Shorter Contributions to General Geology 1966 GEOLOGICAL SURVEY PROFESSIONAL PAPER 554 Tflz's volume was paé/z'saea’ as separate c/zapz‘ers fl—G (A) (B) (0) (D) (E v (F) (G) CONTENTS [Letters designate the separately published chapters] Minor elements in alluvial magnetite from the Inner Piedmont Belt, North and South Carolina, by P. K. Theobald, J r., W. C. Overstreet, and C. E. Thompson. Big Snowy and Amsden Groups and the Mississippian-Pennsylvanian boundary in Montana, by Edwin K. Maugham and Albert E. Roberts. The nature of batholiths, by Warren Hamilton and W. Bradley Myers. Cenozoic volcanic rocks of the Devils Postpile quadrangle, Eastern Sierra Nevada, California, by N. King Huber and C. Dean Rine‘hart. Petrology and structure of Precambrian rocks, Central City quadrangle, Colorado, by P. K. Sims and D. J. Gable. The internal magnetization of seamounts and its computer calculation, by Bernardo F. Grossling. Pennsylvanian and associated rocks in Wyoming, by William W. Mallory. 046 wag... u- I... .u,. “v -—.. - ‘ U “mub‘v - an. AD l“;"l' ~11,ku ._-__ v ya: ""u 1""? \Blé 7 DAY; inor Elements in Alluvial Magnetite from the Inner Piedmont Belt, North and South Carolina VGEOLOGICAL SURVEY PROFESSIONAL PAPER 554—A ’ Prepared on ée/mlf of tfie U.S. Atomic Energy Commission Rmh BudQ-BIULQ» &E‘75’ P6 _ u554 EARTH SCIENCES LIBRARY Minor Elements in Alluvial Magnetite from the Inner Piedmont Belt, North and South Carolina By P. K. THEOBALD, JR., W. C. OVERSTREET, and C. E. THOMPSON SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—A Preparedl on ée/mlf of Me U.S. fltomic Energy Commission UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1967 UNITED STATES DEPARTMENT OF THE INTERIOR STEWART L. UDALL, Secretary GEOLOGICAL SURVEY William T. Pecora, Director EARTH SC'FNCES For sale by‘t'he Superintendent of Documents, US. Government Printing Office Washington, DC. 20402 CONTENTS Pace Relations among the elements—Continued Abstract ——————————————————————————————————————————— A1 Ionic-radius grouping of elements _________________ Introduction ....................................... 1 Prominent highs ________________________________ PIEViOIlS WOIk ---------------------------------- 1 Vanadium and barium ___________________________ Geographic setting —————————————————————————————— 2 Relation of trace elements in magnetite to the distribution Geologic setting ________________________________ 2 of magnetite _____________________________________ Collection Of magnetite —————————————————————————— 4 Geologic interpretation ______________________________ Analytical procedures ——————————————————————————— 5 North-south dichotomy __________________________ Reliability 0f the data ——————————————————————————— 5 Distribution of the elements ______________________ Distribution of the elements __________________________ 7 Manganese, chromium, and titanium __________ Manganese, chromium, and titanium ______________ 9 Copper and zinc ____________________________ Copper and zinc ________________________________ 11 Lead, tin, and beryllium _____________________ Lead, tin, and beryllium _________________________ l4 Vanadium and barium _______________________ Vanadium and barium ........................... l5 Prominent highs ________________________________ Relations among the elements ________________________ 15 Geochemical interpretations __________________________ North-south dichotomy __________________________ 16 Literature cited _____________________________________ PLATE FIGURE TABLE Hwy- 7. 10. 11. 1. . Summary of analyses of magnetite separates from the Inner Piedmont belt of North and South Carolina _______ . Correlation coefficients (Stuart modification of Kendall’s 1) for vanadium and barium in comparison with seven NJ ILLUSTRATIONS . Maps showing sample localities and distribution of manganese, chromium, and titanium in magnetite separates_In pocket . Maps showing the distribution of copper, zinc, lead, tin, and beryllium, and magnetite and sillimanite isograman pocket . Histogram showing the distribution of the tin content of magnetite separates from heavy-mineral concentrates containing less than 1 percent magnetite ___________________________________________________________ . Graph showing the relation of the percentage of magnetite in heavy-mineral concentrates to the proportion of the samples containing more than 10 ppm tin ______________________________________________________ . Graph showing the comparison of erroneously reported antimony with chromium reported for the same samples-_ . Graphs showing the comparison of chemical and spectrographic analyses for lead and copper in magnetite sep- arates _________________________________________________________________________________________ . Histograms showing the distribution of manganese, chromium, and titanium in magnetite separates __________ . Histograms showing the distribution of manganese, chromium, and titanium in magnetite separatesfrom the northern and southern zones of the Inner Piedmont belt _____________________________________________ Histograms showing the distribution of copper and zinc in magnetite separates _____________________________ . Graph showing the cumulative frequency distribution of the zinc content of magnetite separates ______________ . Histograms showing the distribution of lead, tin, and beryllium in magnetite separates ______________________ Histograms showing the distribution of vanadium and barium in magnetite separates _______________________ Correlation coefl‘icients among the minor elements in magnetite separates __________________________________ TABLES Grouping of eight minor elements by valence, coordination, and ionic radii with respect to iron in magnetite- __ other minor elements ____________________________________________________________________________ . Analyses of alluvial magnetite concentrates from the Inner Piedmont belt of North and South Carolina _______ III Page A17 18 18 19 20 20 21 21 22 23 26 26 27 28 Page A5 Page A7 9 19 30 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONTgBELT NORTH AND SOUTH CAROLINA By P. K. THEOBALD, J r., W. C. OVERSTREET, and C. E. THOMPSON ABSTRACT Chemical and spectrographic analyses were made for 15 minor elements in 291 samples of detrital magnetite from the Inner Piedmont belt of North and South Carolina. Lead and copper were determined in all samples; zinc in 283 samples; fitanium, chromium, manganese, and tin in 284 samples; beryllium in 282 samples. These eight elements were detected in a large enough number of magnetite separates to allow the data to be analyzed statistically for geographic and geologic relations. Vanadium and barium were determined in 53 and 85 samples, respectively. Molybdenum and silver, sought in 282 samples, were detected in 8 and 2 samples, respectively. Arsenic, anti: mony, and tungsten, sought in 284 samples, were consistently below the lower limits of detection (1,000, 200, and 100 parts per million, respectively). Manganese, titanium, chromium, zinc, and copper tend to group according to the four-coordinated bivalent, and the six- coordinated trivalent ionic radii, suggesting diadochic substitu- tion in the magnetite lattice. Lead, tin, and beryllium also appear to be incorporated in the magnetite lattice, but their positions are not known. The report area is divisible into southern, middle, and north- ern zones, each characterized by particular associations of minor elements in the magnetite. The middle zone is transitional be- tween the northern and southern zones. In the northern zone, magnetite contains more chromium and less manganese and titanium than does that in the southern zone. Trends of high values for zinc and copper apparent in the northern and south- ern zones are abruptly interrupted at the middle zone, but high values for lead, tin, and beryllium persist throughout this zone. The detrital magnetite is derived from metamorphosed sedi- mentary rocks, which cover about 85 percent of the source area; granitic rocks, which cover about 14 percent of the area; and amphibolite, gabbro, and syenite. Apparent regional variation in the bulk composition of the metamorphic rocks and regional increase in metamorphic grade are the principal factors deter- mining the increase in abundance of manganese, titanium, chromium, and copper in the magnetite. The distribution of specific postkinematic granitic rocks rich in tin and beryllium controls the location of magnetite rich in those metals. Zinc- and lead-rich magnetite also is thought to come from particular granitic rocks, but the relations are not clear. Postkinematic granite and granodiorite in the southern zone contain abundant accessory magnetite that is lean in the minor elements. Where floods of this metal—poor magnetite enter the concentrate, the patterns of distribution of the minor elements in the meta- morphic magnetite are strikingly disrupted. The data demonstrate that the regional distribution of cer- tain minor elements can be determined by analysis of detrital magnetite. INTRODUCTION This report is based on appraisals made during 1951—54 by Overstreet, Theobald, and others of fluvia- tile monazite placers in the western Piedmont of the Carolinas and on studies made in 1956—61 by Theobald and Thompson of trace elements in accessory magnetite from igneous rocks genetically related to ore deposits. During the latter investigation, Theobald decided to investigate the regional distribution of trace elements in magnetite. From the many thousands of heavy- mineral concentrates previously collected in the Care- linas, 289 magnetite separates were prepared by Over- street and analyzed by Thompson. The text is largely the responsibility of Theobald and Overstreet. The purpose of the work is to show the regional vari- ations in the minor elements in magnetite in the western Piedmont of the Carolinas and to suggest how the pro- cedures described here might be used in regional geo- chemical studies. PREVIOUS WORK Trace-element analyses of magnetites separated from igneous rocks or from ore deposits have been described in several recent papers. Wager and Mitchell (1951) and Vincent and others (1954) gave spectrographic and chemical analyses for several minor elements in magnetite concentrates from the Skaergaard intrusive. Theobald and Havens (1960) gave the copper, lead, and zinc contents of iron-bearing minerals and the altera— tion products of these minerals in the Swan Mountain intrusive. Green and Carpenter (1961) reviewed work on the thorium and uranium contents of magnetite from a variety of sources, principally ore deposits. Shri- vastava. and Proctor (1962) gave trace-element analyses of the rock-forming minerals of a quartz monzonite stock in Nevada. All these authors reported'fairly A1 A2 high and fairly variable concentrations of several of the trace constituents. There is an evident partition of several trace constituents that favors magnetite, and this partition causes a more striking variation in trace- metal content of magnetite than that shown by com- parative analyses of the whole rock. These considera- tions lead to the conclusion that regional minor-element variations can be elucidated by study of trace-element distribution in magnetite separates if representative composite separates can be obtained. Concentrates obtained from alluvium seem to offer such a composite sample. A reconnaissance study of the distribution of zinc in alluvial-magnetite separates from the central part of the Front Range, 0010., suc- cessfully outlined the zinc-rich part of the Front Range mineral lbelt (Theobald and Thompson, 1959). Henry Bell 3d (1960), working near Concord, NC, identified alluvium derived from a gabbro—syenite complex on the basis of the zinc content of alluvial-magnetite sep- arates. To our knowledge there are no other published accounts of comprehensive studies of the trace-element composition of alluvial magnetites; however, several exploration and research groups associated with uni- versities or mining companies are investigating the use of magnetite as an indirect medium for geochemical prospecting. These groups reportedly are obtaining favorable results, and their investigations have pro- vided much of the impetus and, indirectly, much of the methodology for our work. In the Front Range, local background appears to be about 200 ppm (parts per million) zinc, and anomalous values are usually greater than 1,000 ppm. Several samples collected near the periphery of that area con- tain less than 100 ppm, which suggests that the regional background for zinc is much less than the local back— ground. To gain broader knowledge of the metal con- tent of magnetite, the authors analyzed a large number of magnetite separates prepared during an appraisal of placer deposits in North Carolina and South Carolina. The general methods of sample collection and prep- aration and the results of the placer appraisal and heavy-mineral studies in the Carolinas have been re- ported in several papers. Overstreet, Theobald, Whit- low, and Stone (1956) described methods for sample collection, sample preparation, mineralogic analysis, and data interpretation. Overstreet, Cuppels, and White (1956) reviewed the monazite distribution in the Southeastern States. Overstreet, Theobald, and Whit- low (1959) gave a more complete summary of the mon- azite resources in western North and South Carolina, the area of this report. Overstreet (1962) reviewed the heavy-mineral reconnaissance and the mineral suites and their distribution. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOGRAPHIC SETTING The report area covers 5,200 square miles in the west- ern Piedmont of North and South Carolina. It extends northeastward for 140 miles from the Savannah River at the South Carolina—Georgia border to the great bend in the Catawba River in west-central North Carolina. The altitude of the area is generally 7 00—1,100 feet. The area slopes gently toward the southeast and is char- acterized by low ridges, narrow valleys, and broad inter— fluves. The local relief is generally 100—200 feet, but it is as much as 1,700 feet at the west margin of the area, in the South Mountains and Blue Ridge. The area is agricultural and industrial and includes some of the most heavily populated counties in North and South Carolina. The population density ranges from 44 to 309 persons per square mile and averages about 115. Because the area is heavily populated, most of it is easily accessible by a network of closely spaced roads. In general, the western part of the area is less populous and less accessible than the central and eastern parts. From 30 to 70 percent of the land in the report area is cleared for farming (cotton, peaches, grain, or pasture), urban development, or industrial use. The uncleared land supports a luxuriant second growth of deciduous trees and conifers. The climate is warm and humid; the annual average temperature ranges from 58.2°F at Morganton, N .C., to 62.1°F at Anderson, 8.0. Annual precipitation ranges from 42.13 inches at Gaffney, 8.0., to 53.18 inches at Greenville, SC. Precipitation tends to increase west- ward; temperature, southward. GEOLOGIC SETTING The report area, a part of the Piedmont physio- graphic province, is underlain by metamorphic and igneous rocks of several ages. Little is known about these crystalline rocks, as surveys have been made of only four 15-minute quadrangles at Gafi'ney, Kings Mountain, Lincolnton, and Shelby, flanking the border between North and South Carolina (Keith and Sterrett, 1931; Sterrett, 1911; Overstreet and others, 1963); of one 30-minute quadrangle at Morganton, NC. (Keith and Sterrett, 1907) ; and of Hart County, Ga., across the Savannah River from Anderson, 80. (Grant, 1958). The part of the area in North Carolina has had better coverage by geologic mapping than the part in South Carolina. Syntheses of the regional geology of both States were prepared by Stuckey and Conrad (1958) and by Overstreet and Bell (1965a, b) . The report area is part of the Inner Piedmont belt as defined by King (1955, p. 352—356) and by Overstreet MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT and Grilfitts (1955, p. 549-577). This belt, one of 10 geologic belts into which the southern Appalachians have been divided (King, 1955, p. 337—338), includes the highest rank metamorphic rocks in the Southeastern States. On the southeast the Inner Piedmont belt is bounded by the Kings Mountain belt, a narrow persistent zone of metasedimentary and metavolcanic rocks that is mainly at the greenschist and albite—epidote-amphib- olite facies. The Kings Mountain belt extends south- westward along, and southeast of, the edge of the report area, from Lincoln‘ton, N.C., to Gafl'ney, SC. South- west of Gafl'ney, the Kings Mountain belt lies a few miles southeast of the study area. The northwest edge of the Kings Mountain belt passes through the town of Laurens, 8.0.; the belt reaches the Savannah River between the town of Abbeville and the Abbeville— Anderson County line. At the northwest edge of the Inner Piedmont belt, there is a narrow zone of blastomylonite and phyllonite called the Brevard belt (Reed and others, 1961). This belt is a fault zone that crosses North and South Caro- lina southwestward from a point about 2 miles west of Marion, N.C., to a point on the South Carolina—Georgia border, about 23 miles upstream along the Tugaloo River from the Anderson—Oconee County line. Thus, the Brevard belt is northwest of the report area. No certain stratigraphic succession or age assign- ments have been made for the rocks in the Inner Pied- mont belt and, indeed, indisputable relic bedding is scarce west of the Kings Mountain belt. Locally, how- ever, some of the sedimentary rocks exposed in the Kings Mountain belt can be traced into the Inner Piedmont. The metamorphosed sedimentary rocks of the Inner Piedmont belt are inferred to consist of three sequences of graywacke, argillite, felsic and mafic tuflaceous argil- lite, tufl‘, and flows with sparse interbedded sandstone and limestone (Overstreet and Bell, 1965a) . Erosional unconformities are postulated to separate the sequences, and to separate the oldest sequence from the basement schist and gneiss of early Precambrian age. This con- tact between the oldest sequence and the ancient base- ment rocks is probably exposed only locally in the Inner Piedmont. Ages of the three sequences of metasedimen- tary rocks overlying the early Precambrian basement are inferred to be late Precambrian and Cambrian, Ordovician to Devonian, and Carboniferous. Of the sequences of rocks above the basement, the older two probably constitute most of the metamorphic rocks in the Inner Piedmont belt. A crude lithologic symmetry in the Inner Piedmont was inferred by Overstreet and Bell (1965a, p. 55—56) A3 to represent repetition of stratigraphic units across a broad and complex anticlinorium. Hornblende gneiss, associated with biotite-muscovite schist and local layers of marble, occupies zones along the southeast and north- west flanks of the belt. Between the flanks the belt is underlain by biotite schist and gneiss, sillimanite schist and gneiss, and migmatite. Rocks in the core of the Inner Piedmont belt are less hornblendic and more highly metamorphosed than rocks on the flanks. The decrease in the hornblendic component from the flanks toward the core probably results from original compositional differences in the sedimentary and pyro- clastic rocks that now form the belt. The rocks in the core probably were deposited on the early Precambrian basement and were originally somewhat more alumi- nous than the rocks on the flanks. The rocks in the core were a stratified succession of shaly to sandy graywackes containing interbedded tufl'aceous argillite and flows of intermediate to mafic composition. Rocks on the flanks of the Inner Piedmont belt were probably deposited on the succession exposed in the core of the belt. This overlying sequence seemingly contained more mafic vol- canic tuffs and flows than did the core sequence. The overlying sequence also contained scattered layers of marl, dolomitic tufi', and dolomitic argillite that are now seen as layers of marble of different degreesof purity. Marble is virtually absent from the core sequence. A northeastward regional decrease in the mafic char- acter of the original sedimentary and pyroclastic rocks in the sequence overlying the core is evident on both flanks of the belt. On the northwest flank of the belt, well beyond the re- port area, massive hornblende gneiss and intercalated marble are exposed in central Oconee and western Pick- ens Counties, S.C. These rocks may be correlative with massive hornblende gneiss and marble on the southeast flank in Laurens and Spartanburg Counties, SC, in the extreme southeastern part of the report area. The hornblende gneiss in Oconee and Pickens Counties gives way northeastward to hornblende-biotite-oligoclase gneiss at the North Carolina—South Carolina line. The hornblende-biotite—oligoclase gneiss is more mafic than the rocks in the core of the Inner Piedmont belt, but less mafic than the hornblende gneiss in western Oconee and Pickens Counties. In Polk, Rutherford, and Mac- Dowell Counties, N.C., outcrops of massive hornblende gneiss are sparse, but the hornblende-biotite-oligoclase gneiss is common eastward as far as central Rutherford County. On the southeast flank of the belt, massive hornblende gneiss decreases northeastward from a point about 8 miles southwest of Gaffney, S.C. 'Staurolite-Ibearing A4 schists, interbedded with subordinate homblendic rocks, are abundant on the southeast flank and extend north- eastward to the Catawba. River. The counterparts of these staurolite schists have not been observed along the northwest side of the core of the belt in North Carolina, possibly because they were cut out by the Brevard fault west of Marion. Thus, there seems to be a regional northeastward de- crease in the original mafic component of the metasedi- mentary and metavolcanic rocks on the flanks of the belt. The attitude of these rocks, inferred to be broadly anticlinorial, tends to expose the more argillaceous se- quence in the core of the belt and the more tufl'aceous sequence on the flanks of the belt. An unconformity probalbly separates the two sequences, but the detailed mapping needed to show the unconformity has not been done in South Carolina, and the maps that have been made in North Carolina do not show the stratigraphic sequence in the Inner Piedmont belt. The metamorphism in rocks of the Inner Piedmont belt increased in grade toward the core of the belt. Rocks along the flanks are typically at the albite—epidote amphibolite facies and at the lower subfacies of the amphibolite facies. In the core of the belt, the rocks are at the sillimanite-almandine subfacies of the amphib- olite facies. Areas where the upper subfacies of the amphibolite ‘facies was attained are outlined on plate 2F by the l-percent isogram for sillimanite. The area of metamorphic climax defined by the dis- tribution of sillimanite is strongly asymmetric and is toward the southeast side of the Inner Piedmont belt. The trend of the zone of rocks at the sillimanite-a1- mandine subfacies is athwart the trends of the strati- graphic units in the Inner Piedmont belt (Overstreet and Bell, 1965a, p. 57). Where the continuity of the 1- percent isogram for sillimanite is interrupted between the North and South Tyger Rivers in Spartanburg County, ‘S.C., the rocks decrease in metamorphic grade to the lower subfacies of the amphibolite facies. From a point just north of Paris Mountain in Greenville County, 8.0., to a point a few miles west of the conflu- ence of the North and South Tyger Rivers, a distance of 20 miles, the metamorphic isograds cut across the trend of foliation and the probable trends of bedding. The metamorphic isograds also cross regional planar struc- tures in northeastern Anderson County, SC, where the 1-percent isogram for sillimanite swings west at the Saluda River. The metamorphic climax in the IImer Piedmont belt is interpreted as having occurred in Ordovician time and as having affected all earlier sequences of sedimen- tary and pyroclastic rocks (Overstreet and Bell, 1965a). SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY A subsequent regional metamorphic event in late Car- boniferous or Permian time is recorded in the Inner Piedmont by the retrogressive effects that it produced. Cycles of intrusive activity are represented in the In- ner Piedmont by plutons of granite and granodiorite and by dikes of pegmatite and gabbro that were em- placed at different times. Felsic plutonic igneous rocks of possible Cambrian, Ordovician to Devonian, and Carboniferous to Permian ages have been recognized. Mafic intrusive rocks are of Cambrian or Ordovician and Permian age. Diabase dikes of Late Triassic( ?) age intrude the other igneous rocks. The plutonic ig- neous rocks of each cycle apparently were emplaced during or after the successive metamorphic episodes. The rocks of the area have been continuously weath- ered and eroded since Cretaceous time. The crystalline rocks are now rather thoroughly weathered and are mantled by saprolite. At most places in the report area, the saprolite is 20—50 feet thick; in a few places it is as much as 180 feet thick; and even in the valleys where erosion has been deepest, at least 85 percent of the rocks have been converted to saprolite to depths of more than 5 feet. The heavy minerals in the streams are derived almost entirely from saprolite, so the magnetite used for the analyses reported herein was weathered for a long time before entering the streams. COLLECTION OF MAGNETITE Magnetite used in this study is from detrital heavy- mineral concentrates collected in 1951~52 by personnel of the US. Geological Survey, supported by the Divi- sion of Raw Materials of the US. Atomic Energy Com- mission, during a study of fluvial monazite placers. D. W. Caldwell, N. P. Cuppels, A. M. White, and J. W. Whitlow prepared many of the concentrates. The heavy-mineral concentrates were prepared by panning samples of gravel or sand dug from riffles in the present channels of small streams. Each sample weighed about 40 pounds. The samples were wet—sieved through a screen made of steel punch plate with 1fig-inch holes, the oversize discarded, and the minus 1A3-inch fraction washed free of clay. Sieved and washed sand was then panned in 16-inch stainless steel pans for 20—30 minutes to provide a concentrate containing as much as 95 percent magnetite and other minerals, of which ilmenite, garnet, monazite, and epidote were the most common. Most of the concentrates contained 1—40 per- cent magnetite. Between 5 and 20 percent quartz was left in the concentrate to avoid loss of heavy minerals sought in the final stage of panning. The procedure was described by Theobald (1957, p. 3-6). The mag- netite used in the present investigation was separated MINOR ELEIWENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT with a hand magnet from the other minerals in the concentrate. ANALYTICAL PROCEDURES The magnetite samples were analyzed by field methods described by Ward, Lakin, Canney, and others (1963). Dithizone was the colorimetric reagent used for lead and zinc determinations, and 2,2’-biquinoline was the re- agent used for copper determinations. The spectro- graphic determinations were made 'by a semiqu‘antitative visual comparison method on the 1.5—meter mobile spectrograph of the US. Geological Survey. Because these methods were developed for the analysis of soils and rocks, they had to be modified to prevent inter- ference from the large amount of iron in the magnetite. In the chemical procedure, the samples were fused with an excess of potassium pyrosulfate for at least 5 minutes. Iron interference was then controlled by adding two or three times the specified volume of the bufl'e'red com- plexing agent. In the spectographic procedure, a series of standards was prepared, using as a base magnetite nearly free of the metals sought. RELIABILITY OF THE DATA Variations among analytical results for a suite of magnetite separates may result from variations in (1) metal content of the magnetite, (2) amount of another mineral included in the magnetite, (3) amount of an artificial contaminant in the separates, or (4) the ana- lytical method. The consistency of the distribution patterns for each of the elements across the report area seems to limit the possibility of artificial contamination of the separates. Artificial contamination is more ap- parent in smaller separates than in larger separates, and there appears to be more metal, particularly tin, in samples containing smaller amounts of magnetite. The tin distribution in the magnetite concentrates is bimodal, indicating that the samples may represent two populations. For those samples where the heavy- mineral concentrate contains less than 1 percent magnetite,1 there is a well-defined break between the two populations at about 10 ppm (parts per million) tin. About 20 percent of these magnetite separates con- tain 10 ppm tin or less; the remainder contain about 30 ppm (fig. 1). To assess the possibility that the higher mode for tin results from artificial contamination, a curve relating the tin content of the magnetite to the magnetite content of the heavy-mineral concentrate is compared in figure 2 with a similar curve calculated 1 Percent is used for convenience. In these concentrates percent and weight of magnetite correlate so closely that both give virtually identical results. 237 4367—67—2 A5 from the distribution in figure 1 by successive additions of tin-free magnetite. The actual decrease in the pro- portion of magnetite separates containing detectable tin as the amount of magnetite in the heavy-mineral con- centrates increases is much greater than the calculated change. Artificial contamination may exist, but there is no evidence for it in the distribution pattern for these samples. '15 NUMBER OF SAMPLES WOOOOO Hv-INM TIN, IN PARTS PER MILLION FIGURE 1.——Tin content of magnetite separates from heavy~minera1 concentrates containing less than 1 percent magnetite. 8°— PERCENTAGE OF MAGNETITE SEPARATES CONTAINING 10 PARTS PER MILLION OR MORE TIN I I 0 0-19 20-39 40-59 60-79 PERCENTAGE OF MAGNETITE IN HEAVY-MINERAL CONCENTRATE 80- 100 FIGURE 2.—-Relation of the percentage of mag- netite in heavy~mineral concentrates to the proportion of the samples containing more than 10 ppm tin. The calculated curve is controlled by the distribution of tin in samples containing less than 1 percent mag- netite, and is projected on a basis of dilu- tion by tin-free magnetite. A6 Natural contamination of the separates by the in- clusion of another mineral in magnetite or the inter- growth of magnetite with other minerals does occur. The unusually metal-rich separates were examined chemically, mechanically, and microscopically in search of a separate phase. The following, for example, were found in the residue from 0.5 g (gram) magnetite from sample 52-WE—109 dissolved in two successive 5—ml (milliliter) portions of hydrochloric acid: 15 mg (milligrams) leucoxene, mostly as boxworks resembling exsolution laminae visible in the original magnetite, 10 mg quartz with minor feldspar, 9 grains staurolite, 3 grains ilmenite, 1 grain zircon, and 1 grain magnetite (protected by a frosted coating of leucoxene) . The separates appear to contain about 99 percent mag- netite, some of which has oxidized to maghemite( ?) or hematite (martite). Virtually all gross errors were eliminated by repeating analyses that originally gave anomalous or unusual analytical results; the repeated analyses were usually made using an independent analytical procedure. One gross error, detected in the analyses of the first group SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY chromium determinations. A minor chromium line in the spectrograms was mistakenly read against the antimony standard and reported as antimony. In fig- ure 3 the erroneous antimony results are compared with the results for chronium in the same samples. The circled anlayses have been checked chemically, and none show antimony in excess of 100 ppm. This comparison indicates the validity of the fourfold subdivision of the chromium results that is used in the following discussion. The product-moment-correlation coeflicient for these paired data (even though one set was read against the antimony standard rather than the chrominum stand— ard) is 0.93 for those samples reported to contain at least 100 ppm antimony. A correlation between chromium and some other element or feature would be extremely poor if it could be masked by analytical variation. Lead and copper were determined both chemically and spectrographically (fig. 4:). The two regression lines represent the best fit of the chemical analyses to the spectrographic analyses and the best fit of the spec- trographic analyses to the chemical analyses. A slight bias is evident for both elements; the chemical analysis is relatively high for lead and relatively low for copper, of samples, provided a measure» of the precision of the compared with the spectrographic analysis. The 15°° I I I I I I I I I I I I 1000— e A) / z / o / z 700— e / -— =' / 2 / f5 / IL m 500— / __ *- / D: g / §_ 300— //<'3/ — >— 5 / E / ’- o z 200— - / < / 7 ° / L|J " / fl: 8 150— . / : Q __ E ' // A/ 100— 3 _ / / / <10 $33): .....3 (23:: /:I:: .I. I I I | I I I 070 ::::° ':::- ::::° -:::- 550 700 1000 1500 2000 3000 5000 7000 10.000' 100 150 200 300 CHROMIUM. IN PARTS PER MILLION FIGURE 3.—Oomparison of erroneously reported antimony with chromium reported for the same samples. Each original spectrographic analysis shown as a dot; those checked chemically are circled. MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT spread of the data in comparison with the analytical variation for both elements justified the subdivision of data used in the following discussion and shown in the illustrations of the geographic distribution of the ele- ments (pls. 1, 2) . The product-moment-correlation coeificients—O.86 for lead and 0.65 for copper—are such that only poor correlations among the elements will be missed because of analytical variation. The methods of determining chromium, lead, and copper span the expected range of precision for the other elements. Several years’ experience with these analytical procedures indicates that the method for spectrographic determination of copper is one of the least precise and that the method for chromium is one of the most precise. Because of their greater sensitivity and somewhat greater precision, chemical determina- tions were used wherever available. Where samples were too small for chemical analysis, spectrographic analyses were substituted. Bias between the two methods is insignificant in relation to the subdivisions of the data we have used. DISTRIBUTION OF THE ELEMENTS Fifteen elements were sought by chemical and spectrographic analyses. The results of analyses of the 291 magnetite separates are tabulated in table 4 at the end of the report. The spectrographic method was used to determine titanium, chromium, manganese, and tin contents of 284 separates and beryllium content of 282 separates. Zinc, because of the poor sensitivity of the spectrographic method, was sought by chemical analysis in 283 of the separates (the remaining 8 con- tain less than 200 ppm). Lead and copper were sought in all samples analyzed by either technique, although the chemical results have been given priority in the discussions. Spectrographic determinations of vana- dium, sought in 53 separates, and barium, sought in 85 separates, are discussed only briefly, as the coverage is considered inadequate. Molybdenum was detected in eight samples, and silver was detected in two samples; data for these elements are insufficient. Arsenic, anti- mony, and tungsten were consistently less than the lower limits of spectrographic detection, 1,000, 200, and 100 ppm, respectively. The eight elements for which adequate data are avail- able—zinc, copper, lead, titanium, chromium, manga- nese, tin, and beryllium—may be conveniently grouped according to their valence states, coordination, and ionic radii with respect to iron in magnetite (table 1). Mag- netite, as a member of the spinel group, is composed of 8 atoms of bivalent four-coordinated iron, 16 atoms of trivalent six-coordinated iron, and 32 atoms of oxygen per unit cell (Palache and others, 1952, p. 688). To A7 TABLE 1.-——Grouping of eight minor elements by palence, coordina- tion, and ionic radii with respect to iron in magnetite [Radii from Green (1959, table 2). Most have been recalculated to tour-coordinated value for bivalent state. Geochemical nature from Rankama and Sahama (1950, tablei4g):]s, slderophile; C, chalcophile; L, lithophile; parentheses, subordinate assoc a on Radius, for indicated valence and coordination Geochemical nature Valence ___________________________ +2 +3 Coordination _____________________ 4 6 Fe (magnetite) __________ 0.70 0.64 S, C, (L) Mn ____________________ . 76 . 66 L Cr _______________________________ . 63 L Ti _______________________________ . 76 L Zn _____________________ . 71 __________ C, (L) Cu ____________________ . 68 __________ C Pb _____________________ 1. 14 __________ C, (S), (L) Sn _____________________ . 88 __________ S, (0) Be _____________________ . 33 __________ allow diadochic substitution, an element must have a size at the valence and coordination state similar to that of the element to be replaced. Rankama and Sahama (1950, p. 122) put an approximate size limitation on diadochic substitution of i 15 percent of the size of the element being replaced. Table 1 is therefore based on four-coordinated bivalent and six-coordinated trivalent ionic radii. Manganese, chromium, and titanium, which can substitute for trivalent iron, make up the first group. The inclusion of titanium in this group re- quires a slight stretching of the 15-percent limit, but the presence of titanium as a trivalent ion in magnetite has been inferred by several authors, and the evidence was recently summarized by Basta (1959, p. 711—717). The size of manganese ions would allow substitution in the bivalent as well as trivalent positions, although the substitution would be more easily accomplished in the latter position. As will become evident, the pattern of distribution of manganese in these samples is more simi- lar to those of chromium or titanium than to those of the other elements. Zinc and copper, which can easily sub- stitute for bivalent iron, constitute the second group. The remaining three elements—lead, tin, and beryl- lium—compose the third group. These elements do not have trivalent forms, and their bivalent four-coordi- nated sizes are not compatible with that of iron. This seemingly prohibits diadochic substitution, but it does not eliminate the possibility that these elements may occur in the framework of a magnetite lattice. Solid solution of other members of the spinel group with magnetite could easily allow for their introduction, either through the solid solution itself or by indirect diodochy of the trace elements with those of the solid solution. Although the extensive solid solution that oc- curs among the members of the magnetite series can be of little help, the limited solid solution that exists be- A8 1000 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY 500 200 100 50 COLORIMETRIC CHEMICAL ANALYSIS, IN PARTS PER MILLION 30 100 1000 SPECTROGRAPHIC ANALYSIS. IN PARTS PER MILLION LEAD FIGURE 4.—Oomparison of chemical and spectrographic analyses (290 duplicate determinations) for lead (above) and copper (on the facing page) in magnetite separates. Each pair of analyses denoted by a dot. Dashed line marks perfect corre- lation; short dashed lines denote spectrographic report intervals on either side of perfect correlation; solid lines are the linear regression curves (curved because of the logarithmic scales). tween magnetite and the spinel or chromite series could be important in terms of trace elements. (For example, limited solid solution of magnetite and the spinel series with minor replacement of aluminum by beryllium could easily account for all the beryllium reported.) Artificial preparations of tin, lead, and beryllium spinels have been reported in both the magnetite and chromite series. (See Mellor, 1931, p. 201; 1934, p. 914.) The possibility that lead and tin are present in the quadrivalent form cannot be disregarded, particularly in View of possible compensating substitutions in the structural modification of the form FeMgFeO4 as de- scribed by Barth and Posnjak (1931). This third group could be further subdivided into two subgroups, since lead and tin have radii considerably larger than that of iron whereas the radius of beryllium is con- siderably smaller. The elements in each of these groups have in common many features of distribution; for this reason, they are considered by groups. Because manganese, chromium, and titanium can best be used to illustrate a major geo- graphic subdivision of the data, they are considered first. The data for all the elements are summarized in table 2, which illustrates the similarity of elements within each group and the breaks between groups for each of the measures of central tendency. The com- parison of mean, median, and mode provides a crude measure of the shape of the distribution for each ele- MINOR ELEMENTS m ALLUVIAL MAGNETITE‘ FROM THE INNER PIEDMONT BELT A9 500 200 100 50 COLORIMETRIC CHEMICAL ANALYSIS, IN PARTS PER MILLION 103 / «0,: 4: -/ 'Ii/ / I I I 50 100 500 SPECTROGRAPHIC ANALYSIS. IN PARTS PER MILLION ment. Each distribution is skewed toward higher values. TABLE 2.—-Summary of analyses of magnetite separates from, the Inner Piedmont belt of North and South Carolina [In parts per million] Element Number oi Mean Median Mode Range analyses 284 I 1, 750 l, 600 l, 600 loo->10, 000 284 1890 300 200 20->10, 000 284 1 6, 000 5, 000 3, 000 1, soc—>10, 000 283 98 75 25- 1, 000 291 40 30 10 <10- 450 291 I 50-64 <25 <25 <25- 1, 000 284 I 32-37 10 <10 <10— 2, 000 282 < 1 < 1 <1 <1- 15 1 Minimum value, values of >10,000 taken as 10,000. ! End points of range represent minimum all values of less than the lower limit of sensitivity taken as zero) and maximum ( values of less than the lower limit 0! sensitivity taken as equal to the lower limit of sensitivity) possible values for the mean. COPPER MANGANESE, CHROMIUM, AND TITANIUM Manganese, chromium, and titanium are the most abundant of the trace elements. Distributions of all three are skewed toward high values and are truncated at the upper limit of spectrographic sensitivity (10,000 ppm in most analyses but'20,000 ppm for some titanium determinations). The distributions are not only skewed on an arithmetic scale but appear to retain a considerable amount of skewing on the logarithmic scale used in spectrographic reporting (fig. 5). There are marked differences in detail among the distributions. The patterns formed by plotting and contouring the data on a planimetric base (pl. 1) exhibit a marked dichotomy. On the basis of this dichotomy, the report area was divided into a northern zone and a southern A10 90-— 80 -1 NUMBER OF SAMPLES o 0000 oooo 0880000 Nmmng ‘HHle-Dho "‘ 500 000 3000 5000 7000 10,000 >10,000 HHN MANGANESE SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY 00 000000 0 o ooooo mogggoooooogo o§ooooo Hmmmnomooooqq LO oooqq Hu-‘Nmmhoo HNmmhoo F‘H Hl—I CHROMIUM /\ TITANIUM /\ FIGURE 5.—Manganese, chromium, and titanium (in parts per million) in magnetite separates. zone. In the northern zone the values for chromium are complex but are generally higher than in the southern zone, where values for manganese and titanium are rela- tively higher. The dividing line between the two appar- ent populations cannot be determined precisely (owing largely to the nature of the sampling and the data), but it is about along the course of the Tyger River. A. third, composite zone, herein called the middle zone (pl. 1A) , has been established to include the breaks in the distribution for these three elements as well as breaks in the distribution patterns of copper and zinc (pl. 2A, 0) and a narrowing in the belts rich in lead, tin, and beryllium (pl. 23, D, E). The southern boundary of the middle zone is defined by the locations of the breaks in data for manganese, chromium, and titanium, which are approximately coincident. The northern boundary of the middle zone is defined by the northernmost pos- sible position for the break in the zinc data. The southernmost possible position for the break in the zinc data and the inferred position of the break for copper, lead, and beryllium are approximately coincident and lie near the middle of the zone. Having thus bracketed the break, we can test the validity of the subdivision. Because the data are truncated, are measured by a logarithmic scale, and have a strongly skewed but unknown distribution, a nonparametric test was used here and for most like data throughout the report. The nonparametric tests are distribution free and usually do not require that the data be measured on an interval scale. These ad- vantages eliminate the necessity for guessing at the form of the distribution and then projecting this form into the unknown realm above and below the sensitivity of the analytical procedures. There is the further ad- vantage of speed gained by avoiding what would un- doubtedly be a cumbersome transformation if the nature of the distribution could be unraveled. With the non- parametric tests there is a slightly greater chance of overlooking an existing correlation, but for the number of samples involved this additional risk is negligible. The chance of finding a false correlation is no greater with a nonparametric test than with the parametric test. A median test (Siegel, 1956, p. 111—116) establishes the dichotomy in the distribution of manganese, chro- mium, and titanium. The hypothesis is that the chro- mium content of the magnetites is greater in the northern zone than in the southern, and that the reverse is true for manganese and titanium. The null hypoth- esis is that the manganese, chromium, and titanium con- tents in the northern and southern zones are approxi- mately the same. The x2 test for two independent vari- ables may be used. The two categories are discrete north and south of the middle zone. Rejecting the null hypothesis at 1 chance of error in 20 (a=0.05) is completely reasonable for these data, but, as will be seen, the chance of error is considerably less. A two- tailed test has been used, although a one—tailed test is permissible under the original hypothesis. A total of 256 samples was tested—93 from the southern zone and MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT 163 from the northern, and the smallest entry in the contingency tables is 18. The critical values of X2 are: a x2 0. 05 3. 84 . 01 6. 64 . 001 10. 83 The results of the median test are: Element x“ D010) Manganese ____________________ ILL 54 <0. 001 Chromium ____________________ 33. 78 <. 001 Titanium _____________________ 7. 13 <. 01 These results indicate that the original hypothesis may be accepted with a high degree of confidence; that is, we will consider the dichotomy- to have been established. Separating the data for the northern and the south- ern zones (fig. 6) provides a clearer picturer of the dis- tribution. The possibility that the individual popula— tions of manganese and chromium have logarithmic distributions must be strongly hedged 'because of the inherent logarithmic scale of spectrographic analysis and the irregular sample distribution for chromium. The arcuate manganese-poor area is the most con- spicuous feature of the northern zone. It generally follows the course of the First Broad River in North Carolina and swings westward in South Carolina to cross Cherokee County and the middle reaches of the Pacolet River in Spartanburg County. The area termi- nates against the middle zone on the south and against the south flank of the South Mountains on the north. The northern and southern parts of the manganese low coincide with titanium lows, and the prominent tita- nium-rich area in Cleveland County, N.C., is broken along the First Broad River. The titanium- and chromium-rich area that includes most of Cleveland County dominates the pattern for these elements in the northern zone. A discontinuous area of low values for both elements lies to the west of the high along the western limit of the report area. An area lean in all three metals occupies a large part of Anderson County in the southern zone. A second area that is low in manganese and titanium trends east- ward from the western limit of the report area south of Greenville. A titanium-rich area surrounds the low in Anderson County, which is flanked by manganese—rich areas on the north, east, and south. COPPER AND ZINC Average zinc values in the magnetite separates are intermediate between those of manganese, chromium, and titanium and those of lead, tin, and beryllium. Average copper values are similar to those of lead. The range for "both elements is about one and a half orders of magnitude. The general distribution for both is A11 strongly skewed toward high values, even when plotted on the geometric reporting scale (fig. 7). Weak bi- modal, or possibly trimodal, distributions are evident for both; the second mode is near 150 ppm for copper and 700 ppm for zinc. The skewing is due at least partly to this bimodal character. Geographic distributions of copper and zinc are shown on plate 2A, 0’. There is a strong pattern of elongate north—trending areas of above- or below-aver- age zinc content. The same pattern is evident for cop- per, though the pattern is less distinctive and seems to break down in the northern zone where the prominent highs trend more eastward This pattern is distinctly broken along a line across central Spartanburg and northern Greenville Counties, 8.0., in the middle zone as defined in the preceding section on “Manganese, Chromium, and Titanium.” The pronounced dichot- omy noted for manganese, chromium, and titanium is not evident for copper and zinc. The medians of cop- per and zinc are the same in the northern zone as they are in the southern zone, and they coincide with the median for the combined data—30 ppm for copper and 75 ppm for zinc. Instead of the dichotomy, there is a break in the data and an apparent offset of about 15 miles in the patterns. In the southern zone both the copper-rich area and the elongate low-zinc area extending southward from Greenville follow approximately along the Saluda River. The high-copper area along the middle reaches of the Enoree River is poor in zinc. At no place in the southern zone does a zinc value in excess of 300 ppm coincide with a copper value in excess of 75 ppm; that is, the second modes described for the overall distribu- tion of these two elements are not geographically co- incident in this zone. Despite this apparent reversal of highs and lows, the two elements are not negatively correlated. Kendall’s correlation coefficient 1' (for dis- cussion of 1' see p. A15) is negative but unusually small (-0.009)—only 6 of 112 correlations of various com- binations of the 8 elements are smaller. We can say with considerable confidence that in the southern zone the two elements do not correlate, and the only simi- larities between them are the type of pattern, their con- centration range, and their expected geochemical behavior. In the northern zone the copper- and zinc-rich areas appear to coincide fairly well. The zinc-rich area ex- tending irregularly northward from the [vicinity of Spartanburg overlaps the copper-rich area north of Spartanburg and a large part of the copper—rich area along the Cleveland—Rutherford County line in North Carolina. The principal area of copper-rich magnetite is along the lower reaches of Buffalo Creek near the A12 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY 5° “ Median Northern zone 163 samples w o I Median NUMBER OF SAMPLES 8888888888 8838888888°8 °° ° 88888° “mmhsaassagg ~~~mmh§2§§§§§§ 28832§§ H H H F. H H MANGANESE A CHROMlUM /\ TITANIUM /\ 40— Southern zone - , 93 sampIes Median Median NUMBER OF SAMPLES O O O O O O O O §§828§§8§888 § H -« N (0 xx 0” 0* o‘ H H l" MANGANESE /\ CHROMlUM /\ TITANIUM /\ FIGURE 6.——Manganese, chromium, and titanium (in parts per million) in magnetite separates from the northern and southern zones of the Inner Piedmont belt. MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT 70 NUMBER OF SAMPLES 0 mo HHN 8888K“ COPPER 80- 'NUMBER OF SAMPLES momemomoo CO Nassaszea§s§§2§ ZINC FIGURE 7.—Copper and zinc (in parts per million) in magnetite separates. North Carolina—South Carolina line. This area is also rich in zinc. Kendall’s correlation coeflicient -r for cop— per and zinc in the northern zone (+0.15) is significant at «=0.01. A pronounced difference occurs in the interrelations between the two elements across the hiatus marked by the middle zone. This difference refutes simple fault displacement, as might be inferred from the uniform 237—367r—67'—v—3 A13 displacement of the north-trending elongate patterns. Some change in process must have been involved. The patterns are generally parallel to zones of regional metamorphism, and the areas of high concentration, particularly for zinc, appear to flank the zones of maximum metamorphism. The copper-rich area along Buffalo Creek at the North Carolina—South Carolina line is the principal feature of the distributions. Included in this area are 8 nearby samples of the 27 that constitute the higher mode for copper; four of the six highest values occur in this group. Two of the samples from the second mode of zinc occur in this same area; this is the only place where samples from the second mode for zinc coincide with those from the second mode for copper. More information is available on the zinc content than on the contents of other minor elements in magnetites. In the two areas from which a considerable amount of data has been accumulated—the Front Range of Colo- rado and the Concord quadrangle, North Carolina— values are similarly distributed. The general back- ground level seems to be below 100 ppm; locally higher levels of concentration generally range from 100 to 200 ppm; and the highest levels, not uncommonly in excess of 1,000 ppm, occur in alluvial magnetite separates adjacent to zinc-rich mining districts or in magnetites separated from igneous rocks inferred to be genetically related to zinc-rich ore deposits. The highest values usually occur in large areas in which the zinc content generally ranges from 100 to 200 ppm. This pattern is nearly identical with that in the Inner Piedmont belt. A cumulative frequency curve (fig. 8) shows that 65 percent of the samples contain 75 ppm zinc or less. The slope of the curve breaks at about 100 ppm to ac- commodate about 30 percent of the samples Whose zinc contents are in the range 100 to 200 ppm. Above 200 ppm the slope of the curve is exceedingly low until the second mode is reached above 500 ppm. The nature of the distribution and the analogy between the Front Range and the Concord quadrangle suggest at least two populations. The concentration of values in the range 100 to 200 ppm could be due to either a third population or a mixture of the two populations. By analogy it can be inferred that magnetite separates containing more than 400 ppm zinc are derived from a rock type that had the potential of genetic association with zinc—rich deposits. This rock must be a minor type occurring in isolated rather small areas scattered along the margins of the high-rank metamorphic rocks. A similar interpretation may apply to the copper distribution. The probability of two populations again seems high. If the copper highs can be assigned to a rock type, the distribution of this rock with respect to A14 100 —— 90— 70— 50— CUMULATIVE PERCENTAGE OF SAMPLES 20- 10 l O O ('0 25 50 *- 75 — 500 — 600 — 700 —' 1000 — | O O V“ 250 — I § 100 — 125 —- 150 '- 175 — ZINC. IN PARTS PER MlLLlON FIGURE 8.—Cumu1ative frequency distribution of zinc content of magnetite separates. the metamorphic zones will be similar to that of the zinc- rich rock, but the zinc—rich rocks and the copper—rich rocks must be separate entities in the southern zone. LEAD, TIN, AND BEEYLLIUM The average amounts of lead, tin, and beryllium found in the magnetite separates are the lowest values obtained for the three groups of elements. The average for beryllium seems to be considerably below the aver- ages for tin and lead. The total range for each is more 177 N (A) ‘h C) 0 NUMBER OF SAMPLES' |\) o 10 0 lnl-OOIOOIOOI-DOOOOO LEAD SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY than an order of magnitude. As less than half the separates contain detectable quantities of these three elements (fig. 9), evaluating the nature of the overall distribution is difficult. Nevertheless, it is clear for at least lead and tin that even on the geometric and logarithmic reporting scales the distributions are strongly skewed toward high values. It seems likely that at least two populations are represented, one in- cluding the samples with detectable metal and the other including samples with average metal content below the limit of sensitivity. Two populations can be demonstrated for a part of the tin distribution when the effect of variation in the quantity of magnetite in the concentrate is removed. (See fig. 1.) The geographic distribution of these minor elements is shown on plate QB, D, E. The patterns shown on these maps are markedly similar. High values are clustered in a relatively narrow belt extending north« eastward from the Saluda River to the North Carolina— South Carolina line. The pattern for beryllium loses coherency in North Carolina, and only a series of isolated spots remains. The patterns for lead and tin generally expand into North Carolina and continue to the northern and eastern limits of the report area. In South Carolina the patterns terminate abruptly against the Saluda River, but there is a slight indication of a trend change southward toward an area of high con— centration along the general course of the Saluda River; this trend change is particularly evident in the pattern for tin. Scattered along the northeast-trending belt are areas of higher concentrations that are generally coincident for all three elements. The largest and most metal—rich of these, not reflected in the beryllium data, is along Buffalo Creek in southeastern Cleveland County, NC. Other metal-rich areas are in South Carolina along the western tributaries to the Pacolet River, along the South Tyger River, and near Greenville. ‘ 256 OOOOOOOOOO NN Inl\ HHNMIONOIDO HHN TIN BERYLLIUM FIGURE 9.—Lead, tin, and beryllium (in parts per million) in magnetite separates. MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT The break in data, so conspicuous for the other ele— ments, is obscure for lead, tin, and beryllium. A general narrowing of the belts rich in all three elements near Spartanburg coincides with the break in the patterns of distribution of zinc and copper. This narrowing in- dicates a slight discontinuity in the data for lead, tin, and beryllium, but the most conspicuous feature of the distribution of these three elements is the continuity of the patterns across the middle zone. VANADIUM AND BARIUM Vanadium and barium were determined in a small number of samples collected from unevenly distributed localities, so interpretation of the data is less reliable than it is for the eight elements already discussed. Data are sufficient, however, to allow certain generalizations to be made concerning the usefulness of these elements for future investigations. The geochemical nature of both elements is lithophile, according to Rankama and Sa- hama (1950, table 4.3, p. 88). Vanadium occurs in both the bivalent and the trivalent states in nature, but ionic radii of both are large enough to inhibit diadochy with iron in the magnetite lattice (0.84 A (angstrom) for the bivalent form in four—coordination and 0.74 A for the trivalent form in six—coordination). The most reason- able substitution would be trivalent vanadium in the bi- valent iron position, as suggested by the data of Barth and Posnjak (1931, p. 255). If this were possible, va- nadium could be grouped with copper and zinc; other— wise it could not be assigned to a group. Barium can be bivalent but not trivalent. The bivalent four-coor- dinated form has an ionic radius of 1.28 A, which places barium with the group of elements (tin and lead) that are too large for diadochic substitution in magnetite. The median and modal concentrations for vanadium are coincident at 1,500 ppm, and the mean concentration is 1,400 ppm. These values are high, suggesting that entry of vanadium into the magnetite lattice is fairly easy. The range of concentrations is 100 to 2,000 ppm, somewhat greater than an order of magnitude and com- parable to the range for the other bivalent elements. The distribution of vanadium may be skewed toward low values on the logarithmic reporting scale (fig. 10). Vanadium is the only element for which negative skew- ing is evident. The median and modal concentrations for barium are coincident at 20 ppm, and the mean concentration is between 32 and 33 ppm. These values are low and are similar to those obtained for lead and tin. The concen- trations generally range from 10 to 300 ppm. There is A15 a, 30— LL! _l n. E U) 20-— LI. 0 5 10— m E D Z O— 8 8 8 8 8 8 H u—c N m to N w 30_ VANADIUM Li] _I D. E < (I) LL 0 0: Lu m E D Z OOOOO [\OWOO HHNM S V BARIUM FIGURE 10.—Vanadium and barium (in parts per million) in magnetite separates. slight evidence, both from the comparison of the median and mean and from the histogram (fig. 10), that the distribution is skewed toward high values. RELATIONS AMONG THE ELEMENTS The relative concentrations of several of the elements vary either directly or inversely with each other. Many of these relations may be seen by comparing the geo— graphic distributions of the several elements. Because few of the correlations among the elements are perfect, it is often difl‘icult to decide Whether they are real features of the distribution or result largely from chance and the desire to find correlations. There are several statistical methods for evaluating the degree of relation between two characteristics of a population; one of these, Kendall’s correlation coefficient (-r), was used to evaluate possible combinations of pairs of elements de- termined in the magnetite separates. Like the median test used in the discussion of the distribution of manga- nese, chromium, and titanium, Kendall’s ‘r is a non— parametric test. It is distribution free, and for these data it requires only that the data may be ordered ac- cording to the concentration of the elements to be com— pared. The Stuart modification (T. G. Lovering, writ- ten commun., 1958) allows easy and rapid handling of large numbers of samples. Kendall’s 'r for combi- nations of the eight principal elements is given in figure 11. Ideally, values of +1 or ——1 indicate perfect cor— relation, whereas 0 indicates complete absence of correlation. A16 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY ALL SAMPLES Cr Ti Mn Cu Zn Pb Sn Be —o.33-—o.29* — —o.35*—o.11—o.39*—o.31*—o.09 22:23:; 026* 027* — 020* 036* — Cr ' 042* 023* 0.16‘ 0.14" 012* — Ti — 019* — — — Mn 0.12‘ 0.39“ 0.30“ — Cu — 0.09 — .Zn 028* — Pb -— Sn SOUTHERN MIDDLE NORTHERN Ti Mn Cu .Zn Pb Sn Be Ti Mn Cu Zn Pb Sn Be Ti Mn Cu Zn Pb Sn Be 031* 0.16 0.17 —- — 019* — Cr 0.33 —' 0.30 — 0.34 0.47' — Cr 0.33" 028* 0.10 023* 039* — Cr 0.47* 020* — —— —— — Ti — 042* — — —~— —— Ti 0.38“ 026* 025* 022* 016* —0.11 Ti ————’-—Mn ———-——-Mn —-0.31*——-—-Mn -— 025* 025* 0.18 Cu — 036* 0.30 — Cu 0.15.‘ 045* 029* — Cu — — -- Zn —- ~—— -- Zn 0.12 018* —— zn 022* —— Pb 042* -— Pb 025* —— Pb 0.14 Sn — Sn —, Sn FIGURE 11.—Correlation coeflicients (Stuart modification of Kendall’s -r) among the minor elements in magnetite separates. All numbers shown are significant at «=0.05; those marked by an asterisk are significant at a=0.01. Correlation not significant at a=0.05 where no entry shown. NORTH-SOUTH DICHOTOMY The prominent break in the data in central Spartan— burg and northern Greenville Counties, S.C., was estab- lished in the discussion of manganese, chromium, and titanium (p. A9—A11). This break is evident in the patterns of the other elements also, though its nature changes from group to group, and profoundly affects the correlations among them. In the southern zone the three possible combinations of pairs of manganese, chromium, and titanium all have significant correlation coefficients. The correlation be— tween manganese and titanium is the strongest measured. That between manganese and chromium is significant at a=0.05 but is the weakest of the three and may result largely from the close association of titanium with both chromium and manganese. In the northern zone the correlations of titanium with both chromium and manganese are highly significant, but the corre- lation between manganese and chromium is exceedingly weak (1-: +0.005). In the middle zone, presumed to contain a mixture Of rocks from both the northern and southern zones, only the correlation between titanium and chromium is significant. The prominent positive correlation of manganese with titanium in the northern and southern zones is not only absent from the middle zone but is replaced by a moderate negative correlation coeflicient (—0.13). The coeflEicient for manganese and chromium is also negative (—0.17). This pronounced reversal in the nature of the association between man- ganese and the other members of the trivalent group also suggests that more than mixing of two populations is involved in the middle zone. The dichotomy is further indicated by a comparison of manganese, chromium, and titanium with the other elements. In the southern zone there are significant correlations of copper with chromium and titanium and of tin with chromium, but these are not exceptional. None of the other correlations are significant. In the northern zone the only correlations that are not sig- nificant «are those of manganese with copper, lead, tin, MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT and beryllium, and of chromium with beryllium. Tita- nium correlates with all the other elements, and chro- mium correlates with all but beryllium. The absence of correlation with beryllium is not surprising in View of the small number of samples in which it was de- tected; it should be noted, however, that beryllium has negative correlation coefficients with all three of these elements in the northern zone and positive correlation coeflicients with all three in the southern zone. In the middle zone the only significant correlation coefficients are those of chromium with copper, lead, and tin, and of titanium with copper. All four significant coeffi- cients in the middle zone are larger than their equiva- lents in either the northern or southern zone. The relations between copper and zinc have been noted, and the relations of these two elements to the other elements follow a similar pattern. Copper corre- lates with lead, tin, chromium, and titanium in both the northern and southern zones, and there is little evidence of any differences in correlation between the zones. In sharp contrast, zinc correlates with all the elements but beryllium in the northern zone but with none of the elements in the southern zone. Tin and lead correlate with each other and with cop- per and chromium regardless of zonal location. These are the most significant correlations involving tin and lead, and they emphasize the general continuity of the distribution patterns for the two elements across the dichotomy. They suggest an association based on the ionic radii and the charge of these four elements across the group boundaries. Somewhat weaker correlations of tin and lead with zinc and titanium in the northern zone are absent from the southern zone, indicating the dichotomy in data for zinc and titanium. The small range of values and the limited number of concentrates containing detectable beryllium preclude strong correlations of this element with the others. In general, the value of -r obtained for beryllium correla- tions with copper, lead, and tin is positive and within the range 0.1—0.2 in the southern and middle zones. In the northern zone, the correlation coefficients are 0.05 or less, and correlations with the other elements are all negative. The limited data for beryllium seem to reflect the dichotomy. The differences in associations between elements across the dichotomy are pronounced. They are large enough to suggest that different processes or conditions controlled the accumulation of minor elements in the magnetites on the two sides of the break. Further, there is some evidence that a third set of conditions controlled the distribution of manganese, chromium, titanium, and copper in the middle zone. The relations among lead, A17 tin, and copper are generally continuous across, and un- affected by, the dichotomy. IONIC-RADIUS GROUPING 0F ELEMENTS The grouping of elements previously established (p. A7) was based on the hypothetical assumption of dia- dochic substitution for iron in the magnetite lattice. This grouping has been convenient in discussing the elements and may be, at least in part, a real feature of the distribution. It seems worthwhile to consider next the features of the measured distributions that tend to support or deny the hypothetical grouping. The manganese-chromium-titanium grouping is based on the possible diadochy with trivalent iron. The three elements have in common many features of dis- tribution, particularly the sharp north-south dichot- omy. Close geochemical aifinity is shown by the highly significant correlations of titanium with both chromium and manganese in both the northern and the southern zones. Correlation of manganese with the bivalent groups, the other possible substitution for this element, is nonexistent in the southern zone and less common in the northern zone than is correlation of the bivalent elements with chromium or titanium. The highly sig- nificant correlation of manganese with titanium and the weaker or nonexistent correlation of manganese with chromium suggest some undetermined feature common to manganese and titanium but not shared by chromium. This suggestion is supported by the reversal of medians across the dichotomy: chromium has a significantly higher median in the northern zone than in the south- ern zone, whereas manganese and titanium have sig- nificantly higher medians in the southern zone than in the northern zone. The most obvious feature common to manganese and titanium but not shared by chromium is a stable quadrivalent form. Copper and zinc can substitute diadochically for bivalent iron. The patterns for the two elements have the same general trend, and the elements have similar median values. As has been noted, they are, however, not consistently related. The correlation coefficient is highly significant in the northern zone but insignificant in the southern zone. Despite their geochemical simi- larity, copper and zinc behaved differently and inde- pendently in different parts of the area. Lead and tin ions are too large for direct substitution in the magnetite lattice. Both have remarkably similar patterns of distribution, both have similar median values, and neither is markedly affected by the dichot- omy. The correlation coefficient of tin with lead is highly significant in each of the zones. A strong asso- ciation seems evident. However, both lead and tin cor- A18 relate even more strongly with copper than with each other. Beryllium is too small for direct substitution in the magnetite lattice. Its pattern and low median value suggest a weak association with lead and tin, but this association is largely confined to the middle and south— ern zones. There is a significant correlation between beryllium and tin in the southern zone, but in the north- ern zone beryllium does not seem to be directly related to the other elements. The only significant correlation is negative and relates beryllium with the trivalent group through titanium. Beryllium’s only other coeffi— cients that approach significance are also in the triva— lent group and are negative: —0.09 with manganese and —0.07 with chromium. Correlations of elements across the group boundaries are common and, in general, seem to follow a distinctive pattern. The strongest and most consistent of these correlations are those of copper with lead and tin. Less strong but consistent are those of copper with chromium and titanium, and of tin with chromium. With the exception of the copper—lead-tin association, these cor- relations between groups are subordinate to those within a group. There is little doubt that the grouping of elements based on ionic radii in coordination states acceptable to the spinel lattice is a real grouping. It provides a practical basis for evaluating the data and indicates associations that can be expected in further studies of the minor—element content of magnetite. The presence of this grouping does not, however, validate the theo- retical base on which the grouping was established. The geochemical similarities of the elements in each group are such that the association could be expected from several combinations of events other than diadochic substitution in, or solid solution with, magnetite. The strong association of copper, lead, tin, and chromium that crosses group boundaries should be further investi- gated, as these elements are commonly associated in base-metal deposits. The possibility that these elements are present in a discrete mineral phase included in, intergrown with, or exsolved from the magnetite cannot be eliminated; nevertheless, polished-section studies of several concen- trates did not reveal such a phase. Tin- and lead-rich concentrates were dissolved in hydrochloric acid in an attempt to chemically isolate a metal-rich phase, but both metals were found in the acid-soluble fraction rather than in the residue. The residue from the acid leach was mostly an open latticework of clean white titanium oxide mimicking exsolved ilmenite that could be seen in polished sections of the magnetite. This residue can account for a large part of the titanium but SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY not for tin and lead. The tin and lead are not present as any of the common oxides or silicates of these ele- ments; these compounds are insoluble in hydrochloric acid. Some of the metals, however, could be present as finely divided sulfides that could be overlooked in pol- ished section and are commonly soluble in acids. The variation in metal content of the magnetite con— centrates undoubtedly results from variation in con— centration of metal in the environment of crystalliza- tion. The observed grouping, regardless of whether direct substitution in a spinel lattice is involved, could result from the association of the metals in the environment. PROMINENT HIGHS The metal content of the magnetite separates varies with the concentration of metal in the environment of original crystallization, regardless of the nature of in- corporation of the metal in the magnetite. This varia- tion suggests that areas rich in metals may be sought indirectly by investigating areas containing metal-rich magnetites. Areas rich in one element have been dis- cussed, but several of these warrant emphasis because they contain high concentrations of several minor elements. A major copper and tin high along the lower reaches of Buffalo Creek is in the southeast corner of Cleveland County, N.C., and adjacent parts of Cherokee County, SC. A major lead high lies a short distance to the northwest, and a zinc high lies along the State line. The eastern limit of these highs is not defined by these data. The highest values obtained for copper, tin, lead, and zinc are in this general area. A second copper and tin high occurs along the western tributaries of the Pacolet River north of Spartanburg, 8.0. There is a small lead high in the same area, and the general levels of zinc and beryllium are above average. A lead and beryllium high extends north from the Enoree River through Greenville, S.C. Highs for tin, zinc, and copper have the same trend and overlap those for lead and beryllium. A weak high extends generally eastward from the Second Broad River into the western tributaries of the First Broad River northeast of Rutherfordton, N.C.; it is best defined by zinc but is shown also by tin and copper. VANADIUM AND BARIUM The abundance of vanadium and barium, determined in 53 and 85 of the magnetite separates, respectively, is not adequately known for direct evaluation of their geographic distribution. Comparison of the limited data for these elements with those for the other elements MINOR ELEMENTS IN ALLUVIAL MAGNETI’I‘E FROM THE INNER PIEDMONT BELT allows some insight into expected distributions. Corre- lation coefficients of vanadium and barium with seven of the other elements are given in table 3. TABLE 3.—Correlation coeflicients (Stuart modification of Kendall’s -r) for vanadium and barium in comparison with seven other minor elements in magnetite separates from the Inner Piedmont belt, North and South Carolina [All numbers shown are significant at a = 0.05; those marked by asterisk are significant at a=0.01. Correlation not significant at a=0.05 Where no entry shown] Cr ’ Ti Mn Cu Zn Pb Sn V ______________________________________ *——0. 39 —O.34 *—0. 35 .......... Ba _________ *+0. 32 ________ —0. 14 "+. 46 —. 18 *+. 30 *+0. 27 Vanadium has marked negative correlations with all the other bivalent elements. Though not significant, even the value of 7 for vanadium and tin is large and negative (—0.22) in contrast to that for the trivalent elements, all of which have an absolute value of 1- less than 0.1. The consistent values of the correlations of vanadium with copper, lead, and zinc in contrast to the weak or insignificant correlations of zinc with copper and lead suggest that vanadium can be used as a general indicator element. A separate containing an unusually low vanadium concentration should at least be checked for unusually high copper, lead, or zinc concentrations. This relationship further suggests that vanadium can be expected to be scarce in the central part of the area, where copper, lead, and zinc are abundant, and abun— dant along the flanks of the area, where copper, lead, and zinc are scarce. Barium has an exceedingly strong positive correla- tion with copper and generally parallels copper in asso- ciation with the other elements. Surprisingly, barium correlates less strongly with lead and tin than does copper. (Compare fig. 10 with table 3.) The positive correlations of barium with copper, chromium, lead, and tin are all strong, however, and together with the significant negative correlations with zinc and man— ganese indicate fairly well the expectable distribution of barium: there should be general north-trending highs in the middle part of the area flanked by general lows. RELATION OF TRACE ELEMENTS IN MAGNETITE TO THE DISTRIBUTION OF MAGNETITE The distribution of magnetite, expressed as percentage of the quartz—free heavy-mineral concentrate, is shown on plate 2F. In general, a large area lean in magnetite occupies the central part of the report area. Nearly all samples from this central area contain less than 10 per- cent magnetite, and most contain less than 5 percent. Nearly all samples collected just west of the central, lean area contain more than 10 percent magnetite, and most contain more than 20 percent. To the east in South A19 Carolina, a similar area of magnetite—rich samples occurs, but the transition is less abrupt than the one to the west. In North Carolina the east boundary of the magnetite-poor area is at best poorly defined. A sharp constriction in the magnetite-lean area generally fol— lows the course of the South Tyger River in western Spartanburg and Greenville Counties; it lies in the middle zone as previously defined (p. A10). The mag— netite-lean area is displaced along this constriction; the area to the south is offset to the west, much as are the patterns for copper and zinc content of the magnetite. Where the magnetite-lean area is much narrowed in the middle zone, the patterns for lead and tin content of the magnetite are also narrowed. The pattern of distribution of magnetite rather strongly resembles the patterns of distribution for the minor constituents of the magnetite. Most of the fea- tures or contrasts in the minor-element patterns that we have considered significant are in the zone lean in mag- netite. For most of the elements the higher values are in the central part of the area and the lower values are on the flanks. Some trace elements (notably copper, lead, and tin) are most abundant along the east edge of the report area and near the North Carolina—South Carolina boundary, where control for the east limit of the magnetite-poor area is lacking. To illustrate these general relations, the 10-percent isogram on plate 2F has been superimposed on each map of trace—element distribution (pls. 1, 2). The relation of trace elements to distribution of magnetite is further defined by the statistical com- parison in the top chart in figure 11. The correlation of seven of the eight elements with the percentage of magnetite in the concentrate is negative and significant at «a=0.05. For five of the elements the correlation is highly significant at a=0.01 (ra20,01=0.10). Man— ganese is the only element not having the general inverse relation of metal content of the magnetite to the per- centage of magnetite in the original heavy—mineral con- centrates; for manganese this relation is nearly random (To: +0.003) . ' This strong consistent relation of the metal content to the amount of magnetite could be expected to lead to consistent positive correlations among the various element pairs. Indeed, virtually all the significant cor- relations among these pairs (fig. 11) are positive. However, two lines of evidence suggest that independ- ent control, rather than cause and efl'ect, produces this relation of the positive correlations among element pairs to the negative correlations of the elements with the proportion of magnetite in the original concentrate: First, the strongest and most consistent correlation among the element pairs is that between manganese A20 (which has no relation to the magnetite distribution) and titanium (which has a strong relation to the mag- netite distribution). Both titanium and chromium are strongly correlated with the magnetite distribution, but the correlation of the two as an element pair is consist- ently poorer than that of manganese and titanium. The second line of evidence is provided by the correla- tion of copper and lead, the elements most closely related to magnetite distribution. Plate 2 shows that areas Where the original concentrates contain more than 10 percent magnetite generally coincide with areas where the contents of copper and lead are low. In the magnet- ite—poor zone, the average metal content of the magnet- ite is higher, but there is considerable consistent varia- tion in the metal content of the magnetite; that is, the patterns of high copper and lead‘ content are within the magnetite—poor areas, and these patterns are truncated by the magnetite-rich areas. As a test of the effect of the magnetite distribution on the association of these two elements, the variation in magnetite content of the original concentrates may be removed by considering only those magnetite concentrates from original heavy- mineral concentrates containing 1 percent mag- netite or less. (Using concentrates with more magnet— ite is impractical, because these magnetites generally contain less than detectable amounts of lead, and the critical patterns for copper and lead are chiefly confined to the magnetite-poor area.) From the northern zone there are 7 4: concentrates containing 1 percent magnetite or less. The lead and copper contents of these magnet— ites are <25—1,000 ppm and 10—250 ppm, respec- tively. The Stuart modification of Kendall’s 1- for the association of the two elements gives 1-0: +0.36. This result is highly significant (ra=0_01=0.20), even though only 43 percent of the samples from the northern group can be used. The value at constant magnetite content, To: +0.36, is, however, considerably less than that for the northern group as a whole, 70= +0.45. The associa- tions of the pairs of elements are apparently independ- ent of, but enhanced by, the association of the elements with the distribution of magnetite. The most obvious cause for the observed variation of the elements with the amount of magnetite would be a contaminant in the magnetite concentrates that pro- vides most or all of the trace constituents and dilution of the contaminant by addition of barren magnetite. This possibility is discussed, and for practical purposes eliminated, in the section “Reliability of the Data.” The data presented tend to substantiate the previous conclusion that contamination, either natural or arti- ficial, is negligible. The major features attributable to the trace-element distributions are either evident in the distribution of magnetite (the north-south dichot- SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY omy) or are extant in a given part of the magnetite distribution and absent from the rest of that distribu- tion (the principal metal highs and the trends they define are in magnetite-poor areas). The more reasonable cause for the observed associa- tions seems to be the basic geochemistry of magnetite and its interrelations with the regional geology. GEOLOGIC INTERPRETATION The source rocks for the magnetite are mostly para- gneiss and paraschist and minor bodies of orthogneiss and nongneissic granitic, gabbroic, and syenitic rocks, in the following approximate abundances: Percent Biotitic paraschist and paragneiss _____________________ 65 Sillimanitic paraschist and paragneiss _________________ 20 Amphibolite _________________________________________ 1 Granitic orthogneiss __________________________________ 10 Nongneissic granite ___________________________________ 4 Gabbro and syenite ____________________________________ Trace The magnetite in at least 85 percent of the rocks crystal- lized when the schists and gneisses formed. The minor- element composition of this magnetite presumably re- flects the bulk composition of the rocks under conditions of the regional metamorphic maximum, possibly some- what modified by a later metamorphic episode. Large local variations in the composition of the magnetite must have been caused by large additions of anomalous magnetite in order to overcome the spatially dominant regional trends. The principal geologic factors con- trolling the distribution of minor elements in the mag- netite were either the bulk composition of the meta— morphic rocks or the grade of regional metamorphism, locally modified for some elements, notably zinc, beryl- lium, and tin, by additions from the igneous rocks. Interpretation is generally limited by the inadequacy of present knowledge of the regional geology. NORTH-SOUTH DICHOTOMY Division of the report area into northern, middle, and southern zones appears to reflect the intensity of regional metamorphism. The broadest expanse of rocks at the sillimanite-almandine subfacies, approxi- mately defined by the 1-percent isogram for sillimanite, is in the northern zone. The belt of high-grade meta- morphic rocks narrows greatly and swings westward for about 20 miles in the middle zone. The intensity of metamorphism apparently decreased southward in the southern zone where the amount of sillimanite de- creases and the belt of sillimanite—bearing rocks nar- rows. Garnet, rutile, and tourmaline, which accompany magnetite in the concentrates derived from the highest grade metamorphic rocks, are most abundant in the MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT northern zone, much less abundant in the middle zone, and least abundant in the southern zone. The middle zone marks a change in direction and in intensity of geotherms. There is no structural evidence that these changes are accompanied by faults approach- ing the apparent displacement in the middle zone, as neither the amphibolites nor the Kings Mountain belt on the east side of the high-grade metamorphic rocks is ofl’set. Lithologic changes superimposed on the metamorphic zonation probably enhance the north-south dichotomy, as the middle zone lies immediately south of pronounced changes in the abundance of mafic components in the metamorphosed sedimentary rocks flanking the belt. The composition of the plutonic igneous rocks changes from granite and quartz monzonite in the central and northern parts of the area to granodiorite in the south- ern part. However, the compositional change of the igneous rocks is slight, gradual, and lies to the south of the middle zone. Zircon, a common accessory min- eral in the plutonic rocks, is the only heavy mineral accompanying magnetite in the concentrates that per- sists continuously across the middle zone. DISTRIBUTION OF THE ELEMENTS MANGANESE, CEROMIUM, AND TITANIUM The distribution patterns for manganese, chromium, and titanium in magnetite seem to result from the orig- inal composition of the sedimentary and pyroclastic rocks and from the energy input during metamorphism. The patterns reflect in a minor way the intrusion of granitic rocks, particularly postkinematic rocks. Analyses of slate, phyllite, schist, gneiss, granite, diorite, amphibolite, and gabbro in the Kings Moun- tain and Inner Piedmont belts, available for South Carolina (Sloan, 1908, p. 250—264) and Georgia (Crick- may, 1952), show a narrow range in abundance of man- ganese and titanium in bulk samples of the rocks (no data are available for chromium). Bulk samples of granite, syenite', and pegmatite from Cabarrus County, N.C., 60 miles east of the Inner Pied- mont belt, contain less manganese, titanium, and chro- mium than do samples of mafic rocks from the same area (Henry Bell 3d, oral commun., 1962). Number of Average, in percent Rock analyses in average MnO T10. Slate and phylllte ...................... 7 0.09 ............... 0. 76 (0.32 for 2)1 .................. Biotite schist and gneiss... ..... 8 .32 . 68 Felsic volcanic rocks.__. _____ 5 .26 ....... 54 Granite ..................... 5 Trace. . .. (0.11 {or 1)!.. Diorite and gabbro ..................... 6 .25 Amphihnlifn 8 .17 ________________ 1 Average of entries given as more than trace. A21 Analyses of rocks in the vicinity of Shelby in the northern zone of the Inner Piedmont belt indicate the distribution of manganese, titanium, and chromium (Overstreet and others, 1963) : [Chemical analyses by Lois Trumbull, Faye Neuerberg, L. C. Peck, and W. J. Blake, Jr.; spectrographic analyses by J. D. Fletcher and P. R. Barnett] Average. in percent Rock Chemical Spectro- graphic MnO T10. Cr Biotite schist ....................................... 0. 07 0. 40 0. 000xt Sillimanite schist ................................... . 08 l. 07 . 01 Horublende-biotite—oligoclase gneiss_. . . l2 . 69 . 002 Diopside-biotite eiss 1 ............................ . 09 . 71 . 01 Toluca Quartz onzonite l ......................... . 03 . 22 . 0002 Pegmatite. . 01 . 15 . 0001 I Average oi two determinations. Norm—The concentrations of the elements determined by semiquantitative spec- trographic analysis are bracketed into oups each oi approximately one-third of an order of magnitude, 31+ indicating the igher portion (10-15 percent); x, the middle i’““3‘?afififigamete:31ifii§$§e°3i§€§ié3§lb§eéfitita833523333,‘é‘n‘t‘é? c emical or spectrographic, show that the assigned group includes the quantitative value in about 60 percent of the analyses. Partition of the minor elements in the rocks is less pronounced than in the accessory magnetite, but their distribution in the rocks somewhat resembles that in the magnetite. Manganese may be more abundant in the schist and gneiss in the southern zone than in the northern zone, and titanium may be somewhat more abundant in biotite schist and gneiss than in other rock types. There may be less manganese, chromium, and titanium in the granites and more titanium in the silli- manite schist than in the other rocks. The sillimanite schist in the northern zone is considerably richer in chromium than are the biotite schist and hornblende- biotite-oligoclase gneiss. The southward-broadening series of manganese-high areas along the core of the belt are interpreted as being remnants of a formerly more extensive pattern derived from sedimentary and pyroclastic rocks that are richer in manganese in the southern zone than in the northern zone. Perhaps the sedimentary rocks immediately overlying the ancient Precambrian basement are most common in the south- ern zone, whereas the overlying sequence is most abun- dant in the northern zone. The older sequence, derived from weathered rocks, contains a more abundant pe- litic fraction and is richer in aluminum and titanium than the overlying sequence (Overstreet and Bell, 1965a; Overstreet and others, 1963); it is therefore reasonable to expect more manganese in the older sequence. The efiects of the metamorphic climax in Ordovician time are evident in the original stratigraphic sequence. Pelites at the staurolite-kyanite subfacies appear to have been the preferred hosts of manganese-rich mag- A22 netite, and at higher metamorphic grade the manganese content of the magnetite tended to decrease. Thus, a notable manganese low was formed in the southern zone in the area bounded by the l-percent isogram for silli- manite, at the sillimanite-almandine subfacies, whereas the largest area of manganese-rich magnetite, in Green- Ville, Laurens, and Abbeville Counties, is at the kyanite- staurolite subfacies. Staurolite is abundant in the northern zone near the manganese high along the Cleve- land—Lincoln County line. The absence of a persistent manganese high between Gaifney and the big bend in the Catawba River, coextensive with the area in which staurolite is most abundant, apparently resulted from the composition of the sediments overlying the lower sequence. The most intense metamorphism, and the least manganese in the magnetite, is in the northern zone, particularly in the arcuate area of manganese lows extending from the Pacolet River through Gafiney and Shelby to the South Mountains. The manganese low in this area is probably superimposed on an original general manganese high. Perhaps, as metamorphism increased above the kyanite-staurolite subfacies, man- ganese changed from the trivalent to the quadrivalent state, thereby facilitating volume reduction and making manganese less acceptable in the magnetite lattice. Such behavior is evident for titanium, which, as the quadrivalent form rutile, is rather abundant in areas outlined by the l-percent isogram for sillimanite. Magnetite is abundant in postkinematic intrusive rocks and is probably low in manganese, as observed by Henry Bell 3d (oral commun., 1962), for Cabarrus County, N.’C. The large manganese low centered at Anderson probably resulted from a flood of manganese- poor magnetite from many small bodies of postkine- matic quartz monzonite and granodiorite. The interpretation of the distribution of titanium in magnetite is similar to that for the distribution of man- ganese. In the southern zone biotite schist and gneiss are apparently somewhat richer in titanium than in the northern zone, probably because of a regional lithologic change. High contents of titanium in detrital magnet— ite occur in a broader area in the southern zone than in the northern zone, a distribution considered to reflect lithologic units. The two prominent titanium highs in the northern zone, the high on the Enoree River, and the three peripheral highs around Anderson are prob- ably residual from a former broad zone of highs. In the northern zone the highs are separated by an arcuate low along the First Broad River. This low closely re- sembles the manganese low and is attributable to the same factor: intense regional metamorphism that re- duced the amount of titanium in the magnetite from maximum amounts incorporated at lower metamorphic SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY facies. The low at Greenville is attributable to this factor. Low-titanium magnetite from magnetite-rich postkinematic granitic rocks flooded the concentrates in Anderson County and produced the low in that area. The same geologic processes that governed the distri- bution of manganese and titanium presumably governed the distribution of chromium. In rocks of the northern zone, chromium is most abundant in sillimanite schist and is scarce in biotite schist, quartz monzonite, and pegmatite. Spectrographic analyses of 124 heavy- mineral concentrates in the belt show that in most the chromium content is within the range 100 to 1,000 ppm. The chromium content of the heavy minerals shows no apparent systematic variation. The difference in abundances of chromium in mag- netite in the northern and southern zones is too great to be explained by differences in metamorphic intensity. Chromium presumably was more abundant in the origi- nal rocks of the northern zone. The most significant high for chromium in the magnetite is along the First Broad River north of Shelby, in the area of maximum regional metamorphism; magnetite presumably was still accumulating chromium at the maximum temperatures and pressures reached. The chromium low in Ander- son County can be attributed to a flood of chromium- poor magnetite from postkinematic granitic rocks. COPPER AND ZINC The abundance of copper in rocks in the Shelby area and in the adjacent Lincolnton and Kings Mountain area is known from quantitative and semiquantitative spectrographic analyses of bulk samples of the rocks (Overstreet and others, 1963; W. R. Grifl'itts, written commun., 1958) : [Analyses by J. D. Fletcher, P. R. Barnett, A. A. Chodos, R. G. Havens, A. T. Myers, and P. J. Danton. For explanation of symbols see table, p. A21] Copper content (percent) Number of samples Quantita— Semiquan- tive titative Biotite schist _________________________________ 1 ____________ 0. 000x- Biotite schist and gneiss ________________ 10 ____________ . 00x Blotlte—oligoclase gneiss. _______________________ . 00x- Hornblende-biotite oligoclase gneiss 0. 0006 ____________ Sillimanite schist ___________________ 1, 4 .006 . 00x+ Dlopside-biotite gneiss ____________ ._ 2 . 002 ____________ Granite gneiss ________________________________ 1 ____________ . 00x- Toluca Quartz Monzonite (Ordovician) _______ 2; 7 . 0006 . 00x- Pegmatite related to Toluca Quartz Mon- zonite ______________________________________ 1 . 0003 ____________ Yorkville Quartz Monzonlte (Permian) ______ 4 ____________ .OOX Cherryville Quartz Monzonlte (Mississip- pian(?) to Permian(?)) _____________________ 8 ............ . OOX‘ Pegmatlte related to Cherryville Quartz Monzonite __________________________________ 2 . 00x The zinc content was below the limit of detection in all bulk samples of rock analyzed. In 124 spectro- graphic analyses of heavy-mineral concentrates, copper is uniformly 0.00x percent and zinc is less than 0.01 MINOR ELEIVIENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT percent. The average values for copper in the rocks are of the same order of magnitude as those for copper in the magnetite. The average values for zinc in the magnetite are considerably greater than those for zinc in the rocks. For copper in the magnetite, lows of 40 ppm or less coincide with low-rank metamorphic rocks along the flanks of the belt, and highs of 50—80 ppm copper or more in broad centrally located areas coincide with higher rank metamorphic rocks. The copper high along the Saluda River and most of the copper highs in the northern zone are within the sillimanite isograd. Seemingly, the magnetite was still taking on copper at the peak metamorphism. The large copper low in Anderson County and the small lows around Lincolnton and Newton appear to reflect a superimposed flood of copper-poor magnetite from postkinematic granite and granodiorite. There is little evidence to relate the copper distribu- tion to copper—bearing deposits. The Graham gold mine, reported by Pardee and Park (1948, p. 76) to contain copper, is in an unsampled basin between the drainage basins numbered 262 and 264. The Cameron lead mine, reported by Keith and Sterrett (1931, p. 12) to have been worked for copper, is 2% miles southeast of Gafl'ney, 8 miles along strike from the principal cop- per high. Copper was not reported by Sloan (1908, p. 81—83) in his description of mineral deposits in the southern zone. There is no unusual amount of copper in magnetite from the vicinity of the several gossanlike deposits of limonite in South Carolina (Keith and Sterrett, 1931, map; Sloan, 1908, p. 102, 106). The general pattern of distribution of zinc in magnetite is defined by zones of 100—200 ppm content that parallel the zones of regional metamorphism, but the relatively higher zinc content is marginal to the metamorphic maximum. Thus, the areas of both maxi- mum and minimum metamorphism are relatively lean in zinc. Interruptions in the general pattern as well as in most zinc maximums apparently result from adulteration of the metamorphic magnetite population by zinc-poor magnetite from granite or by zinc—rich magnetite from syenite. Both granite and syenite are rich in zircon and contain the indicated zinc-rich or zinc—poor magnetite in the Concord area, North Carolina (Henry Bell 3d, oral commun., 1962). Distribution of zircon in the heavy-mineral concentrates favors this explanation for most of the zinc lows sh0wn on plate 20. Magnetite from the syenite in Cabarrus County, N.C., is both rich in zinc and lean in copper (Henry Bell 3d, oral commun., A23 1962), and the syenite in the Inner Piedmont belt in South Carolina probably correlates with syenite in Cabarrus County (Overstreet and Bell, 1965a). The zinc highs above 250 ppm in the southern zone do not correlate with copper highs. Several of these highs are in areas where syenite or syenite pegmatite have been reported (Overstreet and Bell, 1965b; Leiber, 1860, p. 29). Similar relations probably exist for the zinc-rich and copper-poor areas in the northern zone, though only the zinc high southeast of Newton can be correlated with a zircon-rich pegmatite of probable syenitic character located across the Catawba River in Iredell County, NC. (Pratt and Lewis, 1905, pl. 2). The east-northeast—elongated area of zinc-rich magnetite between the First and Second Broad Rivers northeast of Rutherfordton is parallel to, and possibly, a continuation of, a strongly developed system of faults that crosses the northwest quarter of the Shelby quad- rangle from the east end of the high (Overstreet and others, 1963). Neither the faults nor the sillimanite schists in which they occur are known to be mineralized in the quadrangle, and syenite was not observed there. The northeast-trending area of zinc—rich magnetite at the State line northeast of Gaffney is associated with copper in abundances greater than 180 ppm. In this respect, the zinc high is unique. The streams from which the magnetite was taken head in staurolite schist and Cherryville Quartz Monzonite along the west side of the Ordovician to Mississippian Blacksburg and Battle— ground Schists (Keith and Sterrett, 1931, maps), where northeast-trending high—angle faults separate the staurolite schist and quartz monzonite on the west from the sericite schist of the Kings Mountain belt on the east. Neither zinc nor copper has previously been re- ported from near this zinc- and copper-rich area, but the Cameron lead mine, reported to have been originally opened for copper (Keith and Sterrett, 1931, p. 12) is 8 miles to the southwest and nearly on strike with the high. The site of the Cameron mine, 2% miles south- east of Gafi'ney, is not a zinc high, and zinc sulfides have not been observed at the mine, which is in a quartz vein containing siderite, galena, and chalcopyrite. LEAD‘, TIN, AND BERYLLIUM Contents of lead, tin, and beryllium are known for rocks in the Shelby area and adjacent parts of the Kings Mountain, Grafl'ney, and Lincolnton quadrangles (Overstreet and others, 1963; W. R. Grilfitts, written commun., 1958). The analyses are given in the follow- ing table. A24 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY [ ........ , looked for but not found; n.d., not determined. Spectrtigraphie analyses by J. D. Fletcher, P. R. Barnett, A. A. Chodos. R. G. Havens, A. ’1‘. Myers, and P. J. or Dunton. further explanation of symbols see table, p. A21] Percent Number Rock of Quantitative Semiquantltative samples Pb Sn Be Pb Sn Be Biotite schist 1 n.d. n.d n.d. 0. 001: __________ 0. 000x Biotite schist and gneiss. 10 n.d. n.d 0. 0001 . 00x- __________ n.d. Biotite—oligoclase amiss 5 n.d. n.d . 0002 00): __________ n d Homblende-biotlte-oligoclase gneiss .......... .._ l 0. 003 .......... . 0001 n.d. n.d n d Silllmanite schist 1 and 4 . 01 .......... . 0003 and . 0001 Trace .......... n d Diopside-biotite gneiss 2 002 .......... n.d. n.d. n.d n (1 Granite eiss .-.. 1 n d n d n.d. . 00x .................... Toluca uartz Monzonite ..... 2 and 7 006 .......... n.d. . 00x __________ n d 0001:: . 00x .......... n d Pegmatite related to Toluca Quartz Mnnmnite 1 n.d. n.d. n.d. n d Yorkville Quartz Mnnmnlm ______ 4 n.d. 003 . 00x+ .......... n d Cherryville Quartz Monzonite ................. 7 n d n.d. 00055: . 00x 0 00x 11 d Pegmatite related to Cherryville Quartz Monzonite ...................................... 2 n n.d. 0055: . 00x . 00x :1 d The contents of lead in the rocks and in the magnetite (table 3) are of about the same order of magnitude, but contents are locally as much as an order of magnitude greater in the magnetite than in the rocks. The lead- rich sillimanite schist shown in the quantitative analyses is anomalously pyritiferous. Possibly the presence of the sulfide accounts for the unusual abundance of lead, because four samples of sillimanite schist for which there are semiquantitative analyses contain less than 100 ppm lead. Lead is slightly more abundant in the Toluca Quartz Monzonite and related pegmatites. The possible relation of the Toluca to the lead content of magnetite is discussed on page A26. Tin is consistently less abundant in the rocks than in the magnetite. The low beryllium contents of the rocks are paralleled by low beryllium contents of the magnetite. The close correlation between the distribu- tion of tin‘ and beryllium-rich magnetite and the geo— graphic distribution of the tin~ and beryllium-rich Cherryville Quartz Monzonite and related pegmatites indicates that the Cherryville is the specific source for this magnetite. Semiquantitative spectrographic analyses of 124 heavy-mineral concentrates from the report area show a distribution of lead, tin, and beryllium that is in re- markably close agreement with the distributions shown by the analyses of the magnetite separates. The lead contents of 71 concentrates ranged from 100 to 1,000 ppm; these concentrates came from the core of the belt, where magnetite contains at least 50 ppm lead. Concentrates along the flanks of the belt contain less than 100 ppm lead, and magnetite separates from the same area contain less than 50 ppm lead. Concen- trates from streams in the parts of Pickens, Oconee, and Abbeville Counties in the report area and in western, southern, and southeastern Anderson County also con- tained less than 100 ppm lead. The magnetite in this area is also low in lead. Two concentrates from north- eastern Greenville County and one from southeastern Rutherford County contained 1,000—10,000 ppm lead. The abundance of lead in the concentrates from north- eastern Greenville County was not duplicated in the magnetite, but the lead high in southeastern Rutherford County is about 4 miles southwest of the high for lead in magnetite east of Rutherfordton. The tin contents of 39 of the 124 concentrates ranged from 10 to 1,000 ppm. The distribution of these tin- bearing concentrates follows and extends the pattern shown by tin in magnetite (pl. 23). Tin in concen- trates from the southern zone extends the pattern of tin in magnetite along the line between Pickens and Anderson Counties to Oconee County and thence south- ward to the Seneca River in Anderson County. Tin in the concentrates also makes a closure between the east fork of tin in magnetite at the Saluda River and the tin in magnetite at the Little River in Abbeville County. In the northern zone, tin in concentrates from Cherokee County northwest of Gaffney forms an extension of the high values of tin in magnetite at the State line on the west side of the Broad River and brings it to the river east of Graffney. The great highs for tin in magnetite south of Morganton and in Catawba County are not matched by tin in the heavy-mineral concentrates. The beryllium contents of 31 of the 124 heavy-mineral concentrates ranged from 1 to 10 ppm. The distribu- tion pattern for the beryllium-bearing concentrates resembles that for beryllium in magnetite across Chero- kee, Spartanburg, and Greenville Counties. Beryllium- bearing concentrates also form a strong trend from Gafl'ney northeastward through eastern Cleveland County into western Gaston and Lincoln COunties. North-central Lincoln County and east-central Catawba County have several large drainage basins from which beryllium-bearing concentrates were taken. In the northern zone, beryllium is much more common in the concentrates than in the magnetite. The concentrates probably contain other beryllium-bearing minerals, most likely beryl. In the northern zone, the beryllium content of the concentrates has a decided trend along the west side of MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT the pluton of Cherryville Quartz Monzonite and through the areas of beryl-bearing sheet-muscovite peg- matites in Cleveland, Gaston, Lincoln, and Catawba Counties (Griffitts and Overstreet, 1952, fig. 1; Grif- fitts and Olson, 1953, fig. 77; Griffitts, 1954, fig. 1). In the middle zone, however, the distribution of beryl- lium-bearing concentrates and beryllium-bearing mag- netite in western Cherokee and central Spartanburg Counties is not matched by the occurrence of numerous pegmatite dikes containing sheet muscovite (Griflitts, 1953, fig. 104). In the southern zone an extensive area from which beryllium-bearing concentrates were taken extends southward on the east side of ernville from the north- ern part of the county through the high for beryllium in magnetite nearly to the Saluda River. Beryllium in the concentrates also extends the trend of the beryllium in magnetite from Greenville southeastward along the divide between the Enoree and Reedy Rivers into north- western Laurens County. As in the middle zone, the locations of beryllium-bearing concentrates and high values for beryllium in magnetite in the southern zone are not accompanied by numerous sheet-muscovite-bear- ing pegmatite dikes, though a few dikes occur in south- ern Greenville County (Griffitts, 1953, fig. 104). These dikes are far to the southwest of the southwest end of the pluton of Cherryville Quartz Monzonite (Griflitts and Overstreet, 1952, fig. 1) . Relations of the beryllium-bearing concentrates to beryllium-rich magnetite in the middle and southern zones seem to indicate that pegmatites related to the Cherryville Quartz Monzonite occur in these areas, but these pegmatites do not contain exploitable sheet mus- covite. Hence, they have not been mined, and their dis- tribution has not been mapped. Concentrations of beryllium and tin occur where magnetite from other sources is not so abundant that the values of beryllium or tin in the concentrate or in magnetite are depressed to background or undetectable levels. In the middle and southern zones, pressure and temperature during emplacement of these pegmatite dikes may have been generally too low for crystallization of sheet muscovite. Since the presence of sheet muscovite was previously the criterion for determining the distribution of pegmatites, the continuity of their zone of emplacement was assumed to be interrupted where the muscovite was not found. On the basis of new criteria, it can be inferred that the pegmatite dikes related to the Cherryville Quartz Mon- zonite form a zone 50 miles long that cuts across strati- graphic units and metamorphic facies from the south- west end of the pluton at Gaflney to the Saluda River south of Greenville. Lead in magnetite apparently is remarkably sensitive A25 to the abundance of magnetite in the concentrate; lead highs develop only where amounts of magnetite are small. Closure of the pattern of lead highs at the Saluda River closely parallels the flood of magnetite entering concentrates south of the river. The south- east and northwest edges of the pattern for lead closely conform to the place where the magnetite content rises above 10 percent of the weight of the concentrate. Tin is slightly less sensitive than beryllium to the total abun- dance of magnetite, particularly in the extreme north— eastern and southwestern parts of the report area; nevertheless, tin clearly occurs in areas lean in mag- netite. Beryllium in magnetite is very sensitive to the total amount of magnetite in the concentrate, and it de- clines abruptly in magnetite-rich areas like that south of the Saluda River, eastern Greenville County, north- western Spartanburg County, east-central Cherokee County, and Lincoln and Catawba Counties. No un- usual abundances of lead, tin, or beryllium occur in the great magnetite highs in Polk, Rutherford, McDowell, Pickens, Oconee, Abbeville, Greenwood, and Laurens Counties. The shapes, trends, and sizes of the areas high in tin and beryllium, and the similar spatial relations of these areas to the areas containing sparse magnetite suggest that tin and beryllium are most abundant in magnetite from a single source poor in magnetite. Only small amounts of this magnetite are available for concen- tration in the streams; thus, its effect on the concen- tration of minor elements in the magnetite separate is negligible where magnetite from other sources is abun- dant. Trends of the tin-rich and beryllium-rich areas are across the metamorphic isograds and regional struc- ture. Therefore, the source of the tin- and beryllium- rich magnetite must be younger than the metamorphic climax in which the isograds were formed. The source of the tin- and beryllium-rich magnetite may confidently be ascribed to the Cherryville Quartz Monzonite and related pegmatite dikes, because (1) these rocks are uniquely rich in tin and beryllium; (2) they characteristically are lean in minor accessory minerals, including magnetite; (3) they are postkine— matic intrusive rocks emplaced after the metamorphic climax; and (4) the pluton of Cherryville Quartz Monzonite is a crosscutting body on a regional scale, and the bodies of pegmatite are mostly dikes. The lead distribution may not have resulted from the factors noted for tin and beryllium, despite their simi- lar patterns. The available data do not clearly indi- cate Whether the similarity in distribution arises from similar origins of the lead-rich magnetites or from simi- lar conditions permitting detection of the minor amounts of lead, tin, and beryllium. Analyses of bulk samples of the rock show that the Cherryville Quartz A26 Monzonite and its pegmatites contain no more lead—— possibly even less—than the Toluca Quartz Monzonite and its pegmatites, synkinematic intrusives that occur throughout the lead high from the Catawba River at least as far south as Spartanburg. Lead is absent from such typical areas of Cherryville Quartz Monzonite as near Lincolnton. However, lead is particularly abun— dant northeast of Gafl'ney in the part of the Buffalo Creek drainage basin that is dominantly underlain by Cherryville Quartz Monzonite. The area containing above-average amounts of lead is larger than that for tin or beryllium, but lead is more abundant in magnetite than is tin or beryllium. Perhaps both the Toluca and the Cherryville have contributed lead-rich magnetite to the streams, and the present broad lead pattern reflects two sources. Lead-rich magnetite has not come from the low-grade metamorphic rocks outside the core of the Inner Piedmont belt nor from the magnetite-rich post- kinematic granitic rocks and granodiorite 1n Anderson County. Close correlation of lead and tin with chromium and copper appears to result from the mixing of chromium- and copper-rich magnetite of metamorphic origin with lead- and tin-bearing magnetite of igneous origin in the core of the Inner Piedmont belt, where only small amounts of magnetite occur. Loss of above-average amounts of beryllium in the northern zone because of influx of beryllium-deficient magnetite of metamorphic origin may explain the di- chotomy in the data for beryllium in magnetite. The amount of beryllium in the concentrates does not de- crease along the southeast side of the northern zone. VANADIUM AND BARIUM Contents of vanadium and barium have been spectro- graphically determined for the crystalline rocks around Shelby, Lincolnton, Gafl'ney, and Kings Mountain (Overstreet and others, 1963; W. R. Griflitts, written commun, 1958): Spectrographic analyses byJ. D. Fletcher, P. R. Barnett, A. A. Chodos, R. G. Havenls, A. T. Myers, and P. J. Dunton. For explanation of symbols see table, p. A21 Percent Number Rock of Quantitative Semiquantitative samples V Ba V Ba Biotite schist ________________________ 0. 00): 0. 0x+ Biotite schist and gneiss ............. 00x+ . 0x Biotiteroli oclase gneiss ______________ ox- . 0x+ Hornblen e—biotite—oligoclase gneiss. __________________ Sillimanite schist _____ . 0011+ 0x— Diopside- biotite gneiss_. __________________ Granite neiss ___________ . 00x- 0x- Toluca uartz Monzonite ........... 00x- 0X+ Pegmatite related to Toluca Quartz Monzonite ......................... 1 . 001 . 09 .................. Yorkville Quartz Monzonite _________ 4 ................ . 00x .1:- Cherryville Quartz Monzonite ....... 7 ................ . 00x- . 0X Pegmatite related to Cherryville Quartz Monzonite _________________ 2 ________________ 0011- 0x- SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Vanadium and barium were spectrographically deter- mined in 124 heavy-mineral concentrates from the Inner Piedmont belt. One concentrate from the flank of the belt in the northwestern part of the northern zone con- tained 1,000—10,000 ppm vanadium. The vanadium con- tent ranged from 100 to 1,000 ppm in 63 concentrates, predominantly from the flanks of the belt, and from 10 to 100 ppm in 55 concentrates, mainly from the core of the belt. Three of the five concentrates that lacked vanadium are from Anderson County; the remaining two came from the core of the belt in the southern and northern zones. The highest value for barium in the concentrate; was 1,000 ppm in a sample from the core of the belt in the northern zone, near the mutual corner of Cleveland, Gaston, and ”Lincoln Counties. The barium content ranged from 10 to 100 ppm in 87 con- centrates, many of which came from the flanks of the belt. Two samples from the core and one from the northwest flank of the northern zone contained 1—10 ppm barium. Barium was absent from 33 concentrates, mostly from the middle zone and the core of the northern zone. The distributibn of vanadium in the concentrates accords with the postulated behavior of vanadium in magnetite. Vanadium is more abundant on the flanks of the belt than in the core. As currently understood, the main factor controlling the abundance of vanadium in magnetite is progressive regional metamorphism. As metamorphic grade increases, vanadium in magnetite decreases, whereas copper increases. Barium in the heavy- mineral concentrates tends to behave conversely to its postulated behavior in mag- netite. In the concentrates it tends to be most abundant along the flanks of the belt, whereas in magnetite it is most abundant in the core of the belt. Available data are inadequate to resolve these apparent conflicting ob- servations, but the concentrates may contain another barium- bearing mineral, such as barite, which would reverse the postulated trend. PROMINENT HIGHS Anomalous abundances of the metals that may war- rant detailed investigation occur at coincident or ad- j acent highs of several metals in magnetite. The most prominent of these highs are on lower Buffalo Creek, on the Pacolet River, at Greenville, and in the divide be- tween the First Broad and Second Broad Rivers. The prominent highs for lead, tin, copper, and zinc in the lower parts of the Buffalo Creek drainage basin in Cleveland County, N.C., and Cherokee County, 8.0., are localized in a faulted area where Cherryville Quartz Monzonite intrudes rocks ranging in metamorphic grade from the greenschist fvacies to the staurolite- MINOR ELEMENTS IN ALLUVIAL MAGNETITE FROM THE INNER PIEDMONT BELT kyanite subfacies. Small amounts of copper and lead have been mined from a vein a few miles to the south- west, in an area that is not shown by these data to be anomalously metalliferous. The lower Buffalo Creek high is probably the best place in the sampled area for further work, particularly for a detailed geochemical survey. The Greenvi'lle area, from Paris Mountain southward to the Enoree River, contains overlapping highs for lead, tin, beryllium, zinc, and copper. It is a very com- plex region of high-rank metamorphic rocks that are tightly folded, athwart the regional trend, widely in- truded by granite and pegmatite, and doubtless much faulted. It is the second best place in the sampled area for further work, in particular for detailed geologic mapping and for geochemical survey. A small area of high values for copper and tin in the western tributaries of the Pacolet River north of Spar- tanburg, S.C., generally also contains above-average amounts of zinc, lead, and beryllium. The area is north of a major synclinal crosswarp in a region of very high grade metamorphic rocks about which geologic details are unknown. The weak high defined by zinc, copper, and tin that extends eastward from the Second Broad River to western tributaries of the First Broad River northeast of Rutherfordton, N.C., leads into a faulted zone at the west side of the Shelby quadrangle, North Carolina, and is a western projection of that zone. The faulted zone in the Shelby area is not a geochemical high, nor was base-metal mineralization evident in the mapped part of the zone. Therefore, whatever metallization occurs in this zone probably is greater west of the mapped area. Detailed geologic and geochemical map- ping of this fault zone would further define patterns of mineralization in this part of the Piedmont. Some of the tin- and beryllium-rich areas may contain spodumene-bearing pegmatite dikes, because cassiterite and beryl are common accessory minerals in the spod- umene pegmatites of the tin-spodumene belt east of the report area (Griflitts, 1954, p. 3). On the basis of spectrographic analyses of heavy-mineral concentrates made in 1951—52, the occurrence of economic deposits of cassiterite is considered unlikely. {GEOCHEMICAL INTERPRETATIONS A summary of the geochemical features and problems related to the study of the trace-element composition of the magnetite separates includes these pertinent fea- tures: (1) Some of the magnetite separates contain fairly large quantities of titanium, chromium, manga- nese, copper, zinc, lead, tin, beryllium, vanadium, and barium. For most of these elements and separates, the A27 metal content of the magnetite is probably greater than that of the alluvium from which the separates were ob- tained. (2) The content of each metal varies rather widely among the analyzed separates. The range in values obtained for these elements in the magnetite separates probably greatly exceeds the range that could be obtained for the alluvium. (3) The elements are grouped according to the general level of concentration, the geographic pattern of distribution, and the associa- tions among the elements. This grouping is compatible with the chemical properties of the elements, particular- ly with the four- and six—coordinated bivalent and triva- lent ionic radii found in spinel structures. The following conclusions seem warranted on the basis of these facts. A general partition of many of the trace elements favors magnetite over most other rock- forming minerals. The amount of metal attracted and retained by the magnetite will be strongly influenced by the amount of metal available. The availability of the metal is controlled by the absolute quantity of metal in the system, the nature of the system, and the nature of the individual elements. The first of these controls, in large part geologic, was considered under “Geologic Interpretation.” The second cannot be evaluated from these data but seems to be of minor importance. The third results in the grouping of the elements by their chemical properties and involves the nature of incorpo- ration of the metals in the concentrates. The most likely location for the metal in the separates seems, from the limited data available, to be the magne- tite lattice or another spinel in solid solution with the magnetite or an exsolved phase. Though largely specu- lative, this conclusion provides a basis for further work. The possibilities for variety in spinels, beyond those found in natural systems to date, deserve attention. The presence of such a variety of elements interacting in a single relatively simple structure should provide an unusual opportunity to study the vagaries not only of the structure itself but also of the various trace elements. Magnetite, particularly separates from alluvium, is an almost ideal medium for the study of the distribution of some trace elements on a regional scale—a form of geochemical census. Sampling bedrock for such a pro- gram involves technical problems of selective outcrop, gross differences among rock types, and other problems that now seem insurmountable. Sampling of soils in- volves similar problems. For these spot samples, the basic problem is to collect average samples useful for regional comparisons. Therefore, the most successful minor-element—distribution maps produced have been based upon stream sediments. (See, for example, the A28 excellent maps of northern Nova Scotia by Holman and Gilbert, 1959a, b, c, d.) Evidently, one of the most objective and readily obtainable average samples is that obtained from eroded and transported debris in streams. Stream sediments, however, vary in metal content not only with the metal content of bedrock in the drainage basin but also with the erosion, transportation, and sedi- mentary history of the sediment. The effects of this variation are minimized, however, because only a single component of the sediment, magnetite, is used. The use of magnetite has the further advantage of raising the general background of metal content to levels more easily reached by present analytical procedures. There can be no doubt from the data presented that the regional distribution of some trace elements can be studied by the analysis of detrital magnetite. Before general studies of this type can be undertaken and inter- preted, however, stated problems concerning the basic geochemistry of magnetite must be at least partly an- swered. In addition, it will be desirable to have some information on the transformations undergone by mag- netite, and particularly its trace-element constitution, during weathering, erosion, and transportation. For- tunately, such information is becoming available through studies of the magnetite-maghemite-hematite series and of the solubility of trace elements in mag- netite. LITERATURE CITED Barth, T. F. W., and Posnjak, Eugen, 1931, The spinel struc- ture—an example of variate atom equipoints: Washington Acad. Sci. Jour., v. 21, no. 12, p. 255—258. Basta, E. Z., 1959, Some mineralogical relationships in the sys- tem Fe203—F‘e304 and the composition of titanomaghemite: Econ. Geology, v. 54, no. 4, p. 698—719. Bell, Henry, 3d, 1960, A synthesis of geologic work in the Con- cord area, North Carolina, in Short papers in the geological sciences: U.S. Geol. Survey Prof. Paper 400—3, p. 3189— 3191. Crickmay, G. W., 1952, Geology of the crystalline rocks of Georgia: Georgia Dept. Mines, Mining, and Geology BulL 58, 54 p. Grant, W. H., 1958, The geology of Hart County, Georgia: Georgia Geol. Survey Bull. 67, 75 p. Green, Jack, 1959, Geochemical table of the elements for 1959: Geol. ‘Soc. America Bull, v. 70, no. 2, p. 1127—1184. Green, W. D., and Carpenter, David, 1961, The association of uranium and thorium in magnetite: Am. Assoc. Adv. Sci., Southwestern and Rocky Mtn., Div. Mtg. Program, Tempe, Ariz. Griflitts, W. R., 1953, Hartwell district, Georgia and South Caro- lina, pt. 7 of Mica deposits of the southeastern Piedmont: U.S. Geol. Survey Prof. Paper 248-E, p. 293—316. 1954, Beryllium resources of the tin-spodumene belt, North Carolina: U.S. Geol. Survey Circ. 309, 12 p. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Griflitts, W. R., and Olson, J. 0., 1953, Shelby-Hickory district, North Carolina, pt. 5 of Mica deposits of the southeastern Piedmont: U.S. Geol. Survey Prof. Paper 248-D, p. 203—281. Griflitts, W. R., and Overstreet, W. 0., 1952, Granitic rocks of the western Carolina Piedmont: Am. Jour. Sci., v. 250, no. 11, p. 777—789. Holman, R. H. 0., and Gilbert, M. A., 1959a, Geochemistry, zinc in stream sediments, northern mainland of Nova Scotia: Canada Geol. Survey Map 25—1959. 1959b, Geochemistry, lead in stream sediments, northern mainland of Nova Scotia: Canada Geol. Survey Map 26-1959. 1959c, Geochemistry, copper in stream sediments, north- ern mainland of Nova Scotia: Canada Geol. Survey Map 27—1959. 1959a, Geochemistry, heavy metals in stream sediments, northern mainland of Nova Scotia: Canada Geol. Survey Map 33—1959. Keith, Arthur, and Sterrett, D. B. (King, P. 3., ed.), 1907, Geologic map of Morganton quadrangle, North Carolina: U. S. Geol. Survey open-file map. 1931, Description of the Gaflney and Kings Mountain quadrangles, South Carolina—North Carolina: U.S. Geol. Survey Geol. Atlas, Folio 222, 13 p. King, P. B., 1955, A geologic section across the southern Appa- lachians—an outline of the geology in the segment in Ten- nessee, North Carolina, and ‘South Carolina, «in Russell, R. J ., ed., Guides to southeastern geology: Geol. Soc. Amer- ica Guidebook, Ann. Mtg., New York, 1955, p. 332-873. Lieber, O. M., 1860, Report on the survey of South Carolina: South Carolina Geognostic Survey Repts., v. 4, 194. Mellor, J. W., 1931, A comprehensive treatise on inorganic and theoretical chemistry, v. 11, Te, Cr, Mo, W: London, Long- mans, Green & 00., 909 p. 1934, A comprehensive treatise on inorganic and theoreti- cal chemistry, v. 13, Fe (pt. 2) : London, Longmans, Green & Co., 948 p. Overstreet, W. 0., 1962, A review of regional heavy-mineral reconnaissance and its application in the southeastern Pied- mont: Southeastern Geology, v. 3, no. 3, p. 133-173. Overstreet, W. C., and Bell, Henry, 3d, 1965a, The Crystalline rocks of South Carolina: U.S. Geol. Survey Bull. 1183, 126 p 1965b, Geologic map of the crystalline rocks of South Carolina: U.S. Geol. Survey Misc. Geol. Inv. Map L413. Overstreet, W. C., Cuppels, N. P., and White, A. M., 1956, Mon- azite in southeastern United States, in Page, L. R., Stocking, H. E., and Smith, H. 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L., 1951, The distribution of trace elements during strong fractionation of basic magma—a further study of the Skaergaard intrusion, East Greenland: Geochim. et Cosmochim. Acta, v. 1, no. 3, p. 129-208. Ward, F. N., Lakin, H. W., Canney, F. 0., and others, 1963, Analytical methods used in geochemical exploration by the U.S. Geological Survey: U.S. Geol. Survey Bull. 1152, 100 p. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY A30 cHV HV ooo .H ............ HV ooo .H cco HoH ooH oon .H cooV ooH cm on 3. on m .H H H o ooH own on oV HV .................... m ooo ooo .m co com coaV on on E. ooH ooH H .u m a n .coH own on oV HV 1 .................... NV coo ooo .H. 3 com .H oomV om oH E. 3 3V w . m H. a do. «on no nV HV ..................... NV ooH ooo .m oHV cco .H comV oH oHV HRH on 3V w Ho Ho 8 m 63 ooo on oV HV ...................... 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ATOMIC ENERGY COMMISSION PLATE 1 EXPLANATION EXPLANATION ° . (3 Chromium content, I 3g Sample locality . If in parts per million 10—percent-magnetite isogram ' I: 158 ” "/V Arrow indicates direction of increasing mag- netite content. Isogram controlled by about 4000 samples Outline of drainage area above sample locality Dashed where within the drainage area of another sample NORTHERN ZONE MIDDLE ZONE SOUTHERN ZONE \‘ o . . . :5 ' fig: ‘1‘" > ,0“. I , . x . , ~w \.-,_,I-._; .- ' I / E . . . g-I. . x; 1 . . . x ,4 I , \ {fl I WWI I VF " ’"I'Xye‘w/ ‘ i , ; I ,1 "Iv? urge mom 'I- -’ I ' 7 ,I J 7“ III 7, / if \m , /. I I _ @R‘ Kn” y : Kr” V}; \ A ‘7???» . ‘\\“§‘V /’ , ,/,.,fl47 , A; . ,%§&>§%%g I \M\ I} G”? 17/ M§§§®>V J/é/j "‘ x ‘ ‘I \' //<{\\\‘ V 1,. ' 1“ £34 ' ' If} 352 x I «f, If» I» « sea»- I {OENQI m{%\ j I mix} ’ ' / ”VI I -— I. file/N \ _» _, I\/ . N L" \ 2 i 3 "J 3%“ _ of; 1' z‘ 3v 3/ “\ ~__ .x’ "W” \I, I A f I II x (w, ,1 § I \I V. , i so _ / __ _ I; / I \ r , x I «z I -’ __ , w I m . ~~ \ I. l I.‘ , fi' _ i \3 t \«V; V, .- l: A} ' ' x ,2 \3 {} EJJ é ~\r {i I ‘9 « I : ‘ K .5 _. “I “m 7 I “Q; ,’ NW4 , _ / /’ «I; _, v I “ ~\ .....I e I ’ ‘ I C ‘ . I / K / A s I » I II } / 1 / I! I x ‘ J , f 1 I \l , ’ u-u.‘ ‘ x l - 7”" ~ j _- / K V » x ® \ I ‘I ,- “ ”,3, W x ”A: ’ ‘~\., I g’ \k ‘I \ ‘ /.,,r , {I I ( / ,_~‘~.-~- r" ) E x _ k /M‘ \ L‘,‘ 2 / I "‘ I _ 5, A} x I \ , I I I,“ \\L 2 I 4 II 7 f I \I I sf ~~ , ,,.,-.»w”\.rKW» a ‘1 x" x \V f?” «iii/A {S y? i.) N \ AIYK,‘ _/ 2 ”/1 ' '- I I 5/ 2 \ ‘ w “"* ,- , , L/ ‘ V” ) x (I, A” A. SAMPLE LOCALITIES AND GEOGRAPHIC SUBDIVISION OF MAGNETITE SEPARATES B. CHROMIUM EXPLANATION E X P LA N AT I O N Manganese content, ____l__._ i . . . Axe Chromium content, M 8 111 parts per "11111011 10—percent-magnetite isogram _ c; m parts per mlllion 10-percent—magnetite is 0 gr am a? ........ Arrow indicates direction of increasing mag- ' , Arrow indicates direction of increasing mag- -------- netite content. Isogram controlled by about ........ netite content. Isogram controlled by about 4000 samples 4000 samples 700 or less 2000 or less A Jir’ e‘» I amonngw HM? <5“ <\ (\ rr’ 2’ _, ‘ r KW /" / K \ (7 x \ / I ; 1‘ In “« x’l’” I ' K 7 s I) I \. II I a ‘ . > r 90, E / 5 ~\ I l! . .' I. x N g \- \ K a = I I «I h 3 3 _ I \, II ’V I = . '. 7' 4 ”J \ _ \ \ I‘ “ t V . 2 ‘ , . I . . \ ‘ r4» , I Y I I} f , , 3000—5000 ; ~ 1000— 1500 20003000 5000 or more w w, in? we ’ ’ ' ‘ >231Irrara-nt*17n . I ”I“ z r 7’ “' M ”$7,ka I m moo—10,000 P109 ....... ‘Y/ if} l ----- / . ' » II ,2 ,3 > ”I More than 10,000 \_ Elam" Ii/wja’vx 7.» \9 x ««««« MRI , I“ \(Jy (f I / ,n )I I _M f I’ .. / I0 V90 ) * .. I ,' I ’ 0 i I ' ' ’ ‘ )6” “13230397 if“ {K ' i \ , /’ _// . L ' ......\, I \ a , ....... \E_>/__\V,~.m_ \31:~\I ‘ .. 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COPPER \ .............. .\.,, ....................................... 4% x ........... xl/ aw); _ . will; m ......... f, i/ s ...... a; ....... f. x . // \ ............... a. 5 ~ 1V3 / .................. . . /l|ll ................ ................... ....... 10—percent-magnetite isogram 10—percent—magnetite isogram 10—percent-magnetite isogram netite content. Isogram controlled by about netite content. Isogram controlled by about netite content. Isogram controlled by about .\\\\M«h\\ xxwégkmz . \®®\xx§ \\VM.\\ I ,. L"“/3/ru I v , «hm AV /«\\,\>’ a: RT}. 9 ‘V?Ifl/a.//Ir4 , , .. « ”V I ./ ,/ ‘ . \x///fl/////////, W...mmnmm.y\x..&/////%/m t...”“WEEK”All",.w.mg,..4. 4000 samples 1,000 samples 4000 samples Arrow indicates direction of increasing mag— Arrow indicates direction of increasing mag- Arrow indicates direction of increasing mag- EXPLANATION EXPLANATION / \\ ,. x EXPLANATION .....,..\ ...... IIIIIIIIIII 300 or more 50— 100 125—250 N 20—40 90— 160 N 100—200 \ \\ 50—80 \N 15 or less 180 or more Zinc content,in 250 or more Lead content,in parts per million Copper content,in parts per million A Abbevilie fl Abbeville L» _\ ,i ; ‘3‘ ' ‘( *ri}. :5\ GEOLOGICAL SURVEY UNITED STATES DEPARTMENT OF THE INTERIOR Abbeville INTERIOR—GEOLOGICAL SURVEY, WASHINGTON, D.C.—1967—666283 F. MAGNETITE AND SILLIMANITE ISOGRAMS MAPS SHOWING DISTRIBUTION OF COPPER, ZINC, LEAD, TIN, AND BERYLLIUM, AND MAGNETITE AND SILLIMANITE ISOGRAMS 50 MILES 40 5o KILOMETERS 4O 3O 30 IN DEX MAP 20 20 10 10 HHHI—IHI»———-———l 10 I—II—ll—«IHI—ll————I 10 INNER PIEDMONT BELT, NORTH AND SOUTH CAROLINA E. LEAD Base from County road imao's'frdih‘ South Carolina State Highway Department and North'Carolina State Highway and PublicEWor‘ks Commission 057? W , “(553415 Big Snowy and Amsden Groups and the 7 DAY Mississippian-Pennsylvanian Boundary in Montana SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY ‘- GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—B Big Snowy and Amsden Groups and the Mississippian-Pennsylvanian Boundary in Montana By EDWIN K. MAUGHAN and ALBERT E. ROBERTS SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—B fl study of t/ze Upper Mississippian ana’ Pennsylvanian sedimentary formations witn special empnasis on t/zeir regional correlation, variation in litfiology, anaI age assignments UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1967 UNITED STATES DEPARTMENT OF THE INTERIOR STEWART L. UDALL, Secretary GEOLOGICAL SURVEY William T. Pecora, Director For sale by the Superintendent of Documents, US. Government Printing Office Washington, DC. 20402 CONTENTS Page Abstract ___________________________________________ B1 Mississippian—Pennsylvanian boundary ________________ Introduction _______________________________________ 1 Tectonic framework _________________________________ Development of nomenclature ________________________ 1 Early and Late Mississippian (Madison Group)- _ _ _ Mississippian System, Big Snowy Group _______________ 5 Late Mississippian (Big Snowy Group) ____________ Kibbey Formation ______________________________ 5 Early and Middle Pennsylvanian (Tyler Formation Otter Formation ________________________________ 7 and Alaska Bench Limestone) __________________ Heath Formation _______________________________ 7 Middle and Late Pennsylvanian (Devils Pocket, Regional relations ______________________________ 10 Quadrant, and Tensleep Formations) ____________ Pennsylvanian System, Amsden Group ________________ 11 Permian to Jurassic events _______________________ Tyler Formation ________________________________ 11 References cited ____________________________________ Alaska Bench Limestone _________________________ 14 Devils Pocket Formation ________________________ 15 Regional relations ______________________________ 19 ILLUSTRATIONS PLATE FIGURE TABLE i—l [Plates are in pocket] . Columnar sections showing correlation of Big Snowy and Amsden Groups in central and southern Montana. . Columnar sections showing correlation of Big Snowy and Amsden Groups in central and eastern Montana. Paleogeologic map showing distribution of Big Snowy and Madison Groups in Montana. Chart showing stratigraphic distribution and comparative ranges of fossils from the Big Snowy and Amsden Groups in central Montana. . Chart showing development of nomenclature for Mississippian and Pennsylvanian rocks in central Montana- _ _ _ . Index map of Montana and adjacent areas ______________________________________________________________ . Chart showing nomenclature and stratigraphic relations of the Upper Mississippian and Pennsylvanian rocks in Montana, Wyoming, South Dakota, and North Dakota ____________________________________________ . Index map of localities in Big Snowy Mountains and vicinity, Montana ____________________________________ . Diagrammatic section showing relations of Big Snowy and Amsden Groups in central Montana ______________ . Diagrammatic section of erosional unconformity between Heath Formation (restricted) and Stonehouse Canyon Member of the Tyler Formation exposed at west end of Middle Bench, Little Snowy Mountains __________ . Photograph showing unconformable contact of the Stonehouse Canyon Member of the Tyler Formation upon restricted Heath Formation _______________________________________________________________________ TABLES . Comparison of thicknesses of Heath Formation and Stonehouse Canyon Member of the Tyler Formation ______ . Pollen and spores recovered from Mississippian and Pennsylvanian samples from Montana ____________________ III Page B20 23 23 23 24 24 25 25 Page B2 Hm 12 14 Page: B 7 22 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY BIG SNOWY AND AMSDEN GROUPS AND THE MISSISSIPPIAN-PENNSYLVANIAN BOUNDARY IN MONTANA By EDWIN K. MAUGHAN and ALBERT E. ROBERTS ABSTRACT The Big Snowy Group is redefined to include the Ki‘bbey, Otter, and Health Formations, as the group was originally estab— lished. However, the Heath Formation, and consequently the top of the Big Snowy Group, is restricted to strata beneath a Late Mississippian to Early Pennsylvanian regionlal unconform— ity. This restriction limits the Big Snowy to three closely re- lated formations that comprise one sedimentary cycle uncom- plicated by intraformational structural movements. Also, the Big Snowy Group, as restricted, closely approximates a time— stra‘tigraphic unit of Late Mississippian age. The Amsden Group unc‘onformably overlies the Big Snowy or Madison Groups and consists of three formations in central Montana that are stra'tigraphically and lithologically nearly equivalent touthe type Amsden Formation in northern Wyoming. The formations, in ascending order, are the Tyler Formation, the Alaska Bench Limestone, and Devils Pocket Formation. The Tyler Formation is locally divided into a Stonehouse Canyon Member at the base and a Cameron Creek Member at the top. Spores collected from the upper part of the Stonehouse Canyon and Cameron Creek Members are of Early Pennsylvanian age. Fusulinids from the Alaska Bench Limestone suggest a Morrow and Atoka age, and those from the Devils Pocket Formation are of Atoka or early Des Moines age. Regional upwarp in south-central Montana and much of adja- cent ‘Vyoming took place near the end of Mississippian time. The area of uplift was bounded 011 the north by a system of probable faults and monoclinal folds. Rocks of the Big Snowy Group were stripped south of this structural belt, and were tilted northward and beveled north of this belt. Subsequent to erosion, the region gradually submerged during Early Pennsylvanian time. Seas inundated the region and detrital sediments of the basal part of the Amsden Group were deposited unconformably on Upper Mississippian rocks in central Montana, and on older rocks farther south in southern Montana and northern Wyoming. INTRODUCTION Upper Paleozoic rocks of the Big Snowy and overly- ing Amsden Groups of central and western Montana show wide variations in lithology and thickness and have long presented problems of identification, correla- tion, and dating. This paper summarizes the stratig- raphy and presents revisions in nomenclature compati- ble with present understanding of the relations of the rocks comprising the upper part of the Mississippian and lower part of the Pennsylvanian of this region. Study of these rocks was begun in 1957 and particular attention has been given to the problem of the systemic boundary since 1960. The authors have collected sur- face and subsurface stratigraphic information through- out Montana and adjacent States. Surface sections of the west half of Montana have been Visited and studied in detail; subsurface sections in the east half have been studied from sample and geophysical logs. In central Montana special emphasis has been given to detailed stratigraphic correlation, description, and fossil content of individual units. New information re- garding the ages of these rocks is reported, and the equivocal position of the boundary between the Mis- sissippian and Pennsylvanian Systems is resolved. Much of this basic data has been obtained from pub- lished and unpublished reports and credit is given to the original sources throughout this report. However, interpretations of these data are those of the authors and we assume full responsibility for them. The boundary between the Mississippian and Penn- sylvanian Systems, which is believed to be at the un- conformable contact of the Big Snowy and Amsden Groups, is given particular attention. Four groups of detailed columnar sections (pls. 1, 2) are presented to illustrate the unconformity between the Big Snowy and Amsden Groups and to show the regional correla- tion of the stratigraphic units that compose these groups. DEVELOPMENT OF NOMENCLATURE Quadrant Formation was first used in the Three Forks, Mont, area by Peale (1893) (as shown in fig. 1) for strata between underlying Madison Limestone (Mis- sissippian) and overlying Ellis Formation (Jurassic). Quadrant Quartzite was applied by Weed (1896) and by Iddings and Weed (1899) (as shown in fig. 1) at its B1 B2 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY Peale ldd‘ii-ieesdziriggveieed weed Freeman Reeves Scott Sloss (1893) $1899) (1900) (1922) (1931) (1935) (1952) .2 .2 I o 2 Ellis = a! . '3 . . '5 . . '3 . '3 3 . .‘Eg Teton formation 6: Ellis formation 100 feet of black shale 3 Ellis formation 8 Ellis formation 8 Ellis group é: formation 5,: ‘5 5 5 5 a, —. q. a -. LL13 v = Q ‘g Upper .9 _ c = 5 Minnelusa .: (Not described) E E 'E 'u . , N “3 g 5 Undifferentiated E. .E 24 3., g 4- E g “ E Upper Amsden O 0 . - Alaska Bench w 0 A d 0- uadrant uadrant uartzite .2 - “- “- ms en Q . Q q E g E limestone 2 formation formation or sandstone E 2 E 100 feet of 2 Lower Amsden g g g =0 5 _?_ .8'£¥_SlE|£__7_ a W E E _7_ _ ._ 'a Quadrant ._ H th q . g {3 . g ‘5 Tyler sandstone a fan-nation :0 forrfigtion :1 Heath formation s: = 'U to .— B 3 ° ° ‘° Otter shale ‘ '0 o c e E 3 E Otter shale ‘9 = Otter = Otter formation to (v m = E ‘0 formation = V1 "’ ° 8 °' 5 .2» . 5 .20 5 Kibbey .. M '2 °° “my ‘"‘ é “ Kibbey ““8““ .E A w ._ m in g g 'g ,3 Charles formation a = .2 E = - . E E . 2 . E, Mission Canyon Madison Madison o _ , 9 Madison Madison limestone . _ 3 Castle limestone (Not given) 7» _ . S limestone limestone 3 g limestone limestone 3 .2 ‘5 Lodgepole ID 2 E limestone FIGURE 1.—Development of nomenclature for Mississippian and Pennsylvanian rocks in central Montana. Scott (1935) originally placed all the Amsden in the Mississippian but later (1945; 1950, p. 48) placed part in the Pennsylvanian on the basis of fusulinids. type section on Quadrant Mountain in the northwestern part of Yellowstone National Park (fig. 2) for strata between the underlying Madison Limestone and the overlying Permian and Triassic Teton Formation. The name Teton was abandoned and replaced by the Per- mian Phosphoria Formation or its equivalents and the Triassic Dinwoody Formation. The more inclusive concept of Peale’s Quadrant was widely used for a time in western Montana, but these strata have been subse- quently separated in this region into the Big Snowy Group, Amsden Formation, Quadrant Quartzite (con— forming to Weed’s definition), and Phosphoria or Park City Formation or Shedhorn Sandstone. Freeman (1922) subdivided the Quadrant Formation in the east- ern part of the Big Snowy Mountains into the follow- ing: Kibbey Sandstone, Otter Shale, Tyler Sandstone, and Alaska Bench Limestone (fig. 1). Reeves (1931), in his reconnaissance mapping of the Big Snowy Moun- tain‘s, included all rocks between the Madison Limestone and the Ellis Formation in the Quadrant Formation rather than divide them into the units recommended by Freeman. The Big Snowy Group, as established by Scott (1935), consisted of the Kibbey and Otter Formations, previously named by Weed (1900), for units in the Little Belt Mountains, and the Heath Formation which Scott recognized and named as a unit overlying the Otter in the Big Snowy Mountains (fig. 1) . The Heath is not in the vicinity of the type sections of the Kibbey and Otter Formations in the Belt Creek area (fig. 1) where the Ellis Group of Jurassic age lies unconform- ably over the Otter (Easton, 1962, p. 114), but it is pres- ent a few miles east. Farther east, in the Big Snowy Mountains, Scott called red mudstone and limestone conformably overlying his Heath Formation the Ams- den Formation. He assumed, as have many others, that these rocks were a northward extension of a similar se- quence that comprises the Amsden Formation at its type locality (Darton, 1904, p. 396, 397) in the northern Big- horn Mountains, Wyo., and that is well exposed in the Pryor Mountains, Mont., about 95 miles south of the Big Snowy Mountains. The base of the Big Snowy Group was lowered by Seager (1942, p. 864) to include the Charles Formation, which he described as a series of evaporite and dolomite beds lying between the basal “member” of Scott’s Big Snowy Group and the Madison Group. He included strata equivalent to the lower part of the Kibbey For- mation in his Charles. Seager suggested that the Charles possibly should be included with the Madison; but, as he reasoned, a time break was indicated by poros- ity in the underlying limestones in the upper part of the Madison. Therefore, he included the Charles in the Big Snowy Group as the basal “member.” Perry and Sloss (1943) also included the Charles Formation in the Big Snowy Group. The Charles was redefined in the BIG SNOWY AND AMSDEN GROUPS IN MONTANA B3 Mundt Gardner Willis Th‘ (1956a) (1959) (1959) '5 paper ,3 .e .2 g Ellis group Jurassic Ellis group g Ellis group a Ellis Group a ‘5 E m _~W_/; Triassic(?), ' g Tensleep formation Permian(?), and Pennsylvanian(?) g? Quad an F/rma/tio E Undifferentiated Tensleep 2° —: Amsden Pennsyl- Devils Pocket E “m— dolomite E ‘7‘ Devils Pocket 2 formation vanian formation 5 g Amsden member 5 m g. Limestone 5 MW g 2 formation a g 9 “' Alaska Bench Mississip' g- Alaska Bench g -— (restricted) Alaska Bench 2 E ‘i “3.5"" Bench _.7_ formation pan 5 formafion E ; member E __7_ § Limestone PennsyI- > Cameron Creek g Tyler Cameron Creek 5 5: Tyler Cameron CFSEK Tyler vanian 3 formation 2 , member m 2: , Member r formation 3 H [h f t’ formano" lower mbr g g Formatlon Stonehzufigefianyon 5 Heath fm a :3“ orma ion 2 Heath formation *3 g Heath Fm ” ‘ c 5 <5 5 E Otter fm :3 Otter formation g E Otter formation g g g OtterFormation 'a .,=, .c w o z c 7. a .3 Kibbey fm g Kibbey sandstone 5 o .29 Kibbey formation = U :3. FoKrlrggliyon .2 a N E. _ _r_v o no , g 5 Charles fm g a Charles formation .3 {1 Charles formation g g a Charles Formation = '7) = )2 :i II) b :i 1 =90 Mission Canyon é ._ 3 Mission Canyon _3 a Mission Canyon 'g in 3 Mission Canyon = limestone g I: limestone E = limestone E g z Limestone o o o g 8 o 2 ~ —i .2 .6 o .2 '° Lodgepole '5 Lodgepole a . %‘ '° . “ ‘° Lodgepole limestone ‘4: ‘° Lodgepole Limestone E limestone E limestone E g}: E X FIGURE 1.—Continued type well and included in the Madison Group by Nord— quist (1953). With the Charles included, Nordquist (1953, p. 73) lowered the upper contact of the Madison to where the lithology changes between the clastics of the Kibbey Formation and the carbonates and evapo- rites‘of the Madison Group. The Charles Formation, as thus redefined, contained characteristic anhydrites and carbonate lithology similar to that of the under- lying Mission Canyon Limestone. Subsequent workers have established an intertonguing depositional relation between the Mission Canyon and the Charles, and Nord- quist’s assignment of these rocks to the Madison Group is now accepted by most stratigraphers. The Kibbey Formation remains as Scott defined it, the basal unit of the Big Snowy Group. A twofold division of the Heath, as defined by Scott, was made by Beekly (1955). He separated the two units at an unconformity and restricted the name Heath Formation to interbedded black shale and thin limestone strata of marine origin which lie below the unconform- ity. Beekly considered the unnamed upper unit above this unconformity as a basal member of the Amsden. This unnamed member, composed of interbedded sand- stone and black shale, was described by him as “a diverse facies of lagoonal, deltaic, and estuarine deposits vary- ing from marine to nonmarine” deposited on an irreg- ular erosional surface following deposition of the restricted Heath. He divided the remainder of the overlying Amsden into three additional members‘which, in ascending order, are a “Basal (sic) Amsden sand,” “Amsden red beds,” and “Amsden carbonates.” Mundt (1956a) restricted the Heath Formation, as Beekly did (fig. 1), to marine interstratified black cal- careous shale and limestone beneath an unconformity. Rocks between the unconformity and the base of Beekly’s “Amsden carbonates,” Mundt called the Tyler Formation, a restricted use of the name previously in- troduced by Freeman (1922). Overlying the Tyler Formation, as used by Mundt, is limestone he called the Alaska Bench Formation in the same sense that this name was proposed and used by Freeman (1922) (fig. 1). Rocks above the Alaska Bench were called the Ams- den Formation by Mundt (1956a). They include cherty dolomite lying unconformably upon the Alaska Bench but are only part of the carbonate member of the Ams- den Formation of other geologists in the Big Snowy area. The Amsden Formation of Mundt correlates with only the upper red shale member of the Amsden at its type section in the Bighorn Mountains, \Vyo. He be- lieved that lower strata of the formation pinch out northward and are not laterally continuous with any units present in central Montana. According to Mundt, the upper cherty dolomite overlaps these lower strata and constitutes all of the Amsden Formation in the Big Snowy area. Gardner (1959, p. 335—337), who did not recognize the regional unconformity Within the Big Snowy dep- SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY B4 .28; aged?“ 95 553:0? mo amzld 598% _I _ ‘ . _ _ $4.2 oom om o _ , mLTTv . ,, IL, |__I m9 _ \II I1 I II. m SWIM ........... I. m, .m \ \ ‘ , m «go ISI _ s, e n, m , In , 57 .I , I IN . mm? \x\\\ r_ MI m. m E 3,; s ,r “x w m a ,, G 7 xxx ,<;S¢ _ n mieemlkoa am=w I_. \\\. 4 xxx“ . r v.1, WELSWS RIJ ,, >9 \\ .0, II ,_ . n 90 .NO \\\u\::_u~:omw\// .IV 0 . H . \ u I I Ixxt \IImuWK \SI ll I I, HI:- Iévmw 4, I. 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Yellowstone Park Bighorn Mountains Black Hills and . . . a) Central Southwestern Quadrant Mtn Amsden Creek Hartville uplift Williston basm Virgil W A :t c 8 .9 L ’E .9 o) > a 3 a g 5 v _—___ s.— w _ _ __ _ ______ .._. C .9 .E > :5, 5 ' ' ‘ ll Missouri % D'V's'on E (lower part) ——i E 2 E -z _____.__ a __ < i _ g Z \ a < t s > Des Quadrant Q d Tensleep 8. middle _, . Moines Formation ua rant 2 mem er >- Formation Quadrant Sandstone 5 Division III 'm Formation E E Z Devils g .3 5 Pocket upper red 2 E (D \- D- 7 Formation " ‘_ shale membeF‘g ‘— u‘? T c, a é % e I W e E E AtOka D. S ‘2 \ S \ 3 LI- _ N 9 middle g 0 Alaska Bench C . E . , , c _ c g limestone : DiViSions g Limestone ,9 w .E m +6 E member 2 IV and V E _, E carbonate ______ < __ __ _ :5? member T lower : C lower ameion red shale - ~ ~ member -% Creek Mbr g Amsden member DiVision VI E Stonehouse -3 shale flqu— _____ __ _ Morrow E E: member ‘ E \ >\ i— \ Z O. 5 3 LL <3 0- B‘ Sn w <7) Chester E‘ lg o y 9 0 Formation m c m (n — no 2 m Kibbey Formation Formation Charles Formation Charles Formation Charles Formation , and Mission (7) and Mission Mission Can - . and Mission . yon Missmn Canyon Pahasapa _ [gaging Canyon Lime- Canyon Um? Limestone Limestone Limestone Canyon 'j'mF' stone (MiSSiS- stone (M'SS'S' (Mississippian) (Mississippian) (Mississippian) stone (M'SS'S‘ sippian) sippian) sippian) FIGURE 3.——Nomenclature and stratigraphic relations of the Upper Mississippian and Pennsylvanian rocks in Montana, Wyoming, South Dakota, and North Dakota. BIG SNOWY AND unit generally consists of limestone or dolomite although it consists locally of sandstone or anhydrite according to Nordquist (1953, p. 81). This medial limestone of the Kibbey is a widespread marker bed that is easily recognized on geophysical well logs. The Kibbey Formation is generally considered to be of Chester age as in some places it rests unconformably on beds of Meramec or older age. However, in other places (pl. 2) the basal part of the Kibbey interfingers with the uppermost part of the Charles Formation, a fact suggesting a late Meramec age. A Chester age is also indicated by fossils from lower strata in the Otter Formation, which are laterally equivalent to the upper part of the Kibbey, and by the gradational contact with the overlying Otter Formation. O‘TTER FORMATION Weed (1892, p. 307) first used the name Otter Creek Shales for exposures along Belt Creek on the north flank of the Little Belt Mountains near the type locality of the Kibbey Formation. He later (1899, p. 2) re- ferred to these rocks as the Otter Shales and included them as part of the Quadrant Formation. The follow- ing year (1900, p. 295) he placed the Otter Shale as the upper formation of the Quadrant Group. The Otter Formation conformably overlies the Kib- bey Formation in central and east-central Montana and west-central North Dakota. In the Big Snowy Moun- tains the Otter is about 375—47 5 feet thick. Eastward, in the central part of the Williston basin, it has a maxi— mum thickness of about 225 feet. Weed’s (1892, p. 307) Otter Creek at the type locality in the Little Belt Mountains is 198 feet thick; however, at this locality the top of the formation is truncated and is overlain disconformably by the Ellis Group (Jurassic), The Otter Formation consists predominantly of greenish-gray shale; locally it also contains gray, pur- ple, and black shale and interbeds of yellowish-gray argillaceous limestone and dolomite and gypsum. The contacts between the Otter, Heath, and Kibbey Forma- tions are commonly gradational and the three forma- tions intertongue (pl. 2, 0—0’). Fauna collected by Weed (1900, p. 295, 296) were assigned to the Carboniferous, apparently the lower Carboniferous. Fauna studied by Scott (1942) and Easton (1962) indicate a marine environment of Ches- ter age. HEATH FORMATION The Heath Formation was named by Scott (1935, p. 1028) for exposures at Beacon Hill near the town of Heath, Mont, on the north flank of the Big Snowy Mountains (fig. 4). The Heath was defined as the upper unit of his (1935, p. 1025) newly defined Big AMSDEN GROUPS IN MONTANA B7 Snowy Group. Easton (1962, p. 14) revised slightly the contact between the Otter and Heath Formations by extending the base of the Heath downward to in- clude the productid-bearing limestones that Scott placed at the top of the Otter. The Heath Formation is restricted in this paper to exclude sandstone and black shale above a regional unconformity in the same way that it is restricted by Beekly (1955) and by Mundt (1956a). The Heath Formation is unconformably overlain by the Tyler Formation or the Ellis Group (Jurassic). The Heath Formation underlies central and east- central Montana and west—central North Dakota. The restricted Heath at the type section is 76 feet thick, and at three other localities in the Big Snowy Mountains it is 551, 270, and 322 feet thick (table 1). To the east, in the Williston basin, it is no more than about 100 feet thick. TABLE 1.——Comparison of thicknesses of Heath Formation and Stonehouse Canyon Member of Tyler Formation [Adapted from measured sections by Easton (1962, p. 117424)] Heath Formation Stonehouse Canyon (restricted) Member of Tyler Formation Location Beds Thickness Beds Thickness (feet) (feet) Alaska Bench ___________ 28-30 76 18—27 288 Potter Creek Dome _____ 13—26 551 6—12 289 Durfee Creek Dome _____ 33—46 270 297—32 138 Stonehouse Canyon _____ 38-66 322 24—37 101 The Heath Formation consists of interbedded dark- gray to black marine petroliferous limestone and shale and, locally, beds of gypsum. In many places the shale is calcareous and silty. Limestone is increasingly abun- dant from central to western Montana. Freeman (1922, p. 826) was the first to consider the Heath as the source bed for petroleum that is found in adjacent strata, particularly in the overlying sandstone reservoirs of the Stonehouse Canyon Member of the Tyler Formation which are prolific oil producers. Scott (1935, p. .1031, 1032) considered the fauna of the Heath Formation closely related to that of the Brazer Limestone of Idaho and the Moorefield F orma- tion of Arkansas and on this basis assigned to the Heath an age no younger than late Chester. Easton (1962, p. 14), on the basis of abundant fossil collections, also assigned the Heath Formation to the Late Mississippian (Chester). Measured section of the type section. of Big Snowy Group and reference section of Tyler Formation [From Easton, 1962, p. 116—118; modified by Maughan, 1967] Located along the prominent cliff forming the west end of Alaska Bench (locally called “Beacon Hill,” see fig. 5) extending B8 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY 47° 109°30’ 00' ' 109“OO’ 10 15MILES I J FIGURE 4.—Map of localities in Big Snowy Mountains and vicinity, Montana. from the NE% SW14 sec. 25, T. 13 N., R. 19 E., to the SW14 sec. 36, T. 13 N., R. 19 E., and thence continued to the NEW; sec. 1, T. 12 N., R. 19 E., Fergus County, Mont. This section can be reached by going 11.5 miles south of the gypsum plant at Heath, taking the road toward the beacon and stopping on the grade 0.2 mile north of the intersection of the gravel road eastward along Alaska Bench and the unimproved dirt road to the beacon. Most of the upper approximately 400 feet of section through the thick sandstone (bed 20) was meas- ured here. The rest of the section can be reached from above by descending the hillside southwestward below the beacon, or from below by going 12 miles south of the gypsum plant at Heath to a point 0.3 mile south of the boundary of Lewis and Clark National Forest. The Madison Group crops out about 100 yards west of the foregoing point and the rest of the section lies along a line of sight northeast toward the beacon. An unimproved road branches off from the next road going eastward (about one- half mile south of the starting point) and leads toward the base of the cliff. This is the type section of the Big Snowy Group and of the Heath Formation as originally proposed by Scott (1935). PENNSYLVANIAN SYSTEM LOWER PART OF AMSDEN GROUP Alaska Bench Limestone (incomplete section; overlying beds eroded) : 1. Limestone, light-gray, sublithographic ; beds 2—6 in. thick grades upward into brownish-gray 2- to 3-ft limestone beds; cliffmaker; loc-s. 13389, 13399-_- 2. Limestone, yellowish-weathering, fine-grained, massive _____________________________________ 3. Shale, green and red; and yellowish, argillaceous, resistant limestone; 100. 13384 ----------------- 4. Limestone, gray to bufi, sublithographic to fine- grained; beds 1-2 ft thick; crystalline, hard; shale, yellow, green, in partings to beds 1 ft thick; unit forms prominent bluff ; cross sections of productid brachiopods present -------------- Feet 2-1 Total Alaska Bench Limestone measured --------- BIG SNOWY AND AMSDEN GROUPS IN MONTANA Tyler Formation: Cameron Creek Member: 5. Shale, purplish-brown and gray-brown, cal- careous, nonresistant ____________________ 6. Limestone, gray, massive, resistant, fine- grained ________________________________ . Shale and shaly limestone, gray, 6-in. beds, hackly weathering and rather nonresistant; loc. 14220 _______________________________ 8. Limestone, gray, fine-grained, vugular, mas- sive; with laminated wavy minute beds--- 9. Shale, maroon, fissile, calcareous; limestone nodules; loc. 14221 ---------------------- 10. Limestone, gray-brown, fine-grained, wavy laminations; forms resistant ledge; 10c. 14223 from base of unit ; 14222 from t0p_-- 11. Shale, maroon or reddish-brown, calcareous, fissile; nonresistant slope former; base of an 8—in. impure nodular limestone lies 1.5 ft from top of unit; 10c. 14226 just beneath the nodular limestone; loc. 14224 from the uppermost shale ------------------------- 12. Limestone, gray, sublithographic, hard, very resistant ------------------------------- 13. Sandstone, yellowish-buff, stained red on surface, very fine grained, calcareous; prominent clift‘maker; 10c. 13383 in middle of bed ---------------------------------- 14. Siltstone, pinkish to purplish, weathering white, shaly, calcareous; platy and wavy beds; surfaces covered with wormlike pur- plish lines and spots --------------------- 15. Shale, greenish-gray, weathering red, fissile; locs. 13382, 13386 ------------------------ 16. Dolomite, light—gray, very fine grained, cal- careous, hard, resistant; sinuous vertical vugs in lower half; grades up into dark- brownish-gray limestone, very fine grained matrix with pebbly limestone grains, very hard, resistant; loc. 13381 in upper half--- 17. Shale; basal 5 ft calcareous, weathering yel- low; upper 3 ft black, fissile; 4-in. lime- stone 3 ft from top; 10c. 13380 in upper 3 ft- -l Total Cameron Creek Member ___________________ Feet 10 15 21 83 Stonehouse Canyon Member: 18. Limestone, light-gray to buff, very fine grained, seminodular, vugular, very hard, resistant 19. Shale and clay, yellow, greenish-gray, red, very weak (exposure dug out) ; loc. 13387-- 20. Limestone, buff with purple mot‘tling, very fine grained, dolomitic, laminated, platy- weathering; clifif—maker ------------------ 21. Shale, black, fissile; weathers to green clay at base and to red clay at top; plant frag- ments near base; loc. D 3121 A near base-- 22. Sandstone, light-brown, fine-grained; 2-in. beds of alternating clean porous beds and shaly beds; subangular grains; slightly re— sistant 23. Shale, black and dark-gray, fissile; loc D 3121 B near base and D 3121 C near top ________ 03 10 28 B9 Tyler Formation—Continued Stonehouse Canyon Member—Continued 24. Sandstone, buff, weathering yellowish-brown, fine-grained, calcareous, porous, cross- bedded, ripple-marked; 2 in. to 2 ft beds; forms prominent bluff; loc. 13388 in 2 in. ironstone band 5 ft below top ------------ 14 25. Shale, black, brown, and gray, fissile, poorly exposed on steep slope; 100. D 3121 D near Feet middle -------------------------------- 108 26. Sandstone, buff to brown, fine—grained, cross- bedded, friable, calcareous, massive, cliflf- making ------------------------------- 48 (The section continues southwestward from the aerial beacon) 27. Sandstone, brownish, conglomeratic; only basal part exposed; estimated thickness (83 ft) is average of Reeves’ (1926, p. 53, 54) and Scott’s 1935b, p. 1024) total pos- sible thicknesses of associated sandy and covered units less thickness of item 26 above ---------------------------------- 36 Total Stonehouse Canyon Member ________________ 288 Total Tyler Formation -------------------------- 371 MISSISSIPPIAN SYsTEM BIG SNOWY GROUP Heath Formation: Feet 28. Limestone, black, argillaceous, hackly-weather- ing, slightly resistant ----------------------- 10 29. Covered; probably black or dark-brown -shale---- 56 30. Limestone, black, argillaceous, hackly-weather- ing ---------------------------------------- 10 Total Health Formation _________________________ . 76 Otter Formation: 31. Covered; probably greenish-gray sh‘ale --------- 83 32. Shale, greenish-gray; interbedded with argilla: ceous limestone; steep slope ------------------ 42 33. Shale, greenish-gray to black; float of gray lime- stone slabs containing Spirorbis near top ------ 28 34. Covered; probably shale ----------------------- 14 35. Limestone conglomerate, platy _________________ 5 36. Covered; probably shale ----------------------- 28 37. Mostly covered, upper part shale, greenish, some silty; with thin beds of oolitic limestone; lime- stone float to (base of unit ____________________ 90 Total Otter Formation __________________________ 290 Kibbey Formation: 38. Covered; greenish-gray soil; light greenish-gray limestone chips in soil at top of unit; grades downward to reddish-brown soil in lower half; silty _______________________________________ 45 39. Sandstone, brown to reddish-brown, fine-grained, platy to flaggy, calcareous, poorly exposed_-_- 30 40. Covered; soil is reddish ----------------------- 115 Total Kibbey Formation ------------------------ 190 Total Big Snowy Group _________________________ 556 B10 REGIONAL RELATIONS The Big Snowy Group, with one exception, rests upon the Mississippian Charles Formation or the equivalent part of the Mission Canyon Limestone where the Charles is not distinguished. The contact seems conformable in eastern Montana, but an erosional unconformity formed in western Montana prior to deposition of the Kibbey. In the Three Forks area, Scott (1935, p. 1026) noted that the Kibbey rests upon an erosional uncon- formity formed on the Madison Group. A karst surface described by Robinson (1963, p. 43) in this same area has local relief of as much as 100 feet along 300 feet of the basal contact of the Kibbey. At a few places else- where in western Montana, Big Snowy rocks rest on markedly thinned Madison. At an isolated exposure in the Pioneer Mountains southwest of Butte, Mont., Big Snowy rocks rest upon the Lodgepole Limestone (G. D. Fraser, oral commun., 1964). These relations indicate the maximum known erosion and removal of part of the Madison Group prior to deposition of the Big Snowy. The Amsden Group overlies the Big Snowy Group throughout central Montana except for a narrow zone immediately adjacent to the northern flanks of the Belt and Big Snowy Mountains. Here pre-Jurassic erosion removed the Amsden, but not all of the Big Snowy, prior to Jurassic deposition (Perry and Sloss, 1943, p. 1292, and pl. 3). It is not possible to determine the original northern extent of the Big Snowy Group, be- cause a short distance farther north of the truncation of the Amsden by pre-Jurassic erosion the Big Snowy Group is truncated also, and the Jurassic Ellis Group rests upon Madison strata (Perry, 1951, p. 57). The Big Snowy Group changes little in lithology northward in central and eastern Montana, although the thickness varies considerably (Willis, 1959, fig. 3, p. 1945) owing to relief of the Late Mississippian to Early Pennsyl- vanian unconformity. Rocks of the Big Snowy Group were deeply eroded or were removed completely in Late Mississippian time in south-central Montana south of Musselshell, Golden Valley, Sweetgrass, and Park Counties (pl. 3). At least 550 feet of interstratified dark-gray shale and limestone of the Heath Formation (restricted) and a total of 1,189 feet of the Big Snowy Group had been deposited in central Montana (beds 13—50 of Potter Creek Dome section, Easton, 1962, p. 118, 119) prior to this period of erosion. No remnants of the Big Snowy Group are known to be preserved in south-central Moh- tana that would indicate how far south these rocks originally may have extended; instead, these rocks are abruptly truncated, and the lithology of the Big Snowy : Hadley, 1960; and Robinson, 1963). SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Group now preserved immediately to the north does not suggest a shore facies (fig. 5 and pl. 1, A—A’). The length of hiatus represented by the Late Missis- sippian to Early Pennsylvanian unconformity is not known, but it probably was relatively short. The youngest rocks of the restricted Heath Formation are of Chester age (Easton, 1962, p. 23) , whereas the oldest rocks above the unconformity are either of very latest Mississippian or earliest Pennsylvanian. Certainly enough time elapsed prior to deposition of the overlying sediments for the entire Big Snowy Group to be stripped from some areas; yet similarity between shale in the Heath and in the overlying strata suggests that there was no significant change in depositional environment, although conglomerate, sandstone, and plant debris in- dicate a change to a shallower, near-shore environment. Additional evidence for this unconformity is presented in the discussion of the Tyler Formation. Correlation of the Big Snowy Group west and south- west of the Big Snowy Mountains becomes increasingly difficult. Thickness and lithologic changes as well as limited exposures of the stratigraphic sequence do not permit correlation with the individual formations of the type Big Snowy Group. Thus, in parts of south- western Montana where these rocks cannot be mapped separately the Big Snowy is of formation rank (pl. 3 and pl. 1, B—B’). The Big Snowy Formation is com- posed mostly of light-weathering dark-gray limestone and dark-gray shale which very closely resemble the limestone and shale of the Heath Formation, and which has been termed the Lombard facies by Blake (1959). In parts of southwestern Montana and western Wyo— ming, the Big Snowy Formation is composed only of strata equivalent to the Kibbey Formation and is a “red bed sequence” similar to the Amsden Formation. Mis- correlations have understandably been made, and some authors have used “Basal Amsden” or “Lower Amsden” for the Big Snowy Formation or its equivalent to indi- cate that the unit is older than the type Amsden. This misuse of Amsden has contributed to confusion in the literature and has brought the meaning of the term “Amsden” nearly into disrepute. Equivalents or partial equivalents of the Big Snowy Group have been mapped or described in parts of south- western Montana where Ki'bbey and Heath lithologies are apparent (Sloss and Moritz, 1951; Gealy, 1953; Me- Mannis, 1955; Scholten and others, 1955; Blake, 1959; Big Snowy sedi— ments have not been recognized east of Boulder Canyon in Sweetgrass County in south-central Montana. In western Wyoming sandstone and red shale in this in- terval (Rubey, 1958) are similar to those in southwest- ern Montana: however, 50 miles west of the area mapped BIG SNOWY AND AMSDEN GROUPS IN MONTANA SOUTH Ellis Group Devils Pocket Formation G / OUD TJ’Ie, F Orr,7 . Alaska Bench Limestone atlon/ _________________ __ ——’—" L .r‘. :3 " .. Cj‘njeFOH Creek Member“ —- _ — ‘ — —. fl' ‘ \yv V - _-_——_——_———— . FEET 750 500 81, 250 atfor, 5 10 15 MILES APPROXIMATE SCALE tone ouSe‘—n '— ‘15 ‘ Big Snowy Group Madison Group FIGURE 5.—Diagrammatic north-south section in central Montana showing relations of Big Snowy and Amsden Groups. (Modified from Mundt, 1956a, p. 1925.) by Rubey, equivalent strata in the Chesterfield Range of southeastern Idaho are cherty limestones of the upper part of the Monroe Canyon Limestone (Du‘tro and Sando, 1963b, p. 1983, 1984). E. T. Ruppel (written commun., 1964) measured and described 850 feet of mudstone, shale, and limestone in the Beaverhead Mountains near Leadore, Idaho, which he tentatively correlated with the Big Snowy Group. We have briefly studied the strata at this locality and concur with Ruppel’s assignment. This is the western- most extent. of this sequence reported to date; and in the vicinity of Challis, Idaho, equivalent strata are carbon- ates in the White Knob Limestone (W. J. Mapel, written commun., 1963). Branson (1937, p. 650) applied the name Sacajawea and “Lower Amsden” to a dominantly carbonate se- quence in the upper part of the Madison in the Wind River Range in western Wyoming. Sacajawea was applied to strata he believed to lie between the Madison and his “Lower Amsden.” He (1937, p. 653) concluded that Sacajawea is Salem to Ste. Genevieve in age, with a closer affinity to the Ste. Genevieve. He correlated this unit with a part of the Brazer in western Wyoming and with a part of the Big‘Sn-owy Group in central Mon— tana. However, his (1937, p. 651) Sacajawea correlates with the upper part of the Mission Canyon Limestone and his overlying “Lower Amsden” may correlate with the Big Snowy Group. Failure to discriminate Saca- jawea from “Lower Amsden” for younger stratigraphic units in a subsequent paper (Branson, 1939) has made the term ambiguous. PENNSYLVANIAN SYSTEM, AMSDEN GROUP The Amsden Group as used herein for central Mon- tana constitutes a change in stratigraphic rank and consists of three formations, each of which we believe has counterparts in the type Amsden Formation in northern Wyoming. The formations, in ascending order, are the Tyler Formation, the Alaska Bench Lime- stone, and the Devils Pocket Formation. A correlation between the type section of Amsden Creek in north— central Wyoming and equivalent strata in the Big Snowy Mountains is illustrated by the detailed strati- graphic section A—A’ on plate 1. We recognized three lithologic units in the Amsden Formation at the type section: a lower red shale member that locally is a sandstone at the base, a medial limestone member, and an upper red shale member that includes interbedded carbonate rock and sandstone. TYLER FORMATION Interstratified dark-gray shale and sandstone above the Late Mississippian to Early Pennsylvanian uncon- formity and red beds that are laterally and vertically gradational into the shale were named the Tyler For- mation. These gray beds, originally included in the Heath, and the red beds have been called collectively the Tyler Formation by many geologists, including Freeman (1922) and Mundt (1956a). (See fig. 1 of this report.) The reference section of the Tyler Formation is des- ignated in this report as the well—exposed rocks at Alaska Bench in secs. 25 and 36, T. 13 N., R. 19 E., de- scribed in beds 5—27 of the section measured by Easton (1962, p. 116—118), and beds 11 probably through 34 of the section as described by Scott (1935, p. 1024). Both the Big Snowy Group and the Heath Formation are well exposed at this location (Scott, oral commun., 1966) ; although, due to inadequate maps, Scott (1935, p. 1024, 1025, 1028) incorrectly placed these exposures in sec. 6, T. 12 N., R. 20 E. Beacon Hill is also the lo— cation of excellent exposures of the strata that has been B12 suggested for the type locality of the overlying Alaska Bench Limestone (Easton, 1962, p. 15). The Tyler Formation is divided into two members, based largely on color and partly on lithology. The lower member, the Stonehouse Canyon Mem- ber—a new name presented in this report—includes those strata composed of predominantly dark-gray rocks. These are beds 18—26 of Easton’s (1962, p. 117) section at Alaska Bench which are here designated as the type section. Stonehouse Canyon is a name long applied to a stratigraphic section of Upper Mississip- pian and Lower Pennsylvanian rocks in the southeast- ern Big Snowy Mountains (fig. 2). The name was derived from Stonehouse Canyon in secs. 29, 31, and 32, T. 11 N., R. 21 E., Golden Valley County, Mont. and this is its reference section. The rocks at this location, described on pages B17, B18, are moderately well ex- posed; they have also been described by Gardner and others (1945; 1946, p. 51—54), Hadley and others (1945), Mundt (1956a, fig. 9, p. 1932), Gardner (1959, p. 338—342), and Easton (1962, p. 121—124). The Cameron Creek Member is a reduction in nomen- clatural rank, but otherwise it has the same contacts as the Cameron Creek Formation of Gardner (1959, p. 347). Willis (1959, p. 1952, 1953) made a similar pro- posal for the Cameron Creek when he made it a mem- ber of the Tyler Formation (fig. 1). The Cameron Creek Member is composed mostly of red beds that gen- erally comprise the upper part of the formation and are beds 5—17 that Easton included in the Cameron Creek Formation. The division of the Tyler into two members empha- sizes that the members are lithologically similar and FEET 100 \\\_\\\\_\\\\\__\\\\\\\\—\\\\_\\\_\\\\\_\\\_\\\\- \ 5° \\\\\\\\\\\\\\\\\\\\\\\\\\\_ _:Ty|er * _Formation-—_-— SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY that the contact between them is a color change only. The boundary between Stonehouse Canyon and Cam- 'eron Creek is difficult or impossible to pick consistently at the same stratigraphic position from place to place owing to the gradation and intertonguing of one into the other. Easton (1962, p. 13) was aware of this affinity when he stated that “part of the Heath forma- tion is a series of lenticular sandstones in black shale which occur near the top * * * and may in part belong to the overlying beds of the Cameron Creek formation.” Strata equivalent to the Stonehouse Canyon Member, but mostly red, have been included in the Amsden For- mation by Vine (1956, p. 424-434) along the northeast side of the Little Belt Mountains. The unconformity at the base of the Tyler Forma- tion is not readily observed in outcrops owing to poor exposures; however, its location can generally be estab- lished by careful tracing of beds in the rocks above and below. Conglomeratic sandstone, less than half a foot thick at the base of the Stonehouse Canyon Member of the Tyler Formation in Stonehouse Canyon, thickens gradu- ally westward and rests on successively older limestone and shale beds of the Heath Formation. In Stonehouse Canyon this conglomerate is 322 feet above the base of the Heath; but about 21/3 miles west at State Road 25, the conglomerate is 10 feet thick and lies about 60 feet above the base of the Heath, a fact indicating a local westward beveling of about 100 feet per mile. Farther north, Norton (1956, p. 58) described an ex- posure of the Heath-Tyler contact located in secs. 8—17, T. 12 N., R. 20 E., about one-fourth mile east of State Highway 25 (fig. 6). Here the coarse-grained sand- 3 Covered 3» mile _L FIGURE 6.—Section of erosional unconformity between Heath Formation (restricted) and Stonehouse Canyon Member of the Tyler Formation exposed at west end of Middle Bench, Little Snowy Mountains, S1/2, sec. 8 to NE1/4, sec. 17, T. 12 N., R. 20 E., Fergus County, Mont. (modified in part from Norton, 1956, p. 61). BIG SNOWY AND AMSDEN GROUPS IN MONTANA stone of the basal part of the Stonehouse Canyon can be seen overlying and abutting against dark limestone and shale of the older Heath (fig. 7). The sandstone con— tains angular blocks of dark limestone that yield pro- ductids, which are common at most fossiliferous Heath horizons. These limestone blocks probably were broken away by wave action as Stonehouse Canyon seas in- vaded valleys carved in the limestone—dark shale unit of the Heath. Differences in lithology between the Heath Formation and the Stonehouse Canyon Member 'of the Tyler Formation, although generally slight, contribute evi- dence for an unconformity between these two units. The limestone and shale in the Heath are thin bedded, very fine grained, and well sorted and thus indicate that deposition occurred in relatively quiet water. The clastic and carbonate beds of the Stonehouse Canyon are thick bedded, lenticular, very fine to coarse grained, and generally poorly sorted and thus indicate more tur- bulent Conditions that must have existed at the time that these rocks were deposited. The differences in transporting energies suggested by these differences in lithology indicate an abrupt change in depositional con— ditions. At most places where there is sandstone at the base of the Stonehouse Canyon it is conglomeratic. Al- though the contact is poorly exposed, it has been mapped by Douglass (1954) as his Amsden-Big Snowy boun- dary in the southwestern part of the Big Snowy Moun- tains; and the same contact has been extensively mapped on the northeast flank of the Little Belt Mountains by Vine (1956). The Heath seems to have been deposited at consider- able distance from shore. The fossils in the Heath suggest that the Heath sea was of normal salinity; the lithology, fabric, and color of the rocks of the Heath indicate that the water was calm or even stagnant. Such conditions are likely to have prevailed in a broad, open sea and fairly deep water, probably many miles from shore. The shore during deposition of the Stonehouse Can- yon presumably was near the beveled and faulted southern edge of the underlying Big Snowy Group or within about 25 miles of the south flank of the Big Snowy Mountains. Deposition of the Stonehouse Can- yon was in locally turbulent water that was probably shallower than the water during deposition of the Heath. The lithology and fabric of the rocks and the fossils in the Stonehouse Canyon indicate a near-shore marine environment for some strata, whereas other strata, notably the coaly or carbonaceous beds, suggest deposition in swampy or brackish lagoonal environ- ments. Several environments, ranging from normal marine through brackish to possibly fresh water, may B 13 have existed simultaneously in adjacent areas as well as in rapid succession at any one place. The rocks also suggest deposition-by currents ranging from moderately strong in some areas to weak or lacking in others. Gardner (1959, p. 344) stated that “the sediments in the Big Snowy basin of deposition indicate a single es- sentially uninterrupted cycle of deposition rather than two cycles interrupted by a widespread erosional un- conformity. The Big Snowy sequence began with sand- stone [Kibbey], continued with thick shale [Otter, Heath, and Cameron Creek], and ended with a predomi- nantly limestone sequence [Alaska Bench] * * *.” A more detailed inspection of these rocks indicates that the Kibbey, Otter, and the restricted Heath in themselves represent such a cycle of deposition and, characteris- tically, are composed respectively of sandstone, shale, and limestone. Similarly the Stonehouse Canyon, Cameron Creek, and Alaska Bench comprise a similar depositional cycle of sandstone, shale, and limestone. The thickness of the restricted Heath Formation var- ies markedly from one exposure to another in the Big Snowy Mountains. Norton (1956, p. 58) suggested that “Many of the various units of the marine Heath formation can be identified and the sandstones of the Tyler * * * may rest unconformably on any of these.” In table 1 thicknesses of the Heath at four localities are compared with each other and also with thicknesses of the sandy beds that lie above the unconformity. Subsurface examination indicates the relief and ex- tent of the Heath-Tyler unconformity. The uncon- formity is present throughout eastern Montana and western North Dakota, and evidence for it is published by Beekly (1955), Mundt (1956a), Foster (1956), and Willis (1959). Four detailed cross sections (pls. 1, 2) show correlation of the Big Snowy and Amsden Groups and the unconformity between them. In addition, a paleogeologic map of the Mississippian rocks below the unconformity (pl. 3) illustrates beveling of the Heath, Otter, Kibbey, and older Mississipian rocks, contrary to Gardner’s (1959, p. 344) contention that these rocks are not regionally beveled. Mappability of the Tyler Formation and its subdivi- sions has been questioned. Scott (1935, p. 1029) noted that the Tyler sands, which had been named by Free- man, “are neither a lithologic, paleontologic, or mappa- ble unit over broad areas * * *.” Gardner (1959, p. 335) repeated this statement and amplified it by saying that the “Tyler is an identifiable lithic unit only in a few relatively small widely spaced areas of central Montana.” Exception is taken to these statements because we have learned from inspection of exposures of these rocks in the Big Snowy Mountains and elsewhere in central and B14 FIGURE 7.—Unconf0rmable contact of the Stonehouse Canyon Member of the Tyler Formation upon restricted Heath Formation, west end of Middle Bench, NEIA, sec. 17, T. 12 N., R. 20E, Fergus County, Montana, a, unconformable contact; b, limestone beds in the Heath Formation; c, conglomeratic sandstone at base of Stonehouse Canyon Member; d, angular blocks of limestone derived from the Heath Formation. southwestern Montana that several significant lithologic differences can be used to distinguish Tyler from Heath as readily as Heath is distinguished from Otter and Otter from Kibbey. The Tyler Formation thickens northward from the Stonehouse Canyon section, and it thins markedly a few miles south in the subsurface. This thinning is in the same area where the underlying Big Snowy Group thins markedly by truncation, but in the Tyler the thinning seems to be depositional rather than erosional. A thin tongue of the formation extends in the subsurface southward from the Big Snowy Mountains and is con- tinuous into the lower red shale member of the Amsden Formation at its type section as illustrated by strati- graphic section A—A’, plate 1. The Tyler equivalent has been identified in the northern Big Horn Mountains as the “lower elastic zone” by Gorman (1963), and is the Darwin Sandstone and Horseshoe Shale Members of the Amsden of Mallory (1967) . The Tyler generally thins from the Big Snowy Moun- tains eastward across Montana to about the Cedar Creek anticline as illustrated by stratigraphic section 0—6", plate 2. It also thickens east of the Cedar Creek anti- cline in the Williston basin in northeastern Montana SHORTER CONTRIBUTIONS ’I‘O GENERAL GEOLOGY and North Dakota. The formation is sharply trun- cated in the VVilliston basin by pre-Jurassie erosion along an arcuate trace approximately parallel with the Missouri River from northeastern Montana into south- central North Dakota. West of the reference section, the Tyler Formation thickens; on East Buffalo Creek in the Big Snowy Mountains (fig. 4) and on the northeast flank of the Little Belt Mountains it is as much as 700 feet thick. In the latter area it was assigned by Vine (1956) to the lower part of his Amsden Formation. Strata equivalent to the Tyler Formation are included as a part of the Amsden Formation undifferentiated in southwest Montana. In general, these strata are thick immediately northwest of a line trending northeasterly from near Monida Pass in southwestern Montana, through the Gravelly Range to about Judith Gap in central Montana (fig. 2). The equivalent rocks are thin and locally absent immediately southeast of this line, evidence suggesting erosion on the upraised side of a Late Mississippian and Early Pennsylvanian fault or fold system. In central western Montana the Tyler equivalents are relatively thin owing to overlap, proba- ble slower rate of deposition, and beveling at the top. Complex structure, intrusion, metamorphism, and weak resistance to erosion contribute to a dearth of informa- tion about the Amsden Formation in this area. Northwestern Montana includes only a single known remnant of rocks of Pennsylvanian age, which is in the Whitefish Range near the town of Trail Creek (fig. 2). These rocks may be equivalent in part to the Tyler For- mation, but the relations of this small outlier as well as known Pennsylvanian rocks farther north in British Columbia and Alberta to the Pennsylvanian rocks in central Montana are unknown. Some speculations have been offered by Halbertsma and Staplin (1960). ALASKA BENCH LIMESTONE The Alaska Bench Limestone forms prominent hog- backs and dip slopes throughout the Little Snowy Mountains and along the flanks of the Big Snowy Mountains. The formation was named for the Alaska Bench, a broad mesa on the north side of the Little Snowy Mountains (fig. 4). Prominent exposures at the west edge of the bench, at a place locally known as Beacon Hill, the‘ type section, have been described (Easton, 1962, p. 15; Mundt, 1956a, p. 1925—1928; Gard- ner, 1959, p. 347). The section at Beacon Hill is in- complete, however; and the section at Stonehouse Can- yon (p. B17 ) is here given as the reference section. The formation is composed mostly of limestone and some dolomite in 1- to 2-foot-thick beds interstratified with red or gray mudstone. Carbonate rock dominates, BIG SNOWY AND AMSDEN GROUPS IN MONTANA especially at the Beacon Hill section and elsewhere where only the lower 100 feet or so of this unit are pre- served; but in a few places, where there are thicker remnants of the Alaska Bench—as at Judith Gap (pl. 1, B~B’)—the upper part is composed of almost equal amounts of carbonate rock and mudstone. Rocks sim- ilar to the Alaska Bench occur throughout most of Mon- tana wherever Pennsylvanian strata are preserved. In eastern Montana this limestone unit has previously been included in the Minnelusa Formation; it also seems to correlate with Division V, and possibly in part with Division IV, of the Hartville Formation and with the equivalent of part of the lower member of the Minnelusa Formation farther south in eastern Wyoming and west- ern South Dakota. Similar rocks are the limestone of the middle member of the Amsden Formation in south- western Montana and Wyoming and are correlated with the Alaska Bench. These are the same as the strata identified as the “quartz deficient zone” by Gorman (1963) in his study of the Amsden Formation in the Bighorn Mountains and the lower part of the Ran- chester Limestone Member of the Amsden of Mallory (1967). The thickness of the Alaska Bench and its equivalents varies considerably throughout this region and at many places within short distances, owing chiefly to a regional erosional unconformity beneath the overlying Devils Pocket Formation. At some places the formation has been completely removed by this erosion, but at other places the Alaska Bench is as much as 143 feet thick, as at Durfee Creek Dome (Easton, 1962, p. 120), and as much as 290 feet thick as at Judith Gap (pl. 1, B—B’). The Alaska Bench Limestone is gradational with the underlying Cameron Creek Member of the Tyler For- mation. Lateral gradation and interbedding of red beds and limestone at the contact seems evident in the subsurface east of exposures in the Big Snowy Moun- tains (pl. 2, 0—0’). It seems likely that similar inter- tonguing, established in the subsurface by detailed marker horizons provided by geophysical logs, may take place in surface exposures west of the Big Snowy Mountains; but this relationship has not yet been estab— lished because most data are from widely separated measured sections at outcrops of limited lateral extent. DEVILS POCKET FORMATION Gardner (1959, p. 347—348) named the Devils Pocket Formation and designated the type section at exposures in Road Canyon in sec. 31, T. 11 N., R. 21 E., about half a mile west 0f the ranch at the mouth of Stonehouse Can- yon. These strata are 143 feet thick at the type section (p. B17) and 80 feet thick at Durfee Creek Dome (Eas- B15 ton, 1962, p. 119) and are not present farther north in central Montana owing to removal by erosion prior to deposition of the Jurassic Ellis Group. On Galloway Creek in the southwestern part of the Big Snowy Moun- tains (fig. 5), Douglass (1954) described 220 feet of strata lying between the Alaska Bench Limestone and quartzitic sandstone of the Quadrant Formation. These strata resemble the Devils Pocket Formation farther east, and most or all of the interval is equivalent to it. The Devils Pocket is partly truncated beneath Jurassic rocks at its type section; but the formation is gradational into the overlying Quadrant Formation on Galloway Creek. Regional correlation, illustrated on plates 1 and 2, indicates that the complete exposures on the southwest flank of the Big Snowy Mountains are the thickest sections of the Devils Pocket Formation. The Devils Pocket is composed chiefly of dolomite, but it includes abundant red mudstone, mostly in the lower part, and sandstone or quartzite, mostly in the upper part. Equivalent strata of similar lithology ex- tend throughout most of eastern and southern Montana. In eastern Montana they previously have been included within the Minnelusa Formation. They are correlated with the lower part of Division III of the Hartville Formation of east-central Wyoming and equivalent of part of the middle member of the Minnelusa Formation of the Black Hills. The Devil’s Pocket is also equiv- alent to Gorman’s (1963) “upper clasitic zone of the Amsden (upper red shale member in this report) and to Mallory’s (1967) Ranchester Limestone Member of the Amsden (pl. 1, section A—A’). An unconformity of regional extent underlies the Devils Pocket and equivalent rocks in the adjacent areas. This unconformity accounts for much of the variation in thickness and in lithologic sequence of Pennsylvanian rocks in this region, and has caused much of the difficulty that has perplexed students of the Amsden in Montana and Wyoming. Reasons for this unconformity given by Mundt (1956a, p. 1931) are. (a) a sharp lithologic break be- tween the two formations, (b) variable thickness of the underlying Alaska Bench Limestone, (c) the overlap of the Alaska Bench by the Devils Pocket at the base of the Devils Pocket that suggests possible formation in south-central Montana, and (d) locally formed red shale of a lateritic soil on top of the Alaska Bench Lime- stone. These reasons are confirmed by J. G. Mompers (oral commun., 1963) and by the correlations shown on plates 1 and 2. In a few places in southern Montana, lower strata of the Amsden are absent and the Devils Pocket rests upon rocks of the Madison Group. One of these places B16 is near the Montana-Wyoming boundary in the north- ern Bighorn Mountains. About 25 feet of red mud— stone (equivalent to the Cameron Creek Member of the Tyler Formation) overlain by about 80 feet of lime- stone (equivalent to the Alaska Bench Limestone) make up the lower and middle members of the Amsden F or- mation along the Little Big Horn River (Agatston, 1954, p. 569). These rocks are not present approxi- mately a mile north of the State line, where only about 60 feet of cherty limestone and dolomite and some sand- stone of the upper half of the upper member of the Amsden (equivalent to the upper part of the Devils Pocket Formation) rests upon Madison. Evidence that red beds were deposited here above the Madison, as in the exposures not far to the south, is indicated by red mudstone in the fissures and channels that compose the karst zone at the top of the Madison. At Storm Castle Mountain (pl. 1, A—A’) in the Gallatin Range, the Devils Pocket also overlaps the truncated Tyler and Alaska Bench. The top of the Devils Pocket Formation at its type section is an erosional unconformity beneath Jurassic rocks of the Ellis Group. At other places where young- er strata of Pennsylvanian age are preserved, the Devils Pocket Formation and the equivalent upper part of the Amsden Formation grade upward into the Quadrant Formation or Tensleep Sandstone. The boundary be- tween these formations is generally arbitrarily placed between the dominantly carbonate sequence and the overlying dominantly quartzite or sandstone sequence. This contact is not everywhere consistently placed stratigraphically, although a relatively uniform thick- ness of the Devils Pocket seems to persist throughout south-central Montana. Also, the Tensleep grades eastward into dominantly dolomite strata which have been included in the middle member of the Minnelusa Formation in eastern Montana. Here, Devils Pocket is expanded to include these dolomitic rocks that are the equivalent of the Tensleep of central Montana and Wyoming. The Devils Pocket is increasingly sandy westward, and in western Montana equivalent strata are not read- ily separated from the overlying sandstone. In the extreme southwestern part of the State, both Devils Pocket and Tensleep equivalents are included in the Quadrant Formation, but in other parts of the State, especially north and east of the Gallatin Range and Yellowstone National Park, the Devils Pocket is in- cluded as part of the Amsden, and the Quadrant is re- stricted to strata equivalent to the Tensleep farther east. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY Reference section of the Amsden Group in Montana [From Easton, 1962, p. 121—124] Located along a line trending east-southeast from the first prominent red outcrops of Kibbey Formation in road cuts along State Road 25 where it ascends into the Big Snowy Mountains, through Stonehouse Canyon; the section extends from the center of the west line, sec. 25, T. 11 N, R. 20 E., across sec. 30 and into SW14 SW14 sec. 29, thence down the dry canyon running southward across N1/2NW174 sec. 31, T. 11 N., R. 21 E., Golden Valley County, Mont. Most of the section is exposed in Stonehouse Canyon. This can be reached by going 0.5 mile west of the intersection of the Red Hill Road (State Road 25) with US. Highway 12 at a point about 1 mile north of Lavina, then going north along State Road 25 toward the Big Snowy Mountains. Bear left and downhill at the first Y-intersection; go 26.0 miles, always taking the better road when a choice is necessary and making some rightaangle turns, to a road crossing; continue north 1.0 mile to a cattleguard. By bearing right at this point, the road leads to the stone ranch house at the mouth of Stonehouse Canyon; entry to the top of the Paleozoic strata is possible through the fence 011 the west side of Ithe creek 150 feet west of the house. By bearing left at the cattleguard, one can drive onto the Otter Formation. in the upper reaches of Stonehouse Canyon; from the cattleguard bear left and continue roughly northward 0.8 mile to a T-intersection ; turn right, crossing fence line just east of intersection; go downhill and swing northward a total of 0.3 mile; bear left on faint vehicle tracks, going 0.6 mile to northwest corner of fenced field; cross fence line, and immediately turn northward upvalley (called Road Canyon in Gardner and others, 1946, p. 49), going 0.3 mile to fence; cross fence line, bear eastward into dry wash and go a total of 0.4 mile upwash to earth dam; continue around west end of dam, swinging eastward upgrade, crossing low summit, and descend the grade into Stone'house Canyon at another earth dam, a total of 0.9 mile. Most of beds 1—102 may be studied east of the second earth dam or south (downstream) from it. The Otter and the Heath Formations are best exposed to the east; younger beds are best exposed to the south. Beds 1-33 were originally published (Gardner and others, 1946, p. 52-54) as part of the Stonehouse Canyon section; beds 34—102 were originally published (Gardner and others, 1946, p. 46) as part of the State Road 25 section. The sections were originally measured by L. S. Gardner, H. D. Hadley, and C. P. Rogers, Jr. The section was sampled again by \V. H. East'on, J. E. Smedley, and Kasetre Phitaksphraivan. JURASSIC SYSTEM ELLIS GROUP Swift Formation : Feet 1. Sandstone, brown. glauconitic, friable, impure, resistant _______--_____-_______-_____. ________ 5 Total Swift Formation measured _________________ 5 Rierdon Formation: 2. Siltstone, dark—red to brown, sh‘aly, calcareous____ 10 3. Covered; red soil; 200 ft to east, red siltstone is in upper 13 ft of unit ___________________________ 16 Total Rierdon Formation measured _______________ 26 Devils 4. oo 10. BIG SNOWY AND AMSDEN GROUPS IN MONTANA PENNSYLVANIAN SYSTELI AM SDEN GROUP Pocket Formation: Sandstone, White to mottled gray and pink, clean, porous, noncalcareous, poorly bedded; made up of medium-grained, well-sorted, clean, quartz sand; locally quartzitic to cherty, but upper 6 ft is friable and nodular. In Road Canyon, about 1 mile to the west, this interval overlain by 14 ft of breccia with Pennsylvanian fusulinids '. Interbedded sandstone, siltstone, and dolomite; sandstone, white, gray, pink, and purple, fine- to medium-grained, mostly calcareous, friable, unresistant; siltstone, red, unresistant; dolo- mite, white, light-brown, and gray with pink tint; some sandy, some with chert nodules; beds as much as 5 in. thick _________________ . Covered; soil is light purple ___________________ . Dolomite, mottled gray and pink, dense to finely crystalline, siliceous, chalky, brittle, poorly bedded _____________________________________ . Covered ______________________________________ 9. Dolomite, White to gray, sandy textured; beds to 1 ft thick ________________________________ Covered; red soil on steep slope ________________ Feet 18 35 Total Devils Pocket Formation __________________ Alaska 11. 12. 13. 14. . Limestone, gray, some pinkish and purplish, dense 16. 17. 18. 19. 20. 21. 143 Bench Limestone: Limestone, light-gray, dense to finely crystalline; 6-in. to 2-ft beds; top of resistant sequence forming dip slopes; 10c. 13423 from upper 2 ft; colln. 44—37—69 ______________________________ Covered ; soil is red and pink, silty and clayey____ Limestone, gray; beds 2—8 in. thick, some fossils interbedded With mottled pink and gray, silty and sandy dolomite in middle 4 ft ____________ Siltstone, red, shaly, calcareous, locally nodular- to finely crystalline sparsely fossiliferous, re- sistant; beds 2 in. to 2 ft thick; base of shaly interval 4 ft thick is 9 ft from base __________ Covered Limestone, gray, some mottled with pink, stylolit— ic; 3-in. to 2-ft beds; fossiliferous ____________ Shale, purplish-red; a few thin beds of siltstone, red, very calcareous, unresistant ______________ Limestone, mottled gray and pink, finely crystal- line; beds 1—6 in. thick; partings of red, cal- careous shale; sparsely fossiliferous; loc. 13422 ______________________________________ Limestone, dark-gray, dense, massive; veinlets of silica on weathered surfaces _________________ Interbedded limstone and shale, mottled gray and pink ___________________________________ w 20 47 Alaska Bench LimestonehContinued 22. Siltstone, mottled gray and purple, argillaceous, very calcareous; probably with some beds actually of impure limestone, very fine grained; beds mostly 1—6 in. thick, some shaly partings; 10c. 13421, colln. 44—37—83 ____________________ B17 Feet 10 Total Alaska Bench Limestone ___________________ 127 Tyler Formation: Cameron Creek Member: 23. 24. 25. 26. 27. 28. 29. 30. 31. 32. 33. Covered; soil, reddish, silty ________________ Limestone conglomerate, gray, dense, hard, tough, massive; vugs with quartz linings__ Covered; soil reddish _____________________ Shale, red, very calcareous, silty, locally sandy; with thin layers of white to green- ish-gray siltstone and sandstone; unre- sistant _________________________________ Limestone, gray, dense, massive; with calcite veins __________________________________ Interbedded shale and siltstone, red, very cal- careous; with nodules of bluish-gray lime- stone; unresistant; 10c. 13420; colln. 44— 37—80 from siltstone _____________________ Covered; soil is red, silty __________________ Shale, red, fissile to brittle, unresistant; with lenses and laminae of pale-yellow siltstone- Conglomerate, dark-red, very calcareous, massive; clasts angular, of red claystone, ferruginous chert, and siltstone, having diameters as much as 3 in.; matrix silt and clay, calcareous _____________________ Shale, mostly reddish, some purplish-gray, fissile, unresistant; lower 16 ft claystone, silty, gray; outcrops dug out; 2 ft of inter- bedded siltstone and shale with limestone nodules 41 ft from base yielded colln. 44—37—94 _______________________________ Sandstone, white, gray, and brownish, fine- to medium-grained, calcareous; top 6 ft mud- cracked and ripple marked; beds 1/2—6 in. thick; bottom 4 ft massive, porous, cross- bedded; loc. 13418 in basal 4 ft; mostly poorly exposed __________________________ 8% 18 66 Total Cameron Creek Member ___________________ 222 Stonehouse Canyon Member: 34. 35. 36. Covered; soil is dark gray ________________ Shale, dark greenish-gray to very dark brown- ish-gray, slightly calcareous, fissile, poorly exposed ________________________________ Mostly covered; some black, fissile shale, poorly exposed; some dark-gray soil that is presumably weathered shale __________ GI 20 75 B18 Tyler Formation—Continued Stonehouse Canyon Member—Continued Feet 37. Conglomerate, mottled yellow and gray, very calcareous, hard; contains subrounded limestone fragments, 2 in. in diameter, in fine-grained sandstone matrix; poorly ex- posed __________________________________ 1 Total Stonehouse Canyon Member ________________ 101 Total Tyler Formation __________________________ 323 Total Amsden Group ____________________________ 593 MIssISSIPPIAN SYSTEM BIG SNOWY GROUP Heath Formation: Feet 38. Mostly covered; shale, black, fissile, fossiliferous ; alternating with Siltstone, brown to medium- gray, calcareous; upper 51 ft gray soil 011 low slope, probably weathered shale; loc. D 3430 A, 50 ft below top, and D 3430 B, 10 ft below top“ 72 39. Siltstone, brown to gray, very calcareous; grading upward into limestone, black, silty to sandy, earthy, porous, poorly exposed ______________ 3 40. Covered ______________________________________ 5 41. Shale, black, fissile, calcareous, poorly exposed; loc. 13417 _________ l __________________________ 3 42. Covered; soil is dark‘gray, probably shale, black, fissile, unresistant ___________________________ 25 43. Siltstone, mottled gray and yellowish-gray, calcar- eous, massive; contains some plant fragments; overlies shale, black, fissile, poorly exposed, 1 ft thick _____________________________________ 3 44. Siltstone or claystone, dark-gray, weathers yellow- ish-gray, vslightly resistant; locs. 13415, 13416, colln. 44—37—115 _____________________________ 1 45. Covered; soil is gray; digging yields gray-buff shaly siltstone and thin conglomerate at base__ 10 46. Limestone, dark-gray to black, dense to finely crystalline, well-bedded, locally laminated; 2~ft layer 3 ft above base is siltstone, gray calcar- eous, shaly, poorly exposed; locs. 13414; 13424; colln. 44—37—117 from upper 14 1ft ______________ 19 47. Shale, black, weathering gray, fissile, calcareous, poorly exposed ______________________________ 1 48. Limestone. dark-gray, hard ____________________ 2 49. Covered; digging yielded sandstone, light-gray, calcareous, platy ____________________________ 9 50. Limestone, dark-gray to black, dense, brittle; beds as thick as 2 ft; loc. 13425, colln. 44—37—123--__ 6 51. Covered; soil is dark gray _____________________ 20 52. Limestone, dark-gray to black, dense __________ 3 53. Covered; soil is dark gray _____________________ 5 54. Dolomite, black to brownish-gray, dense, massive- 3 55. Shale, dark-gray to black, fissile, silty, poorly exposed ___________________________________ 12 56. Limestone, dark-gray to black, dense, massive, very fossiliferous ____________________________ 2 57. Covered; digging yielded shale, dark-gray to black, fissile, silty ___________________________ 6 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY Heath Formation—Continued 58. Limestone, black to dark-gray, very silty, fossil- iferous; platy beds as much as 14 in. thick, inter- bedded with shale, black, carbonaceous; 10c. 13413 ______________________________________ 59. Shale, black, fissile; upper 33 ft poorly exposed" 60. Shale, dark-gray, fissile, gypsiferous, fossiliferous, poorly exposed ______________________________ 61. Siltstone, greenish-gray, calcareous, poorly exposed ____________________________________ 62. Covered; soil is black and contains flakes of black fissile shale ___________________________ 63. Covered; soil is gray __________________________ 64. Shale, black, fissile, noncalcareous; contains nod- ules and thin lenticular layers of limestone, gray to brown, finely crystalline; loc. 13412, colln. 44—37—138_____; _______________________ 65. Limestone, dark-gray to brown, silty, sandy tex- ture, poorly bedded ; interbedded with greenish- gray, silty shale; 100. 13411, colln. 44—37—138--- 66. Shale, dark—gray or black, noncalcareous, locally limonitic, paper-thin beds; interbedded with greenish-gray mudstone or Claystone; plant fragments _________________________________ Feet C! 48 11 17 0; Total Heath Formation _________________________ 322 Otter Formation: 67. Shale, light-green to greenish-gray, fissile; some layers of limestone, silty, nodular; upper half poorly exposed; lower half with selenite on slope ______________________________________ 68. Shale, green to yellow, fissile, unresistant, poorly exposed; contains basal 1-ft bed and other beds at 8- to 12-ft intervals of limestone, gray, dense, brittle, some laminated; contains some shale, gray, calcareous ____________________________ 69. Covered; soil is light-greenish-gray ____________ 70. Limestone, light-greenish-gray, finely crystalline, silty, porous, locally pebbly; ostracodes ______ 71. Claystone and shale, green, silty, very calcareous ; with thin layers of greenish-gray; finely crys- talline limestone ____________________________ 72. Covered; soil is greenish gray __________________ 73. Siltstone, light-yellowish-green, shaly, unresist- ant; upper 1 ft sandstone, yellow to brown, cal- careous, clayey, poorly sorted; contains clasts of fine-grained sand to angular sandstone _____ 74. Limestone, gray to white, medium crystalline; basal 6 ft interbedded with Siltstone, light- green and white, thin-bedded ________________ . Covered; soil is greenish gray _________________ . Limestone, light-gray to light-brown, dense to finely crystalline, dolomitic; weathers to knob— by surfaces ; forms resistant hogback ; algal( ‘3) _ 77. Covered; soil is gray, silty ____________________ 78. Limestone, medium-gray, medium-crystalline, slabby, fairly resistant; colln. 44—37-156 _____ 79. Covered; soil is gray to greenish-gray, silty, with some float of limestone, gray, fine-grained; forms wide valley floor and rounded hogback__ q .1 a: Cl 10 45 37 11 13 ‘) 10 172 81. 82. 83. 86. BIG SNOWY AND AMSDEN GROUPS IN MONTANA Kibbey Formation : 87. 88. 89. 91. 93. 81% 98. 99. 100. 101. Otter Formation—Continued Feet 80. Limestone, gray to white, finely crystalline, mas- sive to poorly bedded, tufialike; contains some sandy lenses; vugs with calcite crystals ______ 4 Siltstone, light—greenish-gray, shaly, poorly ex- posed; interbedded with stringers of black shale; sandy near top ______________________ 9 Limestone, light-greenish-gray, dense; beds 1&— 1 in. thick __________________________________ 4 Claystone, dark-greenish-gray, silty, shaly at top; agate fragments in the soil probably weather from this unit _______________________________ 16 . Limestone, dark-gray to black, weathering gray to white, dense, shaly-bedded ________________ 2 . Shale, yellowish-brown, silty, fissile, poorly ex- posed ______________________________________ 5 Siltstone, greenish-gray, shaly to platy __________ 13 Total Otter Formation _________________________ 374 Sandstone, gray to yellowish and brown, fine- to medium-grained; calcareous in lower half, partly friable, porous; beds as thick as 3 ft__ 23 Covered ____________________________________ 1 Sandstone, yellow to gray with pink mottling, fine- to medium grained, calcareous; consists of poorly sorted quartz sand with minor im— purities; beds 1 in. to 2 ft thick _____________ 16 . Covered ____________________________________ 12 Sandstone, mottled pink and yellowish-gray, calcareous, porous; beds 6—18 in. thick; me- dium-grained, poorly sorted quartz sand _____ 28 . Covered. ____________________________________ 19 Sandstone, greenish-yellow and brown, calcare- ous, medium-grained; consists of poorly sorted quartz sand with minor impurities; beds 1/2—1 ft thick __________________________ 3 . Covered ____________________________________ 6 . Sandstone, yellow or mottled red, gray, and yel- low; very fine- to medium-grained, friable, porous; some grains rounded and frosted; beds 2-12 in. thick _________________________ 19 . Covered ____________________________________ 10 . Sandstone, yellow calcareous, fine—grained; beds 1—8 in. thick; upper three-fourths of unit grades upward from light-brown to pur- ple, calcareous, shaly siltstone into impure, fine-grained sandstone ______________________ 21 Covered _____________________________________ 6 Sandstone, light-yellow to light-brown, fine- grained, calcareous; mostly with black spots that may be dried oil ______________________ 12 Covered ____________________________________ 6 Siltstone, reddish-brown, calcareous, locally shaly, gray-splotched; becomes progressively more sandy upward; top of unit being re- sistant 1-ft bed ____________________________ 38 Total Kibbey Formation ________________________ 220 Total Big Snowy Group ________________________ 885 B19 REGIONAL RELATIONS The nomenclature proposed here for Pennsylvanian units in central Montana is applicable to subsurface equivalents throughout the eastern part of the State. Some of these units have been extended well into the Williston basin in North Dakota by Willis (1959), but such an extension is beyond the scope of this paper. The more comprehensive term, Minnelusa Formation of the Black Hills area, is used here to include all Pennsylvanian and some Permian strata in North Dakota, South Dakota, and adjacent parts of eastern Wyoming. For the present, the State boundary is arbitrarily designated as a convenient place to sepa- rate the different nomenclatures of these two areas. The correlation illustrated in stratigraphic section A—A’, plate 1, shows, as Mundt (1956a, p. 1929) has noted, that some beds in the lower part of the Amsden pinch out northward and are not present in south- central Montana; but it further shows that homologous units which are present in central Montana pinch out southward from exposures in the Big Snowy Mountain. Gardner (1959, p. 345) noted that the Amsden Forma- tion of northern Wyoming and the Big Snowy [and Amsden] of central Montana “represent the same gen- eral interval of geologic time * * * [and] occupy dis— tinct basins * * * separated by a divide toward which the Amsden rocks thin and vanish from the south and the Big Snowy [and Amsden] thin and vanish from the north.” In western Montana and immediately adjacent parts of Wyoming and Idaho, it is recommended that strata approximately equivalent to the Amsden Group con- tinue to be recognized as the Amsden Formation. The Amsden Formation is widespread in southwestern Montana and western Wyoming. In the Bridger Range it is divided by McMannis (1955, p. 1402) into a basal red silty unit (11—189 ft. thick) and an upper carbon- ate unit (113—185 ft. thick) which grades upward into the quartzose sandstone of the Quadrant (50—165 ft. thick). The lithology of his basal unit suggests cor— relation with the Tyler Formation, and the lithology of the upper carbonate unit suggests correlation with the Alaska Bench Limestone and the Devils Pocket Formation. The Quadrant in the Bridger Range is correlative with the Tensleep farther southeast. How- ever, in the Gallatin Range and farther west, the Ams- den Formation is not as inclusive as farther east. In this part of southwestern Montana the Amsden is re- stricted to equivalents of the Tyler Formation and the Alaska Bench Limestone because the Devils Pocket equivalent here is a part of the Quadrant Formation. B20 MISSISLSIPPIAN—PENNSYLVANIAN BOUNDARY The Mississippian-Pennsylvanian boundary is estab- lished in Montana at the regional unconformity between the Big Snowy and Amsden Groups. It is be- lieved that this unconformity was formed nearly con- temporaneously with that unconformity which, by definition (Chamberlin and Salisbury, 1906; Cheney and others, 1945), separates these two systems in the Mississippi Valley. The paleontologic evidence in Montana seems to confirm this accepted position (pl. 4). Early faunal collections from the Quadrant Forma- tion, which included both Big Snowy and Amsden equivalents in Montana, were regarded by Girty (in Calvert, 1909, p. 17—19; Condit, 1919, p. 112, 116; Reeves, 1931, p. 142—143; and Calkins and Emmons, 1915, p. 8) as of Late Mississippian or of Early Penn— sylvanian age. These early age assignments, although Girty made many of them tenuously, are consistent with the stratigraphic relations and age assignments of this report. Faunal collections to which Girty assigned a Late Mississippian age, wherever we have been able to determine, are from the Big Snowy equivalent; and those assigned an Early Pennsylvanian (Pottsville) age are from the Amsden equivalent. Later workers, possibly influenced by the conclusion, now known to be erroneous, that the Amsden in central Wyoming was of Mississippian as well as Pennsyl- vanian age (Branson and Greger, 1918; Branson, 1937; p. B11 of the present report), considered that the Amsden in Montana also included strata of Missis- sippian age (L. L. Sloss, in Gardner and others, 1945, p. 6—8; Scott, 1935). Most of the faunal collections described and evaluated by Sloss were obtained from locations where Big Snowy rocks are present but were not differentiated from the Amsden. Although many species were recognized and indicated a Pennsylvanian age, Sloss based his age assignment of the Amsden on a statistical analysis and regarded the lower part of the Amsden as Mississippian because of the preponderance of Mississippian over Pennsylvanian forms. Scott gave no paleontological evidence to support his conclu- sion that part of the formation was Mississippian. More recently, Easton (1962) has analyzed the fossil evidence from these rocks in central Montana. The Kibbey, Otter, and Heath (as now restricted) were firmly established by him to be of late Mississippian (Chester) age; but he also regarded the Tyler and Alaska Bench as Mississippian (Easton, 1962, p. 25). Data obtained during the present work and re-evalua- tion of Easton’s collection in light of the present cor- relations indicate a Pennsylvanian age for these rocks. Plate 4 is a rearrangement of Easton’s table 5 into strat- igraphic order and separates fossil collections of the SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY restricted Heath from collections of the Stonehouse Canyon Member of the Tyler Formation. This rear- rangement gives greater clarity and definition of those elements of the fauna regarded as typically Pennsyl- vanian. Many brachiopods restricted to the Chester occur in the Heath and Otter formations and terminate at the unconformity beneath the Tyler. According to Mac— kenzie Gordon, Jr. (written commun., 1965) , diagnostic Mississippian (Chester) fossils in these rocks include Diaphragmus fasciculatus (McChesney) [identified as Productus fasciculatuc by Easton]; Oolediwm obes'wm, (Clark) [Stenocisma 066% (Clark) of Easton], Infla- tia obsoleta (Easton) [“Dictyoclostus” inflatus obsole- tus Easton], Eumetria cf. E. vem (Hall) and Girtyella woodworthz' Clark. Invertebrate fossils from the Tyler Formation are not diagnostic for age. The fauna is composed of gene— ra that have little stratigraphic significance. Plant spores from the upper part of the Stonehouse Canyon Member are of definite Early Pennsylvanian age and this age assignment is extended to the base of the Tyler. Petrocmnia chesterensz's (Miller and Gurley [Omnia cf. 0. modesta White and St. John of Easton] was collected high in this member at the section on Potter Creek dome. This fossil is restricted elsewhere to the. Chester, according to Gordon (written commun., 1965), and could indicate a Late Mississippian age for the lower part of the Stonehouse Canyon Member. However, the occurrence of Petrocmm'a chesterensis in association with Chomtes pseudolimtus Easton which may have a range restricted to the Pennsylvanian, as well as a stratigraphic position of these two brachiopods about equal to the Pennsylvanian pollen spores suggests that P. chesterensz's may have ranged later here than else- where and is not restricted to the Chester in central Montana. An Early Pennsylvanian age is attributed to the Tyler Formation by Willis (1959, p. 1959—1962). He listed 11! argz'nifem haydenensis Girty and M. lasalle‘n- sis? from the lower part of the. Tyler in the Lombard Hills near Three Forks, Mont, and Marginifem mum'- catina Dunbar and Condra from near the top of the Stonehouse Canyon Member at Stonehouse Canyon, and noted that this is an unusually low stratigraphic occur- rence for these Pennsylvanian forms. Willis also listed faunal reports from the Tyler Formation in eastern Montana and North Dakota that were prepared by G. O. Raasch and G. A. Stewart, Canadian Stratigraphic Service, Ltd., Calgary, Alberta, Canada. These reports included many ostracodes, especially Cypridopsz's fabu- limz Jones and Kirkby, which are believed to indicate Early Pennsylvanian age. Mackenzie Gordon, Jr. BIG SNOWY AND AMSDEN GROUPS IN MONTANA (written commun., 1966) suggests that M. waltz-{amino cited by Willis from Stonehouse Canyon is the same as “M.” pla/m'costa Easton from the same location, with the notation that “M.” planicosta was described by Eas— ton as similar to M. muricatina. Collections of plant spores from the Heath and Stone- house Canyon Member of the Tyler have helped sig- nificantly to narrow the zone of uncertainty between the Mississippian and Pennsylvanian in Montana. The spores, identified by R. H. Tschudy (written commun., 1963 and 1964), are listed in table 2; their stratigraphic positions are indicated on stratigraphic sections A—A’ and B—B’, plate 1. Spores, specifically Monoletes, from the upper part of the Stonehouse Canyon Member (bed 21 of the sec- tion by Easton, 1962) at Alaska Bench are of Penn- sylvanian age. Tschudy noted that M onoletes has not been reported below the Namurian B horizon and is not represented in Upper Mississippian strata anywhere in the world (Winslow, 1959, p. 62, and fig. 9, p. 101). The overlying Cameron Creek and Alaska Bench, therefore, are of Pennsylvanian age. An Early Pennsylvanian (Morrow) age for the Cameron Creek and an Early to Middle Pennsylvanian (Morrow to early Atoka) age for the Alaska Bench agrees with the fusulinid evidence according to George Verville (oral commun., 1963). The age of the lower part of the Stonehouse Canyon Member is also considered here as Pennsylvanian, but the evidence is less definite. Spores from beds 23 and 25 in the section at Alaska Bench “may be from a Penn- sylvanian horizon not yet examined, or may represent a transitional flora between the Late Mississippian and the Early Pennsylvanian,” according to R. H. Tschudy (written commun., 1963) . A nearly complete sequence of rocks that span the systemic boundary is exposed on Big Sheep Creek in the Tendoy Mountains, Southwestern Montana (col. 20, 8—15”, pl. 1). Samples from this section indicate an in- creasingly diverse flora upward from the Big Snowy Formation of Late Mississippian age through the lower part of the Tyler equivalent into the upper part of the Tyler equivalent (table 2). Tschudy (written commun., 1964) commented that the flora in this stratigraphic in- terval change abruptly and a complete series of produc— tive reference samples embracing this stratigraphic in- terval must be obtained before any reliable age deter- minations can be provided. Nevertheless, correlation of the strata at Alaska Bench and Big Sheep Creek is definitely suggested by the occurrence of several genera of spores common to both sections—especially Mono— Zetes and Didymosporites. Although there are no taxo- nomic groups reported as limited to the Mississippian or B21 Pennsylvanian, the specimens from the'lower part of the Tyler and equivalent strata have a more Penn- sylvanian than Mississippian aspect according to Tschudy. The lower part of the Alaska Bench equivalent in the Amsden Formation at Amsden Creek has yielded fusu- linids of Morrow age (Gorman, 1963) identified as Paramz'llerellct pinguz's, P. circuli, P. ampla, P. advena, P. sp., Millerella infleota, and M. sp. Scott (1954, p. 1195) and Mundt (1956b, p. 50) collected fusulinids, Millerella marblemis Thompson and Millerella adoena Thompson, from the lower part of the Alaska Bench Limestone near Beacon Hill and believed them to indi- cate a Morrow age. Furthermore, Millerella from the Cameron Creek Member, illustrated by Easton (1962, pl. 3), may be considered Pennsylvanian as much as they may be Mississippian forms (B. A. Skipp, oral commun., 1964). Bairdiacypm’s pumtata described by Scott (1935, p. 153) from the upper 2 feet of his Heath (Stonehouse Canyon Member of the Tyler Formation of this report) was believed by him to be a Pennsylvanian form. His age assignment of this ostracode, which also came from the section at Alaska Bench and probably from strata a few feet above those that yielded the spores of Penn- sylvanian age, is confirmed by the new spore data. The distribution of some important guide fossils is treated in more detail below. Choowtes pseudolimtus Easton (colln. 13396) was con- sidered by Easton (1962) to indicate a Mississippian age. The lowest occurrence in central Montana is from the upper part of the Stonehouse Canyon Mem- ber of the Tyler Formation and was not found in older rocks. This collection is from strata about equivalent to, or higher than, those that yielded the flora of Pennsylvanian age. All other collections of 0. pseudolz'mtus are from younger strata; therefore, this species possibly is restricted to the Pennsylvan- ian. Easton’s (1962, p. 23) remark that the surface ornamentation “bears strong resemblance to that of some Pennsylvanian chonetids * * *,” would seem to further strengthen its assignment to the Pennsyl- vanian. Bradyp/Lyllum (colln. 13420) and Myalz’na (Ortho- myalina) sp. (colln. 13421) are from the Cameron Creek Member of the Tyler Formation and lie above the Pennsylvanian age flora. These genera have long been considered Pennsylvanian although Easton had questioned this age when he found them in the same formation with faunas he believed were of Missis- sippian age (Easton, 1962, p. 22). B22 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY TABLE 2.—Pollen and spores recovered from M ississz'ppian and Pennsylvanian samples from Montana (X, present; A abundant; C, common. Localities indicated on sections A—A’ and B—B’ on pl. 1] Collection locality __________________________ Alaska Bench (Beacon Hill) Big Sheep Creek, Tendoy Mountains Stonehouse ' Canyon Galloway Creek Formation and member ____________________ Stonehouse Canyon Member, Tyler Formation Big Snowy Formation Amsden Formation (Tyler equivalent) Heath Formation Otter Formation US GS Locality N o _________________________ D3121D D3121B D31210 D3121A D3429A D3429B D34290 D3429D D3429E D3429F D3430A D3430B D3431 Lycospom _____________________ Granulatisporites _______________ Convolutispora _________________ Auroraspora? __________________ Punctatisporites A ______________ Leiosphaem ___________________ Densospom'tes __________________ Lycospora of. L. granulatipunctatus ____________ Raistrickia__ ___________________ Schulzospom __________________ Calamospora (small) ____________ Laevigatosporites ________________ Apiculatisporites _______________ Verrucosisporites _______________ N. gen. A _ Reticulatisporites A _____________ Punctatisporites B ______________ Granulatisporites B _____________ Endosporites ___________________ Reticulatisporites B _____________ Triletes _______________________ Crassispora, ____________________ Callisporites ___________________ Didymosporites _________________ Endosporites? n. sp _____________ N. gen. B _____________________ Reticulatisporites C _____________ Knozispom'tes __________________ N. gen. C? Granulatisporites (new) _________ Florinites _____________________ Acanthotriletes _________________ Calamospora (rough) ____________ M onoletes _____________________ N. gen. D _____________________ Grandispom ___________________ Reticulatisporites D _____________ NNNNN NNNNN ____________ Linopmductus nodosus (Newberry) n. subsp. are from the Stonehouse Canyon Member of the Tyler Forma- tion or higher strata. L. nodosus reported in collec- tions 13412, and 13413, which are from strata near the top of the Big Snowy Group, cannot be positively identified according to Mackenzie Gordon, Jr. (writ- ten commun., 1965). According to Gordon, those from USGS localities 13414 and 13424 have spine bases scattered randomly over the pedicle valve and are definitely not L. nodosus. In central Montana, as elsewhere, L. nodosus seems to be restricted to the Pennsylvanian. ,, Dicromyocm'nug granularis Easton (colln. 13423) from the Alaska Bench is well above the flora of Early Pennsylvanian age. The collection is from strata that may be as young as Middle Pennsylvanian (early Atoka). An Early Pennsylvanian age is consistent with the species’ stratigraphic position and its form is “* * * more advanced than Chester species and more primitive than the Moscovian (Middle Penn- sylvanian) genotype” (Easton, 1962, p. 23). Easton (p. 37) preferred a Chester age for this crinoid based upon morphologic evidence, but he quoted from Edwin Kirk (written commun, 1954) that “It is ob- viously much later in the genetic line than the Chester forms and I would be inclined to place it as of at least Morrow age.” “Margim'fem” planocosta Easton (colln. 13420 and 14221) are from Cameron Creek and are above the flora of Early Pennsylvanian age. Easton (1962, p. 23) regarded this as a proemial Pennsylvanian spe- cies. Its stratigraphic position, as well as its very close resemblance to “Marginifem” mmficatim Dun- bar and Condra of Pennsylvanian age, suggests that “M” planoeosm is truly of Pennsylvanian age rather than proeinial. BIG SNOWY AND AMSDEN GROUPS IN MONTANA Neospim'fer pmemmtius Easton (colln. 13369, 13370, and 13409) are all from the restricted Heath. With the exception of a possible similar form in collection 13364 from the Alaska Bench, this species may be re— stricted to the Chester and possibly may serve as a guide to the Heath Formation. This stratigraphic position confirms Easton’s evaluation that this species “is more primitive than the earliest Pennsylvanian Neospz'rifer and is closely allied with forms referred to Spirifer, with which it is associated in strata of Mississippian age.” ‘ Composite], subquadmta (Hall) (colln. 13404 and 13423) and Orbiculoédea wyomingensz’s Branson and Greger (colln. 13399) are in strata of the Alaska Bench Lime- stone well above the Early Pennsylvanian age flora. This occurrence of 0. subquadmta suggests that this fossil, which is an established guide to the Mississip— pian elsewhere, extends well into Early and Middle( ?) Pennsylvanian (Morrow and possibly as late as early Atoka) in central Montana, and has not been reported from Mississippian rocks here. 0. wyomz‘ngensz's may be restricted to the Early Penn- sylvanian or may range longer than previously sup- posed. The Devils Pocket Formation has yielded fusulinids of Middle Pennsylvanian age (Easton, 1962, p. 16—17; Henbest, 1954, p. 50, 51). Correlative strata in the Amsden Formation in other parts of southern Montana and adjacent areas have yielded fusulinids of similar age. These fusulinids have been reported from the Pryor and northern Bighorn Mountains and are sum- marized by Henbest (1954, p. 50, 51). The Devils Pocket and equivalent Amsden strata, on the basis of the occurrence of Profusulinella, were assigned an Atoka age; but Henbest (1956) has suggested that these forms of Profusulz'nella from northern Wyoming and Mon- tana may be as young as early Des Moines age. TECTONIC FRAMEWORK The following summary of late Paleozoic tectonic events places deposition of the Big SnOWy and Amsden Groups into an historical framework of depositional and erosional sequences. These tectonic events are de— duced from the regional distribution of the rocks, their lithologic character, and the presence or absence in some areas of rocks known elsewhere. The summary pre- sented here is based partly 011 evidence presented in the foregoing sections of this paper and partly on addi- tional evidence, deductions, and conclusions that are a part of paleotectonic research not yet completed on the Mississippian and Pennsylvanian Systems of the north- ern Rocky Mountains. B23 EARLY AND LATE MISSISSIPPIAN (MADISON GROUP) Widespread uniformity in thickness and lithology of formations in the Madison Group indicates that the northern Rocky Mountains and adjacent plains area was a stable continental shelf or a moderately sub- merged segment of the craton during Kinderhook, Os- age, and Meramec time. Progressive southward over— lap of Lower Mississippian strata indicates the first gentle uplift in south-central Wyoming and farther south in Colorado of the ancestral Rocky Mountain ranges (Maughan, 1963, p. C26). A similar regional uplift, herein named the Milk River uplift, may have formed at this time north of Montana. This uplift is mostly inferred from the formation later in the Pale- ozoic and early Mesozoic of a definite positive area cen- tered in Alberta and Saskatchewan. Uplifts in Alberta and Saskatchewan may have served in part to inter- mittently restrict the Madison sea southeastward in the Williston basin. Evaporites and dolomites—the Charles Formation which intertongues southwest of the Williston basin into the upper part of the Mission Can- yon that is of Meramec age—formed east of a vast ex- panse of limestone and dolomite deposits in Montana and Wyoming that lie between the Milk River uplift on the north and the ancestral Rocky Mountains on the south. Filling of the shallow seaway between the two uplifts with carbonate sediments probably served to further restrict the sea in the Williston basin from marine water of normal salinity in the Cordilleran geosynclinal sea westward in Idaho. LATE MISSISSIPPIAN (BIG SNOWY GROUP) Regional tectonic movements in early Chester time are believed to have initiated the change from carbonate and evaporite deposits characteristic of the upper part of the Madison Group to the detrital deposits charac- teristic of the lower part of the Big Snowy Group. These tectonic movements extended from the Cordil- leran geosynclinal province eastward into western Montana; parts of western Montana were uplifted and during Chester and Early Pennsylvanian time the Mad- ison Group was subjected to erosion, leading to forma- tion of a karst topography. Central and'eastern Mon— tana seem to have remained stable except for probable slight epeirogenic uplift. Consequently, a sea more shallow than the Madison sea extended across this part of Montana into North and South Dakota at this time. The gradation in the Big Snowy Group from domi- nantly sandstone in the lower part through green and red shale into black shale and limestone in the upper part suggests the gradual deepening of the sea as the region slowly subsided. B24 Regional uplift that began Big Snowy deposition not only reduced the depth of the sea, but also uplifted the bordering land areas higher and produced acceler- ated erosion. The increased detritus is evident in the Kibbey Formation; but the location of bordering lands that served as source areas for these sediments is con- jectural. Only in central North Dakota are the rocks suggestive of deposition proximal to a shore. The Kib- bey here, as elsewhere, is beveled and it is not certain how far east the sea may have extended at this Late Mississippian time. Nevertheless, the continental shield, or Siouxia land area, probably was not far east of this beveled edge in central or eastern North Dakota. Earlier in Mississippian time the shore probably was much farther to the east. In central Montana there is no evidence in the Big Snowy Group to suggest the position of the shore, al- though the Milk River uplift to the north and the ancestral Front Range to the south may have also con- tributed sediments. The Big Snowy Group is more ex- tensively beveled in this area, and the facies of these rocks are laterally very uniform throughout central Montana (pls. 1, 2). If epeirogeny enlarged the land area and decreased the area of deposition, as suggested above, it may be presumed that the incipient ancestral Front Range extended into central, or possibly north: ern, Wyoming. Similarly, the presumed Milk River uplift may have extended into northerii Montana. These assumed land areas and their shorelines were well away from the present limits of the Big Snowy rocks preserved in central Montana (pl. 3) . EARLY AND MIDDLE PENNSYLVANIAN (TYLER FORMATION AND ALASKA BENCH LIMESTONE) Regional upwar , centered about the ancestral Rocky Mountains in Colorado, took place again near the end of Mississippian time. Uplift followed Late Missis— sippian deposition of approximately 1,200 feet of Big Snowy rocks in central Montana. Much of Wyoming and south-central Montana was gently elevated, and Mississippian rocks that had been recently deposited in this area were partlyeroded. The chief area of uplift extended across souther11,Montana and was bounded on the north by an arcuate system of probable faulting in the west and probable monoclinal folding in the east (pl. 3). Big Snowy rocks were com— pletely stripped from the uplifted area south of this structural belt and were tilted northward, beveled, and partly preserved north of the structural belt in a north- ward-thickening wedge (pls. 1, 2). The Milk River uplift either was inactive or its influence at this time was too weak to affect rocks in central Montana. Some beveling of Big Snowy rocks beneath the Tyler Forma— SI-IORTER CONTRIBUTIONS TO GENERAL GEOLOGY tion in northeastern Montana could reflect the influence of the Milk River uplift in this area; but instead, this beveling probably is a local feature marking structural movement along the axis of the Cedar Creek anticline. _ Movement along the Cedar Creek structure, either by faulting or by sharp monoclinal folding, was downward on the east, the Williston basin side, relative to the west side. , The trough—shaped depression in central Montana known as the Big Snowy or Montana trough, formed as a broad seaway during Meramec time, then probably narrowed somewhat during Chester time. The con— figuration of the Montana trough, however, has been em- phasized by subsequent faulting and erosion. The pres- ent zero-edges of the Big Snowy Group were formed by erosion on both the north and the south and are not depositional edges. The present southern limit of these rocks was formed by erosional truncation shortly after their deposition. The northern limit of preservation was defined by a probable system of faults (pl. 3) that formed at a later time and is discussed below. Tectonic stability probably persisted while deposi— tional basins filled with sediments during Morrow and A‘toka time. Sediments that formed the Tyler Forma- tion accumulated in the depositional basin in central Montana north of the extended ancestral Rocky Moun- tain uplift. Gradual regional submergewce followed the initial flood of sediments that had mostly filled the deeper basins. The sea spread southward and soon inundated southern Montana and most of Wyoming. At first, detrital sediments were deposited in this wide, shallow sea as red beds composing the Cameron Creek Member of the Tyler Formation. The source area of these sediments ceased to exist as the adjacent lowlands were inundated further, and deposition of red beds gave way to deposition of the Alaska Bench Limestone. MIDDLE AND LATE PENNSYLVANIAN (DEVILS POCKET, QUADRANT, AND TENSLEEI’ FORMATIONS) Upward epeirogeny and regional erosion is indicated by the unconformity formed after deposition of the Alaska Bench Limestone. Localized warping and probable faulting took place extensively throughout this region during this epeirogeny, These local tectonic features are indicated by the general beveling and local removal of the recently former Lower and lower Middle Pennsylvanian rocks as shown in the cross sections on plates 1 and 2. This regional uplift in Montana coin— cided with the chief uplift in Colorado and southern Wyoming of the Ancestral Front Range in the late Atoka to early Des Moines time. Orogenic uplift also seems to have taken place at this time in the Cordilleran BIG SNOWY AND AMSDEN GROUPS IN MONTANA geosyncline, and the Milk RiVer uplift presumably rose importantly for the first time. During late Atoka, Des Moines, and possibly Mis- souri and Virgil time, older Paleozoic rocks were prob- ably deeply eroded in areas of Middle Pennsylvanian uplift in parts of Idaho, western Montana, Alberta, and Saskatchewan. This erosion presumably included Or- dovician sandstone that seems to be the most likely source rock for sand in the Quadrant and Tensleep. The Ordovician sandstone may have been deposited originally as a continuous sheet throughout most of the northern Rocky Mountain region, but it is preserved only in remnants such as the Swan Peak Quartzite in northern Utah and southeast Idaho, the Kinnikinic Quartzite in central Idaho, the Mount Wilson Quartz- ite in western Alberta, the Winnipeg Sandstone in southeastern Saskatchewan and adjacent areas, and the St. Peter Sandstone in southeastern Minnesota and ad- jacent areas. The principal source area of the Middle and Upper Pennsylvanian sands likely was northwest of present Quadrant exposures in western Montana. Another source area may have been a large area north of Montana in northern Alberta and Saskatchewan from which Ordovician sandstone also could have been eroded, transported southward, and deposited in Mon- tana. Erosion of similar Ordovician rocks southwest- ward in west-central Utah probably contributed also to the Quadrant, Tensleep, and equivalent strata in the Wells, Oquirrh, and Weber Formations (Frank G. Armstrong, oral comm-um, 1965). PERMIAN T0 JURASSIC EVENTS From late in Pennsylvanian time to early in Permian time the Milk River uplift rose further, and the area of uplift extended into central Montana and northern Wyoming. Renewed uplift of the ancestral Front Range highland took place about this time also, and a lowland through central Wyoming separated this high- land on the south from the Milk River uplift on the north. A basin of deposition, centered near the present- day Black Hills, formed east of these land areas and received Early Permian (Wolfcamp) sediments. Sands deposited in this basin were derived mostly from erosion of the Tensleep and equivalent Pennsylvanian rocks of these uplifts. V The Big Snowy and Amsden Groups are abruptly truncated in central Montana north of an arcuate system of probable faults, as shown on plate 3. This probable fault system may have formed as early as Late Penn- sylvanian coincident with Late Pennsylvanian to Early Permian uplift. Erosion north of this fault system at this time, especially of the Kibbey Formation, and dis- tribution of these sediments southeastward probably B25 accounts for scattered grains of medium- and coarse- grained frosted quartz characteristic of Lower Permian rocks throughout eastern Wyoming and the Dakotas (Maughan, 1967). The fault system defines the southern edge of a circu- lar uplift, the Milk River uplift, with an apparent cen- ter in southern Alberta or Saskatchewan. The Milk River (fig. 2) approximately bisects this uplift. The river rises in Glacier National Park, Mont., near the projected western extension of the fault system, it flows eastward across southern Alberta and northern Mon- tana, and it empties into the Missouri River near Fort Peck, Mont., about where the river intersects the eastern part of the arcuate fault system. The term “ancestral Sweetgrass arch” has been used to identify this positive area, but is not used here because there is little or no relationship between the Milk River uplift, a late Paleo— zoic and early Mesozoic structural feature and the Sweetgrass arch, a Laramide structure. Furthermore, the two features are centered in widely separated areas. The Sweetgrass arch is comparatively much smaller and is coincident with only a relatively small part of the southwestern part of the Milk River uplift. Webb (1951, p. 3) included the area of the Milk River uplift as a part of his Alberta shelf. The Permian and early Mesozoic history of the Milk River uplift, identified as a land area centered in Mon- tana, has been summarized by Maughan (1967). Upper Permian and Lower Triassic rocks record a gradual transgressive overlap of sediments that were deposited progressively farther north on the margins of the Milk River uplift. Renewed uplift probably took place in Middle Triassic time, and erosion of Big Snowy and Amsden rocks continued in northern Montana until the Middle Jurassic. By this time the area of the Milk River uplift was stabilized and the Middle Jurassic sea transgressed across Montana from the north. Only “Belt Island” obstructed complete regional submergence and this feature was buried by later Jurassic sediments. REFERENCES CITED Agatston, R. S., 1954, Pennsylvanian and lower Permian of northern and eastern Wyoming: Am, Assoc. Petroleum Geologists Bull., v. 38, no. 4, p. 508—583. Beekly, E. 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Jr., 1945, Mesozoic and Paleozoic in the mountains of south-central Montana: U.S. Geol. Survey Oil and Gas Inv. Prelim. Chart 18. 1946, Stratigraphic sections of Upper Paleozoic and Mesozoic rocks in south-central Montana: Montana Bur. Mines and Geology Mem. 24, 100 p. Gealy, W. J., 1953, Geology of the Antone Peak quadrangle, southwestern Montana: Harvard Univ. unpub. Ph. D. Thesis. Gorman, D. R., 1963, Stratigraphicsedimentational aspects of the Amsden Formation, Big Horn Mountains, Wyoming, in Wyoming Geol. Assoc. and Billings Geol. Soc. First Joint Field Conf. Guidebook: p. 67—70. Hadley, H. D., Gardner, L. S., and Rogers, C. 1’., Jr., 1945, Graphic sections of Mesozoic and Paleozoic rocks that un- derlie the basin areas in south-central Montana: U.S. Geol. Survey Oil and Gas Inv. Prelim. Chart 19. Hadley, J. B., 1960, Geology of the northern part of the Gravelly Range, Madison County, Montana, in Billings Geol. Soc. Guidebook, 11th Ann. Field Conf., West Yellowstone earth- quake area, 1960, p. 149—153. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Halbertsma, H. L., and Staplin, F. L., 1960, The Mississippian- Pennsylvanian boundary from the Peace River area to the Williston basin: Alberta Soc. Petroleum Geologists Jour., v. 8, no. 12, p. 363—373. Henbest, L. G., 1954, Pennsylvanian Foraminifera in Amsden formation and Tensleep sandstone, Montana and Wyoming, in, Billings Geol. Soc. Guidebook, 5th Ann. Field Conf., 1954, p. 50—53. 1956, Foraminifera and correlation of the Tensleep sand- stone of Pennsylvanian age in Wyoming, in Wyoming Geol. Assoc. Guidebook, 11th Ann. Field Conf.: p. 58—63. Iddings, J. P., and Weed, W. H., 1899, Descriptive geology of the Gallatin Mountains, in Geology of the Yellowstone Na- tional Park: U.S. Geol. Survey Mon. 32, pt. 2, p. 1—59. Mallory, W. W., 1967, Pennsylvanian and associated rocks in Wyoming: U.S. Geol. Survey Prof. Paper 554—G. (In press.) Maughan, E. K., 1963, Mississippian rocks in the Laramie Range, Wyoming, and adjacent areas, in Short papers in geology and hydrology: U.S. Geol. Survey Prof. Paper 475-0, p. 023—027. 1967, Eastern Wyoming, eastern Montana, and the Dakotas, in McKee, E. D., Oriel, S. S., and others, Paleo- tectonic investigations of the Permian System in the United States: U.S. Geol. Survey Prof. Paper 515—G. (In press.) McMannis, W. J ., 1955, Geology of the Bridger Range, Montana : Geol. Soc. America Bull., v. 66, no. 11, p. 1385-1430. Mundt, P. A., 1956a, Heath-Amsden strata in central Montana: Am. Assoc. Petroleum Geologists Bull., v. 40, no. 8, p. 1915— 1934. 1956b, The Tyler and Alaska Bench formations [Mon- tana], in Billings Geol. Soc. Guidebook, 7th Ann. Field Conf., 1956, p. 46—51. Nordquist, J. W., 1953, Mississippian stratigraphy of northern Montana in Billings Geol. Soc. Guidebook, 4th Ann. Field Conf., 1953, p. 68—82. Norton, G. H., 1956, Evidences of unconformity in rocks of Car- boniferous age in central Montana, in Billings Geol. Soc. Guidebook 7th Ann. Field Conf., 1956, p. 52—66. Peale, A. C., 1893, The Paleozoic section in the vicinity of Three Forks, Mont: U.S. Geol. Survey Bull. 110, 56 p. Perry, E. S., 1951, The northward termination of the Big Snowy and Amsden strata in central Montana, in Billings Geol. Soc. Guidebook, 2d Ann. Field Conf., 1951, p. 55—57. Perry, E. S., and Sloss, L. L., 1943, Big Snowy Group, lithology and correlation in the northern Great Plains: Am. Assoc. Petroleum Geologists Bull., v. 27, no. 10, p. 1287—1304. Reeve-s, Frank, 1931, Geology of the Big Snowy Mountains, Mon- tana: U.S. Geol. Survey Prof. Paper 165—E, p. 135—149. Robinson, G. D., 1963, Geology of the Three Forks quadrangle, Montana: U.S. Geol. Survey Prof. Paper 370, 143 p. Rubey, ‘V. W., 1958, Geology of the Bedford quadrangle Wyom— ing: U.S. Geol. Survey Geol. Quad. Map GQ—109. Scott, H. W., 1935, Some Carboniferous stratigraphy in Montana and northwestern Wyoming: Jour. Geology, v. 43, no. 8, pt. 2, p. 1011—1032. 1942, Ostracodes from the upper Mississippian of Mon- tana: Jour. Paleontology, v. 16, no. 2, p. 152—163. 1945, Age of the Amsden Formation [lVyoming-Mon- tana] [abs]: Geol. Soc. America Bull., v. 56, no. 12, pt. 2, p. 1195. 1950, (Mississippian) [Mon- The Big Snowy Group BIG SNOWY AND AMSDEN GROUPS IN MONTANA tana], in Billings Geol. Soc. Guidebook, 1st Ann. Field Conf., 1950, p. 47—48. Scholten, Robert, Keenmon, K. A., and Kupsch, W. 0., 1955, Geology of the Lima region, southwestern Montana and adjacent Idaho: Geol. Soc. America Bull., v. 66, no. 4, p. 345—403. Seager, 0. A., 1942, Test on Cedar Creek anticline, southeastern Montana: Am. Assoc. Petroleum Geologists Bull., v. 26, no. 5, p. 861-864. Sloss, L. L., 1952, Introduction to the Mississippian of the Willis- ton Basin, in Billings Geol. Soc. Guidebook, 3d Ann. Field Conf., 1952, p. 65—69. Sloss, L. L., and Moritz, C. A., 1951, Paleozoic stratigraphy of southwestern Montana: Am. Assoc. Petroleum Geologists Bull., v. 35, no. 10, p. 2135—2169. Strickland, J. W. 1956, Mississippian stratigraphy, western Wyoming, in Wyoming Geol. Assoc. Guidebook, 11th Ann. Field Cont: p. 51—57. Vine, J. D., 1956, Geology of the Stanford-Robson area, Central Montana: U.S. Geol. Survey Bull. 1027-J, p. 405—470. Walton, P. T., 1946, Ellis, Amsden, and Big Snowy group, Judith Basin, Montana: Am. Assoc. Petroleum Geologists Bull., v. 30, no. 8, p. 1294—1305. Webb, J. B., 1951, Geological history of plains of western Can- ada: Am. Assoc. Petroleum Geologists Bull., v. 35, no. 11, p. B27 2291—2315; revised, 1954, in Clark, L. M., ed., Western Can- ada sedimentary basin—a symposium: Am. Assoc. Petro- leum Geologists, Ralph Leslie Rutherford Memorial Volume, p. 3—28. Weed, W. H., 1892, Two Montana coal fields: Geol. Soc. America Bull., v. 3, p. 301—330. 1896, Yellowstone National Park; sedimentary rocks: U.S. Geol. Survey Geol. Atlas, Folio 30. 1899, Description of the Fort Benton quadrangle [Mon- tana]: U.S. Geol. Survey Geol. Atlas, Folio 55. 1900, Geology of the Little Belt Mountains, Montana: U.S. Geol. Survey 20th Ann. Rept. 1898—99, pt.- 3, p. 257- 461. Willis, R. P., 1959, Upper Mississippian-Lower Pennsylvanian stratigraphy of central Montana and Williston Basin [North Dakota] : Am. Assoc. Petroleum Geologists Bull., v. 43, no. 8, p. 1940—1966. Winslow, M. R., 1959, Upper Mississippian and Pennsylvanian megaspores and other plant microfossils from Illinois: Illi~ nois Geol. Survey Bull. 86, 135 p. Zieglar, D. L., 1955, Pre-Piper post-Minnekahta “red beds” in the Williston basin [summary], in North Dakota Geol. 'Soc. Guidebook, 3d Ann. Field Cont, p. 49—55; revised and en- larged, 1956, North Dakota Geol. Soc., Williston Basin Symposium 1st Internat., Bismarck, N. Dak., p. 170-178. U.S. GOVERNMENT PRINTING OFFICE : I961 O—ZIJ-IIS UNITED STATES DEPARTMENT OF THE INTERIOR PROFESSIONAL PAPER 554—B GEOLOGICAL SURVEY PLATE 1 / 4* 7 7 13* 17M 18* A Petroleumcgrvoldficiggsl lltilefming Co. Shell Oil Co. L. S. Youngblood Amerada and British 19* A 3* sec. 27, T. 6 8., R. 32 E. 5* 7* 8* 9* 10* 11* 12* Sl‘leiérthgrn gaIIfICRZIIIE Slh73nTnolré Cross :2 American Oil Companies AmeIzgfefegggileT C0- Tidewater Oil Co. Farmers Union 6* Lion Oil Co. GU” OII COTP- Cities Service Aries 0“ CO~ . . P | C Ashland Oil-Zach Brooks I Y i ., I I 16 36C. ' I I, R. E. 25 TBulrke 1 sec. 17' T' 18 N" R' 24 E- 1 Crow Tribal l-A Tribal 1 . Hereford 1 Mackey Ranch 1 Cleveland 1 Spidel 1 PhI'IIDS W0 eum 0- Northern Pacific Railroad 1 Potter Creek Dome E1800 seC- , - 7 N, R. 24 E- Amsden Creek sec. 29,1. 7 3., R. 33 E. sec, 34, 1.33., R. 31 E. Carter07il 00- sec. 1,r.1 M1285, sec. 27, M n, R. 27 e. sec 11 1 4 N R 25 E sec. 12,15 N., R. 23 E. Meme" 1 sec 29 r 8 N R 21 E 14 secs. 5, 8,T. 13 N., R. 21 E. . _ sec. 32' T. 57 N., R. 37 w‘ 7 772900 77 y 7 CTOW Tribal 1 7 _ ' ’ ‘ " ' ' sec. 27. T- 7 N., R- 22 E- ' ’ ' " ' ' Stonehouse Canyon 15 Modified from Easton 7777777-----.WMVere/ssr’” 1 ’ Tensleep thickness from 2 - ‘ " ‘ sec. 17, T. 2 8., R. 29 E. WW ’ 6 - 7 7 7 7-7777 . A2000 sec. 25' T. 11 N., R. 20 E. and Alaska Bench (1962, p. 118419) -§ T7 ? z Agatston (1954, p.569) Lime Bighom Rive, ' ~ - . / ' < 7 secs.29—31,T. 11 N., R. 21 E, secs-7251361T- 13 “-1 R- 19 E1 4400 g 0 E5 Goose Egg Amsden thickness and lithology from sec. 19(7), T. 58 N., R. 39 w, From Easton (1962, p. 121—124) and sec'. '11 T' 12 N.,R.19 E' _ a Piper Form7ati7o7n7 7 3 <3 3 5 Formation P. A. Mundt (written commun., 1955) and Agatston (1954’ p. 569) \\~\.- Modlfled from Easton 7 7777A W W W iii, A . 77¥**’ ”T PM (3 (I) < '5: W. W. Mallory (written commun., 1963) (1962’ p. 116—118) § g g 7 _ AAA-WEE-- -- 'i L i / e .2. ,5 It 9 r: Tensleep i I 5 o n__ Sandstone ,/ g 8 l V) — Z S i 77:5 ‘6 8 S7,". --- - W :l 2‘ g 5 ‘3 ‘ 1 I 3 8 upper red E r. shale member z E E: . Z L .57 31’ LE middle limestone I" g member 13 __ (I) E lower red < shale member 3 E E Madison LINE A—A', FROM AMSDEN CREEK,WYO.,TO NEAR LEWISTOWNMONT. 5 Group 1 7 w . / Q 1 ~ / 2 , z \ a . // 7/]. \ i, \ "034305/ 1’ / 116° 114° 112° 110° 108° 106° 104° \ 1 . / / 1 T l l Alv I l l l l I l ‘ ‘ ‘ ' ,- / // BRITISH COLUMBIA ‘ ._ ALBERTA i SASKATCHEWAN D ( , 7' » / // . 1__.__ CAN—ADA._— -_"_ ----- -——I ————— :TTIT; 53—_—‘—TF-——_ 77 ,\ \\ l // ‘t UNITED STATES .7 7 . 7 . 7 . 7 i- _. 1‘ .. METOEE»: FEOET \\\ \\ Rx ’ .0 . g. . ‘. ' :ZV ' _E \\ \\ \ , 5f ‘10 EXPLANATION _; ( ( 7 o '57 : \i, I, :2 . \ 55 ‘H .« a; 100 I k g; ‘ ,l _ 77J'l/J7n 7 \\ \ 777 j 777 .7. .- . 1\ I1 t- .7 55 ° ,‘,37 Area where upper Paleozoic rocks 50“ \ I, € - . I” probably dep051ted but removed -_ 200 k \\ l ,o by post—Cretaceous erosion — I, l j 125‘ C’ . - l ~ crejjdive ‘ 17 - Measured section -— 300 I \3 5- . 51‘ 100 — I / O | . \ / 50- . I / e K 77. 60 ‘ 7p Borehole \ r / o 4? ‘ 46- (5x 61.1 e400 7 Miles City . / T / \i . 22$}??? 7 \-:77) _ 500 /// \\‘\\E A A :73 /// \- \7\ g;\ 77777 <;\ 0 5 10 20 30 40 MILES / / / >> Ni? z-\77 lW 1 . . l l l l l / // / 77 _ 0) \\\-77777 nOWy l , J: \\ TT‘\~ _l g 77 77 7 > a a /- - - 7 ~0-' ‘ O m g g (I 77777 "harm" . Z ‘2 I :3: Li \- i Z I \'\~ I ‘1 l L” if - .- . QC) 1 T \- 0‘ i R .ffl’ METERS FEET 4 . E 4., 77/77/-/ 0 W7 0 < , é a“ ‘3 l l 2 E I S ‘17 "1500 feet omltted 7-7100 1 1’3 ‘ I >\ \ _ l I I" 50 I ‘ \1 C ‘S 200 I \ .9 - I \\ \ K; i. 13 s— 300 100 A \ // e 400 \ >\ / \ 3 ,/ ‘ O , C , a) / / — 500 E z '3 .20 ,/ ‘V m / 0 5 10 20 3O 40 MILES < E r l i . . . l 1 l W _ l — LE 1 I 0 5 10 20 30 40 KILOMETERS Q. ’2 , / i / I | 1 I 1 I I I I WWWI o. “s ‘1 /, I // >~ / ‘= / — g / // (n g m — .110 .1 U) m I 7/ i (n ‘11 7, — t, ” EXPLANATION 2 fl / /" - Claystone Clayey // // / / 7. i / ,/ _€_ 8 o. / Siltstone Silty U3 3 / '6 e i W g o , / j. ‘. / / Sandstone Sandy // [—7 B ’ l - . 3 .19 A ,1“ 31 32M 33 34 35 J d'th6G 14 Coarse-grained sandstone Dolomitic Sixteenmile Canyon Flank Oil Co. Moss Agate Delpine H0 le Creek 7 U ' ap 15 E Stonehouse Canyon 1 59° 26' T- 5 ”-1 R- 5 5' Gjerde 1 S” 6" I7 ”-1 R- 8 5- sec 14 T 9 N R-10 E- secs 5 8 pi i0 N R 13 E secswm (211942 iilg N23 51nd ' 37 380' ziéTéIer'iI 1120115in T I . -2 sec.24, T.5 N., R. 8 E. . _ ' ' ' " ' l ' ' " ' ' r . secs. — , . ., . . o 7 \‘1 ,7“ Gardner and others (1946, p 28 9) Gardner and others (1946,p 30 33) Gardner and others (1945193335) Gardner and others (1946, p. 36—39) Gardner and others (1946, p. 40742) Galloway Creek Easton (1962, p. 1217124) 5 , ( sec. 7, T. 11 N., R.8 E. a) Conglomerate ( Douglass (1954) 7 Ems Gmu g 3 \ :- \( Dolomite \k 7- Devils Pocket (\ a7) :7: Formation z ITT I \I. I ‘8 % < L: \\ g g: — Nodular dolomite Calcareous dolomite / g E Alaska Bench Q 2 __ .— /i U Limestone g < S 55 > W / g __-7 Limestone Anhydritlc or gypsiferous . S i E “ W 0 ac, >_ '1' \ C 'c a .\ ., Cameron Creek 2 ”l m : 1 \ m E _ \ ~ Member g < z Halitic \‘1, . ‘2 Z :— _ 1 r \ 77* . L m \\ I t ‘ g 0— ' I .. ' 1 Stonehouse Canyon 1— " \ M Member \‘\\ fi‘o \ . _ \ 1 Covered interval or no sample \\ :3" 934298 Poorly exposed interval f \ I 7 Heath \ I Formation Conformable formation boundary \\ ‘ Unconformable formation boundary 1 ‘ ‘~ / WWWWWWWWWW - :1 / 77 Q Z Suggested intraformational ' 3 / 3 correlation line 2 < / ‘5 — 031213 -" / 0. US. Geological Survey / g o. fossil locality / L th /” Otter 5 '— e’f': : i l rs3ztagf 0:96p Formation on m z? 1, . r r :7 . 7 7 w /<\. . . / LINE B—B , FROM NEAR LEADORE, IDAHO,TO STONEHOUSE CANYON, BIG SNOWY MOUNTAINS, MONT. 7 Gammara7y 7-7,, ,’/ (D o , ‘0 i 2-. ; Self pote:t77i7a7l77:( ‘J L . L >' 5 g 'T/ i’ m 8 % Cameron l . . Z E E Creek l < L 4 age LE Member f) E; z :1? Lu / > E 0.. 0L) /:/ . fi 3 Stonehouse T I r g '— Canyon Neil E E . Member “ i . . . l . ‘\ l Heath FL f. 3:: Formation ‘ \X“ // /7 ,//<:; ,/ /. g. . /// / CE: i <2 l / /” (:2 / / g. i : g / //;-/ //L>: ‘ , O /‘ / ;i) i Z L // / <\ <5 / S < // “/75 // if — // “:2 n. / “(3/ o. // 7 > — 3 // O m : Otter / ,E (I) . / LL a) Formation // /’ / _ /// /// // / (f) // / / .20 y// // < (D m { / //// Q~ — / /« ////// EL) // / % E m --—777 //// / (f/ / . // Kibbey /// Formation //' /// Madison Group 53* 54* Carter Oil Co. Zach Brooks Drilling C0. Margaret Nelson 1 Larson 1 sec. 4, T. 37 N., R. 53 E. sec. 2, T. 32 N., R. 55 E. » c 3 7' .9 Ln . < 9 a 2 ~ e E‘: . <:. E, 8.g <1 <2; %.9 t , i> ' _ a N 1 ; > (I) ~'5 if :1 _ ,‘ '/ . r ‘3 (3 f"; (:5: 23 2:. La 7\ m‘ .fli— .- fi'.‘ <:/ 2,- 4, ‘ E) “ . 2:; 4. . Z g ' ‘-:'> . S L i ; R .2 6 _ , p 1 \ (0 ST». 2 :1) fix 2 :3 X «s B _c r o . 4 ~ ; L S I 3 COLUMNAR SECTIONS SHOWING CORRELATION \kf‘fi WALA AvaVn v\/ Mobil , v, (m W .li /\ v Vl“ W WWVV ‘ \ .'»\ /\ “A ’\ Wv R f /\ / 55* Producing Co. F—33—23P sec. 23, T. 29 N., R. 54 E. : gn AAA/\Wv sec. 29, T. 10 N., R. 32 E. sec.5,T.9 N., R. 34 E. WV. 44* Richfield Oil Corp. Cherry Creek Sheep 1 sec. 32, T. 11 N., R. 36 E. 56* Eramont Petroleum and others Northern Pacific Railroad 1 sec. 21, T. 23 N., R. 54 E. l l H “x/ V%x/\ A . M I, [Nail k‘ .N,‘ 1 i Ml ,4” y ..\o/ L/.,,Lm I, Saude Formation of Zieglar, (1955) 45* Phillips Petroleum Co. Northern Pacific Railroad 1 sec. 33, T. 11 N., R. 39 E. 2300: L. \ Coden Exploration Co. Nefsy-Thompson 1 sec. 34, T. 11 N., R. 42 E. Leg/Ml iTFbTmation L \Am/y // Ohio Oil Co.* 57* Stanolind Oil and Gas Co. Northern Pacific Railroad ”E” 1 sec. 21, T. 20 N., R. 52 E. WNW Formation MMWM gmx‘- \RL. leestone T\ T\ Tyler \\ \\ Formation \\\\\ Northern Pacific Railroad 1 / sec. 9, T. IO’MLILSEL/ 46* 47* R. L. Manning-Midwest Oil Co. Northern Pacific Railroad 1 sec. 27, T. 11 N., R. 44 E. E 5800 5999/ EA.” 48* Ohio Oil Co. Government-Cranston 1 sec. 33, T. 12 N., R. 49 E. M ETERS FEET 0 i 100 LIN E ‘P 200 Tm 300 100 ~ — 400 — 500 l ? llo 2'0 30 40 MILES i i i l l o | 110 2‘0 30 4o KILOMETERS 11 l l I I 50* Shell Oil Co. 59* Pine Unit 1 Shell Oil Co. sec. 30, T. 12 N., R. 57 E. Gas City Unit 33X—21 58* sec. 21, T. 14 N., R. 55 E. T‘ E" Texas Co. B. Elpel 1 sec. 35, T. 17 N., R. 53 E. w Shale /V\/ Formation (ma ///" /// i // / / g/ , '\O ///// rtmfl § {\0“ rd“@/// §/// ./ / / ,é/J / //./ /Fof 59/ / //£’ // LINE D-D',IN WESTERN PART OF WILLISTON BASIN THE WILLISTON BASIN 60* Shell Oil Co. Cabin Creek Unit 22—23 sec. 33, T. 10 N., R. 58 E. 49* K Phillips Petroleum Co: 50" Amerada Petroleum Co. Shell OII‘CO. Northern Pacific Railroad 1 Pine Unit 1 sec. 33, r. 12 N., R. 53 E. sec. 30. T- 12 N-. R. 57 E: r 62* W. M. and A. P. Fuller Northern Pacific Railroad 1 sec. 29, T.5 N., R. 60 E. 61* Shell on 00. State 22x36 sec. 36, T. 8 N., R- 59 E- IN 51* Lion Oil Co. Knight 1 sec. 29, T. 14 N., R. 60 E. ‘2 7§oo S IN CENTRAL WESTERN NORTH DAKOTA DI 63* 64* Carter Oil CO. Stanolind Oil and Gas Co. Lewis-Johnson 1 sec. 9, T. 129 N., R. 106 W. A. A. Clark-U. S. A. 1 sec. 11, T. 22 N., R. 2 E. Af/o/x . in / _/ 4‘] - C—C', FROM STONEHOUSE CANYON, BIG SNOWY MOUNTAINS. MONT., EASTWARD TO WESTERN PART OF Spearfish Formatio Opeche Shale n ' ’Wfiéka‘m'; ‘ Limestone upper member PERMIAN middle member lower member Minnelusa Formation PENNSYLVANIAN EASTERN MONTANA OF BIG SNOWY AND AMSDEN GROUPS IN CENTRAL AND EASTERN MONTANA Kibbey Formation Madison Group MISSISSIPPIAN PROFESSIONAL PAPER PLATE 2 C' 52 Shell Oil Co. Northern Pacific Railroad Brown 41—244 sec. 24, T. 142 N., R. 103 W. 554—8 Opeche Shale ‘ z S 2 upper E member a. C .9 S E middle 5 b r LL Z mem e S m E U) 2 5 LL :1) >_ C m E Z 2 E lower *1 member Otter Formation Q. :5 8 z 0 s o. > E 3 (I) g Q Kibbey w 8 . w __ Formation if: 2 Madison Group EXPLANATION Claystone Clayey Siltstone Silty Sandy Coarse—grained sandstone Dolomitic 4x 21:3”. Conglomerate Dolomitic limestone Dolomite Calcareous Calcareous dolomite Anhydritic or gypsiferous 3': Nodular limestone Halitic E Anhydrite or gypsum Cherty />/// X Halite Covered interval or no sample Poorly exposed interval See plate 1 for location of columnar sections Conformable formation boundary Unconformable formation boundary Suggested intraformational correlation line D343OB U.S. Geological Survey fossil locality Type of geophysical log * Lithological information from log by American Stratigraphic 00., Denver, Colo. ** Lithological information from log by Northwest Geological Services, Billings, Mont. lNTERIORiGEOLOGICAL SURVEV, WASHINGTON, D. c.719677666358 UNITED STATES DEPARTMENT OF THE INTERIOR GEOLOGICAL SURVEY 115° 114° 116° PROFESSIONAL PAPER 554—B PLATE 5 111° 110° 109° 108° 107° 106° 105° I I \ \ \s 95 ")0 I I I I EXPLANATION SASKATCHEWAN I 49 ° F Mh 5% Heath Formation 8 (5 >3 3 . g .3, . m Otter Formation .99 Big Snowy Formation CC: ' Mbsh, limestone and shale facies Mk Mbsk, sandstone and limestone facies 48° - - > \ Kibbey Formation 2 Madison Group undifferentiated Charles or Mission Canyon Formations in most places §§g§s¢ b ‘ a \‘fi" \‘3‘9 «M‘5u W ,1... ‘f: '.'- .. 47° H Area where upper Paleozoic rocks probably were deposited but were removed by post-Cretaceous Area where Big Snowy and Madison Groups beveled and truncated beneath Jurassic rocks. In other areas, Big Snowy and Madison Groups overlain conformably by Amsden Group I: Area of no control or beyond limits of investigation .\ >""“‘\.’"‘—__- 46° Contact Dashed where approximately located Probable fault "' 45° Possible fault A 0 Surface measured Subsurface borehole section log Control points 44° “1‘06; . 1 04° 107° 110° 108° {1' 5’. w I .PALEQ_GE0’.IQGIC MAP SHOWIN INTERIOR—GEOLOGICAL SURVEY, WASHINGTON, D. C_—1967—(366358 PS IN MONTANA >1 DISTRIBUTION OF BIG SNOWY AND MADISON GROU SCALE 1:2 500 000 50 o I 50 100 5 5 H I—L I__I 1. L gl~ ”Em j250 MILES 5 O O 50 100 150 200 250 KILOMETERS UNITED STATES DEPARTMENT OF THE INTERIOR PROFESSIONAL PAPER 554—B GEOLOGICAL SURVEY PLATE 4 MISSISSIPPIAN PENNSYLVANIAN BIG SNOWY GROUP AMSDEN GROUP TYLER FORMATION DEVILS POCK FORMATION OTTER FORMATION HEATH FORMATION STONEHOUSE ALASKA BENCH LIMESTONE CANYON CAMERON CREEK MEMBER NAMES OF SPECIES FORAMINIFERS Climacammma ? Roth and Skinner, 1930 Endothyra aff. E. exentralis Cooper, 1947 Endothg'ra sp. Brady/Ma sp. Tetrataxis sp. Ozawainella? sp. Mille'rella aff. M. chesterensis Cooper, 1947 Millerella n. sp. Millerella sp. Pseudostqffella sp. Profusulinella sp. Fusulinella or Profusulinella sp. SPONGE Clionolithes sp. CO RA LS Triplophyllites cf. T. spinulosus (Milne-Edwards and Haime, 1851) Bradypkyllum clam‘gerum Easton ' mplexus contortus Easton Camm'a montane'nsis Easton Kmu'nckophyllum? sp. va tumida Easton C RI N 0 i D Dicromyocrmus granularis Ea ston EC H I N OI D hinocm'nus sp. B RAC H | 0 PO DS Dingula cf. L. carbonaria Sh u m a rd, 1858 Orbiculoidea wyommgensis Bra n son a n d G reger, 1918 Orbiculoidea interlineata Easton Orbiculoidea sp. A Orbiculoidea sp. B cf. C. modesta White and St. John, 1868 Stemscisma obesa (Clark, 1918) Schuchertella sp. pseudolimtus Easton angulata Easton hinoconchus angustus Easton hmoconchus aff. E. altemtus (Norwood and Pratton, 1855) fasciculatus McChesney, 1860 nodosus (Newberry) n. subsp. cranez'si Branson, 1937 ? duodenam'us Easton Dictyoclostu‘s” confluens Easton Dictyoclostus” i’nflatus (McChesney, 1859) ‘Dictyoclostus" inflatus obsoletus Easton Dictyoclostus” inflatus spinolinearis Easton Dictyoclostus”richardsi (Girty, 1927) ntiquatom’a pernodosa Easton ntiqwatom’a n. sp. nivalis Easton Buxtonia”a1izonensis Hernon, 1935 rginifera planocosta Easton carboniferus Girty, 1911 pawulus Girty, 1927 quinqueplecis Easton ”HQ/"er curvatus Easton mfer increbescens Hall, 1858 rifer shoshonensis Branson and Greger, 1918 ' er wellem’ Branson and Greger, 1918 ' er praenuntius Easton sp. rt'im'a” sp. A artim‘a” sp. B Ta superata Easton hirsuta Hall, 1858 aff. C. sublamellosa Hall, 1858 atrypoides Girty, 1910 m'thym's n. sp. laem’s Weller, 1914 posita cf. C. lateralis Girty, 1910 posita ozarkana Mather, 1915 posita sulcata Weller, 1914 subquad’mta Hall, 1858 depressa Easton pidomella n. sp. rifer transversus McChesney, 1859 riina spinosa Norwood and Pratt en, 1855 cf. E. vera Hall, 1858 woodwo'rthi Clark, 1917 ? cf. H.? waltem’ Branson, 1937 .7 circularis Easton PELECYPODS (Palaemmcula) montanensis Easton Mangulata Easton mgodo'rsata Easton sp. A sp. B (Myalina) parallela Easton (Orthomyalina) sp. n.sp. .7 sp. pecten otterensis Easton 1 ~‘ ‘ j _ " ‘. ,_ a \ . ludlovi Whitefield, 1876 um glabratum Easton ' . f 1' I I, _ ¥ , , . , '. _: > ' ‘ ' latifasciatum Branson, 1942 . . . ,, / . sp. sp. . rdellasp. A rdella sp. B inflata Easton walkem' Weller, 1897 .7 sp. altidorsatus Easton GASTROPODS sacajawensis Branson, 1937 rollus (Euomphalus) n. sp.A (Euo'mphalus) n. sp. B ? excavatus Easton (Pseudozygopleura) aff. P. scitula (Meek and Worthen, 1860) cf. N. remex White, 1876 elegans Girty, 1927 SCAPHOPOD ypta subannulata Easton CEPHALOPODS ? sp. sp. ? sp. hespem'um Miller and Furnish, 1940 1 AreviewofLiwproductuandosus by Mackenzie Gordon. Jr., , ~ I V ‘I . i ‘ I‘ s; I ‘ Modified from Easton (1962, table 5); localities for these fossil girittgiesnsgircnirelug nhfsrlé'é‘éfiitezs that these are notL mduosua _ :2 . ‘ ~ : -' .1 EXPLANATION collections are described in detail by Easton on p. 109—114 Rén‘ge bfkspecies which is restricted to the Amsden Group Range of species which is restricted to the Big Snowy Group ;, 243—116 0 - e7 - (In pocket) \ ' , . L -' I MY he Nature of Batholiths GEOLOGICAL SURVEY PROFESSIONAL PAPER 554-C The Nature of Batholiths By WARREN HAMILTON and W. BRADLEY MYERS SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—C 14 survey of features of U. S. oatnolicns leads to Me interpretation Mac c/zese complexes are generally c/zin and Mac c/zey crystallizeaI oeneaz‘n covers consisting largely of Meir own volcanic ejecz‘a UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTONzl967 UNITED STATES DEPARTMENT OF THE INTERIOR STEWART L. UDALL, Secretary GEOLOGICAL SURVEY William T. Pecora, Director For sale by the Superintendent of Documents, US. Government Printing Office Washington, DC. 20402 — Price 30 cents (paper cover) CONTENTS Page Page Abstract ___________________________________________ Cl Regional descriptions and interpretations—Continued Introduction _______________________________________ 2 New England Appalachians ______________________ 013 Regional descriptions and interpretations ______________ 2 Erosion intervals _______________________________ 15 Sierra Nevada batholith _________________________ 2 Strontium isotopes ______________________________ 16 Idaho batholith _________________________________ 5 Origin and emplacement of batholiths _________________ 17 Boulder batholith _______________________________ 6 The environment of batholiths ___________________ 18 Volcanic ash ___________________________________ 9 Origin of granitic magmas _______________________ 18 Tertiary plutons of Cascade Range _______________ 9 Emplacement of batholiths ______________________ 21 Aleutian Islands ________________________________ 11 Batholiths and metamorphism ____________________ 23 Tertiary igenous rocks of Colorado ________________ 11 Batholiths and thrust faulting ____________________ 24 Tertiary igneous rocks of Basin and Range province- 12 Batholiths and younger structure _____________________ 25 St. Francois Mountains batholith _________________ 12 Selected references __________________________________ 26 ILLUSTRATIONS Page Page FIGURE 1. Graph of velocities of compressional waves in FIGURE 5. Geologic and metamorphic map of New a longitudinal section beneath the eastern England _____________________________ 013 Sierra Nevada ________________________ C4 6. Sketch showing metamorphic zones and 2. Geologic and crustal section through the granitic TOCkS Of a part Of southeastern Sierra Nevada of California ____________ 5 New Hampshlre: ----- ;-------;7---8-6-j- 14 3. Geologic map of part of the north end of the 7' Sketch showing lmtlal ratios ,Of St /Sr. {n . Paleozmc and Mesozom batholithlc Boulder bathollth, Montana ___________ 8 . . . _ . . ‘ granitic rocks ________________________ l6 4' Dlagrammatic longitudinal geologic and 8. Diagram showing pressure—temperature phase crustal section through the Boulder batho~ relationships in material of anhydrous lith, Montana ________________________ 10 basaltic composition __________________ 19 III SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY THE NATURE OF BATHOLITHS By WARREN HAMILTON and W. BRADLEY MYERS ABSTRACT A survey of features of batholiths in the United States is interpreted to indicate that batholiths generally are thin, having spread out laterally at shallow depth, and that many of them reach the surface and crystallize beneath a cover of their own volcanic ejecta. It is inferred also that the magmas originate in the lower crust or upper mantle at depths greater than any ever exposed by erosion. Such conclusions agree with those reached by many geologists, but disagree with the concepts that batholiths are masses of great thicknesses, form beneath deep cover of metamorphic rocks, and crystallize from melts mobilized at the levels exposed in gneissic and migma- titic terranes. Successively older Phanerozoic batholiths display in a broad way successively deeper sections into batholithic complexes. Mid-Tertiary batholiths are largely still capped by their vol- canic crusts. The Late Cretaceous Boulder batholith of Mon- tana is only a few kilometers thick; it spread across a floor of prebatholithic rocks, and preserves discontinuously its roof of almost exclusively volcanic rocks, which are contempora- neous and consanguineous with the plutonic ones. The Idaho batholith, largely of middle Cretaceous age, has a few small areas of possible volcanic roof rocks, although the correlative Sierra Nevada batholith has none remaining; but the more than a million cubic kilometers of volcanic ash in Cretaceous strata in the continental interior has no apparent source other than such batholiths. Seismic and gravity data indicate the Sierra Nevada batholith to be probably thin, and gravity data indicate the same for the Boulder batholith. The large Meso- zoic batholiths were unroofed within a small fraction of a geologic period after their formation; since unroofing, they have been incised but not greatly eroded, and shallow depths are indicated. Late Cenozoic uplift correlates with Mesozoic batholithic rock type, apparently increasing with the propor- tion of radioactive components; the present crustal roots of the batholiths are responsible for the uplifts, are more mafic than the exposed rocks, and must somehow have formed be- cause of the overlying batholiths. The uplift is now resulting in selective erosion of Mesozoic batholiths. Batholiths are abundant in Precambrian and Paleozoic ter- ranes, but probably none of these batholiths are as large as the great late Mesozoic batholiths of western North America. One possible interpretation of this contrast is that most pre- Mesozoic batholiths have been selectively eroded away. Gneiss terranes such as the sillimanite “plateau” of New Hampshire may form beneath batholiths, as plutons of magma rise bubblelike and displace heated wallrocks, which flow downward and beneath the plutons and become intensely metasomatized and injected. The metamorphic gradients flanking many gneiss terranes are far too steep to be ex- plained in terms of geothermal heat conducted from the man- tle, and the heat may have been introduced in magmas that largely rose through the crust and coalesced into surficial batholiths. The largest Phanerozoic batholiths are partly in eugeosyn- clinal terranes, and the tectogene hypothesis of melting in downbuckled geosynclines is based on this association. Many Phanerozoic batholiths, however, intrude miogeosynclinal and platform sedimentary rocks and Precambrian basement rocks, and even have formed in the oceanic environment of island arcs. The tectogene hypothesis cannot be applied to such noneugeosynclinal batholiths—if batholiths have a common cause, it cannot be a tectogene. Strontium isotope data indicate that granitic magmas are melted from rocks poorer in ru- bidium than are exposed basement rocks and thus are derived from the lower crust or upper mantle or (in eugeosynclines) from volcanic materials derived in turn from such sources. Batholithic and silicic—volcanic magmas become in general more silicic and more potassic as the continental crust be- comes thicker, so the lower crust may be increasingly involved in melting as its depth increases. The magmas produce batho— liths capped by volcanic fields in some places but produce high-alumina volcanic fields alone in others, depending upon local factors. Laboratory high-pressure data require that high-alumina batholithic and volcanic magmas be equilibrated with crystals above the depth at which basaltic rock under- goes pressure—phase transformation to eclogite; this can be achieved by partial melting, differentiation, or assimilation. Zone melting—whereby volatile components rising in response to pressure gradients within magmas lower the melting tem- perature of the roof while forcing crystallization low in the chamber—can cause great assimilation; indeed, very little of the final high-level magma need represent material present in the initial melt. Much deformation conventionally ascribed to either crustal compression or gravity sliding may be due to the shouldering aside of wallrocks by rising batholith magmas. Batholiths, once formed, resist fragmentation by younger structures and hence greatly influence the subsequent deformation of their regions. Cl C2 INTRODUCTION Batholiths are composite masses of granitic rocks having areas ranging from tens of square miles to tens of thousands of square miles. Some batholiths that out sharply across their wallrocks and that are sur- rounded by contact-metamorphic aureoles clearly formed from magmas intruded from greater depths. Other batholiths are largely concordant and lie within terranes of uniformly high—grade gneisses, and the origin of such batholiths and the source of heat for metamorphism of the associated gneisses are less obvious. We attempt no broad review here of the descriptive features of batholiths; Buddington (1959) has done that ably, and the reader is referred to his work. lVe note features of batholiths and related rocks in the United States which appear to need explanation in any general theory of the origin and emplacement of batholiths. Much interpretation is incorporated with the individual descriptions in the first section of the paper. General synthesis and speculation follow in the second section. The rationale developed is that batholiths form from magmas generated in the upper mantle and lower crust, beneath any depths exposed by erosion; that pluton magmas rise in detached balloonlike forms through the crust, frequently reach- ing the earth’s surface, and coalesce into shallow and fairly thin complexes; and that gneissic terranes form in the zones, beneath the final batholiths, through which the plutons rise, as wallrocks flow beneath the rising magmas and are heated and metasomatized by them. Elements of this synthesis have been sug— gested by other geologists, who are cited in the appro- priate places. The third section of the paper discusses the influence of batholiths, once formed, upon the sub- sequent structural evolution of their regions. Rock names applied on the basis of petrographic criteria to intermediate volcanic rocks generally con- note compositions markedly more mafic and calcic than are connoted by names applied to plutonic rocks of identical chemical compositions. One is likely to think of andesite and diorite, dacite and quartz diorite, and rhyodacite and granodiorite as being of the same compositions, but this is not generally true. For example the postbatholithic andesites of Mount Rainier contain 60 to 64 percent SiO2 and 1.6 to 1.9 percent K20 (Fiske and others 1963, table 2); plu- tonic rocks of the same composition would be classed petrographically as granodiorite or at least quartz diorite. Chemical analyses rather than petrographic names should be employed when comparisons are made between intermediate plutonic and volcanic rocks. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY REGIONAL DESCRIPTIONS AND INTERPRETATIONS SIERRA NEVADA BATHOLITH The Sierra Nevada batholith is known best across its central part (Bateman, 1965; Bateman and others, 1963; Calkins, 1930; Ernst Cloos, 1932, 1935a, 1935b; Durrell, 1940; Hamilton, 1956; Krauskopf, 1953; Macdonald, 1941; Moore, 1963; Rinehart and Ross, 1964; D. C. Ross, 1958; Sherlock and Hamilton, 1958; and others). Bateman, Clark, Huber, Moore, and Rinehart (1963, pl. 1) compiled a geologic map of the central Sierra. The batholith is 55 to 110 kilometers wide, has an exposed length of 650 km, and is a com- posite of plutons of Late Jurassic and Cretaceous ages (Kistler and others, 1965). Large plutons are elon- gate parallel to the northwesterly regional strike and small plutons lie between and within them. Younger plutons in many places cut sharply across structures of older plutons but in some places are nested con- cordantly inside them. Adjacent plutons can be of markedly different compositions within the spectrum quartz diorite-granodiorite-quartz monzonite-alaskite. The larger plutons and the bulk composition tend to become more silicic and richer in alkalis eastward; the dominant rock types are mafic calcic quartz diorite and granodiorite in the west and leucocratic grano- diorite and quartz monzonite in the east, but excep- tions are numerous on both sides. Contacts between plutons and wallrocks and be- tween adjacent plutons are typically so sharp that a hand specimen can be taken across them, although dike—injection zones make some contacts gradational at mapping scales and contact migmatites are extensive along some contacts with metamorphic rocks. Con- tacts may be quite irregular in detail but generally are broad curves at map scale. Thin discontinuous screens of metamorphic rocks locally separate plutons, and large pendants of metamorphic rocks separate groups of plutons. The batholith is bounded by irregular belts of lightly to moderately metamorphosed Paleozoic and Mesozoic sedimentary and volcanic rocks. The pro- portion of mafic to silicic volcanic rocks is higher in the western border belt than in the eastern belt. The rocks of the western border belt in the central Sierra dip steeply toward the batholith and consist of long fault blocks of which those nearest the batholith con- tain the oldest rocks, although in each block the youngest rocks tend to be on the side toward the batholith; the eastern border belt of the central Sierra dips steeply in either direction but its rocks become in a gross way younger westward toward the batholith (Bateman and others, 1963). One pluton of the west THE NATURE OF BATHOLITHS part of the batholith was interpreted by Ernst Cloos (1932, 1935a) to have sent a flat tongue westward over its wallrocks. He (1935a, 1935b) emphasized that the flow—structure domes of some plutons demon- strated the batholith to have spread laterally, pushing its wallrocks aside, across the central Sierra. Assimilation of mafic metamorphic rocks into mo- bile granitic magmas has been demonstrated in many places and probably has contributed much to the more mafic and calcic character of the western plutons, for mafic metavolcanic rocks are abundant in the western border belt, and widespread assimilation is shown along contacts. Static granitization has nowhere been found on more than a very small scale. Both regional and contact metamorphosed rocks show metamorphic grade and intensity decreasing systematically away from contacts: the primary source of heat for meta- morphism to assemblages of higher temperature facies than greenschist was intruded granitic magma (Bate— man and others, 1963; Durrell, 1940; Macdonald, 1941). The flow structures of most plutons show them to have risen as units past their granitic and meta- morphic wallrocks. Injection complexes of gently dipping dikes at some contacts demonstrate vertical stretching of wallrocks (Bateman, 1965, p. 116; Moore, 1963; Sherlock and Hamilton, 1958). Offset belts of metamorphic rocks indicate in some places shoulder- ing aside by rising plutons (Bateman and others, 1963; Moore, 1963; Rinehart and Ross, 1964) but else- where wallrock belts are truncated irregularly by plu- tons. The presence of contact breccias along some contacts indicates that stoping was operative during late stages of intrusion, but the general absence of xenoliths away from contacts seems to be evidence against the process as the dominant mode of emplace- ment. Much detailed mapping apparently demon- strates that the intrusion of the plutons was domi- nantly forcible (Bateman, 1965, p. 115—123; Bateman and others, 1963, p. 44). Evidence for a very shallow depth of crystallization of Cathedral Peak Quartz Monzonite—a large pluton of very coarse grained leucocratic rock—was summa— rized by Evernden (1965). The relation between vari- ations in potassium-argon ages and in elevation led him to conclude that the pluton was emplaced no deeper than 7 km, and possibly as shallow as 4 km. Seismic—refraction data indicate that the Mohoro— vicic discontinuity lies 40 or 45 km below sea level beneath the crestal region of the Sierra Nevada near the 39th parallel (Eaton, 1963) and about 50 km beneath the highest part of the crest farther south (L. C. Pakiser, written commun., 1964). Integration C3 of gravity data with this seismic information requires that the greater part of the crustal thickening repre- sented by the deepening of the Mohorovicic discon- tinuity is in crustal rocks that are markedly denser than the granitic rocks of the exposed batholith. The gravity data published by Thompson and Talwani (1964), for example, when considered with the seismic model, suggest that the Sierra Nevada batholith near the 39th parallel is a tabular structure with a thick- ness of at most 8 km.1 The negative Bouguer gravity anomalies are not sufficiently large to permit both granitic rocks and the entire crust to be thick; as the crust is demonstrably thick, the granitic rocks can- not be. From his preliminary interpretation of explosion seismic waves passing longitudinally through the eastern Sierra Nevada, Jerry P. Eaton (oral commun., 1965) concluded that velocities increase downward in about the manner indicated schematically in figure 1. We are much indebted to Eaton for permission to incorporate this information. Rocks with a velocity appropriate for silicic granitic rocks such as those dominating the surface exposures apparently extend no deeper than about 10 km. We interpret the velocity increase near this depth to occur beneath the thin batholith, and suggest that the underlying high- velocity (6.4 km per sec) rocks are metasomatized schist and gneiss displaced beneath the plutons of the batholith as they rose toward the surface. The depth- velocity fields of common types of igneous rocks are shown in the figure for comparison. Other seismologists have drawn different interpre— tations from seismic data. Mikumo (1965) suggested that low-velocity granitic rocks need extend little deeper than sea level, where they give way downward to denser (Vp=6.3 km per sec) rocks. He achieved close agreement between measured Bouguer gravity and gravity calculated from a model in which the entire crust beneath the Sierra Nevada, from the sur- face to a Mohorovicic discontinuity reaching 46 km, has a density of 2.80 gm per cm3. (Surface rocks lighter than this are, however, exposed throughout much of the region.) Press and Biehler (1964), on the other hand, in- ferred that there is a velocity inversion within the upper crust and that rocks with a velocity appropriate to granite extend to great depth. They studied arrival times of P waves from nuclear explosions in the western Pacific, and found that arrivals at Tinemaha and Reno (both short distances east of the Sierra lThompson and Talwani assumed that the Mohorovicic disconti- nuity lies no deeper than 32 km beneath the Sierra crest. This as- sumption is contradicted by the seismic data; and even so, they could fit no more than about 12 km of granitic rocks to the gravity model. C4 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY [ \\ Batholith \ Metasomatized schist and gneiss 20 — ‘— u) r— ............... O: (D Lu g E r— o -‘ “J O 7 2 -* m 0 o E"- _‘ m C :l 5 ‘5 3 § 3 x — "3 :1 O S ‘9 ~— 2 3 g 0g, 3 Q) _ 5' 5- 0' :‘5 I- ‘D 2. 9: ‘1 r— ” ”" 0 E 3. c; 2:} Lower crust ‘ o '2 3. w (gabbro and amphiboltte) c» a 40 — Er § — 3 FT ‘3 e e. 55 Fr? — ................................... _ MOHOR’OV/CIC """""" DISCONTINUITV Mantle (olivine-rich gabbro) | | l 60 - 5 6 7 8 9 VELOCITY, IN KILOMETERS PER SECOND FIGURE 1.—Ve10cities of compressional waves (heavy lines) in a longitudinal section beneath the eastern Sierra Navadla, from an unpublished interpretation by Jerry P. Eaton. Depth- velocity fields of representative types of igneous rocks are gen- eralized from Birch’s (1960) data. An interpretation of the velocities is shown at the right. Nevada mountain block, but near the axis of the gravity minimum associated with the Sierra and its batholith) were about 0.8 sec later than they would have been if the structure of the crust and upper mantle was the same throughout California as it is at Pasadena. If this delay is due (as they assumed) to low—velocity granitic rocks beneath the Sierra Nevada, then the batholith has a deep root, and a thickness of granite and granodiorite of approximately 37 km is indicated (Press and Biehler, 1964, p. 2987—2988). The interpretation by Press and Biehler is not however in accord with the local seismic—refraction data (such as that illustrated by fig. 1), and the P—wave delays they found can be interpreted alternatively in terms of variations within the upper mantle. To eliminate the possibility that variations in thickness of the lower crust could account for the arrival-time delays, Press and Biehler (1964, equation 10) analyzed the data for fit to these relationships: (Ag) =42Ah( p— p’) z - 12( PB) ozo/ in which (Ag) is the difference in slab—equivalent station (Pasadena) and the seismograph station; Ah is the difference, in kilometers, in depth of the Mohorovicic discontinuity; (p—p’) is the difference in density, in grams per cubic centimeter, above and below the discontinuity; (PD) is the wave delay, in seconds; and a and a’ are the velocities of P waves above and below the discontinuity. Press and Biehler showed that, given reasonable velocities, the expression on the right yields values of (Ag) approximately 50 percent too high. The middle expression, however, provides calculated values agreeing with the observed ones: the reasonable figures of Ah=22 km and (p—p’) = 0.3 gm per cm3, for example, yield (Ag)=280 mgal, the same as that observed. The conflict can be resolved by the interpretation that the wave delay is due to lower velocities in the upper mantle beneath the Sierra Nevada than beneath coastal California, and that the batholith is thin, rather than thick as Press and Biehler assumed. Heat—flow data also indicate the Sierra Nevada batholith to be thin. Thus, granodiorites west of the crest of the central Sierra produce by radioactive Bouguer gravity, in milligals, between the reference decay about 10 microcal per gram per year, yet heat THE NATURE OF BATHOLITHS flow in a deep core hole is only 1.3 microcal per sq cm per see; if the entire flux came from granodiorite like that in the hole, the granodiorite could be only 15 km thick (Lachenbruch and others, 1966; see also Wollen- berg and Smith, 1964). The more mafic granitic rocks of the western part of the batholith produce about 21/2 microcal per gram per year, and heat flow in a deep core hole there is only 1.3 microcal per sq cm per sec, equivalent to the heat production in 30 km of the local rock (Lachenbruch and others, 1966). As much of the heat must in fact come from greater depths, batholithic rocks like those near the surface must be much thinner than these limits. Five heat-flow deter- minations farther north in the Sierra batholith aver— age only 0.9 microcal per sq cm per sec, half the value in the Great Basin to the east (Roy and Blackwell, 1966), leading to the same conclusion that the batho- lith is thin. We conclude that the crustal root beneath the high part of the Sierra Nevada is largely of rocks markedly more mafic and heavier than the exposed quartz mon- zonite and leucogranodiorite, and that the batholith is limited to the upper part of the crust. Figure 2 shows a section through the central Sierra Nevada. The interpretation incorporates seismic and gravity data, and illustrates conclusions developed in this paper. Olivine-rich gabbro of upper mantle VpE 7.9 km per sec (Olivine eclogite lies deeper) W £ Ki // ~—\iZ(//’ ;.I/ f ._> / Gnelsses formed as wallrocks flowed J I / Z.) Gneisses of upper crust ' downward and beneath rising plutons / '? mostly Precambrian C5 IDAHO BATHOLITH* The Idaho batholith of central Idaho has an exposed length of 400 km and a width of 130 km and is sur- rounded by regionally metamorphosed rocks (Hamil- ton, 1963a, b; Larsen and Schmidt, 1958; Reid, 1959; Ross, 1963; Schmidt, 1964; and others). Massive granodiorite and quartz monzonite underlie two main regions, one in the southwest part of the granitic ter— rane and the other in the northeast; elsewhere there is much schist and gneiss interspersed With the gra- nitic rocks of the batholith. Quartz diorite and trondhjemite are Widespread in the gneissic western border zone of the batholith. The culminating intru- sions of the Idaho batholith occurred about the middle of Cretaceous time according to lead-alpha determina- tions on zircons (J afi'e and others, 1959). It is possible that the Idaho batholith formed with- out a roof, its plutons having reached the surface and erupted a volcanic capping, beneath which magma crystallized more slowly. The Casto Volcanics in the east—central part of the region of the batholith can be interpreted speculatively to be remnants of this volcanic cap. The Casto is undated, variably al- tered, and contact metamorphosed but generally non- schistose intermediate lavas and pyroclastics intruded by the Idaho batholith (Leonard, 1962; Ross, 1934). *See note on page 030. Basin and Range province Owens Valley VERTICAL AND HORIZONTAL SCALE (I) 110 210 3'0 40 5'0 MILES I O 50 KILOMETERS L___J__1_l___J__l EXPLANATION a r Upper Cretaceous and Cenozoic sedimentary rocks FIGURE 2.—Geologic and crustal section through the Sierra Nevada of California, along the 37th parallel. Plutons of granitic magma, melted in upper mantle and lower crust, rose through crust and In the Basin and Range Province, Paleozoic sedimentary rocks ilton and Pakiser (1965). coalesced at surface to form Sierra Nevada batholith. , I I \I’ a r sl" Granitic rocks M5 Metamorphic rocks Adapted from Ham- moved along bedding—plane thrust faults, then broke into normal-fault blocks. 232—376—67——-—-2 C6 Deformation of the volcanics is only moderate. Typi— cal dips are about 25°, and the rocks apparently lie unconformably upon highly deformed metasedimen- tary and metavolcanic rocks of which some at least are of late Precambrian age. Ross and Leonard both assumed the Casto to be part of the Paleozoic and Mesozoic eugeosynclinal suite, much older than the batholith, and perhaps of Permian age; if this as— sumption is correct, so is their conclusion that there has been no major deformation or regional metamor— phism since pre-Permian time. In the western border zone of the batholith, however, the fossiliferous Upper Triassic Martin Bridge Limestone was extremely de- formed, highly metamorphosed, and intruded by gra- nitic rocks which in turn were intruded by the main plutons of the Idaho batholith (Hamilton, 1963a, b), and the geometry of the deformation that accom— panied the regional metamorphism indicates a genetic relation to the batholith. The absence of comparable deformation in the Casto Volcanics might indicate that they postdate the emplacement of the batholith—— and yet the volcanics are intruded by the batholith. A possible explanation is that the Casto Volcanics formed by extrusion of lava from the Idaho batholith itself; if so, at least part of the batholith formed with no roof other than a crust of its own ejecta. Leonard (1963) and Leonard and Stern (1966) have, however, presented further evidence which they regarded as indicating that major deformation in cen- tral Idaho long predated the Idaho batholith. Highly deformed upper Precambrian metamorphic rocks were intruded there by syenite whose lead—uranium and lead—thorium calculated age is about 600 million years. (One potassium-argon hornblende age of only 93 my. apparently represents metamorphism by the Cre- taceous batholith.) Leonard considered the syenite to postdate the major folding of the enclosing rocks, and our conjectures in the preceding paragraph are wrong if he is correct. The syenite, however, is grossly con— cordant t0 the structures of its Precambrian wall- rocks, and has been variably crushed and recrystal— lized; so it is possible alternatively that the syenite was intruded into the old rocks before, rather than after, their metamorphism and deformation and thus that the syenite provides a maximum rather than minimum age for that event. BOULDER BATHOLITH The Late Cretaceous Boulder batholith of south— western Montana is a composite mass 100 km long and 50 km wide, and consists of plutons of granodiorite, quartz monzonite, and other granitic rocks (Becraft and others, 1963; Klepper, 1950; Klepper and others, 1957; Knopf, 1963, 1964; Ruppel, 1963; Smedes, SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY 1962). The batholith lies far to the east of the general belt of late Mesozoic metamorphism and batholith formation. The batholith appears to be a floored sheet, roofed only by its own volcanic ejecta. The roof rocks of the batholith are exposed in many areas and consist of dacite, rhyodacite, and quartz latite. (Considerable “andesite” has been reported in the literature also, but published analyses of such rock are clearly of dacite and rhyodacite.) The granitic and volcanic rocks are of the same age within the Late Cretaceous insofar as the radiometric and iso- topic age determinations of the granitic rocks can be compared with the paleontological dates of the vol— canics. The volcanic rocks have a maximum thickness of 2 or 3 km and are mostly pyroclastic. They are broken only by warps and normal faults in most areas, and are but slightly altered except where converted to hornfels near the contact with the batholith. Younger plutons within the batholith are in general more silicic and richer in alkalis than are older ones. A similar irregular compositional age progression within the volcanic roof rocks, silica and alkalis in— creasing upward, has been reported on the basis of field and petrographic studies, although it is not apparent in the few published chemical analyses of the rocks involved. Granitic and volcanic rocks broadly overlap in composition but the granitics are in bulk composition more silicic and more potassic than are the volcanics. The present erosion surface everywhere is probably within 2 km above or below the position of the origi- nal roof of the batholith, for the subhorizontal contact with volcanic rocks is exposed in most regions about the batholith. Presumably the original relief of the roof was still less, as Cenozoic deformation has much affected western Montana, and the initial top of the batholith may have been almost horizontal. Roof- rock lavas lap across the edge of the batholith onto the older wallrocks. A few small masses of contact— metamo‘rphosed sedimentary rocks, probably of pre- volcanic Mesozoic units, are present locally in the roof complex and perhaps represent rafts of floor rocks. ‘ Lower Eocene quartz latite lies upon eroded granitic rocks (Smedes and Thomas, 1965), and shows that the batholith was exposed by erosion very soon after its formation and that magmatism continued into Tertiary time. The volcanic rocks also form the wallrocks for most of the east margin of the batholith. The contact is steep and irregular._ Deformation of the wallrock volcanics has been more severe than that of the roof— rock ones, and the wallrock volcanics have gentle to moderate dips and are broken by many faults. THE NATURE OF BATHOLITHS The north contact of the batholith is semiconcordant to Paleozoic and prevolcanic Mesozoic strata which dip southward beneath the granitic mass (fig. 3; Knopf, 1963; Smedes, 1962). Dips in the wallrocks tend to steepen toward the contact with the batholith. The contact forms in plan three large cusps, concave toward the batholith; the western and central cusps appear in figure 3. The cusps are synclinal sags and are separated by sharp anticlines. The western cusp is 25 km across, and the batholith is in contact with middle and upper Paleozoic beds. The eastern two cusps are 20 and 10 km wide, and the granitic rocks lie against Cretaceous and Jurassic strata. Relatively dense (specific gravity about 2.8) granodiorite lies along the central and eastern cusps and forms an out- crop belt 0 to 4 km wide, south of which is less dense (about 2.71) quartz monzonite. A subhorizontal sheet of still lighter granophyre lies discordantly above the heavy granodiorite, in the same structural position occupied by nearby remnants of volcanic roof rocks (fig. 3). We interpret the general parallelism of the granodiorite-quartz monzonite contact to the margin of the batholith as suggesting that the granodiorite is part of a sheet which dips southward beneath the quartz monzonite. ‘ The south contact of the batholith is against Pre- cambrian crystalline rocks and Paleozoic and Meso- zoic strata. Available data indicate the contact to be complex but are too meager to permit satisfactory generalization beyond the observation that long seg- ments of the contact are semiconcordant to wallrock formations whose tops lie toward the north. The west contact of the batholith is hidden beneath Cenozoic deposits but probably is largely or entirely against Paleozoic and prevolcanic Mesozoic strata. then it is viewed on a broad scale such as that of the geologic map of Montana (Ross and others, 1955), the Boulder batholith is seen to occupy a structural depression. The batholith is surrounded mostly by Mesozoic and Paleozoic strata, whereas Precambrian rocks are extensively exposed in other parts of the same tectonic province elsewhere in Montana. The batholith fills a basin. The roof of the west—central part of the batholith is against an almost constant stratigraphic level in the overlying volcanic rocks through a broad area (Ruppel, 1963, p. 37). Contacts between variant gra- nitic rocks in the batholith are partly sharp and partly gradational but are in general subparallel to the roof contact, and successively lower units tend to be succcessively coarser grained; fine—grained quartz monzonite several hundred feet thick typically lies between the volcanic roof rocks and the coarser quartz C7 monzonite of the interior of the batholith (Ruppel, 1963, p. 32, 37). Lawson (1914) long ago suggested that the Boulder batholith was a floored sheet, intruded between the Cretaceous volcanic rocks and the older rocks beneath. Barrell (1907, p. 166) made similar suggestions still earlier but thought them improbable. Ruppel (1963) suggested that the west-central part of the batholith was a floored sheet because the subhorizontal contacts between plutons are strong evidence for horizontal flow of the intrusive magma. The concordance of the northern contact to the inward-dipping right-side-up Mesozoic and Paleozoic section indicates that sector also to be floored. The near lack of prevolcanic rocks along the east margin of'the batholith suggests that there, too, the granitic rocks lie wholly above the pre- volcanic section. Prevolcanic rocks are virtually lacking in the roof. The Boulder batholith is capped almost exclusively by its own volcanic ejecta and is better regarded as an extrusive complex, of which the volcanic rocks form the upper part and the granitic rocks the lower, than as an intrusive complex. The batholith magma flowed, in effect a gigantic mantled lava flow, across a broad basin whose subsidence may have been due to the withdrawal of magma from depth. Presumably a crosscutting batholith (in the customarily limited sense of the term) within the basin, or several stocks and small batholiths, served as magma conduits. The magma formed volcanic rocks where it erupted to the surface and granitic rocks where it crystallized be- neath an insulating crust of its own ejecta. The cap- ping crust was thickened by eruption over the top, by injection into it of dikes and sills (which are abundant in the volcanic roof), and by crystallization along the bottom of such rocks as granophyre and fine-grained quartz monzonite. Granitic rocks intrude the volcanic rocks wherever the two are in contact, but this does not require the granitic rocks to be wholly younger than the volcanic ones: it indicates that granitic rocks formed where magma crystallized beneath insulating cover. The generally offset compositional ranges of volcanic and granitic rocks—the more silicic half of the volcanic rocks having the same general composition as the less silicic half of the granitic rocks—may indicate that the overlap in age consisted largely of the younger half of the volcanic series and the older half of the granitic series. If this is correct, then thick volcanic rocks formed across the entire basin before much magma spread laterally between volcanics and floor; however, the hidden deep part of the batholith may well consist of relatively mafic rocks of the'same com- SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY NVIWESd CINV ‘NVINVA‘IASNNEd NVIddISSISSIW 'NVINOAEG AHVNUEanO SDOBDVLEHD DISSVHnr NVIUQWVQ NVIHSWVDBHd f_&'\ r—H r—A-a -._ . 3%? § XE: L .2 853-23 8‘ h: 8:“ a. ¥§ a mg: a §-S § ~ ‘9 ~ a: u 353 § 3% g gs: ° :§ 5 Z IEXK $» 3” 9' o W “ :5 3’2 x“ E E “'3 "6 .... w'fiuf- m 3 o 0 0 U39 3 _. o' s l- g fi§§m "236$ 8 2 8 SEE (”w < :3 2s:% “=3 2* E 3* NS § §§ 3 3' eggs: xnzé: gs «2 g =§<~a 2 Ea -~ "were: 1:.«° 0) cu “‘~ 0 =3 0 :th. §p a.) oa§ . _. :5 ”egg airs .2 .§ .2 5:: o W “L = (Swim >g§ a 3 a 32-52? : >< ”1 gmge s m m :12 mos ° LIJ Ex :3 Ga 0Q. 3 ._ g 8 V: a N 5‘“ 93.323 S "S: 5 iafi " an“ a. a) o '3 - g 533$ : E§ w b- 's “(Em V: o~ § °§Et a' -= V: 3 G . (1) “Ex”: :: Buxe‘ Q Q: a W—l snoamnaao “Mn snoamaag sataas nag M90107 6 °‘ L00 00 ‘1'?“ E H ‘z‘ 3 La A Ir} a: . 1/ \ \ 7\ 1:3]— -/\ \l/,\ \ \\ \—/ a _, - )’>.71\/_\' 317% \\/_‘\/ //‘\\/, — ,\\|/\\‘\/ /\/ \I’11'/\\/l\I—\’ / >/,\/\,~\’//\ :/ m, r 501 r“ /_\\:\/\/:< ,Q _\/_\/,\,\ pt \\ /\/\_\\/\/' /’ /__\\/\ \I’ ’I / \ \ \ b /\\/ /.\/’ < H \\— \’/ ‘ /\ \ — o /_\'/ ‘\C~ (I N \‘\\/’,\‘C‘/ \’ .—. /\‘(‘/ /\/ /\ H \ ’//;/\T,\’ \7 [QC/CNS \ C ’<\7\ —’ \/\/: / \, \I\:‘ /\ k types is 10 1‘00 Strike and dip of overtumed beds Syncline Showing trace of axial plane and plume of axis nomenclature of granit 30 Strike and dip of beds Generalized from Knopf (1963) , Montana. that of this paper rather than of Knopf. l 5 MILES 5 KILOMETERS ap of part of the north end of the Boulder bathol oglcm FIGURE 3.—Geol THE NATURE OF position and age as the more mafic volcanics. A rea— sonable picture is that intrusion and extrusion largely overlapped in age and that the volcanic crust was thickened and broadened as granitic rocks crystal— lized beneath it. Repeated injections of magma be- neath the volcanic cover are indicated by the many contacts between plutons. In the west-central part of the batholith, successive plutons flowed laterally. Successive episodes of subsidence are suggested by the geometry along the north margin of the batholith. The granodiorite south of Helena (fig. 3), for exam- ple, may have crystallized after its magma filled the cuspate embayment in the wallrocks, and then both granodiorite and walls may have subsided more. The deepened cusp was then filled by the magma which formed the quartz monzonite to the south and above the granodiorite, if, as we infer, dips between the plutonic units decrease into the batholith. The aggre- gate thickness of the units is probably markedly less than might be inferred from extrapolation of dips of the north contact of the batholith. A crust of volcanic rocks perhaps 2 km thick floated upon granitic magma over a region of about 7,000 square kilometers. Barrel] (1907, p. 166) implied this long ago. Whether or not all the roof floated at any one time is not yet certain. The warping and normal faulting of the volcanic pile may have occurred largely as a result of its vertical and horizontal jostling and stretching while afloat. Many of the points discussed here and in subsequent sections are illustrated by the diagrammatic section through the Boulder batholith (fig. 4). A gravity survey across the north end of the ,' Boulder batholith shows no abrupt decrease in gravity corresponding to the margin of the granitic mass, which thus cannot there have great thickness (Renick, 1965). The data indicate that the surveyed part of the batholith is very thin at its north edge and thickens only gradually southward to a thickness of about 5 km, 15 km south of the north margin. (Renick’s model shows the batholith thickening to 6 or 7 km, but 5 km provides a better fit with the observed gravity.) An unpublished gravity survey of a larger area of the batholith and surrounding region is said to indicate the batholith to be thinner than 15 km (Biehler, 1966). The Boulder batholith lies in a miogeosyncline— not in a eugeosyncline. Most Mesozoic batholiths of western North America occur at least partly within eugeosynclines, but such an environment is obviously not necessary for their formation; and an explanation for the origin of batholiths cannot properly apply to a eugeosynclinal setting alone. Such matters are C9 discussed in the subsequent section on “The environ— ment of batholiths.” VOLCANIC ASH Mesozoic strata of the western interior United States contain 1 or 2 million cubic kilometers of clay altered from volcanic ash blown in from sources nearer the Pacific Ocean. The varicolored siltstones and claystones of the Upper Triassic Chinle Forma- tion of the southern part of the Colorado Plateaus consist largely of montmorillonitic clay derived from latitic or quartz latitic ash and of mixed-layer clay derived from rhyolitic ash, and relic shards and vol- canic minerals are common (Schultz, 1963). Clay in the Middle and Upper Jurassic Carmel Formation and Upper Jurassic Morrison Formation of the C010- rado Plateaus also is largely of volcanic-ash origin, and relic volcanic textures are widespread (Keller, 1962; Schultz and Wright, 1963). Far more volumi— nous still are the volcanic—ash clays (montmorillonite and mixed-layer clay) of the marine Upper Cre- taceous shales of the Colorado Plateaus and Great Plains (Schultz, 1965; Tourtelot and others, 1960; Leonard G. Schultz, oral commun., 1965). The volume of Upper Cretaceous strata of the western interior is approximately 4 million on km (Gilluly, 1963); of this total, perhaps one-fourth or even one-half is altered volcanic ash, to judge by detailed studies of the Pierre Shale and reconnaissance studies of other formations (Schultz, oral commun., 1965). Volcanoes were active throughout much of the Late Triassic and the Jurassic near the Pacific margin of North America, whereas known Cretaceous volcanoes were less widespread and were farther inland (Gilluly, 1965). Granitic rocks of Triassic age are uncommon, those of Late Jurassic and Early Cretaceous age are widespread, and those of Late Cretaceous age are abundant and form large parts of the great Sierra Nevada and Idaho batholiths (Gilluly, 1963, 1965; but see Kistler and others, 1965). Unroofed batho- liths may have provided much of the pyroclastic material preserved in the strata of the western in- terior, and the Late Cretaceous batholiths may have been the source of the greatest part of that material. The numerous middle Cretaceous bentonites of Wyom- ing came primarily from dacitic and quartz latitic sources farther west, presumably volcanoes above the Idaho batholith (Slaughter and Barley, 1965). TERTIARY PLUTONS OF CASCADE RANGE The Cascade Range of Washington is among the regions where middle or late Tertiary plutons are known to have broken through to the surface and pro- duced volcanic piles. Fiske, Hopson, and Waters BATHOLITHS SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY ClO $580320 Ban made 3 nainaaooum 32 33m mm.“ 5 own? 333 gabwhm 25 an.» mama 2583 .95 3.3 we v.8 woo» m: “Ea 5205.3 95. $8352 5325a: nogsom 95 £335 H8303 13950 and 336% gnawing: oEaEEELwEQI.v ”infirm wmmkmionzx on O? om ON OH O m 0“ _ _ _ _____ mug: om ow B o m, S miom 4528on oz< 19E“; :39? 9:96 9.: 58:8 2:80 coszLotcmzwtmoUm .323 3938 can .330 gamma .mBQEQEm 5:383 .0526 3 «53.5.8 62:9: Sun: 83 8Q Ex wkmai 3823.33 E OBSOQOIo: \CLDZCRAOUW mammacw 2052mm ucm BEBEQEN B 259: 88950 .535 $30.. \CLDZEZOQm‘Q QikZOU .8. 2522 m: Em 5:058 2: 2:2 3 SEQ .3395 $8 :2: 3:55: E9: wucmcEw‘. 3E9 m mwmflocm BESEEE ucm 065m 5:058 539% E9; 6 >332. cwmanou .596 Eng: :93 3 32:8 333 GEE. 9.5268 E 35350 982 LooE & fig mxuo. 5.85895 328 am. $69. o_:mu_o> fl wEou :3 8225 So. .o 932 9.2: 0:523 83:5 to 9:23 3 6 c2335 msoc_E=_o> 253.3 Ewes: 3 2:38 62 3 33¢ 635.3320 EELS .23. 353.3 mcScmanoum meB .89. chmo: 3 McEBwhm ucm *oo‘. *0 22a fisEfi>=>So ._oo 3 32:8 $2.30 $5.52 3 32:3 .333 _mE‘_oz .mzz >>wm THE NATURE OF BATI-IOLITHS (1963, p. 40—63) showed that the 250-sq-km Tatoosh pluton of granodiorite and allied rocks in the Mount Rainier area produced explosive eruptions through its roof. Rapid loss of heat and volatiles resulted in distinctive crystallization features—vertically alined vesicles, miarolitic vugs, explosively brecciated rocks, local granophyre, aphanite, and vitrophyre, and tran— sitions between pluton and volcanic plugs—~in the upper few hundred feet of the pluton in vent areas. Rapid crystallization of the entire pluton, despite its granitic texture, is indicated by the absence of flow banding and lineation: “there was no prolonged inter- val during which the material moved as a viscous crystal mush” (Fiske and others, 1963, p. 46). The much larger Snoqualmie Granodiorite batho- lith, farther north in the Washington Cascades, also deroofed itself explosively (R. E. Fuller, unpub. 1925 thesis, as cited by Fiske and others, 1963, p. 59). Pyroxene diorite in cupolas of the batholith grades upward through volcanic plugs into andesite flows and pyroclastics. Here also, resulting dehydration pro- duced chill efi'ects high in the batholith. Miocene plutons of quartz diorite and granodiorite crop out near five of the northern volcanoes of andesite and dacite in the Cascade Range, and their presence sug- gests that the volcanic magmas came from the plutons, which remained active at depth (Hopson and others, 1966). The Tertiary volcanic and intrusive rocks of the Cascades in part were erupted through pre—Cenozoic crystalline complexes but in part have probably formed where only oceanic crust existed in Mesozoic time. In northern Washington and in southern Ore- gon and northern California the Cascade igneous rocks cut and overlie the granitic and metamorphic rocks of the late Mesozoic orogens. These older rocks strike southeastward in the north and northeastward in the south so that northwestern Oregon and southwestern Washington are on the Pacific side of all exposed and projected pre-Cenozoic rocks (Carey, 1958, fig. 56; King, 1959, p. 161). We infer that there is no pre- Cenozoic continental crust within this tectonic em— bayment. The central Cascades (including Mount Rainier) and the Coast Ranges of most of Oregon and Washington have apparently been built upon oceanic crust as a Cenozoic addition to North America. ALEUTIAN ISLANDS The island arc of the Aleutians consists of young volcanoes built upon platforms of middle Tertiary and older submarine lavas and pyroclastics. Andesite dominates both suites, but basalt and dacite are com- mon. Many of the old rocks are variably spilitized, Cll but their prealteration compositions were identical to the young rocks. (The geology of the islands has been described by various authors, under the general title “Investigations of Alaskan Volcanoes,” in the many chapters of U.S. Geological Survey Bulletin 1028.) The older complexes are intruded by Tertiary gra- nitic rocks on a number of islands. These are limited to stocks and smaller masses except on Unalaska Island, where three small batholiths occur (Drewes and others, 1961, p. 610—634 and table 1). About 70 percent of the batholithic rocks consists of grano- diorite. The remainder ranges from quartz gabbro to light—colored quartz monzonite. The granodiorite is chemically the same as andesites and dacites of both older and younger volcanic sequences and presumably is but an intrusive manifestation of the same igneous activity. The Aleutian Islands extend across the North Pacific. No evidence requires that continental crust existed there before the onset of andesitic island—arc volcanism. The mechanism of generation of the magmas from an oceanic mantle is discussed in the second major section of this paper. TERTIARY IGNEOUS ROCKS OF COLORADO Numerous stocks of lower Tertiary granitic rocks dominated by quartz monzonite and granodiorite form a chain trending northeastward across western Colo- rado. The rocks intrude Precambrian plutonic rocks and the thin overlying Paleozoic and Mesozoic plat- form sedimentary rocks. The great middle (and late?) Tertiary volcanic pile of the San Juan Mountains of southwestern Colorado is formed largely of rocks equivalent in composition to quartz diorite and granodiorite. Large calderas and other magmatic collapse structures occur through- out a region of at least 7,000 sq km (Luedke and Bur- bank, 1963; Steven and Ratté, 1963; Thomas A. Steven, oral commun., 1965); Tertiary batholiths must lie hidden beneath the volcanic cover of this area. No data are available to suggest whether the granitic rocks are mostly in a floored complex above the pre-Tertiary rocks. Western Colorado was a stable platform region during late Precambrian(?), most of Paleozoic, and early Mesozoic times, but during Cretaceous and Ceno- zoic times it has been the site of moderate deforma- tion. Western Colorado now has the highest regional elevation in the conterminous United States, and has a thick crust, high heat flow, and low-velocity upper mantle. The heat flow and mantle velocity distinguish the region from the still-stable platform east of the Rocky Mountains, and presumably the abnormal heat flow and mantle velocity, and the crustal thickening. C12 are products of Cretaceous and younger changes and are related genetically to the deformation and igneous activity of western Colorado. A geosynclinal environ— ment is not essential for the generation of granitic magmas: the controlling events occur in the lower crust or upper mantle, not in the upper crust. TERTIARY IGNEOUS ROCKS OF BASIN AND RANGE PROVINCE Silicic volcanic rocks (mostly welded tuffs) and intrusive porphyries and granitic stocks of similar compositions formed throughout broad parts of the Basin and Range province of the Western States dur- ing middle Tertiary time. Presumably the intrusives fed the extrusives. Blank (1963) found a porphyry in southwestern Utah that broke through its roof after partial crystallization and produced quartz latite welded tuf‘fs. Gilluly (1932, p. 69) found that in the Oquirrh Range of central Utah, silicic stocks and porphyries cut the volcanic pile that probably was erupted from the same magma chambers, and similar relationships have been reported elsewhere; but in general, wallrocks of stocks consist of pre-Tertiary rocks. The Tertiary silicic volcanic rocks of Nevada and western Utah have a total volume of approximately 120,000 on km (Mackin, 1960, p. 83). This is equiva— lent to about half the volume of the Sierra Nevada batholith if we are correct in assigning a thickness of approximately 8 km to the batholith. Most of this Great Basin material was melted beneath a miogeo- synclinal terrane, and part of it was melted beneath an unstable platform environment. ST. FRANCOIS MOUNTAINS BATHOLITH The St. Francois Mountains of southeastern Mis- souri expose a Precambrian complex in which a roof of extrusive rhyolite and quartz latite was intruded by leucocratic granite, quartz monzonite, and grano- phyre (Bridge, 1930, p. 59—64; Dake, 1930, p. 26—44; Haworth, 1895; Hayes, 1961; Robertson, 1966; Snyder and Wagner, 1961). Granophyric and granitic rocks are about 1,300 million years old (Allen and others, 1959; Tilton and others, 1962). According to Ander- son (1962; oral commun., 1965), the roof rocks above the plutons of granophyre and granite are chiefly welded tuif and tuffaceous sedimentary rocks and are about 2 km thick. Despite their age and their intru- sion by granite, the volcanic rocks have generally low dips, and are broken only by minor normal faults which developed contemporaneously with volcanism. The volcanic rocks have not been metamorphosed regionally, although some have been weakly silicified and albitized (propylitized?), and they contain well— SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY preserved primary volcanic fabrics such as collapsed pumice lapilli and shards. Granophyre tends to lie with gently dipping contacts between volcanic rocks above and. coarse granitic rocks beneath, and therefore apparently represents a rapidly crystallized roof facies of the batholith. In bulk composition and com- positional variation, granitic and granophyric rocks are similar to the dominant volcanic rocks. Rocks of all three types are mostly red, and are moderately alkalic and low in alumina. Minor sodic dacites present low in the volcanic succession, however, are less silicic than are any of the exposed silicic intrusive rocks (Anderson, 1962). Sheets and irregular intru- sive masses of basalt and diabase cut both intrusive and extrusive silicic rocks. We infer the volcanic rocks to be consanguineous with the intrusive granitic and granophyric rocks. The petrological similarities are great, and the vol- canic rocks lack any evidence of regional deformation and metamorphism to suggest that they represent a wallrock terrane intruded by the batholith. If our inference is correct, then the batholith is roofed only by its own ejecta. The volcanic crust solidified from magma erupted from the molten interior of the com- plex. Granophyre crystallized first beneath the vol— canic crust, and its formation thickened the insulating crust beneath which the granitic rocks then solidified. The cover above the coarse plutonic rocks when they crystallized was probably not thicker than 3 km. The earliest volcanic rocks included types less silicic than the magmas which crystallized at the exposed high levels of the intrusive complex. The red color of most of the granite, granophyre, and rhyolite and their chemical character intermediate between normal calc-alkaline silicic rocks and the moderately alkalic silicic differentiates of lopoliths (Anderson, 1962; compare with Hamilton, 1960) might permit the inference that the St. Francois rocks are the silicic caprocks of a gabbroic lopolith, but gravity surveys indicate that the Precambrian com- plex in the St. Francois Mountains is batholithic rather than lopolithic. The Duluth and Mellen lopo- , liths of the Lake Superior region and the Wichita lopolith of Oklahoma (and the probable lopoliths forming a buried chain trending southwestward from Duluth through Minneapolis and Omaha to Abilene) are marked by great positive Bouguer gravity anoma— lies, with a relief of 100 mgals or more and steep gravity slopes. There is no such anomaly in the St. Francois region. (The lopolithic gravity anoma- lies are perhaps the most remarkable features shown on the United States gravity map by Woollard and Joesting, 1964.) THE NATURE OF BATHOLITHS NEW ENGLAND APPALACHIANS The Paleozoic orogenic terrane of the Appalachian system is best known in New England. Granitic rocks are widespread in this region but there are no great batholiths comparable to the late Mesozoic batholiths of western North America. Some reasons are sug- gested here for the contrasts between the Paleozoic and Mesozoic terranes of opposite sides of the continent. Successive episodes of geosynclinal sedimentation and volcanism, deformation, metamorphism, and gra- nitic intrusion have been superimposed complexly in New England, and many problems remain unsolved. Despite the overlapping of events, the major episodes of metamorphism and intrusion appear in a general way to be younger in the west than in the east and intermediate in the medial belt (fig. 5). Discussion /% V CONN ° 013 here is restricted to that medial belt, whose analysis appears particularly relevant to the topic of this report. We are particularly indebted to Wallace M. Cady and Robert H. Moench for discussions clarifying New England geology. Summary published reports include those of Billings (1956); Billings, Rodgers, and Thompson (1952); and Goldsmith (1964). The medial belt of New England is formed largely of sedimentary rocks and subordinate volcanic rocks, of Silurian and Devonian ages, which were metamor- phosed and widely intruded by granitic masses during the Devonian. In New Hampshire, the belt is a silli- manite-grade “plateau” of highly metamorphosed rocks, injected by granitic plutons which generally are concordant in both structure and mineralogy to the wallrocks, although structurally discordant granites also are present in the northeastern part of the belt. EXPLANATION Granitic rocks Mostly Devonian in New Hampshire and central belt of Maine; mostly Devonian and Carboniferous in Rhode Island, eastern Massachusetts, and south- eastern Maine NEW BRUNSWICK SILURIAN ANO MOSTLY DEVONIAN DEVONIAN AND CARBONIFEROUS Eugeosync ma] rocks Includes gneiss domes of Devonianfl)‘ Oliverian Plutonic' Series in New Hampshire Miogeosynclinal rocks CAMBRIAN AND ORDOVICIAN + + + + 4 + o v + Eugeosynclinal rocks } Basement plutonic rocks PRECAMBRIAN .J._L .l—L J_.L J—L Sillimanite isograd Ticks on side containing sillimanite - __'_. _‘_. _' _. Garnet isograd Dots on side containing garnet 100 MILES 100 KILOMETERS FIGURE 5.——Geologic and metamorphic map of New England. Generalized from Goldsmith (1964). 232—376—67—3 C14 The high-grade belt is bounded on both sides by steep metamorphic gradients (figs. 5 and 6) which cor— respond approximately in some places, although not in others, to the contact zones with older (Cambrian and Ordovician) eugeosynclinal rocks. Outside the metamorphic gradients are less metamorphosed rocks Whose pattern is much complicated by granitic intru- sions of various ages and by superimposed metamor- phisms which produced large areas of garnet—grade rocks but only small areas of sillimanite—grade rocks. The sillimanite plateau of New Hampshire gives way northeastward along the strike in Maine to a terrane of low-grade metamorphosed Silurian and Devonian rocks. These are intruded by crosscutting masses of granitic rocks, mostly Devonian, and middle- and high-grade metamorphism is limited to narrow contact aureoles surrounding the intrusions. The regional biotite isograd, not shown in figure 5, crosses Maine northeast of the garnet isograd shown; most of the medial belt in Maine consists of chlorite- zone rocks. Structures within the high-grade part of the New England medial belt are extremely complex. Litho— SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY logic units swirl, intersect, branch, and pinch out in patterns of great fluidity, and dips typically are gentle to moderate (Billings, 1955; Goldsmith, 1963, 1964). Structures are refolded isoclinally and irregularly (Goldsmith, 1961). Metasomatic changes and granitic injections are widespread. The metamorphic gradients bounding the sillimanitic terrane are locally so steep that biotite and sillimanite isograds are only about 1 km apart (fig. 6; Billings, 1955), although 8 km is a more common distance. Petrologists in general agree that the biotite isograd represents a temperature of 200° to 300°C, whereas sillimanite forms near 600° to 700°C, or at about the temperature of granitic magma. Biotite and silli- manite isograds thus represent a temperature differ- ence of about 400°C, and the temperature gradient in response to which they formed in New Hampshire was as steep as 300°C per km in the plane of the present ground surface. The common assumption that increas- ing metamorphism is due to increasing depth (for example, Turner, fig. 77, in Fyfe and others, 1958) is inadequate to explain such gradients (Hamilton, 1963a, p. 90). No model based upon heat conducted High-grade belt (sillimanite “plateau") of central New Hampshire ~/ \ L_.&EV_V_HAMPSHIRE MASSACHUSS‘EFTS 71° 0 5 MILES O 5 ’KlLOMETERS EXPLANATION Granitic rocks, mostly Devonian FIGURE 6.—Metamorphic zones and granitic rocks of a part of southeastern New Hampshire. Medium grade Low grade (garnet, staurolite) (biotite, chlorite) High grade (sillimanite) Metamorphic rocks, Ordovician to Devonian A steep metamorphic gradient, defined by the sillimanite, staurolite, garnet, and, in part, biotite isograds, separates high-grade and low-grade terranes. Adapted from Billings (1955). THE NATURE OF BATHOLITHS from deep levels in a static system can account for more than a small part of such metamorphism. The apparent alternative is that the heat was intro- duced in magmas from hotter deeper levels. (Barrell, 1921, reached this conclusion long ago.) This require— ment negates the speculation that the gneisses and migmatites of the sillimanite belt represent spontane- ous partial melting due to tectonic depression into levels heated to melting temperatures by heat con- ducted from the base of the crust. The granitic component added to the complex came from markedly greater depths than those exposed. Exposed granitic masses are distributed irregularly through the gneiss belt. Had they alone carried the heat upward, the metamorphic effects should be re— lated concentrically to them, but this is not the situa— tion. The belt is generally bounded by a straight and steep metamorphic gradient, irrespective of the local distribution of exposed plutons within the belt. Some might interpret these relationships as indicating the presence of a continuous great batholith beneath the exposed gneisses, but so steep are the flanking meta- morphic gradients that such a batholith would have to be very near the surface, and it is unreasonable to postulate that the top of the batholith could every- where be at a uniform shallow depth without being exposed. The lmown factors of structure and metamorphism agree with the hypothesis that the gneiss belt of the New England Appalachians formed beneath a batho- lith analogous to that of the Sierra Nevada. Molten plutons may have risen through the gneisses and coalesced above them in a thin, shallow batholith. Granitic melts are much lighter than metamorphic rocks—even solid granitic rocks are lighter than most high-grade metamorphic rocks—and must rise through them wherever the buoyancy of the plutons exceeds the strength of the metamorphic rocks. The extremely plastic and irregular deformation shown by the gneisses indicates that their structures did not form by simple response to systematic regional compression but rather that the rocks flowed about in complex patterns of rising, sinking, underflowing, and over- flowing. This in turn suggests that the dominant causes of flow were gravitational instability and dif- ferences in plasticity, as Rosenfeld (1960) concluded on the basis of detailed analysis of microscopic and macroscopic structures. The heating of the meta- morphic rocks and their copious injection by granitic fluids caused them to flow readily outward, downward, and beneath the rising plutons. Much of the granitic material was enveloped and trapped in the gneisses, but most of it rose through to coalesce higher into Cl5 batholiths, which have since been eroded away. The steep metamorphic gradients bounding the gneissic terrane may mark the margins of the region through which the plutons rose. If this rationale is correct, then the northeastern limit of the sillimanite plateau in the medial belt also represents the limit of initially nearly continuous gra- nitic plutons in the overlying batholithic complex that has since been largely eroded away. Crosscutting granitic masses are more abundant near the along- strike transition from high-grade to low-grade ter- ranes than they are anywhere else in the medial belt (fig. 5) ; this is consistent with the suggestion that a thin batholith, initially continuous to the southwest but since largely eroded away, here gave way north- eastward to a terrane of scattered smaller plutons. The lack of high—grade regional metamorphism in most of the medial belt in Maine is evidence that there is no batholith of regional extent beneath the belt there, and also suggests that there was never a batho- lith of regional extent above the levels now exposed by erosion. Four heat-flow measurements in the Devonian gra— nitic rocks indicated fluxes of only 1.2 to 1.7 micro- calories per square centimeter per second (Birch and Roy, 1965); considered with the probably high radio- activity of these rocks, this low heat flow indicates the granites to be thin. The northern Appalachian region was intruded dur- ing Mesozoic time by the stock and small batholiths of alkalic rocks of the White Mountains magma series. The relation between observed heat flow and measured radioactivity in these White Mountains rocks is such that rocks of high radioactivity, like those exposed at the surface, can extend only a few kilometers downward (Birch and Roy, 1965). EROSION INTERVALS Batholiths have crystallized and then been exposed by erosion within such short intervals of time in a number of places that it is diflicult to visualize the granitic rocks as having formed at depths of many kilometers. Only a small fraction of a geologic period separates rocks deformed and intruded by batholiths from beds resting unconformably upon the same batholiths. The Baja California batholith, for example, intruded upper Lower Cretaceous strata and is overlain unconformably by middle Upper Cre- taceous beds (Allen and others, 1960). Rocks as young as middle Upper Jurassic were metamorphosed and intruded by stocks and small batholiths in the Klamath Mountains of northwestern California, and the resulting crystalline complex is overlain uncon- formably by high Upper Jurassic beds (Irwin, 1960). C16 Zircon of middle and Late Cretaceous age, from batholithic rocks farther west, is abundant in sand— stone throughout the Upper Cretaceous sequence of Montana and Wyoming (Houston and Murphy, 1962, 1965), and thus shows regional exposure of the batho- liths to erosion within a very short interval after crystallization. STRONTIUM ISOTOPES Of the stable isotopes of strontium, some Sr87 is formed radiogenically by the decay of Rbs", whereas the remaining Sr87 and all Sr86 and Sr88 are nonradio- genic. The ratio of radiogenic to nonradiogenic strontium in a magma thus provides an index to the ratio of rubidium to strontium in the source rocks from which the magma was derived; and as these two elements behave quite differently chemically—the ru- bidium being generally associated with potassium, the strontium with calcium—valuable information can be SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY gained from this ratio. The strontium—isotope method of tracing geologic processes was developed by, among others, Hurley and his associates (Hurley and others, 1962), and has been applied by them (Fairbairn and others, 1964a, b; Hurley and others, 1965) and by many more, including Hedge and Walthall (1963) and Zartman (1965). In abundance, Sr87 is com— parable to Sr“, whereas Sr87 is far less abundant than SrSS; hence analyses are generally made for the ratio SIM/Sr“. The young basalts of Hawaii and Iceland have Sr87/Sr8’6 ratios ranging only from about 0.7023 to 0.7045 (fig. 7 ; Hedge and Walthall, 1963; Powell and others, 1965). The mantle sources from which their magmas were melted do not vary much in Sr/Rb ratios. Stony meteorites had an 81‘87/SI'86 ratio of about 0.698 at the time they formed (Hedge and Walt— hall, 1963; Pinson and others, 1965). No igneous rocks of any age yet analyzed have initial ratios that 0-73 l i l l ‘ Most Precambrian rocks have present ratios between 0.71 and 0.72 - - > 0.73. Rocks rich in K-feldspar Change in ratio of grano- diorite formed at 3 x 10‘3 years having Sim/Sr“: 0.702, Rb=0.01 percent. and Sr=0.03 percent INITIAL RATIO. Stall/Sr“ § I 0.70 —— Achondrite region of oceanic basal Mantle source 0.69 I l I Sierra Nevada and British Columbia t changes as Rb“ deca or mica may have much higher ratios; 0.8 is common granitic rocks Nova Scotia granitic rocks ~ Recent continental basalts and most rhyolites Young oceanic basalts (Hawaii and lceland) 5 4 FIGURE 7.——-—Initial ratio of Srs'l/Sr86 in Paleozoic and Meso- zoic batholithic granitic rocks. Because 81-87 includes that strontium produced by radiogenic decay of Rb“, and Sn‘16 and the remaining Srs" are not radiogenic, the ratio Sr37/ Sr86 is a function of the Rb/Sr ratio in the rocks from which the magmas were melted. The initial Sim/Sr86 ratios of granitic rocks, like those of most continental basalts and rhyolites, are a little higher than those inferred for the mantle source regions of oceanic basalts, but much 3 2 TIME BEFORE PRESENT, IN 109 YEARS lower than those of most Precambrian basement rocks. The granitic magmas might represent mantle melts which assimilated crustal rock, or melts of rock whose Sr/Rb ratio was intermediate between that of the mantle and that of Precambrian basement rocks. Data from Fair- bairn, Hurley, and Pinson (1964a, b), Hedge and Walthall (1963), Hurley, Bateman, Fairbairn, and Pinson (1965), Pinson, Schretzler, Beiser, Fairbairn, and Hurley (1965), and Powell, Faure, and Hurley (1965). THE NATURE OF fall as much as 0.001 below the line between the ratios of meteorites and the ratios of the lower limit of young basalts: initial ratios in all igneous rocks fall along the line or are above it by less than 0.005, or at the most 0.008. The slope of this line indicates an Sr/Rb ratio in the source region of oceanic basalt of about 50 :1, and this ratio is the virtual lower limit for the source region of all igneous rocks (Hedge and Walthall, 1963). . This Sr/Rb ratio of 50:1 is similar to, or a little lower than, that of tholeiite. Tholeiite must form by the melting of a large proportion of its mantle source rocks, which thus cannot be ultramafic. The Sr37/Sr86 ratios of Recent continental volcanic rocks overlap those of oceanic basalts and range upward only to about 0.71 (Hedge and Walthall, 1963). The ratios in continental rhyolites and basalts overlap throughout their ranges: the source rocks from which the basalt and rhyolite magmas are derived have about the same Sr/Rb ratios, and there is little basis in these data for the common assumption that basalt magmas are melted from deeper and much more mafic rocks than are rhyolite magmas. Precambrian basement rocks have in general mark- edly higher ratios of rubidium to strontium than have V the mantle source rocks of oceanic basalts, and hence the Sr87/Sr86 ratio of average continental rock has increased more rapidly during geologic time than has the ratio in the mantle. (A line in fig. 7 illustrates the changing ratio in a typical Precambrian grano- diorite whose Rb87 is decaying to Sr87.) Even average Precambrian rocks—markedly more mafic than those most likely to be partially melted to produce granitic magmas—have present Sr87/Sr86 ratios of about 0.720 (Faure and others, 1963). If silicic crustal rocks were melted directly to form any magmas, then the result- ing igneous rocks should instead have initial Sr’37/Sr86 ratios plotting high above the achondrite-oceanic basalt zone of figure 7. Granitic rocks crystallized from magmas whose ini- tial Sim/Sr86 ratios were slightly higher than those of the mantle source regions of oceanic basalts but markedly lower than those of Precambrian plutonic rocks (fig. 7). The initial Sr87/Sr86 ratio in the Paleo- zoic granitic rocks of Nova Scotia averaged 0.708 and only ranged from 0.705 to 0.711 (Fairbairn and others, 1964a), and ratios in the upper Mesozoic granitic rocks of southeastern British Columbia and the Sierra Nevada batholith were mostly from 0.706 to 0.708 (Fairbairn and others, 1964b; Hurley and others, 1965). The average ratio in Precambrian rocks in Paleozoic and Mesozoic time would have been at least 0.715, and many felsic Precambrian rocks would have had ratios of 0.8 to 1.0 and more. Cl7 BATHOLITI-IS Similar relationships hold for granitic rocks of all ages. Thus granitic rocks, 1,000 to 1,120 my. old, of the Llano Uplift of Texas had markedly lower Sr87/Sr86 ratios at the time of their crystallization than did their wallrocks (Zartman, 1965). Tertiary granitic rocks of Colorado and “Washington had initial Sr87/Sr86 ratios of about 0.705, and some Precambrian granitic rocks, 2,400 to 2,800 my. old, of the Pre— cambrian shield had ratios of only 0.701 to 0.703 (Hedge and Walthall, 1963). Some small Paleozoic granites also had very low ratios, such as 0.703 (Czamanske, 1965). As the various authors cited have emphasized, these ratios of radiogenic to nonradiogenic strontium in granitic rocks preclude the possibility that granitic magmas are wholly melted from mantle rock like that which yields oceanic basalts, and also preclude the possibility that they are wholly melted from silicic crustal rocks like those exposed in Precambrian ter- ranes. The source materials for granitic magmas are richer in rubidium (of which the radioactive isotope decays to Sr87) than are oceanic mantle rocks, but they are poorer in rubidium than are exposed crustal rocks. Obviously the magmas either are melted in an environ— ment of intermediate composition, or else are hybrid products 10f melts combined from contrasting rocks. The strontium-isotope data considered alone permit the speculations that granitic magmas are derived in the lower (intermediate) continental crust; that they represent mixed mantle and crustal materials; or that they are melted from eugeosynclinal volcanic rocks and volcanogenic sedimentary rocks. The eugeosyn- clinal-melting hypothesis is attractive for some gra- nitic rocks but obviously is not applicable to the Boulder batholith (which formed in a miogeosyncline) or to the granitic rocks of southeastern California, Colorado, and elsewhere, which were intruded into Precambrian basement complexes. The ratios between the various stable and radiogenic isotopes of lead and their parental uranium and thorium in granitic rocks provide another isotopic method for placing limits on the composition of the source rocks from which the magmas were melted. The lead relationships are being investigated by Bruce R. Doe (written commun., 1965), who finds that the lead isotopes generally support conclusions similar to those reached by the strontium—isotope researchers. ORIGIN AND EMPLACEMENT 0F BATHOLITHS The examples cited in this paper are all from the United States. The local interpretations made could be supported further by other North American exam- ples, and further yet by numerous examples from C18 other continents. The explanation of the examples is accordingly integrated here into a general theory of the characteristics of batholiths. THE ENVIRONMENT OF BATHOLITHS Most large Phanerozoic batholiths lie at least partly within eugeosynclines, and this has led to the popular tectogene hypothesis, whereby batholiths are visual- ized as the products of downbuckling and partial melting of the geosynclinal fillings. Batholiths also form, however, in such other environments as miogeo— synclines (for example, the Boulder batholith), un- stable platforms (as in the San Juan Mountains of Colorado), and oceanic island arcs (including the Aleutian Islands). The Idaho batholith intrudes both eugeosyncline and miogeosyncline. Boundaries be- tween eugeosynclinal and miogeosynclinal suites shift back and forth through time, and in California the boundaries are crossed at high angles by the Sierra Nevada batholith (Gilluly, 1963, 1965). Precambrian basement rocks also are intruded by batholiths, in geosynclinal as well as nongeosynclinal settings. In southern California in the San Gabriel and Orocopia Mountains (Crowell and Walker, 1962), in the West Riverside Mountains (Warren Hamilton, unpub. data, 1966), and elsewhere, large plutons of upper Mesozoic granitic rocks intrude Precambrian basement plutonic complexes as well as Paleozoic and Mesozoic metasedimentary and metavolcanic rocks. The common association of batholiths with eugeo- synclines seems best interpreted to indicate that the batholiths and the geosynclinal volcanic rocks have related origins, rather than that the batholiths form because the geosynclines are present. Batholiths form Wherever temperatures are high enough at depth to melt the needed magmas, and eugeosynclines are only one setting in which these conditions are met. ORIGIN OF GRANITIC MAGMAS The discussions in previous sections indicate to us that all or most granitic magmas are generated at depths below any ever exposed by erosion. Strontium— isotope data preclude the possibility that most gra- nitic (and silicic-volcanic) magmas could form by melting of silicic crustal rocks alone. Sources per- mitted by the isotope data include volcanic rocks and volcanic sedimentary rocks in geosynclines; gabbroic or amphibolitic rocks of the lower continental crust; and mantle rocks, provided some continental rock relatively high in rubidium is assimilated by rising magmas. Geologic, petrologic, and physiochemical reasoning provides much further information on the origin of the magmas. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Some granitic masses are so closely associated with high-alumina volcanic rocks that their magmas must have come from the same source. The granodiorite batholiths of the Cascade and Aleutian provinces share the compositional patterns of the volcanic rocks they intrude and of the volcanic rocks which overlie them, and appear to be but different aspects of con- tinuing magmatism from the mantle. Pregranitic deformation of the older volcanic rocks in these provinces was slight so that major orogeny cannot be postulated to have intervened. Some granitic magmas may form by melting of eugeo- synclinal materials, but many cannot possibly form in this way. In the Boulder batholith and the eastern part of the Idaho batholith, for example, no volcanic piles were available. In other batholiths (for exam- ple, in southern California) the magmas intruded subgeosynclinal basement rocks and hence came from still deeper sources. The rocks formed from such demonstrably noneugeosynclinal magmas cannot be distinguished from those (for example, in the Sierra Nevada) for which a eugeosynclinal origin might otherwise be postulated. The hypothesis that granitic magmas are generated by downbuckling and partial melting of eugeosynclinal materials is certainly inap- plicable in many places and may be Wholly invalid. If there is a single magma—generating mechanism, it must thus be more general than that of geosynclinal downbuckling. The mechanism must produce gra- nitic magmas beneath diverse geologic environments. Quartz diorite, granodiorite, and quartz monzonite are the dominant products required in Phanerozoic terranes, but their relative abundance differs widely from region to region. (In some Precambrian orogens, potassic granite is dominant or at least abundant, but it is lacking or is much subordinate to less potassic rocks in nearly all Phanerozoic suites.) Quartz diorite and trondhjemite are the dominant types in some regions, and gabbro and diorite are present in many assemblages and are abundant in some. Compositions inferred for the lower crust and upper mantle must be incorporated into the explanation. Seismic, heat—flow, and petrologic data provide clues. In the uppermost mantle, seismic waves generally travel faster beneath stable regions of the continental crust than beneath orogenically and volcanically active regions (Herrin and Taggart, 1962; Nuttli, 1963; Pakiser, 1963). Heat flow is in general higher in active regions than in stable ones (Lee and Uyeda, 1965). These relations preclude the possibility that peridotite forms the uppermost mantle in both active and stable regions: pressures in the uppermost mantle are too low to invert olivine (Wentorf, 1959), so that THE NATURE OF BATHOLITHS density-phase transformations cannot be called upon; and neither can the variable hydration of peridotite to serpentine be postulated, for hydration would be suppressed in the high—temperature regions, and hence would correlate with heat flow in the direction oppo- site to that required by the seismic data. Pressure and temperature at the base of the continental crust in stable regions, but probably not in active ones, are appropriate for the eclogite transformation of plagio— clase and low-alumina pyroxene to the denser phases of garnet and aluminous alkalic pyroxene (Yoder and Tilley, 1962, fig. 43). Velocity data determined experi— mentally at high pressure (Birch, 1960) permit the suggestion that the Mohorovicic discontinuity at the base of the continental crust represents a composi- tional change, from basalt above to basalt plus dunite below, and that in the low-velocity mantle and lower- most crust of active regions the basaltic phase is in low-density plagioclase and pyroxene or amphibole. In the high-velocity mantle, and probably also in the deepest crust, of stable regions the basaltic phase is in high—density pyroxene and garnet, but in the low- velocity mantle of tectonically active regions the trans- formation to those dense eclogitic minerals occurs well Within the mantle (Pakiser, 1965). Another possi- bility is that high-velocity mantle contains more olivine than does low-velocity mantle. Such a con- trast might develop as basalt is wrung out of the upper mantle, residual materials being progressively more peridotitic and dunitic. Kimberlite and its dense inclusions provide direct evidence for the existence of mantle rocks of basalt- plus-olivine composition. Kimberlite consists of large corroded crystals of high-pressure minerals—forsterite, jadeitic clinopyroxene and aluminous orthopyroxene, pyrope and almandine, phlogopite, diamond, mag- nesioilmenite—in a matrix of serpentine, chlorite, oli- vine, calcite, and phlogopite. Most kimberlite con- tains inclusions of dense rocks composed of varying combinations and proportions of heavy minerals like those of the corroded crystals (Holmes, 1937; O’Hara and Mercy, 1963; Wagner, 1914; Williams, 1932). Kimberlite has a peculiar bimodal composition of ultramafic components on the one hand and alkalic and volatile components on the other, and is probably a mixture of mantle rock—like the dense inclusions—- with mantle-derived water, carbon dioxide, and alka- lies. The average composition of the dense inclu- sions, and of kimberlite itself minus the 15 or 20 percent of apparently added volatile and alkalic com- ponents, is approximately equal to a mixture of equal parts of tholeiitic basalt (in the mineralogic form of C19 high-pressure pyroxenes and garnet) and magnesian olivine. If these inferences are correct, then the upper mantle and lower crust difl'er primarily in olivine content. As olivine is more refractory than any other major mineral in the rocks, it would not be melted appreciably; therefore, only the melting of the basaltic phase need be considered here, and such dis- cussion may be equally applicable to upper mantle and lower crust. High-pressure laboratory data show that the dense alkali-bearing aluminous pyroxenes and the garnet of basalt-composition pyroxenite and eclogite melt to- gether over a rather narrow temperature range under both anhydrous and hydrous conditions at pressures appropriate to the upper mantle or deep continental crust (fig. 8; Cohen and others, 1966; Yoder and Tilley, 1962, tables 43—47). Tholeiite (basalt low in alkalies, moderately low in alumina, and saturated in silica) and olivine-tholeiite magma could originate by the nearly complete fusion, within such a narrow temperature range, of the pyroxenitic or eclogitic fraction of mantle rock. Partial melting under such conditions to produce granitic magmas is not however possible. No substantial quantity of silicic magma can be generated by direct melting at depths greater than that of the gabbro-to—eclogite pressure transformation if the meltable part of the mantle is of basaltic composition. At pressures corresponding to the uppermost few miles of the crust, pyroxene and plagioclase melt or crystallize together over a limited temperature range 40 l l l I 100— w o l Eclogite cpx+gar Transition zone cpx+p|ag+gar cpx+plag+ gar+|iquid 50-— Liquid PRESSURE. IN KILOBARS S 8 DEPTH, IN Kl LOM ETERS plag+cpx+liquid G a b b ro cpx+p|ag 0 1 1 i 400 600 800 1000 1200 TEMPERATURE, IN “C l I 1400 1600 1800 FIGURE 8.—Pressure-temperature phase relationships in ma- terial of anhydrous basaltic composition. The data plotted are those of Cohen, Ito, and Kennedy (1966), with addi- tions consistent with the data of Yoder and Tilley (1962). Olivine or orthopyroxene or both would be present in rocks of appropriate compositions—the liquidus and solidus curves are in fact those of olivine tholeiite—but are not illustrated here. 0px, clinopyroxene; gar, garnet; plag, plagioclase. C20 (Yoder and Tilley, 1962, figs. 6, 27—30). Crustal tem— peratures, however, preclude melting at such shallow depths. The melting behavior of basaltic rock at pressures appropriate to the lower crust or uppermost mantle, but shallower than the gabbro-eclogite transition, is quite different. Plagioclase at these intermediate pres- sures is a markedly lower temperature mineral than is either pyroxene or amphibole (Cohen and others, 1966; Yoder and Tilley, 1962, figs. 27-30). The tem- perature range of crystallization or melting is broad and increases with pressure. Progressive melting under these shallower-than-eclogite conditions will thus produce highly feldspathic magmas progressively richer in calcic plagioclase and in mafic minerals. Such melting could produce the high-alumina basalt, andesite, and dacite which typify island arcs and eugeosynclines (Hamilton, 1964). This mechanism might provide an explanation for the melting of intermediate and silicic magmas from rock of basaltic composition and thus from the lower crust or mantle. Mantle rocks of the same composi- tion that yield tholeiitic basalt magma by partial melting beneath the eclogite transformation could yield high-alumina basalt, andesite, and dacite magma by less complete melting above it. (High—alumina basalt difi'ers from tholeiite primarily in being richer in the components of calcic plagioclase.) The inti- mate association of tholeiite and high-alumina rocks in some island arcs and also in some continental environments can be explained easily in this way. The contrast between midoceanic volcanism (tholeiite on the sea floor, tholeiite plus alkaline olivine basalt on islands) and island-arc volcanism (tholeiite plus high-alumina volcanic rocks) can then be explained in terms of variations in pressure-temperature rela- tions in the mantle beneath the contrasted provinces, without any need for variations in mantle composi- tions. The curves of actual temperature and of the melting temperature of the basaltic component are probably close together and subparallel within a con— siderable thickness of the upper mantle (MacDonald, 1964, fig. 15), so that such variations in conditions need not be great. Differentiation by partial crystallization of a magma within the crust or upper mantle, but above the eclogite transformation boundary, would produce liquids of the same compositions as would partial melting there, as O’Hara (1965) emphasized. The composition of a magma provides clues as to the depth of its last equilibration with crystals, but not as to whether this equilibration was achieved by par— tial melting or by partial crystallization. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY These mechanisms of deep derivation of magmas by varying degrees of melting satisfy some of the conditions but not all imposed by the data. How, for example, do voluminous granitic magmas form that are almost devoid of mafic components? Why are silicic magmas nearly lacking in the ocean basins except in those places—as, island arcs, and Iceland-— Where abnormally thick volcanic crust is present? Regional relationships yield further data to be con- sidered. Available data suggest parallel regional variations in compositions of eugeosynclinal volcanic rocks and the granitic rocks which intrude them. Magmatism of the type that produced the volcanic piles could also have produced the batholiths. The eastern part of the Sierra Nevada batholith consists largely of light-colored quartz monzonite and granodiorite, and Mesozoic metavolcanic rocks are mostly dacite and quartz latite (Rinehart and Ross, 1964, p. 30—38; Bateman and others, 1963, p. 6). The western part of the batholith consists in general of more mafic and calcic rocks, and the Mesozoic volcanic rocks are cor- respondingly also more mafic and calcic on the average (Bateman and others, 1963, p. 6). Similarly, meta- volcanic rocks along the west side of the Idaho bath- olith are basalt, andesite, and dacite, all low in potas- sium, and the granitic rocks intrusive into them are quartz diorite and trondhjemite (Hamilton, 1963a, table 5 and pls. 1, 2). Metavolcanic rocks are lacking in the eastern part of the batholith, which is of granodiorite and quartz monzonite. The western part of the belt of late Mesozoic bath- oliths of western North America is dominated by quartz diorite, and the central and eastern parts are dominated by granodiorite and quartz monzonite (Moore, 1959, fig. 2). Silica and potassium are higher inland, whereas calcium, iron, and magnesium are higher oceanward. One possible explanation of this difl'erence is that when the batholiths were forming the Mohorovicic discontinuity lay deeper beneath the inland region than beneath the coastal one and that the inland magmas were melted largely within the lower crust whereas the coastal ones came from the upper mantle. An inference proceeding from this postulate and from factors noted previously is that although the lower crust and the meltable part of the upper mantle are both grossly basaltic in composi- tion, the lower crust is richer in potassium and is less mafic. A similar postulate can be made for the melting of magmas during middle and late Cenozoic time throughout the Western United States. The ratio of potassium to sodium in high—alumina rocks of any THE NATURE OF BATHOLITHS given silica content. tends to increase with altitude and hence with crustal thickness. This is demon- strated by the data plotted by Moore (1962, figs. 2, 3) : the K/Na ratio tends to increase as Bouguer gravity becomes increasingly negative.2 (Bouguer gravity in general correlates broadly with regional surface altitude and with crustal thickness.) The bulk com- position of the high-alumina volcanic suites also tends to become more silicic with increasing altitude. This can be seen by considering the many provincial plots of Moore’s figure 1 as frequency-distribution diagrams. These relationships can be explained if the magmas are melted entirely in the mantle where the Mohorovicic discontinuity is relatively shallow and partially or entirely within the lower crust where it is deep. Gradations in magma types reflect grada- tions in crustal thickness and hence might be due either to melting together of both crustal and mantle materials in various proportions or to assimilation of deep-crust material in magmas rising from the mantle. Gilluly (1965, p. 28) made similar suggestions. A mechanism which would permit generation of a magma in the mantle, followed by great assimilation of crustal materials in the rising magma, and which is capable of producing wholly leucocratic rocks, would appear to satisfy all the requirements set forth. Such a mechanism is available if we combine Dickson’s concept of zone melting with that of partial melting within the upper mantle and lower crust. Part of the rise of any magma must be accom- plished by zone melting (Dickson, 1958), and this could result in profound modification of the original magma by contamination. The pressure gradient to which any magma of appreciable vertical extent is subjected must cause migration of the most volatile components toward the top of the chamber. This results in a lowering of the melting temperature at the top and a raising of the melting temperature at the bottom, so that roof rocks are melted and incor- porated into the melt, whereas the basal magma is forced to crystallize. The rising magma becomes progressively enriched in the lowest melting compo— nents of both the initial magma and the rocks through which it passes. The energy needed to melt the roof comes from crystallization at the base. The energy loss as cold rocks are heated is partly compensated for by the lowering of temperatures of fusion as pressure decreases with rise of magma to higher 2 Moore plotted high-alumina and other provinces together. The re- lation between the K/Na ratio and gravity anomalies becomes more regular when only high-alumina provinces are considered, and so prov- inces of tholeiite, olivine basalt, basalt and rhyolite, and highly alka- line rocks—which all may [originate beneath the ecloglte boundary— are logically omitted. 232—376—67—4 C21 levels, and by the rise of magma through a graded or layered crust which becomes less refractory upward. The final magma resulting from such a rise by zone melting could contain only a trifling quantity of material present in the initial melt: an energy envelope has risen through the crust, but only the most volatile of the initial components are present in the final magma. Magma can be mobilized beneath the crust, and yet the final high—level igneous rock can consist largely of components derived from the crust. The correlations between compositions of gra- nitic rocks and the compositions of the columns through which they rose can thus be easily explained, as can the leucocratic character of many granitic rocks and the lead and strontium isotopic relationships in them. The effects of differentiation, both by fractional crystallization and by upward migration of the less refractory components, must further complicate the evolution of the magmas. Zone melting of course represents a combination of assimilation and differ- entiation. The mechanisms suggested here require that the crust be continually growing as material is added to it from the mantle. The general restriction of potas— sic batholiths to terranes more than 1,000 my. old perhaps indicates that the mantle and crust became by that time so differentiated that potassium-rich magmas could no longer be generated in great volume from the mantle. Silicic and high-alumina magmas form primarily in long belts, initially either oceanic or continental, whose tectonic mechanisms are not yet apparent; but the magmas can form anywhere in the high-heat-flow regions of the continents. EMPLACEMENT 0F BATHOLITHS Batholith magmas are melted in the lower crust and upper mantle. The buoyant magma masses rise and probably become completely detached from their zones of melting. Overlying rocks are heated and displaced outward, then sink and flow beneath the rising plutons, becoming intensely metasomatized and injected by granitic material. Much additional gran- itic material is enveloped by the flowing gneisses, joins in their irregular motion, and crystallizes in concordant foliated sheets. Such interpretations have been made by other geol- ogists. Hans Cloos (1923) concluded that many bath- oliths are connected with deep sources by dikelike channels rather than by full-size chambers and that concordance and discordance are determined by the structure of the country rock. Chamberlin and Link (1927) suggested that batholiths are shaped like sheets, tongues, and mushrooms, which have narrow 022 feeders and spread out high in the crust. Lane (1931, p. 823) wrote of batholiths as “intrusions on the surface”—flat mushrooms, reaching the surface in some places and spreading out between basement and overlying sediments in other places, fed by relatively narrow conduits. Bott and Smithson (1966) con- cluded from gravity analyses that granitic plutons extend typically to depths of only about 10 km. The plutons of magma rise until crystallization, forced by loss of heat and volatiles, prohibits further flow, or until they reach the surface. As the plutons stiffen, heating of the wallrocks decreases, metasom- atism virtually stops, and thin aureoles of thermal metamorphism result. Wallrock screens are dragged upward and pushed outward by rising and expanding plutons, and roofs are raised. The deformation of the metamorphic rocks may be due largely to em- placement of the plutons. The larger and more numerous the plutons, the higher they should rise. Many must reach the surface and produce voluminous volcanic eruptions. Any pluton that reaches the surface over a wide area is roofed only by its own volcanic ejecta, and granitic textures develop as the magma crystallizes beneath its insulating volcanic cover. Large batholiths are formed by the coalescence of many plutons and may be less than 10 km thick and may be unroofed over broad areas. The plutons rise a minimum of perhaps 40 km before reaching their site of final crystallization, which is generally within a few kilometers of the surface. The loss of volatiles by eruption through roof rocks must stop the rise of many plutons, but it is difficult to visualize any mechanism which could prevent the broad surface breakthrough of many other large masses of light magma. ' The exposed parts of the Boulder and St. Francois batholiths are covered almost entirely by consanguin— eous volcanic rocks, and such a relationship may be common. Volcanic crusts may float on broad plutons of batholithic magma. The Boulder batholith probably is floored by a downbowed surface of prevolcanic rocks. This shape is indicated by the way the flanking rocks dip right side up beneath the batholith on the north and prob- ably on the south; the shape is also indicated by the horizontal flow shown by internal structure in the west part of the batholith, although the west contact is hidden. The sagging presumably accompanied the withdrawal of magma at depth and its eruption at the surface. Near both east and west margins of the batholith, the volcanic roof rocks dip monoclinally outward, perhaps in response to lifting by the grow- SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY ing batholith. The final major eruptions of magma through the preigneous rocks apparently were primar— ily intrusive; the first eruptions were certainly vol- canic. Extrusion through the volcanic crust and in- trusion beneath it probably occurred simultaneously during much of the history of the complex: new injections of magma beneath the crust fed eruptions at the surface so that crust and batholith thickened simultaneously, although the volcanic rocks spread far beyond the batholith. The sharp contacts between some plutons in the batholith show sporadic intrusion and crystallization and it cannot yet be said how much of the batholith was molten at any one time. The size of the vents through the prevol'canic rocks, which fed the shallow batholith and its volcanic cover, is not known, but the great area of the complex suggests that they were large. The central part of the Sierra Nevada batholith is flanked by steeply dipping metamorphic rocks whose tops face inward toward the batholith. If this struc- ture is related to the batholith, rather than being an old structure fortuitously intruded by the granitic magmas, then conceivably it has an explanation sim- ilar to the structure flanking the Boulder batholith. The Sierran plutons may also have reached the surface and spread laterally as the wallrocks were pushed outward and downward. If this is so, the volcanic cover once above the plutons has been removed by erosion. The overlapping ages of granitic rocks and wallrocks complicate any interpretation, and some of the metavolcanic wallrocks may have formed as roof rocks of early plutons. The largest Phanerozoic batholiths now exposed are of late Mesozoic age. This factpsuggests either that this was a period unique in the earth’s history or else that batholiths are features of the uppermost crust and are so thin that 200 million years’ erosion com- monly suffices to remove them. We prefer the latter alternative. Most gneiss terranes may have formed beneath batholiths since removed by erosion, and the steep metamorphic gradients bounding some gneiss terranes may mark the limits of the belts through which the plutons rose. There are, however, many large Paleozoic and Precambrian batholiths, so either such batholiths were less eroded than their largest correlatives, or else they extended to greater depths than we infer. The rise of plutons can be likened to that of salt domes. Magma and salt both rise because they are lighter than the rocks above them, although the den— sity differential is probably in general greater for magma and metamorphosed wallrocks than for salt and indurated sedimentary rocks. Isolated salt plugs THE NATURE OF BATHOLITHS form where the supply of salt is small, whereas salt megadikes form where the supply is great. The varied salt structures demonstrated by drilling in northern Germany and described by Trusheim (1960) are strikingly similar geometrically to various Late Cretaceous and early Tertiary igneous complexes in western Montana. An inference to be drawn from the salt—dome analogy is that plutons of magma do not mark the locations of hot spots in the mantle, but rather that magma is generated over broad regions at depth and coalesces into masses whose spacing is controlled by the supply of magma and whose posi- tion and shape are controlled by structural features of the crust. Batholiths form where the supply of magma is so great that the masses coalesce and rise toward the surface as large plugs and megadikes. BATHOLITHS AND METAMORPHISM The deformation and heating shown by meta— morphic rocks may be largely products rather than causes of magmatism. The steep metamorphic grad— ients of many terranes of regionally metamorphosed rocks cannot be quantitatively explained as due to conducted geothermal heat, but can be easily under- stood as produced by contact metamorphism on a regional scale. Gneisses and migmatites, ascribed by many writers to anatexis (partial melting in place due to ultrametamorphism), may instead be produced largely by metasomatism and injection by rising plutons. The highly plastic flow patterns of gneisses may form while wallrocks sink and close beneath rising plutons. The contortions in low—grade meta- morphic rocks flanking high—level batholiths also may have been produced largely by rising plutons. Similar suggest-ions have been made by many authors. Sederholm (1919, p. 250) wrote of “Kon- taktmetamorphose regionaler.” Barrell (1921, p. 255) concluded that “batholithic invasion is * * * one of two major factors in dynamometamorphism.” Bud- dington and Chapin (1929, p. 50) used the phrase “contact metamorphism on a regional scale.” The metamorphic rocks of the east-central Sierra Nevada include mineral assemblages typical of the contact-metamorphic hornblende hornfels facies and the regional-metamorphic. almandine amphibolite fac- ies of Turner (in Fyfe and others, 1958). Bateman, Clark, Huber, Moore, and Rinehart (1963, p. D 11) assumed that this combination indicates that the rocks formed at a depth whereat the conditions of contact metamorphism were giving way downward, because of increasing pressure, to conditions of regional meta- morphism; and they accepted Turner’s (fig. 107 in Fyfe and others, 1958) guess that this transition C23 might occur at a pressure corresponding to a depth of about 20 km and specified this as the depth of crystallization of the Sierran plutons. (Interpreta— tions based on granitic rock compositions led Bateman and others to infer a water pressure of 5 kbars, equiv- alent to about the same depth.) The contrasting metamorphic mineral assemblages, however, occur in rocks of different compositions, and we regard this as evidence that the mineralogical contrasts reflect com- positional diiferences rather than some critical pres- sure and that the distinction made between contact and regional metamorphism has no meaning here. The metamorphic grade and degree of recrystalliza- tion of the wallrocks decrease away from granitic contacts (Bateman and others, 1963, p. D 11), and the regional—type metamorphic rocks clearly owe the heat of their metamorphism to the nearby intrusive plutons. The pressure assigned the almandine am- phibolite facies by Turner was based on the assump- tion that rocks of that facies were heated by con- ducted geothermal heat. If such an assumption is accepted, great depth must be postulated even if a steep thermal gradient is also assumed; but as this heat demonstrably came from intruded magmas, and not from deep burial, the assumption is not valid. Hamilton (1963a, p. 89—93) raised numerous objec- tions to the depth-zone correlations and to the over- simplified applications of facies concepts that are still prevalent in the literature of metamorphic petrology. Not all medium- and high-grade metamorphism can be explained as regional effects of batholithic in- trusion, however. The Blue Ridge province of the southern Appalachians, for example, comprises a meta— morphic plateau of general staurolite- or kyanite-zone rocks with local highs of sillimanite-zone rocks (for example, Bryant, 1962, p. 20—23). Upper Precam- brian rocks were metamorphosed progressively to these grades during early or middle Paleozoic time, when the older Precambrian basement rocks (which make up most of the Blue Ridge) were metamor- phosed retrogressively to the same grades. The meta- morphic slope bounding the plateau is rather gentle: the biotite and kyanite isograds are separated by an average distance of 20 km in the Great Smoky Moun- tains (Hadley and Goldsmith, 1963. pl. 3). Known Paleozoic igneous activity in the province is limited to small dikes and to large and small pegmatites in scattered areas. Inasmuch as effects of metasomatism are limited to these areas, passage of plutons upward through the entire exposed terrane cannot reasonably be postulated, nor is there evidence to suggest the presence of a hidden batholith beneath the entire province. Severe deformation accompanied the meta- C24 morphism, and it is likely that the temperatures of deep burial were much increased by deformational heating. BATHOLITHS AND THRUST FAULTING Many thrust faults are temporally and spatially related to batholiths. The relationship might be ascribed to various factors, but one that must be considered in each such relationship is that the in- trusion of a batholith is responsible for the thrusting. Keith (1923, p. 365—375), among others, has proposed such mechanisms, which are worthy of serious evalua- tion even though they are not currently popular. The belt of late Mesozoic batholiths of the Western United States is flanked on both sides by broad ter- ranes deformed by thrust faults during late Mesozoic and very early Cenozoic time. On the east is the Laramide thrust belt, whose east margin is coincident with that of geosynclinal strata and trends south- southeastward through western Montana, southward across westernmost Wyoming, south—southwestward across western Utah and southern Nevada, and inter- sects the crystalline terrane in southeastern California. The belt widens again in southern Arizona and trends into Mexico. The distance of the east margin from the Idaho, Sierra Nevada, and related batholiths ranges from 50 km in southeastern California to as much as 500 km across Nevada and western Utah. (The maximum width of the belt may have been less than 300 km when it formed, however, for the wide sector may have undergone great extension dur- ing the formation of basin-and-range structures and of volcanic piles during Cenozoic time.) The geosyn- clinal assemblage is shingled by thrust faults east- ward from the edge of the batholiths (although inter- pretation is clouded by the widespread late Paleozoic thrust faults present at least in central Nevada). Major thrust faults generally are nearly parallel to bedding of the rocks. Faults in the west and central parts of the belt tend to be subhorizontal or gently folded, and many carry younger strata over older strata, whereas major faults in the east generally dip westward and carry older rocks over younger. Thrust— ing probably preceded folding in most regions, and most of the folding is better regarded as an effect of the sliding that produced the thrusting than as a cause of thrusting. At least in western Wyoming and southeastern Idaho, thrusts in general become younger eastward within the belt (Rubey and Hubbert, 1959, p. 188). The thrust-belt rocks are not metamorphosed except close to the batholiths on the west. Individual faults have horizontal displacements reaching at least 40 km, as proved by older—over-younger exposures in southeasternmost British Columbia, and the aggregate SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY displacement across the belt may reach several times this amount. Basement rocks do not appear in the overthrust sheets. The lack of basement rocks in the overthrust sheets, despite great displacements, requires that the exposed faults are confined to the geosynclinal fill and do not reach the basement. It has been inferred from this relationship, in the Western States and elsewhere, that the strata slid off some highland arch—but the thrust-faulted terrane extends westward to the bath- olithic belt, so such a hypothetical highland cannot be placed anywhere east of the batholiths. Possible explanations are that the blanket slid eastward off the batholiths themselves or that the rising and spread- ing batholiths shouldered the sedimentary rocks east— ward. The lack of basement rocks in the thrust sheets would require that, in the latter explanation, the batholiths expanded eastward above the basement rocks. The thrust belt is particularly irregular in the‘ region of the Boulder batholith and neighboring small batholiths which lie within the belt in southwestern Montana. (See the geologic map of Montana: Ross and others, 1955.) The east margin of the belt swings from a southwestward trend to a southeast- ward trend to define a curve rudely concentric to the Boulder batholith and 50 km or less from it. Trends of thrust faults within the belt are very irregular because early structures have been deformed by later ones which have a tendency toward concentricity about the batholith. Jumbled thrusts trend north- ward a similar distance west of the batholith. One possible explanation is that shouldering by the bath- olith produced much of the thrusting. As the bath- olith thickened, rocks of its inward-sloping floor would be forced plastically outward. Another alter- native is that once formed, the stiff plate of the bath- olith greatly influenced the response of the sedimen- tary terrane to thrusting due to other causes. Both processes may have operated together. We interpret published geologic maps of south- western Montana as indicating that a relatively simple system of east-directed thrust faults had been formed, perhaps by the shouldering action of the Idaho bath- olith, before the eruption of the Boulder batholith and its neighbors at the end of Cretaceous time. The younger batholiths superimposed complex local pat— terns of outward thrusting upon this eastward-thrust terrane. We interpret the structure of the northern Flint Creek Range (McGill, 1965), for instance, as recording the superposition of generations of folds and thrusts in this sequence: (1) Eastward thrusting from the direction of the Idaho batholith to the west, THE NATURE OF BATHOLITHS (2) westward thrusting from the Boulder batholith to the east, and (3) northward crowding by the small Philipsburg and Royal batholiths to the south (see Mutch and McGill, 1962). Thrusting west of the batholiths is of comparable extent but difierent type. Great Late Cretaceous overthrusts carry variably metamorphosed eugeosyn- clinal materials, intruded by stocks and small bath— oliths during Cretaceous(?) and Jurassic, and possi— bly older, orogenies, westward over similar rocks and over little-metamorphosed Upper Jurassic and Cre- taceous deposits. Such overthrusts occur throughout the Klamath Mountains of northwestern California and southwestern Oregon (Irwin, 1964) and in west- ern Idaho (Hamilton, 1963b), and may be present in northwestern Washington (Hamilton, 1963b; Misch, 1952, and Barksdale, 1960, make other interpreta- tions). South of the Klamath Mountains, throughout the length of the Coast Ranges of California, an eastern facies of miogeosynclinal Cretaceous and Upper Jurassic strata has apparently been thrust westward over the eugeosynclinal Franciscan facies of rocks of the same age, the fault extending into the basement rocks beneath the eastern facies (Brown, 1964; Irwin, 1964). Displacements on single faults may exceed 100 km in the Coast Ranges and in VVash- ington. Sheets of serpentine and peridotite lie along many of the faults. , These western thrust faults, broke the basement rocks and tapped sources of ultramafic material, and therefore are very different from the thrusts east of the batholiths. The age of major thrusting approx— imately coincides with the age of major plutonism, so the two processes may have been related genetically, although it is difficult to visualize a satisfactory mechanism linking them. There was much thrusting of sedimentary rocks at least in Nevada during late Paleozoic and early Meso- zoic intervals. No batholiths of corresponding age are known in the region (Gilluly, 1965, p. 23). Great thrust faults can form without assistance from bath- oliths. BATHOLITHS AND YOUNGER STRUCTURE Batholiths, once formed, influence subsequent de- formation because of their mechanical strength. The largest unbroken mountain masses in the Western States—the Sierra Nevada and the mountains of cen- tral Idaho—are carved from those parts of the bath- olithic belt which contain the highest proportion of granitic rocks and the smallest proportion of meta- morphic rocks. The fault zone bounding the Sierra Nevada block on the east cuts obliquely through the C25 batholithic belt, which curves northeastward from the northern Sierra and southeastward from the southern Sierra, but the fault zone approximates the eastern limit of almost continuous granitic rocks within the belt. The crystalline belt contains a higher propor— tion of metamorphic rocks along strike to both the northeast and southeast on the east side of the fault zone, where it is broken into numerous large and small basin-and-range blocks. Similarly Baja California and the Peninsular Ranges of southwestern California are composed of almost continuous batholithic rocks, which here form the west part of the belt; the east part, in southeastern California and western main- land Mexico, contains abundant metamorphic rocks, and is broken into many fault blocks. The large basin—and-range fault blocks striking northwestward into central Idaho abruptly lose structural relief and vanish as they intersect the Idaho batholith. Other fault blocks lie north, west, and east of the batholith, but no large ones break it (Hamilton, 1962). The Coast Range of British Columbia is another massive mountain block consisting largely of batholithic rocks. The history of batholithic mountains indicates that the presence of batholiths markedly influences uplift and erosion. Consider the Sierra Nevada batholith, which was sufficiently raised within Late Cretaceous time that its roof rocks (whether mostly volcanic, as we suggest, or metamorphic, as most geologists would infer) were eroded away before Eocene time. Moun- tains were still present in early Tertiary time, for middle Eocene strata in the foothills lap upon a bedrock surface whose local relief was at least 350 m (Allen, 1929, p. 382), and very coarse Eocene and Oligocene river boulder gravels are preserved in fossil valleys throughout the northern Sierra (for exam- ple, Lindgren, 1911, pls. 4, 21). The west base of the mountain block has remained near sea level through— out Tertiary time, and marine and nonmarine de- posits of various ages lap onto it. The rest of the block has apparently undergone both subsidence and uplift within the Tertiary but not profound erosion. The widespread preservation of lower Tertiary de- posits in the northern part of the Sierra shows that ridge surfaces there are not far from the early Ter- tiary surface, although there has been subsequent uplift and tilting. Lower Tertiary deposits are lack- ing in the central and southern Sierra, but presum- ably there also middle and late Cenozoic erosion has been limited largely to incision. During middle Ter- tiary time the Sierra Nevada was relatively low, and Miocene floras of the northern Sierra crest region indicate an altitude much lower than the present one (Axelrod, 1957). By the end of the Pliocene, C26 local relief in the crest region reached at least 2,000 in (Dalrymple, 1964), so its altitude must have been as high or nearly as high as it is now. Much of the faulting that outlines the present Sierra block has occurred within late Pliocene and Pleistocene time (Dalrymple, 1964), and thus may have consisted pri— marily of the downdropping of blocks (such as Lake Tahoe and Owens Valley) on the east rather than of the upraising of the Sierra block itself. The late Cenozoic uplift of the Sierra Nevada must be related to the deep downbulge in the Mo- horovicic discontinuity beneath it. The range is high because the root is deep. Had this root formed at the same time as the batholith, the range would long ago have reached maximum altitude by rising iso- statically; instead, therefore, much of the downbulge must be a late Cenozoic development. (Christensen, 1966, reached the same conclusion.) Sierran silicic batholithic rocks have a high rate of production of radioactive heat, which increases markedly with po- tassium content (W’ollenberg and Smith, 1964), and perhaps this property is responsible for the root growth and uplift. The highest part of the 650-km- long Sierra Nevada is the 300-km segment, between Sonora Pass and Olancha, in which quartz monzo- nite and silicic granodiorite form most of the east- ern part of the block. The lower crest north and south of this segment, and the middle-altitude region west of the high segment of the crest, consist largely of granodiorite, and the western foothills consist mostly of quartz diorite and metamorphic rocks. Radioactive decay would retard cooling, particularly of the crestal silicic granodiorite and quartz monzo- nite; thus a partial barrier to conduction of heat from the mantle would be formed, and heating of the lower crust and upper mantle would result. Movement of crust or mantle materials, change in pressure-phase mineralogy or hydration state, or par- tial melting might have resulted from this deep heat- ing and so enlarged the Sierra root. Other late Mesozoic batholiths also tend to form mountain masses now standing higher than the sur- rounding nonbatholithic terranes, and the highest re— gions within the batholithic mountains tend to be those with the most silicic and potassic rocks. In the Coast Range of British Columbia, for example, the inland belt of quartz monzonite and granodiorite stands on the average about 1 km' higher than the coastal belt of quartz diorite. Regional or global events unrelated to the batho— liths may of course have caused such observed rela- tionships. It does appear, however, that batholiths are uplifted selectively two or more times, and if so SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY they must also be eroded selectively more than neigh- boring terranes. Batholiths perhaps cause their own destruction. The general scarcity of great pre- Mesozoic batholiths is explicable in such terms. 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Lindgren, Waldemar, 1911, The Tertiary gravels of the Sierra Nevada of California: U.S. Geol. Survey Prof. Paper 73, 226 p. Luedke, R. G., and Burbank, W. S., 1963, Tertiary volcanic stratigraphy in the western San Juan Mountains, Colo- rado, in Short papers in geology and hydrology: U.S. Geol. Survey Prof. Paper 475—0, p. 039—044. Luth, W. C., Jahns, R. H., and Tuttle, D. F., 1964, The granite system at pressures of 4 to 7 kilobars: Jour. Geophys. Research, v. 69, no. 4, p. 759—773. Macdonald, G. A., 1941, Geology of the western Sierra Nevada between the Kings and San Joaquin Rivers, California: California Univ. Dept. Geol. Sci. Bull., v. 69, p. 161—178. MacDonald, G. J. F., 1964, Dependence of the surface heat flow on the radioactivity of the earth: Jour. Geophys. Re- search, v. 69, no. 14, p. 2933—2946. Soc. America THE NATURE OF BATHOLITHS McGill, G. E., 1965, Tectonics of the northern Flint Creek Range, in Fields, R. 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R., 1960, Strati— graphic variations in mineralogy and chemical composi- tion of the Pierre Shale in South Dakota and adjacent parts of North Dakota, Nebraska, Wyoming, and Mon- tana, in Short papers in the geological sciences: U.S. Geol. Survey Prof. Paper 400—B, p. B447—B452. Trusheim, F., 1960, Mechanism of salt migration in northern Germany: Am. Assoc. Petroleum Geologists Bull., v. 44, no. 9, p. 1519—1540. Wagner, P. A., 1914, The diamond fields of southern Africa: Johannesburg, Transvaal Leader, 347 p. Wentorf, R. N., J r., 1959, Condensed systems at high pressures and temperatures: Jour. Phys. Chemistry, v. 63, p. 1934— 1940. Williams, A. F., 1932, The genesis of the diamond: London, Ernest Benn, v. 2, 636 p. Wollenberg, H. A., and Smith, A. R., 1964, Radioactivity and radiogenic heat in Sierra Nevada plutons: Jour. Geophys. Research, v. 69, no. 16, p. 3471—3478. Woollard, G. P., and Joesting, H. R., 1964, Bouguer gravity anomaly map of the United States: US. Geol. Survey. Yoder, H. S., Jr., and Tilley, C. E., 1962, Origin of basalt magmas; an experimental study of natural and synthetic rock systems: Jour. Petrology, v. 1, no. 3, p. 342-532. Zartman, R. E., 1965, Rubidium-strontium age of some meta- morphic rocks from the Llano Uplift, Texas: J our. Petrol- ogy, v. 6, no. 1, p. 28—36. NOTE The discussion of the Idaho batholith in this report (p. 05—06) was based in part upon the interpretation by Ross (1934) that the Casto Volcanics was intruded by the Cretaceous batholith. This interpretation was disproved during the 1966 field season, when all rocks assigned by Ross to the Casto west of the Middle Fork of the Salmon River were mapped by Frederick W. Cater, Warren Hamilton, Benjamin F. Leonard 3d, Raymond L. Parker, and Edwin V. Post. These rocks were found to be that part of the Challis Volcanics (of Eocene age as dated by Axelrod T) which has been altered or converted to hornfels by contact metamorphism by a Tertiary batholith of granite and quartz mon— zonite. (Ross recognized that the young batholith was Tertiary, but he did not see its contacts with the volcanic rocks and did not recognize that fresh Challis grades into the altered rocks which he called Casto.) The “Casto Volcanics” apparently does not exist in the sense intended by Ross. This finding negates the argument in this paper that the Casto might be part of the extrusive cover beneath which the Idaho batholith crystallized—but it is Wholly in accord with the general concepts developed here, for the quartz latite welded tuffs of the Challis form the roof beneath which the Tertiary batholith crystallized. The tuffs presumably formed largely as the ejecta of that batholith. 3‘ Axelrod, D. L., 1966, Potassium-argon ages of some western Tertiary floras : Am. Jour. Sci., v. 264, p. 491—506. 0 Cy‘ IE '7 f 7 DAY P (a if f 55’ ‘71 1D Cenozom Volcanlc Rocks of the Devils Postpile Quadrangle, Eastern Sierra Nevada California GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—D Prepared in cooperation wit/z tne California Division of Mines ana’ Geology OCT 5 1957 ’74)]. ’7' SCIENCE my) CENOZOIC VOLCANIC ROCKS OF THE DEVILS POSTPILE QUADRANGLE, EASTERN SIERRA NEVADA CALIFORNIA Photograph by Gerhard Schumacher. Postpile. Devils 6 h t t a e H». S e d n a .m g .n .U .m 0 :J r a n m u .I. 0 c d e p O 1 e V e d u. m Cenozoic Volcanic Rocks of the Devils Postpile Quadrangle, Eastern Sierra Nevada California By N. KING HUBER and C. DEAN RINEHART SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOLOGICAL SURVEY PROFESSIONAL PAPER 554~D PrepareaI iii cooperation wit/z t/ze California Division of Mines and Geology 'UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1967 UNITED STATES DEPARTMENT OF THE INTERIOR STEWART L. UDALL, Secretary GEOLOGICAL SURVEY William T. Pecora, Director For sale by the Superintendent of Documents, US. Government Printing Office Washington, D.C. 20402 CONTENTS Page Tufif of Beds Meadow—Continued Abstract ___________________________________________ D1 Zonation in the tuft—Continued Page Introduction ——————————————————————————————————————— 1 Upper flow unit ____________________________ D11 Petrologic and Chemical data ————————————————————————— 2 Lower zone of partial welding ____________ 11 Microscopic petrography _________________________ 2 Zone of intense welding ___________________ 11 Chemical data __________________________________ 2 Thickness and lateral extent _____________________ 12 Classification ___________________________________ 5 Age, source, and correlations _____________________ 12 Andesite of Deadman Pass ___________________________ 6 Andesite Of the DeVilS Postpile _______________________ 13 Quartz latite of Two Teats ___________________________ 9 Quartz latite of Mammoth Mountain __________________ 14 Andesite from Deadman—Glass Creeks area ____________ 9 Andesite from Dry Creek area ------------------------ 15 AndeSIte of Pumlce Butte ____________________________ 16 Tuif of Reds Meadow _______________________________ 10 , , , . , , Olivme-bearlng quartz latite _________________________ 16 Zonation m the tut? """""""""""""""" 10 Basalt of the Red Cones __________________________ ’___ 16 Lower flow unit ____________________________ 10 Rhyolite ___________________________________________ 17 Lower zone of no welding ________________ 10 Pumice ____________________________________________ 19 Lower zone of partial welding ____________ 10 Thermal activity ___________________________________ 19 Zone of intense welding __________________ 10 Summary __________________________________________ 20 Upper zone of partial welding ____________ 11 References _________________________________________ 20 ILLUSTRATIONS Page FRONTISPIECE. Photograph showing well-developed columnar jointing in andesite at the Devils Postpile. PLATE 1. Geologic map showing Cenozoic volcanic rocks of the Devils Postpile quadrangle, Sierra Nevada, Calif __ In pocket FIGURE 1. Index map of the Devils Postpile quadrangle__-_ D2 2. Variation diagram of the Mammoth Lakes volcanic suite showing classification according to the “alkali- lime index’ ’ _____________________________ 4 . “Differentiation trend” diagram for the Mammoth Lakes volcanic suite _______________________________ 6 7 3 4. Silica-refractive-index diagram for volcanic rocks of the Mammoth Lakes area _________________________ 5. Diagram showing range in percentage of silica of some of the mapped volcanic units as inferred from refrac— tive indices of fused samples _________________________________________________________________ 7 6. Photograph of basal part of quartz latite of Two Teats on east side of San Joaquin Mountain ____________ 9 7, 8. Photomicrographs of tuif of Reds Meadow— 7. Lower flow unit __________________________________________________________________________ 11 8. Upper flow unit __________________________________________________________________________ 12 9. West-east section in Vicinity of Rainbow Falls showing reconstructed surface of the tufl" of Reds Meadow--- 13 10-13. Photographs of— 10. Polygonal joint pattern displayed by tops of columns in the Devils Postpile ____________________ 14 11. Northernmost of the two Red Cones as viewed from the southern cone_ - - - - - _ _ _ _ - _ _ _- --------- 17 12. Rhyolite dome south of Deadman Creek as seen from Deer Mountain ------------------------- 17 13. Upper surface of rhyolite dome south of Glass Creek showing typical irregular blocky surface_-_- 17 14. Vertical aerial photographs of Deer Mountain area, mounted as a stereopair ___________________________ 18 15. Photograph showing view of Inyo Craters from Deer Mountain -------------------------------------- 19 TABLES Page TABLE 1. Ages of Cenozoic volcanic rocks of the Devils Postpile quadrangle ____________________________________ D1 2. Petrographic summary of the volcanic rocks ------------------------------------------------------- 3 3. Chemical analyses and norms of the volcanic rocks __________________________________________________ 5 V SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY CENOZOIC VOLCANIC ROCKS OF THE DEVILS POSTPILE QUADRANGLE, EASTERN SIERRA NEVADA, CALIFORNIA By N. KING HUBER and C. DEAN RINEHART AB STRACT Cenozoic volcanic rocks of the Devils Postpile quadrangle are of late Pliocene to Recent age and are divided into 11 map units. The suite is alkalic-calcic and ranges in composition from basalt to rhyolite. It includes a rhyolitic welded ash-fl0w tufi which is probably correlative with the Bishop Tuff, although the two units are geographically isolated by the Sierra Nevada drainage divide. The Devils Postpile itself is a classic example of co- lumnar jointing in the lower part of a lava flow. INTRODUCTION The Devils Postpile quadrangle straddles the crest of the Sierra Nevada southeast of Yosemite National Park and contains the Devils Postpile National Monument ( fig. 1). The pre-Cenozoic rocks of the quadrangle in— clude Paleozoic and Mesozoic metasedimentary and met- avolcanic rocks. The metamorphic rocks have been in- truded by various granitoid igneous rocks, which form a part of the composite Sierra Nevada batliolitli (Bate— man and others, 1963; Huber and Rinehart, 1965a) . By late Tertiary time the Sierra Nevada had attained ap— proximately its present overall configuration, except for later modifications caused by faulting, uplift, and in— creased dissection. Regional volcanism began during the late Tertiary and has continued spasmodically to the present. The Cenozoic volcanic rocks of the Devils Postpile quadrangle resulted from several types of eruption throughout late Tertiary and Quaternary time; the forms of eruption include domes, lava flows, ash flows. and extensive pumice falls. The volcanic rocks range in composition from basalt to rhyolite, the oldest being ba- saltic to andesitic and the youngest, rhyolitic; however, most mafic and felsic rocks alternate haphazardly with- out regard to stratigraphic position. For this reason and because many of the volcanic units remain only as scattered erosional remnants, correlations are difficult. Interpretation of the volcanic history has been greatly assisted by potassium-argon age dating. Mapping in the Devils Postpile quadrangle (Huber and Rinehart, 1965a) and in the Mount Merrison quad- rangle to the east (Rinehart and Ross, 1964) allows the division of the Cenozoic volcanic rocks of the Devils Postpile quadrangle into 11 mappable units. Scattered erosional remnants of uncertain stratigraphic position have tentatively been assigned ‘to these units, although some of the remnants may represent additional episodes of volcanic activity not otherwise recognized. Table 1 lists the 11 units and their radiometric ages Where avail- able, and plate 1 shows their distribution. TABLE 1.——-Ages of Cenozoic volcanic rocks of the Devils Postpile quadrangle [See text for source and other data regarding individual age determinations] Potassium-argon age determination 1 Age and unit (million years) Recent: Rhyolite _________________________________________ Basalt oi the Red Cones ......... Pleistocene or Recent: Olivine-bearing quartz latite .......... Pleistocene: Andesite of Pumice Butte _____________ ._ Andesitc from Dry Creek area ............. Quartz latite of Mammoth Mountain-.. Andesite of the Devils Postpile ____________ Tufi oi Reds Meadow ______________________ Pliocene or Pleistocene: Andesite from Deadman-Glass Creeks area ____________________________________ Pliocene: Quartz latite of Two Teats _______________________ Andesite of Deadman Pass- _____________________ ‘ _ 0.37=e0.o4; 0.18:|:0.09 0.94zs0.15; 0.63:t0.35 0.66:|:0.04; (1.1); (1.4) 3.0:l:0.1 3.1:l:O.l; 3.3:t0.1; 3.5:i:0.1 1 The plus-or-minus figure that accompanies these age determinations is not an smartest? baseman::aarzasasaasass-123332 taminated with granite; youngest age is accepted as most reliable. In order to provide a background for description of the individual volcanic units, a general summary of petrologic and chemical data is presented for the entire volcanic suite. Following this, field relationships and general characteristics of each volcanic unit are de-.. scribed in order of decreasing age, rat-her than by pet.- rologic type, to provide a clearer picture of the sequence of events. We wish to thank G. B. Dalrymple for making his potassium-argon age determinations available to us in advance of their being published and Gerhard Schu— macher for the use of some of his excellent photographs of the Mammoth Lakes region. We also express our D1 D2 V \ \ SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY l. , l / L‘ E // Hv/ k DEVILS POSTPILE ’ \ NATIONAL \ / . MONUMENT f] I ALPINE ? \\ I \ I I \ 5 \ 0 l /—\ \ \ ’ \ V \ Area of \ ‘ d \/ \ 5 \\ g % m \ \ ® a» ‘ . \\ / Sonora Pass ‘ \\ AI 9 4f . \ t 62 t Brldgeport \\ \ \ 1\ I o\ J ‘ S‘ ‘Y 0 [mg/era; 6&0 (I f \|\ 4431/9 6);? ” \ 0’ ‘1 TUOLUMNE / g/‘- e x .N \\(1 \ ow \ \a x / I \ f / x \ \ YOSEMITE \ \ \ \ i Tioga Pass ® a r“\ . x i L 9 Mono Ssazfix : : Craters . \ /‘ V \x ( DEVILS j/I\\\\ '/ \ \ ' — / POSTPILE MOUNT MORRISON /, \ ‘~—— kr/ PARK , QUADRANGLE \ OUADRANGLE [I Benton \\> / \ \l "\_, . , L l/ / \L \/\ \ °ng l/ [\(\ \ al/e Casa Diablo II; \ II Hot Spnrigs/z,’ \‘. I / ’ \‘$\ ’9 \~~\, I __\// IItIIIIIIl | | / / MARIPOSA MAMMOTH LAKE CROWLEY / / L Xe LAKES ( x / ~ v Toms Place _/ ‘ L t a . V / \ I5 I{ 37°30’ \ . _.=l // 119nm, ' MADERA F // I _ _ _I / / FRESNO ,/o 10 20 30 MILES > INYO | FIGURE 1.—Location of the Devils Postpile quadrangle and the adjacent Mount Morrison quadrangle. appreciation to R. J. J anda for many stimulating dis- cussions of the interrelations of volcanic and glacial events. PETROLOGIC AND CHEMICAL DATA MICROSCOPIC PETROGRAPHY Data obtained from the petrographic study of the volcanic rocks are recorded in table 2. Percentages of constituents are based upon visual estimates obtained from thin—section study and are little more than rough approximations. The percentages, however, can be ef- fectively used to distinguish major and minor constitu— ents and to reveal the large proportion of indeterminate material occult in the groundmass of most of the rocks. Minerals listed in groups of two or more are given in order of decreasing abundance. In most specimens - range in anorthite content of the plagioclase was deter- mined from twinning relationships. Olivine is typi— cally magnesium rich. CHEMICAL DATA Table 3 presents chemical analyses and norms for seven samples of volcanic rock from the Devils Postpile quadrangle and one from the Mount Morrison quad— rangle. CENOZOIC VOLCANIC ROCKS, DEVILS POSTPILE QUADRANGLE, CALIF. TABLE 2.——-Petrographic summary of the volcanic rocks [An, anorthite; n, refractive index; Fa, fayalite] D3 Phenocrysts Matrix Volcanic unit Color Texture Remarks (Goddard, Average 1948) Mineral range in Percent Material Percent size (mm) Plagioclase 1—5 See re- Glass and impal- See re- (An’b’io). marks. pable dust, marks. Material is dominantly obsidian and minor feldspar, pumice. Where phenocrysts are locally microlites. abundant (as at Deer Mountain), they Light gray to Holohyaline Sanidine. 2—10 occur in approximately the following Rhyolite. black. to vitro— Quartz. 0. 5-2 ratio—plagioclase:sanidinezquartumafic phyric. Biotite. <0.5 minerals: 612:1:1. The n of fused obsid- Oxyhornblende. <0.5 ian from small dome north of Deadman Granular opaque <0. 5 Creek is 1.490. minerals. Plagioclase 1-5 5-15 Piagioclase and 0—75 (Anso-so). potassium feld- Plagioclase phenocrysts are commonly Vitrophyric to spar. zoned and contain irregular patches of Quartz latite Light gray to pilotaxitic, Biotite, oxyhorn- 0.1—2 1-5 Glass and impal- 5-90 glass; potassium feldspar in matrix is of Mammoth nearly dominantly blende, clinopy- pable dust. indicated by stain tests; '12 of natural Mountain. black. hyalopilitic. roxene. granular glass is 1.493—1.500, rock commonly shows opaque min- flow banding. ' erals, minor apatite. Plagioclase 1-5 15—30 Plagioclase. 15-70 (Alias-45)- Quartz latite Light gray to Biotite, oxyhorn- 0.1—1 <5—10 Glass and opaque 15—70 Most plagioclase phenocrysts are embayed, of Two dark or Hyalopilitic. blende, granular dust. riddled with glass. and zoned; made Teats. purplish opaque min- phenocrysts are generally strongly oxi- gray. erais, minor dized and otherwise altered. A apatite and zircon. Plagiociase 2—5 20 -30 Plagioclase, glass, 60—70 (Ants—so)- and minor Plagioclase phenocrysts are commonly Olivine- Hyalopilitic opaque dust. riddled with glass disposed in somewhat bearing Dark gray. to inter- Sanidine. 2-5 5—10 vermicular pattern and have some re- quartz latite. granular. Biotite, oxyhorn- 0.1-0.5 5 verse zoning of plagioclase. Olivine has blende, clinopy- a composition about Fazo according to roxene, orthopy< optical data. roxene, olivine. Many phenocrysts are commonly frag- mented, showing some preserved crystal Variable with Sanidine. 0. 2-2 5-10 Glass and impal- 70~90 faces; exotic rock fragments are common; position Quartz. 0.2—2 = pable dust; degree of welding and formation of eu- within flow; Plagioclase 0. 2—1 = minor devitri- taxitic structure is variable; n of pumice T113 0‘ RBdS light to Vitroclastic. (Ame-25). , fication locally. from base is 1.497; phenocrysts make up Meadow. dark gray Biotite. granular 0. 1—0. 5 <1 10—30 percent of the rock with average of to reddish opaque min- 20 percent. Potassium-sodium ratio in brown. erals. sanidine is about 2:1, as estimated from partial chemical and X-ray difiraction data. Plagioclase 0. 1—3 <5—20 Plagioclase. 50-75 (Anus—70). Olivine and pyroxene are moderately to Andesite from Medium light Trachytic to Olivine. 0.1—1 5—10 Olivine and clino- 10-20 strongly altered to black opaque mate- Dry Creek gray to me- intergranu- pyroxene. rial; olivine is composed of about Faro-20 area. dium dark lar. Clinopyroxene. <0. 1—0. 5 <5 Magnetite and 10—20 according to optical data and also altered gray. other opaque to iddingsite(?) or bowlingite(?). minerals, giass(?). Olivine. 0.15-2 3—5 Plagioclase 50—75 Only about half a dozen highly altered An(as-55). plagioclase phenocrysts occur in thin Andesite from Medium light Clinopyroxene. <0.01-0. 13 5—10 Olivine and clino- 10-25 section; otherwise, alteration in rock is Deadman- gray to me- Pilotaxitic. pyroxene. restricted to rims of iddingsite(?) or Glass Creeks dium dark Magnetite and 5—10 bowlingite(?) around olivine. Sufficient area. gray. minor opaque magnetite is present to affect hand mag- minerals. net. Olivine is composed of about Fazo according to optical data. 249-700 0 - 67 - 2 D4 SHORTER CONTRIBUTIONS TO GENERAL, GEOLOGY TABLE 2.—Pelrographic summary of the volcanic rocks—Continued [An, anorthite; n, refractive index; Fa, fayalite] Phenocrysts Matrix Volcanic unit Color Texture Remarks (Goddard, Average 1948) Mineral range in Percent Material Percent size (mm) ‘ In specimens with appreciable glass, pla- gloclase phenocrysts are typically em- Plagioclase 0.1—3 10—20 Plagioclase. 20—50 bayed and riddled with matrix material; (Ana—m). olivine is composed of about Fain-20 ae- Andesite oi Medium light Merocrystal- Clinopyroxene, 0. 05—0. 3 2—5 Olivine, pyroxene, 30—70 cording to optical data. Specimens from VPunu'ce gray to line to olivine. impalpable small knobs at north edge of outcrop area Butte dark gray. trachytic. dust, opaque appear to be more latitlc than average, a minerals, and conclusion based upon data from fused glass. glass beads; they also contain up to 5 per- cent euhedral plates and wedges of opaque material apparently altered from hornblende. Rock in'unit as mapped ranges from basalt to latite (48—66 percent SiOe) with ande< Medium light Plagioclase 0. 2-5 1—5 Plagioclase. 40—75 site predominant; plagioclase phenocrysts Andesite of the gray to Trachytic to (Ann—70). in andesitic composition range are Anss Devils Post- in dium intergranu- Clinopyroxene, 0.1—2 <5-25 Olivine, pyroxene, 20—50 to A1155; plagioclase and pyroxene pheno- pile. dark gray. lar. olivine. opaque min- crysts are commonly zoned. Olivine, erals, and im- being composed of about Faio—zo accord- palpable dust. ing to optical data, is locally altered to iddingsite(?) or bowlingite(?). Rock is predominantly andesitic but in- cludes some basalt. Percentage of oli- Plagioclase. 50—75 vine and Clinopyroxene decreases with Andesite of Medium light Olivine. 0.1—2 10_25 Olivine, pyroxene, 20—25 increasing silica. One specimen, with Deadman gray to Pilotaxitic to Clinopyroxene. 0.1—1.5 and granular the highest silica content as estimated Pass. medium trachytic. opaque min- from fused beads, contains both ortho- dark gray. erals. and Clinopyroxene. Olivine, being com- posed of about Fazo according to optical data, is locally altered to iddingsite(?) or bowlingite(?). Plagioclase 0. 5-3 20—25 Plagioclase, oli- 70 Basalt of the Medium gray Merocrystal- (Anus-7o). vine, pyroxene, Clinopyroxene and olivine are generally Red Cones. to medium line. Olivine. 0.2—2 <5 granular opaque unaltered, with the composition of oli- dark gray. Clinopyroxene. 0.2—2 <5 minerals, and vine being about Fazo according to glass. optical data. While only the Cenozoic volcanic rocks of the Devils (CLASS'F'CAT'ON 0" PEACOCK) Postpile quadrangle are described in this report, the Alkanc E Acl'a‘f‘c'lf : alfa'ii'c E Calcic volcanic rocks of both the Devils P‘ostpile and Mount 0 (51, l (56, (61) Morrison quadrangles (an area we call the Mammoth 3, I § 9, 8 ,3 :2 . . | a cu a) no -‘ H .4 Lakes area) overlap in time and space and are here con- (0 10 _ 2 g :3 2 2‘ ' ,5 2'; 2‘ 5' ' - . o o w 0 in srdered part of a consanguineous suite which has alka- E C', i o' +K2° ° . . . . . x 110 affinities, at least in the mafic rocks of the suite. The 0 ' . . . . | nature of this suite is illustrated by a. Peacock (1931) § 5 _ i variation diagram (fig. 2). The alkali-lime index, ap- ’2 : . . . . . Lu I prOXimately 54, places the suite in the alkalic-calcm g I . . . LLJ type. Alkalic affinities also are well shown at the mafic a o 50 l 60 70 end of the series, Where the Rittmann p values are less (54) PERCENTAGE 3,02 than 55 (alkaline <55 <10'11 mole per g (average of two analyses); radiogenic Ar”, 56 and 62 percent for the two argon extractions. CENOZOIC VOLCANIC ROCKS, DEVILS POSTPILE QUADRANGLE, CALIF. D13 W s 0 V2 1 MILE E .3 L_‘_‘_1_1_L___J___|_‘_|_J . in Mammoth Summlt .2 is crest Meadow § 6 § 5 is a 10,000’ g fig 10,000, 9200. Q ES ------ 9200' 8400’ _____________________________________________________________________________________________ 8400’ 7600, llllllll 7600, 6800’ 6800’ [Mill Andesite of the Devils Postpile Tuff of Reds Meadow E PrevTertiary granitic rocks FIGURE 9.—West-east section in vicinity of Rainbow Falls showing reconstructed surface of the tuff of Beds Meadow. present, was preceded by the Matuyama reversed- polarity epoch from approximately 2.5 to 0.7 million years ago (Dalrymple, Doell, and Cox, 1965). This suggests that the tuff of Reds Meadow is not appreciably older than 0.7 million years and could well be the same age as the Bishop Tuif. Cobbles of gray ash-flow tuff and pumice pebbles are common in alluvial deposits at Friant, where the ‘San Joaquin River emerges from the foothills of the Sierra Nevada and flows onto the alluvial plain of the San Joaquin Valley. The tuff pebbles are similar to ma- ' terial from the t-uff of Reds Meadow, although some petrographic and chemical differences do exist; no other source of similar tuff is known in the San Joaquin drain- age basin. A potassium-argon age determination on sanidine phenocrysts from pumice pebbles associated with the tuff pebbles gave and age of 0.60:0.02 mil- lion years (J anda, 1965), which is compatible with the age determined for the. tuft of Reds Meadow, N0 source for the tuff of Reds Meadow was found within the present drainage basin of the Middle Fork of the San Joaquin River. Possible vents could be hid— den beneath the younger rocks in the Mammoth Moun- tain area or farther northeast. Past areal contiguity with the Bishop Tuff is a possibility, but one for which direct evidence is lacking. Glacial deposits lie beneath the Bishop Tuff (Put— nam, 1960, 1962; VVahrhaftig, 1965; Sharp, 1965) ; hence if the tuff of Reds Meadow is equivalent to the Bishop Tuff, it also should overlie glacial debris. Be— cause the base of the tuff of Reds Meadow is nowhere exposed, we do not know whether or not the surface on which it rests was glaciated. Indirect evidence suggests pretufl' glaciation, 110w- ever. Numerous pebbles and cobbles of older andesite and granitic and metamorphic rocks from the head- waters of the Middle Fork (Hulber and Rinehart, 1965a) occur in the unconsolidated ash deposit at the base of the tufi near Sotcher Lake. The pebbles are rounded and must have been present as gravels in the Middle Fork valley at the time of eruption of the tuft. R. J. J anda (oral commun, 1965), in his study of the stream regimen of the upper San Joaquin basin, has concluded that these pebbles are too heterogeneous and too large to have been transported to this site 'by modern stream processes; hence the pebbles may represent in- corporated glacial materials. Prior to the eruption of the andesite of the Devils Postpile, a minimum volume of 4 cubic miles of the tufi' had probably been removed from the Middle Fork val- ley during an interval of a few hundred thousand years at most. This feat would seem to require some process, presumably glaciation, in addition to normal stream erosion. Owens River, about 15 miles east of the quad- rangle, receives drainage from a 25-mile segment of the eastern part of the Sierra Nevada, yet in a much longer time has only been able to cut a narrow, steep-walled gorge into the Bishop ’lfiiff. Thus we favor glacial ero- sion as the agent of removal of the tuff of Reds Meadow prior to the eruption of the andesite of the Devils Post- pile; the eruption probably occurred about half a mil- lion years ago. ANDESITE OF THE DEVILS POST'PILE Andesite and basalt are exposed on Mammoth Pass, in the Mammoth Lakes basin, and in the valley of the Middle Fork of the San Joaquin River. This unit is here referred to as the andesite of the Devils Postpile. Two of the three analyzed samples (table 3) are trachy- basalt, and the third is trachyandesite; but inasmuch as data from fused glass beads (fig. 5) give an average composition of approximately 53 percent silica, the map unit is called andesite. Basalt predominates in the bottom of the Middle Fork valley and in the Mammoth Lakes basin; andesite is pre- D14 dominant elsewhere, except in the southern part of the outcrop area where more siliceous quartz latite occurs. A potassium-argon age determination of 0.94:0.16 million years was reported by Dalrymple (1964b) for the andesite of the Devils Postpile. A redetermination on a second split of the same sample yielded an age of 0.63:0.35 million years.2 Because the amount of potas— sium in the sample and the percentage of radiogenic argon relative to total argon are so low in both of these determinations, they indicate little more than that the age of the andesite probably is between a quarter of a million and a million years. It is thus possible to recon- cile these data with an age of approximately 0.7 million years for the tufl’ of Reds Meadow, which the andesite unconformably overlies. The andesite was erupted in large part in the Mam— moth Pass-Mammoth Mountain area, as noted by Mat- thes (1930) and Erwin (1934), from whence it flowed eastward into the Mammoth Lakes basin and westward into the valley of the Middle Fork. There is some evi- dence for an additional local source in the northern part of Pumice Flat, where an apparent andesitic dike cuts a deposit of ash and volcanic rubble. The pattern of silica distribution (pl. 1) also suggests the possibility of separate source areas within the Middle Fork valley and in the Mammoth Lakes basin. Perhaps the samples anomalously high in silica from south of the Lower Falls also came from local vents. Three small outliers south of Lost Dog Lake are also derived from a local source; their correlation with the andesite of the Devils Postpile is uncertain. The Red Cones, considered to be an additional source by Erwin (1934, p. 49), have not been glaciated and are of Recent age rather than Pleis- tocene. Prior to the eruption of the andesite of the Devils Postpile, the tuff of Reds Meadow had been almost com- pletely removed from the central part of the Middle Fork valley (see above), and there the andesite rests directly upon granitic rocks. The andesite was sub- sequently largely removed by glaciation, leaving behind remnants on the valley bottom and slopes. Reconstruc- tion from these remnants suggests an original thickness of at least 600 feet for the andesite in the vicinity of the Devils Postpile. The original extent of the andesite is unknown. The andesite displays conspicuous jointing nearly ev- erywhere, with three distinct types represented. Lo- cally near the base of the flows, particularly along the Middle Fork in the vicinity of Rainbow Falls, nearly 2The analytical data for this determination are as follows (G. B. Dalrymple, written commun., 1965): weight, 10.02 grams; K20, 0.268 percent; radiogenic Arm, 1.1 percent; radiogenlc Arm, 2.50X10-13 mole per g. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY horizontal platy jointing has formed. ‘Most com- monly the jointing is of the orthogonal contraction: crack polygon type, giving outcrops a blocky appear- ance. Least common, but in places very well developed, are joints of the nonorthogonal contraction-crack poly- gon type, as for example, in the Devils Postpile itself (frontispiece and fig. 10). The origin of both orthog— onal and nonorthogonal polygonal jointing has recently been studied by Lachenbruch (1962) and by Spry (1962), and we have discussed it briefly with reference to the Devils Postpile (Huber and Rinehart, 1965b). The columns of the Devils Postpile, some of which are as much as 60 feet long, are polygonal, with hexagonal columns slightly more abundant than pentagonal ones, and with other forms making up 10—20 percent of the total (Hartesveldt, 1952; Beard, 1959). The jointing which defines the columns formed near the base of the flow, and the curved and tilted forms of some of the columns are probably due to irregularities in the iso- therms during cooling, which may have been caused, at least in part, by topographic irregularities on the bed- rock surface over which the lava flowed. QUARTZ LATITE 0F MAMMOTH MOUNTAIN Mammoth Mountain has been described by Mayo (1941, p. 1068) as “the most impressive volcanic edi- fice in the region,” an apt description, because it rises abruptly over 2,000 feet above the surrounding country- side. The eruptive history of Mammoth Mountain in- cludes intrusion of a massive dome or domes, explosive FIGURE 10.——Polygonal joint pattern displayed by tops of columns in the Devils Postpile. Note glacial stria- tions and remnants of glacial polish. Photograph by Gerhard Schumacher. CENOZOIC VOLCANIC ROCKS, DEVILS POS’I‘PILE QUADRANGLE, CALIF. activity, and the outpouring of extensive glassy flows, especially on the northeast side of the mountain. Like the quartz latite of Two Teats, the rock which makes up this unit is extremely variable in color and texture. It ranges from shades of brown and pink through gray to nearly black in the glassiest flows. The rock is typically flow banded and porphyritic; plagio- clase and biotite dominate and vary widely in crystal size and relative amounts. Despite the wide petro- graphic variation, the silica in 17 of 21 samples shows a range of only 4 percent and averages 68 percent, as in— ferred from the index of refraction of fused samples (fig. 5). The other four, with 73—74 percent silica, represent rhyolitic obsidian flows that are locally inter- layered with lithoidal quartz latite flows. Sampling of this unit and analytical work are inadequate to deter— mine whether these obsidian flows are chemically dis- tinct from the quartz latite, although the data available indicate such a possibility. The quartz latite of Mammoth Mountain was con- sidered by Erwin (1934, p. 45) to be of Miocene age and by Matthes (1930, 1960) to be “preglacial” in age. Rinehart and Ross (1964, p. 64) also considered the quartz latite of Mammoth Mountain to be of late Ter- tiary (Pliocene) age, because of its degree of dissection and its earlier correlation with the very similar quartz latite on Two Teats and San Joaquin Mountain, for which a late Tertiary age appears correct. Unlike Two Teats and San Joaquin Mountain, however, Mammoth Mountain does retain appreciable constructional form, and its degree of dissection is not incompatible with a Pleistocene—even a relatively late Pleistocene— age. Although he did not study the quartz latite of Mammoth Mountain in any detail, Gilbert (1941, p. 799 and fig. 2) implied that it is of Pleistocene age. Field relations between the quartz latite of Mammoth Moun- tain and the andesite of the Devils Postpile are equivo- cal as to their relative ages. However, two potassium- argon age determinations on the quartz latite yielded ages of 0.37i0.04 million years (Dalrymple, 1964a, table 1) and 0.18i0.09 million years.3 Mammoth Mountain is therefore considered to be of late Pleisto— cene age and younger than the andesite of the Devils Postpile. The construction of Mammoth Mountain appears to have been complex, involving extrusion of viscous flows of highly varied petrographic character that range from crystal-rich lithoidal types to crystal-poor obsidian. Some of the flows were sufficiently fluid to advance at least 2 miles beyond the north base of the mountain, 3The analytical data for this determination are as follows (R. W. Kistler, written commun., 1961) : biotite, 9.25 g; K, 6.8 percent; radiogenlc Arm, 3.8 percent; radiogenlc ATM), 2.21 X 10'12 moles per g. D15 but most appear to have moved only down the slopes over earlier flows, thus piling up around the vent to form a rather typical cumulodome. The high degree of dissection of the north side of the mountain is prob— ably due to the removal of much material by violent explosive activity and to the subsequent modification by glaciation. In addition to the construction of Mam- moth Mountain, this episode of volcanism produced a satellite dome 1 mile in diameter and 1,000 feet in height, about 1 mile northeast of the base of Mammoth Moun— tain. Glassy flows are discontinuously exposed over an extensive area north of Mammoth Mountain. Minor late-stage activity is manifested by small p-ostglacial phvreatic explosion pits at the north base of Mammoth Mountain and by present-day fumarolic activity at the south base and at the crest. Mammoth Mountain lies on or immediately adjacent to the western perimeter fault that defines the Long Valley volcano«tectonic depression (Paki‘ser and others, 1964, p. 17 and pl. 1). Inasmuch as the location of a volcanic pile the size of Mammoth Mountain would probably be dictated by a major zone of crustal weak- ness, it is likely that this fault provided structural control for the vents from which the quartz latite erupted. In this respect Mammoth Mountain would lie in a structural position similar to the “late rhyolite domes” peripheral to the Valles caldera, New Mexico (Smith and others, 1961). When the structural depres- sion began to form cannot be stated with any degree of certainty because not enough is yet known about the de- tailed relationships of the volcanic rocks in the Long Valley basin and adjacent areas, especially their rela- tive and absolute ages. However, the vertical displace- ment along the western part of the fault, which re- sulted in the Sierran escarpment east of Deadman Pass, must have occurred later than approximately 3 million years ago, the age of the quartz latite of Two Teats. If this fault controlled the location of Mammoth Moun- tain, then it must have been active prior to approxi- mately a quarter of a million years ago, and thus move- ment must have begun between approximately a quarter of a million and 3 million years ago. ANDESITE FROM DRY CREEK AREA In the Dry Creek drainage area, north of Mammoth Mountain, a number of scattered outcrops of andesite and basalt are exposed through the mantle of pumice and alluvium, chiefly along recent faults or in explo— sion pits or cinder cones. Because of the generally poor exposures of this unit, it is difficult to determine its stratigraphic position relative to the other volcanic rocks in the area and, indeed, whether all of the scattered outcrops are correlative. D16 Analysis of fused-bead data (fig. 5) indicates a dis- tinct bimodal composition distribution with maximums at basalt (51 percent Si02) and andesite (56 percent SiOZ). The more silicic rocks are concentrated near the Inyo Craters and along the eastern edge of the quadrangle (pl. 1). At three localities the andesite and basalt can be dem- onstrated to be younger than the quartz latite of Mam- moth Mountain. These three exposures are (1) along a recent fault scarp about 1 mile north of Mammoth Mountain, (2) 0n the north flank of the quartz latite dome northeast of Mammoth Mountain, and (3) on the northeast flank of Mammoth Mountain just west of a recent fracture known locally as the “Earthquake Fault.” The rock in exposures 1 and 3 has been glaci- ated, and therefore the unit is considered to be of late Pleistocene age rather than Recent. Although exact correlations are questionable, all the andesite and basalt in Dry Creek area appear to be approximately the same age and are arbitrarily assigned to the same map unit. In addition, some glaciated andesitic cinder cones near the base of the San Joaquin Mountain ridge are also assigned to this unit. ANDESITE 0F PUMICE BUTTE Pumice Butte is one of the two andesitic cinder cones in the southeast quadrant of the quadrangle. These cones, together with a vesicular andesite that appears to have originated as a flow from the base of Pumice Butte, were not overridden by glacial ice. They are above the elevation limit of the main valley glaciers of Wisconsin age, but a small tributary glacier of presumed Wisconsin age overrode a low shoulder on the north- westernmost of the two cones, and the margins of the flow were glaciated during latest Wisconsin (R. J. J anda, oral commun., 1965). This andesite is therefore thought to have been erupted during a pre-Wisconsin or intra-Wisconsin interglaciation. The rock is typi- cally scoriaceous and appears fresh—quite unlike most of the andesitic flows and rubble south of Deer Creek and east of Pond Lily Lake, which are correlated with the andesite of Deadman Pass. Contacts shown between these units on the map are only approximate, however, and have been distinguished in part by aerial photo- interpretation. Included with the andesite of Pumice Butte is a domelike mass of dacite (at the north edge of the map unit), which appears to be more highly dis- sected and is probably older than the main mass of the andesite. OLIVINE-BEARING QUARTZ LATITE In the extreme northeast corner of the quadrangle, olivine-bearing quartz latite covers an area of about half SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY a square mile. It is exposed somewhat more extensively in the Mount Morrison quadrangle, where it has been described by Rinehart and Ross (1964, p. 57). The three following lines of evidence suggest that the quartz latite is of late Pleistocene to Recent age: (1) There is no evidence that the quartz latite has been glaciated, al- though late Pleistocene glaciations probably did not ex- tend this far from the range front, (2) only slight dis- section is shown in the major area of outcrop, which consists of a relatively flat topped stubby steepfaced flow that can be traced westward into steep-sided domes that retain much of their original form, (3) a north-trending fault scarp in the Mount Morrision quadrangle, along which old glacial till has been displaced, terminates abruptly at the edge of the flow and is probably buried by it. The rock contains abundant white feldspar pheno- crysts 2—5 millimeters long in a dark-gray aphanitic or glassy groundmass. Flow banding is visible locally, and in general appearance the rock is not unlike some of the darker varieties of the older quartz latites. BASALT OF THE RED CONES At the point where Crater Creek tumbles from a bench above the Middle Fork of the San Joaquin River— a bench which Matthes (1960, p. 43) considered part of the Broad Valley Surface—two basalt cinder cones lie on either side of the creek, and associated basalt flows cascade down the slope into the Middle Fork valley. The Red Cones were previously considered contempo- raneous with, and a source of, part of the andesite of the Devils Postpile in the Middle Fork valley. Unlike the andesite of the Devils Postpile, however, the basalt of the Red Cones is typically vesicular with prominent phenocrysts of plagioclase and olivine. Moreover, the Red Cones have obviously not been glaciated, as has the andesite of the Devils Postpile; Matthes (1960, pl. 1) and Erwin (1934, p. 49) attributed this to the elevation of the cones above the Wisconsin glaciers in the Middle Fork valley. However, the associated flows extending into the Middle Fork valley also escaped glaciation—— not only is there a complete absence of glacial erratics, striations, or polish, but the flows also exhibit scoria- ceous and rubbly surfaces of a type nowhere present on the glaciated andesite of the Devils Postpile. The basalt of the Red Cones is therefore considered to be of post- glacial, or Recent, age. The Red Cones still exhibit summit craters (fig. 11) , and a detailed examination of aerial photographs suggests that most of the flow ma- terial was erupted from the western base of the southern— most of the two craters. The flow has been only slightly modified by subsequent erosion. CENOZOIC VOLCANIC ROCKS, DEVILS POSTPILE QUADRANGLE, CALIF. FIGURE 11.—The northernmost of the two Red Cones as viewed from the southern cone. Note breached summit crater. RHYOLITE In the northeast corner of the quadrangle is a series of rhyolite domes that have been considered to be a southern extension of the Mono Craters and have been described in some detail by Mayo, Conant, and Cheli- kowsky (1936). This series consists of six rhyolite domes, alined in a north-south direction over a distance of 6 miles, beginning about 2 miles south of the main mass of the Mono Craters. They tend to be nearly equidimensional and range in size from less than half a mile to nearly a mile in diameter. The four southern— most domes are in the Devils Potpile quadrangle. There is little doubt that these domes are of very re- cent origin; for erosional modification of their forms is negligible and several of them support no vegetation. Each has a multiphase origin, which is best understood by comparison with Panum Crater at the northern end of the Mono Craters, where evidence of the evolution of a dome has been excellently preserved (Williams, 1932; Putnam, 1938). At Panum Crater the initial phase involved the excavation of an explosion pit with the expulsion of vast quantities of pumice and the construc— tion of a pumice-lapilli rim around the perimeter of the pit. This explosive phase was succeeded by the protru- sion of a nearly solid column of pumiceous obsidian and partly crystalline material in the form of a dome on the floor: of the explosion pit. As the dome rose, it tended to fracture and spread laterally, with blocks of pumiceous obsidian spalling off and accumulating in steep talus slopes at its margins. The two larger domes in the Devils Postpile quad- rangle—the one just south of Deadman Creek and the other just south of Dry Creek—had a similar mode of origin. However, these two domes spread laterally dur- ing their protrusion so that they almost completely over— D17 whelmed their original lapilli cones and thus nearly erased all evidence of the earlier explosive phase of their eruption. The dome just south of Deadman Creek (fig. 12) has only a small remnant of its lapilli rim on its southwest side, and the dome south of Glass Creek has none. The top surfaces of each of these domes are very irregular with numerous spires and loosely piled, angular blocks of pumiceousiobsidian (fig. 13). They have a generally concentric structure which is obscure on the surface but readily apparent on aerial photo- graphs (fig. 14). The small tree-covered dome north of Deadman Creek is probably of the same origin, having similar spires and an irregular blocky surface. This FIGURE 12.——Rhyolite dome south of Deadman Creek as seen from Deer Mountain. Dome is about three-quarters of a mile in diameter. A second rhyolite dome, just south of Glass Creek, can be seen in the right middle distance. Photograph by Gerhard Schumacher. FIGURE 13.—Upper surface of rhyolite dome south of Glass Creek showing typical irregular blocky surface. Taken from top of dome in right middle distance of figure 12. D18 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY FIGURE 14.—Vertical aerial photographs of Deer Mountain area, mounted as a stereopair. Main features from south to north are the Inyo Craters, Deer Mountain, and several rhyolite domes. Note remnant of lapilli rim on southern side of rhyolite dome south of Deadman Creek. CENOZOIC VOLCANIC ROCKS, DEVILS POSTPILE QUADRANGLE, CALIF. dome, however, is slightly older than the two larger domes, for it is mantled with debris from later explo- sions and its surface structure consequently obscured. Deer Mountain, the southernmost of the rhyolite domes, has a somewhat more complex history. It may have had an initial origin similar to that of the other domes, but this has been obscured by a late explosive phase, which blasted a summit crater and expelled large quantities of material, including the fine ash that man- tles the mountain and gives it the appearance of a vol- canic cone and not a dome. Several explosion pits near the base of the mountain, the largest two of which are known as the Inyo Craters (fig. 15), further attest to the late explosive phase. A radiocarbon age determination on a log buried within the ejecta that form the parapet of the southern— most of the Inyo Craters yielded an age of 650i200 years (lab. No. W—1431). However, a minimum age of 500 years for the craters, slightly older than the mini- mum radiocarbon age, is required by the fact that 400- year-old pines and firs, whose age is based on ring counts of recently cut specimens, are growing within the craters (fig. 15). Allowing 100 years for trees to become estab- lished after eruption fixes the minimum age of the craters at 500 years. The craters, therefore, must have formed between 500 and 850 years ago (Rinehart and Huber, 1965). lPUMICE Pumice of varying thickness is ubiquitous in most of the Devils Postpile quadrangle east of the Ritter Range, and has obscured many stratigraphic relations between the various Cenozoic units. The pumice cover is most extensive on the east flank of the San Joaquin Mountain ridge and in the lowlands to the east. Much of this pumice was undoubtedly de- FIGURE 15.—View of Inyo Craters from Deer Mountain. Ash- covered rim of summit crater on Deer Mountain can be seen in foreground. D19 rived from explosive eruptions associated with the rhyo- lite domes just described and the Mono Craters farther to the north. It is also probable, however, that large quantities of pumice were ejected during one or more phases in the formation of Mammoth Mountain; much of the pumice in the valley of the Middle Fork of the San Joaquin River and in the vicinity of Pumice Butte, southwest of Mammoth Crest, may have come from this source. Despite its name, Pumice Butte, as has been noted, is an andesitic cinder cone and therefore an un- likely source of pumice in that area, although Birman (1964, p. 7) suggested this possibility. No other likely source of pumice within the San Joaquin drainage basin was noted. A study of the surficial pumice in and around the Mammoth Embayment and Mono Basin should contrib- ute much to the understanding of volcanism in this region over the past few million years, particularly in the light of the recent success of similar studies (Powers and Wilcox, 1964; Czamanske and Porter, 1965). Our knowledge of the volcanic history of the region will not be complete until such a study has been made. THERMAL ACTIVITY Extensive thermal activity in the general Mammoth Lakes area also attests to the recency of volcanic activity. Although hot springs and areas of thermal alteration are more numerous in the Mount Morrison quadrangle to the east (Rinehart and Ross, 1964; McNitt, 1963), Mammoth Mountain itself is a thermal area, and hot springs exist at two other localities in the Devils Post- pile quadrangle. Until at least the‘ late 1950’s, there were a number of active fumaroles near the summit of Mammoth Moun— tain, and indications of thermal alteration are visible over an extensive area. On the south flank of Mammoth Mountain, within a less extensive area of thermal altera- tion, just north of Mammoth Pass, a few small steam vents are still active. The hot spring at Reds Meadow emerges from a grassy slope underlain by the tulf of Reds Meadow. It appears to issue from near the base of the lower zone of intensely welded tuft, which exhibits conspicuous j ointing in nearby outcrops. As the tufl’ is an ash flow whose source we believe to be outside of the Reds Mea- dow area, it is unlikely that the hot spring is genetically related to the tufl". In fact, it cannot be related to any episode of local volcanic eruptive activity. Nor can Fish Creek Hot Springs be related to known volcanism; for they issue from a grassy slope underlain by granitic rocks, 3 miles from the nearest exposure of Cenozoic volcanic rocks. In addition, several cold springs rich in carbon dioxide are in the quadrangle, but their rela— tion to volcanism, if any, is unknown. D20 SUMMARY By late Pliocene time, the rolling upland surface of Matthes’ Broad Valley erosion stage had been consider- ably incised during the Mountain Valley and possibly the Canyon erosion stages. The Middle Fork of the San Joaquin River appears to have had a principal tribu— tary whose source was east of the present Sierran divide; the ancestral channel crossed what is now the drainage divide, probably in the area between Deadman Pass and Two Teats. Upon this surface, somewhat less than 4 million years ago, large quantities of andesite and basalt were erupted from scattered vents to cover an extensive part of the quadrangle. Shortly thereafter, the quartz latite of Two Teats buried a large area in the northeast- ern part of the quadrangle. It is possible that some of this volcanic activity was related to an early stage in the formation of the Long Valley volcano—tectonic depression, although at least some of the movement along the western perimeter fault occurred later. If the relationships are analogous to those at the Valles Caldera, N. Mex, this faulting prob- ably was associated with the eruption of the Bishop and Reds Meadow ash-flow tuf'fs, the next major volcanic episode. About 0.7 million years ago the tuff of Reds Meadow was erupted into the valley of the Middle Fork of the San Joaquin River, which probably had been previously glaciated. In a relatively brief period of time, several cubic miles of the tuff was almost completely removed from the valley, possibly during a glacial advance, and subsequent eruption of the andesite of the Devils Post- pile followed about half a million years ago. By this time it is probable that the Long Valley vol- cano-tectonic depression had formed, and quartz latite was erupted to form Mammoth Mountain 011 the pe- riphery of the depression. Subsequently, other minor andesites and quartz latites were erupted and were over- ridden by glaciers. Two Recent, or postglacial, events were the eruption of the basalt of the Red Cones in the Middle Fork valley and the extrusion of the rhyolite domes in the northeast corner of the quadrangle. Minor phreatic explosions in and near the domes, resulting in craters up to 600 feet in diameter, occurred as recently as 500 to 850 years ago, and thermal activity continues to the present. REFERENCES Axelrod, D. I., 1962, Post-Pliocene uplift of the Sierra Nevada, California: Geol. Soc. America Bull., v. 73, no. 2, p. 183—198. Axelrod, D. I., and Ting, ‘V. S., 1960, Late Pliocene floras east of the Sierra Nevada: California Univ. Pubs. Geol. Sci., v. 39, n0. 1, p. 1—117. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Bateman, P. 0., Clark, L. D., Huber, N. K, Moore, J. G., and Rinehart, C. D., 1963, The Sierra Nevada batholith—a synthesis of recent work across the central part: U.S. Geol. Survey Prof. Paper 414—D, p. D1—D46. Beard, C. N., 1959, Quantitative study of columnar jointing: Geol. Soc. America Bull., v. 70, no. 3, p. 379—381. Birman, J. H., 1964, Glacial geology across the crest of the Sierra Nevada, California: Geol. Soc. America Spec. Paper 75, 80 p. Czamanske, G. K., and Porter, S. C., 1965, Titanium dioxide in pyroclastic layers from volcanoes in the Cascade Range: Science, v. 150, no. 3698, p. 1022—1025. Dalrymple, G. B., 1964a, Cenozoic chronology of the Sierra Ne- vada, California: California Univ. Pubs. Geol. Sci, v. 47, 41 p. 1964b, Potassium-argon dates of three Pleistocene in- terglacial basalt flows from the Sierra Nevada, California: Geol. Soc. America Bull., v. 75, no. 8, p. 753—758. Dalrymple, G. B., Cox, Allan, and Doell, R. R., 1965, Potassium- argon age and paleomag'netism of the Bishop Tuff, Cal- ifornia: Geol. Soc. America Bull., v. 76, no. 6, p. 665—674. Dalrymple, G. B., Doell, R. R., and Cox, Allan, 1965, Recent developments in the geomagnetic polarity [abs] : Geol. Soc. America, 78th Ann. Mtg., Kansas City, Mo., 1965, Program, p.40. Erwin, H. D., 1934, Geology and mineral resources of north- eastern Madera County, California: California J our. Mines and Geology, v. 30, no. 11, p. 7—78. Gilbert, C. M., 1938, Welded tuff in eastern California: Geol. Soc. America Bull., v. 49, no. 12, pt. 1, p. 1829—1862. 1941, Late Tertiary geology southeast of Mono Lake, Cal- ifornia: Geol. Soc. America Bull., v. 52, no. 6, p. 781—815. Goddard, E. N., chm., and others, 1948, Rock-color chart: Wash- ington, Natl. Research Council (repub. by Geol. Soc. Amer- ica, 1951), 6 p. Hartesveld‘t, R. J ., 1952, a brief story of the Devils Postpile National Monument [California] : Yosemite Nature Notes, v. 31, no. 10, p. 138—149. Revised and repub, 1963, Your guide to Devils Postpile National Monument: Yosemite Natl. Park, Caiif., Yosemite Nat. History Assoc, 22 p. Huber, N. K., and Rinehart, C. D., 1965a, Geologic map of the Devils Postpile quadrangle, Sierra Nevada, California: US. Geol. Survey Geol. Quad. Map GQ—437, scale 1: 62,500. 1965b, The Devils Postpile National Monument: Cali- fornia Div. Mines and Geology, Mineral Inf. Service, v. 18, no. 6, p. 109—118. 1966, Some relationships between the refractive index of fused glass beads and the petrologic afiinity of volcanic rock suites: Geol. Soc. America Bull., v. 77, no. 1, p. 101—110. Hudson, F. S., 1960, Post-Pliocene uplift of the Sierra Nevada, California: Geol. Soc. America Bull., v. 71, no. 11, p. 1547— 1573. ' J anda, R. J ., 1965, Quaternary alluvium near Friant, California, in Guidebook for Field Conf. I, Northern Great Basin and California : INQUA (Internat. Assoc. Quaternary Research) Cong, 7th, Boulder, Colo., 1965, p. 128—133. Lachenbruch, A. H., 1962, Mechanics of thermal contraction cracks and ice-wedge polygons in permafrost: Geol. Soc. America Spec. Paper 70, 69 p. McNitt, J. R., 1963, Exploration and development of geothermal power in California: California Div. Mines and Geology Spec. Rept. 75, 45 p. CENOZOIC VOLCANIC ROCKS, DEVILS POSTPILE QUADRANGLE, CALIF. Matthes, F. E., 1930, The Devils Postpile and its strange setting : Sierra Club Bull., v. 15, no. 1, p. 1—8. 1960, Reconnaissance of the geomorphology and glacial geology of the San Joaquin basin, Sierra Nevada, California : U.S. Geol. Survey Prof. Paper 329, 62 p. Mayo, E. B., 1934 Geology and mineral resources of Laurel and Convict basins, southwestern Mono County, California: California Jour. Mines and Geology, v. 30, no. 1, p. 79—88. 1941, Deformation in the interval Mt. Lyell—Mt. Whitney, California: Geol. Soc. America Bull., v. 52, no. 7, p. 1001— 1084. Mayo, E. B., Conant, L. C., and Chelikowsky, J. R., 1936, Southern extension of the Mono Craters, California: Am. Jour. Sci., v. 32, no. 188, p. 81—97. Nockolds, S. R., 1954, Average chemical composition of some ig- neous rocks: Geol. Soc. America Bull., v. 65, no. 10, p. 1007— 1032. Pakiser, L. 0., Kane, M. R, and Jackson, W. H., 1964, Structural geology and volcanism of Owens Valley region, California— a geophysical study: U.S. Geol. Survey Prof. Paper 438, 68 p. Peacock, M. A., 1931, Classification of igneous rock series: Jour. Geology, v. 39, no. 1, p. 54—67. Powers, H. A., and Wilcox, R. E., 1964, Volcanic ash from Mount Mazama (Crater Lake) and from Glacier Peak: Science, v. 144, no. 3624, p. 1334-1336. Putnam, W. C., 1938, The Mono Craters, California: Geog. Rev., v. 28, no. 1, p. 68—82. 1960, Origin of Rock Creek and Owens River Gorges, Mono County, California: California Univ. Pubs. Geol. Sci., v. 34, no. 5, p. 221—279. 1962, Late Cenozoic geology of McGee Mountain, Mono County, California: California Univ. Pubs. Geol. Sci., v. 40, no. 3, p. 181-218. Rinehart, C. D., and Huber, N. K., 1965, The Inyo Crater Lakes— a blast in the past: California Div. Mines and Geology, Mineral Inf. Service, v. 18, no. 9, p. 169—172. D21 Rinehart, C. D., and Ross, D. C., 1964, Geology and mineral deposits of the Mount Morrison quadrangle, Sierra Nevada, California, with. a section on A gravity study of Long Valley, by L. C. Pakiser: U.S. Geol. Survey Prof. Paper 385, 106 p. Rittmann, Alfred, 1952, Nomenclature of volcanic rocks pro- posed for the use in the catalogue of volcanoes, and key- tables for the determination of volcanic rocks: Bull. V01- canol., ser. 2, v. 12, p. 75—103. 1953, Magmatic character and tectonic position of the Indonesian volcanoes: Bull. volcanol., ser. 2, v. 14, p. 45—58. Shapiro, Leonard, and Brannock, W. W., 1956, Rapid analysis of silicate rocks: U.S. Geol. Survey Bull. 1036—C, p. Cl9—C56. Sharp, R. P., 1965, Rock Creek to Owens Gorge, m Guidebook for Field Conf. I, Northern Great Basin and California: INQUA (Internat. Assoc. Quaternary Research) Cong, 7th Boulder, Colo., 1965, p. 97—99. Smith, R. L., 1960, Zones and zonal variations in welded ash flows: U.S. Geol. Survey Prof. Paper 354—F, p. F149—F159. Smith, R. L., Bailey, R. A., and Ross, C. S., 1961, Structural evolution of the Valles caldera, New Mexico, and its hear- ing on the emplacement of ring dikes, in Short papers in the geologic and hydrologic sciences: U.S. Geol. Survey Prof. Paper 424—D, p. D14EHD149. Spry, Alan, 1962, The origin of columnar jointing, particularly in basalt flows: Geol. Soc. Australia Jour., v. 8, pt. 2, p. 191—216. Thornton, C. P., and Tuttle, O. R, 1960, Chemistry of igneous rocks, [pt.] 1, Differentiation index: Am. Jour. Sci., v. 258, No. 9, p. 664—684. Wahrhaftig, Clyde, 1965, Roadcut at Rock Creek, in Guidebook for Field Conf. 1, Northern Great Basin and California: INQUA (Internat. Assoc. Quaternary Research) Cong, 7th, Boulder, 0010., 1965, p. 97. Williams, Howel, 1932, The history and character of volcanic domes: California Univ. Dept. Geol. Sci. Bu11., v. 21, no. 5, p. 51—146. UNITED STATES DEPARTMENT OF THE INTERIOR PREPARED 1N COOPERATION WITH ‘ PROFESSIONAL PAPER 554—D GEOLOGICAL SURVEY -‘ PLATE 1 (MONO CRATEK’S) EXPLANATION SURFICIAL DEPOSITS Surficial deposits undivided Include alluvial, colluoial, and glacial deposits, and pumice that covers extensive areas east of Mammoth Mountain-San Joaquin Moun- tain ridgei Omitted from most of northwest quadrant of map and Recent mr—bfir—h—fi Pleistocene QUATERNARY CENOZOIC VOLCANIC ROCKS Rhyolite < V ‘Qbr A /’>/\v/\ Basalt of the Red Cones or Recent Pleistocene \ Andesite- In Dry Creek area OUATERNARY Quartz latite of Mammoth Mountain + + + + Qap + + + + Andesite of Pumice Butte Age relationship uncertain Pleistocene Andesite 0f the Devils Postpile CENOZOIC Tuff of Reds Meadow Rhyolite ash-flow tuff, in part welded * ‘Tati: ' i if 3 b . Andesite In Deadman- Glass Creeks area. Age relationship uncertain or Pleistocene O R Q U AT E R N A R Y Pliocene (ME‘RCEI‘O PEAK) (Mr MORRHSON) Y TE RT I A R Y I Quartz latite of Two Teats Pliocene Y TE RT I A R Y Andesite of Deadman Pass Includes scattered remnants on Broad Valley erosion surface of Matthes (1960) IGNEOUS AND METAMORPHIC ROCKS pTu Igneous and metamorphic rocks undivided Predominantly of Mesozoic age PRE—TERTIARY Contact _T____.......... Fault Dashed where approximately located; dotted where inferred. Bar and ball on downthrown side 62? Explosion pit T 0.. Thermal springs (T) and fumaroles (F) 57 Percentage Si02 inferred from refractive index of fused sample <—v——— Sampling traverse ‘x// M,‘ . -‘ , (KAISER PEAK) .V INTERIORPGEOLOGICAL SURVEY, WASHINGTON, D. C.—1967v666386 1153“» ,= Geology by N. King Huber and C. Dean Rinehart, 1956— 59; assisted by R. V. Sharp, 1956 'GEOLOGIC MAP SHOWING CENOZOIC VOLCANIC ROCKS OF THE DEVILS POSTPILE QUADRANGLE SIERRA NEVADA, CALIFORNIA SCALE 1:62 500 Base by U.»S. Geological Survey, 1953 4 5 MILES 1 .5 O 1 4 5 KILOMETERS H H H l-leji—“i l-————l APPROXIMATE MEAN CONTOUR INTERVAL 80 FEET DECUNAITON'IQN DATUM IS MEAN SEA LEVEL TRUE NORTH @573" 7 DAY m \/'y Sgt; "E _Petrology and Structure of Precambrian Rocks Central City Quadrangle Colorado GEOLOGICAL SURVEYIPROFESSIONAL PAPER 554—E Petrology and Structure of Precambrian Rocks Central City Quadrangle Colorado By P. K. SIMS and D. J. GABLE SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—E A stuajz of flz'g/z—graa’e metamorp/zz'c and zgfleous rocés wz'tflz'fl Me Colorado mineral oe/t UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1967 UNITED STATES DEPARTMENT OF THE INTERIOR STEWART L. UDALL, Secretary GEOLOGICAL SURVEY William T. Pecora, Director For sale by the Superintendent of Documents, US. Government Printing Office Washington, DC. 20402 CONTENTS Page Page Abstract ___________________________________________ E1 Intrusive rocks—Continued Introduction _______________________________________ 2 Emplacement and origin of the intrusive rocks _____ E41 Geologic setting ____________________________________ 3 Metamorphic facies _________________________________ 43 Rock units _________________________________________ 6 Structure __________________________________________ 44 Metamorphic rocks _________________________________ 6 Terminology ___________________________________ 45 Lithologic succession ____________________________ 6 Folds __________________________________________ 45 Microcline-quartz-plagioclase—biotite gneiss _________ 8 Major folds ................................ 47 Amphibolite ____________________________________ 12 Central City anticline ___________________ 47 Cordierite-amphibole gneiss ______________________ 13 Idaho Springs anticline __________________ 47 Cordierite—gedrite rocks ______________________ 14 Lawson syncline ________________________ 47 Cordierite—biotite rocks ______________________ 16 Intermediate-scale folds _____________________ 48 Hornblende—cummingtonite rocks _____________ 16 Pewabic Mountain syncline ______________ 48 Chemical composition _______________________ 17 Bald Mountain syncline __________________ 48 Cale-silicate gneiss and related rocks ______________ 17 Dumont anticline _______________________ 48 Biotite gneiss ___________________________________ 18 Syncline at Mount Pisgah _______________ 48 Biotite—quartz-plagioclase gneiss ______________ 18 Peeks Flat anticline and adjacent overturned Sillimanitic biotite gneiss .................... 20 folds ________________________________ 48 Cordierite— and garnet-bearing sillimanitic bio- Linear elements related to folds _______________ 49 tite gneiss ----- f j - - - - - - - - - j - -.- - - - - "- ------- 2 1 Idaho Springs-Ralston shear zone _________________ 49 _ 'Chemical composition 0f the bmtlte gne1sses-___ 24 Shearing and associated alteration in Dakota Hill Origin of metamorphic rocks _____________________ 26 . . . area _________________________________________ 50 Gramte gneiss and pegmatlte _________________________ 28 Intrusive rocks _____________________________________ 29 Faults """"""""""""""""""""""" 51 Granodiorite and associated rocks _________________ 30 Jomts ---------------------------------------- 51 Gabbro and related rocks ________________________ 35 Character and environment of deformation _________ 52 Quartz diorite and hornblendite __________________ 38 Geologic history -------------------------------- 54 Biotite—muscovite quartz monzonite _______________ 39 References cited .................................... 55 ILLUSTRATIONS Page PLATE 1. Geologic map and sections of the Central City quadrangle, Colorado ____________________________________ In pocket FIGURE 1. Map showing location of the Central City quadrangle in the Front Range, Colo _____________________________ E2 2. Generalized geologic map of the central part of the Front Range and adjacent areas _________________________ 4 3. Triangular diagram showing variation in composition of microcline gneiss but excluding Quartz Hill layer_-__-__ 9 4. Sketch showing layering in a body of cordierite-amphibole gneiss, Quartz Hill ............................... 15 5. Diagram showing Na20:K20 ratios of biotite gneisses ____________________________________________________ 26 6. Diagram showing FeO: MgO ratios of biotite gneisses ____________________________________________________ 27 7. Triangular diagram showing variation in composition of granodiorite and associated rocks ____________________ 32 8. Contour diagrams of lineations ........................................................................ 46 9. Contour diagrams of joints ____________________________________________________________________________ 52 TABLES Page TABLE 1. Lithologic succession of Precambrian metamorphic rocks, Central City quadrangle ___________________________ E7 2. Modes of microcline-quartz-plagioclase-biotite gneiss _____________________________________________________ 10 3. Modes of amphibolite ________________________________________________________________________________ l2 4. Chemical and spectrochemical analyses and norms of amphibolite __________________________________________ 13 5. Modes of representative varieties of cordierite-gedrite gneiss and associated rocks ____________________________ 14 m IV CONTENTS TABLE 6. Chemical analyses and modes of cordierite-gedrite gneiss and associated rocks _______________________________ 7—12. 14. 15. 16. 17. 18. 19. 20. 21—23 Modes of: 7. Biotite—quartz-plagioclase gneiss ________________________________________________________________ 8. Garnetiferous biotite-quartz-plagioclase gneiss ____________________________________________________ 9. Sillimanitic biotite—quartz-plagioclase gneiss ______________________________________________________ 10. Sillimanitic biotite-quartz gneiss ________________________________________________________________ 11. Garnetiferous sillimanitic biotite—quartz—plagioclase gneiss __________________________________________ 12. Cordierite—bearing garnet-sillimanite—biotite gneiss ________________________________________________ 13. Cordierite-biotite—quartz-plagioclase gneiss _______________________________________________________ Estimated chemical compositions and average modes of principal types of biotite gneiss ______________________ Chemical analyses and modes of cordierite—bearing garnet-sillimanite-biotite gneiss, sillimam'tic biotite—quartz—plagio- clase gneiss, and garnet-sillimanite—biotite-quartz-plagioclase gneiss ____________________________________ Semiquantitative spectrographic analyses of minor elements in biotite gneisses ______________________________ Modes of granite gneiss and pegmatite _________________________________________________________________ Modes of granodiorite and associated rocks _____________________________________________________________ Chemical and spectrochemical analyses and norms and modes of intrusive rocks _____________________________ Chemical and spectrochemical analyses of biotites from intrusive rocks _____________________________________ . Modes of : 21. Gabbro and related rocks ______________________________________________________________________ 22. Quartz diorite and hornblendite ________________________________________________________________ 23. Biotite-muscovite quartz monzonite _____________________________________________________________ Page E15 19 20 21 21 22 22 23 25 26 27 29 32 34 35 37 39 40 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY PETROLOGY AND STRUCTURE OF PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE COLORADO By P. K. SIMS and D. J. GABLE ABSTRACT The Central City quadrangle, in the east-central part of the Front Range, is underlain by Precambrian rocks which are intruded by abundant porphyritic igneous rocks of Tertiary age. The Central City district and parts of the Idaho Springs and Lawson-Dumont-Fall River mining districts, which are integral parts of the Front Range mineral belt, are located at intrusive centers of the Tertiary igneous rocks in the southern part of the quadrangle. The Precambrian rocks are dominantly microcline-quartz- plagioclase-biotite gneiss and migmatitic biotite gneiss. These rocks contain layers and lenses of amphibolite, cordierite- amphibole gneiss, and Gale-silicate gneiss and related rocks and are intruded by generally small bodies of several types of igneous rocks, some of which have been metamorphosed. The metamorphic rocks constitute a well-defined lithologic succession that seems to represent a normal stratigraphic sequence. Three principal layers of microcline gneiss, each several hundred to a few thousand feet thick, are interlayered in alternate sequence with equally thick layers of biotite gneiss. An estimated maximum of 15,000—16,000 feet of strata is exposed in the quadrangle. The amphibolite forms concordant bodies as much as 500 feet wide and 3,000 feet long within and at the margins of major layers of microcline gneiss and local small bodies in the biotite gneiss. The cordierite-amphibole gneiss and calc- silicate rocks occur as small lenses within both microcline gneiss and biotite gneiss. Common mineral assemblages of the metamorphic rocks, listed by rock types, are: Microcline gneiss: Biotite-plagioclase-potassium feldspar-quartz Biotite-muscovite—plagioclase—potassium feldspar-quartz Biotite gneiss: Biotite—plagioclase—quartz Biotite—garnet-plagioclase-quartz Biotite-quartz—sillimanite Biotite-plagioclase-quartz-sillimanite Biotite-plagioclass-potassium feldspar-quartz-sillimanite Biotite-muscovite—plagioclase-potassium feldspar-quartz- sillimanite Biotite-garnet—pla gio cl a s e-potassium feldspar-quartz- sillimanite Biotite-c o r d i e r i t e-garnet—magnetite—plagio cla se-pota s- sium feldspar-quartz-sillimanite Biotite—cordierite-magnetite-plagioclass-quartz Amphibolite : Andesine-hornblende-quartz Cordierite—amphibole gneiss: Biotite—cordierite-garnet—gedrite—plagioclase-quartz Cale-silicate gneiss and related rocks: Clinopyr0xene-garnet-plagioclase-quartz-sphene Garnet-magnetite-quartz Clinopyroxene-epidote—hornblende-plagioclase-quartz The metamorphic rocks are interpreted to be dominantly of metasedimentary origin. The microcline gneiss is thought to represent arkose, and the biotite gneiss, to represent interlayered shale and graywacke. Garnet- and cordierite-bearing varieties of biotite gneiss formed from shale that was somewhat enriched in iron and magnesium and deficient in calcium. Probably the minor metamorphic rocks also were derived mainly from sedi- mentary rocks. The biotite gneisses are migmatized and contain an estimated 15—20 percent by volume of granite gneiss and pegmatite. Other rock types contain lesser amounts of similar material, as streaks or interlacing veinlets. Four types of Precambrian intrusive rocks—granodiorite and associated rocks, gabbro and related rocks, quartz diorite and hornblendite, and biotite-muscovite quartz monzonite—each with associated pegmatites, intrude the layered rocks. From oldest to youngest, they are described as follows: 1. Granodiorite and associated rocks occur as subconcordant folded sheets and small plutons, some of which are phaco- liths. Individual bodies range in composition from mafic quartz diorite to quartz monzonite. The bodies are satel- lite to the batholith of Boulder Creek Granite exposed northeast of the quadrangle. 2. Gabbro and related rocks also form subconcordant bodies; the large Elk Creek pluton probably is a compound phacolith. These rocks are distinguished by their content of both ortho- pyroxene and clinopyroxene and by their range in compo- sition from melagabbro to quartz diorite, diorite being the dominant facies. 3. Quartz diorite and hornblendite grade locally into gabbroic rocks and are interpreted to have formed by retrograde metamorphism of gabbro and related rocks. 4. Biotite-muscovite quartz monzonite forms generally small crosscutting bodies that are peripheral to larger masses at Silver Plume, 0010., and vicinity. In contrast to the older intrusive rocks, it is remarkably uniform in composition. The older intrusives were emplaced, synchronously with the major period of deformation, in the catazone of the crust; sub- sequent to crystallization they were deformed and were largely recrystallized. The biotite-muscovite quartz monzonite has a primary flow structure and is late syntectonic. E1 E2 Progressive metamorphism developed mineral assemblages of the sillimanite zone in all but the youngest intrusive rocks. In rocks of suitable composition, sillimanite and potassium feldspar coexist; muscovite is stable with these minerals in rocks contain- ing sufficient K20. Cordierite is stable in calcium—poor, mag- nesium- and iron-rich biotite gneisses and gedrite-bearing gneisses. Locally, adjacent to the Precambrian Elk Creek pluton the biotite gneiss assemblages are changed to the pyroxene hornfels facies. Adjacent to the largest of the Tertiary intrusive bodies the assemblages are modified also mainly by the conver— sion of highly triclinic microcline to orthoclase. The quadrangle is in a region dominated by northeastward- trending folds; a narrow segment of the major Idaho Springs- Ralston shear zone extends across the extreme southeast corner of the quadrangle. The northeastward‘trending folds are mainly open upright anticlines and synclines that have steeply dipping axial planes and gently plunging fold axes. Closed overturned folds occur in the west-central part of the quadrangle. Lineations that are parallel to the major fold axes (B) and that are nearly normal (A) to them are cogenetic with the folding. The Idaho Springs-Ralston shear zone trends N. 55° E. and is characterized by extreme cataclasis and minor folds that are subparallel to the zone itself. The major folds and associated linear elements were formed during the principal episode of plastic deformation; the younger folding and cataclasis formed during a distinctly later deforma- tion, but still in the Precambrian. This deformation was followed by an episode of faulting, the youngest known mani- festation of Precambrian structural activity in this area. INTRODUCTION Geologic mapping of the Central City quadrangle was undertaken for two principal purposes: (1) To ex- tend to a broader region the knowledge obtained through detailed mapping of the mining districts in the central part of the Front Range and (2) to form a nucleus for a program of quadrangle geologic mapping intended to result in a geologic section across the Front Range. The ultimate objective of this program was to gain a comprehensive knowledge of this segment of the Front Range in order to improve the understanding of the composition and structure of the range and of the factors controlling the localization and extent of the ore deposits of the Front Range mineral belt. This report describes the Precambrian rocks of the quadrangle and places particular emphasis on their petrology and structure. Although the Tertiary por- phyritic igneous rocks and metalliferous veins are shown on the geologic map (pl. 1), Which was origi- nally published by Sims (1964), they are not de- scribed herein because the main features of the rocks are discussed by Wells (1960) and the principal mining districts are described in reports by Sims, Drake, and Tooker (1963), Moench and Drake (1966), and Hawley and Moore (1967). The Central City quadrangle lies astride the boundary between Gilpin and Clear Creek Counties, in the east—central part of the Front Range (fig. 1). SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY 41°—,1Or§— --——//"’19IT—‘_—7 o /// l \/ / LLI // o / i ‘4‘ Z I 4 ! i < . .. . _| x—_— \ 1‘ I- / J J. E" K, B U/ Boulderl . 0°— ” _ __ _ __ 4 ’9 / f . ___t , % DENVER fii CLEAR l ——— ‘ CREEK 5/ l 2_- a _-_. / l E I Breckenrid/e '— ‘ a / ‘ r~ / Z '1 I . f O I I . ‘ \ 0: l / // _1_ / LL__.__|. _ 3'9"— [L 'I/ / LL ! | _ \ / Colqrad » a g 9 , I ICripple _ __ _ ’ Creek | #‘V// / .Canon Cityll \ l /______ A __,,_L O 20 40 MILES APPROXIMATE SCALE FIGURE 1.—Location of Central City quadrangle in the Front Range, Colo. It is about 10 miles east of the crest of the range, in a region characterized by highly dissected, rolling up- land surfaces that slope to the east. Altitudes range from 11,204 feet in the northwestern part of the quad- rangle to 7,650 feet in the valley of Clear Creek at the south boundary. Local relief exceeds 1,000 feet in the Vicinity of the major streams but is less along smaller stream valleys. Bedrock is well exposed throughout most of the quadrangle. A notable exception is the extreme PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. northwestern part, where the upland surfaces are largely mantled by a rubble of the underlying bedrock or are covered by thin deposits of colluvium or glacial materials and where the major stream valleys are largely filled with till left by retreat of the latest valley glaciers. Meraines deposited by these glaciers extend downstream to an altitude of about 9,100 feet in North Clear Creek in the northern part of the quad- rangle and to about 8,600 feet in Fall River in the west-central part of the quadrangle. The Precambrian rocks of the Central City quad- rangle were first mapped early in the 20th century by E. S. Bastin (Bastin and Hill, 1917) at a scale of an inch to the mile. Later, the rocks within a segment of the area were remapped at the same scale by Lov- ering and Goddard (1950) as a part of their compilation of the geology of the entire Front Range. These stud- ies provided an excellent framework for subsequent detailed mapping by the US. Geological Survey in the Central City (Sims and Gable, 1964), Idaho Springs (Moench, 1964), Freeland-Lamartine (Harrison and Wells, 1956), Chicago Creek (Harrison and Wells, 1959), and Lawson-Dument—Fall River (Hawley and Moore, 1967) mining districts. A summary of the stratigraphy and a comprehensive discussion of the structure of the Precambrian rocks in the area of the detailed studies have been published by Moench, Harrison, and Sims (1962). Geologic mapping of the Central City quadrangle was done in two stages. The southern part was com- piled from the more detailed geologic maps (1:6,000) prepared for the Central City, Idaho Springs, and Lawson-Dument-Fall River mining districts, and a small area in the central part was compiled from mapping at the same scale by E. W. Tooker and A. E. Dearth. The rest of the quadrangle was mapped at a scale of 1:20,000 during the field seasons of 1959 and 1960 by the authors and P. D. Lowman, Jr. The areas of responsibility for mapping are shown on the index map on plate 1. Beth stages of the mapping were done under the supervision of P. K. Sims. The laboratory studies were carried on jointly by both authors. D. J. Gable is responsible for most of the quantitative mineralogic data. Several colleagues in the US. Geological Survey assisted in the study by providing mineralogic and chemical data; these indi— viduals are acknowledged at appropriate places in the report. The compositions of plagioclase were determined by oil-immersion methods that determined indices of refraction to an accuracy of i0.003. Modal analyses were made from standard %- by 1-inch thin sections. Sections were cut normal to lineation, and grain count- ing was done by a point counter with spacing of 0.5 E3 millimeter in one direction and 1 mm in the other (Chayes, 1949). On the average, 1,000 points were counted for each thin section. Average grain diameters as reported in the tables were determined from the thin sections from which modal analyses were made. Each thin section was divided into six or eight equal parts, and an average grain diameter was observed and measured for each section. Then these six or eight averages were added and were then reaveraged to yield a grain diameter to represent the entire thin section. Grain sizes given in the text are largely based on megascopic observation and are more meaningful for describing the physical appearance of the rock as a whole than are results from grain-size measurements determined in thin sections. GEOLOGIC SETTING The Front Range is a broad massive mountain unit 30—60 miles wide that extends from the vicinity of Canon City northward about 180 miles to the Wyoming State line (fig. 1). It has a Precambrian rock core which is flanked by steeply dipping Paleozoic and Mesozoic sedimentary rocks and which is locally overlapped by Cenozoic sedimentary and volcanic rocks (fig. 2). The range is crossed at about midlength by the Front Range mineral belt, a narrow northeast- trending belt of porphyritic igneous rocks and associated ere deposits of Laramide age. This belt contains all the important mining districts in the range, except Cripple Creek and the uranium mining areas in Jef- ferson County (Sims and Sheridan, 1964). The Precambrian rocks of the Front Range consist of roughly equal amounts of metamorphic and igneous rocks. The dominant metamorphic rocks are metasedi- mentary biotite gneisses and schists of several types and associated migmatites that were mapped as the Idaho Springs Formation by Ball (1906) and by Levering and Goddard (1950, p. 19—20, pls. 1, 2). Less common are (1) hornblende gneisses and amphib- olites of uncertain derivation, which were mapped as the Swandyke Hornblende Gneiss by Levering and Goddard (1950, pls. 1, 2), (2) quartzite as described by Wells, Sheridan, and Albee (1964), and (3) microcline paragneiss as described by Moench, Harrison, and Sims (1962). The microcline paragneiss previously was considered to be an orthogneiss by Levering and Goddard (1950, p. 23). Estimates of the thicknesses of these formations and their correlations throughout the range cannot yet be made confidently because of complex folding and because of the lack of detailed mapping in many areas. The igneous rocks are of several types and intrude the gneisses and schists. The dominant intrusive rocks E4 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY O 10 :20 3O 4OM|LES It:11|1|11] I I 4| FIGURE 2.——Generalized geologic map of the central part of the Front Range and adjacent areas. Modified from Tweto and Sims (1963, pl. 1). PETROLOGY AND STRUCTURE, PRECAMBRIAN EXPLANATION IGNEOUS AND METAMORPHIC ROCKS >— v v V n: . .. v ls }_ v v v j 0: Lu Volcanic rocks '— >.. g n: O E LLl l— U m < Lu . . . l— *— Laramide 1ntrus1ve rocks , 151.: o . u s Z . . <: Granltic rocks E m \\ \ \ f <2: \\ \ 9% ti \ \ o: 0. Metamorphic rocks Lines indicate general trend offoltatton; star pattern indicates trend not known J SEDIMENTARY ROCKS Mszs Sedimentary rocks and surficial deposits QTS, rocks of Quaternary and Tertiary age Mszs, rocks of Mesozoic and Paleozoic age Contact High-angle fault Dashed where inferred uH—L—A— A — A. _ A .— Thrust fault Dashed where inferred; barbs point toward upper plate ~ vs ~~ ~~ ~’\J ~ Shear zone FIGURE 2.—Continued. 247—419—67—52 ROCKS, CENTRAL CITY QUADRANGLE, COLO. E5 have been named the Boulder Creek Granite, the Silver Plume Granite, and the Pikes Peak Granite (Lovering and Goddard, 1950, p. 25—29); other intrusive rocks of smaller areal extent have been designated by lithologic terms. These lesser intrusive units include quartz dio— rite and hornblendite, which are known to form small bodies in the central. part of the range, and gabbro and related rocks, which were described from the area of this report by Taylor and Sims (1962). The older ig— neous rocks of the intrusive sequence, gabbro and re— lated rocks, granodiorite and associated rocks of Boul- der Creek affinity, and quartz diorite and hornblendite are partly metamorphosed and are interpreted as syn— tectonic intrusives; the rocks of Silver Plume affinity are interpreted as late syntectonic; and the Pikes Peak Granite probably is posttectonic. Pegmatites of gra— nitic composition are related to all intrusive rock types, and aplites are related to some. The metamorphic rocks of the Front Range have min- eral assemblages that conform generally with the alman- dine amphibolite facies of Turner and Verhoogen (1960, p. 544—550). The data from scattered localities indi- cate that the assemblages range from sillimanite-bearing assemblages, as in the region of this report, to lower grade biotite-chlorite—muscovite-quartz assemblages, as in the northern part of the range (W. A. Braddock, in US. Geological Survey, 1964, p. A94). The internal structure of the Precambrian rocks that constitute the core of the Front Range developed largely in Precambrian time. The rocks were folded and re- gionally metamorphosed, were locally sheared, and were faulted and jointed. The regional folding was complex. Reconnaissance mapping, mainly by Levering and God- dard (1950, pls. 1, 2), and later detailed mapping in selected areas indicate that the folds differ in trend from place to place but have systematic patterns locally. Details of the folding and metamorphism remain largely unknown except in the central part of the range, which has been mapped and studied intensively in recent years. (See particularly Moench, Harrison, and Sims, 1962; Tweto and Sims, 1963; Wells, Sheridan, and Al- bee, 1964.) In this area, the main structures in the Precambrian rocks resulted from three successive epi- sodes of deformation. The oldest deformation created broad warps and smaller associated folds whose axes trend northwest; the deformation appears to be partly syntectonic with the Boulder Creek Granite. This de- formation has been recognized (R. B. Taylor, oral commun., 1963) in the Black Hawk and adjacent quad- rangles to the east of the Central City quadrangle, but its full extent beyond this area is not known. A second period of deformation, probably only slightly later than the first episode, developed folds trending north-north- east. This period is the dominant episode of deforma- E6 tion recognized in the Central City quadrangle. In the Central City area this deformation was accompanied by regional dynamothermal metamorphism and migmati— zation and was partly syntectonic with all the intrusive rocks of the area except perhaps the Silver Plume Gran- ite. A third period of deformation began with folding on axes trending east-northeast and progressed into cat- aclasis along shear zones of the same trend. A major zone of shearing related to this deformation impinges on the southeast corner of the Central City quadrangle. The cataclastic deformation was followed by local fault- ing in at least two dominant directions trending north- west and north—northeast (Sims and others, 1963), which produced the initial structures containing the “breccia reefs” of the Front Range. Data on the ages of Precambrian metamorphic and igneous events in the Front Range still are too meager for accurate dating. The determination of reliable ages is hindered by the multiple deformations during the Precambrian and, within the mineral belt, also by the thermal and structural events that accompanied the Laramide revolution. Presently available data indi- cate that the older deformations (episodes 1 and 2, pre- ceding paragraph) probably took place about 1,700— 1,800 m.y. (million years) ago, for George Phair (in U.S. Geological Survey, 1964, p. A95) obtained a zircon- isotope age of 1,730 m.y. for undeformed granitic rock within the Boulder Creek batholith west of Boulder. Rocks of Boulder Creek affinity are thought to be syn- tectonic with all or parts of the older deformations, and thus they should give ages approximating the time of these deformations. The age of the Silver Plume Gran- ite should approximate the third deformation, for the granite is thought to be virtually syntectonic with the plastic stage of this deformation. K—Ar ages of 1,210 and 1,230 m.y. and Rb-Sr ages of 1,360 and 1,350 m.y. for muscovite and biotite were obtained on the Silver Plume Granite by Aldrich, Wetherill, Davis and Tilton (1958, p. 1130). During the Laramide revolution, the core of the Front Range was uplifted and hypabyssal igneous rocks and attendant ores were emplaced within the mineral belt. Mountain building was accomplished without marked internal deformation. Except for the formation of some new fractures in the mineral belt (Sims, Arm- strong, and others, 1963), the rejuvenation of older frac- tures, and the formation of a regional joint set (Harrison and Moench, 1961, p. B5—B12), the Precambrian core apparently was not appreciably deformed during the Laramide. Intrusion of the hypabyssal rocks was largely confined to the narrow strip of ground that con— stitutes the mineral belt. Most of the intrusives were emplaced as dikes, sills, and small stocks, larger stocks were emplaced along the northwest margin of the belt. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY The larger stocks are known, from study of the Tertiary stock at Eldora (Hart, 1964), to have contact meta- morphic halos in which microcline in Precambrian gneisses is changed to orthoclase and isotopic mineral ages are reduced. Comparable mineralogic changes were observed by us adjacent to the stock near Apex in the Central City quadrangle. The full significance and extent of the thermal metamorphism of the Tertiary intrusives is not yet known, and further studies are needed to aid in interpreting the geochronology of the Precambrian rocks. ROCK UNITS The Precambrian rocks of the quadrangle are domi- nantly microcline—quartz—plagioclase-biotite gneiss and biotite gneiss units. These rocks contain small lenses of other metamorphic rocks and are intruded by generally small bodies of granodiorite and associated rocks, gabbro and related rocks, quartz diorite and hornblendite, biotite-muscovite quartz monzonite, and pegmatites of several types. The terminology of the rock units described in this report accords generally with that used previously in the report on the Central City district (Sims and Gable, 1964). The lithologic names are assigned on the basis of quantitative mineral content and on the presence of diagnostic minerals. Where mineral as- semblages are given, the minerals are listed in alpha— betical order without regard to relative abundances. METAMORPHIC ROCKS Metamorphic rocks dominate the bedrock in the quad- rangle. Microcline-quartz—plagioclase—biotite gneiss, hereafter called microcline gneiss in the text, is inter- layered on a gross scale with biotite gneiss to constitute the lithologic framework of the district (pl. 1). Am— phibolite, cordierite—amphibole gneiss, and calc-silicate gneiss and related rocks form small lenses and pods in the microcline gneiss units; and calc-silicate gneiss and associated quartzite and amphibolite occur sparsely in the biotite gneiss units. Internally, the biotite gneiss units are variable in composition and are migma- tized. They consist dominantly of interlayered biotite- quartz-plagioclase gneiss, sillimanitic biotite gneiss, and garnet— and cordierite-bearing sillimanitic biotite gneiss. The garnet- and cordierite—bearing biotite gneiss is distinguished separately from the other biotite gneiss units on plate 1. LITHOLO GIG SUCCESSION The metamorphic rocks constitute a well-defined lithologic succession that seems to represent a normal stratigraphic sequence. Except for rocks in the upper part of the succession, the stratigraphic order has been PETROLOGY AND STRUCTURE, PRECAMZBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. defined previously in the summary report by Moench, Harrison, and Sims (1962, p. 38—39, pl. 2). The succession consists of three principal layers of microcline gneiss, each at least 1,000 feet thick, sepa- rated by biotite gneiss units of comparable thickness, as summarized in table 1. The major microcline gneiss units are, from lowest to highest, the Big Five, the Quartz Hill, and the Lawson layers. In earlier reports the Big Five layer was called the Idaho Springs layer, and the Quartz Hill layer, the Central City layer. The biotite gneiss layers have not been similarly designated. An estimated maximum of 15,000—16,000 feet of strata is exposed in the quadrangle. A biotite gneiss unit exposed on the crest of the Idaho Springs anticline, in the southeast corner of the quadrangle (pl. 1), is interpreted stratigraphically as the lowest unit in the area. The unit is estimated to exceed 1,000 feet in thickness, but its lower part is not exposed in the quadrangle. It consists mainly of sillimanitic biotite gneiss but contains layers of biotite- quartz-plagioclase gneiss and, locally, quartz gneiss. TABLE 1.——Lithologic succession. of Precambrian metamorphic rocks, Central City quadrangle [From highest to lowest, stratigraphically] Estimated maxi- mum thickness Rock and description (feet) Biotite gneiss: Dominantly migmatized interlayered sillimani- tic biotite gneiss and biotite—quartz—plagio- clase gneiss, with lenticular zones of cordier- ite-bearing biotite gneiss, garnetiferous biotite gneiss, and garnetiferous sillimanitic biotite gneiss. Intertongues in lower part with microcline gneiss of Lawson layer. Top of unit not exposed in quadrangle-__- Microcline gneiss (Lawson layer): Ranges in composition from quartz monzonite to granodiorite; contains several small bodies of amphibolite and local lenses of biotite gneiss. Unit thins and is highly folded in the area between Fall River and Peeks Flat. In northern part of quad- rangle, upper part of unit intertongues with overlying biotite gneiss unit ______________ Biotite gneiss: Consists dominantly of migmatized inter- layered sillimanitic biotite gneiss and bio- tite-quartz-plagioclase gneiss with lenses of garnetiferous biotite gneiss and garnetr and cordierite—bearing sillimanitic biotite gneiss. Locally contains lenses of calc—silicate gneiss, amphibolite, and microcline gneiss__ Microcline gneiss (Quartz Hill layer): Average composition is granodiorite; contains several thin layers and lenses of biotite- quartz-plagioclase gneiss and pods of am- phibolite, calc—silicate gneiss, and cordierite— amphibole gneiss. Unit thins to south____ Biotite gneiss: Dominantly sillimanitic biotite gneiss; unit probably pinches to southwest ___________ Microcline gneiss (Big Five layer): Small lenses of amphibolite occur along margins; unit pinches out at depth and to southwest but thickens to east in adjacent quadrangles ____________________________ Biotite gneiss: Dominantly sillimanitic biotite gneiss. Bot- tom of unit not exposed im mapped area__ >2, 500 2, 500 3, 000—4, 000 3, 000 2, 000 1, 000 >1, 000 E7 Above the biotite gneiss unit is a discontinuous layer of microcline gneiss, designated the Big Five layer, which has an estimated maximum thickness of 1,000 feet. This layer is more felsic and generally more massive than the other major layers of microcline gneiss in the area. It pinches out at depth and toward the southwest on the northwest limb of the Idaho Springs anticline but thickens eastward. The Big Five layer of microcline gneiss is overlain by a biotite gneiss unit that has an estimated maximum thickness of 2,000 feet. The biotite gneiss forms a curved outcrop pattern on the crest of the Idaho Springs anticline. It is lithologically similar to the lowest recognized biotite gneiss unit. Above this biotite gneiss unit is the Quartz Hill layer of microcline gneiss, the dominant exposed unit in the eastern part of the quadrangle. It crops out on the axis of the Central City anticline and forms a shieldlike out- crop near Central City and an irregular prong-shaped mass along the steep slopes of the valleys of Clear Creek and Fall River. The structural configuration of the unit is clearly shown on plate 2 of the report by Moench, Harrison, and Sims (1962). The unit is known from studies in the Central City district (Sims, Drake, and Tooker, 1963 ; Sims and Gable, 1964) to have an average composition of granodiorite and a maximum thickness of about 3,000 feet. The Quartz Hill layer is overlain by a biotite gneiss unit that crops out continuously in a 2-mile—wide north- eastward-trending band across the central part of the quadrangle. The maximum thickness of the layer has been estimated from exposures along Clear Creek and Fall River to be about 4,000 feet (Moench, Harrison, and Sims, 1962, table 1). In the vicinity of North Clear Creek (see pl. 1, section A—A’), in the northern part of the quadrangle, the unit seems to be thinner and prob- ably does not exceed 3,000 feet in thickness. This biotite gneiss unit is remarkably diverse in composition, as noted in table 1. A major zone of lenses of garnet- and cordierite-bearing biotite gneiss occurs locally in the middle of the unit, and a thin, discontinuous zone of lenses of the same rock type lies at or near the top. Above this biotite gneiss unit is the Lawson layer of microcline gneiss, originally defined by Moench, Harrison, and Sims (1962) from exposures near Lawson (just west of the Central City quadrangle, in the valley of Clear Creek). The Lawson layer is inferred to extend discontinuously and irregularly northeastward across the quadrangle, its northern extremity in the quadrangle being in the vicinity of Gamble Gulch (pl. 1). According to this interpretation, the layer thins drastically from the west margin of the quad— rangle, is intricately folded between Fall River and Pecks Flat, and apparently pinches out at Peeks Flat on E8 the crest of Peeks Flat anticline. In the vicinity of Peeks Flat it is cut out by the Mount Pisgah pluton of granodiorite and associated rocks but reappears north- east of the pluton. The upper part of the microcline gneiss unit intertongues with the overlying biotite gneiss unit northeast of the pluton; possibly the microcline gneiss that encloses the body of gabbro and related rocks on the axis of the Arizona Mountain anticline is a major tongue of this unit. The unit is estimated from expo- sures near Lawson to exceed 2,500 feet in thickness. Mapping in the Central City quadrangle indicates that the maximum thickness in the vicinity of Blackhawk Peak also is about 2,500 feet. Above the Lawson layer is a layer of biotite gneiss that is interpreted stratigraphically to be the highest unit in the map area. It occupies the northwest quar- ter of the quadrangle. We estimate that about 2,500 feet of strata in the unit is exposed in the quadrangle, but this estimate is less accurate than the estimates for other units because of complex folding and poor exposures—the top of the unit is not exposed. As indicated above, the lower part intertongues with the upper part of the Lawson layer of microcline gneiss. Lithologically, the unit resembles the biotite gneiss unit that underlies the Lawson layer. Garnet- and cor- dierite-bearing biotite gneisses occur locally throughout the unit. The major rock type is sillimanitic biotite gneiss. MICROCLINE-QUARTZ-PLAGIOGLASE-BIOTITE GNEISS Microcline gneiss forms three major stratigraphic layers—~from oldest to youngest, the Big Five, Quartz Hill, and Lawson layers—and scattered smaller lenses and layers within larger masses of biotite gneiss. Although the layers differ somewhat in detail, the gross lithologies are similar, and the rock type is discussed as a unit herein. GENERAL CHARACTER The microcline gneiss is a distinctive rock unit, moderately variable in composition, that can be dis- tinguished by its granitic appearance, generally con- spicuous layering, and well-developed foliation. It is a light— to medium-gray, fine- to medium- grained, equigranular, layered rock. Weathered surfaces are various shades of yellowish gray, gray orange pink, and very pale orange. Layering is the result of alternating layers of slightly different mineral composition and, at places, of regular paper-thin par- allel streaks of biotite. The microcline gneiss contains less biotite and, accordingly, is lighter in color and has a more uniform and straighter layering than the associated biotite gneisses. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY The microcline gneiss within each of the major layers is associated with amphibolite and to a lesser extent with other metasedimentary rock units. The amphi- bolite forms thin, concordant lenses and stubby layers within, and at the margins, of the gneiss bodies. Some lenses, discontinuous along both strike and dip, appear to occur at similar stratigraphic positions within certain gneiss layers and probably constitute stratigraphic marker beds. The Quartz Hill layer, which has been studied in detail, contains small concordant bodies of calc-silicate rocks and cordierite-amphibole gneiss as well as amphibolite. It also has a few layers of biotite gneiss that are remarkably persistent (Sims and Gable, 1964). Contacts of the various rock units with the microcline gneiss generally appear sharp, but gradations across a few inches or a few feet are common. Biotite gneiss grades into microcline gneiss through gradual increases in microcline and diminution of bio- tite across the transition zone; adjacent to amphibolite, the microcline gneiss generally contains hornblende as the dominant mafic mineral rather than the more common biotite. The northern segment of the Lawson layer is partic- ularly heterogeneous. The contact zones are marked by interlayering of biotite gneiss with microcline gneiss and by intergradations, both along and across strike, of the two rock types. The gross interfingering of the two rock types is shown by the several tongues of boi- tite gneiss that extend into the Lawson layer on both sides of North Clear Creek (pl. 1). An intertonguing on finer scale is also present but could not be shown on plate 1. The central part of the layer is relatively homogeneous except in the area north of Blackhawk Peak where it contains numerous beds of amphibolite. PETRO GRAPI-IY The microcline gneiss is generally an equigranular rock of allotriomorphic granular texture, but at a few places it is inequigranular and contains subhedral pla- gioclase and sparse potassium feldspar crystals as much as 5 mm in diameter. As detailed mineralogic data and modal analyses have been given previously for the rock unit within the Central City district, only that part of the region not covered previously is discussed herein; representative modes of the Lawson layer, par— ticularly that part north of Fall River, and modes for other scattered bodies are listed in table 2. The modal data are summarized in the triangular diagram in fig- ure 3. The earlier reports on the Central City district (Sims and Gable, 1964), Idaho Springs district (Moench, 1964), and the Lawson—Dumont-Fall River district (Hawley and Moore, 1967) contain additional information on this rock unit. PETROLOGY AND STRUCTURE, PRECAMBRIAN Quartz Average for Lawson layer Average for scattered bodies (21 modes) Average for Quartz Hill layer (66 modes) Potassium feldspar Plagioclase FIGURE 3.—Variation in composition (volume percent) of mi- crocline gneiss but excluding Quartz Hill layer. 0, Lawson layer; 0, scattered bodies of microcline gneiss. The microcline gneiss within each of the layers that have been mapped is similar in gross aspect to that described from the Quartz Hill layer (Sims and Gable, 1964). Plagioclase, quartz, and potassium feldspar, the dominant minerals, are intergrown in anhedral or rarely subhedral grains. Except locally, the plagio- clase has well-defined albite twinning; Carlsbad twins and pericline twinning are less common. Narrow albitic rims or myrmekite are common at contacts of plagioclase with potassium feldspar. In general, the plagioclase contains a few percent of potassium feldspar as small patches alined parallel to the twin lamellae to constitute antiperthite. The quartz occurs mainly as irregular anastomosing grains, which show strain shadows and which are interstitial to the feld- spar grains, and occurs also as small subrounded inclusions in other minerals or as myrmekitic inter- growths with biotite, plagioclase, and rarely muscovite. The potassium feldspar has both conspicuous grid twinning and fine perthitic intergrowths of plagioclase. The perthitic intergrowths constitute an estimated 10 percent by volume of the potassium feldspar grains. The potassium feldspar was determined in two speci- mens from the Quartz Hill layer to be near maximum microcline and to contain about 80 percent KAlSiaog (Sims and Gable, 1964). The biotite is a distinctly greenish strongly pleochroic type; optical data and chemical analyses of two specimens from the Quartz Hill layer have been reported by Sims and Gable (1964, p. 014). Muscovite has two modes of occur- ROCKS, CENTRAL CITY QUADRANGLE, COLO. E9 rence. Locally it is intergrown with biotite, plagio- clase, and potassium feldspar and appears to be primary, but much of it occurs as patches in microcline and plagioclase and as overgrowths on biotite and appears to be secondary. Magnetite, zircon, and apatite are the most common accessory minerals. Hornblende is a local accessory mineral, occurring especially adjacent to amphibolite bodies. Garnet is a widespread acces- sory mineral in the northern part of the Lawson layer, particularly north of North Clear Creek (see samples 15, 17, 18, and 21 of table 2); it is most abundant and conspicuous along the northwest margin of the layer north of Blackhawk Peak (pl. 1). Typically, altera- tion of the rock is slight; the plagioclase is partly clouded by clay minerals. Chlorite, epidote, and calcite are local alteration products. Typically, the microcline gneiss has a crystallo- blastic texture, and the dominant minerals are inter- grown in a mosaic pattern. Quartz and the feldspars are nearly equidimensional, and biotite forms plates that are slightly elongate parallel to the lineation of the rock. On the east slope of Dakota Hill, east of the stock at Apex, the microcline gneiss in the extreme northwestern part of the Lawson layer is profoundly granulated and altered. Megascopically, the rock is noticeably finer grained and has a more pronounced foliation and lineation than elsewhere. In thin section, the gneiss is seen to be strongly granulated and recrystallized. In contrast to the common mosaic texture, the quartz forms elongate, amebiform aggregates of sutured grains as much as an inch long that form fingerlike projections through a finer grained groundmass of feldspar. Bio— tite is fine grained, typically frayed, and streaked out parallel to the quartz aggregates. The feldspars are strongly altered to clay minerals and sericite. The potassium feldspar has dark shadowy extinction, is microperthitic, and lacks the grid twinning which characterizes it elsewhere. Also, because twin lamellae in plagioclase are partly destroyed, distinction between the two feldspars is very difficult. CHEMICAL COMPOSITION That the microcline gneiss varies in chemical com- position from a quartz diorite to a granite can be inferred from figure 3. 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ASTOOV 591535 25% EN nudism ac ”Emu “32:58 89¢ .33» was .mn .NH .3 .2 .m .w .N. 6 NMV‘D A E12 AMPHIBOLITE OCCURRENCE AND CHARACTER In the quadrangle, amphibolite occurs as small concordant lenses Widely dispersed in each of the microcline gneiss layers and less commonly as local lenses in the biotite gneiss layers. It also occurs as boudins in these rocks. Many of the lenses lie at or near the contact of major layers of microcline gneiss with biotite gneiss and range in width from a few inches to about 500 feet and in length from about a foot to at least 3,000 feet. Most are 1—20 feet thick and a few tens to a few hundred feet long. Accord- ingly, only the larger lenses can be shown at the scale of the geologic map (pl. 1). Megascopically, contacts with the enclosing gneiss appear sharp, but in detail the amphibolite is seen to grade transitionally into the gneiss. Not uncommonly, pegmatite occurs along the contacts and intrudes both the amphibolite and the enclosing gneiss; the intruding pegmatite forms irregular crosscutting veinlets and stringers in the amphibolite and thus produces blocky forms that contrast sharply with the lit-par—lit structure of the migmatized biotite gneisses. Within the Quartz Hill layer of microcline gneiss some bodies of amphibolite are associated with calc-silicate gneisses and cordierite—amphibole rocks. The amphi- bolite bodies in biotite gneiss layers commonly are associated with calc-silicate gneiss or skarn. The amphibolite is a grayish black or medium gray, » predominantly medium grained, generally homogeneous SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY rock. Typical varieties have a uniform salt-and-pepper appearance; other varieties are finely layered as a re- sult of segregation of the minerals into dark and light layers. The rock is more massive in appearance than other rock types in the region but has a moderately well developed foliation and lineation. Lineation is ex- pressed mainly by the alinement of hornblende crystals. PETROGRAPHY Typical amphibolite in the quadrangle contains horn- blende in excess of plagioclase, less than 10 percent quartz, and at places a few percent of clinopyroxene (samples 1, 2, and 5 of table 3). A variety of the rock herein referred to as hornblende gneiss contains about 25 percent hornblende, as much as 70 percent plagio- clase, and generally sparse quartz (samples 6, 7, and 8 of table 3). Hornblende gneiss constitutes some of the mappable bodies within the Lawson layer of microcline gneiss on the east slope of Dakota Hill (pl. 1). All vari- eties of the rock have a dominantly hypidiomorphic granular texture. The petrography of the amphibolite is similar to that described previously from the Central City dis- trict (Sims and Gable, 1964). The plagioclase is dom- inantly andesine and occurs as slightly cloudy anhedral or subhedral crystals. Most grains have well-developed simple twinning. Potassium feldspar occurs as tiny blebs in the plagioclase to constitute antiperthite; it occurs less commonly as small interstitial grains. Hornblende is the common green variety; it tends to TABLE 3.—Modes, in volume percent, of amphibolite [Tr, trace; Nd, not determined; __., not found. Field number is in parentheses after description of sample. For chemical and spectrochemical analyses and norms of samples 1, 2, and 3 see corresponding numbered sample, table 4] Average of Mineral 1 2 3 4 5 6 7 8 19 modes, Central City district Potassium feldspar _____________________________ 0. 6 ________________ 0. 1 0. l ________ 0. 3 0. 4 Plagioclase ____________________________ 30. 5 36. 1 40. 4 13. 1 35. 5 70. 0 70 6 55. 5 43. 0 Quartz _______________________________ 6. 8 7. 1 9. 0 12. 9 1. 9 . 8 ________ 9. 4 4. 8 Biotite _______________________________ . 3 1. O ________ 3. 1 . 1 1. 0 . 1 . 5 1. 7 Homblende ___________________________ 58. 9 51. 7 44. 7 64. 5 59. 4 23. 2 28. 3 28. 9 44. 4 Clinopyroxene _________________________ 1. 9 2. 2 ________________________ 2. 7 ________________ 1. 3 Muscovite ____________________________________________________ 3. 4 ________________ Tr ________ Opaque iron oxides _____________________ . 4 Tr 3 0 . 7 1 5 Tr . 6 4. 6 Zircon ________________________________________ . 1 ________ . 4 ________ . 1 ________ . 1 Sphene _______________________________________ . 6 ________________________ . 4 Tr . 2 4 4 Chlorite ______________________________________________________ 1. 1 . 6 ________ Tr ________ ' Epidote ______________________________________ Tr ________ Tr . 1 . 4 . 3 . 1 Calcite _______________________________ . 9 . 1 2. 6 . 5 . 5 1. 3 ________________ Apatite _______________________________ . 3 . 5 . 3 . 3 3 Tr 1 . 4 Total _____________________________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. O 100. 0 Composition of plagioclase ______________ An“ A1130 An4g Nd Nd Nd Nd Nd Ana; Average grain diameter ___________ mm__ 0. 4 0. 5 Nd 0. 4 0. 7 0. 6 0. 5 0. 5 __________ 1. From mine dump south of the head of Peeks Gulch. (CC—309—B) 2. From thin layer of amphibolite in Central City layer of microcline gneiss, caved adit 0.5 mile northeast of Missouri Falls. (CC—730) 3. From dump of Grand Army shaft. Central City district. (8472—0—53) 4. From south margin of crescent-shaped microcline gneiss layer, nose of ridge south of North Clear Creek and west of Pine Creek. (CC-616—B) 5. From thick layer within the Lawson layer of microcline gneiss, south side 0! Stewart Gulch 0.5 mile west of Missouri Gulch. (CC—1001) 6, 7. Spotted plagioclase-hornblende gneiss, from layer within Lawson layer of mi- crocline gneiss, crest of hill north of Stewart Gulch. (CC—9844, 00-985) 8. Plagioclase-hornblende gneiss, from west margin of the Lawson layer of micro- cline gneiss, saddle west of Stewart Gulch. (CC—996) PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. be poikilitic and is intergrown with plagioclase to form a mosaic pattern. In a few samples clinopyroxene is intergrown with the hornblende; it is a very pale green, slightly pleochroic variety and is altered to hornblende along cleavage planes and grain boundaries. Quartz forms anhedral interstitial grains. Biotite is local in occurrence and is intergrown with but mainly secondary after hornblende. The biotite is partly altered to chlo- rite. Other alteration minerals include calcite, epidote, and muscovite. Opaque iron oxides, apatite, zircon, and sphene are common accessory minerals. The amphibolite has a typical crystalloblastic texture, indicative of virtually contemporaneous crystallization of the dominant minerals—hornblende, plagioclase, and quartz. Pyroxene crystallized locally with the horn- blende but subsequently was partly altered to horn— blende. Biotite also apparently is secondary. CHJEMICAL COMPOSITION Chemical and spectrochemical analyses of three typical specimens indicate some variation in chemical composition of the amphibolite (table 4). In general, the amphibolite is similar in composition to plateau basalt (Washington, 1922, p. 774), but it contains slightly more silica and alumina and less magnesia and titania; also, it overlaps the range in chemical composition of spilites (Reed, 1957, p. 37) but has less soda than does the typical spilite. The amphi- bolite contains less chromium and nickel than do most analyzed mafic igneous rocks (Engel and Engel, 1962, p. 65). The amphibolite from Central City is closely similar chemically to the para-amphibolites from North Carolina (Wilcox and Poldervaart, 1958, p. 1351). Variations in alumina content of the rocks correlate directly with changes in the amount of modal plagio- clase; the relatively low amount of calcium oxide in sample 3 (table 4) reflects a relatively small proportion of hornblende in the amphibolite. TABLE 4.—Chemical and spectrochemical analyses and norms of amphibolite [Laboratory number given in parentheses below sample number. Results of chemi- cal analyses given in weight percent and of spectrochemical analyses, in parts per million. Normative composition or mesonorm computed by method of Barth (1959 1962). For mode and sample description and locality, see corresponding num cred sample, table 3. Dorothy Powers and P. R. Barnett, analysts] Sample 1 2 3 (03102) (G3103) ((33104) Chemical analyses Si02 _____________________ 50. 01 49. 43 48. 54 A1.o, ____________________ 13. 86 15. 02 17. 18 Fe103 ____________________ 2. 95 3. 11 3. 61 FeO _____________________ 10. 35 8. 73 10. 61 MgO _____________________ 5. 99 5. 88 5. 42 CaO _____________________ 10. 33 10. 18 7. 87 NazO ____________________ 2. 57 3. 38 3. 14 24 7—4.1 9—67~—-.3 E13 TABLE 4.—Chemical and spectrochemical analyses and norms of amphibolite—Continued Sample 1 2 3 (G3102) (03103) ((33104) Chemical analyses—Continued K20 _____________________ . 57 1. 06 34 H20+ ___________________ 1. 18 . 95 1 46 20* ___________________ . 08 . 06 10 T10, _____________________ 1. 19 1. 33 1 05 205 _____________________ . 12 . 20 13 MnO ____________________ . 22 . 21 29 002 ______________________ . 33 . 27 27 Cl _______________________ . 02 . 05 01 F ________________________ 07 . 09 09 S ________________________ 10 . 01 04 Subtotal ______________ 99. 94 99. 96 100. 15 Less 0 ___________________ . 08 . 06 . 06 Total ________________ 99. 86 99. 90 100. 09 Bulk density ______________ 3. 02 2. 95 2. 96 Powder density ___________ 3. 06 3. 04 3. 01 Spectrochemlcal analyses Co _______________________ 50 39 34 Cr _______________________ 60 200 16 Cu ______________________ 73 4 16 Ga ______________________ 23 23 2 La _______________________ < 100 < 100 < 100 Ni _______________________ 44 49 1 Pb _______________________ <30 <30 <30 Sc _______________________ 77 55 67 Sr _______________________ 150 250 180 V ________________________ 410 320 410 Y _______________________ 50 50 40 Yb ______________________ 5 4 4 Zr _______________________ 90 140 100 Normative compositions Quartz ___________________ 0. 52 __________ 2. 66 Potassium feldspar ________ 3. 45 6. 50 __________ Plagioclase : Albite _______________ 23. 70 25. 12 28 90 Anorthite ____________ 5. 65 5. 85 __________ Biotite _______________________________________ 3. 28 Hornblende: Actinolite ____________ 58. 80 ____________________ Edenite ________________________ 17. 84 __________ Hornblende _____________________ 34. 43 50 62 Pyroxene: Diopside _____________ . 52 2. 84 __________ Hypersthene ______________________________ 1. 06 Magnetite ________________ 3. 16 3. 28 3. 85 Sphene (total titanium) ____ 2. 58 2. 82 2. 25 Apatite __________________ . 27 . 43 . 27 Pyrite ____________________ . 27 . 03 . 09 Corundum ____________________________________ 6. 27 Calcite ___________________ . 84 . 68 . 68 O ORDIERITE-AMPE‘IB OLE GNEISS Cordierite- and amphibole-bearing gneisses occur as scattered stubby lenses predominantly within the Quartz Hill layer of microcline gneiss. The lenses are elongate parallel to the regional foliation and range from a few tens of feet to 400 feet in length and from a few feet to about 125 feet in width. Most lenses are in E14 direct contact with the enclosing gneiss or a pegmatite, but a few are adjacent to amphibolite, calc-silicate gneiss, or to hornblende-calcic plagioclase—quartz gneiss. The larger lenses are shown on plate 1 of this report but are given in more detail on the geologic map of the Central City district of Sims and Gable (1964). One small lens in biotite gneiss was mapped on the hill west of Missouri Lake (pl. 1). The gneisses are variable in composition but consist mainly of three principal rock types, which in order of decreasing abundance are cordierite-gedrite rocks, horn- blende- and cummingtonite-bearing rocks, and cor- dierite-biotite rocks. Previously (Sims and Gable, 1963) the cordierite-gedrite rocks were called cordierite- anthophyllite rocks. Commonly, the cordierite- gedrite rocks and the hornblende- and cummingtonite- bearing rocks are interlayered on different scales; cor- dierite—biotite rocks locally are intercalated with them. the complexity of the interlayering is shown schemat- ically by the sketch in figure 4. In detail, the layering is still more complex, for each layer has a finer scale layering resulting from segregation of the constituent minerals in different proportions. CORDIERITE-GEDRITE BOOKS The cordierite-gedrite rocks are dark—gray or medi- um—gray and generally medium grained gneisses; but SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY fine-grained and, locally, very coarse grained varieties are present. Freshly broken surfaces have a distinctly greasy, lustrous appearance, whereas weathered surfaces are grayish brown or reddish brown and are commonly ribbed. The layers rich in gedrite (for example, sam— ple 1, table 5) tend to be coarse grained and to have a blocky appearance; the gedrite forms either radiating bundles or columnar aggregates as much as 5 centime- ters long that radiate from a common center which con- tains abundant cordierite. The quartz-rich layers are finer grained and more inequigranular. Garnet and, to a lesser extent, cordierite typically are porphyroblastic and form crystals as much as 3 cm in diameter. Except for the fibrous gedrite, the tabular and fibrous minerals generally impart a conspicuous foliation and lineation to the rock. The rocks typically have a granoblastic texture but locally are granulated. Where the rocks are granulated, quartz forms elongate plates parallel to the gedrite laths and has serrate grain boundaries against cordierite and plagioclase. The cordierite—gedrite gneiss contains quartz, garnet, biotite, and plagioclase as well as the gedrite and the cordierite as major minerals, but the amounts and pro- portions of each of the minerals differ substantially from layer to layer (table 5). Magnetite—ilmenite, apatite, TABLE 5.—-Modes, in volume percent, of representative varieties of cordierite-gedrite gneiss and associated rocks [Tr, trace; Nd, not determined; __-., not found. Field number is in parentheses after description of sample] Cordierite— Homblende- and cummingtonite-bearing Cordierite—gedrite rocks biotite rocks Mineral racks 1 2 3 4 5 6 7 8 9 10 11 12 Quartz ______________________________ 44. 4 24. 4 4 O 1. 7 51. 5 52. 7 66 6 20. 9 12. 8 8. 7 3. 0 Plagioclase ________________________________ . 3 9 0 . 4 ______ 14. 5 ________ 16. 7 31. 7 54. 1 52. 2 Cordierite _____________________ 3. 8 29. 6 29. 7 29 3 38. 6 27. l 29. 6 25. 5 ______________________________ Gedrite _______________________ 90. 0 2. 2 23. 4 37. 9 44. 3 10. 4 1. 3 ______________________________________ Cummingtonjte-hornblende ________________________________________________________________________ 30. 2 40. 7 Homblende ______________________________________________________________________ 47. 2 52. 0 ______________ Magnetite—ilmenite _____________ .x 1 . 7 . 2 1. 6 ______ . 1 Tr . 1 2. 8 3. O 2. 0 3. 5 Biotite ________________________ 5. 4 8. 8 4. 5 10. 2 13. 8 8. 0 1. 9 7. 5 ________ . 2 ______________ Apatite _____________________________ Tr ...... Tr __________________________ 1 . 3 .9 6 Zircon ________________________ . 4 . 1 Tr . 1 . 1 Tr Tr Tr ______________________________ Chlorite _____________________________ . 3 ______ Tr 1. 1 2. 7 ______ 2 4. 3 ________ 2. 6 ______ Spine] _________________________ . 3 ____________ Tr ________________________________________________________ E idote _________________________________________________________________________ 1. 2 ________ . 4 ______ uscovite (secondary) ________________________________________ 2 Tr 1 5. 1 Tr 1. 0 ______ Garnet ______________________________ 13. 9 17. 5 7. 9 __________________________ 1. 7 Tr . 1 ______ Sillimanite--- _ _ _« _____________________ Tr ________________________ Tr Tr ______________________________ Corundum _____________________________________________ Tr __________________________________________________ Total _____________________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. O 100. 0 100. 0 100. 0 100. 0 Composition of plagioclase ___________________ Nd Nd Nd ______ A1125 ________ Ammo Anus-” Ann-“ Nd Average grain diameter- ___mm-_ Nd 0. 5 1. 7 0. 6 0. 6 0. 5 Nd 0. 6 0. 3 0. 4 0. 3 0. 6 1. Cordierite-gedrite-biotite gnless, from small body on crest of Negro Hill. (S 755- 6. Cordierite-gedrite-quartz-biotite gneiss, same locality as 5. (S 257—7—52) 2. Cordieritevgamet-gedrite—quartz gneiss, same locality as l. (S 755-]3“—53) 3. Cordierite—gedrita—gamet-quartz-biotite gneiss, from small lens at margin of gedrite body on crest of Central City anticline, roadcut on north side of North Clear Creek. (B—17-1) . 4. Cordierite—gedrite—gamet-biotite-plagioclase—quartz gneiss, from small body near crest of Quartz Hill. (S 53-A-52) 5. Cordlerl edrife-biotite gneiss, from small body 1,800 ft south of the Patch, Quartz ill. (S 257—1—52) 7. Cordietite-gedrlte-plagioclase quartz gneiss, same locality as 5. (S 257—8-52) 8. Cordierite-quartz—biotite gneiss, from small body 1,500 ft south of the Patch, Quartz Hill. (S 251-B—52) 9. Garnet-homblende—plagioclase‘quartz gneiss same locality as 5. (S 257—6—52) 10. Homblende-plagioclasequartz gneiss, same locality as 1 and 2. (S 7541-1—53) 11. Cummingtonite-hornblende-plagioclase-quartz gneiss, same locality as 5. (S 257— 12. Cgfirmggonlte-hornblende-plagioclase-quartz gneiss, same locality as 11. (S PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. TABLE 6.—C’hemical analyses and modes of cordiert‘te-gedrt'te gneiss and associated rocks [Serial number given in parentheses below sample number. Results 01 chemical analyses given in weight percent and modes, in volume percent. Tr, trace; Nd not determined; _.__, not found. Analysis of sample 6 by C. L. Parker; another analylsels by E. L. Munson. Field number is in parentheses after descriptlon of samp e Cordi— Cum- 5 k till; 3133‘? Cordierite- ite me 5 o e o e- 86 rocks bearing rocks 1 2 3 4 5 6 7 (H3461) (H3457) (H3458) (H3460) (H3462) (14196) (H3459) Chemical analyses 48. 09 65. 18 52. 64 76. 86 76. 94 74. 02 52. 64 18. 75 14. 97 22. 20 7. 40 7. 87 11. 03 18. 23 2. 04 . 67 1. 44 1. 00 1. 03 . 63 1. 23 17. 55 7. 76 9. 73 6. 11 7. 27 4. 50 11. 12 8. 76 6. 75 7. 85 5. 85 4. 68 5. 04 4. 86 .58 .31 .48 .06 .04 .07 6.53 .35 .25 1.91 .32 .41 . 1.60 .29 .12 .83 .22 .18 1.74 .40 . 53 . 13 . 09 . 04 . 04 . . 79 1. 59 2. 77 1. 88 1. 56 . 95 1. 83 1. 12 .16 .15 .15 .29 .13 .17 .12 105 .76 .37 .09 .13 .19 117 1 . 15 01 . 01 . 01 . 01 32 01 . 02 01 . 00 . 01 . . . 01 . 18 __________ 99. 89 99. 99 99. 59 99 81 99.69 99. 75 100. 16 Mode: 20 55 61 59 20 40 19 13 21 .......... 20 24 24 .......... 17 9 2 2 17 1 Tr 19 6 2 43 5 ________________ Tr __________ Tr Tr Tr Tr Magnetite- ilmenite __________ 2 2 Tr Tr Tr Tr Tourmaline. Tr Tr Alteration minerals ......... Tr 3 Tr Tr Tr __________ Tr Pyrrhotite(?) .............. Tr Zircon. . Tr Tr Tr .................... Pyrite. . _ Tr , Total accessory minerals ......... Tr Tr Tr Tr Tr 1 T1- Total __________ 109 100 100 100 100 100 100 Composition of plagioclase ....................... Nd ________________ Nd An" 1. Cordieritegamet-gedrite-quartz gneiss, from small body on south slope of Quartz Hill, near Leavenworth mine. (863—A—52) 2. Cordierite-gedrite-quartz gneiss, same locality as 1. This phase is interlayered with 1. (S63—C—52) 3. Cordierite- edrite-plagioclase—quartz—spinel—biotlte gneiss, from small body on west-sout west slope of Negro Hill at about the 8000-“ contour. (8682—11—53) 4. Coggisirike-ggdrite-quartz-blotite gneiss, from small body on crest of Negro Hill. 5. C(gdigrite-gedrite-quartz—biotlte gnelss, irom small body same locality as 4. (8755- —5 6. Biotite—cordierite-quartz gneiss, from small lens 0.25 mile south of the Glory Hole on east side or road. (S251-A—52) 7. Cummingtonite— et—plagioclasequartz gneiss, from small body near crest of Quartz Hill. ummingtonite is locally intergrown with hornblende. ($53—52) zircon, spinel, sillimanite, and corundum are sparse local minerals. Cordierite occurs as equant grains about 0.3 mm in diameter that poikilitically include tiny subrounded crystals of quartz, magnetite-ilmenite, zircon, and spinel. It is nearly colorless or very pale blue in thin section but locally is deep blue and has a violet cast adjacent to spine] grains. Twinning consists both of interpene- tration and of simple forms. Cordierite is partly . chroic halos. E15 ++++++ Mlcrocllne-quartz-plagioclase-blotite gnelss Cummingtonite-hornblende—plagloclase-quartz rock (mode 12) Cordierite-gedrite-quartz-biotite rock, coarse- grai'ned, locally strongly porphyroblastic; local lenses of green biotite schist at margins Cummingtonite-hornblende-plagioclase-quartz rock, fine-grained (mode 11) Garnet-hornblende-plagioclase-quartz rock, lens 15 ft long (mode 9) Cordierlte-gedrile-quartz-biotite rock (mode 6) Cordlerite-gedrite-plagioclase-quartz-blotite rock (mode 7) Cordierite-gedrite-plagioclase-quartz-biotlte rock No outcrops O 4 8 I_ l l APPROXIMATE SCALE 1'2 FEET FIGURE 4.—Layering in a body of cordierite-amphibole gneiss, Quartz Hill. Mode number refers to table 5. altered to chlorite and sericite (pinite) along grain boundaries. The zircon inclusions have strong pleo- Cordierite embays and corrodes biotite where the two are in contact. On eight grains of cordierite from different rock samples, n, ranges from 1.541 to 1.548i0.003 and averages 1.545i0.003. It is biaxial positive or negative and has a very large but variable 2V. Gedrite forms subhedral or euhedral grains that tend to be slightly larger than the cordierite grains; for the most part it is oriented parallel to the lineation, but some occurs as radiating fibers. It is moderately or weakly pleochroic, and X=pale yellow, Y= pale yellow or pale greenish yellow, and Z=pinkish gray or greenish gray. It is positive and has a large 2V. Three samples gave values for refractive index, n,, as follows: 1.665 130.003, E16 1.674i0.003, and 1.676 i0.003. Chlorite is a local alteration product. Commonly the gedrite is embayed by cordierite and quartz. Garnet is local in occurrence and forms pale pinkish- orange or very pale lavender equant poikilitic grains, mainly containing inclusions of quartz and magnetite- ilmenite but also of sillimanite, cordierite, plagioclase, and biotite. Some crystals are skeletal. Characteris- tically the crystals have closely spaced intersecting fractures. The garnet has been determined from chemical analysis to be an ahnanditic variety; the range of n for five samples is 1.787 21:0.003 to 1.800 -_|-0.003. Biotite forms small ragged subhedral flakes and ranges from grayish yellow to orange brown. Many crystals have strong pleochroic halos around zircon inclusions. The biotite is embayed by both cordierite and gedrite. Plagioclase (calcic oligoclase) is subhedral, is twinned according to the albite and pericline laws (locally having complex twins), and is strongly embayed by cordierite and quartz. A later plagioclase that commonly encloses other minerals in the same sections is more sodic and anhedral and has weak albite twinning. Green spinel occurs locally within cordierite crystals, mainly in association with magnetite-ilmenite. Mag- netite—ilmenite forms anhedral grains in quartz, gedrite, and cordierite and between other mineral grains. Corundum was observed only in one section. Silli- manite occurs as inclusions in plagioclase and garnet and apparently is a relict mineral. The apatite in one specimen (S755—F—53) under crossed nicols has bright- blue polarization colors; as inferred by Vasileva (1958) , it may be a manganese-rich apatite, for the interference color of apatite increases as Mn+2 increases. CORDIERITE-BIOTITE ROCKS Cordierite—biotite rocks are light gray, equigranular, fine to medium grained, and generally mottled because of the segregation of the light and dark minerals into clots. They constitute the bulk of one small body, and representative modes are given in table 5 (sample 8) and table 6 (sample 6). At places in this body a few percent of gedrite occurs in the cordierite-biotite rocks (Sims and Gable, 1964, table 20, sample 4). In another body described in this same report (see table 20, samples 5, 9, 10, 11, and 12), cordierite-biotite rocks are intercalated with thin layers of cordierite-gedrite rocks and cummingtonite— and hornblende-bearing rocks. The cordierite-biotite rocks are granoblastic and, like the cordierite—gedrite rocks, are quite variable in composition. Typically, they consist mainly of quartz, cordierite, and biotite; some varieties contain more than 90 percent biotite and only traces of quartz and SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY cordierite. The biotite has pleochroic colors ranging from light greenish orange brown to moderate reddish brown and is locally skeletal. In one specimen the biotite is olive green and nx=1.597:|:0.005, n,=n,= 1.612 i0.005, and X>Y=Z. Cordierite forms anhedral grains, has parallel twin- ing, and is locally intergrown myrmekitically with quartz. Zircon inclusions in the cordierite have strong pleochroic halos. In one specimen of cordierite n, equals 1.549;}:0.003 ; the mineral is negative and has a very large 2V. Sillimanite occurs locally in cordierite grains and is strongly embayed by it. HORNBLENDE-CUMMINGTONITE ROCKS Hornblende— and cummingtonite-bearing rocks, which occur as distinct layers intercalated with cordier— ite-gedrite gneiss, are of two types—~one consisting of hornblende, calcic plagioclase, and quartz and the other of cummingtonite (with or without intergrown hornblende), calcic plagioclase, and quartz. Both hornblende— and cummingtonite-bearing rocks locally contain biotite. The hornblende—plagioclase-quartz facies is a dark- gray fine— to medium-grained homogeneous rock. It resembles amphibolite but difiers from it in containing more than 10 percent quartz and in having a much more calcic plagioclase. It has a typically granoblas- tic texture. The plagioclase is moderately to strongly zoned and has cores of labradorite or bytownite and rims of andesine or more sodic plagioclase. The hornblende is subhedral or euhedral and has a variable pleochroism. Commonly X=brownish yellow, Y=bluish green, and Z=dark olive green. In one specimen, n, was 1.668:l:0.002. Because the mineral is negative and has a large 2V, it is probably hastingsite. In some sections a little cummingtonite is intergrown with the hornblende. The cummingtonite-bearing rocks are typically greenish gray, equigranular, and fine to medium grained and have a granoblastic texture. The rocks are characterized by intergrowths of cummingtonite and hornblende; the cummingtonite, which usually constitutes about 80 percent of the intergrowths, occurs in structural continuity with the hornblende and either surrounds it or forms irregular patches Within it. In those rocks which contain cumming- PETROLOGY AND STRUCTURE, PRECAMIBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. tonite without intergrowths of hornblende, the cum- mingtonite is peppered with inclusions of ilmenite. The cummingtonite is positive (ZAcz20°); n, was determined in one specimen to be 1.664 i0.003. Garnet, which occurs sparsely, is pale pinkish orange and forms small clear anhedral grains or large irregular porphyroblasts with abundant inclusions of quartz and magnetite. Two separates of sample 853—52 that were X-rayed gave A0=11.557:l:0.0005A and 11.532:{: 0.002A. Quantitative spectrochemical analyses by N. M. Conklin gave 2.4 percent manganese, 0.97 percent calcium, 3.4 percent magnesium, and 23 percent iron. The plagioclase is similar to that in the hornblende rocks described above. Biotite is rare except locally; it has a pleochroism almost identical to that in the associated rocks. In one specimen (853—52), n, was 1.629:l:0.003. CHEMICAL COMPOSITION The rocks belonging to this unit vary in composition from layer to layer, and a chemical analysis is mean- ingful only with respect to a layer which has a specific mineralogy. Chemical analyses that represent the principal mineral associations are given in table 6. Five of the analyses in table 6 represent cordierite- gedrite rocks, one represents cordierite-biotite rock, and one represents cummingtonite-bearing rocks. Empha- sis, therefore, is given to cordierite-bearing mineral associations. The cordierite-gedrite rocks (analyses 1—5, table 6) vary widely in composition but are characterized by low content of CaO, Na20, and K0 and moderate to high content of FeO and MgO. Cordierite-gedrite— quartz rocks, with or without garnet, have FeO+MgO ~20 CaO+Na20+K20~ With increasing amounts of N a20, sodic plagioclase forms (sample 3, table 6), and as K20 increases, biotite forms in substantial amounts (sample 6, table 6) . Garnet forms and both gedrite and cordierite decrease in those rocks which have an uncommonly high ratio of FeO to FeO+MgO (sample 1, table 6). In chemical composition, the rocks are similar to cordierite-antho- phyllite rocks from many other localities. (See, for example, Tilley, 1937 ; Prider, 1940, p. 374.) The cordierite-biotite rocks differ from the cordierite- gedrite rocks in having greater amounts of potassium and somewhat less magnesium and ferrous iron. The potassium favored the formation of biotite and inhib- ited the formation of gedrite as the available magnesium was used in the formation of cordierite and biotite. E17 The cummingtonite—bearing rocks intercalated with the cordierite-gedrite rocks have FeO+MgO <2 Ca0+ N a20+ K20 and a high ratio of CaO to K20+Na20. The high content of CaO, N210, and K20 favored the formation of plagioclase. Some of the available 090 contributed to the formation of hornblende that is intergrown with the cummingtonite. CALC-SILICATE GNEISS AND RELATED BOOKS Rocks containing dominant calcium-iron silicates occur sporadically throughout the region as scattered small pods and discontinuous layers in all major layers of biotite gneiss and microcline gneiss. With few excep- tions the bodies are a maximum of a few feet in thickness and a few hundred feet in length; most are too small to show at the scale of plate 1. The rocks are extremely variable in composition but consist mainly of two types—skarns and calc-silicate rocks. Amphibolite is intercalated with many bodies throughout the quadrangle, and quartzite is associated with some bodies at Pewabic Mountain in the Central City district and on the northeast slope of Mount Pisgah. Skarn is used in this report to designate dark-colored aggregates of calcium, magnesium, and iron silicates that resulted from metamorphism of interlayered impure calcareous and siliceous beds. The name embraces a group of rocks analogous to the reaction skarns of some Fennoscandian geologists (see discus- sion in Leonard and Buddington, 1964, p. 23—24) and does not imply an origin through metasomatism of carbonate rocks by emanations from a nearby cooling granitic body. The light-colored aggregates of similar mineralogy are referred to as calc-silicate rocks al- though some writers would also apply the term “Skarn” to such masses. The skarns have the following characteristic assem- blages: Clinopyroxene-garnet Clinopyroxene—garnet—quartz Clinopyroxene—epidote-hornblende Garnet-magnetite—quartz Lighter colored calc-silicate rocks commonly have the following assemblages: Clinopyroxene-garnet—plagioclasequartz—sphene Calcite-clinopyroxene-garnet—quartz Clinopyroxene-epidote—hornblende-plagioclase— quartz Clinopyroxene-epidote-hornblende—quartz-scapolite Garnet-plagioclase—quartz E18 Epidote-opaque iron oxides-quartz—sphene Epidote—hornblende—quartz—sphene Epidote—plagioclase-quartz—sphene Characteristically the rocks have crystalloblastic tex- ture, indicative of virtually contemporaneous crystal- lization. Poikilitic textures are common in the garnets and less common in clinopyroxene and hornblende. In an earlier report on the Central City district (Sims and Gable, 1964), garnet in the calc-silicate gneisses was reported in molecular percentages. Initial work on these garnets was completed in 1957—58, and at that time it was believed reasonable that percentages for garnets could be determined from combined X—ray and physical data. We now know that this assumption was erro- neous. As no further work has been carried out on the garnets in calc-silicate gneisses since completion of the Central City district report, the type of garnet in these rocks is not designated in this report. BIOTITE GNEISS Biotite gneiss, the most abundant rock unit exposed in the quadrangle, occurs as thick layers intercalated on a gross scale with microcline gneiss. The layers have maximum thicknesses ranging from about 1,000 to 4,000 feet. Smaller bodies, from a feW feet to a few hundred feet thick, occur at places within the major layers of microcline gneiss. Although the lithologies of each of the major biotite gneiss layers differ in detail, each layer is grossly similar, and accordingly the‘biotite gneisses are discussed as a unit herein. The biotite gneiss can be divided on the basis of mineralogy into three principal rock types that are distinctive but to some extent gradational: (1) Biotite- quartz-plagioclase gneiss, (2) sillimanitic biotite gneiss, and (3) cordierite- and garnet-bearing sillimanitic biotite gneiss. All types of biotite gneiss contain moderate amounts of granite gneiss and pegmatite, either as thin streaks and concordant layers to constitute migmatite, or as larger discrete bodies. Sillimanitic biotite gneiss and biotite-quartz-plagio- clase gneiss are the dominant rock types and are intercalated in all the major biotite gneiss layers. Individual interlayers range in thickness from about an inch to several feet. Because of the intimate interlay- ering, mapping of the separate bodies of the two rock types was not practical at the scale of plate 1. Cordierite— and garnet-bearing sillimantic biotite gneiss is mapped separately on plate 1. Contacts of the bodies of garnet— and cordierite-bearing gneiss are not shown, however, because the rock is gradational into sillimanitic biotite gneiss and it is difficult to de- lineate the different rock types in areas of sparse ex— posures. The garnet— and cordierite-bearing biotite gneiss occurs as discontinuous layers within the major SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY biotite gneiss layers that lie above and below the Law- son layer of microcline gneiss. The largest body, about 1,000 feet thick and at least 2% miles long, is strati- graphically near the middle of the lowermost of the two biotite gneiss layers. Discontinuous lenses occur in the upper part of the same layer and locally within the biotite gneiss layer overlying the Lawson layer. The composition of the three principal types of biotite gneiss varies. Because abundant chemical data are lacking, the variations are shown in the text that follows by means of numerous modal analyses. The three types are subdivided for descriptive purposes into nine mineralogic groups. Both the biotite-quartz-plagioclase gneiss and the sillimanitic biotite gneiss have been described in some detail from the Central City district (Sims and Gable, 1964), as well as from adjacent areas (Harrison and Wells, 1956, 1959; Moench, 1964), and are discussed only briefly on the following pages. The garnet- and cordierite-bearing rocks have not been described except in a short preliminary report (Sims and Gable, 1963) and accordingly are discussed more fully herein. A detailed report that will present additional data on the petrology and geochemistry of the cordierite—bearing rocks is now in preparation by the authors. BIOTITE-QUARTZ-PLAGIOCLASE GNEISS GENERAL CHARACTER Biotite-quartz—plagioclase gneiss is a medium- or light-gray fine— to medium—grained equigranular rock. It weathers gray or brownish gray. Rarely, the rock is inequigranular with clots of felsic minerals as much as half an inch in diameter. Most varieties have a conspicuous compositional layering and a marked pre- ferred planar and linear orientation of biotite and other tabular minerals. The mafic and felsic minerals are partly segregated into distinct layers. In many out- crops the biotite-rich layers are paper thin and impart a fissility to the rock, but in others the biotite is evenly dispersed. Some varieties resemble the more massive types of microcline gneiss but can be distinguished from the microcline gneiss because they are darker, finer grained, and richer in biotite. Garnetiferous biotite-quartz-plagioclase gneiss, a local variety gradational into the major type and distinct from the garnetiferous sillimanitic biotite gneiss, forms scattered small bodies, especially in the northwestern part of the quadrangle and in the Central City district (Sims and Gable, 1964). The largest bodies that have been delineated are about 500 feet long and 100 feet wide. The gneiss has a conspicuous layering, the garnet tending to occur within the biotite-rich layers. It re- sembles the more common biotite—quartz-plagioclase gneiss megascopically except that is is darker and con- PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. tains conspicuous garnet porphyroblasts. Commonly, weathering of the exposures forms a reddish stain that aids in distinguishing the rock. PETROGRAPEY The biotite-quartz-plagioclase gneiss contains pla- gioclase, biotite, and quartz as major minerals and a few percent of accessory minerals (table 7). It has an allo- triomorphic granular texture. The plagioclase is dom— inantly sodic andesine but locally is labradorite; it varies more in composition than it does in other types of biotite gneiss. Most grains have narrow rims of albite. Twinning, of which the most prevalent is albite, is com- mon but not ubiquitous, for about 15—20 percent of the grains in the sections examined appear to lack twinning. In sections that lack potassium feldspar the plagioclase grains tend to be poikilitic and to contain tiny inclu- sions of anhedral quartz; where potassium feldspar is present, the plagioclase is slightly antiperthitic and con- tains tiny irregular patches of microcline. Alteration of the plagioclase is slight and consists mainly of clay minerals oriented along twin lamellae and grain bound— E19 and bleb perthite), and tends to be molded around pla- gioclase and quartz grains. Myrmekite occurs locally between the potassium feldspar and plagioclase grains. Quartz forms anhedral grains that show ubiquitous strain shadows; it contains abundant inclusions, at least locally, of magnetite-ilrnenite, zircon, and plagioclase; some of the included magnetite has rims of musco- vite(?). Strongly pleochroic biotite, ranging from pale yellow to deep reddish brown, typically occurs as fresh subhedral crystals having ragged terminations; zircon inclusions have pronounced pleochroic halos. At places the biotite is altered to chlorite and magnetite or to muscovite. Magnetite—ilmenite, apatite, and zircon are nearly ubiquitous accessory minerals. The magnetite locally is altered to hematite. Hornblende, which is a strongly pleochroic olive-green variety, is a local accessory min- eral that occurs mainly adjacent to amphibolite bodies. It forms anhedral grains, at places closely associated with biotite. Not uncommonly, traces of calcite occur with the hornblende. The garnetiferous variety of biotite—quartz-plagi- aries. Potassium feldspar occurs as grains smaller than oclase gneiss differs modally (table 8) from the more those of plagioclase, is typically slightly perthitic (film abundant variety, mainly in containing garnet. The TABLE 7 .—M odes, in volume percent, of biotite-quartz—plagioclase gneiss (Tr, trace; Nd, not determined; _-__, not found. Field number is in parentheses after description of sample] Average of 18 modes, Mineral 1 2 3 4 5 6 7 8 9 10 11 Central City district Potassium feldspar _________________________________________________ 10. 5 4 0 2. 4 Tr 0. 8 1. 0 ________ Plagioclase __________________________ 37. 4 56. 0 51. O 44. 4 43. 6 33. 6 39 3 19. 3 55. 0 63. 2 48. 3 43. 0 Quartz ______________________________ 31. 7 22. 6 32. 2 17. 7 24. 8 30. 8 51 0 66. 0 17. 5 18. 6 22. 0 40. O Biotite ______________________________ 20. 0 20. 4 7. 27. 2 16. 5 21. 3 4 2 10. 4 10. O 4. 0 25. 0 14. 5 Muscovite ___________________________ 1. 6 . 5 ______ Tr ____________ 5 . 7 ____________ . 8 . 5 Magnetite-ilmenite ____________________ 7. 1 . 3 3. 8 2. 6 7. 8 1. O 1. 0 1. 2 2 1 4. 4 2. 3 1. 5 Garnet ______________________________________________________________________________ Tr Tr ______ Chlorite _____________________________ Tr Tr ______ 1. 2 . 3 Tr Tr ________________________ Epidote _____________________________ Tr ____________ Tr ________________________ 1 ____________ Apatite _____________________________ 1. 6 . 2 . 4 . 7 . 3 4 ____________ 3 . 5 6 Zircon ______________________________ . 1 Tr Tr Tr . 1 Tr Tr Tr Tr Tr ...... 5 Calcite __________________________________________ .1(?) ______ 1. 6 ____________________________________ ' Clinozoisite __________________________ . 5 ____________ Tr __________________________________________ Hornblende ______________________________________ 5. 5 6. 1 4. 9 2. 0 ____________ 15. 0 8. 2 ______ Sphene ______ r _________________________________________ . 1 ______ . 4 ____________ Tr . 3 ______ Allanite _____________________________________________________ 1 Tr __________________ Tr ______ Total ___________________________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. O 100. 0 Composition of plagioclase _____________ An54 Nd An“ Ana.) An“ Nd Nd Nd Anal An“ Nd ________ Average grain diameter __________ mm_- 0. 3 0. 5 0. 5 0. 3 0. 2 0. 6 O. 4 0. 4 0. 4 0. 3 0. 5 ________ 1. Biotite gneiss, stron 1y foliated, homogeneous, from southwest flank of Mount Pisgah, between amlin Gulch and Woodpecker Gulch. (CC—155) . Biotite gneiss, thick layer, in migmatlzed biotite-sillimanite gneiss, from north side Fall River, 1.1 miles from west edge of quadrangle. (CC—162A) . Biotite gut-ass, taken from 10-it-thick layer adjacent to amphibolite in Lawson layer of mlcrocline gneiss, north side of Fall River, 0.6 mile from west edge of quadrangle. (CC—21—1A) . Gneiss, from outcrop opposite the Fall River power-plant. (0 0-103) . Biotite gneiss fine-grained, interlayered with microcline gneiss, taken from near junction of éilver Creek and North Clear Creek. (CC-399-B) . Biotite gneiss, strangliy foliated, from contact zone against microcline gneiss layer, head of Freeman ulch. (CC—315—A) WM GOIA 7. Biotite-quartz—plagioclase gneiss, interlayered with microcline. gneiss and silli- manitic biotite-quartz gneiss, from east slope of ridge, near head of Chase Gulch. (CC-443-A) 8. Biotite gneiss interlayered with granite gneiss and pegmatite, from hill southeast of Freeman Gulch, south of North Clear Creek. (CC—3674B) 9. Biotite gneiss, from infolded lens about 100 ft from contact of microcline gneiss body in a saddle 1 mile east of Oregon Hill, on north side of Stewart Gulch. A small body of amphibolite is exposed about 200 It to north. (C 0—994—2) 10. Biotite gneiss, from layer within body of amphibolite on east slope ol ridge between Stewart Gulch and Pickle Gulch. rades transitionally into amphi- bolite. (C 0-1009) ll. Biotite gneiss, from north slope Bald Mountain. (EWT-66B—54) E20 garnet forms porphyroblasts as much as 7 mm in diam- eter but more commonly occurs as grains about 2 mm across. Many grains are almond shaped, flattened in the plane of foliation, and elongated in the direction of dominant lineation. Many grains are broken or frac- tured roughly at right angles to the lineation. Biotite commonly forms thin sheaths around garnet porphy— roblasts. The garnet is strongly poikilitic and contains abundant quartz and plagioclase and less biotite, mag- netite-ilmenite, and apatite. TABLE 8.——Modes, in volume percent, of garnetiferous biotite- quartz—plagioclase gneiss [Tr, trace; Nd, not determined; ______ , not found. Field number is in parentheses after description of sample] Average of 10 modes, Mineral 1 2 3 4 5 6 Central City district 2. 3 ____________ 0. 1 ________________________ 49. 2 36. 4 39. 5 25. 9 43. 4 48. 0 29. 0 35. 3 45. 4 34. 2 50. 4 15. 6 41. 0 38. 0 9. 2 16.6 12. 4 16. 3 34. 2 5. 7 20.0 3. 7 . 2 . 2 ...... . 3 __________________ Tr .8 3.3 4.2 3.6 2.9 1.0 .2 .4 10.3 2.5 2.9 1.0 10.0 . 1 __________________ Tr ______ Tr . 2 . 1 . 5 Tr . 4 2 0 Zircon _________________________ Tr Tr Tr . 1 Tr Tr Amphibole (cummingtonite)..- ______________________________ 1. 0 Total ______________________ 100. 0 100. 0 100.0 100 0 100. 0 100 0 100.0 Composition of plagioclase _____ Nd Average grain diameter. .mm" 0.5 1. Biotite gneiss, thin layer, associated with granite gneiss and pegmatite in Lawson layer of microcline gneiss; sample taken just north of Fall River, about 0.5 mile from west edge of quadrangle. (CC—24) 2. Migmatitic biotite gneiss intercalated with microcline gneiss; sample taken just below base of microcline gneiss layer, south slope of Blackhawk Peak near North Clear Creek. (CC—363—1B) 3. Migmatitic garnetiferous biotite gneiss; sample taken about 800 ft west of locality of sample 2. (CC—366—2) 4. Garnetiferous biotite gneiss, from ridge on north side of Miners Gulch, 1 ,100 ft from west edge of quadrangle. Gneiss is interlayered with biotite—quartz-piagioclase gneiss and sillimanitic biotite-quartz gneiss. (CO—677—10) 5. Garnetiferous biotite gneiss, from thin layer of biotite gneiss in granodiorite body roadcut near head of Missouri Gulch. Garnetiferous biotite gneiss is interce- 1:013:51 sgitgsiillimanitic biotite-quartz gneiss, which ilocally is gatnetiferous. 6. Migmatitic garnetiferous biotite gneiss, adjacent to blunt layer of microcline gneiss; sample taken from steep westward-facing slope of Michigan Hill, 1,000 ft north- east of bend in North Clear Creek. (CC-889). The rocks have a typical crystalloblastic texture; the plagioclase, quartz, and biotite are intergrown and show a well-defined mosaic texture. Hornblende and garnet, where present, appear to have crystallized almost simultaneously with the major minerals. Potassium feldspar apparently crystallized later. The albitic rims on plagioclase, and locally on myrmekite, indicate some modification subsequent to crystallization of most of the rock. The nearly complete absence of potassium feldspar, except in strongly migmatized biotite gneiss, and the apparently late paragenetic pos1tion are in— terpreted to indicate that the potassium feldspar was largely introduced during migmatization. The assemblages of the biotite—quartz—plagioclase gneiss are: Biotite-plagioclase—quartz Biotite-plagioclase—potassium feldspar-quartz SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Biotite-magnetite-plagioclase-potassium feldspar- quartz Biotite-hornblendepmagnetite—plagioclase-quartz Muscovite is a local stable mineral in the preceding rock type as well as in the following assemblages of garnetiferous biotite—quartz-plagioclase gneiss: Biotite-garnet—plagioclase-quartz Biotite-garnet-magnetite—plagioclase—quartz SILLD/IANITIC BIOTITE GNEISS GENERAL CHARACTER Sillimanitic biotite gneiss consists of two varieties, sillimanitic biotite—quartz gneiss and sillimanitic biotite- quartz-plagioclase gneiss, which are gradational and inseparable in the field but differ somewhat in quantita- tive mineralogy. The rocks are light gray or bluish gray fine to medium grained and generally equigranular. Inequigranular porphyroblastic varieties occur locally. Weathered exposures are brownish gray and at places have a conspicuous silvery sheen. In most exposures the gneiss has a distinct layering consisting of alternating layers a fraction of an inch to a few inches thick of slightly different lithology or texture. Sillimanite is distinctive and conspicuous and occurs as alined needles and as aggregates of fibers as much as an inch in length, or less commonly as tabular discoid masses of about the same length. Parnoenarnr The gneiss consists mainly of quartz, biotite, sillima- nite, and plagioclase; minor minerals include potassium feldspar, magnetite—ilmenite, and muscovite. Sillima- nitic biotite-quartz—plagioclase gneiss (table 9) is distinguished from sillimanitic biotite-quartz gneiss (table 10) in that it contains more than 15 percent plagioclase; it also has less biotite and more potassium feldspar. However, as can be seen by reference to the two tables of modes, each variety’s mineral content varies considerably. Both varieties of gneiss have an allotriomorphic granular and, locally, a lepidoblastic texture. Grains average about 0.4 mm in diameter; porphyroblasts are commonly 2.5—3.0 mm in diameter. Plagioclase of relatively uniform composition (Ann—29) forms anhedral or rarely subhedral grains and has three modes of occur- rence: (1) Grains intergrown with the dominant min— erals, (2) grains interstitial between larger crystals of quartz and biotite, and (3) poikilitic grains, generally untwinned, that poikilitically contain inclusions of bio- tite, magnetite, quartz, sillimanite, and zircon. In many sections nearly every plagioclase grain is twinned, whereas in other sections only about half the grains are twinned; albite twinning is dominant; Carlsbad and pericline twins are less common. Adjacent to potas- sium feldspar, the plagioclase has narrow rims of albite PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. E21 TABLE 9.—-Modes, in volume percent, of sillimantttc biotite-quartz—plagioclase gneiss [Tr, trace; Nd, not determined; ._-., not found. Field number is in parentheses after description of sample] Mineral 1 2 3 4 5 6 7 S 9 Potassium feldspar _______________________ 7. 8 13. 3 5. 6 7. 4 0. 6 8. 1 7. 6 ________ 9. 5 Plagioclase ______________________________ 16. 2 19. 6 30. 1 15. 4 28. 9 16. 5 21. 5 38. 1 19. 8 Quartz _________________________________ 49. 1 54. 2 47. 5 53. 2 57. 6 52. 1 47. 2 23. 4 54. 6 Biotite _________________________________ 16 5 9. 0 13. 5 13. 9 8. 7 14. 3 16. 0 34. 1 10. 8 Muscovite ______________________________ 1 3 . 4 . 3 . 9 1. 5 3. 3 . 4 1. 8 1. 0 Magnetite—ilmenite _______________________ 1. 3 1. 1 2. 7 1. 5 1. 7 2. 2 . 8 . 2 1. 9 Sillimanite ______________________________ 7. 8 2. 2 Tr 7. 7 l. 0 3. 5 6. 5 2. 2 2. 4 Chlorite ________________________________ Tr Tr _______ Tr ________________ Tr Tr ________ Apatite _________________________________ Tr _______________________________________________ . 2 ________ Zircon __________________________________ Tr . 2 Tr Tr Tr Tr Tr Tr Tr Calcite _________________________________________________ Tr ________________________________________ Total _______________________________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 Composition of plagioclase ________________ Nd Nd A1129 Ann Nd Ann A1125 Nd An,3 Average grain diameter _____________ mm__ 0. 4 0. 4 0. 4 Nd 0. 4 0. 3 0. 3 0. 4 0. 4 1. Sillimanltic biotite-quartz—plagioclase gneiss interlayered with granite gneiss and pegmatite; sample taken [torn outcrop north side of Fall River, 1.1 miles from west edge of quadrangle. (CC—226) _ 2. Sillimanitic biotite-quartz-plagioclase gneiss lnterlayered wrth biotite-quartz-plag— ioclase gneiss; outcrops on steep west slope of Michigan Hill opposite mouth of Freeman Gulch. (CC-243) 3. Gneiss from mine adlt in eastward-trending gulch located between North Clear Creek and Freeman Gulch; rock is interlayered with rock represented by sample 4. (CC—611—A) . 4. Sillimanitic biotite-quartz-plagloclase gneiss, from same locality as that of sample 3. (CC—611—B) TABLE 10.—Modes, in volume percent, of sillimam’tic biotite- quartz gneiss [Tr, trace; Nd, not determined; ._.., not found. Field number is in parentheses after description of sample] Average of 18 modes, Mineral 1 2 3 4 5 6 Central City District Potassium feldspar ___________________ 0.9 8. 0 ...... 0.6 ______ 0. 5 Plagioclase _________ 10. 0 5 7. 8 10. 2 9. 2 5. 3 14. 0 . 3. 0 34. 0 .2 38. O 49. 0 . 4 .3 .8 31. 5 3. 5 r . 6 . 5 . 6 0 . 1 . . 4 . 1 . 0 . 4 .3 . 4 . 0 Composition of plagioclase _____ Arm Averagegraindiameter-_mm_- 0.4 Nd Nd Nd 0.3 0.2 1. Sillimanitic biotite-quartz gneiss interlayered with biotite-quartz-plagio— clase gneiss, sample taken from outcrop on north side of Fall River, about 1.1 miles from west edge of quadrangle. (CC—162—B) 2. Biotite- and Sillimanite-rich variety of gneiss; thin layer in granodiorite body at head of Missouri Gulch; interlayered with garnetiferous biotite- quartz-plagioclase gneiss. (CC—855—A1) 3. Typical sillimanitic biotite-quartz gneiss from Missouri Gulch area. (J G— 2 4, 5, 6. Samples from north slope of Bald Mountain and vicinity. (EWT—73-54. EWT—87-54, EWT—96—54) and locally is myrmekitically intergrown with quartz. Some grains show a few percent of potassium feldspar in oriented patches, to constitute antiperthite. The potassium feldspar is dominantly microchne; it is slightly perthitic (string and film perthite), generally poorly twinned, and dominantly interstitial. Quartz forms anhedral grains and has strong strain shadows. The biotite is strongly pleochroic, ranging from light yellowish brown to dark reddish brown, and occurs as 247—419—67—4 5. Gneiss from conspicuous outcrop near road on south slope of California :(lgtbmibaglnlillfl north of Nugget; interlayered vn'th rock like that of sample 6. 6. Gneiss from same locality as that of sample 5. (CC—1051—B) 7. Gnelss from roadcut, west of entrance to Cold Spring campground, just east of quadrangle. (J 0—84) 8,9. Samples from north slope of Bald Mountain and vicinity. (EWT—53—54, EWT—89—54) small stubby, ragged laths; at places it contains con- siderable Sillimanite. Most sections show some alter- ation of the biotite to chlorite and magnetite. Optical and chemical data on a sample of biotite from the Cen- tral City district are reported by Sims and Gable (1964). Sillimanite occurs in sheaths and stringers and is gen- erally associated with quartz and biotite. Muscovite forms overgrowths on and is intergrown with biotite, and it rims magnetite. Magnetite-ilmenite and zircon are common accessory minerals. Many sections show some alteration of magnetite to hematite. Biotite, quartz, potassium feldspar, muscovite, plagioclase, and Sillimanite are intergrown and have sharp contacts with one another. Any one of them can be found in direct contact with any one of the others. Muscovite is dom— inantly a primary mineral. Sillimanite and potassium feldspar occur together and only locally have inter- vening muscovite. Chlorite and the clay mineral alteration of plagioclase are secondary in origin. CORbIZERITE- AND GARNET-BEARING SMANITIC BIOTITE GNEISS GENERAL CHARACTER The rocks mapped as cordierite- and garnet—bearing sillimanitic biotite gneiss are distinguished from other types of biotite gneiss in that they contain cordierite and (or) garnet as well as Sillimanite. The cordierite- and garnet-bearing biotite gneiss varies in composition but consists mainly of three va- rieties: (1) Garnetiferous sillimanitic biotite-quartz- plagioclase gneiss (table 11), (2) cordierite—bearing E22 garnet-sillimanite—biotite—quartz—plagioclase gneiss (table 12), and (3) cordierite—biotite—sillimanite—quartz- plagioclase gneiss (table 13). The varieties are gra- dational and were not distinguished separately in the SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY field. Varieties l and 2 are difficult to distinguish meg- ascopically except Where cordierite occurs as distinct porphyroblasts; variety 3 is sparse and is generally intimately intermixed with variety 2. TABLE 11.—M odes, in volume percent, of garnetiferous sillimanitic biotite-quartz-plagioclase gne’iss [Tr, trace; Nd, not determined; ._.-, not found. Field number is in parentheses after description of sample] Mineral l 2 3 4 5 6 7 8 9 Potassium feldspar _______________________ 26. 6 l. 4 9. 5 0. l(?) 1. 2 Tr(?) 18. 5 3. 0 ________ Plagioclase ______________________________ 18. 5 23. 8 40. 8 2. 5 60. 4 41. 8 31. 5 22. 9 11. 8 Quartz ................................. 27. 6 53. 9 14. 5 59. 5 7. 2 42. 0 35. 8 64. 1 48. 4 Biotite _________________________________ 16. 3 14. 3 26. 0 24. 0 28. 5 13. 6 12. 2 7. 7 8. 5 Muscovite ______________________________________ . 1 ' . 2 ________ . 4 . 5 . 6 ________ 1 0 Magnetite—ilmenite _______________________ . 9 2. 6 8 1. 0 . 3 2. 0 1. 2 . 9 3 9 Garnet _________________________________ 3. 3 Tr . 5 1. 9 . 4 . 1 Tr . 6 25 9 Sillimanite ______________________________ 6. 5 3. 9 7. 7 11. 0 1. 6 Tr 1 . 8 4 Apatite ________________________________________________________ Tr Tr Tr ........................ Zircon __________________________________ 3 Tr Tr Tr Tr Tr l Tr 1 Total _______________________________ 100. 00 100. 00 100. 00 100. 00 100. 00 100. 00 100. 00 100. 00 100. 00 Composition of plagioclase ________________ Ana, Nd Ann An“ An” An24 Nd Nd An” Average grain diameter; _____________ mm__ 0. 2 0. 3 0. 2 0. 4 0. 5 0. 4 0. 2 0. 4 0. 6 1. Gneiss layer in microcline Engisgjglill north of Pecks Flat, about 0.3 mile south or . Gneiss, about 1,000 ft southeast of locality of sample 3. (0 0-574) North Clear Creek. (C 2. Gneiss, from upper part of biotite gneiss layer, about 500 it east of contact with microcline gneiss; ridge between Peeks Gulch and Chase Gulch. (CC-383—1) 3. Gneiss from dump of caved adit on east side of Miners Gulch, just west of Elk Creel: gabbro pluton. (00-681) 4. Quartz-rich phase of gneiss about 1,0001t southwest of locality of sample 2, near head of Chase Gulch. (é C—367—B) Om \1 03m . Slightly migmatized gueiss, from ridge crest 0.5 mile south—southwest of Sheridan Hill. (CC—460 . Gueiss interlayered with biotitequartz-plagioclase eiss, from ridge overlooking North Clear Creek, east of mouth of Freeman Gu ch. (0 C—425-1) . Gneiss, from south flank of Michigan Hill. (CC-911—2A) . Layered possibly sheared gneiss, from outcrop on east slope of Oregon Hill, 0.5 mile east of crest. (CC—1077—1B) TABLE 12.—M odes, in volume percent, of cordierite—bearing garnet-sillimanite—bz'otite gneiss [Tr, trace; Nd, not determined; ________ , not found. Field number is in parentheses after description of sample] Mineral 1 2 3 4 5 6 7 8 9 10 Potassium feldspar ______________ 3. 9 14. 2 . 3 Tr Tr(‘?) Tr ........ 29. 7 3. 3 7 7 Plagioclase _____________________ 2. 2 2. 2 6. 0 12. 1 11. 9 5. 7 . 2 10. 6 8. 0 3 9 Quartz _________________________ 30. 4 24. 5 27. 0 27. 3 40. 9 52. 7 22. 2 1. 2 29. 6 26. 7 Cordierite ______________________ 26. 7 19. 2 27. 9 5. 9 11. 2 l. 9 39. 9 10. 0 4. 3 11. 7 Sillimanite ______________________ 8. 2 7. 3 6. 1 . l 7. 8 10. 6 17. 5 20. 1 5. 7 3. 2 Garnet _________________________ 3. 9 3. 6 3. 3 33. 1 Tr 7. 1 9. 7 1. 2 20. 0 17. 4 Biotite _________________________ 22. 0 25. 0 24. 6 l5. 2 25. 0 21. 0 7. 0 20. 8 28. 9 27. 7 Muscovite (secondary) ___________________ Tr _______________________ . 2 ________ 5. 6 Tr ________ Magnetite—ilmenite ______________ 2. 7 3. 9 4. 6 6. 3 3. 2 . 8 3. 3 . 5 . 1 1. 4 Zircon __________________________ Tr Tr . 1 Tr Tr Tr . 2 . 3 Tr Tr Apatite ________________________________ 1 . 1 Tr Tr ________________________________ Tr Spinei __________________________________________ Tr Tr ________ Tr _________________________________ Chlorite _______________________________________________________________________________________________ . 1 Andalusite ______________________ Tr Tr Tr _______ Tr ________________________________ . 2 Total ______________________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 Composition of plagioclase ________ Nd Nd Anne-32 Anss An” A1133 Nd Nd A1130 Nd Average grain diameter _____ mm-_ 0. 7 0. 3 0. 4 0. 3 0. 4 0. 3 Nd 0. 4 Nd Nd 1. Migmatitic gneiss, outcrop 0.25 mile up Missouri Gulch road from State High- way 119, east side of Missouri Creek. (J G—12a) 2. Gneiss, from ridge, 0.5 mile northwest of Missouri Lake. (J G-30) 3. Gneiss, from outcrop in gulch, 0.5 mile West of Missouri Lake and 0.2 mile south- west of locality of sample 2. (J G—38a) 4. Biotite gneiss, adjacent to small lens of mierocline gneiss, east slope of Yankee Hill (CC-530—1 5. Layerwithin interlayered migmatitic biotite gneisses, from east slope of hill at altitude of about 10,800 ft, 0.7 mile north of Yankee Hill. (CC—677—1B) 6. Layer within folded migmatitic biotite gneisses, from east of body of quartz diorite and hornblendite that crosses Pecks Gulch south of North Clear Creek, central part of quadrangle. (00483-4) 7. Biotite gneiss, from mine dump in eastward-draining gulch midway between Oregon Hill and Michi an Hill, 0.3 mile west of Silver Creek. (CC—789—A) 8. Gneiss, from southwest ank of Michigan Hill, 900 feet east of junction of North Clear Creek and Pine Creek. (00—887) 9. Gneiss, from roadcut on north slope of hill, 1.5 miles west of the Bald Mountain Cemetery. (EWT—QO) 10. Gneiss, from east side of small hill south of Chase Gulch and 1.5 miles north of Bald Mountain Cemetery. (D—l68) PETROLOGY AND STRUCTURE, PRECANFBRIAN TABLE 13.—Modes, in volume percent, of cordierite biotite-quartz- plagioclase gneiss Tr, trace; Nd, not determined; __.., not found. Field number is in parentheses after description of sample] Mineral 1 2 3 Potassium feldspar ______________ Tr _ _ _ _ _ _ _ _ ________ Plagioclase _____________________ 18. 3 0 1 15. 3 Quartz _________________________ 24. 3 Tr 37. 2 Cordierite ______________________ 24. 1 89. 0 31. 7 Biotite _________________________ 28. 0 4. 3 8. 3 Magnetite—ilmenite _______________ 2. 7 3. 6 6 0 Rutile __________________________ Tr ________________ Apatite ________________________ . 6 ________ Tr Zircon _________________________________ Tr Tr Muscovite ______________________ . 1 1. 3 ________ Sillimanite ______________________ 1. 9 1. 7 l. 5 Spine] __________________________________ Tr Tr Garnet _________________________________ Tr Tr Staurolite ______________________ Tr ________________ Andalusite ______________________ 1. 9 ________________ Total ______________________ 100. 0 100. 0 100. 0 Composition of plagioclase ________ Nd Nd An26 Average grain diameter _____ mm__ 0. 5 0. 9 0. 3 1. North slope Bald Mountain, about 2,200 it west-southwest of sharp curve in road , upper Eureka Gulch. (EWT—fiGb—54) 2. Gneiss from top of ridge, 0.45 mile west-southwest of Missouri Lake. Sample taken 1,000 ft southeast of location of sample JG-3Sa, table 12. (J G—69) 3. Cordien'te-biotite gneiss adjacent to hornblendite body, just north and east of junction of Mount Pisgah road with Forest Service road to experimental tree plot and dam. (CC—360—1) The rocks are medium or dark gray, fine to medium grained and extremely variable in structure. They range from nearly massive, homogeneous-appearing rocks to strongly layered gneisses. The more massive rocks tend to be nearly equigranular, but locally they have conspicuous porphyroblasts of either intergrown garnet and sillimanite or cordierite. The strongly layered rocks tend to be coarser grained and inequi- granular. Not uncommonly, the layered rocks have a pronounced crinkly foliation. All varieties contain some pegmatite that has potassium feldspar and that occurs as streaks or conformable pods. Am- phibolite and calc-silicate gneisses are commonly found adjacent to the cordierite-bearing gneisses. PETROGRAPHY The principal minerals—cordierite, quartz, biotite, plagioclase, garnet, sillimanite, and potassium feld— spar—occur in different amounts and proportions, as can be seen by comparing the modes in tables 11, 12, and 13. Garnet occurs as porphyroblasts, from less than 1 mm to about 5 cm in diameter, that are strongly poikilitic and contain inclusions of biotite, sillimanite, plagioclase, quartz and, rarely, potassium feldspar. The porphyroblasts have two modes of occurrence: (1) Megacrysts that truncate biotite laths and (2) megacrysts that are enclosed by conformable biotite- sillimanite sheaths. Although generally subequant, ROCKS, CENTRAL CITY QUADRANGLE, COLO. E23 some garnet megacrysts are elongate in the plane of foliation. In some cordierite-bearing varieties, the garnet is strongly embayed and corroded by cor- dierite and is surrounded by well-defined coronas of it; outside the coronas, biotite and sillimanite are profusely intergrown with quartz and plagioclase. A reddish-brown biotite has the pleochroic formula in which X=straw yellow, Y= deep orange brown, and Z=dark reddish brown. It forms stubby, ragged laths, generally having a strong preferred orientation. The crystals are embayed by quartz, plagioclase, cordierite, and garnet and have grain boundaries that are fuzzy against cordierite. Myrmekitic intergrowths of quartz in biotite are as- sociated With alteration of the biotite to andalusite. Zicron inclusions have strong pleochroic halos. Sillimanite is generally associated with biotite, mainly as capillarylike needles and short stubby crystals in felted masses. Except rarely, sillimanite in contact with potassium feldspar lacks intervening muscovite. Quartz, the dominant mineral in the rocks, forms anhedral grains that are slightly larger than all other minerals except garnet. The quartz contains a few tiny inclusions of sillimanite, plagioclase, biotite, and magnetite-ilmenite. Plagioclase (oligoclase-andesine) is anhedral, and is embayed by quartz, potassium feldspar, and biotite. In a few sections it is slightly antiperthitic. The plagioclase is twinned according to the albite and pericline and, rarely, Carlsbad twin laws. Adjacent to potassium feldspar it commonly has albitic rims as much as 0.02 mm thick and locally has myrmekite rims. Typically, it is altered slightly to clay minerals. Cordierite occurs as nearly colorless or as pale-blue equant grains of about the same diameter as those of plagioclase, as porphyroblasts as much as 5 mm in diameter, and as coronas around garnet, as previously noted. Nearly all grains are poikilitic and have in- clusions of magnetite, biotite, quartz, garnet, and sillimanite. In some sections the cordierite has myr- mekitic intergrowths of quartz and, rarely, potassium feldspar. The cordierite can be distinguished by its diagnostic alteration to pinite, twinning of both simple and interpenetration type, low birefringence, sponge- like appearance, and by its content of abundant minute opaque(?) minerals and zircon which has strong radiohalos. Where cordierite coronas are well formed around garnet, they generally lack inclusions of other minerals except quartz and rarely potassium feldspar. These coronas separate the garnet from parts of the rock that are rich in biotite and sillimanite. The E24 cordierite and associated quartz embay the garnet. The cordierite is biaxial negative (rarely biaxial posi- tive) and has a 2V greater than 65°. Seven determina- tions of n, ranged from 1.548:I:0.003 to 1.554i0.003 and averaged 1.550:|:0.003. The potassium feldspar consists of grid-twinned mi- crocline, orthoclase which has fine hairlike albitic twin lamellae, and vein and film microperthite. It forms small anhedral grains and, rarely, large poikilitic por- phyroblasts which have inclusions of sillimanite (un- altered), clear quartz, plagioclase, and biotite. Both monoclinic and triclinic feldspars, as determined by X-ray, coexist in the cordierite—bearing gneiss units. Monoclinic feldspar is dominant in the lens that is con- tinuous with the Dumont anticline in the northeastern part of the quadrangle, whereas triclinic feldspar is dominant in the lens that is due south of Michigan Hill. A small percentage of monoclinic orthoclase micro- perthite occurs with the triclinic feldspar. The triclinic feldspars have a distinct microcline grid twinning and contain about 75i1 percent KAISiaos by weight. P0- tassium feldspars in the other small, scattered, cordi- erite-bearing biotite gneiss bodies within the quad— rangle, as observed in thin section, resemble those at Michigan Hill. The minor minerals include magnetite-ilmenite, which is ubiquitous, and lesser amounts of andalusite, stauro- lite, spinel, apatite, and zircon. Magnetite-ilmenite occurs as poikilitic anhedral grains that have inclusions of quartz, biotite, sillimanite, and spinel or occurs as smaller anhedral grains closely associated with cordi- erite. Narrow coronas of slightly pleochroic, pale-green to pale-brown newly formed biotite were found in a few samples around magnetite in contact with cordierite grains. In those sections containing spinel, the spinel commonly is closely associated with the mag- netite, and it may be surrounded by coronas of biotite similar to those around the magnetite. The spinel is an apple-green magnesian hercynite, as determined by chemical analysis. Andalusite occurs sparsely in many sections that contain cordierite. It forms irregular. parallel intergrowths with biotite in a few sections and embays and corrodes the biotite. It is negative and can be distinguished by its high relief, cleavage, low bire- fringence, very large 2V, and weak dispersion. Most andalusite grains have vermicular rims adjacent to al- tered biotite. The biotite adjacent to the andalusite also is vermicular and commonly is altered to a greenish brown. Rounded relict grains of staurolite occur rarely as inclusions in both plagioclase and cordierite. Alteration minerals that occur within or corrode and embay the major minerals include chlorite, mus- covite, and clay minerals. The pinitic alteration of SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY cordierite appears to consist of an intimate mixture of sericite and chlorite. The textures of the cordierite-garnet gneisses are more complex than those in other biotite gneisses in the area. The presence of porphyroblasts and coronas and the rather common embayment and corrosion of minerals indicate some modification of the rock sub- sequent to original crystallization. Cordierite, garnet, and, less commonly, potassium feldspar and magnetite-ilmenite occur as porphyro- blasts in the rocks. The porphyroblasts nucleated and grew after some recrystallization had taken place. Garnet grew both by pushing aside surrounding min- eral grains and by replacement. The common elon- gation in the plane of foliation indicates that the garnet grew during deformation; some of the garnet apparently continued to grow after the peak of maximum defor- mation, as garnet megacrysts truncate biotite laths. Cordierite evidently formed later than the garnet, for cordierite porphyroblasts commonly contain garnet, whereas cordierite has not been observed in garnet porphyroblasts. Cordierite occurs both as nearly equant grains and as coronas around garnet. Probably the cordierite in the coronas formed at the expense of garnet by the following reaction: Garnet+biotite +sillimanite +quartz +02—> cordierite +magnetite +potassium feldspar +H20. This reaction has been suggested previously for similar rocks by Schreyer and Yoder (1961, p. 150). The virtual absence of biotite and sillimanite in the cor- dierite coronas supports the concept that this reaction is the actual one. The rare occurrence of both spinel and quartz in the same thin section is indicative of local dis- equilibrium. The spinel evidently formed by the breakdown of magnetite and did not react with quartz because it was separated from it by intervening min- erals. This association of spinel and quartz in the same rock is not uncommon and has been described, for example, from a gneiss north of Great Slave Lake, Northwest Territories, Canada (Folinsbee, 1940, 1941). Andalusite, which so far as known is restricted to cordierite-garnet gneisses, seems to have formed at the expense of biotite, probably by the following reaction: Biotite + Ogaandalusite + magnetite-ilmenite + quartz + biotite (new). The quartz released by the reaction formed vermicular intergrowths with both the biotite and the andalusite; magnetite-ihnenite formed discrete crystals. CHEMICAL COMPOSITION OF THE BIOTITE emssns Because of the wide variation in mineralogy of the biotite gneisses, we have calculated approximate PETROLOGY AND STRUCTURE, PRECAMZBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. chemical compositions of the major rock types from the modes (table 14). The technique was described by Sims and Gable (1964) in the report on the Central City district. Although the samples are insuflicient in number to give statistically meaningful bulk chem- ical compositions, they are adequate to give a reliable order of magnitude. Chemical analyses of five repre- sentative samples of biotite gneisses supplement the calculated analyses and are given in table 15. The biotite gneisses are of two distinct chemical types, indicated mineralogically by the presence or absence of sillimanite. The biotite-quartz-plagioclase gneiss and the related garnetiferous variety are char- acterized by having an excess of Na20 over K20 and A1203 1 Fean+FeO+MgO+CaO+Na20+K20< ’ whereas the sillimanitic biotite gneisses have an excess of K20 over Na20 and A120” >1 or z 1 F6203+F90+MgO+CaO+N320+K20 . E25 Also, the sillimanitic biotite gneisses have a low CaO content. The close dependence of rock type on the NazoszO ratio is shown graphically in figure 5, a plot of 23 analyzed samples. The N a20:K20 ratios of biotite-quartz—plagioclase gneiss range from about 1:1 to slightly more than 2:1; the ratios of sillimanitic gneisses range upward to about 1:3. Garnet forms in the biotite-quartz-plagioclase gneiss when FeO exceeds the requirements for biotite. The available data indi- cate that it forms in those rocks having FeO+MgO( +MnO) CaO+Na20+K20 >1 and a high ratio of Fe0+MgO to A1203. In general, the sillimanitic varieties of biotite gneiss are characterized chemically by having FeO+MgO >1 CaO+Na¢O+K20 TABLE 14.—Estimated chemical compositions and average modes of principal types of biotite gneiss [Tr, trace; -.--, not found. Estimated chemical composition given in weight percent and average modes, in volume percent] 1 2 3 4 5 6 7 8 9 Estimated chemical composition 61. 9 70. 5 65. 1 63. 5 57. 9 68. 2 72. 7 67. 3 56. 7 15. 4 13. 0 14. 8 13. 3 23. 0 15. 3 12. 8 14. 9 19. 6 4.2 2.3 3.5 1.9 2.4 2.1 2.2 2.2 3.8 5.0 3. 5 6.2 9.4 7.0 5.4 3.9 5.8 11.1 2.0 1.6 1.4 2.1 2.8 2.1 1.4 1.8 3.6 4.4 2.5 3.3 1.6 .3 .7 1.1 1.5 .5 29 3.6 2.7 2.3 .7 1.2 2.0 2.4 .8 K90 ____________________________________ 1 9 1. 6 1. 7 1. 9 3. 5 2. 8 2. 8 2. 6 2. 8 H20 (total) _____________________________ 6 . 5 . 6 . 7 1. 2 1. 0 . 5 . 6 . 9 TiOz ___________________________________ .6 .5 4 .5 1.1 .8 .5 .6 .7 Total _______________________________ 98. 9 99. 6 99. 7 97. 2 99. 9 99. 6 99. 9 99. 7 100. 5 Average modes Potassium feldspar _______________________ 1. 7 0. 1 0. 4 ........ 1 6 0. 5 6. 7 6. 7 5. 9 Plagioclase ______________________________ 44. 7 43. 1 40. 4 29. 0 7 2 13. 9 22. 9 28. 2 6. 3 Quartz _________________________________ 30. 4 40. 1 37 0 38. O 35 6 48. 8 48. 8 39. 2 28. 3 Cordierite____________; _________________________________________________________________________________ 15. 9 Biotite _________________________________ 15. 1 14. 4 15 8 20. 0 32 5 23. 5 15. 2 16. 8 21. 7 Muscovite _________________________ . 4 . 4 8 ________ 9 3. 7 1. 2 . 3 . 6 Amphibole _______________ 3. 8 _______________________________________________________________ Opaque iron oxides- 3. 0 1. 5 2 5 1. 0 1 5 1 3 1 5 1. 5 2. 7 Tr ________ 2 9 10. O ________________________ 3. 6 9. 9 Sillimanite _____________________________________________________________ 20. 6 8. 2 3. 7 3. 6 8. 7 Other minerals __________________________ 9 . 4 2 2. 0 . 1 . 1 Tr . 1 Tr Total _______________________________ 100. O 100. 0 100. O 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 Composition of plagioclase (average) _______ An“; Anzs An” A1123 AD27 A1125 An“ A1123 Ans; 1. Biotite—quartz—plagioclase gneiss. Analysis of biotite in 8622—53 used for compu- tatilon; composition of hornblende estimated. Average of 11 modes listed in tab e 7. . Biotite—quartz- lagioclase gneiss, from Central City district (Sims and Gable, 1964. table 11 . Average of 18 modes. 3. Garnetiferous biotite-quartz-piagioclase gneiss. Analysis of garnet in EWT-90 used for computation. Average of 6 modes listed in table 8. 4. Garnetilerous biotite-quartz-plagioclase eiss, from Central City district. Aver- age of 10 modes listed in Sims and Ga 10 (1964, table 14). 6. Sillimanitic biotite-quartz gneiss. Analysis of biotite in 8378—53 used for compu- tation. Average of 6 modes listed in table 10. M 6. Sillimanitic biotite—quartz gneiss, from Central City district (Sims and Gable, 1964, table 13). Average of 18 modes. . 7. Sillimanitic biotite-quartz-plagioclase gneiss. Analysis of biotite in 8378—53 used for computation. Average of 9 modes listed in table 9. . 8. Garnetifcrous sillimanitic biotite-quartz—plagloclase gneiss. Analysrs of biotite in J GlZ—A, and analysis of garnet in EWT—90 used for computations. Average of 9 modes listed in table 11. _ _ 9. Cordierite-bearing garnet—siliimmite-biotitequartz gneiss. Analysrs for cordierite and garnet in EWT—90 and analysis of biotite in J G12—A used for computation. Average of 10 modes listed in table 12. E26 TABLE 15.—Chemical analyses and modes of cordierite-bearing garnet-sillimanite—biotite gneiss, sillimam'tic biotite-quartz—pla— gioclase gneiss, and garnet—s71limanite—biotite—quartz—plagioclase gneiss [Serial number given in parentheses below sample number. Results of chemical analyses given in weight percent and modes, in volume percent. Tr, trace; Nd not determined; ________ , not found. Analyses of samples 1—4 by C. L. Parker, 5 by E. S. Daniels. Field number is in parentheses after description of sample] 1 2 3 4 5 (14157) (14156) (14158) (I4155) (D 100276) Chemical analyses 58. 06 57. 28 63. 23 62. 67 57. 93 20. 76 21.13 17. 39 19. 12 21. 82 1. 55 2. 22 3. 03 2.11 1. 08 9. 72 7. 90 5. 95 5.13 7. 02 3. 23 3. 04 2. 67 2. 44 2. 91 . 33 . 65 . 36 . 29 . 46 . 39 . 97 . 62 . 72 . 94 3. 25 3. 56 3. 69 4. 50 5.17 . 10 . 14 . 09 . 04 . 05 1.10 1.47 1. 30 1.48 1.21 . 17 . 20 . 20 . 17 . 10 1.00 1.02 1.08 95 1.09 . 07 . 07 . 06 05 . 07 01 02 . 01 02 01 12 11 . 11 11 02 01 01 . 01 02 14 Subtotal __________________ 99. 87 99. 79 99.80 99. 82 100. 02 Less 0 ________________________ . 03 . 03 . 03 .03 . Total ..................... 99. 84 99. 76 99. 77 99. 79 99. 96 Powder density _______________ 3. 01 2. 94 2. 88 2 84 Nd Modes Potassium feldspar ____________ 1. 6 3. 3 7.0 14. 3 15. 6 Plagioclase ________ 2.0 8.0 7. 1 5. 4 7. 4 26.1 29. 6 41. 2 31. 2 31.2 16. 4 4. 3 ______________________________ ll. 7 5. 7 10. 3 16. 3 10. 7 17. 3 20. 1 . 5 __________ 4. 8 23. 6 28. 9 32. 0 32. 1 29. 1 __________ Tr . 2 Tr . 8 1. 3 . 1 1. 7 . 7 . 4 Tr Tr Tr Tr Tr Tr Andalusite.____________.__ZZZ: _________ 1 Total _____________________ 100.0 100. 0 100.0 100. 0 100.0 1. Cordierite-bearing garnet-sillimanite—biotite gneiss; at bottom of gulch 0.5 mile west of Missouri Lake and 0.3 mile west of fork in stream. (J G—38b) 2. Cordierite—bearing garnet-sillimanite-biotite gneiss; taken from outcrop 1.5 miles west-northwest of Bald Mountain Cemetery on Bald Mountain road. (EWT— 90) 3. Sillimanitic biotite—quartz-plagioclase gneiss; taken from outcrop bordering locality of EWT-90 on the west. (EWT—90b) 4. Sillimanitic biotite~quartbplagioclase gneiss; taken from outcrop on north side of North Clear Creek road 3,400 ft west of bench mark (alt 8,513 ft) and adjacent to garnet: and cordierite-bearing sillimanitic biotite-quartz gneiss. (CC—646) 5. Garnet-srllimanite—biotite—quartrrplagioclase gneiss irom vicinity of Petunia mine, northeast corner of quadrangle. (CC—871—A) and A1203 ~2 F‘ e'0+M“g0 ' The ratio A12032FeO—l—Mg0 is slightly less for cor- dierite—bearing varieties than for varieties lacking cordierite. Garnet and cordierite occur in those sillimanitic rocks that contain high molecular pro- portions of MgO and FeO. Cordierite is dependent on a substantial content of MgO. Judged from the available analyses, it is restricted to those highly aluminous rocks that contain 32 percent or more SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY 6 (A) 41> Na20, IN WEIGHT PERCENT I\) K20, IN WEIGHT PERCENT FIGURE 5.——Na20:K20 ratios of biotite gneisses. O, biotite- quartz—plagioclase gneiss; O, sillimanitic biotite-quartz gneiss; e, garnetiferous sillimanitic biotite—quartz gneiss; X, garnet-cordierite-sillimanitebiotite-quartz gneiss. Flame photometer analyses by J. B. McHugh. A1203+FeO+MgO, 11 percent or more FeO+MgO, and that have A1203+FeO+MgO> 6 CaO+Na20+K20 The control of the variables FeO and MgO on the mineralogy of the gneisses is summarized in figure 6. The occurrence of garnet and garnet plus cordierite is limited to rather well-defined fields. Data on the minor element content of the various types of biotite gneisses are tabulated in table 16. In general, abundances are comparable in the different types, but barium, manganese, and strontium contents are somewhat greater and titaniumcontent somewhat less in the biotite-quartz-plagioclase gneisses than in the sillimanitic biotite gneisses. The high manganese con- tent of garnetiferous biotite-quartz—plagioclase gneiss reflects the presence of the spessartite molecule in the almanditic garnet. Zinc is somewhat more abundant in the garnetiferous biotite gneiss than in other rocks. ORIGIN OF METAMORPHIC ROCKS Data gathered during this study support our earlier interpretation (Sims and Gable, 1964) that the meta- morphic rocks are dominantly of metasedimentary origin. The major rock units in the layered succession have the wide areal extent and the lithologic variations across layers that characterize many sedimentary rock sequences. The principal rock units, microcline gneiss and biotite gneiss, are thought to represent respectively arkose and interlayered shale and graywacke. Aside PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. 4 l I I l I x8 — Cordierite and — \ garnet \\ x0 p.— Z 3— \ xc .— Lu 0 \~‘_ g \e if 00 \ '5 ° \ <5 _ l Garnet “ I Go a l 3 Biotitesillimanite o / A 902 — (no garnet) . / _. E / 9 / o / _ / _ o / A / 1 I I I I I 0 2 4 6 8 10 12 FeO, IN WEIGHT PERCENT FIGURE 6.—Fe0:MgO ratios of biotite gneisses. O, biotite- quartz-plagioclase gneiss; A, gametiferous biotite—quartz— plagioclase gneiss; O, sillimanitic biotite-quartz-(plagioclase) gneiss; 69, garnetiferous sillimanitic biotite-quartz-plagioclase gneiss; X, cordierite—bearing garnet-sillimanite-biotite— quartz-plagioclase gneiss; 0, chemical analysis. Data from tables 14 and 15. E27 from amphibolite, whose origin is uncertain, the lesser rock units are considered to be metamorphosed sedi- mentary rocks. Metamorphism appears to have been virtually isochemical, although migmatization resulted in the local addition of some materials. The biotite gneisses of the area have been interpreted, from lithologic and chemical resemblances to unmeta- morphosed sedimentary rocks, to represent interlayered shales and graywackes. The biotite—quartz—plagioclase gneiss, which contains an excess of Na20 over K20, is inferred to have been derived from original graywacke sediments, whereas the sillimanitic biotite gneisses, which have K20 in excess of Na20, are presumed to have formed from original shales. Small differences in the chemical composition of the original sediments caused the variations in mineralogy and chemical composition observed in the metamorphic rocks. Those graywacke sediments that were somewhat rich in lime yielded hornblende-bearing biotite gneisses, as represented by sample 1, table 14, and those rich in iron (samples 3 and 4, table 14) yielded garnetiferous varieties. In the same way, the dominant facies of the shale yielded sillimanitic biotite gneiss; iron- and magnesium-rich and calcium-poor facies yielded garnet— and cordierite-bearing sillimanitic gneisses. Probably this facies was slightly more impoverished in feldspar than the normal shale; the relatively high iron and TABLE 16.—Semiquantitative spectrographic analyses, in parts per million, of minor elements in biotite gneisses [._.. not found; Uteana Oda, analyst] Sample Slegial Ba Be 00 Cr Cu Ga La Mn Mo Nb Ni Pb Sc Sn Sr Ti V Y Zn Zr 0. Biotite-quartz-plagioclaae gneiss 60-2832 700 2 20 200 7, 000 200 15 <150 200 60-2833 300 2 10 150 5, 000 70 30 <150 300 60-2834 200 2 <10 150 5, 000 70 30 <150 300 60—2835 1, 500 1 15 50 5, 000 100 10 <150 200 60—2836 150 1 <10 150 5, 000 70 30 <150 300 60~2837 1, 500 1 20 30 5, 000 100 20 (150 300 60-2839 300 1 <10 <10 1, 500 <10 100 <150 200 61-1308 500 <1 <10 3, 000 15 30 (200 500 61—1309 1, 000 1 30 100 7, 000 150 50 <200 200 60—2844 500 2 30 100 30 20 <50 5, 000 5 10 50 15 15 <10 30 7, 000 150 20 150 150 60—2845 600 <1 <10 <10 30 15 <50 3. 000 5 <10 5 <10 20 <10 50 3, 000 10 70 300 200 60—2846 500 l <10 <10 10 15 <50 2, 000 <5 <10 5 10 20 <10 100 3, 000 <10 70 150 300 61—1301 300 2 20 70 70 15 70 3, 000 5 ______ 30 20 20 <10 200 3, 000 100 70 <200 200 Sillimanitic biotite-quartz gneiss 61—1302 500 <1 15 100 5 15 <50 300 <2 <10 50 20 10 <10 50 7, 000 100 30 200 300 61—1303 300 <1 15 10 5 20 100 300 <2 <10 30 20 15 <10 50 7, 100 100 200 700 60—2840 300 5 15 200 30 15 <50 300 <5 15 50 <10 15 <10 <20 15, 000 100 20 200 300 60—2841 300 1 10 150 30 5 <50 300 <5 10 20 10 10 <10 300 500 50 20 <150 300 _ 60-2842 200 1 20 200 70 15 70 300 <5 15 70 <10 20 <10 70 10, 000 100 50 <150 200 8379—21—53 _________ 60—2843 500 <1 50 200 10 20 150 300 <5 20 100 <10 30 <10 50 15, 000 150 20 200 500 Garnetiferous sillimanitic biotite-quartz-plagioclase gneiss 00—383—1 _________ 61—1304 300 <1 10 200 50 10 <50 200 <2 ______ 30 15 10 <10 150 7, 000 70 10 <200 700 00—858 ........... 61—1307 1, 000 1 30 150 7 50 100 1, 000 2 ______ 70 30 20 <10 100 7, 000 200 50 300 200 Cordierite-bearing garnet-sillimanite-biotite gneiss C C—789—A ......... 61-1305 1, 500 <1 30 200 5 50 <50 700 3 ______ 150 70 30 <10 200 10, 000 300 50 300 500 CC—783—4 _________ 61-1306 700 <1 15 100 3 15 <50 500 <2 ______ 50 20 15 <10 70 7, 000 100 30 <200 200 E28 magnesium content can be attributed to a lack of prolonged weathering. The high Fe02F6203 ratio compared with that of average shale probably resulted from reduction of the original Fe203 during metamorphism. The lithologic layering and the heterogeneity and wide areal extent of the major layers suggest a meta- sedimentary origin for the microcline gneiss, as proposed earlier (Sims and Gable, 1964). Further support for this proposed origin is given by the intimate interlayer- ing of biotite gneiss with microcline gneiss along both margins of the Lawson layer in the vicinity of North Clear Creek and Blackhawk Peak (pl. 1). Gradations both along and across strike near the contacts indicate gradual changes in lithology that can only result from original deposition. In addition to the gradual changes in sedimentation, repetition in the deposition of lithologically similar sediments occurred. Special chemical and mineralogical studies were not made of the calc-silicate rocks, but from their gross composition, heterogeneity, and local association with quartz gneisses or quartzite, one may conclude that they were derived from impure carbonate rocks. The skarns may in part have resulted from iron metaso- matism, as some are too rich in iron to have formed directly from any common carbonate rock, but this derivation is problematic and should be investigated further. The amphibolite has been interpreted as mainly of sedimentary origin because of its association locally with calc-silicate gneiss and quartz gneiss and its gradation into microcline gneiss (Sims and Gable, 1964). How- ever, it may in part be metamorphosed mafic igneous rocks, possibly spilite. The association of spilites with graywacke and shale sediments is not uncommon in eugeosynclinal terranes, as for example in the Welling- ton district of New Zealand (Reed, 1957). The origin of the cordierite-amphibole gneiss is un- certain. If metamorphism was nearly isochemical, as seems probable from the data obtained on other meta- morphic rocks in the sequence, the gneiss has no com- mon unmetamorphosed equivalent. It has abnormally high percentages of both ferrous iron and magnesium and abnormally low amounts of calcium, sodium, and potas- sium compared with those in known sedimentary or igneous rocks. It most closely approximates iron- formations in composition, as these rocks are exceed- ingly low in alkalic content, but it contains substantially more magnesium than do typical iron-formations. Pos- sibly it represents an unusual magnesium- and iron-rich chemical sediment of an origin similar to that of iron- formations. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY GRANITE GNEISS AND PEGMATITE Granite gneiss and pegmatite constitute the felsic material in migmatites and less commonly form discrete bodies a few feet to several feet wide. They are esti- mated to form 15—20 percent by volume of the mig- matitic biotite gneisses, and in addition they occur as a network of veinlets in some amphibolite bodies and as irregular, generally conformable bodies in microcline gneiss. In general, granite gneiss and pegmatite de- crease somewhat in abundance from the southern part of the quadrangle to the northern part, and concomitantly become coarser grained northward. Because of the generally small size and the abundance of the bodies, the granite gneiss and pegmatite were not mapped sepa- rately on plate 1; instead, they were included with the dominant rock units. GENERAL CHARACTER Granite gneiss and pegmatite constitute a light-gray or yellowish-gray medium- or coarse-grained rock that is somewhat inequigranular. They resemble microcline gneiss but generally can be distinguished from microcline gneiss both by their mode of occurrence and by their megascopic appearance. Granite gneiss and pegmatite typically occur as thin stringers and tabular bodies in biotite gneisses to constitute migmatite, whereas the microcline gneiss generally forms larger, more distinct layers. In appearance, most exposures of granite gneiss and pegmatite are leucocratic, nearly massive, and gen- erally homogeneous, except for local wisps and streaks of biotite-rich gneiss, whereas the microcline gneiss characteristically has a conspicuous even layering. Contacts with adjacent rocks appear to be sharp mega— scopically but in detail are seen to be transitional. PETRO GRAPHY The rock consists mainly of potassium feldspar, pla- gioclase, and quartz and contains biotite, magnetite-il- menite, muscovite, zircon, sillimanite, garnet, and sphene as accessory minerals (table 17). Typically it has a granoblastic texture. Potassium feldspar is perthitic and contains a few percent of plagioclase as film, string, and patch inter- growths. About two-thirds of the grains have a con- spicuous grid twinning. In any given sample, crystals of potassium feldspar tend to be the coarsest grains in the rock. Plagioclase (oligoclase) is slightly antiperth- itic and is twinned according to the albite and Carlsbad twin laws. Albitic rims as much as 0.03 mm wide are common where plagioclase is in contact with potassium feldspar. Myrmekite is present in many sections. Quartz that has conspicuous strain shadows forms ir- regular anastomosing grains and smaller subrounded PETROLOGY AND STRUCTURE, PRECAIVIBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. grains. Biotite is nearly ubiquitous but sparse and commonly is intergrown with muscovite. It is a green- ish-brown pleochroic variety, generally somewhat al- tered to chlorite, and in appearance resembles the biotite in the biotite gneiss units. Sillimanite and garnet are local accessory minerals. The sillimanite forms sub- hedral grains in stringers and clots; it is associated With muscovite and at places where in contact with po- tassium feldspar is separated from it by sheaths of muscovite. Zircon forms tiny anhedral grains, mainly in plagioclase and quartz. Xenotime and monazite occur locally in related pegmatites in the eastern part of the Central City district; their occurrence has been described previously by Young and Sims (1961). Specimens from Dakota Hill and vicinity difl“er from typical samples in being finer grained and in having a cataclastic texture. Hand specimens appear sheared; the mafic minerals are in streaks and wisps and quartz occurs as discontinuous veinlets. In thin section the quartz is seen to form elongate aggregates, subparallel to cataclastic zones in Which the plagioclase, in par- ticular, is granulated. Most plagioclase grains are partly altered to clay minerals. The potassium feldspar has strain shadows and lacks grid twinning. Biotite is fresh and lies dominantly in the cataclastic zones. Potassium feldspar from typical pegmatitebelonging to this granite gneiss and pegmatite unit has been determined from two localities in the Central City E29 district to contain about 78 percent KAlSiaog by weight after homogenization; it has a triclinicity index of 0.75 (Sims and Gable, 1964). An estimated chemical composition of the pegmatite was given in the same report. Flame photometer analysis by J. B. McHugh of three samples gave the following content, in percent: Sample NaO K10 CC—1045—A _____________________________ 3. 0 7. 6 CC—1124—1 _____________________________ 3. 9 4. 7 D— ____________________________________ 2. 6 7. 6 ORIGIN The origin of granite gneiss and pegmatite was discussed in the report on the Central City district by Sims and Gable (1964) and need not be discussed in detail here. In brief, we infer that the granite gneiss and pegmatite formed from a fluid that was able to penetrate readily along the foliation planes of the biotite gneiss country rock, part of which it replaced. A source for the fluid through ultrametamorphism is favored, but an origin from a silicic melt cannot be dismissed from consideration. If the muscovite in the rock is primary, as seems probable from its textural relations, the pegmatite can be inferred from the experimental data of Yoder and Eugster (1955, p. 267) to have formed at more than 1,500 atmospheres of water pressure. INTRUSIVE ROCKS Four types of intrusive rocks—granodiorite and associated rocks, gabbro and related rocks, quartz TABLE 17.—Modes, in volume percent, of granite gneiss and pegmatite [Tr, trace; Nd, not determined; ---- , not found. Field number is in parentheses after description of sample] Mineral 1 2 3 4 5 6 7 8 9 10 ll 12 13 Potassium feldspar ............... 27. 8 36. 61. 52. 1 37. 5 35. 4 37. 9 30. 3 54. 2 28. 7 26. 5 50. 6 40. 0 Plagioclase ______________________ 28. 9 25. 4 18. 6 19. 5 7. 3 33. 3 29. 5 32. 0 14. 4 34. 5 32. 4 22. 2 26. 8 Quartz _________________________ 39. 8 35. 7 20. 0 27. 3 55. 0 27. 4 28. 8 E30. 0 23. 9 30. 2 31. 6 23. 6 32. 6 Biotite _________________________ . 8 Tr Tr . 1 ______ 1. 9 1. 5 1. 8 . 8 . 6 1. 3 1. 6 . 1 Magnetite—ilmenite _______________ 1 . 7 ...... . 2 Tr Tr . 7 . 6 . 3 . 7 . 6 Tr . 1 Muscovite ______________________ 1. 3 2. 2 . 4 . 7 . 2 1. 0 1. 6 4. 2 6. 2 4. 0 5. 5 1. 9 . 3 Zircon __________________________ r . 1 Tr Tr Tr . 1 Sillimanite ______________________ . 5 . 1 1. 3 2. 1 . 1 ...... Chlorite ________________________ . 6 ______ Tr __________________ Garnet _______________________________________________________ Sphene _________________________________________________________________________________________________ Tr Total _______________________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 Composition of plagioclase ________ Nd Nd Nd Nd Nd Nd Ann An“, AD" An“, An" Nd Nd Average grain diameter ______ mm__ 0. 7 1. 3 1. 3 1. 7 1. 4 Nd Nd Nd Nd Nd Nd Nd 0. 5 1. Interlayered with sllllmanitlc biotite uartz gneiss- collected north of Fall River 7. From outcrop on north side of road. just below junction of North Clear Creek 0.9 mile from west edge of quadrangle. (CC—615 and Freeman Gulch. (CC—426A) 2. Interlayered with biotite gneiss, taken from nose olridge north of Fall River and 8. Interlayered with biotite-quartz-plagioclase gneiss, on old road along slope east of Hamlin Gulch. (CC—110) between Peeks Flat and North Clear Creek. (CC—443B) 3. From outcrop east side of Mount Pisgah. (CC—175) 9. Interlayered with biotite gneiss, taken 0.75 mile east of Yankee Hill. (00—535) 4. Leucocratic granite gneiss and pegmatite from contact zone between granodiorite 10. Sillimanitic granite gneiss and pegmatite in biotite gneiss collected 0.5 mile south and migmatitic sillimanitic biotite—quartz gneiss, northwest tip of Mount Pisgah pluton. (CC—313A) . From just northwest of locality ol saméale 4; contains layers and streaks of silllmanitic biotite-quartz gneiss. (C —316—1 . Sillimanitic granite gneiss and pegmatite, from outcrop along North Clear Creek on nose of Blackhawk Peak. (CC—366-1A) 050! (00667—2) From outcrop just west of locality of sample 10. (CC-669) FIE-Eli: i162;- An)orth of American City, northwestern part of quadrangle From ridge on west side of Dakota Hill, 0.5 mile northeast of Apex. Rock is sheared and granulated. (CC—1124—1) 11 of Kingston, west margin of quadrangle. 12: 13. E30 diorite and hornblendite, and biotite-muscovite quartz monzonite—each with associated pegmatites, intrude the layered rocks of the quadrangle and constitute about 15 percent by volume of the exposed bedrock. Each intrusive rock type mapped in the quadrangle, except gabbro and related rocks, is widespread throughout the central part of the Front Range, and now the types can be correlated with confidence from one area to another. The granodiorite and associated rocks of this report are equivalent to the intrusive rocks of the Boulder Creek batholith that are exposed just west of Boulder (Lovering and Goddard, 1950, pl. 2), the type area of the Boulder Creek Granite. The gabbro and related rocks are correlated with gabbro found southwest of Fraser in the Vasquez Mountains (Taylor and Sims, 1962). They are linked to the Vasquez Mountains occurrence by lithologic similarity and by age relative to known Precambrian events. Quartz diorite and hornblendite form numerous small scattered bodies in the region but have not been formally named. The biotite-muscovite quartz monzonite is equivalent to the Silver Plume Granite at the type area at Silver Plume, 0010., and to the biotite-muscovite granite recently described from nearby areas in Clear Creek County (Harrison and Wells, 1956, 1959). GRANODIORITE AND ASSOCIATED ROCKS DISTRIBUTION AND CHARACTER Rocks that range in composition from a mafic quartz diorite to quartz monzonite but have a com— position similar to granodiorite form scattered small or moderate-sized bodies in the quadrangle. The bodies are more abundant toward the northeast, and thus appear to be satellitic to the batholith of Boulder Creek Granite (Lovering and Goddard, 1950, pl. 2). Although the rocks vary rather widely in composition and overlap other intrusive rocks of the region, they constitute a diagnostic type that can be distinguished readily from biotite-muscovite quartz monzonite and from gabbro and related rocks. The granodiorite and associated rocks typically form lenses in the apical areas of folds and sheets on the limbs of folds. Several of the bodies are notably thicker in the axial areas and are interpreted as phaco- liths. The bodies dominantly lie within thick biotite gneiss units, but a few are in the contact zones between biotite gneiss and microcline gneiss units and at least one is at the contact between biotite gneiss and the Elk Creek pluton of gabbroic rocks. The contacts against the country rocks are sharp and only rarely are marked by a zone of interlayering a few feet wide. As a generalization, the contacts are grossly con- cordant with the gneissic structure of the country SHORTE‘R CONTRIBUTIONS TO GENERAL GEOLOGY rocks, but in detail they can be seen to crosscut, and on the scale of the map (pl. 1) they locally appear to cut across some of the larger units, as in the vicinity of Peeks Gulch. The internal structures of the bodies vary as widely as does the composition. All small bodies have a wholly gneissic structure. The larger bodies, how- ever, tend to have weakly foliated interiors and mod- erately or strongly foliated borders. The foliation and lineation within the rocks is parallel to that in the metasedimentary country rocks and is interpreted to have resulted from syntectonic enplacement. The largest body, exposed in the vicinity of Mount Pisgah (pl. 1), and herein called the Mount Pisgah pluton, has a surface area in excess of 1 square mile. It lies in the contact zone between the Lawson layer of microcline gneiss and the underlying unit of biotite gneisses. In most exposures the granodiorite is vir- tually conformable with the gneisses, but on a gross scale its east margin cuts across the internal gneissic structure of the microcline gneiss unit. In detail, prongs such as the one mapped on the south slope of Mount Pisgah (pl. 1) clearly transect the adjacent gneisses. The pluton is moderately thick in the trough of the synchne at Mount Pisgah and in the crest of the adjacent Pecks Flat anticline, but it pinches out abruptly on the limbs and accordingly is inter- preted as a compound phacolith. The borders are strongly foliated and lineated, whereas the interior is nearly massive. Inclusions which are tabular lenses or pods and which conform to the internal structure of the body tend to be elongated parallel to the lineation and are more abundant than in other intrusive masses in the quadrangle. The inclusions are mainly biotite gneiss but include interlayered quartzite and calc—silicate gneiss. The body is cut by a few small dikes of biotite-muscovite quartz monzonite that are too small to be shown at the scale of plate 1. The pluton is more felsic than other bodies in the quadrangle (table 18). Apparently because of its relatively felsic composition, the body was mapped previously as Silver Plume Granite by Bastin and Hill (1917) and by Lovering and Goddard (1950, pl. 2). Another body of moderate size is exposed on the _ upper slopes of Bald Mountain (pl. 1), within the bio~ tite gneiss unit that lies between the Lawson and Quartz Hill layers of microcline gneiss. It contrasts sharply with the Mount Pisgah pluton in being more mafic in compositon (table 18) and in having a wholly gneissic structure. The body is either a folded sheet or a phaco- lith estimated to be about 700 feet thick. (See section 0—0”, pl. 1.) It was mapped earlier by Bastin and Hill (1917) as granite gneiss, for they considered it to be PETROLOGY AND STRUCTURE, PRECAMIBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. equivalent to the granitic gneiss that we have named microcline—quartz—plagioclasebiotite gneiss (Sims and Gable, 1964). Lovering and Goddard (1950, pl. 2) showed the body to be quartz monzonite gneiss. Sub- sequently, in detailed mapping of the Lawson-Dumont— Fall River district, Hawley and Moore (1967) mapped the body as quartz diorite gneiss. A thin layer of quartz diorite gneiss, considered to be a part of the granodiorite unit of this report, was mapped by Hawley and Moore (1967) in the Lawson area but is omitted from plate 1 because of its small size. Several other bodies, mostly of small dimension, crop out in the quadrangle. A folded, subconcordant, com- posite sheet estimated to be about 400 feet thick is exposed south of Central City. It is strongly differen- tiated and ranges in composition from a quartz diorite to a fluorite-bearing quartz monzonite (Sims and Gable, 1964). A few moderate-sized conformable bodies and small lenses are exposed in the northeastern part of the quadrangle, within the biotite gneiss unit intruded by the Bald Mountain pluton. On the whole these bodies are relatively uniform in composition and contain very few parts as felsic as quartz monzonite. Some, as for example the curved body exposed along and adjacent to State Highway 119 at the east edge of the quadrangle, contain substantial amounts of biotite-hornblende quartz diorite. In the northwestern part of the quad- rangle several long, thin, conformable sheets, some of which are probably folded, crop out Within biotite gneisses, and a few tightly folded lenses or phacoliths occur in the biotite gneisses and at the contacts of bio- tite gneiss with microcline gneiss. These bodies are dominantly granodiorite and quartz diorite in compo- sition and, like the small bodies in the northeastern part of the quadrangle, appear to be relatively uniform throughout. PETROGRAPHY The granodiorite and associated rocks are gray, mot- tled black and white, medium-grained, generally gneissic rocks that vary widely in composition and appearance. The more gneissic facies of the granodiorite is generally finer grained. In addition to the dominant facies—a medium—gray biotite granodiorite—a biotite-rich facies containing 20—30 percent biotite and a biotite— and hornblende-rich facies containing 30—50 percent dark minerals occurs. Where the field relations are clear cut, the dark facies are interlayered conformably with the light facies. In general, the dark facies has a more strongly developed foliation than does the light facies; the foliation is due to a subparallel alinement of biotite and hornblende crystals and to a weak compositional layering. Exposed bodies of the dark facies generally are associated with a coarse-grained felsic pegmatite. E31 Although the granodiorite apparently is medium grained and nearly equigranular, local coarse-grained facies contain microcline and plagioclase laths, which are as much as 2.5 cm long, in a finer grained matrix. A few small lenses just east of Yankee Hill contain scat- tered blue-gray plagioclase crystals as much as 10 cm long. The granodiorite and associated rocks have a hypidio- morphic or less commonly an allotriomorphic granular texture. The most abundant mineral, plagioclase, ranges in composition from oligoclase to calcic andesine and rarely to labradorite and generally contains small patches and blebs of potassium feldspar. Commonly the crystals have a weak gradational normal zoning. The plagioclase forms subhedral or anhedral laths that have well-developed twinning—albite, Carlsbad, and albite-pericline twins being the most common. Albitic rims of an average width of 0.025—0.05 mm occur on some plagioclase crystals adjacent to microcline. Myrmekitic intergrowths of quartz and oligoclase— andesine are common. The potassium feldspar is perthitic microcline, which contains 30—50 percent by volume of plagioclase, or microcline that has a fairly well-developed grid twinning. Locally, however, as at Bald Mountain, it lacks visible microcline grid twinning and is diflicult to identify unless it is stained. Quartz generally has conspicuous strain shadows. In some aggregates, quartz grains in contact with one another have serrated borders. Inclusions of subrounded clear quartz grains are common in the feldspars. The biotite is pleochroic, ranging from straw yellow to olive greenish brown. It forms laths that tend to be clustered with other mafic minerals, including magnetite-ilmenite, hornblende, allanite, sphene, and epidote. It cuts and embays hornblende. Hornblende forms euhedral to anhedral crystals and generally is altered somewhat to magnetite and biotite. Lamellar twinning is evident where the hornblende is bleached and altered. The pleochroism of the hornblende varies from pale brownish yellow to a dark olive green or dark bluish green. Allanite is a persistent and characteristic accessory mineral but never exceeds 1 percent by volume. Crystals of the mineral generally range in maximum dimension from 0.1 to 0.4 mm; in the Mount Pisgah body, however, one observed crystal was 3.1 mm in diameter. Commonly the allanite is zoned, with an outer dirty greenish-brown zone that grades inward to reddish brown or brown. The crystals are both positive and negative, have a large 2V, and a moderate pleochroism ranging from light tan to a reddish brown. Magnetite-ilmenite commonly is rimmed by sphene Where abundant, sphene embays biotite. Alteration minerals in the granodiorite and associated rocks formed subsequent to consolidation and include E32 epidote, chlorite, magnetite, sericite, clay minerals, and muscovite. Mineralogical variations of the rock unit in various bodies within the quadrangle are shown by the modes (in volume percent) in table 18. A summation of the mineralogic variations is given in figure 7. In brief, the Mount Pisgah pluton ranges in composition from a quartz diorite to felsic quartz monzonite and has an average composition, as determined from 18 modes, of 27 percent quartz, 38 percent plagioclase, 17 percent microcline, and 16 percent total mafic minerals. The Bald Mountain pluton has a similar range in mineralogy but is dominantly quartz diorite. The other smaller bodies in the quadrangle have comparable ranges in composition. CHEL’IICAL COMPOSITION Analyses of a sample of the quartz monzonite facies and of a hornblende-quartz diorite facies of the gran- odiorite unit are given in table 19. Analyses of two samples, a granodiorite and a quartz monzonite facies, from the Central City pluton were presented earlier SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY Quartz Granite Quartz monzonite Potassium feldspar Plagioclase FIGURE 7.—-Variation in composition (in volume percent) of granodiorite and associated rocks. I, in the Mount Pisgah pluton; O, in the Bald Mountain pluton; and A, in scattered small plutons (53 plots). TABLE 18.—M odes, in volume percent, of granodiorite and associated rocks [Tr, trace; Nd, not determined; ______ , not found. Field number is in parentheses after description of sample] Mount Piszah pluton Mineral 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 Average (1—18) Potassium feldspar ....................... 27. 4 13. 1 34. 4 2. 5 16. 7 Plagioclase _______ 28. G 39. 1 32. 9 56. 4 38. 0 uartz. _ 24. 7 25. 9 18. 3 29. 1 27. 2 Biotite. . 13. 3 19. 9 11. 9 7. 8 l3. 1 Muscovi ..__ 1.5 .9 1.0 1.4 1.5 Magnetite-ilmenite. 2. 6 . 5 . 6 1. 6 2.5 Apatite .......... . 9 . 2 . 4 ______ . 4 Sphene__ 1. 0 ______ Tr . 2 Ca1cite...- '3 1. 2 . 1 Chlorite... Tr ______ . 3 Zircon _______ 2 Tr Tr Allanite ................. Tr Epidote __________________________ Tr Total _________________ 100. 0 100. 0 99. 9 100. 0 99. 8 99. 1 100. 0 100. 0 99. 9 100. 0 100. 0 100. 0 Composition of plagioclase ................ Amv A1123 A1127 Anzs A1121 A1127 An27 Aug-1 Ann A1133 A1130 Anzac Nd Ann Average grain diameter ............. mm.- 0.6 0.8 1. 5 1. 2 0.7 1. 0 1. 5 0.9 0.6 0.4 0. 4 Nd Nd 0.9 Bald Mountain pluton Mineral 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 I}?! e 1 —37 Potassium feldspar _________________ 7. 2 25. 6 ...... 0. 2 0. 1 _________________ 3. 0 2. 3 14. 0 Tr ______ 22. 7 ______________________________ 4. 0 Plagioclase ________ 43. 5 33. 5 54. 6 49. 7 56. 5 52. 2 39. 5 54. 52. 0 39. 4 49. 0 51. 1 40. 7 35. 3 53. 3 54. 5 52. 4 88. 7 45. 4 47. 0 Quartz ________ 30. 0 28. 1 18. 5 11. 9 24. 7 28. 1 11. 3 22. 31. 4 38. 4 26. 3 21. 5 51. 6 28. 0 18. 5 17. 5 13. 5 10. 9 24. 4 24. 0 Biotite ________ 16. 8 10. 4 22. 6 13. 2 17. 2 18. 5 19. 0 21. 12. 5 16. l 9. 8 . . 5 1 14. 6 3. 2 18. 5 21. 8 12 1 l5. 3 Muscovite ________ - 1.8 1.9 .1 Tr .4 .2 ______ . .5 .6 .1 .1 .5 ________________________ .4 Magnetite—ilmenite . 5 . 3 2. 5 1. 2 . 4 . 5 Tr 1. . 4 . 3 . 2 . 9 . 1 . 9 Apatite ___________ Tr . l 1. 0 Tr . 5 2 . 5 . . l ______ . 2 ______ . 1 . 3 Sphene ........... .-_ Tr .................. Tr _____ Tr Chlorite __________ _ . 1 Tr . 1 ’I‘r Tr Tr ______ 2. 8 . 2 7. 2 a Tr . 2 100. 0 Composition of plagioclase _________ Ange A1131 Ange A1134 Ann A1134 Am: Ann A113! A112: Ann Ann, A1134 Auzs An44 Nd A1144 A1135 Average grain diameter ...... m m_. 0. 8 0. 5 0. 6 0. 5 0. 6 0. 5 0. 4 0. 6 0. 8 Nd Nd 0. 3 Nd 0. 6 0. 5 0. 6 0. 4 0. 4 0. 6 0. 5 See notes at end of table. PETROLOGY AND STRUCTURE. PRECAlVIBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. E33 TABLE 18.—Modes, in volume percent, of granodiorite and associated rocks—Continued Smaller scattered bodies Mineral 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 Aver 3 (38-58 Potassium feldspar... . 2 9. 7 ____________ 10. 4 1. 4 8. 0 9. 0 ...... 22. l 6. 5 ______ 7. 7 Plagloclase._.____ . 0 38.1 38.2 51. 5 43.9 45.1 50. 4 52. 9 43. 2 52. 9 49.7 35. 3 38. 8 51. 4 43. 8 Quartz- - . 2 8.0 15. 4 25.0 17. 8 5. 4 19. 9 26. 0 30.0 19. 4 30. 6 26. 9 27. 5 17. 7 23. 6 Blotite- _ .8 4. 8 14. 9 12. 5 20.0 11. 6 6.6 17. 9 10. 7 9. 2 15. 2 .1 17. 8 12. 9 13. 7 Muscovit .__ __ 7 ...... .7 .5 3.3 ____________ 1.2 1.8 Tr .3 .4 ____________ .6 Magnetite-ilm ni _ 3 1 5 1.6 .3 l. 2 .6 3. 2 .4 5. 5 4.0 3.7 .1 3.5 5.9 2. 2 Apatite ______________ .1 .2 .1 .5 .5 r 1.9 ______ .8 1.0 .5 .9 .9 1.0 .6 Spheue. Tr .................. 1. 9 2. 2 1. 8 2. 7 3. 4 . 7 Calcite ________________ . 5 __________________ 4 . 1 ' Tr 1. 0 2. 3 Tr Tr . 1 . 2 ______ 45.9 24. 1 ____-- 12. 5 35.3 5.0 ____-, -_--__ _--_._ __-_-. Tr 9 7.6 6.5 Tr Tr Tr Tr Tr Tr Tr 1 Tr Tr Tr . 1 Tr Tr _____ - . _ Tr __________________ Tr Tr Tr ______ Tr Tr Tr ______ . 4 Tr 1 Epodote ______________ Tr ...... 2. 6 ...... 8 ______ 2 1 Tr . 1 Tr 3 1. 0 ______ . 2 Total _____________ 100.0 100.0 99. 9 100.0 100.0 100.0 100.0 100. 0 100.0 97. 8 100.0 100. 0 100.0 100. 0 100. 0 Composition of plagloclase __________ A1131 A1114 A1145 All“ A1145 Ami Ania Nd A1150 An“ Anal; Arm Ania Allao Nd Nd A1132 Nd A1115 All“ Amt A1133 Average grain diameter (mm) _____ 0.4 Nd 0.6 0.6 0.6 0.9 0.6 0.5 0.3 0.3 0.9 0.4 0.6 1.0 0.6 0.3 0.7 Nd Nd Nd Nd 0.6 1. Quartz monzonite, from crest of Mount Pisgah. (1) . 32. Granodiorlte, irom outcrop 600 it due east of peak of Bald Mountain. (EWT- 2. Granodmrite, from outcrop along south edge of pluton, 600 it from contact wrth 70—54 srlllmanlte-blotlte-quartz gneiss. ) _ 33. Homblende quartz dion'te, from pit in the southeasternth part of the bulge of 3. Fresh quartz dlorlte, from plt adjacent to roadcut in east edge of pluton. (13) the pluton. (EWT—63—54) 4. Quartz diorlte, 500 ft northwest of locality of sample 3, adjacent to a sillimanite— 34. Homblende quartz diorite, from outcrop 600 it north of locality of sample 33. biotite-quartz lens. (13b) _ (EWT—71—54) 5. Quartz monzomte, from southeast slope, Mount Plsgah, 80ft below crest. (18b) 35. Homblende quartz diorite, fine-grained, from it south of Eureka Gulch;adjaoent 6. Quartz monzonite, from outcrop 500 ft southeast of locality of sample 2, near to migmatitic biotite—quartz gneiss lens. EWT—80a—54) blotite gnelss contact. (30) . . 36. Biotite-hornblende quartz diorite, fine-grained, irom outcrop adjacent to biotite 7. Quartz monzonlte, from outcrop just south of Lake Plsgah and west of the Pisgah gneiss, about 250 ft west-northwest of locality of sample 35. (EWT—82—54) road. (L1) . ' 37. Homblende quartz diorite, from outcrop 2,500 ft northwest of peak of Bald Moun- 8. Quartz monzonlte, from knob due east of Lake Pisgah and the Plsgah road. (L8) tain and just south of Eureka Gulch. (EWT—83—54) 9- Quartz dlorlte. from outcmp 1.000 ft west of Lake Pisgah- (L10) 38. Quartz diorlte, lens on ridge between Miners Gulch and North Clear Creek, 10 northwest corner of mapped area. (CC—664) . Quartz monzonite, from outcrop on west slope, Mount Pisgah, 200 ft below crest. ( 11. Granodiorite, from outcrop 600 ft southwest of locality of sample 10. (R98) 12. Quartz mcnzonite, from outcrop due west of mid int between Lake Pisgah and the top oiMount Plsgah, just above flume. ( 12) 13. Grangdiorite, from outcrop 1,500 it north of Lake P'sgah and 500 ft west of Pisgah roe. . 14. Quartz diorite, from outcrop on prominent knob northeast of Lake Pisgah. ($6) 15. Quartz diorite, from knob on Peeks Flat. (CC-330) 16. Quartz diorite, from outcrop 2,000 ft northeast of locality of sample 14. (CC—375) 17. Qiarfitazfiorite, from outcrop west of road, east slope of Mount Pisgah. (S373 18. Qlfiarst; monzonite, from outcrop west of road, east slope of Mount Pisgah. (S374 19. Quartz diorlte, medium- to coarse-grained, from outcrop west or peak 01 Bald Mountain at an altitude of 9,750 ft. (CCH—24—B) 20. Granodiorite, from outcrop 1,000 ft dowuslope southeast from peak of Bald Moun- tain, adjacent to silllmanite-biotite-quartz gnelss contact. (OCH—91 21. Biotite quartz diorite, from outcrop 1,000 it due east of peak of Bald Mountain, and just north of well-defined east-west Tertiary dike. (ET—129) Hpfinqblgpélg quartz diorite, from outcrop 1,000 ft south of locality of sample 21. 23. nggz diorite, from outcrop 2,500 ft northeast of peak of Bald Mountain. (ET— 2 24. uartz diorite, from outcrop 500 it southeast of locality of Sample 20. (ET—220) 25. omblende quartz dion'te, fine-grained, from outcrop 500 ft northwest of locality of sample 24 in the vicinity of biotite-muscovite quartz monzonite. (ET—296) 26. Biotite guartz diorite, from outcrop in the area of sample 19. (OCH-2+0) 27. Quartz ion’te, medium- to coarse—grained, from outcrop 250 it east-southeast of localit of sample 19. (CCH—25—A) 28. Quartz lorite, from pit along northeast edge of pluton, 0.5 mile north of Eureka Gulch. EWT—45—54) 29. Granodiorite, from outcrop near center of pluton, just north-northwest of locality of sample 28. (EWT—46—54) 30. Biotite quartz diorite, fine-grained, irom outcrop near center of pluton and north of Eureka Gulch. (EWT—52—54) 31. Quartz diorite, from outcrop adjacent to granite gneiss and pegmatite, 1,000 ft north of locality of sample 28. (EWT—58—54) N: 5° 39. Quartz diorite, southernmost tip of above lens. (CC—536) 40. Biotite-hornblende quartz diorite; cuts across Missouri Gulch road 1.5 miles north of junction of road with State Highway 119. (J G-9a) 41. Granodiorite, 0.25 mile north of locality of sample 40, associated with quartz diorite, homblendite, and pegmatite. (J 6—21) 42. Gggodgqgite, southwest prong of large body in northeast corner of mapped area. —1 43. Qllnlagtz 38113336), peak 10,383 ft north of bench mark 9101 on State Highway 44. Qrzgréz 8di3rite, northwest shoulder of peak from which sample 43 was taken. — 1 45. Quartz monzonite, taken 0.25 mile east of locality of sample 44. (CC-818a) 46. Biotite-hornblende quartz diorite, lens west of Colorado Creek, northeast edge of mapped area. (CC-855—B) . 47. Biotite-homblende quartz diorlte, along State Highway 119, and 0.25 mile north- east of Missouri Lake. 10—148) 48. Quartz dion‘te, along State Highway 119, Cold Spring campground, near bench mark 9101. (J G—166) 49. Biotite-hornblende granodlorite; sample taken adjacent to locality of sample 48. (JG—166—1) 50. Biotite-homblende quartz diorite, taken midway between localities of samples 47 and 48. (J G—172) 51. Biotite-hornblende granodiorite, taken midway between Blackhawk Peak and Oregon Hill. (CC—789—B) 52. Quartz diorlte; same locality as that of sample 51. (CC—790—A) 53. Gpgpogiorite, taken just east oilocality of sample 52, along Silver Creek. (CO- 54. Glgengdlipgge, lens just north of Montana Creek, northeast edge of mapped area: 55. Quartz diorite, Michigan Hill area. (5381-11—53) 56. Quartz mcnzonite, small outcrop west of northernmost prong of Mount Pisgah along North Clear Creek road. (CC—431) 57. Biotite-hornblende granodiorite, from Sheridan Hill area. (CC-495) 58. Biotiée-gitflaiinblende quartz diorite, from small body 1 mile east of Yankee Hill. (Sims and Gable, 1964). The few chemical analyses reflect the same range in composition as that indicated by the modal analyses. Judged from a single analysis of quartz diorite from Bald Mountain (Sample 2, table 19), the rock differs chemically from the quartz diorite facies in the Elk Creek pluton mainly in its higher aluminum and magnesium content and lower titanium and total iron content. Apparently the quartz diorite facies of the granodiorite unit. can be distinguished chemically from the quartz diorite facies of the gabbro unit in the Elk Creek pluton by its significantly lower titanium content. Chemical and spectrochemical analyses of a sample of biotite from the quartz monzonite phase at Mount Pisgah are given in table 20. The biotite differs slightly from a biotite in biotite-muscovite quartz monzonite, mainly in its higher aluminum and calcium content and lower total iron content. It contains more calcium than any biotite that has yet been analyzed from the district. E34 TABLE 19.—Chemical and spectrochemical analyses and norms and modes of intrusive rocks [Laboratory number given in parentheses below sample number. Field number is in parentheses after description of sample. Results of chemical analyses given in weight percent and of spectrochemicalanalyses, in parts per million; modes given in EolumfiJpercent. Tr, trace. Chemical analyses by Dorothy Powers and P. R. ame Sample 1 2 3 4 (6-3098) (Gr—3099) (G—3100) (G—3101) Chemical apalyses SiOa ___________________ 64. 37 54. 41 49. 56 54. 23 A1303 __________________ 15. 86 17 12 13. 70 15. 38 1. 78 3 14 6. 17 4 66 3. 04 5 60 9. 27 6. 84 1. 69 5 18 4. 09 3 70 2. 37 7 25 7. 46 6 14 3. 09 2 98 2. 46 2 97 5. 00 1 40 l. 22 1 35 . 05 15 . 17 16 . 52 1 23 . 79 95 . 08 07 . 06 08 72 85 3. 75 2 26 32 07 . 41 39 23 . 08 . 29 36 03 07 . 02 03 12 07 . 12 11 14 10 . 36 18 23 . 06 . 07 11 Subtotal ___________ 99. 64 99. 83 99. 97 99. 90 Less 0 _________________ . 13 . 10 . 23 . 15 Total ______________ 99. 51 99. 73 99. 74 99. 75 Bulk density ___________ 2. 66 2. 83 3. 00 2. 81 Powder density _________ 2. 73 2. 89 3. 06 2. 92 Spectrochemical analyses Co ____________________ 8 34 51 33 Cr ____________________ 26 220 48 66 Cu ____________________ 100 73 120 77 Ga ____________________ 23 25 28 26 La ____________________ 140 <100 <100 <100 Ni ____________________ 20 85 44 58 Pb ____________________ 50 <30 <30 <30 Sc _____________________ 13 28 40 22 Sr _____________________ 620 540 500 500 V _____________________ 80 260 600 260 Y _____________________ 50 30 50 50 Yb ____________________ 4 3 4 4 Zr _____________________ 700 130 280 360 Norms Quartz _________________ 19. 44 7. 20 10. 20 13 26 Orthoclase _____________ 29. 47 8. 34 7. 23 7. 78 Albite _________________ 26. 20 25. 15 20. 96 25. 15 Anorthite ______________ 10. 29 29. 19 22. 52 24. 74 Diopside______---______ _; ...... 4.73 7.85 .68 Hypersthene ______ _ 6 34 17 04 11.56 13 85 Magnetite ______________ 2. 55 4. 64 ‘ 9. 05 6. 73 Ilmenite _______________ 1. 37 1. 67 7. 14 4. 26 Corundum _____________ 1. 63 ________________________ Apatite ________________ . 66 ________ . 96 . 96 Pyrite _________________ . 48 . 36 1. 32 . 72 Fluorite ________________ . 16 . 16 . 16 . 16 Calcite (secondary) ______ . 50 . 20 . 70 . 80 See notes at end of table. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY TABLE 19.—C’hemical and spectrochemical analyses and norms and modes of intrusive rocks—Continued I Sample 1 2 3 4 (Ci-3098) (G-3099) (Ci—3100) (G-3101) Modes Quartz _________________ 18. 3 11. 9 12. 7 13. 2 Potassium feldspar ______ 34. 4 0. 2 1. 5 ________ Plagioclase _____________ 32. 9 49. 7 46. 1 55. 1 Biotite _________________ 11. 9 13. 2 10. 5 11. 4 Muscovite _________ __ 1. O Tr ________ Tr Hornblende ____________________ 23. 3 2. 0 13. 8 Magnetite—ilmenite ______ . 6 1. 2 10. 0 3. 7 Epidote ________________________ . 5 . 7 ________ Allanite ________________________ Tr Tr ________ Apatite ________________ . 4 Tr . 2 . 3 Zircon _________________ . 2 Tr Tr Tr Chlorite _______________ Tr Tr 1. 6 1. 2 Orthopyroxene __________________________ 4. 3 Tr Clinopyroxene __________________________ 10. 4 ________ Sphene ________________________________ Tr ________ Calcite ________________ . 3 ________ Tr 1. 3 Composition of plagio— clase ________________ Ann Am; An“ A1133 1. Quartz monzonite phase of granodiorite rock unit, dump east side of Mount Pisgah. Mode is average of three sections; total count is 3,000. (ST—8) 2. Quartz diorite phase of granodiorite rock unit, dump of small pit, east slope of Bald Mountain. Rock has strong follation. (ET—219) 3. Pyroxene quartz diorite phase of gabbro rock unit, dump east side of Elk Creek pluton. (C C—1168-A) 4. Quartz diorite phase of gabbro rock unit, same locality as 3. (CC—1168—B) Potassium feldspar from the same quartz monzonite facies of the Mount Pisgah pluton from which biotite was separated for chemical analysis was analyzed with the X-ray diffractometer by E. J. Young, who used the (201) method of Bowen and Tuttle (1950). The method of homogenization and X—ray investigation used in our study was described in the earlier report on the Central City district (Sims and Gable, 1964). The composition of the feldspar before heating, expressed as weight percent KAlSiaog, was 92:1:1 percent; the composition after homogenization, expressed as weight percent KAISlgOg, was 81 :1: 1 percent. The triclinicity index of the feldspar determined by measuring the difference between 260111;“ (130) and 260m“ (130) is 0.75 i002. Maximum microcIine has an index of 0.84 (MacKenzie, 1954). Quantitative spectrographic analyses by J. C. Hamilton of certain elements in the same sample were determined as follows, in percent: calcium, 0.25; barium, 0.98; strontium, 0.13; rubidium, 0.035; iron, 0.047; and lead, 0.012. PEGMATITE The pegmatite characteristically found associated with the granodiorite, and presumably genetically re- lated to it, forms homogeneous bodies which range in width from several feet to several hundred feet and which form ridges and knobs that tend to stand above the surface of the granodiorite. The pegmatite is PETROLOGY AND STRUCTURE, PRECAMZBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. TABLE 20.—Chemical and spectrochemical analyses of biotites from intrusive rocks [laboratory number given in parentheses below sample numbers. Field number is in parentheses after description of sample. Results of chemical analyses given in weight percent and of spectrochemlcal analyses, in parts per million. Chemical analyses by E. L. Munson and P. R. Barnett] Sample 1 2 (H3381) (H3382) Chemical analyses 2 28 3. 99 17 68 19. 31 9 78 9. 09 18 . 00 22 . 23 9 22 8. 26 .04 I 17 2 96 2. 55 .04 .03 59 . 93 99. 87 99. 90 Less 0 __________________________________ . 25 . 39 Total _______________________________ 99. 62 99. 51 Spectrochendealanalysea Ba _____________________________________ 1, 500 630 Go _____________________________________ 30 40 Cr _____________________________________ 100 30 Cu _____________________________________ 41 24 Ga _____________________________________ 50 60 Nb _____________________________________ 90 80 N1 _____________________________________ 80 30 Se _____________________________________ 60 40 Sr ______________________________________ 20 10 V ______________________________________ 280 180 Y ______________________________________ 30 40 Zr _____________________________________ 380 320 1. Qq’artz #0163192; variety of granodiorite rock unit, dump east side of Mount isga . — 2. Biotite-muscovite quartz monzonlte, Lawson area, just west of Central City quadrangle. (ST—11) light gray, locally iron stained, coarse grained and biotitic, but the feldspar laths rarely exceed 2 inches in length. The bodies are composed dominantly of feld- spar and commonly contain about 5 percent each of quartz and biotite. Microcline greatly exceeds plagi- oclase, which is randomly dispersed. The biotite tends to occur in small books scattered throughout the pegmatite; in appearance and composition the pegma- tite is similar to that associated with biotite—muscovite quartz monzonite. GABBRO AND RELATED ROCKS OCCURRENCE AND GENERAL CHARACTER The classification “gabbro and related rocks” is used in this report for rocks that contain an intermediate plagioclase and both orthopyroxene and clinopyroxene E35 and range in composition from melagabbro to quartz diorite. Diorite is the most common rock type within the quadrangle. In composition the rocks overlap mafic phases of the granodiorite and associated rocks and certain phases of the quartz diorite and hornblend- ite unit. They can be distinguished from these units, however, by their megascopic appearance and by their diagnostic content of pyroxenes. The gabbro and related rocks occur in the north- western and central parts of the quadrangle. All known bodies lie northwest of a diagonal line drawn from the southwest to the northeast corner of the quad- rangle. The largest body, referred to in this report as the Elk Creek pluton, is hook shaped and is about 2 miles in total length and three-fourths mile in maximum width; it is interpreted as a complex phacolith. Several smaller bodies less than 1,000 feet in maximum dimension occur near the Elk Creek pluton, particularly to the northwest, and masses a few tens of feet in Width are associated with some of the quartz diorite and horn- blendite bodies that were mapped sporadically north- eastward from the Vicinity of Mount Pisgah (pl. 1). Contacts of the gabbro and related rocks against the older layered rocks are sharp. Transitional phases are absent except at one locality on the south slope of Ari- zona Mountain, near the axis of the Arizona Mountain anticline (pl. 1), where gabbroic rocks intertongue with microcline gneiss across a width of a few feet. Gabbro appears to crosscut a granodiorite dike at one locality in the northwest corner of the quadrangle; accordingly, the unit is interpreted to be younger than the grano- diorite. Contacts between these units are poorly ex- posed elsewhere, however, and this age relation was not confirmed for other bodies. Gabbro is closely associ- ated with and appears gradational into quartz diorite and hornblendite at a few localities. The gabbro and related rocks are dark gray, generally medium or coarse grained, equigranular, and homoge- neous ; locally they are light gray or olive gray and mot- tled, are inequigranular, and contain feldspar crystals as much as 2 inches in diameter. The rocks tend to weather spheroidally to subrounded boulders, to coarse grus, or under forest cover, to a sticky dark-brown soil. Surfaces 0f the weathered rocks are commonly dark olive gray. Typically the gabbro is massive and has interlocking plagioclase and pyroxene crystals. Local parts of the Elk Creek pluton, however, are foliated and lineated. On the southeast nose of Idaho Hill (pl. 1) a few sub- angular blocks or lenses of strongly foliated gabbro, as much as 25 feet in width and 50 feet in length, are enclosed Within massive homogeneous gabbro of the pluton. The blocks occur near the outer margin of the pluton and have sharp contacts against the massive E36 gabbro. In one block about 25 feet from the contact, the foliation within the block is subparallel to the outer contact of the pluton; in another the foliation within the block is virtually parallel to the contact with a large rotated inclusion of microcline gneiss. In contrast to the other intrusive rocks, the country rocks adjacent to and included within the Elk Creek pluton are altered to hornfels through recrystallization and reconstitution and through shearing. The biotite gneisses in the contact zone across a width of a few tens of feet commonly are porphyroblastic and are coarser grained than the unmodified biotite gneisses. Silli— manite-bearing gneiss is changed to inequigranular rocks which contain discoidal sillimanite aggregates. Augen in one specimen consist of sillimanite, a phlogo- pitic mica, and a green spinel, which are surrounded by sericite. Biotite—quartz-plagioclase gneiss is reconsti— tuted, at least in part, to an orthopyroxene-biotite- quartz-plagioclase rock of felted texture. Microcline gneiss in contact with the gabbro of the Elk Creek pluton on Arizona Mountain is altered to a flaser gneiss and is bleached for a distance of nearly 50 feet from the contact; quartz forms lens-shaped, strongly lineated aggegates, and biotite is altered to chlorite. PETROGRAPHY Composition of the gabbro and related rocks varies greatly, from a melagabbro to a quartz diorite, but it is dominantly diorite. In general, the small bodies scat- tered throughout the quadrangle and particularly those masses associated with the quartz diorite and horn- blendite bodies are more mafic than the larger Elk Creek pluton. (See table 21.) The rocks are hypidiomorphic or allotriomorphic granular and contain intermediate plagioclase, ortho— pyroxene, clinopyroxene, hornblende, biotite, quartz, and opaque iron oxides as principal constituents and potassium feldspar, muscovite, apatite, sphene, zircon, chlorite, epidote, allanite, calcite, and clay minerals as minor constituents. Plagioclase and pyroxene com- monly are intergrown in interlocking grains, but at places they occur in aggregates with a synneusis texture. Plagioclase, mainly calcic andesine but locally lab- radorite, is ubiquitous, constituting from about 2 per- cent to more than 70 percent of the rocks. It is dominantly anhedral, slightly clouded with alteration products, and twinned according to pericline and com- bination Carlsbad—albite twin laws. Most crystals appear homogeneous and slightly antiperthitic; a few show normal concentric zoning with a small range in composition from one zone to another. Dustlike in- clusions, commonly oriented, of a dark-gray and red undetermined material are present in most grains. Myrmekitic intergrowths with quartz are fairly com- SI-IORTER CONTRIBUTIONS T0 GENERAL GEOLOGY mon, especially against biotite and potassium feldspar grains. Some plagioclase crystals contain subrounded blebs of quartz. At places the plagioclase has strain shadows; at others the laths are bent or fractured, and twin lamellae are discontinuous or are actually offset. Both orthopyroxene and clinopyroxene occur in most sections, and at least one of them is present in all. The orthopyroxene occurs as dominantly sub- hedral crystals having a conspicuous schiller structure resulting from the alinement parallel to the optic plane of thin lamellae of brown, translucent, unidentified minerals and opaque iron oxides; other lamellae may be clinopyroxene or hornblende. Because augite la- mellae at a small angle to the (101) plane in the ortho- pyroxene have not been noted, it is inferred that the original form of crystallization was orthopyroxene rather than pigeonite. Most grains are altered, and Skeletal crystals have cores nearly completely changed to opaque iron oxides, hornblende, biotite, and quartz or to a fine aggregate of epidote, calcite, quartz, and chlorite; some crystals poikilitically contain blebs of quartz, opaque iron oxides, and hornblende. With few exceptions, the orthopyroxene has rims of, or is mot- tled by, green hornblende. The orthopyroxene is dom- inantly bronzite. It has a large 2V; X=moderate orange pink, Y=colorless, and Z=light greenish gray. In three samples n, ranged from 1.684 to 1.694. The clinopyroxene is a nearly colorless, slightly pleochroic variety of augite, which occurs in subhedral crystals, is generally skeletal, and has inclusions of opaque iron oxides. Its 2V is variable, ranging from about 40° to 60°; c/\Z ranges from about 38° to 46°. In one sample n,=1.684i0.005. With few exceptions the clinopy- roxene is partly replaced by hornblende that lies along grain boundaries or forms a cuneiform intergrowth which is inside the grain and is controlled by cleavage. Hornblende is present throughout all the bodies. It is variable but seems to be dominantly of two types. One type that forms rims on pyroxene with straight, sharp contacts is green and moderately pleochroic. The pleochroism is generally X: grayish yellow, Y=olive green, and Z=dark olive green or dark olive brown. The hornblende tends to be poikilitic and contains blebs of quartz and locally, biotite, opaque iron oxides, and apatite. The other type of hornblende varies from a nearly colorless to a blue-green variety. It typically is anhedral with ragged outlines, embays pyroxene along cleavages, and forms blebs parallel to cleavage. 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Iii-EuacG u .3 « .NA. m .8 m .AN a A o .a A .ms a do N A6 N .3 N .A m .3 A .3 A .3 o .3 a Na N .N@ A .3 o .E c .$ n An m .3 c .3 c .NA. w .2. - . .NmaAqumaAnA o .A a .A .AA. ------------------------ n A m d ------ an. 1-39.38 gammauom nN AN MN NN AN 8 NA NA NA 3 3 «A NA NA AA A: a w A. c a v r. N A A8232 8:5; .850 A833 A920 MAMA $2353 .0 nosatomeu ASE $355.39 5 E .5983. 22m $3555 M36 «0 38.3 053: mA 565 .830 32: @832 ES 923% Me .2823 3:39. E .auASEIAN H455”. 6:58 Go: .- ”cognac con 62 83: FE E38 Red-brown biotite also is ubiquitous but variable in amount. It is closely associated with the hornblende and pyroxene as subhedral laths and as ill-defined rosettes that contain tiny inclusions of apatite, zircon, allanite, and opaque iron oxides. At places it sur- rounds and occurs along the cleavages of pyroxene and hornblende. The pleochroism is as follows: X=very pale orange, Y=light brown, and Z=moderate reddish brown. Potassium feldspar occurs sparsely as interstitial, anhedral grains clouded with alteration products. It is untwinned and slightly perthitic. Quartz is com- mon, locally constituting as much as 20 percent of the rock. It occurs as anhedral interstitial grains, as blebs poikilitically included in the more abundant minerals, and as myrmekitic intergrowths with pla- gioclase and biotite. Allanite forms pleochroic sub- hedral crystals ranging from light brown to dark reddish brown; some crystals are zoned. Sphene is associated with biotite, hornblende, epidote, and opaque iron oxides. Apatite, as small subhedral crystals, is dispersed throughout the rock. Granular masses of epidote ranging from a very pale yellow to a pale yellow green occur with the mafic minerals, especially biotite; some crystals are zoned. Calcite typically occurs as small granular aggregates associated with epidote in grains of plagioclase and pyroxene. Chlorite, apparently of variable composition, locally surrounds hornblende and pyroxene, embays pla- gioclase along fractures and cleavages, and embays and replaces biotite. Muscovite and clay minerals occur as alteration products in intensely altered rocks and especially are associated with plagioclase. CHEMICAL COMPOSITION Chemical analyses of two typical samples from the Elk Creek pluton are given in table 19. The rocks differ from the quartz diorite phase of the granodiorite unit in containing more T102, Fe203, and F60 and less A1203. In particular, a high titania content seems to be diagnostic of the rocks in the Elk Creek pluton. This content is reflected in the notable amounts of contained magnetite-ilmenite. PEGMATITE A distinctive pegmatite is associated with all the larger bodies of gabbroic rocks and occurs in the vicin- ity of many smaller bodies. The pegmatite forms irregular, crudely zoned bodies as much as a few hun- dred feet in maximum diameter. The bodies are coarse grained and consist generally of a milky quartz core and a perthite—biotite or perthite—muscovite border zone. Crystals of quartz and feldspar are as much as 12 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY inches long. The bodies, which contain abundant milky quartz, form conspicuous knobs that stand out promi- nently in the terrain. At a few localities syenitic peg- matite, which consists dominantly of plagioclase and biotite, forms inch-thin discontinuous tabular bodies along joints in the gabbro. QUARTZ DIORITE AND HORNBLENDITE OCCURRENCE AND GENERAL CHARACTER The intrusive rocks that dominantly contain horn- blende and plagioclase in varying amounts were mapped as quartz diorite and hornblendite. They occur in all parts of the quadrangle except the northwestern part within major units of both biotite gneiss and microcline gneiss. Most of the mappable bodies are alined at ir- regular intervals along a narrow zone which trends N. 45° E. from the southwest corner to the northeast cor- ner of the quadrangle (pl. 1). Several bodies in the Central City district, which were mapped at a scale of 1 :6,000 (Sims and Gable, 1964), were too small to show at the scale of plate 1 of this report. The bodies are generally blunt lenses a few hundred feet in maximum dimension or are, less commonly, tab- ular bodies as much as 4,000 feet long and 500 feet wide. Contacts are poorly exposed, but observations at various points of the contact zones and the map patterns of separate bodies indicate that the bodies are subconformable to the older rocks. The age of the in- trusive bodies relative to the ages of other intrusive rocks in the quadrangle has not been determined une- quivocally. However, rocks of nearly identical lithol- ogy, texture, and structure in areas a few miles to the south (Harrison and Wells, 1959, p. 15) clearly intrude granodiorite and in turn are intruded by biotite-musco- vite granite, the equivalent of the biotite-muscovite quartz monzonite of this report. Within the Central City quadrangle, several of the bodies of quartz dio- rite and hornblendite, especially in the northeastward- trending zone of intrusions that extend northeast from the vicinity of Mount Pisgah, are associated with small masses of gabbro. Small irregular masses of gabbro of a composition represented by samples 20, 21, and 22 (table 21) occur within and are intimately associated with quartz diorite and hornblendite in the bulbous mass on the east side of Mount Pisgah (pl. 1). Else- where, recognizable bodies of gabbro associated with quartz diorite and hornblendite bodies are less common. The field relations are best interpreted as indicating that the bodies of quartz diorite and hornblendite grade into gabbro. The hornblendite is megascopically similar to that described from the Central City district (Sims and Gable, 1964). It is a black or greenish-black, medium- PETROLOGY AND STRUCTURE, PRECAMBRIAN grained or rarely coarse grained, generally massive, equigranular rock. The quartz diorite is a dark-gray, fine- to medium— grained, generally massive and equigranular rock. Lo- cally it is inequigranular and contains clots of biotite as large as half an inch in diameter. Some of the smaller tabular bodies have border phases that are finer grained than the interiors. The quartz diorite lacks the strong foliation that characterizes the amphibolites of the area and thus is readily distinguishable from these rocks. PETROGRAPHY The quartz diorite and hornblendite of the quadrangle are similar mineralogically to that unit described pre- viously by Sims and Gable (1964) from the Central City district but differ in that they contain substantially more pyroxene, which is partly altered to hornblende. Because of this difference and its significance in inter- preting the origin of these rocks, some aspects of the petrography are discussed as follows. The rocks contain dominant hornblende and plagi- oclase of intermediate composition, variable amounts of biotite and pyroxene, and lesser amounts of quartz, opaque iron oxides, and a variety of accessory minerals (table 22). Plagioclase (calcic andesine) forms sub- hedral and anhedral grains dominantly intergrown with pyroxene or, where pyroxene is absent, with hornblende or biotite. It is twinned according to the Carlsbad, albite, and pericline twin laws and shows weak concen- tric zoning. Both clinopyroxene and orthopyroxene occur sparsely in the rock. The orthopyroxene forms subhedral crystals that have a conspicuous schiller structure resulting from thin plates of an unidentified brown translucent material oriented parallel to the optic plane. Other lamellae of the same alinement may be clinopyroxene or a nearly colorless amphibole. With few exceptions, the orthopyroxene is embayed and rim- ’ med by green hornblende, and contact between the minerals is sharp and ragged. Blebs of opaque iron oxides oriented in a crystallographic or a cleavage di- rection occur in the pyroxene that is partly altered to hornblende. The clinopyroxene occurs as irregular, shredded crystals that are partly altered to hornblende. The hornblende forms irregular patches and blades and wedges oriented in the plane of the cleavage and com- monly is optically continuous with the clinopyroxene. Alined grains of magnetite occur in the altered clinopyroxene. The hornblende is a green, strongly pleochroic variety having the following general pleochroic formula: X=moderate greenish yellow, Y=dark greenish yellow, and Z=moderate olive brown. ROCKS, CENTRAL CITY QUADRANGLE, COLO. E39 TABLE 22.—Modes, in volume percent, of quartz diorite and hornblendite [Tr, trace; Nd, not determined; __-_, not found. Color index is volume percent of dark minerals. Field number is in parentheses after description of sample] 1 2 3 4 5 6 7 Potassium feldspar ..... Tr ................ 3. 5 ________________________ Plagloclase ............. 51. 8 43. 5 21. 8 39. 5 52. 7 33. 6 0. 2 Quartz _________________ 3. 4 1. l 9. 5 9. 5 2. 8 5. 3 . 5 Biotite _________________ 16. 4 ________ 21. 4 22. 5 8. 1 17. 1 ........ Musoovite. . . 1 _ _ Opaque iron oxides_-__ 4. 9 3. 6 1. 1 4. 6 6. 8 4. 2 Tr Amphibol _____________ 15. 5 47. 8 43.1 16. 1 3. 3 37. 9 99. 2 Orthopyroxene ......... 2. 3 ________ l. 1 ........ 14. 3 1. 4 ........ Clinepyroxene.. _ 5. 0 . 4 1. 6 . 5 7. 4 _ Apatite _________ . 3 Tr . 4 2. 0 1. 3 . 2 . 1 Sphene- - - . _______________ Tr ________ . ________________ Tr Zircon ____________ Tr Tr Tr Tr Chlorite. _. ....... 2. 7 _- 1.0 . 3 ________________ E idote ................ . 1 . 8 ________ Tr 3. 0 3 ........ A lanitp .3 ________________ - (‘nlr‘ifp Tr ________________________ Total ______________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 Composition of plagia- clase ................. Ana Nd Am: Nd Ami Ame Nd Color index ____________ 45 55 68 55 43 61 99 Average grain diame- ter (mm) ............. Nd Nd 0. 5 0.3 0. 1 0. 2 Nd 1. Diodte, from lens-shaped body that cuts across Missouri Gulch near Stewart Gulch. (CC—1209) 2. Fine-grained phase ofsample 1. (CC—12094) 3, 4, 5, 6. Succession of sam les taken across strike of body exposed on south side of North Clear reek, just east of mouth of Peeks Gulch. Sample 3, diorite that contains biotite clots; sample 4, similar to 3 but finer grained; sample 5, phase finer grained than 4; sample 6, fine-grained hase that occurs as ov01d blocks in normal phase. (CC—362—A; C—362—B; CC—362-C; CG—362—D) 7. Homblendite taken from small unmapped body on north slope of Bald Mountain. (EWT-54—54) In addition to occurring as rims on pyroxene, it forms anhedral irregular grains, generally slightly larger than those of the pyroxene, that poikilitically include quartz, opaque iron oxides, and rarely apatite. Red-brown biotite is distributed irregularly. It has the following pleochroic formula: X=grayish orange or very pale orange, Y=dark yellowish orange or dark reddish orange, and Z=moderate reddish brown. It forms moderately large subhedral, generally poi- kilitic crystals. In general, a substantial proportion of quartz is associated with the biotite. The quartz is mainly in the interstices of the dominant minerals. Some sections which contain large amounts of biotite also contain myrmekitic intergrowths of quartz and potassium feldspar. PEGMATITE The pegmatite that is associated with bodies of quartz diorite and hornblendite is similar to that associated with the gabbro. It is coarse grained and forms crudely zoned bodies. White quartz cores typically grade laterally toward the walls into perthite-quartz-mica pegmatite. Sillimanite occurs locally in the pegmatite in areas where sillimanitic biotite gneisses occur. BIOTITE-MUSCOVITE QUARTZ MONZONITE Scattered small bodies of biotite-muscovite quartz monzonite, generally either in dike form or in crescentic E40 masses thickened in apical areas and thinned or pinched on the limbs of folds, occur in the southwestern and central parts of the quadrangle. The bodies are peripheral to the large pluton of Silver Plume Granite exposed at and near Silver Plume, Colo. (Levering and Goddard, 1950, pl. 3) and also to outlying smaller plutons such as at Alps Mountain in the Freeland- Lamartine and Chicago Creek areas (Harrison and Wells, 1956, 1959). Nearly all the bodies exposed in the quadrangle lie within an ill-defined, relatively narrow, northeastward- trending zone that extends from the southwest corner of the quadrangle to the vicinity of Bald Mountain. The bodies are more numerous and larger in the south- western part of the quadrangle than in the central part. They range in size from three-quarters of a mile to a few tens of feet in length and 1,000 feet to a few feet in width. They occur in biotite gneisses and to a lesser extent in microcline gneiss and granodiorite and asso— ciated rocks. Contacts of the quartz monzonite bodies with the country rock are sharp; however, several inches or feet of pegmatite may occur along the contacts. The quartz monzonite is nearly massive except near the borders where it has weak foliation imparted by oriented biotite flakes and tabular feldspar crystals. The foliation is subparallel to the contacts, even where the contacts are discordant. Accordingly the foliation is interpreted as a primary flow structure. The biotite-muscovite quartz monzonite is a light- SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY or medium-gray rock that weathers to yellowish gray, pinkish gray, or buff. In this area it is dominantly fine or medium grained and virtually equigranular, but parts of the larger bodies are medium grained and se- riate porphyritic in texture. The rock is closely sim- ilar in composition, fabric, and texture to the fine- grained varieties of the type area and of the nearby Freeland-Lamartine and Chicago Creek areas, but it has a slightly lower color index. The rock contains roughly equal amounts of potas— sium feldspar, plagioclase, and quartz as essential min— erals and biotite and muscovite as diagnostic varietal minerals (table 23). The proportions of the essential minerals vary somewhat, but the rock is dominantly quartz monzonite, which has a hypidiomorphic gran- ular texture. As the petrography of the rock has been described previously from the Central City district (Sims and Gable, 1964) and the Lawson-Dumont-Fall River district (Hawley and Moore, 1967), only a brief description is given here. Typically, potassium feld- spar, quartz and plagioclase form crystals that are 3—4 mm in diameter, but some tabular potassium feldspar crystals are as much as 6 mm in width. Interstitial to these grains are smaller crystals of biotite, mus- covite, magnetite, and some of quartz and plagioclase. As a generalization, the quartz grains tend to be clus- tered at random. The plagioclase is typically cloudy and is partly altered to muscovite, sericite, and clay minerals. The potassium feldspar, however, tends to TABLE 23.——M odes, in volume percent, of biotite-muscovite quartz monzonite [Tr, trace; Nd, not determined; ........ , not found] Lawson- Central Quarry, Chicago Scattered small bodies in Mount Pisgah and Bald Mountain areas Dumont- City Silver Creek Fall River district 3 Plume 3 area 4 district 1 Field No ......................... S—lZA-52 CC-296—2 CC-300A EWT-50-54 EWT—81-54 EWT—95—54 Average ________________________ SPQ—l ____________ Potassium feldspar ...... 27. 2 32. 0 39. 4 35. 6 41. 0 21. 5 32. 8 30. 0 25. 8 31. 7 34. 4 Plagioclase _____________ 36. 0 29. 0 27. 3 28. 1 27. 7 39. 0 31. 3 32. 5 34.0 29. 8 27. 8 Quartz _________________ 24. 9 26. 3 23. 4 26. 0 20. 9 26. 5 24. 6 29. 0 30. 1 24. 0 29. 6 Biotite _________________ 9. 4 1.0 4.0 6. 3 4. 3 5.1 5.0 5 0 4. 9 9. 2 3. 4 Magnetite-ilmenite ...... 1. 2 1. 5 1. 6 1. 3 . 2 .4 1. 0 1.0 . 7 1. 4 . 6 Muscovite _____________ 1. 3 3. 3 1. 6 2. 5 5. 4 7. 2 3. 6 2. 5 4. 5 3. 1 3. 2 Apatite ________________ Tr . 2 . 3 . 2 . 3 Tr . 2 Tr Tr . 6 Tr , Zircon and monazite__ _ _ Tr Tr Tr Tr Tr . 3 Tr Tr Tr . 2 . 4 Fluorite ______________________________________________________________________________________ Tr ________ Allanite ________________________ . 4 ........ Tr _______________ Tr Tr ________________________ Epidote ________________________ 1. O 7 _______________________ . 3 ________________________________ Calcite _______________________________________________________________________________________ Tr ________ Chlorite _______________________ 5. 3 1. 7 Tr . 2 ________ 1. 2 ________________________________ Rutile _________________________________ Tr _______________________________________________ Tr ________ Total ______________ 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100. 0 100.0 100.0 99. 4 Composition of plagioclase ___________ A1130 Nd A1129 Ann A1125 Nd An” Nd A1125 A1123 Nd Average grain diameter mm Nd 0. 6 0. 6 0. 6 0. 4 O. 3 O. 5 Nd Nd 0. 6 Nd 1 Average of 12 modes (C. C.Haw1ey, written commun., 1961). 3 Average 015 modes (Sims and Gable, 1964). 3 Type section of Silver Plume granite 4 Average of 29 modes (Harrison and Wells, 1959, p. 19). PETROLOGY AND STRUCTURE, PRECAMBRIAN be clear and little altered. The biotite is dominantly the brown variety; the chemical composition of a typ— ical sample is given in table 20. Associated with the biotite are traces of allanite, zircon, and monazite, which all produce pleochroic halos. Muscovite gen- erally appears shredded and occurs as overgrowths on feldspars and biotite and locally as aggregates and stringers between the larger quartz and feldspar grains. As zircon and monazite are virtually identical in ap- pearance in thin section, they are grouped together in the table of modes. Monazite is sufficiently abun— dant to produce anomalously high radioactivity. Potassium feldspar from the same rock (ST—11) from which biotite was separated for chemical analysis (table 20) was analyzed with the X—ray diffractometer by E. J. Young, who used the (§OI) method of Bowen and Tuttle (1950). The composition of feldspar before it was heated, expressed as weight percent KAlSi308 was 91.5i1 percent; the composition after it was heated, expressed as weight percent KAlSi308 was 82:1;1 percent. The triclinicity index of the feldspar determined by measurement of the diflerence between 29m... (150) and 29cm. (130) was 0.84:};002; this index is equivalent to maximum microcline as defined by MacKenzie (1954). Quantitative spec- trographic analyses by J. C. Hamilton of certain elements were determined on the same material and were, in percent—calcium, 0.15; barium, 0.37; stron- tium, 0.049; rubidium, 0.046; iron, 0.075; and lead, 0.017. Flame photometer analyses by J. B. McHugh of representative samples of biotite-muscovite quartz monzonite gave the following compositions, in percent: Sample N820 K20 CC—296—2 ______________________________ 2. 8 5. 5 2. 7 5. 9 3. 2 6. 0 3. 1 6. 4 3. 1 5. 6 All samples except SPQ—l were collected in the Central City quadrangle; SPQ—l was from the Silver Plume Granite quarry in Silver Plume, Colo. Pegmatite that is associated with biotite-muscovite quartz monzonite is unzoned and consists dominantly of perthitic microcline, sodic plagioelase, quartz, and biotite. It is characterized by the presence of biotite as irregularly shaped books that have little or no preferred orientation and that tend to be clustered. Uraninite-bearing pegmatites within a northeastward- trending belt that extends across Virginia Canyon in the southeastern part of the quadrangle possibly are related to the granite; they have been described in some detail in an earlier report (Sims, Armstrong, and others, 1963, p. 10—12). Flame photometer analyses by J. B. McHugh of two samples from the Highlander ROCKS, CENTRAL CITY QUADRANGLE, COLO. E4]. claim in Virginia Canyon gave the following composi- tions, in percent: Sample N320 K20 H—l ____________________________________ 2. 8 11. 8 A—13-9 _________________________________ 3. 0 8. 0 EMPLACEMENT AND ORIGIN OF THE INTRUSIVE ROCKS The Precambrian intrusive rocks in the Central City quadrangle part of the Front Range were emplaced synchronously with the major period of folding and metamorphism in the catazone of the earth’s crust. The texture, structure, and lithology of the rocks are interpreted to indicate that the intrusives crystallized largely from magma. The sequence of emplacement of the major intrusive rocks in the central part of the Front Range has been established with certainty by the previous detailed geologic studies of Harrison and Wells (1959), Sims and Gable (1964), and Moench, Harrison, and Sims (1962). The order from oldest to youngest, as determined by crosscutting relations and by internal structures of the intrusive bodies, is granodiorite and associated rocks, quartz diorite and hornblendite, and biotite-muscovite quartz monzonite. Our study has established that the gabbro and related rocks, which were not noted in the adjacent area mapped previously by Harrison and Wells (1959), are intermediate in the succession—between the granodiorite and associated rocks and the biotite- muscovite quartz monzonite. It has also established that the quartz diorite and hornblendite bodies which are so profusely scattered through the adj acent terranes probably were formed by the retrograde metamorphism of the gabbro and related rocks. The oldest intrusive bodies, granodiorite and asso- ciated rocks, were intruded early in the episode of plastic deformation. They were emplaced subsequent to the development of most of the migmatite, for inclusions of migmatized biotite gneiss are found scattered through the granodiorite bodies, and were emplaced after folding had begun, for inclusions commonly are folded more intricately than the enclosing granodiorite. The subconcordant contacts and local phacolithic forms are interpreted to indicate that the rock formed from a magma or a fluid that welled up between and along foliation planes and that followed preexisting structures. Magma moved into relatively low pressure sites in the crests of anticlines and the troughs of synclines as folding progressed, and phaco- liths resulted. Except locally, the magma was not able to penetrate for long distances along the limbs of the folds, and as a result the limbs of phacoliths are short and the sheets generally are lenticular. Crystalli- zation took place under directed compressive stresses, and the result was a foliation and lineation subcon- E42 cordant to that in the country rock. Continued stress after consolidation caused recrystallization in the borders of the bodies, and a secondary foliation and lineation resulted. A concept of magmatic origin for all or most of the rock is supported by the crosscutting contacts on both a small scale and a regional scale, by sharp contacts, by a composition consistent with that expectable from a differentiating magma, and by a probable flow structure marked locally by tabular inclusions oriented parallel to crosscutting contacts (Harrison and Wells, 1959, p. 13—15). Gabbro and related rocks are inferred to be younger than the granodiorite because of the apparent cross- cutting relations in the northwest corner of the quad- rangle (pl. 1) on the northwest slope of the Montana Mountain. This interpretation of age is supported by marked differences in the internal structures of the bodies of the two rock types; the gabbroic rocks have a more massive structure than the granodiorite, and thus they probably crystallized later in the epi- sode of deformation under less intense deforming stresses. The Elk Creek pluton, the principal body of gabbroic rocks, is interpreted to be a compound phacolith. In some aspects, it resembles the “pine tree” structures described by Boos and Boos (1934) from the granitic rocks in the Longs Peak-St. Vrain batholith in the northern part of the Front Range. Smaller bodies are stubby sheets or lenses, wholly conformable to the country rock. Presumably the mechanism of emplacement was similar to that of the granodiorite bodies—the magma was injected under directed compressive stresses but except locally was not recrystallized sufficiently to produce a secondary foliation and lineation. The blocks of foliated gabbro observed near the east margin of the Elk Creek pluton suggest that an early consolidated phase of the magma, which crystallized under directed compressive stresses, was broken and engulfed within still-fluid magma. At least one of these xenoliths was oriented, presum- ably by flow of the magma, parallel to the pluton walls. The gabbroic rocks are interpreted to have formed from a normal gabbroic magma. A temper- ature difference between the magma and the country rocks, which may be estimated at possibly as much as 500°C., was sufficient to produce a narrow halo of metamorphism in the country rock and, possibly, a chilling of the margins of the bodies. After crystal- lization, green hornblende formed at the expense of pyroxene and plagioclase. Later biotite and probably other minerals replaced the older minerals. We are not able, however, with our present knowledge to distinguish confidently all phases of igneous or deuteric and metamorphic stages of crystallization. Bodies of quartz diorite and hornblendite are widely SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY distributed in the region, and their equivalence has been established by detailed studies in several areas. All the bodies are small and either podlike or lenslike. That some have border phases that are finer grained than the interiors indicates probable chilling during intrusion. The bodies intrude granodiorite and asso- ciated rocks and in turn are intruded by biotite- muscovite quartz monzonite. An intermediate posi- tion in the intrusive sequence, therefore, is firmly established for the rock unit. The regional variations in the mineralogy and struc- ture of the quartz diorite and hornblendite rock unit and its close association with and gradation into gabbroic rocks in the central and northwestern parts of the Central City quadrangle are interpreted to indicate that the unit was derived from gabbro through retro- grade metamorphism. From localities south of the Central City quadrangle and extending northward into the quadrangle there is a systematic change in the character of the quartz diorite and hornblendite bodies. In the area at and south of Idaho Springs, the bodies are foliated and lineated and contain hornblende and biotite as the major mafic minerals. The strongly foliated masses contain as much as 20 percent micro— cline and plagioclase in the range Ann—Ann; the less foliated ones have virtually no microcline and contain calcic andesine or labradorite (Harrison and Wells, 1959, p. 16). In the Central City district, on the east side of the Central City quadrangle, bodies of the same rock are weakly foliated and contain a few percent of clinopyroxene as well as the principal constituents hornblende and andesine (Sims and Gable, 1964). Similar bodies in the central part of the quadrangle, and particularly those extending northeastward from Mount Pisgah, are weakly foliated or nearly massive and con— tain some orthopyroxene as well as clinopyroxene (table 22). The hornblende appears to have formed largely at the expense of pyroxene; biotite, too, is a secondary mineral. The hornblende is similar in color and optical properties to that formed after pyroxene in the gabbroic rocks. These mineralogic data, together with the ap— parent gradation in the field between gabbroic rocks and quartz diorite and hornblendite, are consistent with a View that the quartz diorite and hornblendite formed in some manner from the gabbroic rocks. Evidently the smaller bodies of gabbro were more susceptible to change than were the larger bodies, and some were so completely modified that all remnants of the original texture and mineralogy were destroyed. The regional pattern of distribution of gabbro and quartz diorite and hornblendite is consistent with an interpretation that original gabbroic rocks were modified by the deforming stresses, the younger episode of Precambrian defor- mation (see p. E49) that produced the Idaho Springs— PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. Ralston shear zone. This deformation was intense in the shear zone and weak elsewhere. Probably all bodies of original gabbro within the shear zone would be modified by retrograde metamorphism because the pressure—temperature environment was one of cat- aclasis with little recrystallization. Outside the zone, modification by the new metamorphic environment was less intense, and some of the primary features of the gabbro were not destroyed. The same episode of meta- morphism can account for the mineral changes in the Elk Creek pluton of gabbro and related rocks, as discussed previously. The biotite-muscovite quartz monzonite, the young- est of the Precambrian intrusive rocks, was emplaced at about the same time as or after the cessation of the deforming stresses that produced the major folds of the region. Evidence for a late-syntectonic origin is pro- vided by occurrence of the quartz monzonite as phaco- liths in small folds, previously documented for the Chicago Creek area (Harrison and Wells, 1959, p. 20), a few miles south of the quadrangle. Within the quad- rangle, the rock occurs as small subconcordant bodies or more commonly as crosscutting dikes. Some of the subconcordant bodies along Clear Creek are possibly phacolithic. The mineralogic homogeneity, crosscutting contacts, and primary foliation of the rocks support the theory for a wholly magmatic origin. The Precambrian intrusive rocks were emplaced in the catazone, as defined by Buddington (1959), in an environment of intense pressure-temperature con- ditions. The thickness of cover can be presumed, from the estimates of Buddington for the catazone, to be in the range of from 7 to 12 miles. The earliest intrusive rocks—granodiorite and associated rocks—crystal- lized after or near the thermal maximum, which was equivalent to the upper sillimanite grade of metamor- phism of the older, major Precambrian deformation. By this time regional migmatization of the country rock had been largely accomplished. Later in the defor- mation the gabbroic rocks were emplaced under similar pressure-temperature conditions. Temperatures in the Elk Creek pluton of gabbro, the largest intrusive body in the region, were sufficiently greater than the country rock to form a narrow metamorphic halo. Evidently a substantial interval of time preceded emplacement of the biotite-muscovite quartz monzonite. The de- forming stress had lessened and probably a substantial thickness of cover had been eroded. Although mild compressive stresses were still active, emplacement was controlled to a considerable degree by cross fractures. The physical conditions of emplacement insome re- spects were transitional to the mesozone environment. Metamorphism under decreasing pressure—temperature conditions may have begun soon after emplacement of E43, the biotite-muscovite quartz monzonite but more likely took place concomitantly with the younger Precambrian deformation. METAMORPHIC FACIES Progressive metamorphism of the rocks of the region produced mineral assemblages characteristic of the silli- manite grade. In rocks of suitable composition, potas- sium feldspar, sillimanite, and muscovite are stable; this stability indicates that the rocks are above the silliman- ite—potassium feldspar isograd, as defined by Guidotti (1963). Judged from the mineral assemblages, regional meta- morphism was virtually uniform in the quadrangle. The assemblages can be related to the metamorphism that accompanied the episode of dominant regional de- formation; the younger Precambrian deformation ap- parently did not appreciably modify the assemblages, except locally in the southeastern part of the quad- rangle, within the Idaho Springs-Ralston shear zone. The younger episode of deformation mainly resulted in cataclasis and the formation of new structures. The assemblages of the regionally metamorphosed Precambrian rocks adjacent to the Tertiary stock near Apex were modified in a manner similar to that de- scribed recently by Hart (1964) from the adjacent Nederland quadrangle. The mafic metamorphic rocks of the quadrangle, re- presented by amphibolite, dominantly contain the as- semblage andesine-hornblende-quartz; clinopyroxene is a local constituent. These minerals crystallized near equilibrium during the highest grade of metamorphism in the area. Some of the hornblende and biotite may have resulted, however, from a later retrograde meta- morphism. Apparently the rocks formed under con- ditions of progressive metamorphism that approximate the transition in which sphene disappears and clinopy- roxene appears, as represented by the amphibolites in the zones of progressive metamorphism in the northwest Adirondack Mountains, New York (Buddington, 1963, p. 1163; Engel and Engel, 1962, p. 69). The horn— blende gneiss in the Lawson layer of microcline gneiss, in the northeastern part of the quadrangle, represents rocks which had an original bulk composition that was different from that of the normal amphibolite. Microcline gneiss contains the common assemblages biotite-plagioclase—potassium feldspar-quartz and bio- tite-muscovitaplagioclase—potassium feldspar-quartz. Garnet, sillimanite, and hornblende are local constitu- ents. None of the assemblages are diagnostic of a particular grade of metamorphism. Metamorphism of the pelitic rocks produced several mineral assemblages, the composition of which depends upon the original bulk chemical composition. The E44 principal assemblages of the dominant biotite gneisses—— biotite-quartz-plagioclase gneiss and sillimanitic biotite gneiss—are as follows: Biotite-plagioclase—quartz Biotite-garnet-plagioclase-quartz Biotite-quartz-sillimanite Biotite-plagioclase—quartz—sillimanite Biotite—plagioclase—potassium feldspar-quartz-silli— manite 6. Biotite—muscovite-plagioclase—potassium quartz-sillimanite Muscovite formed in those rocks that had a sufficiently high K : Na ratio. The plagioclase is consistently An23_29; the average composition is about An”. Rarely muscovite apparently formed also as a product from the reaction between sillimanite and potassium feldspar. QPP‘DF’E‘ feldspar- Metamorphism of pelitic sediments rich in A1203+ FeO+MgO produced the following common assem- blages: 1. Biotite—garnet-magnetite-plagioclase—potassium feld- spar—quartz-sillimanite 2. Biotite—cordierite-garnet—magnetite—plagioclase—potas— sium feldspar-quartz-sillimanite 3. Biotite—cordierite-rnagnetite-plagioclase—quartz Assemblage 1, characteristic of the gneisses listed in table 11, locally lacks potassium feldspar, magnetite ilmenite, and sillimanite. Assemblage 2 may lack or show only traces of magnetite, potassium feldspar, and plagioclase (table 12). Cordierite commonly forms coronas around garnet, as indicated on page E23. The potassium feldspar is dominantly microcline and microperthite but locally is orthoclase. Andalusite— biotite-quartz and biotite—cordierite-magnetite-silliman- ite—spinel form local subassemblages. Assemblage 3, characteristic of the rocks listed in table 13, may con— tain garnet or sillimanite or both. It apparently is more restricted in occurrence than the other two major assemblages. The magnesian cordierite-amphibole gneiss consists of the following several stable mineral assemblages, the composition of which depends upon the bulk chemical compositions : 1. Biotite—cordierite—gedrite 2. Biotite-cordierite—gedrite—quartz 3. Biotite-cordierite—garnet-gedrite—quartz 4. Cordierite—plagioclase-quartz 5. Biotite—cordierite—quartz Except for the occurrence of quartz and spinel in the same rock, evidences of disequilibrium are lacking. The calc-silicate rocks contain a variety of assem- blages. (See p. E17.) They contain epidote as an apparent stable mineral phase and thus difl’er from the common calcareous assemblages in the sillimanite- SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY almandine subfacies (Fyfe, Turner, and Verhoogen, 1958, p. 231). In a narrow zone surrounding the Elk Creek pluton the assemblages of the biotite gneisses were changed to the pyroxene hornfels facies. The biotite gneisses were reconstituted and recrystallized to a coarser grained hornfels. Sillimanitic biotite gneisses were changed to inequigranular rocks which contain discoidal aggregates of sillimanite, phlogopitic mica, quartz, plagioclase, and spinel; biotite-quartz—plagioclase gneiss was converted to orthopyroxene-biotite-quartz—plagioclase rocks of felted texture. A small amount of clinopyroxene formed with the orthopyroxene. The thermal metamorphism related to the larger Tertiary intrusive bodies in this area cannot be entirely separated from the hydrothermal effects; further studies of the contact metamorphic aureoles are needed. STRUCTURE The main structures of the Precambrian rocks of the Central City quadrangle are interpreted to have resulted from two episodes of Precambrian deformation. Northeastward—trending folds and associated linea- tions—the dominant structures of the area—formed during an older plastic deformation; and shearing and related folds and lineations trending east-northeast formed during a distinctly younger deformation. The older deformation was pervasive and virtually con- temporaneous with emplacement of most of the Pre- cambrian intrusive bodies, whereas the younger one was more local in extent, and megascopically visible effects of it are mainly confined to the southeastern part of the quadrangle. The principal manifestation of the young- er deformation is the Idaho Springs Ralston shear zone (Tweto and Sims, 1963). This and other structures related to the younger deformation have been described in considerable detail previously (Moench and others, 1962) and are only discussed briefly in this report. A still older deformation that has been recognized in similar rocks to the east of Central City (p. E5) has not been distinguished in the Central City quadrangle. In accord with previous reports on the region (Moench, Harrison, and Sims, 1962; Sims and Gable, 1964), we refer the linear elements associated With the fold systems to directional coordinates. In this report, B refers to the major fold axes and to linear elements parallel to them, and A refers to the minor fold axes and associated linear elements that are nearly at right angles to the axes of the major folds. A and B are, therefore, used in a geometric sense. However, B conforms to established petrofabric terminology in which 1) refers to the axis of internal rotation, which commonly is parallel to fold axes and normal to a, the PETROLOGY AND STRUCTURE, PRECAMBRIAN direction of tectonic transport (Fairbairn, 1949, p. 6). A has the required geometric relations of an a fabric direction to b but is also a direction of folding. In the commonly established petrofabric terminology, there- fore, it is a b fabric direction and might be termed accordingly I)1 or 62. TERMINOLO GY The Precambrian rocks have a well-defined foliation and lineation that is analagous to the “gneissic struc- ture” of high-grade metamorphic rocks throughout the world. As these structures have been described pre- viously for rocks in this region (Sims and Gable, 1964; Moench, 1964), only brief discussions of the terms are given herein. Foliation—Foliation is the term used for a preferred planar mineral orientation as well as for a composi- tional layering. In the metamorphic rocks the min- eral orientation is subparallel to the compositional layering, except for a few scattered outcrops where a planar mineral orientation is parallel to axial planes of small folds and is at a large angle to the lithologic layering. Except for the exceptions noted above, the foliation is secondary and formed prior to the culmin- ation of folding. In the intrusive rocks the foliation is dominantly a planar mineral orientation and is in part secondary and in part primary. Foliation in the granodiorite and associated rocks, and to a lesser de- gree in the gabbro and related rocks and in the quartz diorite and associated hornblendite, is largely second- ary, for it conforms to and is continuous with the foli- ation in the layered gneissic country rocks. It formed during the regional plastic deformation, subsequent to crystallization. However, as these rocks were em- placed during the deformation, an original primary foliation probably was also induced, and this foliation probably was modified by postconsolidation recrystal- lization. The foliation in the biotite-muscovite quartz monzonite is wholly primary and formed as a flow structure, for it is parallel to the walls of the intrusive bodies even where they crosscut the gneissic structure of the country rock. In the sheared rocks of the Idaho Springs-Ralston shear zone, foliation at places is due to a subparallel mesh of closely spaced fractures, with or without a planar mineral orientation. Lineation.—Lineation is defined (Cloos, 1946, p. 1) as “a descriptive and nongenetic term for any kind of linear structure within or on a rock.” In this area lineation is expressed by the axes of small folds, elon- gate minerals and mineral aggregates, boudinage, and rarely slickenside striae and rodding. With a few ex- ceptions the observed lineations were formed by sec- ondary flowage that accompanied the older, plastic deformation of the region. Consequently the linea- ROCKS, CENTRAL CITY QUADRANGLE, COLO. E45 tions are parallel to the major fold axes (B) or are nearly normal (A) to them. Slickenside striae and rodding and rarely other linear elements are related, however, mainly to the cataclastic deformation that pro- duced the Idaho Springs-Ralston shear zone. These lineations are parallel to folds of the younger deforma- tion or are nearly normal to them. To summarize the lineation data for the quadrangle, lineations measured in surface exposures were plotted on the lower hemisphere of Schmidt equal-area nets. Four diagrams were constructed according to the meth- od described by Billings (1942, p. 119—121) for the poles of joints, except that the lower hemisphere was used. Figure 8A represents measurements made in the west- ern and northwestern parts of the quadrangle, in an area roughly bounded on the east by a straight line passing through Oregon Hill and Mount Pisgah. Most measurements in this area were made north of Fall River; lineation data on the area south of Fall River are given in the report on the Lawson-Dumont—Fall River mining district by Hawley and Moore (1967). Fig— ure 8B represents measurements in the northeastern part of the quadrangle, in the area north of North Clear Creek and east of Blackhawk Peak. Figure 80 covers the east-central part of the quadrangle; it includes meas- urements from Bald Mountain and vicinity northward to North Clear Creek, west of the Central City district. Figure 8D represents measurements made from the southern and central parts of the Central City mining district, in an area centering at Quartz Hill. Many additional data are given for this area in the report on the Central City district by Sims and Gable (1964). Lineation data for the extreme southeastern part of the quadrangle are not presented herein as these are given in the report on the Idaho Springs district by Moench (1964) and in summation, in the report by Moench, Harrison, and Sims (1962, p. 43). FOLDS The northeastward—trending folds are the most con- spicuous structural feature of the region and provide the structural framework of the quadrangle (pl. 1). They are dominantly open, upright anticlines and syn- clines which have steeply dipping axial planes and gently plunging axes. Closed, overturned folds of in- termediate size occur in the west-central part of the quadrangle in the vicinity of Hamlin Gulch. According to our interpretation of the lithologic succession, the dominant folds of the quadrangle are the Central City anticline in the east and the Lawson syncline in the west. The Idaho Springs anticline in the extreme southeastern part of the quadrangle is similarly a major structure, but it does not affect the distribution of the rocks Within the quadrangle to the same extent as do E46 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY A. WESTERN AND NORTHWESTERN PARTS; 641 LlNEATIONS B. NORTHEASTERN PART; 416 LlNEATIONS e C. EAST-CENTRAL PART; 479 LlNEATIONS D. QUARTZ HILL AREA; 1505 LlNEATIONS FIGURE 8.—Contour diagrams of lineations in the Central City quadrangle. Lower hemisphere plots, contoured in percent. PETROLOGY AND STRUCTURE, PRECAMBRIAN the other two folds. In general the strata dip westward from the axis of the Central City anticline to the axis of the Lawson syncline; a structural rise of more than 2 miles occurs from the Lawson syncline southeastward across strike to the southeast corner of the quadrangle. MAJOR FOLDS CENTRAL CITY ANTICLINE The dominant fold in the eastern part of the quad- rangle, the Central City anticline, is a broad, open, nearly symmetrical, upright fold outlined by the Quartz Hill layer of microcline gneiss and the overlying biotite gneiss unit (pl. 1). It has been traced along its axis for a distance of about 10 miles and extends 5 miles northeastward beyond the east border of the Central City quadrangle. Near Central City the fold has a breadth in excess of 3 miles; southwestward it decreases in both breadth and amplitude and near Clear Creek becomes a gently dipping, westward—facing monocline. The axial plane is interpreted from both surface and subsurface data in the Central City mining district to-dip about 85° SE. (Sims and Gable, 1964). The trace of the axial plane trends N. 40° E. on the average but is sinuous in detail. The deflections in the trace of the axial plane account largely for the spread in the maximums of lineations in the northeast and southwest quadrants of the diagram in figure 8D. In the same way the plunge of the axis varies. On the whole the plunge is southwestward within the quadrangle, but locally it is northeastward; small reversals in plunge along the axis are characteristic and can be seen in section E—E’ of plate 2 in the report on the Central City district by Sims and Gable (1964). The limbs dip 40° away from the crest on the average, but in detail they are corrugated by minor folds of several types. The principal minor folds are open, upright anticlines and synclines that have breadths that exceed heights. Many are nearly symmetrical; others are asymmetrical and have the normal relations of drag folds. Fold axes for the most part are nearly parallel to the axis of the Central City anticline. Locally, tight folds whose heights exceed breadths occur on the limbs, particularly near the crest. One such fold belt is marked by the zigzag pattern of the calc-silicate rocks on the northeast slope of Quartz Hill. Less commonly small recumbent folds whose axes are subparallel to the major fold axes occur. The folds that were observed during detailed mapping of the Central City district are at most a few feet across (breadth) and a few tens of feet from crest to crest. Without exception the axial planes of the recumbent folds are subparallel to the foliation planes of the overlying and underlying rocks. Small-scale folds that trend nearly at right angles ROCKS, CENTRAL CITY QUADRANGLE, COLO. E47 to the major fold axis are uncommonly abundant on the flanks of the Central City anticline. They are small undulations that do not modify the structural framework at the scale of plate 1 but warp or crinkle the foliation planes and the B lineation. Rarely, in biotite-rich rocks, these folds are accompanied by a subparallel mineral alinement (A). The folds differ in form within rocks of difl’erent lithology. Within the relatively competent microcline gneiss, they are typically open, low-amplitude warps and irregular undulations that range in breadth from a few inches to at least several tens of feet. They are commonly nearly symmetrical but at places are strongly asym- metrical; the axes are poorly defined and discontinuous and generally are arranged en echelon. At places pegmatite streaks and boudins or incipient boudinage are subparallel to the fold axes. The folds within relatively incompetent biotite gneiss, however, are more sharply contorted and commonly are more strongly asymmetrical. They range in breadth from tiny crinkles to folds several feet across; commonly their heights are nearly as great or are greater than their breadths. The crinkles are mostly of chevron type and are most common in biotite-rich layers. They bend the biotite flakes and sillimanite needles; rarely the biotite flakes are broken at the crest of the chevrons. The axial planes of the crinkles and of other small-scale folds typically converge upward toward the axes of larger anticlines trending in the A direction. IDAHO spnmes ANTICLINE The Idaho Springs anticline, in the southeast corner of the quadrangle, is strongly asymmetrical and has a steep northwest limb and a gentle southeast limb. Its axial plane dips steeply southeastward as noted in section D—D’, plate 1. The axis plunges 25°—50° NE., which is steeper than the plunge of the axes of most other folds in the quadrangle. The fold has been traced with certainty southwestward from the quad- rangle for a distance of about 2 miles and, judged from the regional map of Lovering and Goddard (1950, pl. 2), extends still farther south; probably the bulbous, northeastward-projecting mass of Boulder Creek Granite mapped by Lovering and Goddard (1950) along Barbour Fork is on the axis of this anti— cline. Details of the fold, including the superposed folds imposed by the younger deformation within the Idaho Springs-Ralston shear zone, are described by Moench (1964). LAWSON swarms The Lawson syncline, the dominant fold in the western part of the quadrangle, is a complex, broad, open, upright structure outlined by the Lawson layer of microcline gneiss and adjacent units of biotite gneiss E48 (pl. 1). It has been traced along its axis with certainty for about 5 miles, from south of Clear Creek northeast- ward to the thick lens of microcline gneiss east of Yankee Hill (pl. 1). The axial plane dips nearly vertically; the trace of the axial plane trends N. 15°— 20° E. and is moderately straight. The axis plunges northeastward at a moderate angle, except for local small reversals. Judged from measurement of minor folds and related lineations at and near the major fold axis, the axis plunges on the average about 25° NE. (fig. 8A). The limbs dip gently inward toward the axis. According to our interpretation of the lithologic succession, the east limb, marked by micro- cline gneiss, extends irregularly northward into the adjacent Nederland quadrangle. The west limb lies outside the quadrangle in the Empire quadrangle; W. A. Braddock (written commun, 1963), who mapped this quadrangle, reported that the microcline gneiss on the west limb pinches out abruptly a short distance west of the boundary of the Central City quadrangle, possibly as a result of a stratigraphic pinchout. Within the trough and on the east limb of the syncline, between Clear Creek and Fall River, the microcline gneiss and adjacent biotite gneiss are warped by a series of small, subparallel folds that plunge gently either about N. 70° E. or S. 70° W. The folds are accompanied by a mineral lineation and boudinage. A second lineation direction marked by small-scale folds and a mineral alinement plunges about N. 20° W. or S. 20° E., approximately at right angles to the other fold axes. Folds in the N. 20° W. direction locally contain elongate pods and larger phacolithlike bodies of biotite—muscovite quartz monzonite; at other local- ities these folds are crosscut by dikes of biotite-mus- covite granite and were apparently formed prior to the emplacement of the intrusive rock. The axis of the Lawson syncline can be traced from the Lawson layer of microcline gneiss northward into biotite gneiss but apparently terminates abruptly in the area of the small body of microcline gneiss on the ridge on the east flank of Yankee Hill (pl. 1). The reason for the abrupt termination is not known. From what can be seen at the sparse outcrops the fold axis is interpreted to pass into a slightly overturned fold in biotite gneiss which contains thin stringers of micro- cline gneiss, or it dies out in these beds. Equally un- certain is the reason for the layers of microcline gneiss which trend north-northwest on the north slope of Sher- idan Hill. Although the layers conform to the foliated biotite gneiss, on a gross scale they appear to transect stratigraphic units within the biotite gneiss. The un— certainties in the structure of the Yankee Hill area are reflected by the very generalized form lines within bio- tite gneiss in section B—B’ of plate 1. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY INTERMEDIATE-SCALE FOLDS Several folds of intermediate size, subparallel to the major folds, occur between the axes of the larger folds. These are described as follows in order of their succes— sion from southeast to northwest. PEWABIC MOUNTAIN SYNCLINE The Pewabic Mountain syncline, a relatively small, virtually symmetrical flexure midway between the Idaho Springs and Central City anticlines, has been traced along its strike for about 5 miles. Throughout most of its length it is an open, upright syncline, which plunges gently in either direction. Southwestward, on the east flank of Bellevue Mountain, the fold tightens and locally is overturned to the southeast. BALD MOUNTAIN SYNCLINE The Bald Mountain syncline is a relatively narrow fold that is outlined in part at Bald Mountain by a folded sheet or phacolith of granodiorite. The axis of the syncline trends about N. 40° E. and plunges gently or moderately northeast on the average. Associated small folds and B lineations plunge variably either northeast or southwest (fig. 80). Although the syn- cline is dominantly an open fold, tight folds plicate the limbs, as shown by the detailed map in the report on the Lawson-Fall River-Dumont district (Hawley and Moore, 1967). DUMONT ANTICLINE The Dumont anticline is a tight, upright fold in bio- tite gneiss about midway between the Quartz Hill and Lawson layers of microcline gneiss. It is inferred to extend from Clear Creek nearly to lVIissouri Creek, but because of widely scattered exposures and the absence of definitive stratigraphic markers in the biotite gneiss unit, its validity as a single fold is not certain. Obser- vations of small-scale folds and related lineations indi— cate reversals in plunge along the length of the fold (fig. 80). The crest of the fold is marked nearly every— where by abundant, tight, nearly symmetrical folds of small size. As mapped, phacolithlike bodies of biotite- muscovite quartz monzonite occur locally along the fold axis. smomm AT MOUNT PISGAH The syncline at Mount Pisgah is relatively tight and contains a phacolithic body of granodiorite and associ- ated rocks. The fold is well defined on Mount Pisgah but apparently dies out within a mile to the southwest; it can be traced only a short distance into the pluton. Judged from lineation measurements within the pluton and the biotite gneisses, the fold plunges about 30° NE. PECKS FLAT ANTICLINE AND ADJACENT OVERTURNED FOLDS Between the syncline at Mount Pisgah and the major Lawson syncline to the west are several small, tight, PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. overturned folds that are spaced about 1,000 feet apart and that produce sharp plications in the Lawson layer of microcline gneiss and the underlying and overlying units of biotite gneiss. The folds are dominantly closed and slightly asymmetrical and are overturned to the west. Details of the folds are not well known; they occur in a structurally complex area, which is compli- cated by drastic thickening and thinning of the Lawson layer of microcline gneiss, the dominant stratigraphic marker of the area. The Peeks Flat anticline, the easternmost of the folds, is a complex tight fold in Hamlin Gulch and becomes more open and upright northward along its axis. The axial plane dips steeply eastward, and its trace trends north. The axis plunges about 30° NE. on the average. At the head of Hamlin Gulch the granodiorite that constitutes the western part of the Mount Pisgah pluton is substantially thickened on the crest of the anticline and greatly thinned on the west, overturned limb. About 1,000 feet west of the Peeks Flat anticline, an irregular, narrow tongue of microcline gneiss (Lawson layer) appears to mark the trace of a closed, overturned syncline, the limbs of which probably dip about 50°—60° SE. A parallel, overturned anticline is inferred to lie just west of the syncline; it bisects the narrow tongue of biotite gneiss that lies west of the microcline gneiss. Two other folds, which are dominantly open and up- right, have been mapped on the north side of Fall River, just west of the overturned folds (pl. 1). They appear to have relatively small amplitudes and to be of re— stricted length. The syncline on the east is open in the microcline gneiss but becomes closed southward along the fold axis in the biotite gneisses. This change in the character of the fold is a manifestation of disharmonic folding. LINEAR ELEMENTS RELATED TO FOLDS Linear elements that are subparallel to the major fold axes (B) are ubiquitous in the metamorphic rocks and are present locally in all intrusive rocks except the biotite-muscovite quartz monzonite. Less commonly lineations are oriented nearly at right angles (A) to major fold axes. More than 90 percent of the visible lineations in the Precambrian rocks are subparallel to the major fold axes, or the (B) axis. These are mainly mineral aline- ments and the axes of small-scale folds but include boudinage. The lineations are represented by the strong maximums in the northeast quadrants of the lineation plots and by the weaker maximums in the southwest quadrants (A, B, 0, and D of fig. 8). As the statistical plots indicate, the lineations in the B direction vary somewhat both in orientation and angle of plunge throughout the quadrangle. In the north- E49 western and western parts of the quadrangle (fig. 8A) , the lineation “high” is rather sharply defined and sym- metrical. The maximum is 28° N. 17° E.; a moderate spread in bearing from about N. 10° W. to N. 50° E. and a substantial variance in angle of plunge occur. In the northeastern part of the quadrangle (fig. SB) the maximum is 12° N. 32° E. However, the bearing of the measurements varies in orientation about 90°, but the variance in the angle of plunge is less than in the western part of the mapped area (fig. 8A). The wide spread in the lineations is accounted for by fold axes that range in bearing from about north to east, as described in the Central City district report (Sims and Gable, 1964) for the area on the north side of North Clear Creek. Except for local areas near North Clear Creek and west of State Highway 119, it is doubtful that lineations related to the younger deformation can account for the east-northeast bulge in the northeast segment of the diagram (fig. SB), for other than in these areas, superposed folds have not been recognized. Lineations in the east-central part of the quadrangle (fig. 80) are comparable in bearing and angle of plunge to those in the northeastern part of the quadrangle but have less spread in orientation. Lineations in the Central City district, as indicated by the measurements plotted on figure 8D vary substantially in bearing; the mean value is 10° N. 27° E. The spread in orientation is mainly a reflection of a marked spread in the bearing of the major fold axes which can be seen on the geologic map of the Central City district (Sims and Gable, 1964, pl. 1); the 1— to 2-percent bulge at N. 60° E., however, mainly reflects lineations related to the younger deformation. Lineations in the A direction, oriented nearly at right angles to the major fold axes, occur sporadically throughout the quadrangle but are statistically suf- ficiently abundant to be shown by the contour lines only on figures 83, 80, and 8D. Lineations in the A direction are mainly small folds but include rare mineral alinements, boudinage, and slickenside striae. In figure 8D, the maximums in the northwest quadrant reflect measurements made on the northwest limb of the Central City anticline, and the maximums in the southeast quadrant reflect measurements made on the southeast limb, for the plunge, of course, is controlled by the dip of the limbs. IDAHO SPRINGS-RALSTON SHEAR ZONE The Idaho Springs-Ralston shear zone, which im- pinges on the southeast corner of the quadrangle, is the largest of several Precambrian shear zones known in the Front Range (Tweto and Sims, 1963, p. 998— 1000). It extends from a point a few miles south of the quadrangle northeastward to the mountain front, E50 a distance of about 23 miles. To the northeast it disappears beneath the Fountain Formation of Penn- sylvanian and Permian age; to the southwest it dies out, apparently in a large pluton of granodiorite of Boulder Creek affinity. The northwest boundary of the shear zone, which is southeast of the crest of Bellevue Moun— tain, marks the approximate limit of intense cataclastic deformation (pl. 1). Within the quadrangle the shear zone is characterized by extreme cataclasis, minor folds and related linear elements, and weak local recrystallization. The rela- tively competent rock units, as microcline gneiss and granite gneiss and pegmatite, are nearly pervasively cataclastically deformed, whereas the less competent rock units, although sheared parallel to their foliation, are folded. The folds and cataclasis are superposed on the folds formed by the older, plastic deformation. The folds within the shear zone trend about N. 55° E., subparallel to the zone itself, and plunge at various angles, depending upon the attitude of the older, larger folds. They range from about a foot in breadth to about 400 feet and consequently are too small to show at the scale of plate 1. The axial planes dip steeply southeastward and are remarkably straight and persistent. The folds tend to be sharp crested and are either simple or complex structural terraces or closed chevron folds; with few exceptions they are strongly asymmetrical and have steep, generally long northwest limbs and short, crumpled southeast limbs. The form depends largely on the geometry of super- position and to a lesser extent on the type of movements that produced the folds. The distribution and charac- ter of the folds are shown in the report on the Idaho Springs mining district (Moench, 1964), and the geometry of superposition is discussed in the report by Moench, Harrison, and Sims (1962, p. 49—55). Cataclasis is nearly pervasive in the shear zone and takes several forms, each form depending on the intensity of shearing and the nature of the rock. The effects of shearing range from incipient granulation and mortar textures, visible only in thin sections, to flaser structures and, locally, to mylonite. In general the shearing was localized along preexisting foliation surfaces, especially in the biotite gneisses. Biotite tends to be smeared out and locally chloritized in such rocks, and quartz tends to be elongate or amoebiform. On the crests of minor folds the micas are bent and sillimanite needles are broken. In the more competent and homogeneous rocks the original foliation planes are transected in places and locally obliterated by shear planes which have produced a new foliation—a meshwork of subparallel, interconnecting fractures. In general these shear planes dip roughly parallel to the contact of the Big Five layer of microcline gneiss and SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY the overlying biotite gneiss unit—about 35°—50° NW. Linear elements on the shear surfaces include slickenside striae, mineral streaking, and rodding, and all are oriented northwest, nearly at right angles to the associated fold axis. SHEARING AND ASSOCIATED ALTERATION IN DAKOTA HILL AREA Adjacent to the Tertiary intrusive porphyry at Apex, the Precambrian rocks are sheared as well as altered. Details of the shearing are obscure because of poor exposures, but the general character of the shearing is known from scattered surface exposures and from openings in the tunnel of the Nye-Mathews molyb- denum prospect on the west flank of Dakota Hill. On the west slope of Dakota Hill the rocks are sheared and cataclasized for a distance of at least 2,000 feet from the porphyry contact; shearing is most intense within a few hundred feet of the contact and decreases westward. Adjacent to the contact the Precambrian rocks are profoundly altered as well as sheared and are cut by quartz veins that locally contain molyb- denite. Some quartz veins near the contact and within the porphyry, which similarly is sheared, are as much as 8 inches thick, but most are less than an inch. On the crest and east slope of Dakota Hill shearing and veining is somewhat less intense except near the contact. Megascopically, the sheared rocks are bleached to a grayish white and are locally veined by milky quartz. They have a well-developed gneissic structure and a distinctly granulated appearance, and thus individual rock types are difficult to distinguish separately. The predominant textural changes are a decrease in average grain size and a destruction of the common granoblastic texture by cataclasis and recrystallization. Quartz forms sutured, elongate aggregates of grains that have a marked linear fabric in a fine-grained matrix composed mainly of feldspars. The close association of both shearing and alteration with the Tertiary stock suggest that both were produced by the emplacement and attendant thermal effects of the igneous body. However the stock may have been intruded into a previously deformed zone, for it occurs at the approximate junction of the northward-trending Apex fault and the northwestward—trending Blackhawk fault—faults believed to have originated during the Precambrian (Sims, Drake, and Tooker, 1963, p. 20—22). The shearing may, therefore, be partly of Precambrian age. Some of the deformation is of definite Tertiary age, however, for the stock itself is locally broken by a complex set of fractures, some of which contain quartz and molybdenite. Further studies of the Tertiary intrusive bodies are needed. PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. FAULTS Faults that formed initially in the Precambrian, subsequent to the development of the Idaho Springs- Ralston shear zone, are widely spaced in the quad- rangle. These are northwestward-trending faults, which belong to the group of faults known in the Front Range as breccia reefs (Lovering and Goddard, 1950), and faults trending north-northeast. Other faults, which are far more abundant and which contain most of the ore deposits, formed later during the Laramide orogeny; these form a meshlike pattern in the Central City district (Sims, Drake, and Tooker, 1963). The northwestward-trending faults are character- istically long and persistent; some have been traced across the full width of the mineral belt and for many miles on either side. They are spaced 1—6 miles apart. Within the quadrangle the faults have a left-lateral displacement of as much as 600 feet; they commonly contain porphyry dikes of Laramide age along a part of their length. In detail the faults are complex sheared zones consisting of one or more subparallel major strands and intervening branching and inter- connecting fractures. Most have a complex history of movement. The Idaho Springs and Blackhawk faults, the principal faults of this set in the quad- rangle, show predominant strike-slip movement of several generations, probably Precambrian, and rela- tively late vertical and oblique movements of Lara— mide age. Local concentrations of base-metal sulfides in the faults are related to openings formed by recur— rent movements in Laramide time (Sims, Drake, and Tooker, 1963; Sims, Armstrong, and others, 1963). The faults trending north—northeast are likewise long and persistent but are more widely spaced than are the northwestward-trending faults. The only significant fault of this trend in the quadrangle is the Wild Wagoner-Apex fault, a conspicuous fracture zone that extends northward through the western part of the quadrangle. It is poorly exposed but appears to consist dominantly of a single, major strand and subordinate subparallel fractures. It is filled mainly by gouge and breccia but locally by siliceous breccia and quartz-sulfide veins. At places it contains por— phyry dikes. Movement along the fault is right lat- eral; also the east wall is dropped relative to the west wall. The apparent displacement is determined mainly by the offset of the Tertiary bostonite porphyry dike in the Pecks Flat area (pl. 1). Evidence of a Precambrian origin for both sets of faults is strong in the Front Range as a whole (Tweto and Sims, 1963) but is not definitive within the Central City quadrangle itself. Within the quadrangle the Idaho Springs fault equally displaces the young folds and the older features of the Idaho Springs-Ralston E51 shear zone; this indicates that the displacement is younger than the cataclasis and the younger folding. The Precambrian age is inferred from the local presence of mylonite, a product of a relatively deep-seated environment, and of Precambrian aplite and pegmatite dikes along faults of similar attitude and habit in adjacent areas to the north (Lovering and Tweto, 1953, pl. 1) and east. Many of the faults disappear beneath the Paleozoic and sedimentary rocks at the mountain front, and even those that were rejuvenated in Laramide time apparently die out rapidly in the sedimentary rocks (Lovering and Goddard, 1950, pls. 1—2; Boos and Boos, 1957, figs. 3—10). JOINTS Joints were measured in the quadrangle to supple- ment the data compiled earlier by Harrison and Moench (1961) for the Central City-Idaho Springs area. The joints were plotted on Schmidt equal-area nets to show statistically their orientation. Figure 9A includes meas- urements from the northwestern and western parts of the quadrangle and covers the same area as figure 8A, a plot of lineations; figure QB includes measurements from the northeastern part and covers the same area as figure SB. Diagrams are not included for those parts of the quadrangle that were mapped previously at a more detailed scale; these are given in the earlier sum- mary report by Harrison and Moench (1961). The joint diagrams were constructed according to the method described by Billings (1942, p. 119—121). In such diagrams each pronounced maximum represents the approximate attitude of the poles of many joints. The most pronounced and abundant joint set in both the areas (A and B of fig. 9) strikes northwestward and dips steeply either to the northeast or to the southwest. That the maximum for this set is broad and irregular in both diagrams indicates a substantial spread in both strike and dip. In the northwestern part of the quad- rangle (fig. 9A), the joint set has an average strike of N. 50°—65° W., and dips 85° NE. to 87° SW.; in the northeastern part it has an average strike of N. 45°—55° W. and dips 80° NE. to 85° SW. The second most abundant joint set trends roughly east and dips steeply either to the north or the south. The concentration of poles in figure 9A indicates an average attitude of N. 85°—90° E., vertical; similarly, the average attitude in figure 9B is N. 85°—90° W., 85° S. to 85° N. Other less conspicuous joint sets, judged from the contour diagrams, are not equally developed in both parts of the region. A joint set that strikes N. to N. 5° W. and dips vertically is represented by the 2.5— to 3-percent concentration of poles at the east and west axes of the contour diagram for the northwestern E52 part of the quadrangle (fig. 9A); the set is not evident in the northeastern part of the map area (fig. 9B). A joint set that has an average strike of N. 45° E. and dips 70° NW. is indicated by the 3—percent maximum in figure 9B; it probably is represented by the bulge in the 2-percent concentration of poles at the same location in figure 9A. A joint set that is conspicuous in the northeastern part of the quadrangle has an average strike of N. 70°—80° E. and dips 70°—80° NW. but apparently is not present in the northwestern part, unless it is represented by the westward bulge of the 2.5— to 3—percent contour at the top of figure 9A. If so, it probably overlaps the more conspicuous set striking N. 85°—90° E. and dipping vertically. Another joint set striking N. 32° W. and dipping 72° NE. is represented by the 2.5- to 3-percent “bulls eye” in figure QB; this set is not evident in figure 9A. The joint pattern in the Precambrian rocks of the region can be interpreted to have resulted mainly from stresses related to the major episode of Precambrian folding and to the uplift of the Front Range highland in the Laramide orogeny (Harrison and Moench, 1961). The prominent northwestward-trending joint set can be inferred to be a cross joint to the major folds. Com- parison of figures SB and 9B shows that the position of the joint set striking N. 45°—50° W. corresponds exactly to the theoretical position of a cross joint to the major folds; that is, the joint set is perpendicular to the SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY average plunge of the fold axes as represented by the lineation “high” in figure 8B. The northwestward— trending joint set in the northwestern part of the quadrangle is not quite at right angles to the mean value for the plunge of the fold axes, as indicated in figure 8A; instead it is steeper and differs in strike by a few degrees. The rest of the joint sets represented by the concen- trations of poles in the diagrams can be related to the regional joint system distinguished by Harrison and Moench (1961, p. B12—B14). The north-south joint set in the northwestern part of the map area (fig. 9A) has approximately the relationship of a regional longi- tudinal joint; the set striking N. 70°—80° E. in the northeastern part (fig. 9B) perhaps represents a regional cross joint, and the set striking N. 45° E. represents a regional diagonal joint. Although these interpretations may be speculative, they are consistent with the earlier conclusion of Harrison and Moench (1961), which was derived from a study of the joint data in the southern part of the quadrangle and adjacent areas to the south. Their conclusion was that the regional joint system was formed during the uplift of the Front Range highland. CHARACTER AND ENVIRONMENT OF DEFORMATION The gneissic structure, granoblastic texture, and folds and related lineations can all be related to the episode of plastic deformation that affected the central part of A. WESTERN AND NORTHWESTERN PARTS; 522 POLES FIGURE 9,—Contour diagrams of poles to joints in the Central City quadrangle. B. NORTHEASTERN PART; 386 POLES Upper hemisphere plots, contoured in percent. PETROLOGY AND STRUCTURE, PRECAMBRIAN ROCKS, CENTRAL CITY QUADRANGLE, COLO. the Front Range. Deformation occurred in an envi- ronment in which metamorphic temperatures over- lapped the magmatic range. The Idaho Springs- Ralston shear zone, which resulted from a younger deformation, formed in a less intense pressure-tempera- ture environment. In contrast to the earlier pervasive deformation, this deformation was local in extent and resulted from shearing along a linear zone subparallel to the Colorado mineral belt (Tweto and Sims, 1963). The parallelism of foliation with original bedding, evident in nearly every outcrop in the region, is char- acteristic of high-grade metamorphic rocks in all parts of the world. The parallelism cannot be accounted for by an hypothesis of load metamorphism, for the fabric resulting from such metamorphism should show axial symmetry about the axis of loading, whereas the ob- served fabric shows symmetry about the axes of folding. The origin of the fabric, therefore, must be directly connected with the regional stresses that produced the folding. To account for the foliation and lineation we suggest that initially compacted rocks, buried under a substantial thickness of cover, recrystallized under high temperatures and pressures along slip planes that fol- lowed original bedding. Both foliation and lineation were formed by penetrative movement along the planes, probably involving both external rotation and micro- folding about the B lineation direction. Microfolding was a significant mechanism only in the relatively in- competent biotite gneisses. Studies of the space-lattice orientation of mineral grains were not made specifically to determine the mechanism of deformation, but petro- fabric diagrams of quartz and xenotime from unoriented samples made for other purposes (Young and Sims, 1961, p. 292—293) indicate maximums in the foliation plane. This suggests that movement was in the plane of the foliation. The relatively uniform character of foliation and lineation in all the metamorphic rocks of the region, regardless of the intensity of the associated folds, shows that movements along the slip planes need not be large to produce strong foliation. The relative abundances of platy and tabular minerals are more sig- nificant than the intensity of folding in determining whether a specific rock is weakly or strongly foliated. The axial-plane flow cleavage noted locally in the area formed only in incompetent rocks. At the previ— ously described locality at the mouth of Silver Creek, in the eastern part of the Central City district (Sims and Gable, 1964), well-oriented sillimanite and mica, in particular, formed in shear planes that were parallel to axial planes of small folds in biotite gneiss. These shears do not cut the more massive rock layers. Evi- dently the compression that produced the folds in these incompetent rocks also produced parallel shear planes that were perpendicular to the greatest stress axis and E53 an extension parallel to the least stress axis. In more competent, intercalated rocks it produced concentric shears and slippage along bedding planes. The folding was accomplished largely by fiexure slip. The thick layers of relatively competent microcline gneisses, which were separated by relatively incom- petent biotite gneisses, favored the formation of broad open folds. Tight and locally overturned folds of short wavelength formed mainly within the relatively incompetent biotite gneisses. Throughout the interval of folding the competent layers, especially the micro- cline gneiss layers, probably yielded dominantly by concentric folding. Elastic bending was accompanied by the formation of subparallel concentric shearplanes on the flanks and, to a lesser extent, on the crests. The less competent biotite gneisses in the succession also yielded by slip movements along concentric shear planes. In contrast to the competent microcline gneisses, however, the biotite gneisses were deformed partly by small~scale folding that produced drag folds, the attitudes of which reflect the direction of relative displacement of the overlying and underlying more competent layers. Transfer of material from the limbs to the crests in the incompetent layers was accomplished largely by this type of folding. Possibly to a certain degree it was accomplished also by solution and redep- osition. In addition the potential open spaces in the crests of several folds of intermediate scale were filled by syntectonic intrusive rocks. To a considerable degree the folding is disharmonic. Disharmonic folds of different types are visible on a small scale at scattered localities throughout the region and can confidently be inferred to exist on a larger scale. A small—scale example of a strongly disharmonic fold characterized by extreme attenuation of a rela- tively incompetent layer within amphibolite has been illustrated by Moench, Harrison, and Sims (1962, pl. 3, fig. 2). Larger scale folds of a similar type are mani- fested by the local belts of tight folds within broader areas of open folding, as shown, for example, by map patterns in the Central City district (Sims and Gable, 1964, pl. 1). Within the quadrangle pronounced dis- harmonic folds on a relatively small scale can be observed in the syncline along Fall River, just west of Hamlin Gulch. As the fold axis is traced southward from microcline gneiss into underlying biotite gneisses, the fold tightens and the axial area is characterized by abundant small chevron folds as much as a few feet in width. Insofar as known, the contact between the microcline gneiss and the biotite gneiss, which must be a detachment zone, is relatively smooth and even. Larger scale disharmonic folds of a similar type are suggested by the map patterns (pl. 1). The northeastward-trending fabric of the Precam- E54 brian rocks has been interpreted to have resulted from a horizontal couple (Sims and Gable, 1964). Such a stress pattern can account for the bends in the fold axes and for the minor warps in the A direc- tion as well as for the major fold pattern. Knowledge of the folds within a larger area is needed, however, to confidently interpret the stress pattern. The Idaho Springs-Ralston shear zone is one of several northeastward-trending shear zones in the Front Range. The folding within the shear zone has many characteristics of slip or shear folds, as defined by Turner (1948, p. 165—174) and others but lacks an associated distinct fracture cleavage (Moench, Harrison, and Sims, 1962, p. 54). Associated cataclasis indicates deformation under less intense pressure-temperature conditions than for the older, plastic folding and meta- morphism. The shearing is a manifestation of regional stresses that produced major shear zones which are approximately coextensive with the Colorado mineral belt. The shear zones are interpreted as a dominant controlling structure responsible for localizing the mineral belt (Tweto and Sims, 1963). GEOLOGIC HISTORY The natural history recorded by the rocks in the central part of the Front Range is moderately well known as the result of investigations carried on in recent years. Nevertheless many details are lacking, and knowledge particularly of events that preceded the major episode of Precambrian metamorphism and de- formation is lacking. Accurate knowledge of the ages of deformation and attendant metamorphism as well as of the intrusive igneous rocks still is meager. Until reliable information secured by dating methods is available, correlation of events in this area with those in other parts of the Rocky Mountain region is not feasible. The first event that was recognized in the area was deposition of the sediments that subsequently yielded the metamorphic rock succession. Thick units of inter— bedded graywacke and shale and of feldspathic sand- stone were deposited repeatedly. An estimated 15,000 feet of sediments was formed, apparently without a major break in deposition, for major unconformities have not been recognized in the succession. Lesser bodies of carbonate(?) sediments and clean quartz sands were deposited within the major sedimentary deposits. Judged from the layered rocks now visible, the character of sedimentation did not change materially with time, for units of the two major rock types throughout the succession are similar both in composition and in struc- ture. The nature and source of the sediments are conjectural. Inferrence can be drawn from their . present compositions that the source materials were SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY either felsic metamorphic rocks or felsic or intermediate intrusive rocks, but an older succession of this lithologic character has not been identified within the Front Range. Regardless of source it is probable by analogy with more recent sediments resembling wacke that the sediments were deposited in a eugeosynclinal environ- ment. Possibly the area was transitional to a mic- geosynclinal environment, for the rocks preserved in areas immediately to the east are more typical of those formed in a miogeosyncline (D. M. Sheridan, written commun., 1965). Subsequent to deposition the sediments undoubtedly were buried under a substantial cover of younger strata, probably to depths in the range of 7 to 12 miles. As a consequence of the superincumbent load, the materials must have been largely dewatered. Original bedding probably was accentuated by the for- mation first of clay minerals, then of micas, and per- haps also by incipient growth of other platy or tabular minerals. No evidence has been found to indicate that the strata were deformed by other than mild load metamorphism during this interval; apparently at the onset of regional deformation the succession was sub- horizontal and lacked any notable cleavage. The episode of regional metamorphism and plastic deformation that impressed upon the rocks most of the features visible today took place at high temperatures and pressures in the catazone of the crust as defined by Buddington (1959). The foliation and lineation and the mineral phases characteristic of the sillimanite zone were accomplished relatively early in the deformation. Migmatites formed at or near the culmination of the thermal cycle, in the interval during which metamor- phic temperatures overlapped the igneous range of temperatures. Presumably they formed from the rel- atively low melting components of the country rocks, which were mobilized in the deeper parts of the se- quence and moved upward and outward as a fluid phase along the foliation of the biotite gneisses and, to a lesser extent, into the other metamorphic rocks. During the episode of deformation the intrusive igneous rocks were emplaced. Granodiorite intruded early in the episode, probably near the thermal max- imum, but after most or all of the migmatization. The granodiorite moved upward from depth, probably as a magma, and along foliation planes and into the crests and troughs of active folds. Intrusion appar- ently was largely passive; the magma entered potential open spaces in the apical areas of the folds without notable disruption of the walls. Continuing stresses after consolidation of the granodiorite produced a sec- ondary foliation and lineation, subparallel to the pri- mary flow structures, in the borders of the intrusive bodies. The smaller bodies, and probably also those PETROLOGY AND STRUCTURE, PRECAMBRIAN emplaced relatively early in the intrusive interval, were completely recrystallized and folded. After the thermal maximum was reached, gabbroic magma was emplaced, also largely by a phacolithic mechanism. The rocks in a narrow halo around the larger Elk Creek pluton were progressively metamorphosed, but the latent heat in the smaller bodies was not sufficient to noticeably change the mineral compositions of their wallrocks. After cessation of the major deforming stresses, biotite—muscovite quartz monzonite was em- placed mainly along crosscutting structures or across the gneissic structure of the wallrocks to yield mark- edly crosscutting bodies, but to a small degree it was emplaced by a phacolithic mechanism. The quartz monzonite was emplaced sufficiently late in the de- formation to escape crushing and recrystallization; it has a wholly primary flow texture and was probably the biotite-muscovite quartz monzonite crystallized from a magma. Probably a period of relative quiescence during which the region was undergoing erosion preceded the next event that has been recognized—cataclastic deforma— tion. The Idaho Springs-Ralston shear zone and scattered lesser shear zones were formed as a result of shearing along linear belts which were coincident with the present site of the Colorado mineral belt. Judged from the pervasive cataclasis and minor recrys- tallization, the deformation took place under much lower pressure-temperature conditions than did the earlier regional deformation, perhaps at a significantly lesser depth. Possibly the deformation caused the retrograde metamorphism evident in many parts of the region and even outside the shear zone. Later, but still in the Precambrian, faults trending northwest and northeast and related breccia reefs were formed. The post-Precambrian history of the central part of the Front Range has been discussed by several geolo- gists (Lovering and Goddard, 1950; Sims, Armstrong, and others, 1963; Sims, Drake, and Tooker, 1963; and Harrison and Moench, 1961) and except for the events of the Laramide orogeny need not be reiterated here. Uplift of the Front Range along an arm's trending north— northwest began in Late Cretaceous time and continued, probably intermittently, into the Tertiary. In early Tertiary time several varieties of porphyritic igneous rocks (Lovering and Goddard, 1938; Wells, 1960) were emplaced as small plutons and dikes within and adja- cent to the present mineral belt. In large part the porphyries were intruded, probably passively, into joints, faults, and other planes of structural weakness in the Precambrian country rocks, but in part the intru- sions appear to have been accompanied by explosive activity to yield breccias. Near the close of the hypabyssal igneous activity, three main sets of faults ROCKS, CENTRAL CITY QUADRANGLE, COLO. E55 that make a complex intersecting network formed with- in the mineral belt, and the older (Precambrian) faults were rejuvenated. Mineralization of these fissures yielded gold- and silver-bearing base-metal sulfide ores. These are described and discussed in the report by Sims, Drake, and Tooker (1963). Subsequent uplift, erosion, and weathering have altered the veins to shallow depths, and locally rich supergene ores and placers have resulted. REFERENCES CITED Aldrich, L. T., Wetherill, G. W., Davis, G. L., and Tilton, G. R., 1958, Radioactive ages of micas from granitic rocks by Rb- Sr and K-A methods: Am. Geophys. Union Trans, v. 39, no. 6, p. 1124—1134. Ball, S. H., 1906, Pre—Cambrian rocks of the Georgetown quad- rangle, Colorado: Am. Jour. Sci., 4th ser., v. 21, p. 371~389. Barth, T. F. W., 1959, Principles of classification and norm cal- culations of metamorphic rocks: Jour. Geology, v. 67, no. 2, p. 135—152. 1962, A final proposal for calculating the mesonorm of metamorphic rocks: Jour. Geology, v. 70, p. 497—498. Bastin, E. S., and Hill, J. H., 1917, Economic geology of Gilpin County and adjacent parts of Clear Creek and Boulder Counties, Colorado: US. Geol. Survey Prof. Paper 94, 379 p. Billings, M. P., 1942, Structural geology: New York, Prentice- Hall, Inc., 473 p. Boos, C. M., and Boos, M. F., 1957, Tectonics of eastern flank and foothills of Front Range, Colorado: Am. Assoc. Petro- leum Geologists Bull., v. 41, no. 12, p. 2603—2676. Boos, M. F., and Boos, 0. M., 1934, Granites of the Front Range—the Longs Peak-St. Vrain batholith: Geol. Soc. America Bull., v. 45, no. 2, p. 303—332. Bowen, N. L., and Tuttle, O. F., 1950, The system NaAlSiaog- KAISiaoa-Hzoz Jour. Geology, v. 58, no. 5, p. 489—511. Buddington, A. F., 1959, Granite emplacement with special ref- erence to North America: Geol. Soc. America Bull., v. 70, no. 6, p. 671—747. 1963, Isograds and the role of H20 in metamorphic facies of orthogneisses of the northwest Adirondack area, New York: Geol. Soc. America Bull., v. 74, no. 9, p. 1155—1182. Chayes, F. A., 1949, A simple point counter for thin-section analysis: Am. Mineralogist, v. 34, nos. 1—2, p 1—11. Cloos, Ernst, 1946, Lineation, a critical review and annotated bibliography: Geol. Soc. America Mem. 18, 122 p. Engel, A. E. J., and Engel, C. G., 1962, Progressive metamor- phism of amphibolite, northwest Adirondack Mountains, New York, in Engel, A. E. J., James, H. L., and Leonard, B. F., eds. of Buddington volume, Petrologic studies: New York, Geol. Soc. America, p. 37—82. Fairbairn, H. W., 1949, Structural petrology of deformed rocks [2d ed.]: Cambridge, Addison-Wesley Press, 344 p. Folinsbee, R. E., 1940, Gem cordierite from the Great Slave Lake area, N.W.T., Canada [abs]: Am. Mineralogists, v. 25, p. 216. 1941, The chemical composition of garnet associated with cordierite: Am. Mineralogist, v. 26, no. 1, p. 50—53. Fyfe, W. S , Turner, F. J , and Verhoogen, Jean, 1958, Metamor- phic reactions and metamorphic facies: Geol. Soc. America Mem. 73, 259 p. E56 Guidotti, C. V., 1963, Metamorphism of the pelitic schists in the Bryant Pond quadrangle, Maine: Am. Mineralogist, v. 48, p. 772—791. Harrison, J. E., and Moench, R. H., 1961, Joints in Precambrian rocks, Central City-Idaho Springs area, Colorado, in Shorter contributions to general geology: U.S. Geol. Survey Prof. Paper 374—B, p. B1—B14. Harrison, J. E., and Wells, J. D., 1956, Geology and ore deposits of the Freeland-Lamartine district, Clear Creek County, Colorado: US. Geol. Survey Bull. 1032-B, p. 33—127. 1959, Geology and ore deposits of the Chicago Creek area, Clear Creek County, Colorado: U.S. Geol. Survey Prof. Paper 319, 92 p. Hart, S. R., 1964, The petrology and isotopic-mineral age rela- tions of a contact zone in the Front Range, Colorado: Jour. Geology, v. 72, no. 5, p. 493—525. Hawley, C. C., and Moore, F. B., 1967, Geology and ore deposits of the Lawson-Dumont-Fall River district, Clear Creek County, Colorado: US. Geol. Survey Bull. 1231 (in press). Leonard, B. F., and Buddington, A. F., 1964, Ore deposits of the St. Lawrence County magnetite district, northwest Adiron- dacks, New York: US. Geol. Survey Prof. Paper 377, 259 p. Lovering, T. S., and Goddard, E. N., 1938, Laramide igneous sequence and differentiation in the Front Range, Colorado: Geol. Soc. America Bull., v. 49, no. 1, p. 35—68. 1950, Geology and ore deposits of the Front Range, Colo- rado: U.S. Geol. Survey Prof. Paper 223, 319 p. [1951]. Levering, T. S., and Tweto, Ogden, 1953, Geology and ore deposits of the Boulder County tungsten district, Colorado: US. Geol. Survey Prof. Paper 245, 199 p. Mackenzie, W. S., 1954, The orthoclase—microline inversion: Mineralog. Mag, v. 30, no. 225, p. 354—366. Moench, R. H., 1964, Geology of Precambrian rocks, Idaho Springs district, Colorado: U.S. Geol. Survey Bull. 1182—A. p. A1—A70. Moench, R. H., and Drake, A. A., Jr., 1966, Economic geology of the Idaho Springs district, Clear Creek County, Colorado: US. Geol. Survey Bull. 1208, 91 p. Moench, R. H., Harrison, J. E., and Sims, P. K., 1962, Precam- brian folding in the Idaho Springs-Central City area, Front Range, Colorado: Geol. Soc. America Bull., V. 73, no.1, p. 35—58. Prider, R. T., 1940, Cordierite-anthophyllite rocks associated with spinel-hypersthenites from Toodyay, Western Aus- tralia: Geol. Mag. [Great Britain], v. 77, no. 5, p. 364—382. Reed, J. J ., 1957, Petrology of the lower Mesozoic rocks of the Wellington district [New Zealand]: New Zealand Geol. Survey Bull. 57, new ser., 60 p. Schreyer, W. F., and Yoder, H. 8., Jr., 1961, Petrographic guides to the experimental petrology of cordierite: Carnegie Inst. Washington Year Book 1960—61, p. 147—152. Sims, P. K., 1964, Geology of the Central City quadrangle, Colorado: US. Geol. Survey Geol. Quad. Map GQ—267. Sims, P. K., Armstrong, F. C., Drake, A. A., Jr., Harrison, J. E., Hawley, C. C., Moench, R. H., Moore, F. B., Tooker, E. W., SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY and Wells, J. D., 1963, Geology of uranium and associated ore deposits, central part of the Front Range mineral belt, Colorado: U. S. Geol. Survey Prof. Paper 371, 119 p. Sims, P. K., Drake, A. A., Jr., and Tooker, E. W., 1963, Economic geology of the Central City district, Gilpin County, Colo- rado: U.S. Geol. Survey Prof. Paper 359, 231 p. Sims, P. K., and Gable, D. J ., 1963, Cordierite-bearing mineral assemblages in Precambrian rocks, Central City quadrangle, Colorado, in Short papers in geology and hydrology: U.S. Geol. Survey Prof. Paper 475—B, p. B35—B37. 1964, Geology of Precambrian rocks, Central City district, Colorado: U.S. Geol. Survey Prof. Paper 474—0, p. 01—052 [1965]. Sims, P. K., and Sheridan, D. M., 1964, Geology of uranium deposits in the Front Range, Colorado: US. Geol. Survey Bull. 1159, 116 p. Taylor, R. B., and Sims, P. K., 1962, Precambrian gabbro in the central Front Range, Colorado, in Short papers in geology, hydrology, and topography: U.S. Geol. Survey Prof. Paper 450—D, p. D118—D122. Tilley, C. E., 1937, Anthophyllite—cordierite—granulites of the Lizard [Cornwall, England]: Geol. Mag. [Great Britain], v. 74, no. 7, p. 300—309. Turner, F. J., 1948, Mineralogical and structural evolution of the metamorphic rocks: Geol. Soc. America Mem. 30, 342 p. Turner, F. J., and Verhoogen, John, 1960, Igneous and met- amorphic petrology [2d ed.]: New York, McGraw—Hill Book Co., Inc., 694 p. Tweto, Ogden, and Sims, P. K., 1963, Precambrian ancestry of the Colorado mineral belt: Geol. Soc. America Bull., v. 74, no. 8, p. 991—1014. US. Geological Survey, 1964, Geological Survey Research 1964: US. Geol. Survey Prof. Paper 501-A, 367 p. Vasileva, Z. V., 1958, On the role of manganese in apatites: Akad. Nauk SSSR, Vses. Mineral. Obshch. Zapiski, v. 87, no. 4, p. 455—468. Washington, H. S., 1922, Deccan traps and other plateau basalts: Geol. Soc. America Bull., v. 33, p. 765—804. Wells, J. D., 1960, Petrography of radioactive Tertiary igneous rocks, Front Range mineral belt, Gilpin and Clear Creek Counties, Colorado: US. Geol. Survey Bull. 1032—E, p. 223—272. Wells, J. D., Sheridan, D. M., and Albee, A. L., 1964, Rela- tionship of Precambrian quartzite-schist sequence along Coal Creek to Idaho Springs Formation, Front Range, Colorado: US. Geol. Survey Prof. Paper 454—0, p. 01—025. Wilcox, R. E., and Poldervaart, Arie, 1958, Metadolerite dike swarm in the Bakersville—Roan Mountain area, North Carolina: Geol. Soc. America Bull., v. 69, p. 1323-1368. Yoder, H. S., and Eugster, H. P., 1955, Synthetic and natural muscovites: Geochim. et Cosmochim. Acta, v. 8, nos. 5—6, p. 225—280. Young, E. J., and Sims, P. K., 1961, Petrography and origin of xenotine and monazite concentrations, Central City district, Colorado: US. Geol. Survey Bull. 1032—F, 273—299. V) - w /. flu ex 0.2:; 49 “o C)» ‘0 o .5 UNITED STATES DEPARTMENT OF THE INTERIOR We PROFESSIONAL PAPER 554+: GEOLOGICAL SURVEY ‘ PLATE 1 105°37'30” , 39 {‘2ij , R. 74 W. R. 73 W, . , ‘ 35/ (NEDERLAND) 105 :0 . H E X P L A N A T l O N count-nos,» gouache?“ ~ ,-» , . .. I , « . .. I . 39 5230 0°2°uflconzzf° Iago 0°°°o-n n , I ‘ » L ll ( 33 >- °:o°°°o\o alocobog u°°€°°° i/g‘ Q: 75 O K II Izzy» I 2 /\\--/~ I o o ° 0 ° ° N 0° ‘2” El D h d h Contafti ted h t d h d vein’ Showmg dip “ as e w ere approxima e y oca ,' s or as e . . . . . , erei erred,d tedwh ,. E BlOtlte gnelss where inferred; dotted where concealed Dashed Wk "f at 6W pom-W71", \ 2) Gray, medium-grained, migmatitic biotite gneisses O gnb, dominantly sillimanitic biotite-quartz gneiss with ~\“ I] interlayers of biotite-quartz-plagioclase gneiss and / 8S0 ~.... . * locally garnetiferous biotite-quartz-plagioclase gneiss - - 7 , Intrusion breCCia gnbc, garnet- and cordierite-bearing phase of silli— Faulthshowmg dlp Disseminated ore I .. g 2 : , .. Altered f t . u . . manitic biotite-quartz gneiss- grades into sillimanitic Dashed where approximately located; short dashed , I, . Tame" S of (Home gneiss, pegmatite, and b‘ t't _ t - I ’ d t e t't where inferred; dotted where concealed :1 ‘ ' "£10700l’me-quartz-plagioclase—biotite gneiss in a io i e quar z gneiss. ncludes abun an I gma i e f ‘— 7 . matrix of quartz monzonite (g) and granite gneiss, both as discrete bodies and as __ Sha t . I , , ' thin wisps and layers 3(” \T~~.-.- i Q ‘ ‘Z r / - - /L I \ I I. go Ant1c11ne Portal of adit or tunnel . 1. . by vi =6 Showing approximate trace of axial plane, Short d . 'I 7 .37" , 0 lg dashed where inferred; dotted where concealed / Th? ,V, ‘ g It I f: i \; — , _ ‘ . ‘ I l Bostomte group 2 Cale-Silicate gneiss and related racks . *’ ‘\\‘\ _____ Trace of tunnel workings _ j . y \ M g g " I» ' $3 .> "to/“dis gluartz botstonite 190717,”va E 45’ Green, brown, white, and red, medium- to coarse- / _ l . 7'. \ 4 ) JO ‘ < b “to at” 97717;: e porphyry, “M < E grained, layered rocks of variable composition. Synclme fff—r-‘r' ’ l. / a I gnm 4; _. , “ 08 am e p Mp yry g — 8 Consist mainly of andraditic or grossularitic garnet, <2: Showing approximate trace of axial plane. Short —-\—-r“l’_" ‘l J ' ‘ ‘ x " /m'i f/ ,1 1/ ' ‘ * . »_ , . ' ' .. , E '8 clinopymxe’ne, epidote, feldspar, quartz, 36111101719. — das’bed Where ”inferred; ‘10th Where concealed Approximate boundary of zone of cataclasis and I - \ ,zl \ 47 z i . if _ ,_ {30 ‘ . , Hornblende granodlorite group DJ 56 and amphibole; locally interlayered with amphibolite g younger folds __ y' I ‘ ‘ ' . I, - . . . *\ / .1 . y W \ .. \ , . ‘ . . Includes biotite-quartz—latite por- i— "g \\‘ H . ‘ ///‘ ‘QII K“ 03?“ ' ~ , I. i phyry, biotite-quartz monzonite ,. i < > <2: Xfiao ——- “Chm‘es 1’0"” toward ”me >/ _ l— _‘ f . , \ ~y , . j , I. . , porphyry, biotite granodiorite'por— - Quartz monzonite group 8 U Overturned anticline 4/ 1 “ El ‘ . ’ I _ -‘ I phyry, and hornblende granodiorite Includes alaskite porphyry granite & LL] Showing approximate trace of axial plane, direction of /I/ l i l \ \ ' ‘ ‘ ' porphyry. Biotite-quartz latite por h d t , ' 4.» - - - - 0: dip of the limbs and bearing and plunge of axis . . . . . 5 ,i K/ ,, ;;.i’ , _ , 5 . porphyry is younger than bostonite 10h yry, an quar z monzonite fi Cordlerlte-amphlbole gnelss 0. short dashed where inferred ’ Note: Fold axes are based upon attltude of foliat1on; fol-lation _. I‘ ' ' ‘ group. Biotite granodiorite por- p orp yry g Olive-gray, medium— to coarse-grained, layered rocks; is presumed to be parallel to original sedimentary beddlng 26, c . > * I phyry and hornblende grano- 3 contain cordierite, garnet, gedrite, cummingtonite, /T‘\\\— diorite porphyry are older than S quartz, and plagioclase /Q( 25 . l . . . . 5" eucocratic granodiorite group ‘ E“ Overturned synCllne Leucocratic granodiorite group E) Showing approximate trace of axial plane, direction , _ c: of dip of the limbs, and bearing and plunge of axis Includes alkali syenite porphyry, "‘ albite granodiorite porphyry, and Amphibolite \ leucocratic di 't h . . . . \ grano am e porp yry Dark-gray or black, medium-grained rocks containing \ 30 . \ hornblende and andesine with minor amounts of Plunge 0f f01d 3X13 A p ‘ . mb quartz, Includes both massive and layered varieties; f . . I q locally intercalated with calc-silicate gneiss —€—> 3 J ' “ . . . , Plunge of minor anticline , - B K Blotite—muscowte quartz monzonite 23 xx . , . . ' . ' y , \ . I‘ » , ' ' 2 Gray or pink, fine- to medium-grained, weakly foliated gnm 3‘ x ‘ 7 , 0-“ Blackhawk Peak , , , . f p. \ . , . . ‘ \ j “A _ W intruswe quartz monzonite;fine-grained phase is . 5° \5 _ 65/40'6' ‘ f . , ‘ ’ ' , ‘ , , ‘ , F \J/ 3 "9 eqmngflgle} ”lbegmm'gmmed Phase is sen-ate Microcline-quartz-plagioclase-biotite gneiss Plunge 0f minor drag fOId P D Lowman Jr with additions , ., has . , ‘ ,1 _ ‘ , , y w I ‘ \, to _ M‘ Q 950 porp y” w no u es some pegma ite Light-gray or yellowish-gray, medium-grained rock; 69‘ ' ' by p'. K_ Sims .‘ , .. , ‘. 27; . ‘ , ‘ . , I , ‘ I ‘ -, I , , T -‘ 8 § , generally has a well-defined layering; has a granitic 2 / x' 962/ [a _ 39/1/ 5." a , . ” .I j , 1 ’ , . ‘. . ' . l 9‘ ' :6 \ appearance 1 Plunge of minor recumbent fold TILE \ / “ I ‘ 4 . '_ \ . I a 7‘ .‘ i I - 9% Q d’ ' d " bl K45 0' J' Gable , , y ” , -. 52 I 30 . . . . 3 J . ‘ . I. ‘ r , 3.4 uartz write an horn endite - - ‘ ' *I , I. ~ ’ . .. , . . . . . enerahz n of crenulated foliatlon \ . , sum, 34/ \.: zJ , ‘ y , . , ‘ ' 7' . ., , ’._, , 7 . > . 8 Black or mottled black and white, medium— to coarse- G ed strlke a d plunge \ ’ , . \V , , . . , , ‘1,” 4 . -. . L 9, ‘ ., . > I» ' . ' . ‘25 . ' ‘ y» . \0‘ _‘ G grained, nearly massive intrusive rocks. Includes <2: \\65 50, ‘ . .p I 2 . P l ‘ ’ ' . l ‘ 5 a °,a;°’o‘ ,4 , ' “We .7 . ’ ‘ A. I‘ - ,, , . 4O ' . j _ " ‘ ’7'“ ‘v 1,1,3," E? small local bodies ofgabbro and associated rocks, —- _ _ _ _ ’ ‘ ' , ' \ i ‘ , , , » ’ , , a 50‘s" , r. “ 3 7 , , 5I _ I “ . . " ' , , ,. “*50’ E granodiorite and associated rocks, and pegmatite FD Strlke and dlp 0f £011at10n ' 5-1 ' ' 0) 5° :2 >3 \ , g" 1: g Q Strike of vertical foliation and 8 77’ g E? \ A. E. Dearth .E“ Gabbro and related rocks I]. a P. K. Sims, 30 S < Dark-gray or brownish-gray, medium- to coarse— Bearing and plunge of lineation A. A. Drake, Jr., , $ grained, massive, equigranular, intrusive rocks and Eklvgégfomr j E ranging from melagabbro to quartz diorite. Includes 55527 [1? _ , k _ 6 8 abundant Pegmame Strike and dip of foliation and bearing and plunge \ :_/ ’9 ‘5 _ I, I_ o\\ ‘5 of lineation T 2 S fry? 755/90_Missouril ,, :l ' ' _ , ' '25.! * «~‘Il\‘+90Falls‘ 335481;? 8 \f0 TBS, _ o awe 803;, IN”: «WI: ”1 CCH l dFBM o I ~ 35 - , " 2,; ' ‘ ' Strike and di of foliation and horizontal lineation ' ' aw ey 3“ ' - °°'e on 03:33; 3\ 35' 39/ “1/ % Granod10r1te and associated rocks p R- H- M°e“°h and n ’ / ' 5' . . . , _ 913:9 6:200:10 50V * ’8 , . , 8 Gray, medium—grained, nearly equigranular, intrusive A' A‘ Drake' Jr' °° ,/;°°°°:°‘»'°°°, “>9 rocks ranging f7‘0m quartz diorite to quartz mon— 13 ° Q 6°°°°°°°°Qa°o°° i / ‘55 zonite. Dominantly foliated, but interior parts of 9 ° [69°u°°°° 0°? 0:5: '. E larger bodies are nearly massive. Includes some top . ”a ”'3 °° ‘3 ”3 l ‘E \ pegmatite INDEX TO GEOLOGIC MAPPING "a, ’. accountant; on 1\ D—l an o OVA o°u ° ' J’\ 90$ 0 5 , 9: 0° 2 0 , ‘l i °°°°8 IO o goo, \ I12 / a o °3 ‘ ' ‘ a 6 5° . 4 ‘ .359“. . I,» _ 37‘” flanHlH g “I. o , . , \ Lu 3 ' _.__ ”I I ’ _ , I, n.38 z 0 I: . A , . 4 , _, 7‘ A j .2— g I: LL] 5 ~_\\\AI LU ’ ». f V ' i gnbc % O I: E Z 3 “ is l ‘ 3/ ; 11,0001 ,, gnb > 5 0 F 8< ~ 11,000' ; Ymg gnm i I 0’ miii ELK CREEK 3 < L'- > g > 7 25 (X) PLUTON LL g X t m 3 . 32 ’8 , 3 g; obi 3% £9 10,000 ! 0 :’ < / g9 — 10,000' i gnb E b 3 ~ / E; ' ' ‘ gnb gnm gd gnb 8d x 8< 9000' ‘ H i “ Qa gnm 9000' ' gnb qdh Qa gnb gnb gnb d gnm gnb Tqm am . l x . - g /j/ gnm m gnm 8000’ / - 8000' .N Z i B l B l I a; _ , - , ._ n~ "t 12’000 g Lu: m 12,000 : Centratimh {I Yankee Hill 3 Lu )- 53 ”J E I Raj/”x nb 0 Z < L” o < z J ." 83‘”! “lb ”A * , g i j d2 5 6? 0m 2 f. ft ‘ T * ‘ X I/ (l3 -. . U’ i Lu 2 _l 0' < ’/ > i O L KL.) “3gnrri ‘K °-< _< ’— ' t 65 j 7 l ‘ ' 10 000 ’ i ' A ‘ ‘\ am ‘ g E U rib 2, ._: T, I , . . ’ ‘ I‘ v . I If? :34? i . - ~ MT. PISGAH PLUTON A 4 — 10,000 g H49 ,. Qatari l i ‘ * * « ,i \ PQJQ¥Q\§ ‘ ' ’ ‘ gd I; gnb gd . i . gd 5 ’ ’ E .7 y (DIE I. ,‘ ‘ . . ~ , , VJ\/\f\ _ . _ . ‘ . Tb Q .— 47/3o~~ gnb / ' 3n ,2)" « 47,30” \ ” , ' ' ‘ Z . ’ 65 ‘ ' “J 9 .5 . , Tb gnm am 0 ' 76 . ,1” I‘ 9000 l 5 l gnc — 9000’ ' j ; i l Tb gnm ‘ a /‘ am 8000’ i i I! _ 8000’ 7000' _ 7000, 0 , 0/ > E m z L‘ In 3 z 2 0 g o 2 2% id C E . g g g E P E C ’ \ 11,000! 2 5F 2 22 53 .‘5 —11,000' 0 z _I < 0 5 0 z N am m 0 3 '2 2' 8 E ' gnm E 2 ‘1: g m 2 Ci 10,000' " am 5 [>3 3 Ed 1/ r . g 5 —10,000' D n_ - , ,/ —— < gnm :‘ < E ' , i ; gnb a , 9000 .. ‘ 7 9000 8000’ 8000’ 2 Lu _ g I Z If a s a e; ‘2) z '3 l: 5 O 3 >- o 3 Z 3 > 2 0 m 0 < < < t In I) z D 2 < U- l- LL E u g a 5 m D , I 0 ; Z m < Lu 4 _l < 0 10,000— m E O 0 '22 < (_) g z m 10.000’ 3 0 a E E D j E '— L“ E E j :1 < a CE 0 U 2 E '1 CL .J gnm 3 a- E Z N gnb U) £_) m D > U , o l- 9000,_m O J u) I Z 9000, I < . < < < m ‘I 0 45 9 _ . — 5/ """ gnm 52 8000— Qa 8000/ . 7000’— 7000' ‘ ' ‘ ‘ ' m i 39 p < > F ’ . ‘ 045’ 105" 31'»; R. 74 W R. 73 W. (/DAHO SPR/NGS) 0/105: 30’ INTERIoR—GEOLOGICAL SURVEY, WASH'NGTON, D. c.—1967—056459 PH Base by U.S. Geological Survey, 1942 ‘i ”’5’ , SCALE 1:24 000 Geology mapped in 1952—54 and 1959—60 by P. K. Sims. (300 E 1 in 1952—54 by A. A. Drake,Jr.,C,C.Hawley, R. H. 17% E § ECZEfiZ—fir ]__, O 1 M‘LE Moench, F. B. Moore, E. W. Tooker, and J. D. Wells, 4 j s9 and in 1958—60 by D. J. Gable and P. D. Lowman, Jr. 75;; I z u .- <5 , g 1 .5 O 1 KILOMETER .- APPROXIMAIE MEAN W DECLINAYION. 1967 CONTOUR lNTERVAL 50 FEET DATUM IS MEAN SEA LEVEL GEOLOGIC MAP OF THE CENTRAL CITY QUADRANGLE, COLORADO “C SEP121967 ” ’7 $9 \Euffib‘?’ The Internal Magnetization of Seamounts and Its Computer Calculation GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—F - ) 1,: ’9 The Internal Magnetization of Seamounts and Its Computer Calculation By BERNARDO F. GROSSLING SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—F The triple-field method of calculating the internal magnetization of a seamount, or other geologic body, from its shape and its magnetic anomaly UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTONzl967 UNITED STATES DEPARTMENT OF THE INTERIOR STEWART L. UDALL, Secretary GEOLOGICAL SURVEY William T. Pecora, Director For sale by the Superintendent of Documents, US. Government Printing Oflice Washington, DC 20402 — Price 75 cents (paper cover) CONTENTS Page Symbols ___________________________________________ FIv The computer method—Continued Abstract ___________________________________________ 1 Biased-field anomaly ____________________________ Introduction and acknowledgments ___________________ 1 Magnetic field expressed in x, y, and z polarizations- _ Seamounts and their magnetization ___________________ 2 The triple-field method of determining J ___________ The nature of seamounts ________________________ 2 Total and remanent magnetizations _______________ The magnetization of seamounts __________________ 3 Input and output of the program _________________ Thermoremanent magnetization of lavas _______ 3 The magnetizations of Maher, Boutelle, and Hoke Sea- The magnetization of a submarine volcano _____ 4 mounts __________________________________________ Virtual paleomagnetic poles __________________ 5 The setting of Maher, Boutelle, and Hoke Seamounts_ Importance of the magnetization of seamounts__ 5 The computer calculation of the magnetization of The computer method _______________________________ 6 Maher, Boutelle, and Hoke Seamounts __________ General features of the method ___________________ 6 Discussion of results ____________________________ Scheme of mass approximation ___________________ 6 Proposed procedure to interpret the data of many sea.- The field of the elementary prism _________________ mounts ______________________________________ Determination of the magnetic field with the Gravity 9 Conclusions ________________________________________ Slave ________________________________________ 9 References _________________________________________ ILLUSTRATIONS [Plates are in pocket] PLATES 1—3. Contour maps showing bathymetry, computer model, and hypothetical and observed magnetic fields for the following seamounts, northeastern Pacific Ocean: 1. Maher Seamount. 2. Boutelle Seamount. 3. Hoke Seamount. FIGURES 1—11. Diagrams showing— Page 1. Method of locating virtual pole ______________________________________ F5 2. Grid of stations and prism array _____________________________________ 7 3. Prism array and mass description in row ______________________________ 8 4. Small magnetized element __________________________________________ 10 5. Biased-field anomaly A _____________________________________________ 12 6. Main steps in the calculation of A ____________________________________ 12 7. Small magnetized element and system of reference _____________________ 13 8. Relations between J, K5, and J," when 6—B< 180° and J is east of a"--- 15 9. Fracture zones and locations of Maher, Boutelle, and Hoke Seamounts--- 16 10. Loci of north virtual pole for Maher and Hoke Seamounts ______________ 21 11. Locus of north virtual pole for Boutelle Seamount _____________________ 22 TABLES Page TABLE 1. Spherical harmonics and specialization of the body _____________________________ F9 2. Computer solution of the total magnetization of Maher, Boutelle, and Hoke Seamounts ______________________________________________________________ 19 3. J ,. and locus of north virtual pole ____________________________________________ 20 III Page F11 12 13 15 16 16 16 18 20 23 24 25 3.939;.99-sua {owe gqscmwawue s §§§ “NI- ‘fi (5) e mmnmssm@>xk assets. Pry”? SYMBOLS x coordinate of field point P. y coordinate of field point P. a: coordinate of center of gravity Q of a prism. y coordinate of center of gravity Q of a prism. z coordinate of center of gravity Q of a prism. Field point. Center of gravity of a prism. Distance QP; Gravimetric potential. Magnetic potential. Universal gravitational constant. Magnetic susceptibility. Intensity of biasing field. Angle of inclination of the biasing field, positive below the horizontal plane. Angle of declination of the biasing field, positive east of north. Total magnetization vector (shown in lightface italic with an arrow above on illustrations). Angle of inclination of the total magnetization, positive below the horizontal. Angle of declination of the total magnetization, positive east of north. Remanent magnetization vector. Angle of inclination of the remanent magnetization, positive below the horizontal. Angle of declination of the remanent magnetization, positive east of north. Ratio J n/ (KS2). Elementary virtual displacement vector. x cosine director of As. y cosine director of As. 2 cosine director of As. Field intensity of the anomaly field. :2: component of f, similarly for Y and Z. a; component of I when the total internal magnetization is parallel to s, similarly for Y“) and Z w. Biased-field anomaly. Grid spacing between stations. a: dimension of a prism. y dimension of a prism. x coordinate of the center of symmetry of a prism array. y coordinate of the center of symmetry of a prism array. Center of top of prism array. 2 coordinate of the top of a prism array. Thickness of [0 layer; k=1, . . ., 10. Density of k layer; k=1, . . ., 10. Virtual density of k layer in magnetic problems; k=1, . . ., 10. Layer number; k=1, . . ., 10; Ic=1 for top layer. Row number; n=1, . . ., 20. In a given layer, row 1 is the first row in the positive y direction. The symbol n is used also as Sequential index as specifically indicated in text. Prism number; l: l, . . ., 20. In a given row, prism 1 is the first prism in the positive a: direction. First prism number of the first segment in a given row. Last prism number of the first segment in a given row. First prism number of the second segment in a given row. Last prism number of the second segment in a given row. Subscript, denotes number of a column in the grid of stations. Column 1 is that with algebraically smallest a: of the grid. The symbol iis also used to denote inclination of the total magnetization. Subscript, denotes number of a row in the grid of stations. Row 1 is that with algebraically smallest y of the grid. Order number of filled prisms. Number of filled prisms. Order number of stations in grid. Number of stations in grid. Order number of a region. Temperature. Time. Latitude. Longitude. Sum of squares of errors, in adjusting calculated to observed fields. IV SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY THE INTERNAL MAGNETIZATION OF SEAMOUNTS AND ITS COMPUTER CALCULATION By BERNARDO F. GROSSLING ABSTRACT Seamounts are veritable markers of the paleomagnetic field. Their magnetizations can be determined by an analysis of the magnetic anomaly in relation to the seamount shape. Probably most seamounts are thermoremanently magnetized. The method presented, which is not limited to seamounts, gives the total magnetization vector froma comparison of the observed field with three hypothetical fields obtained by assuming unit magnetiza- tions in three orthogonal directions. A least-squares fit of a linear combination of the three fields to the observed one gives the magnetization components. By examining the relationship between J and J" as a function of K, limits can be established for J ,, ; a locus for the virtual pole can be calculated as a function of K. For the computer model, ensembles of rectangular prisms are used. Four of the harmonic functions of the potential are taken into account. The derivatives of the magnetic potential are computed by a numerical differentiation of the corresponding gravimetric derivatives. The concept of “biased- field anomaly” as an algorithm built in the program gives the option of obtaining either (a) an arbitrary component of the anomaly field intensity, (b) the conventional total-intensity anomaly, or (c) the magnitude of the anomaly field intensity. As an illustration, the magnetizations of Maher, Boutelle, and Hoke Seamounts, located in the northeastern Pacific, are determined. In Maher, the magnetization is predominantly east; in Boutelle, 126° W.; and in Hoke, 17° E. With the deviations found, any rotation of crustal blocks in Maher and Boutelle Seamounts, due to crustal shearing along the Murray and Mendocino fracture zones, respectively, would be con- sistent as to direction of rotation. Alternative explanations of the deviations are also examined. A systematic analysis of the magnetization of seamounts in the northeast Pacific may be useful in deciphering crustal displacements and rotations. INTRODUCTION AND ACKNOWLEDGMENTS This paper discusses the magnetization of seamounts and describes its determination, using computer meth- ods that I developed in 1959—1960 While with the Chevron Research Co. (formerly the California Re- search Corp.). The actual analysis of the problem of the seamounts and the paper itself were prepared later, after I joined the U.S. Geological Survey. I wish to thank the Chevron Research Co. for its permission to publish this work and for its kind cooperation in the preparation of many of the illustrations. The bathym— etry and the total-intensity magnetic anomalies on Maher, Boutelle, and Hoke Seamounts were furnished to me by the Scripps Institution of Oceanography and are small detailed parts of the surveys by Mason and Raff (1961) and Vacquier, Raff, and Warren (1961). The bathyrnetry was collected in the course of several expeditions wholly supported under contract with the Office of Naval Research. The criticisms received from J. R. Balsley, R. G. Henderson, M. K. Hubbert, H. W. Menard, and F. J. Vine are sincerely appreciated. Finally, I am grateful to Professor Victor Vacquier for suggesting the problem of the seamounts. The work originated from a computer program for the calculation of the gravity field produced by a body of arbitrary shape. In this program, nicknamed the Gravity Slave, the body is approximated by sets of rectangular prisms. Taking advantage of the inher- ently high precision of a digital computer, I extended the method to the calculation by numerical approxima— tion of the derivatives of the field components. In particular, this permitted the calculation of the mag- netic field produced by a body uniformly magnetized. A variant of the Gravity Slave program aimed especially at aeromagnetic interpretation was then prepared. This program can calculate the anomaly field under different assumptions about the direction of the internal magnetization. Taking advantage of this feature, I then developed a method—named the triple-field method—to determine the direction and magnitude of the internal magnetization when, in addition to the shape of the body, the anomaly field is known and a uniform internal magnetization is assumed. In geology two different techniques have been used to determine the remanent magnetization. The one most widely used involves an actual measurement on a sample of the rock. The other—which in a certain way is an outgrowth of the interpretation of magnetic anomalies in prospecting—is based on the analysis of the magnetic anomaly produced by the geologic body examined. It consists of the calculation of the mag- netic anomaly for a model of the geologic body, under F1 F2 different assumptions about the internal magnetization, with the intent of reproducing the anomaly observed. The first technique cannot as yet be applied to sea- mounts because of the difficulty of obtaining oriented core samples in deep ocean waters. On the other hand, the second is facilitated by the existence, for certain areas of the oceans, of air- or ship-borne magnetic surveys and of detailed bathymetric maps. The results from the two techniques are not directly comparable. The measurement on a sample gives the remanent magnetization, whereas the anomaly analysis gives the total magnetization; that is, the remanent plus the induced magnetization. Further, the anomaly analysis provides a kind of average of the direction of the geomagnetic field throughout a geologic body. The direction and intensity of the remanent magneti- zation usually vary throughout geologic bodies. Slump- ing, flow, tectonic deformation, and other factors may cause such variation. To obtain an accurate value of the average direction for a complex body, such as a seamount, would require a large number of samples taken not only near the surface but also in the interior of the body. The practice of removing unstable com- ponents of the remanent magnetization in the laboratory is another source of discrepancy between the two tech— niques, for the analysis of a magnetic anomaly refers to the condition as it occurs in nature. The average obtained by anomaly analysis would be most meaningful when the main component of the geomagnetic field has varied only slightly in direction during the period of acquisition of the magnetization. If the remanent magnetization changes, especially in direction, in an erratic or too complex a manner through- out the geologic body, the average may not be signifi- cant. If the geomagnetic field has been reversed numerous times during the period of acquisition of the remanent magnetization and if the body consists of several layers or regions—such as a volcano—~which are magnetized in alternate directions, then the anomaly field would be weak and indistinct; for the fields of successive layers or regions would practically cancel each other, and the anomaly analysis would fail to provide a meaningful average magnetization. But if the anomaly is distinct and can be reproduced accurately by the modeling, then it can be inferred that the magnet— ization used in or derived from the calculation is prob- ably meaningful. Several papers have been published on the computer calculation of the magnetic anomaly produced by a given body when the direction of magnetization is also given. Thereby the direction may be determined by a trial—and-error procedure. We will mention only some of the more recent ones. Bott (1963) calculated the magnetic anomaly by two alternative methods, one SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY using a surface integration and the other a volume integration. Morgan and Grant (1963) calculated the gravity and the magnetic anomaly for two-dimensional bodies by approximating the cross section by a polygon. Talwani and Heirtzler (1964) used the closed expression for the magnetic anomaly caused by a two-dimensional polygon. Bhattacharyya (1964) used a closed mathe- matical expression for the total-intensity anomaly corresponding to a rectangular prism of infinite vertical sides. Several papers have been published on the investiga- tion of oceanic features with the magnetic method. Press and Ewing (1952) computed theoretical magnetic anomalies for typical two-dimensional oceanic struc- tures. Laughton, Hill, and Allan (1960), using a one- prism model and assuming magnetization parallel to the present geomagnetic field, analyzed the magnetic anomaly of a seamount north of Madeira. Also, Gerald Van Voorhis (in US. Naval Oceanographic Ofiice, 1962) and Van Voorhis and Walczak (1963) investigated the magnetization of seamounts with the computer program utilized by Vacquier (1962b). In the triple-field method of analysis of an anomaly, the magnetization is directly calculated, that is, without trial-and-error. Vacquier (1962b) and Henderson and Allingham (1964) used it to determine the magnetiza- tion of two buttes in Montana. Their model is a set of rectangular prisms. The magnetization of each prism is approximated by magnetic particles located at the center of opposing faces of the prism. By comparison, in the method to be described in this paper, the first four terms of the potential expansion are used, which allow the use of relatively larger prisms for the same relative error. Another difference is that the magnetic components will be calculated by a numerical differen- tiation of “virtual” gravimetric components. This paper begins with an analysis of how a seamount may become magnetized. Then the Gravity Slave program is briefly described, and the theory of the triple-field method is presented in extent. As an illus- tration, the magnetizations of three seamounts off California are determined. Finally, the discrepancies between the direction of the internal magnetization in these seamounts and the geomagnetic field are discussed in relation to the geology of the northeast Pacific, where the three seamounts are located. SEAMOUNTS AND THEIR MAGNETIZATION THE NATURE OF SEAMOUNTS The oceans are dotted with submarine mountains. Inside the andesite line, the bedrock in all of them is basalt. In the Pacific Ocean the total number of these submarine mountains is estimated at about 10,000 (Menard, 1959). About 2,000 have already been found, INTERNAL MAGNETIZATION 0F SEAMOUNTS AND ITS COMPUTER CALCULATION of which about 1,000 are in the Baja California sea- mount province, delimited by the Murray and Clarion fracture zones (Menard and Ladd, 1963). Some of them, designated as seamounts, have conical shapes, and are either active volcanoes or believed to be extinct ones. Some seamounts have flat tops, which Hess (1946) interpreted as wave erosion at a time when they stood with their tops near the ocean’s surface. Since then, these flat-topped seamounts, or guyots as desig- nated by Hess (1946), would have subsided. Nayudu (1962) suggests that some of the flat tops may be primary features. If this were true—that is, that some guyots were not subsided seamounts—then the oceanic crust would have a long-term strength at least sufficient to support them. Because the evidence about their age is scant, the limits for their time of formation are too broad. The intense weathering of oceanic basalts has precluded radioactive age determinations. Heezen and Menard (1963) indicated that Cretaceous fossils are the oldest sampled from seamounts. Menard and Ladd (1963) believed that either all seamounts are post-Paleozoic or that in Paleozoic time the seamount volcanism was relatively weak. This volcanic activity would have culminated in late Paleozoic to early Cenozoic times. Erben Guyot (lat 32°50’ N., long 132°32’ W.), which became submerged by Miocene time (Carsola and Dietz, 1952), is about 300 nautical miles northwest of Hoke Seamount, one of the seamounts studied in this paper. Hamilton (1956) discussed guyots some 600— 1,100 miles west of Hawaii and indicated that they formed a chain of islands in Cretaceous time. THE MAGNETIZATION OF SEAMOUNTS THERMOREMANENT MAGNE’I‘IZATION 0F LAVAS Because of certain properties of their constituent lavas, seamounts are actual markers of the geomagnetic field which existed at the time of their formation. The magnetization of a lava is seated in the iron oxides and sulfides which exhibit ferromagnetism. Their mag— netization disappears at the Curie point temperature. Above the Curie point the mineral becomes paramag- netic. The most important Curie point is that corre- sponding to magnetite, namely 578°C. Some lavas exhibit a second Curie point at about 250°C. A lava placed in a steady magnetic field 3 acquires, during its cooling below the Curie temperature, a remanent magnetization Jn which is designated as thermo- remanent magnetization (TRM). This has been shown both by laboratory experiments and by measurements on field samples (Nagata, 1953). For example, recently cooled lavas acquire magnetizations which are accurately parallel to the direction of the F3 present geomagnetic field. Measurements on recent igneous rocks, including lavas, show that their remanent magnetizations are parallel to the direction of the axial geocentric dipole, thus the secular variation is averaged out. The TRM is of great stability. The TRM is progressively acquired as the tempera— ture falls below the Curie point or points. Nagata (1953, p. 142) defined the TRM as “the remanent magnetization after field cooling throughout in a weak magnetic field from T to 0°C.” Below about 100°C, however, the magnetization is increased only by negligible amounts. Furthermore, most of the mag- netization is acquired when the temperature is in the interval from the Curie point to the Curie point minus about 100°C (Nagata, 1953, p. 146). Nagata (1953) defined a function P=§2415 <1) designated as “characteristic function of TRM,” which is the rate of production of remanent magnetiza- tion per degree centigrade per unit field applied. Here H denotes the field intensity, J the thermoremanent magnetization, and T the temperature. In this treatment, T is taken as the temperature increasing with time, for Nagata is assuming reversibil— ity with temperature of the magnetization process. In the cooling process, however, T decreases with time. We assume that not only T is a function of time (t), but that H is also. If we replace T by T’ =Tc— T, that is, if we measure the temperature from the highest Curie point downwards, then we have the same sense of the t scales for the T' and H variables. Because of this, we can rewrite equation 1 in the form , -i 31. HT )—H OT, (2) Nagata found experimentally that the remanent magnetizations acquired in definite temperature inter- vals are additive; hence when H is constant throughout the cooling, T, J=Hfl P(T’)dT’. (3) He does not indicate whether this additive law would be valid for changes in the direction of H, which may occur during the cooling of the seamount. That is, we don’t know whether the following generalization of equation 3 J=LTIHP(T’)dT’ (4) is valid or not. F4 If H varies only in magnitude during the cooling, then T . ,2 J=—f H(t)P{Tc——T(t)} %dt. (5) l 1 Because % is negative and P(T) is positive, the direction of J is the same as that of H. Therefore, the TRM acquired by a rock is not simply determined by the time-average value of H during the period of acquisition of the TRM, but H is weighted by the characteristic function P(T). Many factors—apparently depending on obscure petrologic details—determine the function P(T). Hence, without having a sample of the rock, it would be difficult, if not impossible, to deduce which intensity the paleomagnetic field had when the TRM was acquired. THE MAGNETIZATION OF A SUBMARINE VOLCANO Because of its complexity, we can only visualize the major stages in the magnetization of a submarine volcano; that is, of a seamount. The following dis- cussion is intended as a sketch of the phenomena involved. Present-day magmas reach the earth’s surface with temperatures of about 900°—1200°C, depending on their composition. Hawaiian tholeiitic basalt magmas reach the surface with temperatures of 1050°—1200°C (MacDonald, 1963). At these temperatures a basaltic magma behaves like a viscous fluid, its viscosity being about 102—105 times that of water (Wentworth and others, 1945). As the temperature decreases the viscosity increases, and fluidlike motion ceases at about 800°C. Hawaiian basaltic lava shows a small amount of movement at about 760°C (MacDonald and Finch, 1950). By the time the highest Curie point, namely 578°C, is reached, a basaltic lava is already a solid. In the early stages of a submarine volcano, the structure and temperature distribution are probably of great complexity. , The outer layers of magma oozing at the surface of the incipient volcano would be quickly chilled and thus form a resistant stretchable skin. Flow may continue inside, with the formation of pillow lavas (Rittmann, 1962, p. 70—74). Eventually these outer layers are fractured and are repeatedly incorporated into the fluid by continued flow. The overall magnetic field of the volcano would be weakened by the disordered mingling of shell fragments and by the demagnetization caused by reheating of the frag- ments. The magma injected into and retained in the shallow volcanic chamber would bulge and distend the upper crust, and produce tear breaks followed by water invasion and magma outflow. The relative SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY importance of these two kinds of mechanisms of growth, namely chamber injection and outpouring of magma, is not known; however, the first might be the more important in submarine volcanoes because of the formation of a strong outer shell when magma is chilled by water. As the major eruptive phase comes to an end, the local irregularities in the temperature distribution would begin to disappear and the local pockets of magma would solidify. The neck, or dike-type feeder, would have solidified or have been severed and thus ceased the transport of hot magma from below. Thereafter, the flow of heat in the volcanic mass would be essentially controlled by the outer boundary condi— tions. At this stage in the history of the volcano, the points with temperature equal to the 578°C Curie point would form a bulber surface, inside of which the material is paramagnetic and thus has a weaker magnetization than the outer thermoremanently mag- netized zone. This zone would increase in thickness with further cooling. .The material still fluid would form a chamber surrounded by the shell of paramag- netic solidified rocks enclosed by the Curie surface. As the cooling proceeds this chamber would contract and finally disappear. With still further cooling, the Curie surface would continue to contract, until it would be reduced to a point and then disappear. An important consequence of the process described is the remagnetization of the crustal rocks underlying the volcanic mass. Any earlier thermoremanent magneti- zation of these rocks would be lost to the extent that they are reheated. On reaching the Curie point the earlier TRM would be completely lost, and then TRM would be reacquired when the temperature falls again below the Curie point. Hence in making a computer model of a seamount, allowance should be made for such a root; however, we have not yet taken this effect into consideration in our calculations. The time required to dissipate the heat of a sub- marine volcano is but a brief episode in comparison to geologic time, which encompasses billions of years. A simple model can serve to set an upper estimate for the order of magnitude of the cooling time. Let us con- sider a homogeneous infinite medium in which, at t=0, the temperature is T: To in a spherical zone of radius R of the medium, and T=0 outside the zone. Such a model will overestimate the time required for the cool— ing of a pocket of lava or of a magma chamber. First, the presence of the earth’s free surface shortens the cooling time with respect to an infinite medium. Sec- ond, heat is transported while the lava is fluid not only by conduction but also by convection, which is a more vigorous mechanism. In the model just described the center of the spherical zone remains as the highest INTERNAL MAGNETIZATION 0F SEAMOUNTS AND ITS COMPUTER CALCULATION temperature point through the entire cooling process. The temperature Ta at the center, at time t, is given by the following equation (Ingersoll and others, 1954, p. 141): Ta: To[d>{ a5/(2x/ozt) }-2wZ(x)], (6) where x=R/(2‘/c_&), a is the thermal diffusivity, ¢ is the probability integral, and Z(x) is the normal prob- ability density function. In equation 6 the nomen- clature of Ingersoll, Zobel, and Ingersoll (1954) is retained, although some of the letters are used else- where in this paper for other purposes. As for the thermoremanent magnetization, it is not necessary to consider cooling below about 100°C, that is, Ta=0.1To because T0=1050°—1200°C. The equa- tion gives Ta=0.1To for x=0.765. If we assume a: 0.010 in cgs (centimeter-gram-second) units, then t= 790,000 years for R=1 kilometer, t=7,900 years for R: 100 meters, and t=79 years for R=10 meters. This demonstrates that the heat of small pockets of lava would be dissipated, at most, within a few decades and that, even for a magma chamber, not more than about 1 m.y. (million years) have to be considered. Because the ages of the three seamounts which we have been studying are probably of the order of tens or even hundreds of millions of years, their complete masses should be well below the 578°C Curie point and thus can be thermoremanently magnetized throughout. This would be true even if the volcanic activity of the seamounts had extended for a long period of geologic time, as long as activity had ceased tens of millions of years ago or more. On the other hand, the time of cooling—when the thermoremanent magnetization is acquired—is long in comparison with the secular vari- ation. Thus this variation may be averaged out. Moreover, the fact that the magnetic anomalies of some seamounts may be accounted for by assuming uniform magnetization throughout indicates that the main di- pole field has been relatively stable during those inter- vals of time when the seamounts were cooling between the Curie point and about 100°C. VIRTUAL PALEOMAGNETIC POLES To utilize the direction of remanent magnetization of a seamount, it is necessary to relate this direction to the morphology of the geomagnetic field at the time when the seamount became thermomagnetized. At present, the nondipole part of the field is, changing rapidly. It is conceivable that at some time in the past the contribution of the nondipole part of the field may have been larger than that of the dipole part. As a first step, however, it is customary to calculate the positions of a north or a south magnetic pole under the assump- tion that the field was dipolar, an assumption 243—571 0—-‘67#——2 F5 neglecting the nondipole part. In the specific case of the seamounts, the features of the geomagnetic field with a life span shorter than about 0.1 m.y. would be averaged out and not be reflected in the overall magneti- zation of the seamount. Under this assumption and with the aid of spherical trigonometry (fig. 1), we can obtain the position of the virtual poles. The latitude ¢’ of the virtual pole is given by sin ¢’=sin <1: cos p+cos ¢ sin 12 cos 6", (7) and its longitude N by sin (A—A’)=sin 1) sin (Sn/cos ¢’, (8) where d) and A are the latitude and longitude, re- spectively, of the seamount; 6,, is the declination of the remanent magnetization; and p is the angular distance between the seamount and the virtual pole and, for a dipole field, is given by tan in=2 cot p, (9) where 12,, is the inclination of the remanent magnetiza- tion (Chapman and Bartels, 1940, p. 22, formula 35). IMPORTANCE OF THE MAGNETIZATION OF SEAMOUNTS A‘knowledge of the magnetization of seamounts can be valuable in the study of the relative movements of segments of the earth’s crust and of the crust as a Whole with respect to the earth’s magnetic field. The facts that seamounts are very numerous and that they are scattered throughout most of the oceans indicate the potential wealth of information which may be available. As seamounts are frequently distinctive features North Greenwich meridian FIGURE 1.——Method of locating virtual pole. F6 rising above the ocean floor, their associated magnetic anomalies can be analyzed in relation to their topog- raphy. This permits the determination of their total magnetization—that is, of the vector sum of the induced and of the remanent magnetizations—in the manner described in this paper. Changes since the time of cooling, in the position of a seamount with respect to the geomagnetic field, create discrepancies between the direction of the total and induced magnetizations. Seamounts might have formed at different geologic times. In the meantime, crustal movements—such as overall shifts of the earth’s crust with respect to the geomagnetic field, continental drift, and smaller scale tectonic displacements—may have taken place. An area of particular interest for an investigation of crustal movements using the magnetization of sea- mounts is the Pacific Ocean west of North America. The three seamounts investigated in this paper are in this area. Vacquier, Raff, and Warren (1961) have found strong indications of very large displacements along certain linear fractures of the oceanic crust. To reconcile such large displacements in the oceanic crust with the absence of, or with only small, displacements in the neighboring continental block may require either the postulating of other types of displacements, such as rotation of crustal blocks, or an entirely new conception about the origin and evolution of the oceanic and continental crusts. THE COMPUTER METHOD GENERAL FEATURES OF THE METHOD For the determination of the internal magnetization of the three seamounts, a computer program that I developed for gravity and magnetic calculations was used. Only the main characteristics of this program Will be presented here. The Gravity Slave program determines the gravi- metric or magnetic field produced by bodies of arbitrary shape. The particular variant of that program used has optional channels for the calculation of the vertical component of the field intensity, a quantity designated as “biased-field anomaly,” and the second vertical derivative of the biased-field anomaly. The biased- field anomaly can be readily specialized into either (a) the conventional total-intensity anomaly (Vacquier and others, 1951), (b) the component of the anomaly field on an arbitrary direction, or (c) the magnitude of the anomaly field. The field values are calculated at the intersections of a square grid of stations lying on a horizontal plane (fig. 2). This plane may represent the earth’s surface. The density, fer the gravimetric field, and both the SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY magnetic susceptibility and the remanent-magnetiza- tion intensity, for the magnetic field, can change step- Wise with depth. The internal magnetization can have arbitrary direction and intensity. The blocks of results are compactly stored in the computer. Despite that, various lengths of results can be handled. Thus in core storage the storage used expands and contracts like an accordion. As presently coded, the program can be run on an IBM—704, —7090, or —7094 computer. The program was designed to relieve the user of as much routine as possible. For example, it handles the conversion of units, and the results are plotted automatically. SCHEME OF MASS APPROXIMATION To approximate the body in the Gravity Slave, sets of rectangular prisms are used. To facilitate the coding of such sets of prisms, they are organized in rectangular prism arrays of 4,000 elementary prisms, which contain 10 layers each with 20x20 prisms. As many prism arrays as necessary, each of different dimensions and positions, may be used. If the approximation for a given number of prisms is to be most effective, different- sized prisms locatable at arbitrary positions should be handled. The dimensions E and F of the prisms, the thicknesses Hk, and densities (1,, of each layer can be chosen at will. The position of the prism array is specified by means of the coordinates A, B, and H, of the center G of the top of the prism array (fig. 2). The given mass distribution is approximated by specifying which prisms of the arrays are filled with matter. A particular prism is allowed to be either filled or empty. If there are no filled prisms in a layer, its density is made equal to zero. Then, for the layers containing some filled prisms, the “rows” containing filled prisms are identified. A “row” is defined as a set of 20 prisms lying side by side in the x direction (fig. 3). Each layer contains 20 rows. In a row, only one or two contiguous sequences of prisms may be occupied. N 0 loss of generality results from this restriction because as many prism arrays as required may be used. This system allows the representation of bodies with cavities and reentrants. We may think of each row as a skewer with meat pieces, such as shish kebab. Up to two sections of adjacent pieces of “meat” are allowed per “skewer,” with no more than 20 pieces per “skewer.” For the study of the problem and its computer handling, an 1312 system of coordinates is used, in which the x axis is horizontal and points eastward. The 3; axis is horizontal and points northward, and the z axis is vertical and points downward. The my plane is taken to correspond to the plane on which the magnetic observations have been made or to which INTERNAL MAGNETIZATION 0F SEAMOUNTS AND ITS COMPUTER CALCULATION F7 Y (50,50) /— _____ — _——7——7 / H / [PK/,1) / / (d1) H1 (d2) H2 (0'3) H3 (d4) H4 (d5) H5 (0’6) ((17) (d8) (d9) (0'10) H I; 20;.r JL FIGURE 2.—Grid of stations and prism array. F8 SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY \ \ ROW k, n EXAMPLE: (V1) (W1) V1 W1 = 006021 \ \ \ \\ \ ‘ \\ §\ \ \ \ \\ EXAMPLE: (V1) (W1) (V2 ) (W2) I/1 W1 l/2 W2 = 003010013023 FIGURE 3.~—Prism array and mass description in row. INTERNAL MAGNETIZATION OF SEAMOUNTS AND ITS COB/[PUTER CALCULATION they have been reduced. Although this system of coordinates is left handed, we have chosen it because this system avoids coding negative depths and because we preferred to lay out the :1: axis and y axis parallel to the east and north directions, respectively. THE FIELD OF THE ELEMENTARY PRISM To calculate the potential of the rectangular prism, the potential expansion in spherical harmonics Hn expressed in rectangular coordinates can be used, namely - U=§H=$03a1—+,, ff dhndsdndg“, (10) (v) where h, is a homogeneous polynomial of degree n in x, y, and z; p is the radial distance to the origin; and d is the density (Kellogg, 1929, p. 135). Here H” is not to be confused with Hk, which is used later in this paper to designate the thickness of the k-layer of an array of prisms. By using the potential expansion and not merely the point-mass approximation, the prisms may have a relatively large size, which thereby decreases the number of prisms required and the com- putation expense without introducing serious errors. The expansion of the potential of a point mass or charge in powers or inverse powers of the distance is well known in mathematical physics and geophysics. For example, Grant (1952) used it to interpret gravity data. He uses “reduced 2l-pole moments,” to deter- mine approximately the size and shape of a three- dimensional mass distribution required to produce a given gravitational field. As the degree of symmetry of a body increases, an increasing number of the harmonics are equal to zero. We may start by considering a body of arbitrary shape and then successively specialize its form. The more lower order harmonics there are that are equal to zero, the more accurate is an approximation using only the first ones which are not. In table 1 is summarized the situation for the first five harmonics, namely Ho—H4. The first harmonic is simply equal to M/p; that is, the mass divided by the radial distance to the origin. The second harmonic, namely H1, is zero when the origin is at the center TABLE 1.—Spherical harmonics and specialization of the body H0 H1 H2 H3 H4 Arbitrary body and arbitrary origin ..... M/p #0 #0 #0 #0 Arbitrary body and origin at center of gravity ................................ M/p 0 #0 #0 #0 Rectangular symmetry and origin at center of gravity ...................... MI}: 0 740 0 #0 Cubical symmetry and origin at center of gravity ............................. M/p 0 0 0 #0 F9 of gravity. Moreover, in addition to selecting the origin at the center of gravity, H220 if the body has cubical symmetry, and H3=0 if the body has rectangu- lar or cubical symmetry. The mass element used in the program is an homo- geneous rectangular prism, for which Ho=M/p, H1: H3=0, HfiéO, and H4¢0. The program calculates Ho and also H2 when certain circumstances indicated below are met. Thereby, the first four harmonics of the series expansion for the homogeneous rectangular prism are in effect included, namely U=A74+éé {E2(2:cz—-y2—22)+F2(2y2—x2—22) +H2<222—z2—y2> }, (11> where E, F, and H are the ac, y, and 2 dimensions of the prism. Because the calculation of the third harmonic entails many extra arithmetical operations, it is wise to avoid laying out rectangular prisms for which EzeH, when a cube should be used instead. Furthermore, when a prism is deep or offset enough With respect to a field point, the calculation of the third harmonic is shunted off and thereby falls back to the point-mass approximation. Specifically, it is safe to use the point-mass approximation when {:5 max (E', F, H), where g‘ is the depth of the center of the prism. DETERMINATION OF THE MAGNETIC FIELD WITH THE GRAVITY SLAVE The procedure employed to calculate the magnetic field applying the Gravity Slave program will now be explained because this program basically determines the gravimetric field. The magnetic potential W at a point P corresponding to a magnet of finite dimensions can be expressed, using vector notation, as the sum of a surface and a volume integral (Maxwell, 1904, p. 10, second equation), namely _ Lizdw+ divJ (a) 7‘ (a) do. (12) Here J is the magnetization vector at an interior point Q of the magnet; 71 is a unit vector parallel to the outward normal to the surface element dw; do is a volume element; r is the distance QP. The surface and volume integrals extend over the surface and volume of the magnet, respectively. The signs in equation 12 are the opposite of those given by Maxwell; by following Kellogg (1929), the potential of the unit—positive magnetic pole is taken in this paper to be — l/r instead of +1/r. F10 As usual in magnetic interpretations it will be assumed that the magnetization is uniform throughout the magnetized region being considered and that the fields produced by difierent bodies are superposable. When J is uniform, then div J=0, and the potential W reduces to the surface integral in equation 12. That is, the effect of the magnet or of the region being considered is equivalent to a surface distribution of a density equal to —J-7,. Let us now draw at some interior point 0 of the body a small surface Aw, the normal of which forms an angle 11 with the magnetization intensity J (fig. 4). Consider then the elementary volume Av defined by the tube of force drawn through the perimeter of Aw and by another section Aw’ at a point 0’ distant As from 0. The external fie‘d produced by this volume element is equivalent, from what has been said previously, to the field produced by a pole of strength Am=JAw’ cos 1” distributed on Aw’, and another one —Am distributed on Am. Hence the magnetic potential AW at a point P distant r from 0, corresponding to the element Av, is 1 r+Ar A1) 1 . AW—J E< —;)+h1gher order terms. The potential for the body is obtained by integrating over the volume, namely (13') a 1 : (v)Jb—8,(;)dv- Am=JAw'cos u’ —Am=-Jch05v FIGURE 4.—Small magnetized element. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY The integration in the previous equation extends over the variables 5, n, and 5' corresponding to a point Q of the element Av, which may be its center of gravity. On the other hand, the variables 2;, y, and z of the field point P are constant throughout the integration. 5 . . . The operator W in the integrand involves the coordi- nates E, 17, and 5°, because As is here an elementary displacement vector of the point Q. That is Leg- «Lug J: a as"? as *J 5.7+? a? (14) where Jx/J, Jy/J, and Jz/J are the direction cosines of J. As we have 6— J: a J 6 J2 a aT—("J‘ 275+?” 671+ 76—2 ’ (15’ then in the derivative 5—: <%) in the integrand of equa- tion 13 the differentiation can be carried out with respect to the x, y, and 2 variables, instead of E, n, and g‘, merely by changing the sign. By so doing, the differ- entiation operator can be taken outside the integral sign, namely 2) d1) W=—— J— 68 (a) 7' (16) The prime index has been dropped, for now the 6/68 operator refers to the x, y, z coordinates. It is now assumed that the body instead of being magnetized is occupied by a fictitious gravimetric mass distribution of density J/‘y, where 7 is the universal gravitational constant. The gravimetric potential U of this distribution is U=fJ@ an 1- which, after comparison with equation 16, gives bU which is known as Poisson’s theorem. A relationship similar to equation 18 is valid for any other derivative of the potential because aWn+m+l a bUn+m+l 9 bx"0y"‘bz‘=_5~s— bx"by"'bzl (1 ) which is obtained by successive differentiations of equation 18. Instead of carrying out analytically the partial differentiation of U with respect to s to obtain W, we INTERNAL MAGNETIZATION 0F SEAMOUNTS AND ITS COMPUTER CALCULATION carry it out numerically by means of a finite difference approximation, namely Wz—(Ur+As—Ur)/As. Furthermore, by including the divisor As in the virtual density, namely (1’ =J/(7As), the magnetic potential becomes WzUr—Urfla- (20) With respect to signs, the customary convention in potential theory (Kellogg, 1929, p. 52—53) has been followed; namely, the potential of the unit-positive gravimetric pole is taken as +1 /1'. That is, in a gravimetric field the potential U is equal to the work function and to the negative of the potential energy. For the magnetic potential W the relations are similar but have opposite signs. Moreover, by considering attraction to be positive, the field intensity is f= grad U for gravimetric problems, and f = —grad W for magnetic problems. At this point the problem needs to be redefined. Let us consider a body A in an a}, y, z framework of ref- erence, having a density distribution d’=d’(£, n, f). Its gravimetric potential U(a:—xo, y—yo, 2—20), ob- tained by integration on the 5, n, and g“ coordinates, is known. Some reference point 20, yo, 20, fixed to the body, B, is taken to represent the position of the body in the coordinate system. The magnetic potential W corresponding to the body A when magnetized with an intensity J ='yd’sl where s, is a unit vector parallel to J, is given by equation 18. The 0/08 operator, namely 6 b b b a—s=u1a+u2 6—y+u3 BE (21) involves the x, y, and z coordinates; that is, it implies shifting the field point. The values ul, uz, and ua are the direction cosines of 31. If now the 6% operator is to involve the x0, yo, 20 coordinates—that is, implying a shift of the body as a whole—then b D b b fbs”=_ul b—xo_u2 *bE—ua 52—0 (22) The double prime sign on the 8 denotes that the dif- ferentiation involves x0, yo, and 20. As an example, let us consider the calculation of the vertical component of the magnetic intensity, namely DU :Zgrav. Oz Differentiating with respect to s and considering equa- tion 18, we have F11 OW a mag=____=__ grav Z 52 68 Z . (23) If the 538 operator is to involve :50, yo, and 20, then Zmag_ a ZZI'EV 24 _ _ as; I y ( ) or by numerical approximation ZmzzraV—zm (25) In summary, if some kind of a device or software is available which can provide a gravimetric quantity such as the potential or one of its derivatives, the procedure to obtain the similar magnetic quantity is as follows: A fictitious mass is assumed of density d’ =J/(7As) occupying similar volume and position as those of the magnetized body. Then two positions, say 1 and 2, of the fictitious mass are considered, one being displaced —As/2 and the other +As/2 with respect to the true position. The desired magnetic quantity is the value of the similar gravimetric quantity for position 1 minus its value for position 2. The displacement vector As is chosen parallel to the magnet- ization J. For the magnitude of As a small value should be chosen, so that the variation of the geometry because of the shift is negligible and yet sufficiently large so the device can still sense with adequate accuracy the corresponding variation of the quantity. A digital computer is a most suitable device for such a purpose because of its high precision. If all the quantities are measured in the cgs system, the magnetic anomaly will be given in oersteds, which can be converted to gammas by multiplying by 105. But a scaling factor of 105 is already built into the Gravity Slave program because all the lengths in the input data are to be given in meters, and the results produced by the program are in milligals instead of gals. Hence, if the cgs system is used in the calculation of the virtual density and all the lengths in the input data are given in meters, then the Gravity Slave for magnetic problems gives the answers in gammas. BIASED-FIELD ANOMALY At this point I will introduce the concept of biased- field anomaly. Let us assume that the anomaly field f _, produced by the body V is immersed in a field 9 of the same nature, but of constant intensity and direction which acts as a bias (fig. 5). The definition of the biased-field anomaly is A=l§+fI—fl, (26) from which is obtained F12 BODY FIGURE 5.——Biased—field anomaly A. A: (92+ 12+ 29) cos (9)1/2—9, (27) where cos 0: (QzX—i— Q,Y+SZZZ)/(S2f). _> Here 0 is the angle between 9 and f, and X, Y, and Z are the components of f. The 2;, y, and 2 direction _, cosmes of (l are S2,/Q=cos a sin 6, Q,/Q=cos 0: cos B, (28) and Qz/Q=Sin a, where a is the angle of dip (positive downwards), and [3 is the declination (positive east of north) of 3 The biased-field anomaly can be readily specialized, becoming equal to several geophysically significant quantities. First, it becomes equal to the component of f parallel to the biasing field ii when 9 is large. Second, when (l is the geomagnetic intensity, and S2 >>f, the biased-field anomaly A becomes equal to the total intensity anomaly AT used in aeromagnetic prospecting (Vacquier and others, 1951). This can be seen by expanding in series the first term in equation 27, namely SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY A=f cos «Hg—5+ (29) As 9 increases, A—>f cos 0; (30) that is, A becomes equal to the component of f parallel to 5. Third, the biased-field anomaly becomes equal to f when 9:0, as can be seen from equation 27. Finally, when f is not negligible in comparison with Q, then A has to be used rather than the approximation AT=f cos 0. In the computer calculation of the biased-field anomaly A, corresponding to a magnetic problem, it is necessary to retain in memory the total value of each of the components of f for each of the two positions of the fiducial mass. The algorithm built in the program to do this is shown diagrammatically in figure 6. The quantities X”, Y T, and Z”, which represent the com- ponents of f for the set of prisms, are calculated for positions 1 and 2 of the fiducial mass. The components of the magnetic anomaly field are the differences of these values for the two positions. The program de- termines the direction cosines of S2 and fume, then cos 0, and finally A. MAGNETIC FIELD EXPRESSED IN x, y, AND 2 POLARIZATION 8 Let us consider a small element Av drawn about a point Q and magnetized in a certain As direction (fig. 7). The as component of the magnetic field at a point Position 1, gravitational Position 2, gravitational X’, W, 27 |*_ XT Y7 ZT I Magnetic X7 YT, 2’ “1r ”2: ”3 cos 9 A FIGURE 6.—Main steps in the calculation of A. INTERNAL MAGNETIZATION 0F SEAMOUNTS AND ITS COWUTER CALCULATION I ”P (X,_y,Z) // / / / / / A? / / @ 0 (Ear) ° y X FIGURE 7.—Small magnetized element and system of reference. P, obtained by diflerentiation in the A8 direction of the as component of the gravimetric anomaly, is X=JAv %——3(x5—$) {u1(x—£)+u2(y—n)+ua(z-§) }] r (31) and similarly for the y and 2 components, where J is the magnetization intensity; :0, y, 2 are the coordinates of P; E, 1,, and y those of Q; and ul, U2, and m, are the direction cosines of As. If the magnetization I were parallel to the :1: axis (that is, u1=1, u2=u3=0), then we would have Xm=JAv {1—W}i 1'3 1' 3(11— 1000—5) ——rT——’ (3?) Y”) = _JAU and Z“) =—JAU 3(2-{2‘523—5) The subscripts a), y, and z in parentheses indicate polarization parallel to either the x, y, or z axis, respec— tively. The comparison of equation 31 with equations 32 shows that X is a linear combination of X”), Ym, and Z“). This relation, together with the similar ones for Y and Z, are X(s)=u1X(1)+u2X(ll) +u3Xm ) Y(s)=u1Y(z)+u2Y(y)+u3Y(z), (33) and Z0) =u1Z(,)+u2Z(v,+u3Zm. F13 The subscript s in parentheses, on the X, Y, and Z components on the left side of equations 33, denotes magnetization parallel to As. Now the biased-field anomaly A can be expressed in terms of the x, y, and 2 polarizations. When the —) magnitude of the biasing field 9 is much larger than the magnitude of the anomaly field I, the biased-field anomaly reduces to equation 30, namely A=f COS 0=(92X(s)+QyY(3)+Q;Z(3))/Q. (34) If instead of one prism, we have a set of N prisms, we have A: (:2. ‘74 X8?+9u ; 133%.; 222;): (35) where the superscript n in parentheses denotes the n-th prism, thereby we assume that the prisms are num- bered 1, 2,. . ., n,. . ., N. As each of the nine components for a prism—namely X“), X m, . . ., Z(,,—are additive for a set of prisms, so are X (3,, Y“), and Z“) which are linear combinations of the first group. Hence each of the summation terms in equation 35 can be expressed in terms of summations of terms corresponding to x, y, and z polarizations. By doing this and grouping together the components for each polarization we have A=§,‘—l{sz. g X223+ 9. g Y223+922n1 2:23} +§{a. g Xzza+a ; Yaw. g 223:} +%{ :2. gxw s2. gnaw. ; 2:2: } <36) The three quantities within the braces in the above expression are the values attained by A when there is only x, y, or z polarization. Hence A=JzA(z)+JuA(v)+JzA(z); (37) where Am, Am, and A”, are the values of A for unit polarizations parallel to the three coordinate axes. It is convenient to define these quantities in terms of a unit polarization, or a reference polarization, because in that way their calculation does not require a previous knowledge of J, which is an unknown. THE TRIPLE-FIELD METHOD OF DETERMINING J Let us now consider the problem of determining the internal magnetization of a given body when its biased-field anomaly A“) has been measured at a grid _, of stations. The biasing vector 9, in this problem the F14 geomagnetic intensity, is given. It is assumed, first, that the body is uniformly magnetized throughout, J being the intensity of magnetization. Let X“) denote the measured value of A“), Where the superscript v in parentheses denotes the v-th station of the grid. We assume that the stations have been numbered 1, 2,. . ,v,. ..T. The quantity A0"; should satisfy equation 37. That is, Am is a linear combination of the Am, A0,), and A”) fields, the values of which can be calculated at each of the stations of the grid from the geometry of the body. Three quantities are unknown—namely J 3, J”, and J2. Hence, if the equation were exactly fulfilled, three observations of A would be enough to determine the unknowns. But in practice the equation will not be SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY satisfied exactly because of measurement and other kinds of errors. Furthermore, many more than three observations are usually available, which provide more equations than required. These considerations lead to the choice of the method of least squares for determining the three unknowns. The error at the v—th station is e.=X—A.—J.A22 —J.A22.—J.A22. (38) It is assumed that in addition to any error at each station there is a datum adjustment A0 to be made on the measured data TV“. The condition of least-squares errors provides four linear equations for the determi- nation of J,, J,, J,, and A0, namely A.T+J. Z‘. A22+J. 2 A22.+J 2 A22 =21“) A02 AEB+JzZl {Am }2+Ju2 A ESAE¥A+J2 Z) 1183A 22 =2 A22122 (39) A0$Ai1$+JxZAgiAiil+Jyz{Ag/lithe] 2A22.A 22=2A22A<2 AoZlA 83+Jz 2118M E§§+Jy 2A The magnetization intensity in terms of J,, J”, and J., is J = (Ji-l-JE-l-Jf)“, (40) its declination 6 is given by tan 6=J,/J,,, (41) and its inclination i is given by tan i=Jz/(Jf-I—Jfill/2. (42) The procedure can be readily generalized to the case when the body consists of 0’ regions of known form. Each region is assumed to be uniformly magnetized with certain magnetizations J,, which may differ both in directions and magnitudes. Let A23. denote the value at the v-th station of the biased-field anomaly produced by the a-region under the assumption of unit polarization in the x direction, and similarly for Au?) , and 118;... These values can be obtained by constructing a prism model for each region and by using the computer method previously described to calculate Am, Am, and A”). There are 3a+1 param— eters to be determined—namely A0, and J“, J”, 85A E23+Jz Z {A83}2= 23 A22?» ll JAM - - '1 sz Jum and Jav- of the errors is The sum of the squares ®=Z{X‘")—Ao—Z (J...A22.. 2 +JI/.0A2.13)).a+Jz.oAgl.v>} ' (43) The condition of least-squares errors provides (30+1) linear equations for their determination, namely bi) 2.13:0 a<1> b Exp _0 2.1“: 5.1.1: an. (44) be _ a<1> _ 64> _0 aJ,,._aJ.,.—E>J.,.‘ ' INTERNAL MAGNETIZA’I‘ION 0F SEAMOUNTS AND ITS COMPUTER CALCULATION TOTAL AND REMANENT MAGNETIZATIONS The triple-field method gives the total magnetization J, which is the sum of the remanent magnetization J" _) plus the induced magnetization KS2 (see fig. 8). Hence the determination of the remanent magnetization _, In: J -K9 (45) requires a prior knowledge of K. When J" is of thermoremanent origin, it is parallel to the magnetic field 5 at geologic time t, and its mag— nitude depends on the magnetic minerals present. It is also proportional to Q as long as S2 is not higher than about 1 oersted; the increment AJn, for a given incre- ment A9, decreases gradually with increasing 9. The ratio Qn=Jn/(KSZ) is usually different from 1 and fre- quently higher than 10, sometimes exceeding 100 (N a- gata, 1953). Hence KS2 is not always negligible in comparison of J". Now let us examine what information can be gleaned about Jn and Q” when viewed as functions of K, the _, quantities J and 9 being given. The magnitude of J", deduced from equation 45, is J,=(J2+K292—2K9J cos to”? (46) Here co is the angle betweena and J, given by cos w=(u192+1429y+’U/392)/9, (47) where (21/9, Sly/9, and mm, and u], U2, and U3 are the _) direction cosines of (z and J, respectively. The function Jn reaches the minimum value (Jn)mln=J Sin 0’ (48) for K=K,=(J/S2) cos 0:, and the function Qn=Jn/(KQ) reaches the minimum value (Qn)mln=SiIl a.) (49) for K=K2=J/(§2 cos 0)). These minimums are ob- tained from the equations dJfi/dK=O and dQEL/dK=0 because the squares of J, and Q” are minimized con- currently with J” and Q”. It may be possible, by considering petrologic infor- mation or magnetic measurements on samples, to set limits for the range of likely values of K, namely 0Kly then JnS(J2+K§§22—2K,,J cos w)”2 (53) when %(Ka+Ko)>K1, and Jug (J2+K§522——2KaJ cos w)!” (54) when %(Ka+Ko)K1, we have 2 2J cos 0) “2 Qns<1+ sz—‘m—sz‘ <1 (55) or J 2 2J c “2 Q"S(1+K_:QT_K:—w ’ (56) whichever is smaller. F16 Now let us study the direction of In using the unit sphere. In figure 8, the points P(fl), P(J), and P(Jn) represent the directions of Q, J, and J", respectively. The inclination in and the declination 6,. of J” are given by sin in=cos \p sin i+sin ¢ cos 7: cos c (57) and sin (6fl_5)=sm rlz cos a s1n(6—fl), (58) sin 0) cos in where a and B are the inclination and declination of Q, i and 5 are the inclination and declination of J, and e is given by _sin (6—6) cos a sm 6 sin w (59) and sin a: K——9 J31“ “’- (60) With the previous formulas, In and Q, can be calculated for given values of K or limits established to their admissible values. INPUT AND OUTPUT OF THE PROGRAM The different capabilities of the program are con- trolled by a simple metalanguage consisting of eight mnemonic statements. The name of each statement defines the capability referred to, and an accompanying number indicates what to do about it. The statements and the values of the parameters are recorded in sym- bolic cards, that are loaded into the computer together with the program deck. The input data consist of (a) description and position of the mass distribution in terms of the scheme of prism arrays previously described and (b) magnitude, inclination, and declina- tion of the biasing field. Also, in magnetic problems, the parameters A, B, Ho corresponding to the two assumed positions of the set of prism arrays have to be given. The shifting of the virtual mass is accomplished simply by ordering the program to compute the gravity effect for positions 1 and 2 and then by subtracting the effect for position 2 from that for position 1. The results are printed as maps, under the full command of the program, on an on-line or off-line printer. THE MAGNETIZATIONS OF MAHER, BOUTELLE , AND HOKE SEAMOUNTS THE SETTING OF MAKER, BOUTELLE, AND HOKE SEAMOUNTS As an illustration, I have applied the triple-field method to Maher, Boutelle, and Hoke Seamounts in the northeastern Pacific Ocean. These seamounts were selected after examining bathymetric and magnetic SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY maps at Scripps Institution of Oceanography. The determination of their directions of internal magnetiza- tion may be of significance to the study of several tectonic features of a rather unique character that have been found in the area where they are located. By studying the bathymetry, Menard (1955) dis- covered a system of linear belts of faulting which he denotes “fracture zones” (fig. 9). They closely follow arcs of great circles and are known to extend in an approximately east-west direction for 1,400—3,300 miles, from the western edge of the North American continent to about the longitude of Hawaii. Some of the belts recognized to date are Mendocino, Pioneer, Murray, Molokai, Clarion, and Clipperton. These belts have been compared to the Liiders lines which appear on sheared metals. Menard (1955) believed that they are of Mesozoic to Cenozoic age. On the other hand, Vacquier, Raff, and Warren (1961) believed that the Mendocino, Pioneer, and Murray fault belts have been quiescent since some time in the Paleozoic. An appreciable change of the regional depth of the ocean floor occurs at the Mendocino and Murray fracture zones. South of the Mendocino fracture zone 130°W 120°W T 140° w T «150°N MENDOCINO FRACTURE ZONE —\ o Boutelle Seamaunt , 440° N PIONEER FRACTURE ZONE NF. FRAcTURE 7'0 Hoke AY MU RR Sea mou nt . Maher Seamount MOLOKAI FRACTURE ZONE -420°N ONE A CLAR\0N FRACTURE Z FIGURE 9.—Fracture zones and locations of Maher, Boutelle, and Hoke Seamounts. From “Marine Geology of the Pacific” by H. W. Menard. Copyright 1964 by McGraw- Hill Book Co. Used by permission. INTERNAL MAGNETIZATION 0F SEAMOUNTS AND ITS COMPUTER CALCULATION the sea floor is about one-half mile deeper than to the north, and north of the Murray fracture zone it is about one-quarter mile deeper than to the south. Furthermore, a remarkable pattern of magnetic anomalies—in the form of virtually linear, alternately positive and negative anomalies—has been discovered (Mason, 1958; Mason and Raff, 1961; Menard and Vacquier, 1958). These anomalies extend in an ap- proximately north-south direction, maintaining a great regularity of direction and character between the major fracture zones. They have amplitudes of up to several hundred gammas and a ridge-to-ridge distance of about 10—20 miles. Anomalies with a similar regularity of pattern and covering areas as large have not been discovered in any continental area. The linear anomalies probably extend over wider areas of the Pacific than established until now. Their existence recently has been demonstrated by the US. Coast and Geodetic Survey in the north Pacific, south of the Alaska trench (Peter and Stewart, 1965). The pattern of the magnetics is clearly offset along some of the fracture zones, a condition suggesting a relative displacement of the two bounding crustal blocks. Because of the constancy of their shape along the linear trend, the anomalies can be correlated across the fracture zones (Menard and Vacquier, 1958; Vacquier and others, 1961; Vacquier, 1962a). Thereby Vacquier, Raff, and Warren (1961) conclude that there is left—lateral displacement of about 1,160 km across the Mendocino fracture zone, a left-lateral displacement of about 265 km across the Pioneer fracture zone, and a right-lateral displacement of about 154 km across the Murray fracture zone. A subsequent survey along the Murray fracture zone (Rafi, 1962) revealed that the displacement varies along the fracture zone and at places may be as much as 630 km. As for the origin of the magnetic pattern, which is important for the study of the seamounts and their magnetizations, two possibilities may be explored: (a) the pattern was formed after the oceanic crust came into existence or (b) the pattern and the oceanic crust or a certain layer of it originated at about the same time. In line with the first possibility, three alternatives have been proposed (Mason and Rafl’, 1961; Raff and Mason, 1961; Bullard and Mason, 1963) about the bodies causing the magnetic pattern— namely, (a) isolated bodies of magnetic materials within the “second layer,” (b) block faulting of the main crustal layer, and (c) intrusions of magnetic material. ‘ Along this first line of thought, the sequence of major events could have been as follows: The initial deforma- tion, which defined the locations of the anomalies, could have been either fracturing or folding by buckling F17 of the oceanic crust. Because of the regularity of the pattern, the oceanic crust should have been very uni- form mechanically at the time of the initial deforma- tion, and the causative stresses should have been uniform over the wide areas involved. The evidence about the age of this initial deformation is indefinite. It could have taken place any time before the Mesozoic or early Cenozoic and as far back as the Precambrian. If the cause of the magnetic anomalies is not merely block faulting of the main crustal layer, then either magnetic material (a) was intruded into the second layer, (b) extruded at some time during the deposition of the sedimentary components of the second layer, or (c) extruded on the first sedimentary layer. Submarine volcanoes could have been formed during this period of volcanic activity. Next, the fracture zones would have appeared, with the horizontal displacement beginning some time after the shear fracturing of the crust. Recently Vine and Matthews (1963) proposed the ingenious hypothesis that the magnetic pattern and the main oceanic crustal layer originated together. This layer (seismic layer 3) would have been formed above the center of an oceanic ridge by some kind of segrega- tion of mantle material, spreading laterally by the drag of an underlying convection current. Furthermore, they assumed that at the time the layer was being formed the magnetic field was reversing periodically, and thus the newly created crust acquired magnetiza- tion in strips of alternating directions as it spread away from the oceanic ridge. 7 The structural role of the fracture zones, when the Vine—Matthews spreading-floor hypothesis is accepted, becomes clear with the concept of transform fault introduced by Wilson (1965a, b, c). Transform faults would exist where the crust is absorbed into the interior and formed elsewhere and thus would explain the abrupt ending of many dislocations. Yet the offsets of the midoceanic ridges (Menard, 1960; Heezen, 1962) would be merely an expression of the shape'of the initial crustal breaks and not the result of strike-slip faulting. Talwani, Le Pichon, and Heirtzler (1965) and Vine and Wilson (1965) have further elaborated on the Vine-Matthews hypothesis. According to their views, the magnetic anomalies would be genetically related to the formation of the midocean ridge system in a manner that the magnitude of the anomaly at a given location would depend on the distance to the ridge axis. Hence, the anomalies should be symmetrical about the axis, and the pattern should be linear and parallel to it. A preexisting offset of the ridge would cause, With the spreading of the crust, an offset of the magnetic pattern. In the area of the northeast Pacific, Talwani, Le Pichon, and Heirtzler (1965) postulated that the axial area of the East Pacific Rise south of F18 the Mendocino fracture zone underlies the Basin and Range province and the Colorado Plateaus. Maher Seamountl is about 10 miles south of the Murray fracture zone and about 1,350 nautical miles from California (fig. 9). On the U.S. Coast and Geodetic Survey Chart BC—1506N (1955) it is shown as a double-peaked feature centered at about lat 29°30’N. and long 148°49’W. Boutelle Seamount2 is between the Mendocino and Pioneer fracture zones, about 300 nautical miles offshore (fig. 9). On the U.S. Coast and Geodetic Survey Chart BC—1407 (1953) it is shown at lat 39°1’N. and long 131°5’W. Hoke Seamount3 is about 100 nautical miles south of the Murray fracture zone and about 310 nautical miles offshore (fig. 9). On the Bureau of Commercial Fisheries Topographic Chart 2 (1964) it is shown at lat 32°8’N. and long 126°59’W. ' THE COMPUTER CALCULATION OF THE MAGNETIZATION OF MAKER, BOUTELLE, AND HOKE SEAMOUNTS For each of three seamounts—Maher, Boutelle, and Hoke—a model was built with rectangular prisms defined by prism arrays in the manner described in this paper. For each seamount the bathymetry obtained from Scripps Institution of Oceanography charts differs somewhat from that given in the published charts mentioned above. To construct a model, the seamount is sliced by a number of horizontal planes at elevations which correspond to given contour lines. Each slice is then approximated by an assembly of rectangular prisms with their top and bottom horizontal faces on the upper and lower planes defining the slice. The prism faces parallel to the x and y directions were set parallel to the magnetic east and magnetic north directions. The number of prisms per layer decreases upwards from layer to layer because of the shape of the seamounts. In this manner a staircase surface was constructed closely matching the actual surface. Because of the shape of the seamounts, most of the prisms used are nearly flat. For example, for Maher Seamount some prisms are 183 m (100 fathoms) high in comparison to a horizontal dimension of 2,000 m. The shallowest of these lies with its top surface at a depth of 3,109 m (1,700 fathoms) below the plane of stations. That is, all locations where field values are calculated lie on or beyond a sphere of radius equal to 2.26 times the radius of the minimum enclosing sphere I Named after Captain T. J. Maher, Commanding Officer of the U.S. Coast and Geodetic Survey ship Guide. In 1927 he surveyed Seamounts off the Hawaiian Islands. 5 Named after Charles 0. Boutelle, an Assistant in the U.S. Coast and Geodetic Survey in the 19th century. He worked on geodetic surveys, tides, and currents. 3 Named after Willliam E. Hoke, inventor of a system of navigation to provrdecompass corrections. He died in the 1920’s. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY of the prism. Yet for horizontally flat prisms, the percentage of error of the approximation scheme used is admissible even when the field locations are on a sphere of radius as small as about 0.5 of that of the minimum enclosing sphere. ‘ For the calculation of the hypothetical fields, the following values were taken: J=5>_<10‘4 in cgs units, which corresponds to 0:05 oersted, and K=0.001 in cgs units; 7=6.67><_ 10‘8 in cgs units; and As: 1,000 cm. This gives d’=J/(7As) =75 in cgs units. The values of 0 indicated in plates 1—3 are the actual values of the geomagnetic intensity at the corresponding locations. Values of the total-intensity anomaly (the biased- field anomaly in the theoretical part of this paper) were calculated for a grid of 16X16 stations centered over the model of each seamount, With a unit spacing 8:4,000 m in Maher and Hoke Seamounts and 8:8,000 m in Boutelle Seamount. The calculation, in an IBM—704 computer, of the four hypothetical fields—that is, for polarizations parallel to the magnetic north, magnetic east, vertical and to the present geomagnetic field—required 43, 56, and 30 minutes for Maher, Boutelle, and Hoke Seamounts, respectively. The magnetic anomalies over the three seamounts, which are used in this paper, were measured with total-field flux-gate magnetometers by ships of the Scripps Institution of Oceanography. Regional effects and short-time fluctuations have been removed. Maher Seamount (pl. 1A) is about 30 km long and 12 km Wide and rises about 600 fathoms above the ocean floor. Its model contains 73 rectangular prisms (pl. 13). As for horizontal prism dimensions, three sizes were chosen: 2,000><2,000 m, 2,000><4,000 m, and 4,000><4,000 m. The prisms occupy positions in six layers bounded by seven datum planes—namely 2,300, 2,200, . . ., and 1,700 fathoms. A simple comparison of the hypothetical and ob- served anomalies gives a first qualitative understanding of the manner of magnetization of the seamounts. The hypothetical anomalies for Maher Seamount, shown in plate 10—F, should be compared with the observed anomaly (p1. 1H). Plate 10 corresponds to a least-squares fitted field for Maher Seamount, which will be discussed later. The fact that the range of values for the observed anomaly is wider than for the hypothetical anomaly indicates that the apparent sus- ceptibility is larger than the value K = 0.001 in cgs units taken for the model. Because the shape of the ob- served anomaly (pl. lH) is different when a magnetiza- tion parallel to the present geomagnetic field is assumed (p1. IF), the direction of J should differ substantially _, from that of o. On the other hand, the observed and the east-magnetization anomalies, parts H and E of INTERNAL MAGNETIZATION OF SEAMOUNTS AND ITS COMPUTER CALCULATION plate 1, respectively, are more similar. Hence, in Maher Seamount the internal magnetization should have a strong magnetic-east component. Yet, no linear combination of the anomalies calculated, with the model assumed, can explain the sharp troughs on the north and south sides of the seamount. Boutelle Seamount (pl. 2A) is ovallike, about 42 km long and 22 km wide; it rises about 1,400 fathoms above the ocean floor and reaches within about 900 fathoms of the surface. Its model (pl. 23) contains 100 prisms. As for horizontal prism dimensions, the following sizes were chosen: 1,000><1,000 m, 2,000X2,000 m, 4,000>< 4,000 m, and 8,000>_<8,000 m. The prisms occupy positions in seven layers bounded by eight datum planes—namely 2,300, 2,100, . . ., 1,100, and 900 fathoms. The hypothetical anomalies for Boutelle Seamount are shown in plate 20-F. Plate 20 corresponds to the optimum-fit field, to be discussed later. The trough to the southwest of the seamount, in the observed anomaly (pl. 2H), is more pronounced than that to the north- west, yet the opposite occurs when the magnetization is parallel to the present field (pl. 2F). This means that J has an appreciable southwest component. Hoke Seamount (pl. 3A) is conelike, is about 20 km in diameter, rises about 1,500 fathoms above the ocean floor, and reaches within about 700 fathoms from the surface. Its model contains 67 prisms. As for hori- zontal prism dimensions the following sizes were chosen: 2,000><2,000 m, 2,000><4,000 m, 4,000x4,000 m, and 8,000X8,000 m. The prisms occupy positions in eight layers bounded by nine datum planes—namely 700, 800, 1,000, 1,200, . . ., and 2,200 fathoms. The situation for Hoke Seamount is presented in plate 3B—E, G. Plate 3F corresponds to the optimum- fit field, to be discussed later. Judging by the range of values, the apparent susceptibility of this seamount is about eight times the value K=0.001 in cgs units taken for the model. Because the observed magnetic high is about 5 km southward of its position when the magnetization is parallel to the present field, the angle of inclination of J should be smaller than that of 6, but both these vectors should have about the same declination. Subsequently, the internal magnetizations of Maher, Boutelle, and Hoke Seamounts were calculated with the triple-field method for arrays of 69, 57, and 64 fitting points, respectively. The boundaries of these arrays are shown with segmented lines in the figures. A Burroughs B220 program for multiple regression analysis was used for this calculation. Several factors had to be considered in laying out these grids. Obviously they could not extend much F19 beyond the area defined by the actual observations. Main highs and lows of the hypothetical and observed anomalies were to be included, whereas effects clearly attributable to neighboring anomalies were to be omitted. The fitting points were selected for each seamount from among the stations at which the three hypothetical anomalies had been calculated; thus the need for interpolation was avoided. In the contour map of the observed anomaly, the value was read off at each fitting point. For Hoke Seamount the laying out of the grid offered no difficulty, for its anomaly is distinct and well isolated (pl. 36’). For Boutelle it is difficult to leave aside the effect of the extraneous anomalies which appear to the northeast and southwest of the seamount (p1. 2H), and the grid had to be trimmed in these two directions. The difficulty is greater for Maher Sea- mount because the anomaly apparently merges to the northwest and to the east with other anomalies. The grid chosen (pl. 1H) is a compromise between the need to leave them out and the need to include essential features attributable to the seamount. The findings about the internal magnetization are as follows (table 2): The apparent magnetization is 3.85, TABLE 2.—Computer solution of the total magnetization of Maher, Boutelle, and Hoke Seamounts Quantity Maher Boutelle Hoke Seamount Seamount Seamount J (x104 cgs emu) 1. 89 0. 74 3. 91 Km, (XIII—3 cgs emu) 3. 85 1.48 8. 32 a, declination of] 91°30'E. 126°15’W. 17°10’E. —> findeclination of (1 13°50' E. 18°50’E. 15°40’ E. i, Inclination of J +46°15’ +24°30’ +30°55’ _) a, inclination of Q +58° +61° +56° —> an, angle between J and 9 46°20’ 89°55’ 25°5’ K1 (X10—3 cgs emu) 2. 66 1, 485 7. 53 (Jn)min (x10-3 cgs emu) 1.36 0. 74 1.66 K2 (>(10‘3 cgs emu) 5. 58 1. 44 9.18 (0n)min 0. 72 1.00 0. 42 1.48, and 8.32 in 10‘3 cgs units in Maher, Boutelle; and Hoke Seamounts, respectively. For Maher Seamount the computed J has a declination which is rotated about 78° E. and the inclination about 8° smaller than _, for Q. The vector J forms an angle of about 46° with _) Q. In Boutelle Seamount J has a declination which is rotated about 145° W. and an angle of inclination about 37° smaller than for 5. The vector J forms an —> angle of about 90° with $2. In Hoke Seamount J has _) about the same declination as 9, but its angle of inclina— n I —) u I tlon is about 25° smaller than for 0. Lower 11m1ts were calculated for J, and Q” by analyzing the vector _, relationship between J, J", and (2 as a function of K in F20 the manner described in this paper. Taking the results as a group, one sees that the Jn cannot be smaller than 0.00074 in cgs units, and the (2,, not smaller than 0.42. For Hoke Seamount the anomaly obtained by least-squares fitting (pl. 3F) agrees well in shape, magnitude, and position with the observed one (pl. 36'). In fact, it gives the best reproduction of the observed field of the three seamounts. For Boutelle Seamount the agreement is reasonably good as to the shape and magnitude of the anomalies (pl. 20, H), although the calculation gives negative troughs which are not as pronounced as the observed ones. For Maher Sea— mount the difficulty of isolating the anomaly of the seamount itself impairs the reproduction of the observed field, as can be seen by comparing plates 16' and 11-1. The sharp troughs on the north and south of the magnetic high are not reproduced. Probably a better result would be obtained for Maher Seamount if the seamount were assumed to consist of two adjoining bodies, side by side on an east-west direction, and to have different directions of magnetization. TABLE 3.——J,J and locus of north virtual pole K (X10-~1 cgs emu) J,.(X10-a 6.. i,I 42’ N cgs emu) Muller Seamount 1. 89 91°30’ E. +46°15’ +12° 83°55’W. l. 40 114°30’ E +22°20’ —l4°50’ 81°30’ W. 1. 36 121°45’E +10°55’ —24°14’ 80°35’ W. 1. 51 134°35’ E. —11°20’ —4l°10’ 78°45’ W. 2. 95 158°30’E. —4l°40’ -70° 5’ 69°35’ W. Boutelle Seamount 0. 74 126°15’W. —24°30’ —18° 173°15’E. . 76 129°30’W. —l3°30’ —24°35’ 171°35’E. . 89 135°15’W. +8° 20’ —36°40’ 167°55’ E. 1.24 140°45’ W. +27° 5’ —47°40’ 163°25’ E. Hoke Seamount 3. 91 17°10' E. +30°55’ +68° 5’ 3°50’E. 2. 70 17°35’E. +18" 5' +61°65’ 13°40’ E. 2. 04 18° 5’E. +1° 40’ +54°20’ 20°50’ E. 1.66 19°15’E. —33° ’ +36“ 30°15’E. 2. 02 20° 5’E. —49°10’ +24°55’ 33°55’E. Next, for each seamount the vector 1,, and the locus of the north virtual pole was calculated as a function of K. A suitable set of values of K was chosen, which includes the value K1 that minimizes J,,. The results of these calculations are shown in table 3 and the loci of the virtual poles in figures 10 and 11. DISCUSSION OF RESULTS Because of a seamount’s volcanic nature, the remanent magnetization of a seamount is probably thermo- remanent. The effect upon the magnetization of near— surface alteration is of little significance for the bulk of SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY a seamount’s mass. Hence, it can be assumed that the remanent magnetization is parallel to the geomag- netic field which prevailed at the time of the cooling below the Curie point. The values found for the total intensity of mag- netization J—namely, 0.74, 1.89, and 3.91, in 10‘3 emu (electromagnetic units) in Boutelle, Maher and Hoke Seamounts, respectively—appear reasonable for oceanic basalts. These electromagnetic units, and as used later in this paper, are in the cgs system. By studying the magnetic anomalies, Van Voorhis and Walczak (1963) found a value of 10X10‘3 emu for Kelvin Seamount in the northwest Atlantic. Vine and Matthews (1963), by studying the magnetic anomalies of volcano-like features in the Indian Ocean, have established a value for J of 5X10‘3 emu, and Km, =13.3><10‘3 emu. The lower limits calculated for the remanent magneti- zation J" (table 2) are 0.74, 1.36, and 1.66, in 10‘3 emu in Boutelle, Maher, and Hoke Seamounts, respec- tively. They are consistent with other observations. On dredged material from a seamount north of Madeira Laughton, Hill, and Allan (1960) measured a value for J" of 5X 10~3 emu. Matthews (1961), on lava samples from a small abyssal hill on the North Atlantic, measured a median value for J" of 5><10‘3 emu. Bullard and Mason (1963), for basalt dredged from the Mendocino fault, indicated that J" may reach the exceptionally large value of 0.3 emu. In Maher, Boutelle, and Hoke Seamounts there is a great discrepancy between the directions calculated for the total magnetization (table 2) and the directions of the present geomagnetic field. To investigate the position of the north virtual pole, its locus as a function of K has been calculated for the three seamounts in the manner described in this paper. The loci for Maher and Hoke Seamounts are shown in figure 10, and the locus for Boutelle Seamount in figure 11, all on meridi- onal stereographic projections. The loci are wide apart and do not intersect for the assumed range of values of K. The directions of their 1,, are wide apart. No reasonable values of K can be found which will yield directions of 1,, convergent on a single pole for these three seamounts. Therefore, two alternative explanations are to be considered: (a) the seamounts, when they acquired their TRM, were at different positions with respect to the geomagnetic pole than they are at present, and also they are of different ages, or (b) they are of the same age, but relative changes and local rotations have occurred. Several investigators have proposed reconstructions of the track of the geomagnetic pole for different continents through geologic time, based on paleo- magnetic data. The tracks for Australia and North INTERNAL MAGNETIZATION 0F SEAMOUNTS AND ITS COMPUTER CALCULATION America (Runcorn, 1962) and for South America (Creer, 1964), are shown in figures 10 and 11. No track has been presented as yet for the Pacific area. Granted that there are discrepancies between difierent reconstructions of these tracks, certain features of the phenomenon are well established. The geomagnetic pole increasingly departs from its present position when receding in time from the Cenozoic. The data for // Track for Australia \ 90° w F21 each continent are more internally consistent than the data from continent to continent, and there are large differences between the tracks for the different con- tinents. As the three seamounts analyzed in this paper adjoin the North American continent, it is important to compare the loci for the virtual pole calculated for them with the track for that continent. To account for the deviations between the seamounts’ FIGURE 10.——Loci of north virtual pole for Maher and Hoke Seamounts. Values of K along loci are in 10" emu units. and North America are taken from Runcorn (1962), of the paleomagnetic pole for Australia from 'Creer (1964). A meridional stereographic projection of the hemisphere o Maher antipode Hoke antipode @ The tracks and the one for South America 90°W.—0°—90°E. is used. F22 loci and the track for North America, we can consider different kinds of simple relative motions. Hoke’s locus falls about 90° of arc west of the track for that continent, yet only about 20° of are cast of that for Australia. A northward shift of the locus by about 20° of arc would bring it into practical coincidence with the track for North America. In particular, the point corresponding to K=O would coincide with the Track for South 90° E SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY present pole. On the other hand, an eastward shift of 90°—105° of arc would bring the locus into coincidence with the Jurassic-Silurian segment of the track. Yet, since Hoke and Boutelle Seamounts are only about 7° of are away from the western edge of North America, their feasible westward relative motion is not more than that amount. Otherwise, they would have had to slip under the continent. Boutelle (9 America Hoke ® North America / Track for SP FIGURE 11.—Locus of north virtual pole for Boutelle Seamount. paleomagnetic pole for North America is taken from Runcorn (1962) and the one for South America from Greer (1964). Values of K along locus are in 10"3 emu units. The track of the A meridional stereographic projection of the hemisphere 90°W.—180°W.—270°W. is used. INTERNAL MAGNETIZATION OF SEAMOUNTS AND ITS COLIPUTER CALCULATION Boutelle’s locus falls south of the track for North America, but with its closest point only about 15° of are away. With a northward shift of 15°—50° of are, for the range of K values shown in figure 11, the locus would intersect the track at approximately the Cam- brian segment. On the other hand, an eastward shift of the locus, which would not change the situation ap- preciably, is deemed unfeasible as explained before. Because of the proximity of Boutelle and Maher Seamounts and of the generally eastewest orientation of the intervening fracture zones, any meridional shift should be of generally the same magnitude for both of these seamounts. As explained before, no meridional shift can bring their loci to overlap the track for North America. If it were assumed that the earth’s magnetic field was reversely polarized at the time when Hoke and Maher became magnetized, then the antipode of the locus should be considered instead of the locus itself. There- by, the Hoke and Maher antipodal loci are found to fall near the track for North America. For example, for Hoke Seamount the point corresponding to K: 10X10‘3 emu, and for Maher Seamount that point corresponding to K=0 falls about 20° of are away from the track. The hypothesis proposed by Vine and Matthews (1963) about the origin of the main Pacific crustal layers bears on these interpretations. First, if the oceanic crust had spread eastward away from the mid- Pacific, the seamounts could have been carried eastward by up to approximately 90°. Second, if instead the oceanic crust had spread westward from the East Pacific Rise, the positions in the upper mantle corre- sponding to the seamounts could have been carried westward by up to approximately 20°. In the first alternative the loci for the seamounts would be found eastward of the track for North America, which is opposite of where it occurs. If the second alternative were true, the seamounts could have shifted westward only a maximum of about 7° of arc. Hence, a simple interpretation of these seamounts’ data gives no support to the spreading-floor hypothesis. In the northeast Pacific, because of horizontal dis- placements along the fracture zones, it is conceivable that certain adjoining crustal blocks may have been rotated about vertical axes. Thus, the magnetization vector would be rotated by an amount equal to the rotation undergone by the supporting block after the seamount was built upon it. On the other hand, the angle of inclination would not be affected by such rotation. As the average velocity of slip along a fracture zone may be of the order of a few centimeters per hundred years, such displacements of up to hundreds of kilometers which have been inferred would require tens of millions of years. As the growth and cooling of a F23 seamount can be accomplished in much less than 1 m.y., the rotation during such a period could amount at most to a few degrees. Indeed the magnetic maps of the area of the three seamounts (Mason and Raff, 1961; Rafi and Mason, 1961) suggest that block rotations might have occurred. The magnetic pattern is parti- tioned in blocks, each block being characterized by a certain linear trend which changes direction from block to block. In between there are zones with a disorderly pattern. Some of these changes of direction are ap— preciable. For example, the linear pattern turns by about 20° at about lat 44°N. and long 130°W. (Raff and Mason, 1961, pl. 1). Maher Seamount is about 10 nautical miles to the south of the Murray fracture zone (fig. 9), where the right-lateral displacement of 154 km has been proposed (Menard and Vacquier, 1958; Mason, 1958). The shearing action implied by such displacement, upon the crust on the south side of the fracture, could produce a clockwise rotation. Boutelle Seamount is between the Mendocino and Pioneer fracture zones. The proposed '7 left-lateral displacements of 1,160 km along the Men- docino and of 265 km along the Pioneer fracture zones (V acquier and others, 1961) indicate a shearing action which could produce counterclockwise rotation in the area of Boutelle Seamount. Thus, the east to southeast declination of 1,, in Maher Seamount and the south- west declination in Boutelle Seamount as shown in table 3 could be attributed, in part, to such block rota- tions. To compensate for them, the loci of Maher and Boutelle Seamounts may be rotated anticlockwise and clockwise, respectively, about the corresponding sea- mounts. The arcs thus described by a point of the loci of Maher and Boutelle Seamounts are shown in figures 10 and 11, respectively. It is observed that with this correction the loci approach the tracks of the geomag- netic pole shown in these figures. PROPOSED PROCEDURE TO INTERPRET THE DATA OF MANY SEAMOUNTS I shall now describe the procedure to be used if the data of many seamounts were available. The un- certainties in the interpretation for the three seamounts arise mainly because three is too few. The simplest situation is that of a group of seamounts formed at the same geologic time and when there have been no relative rotations or displacements of one seamount with respect to the others. Then the loci for the virtual pole for the different seamounts should intersect at one point. The value of the susceptibility K for each seamount could be read off, along each locus, at the point of intersection. Next we may consider a group of equal-age seamounts when relative rotations may have occurred. The F24 calculated loci of the virtual pole can be rotated about their seamounts. Each point of a locus, corresponding to a particular value of K, would describe an arc of circle about the seamount. The area common to all these arcs for the likely range of values of K defines an enclosure within which the virtual pole would be found. The enclosure thus defined will be called a corral. The closeness in the determination of the value of K for the different seamounts with this procedure would hinge upon the size and shape of the corral, which in turn would depend on the particular configuration of the seamounts. Furthermore, to investigate possible displacements, the seamounts should first be classified in groups of seamounts which are expected to have equal displace- ments, and the analysis of each group would proceed as indicated before. For example, seamounts between two fracture zones could be assumed to have undergone equal displacements. For each group a corral would be defined for the virtual pole. Then it should be possible to bring these corrals into coincidence by linear displacements, from which the motions along the fracture zones could be deduced. For a group of different-age seamounts, with no displacements or rotations between them, the approach would be as follows: The locus for each seamount would be established as a function of K. These loci should intersect the track of the virtual pole as a function of geologic time. If this track is unknown, a smooth line should be sought intersecting all the loci, as a first approximation to the track of the virtual pole. This procedure might give a good definition of the track of the pole if data for many seamounts of widely different ages were available. The most difficult case is when dealing with seamounts of different ages, and when rotations and displacements between them have to be considered. In such a case, groups of seamounts, which should have equal dis- placements, would be analyzed separately; and next, the group results would be analyzed as outlined in the previous paragraphs. CONCLUSIONS The triple-field method of analysis of a magnetic anomaly gives, without trial and error, the total internal magnetization of a given geologic body as to direction, sense, and magnitude under the assumption of uniform internal magnetization throughout the body or throughout each of a number of given regions in which the body may be subdivided. Besides, the examination of the three hypothetical fields permits a semiquantitative analysis of the problem. By taking advantage of the inherently high precision of digital computers, we can readily calculate the hypo- SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY thetical fields for a model of the body by a numerical differentiation of the gravimetric field. The concept of biased-field anomaly, as an algorithm built in the computer program, gives the option of obtaining several quantities geophysically significant, namely (a) an arbitrary component of the anomaly field intensity, (b) the total-intensity anomaly as conventionally defined in aeromagnetic prospecting, and (c) the magnitude of the anomaly field intensity. Seamounts offer a promising field of application for the triple-field method, because of their commonly distinctive anomalies and known shapes. It may not be enough to consider only the mass standing over the ocean floor, but also a root may have to be included to take into account the demagnetization by reheating and the subsequent remagnetization during cooling of the rocks underlying the volcanic mass. The magnetic field observed in certain seamounts can be reproduced very well under the simple assumption that their masses above the ocean floor are uniformly magnetized in a direction which may be different from that of the present geomagnetic field. By examining the vector relationship between J, J", and K3 as a function of K, we find the TRM vector L, can be a function of K, and from the direction of 1,. a locus for the north virtual pole can be established. Such loci, when the data for many seamounts are available, would permit the investigation of the posi- tions of the virtual pole and of possible differential crustal displacements and rotations. The directions of the TRM in Maher, Boutelle, and Hoke Seamounts appear to be substantially different from those of the present field. Yet, under certain assumptions the loci for the virtual pole can be rec- onciled with the track for the paleomagnetic pole which has been proposed for the North American continent. Hoke and Maher Seamounts would have become magnetized at a time when the field was reversed. Boutelle and Hoke Seamounts would be of Cambrian age or older, and Maher Seamount possibly much younger. Since the Boutelle locus and the Hoke antipodal locus apparently lie within the limits of confidence of the track for North America, only none or a small southward relative motion of these Seamounts would have taken place. Alternatively, rotation of Boutelle and Maher Seamounts, to com- pensate for a possible rotation arising from the shearing along the nearby fracture zones, would bring the loci nearer to the tracks proposed for North America and South America. A systematic analysis, in the manner explained, of the magnetization of a greater number of seamounts in the northeast Pacific may be useful. INTERNAL MAGNETIZATION OF SEAMOUNTS AND ITS COMPUTER CALCULATION REFERENCES Bhattacharyya, B. K., 1964, Magnetic anomalies due to prism- shaped bodies with arbitrary polarization: Geophysics, v. 29, no. 4, p. 517—531. Bott, M. H. P., 1963, Two methods applicable to computers for evaluating magnetic anomalies due to finite three- dimensional bodies: Geophys. Prosp., v. 11, no. 3, p. 292—299. Bullard, E. C., and Mason, R. G., 1963, The magnetic field over the oceans, in Hill, M. N., ed., The sea—ideas and observa— tions on progress in the study of the seas: New York, Interscience Publishers, v. 3, The earth beneath the sea, p. 175—217. Carsola, A. H., and Dietz, R. S., 1952, Submarine geology of two flat-topped northeast Pacific seamounts: Am. Jour. Sci., v. 250, no. 7, p. 481-497. Chapman, Sydney, and Bartels, Julius, 1940, Geomagnetism, v. 1: Oxford, Clarendon Press, 542 p. Creer, K. M., 1964, A reconstruction of the continents for the upper Paleozoic from paleomagnetic data: Nature, v. 203, no. 4950, p. 1115—1120. Grant, F. S., 1952, Three-dimensional interpretation of gravita- tional anomalies: Geophysics, v. 8, no. 2, p. 344—364. Hamilton, E. L., 1956, Sunken islands in the mid-Pacific moun- tains: Geol. Soc. America Mem. 64, 95 p. Heezen, B. C., 1962, The deep sea floor, in Runcorn, S. K., ed., Continental drift: New York, Academic Press, p. 235—288. Heezen, B. C., and Menard, H. W., 1963, Topography of the deep—sea floor, in Hill, M. N., ed., The sea—ideas and observations on progress in the study of the seas: New York, Interscience Publishers, v. 3, The earth beneath the sea, p. 233-280. Henderson, R. G., and Allingham, J. W., 1964, The magnetiza— tion of an inhomogeneous laccolith calculated on a digital computer, pt. 2 of Computers in the mineral industries: Stanford Univ. Pubs. Geol. Sci., v. 9, no. 2, p. 481-497. Hess, H. H., 1946, Drowned ancient islands of the Pacific Basin: Am. Jour. Sci., v. 244, no. 11, p. 772—791. Ingersoll, L. R., Zobel, O. J., and Ingersoll, A. C., 1954, Heat conduction with engineering, geological and other applica- tions: Madison, Univ. Wisconsin Press, 325 p. Kellogg, O. D., 1929, Foundations of potential theory: New York, Frederick Ungar Publishing Co., 384 p. Laughton, A. S., Hill, M. N., and Allan, T. D., 1960, Get» physical investigations of a seamount 150 miles north of Madeira: Deep-Sea Research, v. 7, no. 2, p. 117—141. MacDonald, G. A., 1963, Physical properties of erupting Hawaiian magmas: Geol. Soc. America Bull., v. 74, no. 8, p. 1071-1077. MacDonald, G. A., and Finch, R. H., 1950, The June 1950 eruption of Mauna Loa: Volcano Letter, no. 509, p. 1—6. Mason, R. G., 1958, A magnetic survey off the west coast of the United States: Geophys. Jour. [London], v. 1, no. 4, p. 320—329. Mason, R. G., and Rafi, A. D., 1961, Magnetic survey off the west coast of North America, 32°N latitude to 42°N latitude: Geol. Soc. America Bull., v. 72, no. 8, p. 1259—1266. Matthews, D. H., 1961, Lavas from an abyssal hill on the floor of the North Atlantic Ocean: Nature, v. 190, no. 4771, p. 158—159. Maxwell, J. C., 1904, A treatise on electricity and magnetism, v. 2: Oxford, Clarendon Press, 500 p. F25 Menard, H. W., 1955, Deformation of the northeastern Pacific basin and the west coast of North America: Geol. Soc. America Bull., v. 66, no. 9, p. 1149—1198. 1959, Geology of the Pacific sea floor: Experientia, v. 15, no. 6, p. 205—214. 1960, The East Pacific Rise: Science, v. 132, no. 3441, p. 1737—1746. 1964, Marine geology of the Pacific: New York, McGraw- Hill Book Co., 271 p. Menard, H. W., and Ladd, H. S., 1963, Oceanic islands, sea- mounts, guyots, and atolls, in Hill, M. N., ed., The sea— ideas and observations on progress in the study of the seas: New York, Interscience Publishers, v. 3, The earth beneath the sea, p. 365-385. Menard, H. W., and Vacquier, Victor, 1958, Magnetic survey of part of the deep sea floor 03' the coast of California: U.S. Office Naval Research, Research Rev., June, p. 1—5. Morgan, N. A., and Grant, F. S., 1963, High speed calculation of gravity and magnetic profiles across two-dimensional bodies having an arbitrary cross section: Geophys. Prosp., v. 11, no. 1, p. 10-15. Nagata, Takesi, 1953, Rock-magnetism: Tokyo, Maruzen Co., 232 p. Nayudu, Y. R., 1962, A new hypothesis for origin of guyots and seamount terraces, in MacDonald, G. A., and Kuno, Hisashi, Crust of the Pacific Basin: Am. Geophys. Union Geophys. Mon. 6, p. 171—180. Peter, George, and Stewart, H. B., 1965, Ocean surveys: The systematic approach: Nature, v. 206, no. 4988, p. 1017—1018. Press, Frank, and Ewing, Maurice, 1952, Magnetic anomalies over oceanic structures: Am. Geophys. Union Trans, v. 33, no. 3, p. 349—355. . Rafl’, A. D., 1962, Further magnetic measurements along the Murray fault: Jour. Geophys. Research, v. 67, no. 1, p. 417-418. Rafi, A. D., and Mason, R. G., 1961, Magnetic survey off the west coast of North America, 40°N latitude to 52°N lati- tude: Geol. Soc. America Bull., v. 72, no. 8, p. 1267—1270. Rittmann, Alfred, 1962, Volcanoes and their activity: New York, John Wiley & Sons, 305 p. Runcorn, S. K., 1962, Paleomagnetic evidence for continental drift and its geophysical cause, in Runcorn, S. K., ed., Continental drift: New York, Academic Press, p. 1—39. Talwani, Manik, and Heirtzler, J. R., 1964, Computation of magnetic anomalies caused by two-dimensional structures of arbitrary shape, pt. 1 of Computers in the mineral industries: Stanford 'Univ. Pubs. Geol. Sci., v. 9, no. 1, p. 464—479. Talwani, Manik, Le Pichon, Xavier, and Heirtzler, J. R., 1965, East Pacific Rise—The magnetic pattern and the fracture zones: Science, v. 150, no. 3700, p. 1109—1115. US. Naval Oceanographic Office, 1962, A marine magnetic survey south of the Hawaiian Islands: US. Naval Oceanog. Oflice, Tech. Rept. TR—137, 47 p. Vacquier, Victor, 1962a, Magnetic evidence for horizontal displacements in the floor of the Pacific Ocean, in Runcorn, S. K., ed., Continental drift: New York, Academic Press, p. 135—144. 1962b, A machine method for computing the magnitude and the direction of magnetization of a uniformly magnet- ized body from its shape and a magnetic survey, in Nagata, Takesi, Benedum earth magnetism symposium: Pitts- burgh, Univ. Pittsburgh Press, p. 123-137. F25 Vacquier, Victor, Ralf, A. D., and Warren, R. E., 1961, Hori- zontal displacements in the floor of the northeastern Pacific Ocean: Geol. Soc. America Bull., v. 72, no. 8, p. 1251—1258. Vacquier, Victor, Steenland, N. 0., Henderson, R. G., and Zietz, Isidore, 1951, Interpretation of aeromagnetic maps: Geol. Soc. America Mem. 47, 151 p. Van Voorhis, Gerald, and Walczak, James, 1963, Summary of magnetization computations for Kelvin Seamount: US. Naval Oceanog. Oflice, Informal Manuscript Rept. M—8—63, 19 p. Vine, F. J., and Matthews, D. H., 1963, Magnetic anomalies over oceanic ridges: Nature, v. 199, no. 4897, p. 947—949. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY Vine, F. J., and Wilson, J. T., 1965, Magnetic anomalies over a young oceanic ridge 03‘ Vancouver Island: Science, v. 150, no. 3695, p. 485-489. Wentworth, C. K., Carson, M. H., and Finch, R. H., 1945, Discussion on the viscosity of lava: Jour. Geology, v. 53, no. 2, p. 94—104. Wilson, J. T., 19659., A new class of faults and their bearing on continental drift: Nature, v. 207, no. 4995, p. 343—347. 1965b, Submarine fracture zones, aseismic ridges, and the International Council of Scientific Unions Line— Proposed western margin of the East Pacific Ridge: Nature, v. 207, no. 5000, p. 907—911. 1965c, Transform faults, oceanic ridges, and magnetic anomalies southwest of Vancouver Island: Science, v. 150, no. 3695, p. 482—485., PROFESSIONAL PAPER 554—F PLATE 1 UNITED STATES DEPARTMENT OF THE INTERIOR GEOLOGICAL SURVEY )2 N K=0.001 in cgs units and 9:0.49 oersted are as- sumed. Field contour values In gammas. Area bounded by hachures corresponds to the base of the seamount. PolygonaI segmented line corre- sponds to the boundary of the grid of fitting points. The track lines of the ship-borne magnetometer are shown in figure H /)00 /)00 —2000 _.2300 B. MODEL OF SEAMOUNT CONTOUR INTERVAL 100 FATHOMS, EXCEPT FOR THE 1650-FATHOM CONTOUR A. BATHYMETRIC MAP CONTOUR INTERVAL 100 FATHOMS N MN N , MN fl\o,5 /O.25 [I W / / \ llu IIII O \ \ \\/ \ \0.5__|\ \HHIIIIIIIHH ‘\ / 0 \E \ \ \IIIIIII 20 \\ \/ IIIIIII‘I IIlII\ r“\I ‘0. 25 E. FIELD FOR A MAGNETIC-EAST MAGNETIZATION C. FIELD FOR A MAGNETIC-NORTH MAGNETIZATION D. FIELD FOR A VERTICAL MAGNETIZATION N MN \ \IIIIII III/////// /) F \\\\ \\\\ \\ \/\\\ / G. A LEAST-SQUARES FITTED FIELD H. OBSERVED FIELD CONTOUR MAPS SHOWING BATHYMETRY, COMPUTER MODEL, AND HYPOTHETICAL AND OBSERVED MAGNETIC FIELDS FOR MAHER SEAMOUNT, NORTHEASTERN PACIFIC OCEAN OL 1'0 210 3'0 MILES F. FIELD FOR A MAGNETIZATION PARALLEL TO THE PRESENT FIELD (I) 1'0 2'0 3'0 KILOMETERS 243-571 0 - 67 (In pocket) UNITED STATES DEPARTMENT OF THE INTERIOR GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—F PLATE 2 K=0.001 in cgs units and 9:0.49 oersted are as- sumed. Field contour values in gammas. Area bounded by hachures corresponds to the base of the seamount. Polygonal segmented line corre- sponds to the boundary of the grid of fitting points. The track lines of the ship-borne magnetometer are shown in figure H A. BATHYMETRIC MAP I B. MODEL OF SEAMOUNT CONTOUR INTERVAL 100 FATHOMS CONTOUR INTERVAL SHALLOWER THAN 1100 FATHOMS IS 100 FATHOMS: DEEPER IS 200 FATHOMS \\\\II//// / // D. FIELD FOR A VERTICAL MAGNETIZATION "-1".Crf'E-LD'fi-FOR A MAGNET'C’NORTH MAGNETIZAT'ON , E. FIELD FOR A MAGNETIC-EAST MAGNETIZATION :14 V ~7L ‘F‘ 9‘ . .. ‘ .a......_.,.._»...._,. m»-.. ; , . (nu-7.». {1 2 MN ti“. INA ' I N MN m p" - L1,; : a“ CONTOUR MAPS SHOWING BATHYMETRY, COMPUTER MODEL, AND HYPOTHETICAL AND OBSERVED MAGNETIC 2. ' FIELDS FOR BOUTELLE SEAMOUNT, NORTHEASTERN e PACIFIC OCEAN o 10 20 30 4o 50 MILES I I l I I I \l o 1‘0 2'0 310 4'0 510 KILOMETERS // F. FIE-1D FORi‘A M'AiGNETIZATION‘ PARALLEL . { I’ G. OPTIMUM—FIT FIELD H. OBSERVED FIELD " "To THE PRESENT FIELD -; 243—571 0 - 67 (In pocket) » .1 :bg- fru-‘fiadmfi' ~ ‘ 4v; . '7 £5,“ ., UNITED STATES DEPARTMENT OF THE INTERIOR PROFESSIONAL PAPER 554—1: GEOLOGICAL SURVEY PLATE 5 MN MN 20° 2000 \\ Z / 000— 50— 100 \ llIIII/ll/ :III“\\ \ Q 0 V\ \IlI/I/ A. BATHYMETRIC MAP CONTOUR INTERVAL 200 FATHOMS C. FIELD FOR A VERTICAL MAGNETIZATION B. FIELD FOR A MAGNETIC-NORTH MAGNETIZATION K:0.00l in cgs units and (2: 0.47 oersted are as- sumed. Field contour values In gammas. Area bounded by hachures corresponds to the base of the seamount. Polygonal segmented line corre- sponds to the boundary of the grid of fitting points. The track lines of the ship-borne magnetometer are shown in figure G E. FIELD FOR A MAGNETIZATION PARALLEL F. OPTIMUM FIT FIELD TO THE PRESENT FIELD Z MN P \ \ / \\ \\\ / ‘\ 0 \L \ / ' N“ / \ \\ x / \\\ \\\ ’0 00 \ \\ CONTOUR MAPS SHOWING BATHYMETRY, AND HYPOTHETICAL AND OBSERVED MAGNETIC FIELDS FOR HOKE SEAMOUNT, NORTHEASTERN PACIFIC OCEAN \4 my I” H ”MIC \ om HH\\\‘\\\ ‘005' O 10 20 30 MILES J 10 20 3O KILOMETERS I TO G. OBSERVED FIELD 243-571 0 - 67 (In pocket) Pennsylvanian and Associated Rocks in Wyoming GEOLOGICAL SURVEY PROFESSIONAL PAPER 554-G Pennsylvanian and Associated Rocks in Wyoming SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY GEOLOGICAL SURVEY PROFESSIONAL PAPER 554—G - ”A regional staa’y of t/ze extent, tnicéness, lit/zology, ana’ stratigrap/zic relations of t/ze Tensleep, Ains— a’en, Casper, ana’ Fountain Formations, witn em- p/zasis on origin, paleogeograp/zy, ana’paleotectonie implications UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1967 UNITED STATES DEPARTMENT OF THE INTERIOR STEWART L. UDALL, Secretary GEOLOGICAL SURVEY William T. Pecora, Director For sale by the Superintendent of Documents, US. Government Printing Office Washington, DC. 20402 CONTENTS Page Page Abstract __________________________________________ G1 Casper Formation __________________________________ G17 Introduction _______________________________________ 1 Fountain Formation ________________________________ 18 Location and extent of the area ___________________ 1 Minnelusa and Hartville Formations __________________ 18 Scope and objectives ____________________________ 1 Morgan Formation and Weber Sandstone ______________ 20 Previous work and acknowledgments ————————————— 2 Rocks of Morrow age in the Minnelusa, Hartville, Casper, MgthOdSPf study___.__________________.____:___'_ 2 and Morgan Formations ___________________________ 20 Tectonic settlng of. the Middle Rocky Mountain region 1n T ensl e ep Sandstone _________________________________ 21 Pennsylvaman time """"""""""""""""" 3 Lithology and thickness 21 Definition of the Pennsylvanian System in Wyoming__ _ _ 4 """"" ' """""""" Extent and thickness of the Pennsylvanian System ______ 5 Age """""""""""""""""""""" 21 Lower boundary of the Pennsylvanian System __________ 6 Quadrant Quartzite --------------------------------- 22 Character of the disconformity ___________________ 6 LithOIOSY and thickness ————————————————————————— 22 Time span of the hiatus in Wyoming ______________ 6 Age ___________________________________________ 22 Paleogeomorphology of the disconformity in north- Rocks of Des Moines, Missouri, and Virgil age _________ 23 western Wyoming ............................ 7 Des Moines rocks _______________________________ 23 Interregional significance of the disconformity on Missouri rocks __________________________________ 23 the craton """""""""""""""""" 7 Virgil rocks ____________________________________ 24 Amsden Formation """"""""""""""""""" 8 Upper Boundary of the Pennsylvanian System _________ 25 Lithologic units ________________________________ 8 Age ___________________________________________ 10 Source of sand ______________________________________ 25 Nomenclature and the Amsden problem ___________ 10 Tabulation 0f control data """""""""""""" 27 Darwin Sandstone Member ______________________ 11 Silrface ---------------------------------------- 27 Horseshoe Shale Member ________________________ 12 Subsurface ------------------------------------- 28 Ranchester Limestone Member ___________________ 14 References cited ____________________________________ 30 ILLUSTRATIONS [Plates are in pocket] PLATE 1. Correlation diagrams (stratigraphic sections) of Pennsylvanian and Mississippian rocks in the Middle Rocky Mountain region, Wyoming, showing relation of lithologie units and approximate time boundaries. 2. Maps showing extent, thickness, and lithology of members of the Amsden Formation of Mississippian and Pennsyl- vanian age, and correlative formations in the Middle Rocky Mountain region. 3. Maps showing extent, thickness, and lithology of Tensleep Sandstone (except that part of Permian age) and total thickness of the Amsden, Tensleep, and correlative formations in the Middle Rocky Mountain region. Pale FIGURE 1. Index map showing location of report area _____________________________________________________________ G2 2. Index map showing position of Ancestral Rocky Mountains relative to existing ranges and basins in Wyoming- _ 4 3- Correlation Chart of Pennsylvanian and closely related strata in Wyoming _________________________________ 5 4. Index map of Wyoming showing areal extent of Pennsylvanian formations and exposed Precambrian basement rock in existing mountain ranges __________________________________________________________________ 5 5. Photograph showing erosional disconformity between Amsden Formation and Madison Limestone in Wind River Canyon, south of Thermopolis, Wyo _________________________________________________________ 6 6- Type and reference sections of Amsden Formation ______________________________________________________ 9 7- Photographs showing crossbedding in Darwin Sandstone Member, Tensleep Canyon _________________________ 12 8. View of north wall of Tensleep Canyon near trout—rearing station showing the top of the Madison Limestone, the three members of the Amsden Formation, and the lower part of the Tensleep Sandstone ____________ 14 9. Paleogeologic map showing areal geology of Pathfinder uplift at end of Atoka time and thickness and lithology of Ranchester Limestone Member of Amsden Formation in area surrounding the uplift __________________ 16 10. Photograph of the red marker on east fork of Wagonhound Creek, west of Arlington, Carbon County, Wyo____ 18 11. Geologic section showing generalized correlation of Pennsylvanian and Permian rocks across southern part of Laramie basin and Laramie Mountains, Wyo. and Colo ______________________________________________ 19 III IV CONTENTS Page FIGURE 12. Photograph of Tensleep Sandstone in Sinks Canyon of Middle Popo Agie River ____________________________ G21 13. Photograph showing giant crossbedding in the Tensleep Sandstone, Bull Lake Canyon, Wind River Range____ 22 14. Map showing western limits of Missouri and Virgil rocks in central Wyoming ______________________________ 24 15. Formations which overlie Pennsylvanian strata in Wyoming _____________________________________________ 25 16. Map showing restored regional drainage patterns for Carboniferous time in northern United States and southern Canada ________________________________________________________________________________________ 26 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING By WILLIAM W. MALLORY ABSTRACT The Pennsylvanian System in the Middle Rocky Mountains of Wyoming is represented principally by the Amsden and Ten- sleep Formations. These grade eastward into the Minnelusa, Hartville, Casper, and Fountain Formations, and southward in- to the Weber and Morgan Formations. The Amsden Formation has three members: the Darwin Sandstone at the base, the Horse- shoe Shale (new name) in the middle, and the Ranchester Lime- stone (new name) at the top. The Darwin is of Chester (Mis— sissippian) age; the Horseshoe is of Chester and Morrow ages; and the Ranchester is of Morrow age at the base but is mostly of Atoka age. The Tensleep Sandstone (and its equivalent, the Quadrant Quartzite, in the Yellowstone Park area) is entirely of Des Moines age in the Bighorn and Wind River Basins. In central Wyoming, however, the Tensleep contains beds of Missouri, Vir- gil, and Early Permian ages. The Minnelusa and Hartville For— mations in the eastern part of the State contain interbedded sandstone, shale, and carbonate rock ranging in age from Morrow to Early Permian. The Casper is composed of sandstone and car- bonate rock and grades southward into red arkosic conglomerate of the Fountain Formation in the Laramie basin and adjacent areas. The Casper and Fountain Formations also range in age from Morrow to Early Permian, but the older strata are missing in a wide area in central Wyoming where the Pathfinder uplift, the northernmost element of the Ancestral Rocky Mountains, dominated the paleogeographic scene. In Late Mississippian time the Mississippian Madison Lime- stone was exposed throughout the area except in the Laramie basin and adjacent mountains, where the Precambrian basement rocks were exposed. A karst surface formed in many areas on the Madison in Late Mississippian time. In late Chester and per- haps early Morrow time, marine waters spread eastward across Wyoming from the geosyncline in southern Idaho and invaded the major valleys first and later covered nearly all the Madison terrane. The Darwin Sandstone Member of the Amsden, probably derived from an older sandstone in south-central Canada, was deposited in the resultant bay. Continued submergence and a change in the source and type of elastic sediments carried into the area resulted in the deposition of a nearly tabular layer, the red Horseshoe Shale Member of the Amsden. During very late Morrow and Atoka time, limestone was precipitated from a clear sea to form the Ranchester Limestone Member of the Amsden. In southeastern Wyoming, coarse red arkosic sediments shed from the north end of the Ancestral Front Range uplift and the southern part of the Pathfinder uplift were deposited in the Laramie basin area as the Fountain Formation. In Des Moines time the Tensleep Sandstone was deposited on a broad marine shelf over most of the area, while limestone, sand- stone, and shale of the Hartville and Minnelusa Formations were deposited in eastern Wyoming. Arkose was deposited in dimin- ishing quantities in the Laramie basin area during Atoka, Des Moines, Missouri, and Virgil times. In Missouri and Virgil time, western Wyoming was emergent, and some sand derived from the Des Moines Tensleep there was removed and redeposited in central Wyoming. The Pathfinder uplift had its greatest extent and elevation in Atoka time and was progressively submerged later in the Pennsylvanian. It ceased to exist in late Virgil or Early Permian time. INTRODUCTION LOCATION AND EXTENT OF THE AREA The report area is the State of Wyoming except six counties on the east side of the State (Campbell, Crook, Weston, Niobrara, Goshen, and most of Laramie County) and the thrust belt along the west side (much of Lincoln County and parts of adjacent counties) (fig. 1). The area spans 4° of latitude (ALP—45° N.) and 6° of longitude (105°—111° W.) and encompasses about 70,000 square miles. Physiographically it is dominated by the Middle Rocky Mountains. SCOPE AND OBJECTIVES The stratigraphy of the Pennsylvanian System in the Middle Rocky Mountain region is generally simple, but understanding of relationships has been complicated by several problems, most of which concern the Amsden Formation. The objective of this study was to achieve a regional synthesis of Pennsylvanian strata (and closely associated rocks of Chester age) in an effort to resolve inconsistencies in the literature. The study dis- closed that when these strata are seen in regional per- spective, an uncomplicated pattern is apparent, and their relation to strata of greater, lesser, and equivalent G1 G2 112° 110° 108° 106° 104° 46° MONTANA I CROOK LOHVG HlflOS 44° CAMPBELL 42° COLORADO O 100 200 MILES FIGURE 1.—Location of report area. age can be demonstrated (pls. 1—3) with reasonable clarity. PREVIOUS WORK AND ACKNOWLEDGMENTS Previous work upon which correlation was based in- cludes reports by Thomas, Thompson, and Harrison (1953), Love (1954; written commun., 1956), Henbest (1954, 1956, 1958a, b), Maughan and Wilson (1960), and Verville (1957). Thomas, Thompson, and Harrison (1953) provided a clear and detailed exposition, in age sequence, of the lithology and distribution of the Pennsylvanian System in the Laramie Mountains area. Their plate 9, a series of correlated columnar sections showing fusulinid data, provides a key to correlation and stratigraphic mapping throughout southeastern and south-central Wyoming. Love’s (1954; written commun., 1956) correlation dia- grams, similar to those of Thomas, Thompson, and Har- rison (1953), provide a correlation reference framework for western, central, and northern Wyoming. Henbest’s papers on dating of the Amsden and Tensleep were help- ful, particularly in the Bighorn Basin. Maughan and Wilson’s (1960) sections of the Casper and Fountain Formations are a guide to the interrelation of these two formations. Verville’s (1957) work in the Mayoworth area was significant in locating the Missouri, Virgil, and Wolfcamp subcrop limits in central Wyoming. Verville and Momper (1960) helped explain problems of the deep subsurface of southwestern Wyoming. Parts of the correlation diagrams (stratigraphic sec- tions, pl. 1) were taken directly from the work of these authors and were combined with measured sections col- SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY lected from many sources and with sample logs of deep wells. Most of the well logs used were prepared by the American Stratigraphic Co. J. D. Love prepared sev— eral logs which were helpful. The sources of data are listed in the section, “Tabulation of Control Data.” In the summer of 1963, Ernest E. Glick assisted the author in measuring the type section of the Amsden Formation at Amsden Creek in the northeastern part of the Bighorn Mountains and the reference section of the Amsden at Tensleep Canyon. METHODS OF STUDY In the summer of 1959 the author examined all sig- nificant outcrop areas of Pennsylvanian rocks to gain firsthand familiarity with the strata and to augment study of the literature, measured sections, and sub— surface data. Problem areas or areas of special geologic interest were examined in detail later. The correlation of strata was based on similarity of lithology and vertical sequences of lithology; paleon- tologic data where available were used in dating and as a supplemental aid in correlation. Post—Laramide erosion has exposed Pennsylvanian strata in many areas on the flanks of the Middle Rocky Mountains, and the presence of oil in the region has encouraged the drilling of many wildcat and pool wells in the sedimentary basins. The abundance of surface and subsurface control makes it possible to correlate the regional stratigraphy with confidence. In areas of facies change an effort was made, in constructing correlation diagrams, to select well logs or surface sections that contain the transition facies. Correlation of lithologic units within the Tensleep Sandstone is more difficult than correlation of the lithologically distinct members of the Amsden. Correlation is easiest in the Bighorn and Wind River Basins because there the Tensleep Formation contains strata of Des Moines age only. Farther east the Tensleep (and the equivalent Casper Formation) contains strata of Missouri, Virgil, and Wolfcamp age. However, careful comparison of sur- face sections that contain useful fossil evidence with well logs from adjacent basins permits correlations that give reasonable results. To achieve consistent lithologic and time correlation, a network of correlation diagrams (stratigraphic sec- tions) was constructed showing thickness, lithology, and age for control wells and sections. The cutoff date for incorporation of data into maps and sections was J an- uary 1, 1961. Areas where rocks older than Pennsyl- vanian are exposed are shown on the base map by ab- sence of color or symbol. All maps are intended to be objective in that they show existing geologic relations and features instead of inferred original relations and PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING features. The text discussion of each map contains comments, where appropriate, on inferred original depositional features. Objectivity in subsurface isopleth maps is difficult to attain. Complete objectivity might be achieved by me- chanically spacing all isopleths between available con- trol. But, if additional control points were later incor- porated, the new map would probably look radically different. Thus, the so-called objectivity would be false because the distribution pattern of the control points is treated as a mapping parameter. Another method is to use the available control to detect trends or patterns which seem to be present and then to allow these suspected trends or patterns to in- fluence the actual construction of the map. Such a map would be objective in that it showed a selected param- eter, such as the thickness of a limestone. The map would be subjective to the extent that it contained trends or patterns that the mapper interprets. In this report the author makes substantial use of interpretation of trend and pattern. Control points and parameter values are shown on all maps, so the reader may critically evaluate the author’s interpretations or even make his own maps if he wishes. The maps of western Wyoming accompanying this report are of lithologic units except for those units that encompass several ages of geologic time. The Ams- den Formation consists of three members of distinctly different lithology, and each is mapped separately. The age of the lower (Darwin Sandstone Member) is Ches- ter, that of the middle (Horseshoe Shale Member) is Chester and Morrow, and that of the upper (Ranches— ter Limestone Member) is Morrow and Atoka. The Tensleep Sandstone, on the other hand, is nearly homogeneous lithologically, but as it contains beds of Des Moines, Missouri, and Virgil age, the formation is divided on sections and maps as accurately as possible by age divisions.,Formations in eastern Wyoming span Pennsylvanian time and are therefore mapped accord- ing to age, on appropriate plates. Conventional colors are used to indicate lithology; yellow indicates sandstone (80 percent or more quartz sandstone), blue indicates carbonate rock (80 percent or more limestone or dolomite), and red indicates shale (80 percent or more shale or fine siltstone) . Mixtures are shown by appropriate color combinations. For example, bluish-green indicates sandy carbonate rock, and yellow- ish-green indicates calcareous sandstone; browns indi- cate mixtures of the three end-members—carbonate rock, sandstone, and shale. (See explanation on plates.) In the selection of the colors to be used, the proportions of pure rock types at each control point for each unit mapped were quantitatively determined in 20 percent G3 increments from pure end-members. A judgment factor of about 5 percent was introduced where reasonable to allow some leeway in drawing color-band boundaries. TECTONIC SETTING OF THE MIDDLE ROCKY MOUNTAIN REGION IN PENNSYLVANIAN TIME Throughout Paleozoic time the two major tectonic provinces in the western conterminous United States were the North American craton and the Cordilleran orthogeosyncline. Stille (1936, 1940) recognized within the geosyncline an inner belt adjacent to the craton (miogeosyncline) and an outer belt (eugeosyncline). During Pennsylvanian time most of what is now the State of Wyoming lay on the western margin of the craton; the extreme western edge of Wyoming extended a short distance into the miogeosynclinal belt (fig. 2). Cratonic Western United States in Mississippian time was an unusually stable shelf over which marine waters advanced unimpeded. Tectonic activity in Penn- sylvanian time, however, was greater in this region than at any other time between the Precambrian Era and the Laramide orogenic episodes of latest Cretaceous and Tertiary time. A chain of mountain ranges, the Ancestral Rocky Mountains, arose from the featureless Mississippian surface early in the Pennsylvanian Pe- riod in a belt which extended from central Wyoming to southern Oklahoma. Four separate uplifts composed the Ancestral Rocky Mountains of the Western United States (fig. 2). The northernmost and smallest was the Pathfinder uplift in central Wyoming, which attained its widest areal extent in Atoka time (Mallory, 1963, p. E58). The large Ancestral Front Range uplift of central Colorado (Mal- lory, 1960, fig. 3) was part of a mountain chain that bifurcated in the subsurface into central New Mexico and southern Oklahoma (Rascoe, 1962, fig. 2). The Un— compahgre uplift which centered in southwestern Colorado may have attained along its southwest margin the highest elevation of all four uplifts. The Zuni- Defiance uplift in New Mexico and Arizona was a broad upwarp whose limits are not clearly defined. The report area lies at the north terminus of the (Pennsylvanian) Ancestral Rocky Mountain tectonic complex. It includes the Pathfinder uplift, the northwest terminus of the Front Range uplift, and the cratonic Wyoming shelf, northwest of the Ancestral Rocky Mountains. This shelf was only mildly depressed in Pennsylvanian time in contrast with the strongly down— warped 0r downfaulted intermontane troughs and basins of Colorado and Utah. SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY A r’ r' I . L i I i l """""""" l ‘ "7 ._ ----- J; . | l 'MONTANA l __________ | ___________ .- -—--—-.—--L-.1. ................. -I--l . '02 I l i :2 ¥ i ’ lI-a -\. i ! $5.. — — “1' I l:- ! 'I I i3; BIGHORN 'LBASIN PSIDWDER R'VERI53_.S_._ .. _"l? I """""""" . , . ""I. . l ' ' . - L. I i I L . 1-1.._ i I i i L.-. ' ' I !_,‘_ O‘v/ L.I «a l ! I : {I _Creek_ I _l _____________ 1L'—'—'—'-'-I'T ......... —I ----------- Q I i ! I WIND RIVER i i l “5'“ W Y OIM I N G i .___,__.|_.._ SWEETWATER i i 4/: l :1 UP . - \ - LIFLH ! (f, I . f I i _ _: I I as I: a - G . :1 =w ‘I‘ "a,m\ A; I55 l - ___ ___L _______ __ 41' :> I a ‘ “ s I I: H GREAT DIVIDE ! Fm - —‘-‘I- BASIN ! , i I ' - . |____ _ __4 .......... -! : l ROCK SPRINGS : 1- 11- UPLIFT ._._ _ . l ,r I WASHAKIE' I ! BASIN - ___________ _L--—--.—-‘— — """" a £0er I" I i I l ___ ._ _J._ I ! 'r- r'J i i .l 1 . (I) 50 100 15:0 MILES FIGURE 2.—Position of the Ancestral Rocky Mountains relative to existing ranges and basins in Wyoming. Existing mountain ranges are shaded; Pennsylvanian ranges are finely cross- ruled. The east limit of the Cordilleran geosyncline in western Wyoming (coarsely cross ruled) is the limit of intense Laramide thrust faulting. In this province Pennsylvanian strata thicken markedly westward. DEFINITION OF THE PEN NSYLVAN IAN SYSTEM IN Wyoming. The extent of the regions Where each of these WYOMING names is commonly used is shown in figure 4. Nearly everywhere in the State, Pennsylvanian and closely related older strata rest disconformably on Mis- sissippian or Precambrian rocks, and in many places the contact is exposed. In the Bighorn and Wind River Basins, where the Phosphoria Formation (of Permian age) commonly rests paraconformably (as defined by Dunbar and Rodgers, 1957, p. 119) on Middle Pennsyl- vanian Tensleep Sandstone, the upper boundary of the Pennsylvanian System is usually distinct owing to lithologic contrast. Elsewhere the Pennsylvanian-Per- mian boundary lies within the Tensleep Sandstone and The Pennsylvanian System in Wyoming comprises the following formations: Amsden, Tensleep, Quadrant, Casper, Fountain, Hartville, and Minnelusa (fig. 3). The uppermost parts of the Casper, Tensleep (in some areas), Hartville, and Minnelusa Formations are Early Permian (Wolfcamp) in age. The lower part of the Amsden is of Late Mississippian (Chester) age in at least part of Wyoming. The names Weber Sandstone and Morgan Formation have been used extensively in Utah and Colorado and are also appropriate in southern Sweetwater County, where these formations extend into PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING G5 POWDER AREA BIGHORN, WIND RIVER, AND GREAT DIVIDE BASiNS WASHAKIE BASIN LARAMIE BASIN HARTVILLE AREA Ski/SE"? Overlying units Permian Park City and Phosphoria Formations Division | Virgil J; Weber Division 2 Missouri Sandstone Casper % ii = s E E g a _ Quadrant L; g . . . H- g E Des Momes Quartzite Tensleep Sandstone Formation g “c; Dwrsron III 5:: LU if E “- é: ' E 3 s — e E E E . . . 5 Atoka = Ranchester Limestone Member MW“ D'V‘s'm‘s IV! V .2 E Q .2 . E Morrow E Horseshoe Shale Member Formation Division w a e ___ 3 ___ J ,, — h t —_ __ g C es er Darwin Sandstone Member MM .LL,\,,\, . , Mississrppian Underlying units Mississippian Madison Limestone Precambrian basement Guernsey Formation Madison (part) Limestone FIGURE 3.—Correlation chart of Pennsylvanian and closely related strata in Wyoming. the Casper, Minnelusa, and Hartville Formations and is more difficult to determine. EXTENT AND THICKNESS OF THE PENNSYLVANIAN SYSTEM Rocks of Pennsylvanian age were deposited across all Wyoming except in the southeastern part occupied by the Ancestral Rocky Mountains. These rocks have been eroded from the crests of the Laramide uplifts but underlie the remainder of the region. Included among Laramide uplifts is the Sweetwater uplift (Cohee, 1961) in Natrona, Carbon, and Fremont Counties, where the absence of Pennsylvanian strata is not so obvious as on the major mountain ranges. Here Miocene strata rest on Precambrian basement rocks and partly bury the Granite Mountains. The approximate mean thickness of the Pennsyl- vanian System and associated rocks of Mississippian age in the report area is 600 feet. The rocks are less than 500 feet thick in two areas (p1. 3(0)) : the margins of the Ancestral Front Range and Pathfinder uplifts, and the Bighorn Basin—Montana boundary area, Where these rocks thin into Montana and pinch out. Elsewhere they are 500—1,000 feet thick, except south of Jackson in northwestern Wyoming where they exceed 1,000 feet (thickening westward into the geosyncline) and in Sweetwater County Where they thicken into the Eagle trough in Colorado and Utah. The area of thick Penn- is. i l TensieepiSandstone ._.l, """""""" A . ' : 1-/"\ J"J L11... 1 : ' I V and "We ‘ r : ""ess. .i. ............. .l ......... --------- -! “[1350 i i '- ‘msden Formation 3. Weber Sandstone and . ----- ! Morgan Formation i o9 C; I § ! i l 0 50 100 150 MILES FIGURE 4.—Areal extent of Pennsylvanian formations in Wyoming. Boundaries are approximate. Exposed Precambrian rock in existing mountain ranges shown by shading; thrust belt shown by cross ruling. sylvanian strata and strata of Chester age at Jackson extends eastward into the northern Wind River Range area in two elongate troughs. Near Laramie, between the Pathfinder and Front Range uplifts, a trough per- sisted after Morrow time and connected the Laramie and Denver basin areas. G6 LOWER BOUNDARY OF THE PENNSYLVANIAN SYSTEM CHARACTER OF THE DISCONFORMITY Nearly everywhere in Wyoming the Amsden, Casper, and Fountain Formations rest disconformably on the Madison Limestone. Locally in the Ancestral Rocky Mountains they rest nonconformably on Precambrian basement rocks. Typically, the buried Madison surface shows valleys, sinkholes, and caves. In Sinks Canyon, for example, a short distance west of Lander, Wyo., the Middle Popo Agie River flows for several hundred feet underground in a partly exhumed cave in the upper part of the Madison Limestone. In Wind River Canyon, south of Thermopolis, the highly irregular erosion sur— face on the top of the Madison is visible for several miles (fig. 5). The contact can also be seen in many of the FIGURE 5.—Erosional disconformity between the Amsden For- mation and the Madison Limestone in Wind River Canyon, south of Thermopolis, Wyo. The irregular heavy line empha- sizes solution features of ancient karst topography being exhumed by Recent mass-wasting of the valley wall. canyons on the flanks of the Bighorn Mountains. It is prominent and easily accessible at the entrance to Shell Canyon near Greybull and in many other places. In the northern part of the Wind River Range, however, the upper surface of the Madison is paraconformable with the overlying Darwin Sandstone Member, but. a zone 40—50 feet below the top of the limestone contains abundant solution cavities now filled with material clearly derived from the Darwin and Horseshoe Mem- bers of the Amsden Formation. In several localities on the flanks of the southern part of the Laramie Mountains, chert cobbles, probably de- rived from the Madison Limestone, were reported by Maughan (1963, p. 023). This observation suggests that the Madison extended over much or all of this area and was stripped away prior to deposition of the Fountain and Casper Formations on basement rocks. The findings by Chronic and Ferris (1961, 1963) increase the prob- ability that other rocks of Paleozoic age older than SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY Pennsylvanian were once widespread in southeast Wy- oming and were removed by erosion prior to deposition of Pennsylvanian strata. Chronic and Ferris described two small structurally complex outliers (probably in diatremes) of unquestioned Ordovician and Silurian strata surrounded by a Precambrian granitic terrane of the Laramie Mountains near the Colorado border south of Laramie. The character of the disconformity in the subsurface, however, cannot be inspected, but the gross configura- tion of the disconformity can be mapped indirectly from the thickness map of the Darwin Sandstone Member, which is based on data from deep wells. Because the Horseshoe Shale Member in the Wind River and Big- horn Basins has a nearly constant thickness over a wide area, it must closely approximate a tabular layer with nearly flat upper and lower surfaces. If the base of the Horseshoe is nearly flat, the top of the Darwin, on which it lies, must also be nearly flat. Hence, a thickness map of the Darwin must reflect irregularities of the Madison surface on which it rests. Abundant evidence indicates that uplift of the Front Range and Pathfinder elements of the Ancestral Rocky Mountains in southeastern Wyoming was strong enough (particularly in Atoka time) to cause deposition of feldspathic fanglomerate in some adjacent areas of sub- sidence. Nevertheless, the author knows of no place where angular discordance between Pennsylvanian and Mississippian strata can be seen in the field. Absence of this relationship can be partly explained by the fact that in most of the areas where strong uplift took place Pennsylvanian strata rest directly on Precambrian igneous and metamorphic rocks. On the west side of the northern part of the Laramie Mountains, strata of the Casper Formation of Missouri age rest on Madison Limestone without visible angular discordance. TIME SPAN OF THE HIATUS IN WYOMING On the Wyoming shelf the hiatus beneath Upper Mississippian and Pennsylvanian rocks seems to repre- sent a greater time interval in the eastern part of the State than in the western part near the geosyncline, as would be expected. In Bull Lake Canyon in the Wind River Range the age of the upper part of the Madison Limestone was determined by C. C. Branson (1937, p. 651, 653) . Although he referred to a cavernous zone near the top of the Madison as the Sacajawea Formation, a fauna collected from this zone establishes its equiva- lence to rocks of late Osage or early Meramec age in the upper Mississippi Valley (W. J. Sando, written commun., 1965) . Forty-five miles southeast, in the south- ern part of the Wind River Range, and at several other localities in western Wyoming, the oldest beds of the PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING Amsden Formation are of Chester age. (See p. G13 for a discussion of the age of the Horseshoe Shale Member at Cherry Creek.) Hence, evidence indicates that the disconformity at the base of the Amsden is intra- Mississippian. In the Laramie Mountains area the upper part of the Madison Limestone is of Osage age or older (Maughan, 1963, p. 026) . In the subsurface there, rocks of Morrow or younger age rest on the Madison. In exposures in the northern part of the Laramie Mountains, strata of Mis- souri age or younger locally rest on the Madison (Thomas and others, 1953, pl. 9). In the southern part of the range, Pennsylvanian strata rest on rocks of Precambrian age. Hence, rocks of Morrow to Missouri age rest on rocks of Mississippian or Precambrian age in southeastern Wyoming. PALEOGEOMORPHOLOGY OF THE DISCONFORMITY IN NORTHWESTERN WYOMING In the Bighorn and Wind River Basins of northwest Wyoming, tectonic emergence in immediate post-Madi- son time allowed the forces of mass-wasting and erosion to assume the major role in shaping the interface be- tween the Madison Limestone and overlying rocks. Field inspection of the surface of disconformity at the top of the Madison in Wind River Canyon, Shell Can- yon, and other canyons in the region suggests that a surface having tens of feet of relief is present over a wide area. The narrow dendritic areas which resemble stream valleys on the Darwin isopach map (pl. 2(A)) contain thick Darwin, and the broad areas that resemble hills or interfluves contain thin Darwin. The isopach index numbers show that the thickness of rocks of Chester age in this region generally ranges from a few tens of feet to 100 feet. The topographic pattern of Darwin isopachs in northwestern Wyoming is therefore ascribed to deposition on the dissected upper surface of the Madison Limestone. If the isopachs on plate 2(A) are a reasonably valid representation of the regional topography developed in Late Mississippian time on the Madison Limestone in the Bighorn—Wind River region, then the most promi- nent topographic feature in the area at that time was a broad valley trending nearly east in the west-central part of the State. North and south of this valley are areas which apparently were so dissected as to represent late youth or early maturity. (If adequate control data were available for a local area, mesas formed by the original top of the Madison Limestone might be dis- covered. Such mesas, for example, are present on the dolomite of the Arbuckle Group beneath the Cherokee Shale in central Kansas.) Available evidence indicates that dissection resulted both from surface weathering G7 and erosion and from solution at depth, forming caves. Ancient caves in the Madison filled with material de- rived from the Darwin and Horseshoe Members are found in several localities in the Rocky Mountain region. For convenient reference, the ancient stream which flowed in this east-trending valley is herein termed the “Wyoming River” (pl. 2 (A) ) . This river apparently en- tered the area in southern Johnson County and flowed west as a consequent stream down a regional slope pre- sumably caused by broad epeirogenic arching of a wide region east of the area. Thickness values for the Darwin Sandstone Member in northwestern Fremont and south- ern Teton Counties indicate that the Wyoming River flowed in a valley 1—2 miles wide and about 200 feet deep. Farther upstream, control data are too sparse to define precisely the location and shape of the actual valley. In pre—Darwin time the Wyoming River flowed across the present Wind River Basin, the north end of the Wind River Range, and the Jackson Hole country, and debouched into the ocean in the Cordilleran geosyn- cline probably in eastern Idaho. Tributaries north and south of the river presumably formed a dendritic pat- tern in the flat-lying Madison Limestone. The texture of the topography as shown on plate 2(A) reflects only the density of the control net available for mapping. Actual texture can probably best be estimated by in- spection of the disconformity in canyon walls, a surface Which suggests that local features were small and simi- lar to those of a karst topography (fig. 5). Later the sea advanced eastward into the estuary or bay formed by the valley of the Wyoming River. INTERREGIONAL SIGNIFICANCE OF THE DISCONFORMITY ON THE CRATON A comprehensive grouping and classification of cratonic sedimentary strata in North America was made by Sloss (1963). He defined (p. 95) and discussed the term “interregional unconformity” as follows: Only a very small number of the unconformities observable in the cratonic interior of North America can be shown to be truly interregional. When studied in the field or in the sub- surface these exhibit no obvious characteristics Within a limited area of observation which make it possible to separate them from other unconformities. * * * The most important criteria for the recognition of an interregional unconformity are the magni- tude of its geographic scope and its persistence in previously and subsequently subsiding basins. * * * Only when the outcrop belts are integrated through the application of subsurface data and by the recognition of Widespread lithic units of time-strati- graphic significance on a regional and interregional scale is it possible to identify interregional unconformities. Abundant surface and subsurface evidence through- out the Central and Western United States indicates that the diwonformity at the base of the Amsden, Cas- G8 per, Hartville, Minnelusa, and Fountain Formations in Wyoming has interregional significance (Sloss, 1963, p. 102—104, figs. 3, 6; Henbest, 1958b, p. 37—38; Mallory, 1960, p. 23; Kellett, 1932; Walters, 1946, 1958). This hiatus was clearly defined on a continental scale by Sloss (1963, p. 102—3, fig. 6). He referred to it as the uncon- formity at the base of his Absaroka sequence, a craton- wide grouping of sedimentary rocks which includes upper Mississippian, Pennsylvanian, and younger rocks. Pennsylvanian strata form the base of the sequence over wide areas in the Central United States. The discon- formity is present in the Mississippi Valley, midconti— nent, Rocky Mountain, and Colorado Plateaus regions. In much of this vast cratonic territory, the rocks beneath Pennsylvanian strata are limestone (widely of Missis- sippian age but commonly older) upon which a surface with relief of scores or even hundreds of feet was de- veloped. The form of the surface on carbonate strata is commonly karst in a stage of late youth to early maturity. (See Walters (1946, fig. 7, pl. 1) for a typical example in central Kansas.) Although the erosional surface undoubtedly was once physically continuous from central Ohio, Kentucky, and Tennessee to Mon- tana, Wyoming, and Colorado, the duration of subaerial weathering in the cratonic United States was not every- where the same. In the orthogeosynclines (Appalachian, Ouachitan, and Cordilleran), deposition seems to have been con- tinuous. In most areas, establishing the Mississippian- Pennsylvanian contact is difficult. Marine waters began their transgression from the orthogeosynclines over the continental interior during Chester and Morrow time and spread gradually over the peripheral parts of the craton. By Middle Pennsylvanian time most of the United States had been inundated. The higher, central part of the region was submerged last, allowing pre- Pennsylvanian rocks in these areas to weather longest. I11 parts of Kansas and Nebraska, for example, where strata of Des Moines age lie on rocks of Mississippian to Precambrian age over wide areas, weathering and ero- sion lasted through Atoka and Morrow time and prob- ably much or all of Chester time. In the Wyoming thrust belt and eastern Idaho, Missis- sippian fossils are found above sandstone which has been correlated with the Darwin of the Wind River Range area (Wanless and others, 1955, p. 31—34; Sando and Dutro, 1960, p. 125). AMSDEN FORMATION LITHOLOGIC UNITS Darton (1904, p. 396—397) named the Amsden Forma- tlon for exposures on Amsden Creek, a small tributary SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY of the Tongue River northwest of Dayton (sec. 32, T. 57 N., R. 86 W.), Sheridan County, Wyo. In 1906 Darton (p. 31) described the Amsden in greater detail as follows: * * * Throughout the Bighorn Mountains there is at or near the base of the formation a deposit of red shale, which lies directly on the upper limestone of the Madison formation, ex- cept locally to the south, where it is separated by a bed of brown or gray sandstone. Above the red shale there is a variable suc- cession of pure, white, fine-grained, compact limestone, gray to brownish sandstones, red shales, and cherty limestones. At the top there usually are slabby sandstones, especially to the south- west where they take the place of the greater part of the lime— stone beds. The local basal sandstone above referred to attains its maxi- mum development in the canyons of Otter Creek, where it has a thickness of 80 feet. In the vicinity of Tensleep Canyon its thickness is over 40 feet. It is coarse grained, of gray color, and cross-bedded * * * The basal sandstone of the lower unit in northwestern Wyoming was named the Darwin Sandstone Member of the Amsden by Blackwelder (1918, p. 422—423) after Darwin Peak in the Gros Ventre Range, which is capped by this sandstone. (In Blackwelder’s paper the name is spelled “Dorwin,” presumably a typographical error.) The nomenclature of Darton and Blackwelder is well suited to the Amsden Formation throughout the Middle Rocky Mountain region and can be applied even to superficially atypical sections like those at Bull Lake and Dinwoody Canyons. The three lithologic divisions of the Amsden described by Darton have been referred to as the Darwin Sandstone Member, the middle red shale unit, and the upper cherty limestone unit. To facilitate reference to the latter two units, and to reinforce the identity and stratigraphic relations of the units within the Amsden as established by Darton, new names are here introduced for the two unnamed units of the Amsden Formation above the Darwin. The name Horse- shoe Shale Member is proposed here for the middle red shale. This name is taken from Horseshoe Mountain (sec. 27, T. 56 N., R. 87 W.) 7 miles southwest of Day- ton. The type section of the Horseshoe Shale Member is in the SE14 sec. 33, T. 57N., R. 87 W., on Amsden Creek 4 miles west of Dayton. The name Ranchester Limestone Member of the Amsden Formation is pro- posed here for the upper cherty limestone unit of the Amsden. The name is taken from the town of Ranchester (secs. 18—19, T. 57 N., R. 85 W.) 6 miles northeast of Dayton. The type section of the Ranchester Member is at the same location as that of the Horseshoe. A graphic section of these two members measured at Amsden Creek is shown in figure 6. Because the Darwin Sandstone Member is locally absent at Amsden Creek, the Horseshoe and Ranchester Members compose the total Amsden Formation at the type section. PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING m o. a: mg 0.: 24-! ”o m‘” 24" CE “"3 a)“ 5: Z 0—0) l—‘g S E 2 > 3.. .1 L00 5 “’9 z Em z -"\ E 53': 5.0 mE z 5% < c _ m z a: < > 5 -J .o 5 E z 0 Z 2 Lu 3 g 0. Z a)“ ‘33? <5 Egg G.) 22 E: s; ‘38 3'; n._, °.—. 5 0.). g 3:; 0 5") ‘5 ‘81: w = g "2% ‘E" as U) on.) D 3 '5 <0 (03.1 ‘5 I2 0 m c D E L a: E a L” I: 0 : ‘—“ 5 E g __._. -,—.=-m—-_-_.__ z L; E u, 4.1111m1-a.a-::..: 8 ‘1’ z ‘E: n: '9, o O < < O 2:; 2 Z a)“ 13‘“ << fig :2 _ 2% «1 2% we (I); .<_> 430 CD fl._| 0" ._E E> 5213 2 Ed, ‘6’)? 32 E 32 -z .0 a. mm o __ 3 <12 (I) ‘2 E E _ a) w Q. C: o. 58 3° _ .9“ _-S 3 '03 '00 (I) 5E 2.- Q '3 _J E FIGURE 6.—T‘ype and reference sections of the Amsden Formation. A, Type section of the Amsden Formation and of the Horseshoe Shale and Ranchester Limestone Members at Amsden Creek, northern Bighorn Mountains, Sheridan County, Wyo. B, Reference section for the Amsden Formation and for the Darwin Sandstone, Horseshoe Shale, and Ranchester Limestone Members at Tensleep Canyon, west, flank of Bighorn Mountains, Washakie County, Wyo. G9 G10 A reference section for the Amsden Formation is also proposed. This section is near the lower end of Ten- sleep Canyon, near Tensleep, about half a mile upstream from the Wigwam Trout Rearing Station on the north side of US. Highway 16. This reference section is more satisfactory than the type section at Amsden Creek for several reasons: 1. All three members of the formation are present and well exposed. 2. Lithology of each member is typical of the Amsden in Wyoming. 3. Contact with the underlying Madison Limestone is marked by regional disconformity. 4. Contact with the overlying Tensleep Sandstone is at a sharp local disconformity. 5. Section is visible from the highway and easily ac- cessible for study. Graphic sections of the Amsden Formation at Tensleep Canyon and at Amsden Creek are shown in figure 6. AGE Controversy regarding the identity and the age of the Amsden Formation has existed for many years. Much of the uncertainty can be traced to changes in nomen- clature and to miscorrelations (Branson and Greger, 1918; Branson, 1937, 1939; Branson and Branson, 1941; Wilson, 1962) in which the Amsden Formation was con- fused with the Madison Limestone or merged with the Tensleep Sandstone. Recent work shows that the Amsden Formation is Pennsylvanian and Late Missis- sippian (Chester age) in western Wyoming (Shaw and Bell, 1955; W. J. Sando, written commun.,-1965). Fossils allegedly collected from the Amsden Forma- tion since Darton named and described it in 1904 can be divided into two major groups: 1. Those collected from the Madison Limestone which are Meramec in age and, therefore, have no bear- ing on the age of the Amsden Formation as it was defined by Darton. 2. Fossils actually collected from the Amsden F orma- tion, most of which are Pennsylvanian but some of which have a Mississippian or Mississippian and Pennsylvanian aspect. In 1954 Burk systematically studied the literature on the Amsden Formation and did additional work in col- laboration with C. A. Biggs and others. He stated (p. 4) : There is no previously described fauna of known age which is comparable to the assemblage of fossils which has gradually been taken from the Amsden. In order to establish the age of the Amsden, it is necessary, therefore, to examine the known ranges of all those forms which have been collected from beds above the Darwin sandstone. All of the Madison fossils, mis- takenly assigned to the Amsden, are omitted from consideration. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY as well as those forms which were identified with hesitation. New species taken from the Amsden are also excluded from this tabulation. He concluded (p. 5) : The greatest number of species are Pennsylvanian in age. There are no identifications which would conflict with a Penn- sylvanian age for the Amsden, and there is none which sup- ports an exclusively Mississippian age. If all the forms now generically identified, having a range including both the Mis- sissippian and Pennsylvanian, were found to have been taken from the lower part of the Amsden it might suggest a late Mis- sissippian age for this part of the formation, but an examina- tion of all the faunal lists and the present collection shows that there are as many exclusively Pennsylvanian species known from just above the Darwin sandstone as there are genera and species which have a range including both the Mississippian and Pennsylvanian * * *. The faunal evidence pointed out in this paper makes it necessary to abandon the Mississippian age pre- viously accepted for the Amsden, and confirms its Pennsylvanian age. The fossil collections reviewed and collected by Burk came from the Wind River, Bighorn, Gros Ventre, and Absaroka Mountains. If Burk’s conclusions are valid, that part of the Amsden Formation above the Darwin Sandstone Member in these areas is of Pennsylvanian age. Gorman (1962) collected and identified (with the collaboration of M. L. Thompson) fusulinids from the Amsden in the northern part of the Bighorn Mountains; his conclusions agree with Burk’s. The work of Shaw and Bell (1955) and W. J. Sando (written commun., 1965) indicates, on the other hand, that the lower part of the Amsden is of Late Mississippian age (Chester) in the Wind River Range and Washakie Mountains and westward. Fossils from the Amsden Formation at Darwin Peak in the Gros Ventre Mountains have recently been stud- ied by W. J. Sando and J. T. Dutro (written commun., May 1965). The sequence of faunas is similar to that at Cherry Creek. The age significance of the Cherry Creek section is probably valid for the entire Wind River Range—Gros Ventre Mountains area. Hence, it seems that the Darwin Sandstone Member is of Late (Chester) Mississippian age; the Horseshoe Shale Member is of Late (Chester) Mississippian and Early (Morrow) Pennsylvanian age; the Ranchester Limestone Member is of Early and Middle (Morrow and Atoka) Pennsyl- vanian age. NOMENCLATURE AND THE AMSDEN PROBLEM As stated earlier, Darton (1904, p. 396—397) recog- nized two major lithologic units in the Pennsylvanian System of central Wyoming. He applied the name Ten- sleep to an upper massive sandstone unit and the name Amsden to a lower sequence of limestone, red shale, and basal sandstone that lies on the Madison Limestone. In PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOWG 1906 (p. 31—34) Darton more explicitly described the Amsden. In 1918 (p. 422—423) Blackwelder designated the basal sandstone unit described by Darton as the Darwin Sandstone Member of the Amsden Formation. Also in 1918, however, E. B. Branson and Greger (p. 310—311) discussed strata which they called Amsden at Bull Lake Canyon, where the Amsden and the Madison are well exposed in the canyon walls, and at Cherry Creek, near the Canyon of the Little Popo Agie River (sec. 19, T. 31 N., R. 99 W.) southwest of Lander, where poorly exposed red sandy siltstone and shale of the Horseshoe Shale Member of the Amsden Formation lie on the Madison Limestone. Branson and Greger’s text (p. 310—311) indicates that part of the Cherry Creek fauna was apparently collected from a ledge in the Horseshoe Shale Member and part from purple lime- stone float. The Bull Lake fauna was collected from the upper 76 feet of the Madison Limestone, which they called Amsden. The Darwin Sandstone Member at the base of the Amsden overlying this zone in the Madison is referred to by them as the Tensleep Sandstone. These authors, therefore, correlated the middle red shale (Horseshoe) of the Amsden at Cherry Creek (where the Darwin Member is absent) with the upper part of the Madison Limestone at Bull Lake Canyon despite gross differences in lithology and stratigraphic position. This miscorrelation and the confusion which it engendered are responsible for much of the contradiction in the literature. In a 1937 paper discussing the Amsden Formation and the Madison Limestone at Bull Lake Canyon, C. C. Branson (p. 650—652) also failed to recognize the pres- ence of Blackwelder’s Darwin Sandstone Member of the Amsden Formation, and followed E. B. Branson and Greger’s terminology in that be correlated the Darwin with the Tensleep and extended the name Amsden down into the Madison Limestone. C. C. Branson called (1937, section, p. 651) Darton’s Amsden Formation the “Up- per Amsden” and the uppermost 18 feet of the Madison Limestone the “Lower Amsden.” Below Branson’s “Lower Amsden” is 43 feet of Madison Limestone with a reported Meramec fauna and 2 to 11 feet of solution— breccia limestone which he named the Sacajawea For- mation. In 1939 (p. 1202) C. C. Branson expressed dis- satisfaction with Darton’s dual Pennsylvanian desig— nations——Tensleep Sandstone and Amsden Formation— and proposed that the name Tensleep be redefined to include all of Darton’s Amsden Formation and that part of the Madison Limestone which Branson termed “Lower Amsden” in 1937. Thus, by Branson’s definition the revised Tensleep Formation rests on his Sacajawea Formation. In 1941, E. B. Branson and C. C. Branson (p. 131—132) used the same nomenclature as C. C. Bran- G11 son had in 1939 and proposed that the name Amsden be abandoned. The confusion has recently been intensified by Wilson (1962). For example, in Wilson’s figures 7 and 8 the name Sacajawea (a zone in the upper part of the Madison Limestone) is applied to the Darwin Sandstone Member, and concentric age patterns show that the Ran- chester Limestone Member of the Amsden Formation ranges in age from Morrow to Virgil and grades later- ally into the Darwin and Tensleep. Todd (1964) dis- cussed petrology but took his stratigraphy from Wilson. Because the Amsden Formation is a simple tripartite sedimentary unit with only modest regional facies vari- ation over a wide area, the nomenclature proposed by Darton fits well with observed facts nearly everywhere in central and western Wyoming and has priority over the later proposals by Branson and Branson (1941). DARWIN SANDSTONE MEMBER Lithology and thickness.—The Darwin Sandstone Member is a gray, white, or cream-to-salmon sandstone; locally it is brick red or has brick-red blotches or specks. Its most conspicuous features, aside from color, are crossbedding, a high degree of sorting, and purity of composition (fig. 7). Minute quantities of arkose were noted by Bishop (1957, table 3) in the D‘uBois area in the northern Wind River Range of Wyoming, but gen- erally the member is composed almost entirely of fine to medium quartz sand. The cement is silica or, locally, calcite (Agatston, 1954, p. 516) . A high degree of sorting and a paucity of heavy minerals suggest that the Darwin was derived from preexisting sandstones (Bishop, 1957). The thickness of the Darwin ranges from 0 to 145 feet and changes abruptly from place to place (pl. 2(A) ). Age.—No fossils are known from the Darwin Sand- stone Member, but its age is late Chester in the Wind River Range. At Sinks Canyon, near Lander, the Dar- win is overlain by the Horseshoe Shale Member; but 10 miles southeast, at Cherry Creek, the Darwin is absent owing to pinchout or gradation into the lower part of the Horseshoe, which there lies directly on the Madison Limestone. (See p. G13.) Depositional history—As geosynclical marine waters transgressed eastward onto the Wyoming shelf in late Chester time, the Wyoming River system became a shal- low estuary or bay. Sand from an exotic northern or eastern source was deposited in the bay in earliest Dar— win time and continued to accumulate as the Madison topography was inundated. Because relief on the Mad— ison surface averaged about 150 feet over thousands of square miles, relatively small rise of sea level could have produced widespread inundation. Alternating pe- G12 B FIGURE 7.——Crossbedd.ing in the Darwin Sandstone Member, Tensleep Canyon, Bighorn Mountains, Wyo. A, large-scale crossbedding; B, small-scale crossbedding. Large-scale cross- bedding is more common. riods of emergence and submergence, possibly caused by eustatic changes, may have allowed wind and waves to build a complex network of crossbedded dunes, beaches, and bars. The extent and thickness of the Darwin in the Big- horn Mountains area deserve special comment. The pres- ence of an unusually irregular featheredge of the Dar- win on the west flank of the Bighorn Mountains (pl. 2(A) ) is established by the control net in the area; the presence of an unusually thick local patch of the Dar- win in central Johnson County near Buffalo needs ex- planation. Apparently, in Darwin time the Bighorn Mountains area was uplifted as an asymmetrical arch or fault block with a minimum displacement of about 150 feet. The Wyoming River, flowing south through Johnson County and carrying sand from an exotic source, apparently was antecedent to the arch or block SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY and was ponded for a time by this tectonic event. In the resulting lake, the Darwin accumulated to a maximum thickness of about 150 feet. Another local patch in cen- tral Sheridan County may have had a similar origin. On the west side of this arch or block near Mayoworth, the irregular limit of the Darwin may have been ac- centuated by erosion of the Darwin shortly after deposi- tion. Support for the assumed existence of a fault or sharp flexure in the Bighorn Mountains area near Mayo- worth is provided by the existence later in Morrow time of the Casper peninsula of the Front Range uplift (pl. 2 (B) ), whose trend generally paralleled the Mayoworth fault or flexure. Maughan (1967) noted the existence in Permian time of a tectonic alinement extending from the northeast flank of the Laramie Mountains into the Bighorn Mountains, and this may possibly have been a rejuvenation and southward extension of the Mayo- worth alinement. HORSESHOE SEALE MEMBER Lithology and thickness.—The lithologic unit above the Darwin Sandstone Member is typically red shale and siltstone. It is widespread in northwestern Wyom- ing (pl. 2 (B) ), and its striking color is a useful criterion for identifying the Amsden Formation in outcrop and in well cuttings. Where the Darwin Sandstone Member is absent, the Horseshoe Shale Member rests directly on the disconformity at the top of the Madison Lime- stone. Beds of sandstone, limestone, and dolomite lo- cally compose the bulk of the member as in the Cities Service Sprecher 1 well, sec. 22, T. 36 N., R. 82 W., Natrona County, and in Dinwoody and Bull Lake Can- yons, but red shale predominates nearly everywhere else. Mudstone, claystone, shale, siltstone, and thin sand- stone and carbonate beds are common auxiliary con- stituents. The shale has paper-thin to massive struc- ture, and thin and regular bedding, and commonly has chunky or blocky shale. The color is usually bright, commonly ranging from purple or maroon to brick red'; locally, yellowish and light—pinkish-gray shale is present. At Dinwoody and Bull Lake Canyons the shale is gray green and is copiously interbedded With thin beds of dolomite and sandstone. A distinctive feature of the Horseshoe Member, described in detail by Mundt (1955, p. 68—73), is red to black ferruginous pisolites in the middle of the unit. This “buckshot pellet” zone is widespread in Wyoming and is identifiable on surface exposures in the southern part of the Laramie Mountains, on both sides of the Bighorn Mountains, on the east side of the Wind River range, and elsewhere in the intervening area. Individual pellets are typically about 10 mm in diameter, deep red to black, spherical to ellipsoidal, and in some areas PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING devoid of discernible internal structure. J. F. Murphy (oral commun., 1964) reported that in the Dinwoody— Bull Lake Canyons area well-developed concentric structure is common. A trace of the ferruginous pisolitic zone has been identified in a well at Worland dome, T. 48 N., R. 92 W., Washakie County, on the east side of the Bighorn Basin (Agatston, 1954, p. 518). The isopach map of the Horseshoe Member of the Amsden (pl. 2 (B) ) contrasts markedly with that of the Darwin Member. The Horseshoe is a nearly tabular layer, averaging 7 5 feet in thickness, that was deposited across a relatively flat surface. Locally the thickness exceeds 100 feet; on the margins of the area, 150 feet. The Horseshoe is not so extensive areally as the Darwin. (See pl. 2(3) , northern Natrona and southern Johnson Counties.) Age.—A poor exposure at Cherry Creek was visited by Shaw and Bell (1955), who collected fossils from two ledges in the Horseshoe Shale Member. Here the Madison Limestone forms a prominent ledge; the Dar- win Sandstone Member is absent. Directly above the Madison a few tens of feet of red beds are discernible in the hillside. Forty-seven feet above the base of the Amsden is a sandy limestone 1 foot thick from which Shaw and Bell collected a shell fauna to which they assigned a Chester age. They also reported the presence of a calcareous silty sandy zone 18 feet above the sandy limestone. From this sandy zone they collected a brachiopod fauna to which they assigned a Pennsyl- vanian (Morrow or Atoka) age, partly on the basis of evolutionary changes. Shaw and Bell’s faunal lists were reviewed by J. T. Dutro, Jr. (oral commun., 1963), who would also ascribe a late Chester date to the forms listed from the lower of the two beds and a transitional Mississippian and Pennsylvanian age to the fauna listed from the upper one. The fossils listed by Shaw and Bell (1955, p. 334) from the lower of the two beds are as follows: 00m- posz'ta trinwclea; Oom‘posita sp. indet.; Eumetria oer- neuz'liama; Eumetria sp. indet.; Gflfl‘ithides moor-66; Linop’r'oduictus cronez’si; Mfg/(Iliad sanctiludom'oi; “07"- thotetes” sp.; ?Palaeonez'lo amsdenensis; Produatus phéllz'psz'?; “Pustula genem’evemis”; Spirz'fer wellem'; spirifer Sp. indet.; “Schizophoria swallow”; “Spiri- fem'na” browni; and Torym'fer cf. T. setigem. The fossils listed from the upper bed are as follows: Composz'ta tm'mwlea; Composite; sp.; Linoproductus sp. indet. ; Spar-{fer opimrus; Bellerophontid gastropod ; Tet- racoral; Composz'ta ooata; Composita subtilita; 00m- posz'ta sp. indet.; Eumetm'a sp. indet.; Spirifer sp. indet.; Wellerella sp. indet.; Dictyoclostus sp. indet.; E mnetrz'a sulcata; and Wellerella osagemis? 269—091 0—67—2 G13 W. J. Sando (written commun., 1965) recently col— lected, from beds above the Darwin in the Washakie Mountains, fossils which he considered to be of Late Mississippian age. Hence, the lower part of the Horse- shoe Shale Member in this area seems to be of Chester age, and the upper part, Morrow. Environment and provenance.—The nearly constant thickness and lithology of the Horseshoe Shale Member in the Bighorn—Wind River region suggests that in post-Darwin Morrow time this area was a stable shelf or platform. Because the area was bounded on the west by the Cordilleran geosyncline, on the south by the tectoni- cally active Ancestral Rocky Mountains and associated troughs, and on the east by a shallow basin in north- eastern Wyoming, the term “platform” seems appropri- ate. For this reason the term “Bighorn—Wind River platform” is used in this report as a convenient designation. The most obvious source for the red material would seem to have been the area of exposed Precambrian rocks in the Ancestral Rocky Mountains in southeastern Wyoming. If this were the source, however, strata of the Horseshoe Member should become markedly coarser grained in that direction. Instead, the rocks of equiva- lent age near the Front Range uplift (Casper Forma- tion) are mostly sandstone and limestone (pl. 2(8)). The coarse arkosic conglomerate typical of rocks younger than the Horseshoe in this area is absent. Because Pennsylvanian seas covered only parts of Montana and did not extend very far into south-central Canada (Sloss and others, 1960, p. 34), a wide expanse of Mississippian and older limestone was exposed to weathering and solution in Chester and early Morrow time. A red residual soil may have formed on this lime- stone terrane and then been transported in Chester and Morrow time southward into Wyoming and deposited on the Bighorn—Wind River platform. If this weather- ing and transportation occurred, red soils may also have formed on the weathered surface of the Madison farther south in Wyoming and should now be in place at the base of the Darwin Sandstone Member; but no evidence of this has been found. The thin regular-bedded aspect of the Horseshoe Member beds in many areas, the lack of ripple markings, and the local presence of carbonate beds indicate that the member was deposited in quiet water of moderate depth under unusually stable tectonic conditions. If the source of both the Darwin and the Horseshoe sediments was exotic and distant, an event in the prove- nance area could have been responsible for the change of material being carried into the region. That the origin of red color in sediments is not entirely understood fur- ther complicates the problem. At least it can be stated G14 that at some time the red beds of the Amsden had access to abundant oxygen. Perhaps the material was oxidized in the source region, or possibly the material was in a reduced state at the source and oxidized during trans- portation. The color may have been derived during the process of deposition by shallow inundation alternating with exposure to the atmosphere on a broad depositional shelf. Some combination of all three processes must also be considered. Local areas of green shale must have had a different history. RANCHESTER LIMESTONE MEMBER Lithology and thickness.——The Ranchester Limestone Member of the Amsden Formation is composed pre— dominantly of carbonate rock that is usually gray, tan, pink, or purple, dense or finely crystalline, massive, and cherty. It is characterized by pink to dark-red shale partings or shaly limestone beds that crop out as a series of massive ledges, and by c‘hert that commonly litters the red shale slope of the Horseshoe below (fig. 8). Red pig- ment from overlying iron-rich shales stains the lime- stone ledges to give the Ranchester a pink to red tinge in many areas. Locally sandstone is the dominant lithology in the upper part of the member. FIGURE 8.—North wall of Tensleep Canyon near trout-rearing station showing the top of the Madison Limestone (M), the Darwin Sandstone (D), Horseshoe Shale (H), and Ranchester Limestone (R) Members of the Amsden Formation and the lower part of the Tensleep Sandstone (T). For comparison see Darton (1906, pl. 16B). The carbonate beds of the Ranchester range in com- position from limestone through dolomitic limestone to dolomite. Generally the carbonate beds of the unit are separated by white to gray sandstone, siltstone, and clay- stone, which commonly grade into argillaceous, silty, or sandy limestone. Thin beds of pink, purple, red, lav- ender, or light-green shale are also common. Mundt SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY (1955, p. 75) noted that there seems to ‘be no uniform occurrence of dolomite and limestone within this mem- ber in Wyoming, but he reported that Nieschmidt (1953) showed 24 subsurface and 2 surface sections in central Montana where an upper dolomitic part overlies limestone. Chert occurs as irregular nodular masses, as fracture fillings, and as banded beds; it may be red, orange red, white, gray, or tan (Agatston, 1954, p. 519). Bedding in the carbonate strata ranges from thin to massive; shale partings are common; stylolites and local intraforma- tional disconformities are present but not abundant. Intergranular porosity in the dolomite facies, pseudo— oolitic texture, and vugular porosity are discernible lo- cally on the outcrop and, if in suitable association with other factors in the subsurface, may form petroleum or natural-gas reservoirs. The thickness of the Ranchester Limestone Member generally ranges from 0 to 250 feet (pl. 2( 0) ) and aver- ages about 100 feet. In a belt trending northwest from Casper to the Montana line, the member is less than 100 feet thick; furthermore, isopach patterns of this belt show markedly northwest trends. During deposition of the Ranchester this belt may have been a platform or gently uplifted block related to the fault or flexure southwest of Mayoworth, which was active in Darwin time. The thickness of the Ranchester in the western part of the Bighorn Basin and in the Wind River Basin averages about 150—200 feet; there, a northwest trend is also apparent in the isopach pattern. Age.—The carbonate beds of the Ranchester Lime- stone Member have yielded abundant microfossils which have been useful in establishing its age. At Amsden Creek, in a light-gray dense limestone bed 5 feet above the base of the Ranchester, Gorman (1962, p. 25, 28; 1963, p. 69) collected a fusulinia fauna that he identified as Millerella sp., M. inflecta, Paramillerella pingmis, P. circuli, P. ampla?, P. advena, P. sp., and Ncmkinella, and to which he assigned a Morrow age. From a com— parable horizon along U.S. Highway 14 near Tongue River Canyon (a few miles south of Amsden Creek), Gorman collected Paramillerella advena and N (m- kinella?. A macrofauna was collected by Eliot Blackwelder (Love, 1954) at Darwin Peak from a horizon at the base of the Ranchester. The forms were identified by G. H. Girty as Early Pennsylvanian. Henbest (1954, p. 50—51) identified a protozoan fauna collected at five localities from the upper part of the Amsden Formation in south-central Montana, just north of the report area. He listed the following forms: Endothym sp., Bradyina, sp., Millerella? sp., Profusu- PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING linella sp., Olimacammina sp., Tetratam's sp., and Pseudos taflela Sp. He commented as follows: The species of C’limacammimz, Bradyma, Endothym, and Tetra- tam‘s in these collections have not been clearly defined and are not known to be restricted to the Early or Middle Pennsylvanian, but an association in such abundance is common in rocks of Atoka and Des Moines age and rarely found in older and in younger rocks. All those genera probably range from Mississip- pian to Permian * * *. The Amsden species of Pseudostajfella is minute but has regular growth form. The earliest record of this genus is in rocks of Atoka age. The most characteristic foraminifer in these collections of the Amsden is a highly specialized species classed with question as Profusulinella. No species of fusulinids resembling this form have been recorded in rocks known to be older or younger than the Atoka. All of the earliest Atokan relatives of this species are decidely more primitive, but equally complex fusulinids ap- pear before the end of Atoka time * * *. No evidence of Mississippian age was recognized in these six collections. Atoka, Pennsylvanian, age seems to be rather well substantiated. Nevertheless, the foraminiferal fauna is so pe- culiar and the knowledge of Late Mississippian and Early Pennsylvanian foraminifers so incomplete that this age deter- mination should not be accepted without reservation. In addi- tion, it must be emphasized again that this conclusion on the age of the Amsden applies only to the parts of the formation represented by these collections. Love (1954, col. 14) listed a macrofauna from a lime- stone interbedded with sandstone, which may be upper- most Amsden or lowermost Tensleep. He cited an un- published manuscript by Eliot Blackwelder as source of the information and indicated that the fossils were collected by C. L. Breger at Dinwoody Lakes, in the northwestern part of the Wind River Range, and iden- tified by Gr. H. Girty, who considered the age of this assemblage as Atoka. The available evidence seems to indicate that the lower part of the Ranchester Limestone Member is Morrow and that the remainder is of Atoka age. Tectonic activity during Ranchester time—Tongues or lenses of elastic rock in the Ranchester Limestone Member probably derived their material from four source areas: the Pathfinder uplift, the Front Range uplift, an apparent source on the site of the Wind River Range, and an apparent source in Montana, north of the report area. Of these source areas, the Pathfinder is the most completely documented and had greatest areal extent and elevation in Atoka time (Mallory, 1963). In 30 wells and surface sections on the site of the Path- finder uplift, the Ranchester Member is missing; an ad- ditional 35 control points in the surrounding area show that the Ranchester thickens away from the uplift nearly everywhere at a uniform rate of 10 feet per mile (fig. 9, pl. 2(0) ). In southern Albany County the rate of thickening of rocks of equivalent age increases to 65 feet per mile south of Laramie. The areal geology indicates that the Pathfinder uplift G15 consisted of four distinct parts in Atoka time (fig. 9; Mallory, 1963) : 1. A northeastern segment where Madison Limestone of Mississippian age formed the surface rock. 2. A central belt which bore a cover of carbonate rock, shale, and sandstone of Morrow age (the lower part of Casper Formation) (pl. 2(B), fig. 9). 3. A southern extension Where igneous and metamorphic rocks of Precambrian age were exposed. 4. A western extension where Madison Limestone prob- ably formed all or much of the surface. At present, Tertiary rocks lie on Precambrian over much of the western extension. However, the fact that Pennsylvanian strata rest on the Madison Limestone in many places at the edge of the Sweetwater uplift sug- gests that Mississippian strata extended partly or en— tirely over the western part of the Pathfinder uplift at the close of Atoka time and were removed by erosion following the Laramide deformation. Figure 9, there— fore, shows Madison Limestone on the western extension of the Pathfinder uplift; local inliers of older Paleozoic and basement rocks may have been present in stream valleys. The composition and location of the tongues or aprons of elastic rock marginal to the Pathfinder uplift are related to both the areal geologic pattern of the uplift in Atoka time and the degree of tectonic activity in each of the four divisions. Moderate uplift of the western extension caused sandstone and shale of the Horseshoe Member to be removed and redeposited as tongues of clastic material in the Ranchester sea. Strong uplift of the southern extension, in Albany County, allowed igneous and metamorphic rocks of Precambrian age to be actively weathered, eroded, and redeposited on the southern shores of the uplift as arkose of the Fountain Formation. The northeastern segment was moderately uplifted, causing beds of Morrow age to be removed and redeposited in the Atoka sea, but uplift was not adequate to cause the Madison to be stripped away. The central belt must have been a lowland nearly awash in the sea, because clastic strata of the lower part of the Casper were not removed and adjacent clear marine waters were therefore uncontaminated by mud and sand. Erosion of the Front Range uplift contributed abundant arkosic gravel and sand to the trough in the Laramie basin area. This material mingled with similar material eroded from the south end of the Pathfinder uplift. Arkosic conglomerate of the Fountain Forma- tion of Atoka age is present as far west as the vicinity of Rawlins in the Sinclair Oil and Refining Unit 1 (sec. 28, T. 21 N., R. 86 W.) (pl. 1, B—B’; pl. 2(0)). In G16 SHORTER 108 ° 107° CONTRIBUTIONS TO GENERAL GEOLOGY I 106° '0 '85 105° 09% CONVERSE 42°.— \_/;P\1 / / \> \ \_ /l \ _\ we my: \ C \/\’/ ‘ // \ EXPLANATION AREAL GEOLOGY OF PATHFINDER UPLIFT Lower part of Casper Formation of Morrow age VII/I Madison Limestone of Mississippian age 7 4 - a"... o°'an \ \ \ l / R \ / \: _/’\ P /\/—\?t\ />\\/\\/\ l\\/\0 0 V)‘ no , l\ ” Precambrian granitic and metamorphic rocks LITHOLOGY OF RANCHESTER LIMESTONE MEMBER OF AMSDEN FORMATION .u'u: " o \/ 'n Arkosic congloerate, sandstone, and siltstone 41 ° Sandstone . Approx1mate contact \.og’\ Shaly sandstone 25 Isopachs 4 Showing thickness, in feet, of Rancheste'r O 10 20 30 0 MILES calcareous sandstone Limestone Member. Interva': is 50feet .125 A 180 E3 Well Surface section Shale Control points Carbonate rock Number is thickness QfRanchester Limestone Member. P, paleontologic data are available from Pennsyl- vanian rocks to aid in correlation FIGURE 9.—Areal geology of Pathfinder uplift at end of Atoka time and thickness and lithology of Ranchester Limestone Member of Amsden Formation in area surrounding the uplift. Modified from Mallory (1963, fig. 195.2). PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING southwestern Carbon and southern Sweetwater Coun- ties, scattered control suggests that a belt of sandy calcareous shale fringes the north and west margins of the Front Range uplift. Because coarse material is scarce, uplift in this area presumably was less intense than in the vicinity of Laramie. The presence of rocks of Atoka age along the south edge of Sweetwater County is questionable. Verville and Momper (1960) reported only 67 feet of strata of Atoka age in the Mountain Fuel Unit 1 well in sec. 21, T. 16 N., R. 104 W. Although some workers (Thompson, 1945, p. 26—27; Baker and others, 1949, p. 1181—1182) cited the possible occurrence of Atoka rocks along the south side of the Uinta Mountains, Sadlick (1955, p. 58) found that strata containing Des Moines fusulinids lie disconformably on strata containing Morrow fusulinids along the north edge of the range. Thus, Atoka rocks are thin or absent along the Colorado- Wyoming boundary, and the Front Range uplift may have extended westward into this region during Atoka time. Shale, sandstone, and shaly or sandy carbonate rocks in the Ranchester Limestone Member around the Wind River Range (pl. 1, A—A’; pl. 2(0)) suggest that an island may have existed here in Atoka time, perhaps as a detached segment of the Pathfinder uplift and, hence, the northwesternmost element of the Ancestral Rocky Mountains. If so, the Ancestral Rocky Mountain tec- tonic trend extended for 1,000 miles in a nearly un- broken chain from the cratonic margin of the Cordil- leran miogeosyncline near Jackson Hole southeastward to the cratonic margin of the Ouachita geosyncline near Ardmore in southern Oklahoma. Tongues of calcareous elastic rock in the northern part of Wyoming (pl. 1, B—B’, 0-0’; pl. 2(0)) were apparently derived from a source or sources north of the report area. The materials in these tongues may have been derived from rivers flowing southwest from south- ern Canada and been deposited as deltas and turbidites in the otherwise clear Atoka seas of Wyoming and southern Montana. CASPER FORMATION The Casper Formation, which occurs in the Laramie basin and on the flanks of the Laramie Range in the southeastern part of the State, was named by Darton in 1908 (p. 418—430). He proposed the name for a massive limestone and sandstone sequence which, with the F oun- tain Formation, composes the Pennsylvanian System in the Laramie Mountains area. These rocks are the southeastward extension of the Amsden and Tensleep Formations but differ enough lithologically to warrant another name (Darton, 1908, p. 418). G17 Like many Pennsylvanian formations in the Western United States, the Casper Formation differs markedly in lithology from bed to bed and from area to area. Hence, a detailed description of its lithology at one place does not serve as a competent identification at another. In general, however, the Casper is predomi- nantly composed of carbonate rock and sandstone, com- monly in massive beds. The color is usually gray, bufl", or white, but shades of pink and red are common. The carbonate layers are typically gray, tan, pink, or white, dense to finely crystalline, massive, locally oolitic lime— stone, dolomitic limestone, or dolomite. Locally, they are sandy, vuggy, and cherty. The chert is red, pink, tan, brown, black, white, or gray. Locally, carbonate beds are purple to lavender, dense to lithographic, and thin bedded. The sandstone layers are red, pink, white, buff, or gray, commonly crossbedded, friable, and locally quartzitic. The quartz grains range from coarse to fine. A detailed description of the Casper Formation was given by Agatston (1954, p. 536—548). Crossbedding in the Casper Formation is well ex- posed in the valley of Sand Creek, southwest of Laramie, in a series of spectacular monuments and pil- lars. The term “festoon cross-lamination” was intro- duced by Knight (1929, p. 56; 1953) to describe the primary structures at this locality. Unusually good ex- amples of primary crumpling of sedimentary strata caused by subaqueous gliding are also there (Knight, 1929, p. 74-78). In the Casper area a red mudstone, locally containing sand and conglomerate and generally about 10 feet thick, is found in wells and at the surface within the upper part of the Casper Formation. This bed, informally named the “red marker,” is considered the lowermost unit in the Permian System. The following discussion is quoted from Maughan (1967 ,p. 132) : The base of the Permian System in Wyoming is placed at the lower contact of a mudstone known as the red marker, originally recognized in the subsurface of the Lance Creek field in central eastern Wyoming but now known in most of southeastern Wyo- ming and southwestern South Dakota. This very distinctive unit, composed of red mudstone with minor dolomite and sulfates, is recognized in the subsurface over a very large area. It is also exposed in the Hartville Formation in the Hartville uplift (J. W. Strickland, oral commun., 1958) , in the Minnelusa Formation of the southern Black Hills (C. G. Bowles, oral commun., 1959), and somewhat less certainly, in the Casper Formation of the Lar— amie Range. The red marker does not extend westward into central Wyoming beyond Casper nor northward into northeast- ern Wyoming, northwestern South Dakota, or beyond. On the east fork of Wagonhound Creek, Carbon County, Wyo., the red marker forms a prominent bench in the Tensleep Sandstone (fig. 10). There the marker is about 5-20 feet thick and resembles the Fountain G18 FIGURE 10.—'The red marker (bracketed) on east fork of Wagonhound Greek, west of Arlington, '1‘. 19 N., R. 79 W., Carbon County, Wyo. Photograph by E. K. Maughan. Formation. It is a mixture of red siltstone, sandstone, and conglomerate, is crossbedded, contains cut-and-fill bedding suggestive of fluvial origin, and has numerous thin but prominent lenses of quartz and feldspar pebble conglomerate. The upper surface of the sandstone that underlies the red marker at Wagonhound Creek is highly irregular and contains solution pits. That part of the Casper above the red marker is well sorted crossbedded bufl' quartz sandstone. Scattered within this unit are coarse grains of quartz several times larger than those of the matrix. That part of the Casper Formation that lies below the red marker and is of Pennsylvanian age ranges in thick- ness from about 100 to 600 feet (pl. 3(D)). It is the lateral equivalent of the Tensleep Sandstone, the upper two members of the Amsden Formation, and possibly the upper part of the Darwin Member and is Morrow, Atoka, Des Moines, Missouri, and Virgil in age, al- though at any one locality rocks of one or more of these ages may be missing (pls. 2(8, 0) and 3(A, B, 0); Thomas and others, 1953). Those parts of the Casper that are of Morrow age and nearly all the parts of Atoka age occur only in the subsurface (pls. 1 and 2(B, 0)). Surface sections along the flanks of the Laramie Moun— tains and adjacent parts of the Laramie basin are Des Moines and younger as indicated by an abundant F usu- lim-Tritz'cz'tes fauna. Exceptions exist in Deadhead basin and at Farthing on the east flank of the Laramie Mountains where an Atoka fauna has been reported (Pan American Petroleum Co., oral commun., 1960) . A correlation of Casper strata into subsurface areas adj a- cent to the Laramie Mountains was made by comparing subsurface sections in the Laramie basin with surface sections studied by Thomas, Thompson, and Harrison (1953, pl. 9). SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY FOUNTAIN FORMATION The name Fountain Formation was given by Cross (1894) to a “series of red sandstones, grits, and con- glomerates, a part of the so-called ‘Red Beds,’ found in typical development on Fountain Creek below Manitou Springs (Colo.).” In its type locality, Cross described the Fountain as “chiefly coarse-grained, crumbling arkose sandstones, in heavy banks showing cross-bedding. They are locally conglomeratic, mottled with gray and various light shades of red, through irregular distribution of coloring matter. Near the base and at intervals throughout the series are very dark red or purplish layers of arenaceous shale in fine- grained sandstone.” This description of the Fountain Formation serves well for the entire length of its expo- sure along the east flank of the Front Range in Colorado, the Laramie Mountains in Wyoming, and for exposures in the Laramie basin. Locally cobbles and boulders of granitic and metamorphic rock and chunks of older Fountain sandstone material are observed in torrenti- ally crossbedded sequences. In the northern part of the Front Range on both flanks of the Laramie Mountains and in the Laramie basin, Maughan and Wilson (1960, p. 34—35) recognized two lithologically distinct parts of the Fountain Formation (fig. 11) : The upper part of the Fountain differs from the lower part in that it contains many sandstone units with scattered, rounded, coarse grains of quartz and feldspar. Siltstone strata, common in the lower part, are uncommon and only locally present in the upper part. Although the lower part of the Fountain con- tains some strata of fine-grained sandstone similar in appear- ance to those of the upper part, none contain scattered coarse grains. The contact between the two parts of the Fountain is placed at the base of the lowest sandstone containing these scattered, well-rounded coarse grains. This sandstone appears to be persistent * * * and is a good horizon marker. The Fountain Formation is Atoka, Des Moines, Mis- souri, and Virgil in age. Its areal extent in Wyoming is limited to the southern part of the Laramie basin and the flanks of the southern part of the Laramie Moun- tains, and it intertongues with the Pennsylvanian part of the Casper Formation to the east, north, and west (pl. 1). A few fusulinids collected from thin limestone and sandstone beds by Maughan and Wilson (1960) and identified by Henbest assisted the author in zoning and correlating the Fountain with the equivalent Casper Formation (fig. 11). According to Maughan and Wilson (1960, fig. 2), the upper part of the Casper, Wolf- camp in age, lies unconformably on the Fountain For- mation in the Laramie basin (fig. 11) . MINNELUSA AND HARTVILLE FORMATIONS The name Minnelusa was applied by Winchell (1875, p. 38) to a sandstone in the Black Hills that is G19 PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING .fi .3 68: 82E is. gas“: 58.5 630 5:. .9»? 653582 355.3 5:." 525 25.355 we fig F8553 among wax: damn—Em 55 ding—gem we 50358.80 cosfifioawléfi 555»: : L“ a. a. :7) 2 S uoneuuog lenses zusgangnba u mu no J J0 ued 18M01 1ua|engnba _ 0 _ hm m. u— 000 Dow CON 0 m4 _ _ _ m.__>m::on_ din—2m >>mwz quo. cmtnEmuwnn— 6335 EMS 022,200.. C... van—2. 28am on K «N m» «B mm «K K as o— a: awoa \ , , \ \ \ w M \ “we \ w w H \ \ mcozuom 3:.“qu 3 «two 3398 m:o_u «383 635.85...» E; 53:52.3 m_ 3:... 8022:: S 5:39.. .555; 3523.: .o 3:: 2w 3:: .335; >£Ea¢E8¢D {151 253.228 E.» 2225.3 23.3 2182;. 2.235% v.09. 35.2.30 5:259“. 535:: \ \ \ \ \ \ \\ \ \ \ \ A \\ ‘ l A M \ \ ‘ ‘W m‘ \ \ \ \ ‘ ‘m‘ ‘ l y“ o=ofiu=~m 352:5 \ \ \ “\‘M w \‘m \\ W H w ‘l \ ‘ ‘ m.‘ w,‘ ’5va v\\‘ Y‘l‘ ‘l\\ ‘m \‘\‘x mm \y‘\ x x y \ \ H m N” l\\ \W \ \ , \ \ A W“ \ N \ 1 \\x\ \\\ , \ \ \\ \\\ \v \ \ .200 . .3530 55:3 \ \ \‘Vv v \ \ \ M \ \ \\ \ v \ \ \ \ \ \.\ .39“. .320 Him .253 H .5550 :32 , La_a0 .3580 .ofitfi £03. .255 N G20 “nearly white, crystalline, subsaccharoidal and coarsely granular when weathered and hard,” from which he collected specimens of Streptorhynchus, Athym's, and Zaphrentis. In 1901, Darton (p. 510) used the name to include all sandstones and limestones in the Black Hills region lying between the well-defined limits of the Mis— sissippian Pahasapa Limestone below and the deep-red sandstone and shale of the Permian Opeche Formation above (the entire Pennsylvanian System and the lower part of the Permian System). The Minnelusa is present in the subsurface in the Powder River basin in the northeastern part of Wyoming (not included in this study) and in the Hartville uplift of southeastern Wyoming, where its equivalent is better known by the name Hartville Formation. The Minnelusa Formation in the report area is composed of interbedded limestone, sandstone, and shale. These formations are present only at the east margin of the area. The Minnelusa and Hartville Formations contain strata of Morrow, Atoka, Des Moincs, Missouri, Virgil, and Wolfcamp age, and are the lateral equivalents of the Casper Formation. The thickness of these forma- tions averages about 600 feet (pl. 3(D) ). MORGAN FORMATION AND WEBER SANDSTONE In the Rock Springs uplift, central Sweetwater County, Wyo., only one deep well, the Mountain Fuel Co. Union Pacific Railroad 4, sec. 11, T. 19 N., R. 104 W., penetrates the entire Pennsylvanian System. Formations at depth include the Morgan Formation and the Weber Sandstone. The Morgan consists of cherty fossiliferous carbonate strata with intercalated red shale beds in the lower part, and interbedded sandstone, red shale, and carbonate beds in the upper part. The Weber Sandstone is similar to the Tensleep ; it is a massive fine- to medium- grained light-colored crossbedded sandstone with scat- tered carbonate lenses (American Stratigraphic Co., log CW—1402). The Morgan and Weber are present only in the extreme south-central part of Wyoming. The Morgan is 0 to about 550 feet thick (pls. 2(3) and 3(A) ; the Weber is 0 to about 850 feet thick (pl. 3(3, 0) )- The age of the Morgan is Morrow, Atoka, and earliest Des Moines in the Rock Springs Uplift (Verville and Momper, 1960, John Chronic for American Strati- graphic Co., log CW—1402) ; the Weber was considered of Des Moines, Missouri, Virgil, and )Volfcamp ages by Bissell and Childs (1958, pl. 2). ROCKS OF MORROW AGE IN THE MINNELUSA, HART- VILLE, CASPER, AND MORGAN FORMATIONS Rocks of Morrow age (and locally of Chester age) occur in Wyoming as the Horseshoe Shale Member of SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY the Amsden Formation and the lower parts of the Minnelusa, Hartville, Casper, and Morgan Formations (pl. 2(8) ). The sandstone east of the Laramie Mountains is part of division 6 of the Hartville Formation or the Fair- bank Formation of Condra, Reed, and Scherer (1940). The carbonate strata and the sandy and shaly carbonate strata north of Casper are part of the basal units of the Minnelusa Formation. The sandstone, sandy limestone, and sandy shale in Carbon and Sweetwater Counties are the basal strata of the Casper and Morgan Formations. Several inferences may be drawn from plate 2(3). The northwest arcuate trend of isopachs of the Casper and Morgan Formations in Carbon and Sweetwater Counties reflects the northwest plunge of the terminus of the Front Range uplift, which apparently was in existence by early Morrow time. The northeast trend of isopachs in the Laramie Mountains area, however, is transected by the Casper peninsula of the Front Range uplift, a promontory which extended northward to the city of Casper and contained within its limits nearly all the present-day Laramie Mountains. The discordance between the trend of the Casper Peninsula and the northeast isopach trend nearby leads to the conclusion that Morrow sediments were deposited across the area of the Casper peninsula and removed as a result of up- lift in latest Morrow or earliest Atoka time (Mallory, 1963). Abrupt thinning of Morrow strata in the Casper Formation—from 238 feet in the Tidewater Associates Lawn Creek 81—22 well and 301 feet in the J. J. Lynn Burk 1 well (southeastern Natrona and northeastern Carbon Counties) to zero in outcrop sections a few miles away on the west side of the Laramie Mountains—sup- ports this concept. If the Casper peninsula formed be— fore or during Morrow deposition in this area, isopachs should parallel the margins of the peninsula. The greater thickness of Morrow strata in the Laramie basin compared with the thickness on the east side of the Laramie Mountains indicates that greater subsidence occurred in the Laramie basin, or that some upwarping of the east flank of the Laramie Mountains occurred after deposition (perhaps as a corollary of the uplift of the Casper peninsula), or that both subsidence and upwarping occurred. The predominance of sandstone and sandy limestone in the lower parts of the Casper, Morgan, and Hartville Formations cannot readily be explained. The sand may have been derived from older sediments on the rising Ancestral Front Range uplift, but this assumption is questionable because a wide region in south-central Wyoming and north-central Colorado was periodically emergent in pre-Pennsylvanian time and may or may not have borne a sandstone cover in early Morrow time. PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOLIING Noteworthy is the absence of arkosic conglomerate of demonstrable Morrow age in the Laramie basin and vicinity. TENSLEEP SANDSTONE LITHOLOGY AND THICKNESS The Tensleep Sandstone (fig. 12) was named by Darton (1904, p. 397). The type locality is in Tensleep Canyon, 7 miles east of Tensleep, Washakie County. The formation is commonly white or cream sandstone, but is tan or pink in a few places. It is fine to medium grained and crossbedded (fig. 13), usually in thick ledges. (See also pl. 16A (facing p. 34) of US. Geol. Survey Prof. Paper 51.) The quartz is well sorted, subrounded, and generally frosted. Agatston (1954, p. 522) reported that 59 measurements of crossbedding on the east and west flanks of the Bighorn Mountains indicated no preferred direction of dip, although south and west components were particularly common.’ Tan, white, and pink finely crystalline locally dense limestone and dolomite beds are commonly interbedded with the G21 massive sandstone ledges of the Tensleep, but usually these carbonate beds compose less than 20 percent of the formation. The carbonate beds are broadly lenticular and are, therefore, of limited use in correlation; later- ally they pinch out or grade into sandstone. Through— out much of the area the contact of the Tensleep on the Amsden is sharp, and the two formations are conform— able. Locally, where the Ranchester Limestone Mem- ber of the Amsden contains many sandstone beds, the contact is obscure. The Tensleep Sandstone grades west- ward into the Quadrant Quartzite (Scott, 1935, p. 1018). The Tensleep is best exposed in the Wind River and Bighorn Basins, where it is about 50-350 feet thick. AGE Within this broad area abundant evidence indicates that the age of the Tensleep is Des Moines, and locally very latest Atoka (p1. 3(A)). Henbest (1954, p. 52) listed seven collections of fusulinids from Fremont and Big Horn Counties, Wyo., and from Big Horn and FIGURE 12.—Tensleep Sandstone in Sinks Canyon of Middle Popo Agie River near Lander. Photograph by W. R. Keefer (Keefer and Van Lieu, 1966, fig. 19) . G22 FIGURE 13.—Giant crossbedding in the Tensleep Sandstone, Bull Lake Canyon, Wind River Range. Carbon Counties, Mont., that indicate that the Tensleep of this region is the age equivalent of the lower half of the Des Moines Series. Forms represented include: Wedekvmdellz'na euthysepta, W. egocentric-a, Fus'wlz'na distenta, F. girtyi, F. Zeez', F. tregoensis?, F. rocky/mon- tana?, Bradyina sp., Olimowammim sp., Pseudostafl’ella sp., and others. According to Henbest (1954, p. 53), this Fuwlina-WedekiMeflina fauna is one of the most easily recognized and widely distributed assemblages of Pennsylvanian foraminifers. It is typical of the lower half to two-thirds of the Des Moines Series in the midcontinent region and of that age equivalent else- where. In 1956 (p. 59—62) Henbest examined 11 addi- tional collections of Tensleep fusulinids from Fremont, Teton, Sublette, Washakie, Sheridan, and Johnson Counties and from Grand Teton National Park. Forms represented, in addition to those just listed, are: Calci- tomella sp., Endo-thym sp., Tetratamis millsapemis?, Globz'vahjulina sp., Fugulz'nella gephymea?, and Fusulina or Fusulz'nella sp. An age from very latest Atoka through the early half of the Des Moines is indicated. The specimen of Fusulz'nella gephymea?, is from US. Geological Survey loc. f9791 (Henbest, 1956, SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY p. 61), and is the only specimen of Fusulinella, of which Henbest was inclined to attempt specific identification. The f9791 collection was taken from a zone of inter— bedded dolomite, shale, and sandstone at a horizon that is either at or near the contact of the Amsden Formation and the Tensleep Sandstone at Bull Lake Canyon (Murphy and others, 1956). A transitional fauna and lithology with both Atoka and Des Moines aspects at this zone corroborates the ages assigned by Henbest to the upper part of the Amsden (Atoka) and to the Ten- sleep (Des Moines) in this area. In the Jackson Hole area in extreme western Wyo- ming the Tensleep may include beds that are younger than those in the Bighorn Basin. Love (1954, col. 1) indicated that Tritim'tes occurred in some part of the for- mation; Wanless, Belknap, and Foster (1955, p. 35) re- ported the discovery of a specimen of Tm'tz'cz'tes, but they were not certain as to the horizon represented. At the margin of the Cordilleran miogeosyncline the Tensleep Sandstone may, therefore, range as high as Missouri or Virgil in age. Similarly, in a narrow belt trending roughly north— east through the center of the State, the Tensleep Sand- stone contains strata of Missouri, Virgil, and Wolf- camp ages (fig. 14). Along the east margin of this belt, the Tensleep grades by facies change into the upper parts of the Minnelusa, Hartville, Casper, Morgan, and Weber Formations. QUADRANT QUARTZITE LITHOLOGY AN‘D THICKNESS The Quadrant Quartzite is primarily a well—bedded white to pink fine- to medium-grained quartzite. Locally, where the cement is weak, the rock is friable and read- ily weathers to sand. At Quadrant Mountain in Yellow- stone National Park, the upper part of the Quadrant contains thin beds of calcareous limestone (Scott, 1935, p. 1018). The Quadrant Quartzite is the lateral equiva- lent of the Tensleep Sandstone. According to Scott (1935, p. 1013—1014) : “The Quad- rant formation was named by Iddings and Weed (1899) while working out the geology of the Gallatin Range (1883—93) from its exposure on the southeast corner of Quadrant Mountain in * * ”‘ Yellowstone National Par .” At Quadrant Mountain the Quadrant Quartzite is 243 feet thick. AGE At Quadrant Mountain, Thompson and Scott (1941, p. 350—351) reported a W edekindellina emcentrica—W. euthysepta ?—Fusulina sp. fauna. From this evidence they concluded that the age was clearly Des Moines al— PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING though they referred to the formation as “Lower Penn- sylvanian” in the title of their paper. ROCKS 0F DES MOINES, MISSOURI, AND VIRGIL AGE Strata of Des Moines, Missouri and Virgil age also occur in the Minnelusa, Hartville, Casper, Fountain, Morgan, and Weber Formations. Plate 3(A, B, 0) shows the extent, thickness, and lithology of strata in the listed formations that are of Des Moines, Missouri, and Virgil ages, respectively. DES MOINES ROCKS Rocks of Des Moines age are present everywhere in the area except on the Pathfinder and Front Range uplifts (pl. 3(A)). The rocks of Des Moines age are absent at the margins of these uplifts and are at least 350 feet thick in the Wind River Basin and more than 600 feet thick in Sweetwater County. In the southeast quarter of the State the thickness was controlled by the configuration of the adjacent uplifts; the Pathfinder uplift had diminished appreciably in size since Atoka time although it extended a short distance farther north. In Platte, Albany, and Carbon Counties, the rec- tangularity of the isopach pattern in northeast and northwest directions is apparent. The northeastern trends are similar to the Morrow trends in the area; the northwestern ones are probably related to the dominant northwestern alinement of the Front Range uplift. throughout Pennsylvanian time. The resultant rectan- gular pattern suggests that block faulting was a con- trolling tectonic mechanism. In northwestern Wyoming the Des Moines isopachs trend northeast in contrast with the equally marked northwest trend of Atoka isopachs. The Des Moines trends are unusual in that the linear areas of thin strata are wide and that linear “thicks” are exceedingly nar- row. Inasmuch as strata of late Early Permian age in northwestern Wyoming rest on strata of Middle Penn- sylvanian (Des Moines) age, the existence of a hiatus at this horizon suggests that simple valley cutting was responsible for the abnormal isopach pattern. Yet valley cutting cannot be the reason because the long narrow trends which superficially resemble valleys are areas of thick Des Moines strata. Valleys eroded into the Ten- sleep Sandstone would be “thins.” Perhaps broad, gently arched northeast-trending anticlines were active during or immediately after Des Moines deposition, and the resultant gentle emergence caused unconsolidated sand to be washed or blown from the crests of tectonic—topo- graphic high areas into adjacent troughs. Lithologic associations of Des Moines strata are un- complicated. Because the Quadrant Quartzite and the Tensleep Sandstone contain most of the strata of Des G23 Moines age, plate 3 (A) shows mainly sandstone. Nearly all northwestern Wyoming was probably a broad plat- form sloping gently southwest. Its southeast terminus was the Pathfinder uplift which was flanked on the north and east by sandstone that grades into limestone in the Casper and Hartville Formations. A tongue of sandy carbonate rock (in north-central Carbon County) grades southwestward into the lower part of the Morgan Formation, a widespread limestone in the Eagle basin of northwestern Colorado. In southern Albany County calcareous sandstone and arkose make up the Casper and Fountain Formations. A marked decrease in the area of arkose deposition from that in Atoka time implies that the north end of the Front Range uplift was less active in Des Moines time and supplied a smaller quantity of coarse detritus. A small tongue of calcareous and shaly sandstone in the Jackson Hole area in Teton County is an extension of geosynclinal lithologies farther west. On the platform the Des Moines sea may have been so shallow that the depositional interface lay continually in the zone of wave action. Intermittent emergence and submergence of the platform (related to cyclothems in the Mississippi Valley?) would have allowed wind and waves to continually re-sort the sand grains and arrange them into crossbedded dunes and bars. The thin lime- stone beds in the Tensleep, Quadrant, and Weber For- mations may be records of either deeper—than-usual local subsidences for short periods or temporary quiet-water conditions without influxes of sand. MISSOURI ROCKS Strata of Missouri age (pl. 3 (B )) are absent from all but the easternmost part of the Bighorn and Wind River Basins and are probably absent from much of the Green River basin. The cause is probably lack of deposi- tion, but some postdepositional marginal beveling may have occurred. It seems reasonable that the Bighorn— Wind River platform progressed from periods of shal- low submergence alternating with periods of emergence in Des Moines time to complete emergence in later Penn- sylvanian time. The position of the present west limit of Missouri strata, however, is probably the result of ero- sional beveling, the latest episode of which probably occurred in post-Virgil time. This conclusion is sup- ported by the parallelism of the Missouri and Virgil western limits as suggested by the available surface and subsurface control (fig. 14). Critical evidence in locat- ingthe western limits of the Missouri, Virgil, and VVolf— camp rocks in northern Wyoming was provided by Verville (1957), who reported the occurrence of Schwagem'na and T'm'tz'cites of early VVolfcamp age in beds he assigned to the uppermost part of the Tensleep Sandstone at Mayoworth. K I 0 '''' I\ ‘2’ 1% 425' _ I I "o i l . w Y‘,0\M 1 N G f 'a 1.“ l i | | | lei m T4. ‘ ikr Sr 5 J: k \ g m' ' ill-2 “A ’4' la... \E’ ' —- L515 EV’ : // "7°35“. " a; .1 ..... 4 ’ / y‘aé I \ 1 _ \i_ \ \ \ 0 50 100 150 MILES FIGURE 14.—Western limits of Missouri and Virgil rocks in central Wyoming. Data from which these lines [are drawn are sparse but suggest that the parallelism of the two lines results from post-Virgil pre-Wolfcamp beveling. Shaded areas indicate Precambrian outcrops in present major mountain ranges. The Missouri rocks, averaging about 100 feet in thick- ness, are thinner than the Des Moines. They range from 0 to 200 feet in thickness in the Laramie area and reach a maximum of 250 feet in Sweetwater County. At the east margin of the Bighorn—Wind River platform, a broad belt in Johnson, Natrona, and Carbon Counties was tectonically transitional from a platform margin to a marine shelf in eastern Wyoming, where the rate of subsidence moderately increased eastward. The thick- ness of Missouri rocks in this belt ranges from 50 to 150 feet. Although isopachs in the transition belt are gen- erally sinuous, isopachs are amoeboid in the vicinity of the Pathfinder uplift, which was limited to a small area at the north tip of the Laramie Mountains. Perhaps in later Des Moines or early Missouri time the uplift gently subsided. However, the rectangular isopach pattern in Albany County (pl. 3(3)) suggests an alternate hypothesis—that foundering of the Path— finder uplift took place by progressive northward block faulting. In Missouri time a tongue of carbonate strata more than 150 feet thick in southeastern Natrona County occupied the site of the northern part of the Pathfinder uplift of Des Moines time. The limited extent of arkose deposition in Albany County reflects the progressively diminishing uplift of the north end of the Front Range uplift after Atoka time. The lithologic patterns in the Missouri rocks are generally similar to those of the Des Moines because sandstone is present on the Bighorn—Wind River plat- SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY form (but only its extreme east margin) and grades eastward into carbonate rock in a short distance along a north-trending belt. The sinuous pattern of the facies transition zone is most marked in Natrona and Carbon Counties, where the long narrow tongue of shelf carbonate rocks isolates the residual Pathfinder uplift from the Bighorn—Wind River platform sandstone province. After allowance for minor local variations, the coincidence of the narrow quartz-to-carbonate facies transition belt with the tec- tonic transition from the Bighorn—Wind River plat- form to a shallow basin in eastern Wyoming is note- worthy. The coincidence shows that sand, probably derived from the Des Moines Tensleep of the Bighorn— Wind River platform, was being carried eastward into the sea, where it was transported laterally by longshore currents. At the same time, nearly pure carbonate was deposited a short distance offshore in eastern Wyoming. Control on the southwest flank of the Pathfinder up- lift, though meager, suggests the presence of local sand— stone bodies surrounded by calcareous sandstone. Their existence offshore from points of land projecting south- west and south from the Pathfinder uplift in the sus- pected presence of strong longshore currents suggests that these sands are spits similar in magnitude to those at Cape Cod, Mass. A liberal interpretation was used in indicating the spits on plate 3(3). The sand was presumably swept southward around the east and west flanks of the Pathfinder uplift by longshore currents and redeposited as local sand bodies within sandy lime- stone (Casper Formation). The interpretation regard- ing southward movement of currents at that time agrees with the conclusions of Opdyke and Runcorn (1960, fig. 1, p. 961), who studied crossbedding in the Tensleep and Casper Formations. In the southernmost part of the Laramie basin, small volumes of arkose derived from the Front Range uplift were deposited as tongues in the Casper Formation. The carbonate rock in the Campbell—Converse-Platte County area in eastern Wyoming is part of the Minne- lusa and Hartville Formations. The sandstone in Sweet- water County is an extension of the Weber Sandstone of the Uinta Mountains area. VIBGIL ROCKS Conditions in Virgil time were similar to those dur- ing the preceding Missouri. The Bighorn—Wind River platform was emergent; the J ohnson-Natrona-Carbon- Albany County belt contained the facies transition from sandstone to limestone in northeastern lVyoming; the Pathfinder uplift was almost completely submerged; and the trough in Natrona County persisted but was much reduced in size (p1. 3(0)). The west limit of the PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOIVIING Virgil closely parallels the limit of Missouri strata a few miles farther west, partly owing to nondeposition on the platform but probably mainly owing to post— Pennsylvanian beveling. Thickness of the Virgil in the area ranges from 0 to 300 feet near Casper, but averages about 100 feet. The Virgil isopach pattern differs somewhat from that of the Missouri, however, owing to post-Virgil pre- Wolfcamp warping and erosion in the vicinity of the Laramie Mountains. As a result, rocks of Virgil age are absent in parts of northeastern Natrona County, south- eastern Converse County, central western Carbon County, and an irregularly shaped area east of Laramie. Lithologic relations in the Virgil are also a continua- tion of those in the Missouri. Quartz sandstone of the Tensleep on the extreme east margin of the Bighorn— Wind River platform grades eastward and southward into carbonate rock of the Minnelusa, Hartville, and Casper Formations. Longshore currents moving south- ward bifurcated at the remnants of the Pathfinder up- lift and deposited sand along the shore, where it mingled with carbonate mud. Thin lenses of arkose were depos- ited near Red Buttes, south of Laramie. In southeastern Wyoming the red marker was probably deposited dur- ing a brief period of emergence at the close of Pennsyl- vanian time. UPPER BOUNDARY OF THE PENNSYLVANIAN SYSTEM The upper boundary of the Pennsylvanian System is readily identifiable nearly everywhere in western and central Wyoming. The contrast in lithology and color between the white well—sorted sandstone of the Tensleep and the greenish-gray or red siltstone, shale, or car— bonate of the Phosphoria and Park City Formations is distinctive (fig. 15). Although the Phosphoria and Park City appear to lie in simple conformity on the Tensleep Sandstone at many places in the Bighorn and Wind River Basins, a hiatus is present at the contact. Keefer and Van Lieu (1966) stated, however, that a discon- formity of a few inches to several feet is observable in the Wind River Basin. Also, Eliot Blackwelder (unpub. data, 1911) reported a 4-foot~thick basal conglomerate, with an uneven lower surface, at the base of the Phos- phoria Formation in Dinwoody Canyon. The Tensleep in these areas is Des Moines, whereas the Phosphoria and Park City are mostly or entirely Guada- lupe and Leonard. In the northwestern part of the Wind River Basin, however, J. D. Love (oral commun., 1962) collected macrofossils from the Grandeur (lowermost) Member of the Park City Formation that are reported to be of Wolfcamp age. The strata from which they were _ _ _ _..l_ _ _ .1... ...| cw“ if 3 I n“ I ! 1‘ X I . 1.. , ._._.;4._\. ! ! u, .1 \ i ! _ _— _.J '''''' ‘7' PhosphorLa_l"'\'-'------—" {I ~ ’ i if I L. ,_.| L . . . l . . In '4 L ‘ ! _ 5 {J \K'J 1.,_ §\ .‘| )0. i U312; glaurstaof l ! "W‘s ‘ if !_._._._F.erm93i9_n ______ _‘ “i” "FAQ—i " '''''' "I l l i I ‘ WYO‘iMIINGf i ' | I i I and rear/w I , «a ' Ii J i' Upper part ___.L ......... \31 fr I 0 Q9 " Divisi‘on 101 l 2% l ’ . I ‘ ‘5?" ‘ , Hartville i «— ---.-—.—. ._._....._._. _ 'I ' (I n l §§ ! Park 0in 1 - ' Fm? '° I .g I \ ! I 4. . I . . \7 ! | ' I at] ' I ! " l - _______ . §§’M' , I ' I I . § : Formations i i ' _.._ L _.' 0 50 100 150 MILES FIGURE 15.—-Formations which overlie Pennsylvanian strata in Wyoming. Shaded areas are Precambrian outcrops in present major mountain ranges. After Maughan (1967). collected may be an outlier or part of a thin tongue ex- tending eastward from the Cordilleran miogeosyncline. The hiatus, therefore, accounts for Missouri, Virgil, and much or all of Wolfcamp time. In eastern Wyoming, the systemic boundary is with- in the Casper and Minnelusa Formations and is ob- scure. Locally, as in the Hartville area, the red marker is helpful in distinguishing between beds of Pennsylva- nian and Permian age; but where the marker is absent, paleontological information is necessary. Careful study of the sequence of lithologies is useful, locally, in pro- jecting the boundary from sections where it can be iden- tified from fossils into nearby sections where neither fossils nor the red marker bed is present. SOURCE OF SAND The sources of the sand of the Darwin Sandstone Member are unknown. In many parts of the Rocky Mountain and midcontinent regions, the lithologic unit directly overlying rocks that were exposed in Chester and Early Pennsylvanian time is typically red mudstone or detritus of mixed lithology. Commonly, the materials in this unit can be shown to have had local origin. In Chester time, rocks exhibiting a wide variety of lithol- ogies (particularly in southeast Wyoming and C010- rado) were exposed to weathering and erosion in the region adjacent to the Darwin depositional trough. The Darwin, however, is a remarkably homogeneous well sorted crossbedded quartz sandstone. The grains are rounded and the heavy-mineral content is low. The pu- rity of composition of the Darwin is in marked contrast, G26 therefore, to the relative impurity of other rock units in the Western United States which occupy a similar strati— graphic position. This circumstance suggests that the Darwin was derived from a widespread older sandstone. Because a large part of North. America was emergent in Chester time, long river systems similar to those of the present time probably delivered a variety of materials to the marginal epeiric seas. An exotic source for the Darwin, therefore, seems rea- sonable. The isopach map of the Darwin Sandstone Member (pl. 2(A) ) suggests that the headwaters of the Wyoming River in Chester time lay north or northeast of Wyoming, in Montana, North Dakota, or perhaps as far removed as Saskatchewan, Manitoba, Ontario, and the Hudson Bay area. This whole region was cratonic and contained Paleozoic sandstones older than Pennsyl- vanian. If one of these sandstones were more extensive in Pennsylvanian time than it is today and was not covered in Pennsylvanian time by strata of intermediate age, it could have served as the source of the sand in the Darwin. Two such sandstones are the basal sandstone formations of the St. Croixan Series of Late Cambrian age and the sandstone formations equivalent to the Chazy Group of Middle Ordovician age. Cambrian sandstone is present in Saskatchewan and may have once extended farther east (Sloss and others, 1960, p. 1). Ordovician limestone overlaps the Cambrian sandstone in Saskatchewan, however, for at least 250 miles (Macomber, 1960, p. 4); hence, the St. Croixan sandstones could not have been exposed to erosion and redistribution in Chester or in Pennsylvanian time. Sandstone of Chazy age, much of which is also well rounded and sorted, seems to satisfy the necessary condi- tions. Kistler (1960, p. 6) showed that strata of Chazy age in Central and Western United States and south- western Canada occur in widely separated patches: Harding Sandstone in Colorado; Kinnikinic and Swan Peak Quartzites in Idaho; Eureka Quartzite in Utah and Nevada; the Harding equivalent, the Lander Sand- stone Member of the Bighorn Dolomite in northern Wyoming; the basal sandstone unit of the Winnipeg Formation in Manitoba, Saskatchewan, Montana, and North Dakota; the quartzite unit of the Wonah Forma- tion in the southern part of the Alberta-British Colum- bia area; and a sandstone in extreme northeastern Mani— toba on the shore of Hudson Bay (Kistler, 1960, p. 6; Baillie, 1952, table 2). Neither Kistler nor Baillie iden- tified this sandstone by name; Kistler showed its areal distribution, and Baillie tabulated an unnamed sand-i stone in the Hudson Bay area as equivalent to the Win- nipeg Formation. In many of these areas, isopachs abut the limit of the patch, indicating that the present limit is the result of SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY erosional beveling. Possibly, then, most or all of the isolated patches in North America were originally part of a continuous sedimentary unit. Isopach maps of the Silurian, Devonian, and Mississippian Systems in south— ern Canada (Sloss and others, 1960, p. 12, 14, 23) sug— gest that the present limit of these systems is not far, regionally, from the original limit. Since these systems do not now extend as far as Ordovician rocks pre- sumably did originally, a belt of Chazy sandstone some 500 miles wide may have been exposed during Chester time (fig. 16). 1000 MILES FIGURE 16,—Restored regional drainage patterns for part or all of Carboniferous time in northern United States and southern Canada; present and inferred ancient extent of sandstone of Chazy (Ordovician) age. Stippled areas show present extent of sandstone of Chazy age (after Kistler, 1960) ; shaded areas show inferred former extent of this sandstone stratum shortly after deposition. Hachured line is present limit of strata younger than Chazy. The broad belt southwest of Hudson Bay where basement rocks of Precambrian age are exposed is the postulated locus of the headwaters of Wyoming River in Chester time. The Ontario River of Pepper, de Witt, and Demarest (1954) of Early Mississippian time is compared with the Wyoming River of Late Mississippian time. If a gentle doming in the Hudson Bay region occurred in Pennsylvanian time, a major consequent river system with well-developed subsequent tributaries could have acquired sand of Chazy age, carried it uncontaminated across limestone terranes of Silurian, Devonian, and Mississippian age, and deposited it in Wyoming. There marine currents on the Bighorn—Wind River platform transported it seaward. Support for this hypothesis was provided by the work of Pepper, de Witt, and Demarest (1954), who postu- lated the existence in Early Mississippian time of the Ontario River of eastern Canada, which provided the material for the Berea Sandstone and associated shales in Ohio and Michigan. Their plate 12 shows the head- waters of the Ontario River to have been in Labrador, the Maritime Provinces, and the region east of Hudson Bay; the delta was in Ohio. The sediment source was preexisting immature sedimentary and crystalline rocks PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING of the craton and the northern part of the Appalachian geosyncline (Pepper and others, 1954, p. 99, 100‘; Potter and Pryor, 1961, p. 1226). Potter and Pryor (p. 1227) favored a “recycling hypothesis” for many of the sand— stones of the upper Mississippi Valley and cited as ex- amples the Chester sandstones of the Illinois basin (p. 1225), Whose maturity suggested a contribution from lower Paleozoic sandstones. The Ontario River is plotted in figure 16 to provide visual comparison with the Wyoming River. Both rivers apparently were of the same order of magnitude (at least 1,000 miles long), and both flowed generally south- west on opposite sides of the Transcontinental arch, a gently positive tectonic trend of major areal proportions active in pre-Pennsylvanian time (Eardley, 1951, pls. 3—5). Their ages, however, are dissimilar. The Ontario River is Early Mississippian according to Pepper, de Witt, and Demarest (1954, p. 13). The Wyoming River is assumed to have supplied the sand for the Darwin Member in Chester time. Rocks of the Big Snowy Group were being deposited in central Montana in Chester time. Some question exists as to whether the postulated Wyoming River would have been diverted into the Big Snowy marine embayment. It is not known why sandstone deposition in Wyoming in Darwin time was interrupted in later Morrow and Atoka time and then resumed in Tensleep time, unless stream capture or other unrecorded events in the prove- nance area were responsible. Perhaps, temporarily in Horseshoe time, the source of red clastics was a laterite from one of many possible limestone provenance areas or was an iron-rich Precambrian terrane. A tributary in the Lake Superior region, for example, may have been supplying iron-rich fine clastics to the Ontario River system in Early Mississippian time which were depos- ited in Ohio as the red Bedford Shale (Pepper and others, 1954, fig”, 60, p. 100). Capture of the upper reaches of this tributary by the Wyoming River system in Chester time may have diverted the same iron-rich material to Wyoming, Where it was deposited as the Horseshoe Shale Member of the Amsden with its oolitic iron zone. The problems involved in attempting to re- solve the questions which arise in such a postulate are complex and lie largely in the realm of speculation. During Ranchester time little or no material from an exotic source was introduced into the area. G27 TABULATION OF CONTROL DATA SURFACE No. Location 01251 Locality Source of data p . 1—3 T. R. Sec. Albany County 1 12N 75W 23-24 Camel Rock.....- .. .. . Malaggan and Wilson 2 14N 72W 4 Gilmore Canyon. _. .. . Thomas Thompson, and Harrison (1953). 3 14N 73W . .. .. ... . Red Buttes....... .. .. . Maggan and Wilson 4 MN 77W 17 Sheep Mountain. .. .. . Mallory (unpub. data, 1939). 5 15N 72W 14—18 Telephone Canyon... . Maughan and Wilson 6 17N 72W 5 Rogers Canyon ________ Thomas Thompson, and arrison (1953). 7 18N 72W 6 Wall Rock Canyon.._. Do. 8 19N 70W 31 Farthin (Iron Mtn.). Do. 9 20N 7&72W _ ... .. .. . Sybille pring--_- -- .. . Hensley (1956). 10 23N 73W 29 Wheatland Reservoir. Thomas Thompson, and arrison (1953). 11 24N 74W 11 Garrett...--.. .. .. .. .. . Do. Blg Horn County 18 50N 89W 24 Paintrock Creek _______ Love (written commun., 1956). 19 53N 90W 16 Shell Canyon .......... o. 20 56N 93W 6 Cottonwood Creek... _ Agatston (1954). Carbon County 35 MN 87W 20 Big Sandstone Creek. _ Ritzma (1951). 36 17N 82W 10 Saratoga area section. . Mallory and Meaghan (unpub. data, 1960). 37 19N 79W 32 Wagonhound Creek... Bauer (1952). 38 24-25N 82W 2, 35-36 Medicine Bow River. - Wilson (1954). Converse County 61 29N 72W 23 La Bonte Creek.._. -.- Wilson (1954). 62 29N 77W 21 Little Medicine Creek. Thomas, Thompson, and Harrison (1953). 63 32N 72W 33 Bedtlck.. .. .. .. .. .. .. . Love (1954). 64 32N 74W 6 Boxelder Canyon ...... Do. 65 32N 74W 13 East Fork Little Do. Boxelder Canyon. 66 32N 77W 26 Deer Creek Canyon_. . Do. Fremont County 70 1S ' 2W 33 Trout Creek..-. .. _. _. Love (1954). 71 1N 3W 28 Pevah Creek.. ........ Biggs (1951). 72 2N 4W 2 Bull Lake Canyon.. ._ Love (1954). 73 4N 5W 6 Dlnwoody Canyon_._. Do. 74 6N 4W 1 Black Mountain... _. . Love (written com- mun., 1956). 75 31N 99W 19 Cherry Creek._-. .. .. . Shaw and Bell (1955). 76 32N 94W 23 Long Creek Canyon. . Love (1954). 77 32N 100W 18 Middle Fork Popo Do. Agle River. 78 33N 101W 5 North Fork Popc Do. Agle River. 79 41N 107W 15 Liétle 1Warm Spring Do. ree . 80 42N 105W 7 Wiggins Fork Canyon. Love (written com- mun., 1956). Hot Springs County 98 7N 6E 33 Wind River Canyon_. Love (written com- mun., 1956) 99 43N 93W 28 Red Spring..... .. .. .. Do. G28 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY No. Location No. Location on Locality Source of data 017.; Company Well name Source of data pis. D - 1—3 T. R. Sec. 1—3 T. R. Sec. h Co Jo neon unty . Albany County—Continued 110 42N 84W 20 Mgmle Fork Powder Agatston (1954). i.ver ._ 111 fig 333% 35-36 Ease Cések' "'C" “1;: B“ 13 13N 76W 15 C(filiston and Schmidt 1 ..-. A3211}??? CSt’ratl 112 razy oman ree . o. 1 _ 113 49-505: saw _ -_ -- _ North Fork Crazy Wilson (1954). 14 ”N 75W 36 MRsisvgrliyuel. State 1" 0 Woman Creek and 15 17N 76W 18 Call alifornia-- .- . Wilson 8 ________ Woming 6001 > Dry Kelly Creek. 380‘}. (1956) 114 52N 84W 25 Rock Creek" -- —- —- -- - DO- 16 19N 76W 9 U.S. Smelting Unit 19-76 American Strati- and Refining. 1-9. ugh 10 Co. Lincoln County 17 24N 76W 25 Wasatch Oil ....... Swan 1—25... _. . o. 121 38N new 28 Martin Creek. . .. ... . Lgalg‘vvlligtétg? com- Big Horn County 122 38N 118W 22 South Indian Creek.- . Love (1954). 21 49N 89W 19 Geo. Nolan ________ Government 1.. American Strati- Natrona County gra hie Co. 22 49N 91W 2 Shell Oll-..-..-- -. . Government 1-- o. 123 30N 78W .... __ .. . Bates Creek Reser- Thomas Thompson, 23 50N 92W 2 Kerr MCGee- Uni ..... .- ..- Do. N W Shvo .C k and gI-I'arrison (1953). ggggigm 124 30 79 .--.--__- ee ree -.-._-.-__ 125 30N 83W 17 Nort west Alcova ..... Love D(01954). 24 51N 90W 33 avis 011 .......... Government 1.. Do. 126 32N 73w 22 Hat Six Canyon___.. __ 25 51N 93W 24 Stanollnd 011-. ._ Orchard Unit 1. Do. 127 32N 79W 9 Casper Mountain.. __ _ Do 26 52N 92W 10 Texas ______ Linderman 1.. . Do. 128 33N s7—ssw _ __ -- _ Garfield Peak com- Do. 27 52N 93W 20 Osbome-..————. —. - Kruezer 1-.. .. Do. bined with Rattle— 28 MN 95W 11 Mule Creek 011 Umt 2....- __ ._ - Do. snake gtnc‘iifitlantic 129 33N 89W 23 East CanHyon Creek" _ Do. 6 118- ______ __ _ B __ __ __ __ _ A at to 1954 _ 29 55N 92W 15 Continental 011.. . Government 1.. Do. 13° ”N 85W um" creek 3 s n ( ) so 55N 97W 26 Stanolind 011 ______ J. E. Pepper 1.- Do. 31 56N 96W 12 ...... -Community 1. . Do. Plum County 32 56N 97W 28 Ohio 011..-___.- -_ . G. Easton American Strati- Uni . graghic Co. 162 mm 68W 19 Deadhead basin... __ __ Thomas, Thompson 33 57N 97W 7 801110 Petroleum Dorothy Fox 2. o. 63 27N 66W 1 G L k Land Eigiflson (1955' $th Bamsaan 1 0 uernse a e..._- .-- ove . - 164 ng 70w 15 110395113; Creek.- _ DE). ) 34 58N 96W 34 Mohawk Government 1.. Do. 165 29N 67W 28 Sand Canyon... .. .. ._ Do. Petroleum. Sheridan County Carbon County 123 2211 2213 2 21552132211211.1121)? - e ongue ver. . son __ __ _ _____ __ __ _ 1 s 1. 170 57N 87W 33 Ameden Creek. -- _- -- - Love (wriigtBten com- 39 13N 89W 17 carter 0” unit 1 Afiififi‘fi cg“ mun., . _ __ _ ______ 1 171 58N 89W 6 Little Bighorn Riven. Agatston (1954). 40 MN 83W 23 ”pm“ """" Peters” 1 P131395 Pet” em 41 16N 84W 10 Continental 01].. - Unit 1....- __ _. - Ameflfinc Strati- S i no ra o 00. “h e 0mm" 42 16N 84W 24 Continental 011.. - Unit 2...--- -- - ps Petroleum 175 SW 109W 36 Sheep Mountain ------- Love 0954)- 43 17N 85W 31 Aurora-Kingwood. Unit 1..... -- -_ _ Ameritfin strati- , gra c o T to 0., t 44 17N 88W 27 Shell 011 ........... Rawlins l- -- -- - e n u“ y 45 17N 89W 23 Wasatch 011 _______ Giggimt 1ment— Do. 179 38N 115W 3 Hoback Canyon ------- ngflgtggsfi) 4e 19N saw 2 Ohio 011 ___________ Government 4.. Eitzma (1951). . 180 40N 112w 28 Darwin Peak - Love (1954). ' ‘7 ZON 81W 1‘ Tex“ """"""" n '- -- ' 332%,? sf?“ 181 401“ 116W 16 Jackson-n 130- 4s 20N 78W 35 Ohio 011 ........... Harrison Wyoming Geoi. 182 42N 113W 34 Crystal Cr _ D0. Ucooper 15 Assoc. (1956). 183 421" 115W 2 Flat Creek— D01 49 21N 86W 28 Sinclair 011 and n11..... .. _. _ American Strati- Sheep Creek. Re e.flmng gra 1:110 00. 50 22N 78W 26 California- . -. -. .. . Hodges 1-. __ _. - o. Waghgkie county 51 23N 78W 34 MoCullooh Oil. __ - tha‘zemmeilt- Do. acson . 3; fig aw g I’II‘rout towel: ASBS ton (1954). 52 24N 80W 13 Neatéogai’fis'oc. U.P.R.R. 1 ..... Do. amp on an o. 189 “N saw 9 Otter Creek“ __ _ Do. 53 25N 84W 25 Mfilaslsippi River Unit 1.... -- -_ _ Do. 190 47N 88W 1 Tensleep Canyo Love (written 54 25N 86W 34 Atlantic Refining Unit 3...... -- - Do. commun. 1956). and Fremont 55 26N 73w 15 .T gegomum' B k 1 D Y 11 etc N n 1 Par - - ynn...-._ .-. ur --..-.... 0. ° °" “° ' “m k 56 26N sow 17 Amerada Petrol- Sullivan 00. 1.. Do. 197 lat 44°55' N; long now w Quadrant Mountain.. Scott (1935). g‘imgggfiw 57 26N 89W 24 Sun11 Oils and Government- American Strati- 58 27N 86W 4 s 1 eir 011 d Ulltnlml' era hm 00' inc 8 an I1 ...... ._ ... O. SUBSURFACE Gas 59 28N 81W 22 Stanollnd Oil ..... Unit l-A-_-- _- . Do. 60 12N 92W 10 Phillips Petro— Unit 8..... ._ __ _ Do. No. Location leum‘ olns' Company Well name Source of data 11—3 'I‘. R. Sec. Converse County Alb-“Y County 67 31N 69W 20 California" __ __ -- . Nylen- American Strati- 68 33N 69W 13 c t 011 11011156153“ 1' gm me 00' 12 13N 73W 8 Western Oil Klink 1 ......... M h 3r 91“ - ---- -- -— - 039 e manl- 0- Fields. au_ swingggo) 69 33N 76W 9 Continental 011.. . Whitesides 60. . Love (1954). PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING G29 No. Location No. Location on Company Well name Source of data 011; Company Well name Source of data ls. p . 11—3 T. R. Sec. 1—3 T. R. Sec. Fremont County N atrona County—Continued 81 ls 4E 36 Atlantic Refining. Tribal 5--. _- ._ . Agatston (1957). 144 40N 79W 35 Stanolind 0i1--.. . Government 1.. Love (written 82 2S 2E 19 Stanolind Oil ..... Terry 3. - _ .. .. . American Strati- commun, 1960). graphic Co. 145 40N 80W 31 Shell Oil. .- .. .. .. - Unit-Govern- American Strati- 83 7N 1E 30 Carter Oil.... -. .- . Shoshone- Love (written ment Hill 1. graphic Co. Madden 1. commun.,195_6). 84 IN 1W 27 Continental Oil.- - Sage Creek 2. .. American Strati- _ graphic Co. Park County 85 2N 1W 19 Stanolind Oil ..... Tribal 9—A-- -- _ Love (1954). 86 3N 1W 6 British American. Tribal E—6-- -- . Wyoming Geol. Assoc. (1956). 146 47N 100W 28 General Petro- Banner 1.. .- .. - Love (written 87 6N 2W 15 Continental Oil.. . Chatterton 20-_ Love (written eum. . commun..1960). commun. 1956). 147 47N 102W 20 Continental 011.. . Government 1-. American Strati- 88 27N 95W 18 Gulf Oi1.--. _. ._ ._ . Government 1-. Agatston (1957). . gpra hic Co. 89 28N 94W 2 Califomia.. .. .- ._ . Unit-Govern- American Strati- 148 48N wow 34 Stanolind Oil. --- . Umt 2- .-- .. .. - Agatston (1954). ment 3. graphic Co. 149 48N 101W 4 California.. .- -- -- . Rawhide 1. . American Strati- 90 29N 96W 20 Carter 011.... .. .- . Yellowstone Do. graphic Co. Sheep Co. 1. 150 48N 103W 31 Wiishire-Atlantic- State 1.... .. .. - Agatston (1954). 91 30N 95W 14 California.- .- .. -_ - Government 1.. Do. 151 49N 102W 13 Texas.. ._ .. ._ .. .- - Unit 3. --_ .. ._ . American Strati- 92 31N 94W 5 British American. Government 1._ Do. graphic Co. 93 31N 98W 4 W. A. Barber and Government 1.. American Strati— 152 5ON 100W 13 Sohio Petroleum. - Simmons-Gov- Do. others. gra hic Co. ernment 1. 94 82N 95W 14 Sinclair— Unit 9.-_-- _. _. - o. 153 50N 101W 6 Richfield... .- .. .- . Unit 1- - ._ .- .. - American Strati- Wyoming gra hic Co. 95 33N 92W 1 Sinclair Oil and Unit 2—C--_. .. . Love (1954). 154 50N 105W 25 Continental Oil.. - Unit 1- --_ -. .. _ 0. Gas 155 51N 100W 29 Kirk-Pacific Connagham 5-- Do. 96 33N 96W 10 Stanolind Oil ..... Unit 11--___ -_- . Do. Western. 97 42N 107W 6 Califomia-. .- .. .. . Langguth 1- - - - American Strati- 156 52N 101W 31 Su erior Oil.. Do. graphic Co. 157 53N 100W 27 Ca 'tomia... Do. 158 55N 100W 8 Stanolind Oil. Do. 159 57N 98W 8 Continental Oil-. _ Do. Hot Springl County 160 57N 101W 7 0:31:31 Petro— Do. 161 58N 100W 33 Seaboard 011 -- -- - Do. 100 8N 3E 6 Candana Southern. Tribal 1._. .. .- . Amerigan (SJtrati— ra 1c 0. 101 41N 91W 27 Sohio Petroleum. - Picard 1..... ._ . g 0. PM“! Cum” 103 42% 96W 2 ghell Ci 1-..-. -- ._. Unit 1--.. - Do. 0 43 91W 6 am” 5 mm" ‘ Murphy """ ' ngfinggfitei‘m, 166 25N 65W 29 Seaboard 011. _- -- - Wilson 1..-- -- - American Strati- 104 43N 92w 21 Pacific West ...... Kirby Creek 1.. Do. . . graphic.“- 105 44N 95w 23 Continental 011- - Gebo 42—T- - American Strati- 167 30N 57W 2 N “101131 ASSOCI- Government 1-- Love (written graphic Co. ated Petroleum. commun., 1960). 106 44N 97W 12 Husky Oil-._ - Unit5- .- .--- Do. 107 45N 100W 26 Continental 0 Skelton Unit 1- Do. . 108 46N 98W 25 Texas_- .- .- ._ ._ .- _ Govemment- Do. Sheridan County 109 46N 100w 25 0111 01 d snwalton 1' Stcanollizn 011 mp 1"” " ’ D°' 172 55N 85W 22 shell 011-.. -- -- -- _ Demple 1 ...... American Strati— d ' graphic Co 173 57N 78W 11 Shell Oil._. -. .- -. _ Clear Creek 1-_ Do. . Johnson County 174 58N 84W 30 Shell 011. .. _. .. .. - Buszklesevic l - Do. 115 42N 83W 4 Chicago Corpora- Harlan 1-.-. .. - American Strat - Sweetwnter County tion and Re- graphic Co. 5132110 N Mural 176 16N 104W 21 Mountain Fuel- -. Unit 1.-.. .. .. - Veg/Fine 3111:1960) -. . omper . 116 “N 81W 19 S tanolmd 011 ---- G‘figfig‘fut' Do. 177 19N 102W 35 Cities Service.... . Union Pacific" American Strati- 117 45N 83W 24 Stanolind Oil _____ Brock 1._.‘...- - Do. - . gm" ° ° 118 46N 81W 5 Shell Oil..- __ _ _ _ Government 1- Foster (1958). 178 19N 104W 11 Mountain Fuel. -- Union Pacific 4. Anigrigagggrati- 119 48N 76W 20 Pure 011..... .. .. . Unit -- _. - American Strati- g p . graphic Co. 120 48N 82W 17 Carter Oll-_.- -- .- . Rider 3 ......... Do. Teton County Nan-om: County 184 42N 114W 1 Carter 011-... .. .. - Marie Treg- American Strati- 135 44N 111w 25 c 111 ‘ U1??? 1' griamc 00' a ornla-- -- .. -_ - - ___ __ __ _ o. 131 29N 80W 22 Tidewater Asso— Lawn Creek American Strati— n ciates. - . graphic Co. 132 31N 81W 24 Paézlfic Western Oborne 1__ .- .- . Love (1954). Uinta County 133 31N 84W 9 Atlantic Refining. Foster Unit 1._ American Strati- gra hic Co. 186 16N 117W 32 Shell Oil... .. .. .. . LeRo Unit 1._ Verville and 134 32N 81W 16 Kemmerer and State 1._ ._ .. .. . Dpo. y Momper (1960). Kemmerer. 135 33N 80W 20 M. E. Morton-.-. Johnson 1—A__- Do. 136 34N 82W 10 Skelly Oil.. .. ._ .. - Wallway-Gov— Do. Washakie County emment 1. 137 35N 77W 21 Socony—Vacuum- _ G 32-X—21_ Do. 138 36N 81W 4 Amerada Petro- Unit 1- . __ Do. 191 43N 89W 31 Hiawatha-- -- ._ _. . Government American Strati- leum. 1—31. graphic Co. 139 36N 82W 22 Cities Service" Sprecher 1 Do. 192 45N 92W 19 G & G Drilling..- Unit 1. -.- ._ -. . Do. 140 37N 82W 36 Pure Oil. - .. Um . - -. Do. 193 46N 91W 26 General Petro- 43-26-G..- .. -- . Love (written 141 37N 85W 3 Trigood Oil_--. .. . Goovernment Do. leum. communq 1956). 194 46N 96W 27 Philli s Petro- Unit 1. --. -. _. . American Strati- 142 38N 78W 10 US. Navy. .- -_ .. . NPR3 1—G-10.. Do. leurg. graphic Co. 143 39N 83W 18 Sohlo Petroleum. - Government- Do. 195 48N 89W 31 Gulf 0il.... .. _. _. . Mills-Federal 1. D0. Evans 1. 196 48N 92W 18 Pure Oil..... .- ___ Unit _.-. .- ___ Do. 269—091 0—67—3 G30 REFERENCES CITED Agatston, R. S., 1954, Pennsylvanian and Lower Permian of northern and eastern Wyoming: Am. Assoc. Petroleum Geologists Bull., v. 38, no. 4, p. 508—583. 1957, Pennsylvanian of the Wind River Basin, in Wyo- ming Geol. Assoc. Guidebook 12th Ann. Field Conf., 1957: p. 66—69. Baillie, A. D., 1952, Paleozoic stratigraphy of the outcrop area in Manitoba, in North Dakota Geol. Soc. Guidebook 1st Ann. Field Conf., 1952: p. 32—40. Baker, A. A., Huddle, J. W., and Kinney, D. M., 1949, Paleozoic geology of north and west sides of Uinta Basin, Utah: Am. Assoc. Petroleum Geologists Bull., v. 33, no. 7, p. 1161—1197. Bauer, E. J ., 1952, Wyoming geologic map and structure sections of the Wagonhound Creek area, Carbon County, Wyoming: Wyoming Univ. unpub. Master’s thesis. .Biggs, C. A., 1951, Stratigraphy of the Amsden Formation of the Wind River Range and adjacent areas in northwestern Wyoming: Wyoming Univ. unpub. Master’s thesis. Bishop, W. F., 1957, The petrology of the Darwin Sandstone member of the carboniferous Amsden Formation of west central Wyoming: Miami Univ. Master’s thesis. Bissell, H. J., and Childs, 0. E., 1958, The Weber formation of Utah and Colorado, in Symposium 'on Pennsylvanian rocks of Colorado and adjacent areas: Rocky Mtn. Assoc. Geologists, p. 26—30. Blackwelder, Eliot. 1918, New geological formations in western Wyoming: Washington Acad. Sci. Jour., v. 8, no. 13, p. 417—426. Branson, C. C., 1937. Stratigraphy and fauna of the Sacajawea formation, Mississippian of Wyoming: Jour. Paleontology, v. 11, no. 8, p. 650—660. ~-—— 1939, Pennsylvanian formations of central Wyoming: Geol. Soc. America Bull., v. 50, no. 8, p. 1199—1225. Branson, E. B., and Branson, 0.0., 1941, Geology of Wind River Mountains, Wyoming: Am. Assoc. Petroleum Geologists Bull., v. 25, no. 1, p. 120—151. Branson, E. B., and Greger, D. K., 1918, Amsden Formation of the east slope of the Wind River Mountains of Wyoming and its fauna: Geol. Soc. America Bull., v. 29, p. 309—326. Burk, C. A., 1954, Faunas and age of the Amsden formation in Wyoming: J our. Paleontology, v. 28, no. 1, p. 1—16. Chronic, John, and Ferris, C. S., J r., 1961, Early Paleozoic outlier in southeastern Wyoming, in Symposium on lower and mid- dle Paleozoic rocks of Colorado, 12th field conf.: Denver, 0010., Rocky Mtn. Assoc. Geologists, p. 143—146. 1963, Two Early Paleozoic outliers in the southern Laramie Range, Wyoming, in Geology of the northern Denver basin and adjacent uplifts, Rocky Mtn. Assoc. Geologists Guidebook 14th Field Conf., Colorado, Wyoming, Nebraska, and South Dakota 2 p. 23—30. Cohee, G. V., chm., 1961, Tectonic map of the United States, ex- clusive of Hawaii and Alaska: U.S. Geol. Survey and Am. Assoc. Petroleum Geologists. Condra, G. E., Reed, E. C., and Scherer, O. J., 1940, Correlation of the formations of the Laramie Range, Hartville uplift, Black Hills, and western Nebraska: Nebraska Geol. Survey Bull. 13, 52 p. Cross, 0. “7., 1894, Description of the Pikes Peak sheet [Colo- rado] : U.S. Geol. Survey Geol. Atlas, Folio 7. Darton, N. H., 1901, Preliminary description of the geology and water resources of the southern half of the Black Hills and adjoining regions in South Dakota and Wyoming: U.S. Geol. Survey 21st Ann. Rept., pt. 4, p. 489—599. SHORTER CONTRIBUTIONS T0 GENERAL GEOLOGY Danton, N. H., 1904, Comparison of the stratigraphy of the Black Hills, Bighorn Mountains, and Rocky Mountain front range : Geol. Soc. America Bull., v. 15, p. 379—448. 1906, Geology of the Bighorn Mountains: U.S. Geol. Survey Prof. Paper 51, 129 p. 1908, Paleozoic and Mesozoic of central Wyoming: Geol. Soc. America Bull., v. 19, p. 403—470. Dunbar, C. 0., and Rodgers, John, 1957, Principles of strati- graphy: New York, John Wiley & Sons, Inc., 356 p. Eardley, A. J ., 1951, Structural geology of North America: New York, Harper & Bros., 624 p. Foster, D. I., 1958, Summary of the stratigraphy of the Min- nelusa formation, Powder River Basin, Wyoming, in Wyo- ming Geol. Assoc. Guidebook 13th Ann. Field Conf., 1958: p. 39-44. Gorman, D. R., 1962, A study of the Amsden Formation: Illinois Univ. unpub. Ph. D. thesis. —— 1963, Stratigraphic-sedimentational aspects of the Amsden Formation; Bighorn Mountains, Wyoming, in Wyo- ming Geol. Assoc. Guidebook 18th Ann. Field Conf., 1963: p. 67—70. Henbest, L. G., 1954, Pennsylvanian foraminifera in Amsden formation and Tensleep sandstone, Montana and Wyoming, in Richards, P. W., ed., Billings Geol. Soc. Guidebook 5th Ann. Field Conf., 1954: p. 50—53. 1956, Forminifera and correlation of the Tensleep sand- stone of Pennsylvanian age in Wyoming, in Wyoming Geol. Assoc. Guidebook 11th Ann. Field Conf., 1956: p. 58—63. 1958a, Geologic and ecologic significance of the upper Paleozoic foraminifera in the Hartville area, Wyoming, in Wyoming Geol. Assoc. Guidebook 13th Ann. Field Conf., 1958: p. 127-131. 1958b, Significance of karst terrane and residuum in Upper Mississippian and Lower Pennsylvanian rocks, Rocky Mountain region, in Wyoming Geol. Assoc. Guidebook 13th Ann. Field Conf., 1958 : p. 36—38. Hensley, F. S., J r., 1956, Some microfossils of the Pennsylvanian- Permian Casper Formation along the west flank of the Laramie Range, Albany County, Wyoming: Wyoming Univ. unpub. Master’s thesis. Hoare, R. D., and Burgess, J. D., 1960, Fauna from the Tensleep sandstone in Wyoming: Jour. Paleontology, v. 34, no. 4, p. 711—716. Iddings, J. P., and Weed, W. H., 1899, Descriptive geology of the Gallatin Mountains: U.S. Geol. Survey Mon. 32, pt. 2, p. 1—59. Keefer, W. R., and Van Lieu, J. A., 1966, Paleozoic formations in the Wind River Basin, Wyoming: U.S. Geol. Survey Prof. Paper 495—B, 60 p. Kellett, Betty, 1932, Geologic cross section from western Mis- souri to western Kansas, showing detailed correlation of Permian Big Blue series and Pennsylvanian, with Moore, R. C.», and Condra, G. E., Classification of Lower Permian and Pennsylvanian systems of Kansas and Nebraska, 2d ed., prepared for Kansas Geol. Soc. 6th Ann. Field Conf., 1932: Kansas Geol. Soc.; also in Eardley, A. J., 1951. Kistler, J. 0., 1960, St. Peter—Simpson Group studies, United States and Canada, isopach-lithofacies map, Simpson—St. Peter-Plateville interval, in Sloss and others, 1960. Knight, S. H., 1929, The Fountain and the Casper formations of the Laramie Basin; a study on genesis of sediments: W'yo— ming Univ. Pub. Sci. Geology, v. 1, no. 1, p. 1—82. 1953, Cross-lamination and local deformation in the Casper sandstone, in Wyoming Geol. Assoc. Guidebook 8th Ann. Field Conf., 1953: p. 26—28. PENNSYLVANIAN AND ASSOCIATED ROCKS IN WYOMING Love, J. D., 1954, Tentative diagrammatic correlation of Ten- sleep, Amsden, Casper, and Hartville Formations in Wyo- ming, in Wyoming Geol. Assoc. Guidebook 9th Ann. Field Conf., Casper area, Wyoming, 1954 : in pocket. Macomber, R., compiler, 1960, Ordovician System, isopach- lithofacies map, in Sloss and others, 1960. Mallory, W. W., 1960, Outline of Pennsylvanian stratigraphy of Colorado, in Rocky Mtn. Assoc. Geologists, Geol. Soc. Amer- ica, and Colorado Sci. Soc, Guide to the geology of Colorado : Denver, p. 23—33. 1963, Pathfinder uplift of Pennsylvanian age in southern Wyoming, in Short papers in geology, hydrology, and topog- raphy: U.S. Geol. Survey Prof. Paper 450—E, p. E57—E60. Maugham, E. K., 1963, Mississippian rocks in the Laramie Range, Wyoming, and adjacent areas, in Short papers in geology and hydrology: U.S. Geol. Survey Prof. Paper 475—0, p. 023—027. 1967, Eastern Wyoming, eastern Montana, and the Da- kotas, m McKee, E. D., Oriel, S. S., and others, Paleotectonic investigations of the Permian System in the United States: U.S. Geol. Survey Prof. Paper 515. Maughan, E. K., and Wilson, R. F., 1960, Pennsylvanian and Permian strata in southern Wyoming and northern Colorado, in Rocky Mtn. Assoc. Geologists, Guide to the geology of Colorado: Denver, p. 34—42. Mundt, P. A., 1955, A regional study of the Amsden Formation: Stanford Univ. unpub. Ph. D. thesis; also in [abs] Dissert. Abs., 1956, v. 16, no. 1, p. 101—102. Murphy, J. F., Privrasky, N. C., and Moerlein, G. A., 1956, Geol- ogy of the Sheldon-Little Dome area, Fremont County, Wyo- ming: U.S. Geol. Survey Oil and Gas Inv. Map OM—181. Nieschmidt, C. L., 1953, Subsurface stratigraphy of the Heath shale and Amsden formation in central Montana : U.S. Geol. Survey Oil and Gas Inv. Chart 00—50. Opdyke, N. D., and Runcorn, S. K., 1960, Wind' direction in the western United States in the late Paleozoic: Geol. Soc. America Bull., v. 71, no. 7, p. 959—971. Pepper, J. F., de Witt, Wallace, Jr., and Demarest, D. F., 1954, Geology of the Bedford shale and Berea sandstone in the Appalachian basin: U.S. Geol. Survey Prof. Paper 259, 111 p. [1955]. Potter, P. E., and Pryor, W. A., 1961, Dispersal centers of Paleozoic and later cla'stics of the Upper Mississippi Valley and adjacent areas: Geol. Soc. America Bull., v. 72, no. 8, p. 1195—1249. Rascoe, Bailey, J r., 1962, Regional stratigraphic analysis of Pennsylvanian and Permian rocks in western Mid-Conti- nent, Colorado, Kansas, Oklahoma, Texas: Am. Assoc. Petroleum Geologists Bull., v. 46, no. 8, p. 1345—1370. Ritzma, H. R., 1951, Paleozoic stratigraphy, north end and west flank of the Sierra Madre, Wyoming-Colorado, in Wyoming Geol. Assoc. Guidebook 6th Ann. Field Conf., 1951: p. 66—69. Sadlick, Walter, 1955, Carboniferous formations of northeastern Uinta Mountains [Colo-Utah], in Wyoming Geol. Assoc. Guidebook 10th Ann. Field Conf., 1955 : p. 49—59. Sando, W. J., and Dutro, J. T., Jr., 1960, Stratigraphy and coral zonation of the Madison group and Brazer dolomite in northeastern Utah, western Wyoming, and southwestern Montana, in Wyoming Geol. Assoc. Guidebook 15th Ann. Field Conf., 1960: p. 117—126. 0 G31 Scott, H. W., 1935, Some Carboniferous stratigraphy in Mon- tana and northwestern Wyoming: J our. Geology, v. 43, no. 8, pt. 2, p. 1011—1032. Shaw, A. B., and Bell, W. G., 1955, Age of Amsden formation, Cherry Creek, Wind River Mountains, Wyoming: Am. Assoc. Petroleum Geologists Bull., v. 39, no. 3, p. 333—337. Sloss, L. L., 1963, Sequences in the cratonic interior of North America: Geol. Soc. America Bull., v. 74, p. 93—114. Sloss, L. L., Dapples, E. C., and Krumbein, W. C., 1960, Litho- facies maps—An atlas of the United States and southern Canada: New York, John Wiley & Sons, Inc., 108 p. Stille, H. W., 1936, Die Entwicklung des amerikanischen Kordil- lerensystems in Zeit und Raum: Preuss. Akad. Wiss. Phys.- Math. Kl. Sitzungsber., nr. 15, p. 134—155. 1940, Einfiihrung in den Bau Amerikas: Berlin, Gebriider Borntraeger, 717 p. Thomas, H. D., Thompson, M. L., and Harrison, J. W., 1953, Stratigraphy of the Casper formation, pt. 1 of Fusulinids of the Casper formation of Wyoming: Wyoming Geol. Survey Bull. 46, p. 5—14. Thompson, M. L., 1945, Pennsylvanian rocks and fusulinids of east Utah and northwest Colorado correlated with Kansas section : Kansas Geol. Survey Bull. 60, pt. 2, 84 p. Thompson, M. L., and Scott, H. W., 1941, Fusulinids from the type sections of the lower Pennsylvanian Quadrant forma- tion [Wyoming] 2 J our. Paleontology, v. 15, no. 4, p. 349—353. Todd, T. W., 1964, Petrology of Pennsylvanian rocks, Bighorn Basin, Wyoming: Am. Assoc. Petroleum Geologists Bu11.. v. 48, no. 7, p. 1063—1090. Verville, G. J ., 1957, Wolfcampian fusulinids from the Tensleep sandstone in the Big Horn Mountains, Wyoming: Jour. Paleontology, v. 31, no. 2, p. 349—352. Verville, G. J ., and Momper, J. A., 1960, Pennsylvanian fusulinids and preliminary series correlation in southwestern Wyo- ming, in Wyoming Geol. Assoc. Guidebook 15th Ann. Field Conf., 1960: p. 127—128. Walters, R. F., 1946, Buried pre-Cambrian hills in northeastern Barton County, central Kansas: Am. Assoc. Petroleum Geologists Bull., v. 30, no. 5, p. 660—710. 1958, Difierential entrapment of oil and gas in Arbuckle dolomite of central Kansas : Am. Assoc. Petroleum Geologists Bull., v. 42, no. 9, p. 2133—2173. Wanless, H. R., Belknap, R. L., and Foster, H. L., 1955, Paleozoic and Mesozoic rocks of Gros Ventre, Teton, Hoback, and Snake River ranges, Wyoming: Geol. Soc. America Mem. 63, 90 p. Wilson, P. C., 1954, Pennsylvanian stratigraphy of the Powder River Basin, Wyoming and adjoining areas: Washington State Coll. unpub. Ph. D. thesis. 1962, Pennsylvanian stratigraphy of Powder River basin and adjoining areas, in Branson, C. 0., ed., Pennsylvanian System in the United States—a symposium: Am. Assoc. Petroleum Geologists, p. 117—158. Winchell, N. H., 1875, Geological report, in Ludlow, William, Report of a reconnaissance of the Black Hills of Dakota made in the summer of 1874 : Washington, p. 21—66. Wyoming Geological Association, 1956, Subsurface stratigraphy of the pre-Niobrara formations in Wyoming, pt. 1 of \Vyo- ining stratigraphy: Casper, 97 p., and columnar sections. UNITED STATES DEPARTMENT OF THE INTERIOR GEOLOGICAL SURVEY PROFESSIONAL PAPER 554~G PLATE I NORTHWEST EXPLANATION 78 SOUTHEAST A North Fork 90 A’ ' WT, V1 W 185 97 79 73 72 P090 Ag'e R'Ve' 77 Caner 89 ’ if1 Unconformlty California California Little Warm Dinwoody Bull Lake 71 70 Middle Fork 75 Y ' ‘ V l ‘ . . ellowstone» California -—~___.o F - . . . . ormatlon or mem er Unit 1 Langguth 1 39””E Creek Canyon Canyon Pevah Creek Trout Creek Popo Agie River Cherry Creek Sheep Co. 1 Unit-Govt.3 Arkos1c conglomerate Sandy shale b boundary Phos hor'a ' ' *Tf——_T_—T— p 1 Formation (Permian) M Tlme boundary \ . . E 3 ac) \\ ISSOUI'I E 1 O E Mtgenses % g 2 Key to geographic and stratigraphic § 5 g Tensleep Sandstone 1 Des <>t Small conglomerate pebbles Carbonate rock occurrence of significant fossils 3 i l- m Moines ; (I: one we” only 0 10 2o 30 4o 0% 1 Q8 I78I .. i_i_i_i_Liiiiii i i\l_%_5IOM'LES E z // z 1: Control pomts 1dent1f1ed E at Atoka E _. ., ' :3”! Atoka 'E E in text table 0 10 20 3O 4O 50 K'LOMETERS z _ . //_ M Z i—UJ_I_LLIJ_LIMi“i\g4 fl 7.: / / 0"r0W __ a Shaly carbonate rock ggMormw E /:// z 8 Sandstone VERTICAL EXAGGERATiON APPROXlMATELY x140 E E if M _ E SE NOTE: On this diagram thickness and O E IIEI Limesmr‘e (MISSisSippl-an) % 0 lithology are true only at control points shown FEET (0 Chester U . 0n the diagram. Between control oints a) m 1, 2. Macrofauna (Hoare and _ . 1_ Wedekmdellt' a s . . . i . Z) . , p O a E Burgess, 1960) (Des Moines) 1' iirfflaafi::?::;glrty 1. USGS loc. f9791 microfauna and Fusulz'na sp n p 1' x;::::figig::wd Effignssfillng'sa; 8 C 1 i d Gypsum and anhydrite correlatlons are shown dlagrammatically by < e I _ . . g 1, 2. Blackwelder-Girty macrofauna (Morrow) gacfifeuna (late Atoka to early (Love, 1954) 2_ Chester macrofauna (Shaw and EJ a careous san Stone stralght lines 10 (Projected 32 miles from Darwin Peak es Oines) Bell, 1955) - 20 section) 1. Dictyoclostus hermostmus Linoproductus prattem'anus O Squamulam'a perplexa V 3 Composita subtilata Red Green 40 (Love, 1954) shale shale , ~ . SOUTHEAST NORTHWEST B! B . .' 42 :~ I Continental Oil : r x ' Unit 2 J 94 92 54 49 160 .4 t 109 83 . - -. 57 Atl t' R f' ' d S‘ lair Oil and 41 . » . - - 'l 86 85 84 82 96 Sinclair- British- 89 an '9 e '"'"g an '"c . . . . Gen. Petroleum 158 -" 157 15.5 149 148 0h'0 9" and _107 . Carter 0' . . . . . . . . . . - - - - ' Continental Oil Badu-ra 2 Stanolind ' California _ Kirk-Pacific Western California Stanolind Oil Stanolind Oil ContinentalIOil Shoshone-Madden 1 British-American Stanolind Oil Continental Oil Stanolind Oil Stanolind Oil WIlIiorniéig Aenericzlin CaliEorIiIitia3 sugoeil-filiritztethlers FremorhtnlI’teIEroleum Ramiilig Unit 1 ‘ Unit 1 _. Unit 1 Connagham 5 Rawhide 1 Unit 2 Sheep 1 Skelton Umtl Strata restored Tribal E-6 Tribal 9-A Sage Creek 2 Terry 3 unit 11 “' OVt- UN 0 - - I Gas er Formafion . Red marker bed (base of Permian System) Tertiary z . . . . Upper part 0 D (Permian) < Q ac: . ' A Phosphoria Formation (Permian) §\ 6 i w __ . \ \x - . z -. .. a o I , 4 a .. . \\\ ‘~\~~\_ Q. E Virgil D - 2 .p \\ ~~_ G) O ‘————_‘_——_—_‘__‘_‘——,—%‘ “\flfl- .- ‘>‘,M°§,:$. ea ‘\*~m\—~~\~LLE % ‘5 ESEeLL Formatlon MissO ,l ~93” (0 $1” ‘ ~ , fl g Tensleep Sandstone 5 ‘g—————-—-——" \~—‘~§ \\U . D , «I ”—‘h_ 0 2 co __ Des Moines ’— 0) Des Moines E, z ’Atoka E —\L 1.1% c .9 \\—\\ E o. Morrow ,2 .5 \\\\_\ E Z _ E" E Casper \F\‘\ Atoka m S N AtOKa c g T,” L—W Orr" \\ 9:: COL- E \_ \\ __I,___LL\\ Ranchester Limestone Member % g ___e———\\\ W I? anon \ \\\ _ \ \———‘________________a————JT—T ———— ‘\__!\\\¥_‘r—___-~___—____.— ll) «\ \\\~———J*’T'TT ,,_,—»/-’ . . , \ ° 3 Chester 5" “We ;\\\L‘\ MEI’LSL-L\\__\‘ ___ hoe Sham Member E g f~~m \\\\\I:‘:,L//»/ LImestone (missISSlpplan) Morrow \‘ — , '0 erN#—~”fiflfl Chester_‘%r*\\\:_fi’ Horses er L N _:W 3 E , Madison L' _ , .7 . ‘ \\_-—- in sandstone Memb M E < Imestone (MlSSISSIlean) Darw N NORTHWEST . ‘ 1» SOUTH EAST C r ‘ 4 C’ 34 ' l _ Mohawk ', ‘ ~II ’ 23 Gm,“ 31 28 V r 27 j; ' v 25 Kerr McGee- 22 195 115 143 __ 139 . 135 127 123 . 62 . _ 10 7 6 5 ,2 Texas Mule Creek-Atlantic Osborne 7 Stanolind Phillips Shell Gulf 190 189 Strata Chicago, Republic Nat. Sohio Cities SerVIce Morton CaSper 124 Bates Creek Little Medicme 11 Wheatland _ 9 . Wall Rock Rogers Telephone Gilmore w 3 “5’ Community 1 Unit 2 Krueger 1 Orchard UnitZ Unit 1 Govt. 1 Mills-Federal 1 Tensleep Canyon Otter Creek restored Harlan 1 Govt-Evans] f Iiprecher 1 Johnson 1-A Mountain Sheep Creek Reservmr Creek Garrett Reserv0ir Sybille Spring Canyon Canyon Canyon Canyon 8 <2»: Des % ‘3 ‘ ' > Phosphoria Formation (Permian) /—’ fl 1 -— Red marker bed (base 0 erm'an syStem) —— Upper part of Tensleep and Casper Formations (Permian) Recent //er/o-si—Oh\ Surface Upper part of Casper Formation 1 1 1 2 Moines : U 1 . . I . 1 12 ’ifirL—II 1 \ . . ~‘_ . _ g :2 +63% * T* | S d t \\\\ \ Vll'gll _ __'______e——————“ \\\\\\ Virgll _”___1_2_,.’—’——"’_’ Casper Formation 2 \\\ Virgil // \ mg” 3 g _ _l ens eep an 5 one __,__— i—a—~———T‘T _‘—T——\~~‘— ———J—T’— ‘\~~ ______ ’J—mT—T— _\\ / / ‘—~‘~~‘LE — o . S 5 Atoka : E A. \ \ _—_—__ Missouri , -Missouri 1 T .t. w 1 3 i M‘isgui\\*T / // \ E E . .9 . t . v , ~ - ‘\ \ _____’_—_———“—___—~§___* ________ , -- _ m wt 8 \\\\\ . i ' ' g E g ‘5 ' . -. \\RanChesterL_ Limestone "flfiflfl' kellyemis 3 \\\\ /// g Fountain FairrE-tQD/I \ Missouri 3 E < m'» "’ E , -- ' 1‘” " 7““: ‘ , \ Membe 1» Triticiws “MEWS 1,2,3. Tritict'tes sp Des MOine$\\\// W— \ 5) Z 0 m. Morrow E L . a. Horseshoe Shale r c 2 Tnmcttes sp ' ' ' ' \ 2 Z O < 0 /Darwin Sandstone Member (Chester) , ~ ~ I . M S ,9\\ ' . I I . . . Precambrian basement 1- Tmm'ftes “flaws“ \ Des Z to l . 1. (Prolected 45 miles from Dal-Win Sand ember '0 +5 —~———— 1. Tmttcttes SD 1' Trztwttes Sp 2' Fusultna presttna \ MOIHES “J E " '- Amsden Creek) - Ston 0’ _ I 2. Triticites 1. T 't’ ‘te ‘ ' 3. Fusulimz sp 0. . FOSSIL—DATA ACKNOWLEDGMENT Millerella mflecta *Afu“5"u‘f;j;fes M°'"es 6 Member (Chem) g g Pathfinder uplift of Pennsylvanian tlme pic/nus 2. thifiiwjffbi’iiiizg‘ff 1 T .t. tea I . \/ Atoka o Tensleep Canyon: Henbest (1956) v Millerella sp r L Morrow 3. Fusult’na rem“ . react s om mcosus . *f: Abundant fusulinid evidence of Paramillerella pingUiS gflitcétes ”at“: (I) 11 Fusulma sp Des Moines age‘ Henbest (1954 1956) ' Paramillerella circuli 1. Des Moines Mad’ m wttes ones us . 1 Schwagem'na SD A d C k_ G ' 1962 1963' Paramillerella ampla(?) uses f5423 'Son le , 2- FIISWW sp/ Ims en ree ' orman( ’. ) Paramt'llerella adrena estone ”’MTg'sTs-sipp‘an) . Location near Chicago, Republic Nat, Paramt'llerella Sp ( Harlan 1: VerVIlle (1957) n 'nella s 1. (Projected 8 miles from Mayoworth) Na In D . All others: Thomas, Thompson, and Schwagemna sp . Harrison (1953) Tritict'tes sp INTERIOR—GEOLOGICAL SURVEY, WASHINGTON, D.ci—1967—t367042 CORRELATION DIAGRAMS (STRATIGRAPHIC SECTIONS) OF PENNSYLVANIAN AND UPPER MISSISSIPPIAN ROCKS IN THE MIDDLE ROCKY MOUNTAIN REGION, WYOMING SHOWING RELATION OF LITHOLOGIC UNITS AND APPROXIMATE TIME BOUNDARIES UNITED STATES DEPARTMENT OF THE INTERIOR GEOLOGICAL SURVEY —§____._ PROFESSIONAL PAPER 554-G PLATE 2 ELLOWSToNE —"i’—". T‘— a DUBOIS Shale Member RIVERTON 81 ' I Formation I ROCK EN RIVER ROCK N RIVER %VANSTON ’ gl/ANSTON 41° 111° 1 ‘ ‘ __'._I l 41 ° Base by U.S. Geological Survey, 1949 109 108° 104° 0 107° 106° 1050 111 109° Base by U.S. Geological Survey, 1949 A. DARWIN SANDSTONE MEMBER 105° 108° 107° 1060 B. HORSESHOE SHALE MEMBER EXPLANATION LITHOLOGY Sandstone Arkosic sandstone Arkosic shaly sandstone ‘ 1 ‘ j i I Shaly sandstone Minnelusa Formation , l ,‘ Th F ‘ ' 7 i i ‘ we Wm” 9* Calcareous sandstone Calcareous shaly sandstone Shale Sany hal Calcareous san y shale Calcareous shale Limestone M MAPS SHOWING EXTENT, THICKNESS, AND LITHOLOGY OF MEMBERS OF THE AMSDEN FORMATION Carb°nate 1’0““ OF MISSISSIPPIAN AND PENN SYLVANIAN AGE, AND CORRELATIVE FORMATIONS IN THE MIDDLE ROCKY MOUNTAIN REGION Sandy carbonate rock SCALE 1:1 000 000 25 O 25 50 75 I——I I—I I—I 100 MILES 25 o 25 50 75 100 KILOMETERS Shaly carbonate rock +——-——-—d I l‘ l —-—200———— . 1 fl - Isopach ,,,,, ‘ i I Dashed where control is distant; short dashed where thickness is restored across present-day uplifts and mountains. Interval 25 or 50feet I . 88 ‘75 . 183 25 T" 97 Subsurface Surface Composite Sections Roman number (1) designates well or surface section listed in text table. Italic number (250) is thickness, in feet, of mapped unit Dash (—) indicates infor- , 7,, ’ ' '7 mation not available. Composite sections were derived from two or more adjacent wells or surface sections Outcrop of pre-Pennsylvanian rocks Probable limit of Pennsylvanian strata beneath Tertiary cover Limit of uplift in existence at time of deposition of mapped unit and restored interpretation of high- land areas Postulated river courses developed on Madison Lime- stone prior to deposition of Darwin Sandstone Member (Map A) A ———A ’ Line of geologic section Shown on plate 1 .39 41° I “ > - - RO T RANG 41° . 111° « h p “. ’ - ‘7 - I 109° 108a Base by U.S. Geological Survey, 1949 o 1050 INTERIOR—GEOLOGICAL SURVEY, WASHINGTON, 0.0. 104 107C 106° 1967—667042 A C. RANCHESTER LIMESTONE MEMBER PROFESSIONAL PAPER 554—G PLATE 5 INTERIOR'GEOLOGICAL SURVEYI WASHINGTON, D.C. 1967 (6704; I "3%“ 1 1 1 1 I More than 1000 Less than 500 EXPLANATION Thickness of strata, in feet More than 500 but less than 1000 ajg'FormatiQnI -‘ ’ 2 111L111 7‘ I ‘0’ IF" [i0 1 _.s.~_a=a=~ [On 51; I" .1111: I I I I I I I I g 7 77 I I I I $1.. ' WC: ——.,I1—-1-— 1.1-,— .9111 1 ‘ ' ‘ 2‘ 1141—1—19. SHERIDAN 1 1 1111 157° 107 I1‘IIILIIA I I I I I- .39 ' \I e \\ _\ 3m B. MISSOURI ROCKS D. COMBINED THICKNESS OF THE AMSDENITENSLEEP,ANDCORRELATIVE FORMATIONS UNITED STATES DEPARTMENT OF THE INTERIOR 0 6\.\_ 1 1 \ NE. 1 IIIIII 1 u. m 1 LII anII 1 \\._1 NE WM 1\ 1 T 0 1 1 “MW; IIIIIII 3 10 IIIII IIIWITII _ 1:1 I I 11. III. 1 1 \J. EI 100 E 1 \1\ 8 1 T 1 1O W IITII\\II1III 1 ,I ...HI11ANH n “I IIIII 11 a \\ I11. I I\I"1m S \II1W I 1 \1\ \— 1 m \ 1 \ I I \I1\ I 1 \ \\1 1 1 ‘ 1‘ 11 S w 1, s _1 E _ Y I 11 h 1 _1 K 1 C 1 I ... _ T I I. 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