Constraints on the Latest Movements On the Melones Fault Zone, Sierra Nevada Foothills, California By]. ALAN BARTOW SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—J Potassium-argon dates for late Tertiary intrusions provide age control for the latest movements on the Melones fault zone in the Sierra Nevada foothills FIGURE 1. 2. TABLE 1. CONTENTS Page Abstract _________________________________________________________________ J1 Introduction ______________________________________________________________ 1 Acknowledgments _____________________________________________________ 1 Age of the Mehrten Formation _____________________________________________ 1 Age of dacite intrusions ___________________________________________________ 3 Discussion _______________________________________________________________ 4 References cited __________________________________________________________ 4 ILLUSTRATIONS Page Geologic map of the area southwest of Mokelumne Hill, Calaveras County, Calif. _______________________ J2 Diagram showing the correlation of some Tertiary stratigraphic units in the western Sierra Nevada, show- ing age control from K-Ar dating of volcanic rocks ______________________________________________ 3 TABLES Page Chemical data on late Tertiary intrusions ___________________________________________________________ J3 Analytical data for K-Ar age determinations ________________________________________________________ 4 III “BEAR? , ml 3;: rtmm Shorter Contributions to Stratigraphy and Structural Geology, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—A-J UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1980 UNITED STATES DEPARTMENT OF THE INTERIOR CECIL D. ANDRUS, Secretary GEOLOGICAL SURVEY H. William Menard, Director Library of Congress Catalog—card No. 80—600016 For sale by the Superintendent of Documents, US Government Printing Oflice Washington, DC. 20402 Stock Number 024—001—03285—6 (A) (B) (C) (D) (E) (F) (G) (H) (I) (J) CONTENTS [Letters designate the chapters] Origin of the rive-r anticlines, central Grand Canyon, Arizona, by Peter W. Huntoon and Donald P. Elston Rock control and tectonism—their importance in shaping the Appalachian High- lands, by John T. Hack The Upper Ordovician and Silurian Hanson Creek Formation of central Nevada, by Rueben J. Ross, Jr., Thomas B. Nolan, and Anita G. Harris The Marble Hill Bed: an offshore bar-tidal channel complex in the Upper Ordovi- cian Drakes Formation of Kentucky, by W C Swadley Paleogene sedimentary and volcanogenic rocks from Adak Island, central Aleutian Islands, Alaska, by James R. Hein and Hugh McLean The Livengood Dome chert, a new Ordovician formation in central Alaska, and its relevance to displacement on the Tintina fault, by Robert M. Chapman, Flor- ence R. Weber, Michael Churkin, Jr., and Claire Carter Intertonguing between the Star Point Sandstone and the coal-bearing Blackhawk Formation requires revision of some coal-bed correlations in the southern Wasatch Plateau, Utah, by Romeo M. Flores, Philip T. Hayes, Walter E. Marley III, and Joseph D. Sanchez New evidence supporting Nebraskan age for origin of Ohio River in north-central Kentucky, by W C Swadley Reconnaissance geologic study of the Vazante zinc district, Minas Gerais, Brazil, by Charles H. Thorman and Samir Nahass Constraints on the latest movements on the Melone-s fault zone, Sierra Nevada foothills, California, by J. Alan Bartow CONVERSION FACTORS Metric unit Inch~Pound equivalent Metric unit Inch-Pound equivalent Specific combinations—Continued Length millimeter (mm) : 0.03.937 inch (in) meter (m) : 3.28 feet (ft) kilometer (km) : .62 mile (mi) Area square meter (m2) : 1076 square feet (ft?) square kilometer (km?) : .386 square mile (mi?) hectare (ha) : 2.47 acres Volume cubic centimeter (cm3) : 0.061 cubic inch (in-'3) liter (L) : 631.03 cubic inches cubic meter (m3) = 35.31 cubic feet (ftfl) cubic meter : .00081 acre-foot (acre-ft) cubic hectometer (hms) : 10.7 acre-feet liter : 2.113 pints (pt) liter : 1.06 quarts (qt) liter : .26 gallon (gal) cubic meter : .00026 million gallons (Mgal or 106 gal) cubic meter : 6.290 barrels (bbl) (1 bbl:42 gal) Weight gram (g) : 0.035 ounce, avoirdupois (oz avdp) gram : .0022 pound, avoirdupois (lb avdp) metric tons (t) : 1.102 tons, short (2,000 lb metric tons 2 0.9842 ton, long (2,240 lb) Specific combinations kilogram per square centimeter (kg/cm?) kilogram per square centimeter cubic meter per second (m‘l/S) l l 0.96 .98 atmosphere (atm) bar (0.9869 atm) cubic feet per second (ft3/s) liter per second (L/s) : .0353 cubic foot per second cubic meter per second : 91.47 cubic feet per second er per square kilometer square mile [(ft“/s)/mi2] [(ma/SVkmz] meter per day (m/d) : 3.28 feet per day (hydraulic conductivity) (ft/d) meter per kilometer 5.28 feet per mile (ft/mi) (m/km) kilometer per hour : 9113 foot per second (ft/s) (km/h) meter per second (m/s) : 3.28 feet per second meter squared per day = 10.764 feet squared per day (ftZ/d) (m2/d) (transmissivity) cubic meter per second 2 22.826 million gallons per day (ms/S) (Mgal/d) cubic meter per minute —264.2 gallons per minute (gal/min) (mfi/min) liter per second (L/s) : 13.85 gallons per minute liter per second per : 4.83 gallons per minute per foot meter [(L/s)/m] [(gal/min)/ft] kilometer per hour : .62 mile per hour (mi/h) (km/h) meter per second (m/s) : 2.237 miles per hour gram per cubic : 62.43 pounds per cubic foot (lb/ft3) centimeter (g/cm3) gram per square : 2.048 pounds per square foot (lb/ftfi) centimeter (g/cmz) gram per square 2 0142 pound per square inch (lb/in?) centimeter Temperature degree Celsius (“0) 1.8 degrees Fahrenheit (°F) degrees Celsius (temperature) If H [(1.8X°C)+32] degrees Fahrenheit Origin of the River Anticlines, Central Grand Canyon, Arizona ByPETER W. HUNTOON and DONALD P. ELSTON SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—A PLATE 1. FIGURE 1. @P‘PWN CONTENTS Page Abstract __________________________________________________________________ A1 Introduction ______________________________________________________________ Acknowledgments ____________________________________________________ Geographic and stratigraphic location _____________________________________ Structure ________________________________________________________________ Low-angle thrust faults __________________________________________________ High-angle kink bands ____________________________________________________ Related anticlines ________________________________________________________ Ancient river anticlines ___________________________________________________ Regional distribution _____________________________________________________ Origin ___________________________________________________________________ References cited __________________________________________________________ H wammmmfimmr—I ILLUSTRATIONS Page River anticlines, central Grand Canyon, Arizona _______________________________________________ In pocket Map showing the location of the Colorado River and its principal tributary canyons in the central Grand Canyon, Ariz _________________________________________________________________________________ A2 Stratigraphic columns of Cambrian rocks in the central and eastern Grand Canyon __________________ 3 Photograph of well-developed river anticlines upstream from 010 Canyon _____________________________ 4 Photograph of a pair of low-angle echelon thrust faults _______________________________________________ 5 Photograph and diagram showing kink bands in rocks along the river ________________________________ 6 Diagram illustrating selected thories used to explain the origin of river anticlines in the central Grand Canyon ______________________________________________________________________________________ 7 III SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 19/9 ,— ORIGIN OF THE RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA By PETER W. HUNTOON1 and DONALD P. ELSTON ABSTRACT A system of anticlines lies along the trend of the sinuous course of the Colorado River for a distance of 97 km in the central Grand Canyon. Similar anticlines occur in some perennially wet side canyons. The anticlines are most abun— dant and well developed along northeast-trending reaches of the main canyon where it is floored by the Cambrian Muav Limestone. Dips of the folded strata are as great as 60°, and the folding locally extends more than 250 m from the river. Low-angle thrust faults in the limbs of the anticlines parallel the river and have formed in response to folding of the com- paratively brittle carbonate strata. High-angle reverse kink bands, along which rocks are displaced up toward the river, also parallel the anticlines and have developed in response to the upward bulging of the canyon floor. The river anticlines are an unloading phenomenon. They result from lateral squeezing toward the river of saturated shaly parts of the Muav Limestone and underlying Bright Angel Shale. The driving mechanism for the deformation is a stress gradient that results from a difference in lithostatic load between the heavily loaded rocks under the 650-m-high canyon walls and the unloaded canyon floor. Saturation ap- pears to weaken the shaly rocks sufficiently to allow defor- mation to take place. River anticlines are not present in the eastern Grand Canyon, where the Cambrian rocks also occur at river level. Their absence is explained by a lack of shaly rocks that could flow when saturated. INTRODUCTION A series of anticlines coincides with the trend of the Colorado River in the central Grand Canyon (fig. 1). Because the anticlines follow the river around bends, they appear to have a genetic rela- tion to the river and its canyon. The generalized axial trace of two of the folds, at and below the mouth of Matkatamiba Canyon, is shown on the geologic map of Maxson (1969). These folds have been called “river anticlines” (Ford and others, 1974). This paper describes the river anticlines and associated faults in the Grand Canyon and relates 1Department of Geology and Wyoming Water Resources Research In- stitute, University of Wyoming. the origin of the anticlines to mass flowage of satu- rated shaly Cambrian rocks in response to hori- zontal stresses related to lithostatic load differences between the rocks under the high canyon walls and the unloaded canyon floor. Fieldwork and mapping leading to this report was carried out during parts of two river trips in 1976. The existence and general characteristics of the river anticlines had been noted during seven previous research trips of Elston, the first of which was in October 1971. During a trip in May 1973, E. M. Shoemaker recognized the preferential devel- opment of river anticlines along northeast-trending stretches of the river, which led him to propose a structural origin for the river anticlines (an origin also proposed for the Meander anticline on the C010- rado River, below its confluence with the Green River in Utah). ACKNOWLEDGMENTS We thank the US. National Park Service for al- lowing us to work and sample in Grand Canyon Na- tional Park and gratefully acknowledge the kind cooperation of David C. Ochsner and Robert Yearout. Funds supporting work by Huntoon were pro- vided by the Office of Water Research and Tech- nology through the Wyoming Water Resources Re- search Institute under matching grant 14—34-0001— 6134. The US. Geological Survey work of Elston was carried out under magnetostratigraphic investi- gations of the Cambrian strata of northern Ari- zona. James I. Drever, Department of Geology, Uni- versity of Wyoming, conducted the X-ray analyses of clay samples collected during the course of this study. Eugene M. Shoemaker and Irving J. Witkind of the US. Geological Survey reviewed the manu- script and made many helpful suggestions. Shoe- maker provided a sketch that led to figure 6. A1 A2 GEOGRAPHIC ANI) STRATIGRAPHIC LOCATION We have observed river anticlines along the 97- km reach of the Colorado River between Fishtail and Parashant Canyons (fig. 1). The anticlines characteristically occur in the lower half of the Cambrian Mauv Limestone and in parts of the un- derlying Cambrian Bright Angel Shale where these rocks crop out at and near river level. The combined thickness of'the Bright Angel Shale and Muav Limestone in the area of the river anti- clines in central Grand Canyon is about 300 m. The nomenclature and composition of these two rock units are summarized in figure 2. Member names are those of McKee (1945). His Spencer Canyon and underlying members are here reassigned to the Bright Angel Shale so as to retain the stratigraphic position of the Bright Angel Shale-Muav Limestone contact as originally defined by Noble (1914, p. 65). In the central Grand Canyon, the Bright Angel Shale consists of interbedded shaly siltstone and sandstone and interbedded dolomite. The Muav Limestone consists of cliff-forming, in part thin bedded dolomite and limestone that contain subor- SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 dinate shale partings. The clay fraction diminishes in both formations toward the east, and shale is rare in the Cambrian section in the eastern Grand Canyon. Regional dip of the Cambrian rocks in the central Grand Canyon is less than 2° toward the north or northeast. STRUCTURE The axes of the anticlines generally follow the trend of the center of the river (fig. 1, plate 1). Be- tween Kanab and Mohawk Canyons, there is a marked tendency for maximum dips to occur where the anticlines trend northeast. Local closures occur along the northeast-trending folds and anticlinal structures are missing or subdued along northwest- trending reaches of the river. The anticlines are best developed along the 21-km stretch between Kanab and Havasu Canyons. Here the canyon is at its narrowest, talus is minimal, and the river level is at or above the top of the Bright Angel Shale. The steepest dips and greatest structural relief associated with the anticlines occur where McKee’s (1945) Peach Springs and overlying Kanab Canyon Members of the Muav Limestone 113015' 113° 112°45' 112°30' 36°30’ l l l O? Whispering ,./ 2 Falls Grotto 0 5 10 KILOMETERS / 55 $70? 36°15’-— ~ ‘95)? , 5‘7 14% ‘ 2% an a ‘i/ 2)- ( so; 1 /' 1 FIGURE 1.—Locati0n of the Colorado River and its principal tributary canyons in the central Grand Canyon, Ariz. Stippled reaches contain river anticlines. Downstream is to left. ORIGIN OF THE RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA EXPLANATION 1 r I I I I I Limestone / / / Dolomite Sandstone-siltstone __._ Shale-claystone Mixed Iithologies indicated by superposition of patterns METERS — I 50 CENTRAL GRAND CANYON MARBLE CANYON- EASTERN GRAND CANYON DEVONIAN \ d 1 Upper member Havasu Member Gateway ac) Canyon % m Member > .E —J > 8 Kanab 2 Canyon Member Peach Springs Member 0. J a E >0 2 2 o c a .9 T) > 2’ < E 2’ £5 1 \ a) C 2 (D 'o C r 8 39 o m a. .9 J 1 ‘ PRECAMBRIAN CAMISEAN FIGURE 2.——Stratigraphic columns of Cambrian rocks in the central and eastern Grand Canyon. Note that the clay content diminishes eastward. Figure and nomenclature are slightly modified from Elston and Bressler (1977, fig. 2). Member names within Muav Limestone are those of McKee (1945). A3 A4 (fig. 2) crop out at river level. Because the Muav Limestone is more resistant than adjacent strata, the canyon is narrow and the river is commonly less than 75 m wide. This constriction extends from about 0.8 km east of the Kanab Canyon to Fern I ‘Glen (fig. 1). Between Kanab and Havasu Canyons, anticlinal dips in the Muav Limestone locally are as great as 60°, although dips are mostly less than 30° (plate 1; fig. 3). A river anticline that trends past the mouth of Matkatamiba Canyon (fig. 1) is partic- ularly well exposed. Here dips are a maximum of 40° at the river and decrease away from the river. Folded strata extend into the canyon wall as far as 250 m from the center of the river. Below Havasu Canyon, the inner gorge widens, and dips on the flanks of the anticlines commonly are about 2° to 5°. In general, deformed strata extend no more than 25—50 m above river level. Higher on the canyon walls, the strata are undeformed. Folding is entire- ly restricted to the Cambrian rocks. The maximum structural relief along the anti- clines is of the order of 10 m, and most of the river anticlines have less than 5 m of structural relief. Total shortening across the structures therefore is only a few meters. LO‘V-ANGLE THRUST FAULTS Two sets of intersecting low-angle thrust faults and fractures parallel the anticlines in the Muav Limestone and dip both toward and away from the SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 river. The faults are particularly numerous and have greater stratigraphic offsets where the lime- stone is steeply folded. However, we have observed low-angle thrusts in limestone beds that dip as little as 2°. The low-angle thrusts may occur alone but are more commonly found in parallel and echelon sets. Individual faults are discontinuous and rarely ex- tend more than 100 m along the strike of an anti- cline. As shown in figure 4, the faults characteris- tically terminate in breccia zones. Parallel to the river, the fault planes curve in conformance with the structural relief along an anticline. The result— ing curvature of the fault planes indicates that the faults encircle the up- and down—stream ends of the anticlines; however, conclusive evidence for this relationship is missing, because the critical outcrops are below river level. Displacements along the thrust planes are dip- slip and rarely exceed a few meters. The faults that dip away from the river generally dip at angles no more than 15° greater than the beds in which they are found. The faults that dip toward the river in- tersect bedding planes at angles as great as 30°. In most outcrops, there has been considerably more movement along the thrusts that dip away from the “river than along those that dip toward the river. The low-angle thrusts are the result of internal shearing within the limbs of the anticlines. They are clearly a response to the folding of the brittle ‘ limestone in the section. The thrusts occur because FIGURE 3.—Well-developed river anticlines upstream from 010 Canyon. View is downstream toward west. ORIGIN OF THE RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA A5 a FIGURE 4.—A pair of low-angle echelon thrust faults (arrows). The fault that trends from left toward lower center terminates in a breccia zone. The scene is about 7 m wide. no differential movement is possible between the individual bedding surfaces, which are interlocked by minor irregularities. Bedding plane irregularities have as much as 2 cm of relief. HIGH-ANGLE KINK BANDS Discontinuous high—angle reverse kink bands 10— cally parallel the river anticlines. As shown on fig- ure 5, the kink bands displace the rocks up toward the river and are as wide as 2 m. The bands ap- proximately parallel the slope of the canyon walls and are obliquely to vertically jointed in the lime- stone. We interpret the kink bands to result from stress buildup accompanying the upward bulging of the canyon floor during the growth of the anticlines. RELATED ANTICLIN ES Small river anticlines occur in the perennially wet reaches of Kanab Canyon and a few of its tribu— taries (fig. 1). The crests of these anticlines are 10- cally exposed in the floor of Kanab Canyon and Whispering Falls Grotto and are characterized by buckled beds. The buckling demonstrates that the anticlines have formed in response to lateral compression. ANCIENT RIVER ANTICLINES The canyon of the Colorado River is uncommonly narrow at a few localities between Kanab and Tuck— up Canyons. At such places, McKee’s (1945) Kanab Canyon and Gateway Canyon Members of the Muav Limestone are preserved close to and several tens of meters above the present course of the river. Some of these outcrops are anticlinally folded par- allel to the river. Because the rocks are deformed and the structures lie above thepresent generation of river level anticlines, we interpret these folded outcrops to be remnants of former river anticlines that originated before the canyon attained its present depth. We consider these remnants to be evidence that the formation of river anticlines is a continuing process that accompanies the downward cutting of the canyon through the Muav Limestone and Bright Angel Shale. The remnants suggest that as the river cuts through and exposes the anticlines, they erode and are removed as the canyon widens. Consequent- ly, most traces of the older generations of anticlines have been removed. REGIONAL DISTRIBUTION The distribution of river anticlines in the Grand Canyon is shown on figure 1 and plate 1. Significant- ly, river anticlines and their associated faults are not present in Marble Canyon, 60 km east of the area of figure 1, and in the contiguous eastern Grand Canyon where the canyon is also floored by the A6 FIGURE 5.—Kink bands in rocks along the river. A, A high- angle kink band that parallels river, exposed by spalling along the band; the length of kink band is approximately 20 m. B, A schematic diagram of a high—angle kink band on the flank of a river anticline that results from a bulging of the canyon floor; arrows indicate relative movement. Muav Limestone and Bright Angel Shale. We at- tribute this fact to the absence of sufficient clay in the Cambrian section to allow plastic flowage, even when saturated. As shown diagramatically in figure 2, the percentage of claystone in the Cambrian sec- SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 tion decreases from west to east. Where the river flows on these units in Marble Canyon and the eastern Grand Canyon, the Bright Angel Shale con- sists almost totally of fine-grained sandstone and siltsone, and clay represents only a few percent of the total thickness. Similarly, the shaly partings in the Muav Limestone become more silty toward the east, and the percentage of clay is correspondingly less. ORIGIN Ford and others (1974, p. 126—127) have sum- marized the hypotheses advanced to explain the river anticlines in the Grand Canyon. A number of explanations have been proposed, three of which are illustrated on figure 6: 1. Some workers have evoked variants on the theme of strain adjustments resulting from the removal of large volumes of overburden and the consequent realinement of stresses and deformation of the rocks near river level (McKee, in Ford and others, 1974, p. 126; Sturgul and Grinshpan, 1975). 2. Ford (in Ford and others, 1974, p. 126) con- sidered that the Bright Angel Shale is mobile and is squeezed out from under the adjacent walls as a result of the stress gradients pro- duced by the large lithostatic load differential between loaded rocks under the canyon walls and the unloaded canyon floors. Hamblin and Rigby (1969, p. 57 and 60) invoked a com- bination of these mechanisms to explain the anticlines. 3. Shoemaker (1973) proposed that the thick plate comprising the Paleozoic rocks above the Bright Angel Shale passively glides on the mobile Bright Angel Shale toward the canyon in response to a gentle northeastward dip. Shoemaker envisioned the plate as being rigid and the horizontal extension Within it as be- ing largely accommodated by tensional separa- tions along joints. 4. Huntoon (in Ford and others, 1974, p. 127) thought that the anticlinal bulge along the canyon floors resulted from swelling of clay in the Muav and Bright Angel Formations. We embrace the hypothesis that the river anti- clines are an unloading phenomenon. We believe that the anticlines represent strain within the shaly parts of the Muav and Bright Angel formations and that the strain results from large stress gradients that exist between the unloaded canyon floor and rocks under the heavily loaded adjacent canyon walls. In ORIGIN OF THE RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA \ lPPs x Mr \ Dtb \ \ / _, €m e I] I \\ / K I I 2' p€ 2' 4 k \ A. Distributed flowage resulting from release of con- fining pressure. lPPs Mr Dtb .. V ->—>—-—>'/\<—<—+ ->€b-O—>€lt—/'/T\ <— <— <- / B. Flowage in incompetent shale layers clue to differ— ences in lithostatic load. (— lPPs ‘— Mr {— Dtb ‘— Cm W— 4— <— <- <—- <— <- Cba €t I C. Gliding of a plate on an incompetent layer. FIGURE 6.—Selected theories used to explain the origin of river anticlines in the central Grand Canyon, Ariz. A, This report. B, Ford and others (1974) and Sturgul and Grin- shpan (1975). C, Shoemaker (1973). Combination of fig- ures A and B, Hamblin and Rigby (1969). Arrows indi— cate movement caused by stress. p6, Precambrian igneous and metamorphic rocks. Ct, Tapeats Sandstone. Cba, Bright Angel Shale. Cm, Muav Limestone. Dtb, Temple Butte Limestone. Mr, Redwall Limestone. IF'Ps, Supai Group, capped by Esplanade Sandstone. A7 _ our model, flowage is initiated in the shaly parts of the Bright Angel and Muav formations close to the river as they become exhumed. The model in part devolves from the observation that fossil river anticlines are preserved strati- graphically high in the Muav Limestone, well above present river level (pl. 1). The fossil anticlines are found where the central gorge of the canyon is very narrow. They do not appear to be physically con- nected to underlying river anticlines found at present river level. We consider this evidence that formation of the river anticlines has been a con- tinuing process and that older river anticlines once existed but have been removed by erosion during widening of the canyon. The primary differences between our hypothesis and that of Ford (Ford and others, 1974, p. 126) are that we recognize that saturation of the shale in the section is necessary for the rocks to become plastic enough to deform under the imposed stress gradients and that the river anticlines are not nec- essarily related to the presence of a closely under- lying Bright Angel Shale. Without saturation and without shale, these same rocks do not deform. The various hypotheses proposed by others and the one we favor can be examined in light of the structure of the anticlines and of other data col- lected during the course of our investigation. One fact is clear from field observations: The strain ob- served in the river anticlines is a combination of plastic flowage and brittle failure within the Muav Limestone and of mostly plastic flowage within the Bright Angel Shale. A concept that implies elastic rebound, such as that proposed by Sturgul and Grinshpan (1975), is negated by the observed non- elastic strain. Huntoon’s swelling—clay hypothesis was based on observations in which clay partings between indi- vidual limestone beds appear to be expanded under the floor of Whispering Falls Grotto. The grotto is eroded beneath a waterfall in beds near the middle of the Muav Limestone. Exposures of the grotto floor indicate a slight but measurable increase in the thickness of the shale partings toward the axis of the fold. To test for the presence of swelling clay, we collected claystone and siltstone samples from shaly and thin-bedded parts of the Muav Lime- stone and Bright Angel Shale in the central and eastern Grand Canyon, in areas with and without river anticlines. X-ray analyses revealed that the clay in all samples is composed of illite and kaolinite or chlorite. Montmorillonite, if present, occurs only in trace amounts. Because of the Virtual lack of A8 swelling clays in all samples, we believe that the swelling-shale hypothesis can be rejected. The ex- panded beds observed at Whispering Falls Grotto thus are undoubtedly the result of physical separa- tions between limestone beds caused by buckling of the strata along the crests of the anticlines. The hypothesis of Shoemaker involves a plate that glides toward the river on the incompetent Bright Angel Shale and is detached from the northeast— trending Sinyala fault that parallels the general trace of the river on the south (pl. 1). In this model, the plate moves northwestward toward the river on a dipping surface that slopes less than 2° through- out the central Grand Canyon. As envisioned by Shoemaker, the top part of the Bright Angel Shale deforms plastically or by shearing. As long as the upper contact of the Bright Angel Shale is near or below river level, the MuaV Limestone and upper part of the Bright Angel Shale would bulge upward along the axis of the river. Deformation would be confined to the northeast-trending segments perpen- dicular to the direction of plate movement. Shoemaker’s hypothesis would explain the tend— ency for the anticlines to aline along the northeast— trending segments of the river. However, we find difficulties with the theory. First, we have been unable to document tensional separations alongjoints parallel to the river in the cross sections along tribu- tary canyons south of the river. Moreover, there are no faults or grabens alined parallel to the river that exhibit tensional separations ( Ford and others, 1974, p. 127). (The northeast—trending Sinyala fault is a deep-seated fault.) Tensional separation should be evident in the brittle and erosionally resistant Permian Esplanade Sandstone that caps the benches surrounding the inner gorge. Secondly, no tear faults, which would be required to accommodate the northwestward motion of the moving plate, have been observed along or parallel to the northwest- trending segments of the river. Thirdly, anticlines that we have observed near the upstream and down— stream ends of the river anticline belt (fig. 1, pl. 1) occur in competent limestone beds in the lower and middle parts of the Bright Angel Shale (fig. 2). This demonstrates that anticlinal deformation has occurred through the full thickness of the Bright Angel Shale and is not necessarily restricted to the upper part of the formation, as would be required by the model depicted in figure 60. We thus con— clude that there is no field evidence for appreciable lateral movement of the Paleozoic section toward the Colorado River. SHORTER CONTRIBUTIONS TO STRATIGR APHY AND STRUCTURAL GEOLOGY, 1979 From field relations, then, flowage of saturated shaly strata involves only small volumes of rock in close proximity to the river. These rocks are folded at places where there is a large stress between heavily loaded zones under canyon walls that are as high as 650 m and an unloaded canyon floor. Shaly strata appear to have been squeezed laterally toward the river, and local lateral flowage, which has amounted to no more than a few meters at the river, has occurred in response to the compressive stress to create the river anticlines. In this model, the thrust faults result from the folding of the strata in the anticlines and they are not primary shears re— lated directly to differential lateral flowage of the rocks under the canyon walls. Saturation of the shale is essential for the defor- mation. The lack of flowage features in shaly out- crops high in the walls of the Grand Canyon or away from perennial streams supports this conclu- sion. Hundreds of kilometers of such dry outcrops exist in the Grand Canyon under load conditions identical to those near the river anticlines, yet anti- clinal deformation has only occurred where the shaly rocks have been or are in nearly constant contact with water. Our hypothesis accounts for the formation of river anticlines but not for a preferred orientation where they might tend to parallel a deep—seated fault and not follow the stream course in detail. West of Fern Glen, the anticlinal trends vary, and some minor river anticlines trend northwest as well as northeast. However, east of Fern Glen, very well developed anticlines trend northeast along northeast— trending stretches of the river and disappear along the short segments of the river that trend northwest (fig. 1, pl. 1). In this area the river is paralleled on the south by the northeast-trending Sinyala fault, which lies about 3—6 km to the southeast. The pos- sibility exists that some small component of north— westward, downdip translation of the Paleozoic sec- tion has occurred between the Sinyala fault and the Colorado River, employing the model of Shoemaker in a restricted sense. Such creep would produce a differential stress that presumably would enhance development of river anticlines along the northeast- trending stretches of the river and suppress their formation along the northwest-trending stretches. The differential stress would augment stresses pro- duced as a consequence of unloading along the river, which in turn suggests that a preferred orientation of stream or river anticlines might be a compara— tively sensitive indicator of differential stresses of tectonic origin. ORIGIN OF THE RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA REFEREN CES CITED Elston, D. P., and Bressler, S. L., 1977, Paleomagnetic poles and polarity zonation from Cambrian and Devonian strata of Arizona: Earth and Planetary Science Letters, v. 36, no. 3, p. 423—433. Ford, T. D., Huntoon, P. W., Billingsley, G. H., Jr., and Breed, W. J., 1974, Rock movement and mass wastage in the Grand Canyon, in Breed, W. J., and Boat, E. C., eds., Geology of the Grand Canyon: Flagstaff, Museum of Northern Arizona and Grand Canyon Natural History Association, p. 116—128. Hamblin, W. K., and Rigby, J. K., 1969, Guidebook to the Colorado River, pt. 2, Phantom Ranch in Grand Canyon National Park to Lake Mead, Arizona-Nevada: Provo, Utah, Brigham Young University Geological Studies, v. 15, pt. 2, 126 p. A9 McKee, E. D., 1945, Stratigraphy and ecology of the Grand Canyon Cambrian, pt. 1 of Cambrian history of the Grand Canyon region: Carnegie Institute of Washington Publication 563, p. 31—168. Maxon, J. H., 1969, Preliminary geologic map of the Grand Canyon and vicinity, Arizona (western and central sec- tions): Grand Canyon Natural History Association (PO. Box 219, Grand Canyon, Ariz.) Shoemaker, E. M., 1973, River anticlines of the Colorado: Paper presented at Symposium on Northern Arizona Geology, Museum of Northern Arizona, August 31, 1973 (unpublished). Sturgul, J. R., and Grinshpan, Z., 1975, Finite-element model for possible isostatic rebound in the Grand Canyon: Geology, v. 3, p. 169—171. Rock Control and Tectonism— Their Importance in Shaping the Appalachian Highlands By JOHN T. HACK SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—B FIGURE TABLE 1. ewews N!“ 9° CONTENTS Page Abstract _____________________________________________________________________ B1 Introduction __________________________________________________________________ 1 Major topographic features and drainage systems _______________________________ 2 Relation of the highland form to bedrock geology _____________________________ 2 The belt of (‘amhrian and Ordovician rocks ___________________________________ 6 Duration of erosion _____________________________________________________________ 1‘ Evidence for late tectonic activity _____________________________________________ 9 Rates of erosion and uplift ___________________________________________________ 12 Development of the Appalachian Highlands in late geologic time _______________ 14 Conclusion ____________________________________________________________________ 15 References cited _______________________________________________________________ 16 ILLUSTRAITONS Page leneralized topographic map of the Appalachian Highlands. showing: the divide between the drainage to the Atlantic and to the Gulfs of Mexico and St. Lawrence ______________________________________________ B3 Subenvelope map showing contours drawn on altitudes 211011;: the major stream channels ________________ 4 Generalized map of major rock units based on their resistance to erosion _________________________________ 5 Profile and outline map of the Cambrian and Ordovician helt from Coosa River to Lake Champlain ________ 7 Generalized relief map of l'iedmont and Blue Ridge provinces __________________________________________ 10 Map of Eastern United States showing Bouguer gravity anomalies _______________________________________ 14 TABLES Page Thickness of sediments penetrated in COST No. B—2 and estimated deposition rate ________________________ B9 Estimates of erosion rates ______________________________________________________________________________ 12 Estimates of tilt rates _________________________________ _ ________________________________________________ 12 III ROCK CONTROL AND TECTONISM—THEIR IMPORTANCE IN SHAPING THE APPALACHIAN HIGHLANDS By JOHN T. HACK ABSTRACT The current interest in contemporary tectonic processes in the Eastern I‘nited States is turning up abundant evidence of crustal movements in late geologic time. Topographic analysis of the highland areas from the southern Blue Ridge to the Adirondack Mountains indicates that most of the landforms owe their origin to erosion of rocks of (lif- ferent resistance rather than to tectonic processes. Most areas of high relief and high altitude have been formed on resistant rocks. The Cambrian and Ordovician belt. containing mostly shale and carbonate rock. on the other hand. forms an exten- sive lowland from Alabama to the Canadian border and girdles the Adirondack Mountains. Differences in altitude can be ex- plained by the presence of resistant rocks outside the belt; these resistant rocks form local base levels on the streams that drain the belt. A few areas may have undergone local uplift at a higher rate than areas nearby— for example. the Piedmont region northwest of Chesapeake Bay. Most estimates of erosion rates. based on the load transported by streams, and of uplift rates. based on removal of the known or in— ferred amount of overburden during a known period of time. are of the same order of magnitude. averaging about 4 x 10’2 millimeters per year. Rates of uplift. based on study of tilted Pleistocene beaches and repeated geodetic traverses. are at least an order of magnitude higher for comparable areas. Tectonic uplift of the highlands has been slow and involves mostly warping or tilting on a large scale. Erosion rates keep up with or exceed the rate of uplift and have been suflicient to mask evidence of faulting or other differential movements. The high rates of uplift that are inferred on tilted water planes in the glaciated regions or that are measured by dif- ferences in repeated geodetic traverses cannot have been sustained for long periods of time. INTRODUCTION Differential erosion of rocks of varying resistance has long been considered a major factor in shaping the landscape of the Eastern United States. Theories to explain the erosional history have been concerned mainly with cyclical landform development. relict sur- faces such as peneplains. and the evolution of the drainage pattern (Thornbury. 1965. p. 72—87). Al- though most geologists agree that uplift did take place in the Eastern I'nited States during late geologic time (since the middle Mesozoic). few have thought that local topographic forms are of tectonic origin. Faults active in late geologic time have been discovered (York and Oliver. 1976; )Iixon and Nowell. 1977’). but it has not been proved that any modern topographic fea- tures owe their origin to contemporary faulting. Recent studies of precise leveling data along geodetic traverses that. have been rerun one or more times in- dicate that many of the monuments have changed in elevation during the intervals between the traverses. Some of these changes may be due to errors in leveling or to disturbances of the monuments. but it is generally agreed that many of the changes are related to actual motion of the ground or of the Earth‘s crust. The dif- ferences between traverses expressed as velocities of vertical motion are commonly quite large (see Brown and Oliver, 1976. for discussion of the method of analysis). The interpretation of the repeated traverses poses many problems. The rates of uplift. or depres- sion are commonly so high that they could not be long sustained without resulting in much larger topographic features than are observed. In spite of the difliculties involved in interpreting evidence for recent crustal movements. there is little doubt that rates obtained from precise leveling data as well as from study of ancient strand lines indicate rapid contemporary crustal movements. Analysis of topography of the Appalachian Highlands shows. however. that diii'erential erosion of rocks of different resistance is the major factor that controlled the de- velopment of today’s topographic forms. including even large features. A few major topographic discon- tinuities apparently unrelated to rock type do exist. however. and can be identified. The picture that emerges from topographic analysis is that the major drainage system of the area has. over a long period of time. become closely adjusted to rock type. In places. it still deviates from lithologic controls. Some of the deviations occur in response to competing major drain- age systems that have ditferent base levels. lengths, and gradients. which are basically caused by rock con- trol within the different systems. Some may be due to tectonic influence. and others may be inherited from the past. B1 132 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 The rate of topographic adjustment because of ero- sion is so rapid that detection of the details of tectonic control is difficult or impossible, even in areas where it is reasonable to expect that such control exists. On the other hand, the sedimentary record preserved in the Coastal Plain and Continental Shelf as well as the grossest topographic relationships are clear evi- dence of tectonism. The intent of this paper is to de- scribe the major drainage systems and topographic forms of the Appalachian Highlands between Ala- bama and northern New York and to discuss their ad- justment to bedrock. Some topographic anomalies that may be caused by differential movement of separate tectonic blocks are also considered. I wish to thank C. S. Denny, Richard Goldsmith, Sheldon Judson, and “'ayne Newell for their thought- ful criticism. MAJOR TOPOGRAPHIC FEATURES AND DRAINAGE SYSTEMS The general outlines of the topography and drain— age of the Appalachian Highlands are shown in fig- ures 1 and 2. In figure 1, an envelope map originally defined by Stearns (1967), contours are drawn across the ridge crests and drainage divides. Essentially, fig- ure 1 is a map of the high parts of the topography. In this paper, the envelope map is used as a reference map showing the major streams and the continental divide separating the drainage to the Atlantic Ocean from that to the Gulf of Mexico and St. Lawrence River. Figure 2, a subenvelope map (Stearns, 1967; Hack, 1973), is another way to generalize the topog— raphy. In this map, the contours are drawn on alti— tudes along the major streams. It shows the position of the major drainage divides as well as the general configuration of the topography. The degree of gen- eralization is controlled by the spacing of the streams used in the generalization. The streams chosen can be selected precisely, simply by eliminating the upper reaches of a certain length of all the streams. The length of reach chosen for elimination determines the degree of generalization. The subenvelope map is useful for the present purpose because the larger streams are flowing on bedrock, and their channel gradients are adjusted to the bedrock. In preparing figure 2, the contours were drawn on the streams shown on the U.S. Geological Survey’s base map of the United States at 1:2 500 000 scale. Contours were plotted only on streams more than 32 km long. The resulting map shows the altitudes to which the major streams have eroded. The true alti— tudes of the divide areas, of course, are higher than the contours shown. Narrow ranges of mountains that do not correspond to the major divide areas do not have topographic expression on the map. For example, the Blue Ridge of northern Virginia north of Roanoke (fig. 1, BR) contains peaks higher than 1220 m, but, as the entire range averages less than 32 km in width, it does not shown in figure 2. The highland area (fig. 2) from Georgia to northern New York is about 1600 km long and averages about 240 km in width. It is broken by several large re- entrants. The Tennessee River valley at T in the south divides it into two prongs. The New River valley forms a prominent reentrant at N and is the only river system that crosses the highland from south- east to northwest. Large reentrants occur at the Po- tomac River basin (P) and Susquehanna River basin (S). The Mohawk Valley at M separates the Adiron- dack Mountains from the main highland area. Many of the forms are closely related to the kinds of bedrock directly beneath them. This relationship prob- ably accounts in large part for the features just men- tioned, thus obscuring the evidence for the interpre- tation of the Appalachian and Adirondack Mountains as uplifted features. The problem, then, is the extent to which the present topography has been formed by differential erosion as opposed to tectonism. In a general way, the highlands owe their existence to uplift, for the rocks of the highlands pass under the Coastal Plain at the southern end. On the other hand, local topographic features such as the individual ridges in the Valley and Ridge province owe their height to differential erosion. The local importance of differential erosion is evident even in the southern Blue Ridge, an area of. complex geology in which it is more difficult to identify resistant or nonresistant rock (Hack, 197 3, 1976). Were the highlands, including the Adirondack Mountains and the Blue Ridge, uplifted as a single tectonic unit or are they composed of a group of separate tectonic blocks? If the highland is a single unit, has the uplift been differential, and at what scale or wavelength? RELATION OF THE HIGHLAND FORM TO BEDROCK GEOLOGY Figure 3 is a generalized geologic map of the Ap- palachian Highlands that has selected contours from the subenvelope map superimposed on it. The rock units are grouped according to broad lithologic types having different resistance to erosion. Units 1 and 4 contain extensive outcrop areas of resistant rock that tend to form highlands or high relief. Unit 2 con- tains alternating sequences of resistant and nonre- sistant rocks, which, where folded, produce the elon- gate ridges and valleys of the Valley and Ridge ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS B3 88° 84° 80° 76° 7 44° 40° ? zoo 490 MILES or 250 450 KILOMETERS EXPLANATION fig 1000-2000 2000—3000 3000-4000 Over 4000 feet 305-610 610-915 915-1220 Over 1220 meters —————— Continental divide FIGURE 1.—Generalized topographic map (en- velope map) of the Appalachian High- ferred to in text include: A, Adirondack lands, showing the divide between the Mountains; BR, Blue Ridge of northern drainage to the Atlantic and to the Gulfs Virginia. of Mexico and St. Lawrence. Localities re- B4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 O 88° 84° 80 T I 200 400 MILES I l O-O 250 450 KILOMETERS Contour interval 200 feet (61 meters) — —— — Drainage divide FIGURE 2.—Subenvelope map showing con- source are not included. Ch, Chattahoochee tours drawn on altitudes along the major River; T, Tennessee River; N, New River; stream channels. Altitudes on streams less P, Potomac River; S, Susquehanna River: than 32 km long as measured from the M, Mohawk River. province. Unit 3 contains mostly nonresistant shale New York and Pennsylvania and the Permian of West and carbonate sequences. The four units are defined Virginia. The rocks of Pennsylvanian age are elastic as follows largely on the basis of summary by Colton, sequences, both marine and continental, and contain 1970) : coal beds. They include thick sandstone, especially in Unit 1: Includes the entire sequence of rocks of the lower part of the section. These more resistant Pennsylvanian age as well as the Upper Devonian of rocks crop out near the eastern edge of the Appala- ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS B5 88° 84° 800 760 44° 44° 40° 40° 36° 36° 32° 9 290 490 MILES 5 260 460 KILOMETERS Contour interval 400 feet (122 meters) EXPLANATION E Unit l-—Mostly rocks of Late Devonian and Pennsylvanian age Unit 2——Mostly rocks of Silurian to Middle Devonian age Unit 3--Sedimentary rocks of Cambrian and Ordovician age m Unit 4--Rocks of Precambrian age (from King and Beikman, 1974) -—'—' Inner edge of Coastal Plain FIGURE 3.——Generalized map of major rock units based on their resistance to erosion. Se- lected contours from the subenvelope map are superimposed. I56 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 chian Plateau and in the area east of the Susque- hanna River basin. The higher areas of the plateau of northern Pennsylvania as well as the Catskill Moun— tains are underlain by continental elastic rocks of Devonian age. They include coarse conglomerate and sandstone in the units designated as Upper Devonian by King and Beikman (1974). Unit 2: Includes virtually all kinds of rock typical of platform deposits ranging from limestone and shale to sandstone and conglomerate. This unit includes the following chronostratigraphic map units of King and Beikman (1974): Silurian, Lower and Middle Devo- nian, Devonian, undivided, and Silurian and Devo- nian, undivided. The Upper Devonian of King and Beikman (1974) is not included. South of Pennsyl- vania, the unit also includes Mississippian sequences. Unit 3: Includes the sedimentary rock of Cambrian and Ordovician age, mostly carbonate rocks and shale. The carbonate rocks increase in thickness from 180 m in New York to more than 3000 m in Tennessee. Shale, siltstone, and mudstone make up most of the re- mainder; they increase in thickness to the north. The unit omits the basal quartzite south of Pennsylvania but includes it to the north. Unit 4: Includes the rocks designated Precambrian by King and Beikman (197 4). These rocks vary wide- ly in their degree of resistance, but they do contain resistant sequences of great thickness, especially in the south. The Ocoee Supergroup, for example, which un- derlies the Great Smoky Mountains and the Black Mountains, including Mount Mitchell, contains 5200 m of metamorphic quartz-rich sandstone and silxtstone. The narrow Blue Ridge of Virginia is underlain by quartzite, metavolcanic rocks, and coarse-grained granitic rocks. The subenvelope map superimposed on the geologic units (fig. 3) indicates that, except near the southern edge of the Appalachians, the high areas are on the most resistant rocks. The Lower Pennsylvanian se- quence (unit 1), the Devonian elastic sequence (unit 2), and the Precambrian rocks (unit 4) form the highest areas. The bifurcation of the southern high— land occurs where the main highland shifts southeast- Ward to correspond with the large outcrop area of unit 4. The highland on the Pennsylvanian sequence continues to the southwest, forming the. western prong. In southern Pennsylvania, sandstone of Pennsyl- vanian age holds up large plateaulike areas. In the northern part, the Devonian and Mississippian clas- tic rocks underlie the present erosion surface and form the northern glaciated section of the Allegheny Plateau. The Catskill Mountains of New York locally include peaks as high as 1220 m. The Adirondack Mountains form a separate dome underlain by rocks of Precambrian age. In contrast, the large areas of nonresistant rocks found mostly in units 2 and 3 are responsible for the large reentrants in the Appalachian highlands at T, P, and S in figure 2. The large valley at T is drained by the Tennessee River and its tributaries and cor- responds to a broad outcrop area of carbonate rocks (unit 4). At P, the principal tributaries of the Po— tomac are subsequent streams flowing in valleys in carbonate and shale (unit 2). The same is true of the Susquehanna basin at S. THE BELT OF CAMBRIAN AND ORDOVICIAN ROCKS At the crossover of the subenvelope crest from the Precambrian area to the plateau in southwest Vir- ginia, the altitude of the subenvelope (fig. 3) appears to be unrelated to the bedrock. It crosses the nonre- sistant Cambrian and Ordovician sequence at an alti- tude of more than 730 m comparable with or higher than, altitudes at many other places along the crest. At the Hudson River, for comparison, the Cambrian and Ordovician belt is almost at sea level. Study of this belt of rocks shows that the altitudes within it can be explained at least in part by adjustments of the major streams that cross the belt to bedrock down- stream from the belt. These adjustments, in effect, form local base levels at various altitudes that control the profiles upstream. Figure 4A, which illustrates the argument for this hypothesis, is a profile along a series of stream reaches that flow along the strike in the Cambrian and Ordovi- cian rocks. The profile is drawn along the streams as the crow flies and does not include bends or meanders. The altitudes are on the valley floors or flood plains adjacent to the streams. Except for small parts of the Coosa River, the streams are entirely Within the Cam- brian and Ordovician rocks. However, all have tribu- taries or upper reaches that head in more resistant rocks, and they depart from the belt to cross terrains having a great variety of rock types. The divides be- tween the streams within the belt are narrow every— where, so that a continuous longitudinal profile is formed. The profile along the strike of the Cambrian and Ordovician rocks (fig. 4A) is divided into concave- upward segments, each of which is part of the drain- age basin of a large stream that exits the Cambrian and Ordovician belt. The high point is near Rural Re- treat south of the continental divide, between head- waters of the Tennessee and New Rivers. This place is close to Where the Cambrian and Ordovician belt B7 ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS OOO. OOON .3de :3: @3033 we $633 -1“ “:33me WSQESZ .Efimawno 3:3 8 33m 380 Sony :3 5325520 9% anEQENO 9: me Q: Qua wince ”Ex :3 vacamllé Hannah x»/ / :v ,MW, pm o3 o% kiln-3% lifehf , . 4 I H ¢mwv «wow , b mmmkm2043+oo¢ oom ooN oo. o mMJ_z oov oom cow om. \d (90) (0d) HHAIH ovwomoa HHAIH EHVMV'IHCI / gmmuuwm Hmunm (nH) HEAIH NOSGHH (OS) Auwwmv mmuaufiuam NIVTdNVHD axvq HHAIH TTIXTADHOS (n3) HHAIH VNNVHHHGSHS (PD)HEAIH sawvr (OH)HEAIH HXONVOH (em) HHAIH HHSSENNHL (3N)HHAIH MEN wUH>flQ Hmucmcfluzou B8 narrows and is broken up by ridges of more resistant rocks younger than Ordovician. Profiles of streams of small and intermediate size such as those in the belt are generally shaped by the amount of discharge and various factors related to flow resistance of the chan- nel, the most important of which are probably load and size of bed material. The channel slopes of streams are normally inversely proportional to the discharge, and, as discharge and length are directly related quan- tities, channel slopes are generally inversely propor- tional to a function of length. This explains the con- cave-form characteristics of most longitudinal valley profiles. As channel slope tends to be directly propor- tional to a function of size of bed load, the concave profile is modified in steepness and in the details of its form by the resistance of the rocks in its drainage basin. The high point at Rural Retreat on the divide between the Tennessee River drainage and the New River drainage is a striking feature of the composite profile (fig. 4A). It can be explained by differences in the geology of the Tennessee River drainage basin and the New River basin (Hack, 1973). The Tennessee system descends from Rural Retreat via the Holston River as far as Chattanooga through a broad valley in shale and carbonate rocks. The altitude at Chatta- nooga, where the river system exits the Cambrian and Ordovician belt, is only about 210 m. Below this point, it crosses resistant sandstone beds in a short reach below Chattanooga. It crosses northern Alabama through an area underlain mostly by Ordovician and Mississippian limestone. Thus, almost its entire course is in nonresistant rocks. The New River, on the other hand, heads in the Blue Ridge and crosses the Cambrian and Ordovician belt at a narrow point. It then flows northward toward the Ohio River, crossing thick sandstone sequences of Pennsylvanian age. Its gradient is actually steepened in these reaches in spite of the increased discharge. Its steep descent ceases above Charleston where it joins the Gauley River at an altitude of about 200 m. Northeast of the New River, the regional drainage is to the southeast into the Atlantic by streams that head along the resistant rocks of the plateau and flow more directly across the Cambrian and Ordovician belt. Only in the Potomac and Susquehanna basins are there major streams parallel to the regional strike. The Roanoke and James River cross resistant Precambrian rock before reaching the Piedmont lowlands, but the outcrop widths are narrow compared with those of the resistant rocks crossed by the New River. The P0— tomac exits the belt at a place where the resistant rocks are unusually thin (Hack, 1965, p. 28). The Susquehanna and Schuylkill Rivers flow directly SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 across the nonresistant beds of the Newark Group in the Triassic and Jurassic basin. The Delaware crosses only a narrow band of rocks of Precambrian age in the Reading Prong. The absence of a high resistant barrier between the Cambrian-Ordovician rocks and the Piedmont lowlands in Pennsylvania permits even small streams like the Schuylkill to survive without capture by subsequent drainage from the Susquehanna or Delaware. In New York State, the Cambrian and Ordovician belt forms a narrow lowland separating the Appalachian Plateaus province on the west from the New England province on the east. The altitudes are low; the Hudson River is at tide level as far north as Albany. The belt intersects tidewater again at the St. Lawrence River. Adjustments to resistant and nonresistant rocks do not explain why the Hudson River cuts through the Precambrian rocks of the Hud- son Highlands in a preglacial gorge southeast of the Cambrian and Ordovician belt. Many explanations have been offered (Thornbury, 1965, p. 165). The relationships cited above show that the gen- eral topography of the belt can be explained without involving the hypothesis of differential uplift, except, of course, the obvious condition that the entire Ap- palachian system formed as a result of one or more periods of uplift at some time in the past. The height of the divide at the New River crossing is explained by the necessity for that river to cross many miles of re— sistant rocks downstream exposed on the Appalachian Plateau. North of the New River, the lower areas are explained by the fact that the drainage out of the Cambrian and Ordovician belt crosses lesser barriers to reach the lowlands. The altitudes marginal to the Cambrian and Ordo— vician belt can be judged by the spot altitudes of peaks shown on figure 4B. These altitudes are generally high all along the belt except locally in Pennsylvania, where the belt is bordered on the south by weak rocks and on the north by narrow ridges underlain by thin re- sistant beds. DURATION OF EROSION The general evidence for a close adjustment of the surface forms to the rocks, especially obvious in the Cambrian and Ordovician belt, is an indication that the region must have been exposed to subaerial erosion for a long time; however, more positive evidence does exist. A narrow zone of deposits containing lignite, wood, leaf fragments, and other organic material was called the Brandon Formation by Clark (1891), who regarded these occurrences as of Eocene age. The de- posits are now known to range in age from Cretaceous ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS B9 through Tertiary. They extend down the southeastern margin of the Appalachian basin from Vermont as far south as Alabama. These deposits are small and are generally in residuum overlying the Lower and Mid- dle Cambrian carbonate rocks of the Shady and T oms- town Dolomites and their equivalents. The surface topography in the zone of these deposits is in many places pitted by modern sag ponds that commonly con- tain peat and other organic matter of Pleistocene to Holocene age. The close association suggests a common origin for the organic deposits. They have evidently been formed by solution in the underlying carbonate rocks as part of a continuous process operative since the Cretaceous. The literature on the occurrences is scattered, but a discussion of the implications of a Cretaceous locality in Pennsylvania has been presented by Pierce (1965). Some modern sag ponds have been described by Craig (1969). Additional references have been cited by Mathews (1975). As the fossil deposits could have slumped downward several thousands of meters since they formed because of solution of the underlying carbonate rocks, they do not necessarily inform us about the past nature of the relief, but their occurrence and the lack of marine fossils indicate that the Cambrian and Ordovician belt has been exposed to erosion for a long time. EVIDENCE OF LATE TECTONIC ACTIVITY Abundant evidence indicates that tectonism has taken place in late geologic time, that is, during the late Mesozoic and Cenozoic. The evidence includes changes in the rates and kinds of deposition through time on the Continental Shelf, late faults, tilting or warping of features of the coastal zone, differences in height of land, and areas of anomalous relief. The sedimentary record in the Coastal Plain and Conti- nental Shelf has, of course, been used by many to in— terpret the history of the hinterlands, and this record will surely become increasingly useful as exploration of the shelf continues. Analysis of the Coastal Plain and shelf deposits, however, is outside the scope of this paper, though the use of sedimentary volumes to interpret erosion rates on the land should be men- tioned. Mathews (1975) estimated, by analysis of a series of sections off the Atlantic coast, that at least 2,000 m has been eroded during the Cenozoic on the adjacent land, an average rate of 0.027 mm/yr. The data also suggest that erosion was greatly accelerated during the Miocene. Data (table 1) from the COST No. B—2 well on the outer shelf off New Jersey (Scholle, 1977) permit an estimate of sedimentation rates since the beginning of the Cretaceous, though density, compaction, and other factors are not taken into account. As this well seems to penetrate a typical sequence of major units, the sedimentation rates probably are roughly indicative of relative erosion rates in the source area. The data support the idea that rates were high in Early Cre— taceous and Miocene time. Judging by the quick response of the crust by up- lift to the removal of the continental glacier of Pleis- tocene age, changes in erosion rates probably have been balanced closely by changes in uplift rates; the reverse was probably also true. Thus, high uplift rates probably prevailed in the Early Cretaceous and Mio- cene, and lower rates, from Late Cretaceous to Eocene time. Nevertheless, uplift was probably continuous throughout post-Jurassic time. That faulting took place in late geologic time in the Eastern United States is not easy to demonstrate. Nevertheless, actual faults cutting beds of late geo- logic age have been observed at more than 30 places in eastern North America (York and Oliver, 1976). Probably many more late faults will be discovered. Most of them are along the Fall Zone where thin sedi- mentary strata overlap the basement. Two of these faults have been studied in some detail. The Stafford fault zone, west of the Potomac River (Mixon and Newell, 1977), consists of en echelon faults and a parallel monocline 56 km long. The displacements are small, not exceeding 60 m. They involve sediments as young as middle Tertiary, but some movement in the late Tertiary is possible. The sense of motion is reverse. A somewhat similar group of faults that had small displacements in late geologic time has been found near Augusta, Ga. These faults are also reverse and involve Cretaceous and younger sediments, possibly as young as Miocene (David Prowell, oral commun. 1977). The displacements observed are too small to account for the rate of uplift of the area inland from the Coastal Plain as determined by the sedimentary record. Other faults farther inland would be diflicult to de- TABLE 1.—Thiclmess of sediments penetrated in COST No. 3—2 (Scholle, 1977) and estimated deposition rate Du ration Thickness . . . Rate Period or epoch 3?;2éqgs of sin-3mm (mm/yr) Pliocene and 12 130 0.011 Pleistocene. Miocene ___________ 14 860 .061 Oligocene __________ 12 150 .012 Eocene ____________ 16 220 .014 Paleocene __________ 11 60 .005 Late Cretaceous ___- 35 980 .028 Early Cretaceous ___ 36 2470 .068 Total ________ 136 4870 Average _____ .036 B10 tect because of the sparse sedimentary cover younger than the basement. Another tectonic process of late geologic age is the general downward tilt of the coastal zone to the north- east. This tilt can be seen on figure 3, which shows the Fall Zone or inner margin of the Coastal Plain. In Georgia, the Fall Zone is crossed by streams at an altitude of 120 m. At the northern end of Chesapeake Bay, the zone is at sea level, and, in New England it is below sea level. The tilt of the Fall Zone is matched by the tilt of the Continental Shelf. These features imply a broad tilt of the continent down to the north- east. The head of tidewater extends inland different distances in different rivers, as shown by the 0 contour in figure 3, but this phenomenon is not necessarily re- lated to the tilt of the coastal zone. Eardley (1964) explained the tilt of the Coastal Plain and the Con- tinental Shelf as part of a worldwide process of pole- ward rise of sea level related to a long-term decrease in period of the Earth’s rotation, beginning in the 88° SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 Cretaceous. The fact that some river valleys appear to be drowned more deeply toward the north, however, suggests that a process acting more rapidly may be involved. Differences in relief or sharp changes in altitude may be evidence for unequal uplift in the past, pro- vided that relief differences are in rocks of similar resistance to erosion. Figure 5 is a simplified relief map of the Piedmont and Blue Ridge provinces south of the glacial border. The two provinces have extreme differences in relief. The Piedmont relief is compara- tively low over extensive areas, and the map is in- tended to emphasize differences in relief within the Piedmont. Most of the areas of. moderate relief on the outer Piedmont are directly related to resistant bed- rock. For example, the Pine Mountain belt contain- ing quartzite beds forms a moderate relief area at A. Cataclastic rocks of the Brevard zone occur at B. Re- sistant beds of the Kings Mountain belt are found at C. At area D, moderate relief is caused primarily by fel- 800 76° r 1 l EXP LANAT I ON 40 High relief Moderate relie f Low rel ief 32° l _-.__— 36 32° 200 | O—-o I 200 FIGURE 5.—Generalized relief map of Pied- mont and Blue Ridge provinces. The re— lief classes are based on the number of contours within square areas 10 km on a side, as indicated on the 12250000- seale series of topographic maps by the US. Geological Survey. Less than 90— 400 MILES 4| I 1100 KILOMETERS 120 m in low-relief areas and 120—240 In in moderate-relief areas. High-relief areas are generally more rugged and have relief of more than 240 m and as much as 1200 in. Letters indicate areas discussed in the text. ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS sitic rocks of the slate belt. A large area of moderate relief (E) lies between the Coastal Plain and the Cul- peper and Gettysburg Triassic and Jurassic basins. Its highest point exceeds 305 m above sea level near West- minster, Md. This area was referred to as the VVestZ minster anticline by Campbell (1929, 1933), who be- lieved that it represented an upwarped peneplain of Miocene age on the basis of warped stream terraces along the Potomac. Susquehanna, and Schuylkill Rivers. Although the existence of a peneplain of lower relief in Miocene time cannot be proved, the higher relief that exists there now is probably the re- sult of upwarping of the area at a higher rate than other parts of the Piedmont farther south. Several arguments support this idea. The rocks in the area of the ‘Vestminster anticline, or uplift, to use a less specific term, are both igneous and sedimentary, similar to rocks of the Virginia Piedmont farther south where relief is lower. They include gabbro, granite, gneiss, diamictite gneiss, quartzose schist, vol- canic rock, phyllite, and marble. The altitude of the highest parts of the Westminster uplift exceeds 305 In. Total relief within 100—km2 areas was measured on 1 :2-1 GOO—scale maps of the following quadrangles: Downington, Pa. (162 m), and Mt. Airy (125 m), ‘Vinfield (125 m), and Rockville (105 m), Md. In low-relief areas shown on figure 5, total relief is generally less than 90 m in areas 10 km on a side. On its west side, most of the uplifted area is bounded by a low escarpment that overlooks Triassic and Jurassic basins to the west. As these basins are underlain mostly by shale and mudstone, the relative- ly low altitude and low relief are perhaps related to rock control. On the east, the Piedmont margin is overlapped by a sequence of sedimentary deposits ranging in age from Early Cretaceous to late Tertiary (Darton, 1951). The surface of basement and base of the Cretaceous has a slope of about 0.02 in Washing- ton, D.C., whereas the slope of the Miocene in Wash- ington is only 0.005. A river gravel deposit resting on the Miocene, believed to be of late Tertiary age, has an even lower slope, averaging about 0.0025 at the same cross section. Although the slope of the surface on which these sediments were laid down is not known, the Lower Cretaceous grave] is similar in size compo- sition to the upper Tertiary gravel. Thus, the differ- ence between the two slopes (approximately 0.0175) may reasonably be taken as a measure of the tilt that has taken place. Assuming an age difference of 100 million years between the two gravels, the rate of tilt would average about 1.75 X 10—4 (mm/km)/yr. If the present slope tangent of the basement were pro- jected 50 km inland from the Fall ZOne to a point B11 near the center of the highest altitude near VVest- nunster, Md., assuming no erosion and no bending, the altitude of this point would be about 1000 m above sea level or 700 m above the present surface. The aver- age rate of lowering at this point since middle Creta- ceous time would have been 7 X 10—3 mm/ yr. As will be shown in the next section, these rates of uplift and erosion are somewhat lower than average rates calcu- lated for other places and other time spans using a variety of methods (tables 2 and 3). It should be noted that the present slope of the base— ment beneath the Coastal Plain in the Washington area, as indicated on the “Tectonic Map of the United States” (U.S. Geol. Survey and Am. Assoc. Petroleum Geologists. 1962), is greater in the northern Piedmont near Washington east of the Westminster uplift than it is farther south. At the Virginia—North Carolina boundary, the slope tangent of the basement near the inner edge of the Coastal Plain is about 0.005 as com- pared with 0.02 in the “'ashington area. In the north- ern Piedmont. the increased relief is also related to the general depression of the Coastal Plain in the Salis- bury embayment, which brings the edge of the Pied- mont to sea level in the larger river valleys like the Potomac. The Pleistocene lowering of sea level. lowered the base level still more and probably affected the re- lief inland, at least close to the larger streams (Hack, 1975, p. 96). The most dramatic difference in relief unaccounted for by geology is along the Blue Ridge escarpment. As shown by figure 3, a steep escarpment in Georgia and North Carolina crosses the boundary between the Pre— cambrian terrane and the generally younger meta- morphic rocks of the Piedmont. The same escarpment also crosses the Cambrian and Ordovician belt, and its crest forms the continental divide between the Gulf of Mexico and Atlantic drainages. The relief and alti- tudes of the terrain on either side are strikingly dif— ferent. “’ithin areas 10 km wide, the relief averages less than 90 in in the Piedmont. In areas of comparable size in the Blue Ridge, the relief ranges from about 180 m to more than 1000 m, and altitudes rise to more than 2000 m. Davis (1903) explained the escarpment by the dif- ferences in lengths of the rivers on either side in their descent to the sea. White (1950) thought the escarp- ment was essentially a fault scarp. Hack (1973) thought that the escarpment was due to differences in the geology crossed by the streams on either side. In other words, the westward—flowing streams cross a series of local base levels formed on resistant rocks, thus, in effect, maintaining the high altitude of the continental divide. The faulting hypothesis remains a B12 possibility, though clear evidence for a fault or faults has not been found. Nevertheless, over long distances, the rocks forming the present escarpment are not more resistant than those on the adjacent Piedmont at the toe of the scarp. If no faulting is involved, tilting must have taken place on the southeast side of the Blue Ridge. RATES OF EROSION AND UPLIFT Many estimates have been made of erosion and up- lift rates based on a variety of methods. In addition, contemporary uplift or tilting of the land has been measured in the field by repeated geodetic traverses. Some of these estimates that relate to the Appalachian Highlands are summarized in tables 2 and 3. Table 2 deals with erosion and uplift rates measured at a single locality within an area. These rates are absolute values SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 related to a constant but unspecified data base ex— pressed as erosion in millimeters per year or as ve— locities of uplift in Inillimeters per year. The estimates in table 3, on the other hand, are dif— ferences between rates at points along a linear feature or traverse. In other words, they are relative rates or angular tilt rates rather than uplift or subsidence. They are expressed in millimeters per kilometer per year (compare with Brown and Oliver, 1976, p. 17). The first five items in table 2 are based on amounts of material being carried out of one or more drain- age basins over a period of at least a year. Item 1 in the table is from a general paper by Judson and Ritter (1964), which contains a review of earlier literature as well as newer data. The first three estimates deal with fairly large drainage basins. Items 4 and 5 are small watersheds that may not be typical, though the TABLE 2,—E8timatcs of erosion rates Item no. Process (mill/ta) Sources of data 1 Modern erosion, North Atlantic drainage basins--- 4.8 X 10‘2 Judson and Ritter, 1964. 2 Modern erosion, South Atlantic and Gulf basins-_-- 4.1 X 10‘” Judson and Ritter, 1964. 3 Modern erosion, South Fork of Shenandoah 4.3 X 10‘2 Hack, 1965. River watershed. 4 Modern erosion, small watershed, Coweta, N.C. ---- 9 X 10'“ Berry, 1977. 5 Modern erosion, small watershed, Maryland 4.6 X 10‘3 Cleaves and others, 19701 Piedmont. 6 Cenozoic erosion rate based on sediment volumes 2.7 X 10‘2 Mathews, 1975. on Continental Shelf (Atlantic drainage). 7 Phanerozoic erosion rate for US. based on sedi- 1 x 10'2 Gilluly and others, 1970. mentary volumes. (Figure given is minimum; estimated by authors as up to six times too low). 8 Uplift and erosion (some possibly tectonic) based 4.1 X 10'“ Whitney and others, 1976 1 on depth of intrusion of Stone Mountain Granite, Georgia. 9 Uplift and erosion (some possibly tectonic) based 5.3 X 10‘2 This paper, data from Colton, 1970. on thickness of Paleozoic sequence in eastern Pennsylvania. 10 Removal of overburden from Middle Devonian rocks and conodont color in Appalachian basin—— rate calculated since Permian. 4.8 to 8.8 X 10‘2 This paper, data from Epstein and others. 1977. 1The rates reported here were calculated by Hack from data in references cited. TABLE 3.—E3timates of tilt rates 132:“ Tilted feature [ (mflfifi‘; /yl‘] Source of estimate 1 Modern tilt rate in Potomac basin if balanced by 3.5 x 10" This paper. average erosion rate of 4 X 10‘2 mm/ y. 2 Tilt rate of basement rocks from the Cretaceous 1.75 x 10“ This paper. to late Miocene, Washington, DC. 3 Domal uplift of Canadian Shield since 7000 B.P. 2.2 X 10.2 Walcott, 1972, fig. 16.1 (average tilt of southeast part of dome). 4 Downwarp to south of beach of high stand of 9.3 X 10'2 Denny, 1974, p. 42.1 Champlain Sea since 10 000 BR 5 Upwarp to south of level line west of Lake 3 X 10‘2 Isachsen, 19’75.1 Champlain, Rouses Point to Ticonderoga after 18 years. _ 6 Downwarp of rerun level line 100 km long between 5.5 x 10" Brown and Oliver, 1976, fig. 11. Atlanta and Columbus, Ga. l'l‘he tilt rates reported here were calculated by Hack from data in the references cited. ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS methods of analysis used in these two cases are rigor- ous. Items 6 and 7 are estimates based on the volume of sedimentary sequences that have been redeposited after erosion from the land. These estimates are the same order of magnitude as rates in items 1—3. Item 8 (Whitney and others, 1976) is an entirely different kind of analysis based on depth of emplacement of a granite body in metamorphic rocks. By using isotope ratios, the age and the temperature and, hence, by in- ference, the depth of emplacement can be estimated. The rate of removal of overburden can, of course, be determined from these data, though whether all the removal was due to subaerial erosion is not certain. It is interesting that item 8 is comparable in magni- tude with items 1, 2, 3, and 6. Items 9 and 10 are similar in concept to item 8 in that uplift rates are estimated by inferences about age and depth in the Earth before the time of uplift. It is assumed that because the rocks are now at the surface, the average uplift rate and erosion rate have been the same. Item 10 is taken from a study of conodonts (Epstein and others, 1977), which shows that conodont colors in the Appalachian basin con- sistently indicate depth of burial. These data give re- sults similar in magnitude to the others. The two modern erosion rates from small watersheds (items 4 and 5, table 2) are significantly different, but it should be noted that the rates in these two places are lower rates rather than higher. The results obtained from direct measurements of tilt cannot be compared with the uplift velocities or erosion rates given in table 2 without making some sort of conversion that involves assumptions about the distribution of the uplift or erosion rates within an area. To do so, the absolute rates must be converted to tilt rates. Item 1 in table 3 is such a conversion. It is based on the assumption that the Potomac Basin is now being tilted upward at its western headwaters near the continental divide relative to the mouth at sea level and that uplift is balanced by an erosion rate of 0.04 mm/y (from table 2). As the erosion rate varies throughout the basin, the average rate is as- sumed to be centered at a point about halfway (112 km) from the river mouth to the continental divide. On this basis, the- tilt-rate estimate in the basin is 3.5 X 10—4, expressed as millimeters per kilometer per year. This rate can be compared with tilt rates meas- ured along level lines (items 5 and 6, table 3) or along water planes whose age is known (items 3 and 4, table 3). The data show that the measured modern and post- Pleistocene tilt rates are at least an order of magnitude higher than rate number 1, on the basis of the erosion rates estimated in table 2. The same conclusion is B13 reached by comparing erosion estimates with absolute uplift rates of domal features that have been meas- ured. One example is the Canadian Shield, which near its center at Hudson Bay is estimated by VValcott (1972) to be rising at a rate of 20 mm per year, a rate much higher than any of the absolute rates of table 1. In this example, the uplift area has a radius of about 1000 km, and this rapid rate is, of course, compatible with the calculated tilt rate for this area (item 3, table 3). Another example of a high absolute rate is the Adirondack dome, studied by Isachsen (1976). On the basis of differences between two geodetic traverses, assuming 0 base at the south edge of the rising area at Utica, 240 km from the center, the Adirondacks are rising at a rate of 3.7 mm/y. This translates to a tilt rate of 2.4 X 10—2 mm/ yr, approximately the tilt rate of the southern margin of the Canadian Shield. The northern flank of the Adirondack domal uplift also shows on a geodetic traverse along the shore of Lake Champlain (item 5, table 3), but the direction of tilt is the reverse of the postglacial tilt of the water planes parallel to it (item 4, table 3). Therefore, if the tilt is real, it has only recently reversed its direction. A brief consideration of the rate of change of land- forms in postglacial time is useful. Glacial deposits of VVisconsinan age in the Central United States, as much as 55 000 yr old (Flint, 1971, p. 559), have a con- structional topography that has undergone little change in form since deposition. On the other hand, glaciated terrain of Illinoian age, more than 100 000 yr old, has been considerably modified, though the deposits themselves are mostly preserved. Nebraskan glacial deposits, which are more than 400 000 yr old, have been eroded extensively, and none of the original topographic form survives. These general conditions are consistent with an average erosion rate of 0.04 mm/ yr or 4 m/ 100 000 yr. The data indicate that erosion rates in the Eastern United States, both modern and Cenozoic in age, are, on the average, low. For large areas, they average as much as about 4 x 10—2 mm/yr. Judging by two anomalously low rates estimated by more detailed analysis of very small watersheds, the true rates may be even lower. All these rates are much lower than uplift rates at the center of large areas like the Canadian Shield, where the rate at the center is 20 mm/yr. However, uplift rates depend on the size of the area being deformed, so it is more reasonable to compare tilt rates than absolute uplift rates. If ero- sion rates are converted to or are considered the inverse of tilt rates, then a large discrepancy remains, amount- B14 ing to at least one order of magnitude. It thus appears that the rates of deformation now observed cannot have been sustained for long periods. The rates of re- bound after the removal of the ice are consistent, however, with measured tilt rates. Rebound can take place rapidly, with only a short lag time between re- moval of a load and compensation by uplift. Much of the data contained in repeated geodetic traverses, how- ever, is not readily explained. For example, if the tilt downward south of Atlanta, Ga., involving a 100-km traverse (item 5, table 3), were sustained for 100000 yr, it would result in a depression 550 m deep. No de- pression exists at this locality, and the topography is a normal erosional topography resembling its sur- roundings. DEVELOPMENT OF THE APPALACHIAN HIGHLANDS IN LATE GEOLOGIC TIME The present topography of the Appalachian High- lands is closely adjusted to the rocks. However, the available evidence shows that the highlands have un-‘ dergone warping and minor faulting in late geologic time. No positive evidence has yet been found that the bedrock at the surface has been broken by differential movements of any great magnitude since at least Cre- taceous time. Though uplift took place, it must have involved broad warping of the crust. Data from the Coastal Plain and Continental Shelf indicate that though the tilting of this area has probably been con- tinuous, it was particularly rapid in early Cretaceous and Miocene to Pliocene time. Assuming these con- ditions, the present topography could have been de- rived by normal erosional processes. The outlines of the present landscape began to form some time in the middle Mesozoic after the breakup of the continent and the establishment of a drainage system toward the opening Atlantic. The first coarse elastic sediments deposited on the Coastal Plain and Continental Shelf are Early Cretaceous in age. The outlines of the hinterlands from which this material came are not known except by inference. The material deposited on the Coastal Plain contains elements from both the Piedmont and the Coastal Plain (Glaser, 1969). Because the coarser gravel contains a high pro- portion of quartzite and finer gravel contains a sig- nificant amount, at least some of the material probably came from the Appalachians west of the Blue Ridge. On the other hand, the heavy minerals and the paucity of chert indicate that, in general, the greatest propor- tion of material probably came from the Piedmont, perhaps indicating that the relief on the Piedmont was high at the time. SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 88° 84° 80° 76° 72° 68° 48° ' _ L. i 44° 44° - 1 40° 40° O 36 36° - O 32 32° 0 490 800 MILES I l 0 400 800 KILOMETERS EXPLANATION n E § I s —120 to —so -so to -40 ~40 to o 0 to +50 Bouguer anomalies in milligals Subenvelope contour at 1000 feet (305 m) _l_l_l_|_.| Inner edge of Coastal Plain FIGURE 6,—Map of Eastern United States showing Bouguer gravity anomalies. Data from Defense Mapping Agency, [Aerospace Center, St. Louis, Mo. Geomorphic evidence concerning the limits of the initial drainage system and the position of the divide has been controversial, and much has been written on the subject (Thornbury, 1965, p. 82). Judson (1975) noted that the continental drainage divide corresponds to a gravity low in the central and southern Appala- chians (fig. 6). The meaning of the gravity low is not understood, but the low may relate either to a thick crust or to the thick sequence of low-density rocks in the eastern part of the Appalachian basin. In any case, the correspondence of the low to the drainage divide is consistent with the idea that the divide has had a stable position for a long time. The correspond- ence of the divide to the low is lost north of lat 40°. In a general way, low anomalies do tend to correspond ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS to high topographic areas, though, in detail, this correspondence breaks down in the Adirondacks and New England (fig. 6). During the Cretaceous, base level of erosion was established at about its present position or higher; it could have varied little more than 300 m in altitude since then. During the Early Cretaceous, when large amounts of coarse sediments were poured onto the Con— tinental Shelf, derived largely from crystalline ter- rane, base level was lower than it is at present. By Late Cretaceous time. the rate of sedimentation had slowed, and base level rose so that the sea probably covered some of what is now the Piedmont. In Oligo- cene time, at least in the north, sea level again was lowered. In Miocene time, base level was generally lowered, though in some areas the sea was higher than now. The positions of base level were never parallel to those at present, as certain areas like the Cape Fear and South Jersey uplifts tended to be higher than others (Owens, 1970), and there has been a general down tilt of the shelf to the northeast. It is assumed that continuing uplift as well as erosion inland were responsible for the high rate of sedimentation on the Continental Shelf. As the abso- lute rate of uplift relative to base level must have in- creased inland, the uplift was in the form of a broad arch, its axis centered somewhere near the present continental divide. Thus, topographic relief that formed as a result of uplift would have been higher inland near the crest of the arch, especially in areas of resistant rocks. lVest of the Blue Ridge, where streams were eroding sedimentary rocks arranged in distinct and contrasting layers, the drainage system became adjusted to differences in the bedrock by multiple piracies, and local differences in relief were marked. This condition was especially true along the Cambian and Ordovician belt in which solution was a factor in the erosion mechanism. In the crystalline rock areas of the Blue Ridge and Piedmont, some rocks, such as the Ocoee Supergroup, felsitic volcanic rocks, or the quartz kyanite rocks of the Kings Moun- tain belt, had great resistance to erosion. As Flint (1963) showed in southern Connecticut, rocks having high quartz content tend to form high relief. Massive, coarsely crystalline rocks tend to be more resistant than sheared or foliated rocks. When erosion was rapid, differences in relief related to rock type would have been large. Under conditions of slower erosionin the crystalline—rock areas, because of the polymineralic nature of most of the rocks, a residue of saprolite would collect at the surface. When the residual mantle became generally thick, differences in relief were largely eliminated over broad areas. B15 At present in the Piedmont and Blue Ridge, sapro- lite underlies most areas that have low relief. It may be more than 90 m thick though it averages about 18 m (Hack, 1976). In areas of high relief, it is thin, lack- ing, or is found mostly in valley bottoms. The shape of the domal uplift that produced the highlands is a significant problem. The steep gravity gradient that parallels the Blue Ridge front is sug— gestive of different crustal blocks that may separate areas having two rates of uplift. On the other hand, as shown by figure 5, the gradient loses its character in New England north of lat 40°, yet great contrasts in topographic relief are found in New England. An analysis by Best and others (1973) based on modeling of various assumed crustal conditions indicates that in North Carolina, at about lat 35°45’, the steep gradient is related to high—density rocks in the Carolina slate belt at high levels in the crust. Although the argument for a tectonic break near the east foot of the Blue Ridge is strong, the present topography could have resulted from a warping over the axis of the Appalachian Highlands. According to this hypothesis, the Blue Ridge escarpment south of Roanoke was tilted eastward (on the southeast side of the upwarp), and the escarpment reached its pres- ent position by retreat. The gentle gradient and presumed rate of tilt of the basement just east of the Fall Zone may provide a clue. At the North Carolina—Virginia boundary, the gradient of the basement is only about 0.005. If pro- jected inland as far as the Blue Ridge, the top of the basement would reach an altitude of only about 1200 in, almost equal to the crest of the present mountain range. \Ve have reason to believe, however, from esti- mated erosion and uplift rates, that the middle Creta- ceous crest of the range would have been at about 4000 111 relative to present sea level. Therefore, if faulting is ruled out, the warping of the Piedmont must have taken place at a higher rate inland toward the Blue Ridge than at the Fall Zone. Farther north in the lVashington area where the Piedmont narrows, the warping of the basement at the Fall Zone is steeper and may account for much of the uplift. CONCLUSION Analysis of the landscape, the major concern of this study, indicates that the topography is quite varied in altitude and relief, but this variation is in general related to differential erosion of the bedrock rather than to differential uplift. Even the Adirondack Moun— tains, which have the form of an uplift, are girdled by a wide belt of nonresistant rock. Much of the Ap- B16 palachian Highlands region, however, can be inter- preted as an upwarp of large proportions, but the en- tire highland cannot be explained that simply. For example, differential uplift, or warping, must be in— voked to explain the sharp break between the Blue Ridge and the Piedmont. This is attested to by the gentle slope of the Piedmont at its margin along the Fall Zone and by the lack of large-scale faulting along that zone. The Blue Ridge escarpment, at least in the south, cannot be explained by differential erosion alone. Thus, the up-arching of the Appalachian High- lands must have been differential, and it could have been accompanied by small-scale movements along fault zones parallel to the structural grain. Data from a variety of sources indicate that the Earth’s crust probably responds quickly to loading or unloading and that erosion rates and uplift rates are probably mutually dependent. The effect of a high rate of uplift is to increase erosion rates gradually, eventually producing topography of greater relief. The converse would also be true, that decreased ero- sion rates would be correlated with lower relief. Local variations in the relief, even of great magnitude, are related to rock control and, in places, to differences in uplift rates as well, thus obscuring the outlines of tectonic blocks that may exist. Data obtained from repeated geodetic traverses in- dicate rapid rates of tilt in many areas. These rapid rates are matched by similar rates along ancient water planes in deglaciated regions. Clearly, rapid changes in elevation or rates of tilt can and do occur locally. Nevertheless, many of those observed cannot have been sustained for long periods of time. The erosion rates that have been estimated are an order of magnitude or more lower, and, if the high rates of tilt or uplift were long sustained, they would result in much larger topographic features than are observed. REFERENCES CITED Berry, J. L., 1977, Chemical weathering and geomorphological processes at Coweeta, North Carolina [abs]: Geol. Soc. America Abs. with Programs, v. 9, no. 2, p. 120. Best, D. M., Geddes, W. H., and Watkins, J. S., 1973, Gravity investigation of the depth of source of the Piedmont gravity gradient in Davidson County, North Carolina: Geol. Soc. America Bull., v. 84, no. 4, 1213-1216. Brown, L. D., and Oliver, J. E., 1976, Vertical crustal move- ments from levelling data and their relation to geologic structure in the eastern United States: Rev. Geophysics and Space Physics, v. 14, no. 1, p. 13—35. Campbell, M. R., 1929, Late geologic deformation of the Ap- palachian Piedmont as determined by river gravels: Natl. Acad. Sci. Proc., v. 15, p. 156—161. SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 1933, Chambersburg (Harrisburg) peneplain in the Piedmont of Maryland and Pennsylvania: Geol. Soc. America Bull., v. 44, no. 3, p. 553—573. Clark, W. B., 1891, Correlation papers*Eocene: U.S. Geol. Survey Bull. 83, 173 p. . Cleaves, E. T., Godfrey, A. E., and Bricker, O. P., 1970, Geo— chemical balance of a small watershed and its geomorphic implications: Geol. Soc. America Bull., v. 81, no. 10, p. 3015—3032. Colton, G. \V., 1970, The Appalachian Basin—its depositional sequences and their geologic relationships, in Fisher, G. W., and others, eds., Studies of Appalachian geology— central and southern: New York, Interscience Publishers, p. 5—18. Craig, A. J., 1969, Vegetational history of the Shenandoah Valley, Virginia, in Schumm, S. A., and Bradley, W. C., eds. United States contributions to Quaternary research: Geol. Soc. America Spec. Paper 123, p. 283—296. Darton, N. H., 1951, Structural relations of Cretaceous and Tertiary formations in part of Maryland and Virginia: Geol. Soc. America Bull., v. 62, no. 7, p. 745—780. Davis, W. M., 1903, The stream contest along the Blue Ridge: Geog. Soc. Philadelphia Bull., v. 3, p. 2134244. Denny, C. S., 1974, Pleistocene geology of the northeast Adirondack region, New York: US. Geol. Survey Prof. Paper 786, 50 p. Eardley, A. J., 1964, Polar rise and equatorial fall of sea level since the Cretaceous: Jour. Geol. Education, v. 12, no. 1, p. 1—11; Am. Scientist, v. 52, p. 488—497. Epstein, A. G., Epstein, J. B., and Harris, L. D., 1977, Conodont color alteration—an index, to- organic metamorphism: U.S. Geol. Survey Prof. Paper 995 27 p. Flint, R. F., 1963, Altitude, lithology, and the Fall Zone in Connecticut: Jour. Geology, v. 71, no. 6, p. 683—697. 1971, Glacial and Quaternary geology: New York, John Wiley & Sons, 892 p. Gilluly, James, Reed, J. 0., Jr., and Cady, W. M., 1970, Sedi- mentary volumes and their significance: Geol. Soc. America Bull., v. 81, no. 2, p. 3534376. Glaser. J. D.. 1969, Petrology and origin of Potomac and Magothy (Cretaceous) sediments, Middle Atlantic Coastal Plain: Maryland Geol. Survey Rept. Inv. 11, 102 p. Hack, J. T., 1965, Geomorphology of the Shenandoah Valley, iVirginia and West Virginia, and origin of. the residual ore deposits: U.S. Geol. Survey Prof. Paper 484, 84 p. 1973, Drainage adjustment in the Appalachians, in Morisawa, Marie, ed., Fluvial geomorphology: Bingham- ton, N. Y., State Univ. New York, Pubs. Geomorphology, p. 51—69. 1975, Dynamic equilibrium and landscape evolution, in Melhorn, W. N., and Flemal, R. C.. eds., Theories of land- form development: Binghamton, N. Y., State Univ. New York, Pubs. Geomorphology, p. 87—102. 1976, Geomorphic framework of the Piedmont and Blue Ridge [abs]: Geol. Soc. America Abs. with Programs, v. 8, no. 6, p. 897—898. Isachsen, Y. W., 1975, Possible evidence for contemporary doming of the Adirondack Mountains, New York, and sug— gested implications for regional tectonics and seismicity: Tectonophysics, v. 29, p. 169481. 1976, Contemporary doming of the Adirondack, Moun- tains, New York: EOS (Am. Geophys. Union, Trans), v. 57, p. 325. ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS Judson, Sheldon, 1975, Evolution of Appalachian topography, in Melhorn, W. N., and Flemal, R. C.; eds, Theories of landform development: Binghamton, N. Y., State Univ. New York, Pubs. Geomorphology, p. 29~44. Judson, Sheldon, and Ritter, D. F.. 1964, Rates of regional denudation in the United States: Jour. Geophys. Research, v. 69, no. 16, p. 33953—102. King, P. B., and Beikman, H. M., compilers, 1974, Geologic map of the United States (exclusive of Alaska and Hawaii): Reston, Va., U.S. Geol. Survey, 3 sheets, 1:2 500 000. Mathews, W. H., 1975, Cenozoic erosion and erosion surfaces of eastern North America: Am. Jour. Sci., v. 275, no. 7, p. 818—824. Mixon, R. B., and Newell, W. L., 1977, Staiford fault system— 'Structures documenting Cretaceous and Tertiary deforma- tion along the Fall Line in northeastern Virginia: Geology, :v. 5, no. 7, p. 437—440. Owens, J. P., 1970, Post-Triassic tectonic movements in the central and southern Appalachians as recorded by sedi- ments of the Atlantic Coastal Plain, m Fisher, G. W., and others, eds, Studies of Appalachian geology—central rand southern: New York. Interscience Publishers, p. 417— 427. Pierce. K. L., 1965, Geoinorphic significance of a Cretaceous deposit in the Great Valley of southern Pennsylvania: U.S. Geol. Survey Prof. Paper 525~C, p. 0152—0156. B17 Scholle, P. A., ed., 1977, Geological studies on the COST No. /B—2 well, US. Mid-Atlantic Outer Continental Shelf area: U.S. Geol. Survey Circ. 750, 71 p. Stearns, R. G., 1967, Warping of the western Highland Rim peneplain in Tennessee by ground-water sapping: Geol. Soc. America Bull., v. 78, no. 9, p. 1111—1124. Thornbury, W. D., 1965, Regional geomorphology of the United States: New York, John Wiley & Sons, 609 p. U.S. Geological Survey and American Association of Petroleum Geologists, 1962,. Tectonic map of the United States, ex- (clusive of Alaska and Hawaii: Washington, D.C., 2 sheets, scale 1:2 500 000. Walcott, R. I., 1972, Late Quaternary vertical movements in eastern North America—quantitative evidence of glacial- isostatic rebound: Rev. Geophysics and Space Physics, v. 10, no. 4, p. 849—884. White, W. A., 1950, Blue Ridge Front—a fault scarp: Geol. Soc. America Bull., v. 61, no. 12, pt. 1, p. 13094346 Whitney, J. A., Jones, L. M., and W'alker, R. L., 1976, Age and origin of the Stone Mountain Granite, Lithonia dis- trict, Georgia: Geol. Soc. America Bull., V. 87, no. 7, p. 1067—1077. York, J. E., and Oliver, J. E., 1976, Cretaceous and Cenozoic faulting in eastern North America: Geol. Soc. America Bull., v. 87, no. 8, p. 1105—1114. The Upper Ordovician and Silurian Hanson Creek Formation of Central Nevada By REUBEN J. ROSS, JR., THOMAS B. NOLAN, and ANITA G. HARRIS SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—C Detailed geologic mapping and collection of fossils has led to revision of lithostratigraphy and dating of the formation in the Belleoue Peak Quadrangle FIGURE 1. TABLE 2. H CONTENTS Page Abstract ________________________________________________________________ C1 Introduction ______________________________________________________________ 1 The Hanson Creek Formation near Eureka, Nevada _________________________ 1 Distribution of fossils and their ages ___________________________________ 5 Age of the Hanson Creek Formation _______________________________________ 8 Evidence of the corals ________________________________________________ 8 Evidence of the bryozoans _____________________________________________ 8 Evidence of the brachiopods ___________________________________________ 8 Evidence of the trilobites _______ . ______________________________________ 8 Combined evidence of the megafossils __________________________________ 10 Evidence of the microfauna: conodonts _________________________________ 11 Conclusions ______________________________________________________________ 15 References cited __________________________________________________________ 15 Map showing outcrop areas of the Hanson Creek Formation and numbered fossil localities in the Belle- vue Peak Quadrangle and elsewhere in central Nevada __________________________________________ Preliminary geologic map and sections of the Wood Cone Peak area, Bellevue Peak Quadrangle, Eureka County, Nev _________________________________________________________________________________ Lithostratigraphic section of the Hanson Creek Formation, measured on the West side of the Mountain Boy Range, Nev ______________________________________________________________________________ Trilobites from the Hanson Creek Formation _______________________________________________________ Trilobites from the Kimmswick Limestone, in Missouri, and the Hanson Creek Formation ________________ Late Ordovician conodonts from the lower dark dolomite member and the middle limestone and dolomite member of the Hanson Creek Formation, Mountain Boy Range, and from the g'r‘aptoilite-bearing platy limestone facies of the Hanson Creek Formation, Monitor Range, Eureka County, Nev ______ Latest Ordovician and (or) earliest Silurian conodonts from the middle limestone and dolomite member of the Hanson Creek Formation, Mountain Boy Range, Eureka County, Nev ______________________ TABLES Trilobites of the Hanson Creek Formation in the Bellevue Peak Quadrangle, Nevada __________________ Distribution of conodonts in collections from the Hanson Creek Formation, Mountain Boy Range, and Wood Cone Peak area, Bellevue Peak 15-minute Quadrangle, Eureka County, Nev _________________ IH Page C3 13 15 17 20 Page 09 19 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION OF CENTRAL NEVADA By REUBEN J. Ross, JR., THOMAS B. NOLAN, and ANITA G. HARRIS ABSTRACT The Hanson Creek Formation southwest of Eureka, Nev., in the Bellevue Peak Quadrangle is composed of three litho- stratigraphic members: (1) a basal dark-gray dolomite, (2) a middle silty thin- to thick-bedded, locally nodular, dark- gray, light-yellow-mottled limestone topped by light-gray dolomite, and (3) an upper dark-gray dolomite, which is herein named the Combs Canyon Dolomite Member. Detailed geologic mapping and accompanying fossil collecting prove that the same lithostratigraphic and biostratigraphic sequence is present in the Mountain Boy Range and 11 km to the south near Wood Cone Peak. Minor differences in lithology and fossil preservation are attributed to intrusion of quartz monzonite and complex structural disruption near Wood Cone Peak. Bighornia sp., Lobocorallium major, Sceptropora facula, Thaerodonta sp., chidocyclus capax, Ceraztrinus icarus, and a new species of Hypodicranotus favor a Maysvillian, pre- ferably a Richmondian, age for the middle member. The pres- ence of the conodonts Amorphognathus ordom'cicus and Pro— topanderodus insculptus supports this preference, as does the occurrence near the top of the member of latest Ordovician and (or) earliest Silurian conodonts, including distomodids, Exochognathus keislognathoides s.f., and Aphelognathus? n. Sp. Conodonts also indicate that the lower member is not older than Edenian. INTRODUCTION Mapping of the Bellevue Peak Quadrangle by Nolan has included the outcroppings from which Hague (1892, 1883 atlas) originally reported Ordo- vician fossils in carbonate rocks above the Eureka Quartzite southwest of the Eureka mining district, Nevada. The combined mapping and paleontological effort has shown these historically important out- croppings to be far more complex structurally than might be supposed from Hague’s (1883) reconnais- sance maps, has demonstrated a stratigraphic sequence that crosses the Ordovician and Silurian boundary within a single formation, and has re- vealed the association of Late Ordovician corals, brachiopods, trilobites, and conodonts that have at times been considered to be of differing ages. THE HANSON CREEK FORMATION NEAR EUREKA, NEVADA The Hanson Creek Formation in the central basin ranges is composed of limestone and dolomite, the lat- eral equivalents of which are widely distributed in Utah and Nevada and have been given other forma- tion names—Fish Haven Dolomite to the northeast and Ely Springs Formation to the east and south. The fauna that characterizes these beds was first found by Arnold Hague, assisted by C. D. Walcott, in the course of their survey of the Eureka mining district in Nevada (Hague, 1892, p. 57—59). They assigned the fauna to the lower beds of their Lone Mountain Limestone and considered it to be of Trentonian, or Early Silurian, age. (At the time of their work, the Ordovician System had not been generally recognized.) Their fossil collections were made near Wood Cone Peak, about 19 km southwest of Eureka, and half of the fossil collections discussed in this report were taken from the same localities. The fauna was not found at the type locality of the Lone Mountain Dolomite on Lone Mountain. The fossils listed by Hague (1892, p. 59) were un- fortunately believed to represent a “Trenton” fauna on the basis of misconceptions dating back to the work of James Hall. “Orthis subquadrata” and ”01”- this plicatella,” for instances, were stated by Hall (1847, p. 122, 126) to be common in the “Trenton Limestone” at Maysville, Ky., and at Cincinnati and Oxford, Ohio. Today, these two brachiopods would be recognized as Plaesiomys subquadrata (Hall) of Richmondian age and Plectm'this plicatella (Hall) of Cl CZ SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 Maysvillian age. When Hague’s faunal list is up— dated to reflect modern taxonomy, it agrees with a Maysvillian—Richmondian correlation as much as or more than with a Trentonian—Edenian correlation. The fossils listed by Hague (1892, p. 59), with modern designations, are as follows: Leptaena sericea :Sowerbyella or Thaerodonta, several possible species :Plaesiomys or Pionorthis, species not determined :Plectorthis plicatella (Hall) (we have not identified this species in our Hanson Creek collections to date) :C7‘yptolithoides sp. :Isotelus; might be Anataphrus Orthis subquadrata Orthis (like 0. plicatella) Trinucleus concentricus Asaphus platycephalus Streptelasma :Bighornia, Lobocorallium, or one of several other genera Rhynchonclla :Lepidocyclus, possibly Rhynchotrema Ceraurus :Ceraurus or Ceraurinus Dalmanites :Calyptaulax Illaenus :Illaenus (we have not identified any Illaenid trilobite ‘in our collections from the Hanson Creek) Mapping in the Bellevue Peak Quadrangle has dis- closed additional occurrences of fossiliferous beds in the Hanson Creek Formation, particularly in the Mountain Boy Range about 4 km due east of Dry Lake Well (fig. 1). As originally defined by Merriam (1940, p. 10—11, table 1), the Hanson Creek Formation was that part of the Lone Mountain Limestone of Hague (1892) that overlay the Eureka Quartzite and underlay a conspicuous chert interval at the base of the Silurian Roberts Mountains Formation. Merriam (1940) con- sidered the Hanson Creek Formation at the type section above Pete Hanson Creek in the Roberts Mountains (fig. 1) to be Middle and Late Ordovician in age; he measured a total thickness of 545 ft and described five units, in descending order, as follows: Fact Limestone, massive, dark-gray, fine-grained 1 __________ 180 Limestone, poorly stratified2 ________________________ 140 Limestone, slabby and shaly, dark-bluish—gray; weathers very light gray; very fossiliferous2 _______________ 140 Limestone, poorly bedded, noncrinoidal; small black chert nodules common3 ___________________________ 45 Limestone, dolomitic, dark-gray; flecked with crinoidal columnals3 ______________________________________ 40 _________________________________________ 545 1 Probably Combs Canyon Dolomite Member of this report. 3 Probably middle limestone member of this report. 3 Probably lower dark dolomite member of this report. Merriam (1940, p. 19, fig. 3) extended the name Hanson Creek southward to Lone Mountain, al— though the lithology was considerably different; at Lone Mountain the entire interval between the Eureka Quartzite and the basal chert of the Roberts Mountains is dolomite. The fossiliferous limestone of the middle of the type section was not recognized. In 1942 Merriam and Anderson (p. 1686) included as the Hanson Creek still a third lithology—a lami- nated, thin-bedded graptolite-bearing limestone—on the basis of exposures on Martin Ridge, in the Moni- tor Range (fig. 1). This graptolitiferous lithology is present as far south as Dobbin Summit and is also present near the crest of the Antelope Range not far from the head of Ninemile Canyon. Nolan, Merriam, and Williams (1956, p. 32—34) followed the practice of Merriam and Anderson (1942) but also noted localities in the Eureka dis- trict where the formation is represented by dark dolomite. The Hanson Creek Formation was modified in 1972 by Mullens and Poole, who recognized a sandy zone within the rocks mapped as Hanson Creek; they reported finding a Silurian fauna above the zone. A Silurian age for an upper member of the Ely Springs and Fish Haven Dolomites has been accepted also. In most if not all of these occurrences, the Silurian fauna is found in an upper dark dolomite. Similarly, the uppermost member of the Hanson Creek For- mation in the Bellevue Peak Quadrangle and in all other outcrop areas shown in figure 2 contains Silurian fossils. Indeed, conodonts collected during this study show that even the upper part of the underlying middle limestone and dolomite unit in the Bellevue Peak Quadrangle is of latest Ordovician and (or) earliest Silurian age. Outcrops of the Hanson Creek Formation are widely distributed in the Bellevue Peak Quadrangle and are in all places tectonically disturbed. They constitute a relatively easily deformed sequence of thinly and moderately thick bedded limestones and dolomites between two resistant and massively bedded sequences. The Eureka Quartzite underlies the formation, in most places with a tectonic rather than a sedimentary contact. Similarly, the overlying Lone Mountain Dolomite of probable Silurian age constitutes an extensive thrust plate that not only rests on different units of the Hanson Creek but also locally has cut through both the Hanson Creek and the Eureka Quartzite and directly overlies Limestones of the Pogonip Group of Early and Middle Ordovician age. THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C3 117°00' 116°00' I TYPE SEQTJQN HANSON CREEK, >~ W" ' '0” 2‘ § \‘ HANSON CREEK FORMATION SECTION USGS COLLECTIONS 8320 00-8330 00,02703 CO, D2704 00,02713 CO, 02714 CO “jLONE MOUNTAIN Ordovician section-g? ‘ BELLEVUE PEAK QUADRANGLE 15' 390m Nuuvw 39° 00‘ Nevada Location 0! figure I 215 3‘0 KILOMETERS 38° : 45' FIGURE 1.—Outcrop areas of the Hanson Creek Formation and numbered fossil localities (dots) in the Bellevue Peak Quadrangle and elsewhere in central Nevada. C4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 As a result there is no single continuous exposure of the Hanson Creek Formation, and the exposed sequences normally exhibit one or more minor thrust faults that may bring different units into juxtaposi- tion. In addition, conglomeratic layers have been recognized in all three of the members and may represent sedimentary breccia flows, erosional breaks, or intraclasts. The two most extensive exposures of the Hanson Creek Formation in the quadrangle are near Wood Cone Peak and in the Mountain Boy Range east of Dry Lake. In the Wood Cone Peak area particularly, the beds are cut by at least three minor thrust faults. The fault planes are folded together with the sedimentary rocks; the thinner bedded limestone units are in many places rather intensely folded (fig. 2). A quartz monzonite intrusion on the south- west side of the area of outcrop has heated and slightly recrystallized the limestone. As a result, the silty limestone at the bottom of the middle member does not weather to form the same rubbly nodular slope that is found in the Mountain Boy Range, although clearly the limestone in both areas was originally identical. Three lithologic units can be recognized in the more extensive exposures. The lowest is a massive dark (nearly black) dolomite; the only recognizable megafossils are tiny (1 mm diameter) white crinoid columnals. At several places in the higher part of the southern Mountain Boy Range appear local in- traclastic conglomerate bands in the dolomite; here too, small masses of black chert are included. In the most nearly complete exposure of the member, north and northeast of Wood Cone Peak summit, the unit lies above a thrust fault that separates the member from the light-gray dolomites and thin-bedded lime- stones that normally are conformably above it. At some places in the Mountain Boy Range, the brec- ciated contact between the Eureka Quartzite and the lower dolomite member exhibits nearly iso- clinal folding. Above the basal dark dolomite, a sequence of thin- and medium-bedded silty limestone and thin-bedded light—gray dolomite is found at all the more exten- sive exposures. In most places limestone directly overlies the basal dolomite, but both north of Wood Cone Peak and in the Mountain Boy Range, one or two meters of the light-gray dolomite occur below B. 1,. - MAPPING 3' , NOT ,‘ OMPLETE , c’\ ,/ . ?« ' ' . .Qd “ ‘\ - X r 6,5 * to “\‘Q I .. I . 2' r 4 C. Vb.“ . 1 fi“+-*,\026;15qup’7’. - , ll ‘ ‘ I [\l\/\, 10100 ZOIOO METERS 1 CONTOUR INTERVAL 200 FT: 1 FT=O.3048 METERS FIGURE 2.—Preliminary geologic map and sections of the Wood Cone Peak area, Bellevue EXPLANATION Stream and piedmont gravel OUATERNARY ~ Quartz monzonite OLIGOCENE l Lone Mountain Dolomite-May include DEVONIANP) some Rabbit Hill Limestone AND SILURIAN Combs Canyon Dolomite Member xxx Limestone and light-gray dolomite-- Hanson Creek SILURIAN AND \\\\\\ Fossiliterous Formation ORDOVICIAN Lower dark dolomite member Eureka Quartzite ORDOVICIAN Pogonip Group CONTACTuDotted where concealed -——— THFtUST FAULT—Dashed where uncertain: dotted where concealed Teeth on upper plate. Arrows show direction of movement 0277.‘ 00 Collection locality, u. 5, Geological Survey Peak Quadrangle, Eureka County, Nevada. Geology by T. B. Nolan, R. J. Ross, Jr., and L. A. Wilson. Base from US. Geological Survey 1162,500 Bellevue Peak (1956). THE UPPER ORDOVICIAN AND SILURIAN the limestone. Locally the limestone is abundantly fossiliferous ; localities where fossil collections have been made are indicated on figures 1 and 2. Higher in the sequence, thin— to medium-bedded light-gray dolomite is interbedded with limestone. In the higher units, the limestone beds contain nodules of black chert that are abundant and distinctive. In the higher limestone beds are also a few local intra- clastic conglomeratic zones. Fragments as larger as 5-cm across were found in an isolated fossiliferous limestone outcrop 1 km southwest of Spanish Moun- tain at the south end of the Mountain Boy Range (fig. 1). Limestone of the middle member is thicker bedded in what is apparently the upper part of the sequence. Most of the more massive limestone, however, is part of a minor thrust plate that rests in places on both the thinner bedded limestones and the light- gray dolomite, as well as on the basal dark dolomite. Fossils have been found in the massive limestone in only one place, about 1.6 km northwest of BM 7201 at the divide north of Wood Cone Peak. At what appears to be the very top of the middle member, both at Wood Cone Peak and in the Moun- tain Boy Range, is a sequence of as much as 30 m of very light gray dense dolomite. The uppermost unit of the more inclusive Hanson Creek dolomite of Mullens and Poole (1972) is a second dark dolomite, which occurs in most if not all areas in which fossiliferous Hanson Creek crops out. In most exposures, this upper unit is characterized by beds in which sections of brachiopods are abundant. In some places, the brachiopods are sufi‘i- ciently well preserved to be recognized as penta- merids. In other places, halysitid corals have been preserved. Intraclastic conglomeratic beds and hori- zons with black chert are present in this member also. North of Wood Cone Peak on the south side of Combs Canyon, the dark dolomite member, which contains pentamerid brachiopods and halysitid corals, apparently conformably overlies the upper- most light-gray dense dolomite of the middle mem- ber. The sandy horizon described by Mullens and Poole (1972) was not recognized here, however. The redefinition of the Hanson Creek Formation by Mullens and Poole (1972) includes this third, or upper, member within the formation. Since it is an easily recognized and mappable unit, we suggest that it be given the name of the Combs Canyon Dolomite Member of the Hanson Creek Formation; its type locality is on the south side of Combs Canyon in the NE cor. sec. 31, T. 18 N., R. 52 E. HANSON CREEK FORMATION, NEVADA C5 (unsurveyed). Here the member overlies the upper very light gray dolomite of the Hanson Creek For- mation and appears to be conformable beneath the typical massive, vuggy, coarsely crystalline Lone Mountain Dolomite. DISTRIBUTION OF FOSSILS AND THEIR AGES The larger collections of fossils have been made from two areas in the Bellevue Peak Quadrangle. One of these areas contains exposures of dark-gray locally thick or thin bedded locally silty beds of the middle limestone member of the Hanson Creek and is north of and within 2 km of Wood Cone Peak (figs. 1, 2). This is the area mentioned by Hague (1892, p. 58—59; 1883, Atlas, Sheet IX). Many of the trilobites reported in this paper come from the same outcrops. Collections made from strata in three small conical hills alined approximately north to south and located east and north of Wood Cone Peak are from the lower part of the middle limestone member of the Hanson Creek Formation. There we have found silicified brachiopods and corals, several genera of which are common to the Maquoketa Shale (Iowa) and the Stony Mountain Formation (Mani- toba; Wyoming) (fig. 3; USGS collns. D2707—CO, D2708—CO, D2741—CO). The second area is on the west side of the Moun- tain Boy Range and is exposed mainly on the north side of a dry wash draining westward into Dry Lake playa (fig. 1) (200 m east of C, sec. 31, T. 19 N., R. 53 E.; UTM grid, zone 11: E. 581,775 In, N. 4,369,500 In) (See fig. 3.) This second area is the site of the most nearly complete section of the Hanson Creek Formation we have found in the Bellevue Peak Quadrangle, although it is poorly exposed at the base; the underlying Eureka Quartzite crops out south of the wash and dips northerly, but its contact with the Hanson Creek is covered by sand and gravel. The light-gray dolomite that overlies the main body of the middle limestone member north of the wash may be above a thrust fault; it contains latest Ordovician and (or) earliest Silurian cono- donts. The dark-gray Combs Canyon Dolomite Mem- ber (Silurian) is exposed about 150 m southeast of the main section (UTM Grid, zone 11: E. 581,880 In, N. 4,369,500 m). Despite its limitations, this section is the basis for the lithostratigraphic sequence found to be usable throughout our mapping (fig. 3). 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(Merriam, 1963b, p. 10—11). The stratigraphic positions of our collections are shown accurately in figure 3 for the Mountain Boy Range section. Because of structural complexity (fig. 2), the stratigraphic position of the collections made 11 km to the south near Wood Cone Peak can- not be precisely determined; nonetheless, we have indicated the position in a general way on figure 3. Merriam and Anderson (1942, p. 1686) assigned the entire Hanson Creek a Late Ordovician age on the basis of graptolites from the Monitor Range locality but not on the basis of shelly fossils from the type section or the Wood Cone Peak area. That assignment was continued by Nolan and others (1956) and by Merriam (1963a, p. 31—33). The cor- relation presented by the latest of these references was based largely on the resemblance of brachiopods, including collections from the Inyo Mountains, to those of the Maquoketa Shale of Iowa, which were compared with Richmondian faunas of Ohio. AGE OF THE HANSON CREEK FORMATION Our collections, made from 1975 to 1977, include not only the expectable megafossils, such as corals, bryozoans, brachiopods, and trilobites, but also a very representative conodont microfauna. The evi- dence that each of these kinds of fossils provides for determining the age of the Hanson Creek For- mation is reviewed in the paragraphs that follow, as well as in figure 3 and tables 1 and 2. EVIDENCE OF THE CORALS The corals of the Hanson Creek Formation (fig. 3) were identified by W. A. Oliver (written commun., March 15, 1976) ; they include a generic assemblage reviewed by Duncan (1956, p. 218—220), who con- cluded that they are of Late Ordovician age. Nelson (1963) also found this assemblage in the Hudson Bay Lowland; the questionably identified Nycto- pom(?) and Saflordophyllumfl) are in the Bad Cache Rapids Group, which Nelson (1963, p. 24—25) correlated with the Red River Group of Manitoba but for which he favored a Richmondian age rather than Edenian or Trentonian. Bighornia is restricted to the younger Churchill River Group, as is Labo- corallium tm‘lobatum major Nelson; according to Nelson, that group is of Richmondian or Gamachian age. Oliver (written commun., Jan. 4, 1978) com- mented on the distinctive Lobocorallium found in USGS colln. D2713-CO: 1. If the Lobocomllium is either of Nelson’s species it is the younger one, i.e. L. trilobatum major. The similarity is close. 2. If Nelson’s trend to increasing distinctness of lobation (Nelson, 1963, p. 34) with time is real, then your speci- mens may be even younger than Nelson’s L. major, as yours is one of the most distinctly trilobate group of specimens that I have seen. Duncan (1957, p. 611) reviewed the stratigraphic range of Bighorm'a, concluding that it was of Rich- mondian age and noting that the genus had never been found in Trentonian (in the old sense) strata. As noted by Nelson (1963, p. 28, 39—43), Bighorm‘a is represented by several species in the Churchill River Group but by none in the underlying Bad Cache Rapids Group. EVIDENCE OF THE BRYOZOANS Five of the six genera of bryozoans were identi- fied by O. L. Karklins (written commun., March 3, 197 6) , who stated that preservation was too poor to permit determination of species and that no conclu- sion as to age within the Ordovician was therefore possible. On the other hand Sceptropom facula Ulrich is readily identified (Ross, 1957, pl. 37, figs. 10, 11). Originally identified from Richmondian strata of Minnesota (Ulrich, 1888, p. 228—229) , this species is present in shale of the: Stony Mountain Formation of the upper part of the Bighorn Group in Wyoming (Ross, 1957, p. 459) and in the: English Head Formation and lower part of the Vaureal Formation of Anticosti Island (Twenhofel, 1927, p. 86—87). It is known from no older formations. EVIDENCE OF THE BRACHIOPODS Most of the brachiopo-ds listed in figure 3 are fragmentary. Several genera are of little use in dif- ferentiating rocks of Middle and Late Ordovician age. Pionorthis is present in the Stony Mountain beds of Manitoba and Wyoming (Ross, 1957i Ma- comber, 1970). Lepidocyclus gigas occurs in the Maquoketa Shale of Iowa as well as in the shale of the Stony Mountain Formation of Wyoming. Thaerodonta is another form found in Maquoketa and younger rocks. Leptellinids like Anoptambom‘tes are found in very late Ordovician strata of eastern Alaska (Ross and Dutro, 1966, p. 10—13) and in the Craighead Limestone of Ayrshire of late Caradocian age (probably early Cincinnatian equivalent) (Wil- liams, 1962, p. 170—172) . EVIDENCE OF THE TRILOBITES The ranges of the 18 Hanson Creek trilobite gen- era are shown in table 1. (See also figs. 4, 5.) The THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C9 TABLE 1.—Tm'lobites of the Hanson Creek Formation in the Belleoue Peak Quadrangle, Nevada [Trilobites identified to family, genus, or species where possible. List of previously reported occurrences may not be all inclusive but is intended to show the known range of each family, genus, or species. Leaders, __, no data] Trilobites and collection no. (USGS) Age range previously assigned Middle to late Caradocian. Selected previously reported occurrences Reference and remarks Whittington, 1952; Tripp, 1965, 1967; Ross and Shaw, 1972; Genus reported in Trenton Lime- stone (New York) ; Prosser Hypodioranotus n. sp ___ D2645a, c, d, e—CO D2707—CO Limestone (Minnesota) ; Viola Ancigin, 1973; Ludvigsen, 1975. D2708—CO Limestone (Oklahoma) ; upper- Middle body of hypostome dif- D2771~CO most Copenhagen Formation fers markedly from those of H. Robergia sp ___________ (Nevada) ; Silliman’s Mount (Batfiin Island) ; Whittaker For- mation (Dist. of MaeKenzie) ; Albany and Stincvhar beds (Ayrshire) ; Tiplilski horizon (central Uralss). Road River Formation (Yukon) ; Llandelian to striatulus (Walcott) and H. missoum’ensis (Foerste). Roy, 1941. Whittington, 1950; Lenz and D2645d—CO Ogygiocaris Shale (Sweden). Ashgillian. Churkin, 1966. Isotelus sp ____________ Chazy Group (New York) ; Lex- Chazyan to Twenhofel, 1927; Ross, 1967; D2645d—CO ington Limestone (Kentucky) ; Richmondian. Shaw, 1968, 1974. D2707—CO Ellis Bay Formation (Anticosti Llandeilian to D270‘8—CO Island). Numerous other oc- Ashgillian. Anataphrus sp ________ D2645a—e—CO cu rrences. Sillman’s Mount (Baffin Island) ; Cape Calhoun Formation Late Mohawkian to Richmond‘ian. Whittington, 1954; Troedsson, 1928; Ross and Shaw, 1972. D270‘7—CO (Greenland); Maquoketa Shale D2708—CO (Iowa) ; uppermost Copenhagen D2741—CO Formation (Nevada). D2771—CO Styginid, possibly Chair of Kildare Limestone Ashgillian ________ Dean, 1974, p. 67; Kielan, 1960, p. Stygina. (Eire) ; Upper Ordovician (P0- 82, pl. 8, figs. 1, 2; Warburg, D2741—CO land) ; Boda Limestone (Swe- 1925, p. 96, pl. 3, fig. 1; Whit- den) ; Killey Bridge (North tington, 1950, p. 547—549. Ireland). Proetid _______________ Numerous localities; widespread Middle Ordovician Owens, 1973. D2707—CO in world. to Mississippian. Otam'on sp ____________ D2645a, c—e—CO Many species in United Kingdom, Estonia, Sweden, Poland, Quebec, Llanvirnian. to Middle Devonian Lamont, 1935; Opik, 1937; Twen- hofel, 1927 (listed as Cyphas- D2707—CO MacKenzie, New York, Okla- (Chazyan through phis) ; Begg, 1939; Weller, 1903; D2741—CO homa, Iowa, New Jersey, Vir- Richmondian). Tripp, 1954, 1967; Ingham, D2771—CO ginia, Anticosti Island. 1970; Warburg, 1925; Whitting- Cryptolithoides sp _____ Viola Limestone (Oklahoma) ; Mid-Caradocian to ton, 1966; Kielan, 1960; Schu- chert and Cooper, 1930‘; Shaw, 1968, 1974; Raymond, 1925; Hu, 1975; Ludvigsen, 1975, 1976a. Whittington, 1941; Ross and Shaw, D2645b—e—CO Copenhagen Formation (Ne- Ashgillian. 1972; Lenz and Churkin, 1966. D2707—CO vada) ; Road River Formation D2741—CO (Yukon). Calymenid (possibly Gravicalymene or Numerous localities in North America and Europe. Llanvirnian to Middle Devonian. Shirley, 1936; Twenhofel, 1927; Dean, 1962; Ross, 1967. Calymene). D2645d, e—CO D270‘7—CO D2708—CO Cerauri’nus cf. 0. Richmond Gp. (Indiana and Ohio) ; Richmondian Ludvigsen, 1977 ; Twenhofel, 1927. wa'rus (Billings). English Head Formation and Maysvillian? D2645e—CO Vaureal Formation (Anticosti Island). Ceraurus sp. 1 _________ Whittaker Fm. (Dist. “Black River— Ludvigsen, 1975, pl. 5, figs. 8, 9. D2645d, e—CO MacKenzie). Trenton”. D2741—CO Cerau’rus, sp. 2 (radial ornament)—n. sp.? D2645a—CO D2741—CO Cemurus sp. 3 (fine ornament). D2645e—CO Acanthoparyphinid, possibly Holia or Youngia. D2645e—CO Not known previously( ?) ________ Not known previously ( ?) ________ Kimmswick Ls. (Missouri and Illinois) ; Lincolnshire and Edin- burgh Formations (Virginia); Esbatattaotine and Whittaker Formations (Dist. of Mac- Kenzie). Arenigian to Ashgillian. Chazyan to Edenian. Llandeilian to Llandoverian. For generic range see Lane, 1971, text fig. 2. Schmidt, 1881; Warburg, 1925; Bradley, 1930; Whirttington and Evitt, 1954; Ludvigsen, 1975, 1976a; Lane, 1971. C10 SHORTER CONTRIBUTIONS T0 STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 TABLE 1-—T7‘i105it98 0f the Hanson Creek Formation in the Bellevue Peak Quadrangle, Nevada—Continued Trilobites and collection no. (USGS) Selected previously reported occurrences Age range previously assigned Reference and remarks Sphaerocoryphe sp ______ Uppermost Copenhagen Fm. D2771—CO (Anticosti Island) ; Chair of Kildare Limestone (Eire) ; Boda Limestone (Sweden); Rhiwlas Limestone (North Wales) ; Chazy Group (New York) ; Ashgill (Cumbria); Balclatchie beds, Drummuck beds (Ayrshire). Ellis Bay Formation to Jupiter Formation (Anticosti Island); Encrinurus sp. (may be 2 sps.) D2645b, d—CO Stile End beds (United King- D2.707—CO dom); Ashgill (Cumbria). D2771—CO Cybeloides sp ___________ Maquoketa Shale (Iowa) ; Edin- D2708—CO burgh Formation and Eflna Limestone (Virginia); Lincoln- shire Limestone (Tennessee) ; Rhiwlas Limestone (Wales). Whitehead Formation, Perce (Quebec); Tyrone Limestone Calyptaulax sp _________ D2645c, e—CO D2707—CO (Kentucky); Galena Dolomite D2‘708—CO and Platteville Limestone (of D2713—CO Mississippi Valley); Middle Ordov1cian of Appalachians (Virginia) ; Copenhagen Forma- . tion (Nevada). Lichid, possibly (Nevada) ; Ellis Bay Formation Chazyan to Ross and Shaw, 1972; Shaw, 1968; Gamachian. Lane, 1971; Warburg, 1925; (Llandeilian to Ingham, 1974; Twenhofel, 1927. Ashgillian). Also in middle limestone mem- ber of Hanson Creek type section. Caradocian? Twenhofel, 1927; Dean, 1963; Ashgillian to Ingham, 1974. Silurian. Chazyan to Slocum, 1913; Cooper 1953; Shaw, Richmondian. 1968; Whittington, 1964. Chazyan to Delo, 1940; Shaw, 1968; Cooper, Richmondian. 1953; Schuchert and Cooper, 1930; Ross and Shaw, 1972. Family range Hemwrges. Ordovician to D2645aeCO Devonian. Odontopleurid, possibly _______________________________ Middle Ordovician _______________________________ Ceratocephalw. Late Devonian. D2645e—C0 D2708—CO silioified trilobites are almost all immature: and there- fore either difficult or impossible to identify as to species. Six taxa can be given only family designation. Late Ordovician trilobites from the Cordilleran and Rocky Mountain regions are poorly known. As a result, many if not most taxonomic comparisons must be made with the Middle Ordovician forms that have been much better monographed. In this case the importance of modern taxonomic assessment of a genus is illustrated by the recent work of Ludvigsen (1977) on Ceraurinus; his work leaves no doubt that C. icarus (Billings) or a very closely related variety is present in the Hanson Creek (fig. 4, s, z) and that its age is Richmondian, possibly as old as Maysvillian. Only Hypodicranotus would seem to contradict the evidence of C. icarus. That genus has previously been considered an “indicator of an age equivalent to that of the Trenton Limestone of New York and the Kimmswick of Missouri. (See table 1.) The nearly equally tripartite subcircular middle body of the hypostome of H. missoum'ensis (Foerste) (fig. 5a) differs markedly from the elongate middle body of the Hanson Creek hypostome (fig. 5h-k). The Tren- ton species H. striatulus (Walcott), as shown by Whittlington (1952) , bears the same middle body as H. missoum’ensis. The species of Hypodicranotus from the Hanson Creek beds clearly differs from these two late Mohawkian (or early Cincinnatian) species and may prove to be an index to upper Cincinnatian strata. No other trilobite‘ taxon provides more definitive evidence as to whether the Hanson Creek is of Rich- mondian or an older age. If future collecting should prove that the sty‘ginid is really Ashgillian Stygina, rather than an older form like Raymondaspis, or that the acanthoparyphinid (fig. 4, aa-ac) Youngz'a, rather than an older form like Holia, little. question could remain about the Richmondian or Gamachian age of the trilobites. COMBINED EVIDENCE OF THE MEGAFOSSILS When the evidence of Bighorm'a and particularly of Lobocorallium major Nelson among the corals, of Sceptropom facula Ulrich among the bryozoans, and of Thaerodonta and Lepidocyclus capax among the brachiopods is joined to that: of the trilobites THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA Cerauri'nus icarus (Billings) and a new species of Hypodicranotus, a Richmondian (possibly a Late Maysvillian) age for the middle limestone member of the Hanson Creek is reasonable and agrees with the late Caradocian to Ashgillian age indicated by the graptolites (Ross and Berry, 1963, p. 65—67). EVIDENCE OF THE MICROFAUNA: CONODONTS The microfauna from the measured section in the Mountain Boy Range consists of conodonts. Nearly 2,300 identifiable elements (table 2) were recovered from the section by using dilute acetic- and formic- acid treatment of limestone and dolomite samples, respectively. Insoluble residues were washed through a 20-mesh sieve nested in a ZOO-mesh sieve; a heavy-liquid (tetrabromethane) separation of the residue retained on the 200-mesh sieve was then made. The heavy-mineral concentrate was picked for conodonts at about X 50 magnification. Amorphognathus ordovicicus Branson and Mehl (fig. 6, a—f) occurs in the basal 5.2 m of the section. This index species of the latest Ordovician North Atlantic conodont zone has a range that begins a little above the base of the Ashgillian (a little above the base of the Maysvillian through the Richmon- dian; see Sweet and Bergstrom, 1976, fig. 4). An- other indicator of slightly younger Maysvillian through Richmond‘ian age, Protopanderodus insculp- tus (Branson and Mehl) s.f., occurs 20.7 m above the base of the section. Other elements of less re— strictive but definitely Late Ordovician age that co- occur with or above A. ordovicicus in the section are Belodina sp. D of Shatzer‘ (1976) (fig. 6,1", 3), Oulodus ulm'chi (Stone and Furnish) (fig. 6,m—p), and Plegagnathus nelsom' Ethington and Furnish s.f. (fig. 6,q). These elements along with many other Ordovician conodonts of less biostratigraphic value (fig. 6) were found within the lower 22.9 m of the measured section. An outcrop of dark-gray thick-bedded limestone 200 m west of the measured section (table 2; USGS colln. 8292—00) also contains elements of Amor- phognathus ordovicicus and therefore lies within the same age range as collections 8320—00 to 8328—00 Iof the measured section. Although we have yet to verify the relationship, this outcrop seems to overlie the middle limestone of the measured section but to underlie the light-gray dolomite. Lithologically, it strongly resembles the limestone near Wood Clone Peak. The uppermost collection (table 2; USGS colln. 8329—00) within the measured part of the section Cll contains conodonts of Late Ordovician (zygognathi- form elements) and Silurian aspect (plectospathodi- form elements). A possible fault separates the measured part of the section from overlying beds. Conodonts from light-gray dense dolomite above this fault (table 2; USGS colln. 9785—SD) but still within the middle limestone and dolomite member of the Hanson Creek are of latest Ordovician and (or) earliest Silurian age (fig. 7). These include spatho- gnathodiform and plectospathodiform elements that are probably part of the same Ozarkodma apparatus (fig. 7, j, k) and Exochognathus keislognathoides Pollock, Rexroad, and Nicoll (fig. 7,1) and disto- modids (fig. 7,h, i) that probably belong with other elements assigned to Aphelognathus? n. sp. (fig. 7a—g). Significantly, A.? n. sp. elements are abuLnd- ant and co-occur with latest Ordovician and earliest Silurian species in the basal beds of the upper part of the Ely Springs Dolomite in the Funeral Moun- tains, Calif. (A. G. Harris and J. F. McAllister, USGS collns., unpub. data) .1 Conodont data from the measured section demon- strate that (1) the lower 22.9 m of the Hanson Creek are of Late Ordovician (Ashgillian) age and no older than early Maysvillian; (2) beds immedi- ately below the possible fault are of latest Ordo- vician and (or) earliest Silurian age; and (3) the light-gray dense dolomite above the fault(?) is of latest Ordovician and (or) earliest Silurian age. Moreover, data about conodonts from this section in conjunction with those about other Hanson Creek conodonts from central Nevada provide new infor- mation on the age of the Hanson Creek. The base of the measured section includes the uppermost beds of the lower dark dolomite member. No complete sec- tion of this member has been found thus far in the 1 Since this report was originally prepared (1977), abundant and very well preserved Conodont faunas have been recovered from the Vauréal and overlying Ellis Bay Formations of Anticosti Island, Québec (Nowlan and Barnes, 1979; McCracken and Barnes, 1979). Representatives of what we have assigned to Aphelognathus? n. sp., distomodids, and Exochagnathus keisloynathoides are among the most abundant elements in the Vauréal and Ellis Bay conodont faunas. McCracken and Barnes (1979) placed these elements in a new genus and species, Gamachignathua enaifer. This species is restricted to the Upper Ordovician Vauréal Formation and the Ordovician part of the overlying Ellis Bay Formation, where it co—occurs with other typical Late .,Ordovician conodonts. According to McCracken and Barnes, however, Gamachignathus co—occurs with Ozarkodina of Silurian aspect in the upper 1 to 2 m of its occurrence. McCracken and Barnes used the upper limit of Gamachignathus and the succeeding prevalence of Ozarkodma without Gamachignathus to mark the Ordovician/ Silurian boundary on Anticosti Island. It would appear from their data that USGS coll. 9785~SD, which contains representatives of both genera (table 2), is equivalent to the interval of overlap of Gamachignathus and typical Silurian Ozarkodina on Anticosti Island and therefore lies im- mediately below the Ordovician/Silurian boundary as determined by McCracken and Barnes. Because no international Ordovician/Silurian boundary stratotype has yet been established, this part of the Hanson Creek Formation is assigned a latest Ordovician and (or) earliest Silurian age. 012 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 THE UPPER ORDOVICIAN AND SILURIAN area of this report, although mapping indicates the member is about 46—61 m thick. The Eureka—Hanson Creek contact, however, is exposed in the Mountain Boy Range at Spanish Mountain, 2.4 km southwest of the measured section (fig. 1). There, a single adenticulate detortiform element of Belodma sp. A of Sweet, Ethington, and Barnes (1971) was recovered from 2 m above the base of the Hanson Creek (fig. 6,g). Elements of this same species, moreover, occur 10 m above the base of the Hanson Creek Formation at its) type locality in the Roberts Mountains (about. 53 km northwest). According to W. C. Sweet (oral commun., 1977), Belodina. sp. A has a known range of middle Edenian to middle Maysvillian in his undescribed conodont collections from the western United States, although he suspects that the range of this species may extend to at least the base of the Edenian (base of the Upper Ordovician). Thus, the base ‘of the Hanson Creek in the Mountain Boy Range and Roberts Mountains appears to be no younger than middle Maysvillian and no older than Edenian. HANSON CREEK FORMATION, NEVADA C13 In the course of our studies, conodonts were also collected from. the lower part of the graptolite— bearing platy limestone facies of the Hanson Creek Formation in the Monitor Range, on Martin Ridge, about 2.5 km south of its north end (Horse Heaven Mountain 15-minute quadrangle) and 35 km south- west of the Mountain Boy Range locality. Collections from 6.1, 16.8, and 24.4 m above the base of the formation all contain Protopanderodus insculptus (Branson and Mehl) (fig. 6,ac) and only a few other conodonts of little biostratigraphic value. Some of the P. insculptus elements have part of a second denticle posterior to the cusp preserved—the diag- nostic characteristic of this middle Maysvillian through Richmondian North American species. In this same vicinity on Martin Ridge (although not the same locality), Ross and Berry (1963, p. 65—67) made several float collections of graptolites from about the lower 9.1 m of the Hanson Creek. At. that time, they believed that a fault separated the Hanson Creek from the underlying Eureka Quartzite. They have since decided that no fault exists and the sec— FIGURE 4.—Trilobites from the Hanson Creek Formation. [a-j. q—s, z, aa—ag, from USGS colln. D26459vCO; k, t—y, ah—aj, from USGS colln. D2645d—CO; l—p, ak, from USGS colln. D2645a—CO; ax, from USGS colln. D2707«CO] Ceraurus sp. 1. a—c, cranidium, dorsal, anterior, and lateral views, X 7, USNM 252215. d—f, cranidium, dorsal, anterior, and lateral views, X 6, USNM 252216. g, h, 1', j, cranidium, dorsal, antero-dorsal, lateral, and anterior views, X 3, USNM 252217. k, hypostome, ven- tral view, X 8, USNM 252218. Cemm‘us sp. 2. l, m, n, cranidium, dorsal, lateral, and anterior views, X5, USNM 252219; note radial ornament around each tubercle. 0, par- tial thorax and pygidium, dorsal view, X 10, USNM 252220; lacks radial ornament of associated crandium. p, hypostome, fragmen- tary, ventral view, X 10, USNM 252221; lacks radial ornament. Ceraurus ? sp. 3. Crandium and pygidium, dorsal views, X 7, USNM 252222, 252223. Ceraum'nus icams (Billings). s, hypostome, ven— tral view, X3, USNM 252224. 2, pygidium, dorsal view, X 3, USNM 252225. Calymenid, possibly Gravicalymene. t, hypo- s’oome, ventral view, X 7, USNM 252226. u, pygidium, dorsal view, X 6, USNM 252227. 1;, cranidium, fragmentary, dorsal view, X 5, USNM 252228. w, x, y, pygidium, posterior, lateral, and dorsal views, X 9. USNM 252229. Acanthoparyphinid, possibly Youngia or Holia. aa, pygidium plus one thoracic segment, dorsal View, X 7, USNM 252230. ab, ac, cranid- ium, dorsal and lateral views, X 4, USNM 252231. w—Ic. l—p. q, T. s, 2. 13—11. act—ac. Odontopleurid, possibly Cemtocephala. ad, ae, partial thorax, probably assignable to this genus, dorsal and lateral views, X 7, USNM 252232. of, ag, cranidium, dorsal and anterior views, X 7, USNM 252233. ad—ag. Encrinums sp. ah, pygidium, fragmentary, dor- sal view, X 5, of a small part of the pleura and axis, USNM 252234. at", aj, cranidium, fragmentary, dorsal and anterior views, X 9, USNM 252235. ax, pygidium, dorsal view, X 8, USNM 252236. ah—aj, ax. ak. Lichid, p05sibly Hemiwrges. Cranidium, dorsal view, X 10, USNM 252237. Anataphrus sp. and Isotelus sp. Because com- plete ontogenetic series are lacking distinction is uncertain. Every specimen bears at least one transitory thoracic segment at the front of the pygidium. al, am, pygidia, dorsal views, X 7, USNM 252238, 252239. cm, (10, ar, as, at, an, pygidia, dorsal views, X 10, USNM 252240, 252241, 252244, 252245, 252246, 252247 re- spectively. up, pygidium, dorsal view, X 6, USNM 252242. aq, av, pygidia, dorsal views, X 5, USNM 252243, 252248. al—au. aw. Calyptaulax sp. Cranidium, immature, dorsal view, X 7, USNM 252249. Cybelm'des sp. Pygidium, fragmentary, dorsal view, X 8, USNM 252250. ay. STRUCTURAL GEOLOGY, 1 9 THE UPPER ORDOVICIAN AND SILURIAN tion is probably continuous. Ross and Berry (1963) concluded that the Hanson Creek graptolites were chiefly late Caradocian forms but had some Ashgil- lian components. This conclusion would be in reason- able accord with the conodont data, which indicate an Ashgillian age for beds 6—24 m above the base of the formation. CONCLUSIONS We find that the Hanson Creek Formation in the Bellevue Peak Quadrangle is of Late Ordovician and Silurian age, confirming the previous conclusion of Mullens and Poole (1972) based on other areas. We suspect that continuous deposition took place in this area from latest Ordovician into Early Silurian time. The combined macro- and micro-faunas indicate that the base of the formation is not older than Edenian and that the lower part of the middle lime- stone and dolomite member is at least as young as Maysvillian and is more likely to be Richmondian. The latest Ordovician and (or) earliest Silurian age of the top of this member, evidenced by conodonts, agrees with such a likelihood. HANSON CREEK FORMATION, NEVADA C15 Any apparent uncertainty or discrepancy in dat- ing of these beds is caused largely by ignorance of both stratigraphic and paleogeographic ranges of many of the taxa. The Ordovician paleoenvironments of Ohio, type area of the Cincinnatian Series, and of Nevada were very different (Ross, 1976, 1977). As a result the two areas have little in common pale- ontologically, even in strata deposited at the same time. REF EREN CES CITED Aldridge, R. J ., 1972, Llandovery conodonts from the Welsh Borderland: London, British Museum (Natural'History) Bulletin, Geology, v. 22, no. 2, p. 127—231. Ancigin, N. 1., 1973, Trilobita, in Ordovician Stratigraphy and fauna of the central Urals: Ministry of Geology SSSR, Ural. Terr. Geol. Uprav., Nedra, Moskva, 166 p., 30 pls. Barnes, C. R., 1974, Ordovician conodont biostratigraphy of the Canadian Arctic, in Aitken, J. D., and Glass, D. J., eds., Symposium on the geology of the Canadian Arctic, Proceedings: Calgary, Geological Association and Canad— ian Society of Petroleum Geology Special Volume, p. 221—240. Barnes, C. R., and Fahraeus, L. E., 1975, Provinces, com- munities, and the proposed nektobenthic habit of Ordovi- cian conodontophorids: Lethaia, v. 8, no. 2, p. 133—149. [References continued on p. CIS] FIGURE 5.—Tribolites from the Kimmswick Limestone, of Missouri, and the Hanson Creek Formation. [a—d, from Kimmswick Limestone, from quarry 1.6 km south of Elsberry, Mo.; lowest bed quarried; USGS colln. 258g—CO. e—g, l—p, Hanson Creek Formation, USGS colln. D2645a-CO. an-ar, Hanson Creek Formation, USGS colln. D2645e—CO. All others, from USGS colln. D2645d—CO] a—d. Hypodicranotus missomiensis (Foerste). All X 3. a, hypostome, ventral view; note almost circular middle body divided almost equally into three parts; USNM 252181. b, c, d, cranidium, lateral, anterior, and dorsal views, USNM 252182. Hypodicranottw n. sp. 6, cranidium, dorsal view, X20, USNM 252183. f, cranidium, smallest specimen recovered, X 20, USNM 252184. 9, m, n, cranidial, dorsal views, respectively, X 10, x20, x10, USNM 252185, 252189, 252190. h, i, hypostome, ventral and dorsal views, X 3, USNM 252186; note that middle body has transversely wide anterior lobe and narrowly elongate posterior lobe and that median furrow (sag.) is distinguished more clearly on interior than on exterior. j, k, hypostome, dorsal and ventral views, X 8, USNM 252187. I, o, 10, free cheeks, left, dorsal views, respectively X 6, x 10, X 10, USNM 252188, 252191, 252192. Robergia sp. Fragmentary cranidium and free cheek, dorsal views, X10, USNM 252193, 252194. Isotolus sp. Immature cranidia, deformed tec- tonically, all dorsal views, X 10. USNM 252195, 252196, 252197, respectively. q ,7. 8-“ . v—ad. Anataphrus sp. "a, w, cranidium, dorsal and an- terior views, X5, USNM 252198. 06, :11, ad, cranidia, dorsal views, respectively X 7, X 10, X10, USNM 252199, 252200, 252204. 2, aa, hypostomes, ventral views, X 10, USNM 252201, 252202. ab, ac, free cheek, anterior and dorsal views, X 7, USNM 252203. ae—am. Otarion sp. ae. free cheek, dorsal view, X 7, USNM 252205. af, ag, ah, cranidium, dorsal, anterior, and lateral views, X 7, USNM 252206. oi, oi, ak, cranidium, lateral, dorsal, and an- terior views, X 6, USNM 252207. al, pygidium, dorsal view, x 20, USNM 252208. am, associ- ated hypostome, ventral view, X 20, USNM 252209. Ceraums sp. 1. an, ac. cranidium, fragmentary, dorsal and anterior views, X 6. USNM 252210. ap, aq,. ar, cranidium, fragmentary, anterior, dorsal, and lateral views, X 5, USNM 252211. (Is—aw. Cryptolithoides sp. as, cranidium, dorsal view, X 7, of immature specimen; note that eye tuber- cles are still present. USNM 252212. at, au, ventral and dorsal views, X 5, of right side of fragmentary cranidial fringe with free cheek in place, USNM 252213. aw, free cheek, frag- mentary, ventral view, X 7, USNM 252214. an—ar. SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 19 9 THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA Cl7 FIGURE 6.—Late Ordovician conodonts from the lower dark dolomite member (a—j) and the middle limestone and dolo- mite member (lo—ab) of the Hanson Creek Formation, Mountain Boy Range, and from the graptolite—bearing platy limestone facies of the Hanson Creek Formation (ac), Monitor Range, Eureka County, Nevada. All speci- mens illustrated on figures 6 and 7, except 69, are re- posited in the U.S. National Museum of Natural History, Washington, DC. (USNM). Susann Braden assisted in the SEM photography. a—f. Amorphogmthus ordovicicus Branson and Mehl, from USGS colln. 8320—00. w—e, X 100; f, x 50. This is the zonal index for the latest Ordovician North Atlantic A. ordovicicus Zone (:upper third of Fauna 11 and Fauna 12 of the midcontinent faunas of Sweet, Ethington, and Barnes, 1971); Mays- vi‘llian through Richmondian (Ashgillian). a, b. Outer lateral views of two holodrontiform elements, USNM 248444 and USNM 248445. 6. Lateral view of trichonodelliform element, USNM 248446. d. Inner lateral view of eoligonodiniform ele- ment, USNM 248448. e. Upper view of ambalodiform element, USNM 248447. f. Upper view of amorphognathiform element, USNM 248449. g. Belodimz sp. A of Sweet, Ethington, and Barnes (1971) (:a new species of Belodina described by Craig (1968) from the lower half of the Fernvale Limestone of northern Arkansas). Postero-lateral inner view of adenticulate detortiform element, OSU 33382 (Orton Museum of Geology, The Ohio State University, Columbus, Ohio), x 50. From 2 m above base of Hanson Creek Formation, Span- ish Mountain, Mountain Boy Range, Bellevue Peak 15-minute Quadrangle (UTM grid, zone 11, E. 583,400 m, N. 4,368,250 m). Specimen courtesy of S. M. Bergstrom. This species is characteristic of middle Edenian (but may extend to the base of the Edenian according to W. C. Sweet, oral. commun., 1977) through middle Maysvillian of the central and western United States (Sweet and Bergstriim, 1976; Shatze-r, 1976) and the Canadian Arctic (Barnes, 1974). It appears to be a component of the “Arctic-Red River” fauna of earlier workers. h,z'. Panderodus yracilis (Branson and Mehl) element, from USGS colln. 8321—00. Inner and outer lateral views, USNM 248450, X 25‘. Elements of this species comprise 71 percent of all conodonts in the definitely Ordovician collections from the Hanson Creek Formation in the Mountain Boy Range and Wood Cone Peak area. j. Belodina dispansa (Glenister) s.f., from USGS colln. 832’1—CO. Outer lateral view, USNM 248451, X 100. k, l. Zygognathiform element, posterior and oblique lateral views of two specimens, USNM 248456 and USNM 248460, X 50; USGS colln. 8323—00. Several dif- ferent zygognathiform elements occur in our col- lections without other associated ramiform ele- r, s. w, x. ab. ments. Accordingly these were not referred to a discrete or multielement taxon. Oulodus ulr’ichi (Stone and Furnish), x 50. m. Posterior View of oulodiform element, USNM 248452; USGS colln. 8321—00. 71. Inner lateral view of prioniodiniform element, USNM 248454; USGS colln. 8326—00. 0. Oblique posterior view of trichonodelliform ele- ment, USNM 248455; USGS colln. 8323—00. 1). Lateral view of cordylodiform element, USNM 248453; USGS colln. 8323—00. Plegagnathus nelsom’ Ethington and Furnish s.f. In- ner lateral view, USNM 248459, x50; USGS colln. 8322—00. P. nelsom', P. dartem' Ethington and Furnish s.f., and Cordylodus robustus Ething- ton and Furnish s.f. form one apparatus that is a diagnostic component of postaearly Edenian Ordo- vician “Arctic-Red River” conodont faunas. P. dar— tom' and C. robustus elements have not been iden- fied in our collections. Belodina n. sp. D of Shatzer (1976), X 100; USGS colln. 8325—CO. Outer and inner lateral views of two difl'erent belodiniform elements, USNM 248457 and USNM 248458. This apparatus includes Belo- dina profumla (Branson and Mehl) of most previ- ously described Late Ordovician conodont faunas. This species is a. Maysvillian-Richmondian com- ponent of the “Arctic-Red River” fauna. Coelocerodontus trigom‘us Ethington s.f. Posterior view, USNM 248469, X 100, USGS colln. 8323—00. Oistodus venustus Staufl'er s.f. Outer lateral view, USNM 248464, )< 100; USGS colln. 8322—00. Drepanoistodus suberectus (Branson and Mehl). Inner lateral view of oistodiform element, USNM 248463, x 100; USGS colln. 8325—00. Elements of this species comprise nearly 10 percent of the conodonts in Ordovician collections from the Han- son Creek Formation in the Mountain Boy Range and Wood Cone Peak area. Pseudoneotodus mitmtus (Moskalenko). (=0isto- dus(?) sp. of Branson and Mehl (1933, pl. 9, fig. 3). Upper views of two specimens, USNM 248466 and USNM 248465, x 50; USGS colln. 8322—00. This is a sporadic element in Middle and Late Ordovician conodont faunas in the western United States. Pseudoneotodus beckmwrmi (Bischoff and Sanne- mann) s.f. Upper view, USNM 248467, x 100; USGS colln. 8323—00. P. beckmarmi ranges from the Late Ordovician through Early Devonian and appears to be the only morphotype in its ap- paratus. Belodina compressa (Branson and Mehl). Innerr lateral view of belodiniform element, USNM 248461, x 100; USGS colln. 8322—00. Prioniodiform element. Oblique anterior view, USNM 248470, x 50; USGS colln. 8327-00. Zygognathiform element, posterior vi‘eW, USNM 248462, x 50; USGS colln. 8327—00. This distinc- tive element with a greatly laterally compressed cutsp bearing a posterior costa and short but high lateral processes appears to be rather common in Cl8 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 FIGURE 6.—Continued. Late Ordovician conodont faunas of the western United States. Protopanderodus insculptus (Branson and Mehl) s.f. Outer lateral view, USNM 248468, x 50; from 6.1 m above the base of the Hanson Creek Formation (graptolite—bearing platy limestone facies), about 2.5 km south of the north end of Martin Ridge, Horse Heaven Mountain 15-min- ute Quadrangle. The extremely elongate posterior process is typical of true North American P. insculptus (middle Maysvillian-Richmondian). ac. This is the most common element in the meager conodont fauna recovered from three 5-kg samples from the lower part of the graptolitebearing platy limestone facies on Martin Ridge. In con- trast, only one specimen of P. insculptus was found among the nearly 2,500 conodonts from the Hanson Creek of the Mountain Boy Range and Wood Cone Peak area. Clearly, P. insculptus is a component of deeper and (or) cooler water cono— dont biofacies. References Cited—Continued Begg, J. L., 1939, Some new species of Proetidae and Otarioni- dae from the Ashgillian of Girvan: The Geological Maga- zine, v. 76, p. 372—382, pl. VI. Bradley, J. H., Jr., 1930, Fauna of the Kimmswick Limestone of Missouri and Illinois: Chicago University, Walker Museum of Paleontology Contributions, v. '2, n0. 6, p. 219—290, pls. 23—30. Branson, E. B., and Mehl, M. G., 1933, Conodonts from the Plattin (Middle Ordovician) of Missouri: Missouri Uni- versity Studies, v. 8, p. 101—119. Cooper, B. N., 1953, Trilobites from the lower Champlainian formations of the Appalachian Valley: Geological Society of America Memoir 55, 69 p., 19 pls. Graig, W. W., 1968, The stratigraphy and conodont paleon- tology of Ordovician and Silurian strata, Batesville Dis- trict, Independence and Izard Counties, Arkansas: Austin, University of Texas Ph.D. thesis, 393 p. Dean, W. T., 1962, The trilobites of the Caradoc Series in the Cross Fall Inlier of northern England: London, British Museum (Natural History) Bulletin, v. 7, no. 3, p. 65—134, pls. 6—18, 5 text figs. 1963, The Stile End Beds and Drygill Shales (Oirdovi- cian) in the east and north of the English Lake District: London, British Museum (Natural History) Bulletin, v. 9, no. 3, p. 47—65, pls. 1—5. 1974, The trilobites of the Chair of Kildare limestone (upper Ordovician) of eastern Ireland, part 2: Palaeon- tographical Society Monograph, p. 61—98, pl. 26—44. Delo, D. M., 1940, Phacopid trilobites of North America: Geo- logical Society of America Special Paper 29, 135 p., 13 pls. Duncan, Helen, 1956, Ordovician and Silurian coral faunas of western United States: US. Geological Survey Bul- letin 1021—F, p. 209—236. 1957, Bighorm'a, a new Ordovician coral genus: Jour- nal of Paleontology, v. 31, no. 3, p. 607~615. Hague, Arnold, 1892, Geology of the Eureka district, Nevada: US. Geological Survey Monograph 20, 419 p., and atlas of geologic maps [1883]. Hall, James, 1847, Paleontology of New York, containing descriptions of the organic remains of the lower division of the New York System, in Natural history of New York: Albany, C. van Benthuysen, v. 1, 338 p. Hu, Chung-Hung, 1975, Ontogeni‘es of four species of silici- fied Middle Ordovician trilobites from Virginia: Proceed- ings of the Geological Society of China, no. 18, p. 115—127, 4 figs, 2 pls. Ingham, J. K., 1970, The Upper Ordovician trilobites from the Cautley and Dent Districts of Westmoreland and Yorkshire: Palaeontographical Society Monograph, pt. 1, p. 1—58, pls. 1—9. [Published as part of v. 124.] 1974, The Upper Ordovician trilobites from the Caut- ley and Dent districts of Westmoreland and Yorkshire: Palaeontographical Society Monograph, pt. 2, p. 59—87, pls. 10—18. [Published as part of v. 128.] Kielan, Z., 1960 (1959), Upper Ordovician trilobites from Poland and some related forms from Bohemia and Scan- dinavia: Palaeontologia Polonica, no. 11, i—vi, 198 p., 36 pls. Kirk, Edwin, 1933, The Eureka quartzite of the Great Basin region: American Journal of Science, 5th ser., v. 26, p. 27—44. Lamont, A., 1935, The D'rummuck group, Girvan; a strati- graphical revision, with descriptions of new fossils from the lower part of the group: Geological Society of Glas- gow Transactions, v. 19, p. 288—332, pls. 7—9. Lane, P. D., 1971, British Cheiruridae- (Trilobita): Paleon- tographical Society Monograph, 95 p., 16 pls. [published as part of v. 125.] Le Fevre, Jean, Barnes, C. R., and Tixier, Michel, 1976, Paleoecology of Late Ordovician and Early Silurian conodontophorids, Hudson Bay Basin, in Barnes, C. R., ed., Conodont paleoecology: Canada Geological Associa- tion Special Paper 15, p. 69—89. Lenz, A. C., and Churkin, Michael, Jr., 1966, Upper Ordovi- cian trilobites from northern Yukon: Palaeontology, v. 9, pt. 1, p. 39—47, pls. 4—5. Ludvigsen, Rolf, 1975, Ordovician formations and faunas, southern MacKenzie Mountains: Canadian Journal of Earth Sciences, v. 12, no. 4, p. 663—697, 5 pls. 1976a, Middle Ordovician trilobite biofacies, southern MacKenzie Mountains: Canada Geological Association Annual Meeting, P. S. Warren Biostratigraphy Sympo— sium, Program with Abstracts, v. 1, p. 39. 1976b, New cheirurinid trilobites from the lower Whit- taker Formation (Ordovician), southern MacKenzie Mountains: Canadian Journal of Earth Sciences, v. 13, no. 7, p. 947—959, 2 pls. 1977, The Ordovician trilobite Cemurinus Barton in North America: Journal of Paleontology, v. 51, no. 5, p. 959—972, pls. 1, 2. McCracken, A. D., and Barnes, C. R., 1979, Conodonts and the Ordovician-Silurian boundary from the Ellis Bay For- mation, Anticosti Island: Canada Geological Survey Bul- letin [in press]. [References continued on p. 021] THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA 019 TABLE 2.—Distribution of conodonts in collections from the Hanson Creek Formation, Mountain Boy Range, and Wood Cone Peak area, Belle’vue Peak 15-minute Quadrangle, Eureka County, N ev. [Collections 8320—8329—C0 _and 9785—813 from measured section 2001 m east of center sec. 31, T. 19 N., R. 53 E., dscribed in text. Collection 8292— . .from dark-gray thick-bedded. limestone about 200 m west of measured section (UTM Grid, zone 11. E. 581 450 m, N. 4369 600 m); mapping 111%]th tth115 outcrop may overlie the middle limestone of the measured section. Collections 833278334—00 from Wood Cone Peak area. Leaders, __, no presen Distance (meters) above base of section _______ 0 to 0.6 5 6.1 14 9 18.6 19 8 20.7 21 3 22.9 27.4 to 28 Collection No. (USGS) ________________________ 8320—00 8321—00 8322—C0 8323—00 8325—C0 8326—00 8327—00 8328—CO 8324—00 8329—00 9785—SD 8292—00 8332—C0 1 8333—00 2 8334—00 3 l“ to a. c)! 3. m 2“ on Sample weight ______________________ kilograms- _ 4.6 an :7! m w on 0: at I“ N 1“ c» N w m e N w on Amorphognathus ordom’cicus Branson and Mehl: Ambalodiform element ____________ 7 Amorphognathiform element ______ 43 Eoligonodiniform element _________ 5 Holodontiform element ___________ 6 Tetraprioniodiform element _______ 1 Trichonodelliform element ________ 8 Aphelognathus? n. sp.: Eoligonodiniform element _________ Pri‘oniodiform element ____________ Spathognathodiform element ______ ___ ___ ___ ___ -_ Trichonodelliform element ________ ___ ___ Zygog'nathiform element __________ ___ ___ ___ ___ __ Belodina compressa (Branson and Mehl) : Eobelodiniform element ___________ 5 1 6 3 5 ___ ___ ___ ___ _- Belodiniform element _____________ 1 ___ 2 1 B. dispansa (Glenister) s.f. ___________ __- 26 12 13 33 14 ___ 1 ___ ___ ___ ___ ___ 3 1 B. n. sp. D elements of Shatzer (1976)__ ___ __- ___ ___ 11 2 ___ ___ ___ ___ ___ 1 _-_ ___ ___ Coelocerodontus trigonius Ethington s.f.- ___ ___ ___ 2 ___ ___ ___ ___ ___ ___ 1 _-_ ___ ___ ___ Drepanoistodus suberectus (Branson and Mehl) elements ____________________ 23 34 32 27 30 22 2 ___ 3 ___ ___ 2 26 22 4 Exochognathus keislognathoides Pollock, Rexroad, and Nicoll s.f. _____________ ___ ___ ___ ___ ___ ___ _-_ ___ ___ ___ 6 -__ ___ ___ ___ Oistodus venustus Staufl’er s.f. _________ ___ ___ 2 ___ 1 _-_ ___ ___ ___ ___ __- ___ ___ ___ ___ Pseudoneotodus beckmanni (Bischofl' and Sannemann) s.f. _______________ ___ ___ ___ Pseudoneotodus mitratus (Moskalenko) _____________________ _ _ _ _ _ _ 3 Oulodus ul’richi (Stone and Furnish) : Cordylodiform element‘ ___________ ___ ___ ___ Cyrtoniodiform element ___________ 1 3 8 Eoligonodiniform element‘ ________ 1 ___ ___ Oulodiform element ______________ ___ 5 4 Prioniodiniform element __________ ___ 1 ___ Trichonodelliform element _________ ___ ___ 3 Zygog'nathiform element __________ ___ 1 1 Pande’rodus gracilis (Branson and Mehl) elements _____________________ 169 405 252 288 197 150 16 9 11 10 ___ 12 54 122 26 Panderodus sp. elements ______________ ___ ___ ___ ___ ___ ___ ___ -__ ___ _ Plegagnathus nelsom‘ Ething'ton and Furnish s.f _________________________ 2 1 1 4 ___ ___ ___ ___ 1 ___ ___ ___ 1 ___ ___ P’rotopanderodus insculptus (Branson and Mehl) s.f. _____________________ _-_ ___ ___ ___ ___ e__ 1 ___ ___ _ Cyrtoniodiform elements ______________ -__ ___ ___ ___ ___ ___ 4 ___ ___ Distomodifor'm elements ______________ ___ ___ ___ _-_ ___ ___ ___ ___ ___ Eoligonodiniform elements ____________ ___ -__ ___ ___ -__ ___ ___ -__ ___ 1 ___ Plectospathodiform elements __________ ___ ___ ___ ___ ___ __- ___ ___ ___ 1 1 Prioniodiform elements _______________ ___ ___ ___ ___ ___ ___ 1 ___ ___ Prioniodiniform elements _____________ ___ ___ -__ ___ ___ ___ ___ ___ ___ 1 ___ Spathognathodiform elements _________ ___ ___ ~__ ___ ___ ___ ___ ___ ___ ___ 2 Trichonodelliform elements with very laterally compressed cusp ___________ ___ _-_ 1 ___ 1 __e ___ ___ ___ ___ ___ _ Zygognathifonn elements with very lateraJIy compressed cusp ___________ -__ 4 6 ___ 3 -__ 7 ___ ___ 2‘ ___ ___ ___ ___ 2 I I I I I | | I I I l | I I | | 00%th I I I I I I l I I I I I WHGU‘QQW 03 03 I I I 1Dark-gray platy limestone. UTM Grid,, zone 11, E. 574 750 m, N. 4359 550 m (=D2707—CO). 2 Dark-gray flaggy skeletal limestone. UTM Grid zone 11, E. 674 800 m, N. 4 359 125 m (=D2708—CO). 3Dark gray fine-grained limestone. UTM Grid, zone 11, E. 574 630 m, N. 4 359 840 m. 4Cordylodiform and eoligonodiniform elements both appear to occupy the Sc position (apparatus nomenclature of Sweet and Schonlaub, 1975) in this species of Oulodus. These collections contain no other apparatus in whichan eoligonodiniform or cordylodizform element can be accommodated. C20 FIGURE 7.—Latest Ordovician and (or) earliest Silurian conodonts from the middle limestone and dolomite member of the Hanson Creek Formation, USGS colln. 9785—SD, Mountain Boy Range, Eureka County, Nev. (except f). a—~g. Aphelognathus? n. sp. The generic assignment is queried because Sweet, Thompson, and Satterfield (1975) diagnosed the apparatus of Aphelognathus as containing a dichog'nathiform or prioniodiniform element in the 01 position. In this species, a prionio- diform element occupies the 01 position. In addition, the denticle pattern and profile designated by Sweet, Thompson and Satterfield (1975, p. 29) as char- acteristic of the genus does not fit elements of this species. The apparatus of A.? n.. sp. was partly determined from conodont collections from near the base of the upper part of the Ely Springs Dolomite, Funeral Mountains, Inyo County, Calif. In these collections, A.? n. sp. elements are abundant, well preserved, and are nearly the only conodont ele- ments aside from distomodids, keislognathodids, Oulodus sp., and panderodids. Aphelognathus is sup- SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 FIGURE 7.——Continued. posedly a Middle and Late Ordovician genus. Pol- lock, Rexroad, and Nicoll (1970), however, desig- nated a 1:) element from Lower Silurian beds in northern Michigan and Ontario, Aphelognathus silum'cus. Moreover, the same form occurs in the B and lower C divisions of the Llandovery in the Welsh Borderland (Aldridge, 1972). The 1) element of our species closely resembles A. silum'cus s.f., but A.? n. sp. has only one adenticulate lateral process, whereas A. siluricus s.f. has an outer and inner adenticulate process. A.? n.sp. may well be ances- tral to A. silum'cus. a, b. Spathognathodiform (1)) element, inner and outer lateral views of two specimens, USNM 248472 and USNM 248473, x 50. Note the adenticulate outer lateral process extending from the anterior edge of the main cusp. This feature demonstrates the consanguinity of this and the prioniodiform element. c, d. Eoligonodiniform element, outer and inner lateral views of two specimens, USNM 248474 and USNM 248475, x 75. e, f. Prioniodiform element, oblique anterior views of two specimens, x 50. e. USNM 248476. Thus far, all prioniodi- form elements recovered from the Hanson Creek Formation are fragmentary and only the adenticulate upper part of the anterior process is preserved. If We had not had complete specimens of this element avail- able to us from the Ely Springs Dolomite of the Funeral Mountains (fig. f), we would have assumed this element to be dichognathiform. f. USNM 248477, from USGS colln. 9628— SD (about 1.5 m above the base of the upper part of the Ely Springs Dolomite, northwest base of spur 4298, southeast end of Funeral Mountains, SW14 sec. 33, T. 26 N., R. 4 E., Ryan 15-minute quadrangle, Inyo County, Calif). Note the denticulate anterior process. 9. Tricvhonodelliform USNM 248478, x 75. Distomodiform elements, outer and inner lateral views of denticulate and adenticulate elements, USNM 248472 and USNM 248471, x 100. These are char- acteristic Early Silurian morphotypes. Spathognathodiform element, outer(?) lateral View, USNM 248481, x 100. Plectospathodiform element, inner lateral view, USNM 248480, x 100. This is a typical post-Ordo- vician morphotype. It probably belongs in the same apparatus as the element shown in j. Exochognathus keislognathoides Pollock, Rexroad, Nicoll s.f. Posterior view, USNM 248479, x 100. Better preserved specimens have denticulate outer lateral and posterior processes and an adenticulate inner lateral process. This is a common element in Early Silurian conodont faunas. element, posterior view, THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C21 References Cited—Continued Macomber, R. W., 1970, Articulate brachiopods from the upper Bighorn Formation (Late Ordovician) of Wyoming: Journal of Paleontology, v. 44, no. 3, p. 416—450, pls. 77—80. Merriam, C. W., 1940, Devonian Stratigraphy and Paleon- tology of the Roberts Mountains Region, Nevada: Geo- logical Society of America Special Paper 25, 114 p. 1963a, Paleozoic rocks of Antelope Valley, Eureka and Nye Counties, Nevada: U.S. Geological Survey Profes- sional Paper 423, 67 p. 1963b, Geology of the Cerro Gordo mining district, Inyo County, California: U.S. Geological Survey Pro- fessional Paper 408, 83 p. Merriam, C. W., and Anderson, C. A., 1942, Reconnaissance survey of the Roberts Mountains, Nevada: Geological Society of America Bulletin, v. 53, no. 12, p. 1675—1727. Moskalenko, T. Z., 1973, Conodonts of the Middle and Upper Ordovician on the Siberian Platform: Academy of Sci- ences of the USSR, Siberian Branch, Transactions of the Institute of Geology and Geophysics, v. 137, 143 p. [in Russian]. Mullens, T. E., and Poole, F. G., 1972, Quartz-sand—bearing zone and Early Silurian age of upper part of the Hanson Creek Formation in Eureka County, Nevada: U.S. Geo— logical Survey Professional Paper 800—B, p. B21—BZ4. Nelson, S. J ., 1963, Ordovician paleontology of the northern Hudson Bay Lowland: Geological Society of America Memoir 90, 152 p., 37 pls. Nolan, T. B., Merriam, C. W., and Williams, J. S., 1956, The stratigraphic section in the vicinity of Eureka, Nevada: U.S. Geological Survey Professional Paper 276, 77 p. Nowland, G. S., and Barnes, C. R., 1979, Late Ordovician conodonts from the Vauréal Formation. Anticosti Island, Quebec: Canada Geological Survey Bulletin [in press]. Opik, A., 1937, Trilobiten aus Estland: Tartu University Geological Institution, Publications No. 52, 163 p., 26 pls. Owens, R. M., 1973, British Ordovician and Silurian Proetidae (Trilobita): Paleontographical Society Monograph, p. 1—98, pls. 1—15. [Published as part of v. 127.] Pollock, C. A., Rexroa-d, C. B., and Nicoll, R. S., 1970, Lower Silurian conodonts from northern Michigan and Ontario: Journal of Paleontology, v. 44, no. 4, p. 743—7 64. Raymond, P. E., 1925, Some trilobites of the lower middle Ordovician of eastern North America: Harvard College Museum of Comparative Zoology Bulletin, v. 67, no. 1, 180 p., 10 pls. Ross, R. J., Jr., 1957, Ordovician Fossils from wells in the Williston Basin, eastern Montana: U.S. Geological Sur- vey Bulletin 1021—M, p. 439—510,, pls. 37—43. 1967, Calymenid and other Ordovician trilobites from Kentucky and Ohio: U.S. Geological Survey Professional Paper 583—B, p. B1—B18, 5 pls. 1976, Ordovician sedimentation in the western United States, in Bassett, M. G., ed., The Ordovician System— Proceedings of a Palaeontological Association Symposium, Birmingham, 1974: Cardiff, University of Wales Press and National Museum of Wales, p. 73—105. 1977, Ordovician paleogeography of western United States, in Stewart, J. H., Stevens, C. H., and Fritsche, A. E., eds., Paleozoic Paleogeography of the Western United States, Pacific Coast Paleogeography, Symposium 1, Pacific Section: Society of Economic Paleontologists land Mineralogists, p. 19—38. Ross, R. J., Jr., and Berry, W. B. N., 1963, Ordovician grapto— lites of the Basin Ranges in California. Nevada, Utah, and Idaho: U.S. Geological Survey Bulletin 1134, 177 p., 13 pls. Ross, R. J., Jr., and Dutro, J. T., Jr., 1966, Silicified Ordoh vician brachiopods from east-central Alaska: Smithson- ian Miscellaneous Collections, v. 149, no. 7, 22 p., 3 pls. Ross, R. J., Jr.,.and Shaw, F. C., 1972, Distribution of the Middle Ordovician Copenhagen Formation and its trilo- bites in Nevada: U.S. Geological Survey Professional Paper 749, 33 p. 8 pls. Roy, S. K., 1941, The Upper Ordovician fauna of Frobisher Bay, Baffin Land: Field Museum of Natural History Publications, Geology Memoirs, v. 2, 212 p., 146 figs. Schmidt, F., 1881, Revision der Ostbaltischen Silurischen Trilobiten, Abteilung 1, Phacopiden, Cheiruriden, und Encrinuriden: Memoires de l’Académie Impériale des Sciences de St. Petersbourg, VIIe Série, Tome XXX, No 1, 237 p., 16 [pls. Schuchert, Charles, and Cooper, G. A., 1930, Upper Ordovician and Lower Devonian stratigraphy and paleontology of Percé, Quebec: American Journal of Science, 5th series, v. 20, p. 161—176, 265—288, 365—392. Shatzer, D. C., 1976, Conodont biostratigraphy of the Fre- mont and Priest Canyon Formations (Upper Ordovician) at Kerber Creek, Saguache County, Colorado: Columbus, The Ohio State University M. S. thesis, 141 p. Shaw, F. C., 1968, Early Middle Ordovician Chazy trilobites of New York: New York State Museum and Science Service Memoir 17, 163 p., 24 pls. 1974, Simpson Group (Middle Ordovician) trilobites of Oklahoma: Journal of Paleontology, v. 48, no. 5, supp. II, Paleontological Society Memoir 6, 54 p., 12 pls. Shirley, Jack, 1936, The British trilobites. of the family Calymenidae: Geological Society of London Quarterly Journal, v. 92, pt. 4, p. 384—422, pls. 29—31. Slocum, A. W., 1913, New trilobites from the Maquoketa beds of Fayette County, Iowa: Field Museum of Natural His- tory, Publications in Geology, ser. 4, p. 43-83. Sweet, W. C., and Bergstriim, S. M., 1976, Conodont bio- stratigraphy of the Middle and Upper Ordovician of the United States Midcontinent, in Bassett, M. G., ed., The Ordovician System—Proceedings of a Palaeontological Association Symposium, Birmingham, 1974: Cardiff, University of Wales Press and National Museum of Wales, p. 121—151. Sweet, W. C., Ethington, R. L., and Barnes, C. R., 1971, North American Middle and Upper Ordovician conodont faunas, in Sweet, W. C., and Bergstriim, S. M., eds, Conodont biostratigraphy: Geological Society of America Memoir 127, p. 163—193. Sweet, W. C., and Sch6nlaub, H. P., 1975-, Conodonts of the genus Oulodus Branson & Mehl, 1933: Geologica et Palae- ontologica, v. 9, p. 41—59. Sweet, W. C., Thompson, T. L., and Satterfield, I. R., 1975, Conodont stratigraphy of the Cape Limestone (Mays- villian) of eastern Missouri, in Missouri Geological Sur- vey, Studies in Stratigraphy: Missouri Geological Survey Department of Natural Resources, Report Investigations 57. pt. 1, p. 1—59. C22 Tripp, R. P., 1954, Caradocian trilobites from mudstones at Craighead Quarry, near Girvan Ayrshire: Royal Society of Edinburgh Transactions, v. 62, pt. 3, p. 655—693, pls. 14. 1965, Trilobites from the Albany division (Ordovician) of the Girvan district, Ayrshire: Palaeontology, v. 8, pt. 4, p. 577—603, pls. 80—83. 1967, Trilobites of the Upper Stinchar limestone- (Or- dovician) of the Girvan District, Ayrshire: Royal Society of Edinburgh Transactions, v. 67, no. 3, p. 43—93, pls. I—VI. Troedsson, G. T., 1928, On the Middle and Upper Ordovician faunas of northern Greenland—Part II: Meddelelser om Grénland, Bd. 72, p. 1—197, pls. 1—56. Twenhofel, W. H., 1927, Geology of Anticosti Island: Canada Geological Survey Memoir 154, 480 p., 60 pls. Ulrich, E. 0., 1888, On Sceptropo'ra, a new genus of Bryozoa, with remarks on Helopora Hall, and other genera of that type: American Geologist, v. 1, p. 228—234. Warburg, Elsa, 1925, Trilobites of the Leptaena Limestone in Dalarne with a discussion of the zoological position and the classification of the Trilobita: Bulletin of the Geological Institutions of the University of Uppsala, v. 17, p. 1—446, 11 pls. Weller, S., 1903, The Paleozoic Faunas: New Jersey Geological Survey, Report on Paleontology, v. 3, 462 p., 52 pls. Whittington, H. B., 1941, The Trinucleidae, with special ref- erence to North American genera and species: Journal of Paleontology, v. 15, no. 1, p. 21—41, pls. 5, 6. SHORTER CONTRIBUTIONS T0 STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 1950, Sixteen Ordovician genotype trilobites: Journal of Paleontology, v. 24, no. 5, p. 531—565, pls. 68—75, 9 text figs. 1952, A unique remopleuridid trilobite: Breviora, Harvard University Museum of Comparative Zoology, no. 4, 9 p., 1 pl. 1954, Ordovician trilobites from Silliman’s Fossil Mount: Geological Society of America Memoir 62, p. 119—149, pls. 59—63. 1959, Silicified Middle Ordovician trilobites—Remo- pleurididae, Trinucleidae, Raphiophoridae, Endymioniidae [Va.]: Harvard College Museum of Comparative Zoology Bulletin, v. 121, no. 8, p. 373—496, 36 pls. 1963—1968, Monograph of the Ordovician Trilobites of the Bala Area, Merioneth, 138 p., 32 pls.: London, Palae- ontographical Society, published in parts, as follows: 1962, v. 116, Part I, p. 1—32, pls. I—VIII.; 1964, v. 118, Part II, p. 33—62, pls. IX—XVIII; 1966, v. 120, Part III, p. 63—93, pls. XIX—XXVIII; 1968, v. 122, Part IV, p. 93—138, pls. XXIX—XXXII. Whittington, H. B., and Evitt, W. R., II, 1954, Silicified Middle Ordovician trilobites [Va.]: Geological Society of America Memoir 59, 137 p., 33 pls. Williams, Alwyn, 1962, The Barr and Lower Ardmillan Series (Caradoc) of the Girvan District, south-west Ayrshire, with descriptions of the Brachiopods: Geological Society of London Memoir 3, 267 p., 25 pls. The Marble Hill Bed: An Offshore Bar—Tidal Channel Complex in the Upper Ordovician Drakes Formation of Kentucky By W C SWADLEY SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—D Prepared in cooperation with the Kentucky Geological Survey FIGURE 1. 95‘?" CONTENTS Page Abstract _________________________________________________________________ D1 Introduction _____________________________________________________________ Previous nomenclature ____________________________________________________ Stratigraphy _____________________________________________________________ Extent and thickness ______________________________________________________ Lithology ________________________________________________________________ Environments of deposition ________________________________________________ Depositional history ______________________________________________________ Summary ________________________________________________________________ References cited __________________________________________________________ oooommusoar-AHH ILLUSTRATIONS Page Map showing location of study area and areas of exposure of the Rowland Member of the Drakes For- mation and the Bull Fork Formation in north-central Kentucky ________________________________ D2 Diagrammatic cross section showing relation of the main elements of the Marble Hill Bed to the enclosing stratigraphic units Within the study area ________________________________________________________ Map showing thickness and stratigraphic position of the Marble Hill Bed within the Rowland Member ____ Cross sections showing position of Marble Hill Bed Within Rowland Member __________________________ Sketch map 0f study area showing interpretation of depositional environments ________________________ ~1th HI SHORTER CONTRIBUTIONS TO STRATIGRAPI-IY AND STRUCTURAL GEOLOGY, 1979 THE MARBLE HILL BED: AN OFFSHORE BAR-TIDAL CHANNEL COMPLEX IN THE UPPER ORDOVICIAN DRAKES FORMATION OF KENTUCKY By W C SWADLEY ABSTRACT The Marble Hill Bed of the Rowland Member of the Drakes Formation consists of a group of linear bodies of biospar- rudite and biosparite that occur near the northern edge of the Rowland Member in north-central Kentucky and south- eastern Indiana. The Rowland Member, a unit of interbedded micritic limestone and mudstone, abruptly thins and pinches out into interbedded fossiliferous limestone and shale of the Bull Fork Formation. The Marble Hill was deposited as an offshore bar and tidal channel complex along the seaward northern edge of the shallow marine platform on which the Rowland was deposited. Deposition was probably initiated by uplift that resulted in a northward transgression of the plat- form environment over the southern edge of the open shelf on which the Bull Fork was deposited. Later subsidence of the area resulted in a southward regression of the platform environment and ended Marble Hill deposition. INTRODUCTION The Marble Hill Bed is a unit of bioclastic lime- stone in the Rowland Member of the Drakes Forma- tion in north-central Kentucky and southeastern Indiana (fig. 1). It consists of several related lime stone bodies along or near the abrupt northern pinchout of the Rowland Member into the Bull Fork Formation. Recent mapping in the area as part of the statewide US. Geological Survey—Kentucky Geological Survey cooperative mapping program has for the first time determined the extent and shape of the Marble Hill Bed and its stratigraphic relations. The depositional environments and history are here- in analyzed on the basis of lithology, field relation- ships, and fa-unal content. This report is based largely on field data obtained during the mapping of six 71/2-minute quadrangles in north-central Kentucky and reconnaissance map- ping in southeastern Indiana. Field descriptions of the rock units were supplemented by the study of thin sections. Thin-section preparation and part of the petrographic analysis were done by Brandon C. Nuttall, a graduate student at the University of Cincinnati. PREVIOUS NOMENCLATURE The Marble Hill was first described by Owen (1859), who named it for exposures in a building- stone quarry at Marble Hill, Ind. The name was used in several reports in the early 1900’s and then apparently fell into disuse (Weiss and Norman, 1960). Much of the later literature on the area as- signs the part of the stratigraphic section that in- cludes the Marble Hill to the Waynesville Shale. In a summary of stratigraphic nomenclature of In- diana, Shaver and others (1970) included the Marble Hill Bed in the Dillsboro Formation. Recent work in north-central Kentucky by Swadley and Gibbons (1976) reassigned the unit as the Marble Hill Bed of the Rowland Member of the Drakes Formation. The name Drakes Formation is herein extended to a small area in southeastern Indiana where the Marble Hill Bed is exposed. STRATIGRAPHY Upper Ordovician strata of the study area consist largely of limestone interbedded with shale and mud- stone. The Marble Hill Bed comprises two elastic limestone bodies within and near the north edge of the Rowland Member of the Drakes Formation (fig. 2). The Rowland consists of interbedded limestone and mudstone. It overlies fossiliferous limestone and shale of the Bull Fork Formation and is overlain by fossiliferous limestone and shale of the Bardstown Member of the Drakes Formation. The Bardstown is overlain by the Saluda Dolomite Member of the Drakes. The Bull Fork Formation is a unit of interbedded limestone and shale that is exposed along two north- D1 D2 86° 85° SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 84" 33° ' | 390.— Madison \ \. Carrout ‘.““ 38° ardstown Stanford o Area of figure 3 / § \X\§;\:\ '\ M\\’X°nh § 0 ?‘ \\\\\/il\hitp\f\g!§ dd \ Louisville \\\’L1t\£ \\\sjt\lillglc:\u\m \ § U/l \ ‘ \®£k\iegositi0n KENTUCKY Richmond 0 100 KILOMETERS I Cincinnati OHIO — River 0 Lexington EXPLANATION Area of exposure of Rowland Member of Drakes Formation Area of exposure of Bull Fork Formation l 37° 1 FIGURE 1.—Map showing location of study area and areas of exposure of the Rowland Member of the Drakes Forma- tion and the Bull Fork Formation in north-central Kentucky trending bands ‘of outcrop flanking the inner part of the central Kentucky region (fig. 1). Along the eastern band, the formation thins from about 60 m at the type section near Maysville, Ky. (Peck, 1966), to a pinchout about 75 km to the south. In the west, the Bull Fork is about 65 m thick near Carrollton, Ky., and thins southward, pinching out 50 km to the south. The formation consists of medium-gray poorly sorted biomicrudite in thin, even to irregular beds that are i‘nterlayered with gray calcareous shale. Well-sorted, fine-grained biosparite in thin even beds and even-bedded calcareous siltstone are minor components. Overall, the limestone and shale con- tents are nearly equal, but limestone is generally more abundant near the base of the unit, and shale is more abundant toward the top. Fossil brachiopods and bryozoans are abundant in the Bull Fork; echxinoderms, trilobites, and gastropods are common; and solitary corals occur sporadically. The Rowland Member of the Drakes Formation is a sparsely fossilferous, thin-bedded, dolomitic to limy mudstone named by Weir and others (1965) in THE MARBLE HILL BED IN THE UPPER ORDOVICIAN DRAKES FORMATION, KENTUCKY N s Saluda Dolomite Member ' 38. Bardstown Member E .2 Marble Hill Rowland Member 8 Bed ‘ ‘ Marble Hill Bed Bull Fork Formation FIGURE 2.-Diagrammatic- cross section showing relation of the main elements of the Marble Hill Bed to the enclosing stratigraphic units within the study area. southeastern central Kentucky. It crops out in an arcuate belt along the south and west side of the inner part of the central Kentucky region (fig. 1). Northwestward from the type area it becomes more limy and more fossiliferous. Near the northwest end of its outcrop belt, in the study area, the Rowland consists of about 12 m of light-gray clayey micrite interbedded with meduim-gray to greenish-gray cal- careous mudstone. The limestone is commonly in nodular to even beds as much as 1 m thick but locally is in thin laminae. Many beds have been extensively burrowed and weather to a knobby surface with numerous vertical finger-size holes. Mudstone is dominant in the lower half of the member and com- monly forms a 1- to 2-m-thick bed at the base. Fossils include common to abundant colonial corals as much as 0.6 m in diameter, common ostracodes, and sparse brachiopods and bryozoans. At its (north- ern limit, the Rowland thins abruptly and pinches out into the Bull Fork Formation (fig. 2). The Marble Hill Bed occurs as a thin linear body at the base of the Rowland and as a thicker, more complex unit along the northern edge, where it locally makes up almost the entire thickness of the member. The Marble Hill is described in more de- tail in following paragraphs. Overlying the Rowland Member is the Bardstown Member of the Drakes, a. unit of fossiliferous lime- stone and interbedded shale named by Peterson (1970) in southwestern central Kentucky. It occurs along a belt of outcrop similar to the western part of the Rowland outcrop (fig. 1). In the study area, it consists of 12 to 18 m of medium-gray limestone interbedded with an equal amount of medium-gray calcareous shale. The Bardstown is lithologically and D3 faunally so much like the Bull Fork Formation that it cannot be differentiated from the Bull Fork north of the pinchout of the Rowland Member (fig. 2). The Saluda Dolomite Member is the upper mem- ber of the Drakes and consists of 20 m of dolomite and dolomitic to limy mudstonb. It overlies the Bardstown Member and the later 11y equivalent part of the Bull Fork. EXTENT AND THIC NESS The Marble Hill Bed is expose in a small area of north—central Kentucky and so theastern Indiana south of Madison, Ind. (fig. 1). It consists of two clastic limestone bodies: a narro , elongate unit at the base of the Rowland Member about 11 km south of the Rowland pinchout and a w'der, more complex body that occurs mostly along the northern pinchout of the member (fig. 3). The basal unit is about 1 km wide and is discon— tinuously exposed for 29 km along an east-northeast trending outcrop. Near the east end it breaks into a line of discontinuous pods or lenses. In cross sec- tion, the body is flat bottomed, convex, and asym— metrical and has the steeper face to the northwest (fig. 4, section B—B’). The basal contact appears conformable. The thickness ave ges about 1.2 m; the maximum observed thicknes is 4 m. The body goes into the subsurface at the west end. Its sub- surface extent is unknown. The larger unit of Marble Hi1 limestone occurs along the northernedge of the owland Member as an elongate body that parallels e basal unit and extends a short distance into Indi na before descend- ing into the subsurface (fig. 3). It is as much as 9 m thick, thinning abruptly to the nbrthwest and more gradually to the southeast (figl 4, sections A—A’ and B—B’). The unit appears to rest conformably on the basal mudstone of the Rowland. Near the west end of the exposed body, the southward-thinning edge occurs as an irregular tongue of Marble Hill strata that extends for 3 km into the main body of the Rowland Member. It is overlain by a northward- thinning tongue that has Rowland lithology (fig. 4, section A—A’). From 1 to 5 km east of the Ohio River, the southward extension occurs at the top of the member (fig. 4, section B—B’). There is a small area about 1.5 km east of the Ohio River (fig. 3) where both tongues are present. At the east end the larger body of the Marble Hill Bed is joined by three subunits The first subunit joins at the eastern tip and exten is for nearly 10 km to the east-southeast. For about half that distance, D4 SHORTER CONTRIBUTIONS T0 STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 85°22'30" 85°15' l | B \\\‘~‘“ 330 ,‘:::§$“{{{ V 37" 033%?“ —|— _ 30-, ‘\\\\\ Ms A M __ stew o " sexte\\ 0 Qwiiia,‘ $§t§§§e§ws \ ~ . ;;‘:""‘\‘\\}‘3“£‘\‘{§\\\\\§‘°\{\\‘\\“‘:§§2“ «ssh (‘\§§\\\\\\ _ V ('3 ’ K in .~‘ \ Q~ 38° 30’ B! Area of this figure _ EXPLANATION Marble Hill Bed at top of Rowland Member Marble Hill Bed within Rowland Member KILOMETERS I ------ Northern limit of Rowland Marble Hill Bed at base of Rowland Member Member (where not coincident V 9 4“ 3:5 With edge of Marble Hill Bed? (,0 0° FIGURE 3.—Map showing thickness and stratigraphic position of the Marble Hill Bed within the Rowland Member. Con- tour interval 1.5 m. Cross sections A—A', B—B’, and C—C’ are shown in figure 4. Inset map gives numbers of U.S. Geological Survey 71/2—minute geologic quadrangles within which the Marble Hill Bed occurs. this subunit occurs at the top of the Rowland Mem- ber; then the upper contact drops gradually in the section as the subunit thins to a featheredge near the middle of the Rowland (fig. 4, section C—C’). The basal contact is sharp and locally discornfom- able. Farther west, two other subunits join the larger body of the Marble Hill near its east end. These are south-trending linear bodies as much as 3 m thick and about 0.8 km wide that extend for as much as 13 km to the south (fig. 3). In cross section each of these two bodies is convex downward and has a flat upper surface coincident With the upper contact of the Rowland Member. The basal contact seems to be conformable; these units grade laterally into and intertongue with poorly bedded to thinly laminated Rowland micrite over a distance of 1 to 2 m. LITHOLOGY The Marble Hill Bed typically consists of gastro- pod biosparrudite and echinoderm biosparite. Bio- micrudite locally is an important component. The THE MARBLE HILL BED IN THE UPPER ORDOVICIAN DRAKES FORMATION, KENTUCKY A A' 30 ' ‘[ Odb 20 r - m 0dr 1o — ‘ Km _ Ob O (I) B B' (I: 30( _ w ’— I“ 2 Odb 20 _ Z a; m 0dr 3 10 m _ Z ¥ 0 Ob I 0 p- C C 30 — j 20— 10— Ob EXPLANATION Odb 0dr Rowland Member m Marble Hill Bed Bardstown Member Drakes Formation Ob Bull Fork Formation FIGURE 4.——Cross sections showing position of Marble Hill Bed within Rowland Member. Location of sections shown on figure 3. biosparrudite is very light gray and weathers very pale orange. It occurs commonly in even beds as much as 1 m thick and is locally thinly crossbedded in sets 0.5 to 1.0 m thick. The major components are abundant whole and broken high-spired gastropods as much as 50 mm long, common bryozoan and brachiorpod fragments, common to sparse echino‘derm D5 fragments, and sparse trilobite fragments. Fossil fragments are abraded, and sorting is poor. The cement is commonly medium to coarse sparry calcite with minor amounts of microspar. Small amounts of collophane occur as fillings in fossil cavities. The biosparrudite is resistant and weathers to bold rounded ledges. The sparry calcite that fills the abundant gastropod shells weathers readily, yielding a pitted surface. Biosparite is light olive gray to very light gray and Weathers yellowish gray. It consists of abundant broken and rounded echinoderm plates, common trilobite fragments, and sparse bryozo-an and brach- iopod fragments. Grains are fine to medium, and sorting is generally good. The cement is medium- to cloarsehgrained sparry calcite, much of which oc- curs as overgrowths on echinoderm fragments. The rock is slightly phosphatic. Biosparite is commonly thinly crossbedded in sets 0.3 to 0.6 m thick. It is less resistant than the biosparrudite but forms con- spicuous ledges. Biomicrudite is light to medium gray and weathers light gray to grayish orange. It consists of abundant whole and broken gastropods, abundant coarse to fine brachiopod and b‘ryozoan fragments, and abundant to common echinoderm plates. Sorting is very poor. The cement varies from micrite to sparry calcite. Beds are even to slightly irregular and generally 0.3 to 0.6 m. thick. A common exposure of the Marble Hill Bed con— sists of a basal unit of crossbedded biosparite over- lain by an equal or greater thickness of thick-bedded gastropod biosparrudite. At some exposures, the two lithologies are interbedded, and infrequently the bio- sparite is dominant. Where the unit is 1 to 2 m thick, the biosparite is commonly absent. Biomicrudite is present in areas where the Marble Hill is less than about 1 m thick, and it is commonly the dominant lithology where the unit is less than 0.5 m. At a few exposures, the Marble Hill includes thin tongues of micrit-ic limestone and calcareous mudstone of char- acteristic Rowland lithology. ENVIRONMENTS OF DEPOSITION The study area prior to the development of the Marble Hill Bed was a shallow marine shelf. The abundant and diverse fauna of the underlying Bull Fork Formation indicates an area of shallow, well- aerated water of normal salinity. The jumble of whole and coarsely broken fossils in a micritic mat- rix probably resulted from current action sufficient to produce some breaking and sorting of the fossils but too weak to Winnow out lime mud. Thus, most of D6 the Bull Fork was probably deposited below normal wave base but within the zone affected by storm waves. This environment is the shelf facies or deep undathem of Wilson (1975, p. 355), who describes the water depth as “tens of meters or even a hundred meters.” The shelf environment existed for tens of kilometers south of the study area before Rowland time and north of the study area before, during, and after Rowland deposition. The greatest southern extent of the pro-Rowland shelf is indi- cated by the southern limit of Bull Fork deposition, shown in figure 1. The dolomitic to limy mudstone of the type Row- land Member in southern central Kentucky was de- posited in very shallow areas along the landward (south) side of a marine platform. The great spars- ity of the marine fauna there indicates restricted circulation, and the dolomitic character of the sedi- mentary rocks suggests intermittent hypersaline conditions. Deposition occurred partly in intertidal areas, as evidenced by mudcracks (Weir and others, 1965, p. D17). The Rowland of the study area was deposited seaward (north) of the type Rowland, in deeper water; it represents a lagoonal area in which argillaceous micrite and calcareous mudsto-ne were deposited. The common corals and ostracodes and sparse brachiopods and byrozoans indicate an area of more normal salinity but one with only moderate circulation and limited food supply. The area sup- ported an abundant infauna, as evidenced by numer- ous burrows- in the limestone. There was little wave aetion and very little Winnowing of lime mud. Wilson (1975, p. 358) describes this type of environment as an open marine platform having a water depth from a few meters to a few tens of meters. The open marine platform on which the Rowland was de- posited covered central Kentucky and extended northward. As shown. by the superposition of the Rowland on the Bull Fork Formation (fig. 1), the platform overlapped the southern part of the shelf on which the Bull Fork was deposited. The Marble Hill Bed was deposited along a nar- row band at the northern limit of the shallow Rowland platform. Immediately to the north lay the deeper water of the open marine shelf on which the Bull Fork was being deposited. The well-sorted biosparite and the abraded biosparrudite of the Marble Hill indicate deposition in a high-energy en- vironment capable of extensive reworking. The abundance of sparry calcite cement in an area bounded on the north and south by micritic lime- stone indicates a narrow zone of high energy levels needed for thorough winnowing of carbonate sands. SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 The Marble Hill formed as a platform-edge sand where the open shelf, at depths below wave base, shoaled to the shallow depth of the Rowland plat- form. The impingement of the wave base on this shoal, probably at depths of 5 to 10 m, produced a zone of turbulence in which skeletal material gene- rated by the abundant fauna of the shelf was broken, abraded, and sorted. The environment was one of well-oxygenated water of normal salinity and good circulation, but it probably did not support a plentiful fauna because of the shifting substrate. Gastropods, a major component of the biosparrudite, are typical of organisms indigenous to this environ- ment (Wilson, 1975, p. 358). The Bardstown Member overlying the Rowland in the study area consists of the same rock types and fauna as the underlying Bull Fork Formation. The Bardstown represents a recurrence of the marine shelf environment that preceded the deposition of the Rowland. DEPOSITIONAL HISTORY Prior to the deposition of the Marble Hill Bed, north-central Kentucky was part of a open marine shelf on which the Bull Fork Formation was being deposited. This shelf extended for more than 20 km south of the study area, where it was bordered by a shallower marine platform. The deposition of the Marble Hill was initiated by a shoaling that prob- ably affected all of central Kentucky. As a result of the shoaling, the shallow, restricted-circulation en- vironment of south-central Kentucky extended northward; thus the lagoonal sediments. of the basal part of the Rowland were deposited on the southern part of the Bull Fork shelf. The shoaling seems to have been more pronounced in the study area than it was farther east, as is indicated by the extent to which the Rowland overlies the Bull Fo‘rk (fig. 1). Where the wave base impinged on the rising sea floor, an offshore bar formed along a. northeast- southwest line and deposited the basal unit of the Marble Hill (fig. 5). The bar commonly rests directly on the Bull Fork and marks the seaward edge of early Rowland deposition. The orientation of the bar suggests a prevailing wave direction from the north- west. Toward the eastern end of the area, the water may have been somewhat deeper, as the bar thinned and formed a line of isolated patches of carbonate sanrd. No evidence has been found to indicate that the bar followed the margin of the Rowland environ- ment where it swings southeast (fig. 3). No cor- responding bar deposits have been reported in the area of Rowland—Bull Fork overlap along the east- THE MARBLE HILL BED IN THE UPPER ORDOVICIAN DRAKES FORMATION, KENTUCKY 85°22'30" 85°15‘ OPEN SHELF 38° \ 33°_ 30' 1 | FIGURE 5).—Sketch map of study area showing interpretation of depositional environments. Solid arrows indicate pre- vailing wave direction on the open shelf. Dashed arrows show inferred path of tidal currents onto the platform. Border between depositional areas dotted where inferred. Area same as figure 3. ern belt of outcrop (Weir and McDowell, 1976; Weir, 1976). The deposition of the bar that forms the basal unit of the Marble Hill Bed was followed by an in- flux of terrigenous elastic sediments that formed a 1- to 2-m mudstone unit on both sides of the bar. This mudstone occurs at the base of the Rowland Member throughout the study area, except at the northern edge and where the basal unit of the Marble Hill is present. Further shallowing accom- panying this influx caused the first bar to be aban- doned and shifted bar deposition about 10 km sea- ward. A second bar formed on the basal mudstone unit, parallel to the first bar and flanked by open shelf to the north (fig. 5). This second bar persisted through the remainder of Rowland deposition and formed the main. body of the Marble Hill at the north edge of the Rowland Member. Skeletal mate— rial carried onto the second bar by incoming waves was reworked in the turbulent zone along the bar and deposited as biosparite and biosparrudite. Gas— tnopods living on the bar contributed large amounts of coarse skeletal material to the biosparrudite. The width of the main bar (fig. 3) indicates that deposi- tion was confined to a narrow zone at the seaward edge of the platform. Deposition: of this bar and the laterally equivalent lagoonal sediments generally kept pace With subsidence of the area. A rapid build- up of sediments would have caused the. Rowland to D7 prograde onto the shelf. Conversely, a too slow rate of sedimentation relative to the rate of subsidence would have resulted in an advance of the deeper shelf environment over the outer part of the lagoon. North of the second bar, interbedded b‘iomicrudite and shale of the Bull Fork Formation were deposited on the shelf. On the shallower platform to the south, partly protected from wave action by this bar, Rowland micrite and interbedded mudstone were deposited in a lagoonal environment. The fine grained calcite deposited on the platform. was prob— ably derived from the disintegration of organisms living in the lagoon and from the fine debris pro- duced by abrasion of skeletal grains on the bar. Two southward extensions of the main Marble Hill bar occur along its south edge, one along and near the Ohio River Within the Rowland Member and the other 1 to 5 km east of the Ohio at the top of the Rowland (fig. 3). These tongues, which con- sist largely of biosparrudite, may be washover fan deposits formed of material swept over the bar onto the platform during periods of storm (fig. 5). After‘the s‘econd ban“ was established, a body of carbonate sand at the east end was spread, probably by tidal currents, to the southeast along the seaward edge of the platform. The presence of corals, orachi- opods, and bryozoans in the lagoonal sediment of the Rowland indicates some exchange of water with the open shelf, but the sparseness of the fauna suggests only weak currents that brought in small amounts of food for bottom-dwelling forms. The bar re- stricted flow from the open shelf to the northwest, so that tidal exchange was concentrated in a chan- nel around the east end of the bar. The southeast- ward direction of the tidal channel from the end of the bar may have been influenced by the first Marble Hill bar. The abandoned bar probably re- mained a topographic high that had enough relief to divert the main tidal flow to the east (fig. 5). Deposition of carbonate sand continued along the tidal channel until the deposit was locally as much as 5 m thick, then this channel was abandoned in favor of channels to the west. The eastern half of the old channel returned to a normal platform en- vironment, and a tongue of Rowland micrite and mudsto-ne was laid down across the channel deposits (fig. 3). The new path of the tidal flow was still around the east end of the main bar, but there it split into two south-trending channels (fig. 5). Carbonate sand deposited in these channels formed two linear bodies that eventually extended as much as 13 km south of the northern edge of the Marble Hill Bed. D8 Both these channels are alined with low spots in the lower Marble Hill bar (fig. 3). These low spots may be the result of scouring by tidal currents, or low spots in the older bar may have controlled the path of the later tidal channels. Tidal channels persisted in this position through the remainder of Rowland deposition. The filling of the channels with biosparite and biosparrudite kept pace with the buildup of lagoonal sediments, as evidenced by the intertongu- ing and gradational nature of the contacts along the edges of the channels. In the final stages of deposition of the Rowland Member in the study area, the amount of ter‘r‘igenous elastic sediments supplied to the seaward part of the platform was reduced. As a result, the upper 1 to 2 m of the Rowland consists of micrite and only minor mudstone interbeds. Depositionof the Rowland and Marble Hill ended with the subsidence of the area that brought a return of the shelf environment and the beginning of the deposition of the Bar‘dstown Member. The sharp basal contact of the Bar‘dstown (Swadley, 1977) suggests that subsidence was rapid. Scattered lenses of carbonate sand occur at the base of the Bardstown south of the study area (Kepferle, 1977; Luft, 1977), but there is no evidence indicat- ing a uniform southward migration of a bar ahead of the advancing shelf. SUMMARY The Marble Bed of the Rowland Member of the Drakes Formation is made up of poorly sorted gas- tropod biosparrudite and well-sorted echinoderm biosparite. It was deposited as a complex of offshore bars, washover fans, and related tidal-channel de- posits in a narrow zone at the seaward edge of the shallow marine platform on which interbedded micrite and mudstone of the Rowland Member were deposited. The Bull Fork Formation, which under- lies the Rowland and is its lateral equivalent to the north, and the Bards-town Member of the Drakes Formation, which overlies the Rowland, were de— posited on an open marine shelf that bordered the Rowland platform. The Marble Hill developed in a high-energy environment, where the wave base im- pinged on the sea floor between the open shelf and SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 the shallower surface of the platform. Deposition of the Marble Hill was initiated by an uplift of the area that caused the platform environment of cen- tral Kentucky to migrate northward over the south- ern edge of the open shelf. Deplosition was ended by subsidence that brought about a return of the shelf environment. REFERENCES CITED Kepferle, R. C., 1977, Geologic map of the Ballard‘sville‘ quad- rangle, north-central Kentucky: U.S. Geological Survey Geologic Quadrangle Map GQ—13'89, scale 1:24,000. Luft, S. J., 1977, Geologic map of the Smithfield quadrangle, north-central Kentucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1371, scale 1124,000. Owen, D. D., 1859, Report of a geological reconnoissance of the State of Indiana made in the year 1837, pt. 1: In- dianapolis, John C. Walker, 63 p. Peck, J. H., 1966, Upper Ordovician formations in the Mays- ville area: U.S. Geological Survey Bulletin 1244-3, 30 p. Peterson, W. L., 1970, Bardstown Member of the Drakes ‘ Formation in central Kentucky: U.S. Geological Survey Bulletin 1294—A, p. A36—A41. Shaver, R. H., and others, 1970, Compendium of rock-unit stratigraphy in Indiana: Indiana Geological Survey Bul- letin 43, 229 p. Swadley, W C, 1977, Geologic map of the Bedford quad- rangle, north-central Kentucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1409, scale 1224,000. Swadley, W C, and Gibbons, A. B., 1976, Geologic map of the Campbellsburg quadrangle, north-central Kentucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1364, scale 1:24,000. Weir, G. W., 1976, Geologic map of the Mount Sterling quad- rangle, Montgomery County, Kentucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1335, scale 1:24,000. Weir, G. W., Greene, R. C., and Simmons, G. C., 1965, Callo- way Creek Limestone and Ashlock and Drakes Forma- tions (Upper Ordovician) in south-central Kentucky: U.S. Geological Survey Bulletin 1224—D, 36 p. Weir, G. W., and McDowell, R. C., 1976, Geologic map of the Preston quadrangle, Bath and Montgomery Counties, Kentucky: U.S. Geologic-a1 Survey Geologic Quadrangle Man GQ—1334, scale 1:24,00‘0. Weiss, M. P., and Norman, C. E., 1960, The American Upper Ordovician standard II, Development of stratigraphic classification of Ordovician rocks in the Cincinnati region: Ohio Division of Geological Survey, Information Circular no. 26, 14 p. Wilson. J. L.. 1975, Carbonate facies in geologic history: New York, Springer-Verlag, 471 p. Paleogene Sedimentary and Volcanogenic Rocks from Adak Island, Central Aleutian Islands, Alaska By JAMES R. HEIN and HUGH McLEAN SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—E CONTENTS Page Abstract _________________________________________________________________ E1 Introduction ______________________________________________________________ 1 Acknowledgments _________________________________________________________ 2 Analytical procedures ______________________________________________________ 3 Andrew Lake Formation __________________________________________________ 3 Stratigraphy and depositional environment _______________________________ 3 Secondary minerals ___________________________________________________ 7 Age _________________________________________________________________ 8 Hydrocarbon potential _________________________________________________ 9 Yakak Peninsula strata (Wedge Point area) _________________________________ 9 Stratigraphy and depositional environment _______________________________ 9 Secondary minerals ____________________________________________________ 10 Age _________________________________________________________________ 11 Hydrocarbon potential _________________________________________________ 11 Finger Bay Volcanics _____________________________________________________ 11 Metamorphism ________________________________________________________ 1 1 Age __--__________, __________________________________________________ 12 Discussion ________________________________________________________________ 12 Paleogene sedimentation _______________________________________________ 12 Source of sediment ________________________________________________ 12 Depositional basin ________________________________________________ 13 Alteration of sedimentary rocks ________________________________________ 13 Regional tectonics _____________________________________________________ 14 Concluding remarks ___________________________________________________ 14 References cited __________________________________________________________ 15 ILLUSTRATIONS Page FIGURE 1. Map showing location of outcrops of granodiorite plutons on Adak and Kagalaska Islands in the central Aleutian Islands, Alaska ______________________________________________________________________ E2 2. Map showing sample locations and topography of Andrew Lake area, northern Adak Island ___________ 3 3. Map showing sample locations and topography near Wedge Point, Yakak Peninsula, southwest Adak Island _______________________________________________________________________________________ 4 4. Composite stratigraphic sections of Andrew Lake Formation and two sections near Wedge Point ____ 5 TABLES Page TABLE 1. Description of rock samples from the Andrew Lake Formation, Finger Bay Volcanics, and Yakak Pe- ninsula, Adak Island, Alaska __________________________________________________________________ E6 2. Foraminifers from the Andrew Lake Formation ____________________________________________________ 8 3. K—Ar data and age for an andesite sill, Andrew Lake Formation _____________________________________ 9 4. Organic geochemistry of Tertiary sedimentary rocks from Adak Island _______________________________ 9 III SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS FROM ADAK ISLAND, CENTRAL ALEUTIAN ISLANDS, ALASKA By JAMES R. HEIN and HUGH MCLEAN ABSTRACT The Andrew Lake Formation on northern Adak Island, here redefined, consists of conglomerate, sandstone, chert, shale, and pyroclastic ejecta of late Eocene age. These strata were deposited in a marine basin no deeper than 500 meters. Nonmarine to shallow marine volcaniclastic rocks, probably correlative in age with the Andrew Lake Formation, crop out in the Wedge Point area of southwestern Adak Island. The sedimentary rocks contain secondary minerals including chlo- rite, vermiculite, smectite, analcime, laumontite, clinoptilolite, wairakite, and the rare zeolite yugawaralite. These minerals reflect a complex history of alteration involving burial dia- genesis, migration of hydrothermal solutions associated with intrusion of granodiorite plutons, and local thermal metamor- phism caused by intrusion of dikes and sills. The late Eocene strata both at Andrew Lake and near Wedge Point overlie the Finger Bay Volcanics, which consists of highly altered inter- bedded flows and pyroclastic rocks that contain secondary minerals (chlorit‘e, albite, actinolite, muscovite, epidote) char- acteristic of greenschist-facies metamorphism. The age ot-the Finger Bay Volcanics is unknown, but because of the contrast in degree of alteration and metamorphism between it and the overlying Andrew Lake Formation, it is believed to be late Paleocene or early Eocene. The late Eocene strata are low in total organic carbon and are therefore not considered a poten- tial source rock for petroleum. Other workers suggest that subduction of the Kula ridge spreading center beneath the Aleutian arc, 30 million years ago, resulted in an episode of regional greenschist-facies meta- morphism throughout the arc. However, absence of meta- morphism in the Andrew Lake Formation indicates either that the Kula ridge was subducted at an earlier time (about 50 my. ago) or that low-grade regional metamorphism did not accompany the subduction of the spreading center. INTRODUCTION Adak Island, part of the Andreanof group of cen— tral Aleutian Islands, consists of late Cenozoic stra- tovolcanoes that overlie uplifted Tertiary volcanic, sedimentary, and plutonic rocks (fig. 1). In general, most of the larger Aleutian islands are characterized by a Late Cretaceous(?) and early Tertiary mafic volcanic basement overlain by Paleocene volcanic and sedimentary rocks. Commonly, as on Adak Island, lower Tertiary rocks of the Aleutians are in- truded by middle Miocene plutonic bodies, mostly granodiorite (Fraser and Snyder, 1959; Marlow and others, 1973; DeLong and others, 1978). Marine volcaniclastic strata that crop out between Andrew Lake and Clam Lagoon on northern Adak Island (figs. 1 and 2) were mapped by Coats (1956) as part of the Finger Bay Volcanics and were tenta- tively dated as Paleozoic on the basis of leaflike im- pressions identified as Annulam’a stellatw. Because rocks of unequivocal Paleozoic‘age were not known from other Aleutian islands, Scholl, Green, and Mar- low (1970) re-examined these strata and found that the “Annmarie” beds are in fact of middle or late Eocene age on the basis of foraminifers, dinoflagel- lates, and vpelecypods. They named the fossiliferous strata the Andrew Lake Formation and reported that it rests depositionally on the Finger Bay Vol- canics, which they assumed to be of slightly older Tertiary age. The Finger Bay Volcanics, defined by Coats (1947), occurs widely on Adak Island and is espe- cially well exposed on southern Adak (Fraser and and Snyder, 1959) . Fraser and Snyder (1959) found that, in general, the Finger Bay Volcanics consists of pervasively altered pyroclastic deposits and ba- saltic and andesitic flows. They deduced from dated rocks exposed on nearby Kanaga Island that the volcanic rocks of southern Adak are probably of Tertiary age. They also included bedded pyroclastic rocks, volcanic wacke, and argillite as part of the Finger Bay Volcanics. One of these sedimentary rocks sections is well, exposed at Wedge Point on the Yakak Peninsula (figs. 1 and 3). E1 E2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 l77°00' 177°30' B E R I N G S E A 52°OO' —- _4 '6 \‘6 ‘6 Garmet Cove . P3)“ , KAGALASKA . ISLAND ADAK ISLAND 51°45’ — —— EXPLANATION V, Outcrop of grano— ‘\/“ diorive pluton ,~-~»_ Contact 735 0 20 . -: l J l' I KILOMETERS FIGURE 1.—Location of outcrops of granodiorite plutons on Adak and Kagalaska Islands in the central Aleutian Islands. Area of figure 2 is indicated by rectangle, Wedge Point area is shown in figure 3. Here we present information on the petrology, mineralogy, stratigraphy, and depositional environ- ments of the Andrew Lake Formation and the sedi- mentary and pyroclastic rocks at and near Wedge Point. We also assess the potential of these rocks as sources of hydrocarbons. We define mineral assem- blages formed by hydrothermal processes as distin- guished from assemblages developed by regional low—grade metamorphism. We speculate on the sig- nificance of the Paleocene history of Adak Island in the regional development of the Aleutian island arc and the early Tertiary plate-tectonic interaction of the Kula ridge with the Aleutian subduction zone. ACKNOWLEDGMENTS We thank Capt. T. P. Driver, commanding officer of US Naval Station, Adak, for permission to work on the Naval Station and Comdr. Elton Himes for providing transportation and logistic support for fieldwork. The US. Naval Special Service Corps provided logistic support for our works on Yakak Peninsula. H. N. Meeks of the National Ocean- graphic and Atmospheric Administration Observa- tory, Adak, provided a vehicle during part of our work. Paul T. Fuller was our able field assist- ant. C. E. Gutmacher provided X-ray diffracto— grams and processed a sample for K-Ar dating; G. B. Dalrymple made the K-Ar age determination. John Barron, Kristin MacDougall, Richard Poore, William Sliter, Fred May, and Louie Marincovich searched, often fruitlessly, for fossils. Kam Leong determined the Fe content of three samples by atomic absorption techniques. A. J. Koch, of Mobil Oil Corp., and George Claypool, U.S. Geo- logical Survey, provided organic geochemical anal- yses and interpretation. We benefited from review PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA E3 176°38' l76°361 l l I 2 EXPLANATION l J KILOMETERS o E g ZE $8 QTV : z :2 , 0‘ E . 8.3 Volcanic <_( 2 £5 rocks 5 3 51°57' “- E O . .2 3 g 8 Ardkrew >- o o w D m Forni'ia‘veion g ANDREW : “C’ 3 c “c’ ~ 0 m; E o o a: m 0 c ’— 3 8 g g g 3 Finger 1““ 1’: |B Boy LAKE CLAM S 5 D :3 ° Volcanics LAGOON .803»600 51°56’ 51°55’— lsoloied outcrop studied 0802-100 Measured section locality /40 Strike and dip of beds _/100\ Topographic contour, in feel ?\_./‘-\3 — Contact inferred FIGURE 2.—Sample locations and topography of Andrew Lake area, northern Adak Island. (Note that No. 803—600, for example, includes all samples from 600 through 699.) Measured sections are keyed to figure 4. Base from Coats (1956). by and helpful comments of R. R. Coats, D. W. Scholl, D. A. Swanson, M. S. Marlow, and A. K. Cooper, U.S. Geological Survey; Yotaro Seki, Sai- tama University, Japan; and A. C. Waters, Univer- sity of California at Santa Cruz. ANALYTICAL PROCEDURES Bulk rock samples and mineral separates were powdered and examined with a Norelco1 X-ray dif- fractometer. Most rock samples were cut into thin sections for textural and mineralogical study. Three samples were analyzed for iron content with an atomic absorption spectrophotometer. ANDREW LAKE FORMATION Scholl and others (1970) defined and briefly de- scribed the Andrew Lake Formation. They suggested 1Any use of trade names in this publication is for descriptive purposes only and does not constitute endorsement by the U.S. Geological Survey. that it is more than 850 m thick, although only about 40 m of section is actually exposed in quarries and low cliffs along the east shore of Andrew Lake (figs. 2 and 4). An additional 30 m of sedimentary and pyroclastic rocks crops out south of the south limit of the Andrew Lake Formation as described by Scholl and others (1970). Although this lower sec- tion is only sparsely fossiliferous and its age is a matter of conjecture, we include it as part of the Andrew Lake Formation. Even with the addition of this section, the total thickness of the formation may not be greater than 800 m (fig. 4). STRATIGRAPHY AND DEPOSITIONAL ENVIRONMENT A composite stratigraphic section of the Andrew Lake Formation (fig. 4) compiled from outcrops located on figure 2 ShOWS that the lower half of the section consists mainly of volcaniclastic sedimentary rocks. Volcanic sandstone and silty sandstone are most common, but they range from unsorted sandy E4 176°58’ SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 176°56' 0 l l l KILOMETERS CONTOUR INTERVAL 100 FEET Gonnet Cove 805—10] " Area of figure 3 51°421 — Index Map EXPLANATION 0 Measured section locality 51°41’ 81lA300- . FIGURE 3.—Samp1e locations and topography near Wedge Point, Yakak Pensinsula, southwest Adak Island. Measured sec- tions are keyed to figure 4. Note location of station 805—101 at Gannet Cove, western Adak Island (inset). Base from Fraser and Snyder (1959). conglomerate to tuffaceous mudstone. Diatoms are rare in these rocks. Intercalated devitrified ash- fall tuff attests to coeval volcanism. Numerous dikes and sills cut this part of the Andrew Lake Forma- tion. Descriptions of the samples studied are given in table 1, in stratigraphic order. Overlying‘ this relatively coarse grained elastic section are the only richly fossiliferous IOWer Ter- tiary strata known on Adak Island (fig. 4). These‘ strata mark the lower contact of the Andrew Lake Formation as defined by Scholl and others (1970). In this section are thin devitrified ash-fall tufl“ beds interbedded with quartz chert, laminated quartz por- celanite, siliceous“ shale, laminated pyritic shale, calcareous chert, and the first recognized occurrence known to us of bedded quartz chert that contains abundant diatoms. The stratigraphically highest beds in the Andrew PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA E5 ANDREW LAKE 803—401 803-403 803—402 803-302 803—301 803-303 803—206 803-200 803-204 803-201 803-203 803-205 803-202 803-207 803—21 3 803—21 I 803-209 802-501 602-602 602-403 802-40I 802-405 802-402 WEDGE POINT SECTION SOUTH OF WEDGE POINT Bil-IO] 811-401 811—302 EXPLANATION SHALE SILTSTONE SANDSTONE CONGLOMERATE, PEBBLY MU DSTONE WY? PYROCLASTIC @ CHERT AND PORCELANITE VOLCANIC FLOW OR SILL 802-301 SAMPLE NUMBER mm. m 810—406 810-403 BIO-404 810-401 810L402 810-502 810-3010 BIO-SOIb 810-302 ' BIO—202 810-20] BIO-I01 Vertical scale FIGURE 4.—Composite stratigraphic sections of Andrew Lake Formation and two sections near Wedge Point. Sections and sample localities are keyed to figures 2 and 3. Thicknesg of covered areas (breaks in sec- tions) is unknown. Top part of Andrew Lake section and upper 7 m of Wedge Point section were meas- ured. Queried line marks contact of Andrew Lake Formation with underlying Finger Bay Volcanics. (a) Chert—porcelanite section, (b) section rich with volcanic detritus, (c) Andrew Lake Formation as defined by Scholl and others (1970), and (d) Andrew Lake Formation as redefined. E6 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 TABLE 1.—Descm’pti¢m of rock samples from the Andrew Lake Formation, Finger Bay Volcanics, and Yakak Peninsula, Adak Island, Alaska [Locations and formations are keyed to figures 1, 2, and 3. Samples are listed in stratigraphic order starting at the base of the section. A, analcime; Ac, actlnolite; Al, albite; Am, amphibole; Au, augite; B, biotite; C, calcite; Ca, chalcedony; Ce, celadonite; Ch, chlorite; Cl, clinoptilolite; Cr, chert; Cz, clmozmsite; E, epidote; F, fossil debris: Fe, iron oxides; H, hypersthene; He, heulandite; Ho, hornblende; I, illite: L, laumontite; M, magnetite; Mn, Mn-dendrites; Mu, muscovite: P, plagioclase; Ph, prehnite; Pr, piemontite; Pr, pyroxene; Py, pyrite; Q, quartz; S, stilpnomelane: Sm, smec- tite; Sp, sphene; St, stilbite; T, tremolite; V, vermiculite; Vg, altered volcanic glass; VRF, volcanic rock fragments; W, wairakite; Y, yugawaralite; Z, undifferentiated zeolites. Tr., trace. He and St occur only in sample 803 —— 701] - Major primary Secondary . Sample No. Location Rock type constituents minerals Fossds Comments 802 — 402 Andrew Lake area_- Sandy pebble con- VRF, Pr, P, Q, Vg A], V, Ch, Q, Fe, L Diatoms ________ Poorly sorted glomerate. up to 9 mm. 802 — 405 _-__do _____________ Silty mudstone and fine- Q, P, VRF, Pr, Vg A), V, Ch, Q, Fe .___do __________ Burrowed, lenses, grained sandstone. contorted bedding. 802 - 401 --__do _____________ Lithic sandstone _______ VRF, P, Vg, Q, I Q, V-Ch, Ce, Cl, Ch, Barren _________ Fe, A1 802 — 403 ____do _____________ Lithic sandstone and VRF, Vg, l, P, Q Ch, L, Q, Fe, V-Ch, Rare fragments" Pods of vitric siltstone. C], S ash. lenses, contorted bedding. 802 — 602 ____do _____________ Porphyritic dike _______ P, Pr, Q V-Ch, L, Q, Fe, A] Not looked for __ 802 — 501 ____do _____________ Lithic sandstone and VRF, Vg, Pr, Q, P Q, V-Ch, Ch, L, A], Fe Barren __________ Lenses of or- sandy and silty mud- ganic matter, stone. burrowed. 802 — 303 ___.do _____________ Pebble conglomerate and Vg, VRF, Q, I, Pr L, V-Ch, Q, Fe _-__do __________ Grains compact- silty mudstone and penetrate. Con- sandstone. glomerate is well graded. Clasts to 12 mm. 802 — 301 ____do _____________ Pumiceous ash-fall tufi‘?_ Vg, P, Pr L, Q, V, Fe ____do __________ Nearly totally re- placed by laumontite. 802 — 204 ..--do _____________ Pebbly sandstone _______ VRF, P, Au, Q, Vg Q, V, Ch, L, AI, Ph, ____do __________ Very poorly S, Fe sorted clasts to 13 mm. 802 — 201 -___do _____________ Fine-grained sandstone VRF, P, Q, Pr, Ac, Q, V-Sm, L, Al, Fe -__-do __________ Sandy layers are and silty mudstone. , , h poorly sorted and crudely graded. 802 - 202 ____do _____________ Fine-grained lithic P, VRF, Au, Ac, E, Q, A1, V, Ch, Fe, L ___.d0 __________ sandstone. C, Mu 802 - 802 --._do _____________ Volcanic sill ___________ P, Q, Au, H, B Ch, V Not looked for __ K-Ar date 14 m.y. 802 — 701 ____do _____________ Volcanic flow? or sill -_ P, Pr, Q, B Q, V, Ch ____do __________ Pillowlike structures. 802 ~ 803 ____do _____________ Altered vitric tufl' ______ P, Vg, Q V, Sm, Q Barren __ _ 802 - 804 Volcanic siltstone __ _ Q, P, Vg Q, Z, Ch, V, C, Fe Fragments _ _ Pods of tuif. 802 — 103 Tufi'aceous mudstone _ _ Q, P, Vg, Pr, B Q, C, Ch Diatoms and radiolarians. 802 — 101 Tufi" ___________________ P, Vg, Pr, Q V-Ch, Fe, A], Q, Ch Benthic Completely al- foraminifers. tered. 14 cm below dike. 802 — 102 Volcanic dike _____ P, Pr, Q V, Ch, Q Not looked for _. 803 — 209 Black chert __ Q Q, Tr, Z Dinoflagellate ___ Mldcl: organic e rxs. 803 — 211 Vitric tuff ______ Vg, P Sm, Cl Friable. 803 -— 213 __-_do _________________ Vg. P, Pr Sm-V, Ce, Fe, Q Do. 803 — 207 Chert and porcelanite __ Q, P, I, F Q, Fe, Ch Late Eocene, upper bathyal. 803 — 202 Calcareous chert _______ Q, P, F Q, C Diatoms, « radiolarians. 803 — 205 Chert and porcelanite __ Q, P, F Q, Py, Mn Microfossil molds- Burrowed, laminated. 803 — 203 ____do _________________ Q, P, F Q, Py, Mn ____do .......... Laminated. 803 — 201 ___- .-_ Q, P, F Q, Py, Mn ____do __________ 803 — 204 Siliceous shale _________ 2’ ? Benthic foram- Cretaceous inifers. through Oligocene. 803 - 200 -_--do _____________ Pyritic silty shale ______ Q, P, I, F Q, C, Py, Ce, V See table 2 ...... Laminated, late Eocene, upper bathyal. 803 - 206 ___-do _____________ Black silty shale _______ ? '! Barren __________ 808 — 303 ____do _____________ Diatom chert and sili- Q, P, F, Am, VRF Q, V-Ch, C, Fe Diatoms, Vermiculite clay ceous shale with chert radiolarians, nodules, shale nodules. foraminifers. is laminated. 803 — 301 --__do ............. _-._do _________________ Q, P, F, VRF Q, C, V-Ch, Fe Diatoms, Laminated shale, radiolarians, hematite foraminifers, pseudomorphs and fish. after pyrite. Lake Formation described by Scholl and others (1970) include aquagene tuff, ash-fall tufi‘ and pos- sibly ash-flow tuff, and volcanic dikes (fig. 4). These pyroclastic rocks differ texturally and mineralogic- ally (table 1) from strata in the lower part of the formation. Mild alteration of framework grains and traces of sideromelane(?) that has not devitrified contrast with the greater alteration shown by under- lying rocks and suggest that this uppermost section here. may be younger than the underlying, more altered (part of the section. Therefore, both the lower and upper contacts of this formation as originally pro- posed by Scholl and others (1970) are redefined We redefine the base of the Andrew Lake Forma- tion exposed along the southeast shore of Andrew Lake (fig. 2) to include the conglomerate at locality 802—400 (figs. 2 and 4); the underlying silicified PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA E7 TABLE 1,—Desc'r'iption of rock samples from the Andrew Lake Formation, Finger Bay Volcanics, and Yakak Peninsula, Adak Island, Alaska—Continued Sample No. Location Rock type Mzdgztiltdiernlzsw Slicigxe‘stfigy Fossils Comments 803 -— 302 Andrew Lake area __ Volcanic dike ......... A1,Ch, V Not looked for __ P, Pr V8, P, VRF, Pr 803 — 403 —..-do ————————————— Aquagene tul'f _________ Q, W, V, Ce, rare C ____do __________ Unaltered min- erals and some glass (?) . 803 — 403 __--do ............. Ash-flow tuft __________ Vg, VRF, P, B, Q, W, Sm, Cl, Ch, C, V _--_do __________ Unaltered min- Au, Cz erals. 803 - 401 ---_do ............. Ash-fall tuff ........... Vg, P, B, Am Q, Cl, W, V, Sm Barren __________ Unaltered min— erals and some glass (7) . 803 - 601 --_-do _____________ Clayey siltstone ________ Q, P, F Ch, V-Ch, Fe, Q Forfilminisers, ec moi . 803 - 602 ____do _____________ Lithic sandstone _______ P, Q, VRF V-Ch, C, Q. Ch, Al Barren __________ Baked by nearby dike. 803 — 604 ____do _____________ -___do _________________ P, VRF, Q V-Ch, C, Q. Ch, Al, Fe _-__do __________ Do. 810 — 101 South of Wedge Welded tufi ____________ Vg, P, Pr Y, Ch, Ce. Fe, I-Mu, L ____d0 __________ Purple with green Point. mottling. 810 - 201 ___-do _____________ Volcanic dike __________ P, Pr, Vg Vesicles: Q. S-Ce— Not looked for __ Ch, L or C; Matrix: Fe. C. Q 810 — 202 ____do _____________ Lithic sandstone _______ VRF, P, Pr Y, Q, Al, Ch, Fe Barren __________ Poorly sorted, compact, grains penetrate. 810 — 302 --__do _____________ Tufi ___________________ ? ? ____do __________ 810 - 301a ____do _____________ Pebbly mudstone _______ Vg, P. Au, VRF Y,LQ, Fe, Ch, V. Tr, ____do __________ Lahar. 810 - 301b ____do _____________ Volcanic flow __________ P, Au Vesicles: W, S, L: Not looked for ._ Matrix: Al, Fe, I 810 — 502 __-.do _____________ Sandy shale ____________ '! ? Barren __________ 810 — 402 ..___do _____________ -_-- o ________ ? '! ____ o __________ 810 — 401 __--do _____________ Ash-flow tufi ___________ Vg, Au, P 1.51211, ¥,Ee. 0, Fe, Not looked for __ Slightly welded. r. . 810 — 404 _..__do _____________ Pebbly mudstone _______ Vg, P, W An, Q, Ch, Ce, Y(?) Barren __________ Pumiceous lahar. 810 — 403 ___-do _____________ Lithic sandstone _______ VRF, P, Pr An, Fe, clays ____do __________ Banded, well sorted, crudely graded. 810 - 406 ____do _____________ ____do _________________ VRF, P. Pr, Q Z, Q, Ch, Fe _-__do __________ Poorly sorted, graded. 812 — 104 ____do _____________ Silicified ash ___________ Vg, P Q. Ch, Fe, C Spicules? ________ 811 — 302 Wedge Point _______ Pebbly mudstnne _______ ? ‘.’ Barren __________ 811 - 401 _-__do _____________ Volcanic dike __________ Au, P, Sp, M Vesicles: Q, Not looked for __ More than 50 per- Ch+S+Ce, L, Q, cent altered to Tr, Pr, E: Matrix: quartz and Q, L, Al, E, Fe, Ch, I laumontite. 811 — 2013 ____do _____________ Pebbly mudstone _______ Vg. P. Pr, M Q, C, Ch, S, Ce, Fe, A Barren __________ Pumiceous. 811 — 201b __-_do _____________ Volcanic dike __________ P. Au, M Vesicles: L, Not looked for __ S+Ch+L. Q: Matrix: Q, Fe, Ch 811 — 101 .-__do _____________ Silty sandstone _________ Vg, P, M. Au, Q Q, Fe, Ce, I—Mu, Barren __________ Poorly sorted, Y, L, A pumiceous. 811 — 201 ____do _____________ Volcanic wacke ________ VRF, P, Pr, Q A], Q, Y, Fe, L, Ch ____do __________ 803 — 101 Southeast of Andrew Lithic sandstone _______ P, Q, B, VFR, Ac-T, E, Q, Fe, Ch, V __-_do __________ Finger Bay Vol- Lake. Ho, Cr canics, com- pact. 805 — 101 Gannet Cove _______ Quartzite and epidosite__ Q, P, Cg, F, B E, C, Q, Mu, Fe, Ch, Ghosts of micro— Finger Bay Vol- Pi? fossils. canics, poorly sorted. 803 — 701 Clam Lagoon _______ Hydrothermal vein ..... W, Ca, V, 01, He, St Not looked for ______ Not looked for __ Vein cuts Finger Bay Volcanics. volcanic rocks are assigned to the Finger Bay Vol- canics. The top of the formation is in the area covered by tundra between outcrop localities 803— 300 and 803—400 (figs. 2 and 4). Fossils (table 2) indicate that the strata of the redefined Andrew Lake Formation were deposited in a marine environment, probably at water depths between 200 and 500 m. Sediments were reworked by bottom currents and by infauna. Lenses con- taining diatoms or devitrified ash are locally abund- ant; however, laminated porcelanite, black pyritic shale, and laminated shale from the middle part of the section suggest that the depositional basin was for a time “starved” or cut off from active terri- genous sedimentation. Consequently, mainly biogenic material accumulated, although an active infauna was not present. Other sedimentary structures and current-direc- tion indicators are rare. They include ripple lamina- tions and poorly developed graded bedding. Sand- stone beds are graded and have flute casts, and sand- stone and shale sequences are rythmically bedded, which suggests deposition by turbidity currents. Rare crossbedding indicates eastward current flow. Deposition of most of the Andrew Lake Formation is ascribed to a combination of biogenic-pelagic, turbidity-current, hemipelagic, and pyroclastic processes. SECONDARY MINERALS Most rocks of the Andrew Lake Formation are silicified and in part altered to clay minerals and iron oxides (table 1). In most samples, plagioclase of intermediate anorthite content is partly altered to albite. In the lower part of the section where vol- canogenic sedimentary rocks, ‘pyroclastic debris, and volcanic dikes and sills are abundant, quartz and E8 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 TABLE 2.—Foraminifers from the Andrew Lake Formation [Fossils identified by Kristin McDougall, R. Z. Poore, and W. V. Sliter] Sample No. 802—101 803—200 Benthic forums Other microfossils Age range Depth of water Bathysiphon sp. _________________________________________________ Cyclammina pacified? ______________________ Echinoid spines ______ Cyclammina sp. ___________________________ Radiolarians Dentalina sp. Epom'des sp. _______________________________________________________________________ Gyroidimz soldam'i d’Orbigny ________________________________________________________ Lenticulina sp ______________________________________________________________________ Valvulineria sp.? ___________________________________________________________________ Bathysiphon eocem'cus ____________; _______ Echinoid spines ______ Late Eocene _____ Bathysipho'n sp ____________________________ Fish debris ____________________________ Cyclammimz pacifica Beck __________________ Radiolarians Dentalimz sp. Eponides sp. _______________________________________________________________________ Gyroidina soldam'i d’Orbig'ny ________________________________________________________ Lenticuh‘na sp. _____________________________________________________________________ Melom‘s umbilicatulus (Montague) ___________________________________________________ Cyclammina sp. ___________________________ Indeterminate _ _ _ Late Eocene _____ Upper bathyal. 803—207 Upper bathyal. 803—601 803—204 Bathysiphon eocenicus ? Rhabdammina eocem’ca? laumontite occur as interstitial cement and replace volcanic debris along with vermiculite, chlorite, smectite2 and vermiculite-chlorite (randomly inter- layered). Some volcanic rock fragments are almost completely altered to iron oxides and clays. Locally calcite, analcime, clinoptilolite, and stilpnomelane replace volcanic debris. Three samples (802—202, —201, —204 in fig. 4) from the upper part of this section rich in volcanic detritus contain abundant actinolite, epidote, chlo- rite, and calcite. These minerals, commonly found in greenschist-facies metamorphic rocks, occur as de- trital minerals in these three samples (table 1). Ad- jacent to dikes and sills, however, secondary preh- nite and epidote have formed (for example, prehnite in sample 802—204, table 1). Because minerologically unstable volcanic debris is less abundant, fewer secondary minerals char- acterize the chert-porcelanite part of the section. Quartz and minor calcite and pyrite are the most important secondary minerals. Quartz indiscrimin- ately replaces most sediment components and fills all available void space; thus abundant siliceous shale and porcelanite are produced. Calcite replaces quartz and is therefore a relatively late stage mineral. Ash-fall tufi" in the Andrew Lake Formation has altered to smectite and minor cl-inoptilolite. In places smectite was subsequently converted to vermiculite. ’Smectite is the internationally accepted group name for the clay minerals that include montmorillonite, nontronite,“ and saponite (Brindley and Pedro, 1975). It is used as a general term. Echinoid spines ______ Intermediate ___ _ Cretaceous through Oligocene. Locally vermiculite, chlorite, hematite, and celado- nite replace volcanic detritus (table 1). Pyroclastic rocks that overlie the Andrew Lake Formation as redefined herein contain unaltered framework grains, but most of the glassy volcanic material is replaced by wairakite and to a lesser extent by clinoptilolite, quartz, and celadonite. Ver- miculite, smectite, chlorite, and iron oxides are present; calcite is rare (table 1). Plagioclase is of intermediate composition. Celadonite in this part of the section is blue, whereas lower in the section it is green. Secondary mineral assemblages or mineral facies in the stratigraphic section can indicate the history of burial metamorphism of the rocks. Accordingly, quartz, vermiculite, vermiculite-chlorite, smectite, chlorite, iron oxides, and celadonite are ubiquitous. Clinoptilolite is present in places but is most com- mon at the top. Stilpnomelane and laumontite are in the lower part; wairakite occurs near the top of the section. Calcite and pyrite are mainly at midsection (table 1). AGE The Andrew Lake Formation was assigned a middle or late Eocene age by Scholl and others (1970) on the basis of microfossils and megafossils. Our collections of benthic foraminifers (table 2) indicate that the Andrew Lake Formation was de- posited during the 'late Eocene, virtually identical to the foraminiferal age assigned in Scholl and others PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA TABLE 3.—K—Ar data and age for an andesite sill, Andrew Lake Formation [Potassium measurements by A. Berry; argon measurement and age calculation by A. Berry and E. Sims] “Anna Calculated . Percent mArum age Sample No. Mineral K20 (moles/g) —___4“Armmi (millions of years) 802 —- 802 Plagioclase 0.079 1.640><10‘12 0.04 14.4i3.5 40K decay constants: )xs : 0.572 X 10‘10/yr, A! : 8.73 X 10—13/yr, M : 4.905 X 10—10/yr. Abundance ratio 40K/K : 1.167 x 10*4 percent atomic. (1970). According to Berggren (1972), late Eocene represents an absolute age of 37.5 to 43.0 m.y. A K-Ar date on fresh plagioclase from an andesite sill cutting the Andrew Lake Formation was 14.4:35 m.y. (table 3). We speculate that the pyroclastic rocks that over- lie the Andrew Lake Formation as redefined, but in- cluded in the formation by Scholl and others (1970), are part of a younger series of volcanic rocks. These strata are structurally concordant with the lower part of the formation, but the unaltered framework grains and only mild alteration of glass shards differ markedly from the relatively high degree of altera- tion of samples from only slightly lower stratigraph- ically. Similar mild alteration is typical of other Neogene volcanic rocks on Adak; for example, the sill (802—802) cutting the Andrew Lake Formation dated by the K-Ar method as middle Miocene (table 3) . Perhaps these little-altered pyroclastic rocks are associated with the intrusion of Miocene granodio- rite plutons and related dikes and pyroclastic de- posits (Fraser and Snyder, 1959). However, erup- tion of the pyroclastic rocks at any time after the Eocene cannot be ruled out. HYDROCARBON POTENTIAL Two samples of shaly siltstone were analyzed for TOC (total organic carbon), EOM (chloroform ex- tractable bitumen), and R0 (v‘itrinite reflectance). These quantities, as well as EOM/TOO, are listed in table 4. Sample 803—204 (fig. 4 and tables 1 and 4) contains 0.41 weight percent of TOC, the highest value recorded for any rocks on Adak Island. This value, however, is still below the 0.50 weight percent quantity generally considered to separate a possible source rock from one With no source-rock potential. The Ro value of 2.1+ indicates that sample 802—804 has been heated well beyond the level of crude oil stability (A. J. Koch, written commun., 1976). A low EOM/TOG ratio can be interpreted as resulting from the “cracking” of organic material; the break- up of organic compounds probably resulted from the E9 TABLE 4.—-—Organic geochemistry of Tertiary sedimentary rocks from Adah Island [Analyses of total organic carbon (TOC) by U. S. Geo]. Survey Organic Geochemistry Laboratory, Lakewood, Colo. Analyses of chloroform ex— tractable bitumen (EOM), vitrinite reflectance (Ru), and EOM/TOC by Mobil Oil Corp., Dallas, Tex.] Sample L a - TQC EOM ” EOM/ 803 — 204 Andrew Lake 0.41 28.6 2.1+ 0.70 803 — 206 __do _______ .04 28.7 7.10 810 — 302 Wedge Point .05 44.3 * 8.90 810 — 402 __do _______ .09 40.2 * 4.50 810 — 502 __do _______ .07 29.3 * 4.20 811 — 302 __do _______ .04 26.4 * 6.60 * Insufficient organic carbon for vitrinite reflectance measurement. heat produced by nearby intrusions such as the numerous dikes and sills observed in outcrop. Sample 803—206 has a very low TOC content, but the EOM/TOIC ratio indicates that it has not been subjected to excessive heat. A low TOC content means that there probably never was a significant quantity of organic material present. Attempts to recover organic residue for measurement of R0 were unsuccessful. YAKAK PENINSULA STRATA (WEDGE POINT AREA) Fraser and Snyder (1959) mapped a section of sedimentary, volcanic, and volcaniclastic strata on the west side of Yakak Peninsula (fig. 3) as part of the Finger Bay Volcanics. On the basis of lithologic composition, degree of alteration and induration, and the types of secondary minerals present (see below), we propose that these strata are temporally equivalent to the Andrew Lake Formation. Study of the mineralogy and petrology of samples from two stratigraphic sections, one immediately south of and one at Wedge Point (fig. 3), complements the reconnaissance work done by Fraser and Snyder. STRATIGRAPHY AND DEPOSITIONAL ENVIRONMENT The section south of Wedge Point (figs. 3 and 4) consists predominantly of interbedded sandstone, shale, and pebbly mudstone, with minor ash-flow and ash-fall tuff (including slightly welded tufi'); v01- canic dikes cut the section. Sandstone beds are pri- marily lithic arenite but include lithic to feldspathic arenite and wacke. Virtually all lithic grains are volcanic rock fragments, generally subrounded. Plagioclase, pyroxene, and locally quartz are other common framework grains. Poorly sorted rocks abound, but graded, layered, and well-sorted beds occur near the top of the section. Sandstone is com- E10 monly cemented by laumontite, yugawaralite,3 iron oxides, or clays. Pebbly mudstone beds contain mostly subangular grains, although a complete range of grain round- ness is present. Again, the framework grains are dominantly volcanic rock fragments and, in some samples, are primarily pumiceous. These poorly sorted rocks are texturally like mudflows and are probably lahars. The mudflow units are as much as several tens of meters thick, and they include rock fragments up to 1 m in diameter. Ash-flow tuff is purplish to gray, commonly with a green mottled surface reflecting replaced volcanic glass globules (replaced collapsed pumice 1apilli?). Texturally, it appears massive to weakly flow banded. Sample 810—101 (figs. 3 and 4; table 1) is mostly glass globules and flattened pumice frag- ments (greater than 3 mm) replaced by yugawara- lite. Sample 810—401 has pumice with an elongation ratio of 20 : 1. Large plagioclase and augite glomero- crysts (Carlisle, 1963, p. 58) are present. The groundmass appears to be collapsed pumice and glass shards replaced by illite or muscovite. The section includes many porphyritic volcanic dikes altered to zeolites, quartz, and iron oxides. Ves- icles and amygdules are filled with zeolites, quartz, and clays. Phenocrysts are mostly plagioclase, partly or wholly altered to albite, and pyroxene that is locally fresh. The section at Wedge Point (figs. 3 and 4) is similar to the one to the south just described but is capped by 7 m of alternating sandstone (volcanic wacke) and shale; two pebbly sandstone beds occur in this section. Wacke makes up 93 percent of this 7-m section and consists of beds 3 to 130 cm thick (average 47 cm), whereas the 7 percent of shale consists of beds 2 to 12 cm thick (average 5.4 cm). In general, beds increase in thickness up section. Framework grains are mostly volcanic fragments, plagioclase, pyroxene, and altered volcanic glass. Samples rich in relict volcanic glass also contain abundant zeolites. The matrix is made up mostly of clays, iron oxides, and fine-grained counterparts of framework grains. Relative to the section south of Wedge Point, in- trusive bodies at Wedge Point are more highly altered. From 50 to 75 percent of the host volcanic rock may be altered to zeolites, quartz, and clays. The presence of ash-flow tufi‘ and boulder lahars and the apparent absence of microfossils suggest 3Yugawaralite, a rare calcium zeolite, has been reported from only one other locality in North America, near Fairbanks, Alaska (Eberlein and others, 1971). It has also been reported from Heinabergsjokull, Iceland (Barrer and Mamhall, 1966) and from Japan (Sakurai and Hayashi, I952; Seki and Okumura, 1968; Sameshima, 1969; Seki and others, 1969). SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 that the rocks at Wedge Point were deposited in a subaerial to shallow marine environment and that volcanism was active near the site of deposition. Fiske (1963) and Fiske and Matsuda (1964) have demonstrated that submarine ash-flow tufl“ is not welded. Rocks at Wedge Point must therefore be, in part, subaerial deposits. Further, the presence of graded lithic arenite and sequences of alternating lithic wacke and shale suggest turbidity-current dep- osition; if so, Wedge Point rocks are in part sub- aqueous deposits (probably shallow marine). No freshwater fossils or lacustrine deposits were found at Wedge Point. SECONDARY MINERALS Porphyritic basaltic and andesitic rock fragments are altered to iron oxides (mostly hematite) and clay minerals (dominantly chlorite and illite). Plagi- oclase is altered to albite that has approximately parallel extinction and little or no relict zoning. In contrast, clinopyroxene is commonly unaltered. Much pumice is replaced by vermiculite and chlorite. Other glassy volcanic fragments are replaced domi- nantly by zeolites such as yugawaralite, analcime, and laumontite, and to a lesser extent by quartz, illite, chorite, celadonite, hematite, vermiculite, and smectite. The rare zeolite, yugawaralite, is the most common zeolite in these rocks on Yakak Peninsula. Sandstone is cemented by analcime, yugawaralite, laumontite, chlorite, illite, and hematite. Thin (0.005—0.1 millimeter) clay films coat all grains in some beds of arenite. The framework grains are compact and penetrate adjacent grains, and the re- maining pore spaces (commonly very small, maxi- mum 0.4 mm2 in cross section between grains) are filled with a zeolite such as analcime in sample 810— 403 (figs. 3 and 4; table 1). These observations mean that the grains acquired their clay rim after deposition and were subsequently compacted (prob— ably by deep burial) before cementation by zeolite. The zeolite cement is a relatively late occurrence. Galloway (1974) suggested that these diagenetic changes could occur after 300 to 1200 m of burial. He showed that the surface coatings of authigenic clay were the result of mobilization of silica and aluminum from the volcanic debris. The groundmass of dikes and sills is replaced by quartz, laumontite, analcime, stilpnomelane, albite, iron oxides, chlorite, illite, and calcite; calcite is the latest mineral. Minerals found in vesicles and amyg- dules suggest the following paragenetic sequence: quartz (occasionally replaced by laumontite), granular material (unidentifiable), analcime, phyl- PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA losilicates (stilpnomelane, chloride, celadonite, rare- ly vermiculite or hematite), laumontite, phyllosili- cates again (minor), and rarely quartz, calcite, pyrite, or epidote. The filling of an amygdule may begin any place in this sequence, and four or five different minerals may occur in one amygdule. The granular material listed above is a very fine grained, high-relief, highly birefringent mineral that may be calcite or epidote. It occurs as a thin and at places discontinuous band separating the first and second minerals formed in the vesicles. Laumontite and terminated quartz crystals more than 2 cm long suggest that these minerals formed from hydro- thermal solutions. AGE No fossils were recovered from the strata on Yakak Peninsula. From the lithologic composition, and degree of alteration and induration, we infer that these strata are probably temporal equivalents of the Andrew Lake Formation. HYDROCARBON POTENTIAL Four samples from the Wedge Point area (table 4) show consistently low values of total organic carbon and are not considered to be potential source rocks for petroleum. No organic residue for deter- mination of vitrinite reflectance was removed from any of these samples. FINGER BAY VOLCANICS Coats (1956) and Fraser and Snyder (1959) de- scribed the Finger Bay Volcanics as pervasively altered to chlorite, albite, epidote, and silica. We examined two samples (803—101 and 805—101) of sandstone from the Finger Bay Volcanics (fig. 2; fig. 3, inset) for comparison with sedimentary rocks of the Andrew Lake Formation and of Yakak Peninsula. Sample 803—101, collected from a quarry southeast of Andrew Lake, is a compact, poorly sorted feld— spathic arenite. The main framework grain is plagio- clase with accessory quartz, biotite, volcanic rock fragments, and chert. Silica and less abundant clays cement the rock. Abundant epidote, quartz, and actinolite-tremolite replace pyroxene(?) and feld- spar grains and fill veins (table 1). Plagioclase is altered to albite. Chlorite, vermiculite, and hematite are less abundant secondary minerals. Quartz sandstone, quartzite, and minor epidosite are found at Gannet Cove on the west coast of Adak Island (sample 805—101, fig. 3, inset). Quartz, epi- dote, and muscovite are secondary minerals that now E11 compose the bulk of the rocks. Piemontite, calcite, chlorite, iron oxides, and illite are minor secondary minerals. Granular quartz and deeply corroded and replaced plagioclase are probably the only primary grains remaining. Faint structures are reminiscent of glass shards, collapsed pumice, and microfossils. Rare chlorite spherulites occur. Although the Finger Bay Volcanics is highly altered, it is not penetratively deformed. Open folds with dips gen- erally less than 40° occur (Fraser and Snyder, 1959). METAMORPHISM Secondary mineral assemblages (Fraser and Snyder, 1959) suggest that the Finger Bay Vol- canics was subjected to regional greenschist-facies metamorphism. The diagnostic mineral assemblages range from actinolite- or tremolite-epidote-chlorite- albite-quartz to epidote-quartz-muscovite-chlorite- albite (table 1). Prehnite and pumpellyite mineral assemblages, indicative of lower temperature grades than actinolite-greenschist facies (Coombs, 1953; Seki, 1969, Coombs and others, 1970) , appear in the Finger Bay Volcanics (Fraser and Snyder, 1959), although the reconaissance nature of the work by Fraser and Snyder precludes delineation of a co- herent regional pattern of metamorphic facies. Cer- tainly, the actinolite and epidote greenschist assem- blages appear to be dominant on Adak Island. It is not clear whether emplacement of grano- diorite plutons contributed significantly to meta- morphism (contact metamorphism) of the Finger Bay Volcanics or whether metamorphism was domi— nated by a regional thermal event. Fraser and Snyder (1959) described only a thin zone of con- tact-metamorphic hornfels adjacent to the plutons. It is worth noting, however, that although some outcrops of the Finger Bay Volcanics and the Andrew Lake Formation are equidistant from ex- posed plutonic rocks, these formations show signifi- cant differences in metamorphic mineral assem- blages. More fieldwork is needed to fully distinguish regional patterns from contact metamorphism. In contrast to the Finger Bay Volcanics, sedimen— tary and pyroclastic rocks on Yakak Peninsula and the Andrew Lake Formation at Andrew Lake have not been subjected to greenschist- or even zeolite- facies regional metamorphism. These rocks have been moderately altered by low-temperature super- gene and hydrothermal processes of the zeolite facies and nowhere show greenschist-facies metamorphism. Thermal metamorphism associated with emplace- ment of dikes and sills, together with the supergene E12 and hypogene mineralization, has created a complex milieu of secondary minerals. AGE Except for sedimentary rocks of the Andrew Lake Formation, Paleogene rocks of Adak Island are ap- parently devoid of fossils. Fossil findings indicate that the Andrew Lake Formation accumulated dur- ing the late Eocene (37.5—43 m.y. ago; Scholl and others, 1970; table 2). The Finger Bay Volcanics is estimated to be as old as the initial formation of the Aleutian ridge and no younger than the overlying Andrew Lake Formation. The Finger Bay Volcanics therefore formed sometime before the late Eocene (before about 40 m.y. ago) but probably after latest Cretaceous (about 65 m.y. ago) (Marlow and others, 1973; Scholl and others, 1975) . The Finger Bay Vol- canics and associated sedimentary rocks evolved through a sequence of deposition, burial, regional greenschist-facies metamorphism, uplift, and ero- sion before the Andrew Lake Formation was de- posited. We therefore favor an age representative of the older part of this (40—65 m.y.) timespan, perhaps late Paleocene or early Eocene (about 50 m.y. ago), but rocks may be as old as early Paleo- cene (60 m.y.). Certainly the episode of regional metamorphism must have ended at least 45 m.y. ago. Unfortunately, it may not be possible to obtain reliable radiometric dates from the Finger Bay Vol- canics because of the thermal effects associated with the intrusion of plutonic rocks. The granodiorite on adjacent Kagalaska Island is Miocene (dated as 13.2 and 13.7 m.y.; Marlow and others, 1973; DeLong and others, 1978), approximately the same age as an andesite sill (table 3) cutting the Andrew Lake Formation, and is probably the same age as plutons on Adak. Available published data (Fraser and Barnett, 1959; Powers and others, 1960; Lewis and others, 1960; Carr and others, 1970 and 1971; Gates and others, 1971) suggest that the oldest exposed rocks on the western Aleutian Islands (Attu, Agattu, Shemya, Amchitka, Rat, Amatignak, Ulak, Tanaga, Kanaga) have undergone variable but mild altera- tion. Variability of alteration, presence of fresh calcic plagioclase, only weakly altered volcanic glass, and the occurrence of a variety of temperature-sen- sitive zeolites argue against regional greenschist- facies metamorphisms of the exposed rocks on these islands. Accordingly, the Finger Bay Volcanics on Adak Island appears to be unique among the rocks that crop out on the western Aleutian Islands and may represent the oldest rocks described to date SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 from these areas. Alternatively, the basement rocks of other Aleutian islands may be the same age as, but were not as deeply buried as the Finger Bay Volcanics on Adak Island. DISCUSSION PALEOGENE SEDIMENTATION Fraser and Snyder (1959) estimated that the ex- posed section of the Finger Bay Volcanics includes about 70 percent pyroclastic, 20 percent flow, and 10 percent sedimentary rocks and that most of this 2400-m-thick section was deposited in a marine en- vironment. We infer that this occurred in the late Paleocene or early Eocene. Coats (1956) provided evidence for a minimum thickness of about 600 m and speculated that the maximum is 4600 m. These observations suggest that the Finger Bay Volcanics is part of the initial series rocks, the Aleutian ridge basement complex (See Jake and White, 1969; Mitchell and Bell, 1973; Marlow and others, 1973.) SOURCE OF SEDIMENT The Finger Bay Volcanics was deposited and metamorphosed before deposition of the next recog- nizably younger strata, the Andrew Lake Forma- tion. Probably by middle Eocene time, the growing Aleutian ridge had nearly reached sea level, and subaerial volcanoes contributed debris to surround- ing basins. At times, the Finger Bay Volcanics may have contributed sediment to the Andrew Lake Formation, as clasts in some samples (samples 802— 201, 802-202, table 1) are lithologically similar, but the overall amount of sediment derived from the Finger Bay Volcanics appears to be relatively small. The main source of Andrew Lake detritus was most likely contemporanous volcanism. The host volcanic centers were eventually deeply eroded and perhaps in part covered by younger debris. Because hydrothermal silicification of the Andrew Lake Formation was intense, it is difiicult to iden- tify the origin of the silica in the chert beds. It is not clear whether the silica in the chert-porcelanite sec- tion of the Andrew Lake Formation is released by the dissolution of siliceous biogenic debris, deposi- tion from hydrothermal solutions, or both. Diatoms and minor radiolarians from the quartz chert occur in all states of preservation, from ghosts to speci- mens that retain frustule ornamentation. This sug- gests that at least part of the silica was derived from dissolution of siliceous microfossils. The cal- cite in this section is probably redeposited carbonate PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA released when foraminifers were replaced by quartz. DEPOSITIONAL BASIN The dimensions of the depositional basin of the Andrew Lake Formation are not known. Scholl and others (1970) speculated that the Andrew Lake strata accumulated in a fairly deep (500 m) basin along an early Tertiary Aleutian ridge. If the Yakak Peninsula strata (Wedge Point) are tem- porally equivalent to the Andrew Lake Formation, then they possibly represent the subaerial and shal- low-marine facies of the deeper water Andrew Lake strata. The basin seems to have been no deeper than 500 m (fossils suggest 200—500 m), a situation very much like the present Aleutian Islands and the adjacent 200-m-deep Aleutian ridge platform. We infer that these rocks were deposited on the sub- aerial flanks of a volcanic complex and in adjacent offshore shelf and slope environments. The presence of a small enclosed basin, one isolated from turbid- ity-current deposition, is evident in the lami- nated chert and porcelanite [of the Andrew Lake Formation. ALTERATION OF SEDIMENTARY ROCKS Burial diagenesis, local thermal metamorphism by dikes and sills, and hydrothermal activity contrib- uted to the alteration of Eocene rocks. Burial diag- enesis was an important process in the early stages of alteration of these rocks, primarily because they contain a large fraction of unstable mafic to inter- mediate volcanic rock fragments. Burial caused the transformation of the glassy parts of volcanic rock fragments to clays. These structurally weak rock fragments, upon further burial, decomposed to form a sedimentary rock consisting of framework grains of plagioclase, pyroxene, and rock fragments in a clay matrix. All the original glass shards and glass globules were altered or replaced during burial. After uplift and some erosion, two additional stages of alteration strongly affected the character of the sedimentary rocks: 1. Numerous late Tertiary dikes and sills intruded and thermally metamorphosed adjacent wall- rock. Sedimentary rocks adjacent to dikes were baked, and locally epidote and prehnite formed. More commonly chlorite, hematite, quartz, and vermiculite mixed-layer clay phases formed next to dikes. 2. A more significant episode of alteration occurred in conjunction with the intrusion of Miocene plutonic rocks, when extensive formation of zeolites occurred. E13 Many observations favor hydrothermal fluids rather than zeolite-grade regional or burial metamorphism as the mechanism of alteration of the Andrew Lake and the Yakak Peninsula rocks: 1. At Andrew Lake, wairakite, the highest tempera- ture zeolite, stratigraphically overlies laumon- tite and clinoptilolite, minerals characteristic of a relatively lower temperature metamorphic facies (Coombs, 1961; Harada, 1969; Seki and others, 1969). 2. Nonequilibrium mineral assemblages are com- mon; for example, smectite is associated with laumontite. Low-temperature zeolites occur in close association with higher-temperature forms; for example, laumontite, clinoptilolite, and wairakite at Andrew Lake and laumon- tite, analcime, and yugawaralite at Wedge Point. (See Coombs and others, 1959; Seki, 1969; Kossovskaya, 1975.) 3. There is no apparent stratigraphic or spatial variation in the metamorphic grade of zeolites. The distribution of zeolites does not show zonal relations. 4. A wide variety of secondary minerals is asso— ciated with the zeolites. 5. Calcium zeolites (yugawaralite, laumontite, and wairakite) greatly predominate over sodium varieties (analcime; Kossovskaya, 1975). 6. Some crystals of quartz and laumontite are more than 2 cm long. Less diagnostic but supporting evidence is that (1) ubiquitous quartz silicification suggests deposition from circulating hydrothermal fluids (Fournier, 1973; Coombs, and others, 1959), (2) mixed-layer clays, for example vermiculite-smectite—chlorite in our samples, commonly form in hydrothermal de- posits (Bundy and Murray, 1959; Lovering and Shepard, 1960; Heystek, 1963; Steiner, 1968), (3) secondary mineral assemblages (except in the bio- genic chert-porcelanite section) are independent of original rock type (Sigvaldason and White, 1961), and (4) reported occurrences of yugawaralite (and at most locations, wairakite) are from geothermal areas (Sakurai and Hayashi, 1952; Barrer and Mar- shall, 1965; Harada and others, 1969). Hydrothermal solutions associated with emplace- ment of plutons and with the contemporaneous vol- canic activity apparently permeated the Paleocene rocks and sealed any available pore space by deposi- tion of secondary minerals. Rocks close to the main thoroughfares of circulating fluids were 50-75 per- cent replaced. Deposition of secondary minerals in E14 vesicles probably resulted from several different passes of hydrothermal solutions during which time the temperature decreased and the chemistry of fluids changed. Analysis for iron yielded 3.8, 5.6, and 6.8 weight percent Fe for samples 810—201, 810—301, and 811—201b, respectively, values that are similar to those found for mafic and intermediate volcanic rocks (Turner and Verhoogen, 1960). Cir- culating fluids apparently did not add much iron to the system; rather, the iron in the volcanic rocks was mobilized to form iron oxides and hydroxides, chlorite, celadonite, and stilpnomelane. Alteration of plagioclase and ferromagnesian minerals and ions from the circulating hydrothermal solutions provided abundant calcium for formation of laumon- tite, yugawaralite, wairakite, and minor calcite. Solutions were silica saturated with respect to quartz. Eberlein and others (1971) stated that the conditions for yugawaralite formation include low fluid pressure, ZOO—300°C, and alkaline solutions with silica saturated with respect to quartz. REGIONAL TECTONICS In recent years, there has been much speculation about what effect the subduction of an active oceanic ridge (spreading center) has on an island-arc com- plex (for example, Atwater, 1970; Grow and At- water, 1970; Uyeda and Miyashiro, 1974; DeLong and Fox, 1977; DeLong and others, 1978). It has been proposed by some workers (Atwater, 1970; Hayes and Pitman, 1970; Marlow and others, 1973; among others) that subduction of a ridge will ter- minate ridge spreading. Uyeda and Miyashiro (1974) believed that spreading can continue long after subduction of the spreading center. They also suggested that widespread volcanism accompanies subduction of ridges. Grow and Atwater (1970) equated middle and late Tertiary orogeny in the Aleutian Islands and Alaska to subduction of the Kula ridge beneath the Aleutian-Alaskan part of the North American plate. DeLong and others (1978) speculated that regional greenschist-facies metamorphism in the Aleutian Islands resulted from subduction of the Kula ridge beneath the Aleutian island arc. By the interpretation of Grow and Atwater (1970) and DeLong and others (1978), the Kula ridge entered the Aleutian trench between 20 and 35 my. ago (favored age is 30 m.y.). Our results, however, suggest that the 30-m.y. K-Ar ages on which DeLong and McDowell (1975) and DeLong and others (1978) base their conclusions are initial cooling ages of Aleutian island volcanic rocks and SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 not metamorphic ages. If ridge subduction produces an episode of regional lowgrade metamorphism, as DeLong and others speculate, then because the Andrew Lake and probably correlative sedimentary rocks on Adak Island are not regionally metamor- phosed, subduction of the Kula ridge must have oc- cured before the late Eocene, about 50 my. ago. Models for North Pacific plate motion allow subduc- tion of the Kula ridge at 50 my or 35 my ago de- pending on whether relative motions have been dis- continuous or continuous, respectively (Cooper and others, 1976, fig. 4). Although the plate models are, approximations, there is increasing evidence of dis- continuous motion with faster rates of convergence in early Cenozoic time and slower rates during the middle and late Cenozoic; this evidence, then, favors subduction of the Kula ridge 50 my. ago (Hayes and Pitman, 1970; Hein, 1973; Hamilton, 1973; Scholl and others, 1977; also, see Francheteau and others, 1970; Larson and Pitman, 1972). Therefore, evi- dence for the timing and the effects of ridge sub- duction as proposed by DeLong and others (1978) can be interpreted variously. More likely, early Tertiary (Paleocene or Eocene) metamorphism re- sulted from the depositional burial and tectonic up— lift of more than 4000 m of volcanic and sedimen- tary rocks. Emplacement of plutonic rocks at depth during early development of the arc complex may have contributed to the observed metamorphism. CONCLUDING REMARKS We suggest that sedimentary and volcanogenic rocks in the Wedge Point area are temporally equiv- alent to the upper Eocene, Andrew Lake Formation. However, despite the similarity in lithologic com- position, degree of alteration, and induration, age- diagnostic fossils must be found in rocks in the Wedge Point area to justify including them as part of the Andrew Lake Formation. These rocks ac- cumulated approximately 40 my. ago on the flanks of an active volcanic complex. Subaerial and marine (maximum 200—500 m deep) rocks are represented. Deposits underwent burial diagenesis that signifi- cantly reduced porosity and produced authigenic clay minerals, iron oxides, and possibly quartz. Ad— ditional alteration occurred in conjunction with in- trusion o-f sills, dikes, and especially granodiorite plutons. Secondary minerals formed during these late-stage thermal and hydrothermal events are primarily zeolites (yugawaralite, laumontite, anal- cime, wairakite), clays, iron oxides, and quartz. Most pore spaces that remained after burial diag- enesis were filled during this episode, essentially PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA eliminating any reservoir potential these strata may have had. Organic matter was either initially very low in these rocks, or if present was in places sub- sequently “cooked” by the heat of igneous intrusions. Consequently, these strata are unlikely sources of hydrocarbons. The Finger Bay Volcanics was regionally meta- morphosed to the greenschist facies some time be- fore the late Eocene (possibly 50—55 m.y. ago). These rocks represent the oldest rocks exposed in the western Aleutian Islands. DeLong and McDowell (1975) and DeLong and others (1978) inferred from K-Ar ages that subduction of the Kula ridge spreading center resulted in a regional greenschist metamorphic event along the Aleutian are about 35 m.y. ago (Oligocene). The absence of regional meta- morphism in the Andrew Lake Formation suggests other interpretations. If Adak is typical of other Aleutian islands, either the Kula ridge was sub- ducted about 50 m.y. ago or greenschist meta- morphism need not accompany subduction of a spreading center. REFERENCES CITED Atwater, Tanya, 1970, Implications of plate tectonics for the Cenozoic tectonic evolution of western North America: Geol. Soc. America Bull., v. 81, p. 3513—3536. Barrer, R. M., and Marshall, D. J ., 1965, Synthetic zeolites re- lated to ferrierite and yugawaralite: Am. Mineralogist, v. 50, p. 484—489. Berggren, W. A., 1972, A Cenozoic time sc-alc_Some implica- tions for regional geology and paleobiogeography: Lethaia, v. 5, p. 195—215. Brindley, G. W., and Pedro, G., 1975, Meeting of the Nomen— clature Committee of A.I.P.E.A.: Clays and Clay Minerals, v. 23, p. 413—414. Bundy, W. M., and Murray, II. II., 1959, Arg'illization in the Cochise Mining District, New Mexico: Clays and Clay Minerals, v. 6, p. 342—368. Carlisle, Donald, 1963, Pillow breccias and their aquagene tuffs, Quadra Island, British Columbia: J our. Geology, v. 71, p. 48—71. Carr, W. J., Gard, L. M., Bath, G. D., and Healey, D. L., 1971, Earth-science studies of a nuclear test area in the western Aleutian Islands, Alaska: An interim summa- tion of results: Geol. Soc. America Bull., v. 82, p. 699— 706. Carr, W. J., Quinlivan, W. D., and Gard, L. M., 1970, Age and stratigraphic. relations of Amchitka, Banjo Point, and Chitka Point Formations, Amchitka Island, Aleutian Islands, Alaska: U.S. Geol. Survey Bull. 1324—A, p. A16— A22. Coats, R. R., 1947, Geology of northern Adak Island: U.S. Geol. Survey Alaskan Volcano Inv. Rept. 2, pt. 5, p. 71— 85. 1956, Geology of northern Adak Island, Alaska: U.S. Geol. Survey Bull. 1028—C, p. 045—067. Coombs, D. S., 1953, The pumpellyite mineral series: Min- eralog. Mag, v. 30, p. 113-135. E15 1961, Some recent work on the lower grades of meta- morphism: Australian Jour. Sci., v. 24, p. 203-215. Coombs, D. 8., Ellis, A. J., Fyfe, W. S., and Taylor, A. M., 1959, The zeolite facies, with comments on the interpreta- tions of hydrothermal synthesis: Geochim. et Cosmochim. Acta v. 17, 53—107. Coombs, D. S., Horodyski, R. J., and Naylor, R. S., 1970, 0c- currence of prehnitepumpellyite facies metamorphism in northern Maine: Am. Jour. Sci., v. 268, p. 142—156. Cooper, A. K., Scholl, D. W., and Marlow, M. S., 1976, Plate tectonic model for the evolution of the eastern Bering Sea Basin: Geol. Soc. America Bull., v. 87, p. 1119—1126. DeLong, S. E., and Fox, P. J ., 1977, Geological consequences of ridge subduction: Am. Geophys. Union Ewing Sym- posium, v. 1, p. 221—228. DeLong, S. E., Fox, P. J., and McDowell, F. W., 1978, Sub- duction of the Kula Ridge at the Aleutian Trench: Geol. Soc. America Bull., v. 89, p. 83—95. DeLong, S. E., and McDowell, F. W., 1975, K-Ar ages from the Near Islands, western Aleutian Islands, Alaska: In- dication of a mid-Oligocene thermal event: Geology, v. 3, p. 691—694. Eberlein, G. D., Erd, R. G., Weber, F. R., and Beatty, L. B., 1971, New occurrence of yugawaralite from the Chena Hot Springs Area, Alaska: Am. Mineralogist, v. 56, p. 1699—1717. Fiske, R. S., 1963, Subaqueous pyroclastic flows in the Ohana- pecosh Formation, Washington: Geol. Soc. America Bull., v. 74, p. 391—406. Fiske, R. S., and Matsuda, Tokihiko, 1964, Submarine equiva- lents of ash-flows in the Tokiwa Formation, Japan: Am. Jour. Sci., v. 262, p. 76—106. Fournier, R. 0., 1973, Silica in thermal waters: Laboratory and field investigations, in Proceedings of symposium on hydrogeochemistry and biogeochemistry, Tokyo, 1970: Washington, D.C., The Clarke 00., p. 122—139. Francheteau, Jean, Harrison, C. G. A., Sclater, J. G., and Richards, M. L.,' 1970, Magnetization of Pacific sea- mounts: A preliminary polar curve for the northeastern Pacific: Jour. Geophys. Research v. 75, p. 2035—2061. Fraser, G. D., and Barnett, H. F., 1959, Geology of the Delarof and westernmost Andreanof Island, Aleutian Islands, Alaska: U.S. Geol. Survey Bull. 1028—1, p. 211—248. Fraser, G. D., and Snyder, G. L., 1959, Geology of southern Adak Island and Kagalaska Island, Alaska: U.S. Geol. Survey Bull. 1028—M, p. 371—408. Galloway, W. E., 1974, Deposition and diagenetic alteration of sandstone in northeast Pacific arc-related basins: Im- plications for graywacke genesis: Geol. Soc. America Bull., v. 85, p. 370—390. Gates, Olcott, Powers, H. A., and Wilcox, R. E., 1971, Geology of the Near Islands, Alaska: U.S. Geol. Survey Bull. 1028—U, p. 709—822. Grow, J. A., and Atwater, Tanya, 1970, Mid-Tertiary tectonic transition in the Aleutian Arc: Geol. Soc. America Bull., v. 81, p. 3715—3722. Hamilton, E. L., 1973, Marine geology of the Aleutian Abyssal Plain: Marine Geology, v. 14, p. 295—325. Harada, Kazuo, 1969, Further data on the natural associa- tion of Ca-zeolites: Geol. Soc. Japan Jour., v. 75, p. 629—630. Harada, Kazuo, Nagashma, K020, and Sakurai, Kin-Ichi, 1969, Chemical composition and optical properties of E16 yugawaralite from the type locality: Am. Mineralogist, v. 54, p. 306—309. Hayes, D. E., and Pitman, W. C., 1970, Magnetic lineations in the North Pacific: Geol. Soc. America Mem. 126, p. 291—314. Hein, J. R., 1973, Increasing rate of movement with time be- tween California and the Pacific plate: From Delgada submarine fan source areas: Jour. Geophys. Research, v. 78, p. 7752—7762. Heystek, Hendrik, 1963, Hydrothermal rhyolitic alteration in the Castle Mountains, California: Clays and Clay Min- erals, v. 11, p. 158—168. Jakes, P., and White, A. J. R., 1969, Structure of the Melane- sian arcs and correlation with distribution of magma types: Tectonophysics, v. 8, p. 223—236. Kossovskaya, A. G., 1975, Genetic types of zeolites in strati— fied rocks: Litologia i Poleznye Iskopayemye, no. 2, p. 162—178. Larson, R. L., and Pitman, W. C., 1972, World-wide correla- tion of Mesozoic magnetic anomalies and its implica- tions: Geol. Soc. America Bull., V. 83, p. 3645—3662. Lewis, B. (2., Nelson, W. H., and Powers, II. A., 1960, Geology of Rat Island, Aleutian Islands, Alaska: U.S. Geol. Sur- vey Bull. 1028—Q, p. 555—562. Levering, T. S., and Shepard, A. 0., 1960, Hydrothermal argillic alteration on the Helen claim, East Tintic Dis- trict, Utah: Clays and Clay Minerals, v. 8, p. 193—202. Marlow, M. S., Scholl, D. W., Bufiington, E. C., and Alpha, T. R., 1973, Tectonic history of the central Aleutian Arc: Geol. Soc. America Bull., v. 84, p. 1555—1574. Mitchell, A. H., and Bell, J. D., 1973, Island-arc evolution and related mineral deposits: Jour. Geology, v. 81, p. 381— 405. Powers, H. A., Coats, R. R., and Nelson, W. H., 1960, Geology and submarine physiography of Amchitka Island, Alaska: U.S. Geol. Survey Bull. 1028—P, p. 521—554. Sakurai, Kin-Ichi, and Hayashi, A., 1952, “Yugawaralite,” a new zeolite: Yokohama Natl. Univ., Sci. Rept., sec. 11, no. 1, p. 69—77. SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 Sameshima, T., 1969, Yugawaralite from Shinoda, Shizuoka Pref. Central Japan: Earth Sci. (Jour. Japanese Assoc. Amateur Mineralogists), v. 20, p. 70—78. Scholl, D. W., Bufiington, E. C., and Marlow, M. S., 1975, Plate tectonics and the structural evolution of the Aleu- tian-Bearing Sea region: Geol. Soc. America Spec. Paper 151, p. 1—31. Scholl, D. W., Greene, H. G., and Marlow, M. S., 1970, Eocene age of the Adak ‘Paleozoic(?)’ rocks, Aleutian Islands, Alaska: Geol. Soc. America Bull., v. 81, p. 3583—3592. Scholl, D. W., Hein, J. R., Marlow, M. S., and Buflington, E. C., 1977, Meiji sediment tongue: North Pacific evi- dence for limited movement between the Pacific and North American plates: Geol. Soc. America Bull., v. 88, p. 1567—1576. Seki, Yotaro, 1969, Facies series in low-grade metamorphism: Geol. Soc. Japan Jour., v. 75, p. 255—266. Seki, Yotaro, Oki, Yasue, Matsuda, Tokihiko, Mikami, Keizo, and Okumura, Kimio, 1969 Metamorphism in the Tanya- Wa Mountains Central Japan (11) : Japanese Assoc. Mineralogists, Petrologists, and Econ. Geologists Jour., v. 61, p. 50—75. Seki, Yotaro, and Okumura, Kimio, 1968, Yugawaralite from Onikobe active geothermal area, northeast Japan: Jap- anese Assoc. Mineralogists, Petrologists, and Econ. Geolo- gists Jour., v. 60, p. 27—33. Sigvaldason, G. E., and White, D. E., 1961, Hydrothermal al- teration of rocks in two drill holes at Steamboat Springs, Washoe County, Nevada: U.S. Geol. Survey Prof. Paper 424—D, p. 116—122. Steiner, A., 1968, Clay minerals in hydrothermally altered rocks at Wairakei, New Zealand: Clays and Clay Min- erals, v. 16, p. 193—213. Turner, F. J., and Verhoogen, John, 1960, Igneous and meta- morphic petrology: New York, McGraw-Hill, 694 p. Uyeda, Seiya, and Miyashiro, Akiho, 1974, Plate tectonics and the Japanese Islands: A synthesis: Geol. Soc. America Bull., v. 85, p. 1159—1170. The Livengood Dome Chert, a New Ordovician Formation in Central Alaska, and Its Relevance to Displacement on the Tintina Fault By ROBERT M. CHAPMAN, FLORENCE R. WEBER, MICHAEL CHURKIN, JR., and CLAIRE CARTER SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—F A newly defined formation provides a key to the correlation of lower Paleozoic formations and right-lateral movement on the Tintina fault in east-central Alaska FIGURE 1. 9‘5"???“ CONTENTS Page Abstract _________________________________________________________________ F1 Introduction _____________________________________________________________ 1 Acknowledgments _________________________________________________________ 1 Original description of the Livengood Chert ________________________________ 2 The Livengood Dome Chert ________________________________________________ 2 Areal distribution ____________________________________________________ 2 Lithology and structure _______________________________________________ 6 Reference section in the Lost Creek borrow pit __________________________ 8 Paleontology _________________________________________________________ 9 Regional correlation of rocks in the Livengood and Charley River quadrangles __ 11 Tectonic significance ______________________________________________________ 12 References cited __________________________________________________________ 12 ILLUSTRATIONS Page Generalized bedrock geologic map of the central part of the Livengood quadrangle, showing the Livengood Dome Chert and other major rock units _________________________________________________________ F3 Reference section of the Livengood Dome Chert in the Lost Creek borrow pit ________________________ 4 Map showing the location of the Lost Creek borrow pit and cross section A—A’ ________________________ 6 Geologic cross section A—A' through the Lost Creek borrow pit, showing the structural interpretation ___ 8 Photograph of graptolite-bearing shale beds in the Livengood Dome Chert at the Lost Creek borrow pit 9 Drawings of graptolites from the Livengood Dome Chert ______________________________________________ 10 III « SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 THE LIVENGOOD DOME CHERT, A NEW ORDOVICIAN FORMATION IN CENTRAL ALASKA, AND ITS RELEVANCE TO DISPLACEMENT ON THE TINTINA FAULT By ROBERT M. CHAPMAN, FLORENCE R. WEBER, MICHAEL CHURKIN, _]R., and CLAIRE CARTER ABSTRACT The discovery of Late Ordovician graptolites in the Liven- good Chert plus new data from field mapping and other paleontologic studies in the Livengood quadrangle necessitate major revisions in the Paleozoic stratigraphy in central Alaska. The term “Livengood Chert,” pertaining to a forma- tion consisting of chert, limestone, dolomite, shale, and argil- lite and originally assigned a Mississippian age, is abandoned. A predominantly chert formation, the Livengood Dome Chert, is herein newly defined and is dated by the graptolites as Late Ordovician. It is exposed in an east-northeast-trending belt about 91 km long in the Livengood quadrangle, and the type area is near Livengood Dome. Its structure is complex, and well-exposed thick sections are scarce; therefore the estimated thickness of 300—600 m is uncertain. The Livengood Dome Chert is overlain by an unnamed formation, composed largely of dolomite and limestone, that is provisionally assigned a Middle Silurian to Early Devonian age; the chert is under- lain by Cambrian and Precambrian(?) argillite, slate, quartz- ite, siltstone, limestone, and chert. Extensive chert units similar to the Livengood Dome Chert in lithology and strati- graphic and structural position but paleontologically undated are present in the Tanana, Kantishna River, and Fairbanks quadrangles to the west and southwest. A chert unit that may be correlative with the Livengood Dome Chert crops out in the Circle quadrangle to the east. A correlation is suggested between the Livengood Dome Chert and the Ordovician part of the Road River Formation that lies farther east in the Charley River quadrangle. A correlation between the chert and other rock units of similar ages and lithologies in the Livengood and Charley River areas is significant to under- standing the tectonics of the region because these areas are respectively south and north of the Tintina fault zone. The correlation implies about 300 km of right-lateral offset along this major fault system. INTRODUCTION Recent advances in the stratigraphic and pale- ontologic knowledge of Paleozoic rocks in central Alaska have provided data to allow revision of sev- eral major rock units and their ages and to support significant new regional correlations and tectonic interpretations. A predominantly chert formation, the Livengood Dome Chert, is newly described in this report, and its age is identified as Late Ordo- vician on the basis of graptolites that were discov- ered in it in 1971. A formation consisting mainly of dolomite and limestone that immediately overlies the Livengood Dome Chert is also described but is not named. These formations formerly constituted the Livengood Chert that was described and assigned a Mississippian age by Mertie (1937, p. 105—111), and in this paper the term “Livengood Chert” is abandoned. ACKNOWLEDGMENTS The information presented here is largely the re— sult of geologic mapping and studies in the Liven- good quadrangle by several geologists between 1960 and 1971 (Chapman, Weber, and Taber, 1971). The discovery of graptolites by Donald M. Triplehorn, of the University of Alaska, Michael Churkin, J r., and Claire Carter was a major contribution to this study (U.S. Geological Survey, 1972, p. A51—A52). Re- gional correlations and tectonic interpretations are based also on geologic work in the Charley River quadrangle (Brabb and Churkin, 1969) and the Tanana and Kantishna River quadrangles (Chap- man, Ye‘end, Brosgé, and Reiser, 1975; Chapman, Yeend, and Patton, 1975) . We have benefited also from the geologic investi- gations in the Livengood area and the description of several chert thin sections by Robert L. Foster, from pale‘ontologic and stratigraphic studies in the Liven- good and White Mountains areas by J. Thomas F1 F2 Dutro, Jr., and from geologic mapping by Donald Grybeck, who worked in the eastern part of Liven- good quadrangle with Chapman and Weber in 1968. The geologic report on the Yukon—Tanana region, by Mertie (1937) , who defined the major rock units and many of the geologic problems, was a most help— ful base for the more recent work in the Livengood quadrangle. ORIGINAL DESCRIPTION OF THE LIVENGOOD CHERT The name Livengood Chert was first used in 1926 but Without definition (Mertie, 1926, p. 79). How- ever, Mertie had previously mapped and described this rock unit near Livengood, referring to it as “a stratigraphic series consisting dominantly of chert,” and stated that the valley of Livengood Creek may be considered as its type locality (Mertie, 1918, p. 239—244). In 1937 the Livengood Chert was first defined, and it was described as extending “from a point north of the Sawtooth Mountains to the valley of Beaver Creek, north of the White Mountains, a distance of about 65 miles. The maximum width of this belt, in the vicinity of Livengood, is 81/2 miles” (Metrie, 1937, p. 105). This belt is within the Liven- good quadrangle. Mertie (1937, p. 105—111) identi- fied an eastward extension of this formation in the hills “between the lower valleys of Beaver and Preacher Creeks” (Circle quadrangle) and “in a narrow belt crossing Woodchopper and Coal Creeks a short distance south of the Yukon River” (Charley River quadrangle). He also described westward ex- tensions that occur as isolated beds in the vicinity of Sawtooth Mountain and as “metamorphosed equiv- alents of these rocks” in the Rampart district (Tanana quadrangle). The lithology is dominantly chert, ranging in color from light smoky gray to black, and interbedded minor amounts of limestone, shale, and argillite (Mertie, 1937, p. 105—109). The limestone is com- monly white to cream, crystalline, and in various stages of silicification and is less commonly dark gray, noncrystalline, thin bedded, and mostly un- silicified. Another rock type is chert conglomerate “composed essentially of chert pebbles. in a matrix of chert” that “appears to lie at or near the base of the Livengood chert. Numerous small bodies of basaltic or diabasic greenstone are also found with the sedimentary rocks, but these igneous. members are believed to be largely intrusive and therefore of later origin.” SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 The rocks, according to Mertie, are closely folded, generally strike N. 60° E., and dip steeply south, forming a sequence overturned from south to north in which the same beds are probably repeated several times. Owing to the complex structure, no measure- ment of true thickness is possible, but “a consider- able thickness of beds, perhaps several thousand fee ,” is estimated. The Mississippian age determination was based on one fossil collection, 18AOF8 (Mertie, 1937, p. 110), from a limestone bed on a tributary of Lost Creek that contained crinoid stems Batostomella sp. and Athym's sp. This collection was accepted, with certain reservations, as of Carboniferous, probably Late Mississippian, age, although it lacked the more diagnostic Late Mississipp-ian fossils that were found in collections of the undifferentiated Carboniferous rocks farther north in this region. Mertie states that “it is doubtful if this collection alone, considered without reference to others, even justifies a definite assignment to the Carboniferous,” and “* * * the best estimate of the geologic age of the Livengood chert * * * is that it probably represents the base of the Carboniferous sequence in this region, and it is therefore classified as Mississippian.” THE LIVENGOOD DOME CHERT In recent geologic mapping prior to discovery of the Ordovician graptolites, Mertie’s Livengood Chert was divided into two unnamed units, provisionally assigned an Ordovician to Devonian age: a lower unit that is predominantly chert and minor amounts of interbedded shale and some other rocks and an upper unit of predominantly limestone, dolomite, and minor amounts of chert and shale (Chapman, Weber, and Taber, 1971). The lower, predominantly chert unit in the Liven- good quarangle that is now dated by the Ordovician graptolites is here named the Livengood Dome Chert. The upper, predominantly carbonate unit is here differentiated as a separate but unnamed formation. The term “Livengood Chert” as defined by Mertie is therefore abandoned. AREAL DISTRIBUTION The Livengood Dome Chert crops out in an east- northeast—trending belt (fig. 1) betWeen the Mud Fork and the headwaters of Victoria Creek, a dis- stance of about 91 km. The width of this belt gen- erally is about 5 km but ranges from 1.6 to 9.6 km. Minor amounts of some other rocks, either unrec- ognized or too small to be shown at the map scale, THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT F3 l49°00' 148°30’ 148°00’ 0 5 10 KILOMETERS L____l_l 65°45’ *— “l iwrn “ me: i ive ng ofo‘d 65°30' , l // firea of EXPLANATION '9“ ' CRETACEOUS AND Graywacke, conglomerate, shale, and quartzite JURASSIC(?) TRIASSIC Sedimentary and basaltic volcanic rocks of Rampart Group (Permian?), AND g, and intrusive mafic rocks (Triassic) PERMIANl?) g E Upper DEVONIAN ~ - Graywacke, conglomerate, and shale Devonian OUADRANGLE LOCATION Lower 1. Tanana Dolomite and limestone, in part silicified, chert, shale, siltstone, and Del/30mm ail/00%]: 2- Livengood argillite; also includes some mafic intrusive and volcanic Middle SILURIAN 3. Circle rocks and serpentinite of uncertain ages Silurian 4_ Charley River 5. Kantishna River . . . Upper 6. Fairbanks Livengood Dome Chen Ordovician 0RDOV|C|AN Chert, shale, siltstone, claystone, ruff, and minor limestone CAMBRlAN AND Argillite, slaty shale, quartzite, grit, siltstone, some limestone, and minor chert PRECAMBRIAN(?) /?—\_/ — — — x Contact Inferred fault Site of Lost Creek borrow pit Queried where uncertain FIGURE 1.——Generalized bedrock geologic map of the central part of the Livengood quadrangle, showing the Livengood Dome Chert and other major rock units. Base from US. Geological Survey quadrangle, scale 1:250,000 (1956). F4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 500— METERS EXPLANATION A A A A A A Chert A — q Claystone, clay shale, — ‘l . . — “ sult shale, in part tuffaceous ”””‘ Limestone North margin of pit Maroon shale YYYYYY Tuff and tuffaceous rocks Palagonite tuff,dark—olive-green with white carbonate spots; 450— thin section 7iAWr-570—33A from 450 m. Maroon shale, terruginous, slighty silty, thin~bedded FOSSilIfeFOUS beds Chert, gray, and some cloystone, olive-drab, Graptolite with splintery fracture 8. . . L6 Shale, drab~olive~gray, splintery fracture. Limestone, ¢ ROdIOlOI’lOn —C medium-gray, weathers brown, very fine to line-grained, 3 beds .<_ 30 cm thick. thin white calcite veins. Lithic tuft, '5 Thin section ‘2 moderate yellowish brown, fragmental rock,- thin section A 7lAWr-570-31 B from 405 m, Shale and tuft are soft and t—r— _—_ mostly very thin bedded 400—:— Fault zone; clayey gouge, dark brown Chert, light-gray to medium—dark—gray, with a few reddish-brown bands, very dense and hard, conchoidal fracture, and laminae of clay shale, ochre-stained, waxy luster. Radiolaria in chert thin section 7iAWr-570-30 from 366 m 350 —- Mostly covered; chert, medium-gray, weathers light yellowish brown, and about 10 percent shale Chert, medium- and medium~dark~gray, trace of black chert, some claystone and clay shale, very light gray, a and pale-green clay shale partings; white efflorescence on some clayey bedsi Very soft, round white clay balls, ‘0 2.5 cm in diameter, in one medium—gray cloystone bed‘ All beds are 2.5 to 76 cm thick and evenly layered -C t 0 Z 300 — Shale and chert, dark-gray, in alternating beds 2.5 to 5 cm thick; limestone, dark-gray. sulfide-bearing, rare; thin section 7lAWr-570-27 from 280 m Shale, drab-alive-gray, breaks into tiny splinters, and limestone, mediumdark—gray, weathers brown; beds 5 to 8 cm thick, cut by thin veins of dark-gray calcite Chert, gray, weathers brown, and cloystone, light-gray, weathers yellow in part; alternate in beds 8 cm thick 250 * Continued next page FIGURE 2.—Reference section of the Livengood Dome Chert in the Lost Creek borrow pit (see fig. 3 for leecation). The strati- mined. The amount and direction of movement on the fault are unknown, are included. The exposures are such that a com-1 one place. Within this belt the Livengood Dome plete section and the exact lower contact could not Chert is best exposed in the area between the valleys be found, and the upper contact was seen at only of Lost Creek and the South Fork of Hess Creek, THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT F5 METERS 250 1 A _ 200— 150— 100‘ |>| t> H > bll 50— 0-5 North dips Chert, medium-dark-gray, weathers drab brown, fairly massive Mostly covered; chert, medium-gray, weathers dull brown; a few claystone beds 30 to 60 cm thick Mostly covered; probably largely chert, gray, weathers yellow and brown, massive Clay shale, pale-green and light-greenish-gray, splintery fracture Chert, medium-light-gray, weathers yellow and brown, fairly massive Claystone, very light gray to greenish-gray, weathers yellow, 60 percent; chert, medium-gray, weathers brown, 40 percent; beds 7.5 cm thick Mostly covered; chert, medium—gray, weathers yellowish orange, with minor amount of clay shale interbedded —-/ Claystone, light-gray, in port carbonaceous \f‘hert, medium-gray, weathers yellow, thin clay shale partingsr Claystone, medium-dark-gray to grayish-black, graphitic, greatly sheari/ Poorly exposed; chert, medium-gray, weathers dull brown except for a 4.6 m yellowish-orange band in midsection, fairly massive beds 30 to 60 cm thick, some slightly banded Chert,medium-gray and locally greenish-gray, weathers groy Chert, medium-gray, beds 2 to 8 cm thick; interbedded shale, very light gray to grayish—black, soft, graptolites present in two horizons / Claystone, olive-gray, weathers reddish brown' Claystone, light-grayish-green to gray and brown, weathers yellowish orange with red to maroontints, possibly tuffaceous; chert, gray and greenish-gray, and some Claystone, dark-gray, with light-gray clayey clasts; in thin alternating beds. One bed of black shale 8 cm thick Chert, lightvgray, and interbedded Claystone, very light gray; one bed of chert, grayish-green, 15 cm thick Claystone, dusky-yellow to light-olive-gray, and partings ot clay shale, pale-olive; form 1 cm chips. \ One bed of chert, medium»dark-groy / Clay shale and interbedded tuffaceous siltstone, various shades of gray and yellow, beds 5.. 2.5 cm \ thick; one thin, very soft, brown bed. Chert,medium-gray, in a few thin beds Chert, medium- to medium-dark-gray, iron-stained , beds commonly 2.5 cm thick but one bed at top is 60 cm; some maroon_clay shale partings f Poorly exposed; rubble of clay shale and Claystone, light-yellowish-gray with some medium'gray and some tinged with gray and maroon, finely laminated, hackly fracture; interbedded with chert, medium-dark-gray, very thin bedded, and 2.5 cm beds of felsic tuff, very light gray, to yellowishprange, moderately porous, contains devitrified glass shards and finely divided iron oxide. Thin section 7lAWr ~570-5 from 49 m Mostly covered; chert, gray, weathers pale yellow to dark yellowish orange, beds possibily S. 30 cm thick Clay shale, greenish-gray, soft, breaks into 6 mm chisz_____.____——_ Chert, |ight~gray, weathers ochre, black manganese stains, beds 2.5 to to cm thick; clay shale, ———-\ light-gray, in very thin beds, rare h Chert, greenish-gray, weathers moderate brown, bedded , vitreous, conchoidol fracture, brittle, close jointed South margin of pit graphic orientation and continuity of this section are uncertain because tops and bottoms of beds rarely could be deter- and some beds may be repeated owing to undetected isoclinal folding. tion name is taken from Livengood Dome, which is thickest section, a reference section (fig. 2), was which is herein designated its type area. The forma- l prominent topographic feature in the type area. The l i underlain largely by this chert and is the most uncovered during 197 0—7 1 in a large highway bor- F6 T 9 N. 148°50' T 8N. ' “ , , . \ ‘y ' >\ p \ ‘ V500...\\ ' / ' r: a _‘ 5 ,‘ ‘ fl, 63° ._ . f v 30' R. 7 w. SCALE 1 63 360 l 2 3 L —1 9000 12000 15000 :1 r——l p—l l—{ 2 3 4 4 MILES 18000 I——«1 5 KILOMETERS 21000 FEET p—q FIGURE 3.—Location of the Lost Creek borrow pit and cross section A—A’ (fig. 4). Base from US. Geological Survey Livengood C—4 quadrangle, 1951. row pit 13.6 km west of the town of Livengood and about 2.4 km west of Lost Creek in the SW14 sec. 8, T. 8 N., R. 6 W. (fig. 3). West of the Mud Fork a belt of chert and meta- chert, apparently correlative with the Livengood Dome Chert, is present in the Tanana quadrangle (fig. 1) and extends over a distance of about 73.6 km between Hoosier Creek and Point Tilman on the Yukon River (Chapman, Yeend, Brosgé, and Reiser, 1975). An extensive, predominantly chert unit has been mapped in the northeastern and south-central part of the Kantishna River quadrangle (Chapman, Yeend, Brosgé, and Reiser, 1975; Chapman, Yeend, and Patton, 1975) and in the northwest corner of the Fairbanks quadrangle (Péwé and others, 1966). These cherts outside the Livengood quadrangle are tentatively correlated with the Livengood Dome Chert on the basis of lithologic similarities and stra- tigraphic and structural position. There is no definitive paleontologic evidence that these cherts SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY. 1979 are all the same age. However, the recent discovery of radiolarria-ns of probable early Paleozoic age in samples from both the Livengood Dome Chert and the chert in the south-central part of the Kantishna River quadrangle (D. L. Jones, oral commun., 1978) supports a general correlation of these two chert units. To the east of Victoria Creek in the Circle quad- rangle, the belt of chert mapped by Mertie (1937, pl. 1) between Beaver and Preacher Creeks was ex— amined briefly by Churkin in 1968; no definitive age information was obtained here or in the immediately adjacent areas. Farther east, in the Charley River quadrangle, the Livengood Chert as mapped by Mertie in the Woodchopper and Coal Creeks area, on the north side of the Tintina fault, has been re- interpreted by Brabb and Churkin (1969) as two units: an argillite and chert unit of late Paleozoic age and the Step Conglomerate of Permian age. However, the Road River Formation (Churkin and Brabb, 1967; Brabb and Churkin, 1969) in the southeastern part of the Charley River quadrangle and also on the north side of the Tintina fault has, in part at least, the same age and lithology as the Livengood Dome Chert. LITHOLOGY AND STRUCTURE The Livengood Dome Chert in the type area. con- sists of more than 50 percent chert that commonly ranges from light gray to grayish black, weathers to light and very light shades of gray, green, yellow, reddish brown, and red, and commonly has iron and manganiferous stains and thin coatings. Bedding, which in many outcrops is obscure, ranges from thick and massive (as much as 100 cm) to thin (2— 8 cm) and includes some ribbon chert» units; in part of the unit, banding accentuated by slight color differences is present. The chert and associated rocks are commonly jointed, irregularly fractured, and in places brecciated; thin to hairline white quartz veins are common along fractures and in places form a reticulated boxwork pattern. Very thin layers and partings (a few millimeters to several centimeters thick) of clay shale, argillite, siliceous slaty shale, and siltstone are interbedded with chert, and there are uncommon thin units of tufi', tuffaceous siltstone, limestone, and possibly some graywacke. All these rocks, which form less than 50 percent of any chert section, range from light gray and olive gray to dark gray; included also are a few shaly beds that are grayish red to dark red and reddish brown and some pale-yellow to THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT yellowish-orange lithic tuffs. Because these rocks are thinner bedded and less resistant to weathering than the chert, they are poorly exposed or absent in nat- ural outcrops and rubble patches, and our observa- tions of them have been made chiefly in manmade cuts. A distinctive thin unit of small-pebble con- glomerate, consisting of chert pebbles in a siliceous or chert cement, is present at several places near Livengood and Lost Creek. This chert pebble con- glomerate apparently is in the lower part of the formation, but owing to complex structure and dis- continuous outcrops, its stratigraphic position is uncertain. In thin section the chert is cryptocrystalline but includes some microcrystalline quartz. In part it contains altered remnants of radiolarians (71AWr— 570—30); some sponge spicules are also visible in hand specimens. Minor amounts of sericite and argil- laceous grains are present. Hairline fractures filled with quartz, iron oxides, and mica are abundant. One section shows evidence of at least three sets of frac- tures and a granular quartz that has been invaded by silica. The deformation, fracturing, and rehealing of fractures are indicative of incipient or very low grade metamorphism. A thin section of grayish~black siliceous siltstone, which is interbedded with chert in a borrow pit 6.4 km southwest of Livengood, shows the siltstone to be a moderately sheared and altered volcaniclastic rock containing quartz, much altered albitic plagi- oclase, colorless mica, chlorite, a zeolite (probably laumontite) , and a large amount of What appears ’60 be altered glass. The veinlets are largely quartz or zeolite plus quartz (J. W. Hawkins, written com- mun., 1966) . Tuff is interbedded with the chert in the Lost Creek borrow pit (fig. 2). A very light gray to pale yellowish-orange tuff (71AWr—570—5) in thin sec- ton shows cryptocrystalline to microcrystalline chert, probable devitrified glass shards, and some yellow, red, and brown earthy iron oxides and micro- crystalline white mica; the tuff has a relict clastic or pyroclastic texture, and could be a water-laid sediment. Two other moderate yellowish-brown to dark-olive—green tufi'aceous rocks (thin sections 71AWr—570-313 and —33A) are crystal to vitric tuff and vesicular palagonite tufi', which are gener- ally less extensive and derived probably from basal- tic or andesitic rocks. Several medium-dark-gray limestones, in beds 5—8 cm in thickness and interbedded with shale and chert in the Lost Creek borrow pit, are calcarenites. -A thin section (71AWr—570—27) shows about 85 F7 percent interlocking sand-size carbonate grains that have been recrystallized and 15 percent iron oxides and sulfide that occur as tiny grains and fracture fillings. Well—consolidated medium- to dark-gray and greenish-gray polymictic small-pebble conglomerates and conglomzeratic volcanic graywackes noted in the area between Lost Creek and the headwaters of Erickson Creek are apparently interbedded in the Livengood Dome ’Chert. In thin sections, angular chert pebbles and grains are abundant, and angular pebbles and grains of quartz, feldspar, pyribole, shale, and mafic volcanic rock are common in a very fine grained graywacke matrix. The structure throughout the type area of the Livengood Dome Chert is complex; tight, isoclinal, and overturned folds, joints, irregular fractures, and faults of indeterminate magnitude are common. The regional strike is generally N. 60—7 0° E.; both south and north dips generally exceed 45°. The rocks throughout this area have been subjected to some degree of low-grade metamorphism, and the meta- morphism appears to be controlled by the structural setting and relative competence of the rock type. Foliation is crudely formed in some of the weaker rocks. An accurate thickness for this formation cannot be determined because of the structural complexity, absence of continuous exposures, and concealed upper and lower contacts. The reference section in the Lost Creek blorrow pit has an apparent thickness of about 459 m, but because this section is cut by a fault of unknown displacement and includes dip re- versals, »it has a smaller and unknown true strati- graphic thickness. An estimated thickness of 300— 600 m for the Livengood Dome Chert would be com- patible with its relatively broad outcrop belt in the Livengood quadrangle and with similar size belts of probably correlative chert units in adjacent quad- rangles. The position of the upper contact of the Livengood Dome ‘Chert can. be mapped within narrow limits at many places in the outcrop belt, but the contact zone has been seen. at only one site, on the north side of a hill 1.2 km south of the Lost Creek borrow pit (fig. 4). Here the section of rocks, all dipping 55°— 65° 8., consists of at least 9 m of greenish-gray to black, banded and thin-bedded chert overlain by about 60 m of medium—dark-gray thin—bedded chert wacke and interbedded silty shale, in. which graded bedding indicates that this unit is in upright posi- tion. These rocks are in. turn overlain. by about 15 m of massive fine- to medium-grained sparsely fos- F8 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 3 y, METERS l Lost C eek borrow ' FEET / -f‘ \ pm 600 /’ j l Columnur secllon shown in figure 2 2000 I l \ l 1‘ ‘x \ / I 500 J \ \\\\I\l\ ‘\\\(\Q\\\\\ I, ’ \ 1500 400 ‘. a \ \Zl'l’ll _. ‘ I Bedrock 300 I Ch, n and Shale «\05‘ Y _ ’3}. 49%:‘\\\\‘Inl ‘l‘f/‘CJ concealed lOOO E '/ " ’ . ’ x , J // / x’ T I I ‘I/ \ ’04:"); {of ‘397’ \\\_/'~\ /' 200 Groywocke / I I I "—I l l \ ’c a? 5 I I l I ’5‘ ' 500 conglomevale, I l I \ \ ‘00 pvoiecved // I II Fossiliferous Che" WGCI‘B / \ posilion I limeslone and shale SEA LEVEL SEA [Evil 0 1000 2000 FEET EXPLANATION — —- Contact— Dashed where inferred ° '00 30° 50° Mm“ —v— Foult— Queried where uncerlain Horizontal scale FIGURE 4.—Geologic cmSS section A—A’ through the Lost Creek borrow pit, showing the structural interpretation. See figure 3 for location. siliferous light-gray limestone, the basal 61 cm of which has thin lenticular beds containing some dark- gray rounded chert pebbles. The upper contact of the Livengood Dome Chert is placed immediately below the base of the limestone. Where more fully represented in the Livengood quadrangle, the unnamed formation that overlies the Livengood Dome Chert is composed predomi- nantly of very light gray to light-gray dolomite and limestone, is commonly slightly to nearly completely silicified, includes a minor amount of medium-dark- gray to black chert occurring as thin beds, lenses, and nodules and contains some thin beds of slaty shale, siliceous silrtstone, and argillite. The uniformly dark color of this chert contrasts with the various shades of gray and other light colors that are char- acteristic of the chert in the Livengood Dome Chert. The other noncarbonate rocks in both formations are generally similar. The exact age of this dominantly carbonate rock formation, shown as “dolomite and limestone” on figure 1, is uncertain. Based on exhaustive studies of the several small collections of fragmentary con- odonts, corals, and brachio-pods made in recent years and of collection 18AOF8 (Mentie, 1937, p. 110), the maximum age range of the fauna in the appar- ently basal unit is Middle Ordovician through Middle Devonian (Givetian). The conodont elements have a long range of Middle Ordovician to Middle Devonian (A. G. Epstein, written commun., 1976). The corals have a range of Silurian through Middle Devonian (W. A. Oliver, J r., written commun., 1963). The brachiopods have a possible range of Middle Silurian through Early Devonian, but a Mid- dle Silurian age is considered most reasonable by J. T. Dutro, J r., (written commun., 1976) . Therefore, an age range of Middle Silurian to Early Devonian is provisionally assigned to this unnamed formation. Some mafic intrusive and volcanic rocks and serpen- tinite are, for convenience in mapping, included with this formation. They are closely associated with the carbonate rocks and are considered to be of Silurian and (or) Devonian age; however, their exact ages and relations to the carbonate rocks are uncertain, and a discussion of these rocks is beyond the scope of this paper. The lower contact of the Livengood Dome Chert has not been seen in the outcrop belt. It is, however, closely approximated within a zone of discontinuous outcrops and rubble patches along the northern margin of the outcrop belt. Older lower Paleozoic and Precambrian(?) rocks, including maroon and green argillite and slate, quartzite, siltstone, chert, and some limestone, lie immediately north of this contact zone but were not found in direct contact with identifiable Livengood Dome Chert. Near this contact zone, chert conglomerate and some chert breccia of uncertain origin occur in the Livengood Dome Chert. The nature of the contact is not clear, but we believe that it is unconformable and prob-ably in part a fault contact. REFERENCE SECTION IN THE LOST CREEK BORROW PIT The thickest continuously exposed section of the Livengood Dome Chert known in the type area was examined in 1970 and 1971 in the Lost Creek borrow pit (figs. 2 and 3), which was opened in 1969—70 as a source of rock for construction of the Alyeska pipeline haul road north from Livengood. Subse- quently this pit has been smoothed by bulldozing; in THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT 1976 most of the bedrock section was not well ex- posed, and the graptolite beds were covered. A section having an apparent stratigraphic thick- ness of about 459 m was measured across the floor of the pit from the lower, south edge to the north edge near the top of the hill. About 60 percent of the section is chert, and 40 percent includes thin beds of shale, claystone, siltstone, tuff, and tufl’aceous rocks and a few thin beds of limestone. The beds strike about N. 70° E., dip 50°—70° N. in the south- ern part of the section, and dip about 70° S. in the north part of the section. Graded beds in one out- crop within the north-dipping part of the section suggest that these beds are overturned. Possibly all of the north-dipping section is overturned, but other evidence for this and for the orientation of the beds in the south-dipping part of the section was not found. A fault of unknown displacement, marked by a dark-brown clayey gouge zone, separates the north- and south-dipping parts of the section. A detailed description of this section, including the positions of the graptolite horizon and thin-sec- tion specimens, is given in figure 2. A diagrammatic F9 cross section through the pit and continuing to the south to include a part of the unnamed limestone formation is shown in figure 4. The true stratigraphic thickness represented in this section is not known because the bed orientation, dip reversals, and movement on the fault could not be accurately evaluated and because the section may be repeated by undetected isoclinal folding. Neither the chert conglomerate, which is probably character- istic of the lowest part of the Livengood Dome Chert, nor rocks of other formations above and below the Livengood Dome Chert were found in the pit. There- fore, this section is presumably representative of a part of the formation and does not include either of its contact zones. PALEONTOLOGY Graprtolites were discovered in and collected (col- lection 71ACn—391) from thin shale beds inter- bedded with chert (fig. 5) in the Lost Creek borrow pit by D. M. Triplehorn, Claire Carter, and Michael Churkin, Jr., in 1971. Subsequently, an additional collection (71ACh—287, about 1.5 m above 71ACn— FIGURE 5.—Graptolite-bearing shale beds, indicated under the hammer, in the Livengood Dome Chert at the Lost Creek borrow pit. F10 391) was made from this site by others. Both collec- tions of graptolites present problems in identifica- tion and age determination. The only genera present are of the biserial scandent type, a group that ranges in age from Middle Ordovician to Early Silurian. The usual variety of distinctively shaped genera that characterize most graptolite assemblages is unfor- tunately not found here. Climacogmptus afl’. C. longism‘nus supernus Elles and Wood (fig. 6A) resembles Climacogmptus bicorm‘s tridentatus Lapworth in its three-spined proximal end, but it is narrower and differs also in the arrangement of its proximal thecae and spines. It resembles C. l. supernus in the arrangement of proximal thecae and lateral spines, but it has a much larger rhabdosome and also has an elongated Virgella (middle spine), which C. l. supernus lacks. It is more than 57 mm long and 0.7—1.8 mm. wide. C. l. supernus is an Upper Ordovician species from Great Britain (Elles and Wood, 1901-18, p. 196—197; Toghill, 1970, p. 22; Riva, 1974a, p. 120—125). Climacograptus cf. C. miserabilis Elles and Wood (fig. 6D) is a small species measuring about 15 mm long and about 1 mm wide, having 14—10 thecae per centimeter. The thecal apertures are nearly opposite each other. The Alaskan form differs from C. miser- abilis s.s. in the spacing of its thecae: C. miserabilis s.s. has 10—11 thecae per centimeter. C. miserabih‘s is found in Upper Ordovician and Lower Silurian rocks of Great Britain (Elles and Wood, 1901—18, p. 186—187). It also occurs in Upper Ordovician and Lower Silurian rocks of the Kolyma Basin in the northeastern U.S.S.R. (Koren’ and Sobolevskaya, 1977). Amplexogmptus? pacificus pacificus (Ruede- m‘ann) (fig. 6H) is characterized by a short rhab- dosome and double genicular spines (Riva, 1974b). The Alaskan specimen is about 19 mm long and 1.0—2.1 mm wide; it has 8 thecae in the proximal 5 mm and 13 thecae per centimeter distally. The genicular spines are about 0.5 mm long and occur in pairs on only a few thecae, perhaps owing to im- perfect preservation. Most of the thecae are of the amplexograptid type, which has distinctly everted apertures and outwardly inclined supragenicular walls. T. N. Koren’ (written commun., 1976) has found the Alaskan form to be no different from her specimens of A.? p. pacificus from the Kolym‘a region of the northeastern U.S.S.R. A.? p. pacificus is found in Upper Ordovician rocks in Idaho (Rueh demann, 1947, p. 429; Ros-s and Berry, 1963, p. 125; Riva, 1974b, p. 1457) and in the Cilimacograptus supernus Zone (=Dicellogmptus orndtus Zone) of SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 FIGURE 6.—Graptolites from the Livengood Dome Chert (all X 5). A, Climacograptus aff. C. longispinus supernus Elles and Wood, USNM241051, collection 71ACh—287. B, Clima— cograptus sp., USNM241052, collection 71ACn—391. C, Dia— gram illustrating parts of thecae, as discussed in text: 1, apertural excavation, 2, thecal aperture, 3. supragenicular wall, 4, geniculum, 5, genicular spine. D, Climacoyraptus cf. C. miserab'il'is Elles and Wood, USNM241053, collection 71ACn—391. E, Climacog'raptus sp., USNM241054, collec- tion 71ACn—391. F, Climacograptus sp., USNM241055, col- lection 71ACn—391. G, Amplexograptus? sp., USNM241056, collection 71ACn—391. H, Amplexograptus? pacificus pacifi- cus (Ruedemann), USNM241057, collection 71ACh—287. THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT the Kolyma Basin and Kazakhstan, U.S.S.R. (Koren’ and Soboleyskaya, 1977). The specimen illustrated in figure 6G is question- ably referred to the genus Amplexograptus. It is larger than other species of Amplexograptus (32 mm long, 1.3—3.0 mm wide, 12—10 thecae per cen- timeter) and differs in the details of the proximal end, but it possesses the outwardly inclined supra- genicular walls characteristic of the genus. A few small climacograptids with genicular spines similar to those of A. ? p. pacificus occur in collection 71ACh—287. The remainder of both collections con— sists of numerous large-and medium-size climaco- graptids (figs. 63, E, and F). Some of these may be transitional forms between previously described species, or they may be new species, but they lack any distinguishing features other than rhalodlosome dimensions by which to identify them. Specimen F in figure 6 shows greatly retarded growth of one series of thecae. Only one other specimen having similar biserial-uniserial development was found, but it is a much wider form that has different thecal characteristics from the illustrated specimen. It is possible that this biserial-uniserial development may be due to injury, or it may be a prelude to the uni— serial forms, such as Monograptus. This fauna is Late Ordovician, even though it lacks the dicellograptids and distinctively spined species of Climacograptus that are characteristic of most Late Ordovician faunas of the Cordillera. This age assignment has been confirmed by John Riva (written commun., 1972) and W. B. N. Berry (oral commun., 1972). On the basis of our illustrations (fig. 6), Koren’ (written commun., 1976) has cor- related the Livengood Dome Chert fauna with faunas from the upper part of the Late Ordovician Zone of C. supernus (=Dicellogmptus ornatus Zone) of the Kolyma Basin (Lukav beds) and Kazakhstan (Tchokpar horizon). Nondiagnostic radiolarians and sponge spicules are the only other fossils that were known in the Livengood Dome Chert until early in 1978. Pending final determinations, radiolarians of probable early Paleozoic age have been identified by Brian Holds- worth and D. L. Jones in chert from the Lost Creek borrow pit and from the chert and slate unit of probable Ordovician age (Chapman, Yeend, and Patton, 1975) in the south-central part of the Kan- tishna River quadrangle (D. L. Jones, oral commun., 1978). Conodonts were looked for but not found in the shale irnterbedded with chert in and near the Lost Creek borrow pit. F11 REGIONAL CORRELATION OF ROCKS IN THE LIVENGOOD AND CHARLEY RIVER QUADRANGLES The graptolites discovered in the Livengood Dome Chert indicate that it is coeval with part of the Road River Formation, a similar, but not identical unit that occurs much farther east in the Charley River quadrangle (Churkin and Brabb, 1965) and is widespread still farther east in Yukon Territory (Jackson and Lenz, 1962; Green, 1972). Much of the Livengood Dome Chert is pure chert that con- tains only minor shaly partings. The Road River, in contrast, is mainly graptolitic shale that has only a minor amount of siliceous shale and chert. The Livengood Dome Chert has some green tuff beds not found in the Road River Formation, and the Road River has a few conglomerate beds containing lime- stone and dolomite fragments that have not been recognized in the Livengood Dome Chert. The Road River in the Charley River area unconformably over- lies a thick section of mainly limestone (Brabb, 1967) that in its upper part is rich in Cambrian trilobites (Palmer, 1968). The Road Rive-r underlies the Mc‘Cann Hill Chert, a thin-bedded chert and siliceous shale unit that has a basal limestone unit rich in shelly fossils of Early Devonian age (Chur— kin and Brabb, 1965; Churkin and Brabb, 1967). The fossiliferous Cambrian limestone and distinc- tive carbonate-rich formations of the Tindir Group of Precambrian age that occur below the Road River Formation are apparently absent in the Livengood quadrangle. Instead, the rocks structurally or strat- igraphically below the Livengood Dome Chert along the north side of its outcrop belt are mainly siliceous. However, the presence of a thin fossilifer‘ous lime— stone immediately overlying the chert, chert wacke, and silty shale of the Livengood Dome Chert sug- gests a correlation of this limestone with the basal limestone and shale member of the McCann Hill Chert. Certain other younger formations in the Livengood quadrangle also appear to have close cor- relatives in the Charley River area. For example, the coarse elastic rocks of Late Devonian age near Livengood may be a facies equivalent of the Nation River Formation; the volcanic and elastic rocks of the Rampart Group of probable Permian age may be related to the Circle Volcanics and associated elastic rocks of late Paleozoic age in the Charley River area; finally, the Lower Cretaceous Kandik Group of the Charley River area appears to have its counterpart in the Jurassic ( ?) and Cretaceous gray- wacke formations exposed in the Livengood and F12 Tanana quadrangles. Quartzites, unique to these Mesozoic sections in both the Livengood and Charley River areas, contain Buchia, which further strengthens this correlation. TECTONIC SIGNIFICANCE The tectonic significance of these stratigraphic correlations between the Livengood and Charley River areas is that the two areas of relatively un— metamorphosed Paleozoic strata occur on opposite sides of the Tintina fault zone. The Tintina fault in Alaska, where it crosses the Alaska—Yukon Terri- tory boundary, separates greenschist, quartzite, phyllite, and greenstone of Paleozoic age on the south from essentially unmetamorphosed Precam- brian, Paleozoic, and younger strata on its north side. This contrast in geology across the fault zone can be traced as far west as the western boundary of the Circle quadrangle, where the Tintina appar- ently splays into the several faults that are mapped in the northeastern part of the Livengood quad- rangle (Chapman and others, 1971). On the south side of these faults near the eastern boundary of the Livengood quadrangle, unmetamorphosed rock se- quenc-es like those on the north side of the fault zone in the Charley River area first appear. This geologic correlation strongly suggests a major right-lateral displacement along the Tintina fault system. The distance is about 300 km between the west end of the Road River Formation exposures near Nation in the Charley River quadrangle and the east end of the Livengood Dome Chert exposures near the head of Victoria Creek. Mainly on the basis of data from the Canadian part of the fault, Rod- dick (1967) estimated a 352- to 416-km displace- ment on the Tintina—Rocky Mountain trench sys~ tem, and more recently, Tempelman-Kluit, Gordey, and Read (1976) suggested a displacement of 450 km. A possible explanation for the differences in the calculated amounts of displacement is that the Tin- tina fault has splayed into several faults, and the total offset of an originally widespread unit, such as the Livengood Dome Chert, may be disguised by partial offset of detached, fault-bounded blocks of the unit within the fault system, for example: the Crazy Mountains, Little Crazy Mountains, and the hills between Preacher and Beaver Creeks, which are along the north boundary of the Tintina fault system in the northern part of the Circle quadrangle. Some of the more widespread rock units noted in this area in a brief reconnaissance by Churkin in SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 1968 are chert (probably part of the Livengood Dome Chert), chert-pebble conglomerate, basaltic rocks interlayered with chert (probably part of the Rampart Group), and interbedded quartzite, phyl- lite, and siltstone that in part contains Oldhamia, a fan-shaped trace fossil of either Cambrian or late Precambrian age. The structurally complex se- quences of various rock types, apparently including both lower and upper Paleozoic units, that form these hills are tentatively interpreted, on the basis of limited field data and aeromagnetic data (US. Geological Survey, 1974), as parts of detached, fault-bounded blocks, or a block, within the Tintina fault system. REFERENCES CITED Brabb, E. E., 1967, Stratigraphy of the Cambrian and Or- dovician rocks of east-central Alaska: U.S. Geological Survey Professional Paper 559—A, 30 p. Brabb, E. E., and Churkin, Michael, Jr., 1969, Geologic map of the Charley River quadrangle, east-central Alaska: US. Geological Survey Miscellaneous Geologic Investiga- tions Map I—573, scale 1:250,000. Chapman, R. M., Weber, F. R., and Taber, Bond, 1971, Pre- liminary geologic map of the Livengood quadrangle, Alaska: US. Geological Survey open-file report, scale 1:250,000. Chapman, R. M., Yeend, W. E., Brosgé, W. P., and Reiser, H. N., 1975, Preliminary geologic map of the Tanana and northeast part of the Kantishna River quadrangles, Alaska: US. Geological Survey Open-File Report 75-337, scale 1:250,000. Chapman, R. M., Yeend, W. E., and Patton, W. W., Jr., 1975, Preliminary reconnaissance geologic map of the western half of the Kantishna River quadrangle, Alaska: U.S. Geological Survey Open-File Report 75—351, scale 1:250,000. Churkin, Michael, Jr., and Brabb, E. E., 1965, Ordovician, Silurian, and Devonian biostratigraphy of east-central Alaska: American Association of Petroleum Geologists Bulletin, v. 49, no. 2, p. 172-185. 1967, Devonian rocks of the Yukon-Porcupine Rivers area and their tectonic relation to other Devonian sequences in Alaska, in Oswald, D. H., ed., International Symposium on the Devonian System, Calgary, 1967 : Cal- gary, Alberta Society of Petroleum Geologists, v. 2, p. 227—258. Elles, G. L., and Wood, E. M. R., 1901—18, A monograph of British graptolites: London, Paleontographical Society, 526 p., 52 pls. Green, L. H., 1972, Geology of Nash Creek, Larsen Creek, and Dawson map-areas, Yukon Territory: Geological Survey of Canada Memoir 364, 157 p. Jackson, D. E., and Lenz, A. C., 1962, Zonation of Ordovician and Silurian graptolites of northern Yukon, Canada: American Association of Petroleum Geologists Bulletin, v. 46, no. 1, p. 30—45. Koren’, T. N., and Sobolevskaya, R. R., 1977, Noviy etalon postedovatel’ nosti kompleksov graptolitov na rubezhe THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT Ordovika—Silura (Severo—Vostok S.S.S.R.) : Doklady Aka- demia Nauk S.S.S.R., Tom 236, no. 4, p. 950-953. Mertie, J. B., Jr., 1918, The gold placers of the Tolovana district: U.S. Geological Survey Bulletin 662, p. 239—244. 1926, The Paleozoic geology of interior Alaska [abs.]: Washington Academy of Sciences Journal, v. 16, p. 78—79. 1937, The Yukon-Tanana region, Alaska: U.S. Geo- logical Survey Bulletin 872, 276 p. Palmer, A. R., 1968, Cambrian trilobites of east-central Alaska: U.S. Geological Survey Professional Paper 559— B, 115 p. Péwé, T. L., Wahrhaftig, Clyde, and Weber, F. B., 1966, Geologic map of the Fairbanks quadrangle, Alaska: U.S. Geological Survey Miscellaneous Geologic Investigations Map 1—455, scale 1:250,000. Riva, John, 1974a, Late Ordovician spinose climacograptids from the Pacific and Atlantic fauna] provinces, in Rick- ards, R. B., Jackson, D. E., and Hughes, C. P., eds., Graptolite studies in honour of O. M. B. Bulman: Pale- ontology Special Paper 13, p. 10‘7—126, pl. 7. 1974b, Graptolites with multiple genicular spines from the Upper Ordovician of western North America: Canad- F13 ian Journal of Earth Sciences, v. 11, no. 10, p. 1455—1460. Roddick, J. A., 1967, Tintina trench: Journal of Geology, v. 75, no. 1, p. 23—32. Ross, R. J., and Berry, W. B. N., 1963, Ordovician graptolites of the Basin Ranges in California, Nevada, Utah, and Idaho: U.S. Geological Survey Bulletin 1134, 177 p., 13 pls. Ruedemann, Rudolf, 1947, Graptolites of North America: Geological Society of America Memoir 19, 652 p., 92 pls. Tempelman-Kluit, D. J., Gordey, S. P., and Read, B. C. 1976, Stratigraphy and structural studies in the Pelly Moun- tains, Yukon Territory, in Report of activities, Part A: Canadian Geological Survey Paper 76—1A, p. 97—106. Toghill, Peter, 1970, Highest Ordovician (Hartfell Shales) graptolite faunas from the Mofl'at area, south Scotland: British Museum (Natural History) Bulletin, Geology, v. 19, no. 1, p. 1—26, 16 pls. U.S. Geological Survey, 1972, Geological Survey research 1972: U.S. Geological Survey Professional Paper 800—A, p. A51~A52. 1974, Aeromagnetic map of the Circle quadrangle, northeastern Alaska: U.S. Geological Survey Open-File Report 74—101, scale 1:250,000, 1 sheet. Intertonguing between the Star Point Sandstone and the Coal-Bearing Blackhawk Formation Requires Revision of some Coal-Bed Correlations in the Southern Wasatch Plateau, Utah By ROMEO M. FLORES, PHILIP T. HAYES, WALTER E. MARLEY III, and JOSEPH D. SANCHEZ SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—G Stratigraphic boundary relations/u]; of the Star Point Sandstone and Blackhawk Formation and its bearing to coal correlations CONTENTS Page Abstract _________________________________________________________________ G1 Introduction ______________________________________________________________ 1 Description of the intertonguing and correlation of coal beds ________________ 1 References cited __________________________________________________________ 3 ILLUSTRATIONS Page FIGURE 1. Map of Muddy Canyon-Emery area, Utah, showing the position 0f contact and the zone of intertonguing between the Star Point Sandstone and the Blackhawk Formation, the location of the panoramic photo— graph, and the location of the diagrammatic sections ___________________________________________ G2 2. Composite panoramic photograph and diagrams of the north wall of Muddy Canyon in El/z sec. 36, T. 20 S., R. 5 E., and W1/2 sec. 6, T. 21 S., R. 6 E., showing intertonguing relations between the Star Point Sandstone and the Blackhawk Formation and showing Blackhawk coal-bed correlations of Spieker and revised coal-bed correlations proposed in this paper _________________________________________ 4 3. Diagrammatic section along the Wasatch Plateau front, showing relations among the Star Point Sandstone, the Blackhawk Formation, and tongues of the Star Point and the Blackhawk ______________________ 6 4. Diagrammatic section along a downfaulted block near Emery, Utah, showing relations among the Star Point Sandstone, the Blackhawk Formation, and tongues of the Star Point and the Blackhawk ____ 6 III SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 INTERTONGUIN G BETWEEN THE STAR POINT SANDSTONE AND THE COAL-BEARING BLACKHAWK FORMATION REQUIRES REVISION OF SOME COAL-BED CORRELATION S IN THE SOUTHERN WASATCH PLATEAU, UTAH By ROMEO M. FLORES, PHILIP T. HAYES, WALTER E. MARLEY III, and JOSEPH D. SANCHEZ ABSTRACT Geologic mapping and detailed measurement and descrip- tion of closely spaced stratigraphic sections across the con- tact between the Upper Cretaceous Blackhawk Formation and the Star Point Sandstone of the Wasatch Plateau, Utah have established the presence of a southeast-trending zone of in- tertonguing between these rock units. The contact between the formations drops stratigraphically about 20 m within about 1 km as a result of the intertonguing. This relationship re- quires a revision of correlations of lower Blackhawk coal beds between the two sides of the zone of intertonguing. The Hiawatha coal bed of the east side of the zone is apparently correlative with the Muddy No. 2 coal bed of the west side of the zone and possibly with the Upper Ivie bed of the Con- vulsion Canyon area. Beds referred to as the Hiawath, Upper Hiawatha, and Muddy No. 1 on the west side of the zone of intertonguing pinch out eastward toward the Star Point Sandstone. INTRODUCTION As described by Spieker (1931) , the thickest coal beds of the southern part of the Wasatch Plateau near Emery, Utah, (fig. 1) occur in the lower 30 m of the Upper Cretaceous Blackhawk Formation. According to him, there are two major coal groups in the lower part of the Blackhawk: a lower group consisting of the Hiawatha and Upper Hiawatha coal beds and an upper group consisting of the Muddy No. 1 and Muddy No. 2 coal beds. The Hiawatha coal bed was recognized as immediately overlying the Star Point Sandstone, which was mapped as a continuous, ledge-forming unit. The discontinuous Upper Hiawatha coal bed was con- sidered to be separated from the Hiawatha coal bed by about 6 m of interbedded sandstone, siltstone, shale, and carbonaceous shale. The Muddy No. 1 coal bed was considered to lie 9 to 12 m above the Upper Hiawatha and 8 to 12 m below the Muddy No. 2. During recent geologic mapping (Hayes and Sanchez, 1977 ; Sanchez and Hayes, 1977) and de tailed measurement and. description of closely spaced stratigraphic sections of the upper part of the Star Point Sandstone and lower part of the Blackhawk Formation (Marley and Flores, 1977), a zone of intertonguing between the two formations was ob- served at several localities within a 10-km-long and l-km-wide belt extending south-southeastward from the north wall of Muddy Canyon to a dlownfaulted block near the town of Emery (fig. 1). As a result of this intertonguing, the contact between the two formations is stratigraphically about 20 m higher to the east than it is to the west. The coal-bed corre lations of Spieker (1931) are correspondingly in error. DESCRIPTION OF THE INTERTONGUING AND CORRELATION OF COAL BEDS Figure 2 consists of a composite panoramic photo- graph showing the intertonguing on the north wall of Muddy Canyon and of diagrams of the same area showing the coal-bed correlations of Spieker (1931) and our revised correlations. As shown on the figure, the base of the Blackhawk Formation drops about 20 m westward between Spieker’s sections 463 and 464, owing to the pinching out of the tongue of the Star Point Sandstone. Because Spieker did not rec- ognize the intertonguing and considered the top of the Star Point to be a uniform stratigraphic horizon, he correlated the coal beds immediately overlying the Star Point at the two localities, referring to them G1 G2 111°20' SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 111°15’ I 44000), 39° 00’ WASATCH PLATEAU 38° 55' i EXPLANATION : one of intertonguing Contact between Blackhawk Formation (B) and Star Point Sandstone (SP) + Fault-Dashed where interred. U, upthrown side; D, downthrown side A—-—A’ Line of diagrammatic section ‘X‘ Coal mine 114° 42° 110“ 40"“ UTAH I STUDY AREA 38“ 0 50 100 KlLOMETERS 0 1 2 L__l—l____l 3 KILOMETERS FIGURE 1.—Map of Muddy Canyon—Emery area showing position of contact and zone of intertonguing between Star Point Sandstone and Blackhawk Formation, location of panoramic photograph fig. 2, and diagrammatic sections (figs. 3 and 4). Geology modified from Spieker (1931). both as the Hiawatha. In the revised correlations, the Hiawatha coal bed at section 464 occurs within a sequence of rocks in the Blackhawk Formation that pinches out eastward between lower and upper sandstone ledges of the Star Point. The Muddy No. 1 coal bed of section 464 appears to be at the horizon of the tongue of the Star Point Sandstone but is probably out out by a channel sandstone of the Blackhawk Formation between sections 464 and 463. The Muddy No. 2 coal bed of section 464 is cor- related with the Hiawatha coal bed of sections 461 and 463. Figure 3 and 4 are diagrammatic sections con- structed from closely spaced stratigraphic sections that were measured along the Wasatch Plateau front (A—A’, fig. 1) and in a downdropped fault block near Emery (B—B’, fig. 1). Although different in some details, the lower coal beds in both areas are exposed to the southwest to the pinch out within the Blackhawk Formation or are truncated to the east REVISION OF SOME COAL-BED CORRELATIONS, SOUTHERN WASATCH PLATEAU, UTAH by a tongue of the Star Point Sandstone. The coal bed immediately overlying the Star Point in areas to the northeast is generally traceable southwest- ward into beds that crop out well above the Star Point Sandstone. As a result of the intertonguing, the Star Point—Blackhawk contact drops about 21 m southwest in the exposure along the Wasatch Plateau front and about 16 m in the exposure along the fault block. Although Spieker and Reeside (1925) were aware of intertonguing among various Cretaceous units exposed in the Wasatch Plateau, Spieker (1931), in his excellent bulletin on the Wasatch Plateau coal field, commented on the remarkable “evenness and persistency” of the contact between the Star Point Sandstone and the Blackhawk Formation and appar- ently was unaware of intertonguing between the two rock units within the coal field. Spieker’s con— cept of a very regular contact probably holds true for most localities, but our findings suggest that other examples of intertonguing between the Star Point Sandstone and the Blackhawk Formation may be found within the Wasatch Plateau as more de- tailed geologic work is done. On the basis of our present knowledge, it is neces- sary to revise the terminology in areas to the south and west of the zone of intertonguing described here. For example, the upper coal bed mined in the abandoned mine in Muddy Canyon (fig. 1) , referred to as the Muddy No. 2 coal bed by Spieker (1931), is apparently the Hiawatha coal bed of areas to the northeast of the zone of intertonguing, and a thick (2.5—5 m) coal bed found 12 to 19 m above the Star Point Sandstone in several holes drilled in the Emery West quadrangle west of the zone of intertonguing (Blanchard and others, 1977) may also be the Hiawatha of areas to the northeast. The coal bed mined in. the abandoned Link Canyon mine (fig. 1), identified by Doelling (1972) as the Upper Hiawatha coal bed, is replaced to the east by the Star Point Sandstone tongue and is about 20 m below the strat- igraphic position of the Upper Hiawatha coal bed to the east of the zone of intertonguing. Correlations of coal beds in our study area with those of the Southern Utah Fuel ‘Co. mine in Convulsion Canyon, 14 km to the southwest, are still uncertain, but it may be pertinent to note that the thick bed presently being mined there crops out about 20 m above the top of the Star Point Sandstone and may be equiv- G3 alent to the Hiawatha coal bed of the revised corre- lation. According to Doelling (1972), the coal bed mined in Convulsion Canyon is the Upper Ivie coal bed. The intertonguing noted here, along with data gathered by Marley and Flores (1977 ) , have signif- icant paleogeographic and paleoenvironmental im- plications that are beyond the scope of this report. Some of these implications are presented in pre- liminary form by Marley, Flores, and Cavaroc (1978) and in more detail by Flores (1979) and Marley, Flores, and Cavaroc (197 9) . REFERENCES CITED Blanchard, L. R, Ellis, E. G., and Roberts, J. V., 1977, Lithologic and geophysical logs of holes drilled in the Wasatch Plateau Known Recoverable Coal Resource Area, Carbon, Emery, and Sevier Counties, Utah: U.S. Geological Survey Open-File Report 77-133, 324 p. Doelling, H. H., 1972, Central Utah coal fields; Sevier-San- pete, Wasatch Plateau, Book Cliffs and Emery: Utah Geological and Mineralogical Survey Monograph Series ‘No. 3, p. 1—496. Flores, R. M., 1979, Coal depositional models in some Ter- tiary and Cretaceous coal fields in the U.S. Western In- terior: International Association of Sedimentology, 0r— ganic Geochemistry volume (in press). Hayes, P. T., and Sanchez, J. D., 1977, Preliminary geologic map of the Emery West quadrangle, Emery and Sevier Counties, Utah: U.S. Geological Survey Open-File Re- port 77—822. Marley, W. E., III, and Flores, R. M., 1977, Descriptions of stratigraphic sections, Upper Cretaceous Blackhawk For- mation and Star Point Sandstone in the Emery West and Flagstaff Peak quadrangles, Utah: U.S. Geological Sur- vey Open-File Report 77—833, 257 p. Marley, W. E., III, Flores, R. M., and Cavaroc, V. V., 1978, Lithogenetic variations of the Upper Cretaceous Black- hawk Formation and Star Point Sandstone in the Wa- satch Plateau, Utah: Geological Society of America Ab- stracts with Programs, v. 10, no. 5, p. 233. 1979, Coal accumulation in Upper Cretaceous marginal deltaic environments of the Blackhawk Formation and Star Point Sandstone, Emery, Utah: Utah Geology (in press). Sanchez, J. D., and Hayes, P. T., 1977, Preliminary geologic map of the Flagstaff Peak quadrangle, Emery, Sanpete, and Sevier Counties, Utah: U.S. Geological Survey Open- File Report 77—823, scale 1:24,000. Spieker, E. M., 1931, The Wasatch Plateau coal field, Utah: U.S. Geological Survey Bulletin 819, 210 p. Spieker, E. M., and Reeside, J. B., Jr., 1925, Cretaceous and Tertiary formations of the Wasatch Plateau, Utah: Geo- logical Society of America Bulletin, v. 36, p. 435—454. G4 SHORTER CONTRIBUTIONS T0 STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 VERTICAL EXAGGERATION X 25 464 1 EXPLANATION METERS - 15 Goal Muddy Not 2 { - v Bony coal coal bed E Shale """ Muddy No. 1{ Siltstone coal bed 461 Sandstone Kb _____ I? 0 Correlation line--Queried where extent uncertain __________________________ H‘ h Hiawatha { } cfgleb‘e: coal bed A B FIGURE 2.——A, Composite panoramic photograph of the north wall of Muddy Canyon in E% sec. 36, T. 20 S., R. 5 E., and W% sec. 6, T. 21 S., R. 6 E. Intertonguing relations between Star Point Sandstone (Ksp) and Blackhawk Formation (Kb) are shown. The westward projecting tongue of Star Point (Kspt) overlies the eastward projecting tongue of Blackhawk (Kbt); they have a combined thickness of about 20 m. Channel sandstone (Kbc) in Blackhawk and approxi- REVISION OF SOME COAL-BED CORRELATIONS, SOUTHERN WASATCH PLATEAU, UTAH G5 Hiawatha Muddy No.2 {. . coal bed coal bed Muddy No. 1< coal bed Hiawatha coal bed mate locations of Spieker’s (1931) coal sections 461, 463, and also shown. Outcrop is about 2 km in length and location is indicated on figure 1. B, Blackhawk coal-bed correlations of Spieker (1931). C, Revised coal-bed correlations pro- posed in this paper. (Coal-bed names are those of Spieker.) G6 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 Kb METERS 1o EXPLANATION ’_ Coal-Dashed where ' Sandstone 5 Inferred 0.1 0.2 0.3 0.4 KILOMETERS if: i: Carbonaceous shale l: Shale and siltstone _.J ’7 7’ 7 FIGURE 3.—Diagrammatic section A—A’ (southwest to northeast) along Wasatch Plateau front (see fig. 1) showing rela- lations among Star Point Sandstone (Ksp), Blackh‘awk Formation (Kb), and tongues of Star Point (Kspt) and Blackhawk (Kbt). Section is based on tracing of key beds among 13 detailed stratigraphic sections. METERS 10 EXPLANATION D —— Coal-Dashed where inferred Sandstone 3:5 Carbonaceous shale |:J Shale and Siltstone 0.1 0.2 KILOMETERS 0 FIGURE 4,—Diagrammatic section B—B’ (southwest to northeast) along downfaulted block near Emery (see fig. 1) show- ing relations among Star Point Sandstone (Ksp), Blackhawk Formation (Kb), and tongues of Star Point (Kspt) and Blackhawk (Kbt). Section is based on tracing of key beds among 12 detailed stratigraphic sections. New Evidence Supporting Nebraskan Age for Origin of Ohio River in North-Central Kentucky By w c SWADLEY SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126—H Prepared in cooperation with the Kentucky Geological Survey 3% 3% 5g, ‘ do: .“5 CONTENTS Page Abstract _________________________________________________________________ H1 Introduction ______________________________________________________________ 1 Regional setting __________________________________________________________ 1 Glacial deposits ___________________________________________________________ 3 Discussion _______________________________________________________________ 5 References cited __________________________________________________________ 6 ILLUSTRATIONS Page FIGURE 1. Map showing the preglacial drainage of the northern Kentucky region and the southern limit of glaciation H2 2. Map Showing location of glacial deposits described in this study ____________________________________ 4 III SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 NEW EVIDENCE SUPPORTING NEBRASKAN AGE FOR ORIGIN OF OHIO RIVER IN N ORTH-CENTRAL KENTUCKY By W C SWADLEY ABSTRACT Glacial deposits at four localities in north-central Kentucky support a Nebraskan age for the origin of the Ohio River drainage system. These deposits, including till and outwash, are deeply weathered and must be as old as Kansan. They occur in valleys that deeply dissect the pre-Pleistocene Madi- son divide and that were carved by the Ohio River and its tributaries during a long period of erosion. Thus the glacial diversion that disrupted the Pliocene drainage and created the Ohio River substantially preceded the Kansan Glaciation and must have occurred during the Nebraskan Glaciation. INTRODUCTION Our understanding of the glacial history of north- ern Kentucky and of the glaciated portion of the Ohio River has undergone considerable change in recent years. Prior to about 1950, the glacial de- posits of northern Kentucky were generally con- sidered to be of Illinoian age. In 1956 Rich determined that the upland till in southwestern Ohio and northern Kentucky wars pre-Illinoian, and in 1957 Ray described deposits of deeply weathered till along the Ohio valley in northern Kentucky, to which he assigned a Kansan age. In 1965 Leighton and Ray reported two deeply weathered tills separated by a paleosol on the uplands of northern Kentucky a few miles southwest of Cincinnati. On the basis of the degree of weathering and topographic posi- tion of the deposits, they identified the IOWer till as being of Nebraskan age and the upper as being of Kansan age. In describing the glacial history of the area they attributed the disruption of the preglacial drainage and the formation of the Ohio River to the Nebraskan Glaciation. Swadley (1971) traced the valley of the preglacial Kentucky River northeast- Ward from the area of Carrollton, Ky., and discussed its relation to the development of the Ohio River. In the most comprehensive recent study of the gla- cial history of the region, Ray (1974) described in more detail the evolution of the glaciated Ohio River and again attributed the development of the Ohio drainage system to Nebraskan Glaciation. Ray’s study, based largely on reconnaissance work, has been substantiated by detailed mapping of the joint mapping program recently completed by the US. and Kentucky Geological Surveys. Mapping in north-central Kentucky west of Car- rollton has revealed three small deposits of deeply weathered till and outwarsh that occur at low topographic levels. The low topographic position and deep weathering of these deposits, and their distri- bution along a major preglacial divide lend further support to a Nebraskan age for the initiation of the present Ohio River. REGIONAL SETTING The preglacial drainage system of northern Ken- tucky consisted of three principal streams; the Ken- tucky, Licking, and Manchester Rivers (Ray, 1974, p. 15). These streams flowed northward across the central lowland of Kentucky (fig. 1). This lowland, underlain largely by Ordovician limestones and shales, is bordered on the east and the west by out- ward-dipping escarpments capped by more resistant limestones and dolom‘ites of Silurian age. The pre- glacial Kentucky River headed in southeastern Kentucky and flowed northwestward essentially along its present course to the area of Carrollton, Ky. Near Carrollton the course of the preglacial Kentucky is marked. by an abandoned high-level valley that trends northeast then north from Car- rollton to the area of the mouth of the Great Miami River, about 24 km west of Cincinnati. The river flowed through a broad meander belt that lies largely south and east of the present Ohio River (Swadley, H1 H2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 39° ' KENTUCKY . / 38° EXPLANATION Preglacial drainage --------- Modern drainage._ Shown only where not coincident with preglacial drainage 0 KILOMETERS 50 1 | 1 l FIGURE 1.—Map showing the preglacial drainage of the northern Kentucky region and the southern limit of glaciation. Modified from Swadley (1971) and Ray (1974). 1971) . West of Cincinnati it turned northeast again, flowing generally along the valley of the modern southwest-flowing Great Miami River. The Kentucky continued nonthe‘astward into central Ohio, where it joined the preglacial master stream of the north- central states, the Teays-Mahomet River. The preglacial Licking River headed in east- central Kentucky and flowed essentially along its present course. South of Cincinnati it flowed east of its modern valley, passed to the east of Cincinnati, and continued northward to join the Kentucky near Hamilton, Ohio, about 32 km to the north. The pre- glacial Manchester River headed along the eastern Silurian escarpment and flowed northwest to join the Licking River near Cincinnati. The glacial history of northern Kentucky, as described by Ray (1974), began with ponding of the north-flowing streams by Nebraskan ice moving into Ohio. As the Teays-Mahomet River became dammed by ice, the drainage was diverted to the southwest across the Madison divide. With the ad- vance of ice into northern Kentucky, the valley of the Kentucky River below Carrollton became largely filled with drift, and the drainage over the Madison divide was forced farther to the south (Swadley, 1971). Drift was deposited on the northern Ken- tucky uplands and in many of the existing stream valleys. After the ice front had retreated from its maximum advance (fig. 1), the developing Ohio River established its new valley. The Ohio entered NEBRASKAN AGE ORIGIN OF OHIO the valley of the Manchester River through a break in the eastern Silurian escarpment and then followed that valley to a point north of Cincinnati. North of Cincinnati the Ohio turned southwest along what is now the valley of the Great Miami River, thus re- versing the direction of flow of the preglacial Ken- tucky. From the mouth of the Great Miami it cut a new valley along the edge of the Kentucky River meander belt to the Carrollton area, then flowed westward, probably reversing an east-flowing trib- utary of the Kentucky, across the Madison divide into the valley of the preglacial Ohio. The Aftonian Interglaciation, following the Nebraskan Glaciation, was a time of weathering and erosion in northern Kentucky. The Nebraskan Drift was subjected to intense weathering. In the Madison area, dissection by the Ohio River and its tributaries cut more than 90 m below the drift-fiilled preglacial valleys on the uplands of the divide (Ray, 1974, p. 34). The Kansan Glaciation seems to have been less extensive than the Nebraskan in northern Kentucky. The limit of the Kansan Drift is vague, because the two tills cannot be distinguished except at a few localities, such as those where two tills are found in superposition. The Kansan ice probably blocked the Ohio River drainage temporarily, but no significant drainage changes have been recognized. The second interglacial period (Yarmouth) was also a time of weathering and active stream erosion. The Ohio River system eroded to its greatest depth, the so-called “Deep Stage” (Durrell, 1961, p. 52). Most of the Kansan Drift was removed from the Ohio River valley, except for small areas in pro- tected localities. The high-level Nebrask'an Drift and its partial covering of Kansan deposits were sub- jected to further Weathering and erosion that destroyed any distinctive topographic expression. The advance of the Illinoian Glaciation into Ken- tucky was very limited. Sowtheast of Cincinnati the ice occupied the Ohio valley and only a small upland area. To the southwest it was confined to the Ohio- valley and the lower reaches of a few tributary streams. The only Illinoian Drift recognized in Ken- tucky within the study area is one deposit at Carrollton and a smaller deposit 4 miles to the west (Swadley, 1976) . Permanent drainage modifications affecting northern Kentucky were limited to the Cincinnati area, where the northern loop of the Ohio around Cincinnati was abandoned and the river shifted to its present position on the south side of the city. RIVER, NORTH-CENTRAL KENTUCKY H3 The Sangamon Interglaciation, following the Illinoian Glaciation, was another period of erosion and weathering. Tributaries of the Ohio were prob- ably deepened and extend-ed headward, but erosion along the major stream was limited largely to re- moval of Illinoian Drift and valley train deposits. Ice of the Wisconsin Glaciation, the last of the four glaciations, failed to reach Kentucky. The most pronounced result of the Wisconsin Glaciation on northern Kentucky was extensive alluviation of the Ohio by glaciofluvial deposits and the accompanying pending of tributary streams. GLACIAL DEPOSITS The glacial deposits of this study lie along the Silurian escarpment at the western edge of the cen- tral Kentucky lowland. The escarpment is an area of sharply dissected flat-topped ridges capped by resistant, cliff-forming dolomites and limestones of Silurian age. These formations dip gently westward and form a well-defined cuesta above less resistant limestones and s‘hales of the underlying Ordovician formations. The area (fig. 2) is bordered on the north and west by the Ohio River and on the east by the valley of the Kentucky River. The elevations of the higher ridges range from 260 to 275 m. The area is cut by a number of short, steep streams that flow generally s-outhwestward into the Ohio River at an elevation of 128 m. Streams heading on the east side of the divide flow into the Little Kentucky River, a tributary of the Ohio. Where the Ohio River cuts through the escarpment at Madison, the valley is very narrow with nearly vertical walls. The glacial deposits of the study area (fig. 2) con- sist of till and lesser amounts of outwash. Most are deeply weathered. A high-level till forms thin, dis- continuous caps on many of the flat-topped ridges but exhibits no distinctive topographic expression. The deposits are generally poorly exposed, but where mapped in the Bedford area (Swadley, 1977a, 1977b) they occur at elevations above 245 m. The till consists of reddish-orange to redish-brown clayey silt that contains common to locally abundant gravel and sand. Limonite pellets and nodules are generally abundant. The till is as much as 10 m thick locally, but most ridgetop caps are probably only 3—5 m thick. The gravel portion of the till consists largely of angular fragments of weathered chert of the type producedby weathering of the Silurian dolomites of the area. Other clasts found in the gravel are rounded pebbles, cobbles, and boulders of quartz, quartzite, basalt, granite, gneiss, and schist. H4 SHORTER CONTRIBUTIONS T0 STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 85°30' 85°00' IT» -‘ \. | V. \>/-»;) {2.3 )9 . 1\ f/"* ‘ j/i K i l X \\ I >‘ i <1” “a / , w l N D I A N A Q/ #:1/ .3 Z/ . / 38°_ \ ‘Mi 7 500 _ 45' Q , [Madison ' »*/'\.\j, V>< / r’/ / ‘ 4/16) \ / . / $./ \.\f~a:ngknonf i /Z/,,/ 3/" \OZ’L’fi‘ky / $741; /+” \ ///f \\«/A f/ /’””\\ 3/ 3 j” / / R ail r’ 29/ l 38° 30' EXPLANATION Area of high-level drift along Madison divide in Kentucky Area of low-level drift described in this study @ ®& -------- Approximate southern limit of pre-lllinoian glaciation —> ——- Preglacial Kentucky River. Arrow indicates direction of flow - SOD-foot (152-meter) topographic contour FIGURE 2.—Map showing location of Many igneous rock fragments are thoroughly weathered and crumble easily. In all exposures studied, the drift was completely leached of car- bonate minerals. Locally the drift drapes down along valley sides in what appear to be deposits that filled preglacial valleys at elevations as low as 245 m. Low-level drift occurs at four localities in the study area. Three of these were unreported prior to the Kentucky mapping program. The fourth occur— rence was reported by Ray (1957). At the first local- ity (fig. 2, locality 1), outwash and till occur on low ridges on both sides of Dog Branch in the southern part of the Bethlehem quadrangle (Swadley, 1977b) . K|LOMETERS 10 glacial deposits described in this study The ridge along the north side of Dog Branch is capped by a narrow band of drift for a distance of 1.3 km. The base of the drift descends from an ele- vation of 230 m at the north end to 170 m at the south. The drift is poorly exposed but seems to con- sist largely of weathered sand and gravel that contains no calcareous minerals. The drift includes chert, quartz, quartzite, and igneous and metamor- phic rocks similar to those in the high-level drift. The composition and configuration indicate the deposit is glacial outwash that filled a small, steep valley. Subsequent erosion has cut new valleys on either side, producing a topographic reversal. Dog NEBRASKAN AGE ORIGIN OF OHIO Branch now flows at approximately 152 m elevation in this area. The ridge along the southeast side of Dog Branch is capped for about 0.3 km by two small areas of drift (not shown in fig. 2) at an ele- vation of about 230 m. The drift consists of weathered, noncalcareous silty clay till similar to the high-level till but less oxidized. Boulders, includ- ing some basalt and calcitic dolomite, are common and generally less weathered than those of the higher till. The till is 3 to 5 m thick. The second low-level drift exposure (fig. 2, local- ity 2) is in the Madison West quadrangle (Swadley, 1978) , in the valley of Corn Creek 3.4 km. upstream from where it joins the Ohio River valley. Here drift caps a small ridge along the north side of Corn Creek valley at an elevation of 208 m and extends downward along the nose of the ridge to 152 m. Most of the deposit appears to be graVelly sand out- wash. Near the base, an abandoned borrow pit ex- poses 5 m of reddish-brown, limonite-stained sand and clayey sand containing common to abundant siliceous gravel. This is underlain by 2.4 m of light- brown sandy clay till with scattered chert, quartz, and igneous clasts. The till, which is leached of cal- careous minerals and includes igneous boulders that crumble easily, rests on ledges of Ordovician limes stone. The third occurrence is about 6.4 km upstream from Madison, 1nd, at the base of the south wall of the Ohio River valley (fig. 2, locality 3). This de posit, at an elevation of 152 m, lies 0.3 km west of the point where Hampton Creek debouches into the Ohio River valley (Madison East quadrangle, Gib- bons, 1978). Here drift and loess form, a low spur about 12 m high that j uts northward from the main valley wall. Drift is exposed along the nose and east side of the spur in natural expo-sures as well as along a drainage ditch that cuts across the north end of the spur. A steep face at the end of the spur exposes 5.2 m of medium-brown sandy silt-clay till overlain by 4.6 m of reddish-brown loess. Abundant gravel in the till consists largely of angular fragments of weathered chert but also includes well-rounded quartz and chert pebbles and cobbles of weathered granite, basalt, and gneiss. The till matrix is leached of carbonate minerals and includes no limestone gravel. Along the east side of the deposit an excava- tion for a pond spillway exposes 2.1 m of till resting on limestone. The till is pale-yellowish-brown non- calcareous silty clay streaked with limonite stains. A fourth occurrence of low-level drift in the study area was known prior to the recent mapping. At the mouth of Phillips Branch on the west side of the RIVER, NORTH-CENTRAL KENTUCKY H5 Madison divide (fig. 2, locality 4), drift forms a number of steep, sharply dissected ridges as much as 30 m high. Ray (1957) described the drift in fresh roadcut exposures and assigned a Kansan age based on the weathering profile. Mapping of the Bethlehem quadrangle (Swadley, 1977b) indicates that the de- posit includes both till and outwash. Till capping the ridges is reddish-brown clayey silt containing scattered to abundant chert, quartz, and igneous and metamorphic pebbles and cobbles. Most of the drift body is crossbedded limonitestained sand and coarse, poorly sorted gravel that includes abundant limestone clasts and local calcite cement. The basal calcareous till described by Ray is now slumped and covered. The drift rests on a bedrock bench at an elevation of 137 to 143 m. In marked contrast to the deeply weathered drift are the deposits of drift of Illinoian age in the study area. About 22 km east of Madison, at Carrollton (fig. 2, locality 5), is a body of drift that has long been recognized as being Illinoian in. age (Ray, 1974, p. 48). The deposit occurs in an abandoned meander of the Kentucky River valley at a point where the valley wall separating the Kentucky and Ohio Rivers had been breached by lateral erosion of the two rivers, allowing Illinoian ice moving down the Ohio to push into the Kentucky valley. The till deposited was largely protected from later erosion that removed most of the Illinoian glacial deposits from the Ohio valley in the area. The deposit con- sists of meduim-gray calcareous clay till that con- tains abundant limestone clasts and common to scattered pebbles and cobbles of quartz, chert, and igneous and metamorphic rocks. The elevation of the deposit ranges from 150 to 160 m; the base is not exposed. During the mapping of the Carrollton area (Swadley, 1976) good exposures of till were noted in roadcuts and in a small borrow pit at the north edge of the deposit. DISCUSSION Field evidence in the study area indicates the presence of two tills along the Madison divide. The high-level till is the more deeply weathered and older of the two and occurs at elevations of 245 m and above. Before the Silurian escarpment had been breached, the area was drained by the north-flowing preglacial Kentucky River on the east and by local consequent streams to the west. Near Carrollton the Kentucky flowed at an elevation of 204 m (Swadley, 1971) . The Madison divide, 30 km to the West, was part of a broad upland that stood well above the H6 valley of the Kentucky. The maximum. preglacial dissection along the divide is indicated by remnants of shallow valleys occupied by the high-level till at about 245 m. The deep dissection of the divide by the Ohio River had not begun at the time the high-level till was deposited, as indicated by the following evidence. West of the Madison divide the modern Ohio River follows the valley of the preglacial Ohio (fig. 1). Upstream from Cincinnati it largely coincides with the preglacial Manchester River, but in the area betWeen Carrollton and the mouth of the Great Miami River, the Ohio does not follow the preglacial drainage. As the Madison divide was cut and the ponded waters began draining to the southwest, the Ohio created a new valley, which closely paralleled the preglacial Kentucky River valley and intersected it at several points (fig. 1) but did not occupy it. In order for this to occur some mechanism would haVe been required to make the preglacial Kentucky un- available to the developing Ohio. In a description, of the preglacial Kentucky River between Carrollton and the mouth of the Great Miami, Swadley (1971) reported field evidence in northern Kentucky that indicates the glacial lake ponded against the Madison divide was still intact when the first glacier entered Kentucky. That re— port suggests that the major breaching of the Madi- son divide and the initiation of the Ohio occurred after the first glacier had withdrawn from its maxi— mum advance and after the preglacial Kentucky valley northeast of Carrollton had been filled with drift. The Ohio developed, probably as an ice-margin stream, along the edge of the preglacial Kentucky River meander belt but did not occupy the existing valley because it had been filled with drift. Remnants of this drift still occur along the course of the Old Kentucky at many points. Therefore, drift deposi- tion by the first glacier occurred south of the Ohio in the study area prior to the cutting of the Ohio valley and at elevations equal to or above the pre- glacial drainage. Thus, the elevations would have been about 200 meters within the old Kentucky val- ley and Well above that along the Madison divide. The low-level drift at the Dog Branch, Corn Creek, and Hampton Creek localities was deposited by a later glacial advance. These deposits occur at elevations as low as 150 m, well below the level of the main drainage (204 m at Carrollton) at the start of glacial activity in the region. Therefore, they must have been deposited much later than the high- level till, after a lengthy period of active stream erosion during which the Ohio and its small trib- SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 utaries were entrenched as much as 100 m below the preglacial surface. These exposures, although less weathered than the high-level till, are completely lacking in limestone gravel and appear to be much more extensively leached than either the calcareous Illinoian Till at Carrollton (locality 5) or similar bodies of limestone-rich Illinoian Till further up- stream along the Ohio valley (Swa‘dley, 1973). The Hampton Creek locality, in the Ohio valley 16 km downstream from Carrollton, is considered to be pre—Illinoian only on the basis of the degree of weathering described above. The Dog Branch and Corn Creek localities are considered to be pre- Illinoian because of both the degree of weathering and their topographic position. They occur in sharply incised valleys about 3.2 km upstream from where these valleys join the Ohio. These locations, well back from the Ohio in areas where no indica- tions of Illinoian Glaciation. have been recognized on the surrounding uplands, make it very unlikely that they are Illinoian. Deposition by a tongue of Illinoian ice moving down the Ohio valley, as was postulated for the origin of the till at Carrollton (Ray, 1974, p. 48), would require a minor tongue to have split off at an obtuse angle and have moved more than 3.2 km up a narrow creek valley from the Ohio. Such an explanation seems untenable. On the basis of their topographic position and the degree of weath- ering, the deposits are here assigned a Kansan age. Thus, the presence of drift of Kansan age at low topographic levels within the Ohio River drainage system indicates that the Ohio predates the Kansan Glaciation and, therefore, that the glacial disruption of the Pliocene drainage and the initiation of the Ohio River occurred during the Nebraskan Glaciation. REFERENCES CITED Durrell, R. H., 1961, The Pleistocene geology of the Cincinnati area, in Guidebook for field trips, Cincinnati meeting, 1961: New York, Geological Society of America, p. 47—57. Gibbons, A. B., 1978, Geologic map of part of the Madison East quadrangle, Trimble and Carroll Counties, Ken- tucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1471, scale 1 : 24,000. Leighton, M. M., and Ray, L. L., 1965, Glacial deposits of Nebraskan and Kansan age in northern Kentucky, in Geological Survey research 1965: U.S. Geological Survey Professional Paper 525—B, p. Bl26—B131. Ray, L. L., 1957, Two significant new exposures of Pleisto- cene deposits along the Ohio River valley in Kentucky: Journal of Geology, v. 65, no. 5, p. 542-545. 1974, Geomorphology and Quaternary geology of the glaciated Ohio River Valley—a reconnaissance study: U.S. Geological Survey Professional Paper 826, 77 p. NEBRASKAN AGE ORIGIN OF OHIO RIVER, NORTH-CENTRAL KENTUCKY Rich, J. L., 1956, Pre-Illinoian age of upland till in south- eastern Indiana, southwestern Ohio, and adjacent parts of Kentucky [abs]: Geological Society of America Bul- letin, v. 67, no. 12, pt. 2, p. 1756. Swadley, W C, 1971, The preglacial Kentucky River of north- ern Kentucky, in Geological Survey research 1971: US. Geological Survey Professional Paper 750—D, p. D127— D131. 1973, Geologic map of parts of the Vevay South and Vevay North quadrangles, north-central Kentucky: US. Geological Survey Geologic Quadrangle Map GQ—1123, scale 1:24,000. H7 1976, Geologic map of part of the Carrollton quad- rangle, Carroll and Trimble Counties, Kentucky: US. Geological Survey Geologic Quadrangle Map GQ—1281, scale 1:24,000‘. 1977a, Geologic map of the Bedford quadrangle, north— central Kentucky: US. Geological Survey Geologic Quad- rangle Map GQ—1409. 1977b, Geologic map of part of the Bethlehem quad- rangle, Trimble and Oldham Counties, Kentucky: U.S. Geological Survey Geologic Quadrangle Map GQ—1436, scale 1:24,000. 1978, Geologic map of part of the Madison West quad- rangle, Trimble County, Kentucky: U.S. Geological Sur- vey Geologic Quadrangle Map GQ~1469, scale 1:24,000. Reconnaissance Geologic Study of the Vazante Zinc District Minas Gerais, Brazil By CHARLES H. THORMAN, U.S. GEOLOGICAL SURVEY, and SAMIR NAHASS, MINISTERIO DAS MINAS E ENERGIA, BRAZIL SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126-1 Prepared in cooperation with the Ministerio das Minas e Energia, Brazil The geology and the structural control of the Vazante Zinc deposits in the Paraopeba Formation were determined primarily from aerial photographs FIGURE 1. 2. 3. 9’5"!" .9 CONTENTS Page Abstract _________________________________________________________________ I1 Introduction ______________________________________________________________ Acknowledgments _____________________________________________________ Regional setting __________________________________________________________ Description of geologic units _______________________________________________ Paraopeba Formation _____________________________________________________ Unit A ______________________________________________________________ Unit B _______________________________________________________________ Unit C _______________________________________________________________ Structure ________________________________________________________________ Folds ________________________________________________________________ Lineations ___________________________________________________________ Faults _______________________________________________________________ Mineralization and reserves ________________________________________________ 10 Summary ________________________________________________________________ 11 References cited __________________________________________________________ 11 oooooooo-aoooooooaNNH ILLUSTRATIONS Index map of Brazil and the Vazante region __________________________________________________ Geologic map and cross sections of the Vazante zinc district, Minas Gerais, Brazil ______________ Diagrammatic stratigraphic sections of the late Precambrian Bambui Group (1,000—600 m.y.) for the western side of the Bambui basin and of the Paraopeba Formation in the Vazante district ________ Photograph of thin—bedded, folded phyllite of unit A ___________________________________________ Panoramic view of Serra do Pogo Verde, viewed from the south ________________________________ Diagrammatic examples of ways in which breccia zones may form along step-bedding faults ______ Photograph of resistant ridge of mineralized breccia (primarily hemimorphite, hydrozincite, willemite) on the north end of Serra do Poco Verde __________________________________________________________ III SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 RECONNAISSANCE GEOLOGIC STUDY OF THE VAZANTE ZINC DISTRICT, MINAS GERAIS, BRAZIL By CHARLES H. THORMAN and SAMIR NAHASS 1 ABSTRACT The Vazante district, Minas Gerais, 130 km south of Paracatu, produces most of Brazil’s zinc metal. The district is in the western part of the late Precambrian Bambui basin along the eastern and leading edge of the north-trending Brazilian orogenic belt (ca. 600—500 m.y.). Bedding and low- angle thrust faulting, folding, and low-grade metamorphism dominated the structural development. Trends within the district are northeast, and dips are 20°—45° NW. Three sets of folds developed during the main period of eastward thrust- ing of older Precambrian rocks over the west margin of the Bambui basin. A fourth fold set is transverse to the regional trend. Rocks in the district are tentatively assigned to the Paraopeba Formation of the Bambui Group and are desig- nated units A through C in ascending order. The contacts between units A, B, and C are interpreted to be low-angle or bedding faults; their original stratigraphic positions with respect to each other are unknown. The Vazante deposit is epigenetic. Zinc silicate minerals (hemimorphite and willemite) occur in a folded breccia zone in the lower part of unit B of the Paraopeba Formation. The breccia zone is interpreted to be tectonic in origin, having formed along the step of a step-bedding—plane fault during Brazilian time. The source of the zinc is probably syngenetic in the Paraopeba Formation. Broad uplift and deep Weather- ing of the region and concurrent supergene enrichment took place during late Mesozoic and Cenozoic time. Approximately 110,000 metric tons of zinc metal were produced from 1965 to 1974; reserves may be as much as 3 million metric tons. INTRODUCTION The Vazante zinc district, the largest zinc-produc- ing area in Brazil, is about 130 km south of Paracatu and 120 km northwest of Patos de Minas in the west- ern part of the State of Minas Gerais (fig. 1). Com- panhia Mineira de Metais at Serra do Poco Verde and Companhia Mercantile Industrial Inga at Serra do Ouro Podre (fig. 2), the only companies operating in the district, produced approximately 110,000 1Companhia de Pesquisa de Recursos Mineraia, Ministerio das Minas e Energia, Rio de Janeiro, Brazil. metric tons of zinc metal from 1965 to 1974 (White and Nagell, 1975; Bohomoletz, 1975, p. 176). The purpose of our investigation was twofold: First, a geologic map of the Vazante region was needed to serve as a base for a pilot geochemical study of the Vazante district. This pilot study was the initial phase of Pr‘ojet‘o Bambui, a. reconnais- sance geochemical exploration program of the Bam- bui basin undertaken by the CPRM (Companhia. de Pesquisa dc Recurso-s Minerais.) for the DNPM (Departmento National de Producao Mineral). Second, a better understanding of the structural setting of the district was needed to evaluate the feasibility of looking for additional Vazantertype deposits. This report was prepared as part of the cooperative program of the Ministerio das Minas e Energia (MME) and the U.S. Geological Survey (USGS), sponsored by the Government of Brazil and the U.S. Agency for International Development (USAID) , of the U.S. Department of State. Luciana Jacques de Morais (unpublished data, 1955) presented the first study of Vazante, and Moore (1956), of the USGS, studied the Serra do Pogo Verde. Subsequently, Carvalho, Dequech, and Guimaraes (1962), Branco (1962), Guimaraes (1962), and Cassedanne (1968a,b) reported on the Vazante deposit. These studies Were restricted pri- marily to the deposit and to the narrow cuesta along which mineralization occurs. Amara] (1968) was the first to map the surrounding terrain. Companhia Mercantil e Industrial Inga began production in 1965, and Companhia Mineira de Metais, in 1970. This investigation was of a reconnaissance nature. Approximately three weeks were spent preparing a photogeologic map from 1 :25,000-scale aerial photo- graphs (fig. 2). Five days in March and April 1974 were spent field checking the photogeologic interpre- tations. Il 12 o 46 l 47° \Brasilia l ‘ Paracatu STATE OF MINAS GERAIS Joéo Pinheiro fl WM Belém BRAZIL .Manaus .Porto Velho DEFORMED BELT- .Cuiaba § rasil ia, M ~ Gonaniao ‘2? l‘ Q [ STUDY AREMMU o 85° Paulo. Rio de Janeiro 0 500 1000 km Li-J-LL—I—w 0 \ é _| FIGURE 1.—Index map of Brazil and the Vazante region. Tri- angulo Mineira area (patterned) south of Vazante, studied by Barbosa and others (1970). SHORTER CONTRIBUTIONS T0 STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 Several local names for individual parts of the main cuesta near the base of unit B (subunit B-3) along which mineralization took place are shown on the geologic map. We are not aware of any single name that applies to the entire cuesta. For ease of reference we have adopted the informal term “min- eralized cuesta” to refer to the entire feature. ACKNOWLEDGMENTS We wish to acknowledge the many fruitful discus- sions with Dr. Oscar P. G. Braun, CPRM, that helped us develop our thoughts and ideas on the geology of the region. The aid and support of Dr. Benedito P. Alves, late Chief of the CPRM Agency in Belo Harizonte, was especially appreciated. Drs. Carlos Alberto Heineck and Hugo Peter Steiner, CPRM, took part in our field work, and their as- sistance is gratefully acknowledged. Dr. Peroio Branco, CPRM made several pertinent, helpful sug- gestions during the final preparation of the manu- script. We especially acknowledge the support of Dr. Joao Batista de Vasconcelos Dias, CPRM who was Director and Coordinator of the MME—USAID agreement during our investigation. REGIONAL SETTING The study area is *in a deformed belt (fig. 1) along the western margin of the Bam‘bui basin, a late Pre- cambrian basin (1,000—600 m.y.) filled with silt- stone, marl, conglomerate, sandstone, limestone, and dolomite of the Bambui Group. The clastic and pe- litic rocks were derived from a highland west of the deformed belt. Older Precambrian rocks were thrust eastward over the western. edge of the basin during the Brazilian orogeny (ea. '600—500 m.y.) ; the orogeny formed a major north-trending deformed belt about 1,000 km long, characterized by low- angle thrust faults, bedding-plane faults, large folds, and widespread low-grade metamorphism. Allochthonous rocks in the belt include the follow- ing Precambrian units, in ascending order (Barbosa and others, 1970) ; granite-gneiss complex; mie‘a schist, quartzite, and amphibolite of the Araxa Group; quartzite and phyllite of the Canvastra Group; and calc-schist of the Ibia Formation, a basal unit of the Canastra. The Vazante area is about 30 km east of the Serra dos Piloes, in which the eastern- most major thrust brings Canastra over the Bam- bui (fig. 1). Smaller scale thrusting and broad fold- ing of the deformed belt extend to about 100 km east of Vazante. Only the Bambui Group is present in the study area. RECONNAISSANCE GEOLOGIC STUDY, VAZANTE ZINC DISTRICT, MINAS GERAIS, BRAZIL I3 Previous workers interpreted the structural set- ting at Vazante as one of normal faulting or ten- sional tectonics. A plausible alternative presented here involves compressional tectonics associated with essentially no normal faulting but with thrust, bedding, and step-bedding faulting and related structural complications consistent with the struc- tural environment of the deformed belt within which the Vazante district is located. Barbosa and others (1970) mapped these types of structures in the Triangulo Mineira immediately to the south of Vazante (fig. 1). DESCRIPTION OF GEOLOGIC UNITS Braun (1968; and oral commun., 1974) assigns all the rocks in the Vazante region, herein mapped, to the Paraopeba Formation of the Bambui Group; Amaral (1968) considered the eastern two-thirds of the Vazante region to be underlain by the Sete Lagoa Formation of the Bambui Group and the western one-third, by the Canastra Group. This re- port follows the work of Braun. We divided the Paraopeba Formation into three informal units (A, B, and C, in ascending order). Figures 3A and 3B are diagrammatic stratigraphic sections of the Bam- bui Group of the western side of the Bambui basin and of the Paraopeba Formation in the study area. It is possible that unit A includes some Paranoa Formation and unit C some Tres Marias Formation. The relative stratigraphic positions of units A, B, and C are not clear at this time. The contact between units A and B—the Santa Catarina fault—is based on the marked contrast between the pronounced, pervasive, penetrative deformation and low-grade metamorphism of unit A and the generally local and much weaker low-grade metamorphism of unit B. Beds on opposite sides of the Santa Catarina fault are more or less concordant. The nature of the con- tact between unit B and C—the Vazante fault—is less clear. Along the southern half of the fault’s ex- posure, beds are commonly discordant on a small scale across the contact, whereas beds are concordant across the contact to the north. Sandstone is more abundant than limestone in unit C, whereas limeL stone is more adundant than sandstone in unit B. The Vazante fault, between units B and ‘C, is here interpreted as a bedding fault that locally has brought about discordant stratigraphic relation— ships, thus implying that unit B is older than and grades upward into unit C. In contrast, the Santa Catarina fault appears to be a larger structure, and the stratigraphic relationship between unit A and units B and C is open to question. PARAOPEBA FORMATION UNIT A Unit A is dominantly a silty micaceous phyllite with thin lenses of quartzite. The phyllite varies in color from tan to brown to reddish yellow : the quartz- ite beds are cream colored and have dark red—brown iron—oxide coatings on weathered surfaces. The rocks are commonly laminated (fig. 4). A moderate to well-developed foliation at 20°—90° to the bedding is typical and has often been misidentified as bed- ding. Fracture and crenvulation cleavage and incipient foliation at 20°—40° to the bedding are common in the quartzite lenses. Lineations and small-scale folds are common. Thickness of the unit is unknown, but it probably exceeds 500 m. UNIT B Unit B forms a sinuous northeastatrending belt that ranges in. width from 5 km in the southwest to 2 km in the northeast. The dominant lithology is massive to medium-bedded, cream-colored to light- gray, aphanitic to fine-grained dolomitic limestone. Collem‘a, an algal stromatolite, occurs in some beds (Cassadanne, 1968a), and siliceous layers are locally present. A few thin beds of medium-grained, well- sorted, light-colored quartzite were noted near the crest of Serra do Pogo Verde. Interbedded with the limestone are thin-bedded, in part laminated marl, shale, and siltstone that vary in color from red to bufl’ to brown to green. The massive to thick-bedded limestone underlies cuestas and isolated hills that alternate with elon- gate and irregularly shaped valleys underlain by marl, shale, silstone, and, in some places, limestone. This topography is typical of unit B. A major char- acteristic of the Paraopeba Formation is the rapid lateral and vertical facies change from carbonate to argillaceous rocks (Braun, 1968). A major problem in understanding the geology of the Vazante area is deciphering the relationship of the limestone hills to the shale and marl valleys that are on strike with the hills. At several localities a 50- to 80-m section of limestone will change along strike to a comparable thickness of shale within less than 20 m. In a small pit about 2 km south of the Inga mining operation, approximately 10 m of thick-bedded limestone con~ taining thin argillaceous partings grades laterally into shale and thin-bedded limestone in a distance of about 20 m. I4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 6" ' 77 54 0 1, web 4345' ‘LAGOA I UPPER PRECAMBR/AN ’ PARAOPEBA FORMATION ,/ Siltstone, limestone, sandstone, and marl f Limestone, siltstone, and shale Limestone and shale \\ siltstone, shale, and limestone Limestone, siltstone, and shale Limestone Marl and shale Phyllite and quartzite 18“ 00' o _. —n c.) 5 FIGURE 2.—-Geologic map and cross sections of the Vazante district, Minas Gerais, Brazil. EXPLANATION 3O 8’ RECONNAISSANCE GEOLOGIC STUDY, VAZANTE ZINC DISTRICT, MINAS GERAIS, BRAZIL I5 NORTHWEST SOUTHEAST D D' UNITC UNITB UNITA—> SANTA CATARINA METERS AULT 800 400-; SEA _' , LEVEL‘_‘ . EVEL -400 C UNIT 0’ UNITC UNI B A METERS MINERAL/ZED CUESTA 1200 SANTA CATARINA FAULT 800 ’v 400 SEA SEA LEVEL / LEVEL —400 B B' METERS UNIT C UNIT 3 UNIT A—> MINERAL/ZED CUESTA SANTA CATARINA MINEIRA FAUL FAULT /// /// //// ;"-2// , /.‘/ ¢:;// SEA 3/ .- LEVEL ¢2 ////// / / / A , UNITC ' NITZE S UNITA MET1E223_ MINERAL/ D CUE TA SANTA CATARINA - VAZANTE FAULT MINEIRA FAUL FAULT 800—-~-- «— -~ , / -- - / - /X /’- /" z/ / // 7f // 400—” [Le/<29» / 0/ ; 4 {1//?////é////é/¢/ - ,Lcar' ...¢~ //’ za’ ,/- //r / /a/ ,7 “,1 r’ H7?/ /:/ /// SEA- '*/ 1 " / // / / " // / / MM / SEA LEVEL -//W//,-r , / 4:: ’7 / / 5&0; ” // // LEVEL ’400 -’/-‘ /////"_;.2 ))////v/ ////// //,/// // /// 1 2 3km l_;_J____;__J NO VERTICAL EXAGGERATION __ LITHOLOGIC UNITS 0N CROSS SECTION Sandstone Limestone _ f Siltstone and shale T‘“ Marl and shale Phyllite and quartzite Thrust fault—Arrows show direction of relative movement. Dashed where inferred MAP AND CROSS SECTION EXPLANATION Trace of marker bed from aerial photo— 4—. —~?-h Possible thrust fault along bedding plane LARG E-S.CA LE FOLDS—Showing graphs ._, Tear fault bounded by thrust fault or crest/ine and direction of plunge -—-— Trend or lineament from aerial photo- bedding fault {—5—} Anticline graphs ATTITUDES OF BEDS o—H Syncline Contact—Dashed where approximate .11 Strike and dip H“ SMALL-SCALE FOLDS—Showing : Possible facies change—Photogeo/ogic i5 Strike and dip showing plunge of plunge interpretation lineation ~ ~ ~ ~ Mineralized breccia zone hA— Thrust fault—Dashed where inferred. _I_ Interpreted from aerial photographs xxx Road Teeth on upper plate. Arrows show Stream direction of relative movement FIGURE 2.—Continued. 16 S-HORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 THICKNESS GR UP F ’ O ORMATION LITHOLOGY IN METERS Arkose, micaceous siltstone, graywacke, arkosic sandstone, argillaceous limestone TRES lenses MARIAS 20—400 LAGUNAR CONG LOMERATE SAMBURA CONGLOMERATE ‘5 Pelitic and carbonate rocks including facies m PARAOPEBA named "Sete Lagoas," "Lagoa Jacaré," E (Fig. 3B) "Serra de Santa Helena". Sandstone, 1000—1400 '33 siltstone, argillite, marl, limestone CARRANCAS CONGLOMERATE Ouartzites with interbedded phyllite and metasiltstone PARANOA 100—3800 850 MIGUEL CONGLOMERATE CANASTRA GROUP INFORMAL FORMATION TsL/ICENEESSS’ UNITS, THIS LITHOLOGY REPORT LITHOLOGIES SILTSTONE WITH SHALE 500 C E DOLOMITIC LIMESTONE SANDSTONE 5 E @ IRON FORMATION 0 g 750—1000 B m < MARL Q. ~ ~ PHYLLITE -i€ PHYLLITIC SILTSTONE 500 A ‘ ‘ FIGURE 3.—Diagrammatic stratigraphic sections. A, Stratigraphic section of the late Precam- brian Bambui Group (1,000—600 m.y.) for the western side of the Bambui basin (Braun, 1968). B, Stratigraphic section of the Paraopeba Formation in the Vazante district. Unit A may include some Paranoa Formation and unit C some Tres Marias Formation Unit B has a minimum thickness of approximately 300 m; it is subdivided into six subunits, numbered 13-1 at the base to B—6 at the top (fig. 2) , primarily on the basis of photointerpretation of geologic and geomorphi-c features. Some of the subunits are shown intertonguing with others (fig. 2), a reflec- tion of the abrupt facies changes in the Paraopeba Formation. Subunit B—l consists of marl and shale and forms a lowland occupied in part by the Rio Santa Catarina (fig. 5). Subunit B-2 consists of limestone, which forms a narrow zone of subdued parallel ridges; the limestone appears to grade into the mar] and shale of B—l to the northeast. Subunit B—3, which contains the Vazante ore bodies, forms the prominent “mineralized cuesta” and is limestone along most of its length; shale and siltstone con— RECONNAISSANCE GEOLOGIC STUDY, VAZANTE ZINC DISTRICT, MINAS GERAIS, BRAZIL I7 FIGURE 4.—Thin-bedded, folded phyllite of unit A. These are f1 folds. Note the crenulation cleavage along axial planes. Camera lens cap is 6 cm Wide. stitute a major portion of the subunit at the south end of Serra de Ouro Podre. Subunit B—4 is char- acterized by very low relief and probably is com- posed largely of fine-grained elastic and argillaceous rocks interlayered with some thin-bedded limestone; a few isolated hills are underlain by thin— to medium- bedded limestone. Subunit B—5 is characterized by numerous low cuestas and hills of moderately dip- ping thin-bedded to massive limestone; several hills are traceable on aerial photographs for 3 to 6 km. Subunit B—6 is distinguished by low, isolated hills of gently dipping to horizontal, thick-bedded to mas- sive limestone beds that are discontinuous; the in- tervening areas are flat pasture land underlain by thin- to medium-bedded limestone, siltstone, shale, and marl. Subunits B-l, B—3, and B-4 were not separated in the northern part of the area, because B—3 be- comes subdued and unrecognizable. B—2 and B—3 terminate to the southwest against a tear fault. The rocks on the southwest side of the tear fault are mapped as unit B. Subunits 13—5 and B-6 are not distinguishable immediately northeast of Vazante because of folding and possibly because of stnati- graphic complications. The basal contact of subunit B—5, however, appears to merge with the Mineira fault or to become a bedding fault to the southwest of the tear fault. UNIT C Unit C underlies a terrain of undulating hills and cuestas in the western part of the area. It consists of interbedded siltstone, marl, sandstone, and dolo- mitic limestone. These rocks are lithologically simi- lar to those in unit B, but they have a much higher ratio of siltstone and marl to limestone. Lime-stone underlies most of the prominent, sharp-crested cuestas, whereas sandstone generally underlies rounded, more subdued cuestas. The thickness of unit C probably exceeds 500 m; the top of the unit is not exposed in the map area. FIGURE 5.—Panoramic view of Serra do Pogo Verde, viewed from the south. The valley to the right (northeast) is under- lain by shale and marl of subunit B—l. The ridge is Serra do Pogo Verde, at the southwest end of the mineralized cuesta; the Companhia Mineira de Metais mine is on the opposite side of the ridge. Low hill in the foreground is unit A; hills in the background are unit C. I8 SHORTER CONTRIBUTIONS T0 STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 STRUCTURE The Vazante district displays many structures typical of a deform-ed belt. During the Brazilian orogeny (ca. 600—500 m.y.), older Precambrian rocks were thrust eastward over the western mar- gin of the Bambui basin along a front approximate- ly 1,000 km long. Large-scale thrusting died out west of Vazante; the easternmost major thrust is in the Serra dos Piloes, about 30 km to the west of Vazante (fig. 1). The effects of the orogeny are Widespread at Vazante, and at least tWo phases of tectonism can be documented. First, low-angle, bedding, and step- bedding thrust faults developed, as did. some re- lated folding. Low-grade metamorphism and pene- trative deformation either preceded or was coeval with the thrusting. Second, broad cross folds de- formed all previous structures. Interpretation of the types of structures is based on their characteristics as seen in the field and as interpreted on aerial photographs, as well as on the geological history of the surrounding region. F OLDS Four sets of folds (f1 through f.) are recognized. Set f1 consists of folds that trend N. 20°—50° W. and range in size from micro-folds less than 1 mm to some having wavelengths of 2-5 cm and ampli- tudes of 1—3 cm (fig. 4). This fold set is best de- veloped in unit A, is well developed along the con- tact between units B and C, and is also present in units B and C, where it was observed only in pelitic rocks interbedded with massive limestone and sand- stone beds. Set f2' consists of two northnorthwest-plunging anticlines, :in unit A and in the lower part of unit B. The folds die out upward in the basal marl-shale member of unit B; the overlying cuesta-forming mineralized member is not affected by the f2 folds. A West-plunging anticline, 2.5 km south of Lagoa Feia, is also interpreted as belonging to set f2. This fold is a complex anticline that underwent possibly two stages of folding. A tightly folded axial zone is interpreted to be an f2 fold that was refolded about a larger, more open f4 fold. The f2 folds have wave- lengths and amplitudes of about 0.5 km and are at least 2 km in length. The sets of f1 and f2 folds are parallel. The third fold set, f3, consists of large-scale struc- tures that trend N. 10° W. to N. 40° E., parallel to regional structural trends. All units and f1 and f2 folds are folded about f3 axes. The folds are broad, open, and slightly asymmetrical. They are over- turned to the southeast; wavelengths are 0.5 to 2 km, and. amplitudes are 0.1 to 1.5 km. Set f3 folds are more abundant and more pronounced in units B and C than in unit A. Most of the folds are doubly plunging and discontinuous. Disharmonic folding is dominant, with the more intense folding having occurred in resistant limestone and sandstone beds. Set f4 folds are large west-no-rthwest-plunging folds transverse to the regional trend. All structures in the northern part of the area are folded about f4 axes. Set f1 folds, the oldest, formed early in the Brazil- ian orogeny during the major period of widespread dynamic metamorphism. Lineations in all units are interpreted as being of this same age. Set f2 folds postdate the Santa Catarina fault, which is folded about f2 axes, but appear to predate the Vazante fault and most structures in the lower part of unit B. Both f1 and f2 axes are more or less normal to the direction of Brazilian thrust faulting. Set f3 folds are interpreted as being related to the major period of Brazilian folding and thrust faulting. Transverse folding of all structures by f. folds is the youngest event, postdating all thrust faulting. LINEATIONS Three types of mutually parallel lineations were noted: streaks of fine-grained sericite, f1 fold axes, and traces of fractures that intersect parting and bedding planes. The streaking and, smearing out of sericite appears to be the most widespread form of lineation. The fractures were noted only Where f1 folds were present and are parallel to 1’1 axial planes. Quartz-filled fractures are common. The lineations occur throughout unit A and in most of the thin argillaceous beds in units B and C. In units B and C, lineation is strongest at and near the contacts betWeen argillaceous beds and. sandstone and car- bonate beds. The thinner the argillaceous bed be- tween massiVe competent beds the better developed the lineation. FAULTS Low-angle and bedding faults and related tear faults are believed to be the dominant style of fault- ing. The mineralized breccia zone on the mineralized cuesta may be related to this. faulting. The upper and lower contacts of unit B are herein considered to be thrust faults throughout most of the map area. This faulting predates f3 folding. High-angle normal faulting is probably relatively insignificant. RECONNAISSANCE GEOLOGIC STUDY, VAZANTE ZINC DISTRICT, MINAS GERAIS, BRAZIL 19 The Santa Catarina fault, which separates units A and B, marks a major contrast between the pro- nounced and pervasive penetrative deformation and low-grade metamorphism of unit A rocks and the generally locally developed and much weaker low- grade metamorphism of unit B rocks. Beds on op- posite sides of the fault are generally more or less concordant; however, minor but important discord- ances do occur. In the basal member of unit B, several marker horizons traceable on aerial photo- graphs are truncated toward the north at a. low angle against the fault. The thickness of this basal member is also reduced by at least one half from the southern edge of the map to the point where the marker beds are truncated. The contact between unit B and unit C is thought to be a low-angle thrust and bedding fault, herein named the Vazante fault. We consider the contact to be one continuous fault that is a low-angle thrust southwest of the north-plunging anticline about which it is folded; the thrust climbs up the strati- graphic section to the northeast or east and becomes a bedding fault east and northeast of the anticline. Structures in unit B, south of the anticline, consist of a series of broad, open folds that apparently in- Volve the same beds, indicating that the dip of unit B is very low in this area. In this same area, unit C dips uniformly and more steeply than unit B and is not folded, thus indicating a structurally discordant relationship across» the contact. Strata along both sides of the fault in this interval are discordant and irregular, probably indicating drag folding and re- lated faulting. East and northeast of the anticline, beds in both units are more or less parallel to each other. The contact is a folded surface as evidenced by the general continuity of strata in both units. Transverse high-angle normal faults have been used by previous workers to explain minor offsets or suspected offsets of strata in unit B. However, on aerial photographs numerous marker horizons can be seen to be continuous both above and below some of these offsets. Thus, the faulting that is re- sponsible must be confined by these marker beds. We suspect that multiple bedding faults and step- bedding faults formed, and differential movement of the intervening beds created tear faults that ter- minated at the bounding bedding faults. In this fashion, an uptilted sequence of rocks would have tear faults that offset only a few beds in a limited stratigraphic interval ;these tear faults would have all the characteristics of high-angle normal faults but would terminate abruptly at overlying and un- derlying layers. Two examples of this type of tear-fault relation- ship Will illustrate the point. About 2 km southeast of Vazante along the road to Patos de Minas, a low ridge, on the north side of the road, is underlain by at least 20 m of northwest-dipping limestone. This limestone is not present in the drainage ditch along- side the road, a distance of about 10 m from lime- stone outcrops. Instead, only folded shale is present in the ditch and the roadbed. A second example can be seen in the streambed that crosses the mineralized cuesta about 0.75 km west of the airstrip at Fazenda do Inga. Northeast of the streambed, the cuesta is underlain by 50—80 m of massive to thick-bedded limestone, whereas in, the streambed and to the southwest along the cuesta moderately folded silt- stone and shale are present. The change from lime- stone to siltstone is limited to an interval of about 10 m. In both areas, bounding strata are not offset. Thus, tear faulting bounded by bedding-thrust faults best explains the relationships. Facies changes may have been the dominant factor in controlling the location of such tear faults. The mineralized cuesta has been interpreted as a northwestward tilted fault block by previous workers (Biranco, 1962; Carvalho and others, 1962; Amaral, 1968). If their interpretation is correct, the prominent limestone beds underlying the cuesta might be expected to crop out between the: cuesta and the Santa Catarina fault, which they do not. An alternative explanation, the one favored in this report, is of a poly-folded and thrust—faulted se- quence of interbedded resistant and less resistant strata, which have eroded to ridge-and-valley-type topography. The mineralized breccia. zone, on the dip slope of the mineralized cuesta, has also been considered by previous workers (Moore, 1956; Branco, 1962; Carvalho and others, 1962; Guimaraes, 1962; Amaral, 1968) as being due to normal faulting. The breccia zone ranges from less than 1 m to 200 m in width; the bounding faults are very steep, dipping both to the northwest and southeast. There are sev- eral possible explanations for the formation of the breccia zone. The one favored here is that the breccia formed along the step of a step-bedding fault, herein named the Mineira fault, that originated during the major period of Brazilian tectonism. Both the brec- ci-a zone and the regional dip, which is 30° to 45° southwest, are folded about f.1 fold axes. Removal of the regional dip and f4 folding would change the breccia zone from a curved and near vertical struc- ture to a linear, moderately steeply southwest-dip- ping structure like that of a step~bedding fault (fig. 110 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 / 4 ’/ 4/ FIGURE 6,—Diagrammatic examples of ways in which breccia zones may form along step-bedding faults. A, Tilted, steep step-bedding fault resembling a normal fault. B, Tilted, low-angle step-bedding fault. C, Breccia zone cut by younger low-angle fault. D, Multiple breccia zone along a single fault. E, Two breccia-zone step-bedding faults, one cutting the other. F, Several stages of faulting and brecciation. 6A). This structure would have developed prior to or coeval with f1, f2, and (or) f3 folding but definitely before f4 folding. Some variations in the possible forming of a step-bedding—plane breccia zone are shown in figure 69Mu1tiple steps could form. Or, one or a series of steps might be cut by a low-angle fault, thus planing off and moving the upper por- tion of the breccia zone some distance from the IOWer portion, which in turn would be overlain by unbrecciated “upper pla ” rocks. Other interpreta- tions for the origin of the breccia include: a high- angle fault, as proposed by previous workers; a fault-zone vein that predated f4 folding but was postthrusting, and perhaps was of a rebound ten- sional nature; a faulted and metamorphosed Mis- sissippi-valley type or similar type of “stratabound” deposit; a faulted and metamorphosed massive sul— fide syngenetic deposit. MINERALIZATION AND RESERVES The zinc deposits occur in breccia zones along the Mineira fault, primarily along the dip slope of the mineralized cuetsa (fig. 7). These breccia zones are variable in width and depth; they range from 1 to 200 m in Width and are at least 80 m deep. A wide variety of minerals are present and are listed in decreasing order of abundance: hemimorphite (re- ferred to as calamine in Brazilian reports of the district), willemite, hydrozincite, cerussite, smith- sonite, chalcocite, brochantite, pyromorphite, covel- lite, zincite, cuprite, native copper, malachite, linarite, autrichalcite, acanthite, and native silver (Amaral, 1968). The main ore minerals are hemi- morphite, willemite, and hydrozincite. Secondary zinc silicate minerals extend downward to at least 80 m. Sulfide minerals are recognized at greater RECONNAISSANCE GEOLOGIC STUDY, VAZANTE ZINC DISTRICT, MINAS GERAIS, BRAZIL FIGURE 7.—Resistant ridge of mineralized breccia (primarily hemimorphite, hydrozincite, willemite) on the north end of Serra do Pogo Verde. Trees do not grow over the mineral- ized zone. depth, but their limits are poorly defined, and they apparently occur in irregular small bodies. Proven reserves on the Companhia Mineira de Metais concession are about 6,200,000 metric tons of 16-percent zinc. ore (largely hemimorphite ore) and an additional 800,000 metric tons of 40-percent zinc ore (largely willemite ore), or about 1,312,000 metric tons of zinc metal (White and Nagell, 1975). Reserve data are not available for the Inga deposit but probably are of a similar size. Probable and possible reserves half again as large as proven re- serves seem to indicate that total zinc metal re- serves could be as much as 3 million metric tons. SUMMARY The Vazante district, which occurs along the lead- ing edge of a deformed belt, has been subjected to considerable compression, and the rocks have re- sponded accordingly. Low-grade dynamic meta- morphism of argillaceous rocks, broad folding, and low-angle and bedding faulting and associated tear faulting dominate the structural history. Almost all the structures in the Vazante district were formed probably during the Brazilian orogeny (ca. 600- 500 my.) . The breccia zone along the mineralized cuesta is only a very small part of the area, but is economical- ly important. The breccia probably formed in re- sponse to step—bed‘diing-plane faulting along the base of the carbonate rocks underlying the mineralized cuesta. Localization of the step-bedding faulting may have been controlled by earlier structures, such I11 as high-angle faults or collapse-breccia zones. The step-bedding—plane faulting resulted in the breccia— tion of carbonate and argillaceous rocks. The amount of stratigraphic separation is unknown but it prob- ably was not large. The Vazante deposit is considered to be epigenetic. The source of the mineralization is incompletely understood. We favor a source from syngenetic minerals in host-rock dolomitic limestone of the Paraopeba Formation. Cassedanne (1968b) arrived at such a conclusion for at least part of the min- eralization, pointing out that Collem’a-bearing dolo- mite in the district has an abnormally high zinc and lead content. Amaral and Kawashita (1967, 1968) dated shales at Vazante from the Sete Lagoas For- mation (which we refer to as the Paraopeba For- mation) as 600:22 my. old. Amaral (1968) re- ported an age. of 600:30 m.y. for galena at Vazante. The presence of abnormally high amounts of zinc and lead in certain carbonate facies coinciding with the similar ages for shale and galena in the district strongly supports a syngenetic interpretation for the source of the zinc and lead. Subsequent concentra— tion of the deposit into its present form as an epigenetic deposit probably was made: possible be- cause Brazilian orogenic events formed breccias that provided the “plumbing” system for the later migration of ore-bearing fluids. Redeposition in the breccia zone along the mineralized cuesta formed a deposit that was subsequently oxidized by surface or near-surface waters. During the redeposition and oxidiation phases, possibly during late Mesozoic or early Cenozoic time, the breccia was so altered that whether its origin is tectonic or solution cannot be determined. Solution and collapse features are com- mon in the breccia zone, but they reflect probably only the latest history of the zone. This interpretation suggests that future explora- tion for Vazanteetype deposits should be concen- trated along the deformed belt where folded car- bonate rocks are interbedded with argillaceous rocks. This coincidence of structural and strati- graphic situations is favorable for disharmonic folding, which, in turn, leads to bedding and step- bed-ding faulting. The end result of these events may have been the formation of bireccia zones that were potential depositional sites for orebearing fluids. REFERENCES CITED Amaral, Gilberto, 1968, Contribuicao ao conhecimento dos depésitos de Zn-Pb-Cu-Ag da Serra do Pogo Verde, Va— zante, Estado de Minas Gerais: Sociedade Brasileira I12 Geologia Congresso, 22d, Belo Horizonte, 1968,Anais, p.13—31. Amaral, Gilberto, and Kawashita, Koji, 1967, Idade do g'rupo Bambui: [abs.]. Parananaense Geociéncias B01. 26, p. 39—40. 1968, Idade do grupo Bambui por determinacées Rb-Sr em folhelhos: Sociedade Brasileira Geologia Congresso, 21st, Curitiba, 1967, Anais, p. 214. Barbosa, Octavio, Braun, Oscar P. G., Dyer R. C., Cunha, Carlos, and Rodigues, A. B., 1970, Geologia da regiao do Triangulo Mineiro: Brazil Divisao Fomento Producao Mineral, B01. 136, 140 p. Bohomoletz, M. L., 1975, Brazil’s metallurgical industry— after a slow start it now learns how to run: Engineering and Mining Journal, v. 176, no. 11, p. 176 to 176—D1. Branco, Jose Jaime R., 1962, Principais ocorréncias de zinco em Minas Gerais: Sociedade Intercambio Cultural e Estudos Geologia, Ouro Préto, pub. 2, p. 151—184. Braun, Oscar P. G., 1968, Contribuicao a estratigrafia do grupo Bambui: Soci‘edade Brasileira Geologia Congresso, 22d, Belo Horizonte, 1968, Anais, p. 155—166. SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 Carvalho, P., Dequech, D., and Guimaraes, D., 1962, Jazida plumbo-zincifera do Municipio de Vazante, Minas Gerais: Brazil Divi‘sao Fomento Producao Mineral Bol. 110, 119 p. Cassedanne, Jacques, 1968a, Description du biostrome 5. Col- lenias de la mine de Vazante (Minas Gerais) : Academia Brasileira Ciéncias Anais, v. 40, no. 2, p. 215-225. 1968b, Nota sobre o ambiente die sedimentacao das rochas encaixando a mine‘ralizacao de Vazante (MG): Sooiedade Brasileira Geologia Congrersso, 22d, Belo Hori- zonte, 1968, Anais, p. 33—40. Guimarées, Djalma, 1962, Génese do minério de zinco de Va- zante, Minas Gerais: Sociedade Intercambio Cultural e Estudos Geologia, Ouro Préto, pub. 2, p. 101—147. Moore, Samuel L., 1956, Zinc and copper deposits of the Vazante area, Minas Gerais, Brazil: U.S. Geological Sur- vey open-file report, 16 p. White, M. G., and Nagell, R. H., 1975, Lead and zinc re- sources of Brazil: U.S. Geological Survey Open—File Re— port 75—49, 49 p. SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 CONSTRAINTS ON THE LATEST MOVEMENTS ON THE MELONES FAULT ZONE, SIERRA NEVADA FOOTHILLS, CALIFORNIA By J. ALAN BARTOW ABSTRACT K-Ar dates of 4.12 million years (m.y.) and 4.03 m.y. for two small dacite and rhyodacite intrusive bodies neazr Moke- lumne Hill, Calif, provide important age control for the late Tertiary Mehrten Formation in this area because the Mehrten both is intruded by the dacite and rhyodacite and depositionally overlies the dacite of one intrusion. These dates also establish a maximum age for the latest movement on the Mother Lode fault, a part of the Melones fault zone, because the Mehrten Formation is displaced by this fault and because younger eolluvium containing clasts from one of the intrusive bodies is in fault contact with the Mehrten. INTRODUCTION Four small dacitic to rhyodacitic intrusive bodies near Mokelumne Hill, Calaveras County, Calif., (fig. 1) are closely associated with the Mehrten Forma- tion, a late Tertiary volcaniclastic unit. Inasmuch as the dacite and rhyodacite bodies intrude the Mehrten and are locally overlain by it, dating the in- trusions would give an age within the Mehrten For- mation in this area. This age has added significance because recent work in the Sierra Nevada foothills belt by Woodward-Clyde Consultants for the Pacific Gas and Electric Co. has revealed that the Mehrten Formation is displaced by the Mother Lode fault (fig. 1) near the intrusive bodies. The base of the Mehrten is offset by as much as 30 m (down to the northeast) (J. R. McCleary, oral commun., 1976), and a trench excavated across the Mother Lode fault adjacent to Hamby dome (visited by the author) shows that the highest exposed Mehrten on the ridge crest is juxtaposed with younger eolluvium containing clasts of rhyodacite apparently identical with the rhyodacite of the Hamby dome intrusion. The minimum offset on the eolluvium-Mehrten con- tact is about 5 m (T. A. Grant, oral commun., 1977). ACKNOWLEDGMENTS R. L. Blum, of the Pacific Gas and Electric Co., and J. R. McCleary and T. A. Grant, of Woodward- Clyde Consultants, have kindly shared information resulting from their investigations in the Sierra Nevada foothills fault zones. Dennis Sorg prepared mineral separates, and Marvin Lanphere and J. C. Von Essen were helpful in obtaining the K—Ar dates. The manuscript has benefited from the comments of T. A. Grant. AGE OF THE MEHRTEN FORMATION The Mehrten Formation, a widespread and easily recognized stratigraphic unit in the western Sierra Nevada, is composed principally of andesitic elastic and pyroclastic material. The formation was orig- inally described by Piper, Gale, Thomas, and Robin- son (1939) in the lower Sierra Nevada foothills, west of the area of this report, where the formation is composed mostly of sandstone and minor amounts of conglomerate, mudstone, and breccia. In the Mo- belumne Hill area, the formation consists mainly of coarse andesitic conglomerate and minor interbedded sandstone (Rose, 1959). The Mehrten unconform- ably overlies either rhyolitic tuff and conglomerate of the Valley Springs Formation (late Oligocene and early Miocene) or pre-Tertiary bedrock consisting of deformed metasedimentary rocks, metavolcanic rocks, and plutonic igneous rocks. Piper, Gale, Thomas, and Robinson (1939) con- sidered the age of the Mehrten Formation, on the basis of sparse fossil evidence, to be late( ?) Miocene and Pliocene( ?). More recent work has established that the Mehrten contains a vertebrate fauna of early Clarendonian age below the latite flow of Table J1 J2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 i20°45' 120°42’30” AMADOR COUNTY i \Viw CALAVERAS C(lUNTY 38° ~IthrL/w flew, 333$qu ”0‘ i KILOMETER EXPLANATION 1:1:Tr23'5 Rh odacite , Y Phocene ccia . E 8 III }Pilocene and S N ‘ > . _ _ _ Mehrten Formation Miocene E ?S 38° . V m u; 1 5. ll!” Miocene and l— u Volley Springs Formation Oligocene lone (?) Formation }E°cene / :1 l Granltic rocks ' °‘ ‘-’ ......... JM. . Q ,0 _ . ‘ . Meta- La, 2', ...... other meto- volcanic m 8 sedimentary rocks g 2 ....... rocks a , .............. ‘ . — ' ............... rley Reservozr 8 9 '::'.::'.‘. Com/eras *2“: ....... Formation 0— N ........... c,— — _ ._., x SWIG-3|) j I : I : : I , , . , 3 Contact Fault Trench K-Ar sample ....... locality FIGURE 1.—Geologic map of the area southwest of Mokelumne Hill, Calaveras County, Calif. Geology modified from Rose (1959) and Goldman (1964). Base from US. Geological Survey, Jackson, Mokelumne Hill, San Andreas, and Valley Springs 7.5 minute quadrangle. CONSTRAINTS ON MOVEMENTS, MELONES FAULT ZONE, SIERRA NEVADA FOOTHILLS Mountain, about 15 km southeast of the map area (Stirton and Goeriz, 1942) , and one of late Hemphil- Ian age above the latite (Stirton and Goeriz, 1942, Wagner, 1976) indicating that the Mehrten is Mio- cene and Pliocene. Correlation of K—Ar dates with the North American mammalian chronology by Evernden, Savage, Curtis, and James (1964) indi- cates that the lower part of the Mehrten is at least 11 to 12 m.y. (million years) old and that the top must be as young as 4 to 6 m.y. K—Ar dating of Cenozoic volcanic rocks in the Sierra Nevada (Dal- rymple, 1963, 1964) provides further control on the age of the Mehrten Formation (fig. 2). The date of 9.0 m.y. for the latite of Table Mountain, which occurs as a tongue in the Mehrten Formation, is in good agreement with the vetebrate ages. AGE OF DACITE INTRUSIONS The petrology of the intrusive bodies has been described by Rose (1959) and is not repeated here. Rose classified the larger bodies of Golden Gate Hill and Tunnel Peak (fig. 1) as hornblende-biotite dacite and hornblende dacite, respectively. The smaller two WESTERN FOOTHILLS (Stanislaus and Mokelumne River MIDDLE ALTITUDES (Stanislaus and Tuolumne River drainages) drainagcs) . . _Oakdale launa Andesmc breccia 39‘1“?" (late Hemphillian) ’ , - ~ Lame se Imekn cry 1 , , ’ ’ ’ (Dardanelle facies) '0": s _,_;_:_:__—_—_-; ; : : : Z I _ iolile-augile aliler< 9,0m.y. Lallle of Lalile of 9.2m.y. Table Mountain Table Mounlain —9.0m.y. C — _______________ g —Two-Mile Bar fauna 3 (early Clarendonian) g Andesmc 5 Andesilic '3 u. d' E luff, E’ 58 Imenlary LE) breccia, é rocks E: and b and, E, sedimentary reCCIo '5 rocks 5 W 19-9m-Y:::::::::m V, Upper luff 21 .9m.y Upper luff -21.1m.y. 2’ c_—— ‘ """""""" —— .22 fig Lower luff TE2;8_m;y ___________________ Lower luff —23.1m.y. i3 .2 > W _______________ . Lalile welded luff— 33.2m.y. lone Formalion Aurilerous gravel W W Basemenl rocks Basemenl racks FIGURE 2.—Cou'relation of some Tertiary stratigraphic units in the western Sierra. Nevada, showing age control from K-Ar dating of volcanic rocks (from Dalrymple, 1964). J3 TABLE 1.—Chem1'4:al data on late Tertiary intrusions [X—ray spectrographic analyses by L. F. Espos and G. M. Kawakita. Chemical analyses (*) by M.J. Cremer. Cr, Ni, Ba, Rb, Sr, V, Sc, and Y in parts per million;others in weight percent] Hamby dome Golden Gate Hill GV76-30 GV76-33 Si02 ------------- 67.92 59.97 A1203 ------------ 16.10 15.63 Fe203 ------------ 1.24 5.00 Fe0* ------------- .52 < .1 MgO -------------- .90 4.46 Cao -------------- 2.24 5.07 Na20 ------------- 6.09 4.24 K20 -------------- 2.97 .78 T102 ------------- .33 .81 P205 ------------- .10 .27 14110 -------------- .032 .076 H20+* ------------ .89 1.05 H20-* ------------ .21 .25 co * ------------- .03 .02 S (total) -------- .001 .006 Cr --------------- 55 155 Ni --------------- 26 113 Ba --------------- 960 1250 Rb --------------- 83 64 Sr --------------- 1600 1650 V ---------------- 27 116 Sc --------------- 5 15 Y ---------------- <5 6 intrusive bodies, referred to as Hamby dome and McSorley dome, were classified as hornblende rhyo- dacite. New chemical analyses (table 1) indicate that the Hamby dome intrusion is rhyodacite and the Golden Gate Hill intrusion is dark rhyodacite in Rittmann’s (1952) classification. According to Rose’s (1959) interpretation of the structural relations, the dacite of Tunnel Peak and the two small rhyodacite bodies are clearly intrusive into and therefore younger than the Mehrten For- mation. At McSorley and Hamby domes, the Gala- veras Formation has been dragged up between the rhyodacite and the surrounding Mehrten Formation, and the Mehrten has been deformed adjacent to the intrusive bodies. The dacite of Tunnel Peak is sur- rounded by a tuff breccia, interpreted by Rose (1959) as a vent breccia, that contains Mehrten andesite pebbles. More recent mapping in the area by Woodward-Clyde Consultants has shown that although the Mehrten has been intruded by the dacite of Tunnel Peak, the uppermost part of the Mehrten depositionally overlies part of the dacite (T. A. Grant, written commun., 1977) . Rose (1959) believed that the dacite of Golden Gate Hill, unlike the other small intrusions in the J4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 TABLE 2.—Analytical data for K-Ar age determinations [CV76-30 - Rhyodacite of Hamby dome; GV76-34 — Dacice of Golden Gate H111. Potassium measurements: J. H. Tillman. Argon measurements and age calcu- lation: E. Sims. 0K decay constants: X5 = 0.572 x 10-10 yr'l, 19' = 8.78 x 10-13 yr-l, A3 = 4.962 x 10—10 yr-l. Abundance ratio: “OK/K = 1.167 x 10'“ atom percent.] “DAV‘ HON-rad Calculated 7 Percent rad Ar age Field No. Mineral K20 (moles/g) total (106 years) GV76-30 Hornblende 1.061 6.283X10-12 0.20 4.12:0.11 1.057 owe—31. Hornblende 1.186 6. 900x10'12 o. 33 4.03:0. 26 area, was older than the Mehrten Formation or equivalent to the oldest part of the Mehrten in the area. As evidence, he cited the large angular blocks of dacite “identical to that which forms Golden Gate Hill” (Rose, 1959, p. 14), that were found in the basal part of the Mehrten Formation just southeast of the hill. A sample of dacite from Golden Gate Hill that was presumed to be older than the Mehrten Forma- tion and one of rhyodacite from Hamby dome that was presumed to be younger than the Mehrten were selected for K—Ar dating. The results, summarized in table 2, indicate that both intrusions are only a little more than 4 m.y. 01d and that the ages are analytically indistinguishable. DISCUSSION The age of 4.12 m.y. for the rhyodacite of Hamby dome, together with the structural relations de- scribed by Rose (1959) for Hamby and McSorley domes, suggests that the Mehrten Formation in this area is older than 4 m.y. If all the intrusions are approximately the same age, however, the structural relations at Tunnel Peak indicate that the date must fall within the Mehrten, probably within the upper part. This age for the Mehrten Formation is in fair agreement with‘the upper limit of 4 to 6 m.y. in- ferred from the vertebrate ages. The dacite of Golden Gate Hill (4.03 m.y.) may be considered to be the same age as the rhyodacite of Hamby dome and must also be approximately equiv- alent in age to the upper part of the Mehrten For- mation. The talus blocks of dacite that Rose (1959) interpreted as being within the basal part of the Mehrten might be better interpreted as landslide debris or younger colluvium incorporating dacite talus with debris from the Mehrten Formation. A more important conclusion to be drawn from the data presented here is that the latest movement on the Mother Lode fault, one of several faults mak- ing up the Melones fault zone of the Sierra Nevada foothills metamorphic belt, was more recent than the youngest part of the Mehrten Formation, which is about 4 to 6 m.y. old. Because colluvium contain- ing blocks of rhyodacite from Hamby dome is in- volved in the faulting, the latest movement must be more recent than 4 m.y., but no other time con- straints can be placed on the movement at this site. REFERENCES CITED Dalrymple, G. B., 1963, Potassium-argon dates of some Ceno— zoic volcanic rocks of the Sierra Nevada, California: Geological Society of America Bulletin, v. 74, p. 379—390. 1964, Cenozoic chronology of the Sierra Nevada, Cali- fornia; California University Publications in Geological Sciences v. 47, 41 p. Evernden, J. F., Savage, D. E., Curtis, G. H., and James, G. T., 1964, Potassium-argon dates and the Cenozoic mam- malian chronology of North America: American Journal of Science, v. 262, p. 145—198. Goldman, H. B., 1964, Tertiary fluviatile deposits in the vicinity of Mokelumne Hill, C‘alaveras County, California: Los Angeles, California University MA. thesis, 56 p. Piper, A. M., Gale, H. S., Thomas, H. E., and Robinson, T. W., 1939, Geology and ground-water hydrology of the Make— lumne area, California: US. Geological Survey Water- Supply Paper 780, 230 p. Rittmann, A., 1952, Nomenclature of volcanic rocks proposed for the use in the catalogue of volcanoes, and key-tables for the determination of volcanic rocks: Bulletin of Vol- canology, ser. 2, v. 12, p. 75—102. Rose, R. L., 1959, Tertiary volcanic domes near Jackson, California: California Division of Mines and Geology Special Report 60, 21 p. Stirton, R. A., and Goeriz, H. F., 1942, Fossil vertebrates from the superjacent deposits near Knights Fenry, Cali- fornia: California University Publications in Geological Science, v. 26, no. 5, p. 447—472. Wagner, Hugh, 1976, A new species of Pliotax’idea (Musteli- dae: Carnivore.) from California: Journal of Paleontology, v. 50, p. 107—127. fiU.S. GOVERNMENT PRINTING OFFICE: I980 0—— 3ll-344/24 PROFESSIONAL PAPER 1 126-A PLATE 1 UNITED STATES DEPARTMENT OF THE INTERIOR 112°40' 112035, 1 GEOLOGICAL SURVEY 112°55' 112°50' 112°45’ ' 1 1 113°05' 113°00' / Whispering Fal/s Grotto €mk 6mg ka — 36°25 35°25 — EXPLANATION 75 Q) G “a? 5 Contacts mapped where they intersect river level. Members of Bright Angel Shale and Muav Limestone that exhibit anticlinal structures at and above river level are identified by appropriate symbols. €mk -2 3/ Fish tail Rap/d5 6mg Cm ka Cm” Middle _ Cambrian LOWER AND MIDDLE PARTS OF MUAV LIMESTONE Cmg, Gateway Canyon Member ‘ €mk, Kanab Canyon Member ‘ €mp, Peach Springs Member 1 Cbsc — gonto _ CAMBRIAN Cba ebs “mp €br Middle Upset Rapids 35:51:16? BRIGHT ANGEL SHALE ~ and Lower 30-50 High— and low-angle faults €bsc, Spencer Canyon Member ‘ ’2 _ Cambrian become numerous €bs, Sanup Plateau Member ‘ >2 €br, Rampart Cave Member ‘ >2 ’ Ct "2 €mk20-35 33 / s 75 /\/ 6mg 35-40 TAPEATS SANDSTONE _ L _ ‘ Nomenclature of McKee (1945). 7 This member, included in the Muav Limestone by McKee (1945), is here included in the Bright Angel Shale to retain the base of the Muav Limestone as originally defined by Noble (1914, p. 65) in Bass Canyon, central Grand Canyon. Thrust faults not related to —~ 36°20’ tectonic fold 36°20' ~ Contact % 3. .__.I_ High—angle fault—Dashed where approximately located. Bar / 09 and ball on downthrown side t t i l Low—angle thrust fault closely associated with or on the exten- sion of river anticlines~Dot indicates location of fault where too small to be shown by line in cut. At places thrust faults pass upward into high-angle faults, several more than 10 m Fossil river anticlines about above river 30-60 m above river in €mk and Cmg 8 Reverse fault—R is on upthrown side t\ 5-7 ’1' Y + ' High-angle faults, well above 57/ Monocline river,pass downwardintolow- I - , _ . . 6’ angle thrust faults near river F0551] river anticlines about 9/3.? % River monocljne, displaying sharp bending of strata toward or level 3 30-60 m above river in ka (,0 away from river €mk emp 1% '5 and Cmg $0,?) t ka 0 1 Cmp $52? River anticline and its upriver and down-river limits—Limits / ‘Ktr \ \ 1 I dashed where approximately located t t/\1 / C ’7 r 3 2*: River syncli'ne (So 1 {K36 . . (a \ 3 Strike and dip of beds s 1-2 1 —L~ Inclined «Q, 63 Horizontal REFERENCES CITED Maxson, J. H., 1969, Preliminary geologic map of the Grand Canyon and vicinity, Arizona, western and central sections: ea Cmp Cmp ka/ Grand Canyon Natural History Association. Fe GI . McKee, E. D., 1945, Stratigraphy and ecology of the Grand m e" l \ 4 .. Canyon Cambrian, Part 1 of Cambrian history of the Grand (23 23 5\ National A Canyon region: Carnegie Institute of Washington, Publica- a: 5 ‘— O 5’ tion 563, p. 3-168. C / 5 E V‘ M 36°18 Noble, L. F., 1914, The Shinumo quadrangle, Grand Canyon 2 ‘3, 44 36°15 — DistriCt, Arizona: U.S. Geological Survey Bulletin 549, (3°: 7'2 100 'p. e / g Cbs ‘1' 9 . ‘0 23/ .. V l» R I V?- R 4 2 4‘” Ct Cba O 0 e \ X) 0 9 0 5 § 0 (3° . ~k § ‘17 § § Ct Cba 3 7 8 g / a / 7. / “g / a / O 0‘: / /4 Cbs / Cbr /5 Cbs 9 , 1 \3'5 3-5 / 5 Cmp 3 /. / 3 3 5 gigs l I I ‘ I 1 | 36°10' , r o , ° I 7 ° ' o ' o ' 112°45' 112°40’ 112 35 36 10 1 13005! 1 13 00 112 55 H2 50 InteriorwGeological Survey, Reston, Va.—l9797G79141 Planimetric base from U.S. Geological Survey topographic N14 7 SCALE 1:62 500 - Geology by D. p. Elston and R W. Huntoon, 1976. quadrangles Kanab Point. 1962; Tuckup Canyon, 1962; v 1 5 O 1 2 3 4 5 6 7 8 9 KILOMETERS Location of Sinyala Fault from Maxson (1969). National Canyon, 1962; Grand Canyon National E FrH ,‘4 H H . L r , Monument, W 1/3, 1936 g 1 5% 0 1 2 3 4 5 MILES % i-—l I—i H i—i i—-l r i———————‘{ l—-—-—-———i E ARIZONA Approximate mean declination, 1979 RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA 1979