Shorter Contributions to

Stratigraphy and Structural

J:

GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126-A-J
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, U.S. Government Printing Office

Washington, D.C. 20402

Stock Number 024-001-03285-6
CONTENTS

| Letters designate the chapters]

(A) Origin of the river anticlines, central Grand Canyon, Arizona, by Peter W. Huntoon 

and Donald P. Elston

(B) Rock control and tectonism their importance in shaping the Appalachian Highlands, by John T. Hack

(C) The Upper Ordovician and Silurian Hanson Creek Formation of central Nevada, 

by Rueben J. Ross, Jr., Thomas B. Nolan, and Anita G. Harris

(D) The Marble Hill Bed: an off share bar-tidal channel complex in the Upper Ordovician Drakes Formation of Kentucky, by W C Swadley

(E) Paleogene sedimentary and volcanogenic rocks from Adak Island, central Aleutian 

Islands, Alaska, by James R. Hein and Hugh McLean

(F) 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

(G) 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

(H) New evidence supporting Nebraskan age for origin of Ohio River in north-central 

Kentucky, by W C Swadley

(I) Reconnaissance geologic study of the Vazante zinc district, Minas Gerais, Brazil, 

by Charles H. Thorman and Samir Nahass

(J) Constraints on the latest movements on the Melones fault zone, Sierra Nevada 

foothills, California, by J. Alan Bartow
CONVERSION FACTORS

Metric unit Inch-Pound equivalent

Length

millimeter (mm) 

meter (m) 

kilometer (km)

= 0.03937 inch (in) 

= 3.28 feet (ft) 

= .62 mile (mi)

Area

square meter (m2 ) 

square kilometer (km2 ) 

hectare (ha)

= 10.76 square feet (ft2 ) 

= .386 square mile (mi2 ) 

= 2.47 acres

Volume

cubic centimeter (cm3 ) 

liter (L) 

cubic meter (m3 ) 

cubic meter 

cubic hectometer (hm3 ) 

liter 

liter 

liter 

cubic meter

cubic meter

= 0.061 cubic inch (in3 ) 

= 61.03 cubic inches 

= 35.31 cubic feet (ft3 ) 

= .00081 acre-foot (acre-ft) 

= 810.7 acre-feet 

= 2.113 pints (pt) 

= 1.06 quarts (qt) 

= .26 gallon (gal) 

= .00026 million gallons (Mgal or 

106 gal) 

= 6.290 barrels (bbl) (1 bbl = 42 gal)

Weight

gram (g) 

gram 

metric tons (t) 

metric tons

= 0.035 ounce, avoirdupois (oz avdp) 

= .0022 pound, avoirdupois (Ib avdp) 

= 1.102 tons, short (2,000 Ib) 

= 0.9842 ton, long (2,240 Ib)

Specific combinations

kilogram per square 

centimeter (kg/cm2 ) 

kilogram per square 

centimeter 

cubic meter per second 

( m3/s )

= 0.96 atmosphere (atm) 

= .98 bar (0.9869 atm) 

= 35.3 cubic feet per second (ft3/s)

Metric unit

Specific

liter per second (L/s) 

cubic meter per second 

per square kilometer 

[(mVs)/km2 ] 

meter per day (m/d)

meter per kilometer 

(m/km) 

kilometer per hour 

(km/h) 

meter per second (m/s) 

meter squared per day 

(m2 /d) 

cubic meter per second 

(m.3/s) 

cubic meter per minute 

(m3/min) 

liter per second (L/s) 

liter per second per 

meter [(L/s)/m] 

kilometer per hour 

(km/h) 

meter per second (m/s) 

gram per cubic 

centimeter (g/cm3 ) 

gram per square 

centimeter (g/cm2 ) 

gram per square 

centimeter

Inch-Pound equivalent

combinations Continued

= .0353 cubic foot per second 

= 91.47 cubic feet per second per 

square mile [(ft3/s)/mi2 ]

= 3.28 feet per day (hydraulic 

conductivity) (ft/a) 

= 5.28 feet per mile (ft/mi)

= .9113 foot per second (ft/s)

= 3.28 feet per second 

= 10.764 feet squared per day (ft2/d) 

(transmissivity) 

= 22.826 million gallons per day 

(Mgal/d) 

= 264.2 gallons per minute (gal/min)

= 15.85 gallons per minute 

= 4.83 gallons per minute per foot 

[(gal/min) /ft]

 .62 mile per hour (mi/h)

 2.237 miles per hour

= 62.43 pounds per cubic foot (lb/ft3 ) 

= 2.048 pounds per square foot (lb/ft2 ) 

= .0142 pound per square inch (lb/in2 )

Temperature

degree Celsius (°C) 

degrees Celsius 

(temperature)

= 1.8 degrees Fahrenheit (°F) 

= [(1.8x°C)+32] degrees Fahrenheit
Origin of the River Anticlines, 

Central Grand Canyon, Arizona

By PETER. W. HUNTOON and DONALD P. ELSTON

SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND 

STRUCTURAL GEOLOGY, 1979

GEOLOGICAL SURVEY PROFESSIONAL PAPER 1126-A

CONTENTS

	Page

Abstract _________________________________________------—— Al

Introduction ___________________________________________ — ____ — — _— 1

Acknowledgments ________________________________ — ____— __ 1

Geographic and stratigraphic location _________________ ————— _— 2

Structure ____________________________-__ —— ———— - 2

Low-angle thrust faults ______________________________ — _______ 4

High-angle kink bands ____________________________________ — _ — __ — __ 5

Related anticlines ___________________________________________ — _ 5

Ancient river anticlines __________________________________________ 5

Regional distribution __________________________________-____ 5

Origin ___________________________________________ 6

References cited _________________________________________________ 9

ILLUSTRATIONS

Page

PLATE 1. River anticlines, central Grand Canyon, Arizona _________________________________ — _ In pocket

FIGURE 1. Map showing the location of the Colorado River and its principal tributary canyons in the central Grand

Canyon, Ariz ______________________________________________________________—______ A2

2. Stratigraphic columns of Cambrian rocks in the central and eastern Grand Canyon _____________ — _— 3

3. Photograph of well-developed river anticlines upstream from Olo Canyon _________________________ 4

4. Photograph of a pair of low-angle echelon thrust faults________________________________________ 5

5. Photograph and diagram showing kink bands in rocks along the river ___________________________ 6

6. Diagram illustrating selected thories used to explain the origin of river anticlines in the central Grand

Canyon ____________________________________________________________ 7

in

SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 abundant 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 comparatively 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 appears to weaken the shaly rocks sufficiently to allow deformation 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 relation 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

1 Department of Geology and Wyoming Water Resources Research Institute. University of Wyoming.

the origin of the anticlines to mass flowage of saturated shaly Cambrian rocks in response to horizontal 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 development 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 Colorado River, below its confluence with the Green 

River in Utah).

ACKNOWLEDGMENTS

We thank the U.S. National Park Service for allowing us to work and sample in Grand Canyon National Park and gratefully acknowledge the kind 

cooperation of David C. Ochsner and Robert 

Yearout.

Funds supporting work by Huntoon were provided by the Office of Water Research and Technology through the Wyoming Water Resources Research Institute under matching grant 14-34-0001- 

6134. The U.S. Geological Survey work of Elston 

was carried out under magnetostratigraphic investigations of the Cambrian strata of northern Arizona. James I. Drever, Department of Geology, University 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 U.S. Geological Survey reviewed the manuscript and made many helpful suggestions. Shoemaker provided a sketch that led to figure 6.

Al
A2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

GEOGRAPHIC AND 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 underlying 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 anticlines 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 subordinate 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). Between 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 northwesttrending 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

FIGURE 1.—Location 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 A3

EXPLANATION

CENTRAL 

GRAND CANYON

MARBLE CANYONEASTERN GRAND 

CANYON DEVONIAN

Shale-claystone

Mixed lithologies

indicated

by superposition

of patterns

METERS 

,-150

-100

-50

-0

Upper 

member

Havasu 

Member

Gateway 

Canyon 

Member

Kanab 

Canyon 

Member

Peach 

Springs 

Member

<D 

C 

O 

"<7> 

<D

y E

>

O

o

C/5

g>

CO

Y

Sandstone

Tapea

Q.

e <

Cki

PRECAMBRIAN

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 Dressier 

(1977, fig. 2). Member names within Muav Limestone are those of McKee (1945).
A4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

(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 

"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 particularly 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 entirely restricted to the Cambrian rocks.

The maximum structural relief along the anticlines 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.

LOW-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

river. The faults are particularly numerous and 

have greater stratigraphic offsets where the limestone 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 extend more than 100 m along the strike of an anticline. As shown in figure 4, the faults characteristically terminate in breccia zones. Parallel to the 

river, the fault planes curve in conformance with 

the structural relief along an anticline. The resulting 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 dipslip 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 intersect 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 Olo Canyon. View is downstream toward west.
ORIGIN OF THE RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA A5

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 locally parallel the river anticlines. As shown on figure 5, the kink bands displace the rocks up toward 

the river and are as wide as 2 m. The bands approximately parallel the slope of the canyon walls 

and are obliquely to vertically jointed in the limestone. 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 ANTICLINES

Small river anticlines occur in the perennially wet 

reaches of Kanab Canyon and a few of its tributaries (fig. 1). The crests of these anticlines are locally 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 Tuckup 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 parallel to the river. Because the rocks are deformed 

and the structures lie above the--present 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 

tho 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. Consequently, 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. Significantly, 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 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

A

o

B

FIGURE 5.—Kink bands in rocks along the river. A, A highangle 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 attribute 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 section decreases from west to east. Where the river 

flows on these units in Marble Canyon and the 

eastern Grand Canyon, the Bright Angel Shale consists 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 summarized 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) considered that the Bright Angel Shale is mobile 

and is squeezed out from under the adjacent 

walls as a result of the stress gradients produced 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 combination 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 being largely accommodated by tensional separations 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 anticlines 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 A7

A. Distributed flowage resulting from release of confining pressure.

8. Flowage in incompetent shale layers due to differences in lithostatic load.

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 Grinshpan (1975). C, Shoemaker (1973). Combination of figures A and B, Hamblin and Rigby (1969). Arrows indicate movement caused by stress. pC, Precambrian igneous 

and metamorphic rocks. -Ct, Tapeats Sandstone. -Cba, Bright 

Angel Shale. Cm, Muav Limestone. Dtb, Temple Butte 

Limestone. Mr, Redwall Limestone. PPs, Supai Group, 

capped by Esplanade Sandstone.

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 stratigraphically high in the Muav Limestone, well above 

present river level (pi. 1). The fossil anticlines are 

found where the central gorge of the canyon is very 

narrow. They do not appear to be physically connected to underlying river anticlines found at 

present river level. We consider this evidence that 

formation of the river anticlines has been a continuing 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 necessarily related to the presence of a closely underlying 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 collected during the course of our investigation. One 

fact is clear from field observations: The strain observed 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 nonelastic strain.

Huntoon's swelling-clay hypothesis was based on 

observations in which clay partings between individual 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 Limestone 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 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

swelling clays in all samples, we believe that the 

swelling-shale hypothesis can be rejected. The expanded beds observed at Whispering Falls Grotto 

thus are undoubtedly the result of physical separations 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 northeasttrending Sinyala fault that parallels the general 

trace of the river on the south (pi. 1). In this model, 

the plate moves northwestward toward the river on 

a dipping surface that slopes less than 2° throughout 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 perpendicular to the direction of plate movement.

Shoemaker's hypothesis would explain the tendency for the anticlines to aline along the northeasttrending segments of the river. However, we find 

difficulties with the theory. First, we have been 

unable to document tensional separations along joints 

parallel to the river in the cross sections along tributary 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 northwesttrending segments of the river. Thirdly, anticlines 

that we have observed near the upstream and downstream ends of the river anticline belt (fig. 1, pi. 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 6C. We thus conclude that there is no field evidence for appreciable 

lateral movement of the Paleozoic section toward 

the Colorado River.

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 related directly to differential lateral flowage of the 

rocks under the canyon walls.

Saturation of the shale is essential for the deformation. The lack of flowage features in shaly outcrops high in the walls of the Grand Canyon or 

away from perennial streams supports this conclusion. Hundreds of kilometers of such dry outcrops 

exist in the Grand Canyon under load conditions 

identical to those near the river anticlines, yet anticlinal 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 northeasttrending stretches of the river and disappear along 

the short segments of the river that trend northwest 

(fig. 1, pi. 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 possibility exists that some small component of northwestward, downdip translation of the Paleozoic section 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 northeasttrending stretches of the river and suppress their 

formation along the northwest-trending stretches. 

The differential stress would augment stresses produced as a consequence of unloading along the river, 

which in turn suggests that a preferred orientation 

of stream or river anticlines might be a comparatively sensitive indicator of differential stresses of 

tectonic origin.
ORIGIN OF THE RIVER ANTICLINES, CENTRAL GRAND CANYON, ARIZONA A9

REFERENCES 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 Roat, 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.

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. 3-168.

Maxon, J. H., 1969, Preliminary geologic map of the Grand 

Canyon and vicinity, Arizona (western and central sections) : Grand Canyon Natural History Association 

(P.O. 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).

Sturg-ul, 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

CONTENTS

Page

Abstract ______________________________________________ Bl

Introduction ____________________________________ 1

Major topographic features and drainage systems _________________ 2

Relation of the highland form to bedrock geology ________________ 2

The belt of Cambrian and Ordovician rocks _____________________ 6

Duration of erosion _____________.____________________-___________ 8

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

ILLUSTRATIONS

Pago 

FIGIT EE 1. Generalized 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

2. Subenvelope map showing contours drawn on altitudes along the major stream channels ____—_____ 4

3. Generalized map of major rock units based on their resistance to erosion ___________________ 5

4. Profile and outline map of the Cambrian and Ordovician belt from Coosa River to Lake Champlain _—_____ 7

5. Generalized relief map of Piedmont and Blue Ridge provinces ____________—______ — _____ 10

6. Map of Eastern United States showing Bouguer gravity anomalies ____________— —— __ — ______ —— ——— 14

TABLES

Page 

TABLE 1. Thickness of sediments penetrated in COST No. B-2 and estimated deposition rate ________—_—____ B9

2. Estimates of erosion rates ____________________________________________________ 12

3. 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 United 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 different resistance rather than to tectonic processes. Most areas 

of high relief and high altitude have heen formed on resistant 

rocks. The Cambrian and Ordovician belt, containing mostly 

shale and carbonate rock, on the other hand, forms an extensive lowland from Alabama to the Canadian border and girdles 

the Adirondack Mountains. Differences in altitude can be explained by the presence of resistant rocks outside the belt; 

these resistant rocks form local base levels 011 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 inferred 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 sufficient 

to mask evidence of faulting or other differential movements. 

The high rates of uplift that are inferred 011 tilted water 

planes in the glaciated regions or that are measured by differences 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 surfaces such as peneplains, and the evolution of the 

drainage pattern (Thornbnry, 1965, p. 72-S7). Although most geologists agree that uplift did take place 

in the Eastern United 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; Mixon and Xewell, 1977), but it has 

not been proved that any modern topographic features owe their origin to contemporary faulting.

Recent studies of precise leveling data along geodetic 

traverses that have been rerun one or more times indicate 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 differences 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 depression 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 difficulties 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 differential erosion of rocks of different 

resistance is the major factor that controlled the development of today's topographic forms, including 

even large features. A few major topographic discontinuities 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 drainage systems that have different base levels, lengths, 

and gradients, which are basically caused by rock control within the different systems. Some may be due to 

tectonic influence, and others may be inherited from 

the past.

Bl
B2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

The rate of topographic adjustment because of erosion 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 evidence of tectonism. The intent of this paper is to describe the major drainage systems and topographic 

forms of the Appalachian Highlands between Alabama and northern New York and to discuss their adjustment 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 Wayne Newell for their thoughtful criticism.

MAJOR TOPOGRAPHIC FEATURES AND 

DRAINAGE SYSTEMS

The general outlines of the topography and drainage of the Appalachian Highlands are shown in figures 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, figure 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 topography. In this map, the contours are drawn on altitudes 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 generalization 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 deterhiines 

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 1:2500000 scale. 

Contours were plotted only on streams more than 32 

km long. The resulting map showT s the altitudes to 

which the major streams have eroded. The true altitudes 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 reentrants. 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 southeast to northwest. Large reentrants occur at the Potomac River basin (P) and Susquehanna River basin 

(S). The Mohawk Valley at M separates the Adirondack Mountains from the main highland area.

Many of the forms are closely related to the kinds of 

bedrock directly beneath them. This relationship probably accounts in large part for the features just mentioned, thus obscuring the evidence for the interpretation 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, 1973, 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 Appalachian 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 contains alternating sequences of resistant and nonresistant rocks, which, where folded, produce the elongate ridges and valleys of the Valley and Ridge
ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS B3

400 KILOMETERS

EXPLANATION

1000-2000 

305-610

2000-3000 

610-915

3000-4000 

915-1220

Over 4000 feet 

Over 1220 meters

—— — — — — Continental divide

FIGURE 1.—Generalized topographic map (en- of Mexico and St. Lawrence. Localities revelope map) of the Appalachian High- ferred to in text include: A, Adirondack

lands, showing the divide between the 

drainage to the Atlantic and to the Gulfs

Mountains; BR, Blue Ridge of northern 

Virginia.
B4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

32° .

4QO MILES

0 200 400 KILOMETERS

Contour interval 200 feet (61 meters)

— — — — Drainage divide

FIGUBE 2.—Subenvelope map showing contours drawn on altitudes along the major 

stream channels. Altitudes on streams less 

than 32 km long as measured from the

province. Unit 3 contains mostly nonresistant shale 

and carbonate sequences. The four units are defined 

as follows largely on the basis of summary by Colton, 

1970) :

Unit 1: Includes the entire sequence of rocks of 

Pennsylvanian age as well as the Upper Devonian of

source are not included. Oh, Chattahoochee 

River; T, Tennessee River; N, New River; 

P, Potomac River; S, Susquehanna River: 

M, Mohawk River.

New York and Pennsylvania and the Permian of West 

Virginia. The rocks of Pennsylvanian age are clastic 

sequences, both marine and continental, and contain 

coal beds. They include thick sandstone, especially in 

the lower part of the section. These more resistant 

rocks crop out near the eastern edge of the Appala-
ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS

88° 84° 80°

B5

0 200 400 KILOMETERS

Contour interval 400 feet (122 meters)

EXPLANATION 

Unit 1—Mostly rocks of Late Devonian and Pennsylvanian age

t^tl^tl Unit 2—Mostly rocks of Silurian to Middle Devonian age 

t^V;;!^v!y;!-j Unit 3—Sedimentary rocks of Cambrian and Ordovician age 

R^^j^l 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. Selected contours from the subenvelope map are superimposed.
B6 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

chian Plateau and in the area east of the Susquehanna River basin. The higher areas of the plateau 

of northern Pennsylvania as well as the Catskill Mountains are underlain by continental clastic 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 Devonian, Devonian, undivided, and Silurian and Devonian, undivided. The Upper Devonian of King and 

Beikman (1974) is not included. South of Pennsylvania, 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 remainder; 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 (1974). These rocks vary widely in their degree of resistance, but they do contain 

resistant sequences of great thickness, especially in the 

south. The Ocoee Supergroup, for example, which underlies the Great Smoky Mountains and the Black 

Mountains, including Mount Mitchell, contains 5200 m 

of metamorphic quartz-rich sandstone and siltstone. 

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 sequence (unit 1), the Devonian clastic sequence (unit 

2), and the Precambrian rocks (unit 4) form the 

highest areas. The bifurcation of the southern highland occurs where the main highland shifts southeastward 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 Pennsylvanian age holds up large plateaulike areas. In the 

northern part, the Devonian and Mississippian clastic 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 corresponds to a broad outcrop area of carbonate rocks 

(unit 4). At P, the principal tributaries of the Potomac 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 Virginia, the altitude of the subenvelope (fig. 3) appears 

to be unrelated to the bedrock. It crosses the nonresistant Cambrian and Ordovician sequence at an altitude 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 downstream from the belt. These adjustments, in effect, form 

local base levels at various altitudes that control the 

profiles upstream.

Figure 4^4, 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 Ordovician 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 Cambrian and Ordovician rocks. However, all have tributaries 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 between the streams within the belt are narrow everywhere, so that a continuous longitudinal profile is 

formed.

The profile along the strike of the Cambrian and 

Ordovician rocks (fig. 4J.) is divided into concaveupward segments, each of which is part of the drainage ~ basin of a large o stream that exits the Cambrian

and Ordovician belt. The high point is near Rural Retreat south of the continental divide, between headwaters of the Tennessee and New Rivers. This place 

is close to where the Cambrian and Ordovician belt
Continental Divide

100 200 300 400 MILES

100 200 300 400 KILOMETERS

\

w 

>

H

< 1-1 ;->

: — ' 2 ffi

i O ^

Q

y\ ^ 

IX ^— ^JQ

FIGURE 4.—Profile (A) and outline map (.B) of the Cambrian and Ordovician belt from Coosa River to Lake Champlain. 

titudes of selected high peaks.
B8 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 channel, 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 quantities, channel slopes are generally inversely proportional to a function of length. This explains the concave-form characteristics of most longitudinal valley 

profiles. As channel slope tends to be directly proportional 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. 4.4). 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 Chattanooga, 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 Potomac exits the belt at a place where the resistant 

rocks are unusually thin (Hack, 1965, p. 28). The 

Susquehanna and Schuylkill Rivers flow directly

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 Hudson 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 general topography of the belt can be explained without 

involving the hypothesis of differential uplift, except, 

of course, the obvious condition that the entire Appalachian 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 resistant 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 Ordovician belt can be judged by the spot altitudes of peaks 

shown on figure 4#. 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 resistant 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 deposits 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 Middle Cambrian carbonate rocks of the Shady and Tomstown Dolomites and their equivalents. The surface 

topography in the zone of these deposits is in many 

places pitted by modern sag ponds that commonly contain 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 Continental Shelf has, of course, been used by many to interpret 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 mentioned. 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 Cretaceous and Miocene time.

Judging by the quick response of the crust by uplift to the removal of the continental glacier of Pleistocene 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 Miocene, 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 geologic 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 sedimentary 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 difficult to deTABLE 1.—Thickness of sediments penetrated in COST No. 

B-2 (Scholle, 1977) and estimated deposition rate

Period or epoch

Duration 

(millions 

of years)

Thickness

of sediment

(m)

Rate 

(mm/yr)

Pliocene and 12

Pleistocene. 

Miocene —_————— 

Oligocene _——_—_ 

Eocene ———————— 

Paleocene —_———— 

Late Cretaceous —_ 

Early Cretaceous ___

Total _____ 136 

Average _____

130 0.011

14

12

16

11

35

36

860

150

220

60

980

2470

.061

.012

.014

.005

.028

.068

4870

.036
BIO SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 northeast, 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 northeast. 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 related to the tilt of the coastal zone. Eardley (1964) 

explained the tilt of the Coastal Plain and the Continental Shelf as part of a worldwide process of poleward rise of sea level related to a long-term decrease 

in period of the Earth's rotation, beginning in the

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, provided 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 comparatively low over extensive areas, and the map is intended to emphasize differences in relief within the 

Piedmont. Most of the areas of moderate relief on the 

outer Piedmont are directly related to resistant bedrock. For example, the Pine Mountain belt containing quartzite beds forms a moderate relief area at A. 

Cataclastic rocks of the Brevard zone occur at B. Resistant beds of the Kings Mountain belt are found at 

C. At area D, moderate relief is caused primarily by f el80 76L

40^

36C

32^

EXPLANATION

High relief

Moderate relief

Low relief

36^

32C

FIGUEE 5.—Generalized relief map of Piedmont and Blue Ridge provinces. The relief classes are based on the number of 

contours within square areas 10 km on 

a side, as indicated on the 1:250000- 

scale series of topographic maps by the 

U.S. Geological Survey. Less than 90-

400 KILOMETERS

120 m in low-relief areas and 120-240 m 

in moderate-relief areas. High-relief 

areas are generally more rugged and 

have relief of more than 240 m and as 

much as 1200 m. Letters indicate areas 

discussed in the text.
ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS Bll

sitic rocks of the slate belt. A large area of moderate 

relief (E) lies between the Coastal Plain and the Culpeper and Gettysburg Triassic and Jurassic basins. Its 

highest point exceeds 305 m above sea level near Westminster, Md. This area was referred to as the Wesfminster anticline by Campbell (1929, 1933), who believed 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 result 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 Westminster 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, cliamictite gneiss, quartzose schist, volcanic rock, phyllite, and marble.

The altitude of the highest parts of the Westminster 

uplift exceeds 305 m. Total relief within 100-km2 areas 

was measured on 1:24 000-scale maps of the following 

quadrangles: Downington, Pa. (162 m), and Mt. Airy 

(125 m), Winfield (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 relatively 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 Washington, D.C., whereas the slope of the Miocene in Washington 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 gravel is similar in size composition to the upper Tertiary gravel. Thus, the difference 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~* (mm/km)/yr. If 

the present slope tangent of the basement were projected 50 km inland from the Fall Zone to a point

near the center of the highest altitude near Westminster, 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 average rate of lowering at this point since middle Cretaceous 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 lowTer than average rates calculated 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 basement beneath the Coastal Plain in the Washington 

area, as indicated on the "Tectonic Map of the United 

States1 ' (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 compared with 0.02 in the Washington area. In the northern Piedmont, the increased relief is also related to the 

general depression of the Coastal Plain in the Salisbury embayment, which brings the edge of the Piedmont 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 relief 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 Precambrian terrane and the generally younger metamorphic 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 altitudes of the terrain on either side are strikingly different. Within areas 10 km wide, the relief averages 

less than 90 m 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 differences in lengths of the rivers on either side in their 

descent to the sea. White (1950) thought the escarpment 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 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 uplift 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

related to a constant but unspecified data base expressed as erosion in millimeters per year or as velocities of uplift in millimeters per year.

The estimates in table 3, on the other hand, are differences 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 drainage 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.—Estimates of erosion rates

Item 

HO. Process Rate 

(mm/yr) Sources of data

1 Modern erosion, North Atlantic drainage basins_ 4.8 X 10 2

2 Modern erosion, South Atlantic and Gulf basins__ 4.1 X 10~*

3 Modern erosion, South Fork of Shenandoah 4.3 X 10"2 

River watershed.

4 Modern erosion, small watershed, Coweta, N.C. __ 9 X 10~3

5 Modern erosion, small watershed, Maryland 4.6 x 10"3 

Piedmont.

6 Cenozoic erosion rate based on sediment volumes 2.7 X 10~2 

on Continental Shelf (Atlantic drainage).

7 Phanerozoic erosion rate for U.S. based on sedi- 1 X lO1"2 

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 10~2 

on depth of intrusion of Stone Mountain Granite, 

Georgia.

9 Uplift and erosion (some possibly tectonic) based 5.3 X 10~2 

on thickness of Paleozoic sequence in eastern 

Pennsylvania.

10 Removal of overburden from Middle Devonian 4.8 to 8.8 X 10~ 

rocks and conodont color in Appalachian basin— 

rate calculated since Permian.

Judson and Ritter, 1964. 

Judson and Ritter, 1964. 

Hack, 1965.

Berry, 1977.

Cleaves and others, 1970 1

Mathews, 1975.

Gilluly and others, 1970.

Whitney and others, 1976 1

This paper, data from Colton, 1970.

This paper, data from Epstein and 

others, 1977.

1 The rates reported here were calculated by Hack from data in references cited.

TABLE 3.—Estimates of tilt rates

Item 

no.

1

2

3

4

5

6

Tilted feature

Modern tilt rate in Potomac basin if balanced by

average erosion rate of 4 X 10~2 mm/y.

Tilt rate of basement rocks from the Cretaceous

to late Miocene, Washington, D.C.

Domal uplift of Canadian Shield since 7000 B.P.

(average tilt of southeast part of dome).

Downwarp to south of beach of high stand of

Champlain Sea since 10 000 B.P.

Upwarp to south of level line west of Lake

Champlain, Rouses Point to Ticonderoga after

18 years.

Downwarp of rerun level line 100 km long between

Atlanta and Columbus, Ga.

Rate 

[(m/km)/yr]

3.5 X 10-*

1.75 X 10-4

2.2 X 10'a

9.3 X 10-2

3 X 10-2

5.5 X 10-*

Source of estimate

This paper.

This paper.

Walcott, 1972, fig. 16.1

Denny, 1974, p. 42.1

Isachsen, 1975.1

Brown and Oliver, 1976, fig. 11.

tilt rates reported here were calculated by Hack from data in the references cited.
ROCK CONTROL AND TECTONISM—SHAPING THE APPALACHIAN HIGHLANDS B13

methods of analysis used in these two cases are rigorous. 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 inference, 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 magnitude 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 consistently indicate depth of burial. These data give results 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 inm/y (from table 2). As the erosion rate 

varies throughout the basin, the average rate is assumed 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 measured 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 postPleistocene 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

reached by comparing erosion estimates with absolute 

uplift rates of domal features that have been measured. One example is the Canadian Shield, which near 

its center at Hudson Bay is estimated by Walcott 

(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 landforms in postglacial time is useful. Glacial deposits of 

Wisconsinan age in the Central United States, as much 

as 55000 yr old (Flint, 1971, p. 559), have a constructional 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 erosion rates are converted to or are considered the inverse 

of tilt rates, then a large discrepancy remains, amount-
B14 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 rebound 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 removal of a load and compensation by uplift. Much of 

the data contained in repeated geodetic traverses, however, 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 in deep. No depression exists at this locality, and the topography is 

a normal erosional topography resembling its surroundings.

DEVELOPMENT OF THE APPALACHIAN 

HIGHLANDS IN LATE GEOLOGIC TIME

The present topography of the Appalachian Highlands is closely adjusted to the rocks. However, the 

available evidence shows that the highlands have undergone 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 Cretaceous 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 continuous, it was particularly rapid in early Cretaceous 

and Miocene to Pliocene time. Assuming these conditions, the present topography could have been derived 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 

clastic sediments deposited on the Coastal Plain and 

Continental Shelf are Early Cretaceous in age. The 

outlines of the hinterlands from wT hich 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 proportion of quartzite and finer gravel contains a significant amount, at least some of the material probably 

came from the Appalachians w^est of the Blue Ridge. 

On the other hand, the heavy minerals and the paucity 

of chert indicate that, in general, the greatest proportion of material probably came from the Piedmont, 

perhaps indicating that the relief on the Piedmont 

was high at the time.

48°

44°

88° 84° 76° 72° 68°

44°

800 KILOMETERS

EXPLANATION

-120 to -80 -80 to -40 -40 to 0 0 to +50 

Bouguer anomalies in milligals

Subenvelope contour at 1000 feet (305 m)

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 Appalachians (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 correspondence 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 B15

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 Continental Shelf, derived largely from crystalline terrane, 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 Oligocene 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 absolute rate of uplift relative to base level must have increased 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. West of the Blue Eiclge, 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 Eidge and Piedmont, some 

rocks, such as the Ocoee Supergroup, felsitic volcanic 

rocks, or the quartz kyanite rocks of the Kings Mountain 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 erosion in 

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.

At present in the Piedmont and Blue Ridge, saprolite 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, lacking, 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 suggestive 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 (oil the southeast side 

of the upwarp), and the escarpment reached its present 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 projected inland as far as the Blue Ridge, the top of the 

basement would reach an altitude of only about 1200 

m, almost equal to the crest of the present mountain 

ranse. ~ We have reason to believe, ' however, from estimated erosion and uplift rates, that the middle Cretaceous crest of the range would have been at about 4000 

m 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 

Washington 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 Mountains, which have the form of an uplift, are girdled 

by a wide belt of nonresistant rock. Much of the Ap-
B16 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

palachian Highlands region, however, can be interpreted as an upwarp of large proportions, but the entire highland cannot be explained that simply. For 

example, differential uplift, or warping, must be invoked 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 Highlands 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 erosion 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 indicate 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 tabs.] : 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 

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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 Bellevue Peak Quadrangle

CONTENTS

Page

Abstract __________________________________________ Cl

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

ILLUSTRATIONS

Page

FIGURE 1. Map showing outcrop areas of the Hanson Creek Formation and numbered fossil localities in the Bellevue Peak Quadrangle and elsewhere in central Nevada ________________________-__-__-----_-_--_ C3

2. Preliminary geologic map and sections of the Wood Cone Peak area, Bellevue Peak Quadrangle, Eureka

County, Nev _________________________________________________________________ —— __ 4

3. Lithostratigraphic section of the Hanson Creek Formation, measured on the west side of the Mountain

Boy Range, Nev _____________________________-____________^___-______------_---_-_-----_-_- 7

4. Trilobites from the Hanson Creek Formation _________---__-____________________________________ 13

5. Trilobites from the Kimmswick Limestone, in Missouri, and the Hanson Cre'ek Formation ________________ 15

6. 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 graptolite-bearing 

platy limestone facies of the Hanson Creek Formation, Monitor Range, Eureka County, Nev ______ 17

7. 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 _____________-___----_ 20

TABLES

Page

TABLE 1. Trilobites of the Hanson Creek Formation in the Bellevue Peak Quadrangle, Nevada ___ — ___ — ___— C9 

2. 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 _________ — _____ 19

in

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 lithostratigraphic members: (1) a basal dark-gray dolomite, (2) 

a middle silty thin- to thick-bedded, locally nodular, darkgray, 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., Lepidocyclus capax, Ceraurinus icarus, and 

a new species of Hypodicranotus favor a Maysvillian, preferably a Richmondian, age for the middle member. The presence of the conodonts Amorphognathns ordovicicus and Protopatiderodus 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 Aphelognathusl 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 Ordovician 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 outcroppings to be far more complex structurally than 

might be supposed from Hague's (1883) reconnaissance maps, has demonstrated a stratigraphic 

sequence that crosses the Ordovician and Silurian 

boundary within a single formation, and has revealed 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 lateral equivalents of which are widely distributed in 

Utah and Nevada and have been given other formation 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 unfortunately believed to represent a "Trenton" fauna 

on the basis of misconceptions dating back to the 

work of James Hall. "Orthis subquadrata" and "Orthis 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 Plectorthis plicatella (Hall) of

Cl
C2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

Maysvillian age. When Hague's faimal list is updated 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 

Orthis subquadrata 

Orthis (like O. plicatella)

Trinucleus concentricus 

Asaphns platycephalus 

Streptelasma

Rhynchonella

Ceraurus

Dalmanites

Illaenus

=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) 

=Cryptolithoides sp. 

=Isotelus; might be Anataphrus 

=Bighomia, Lobocorallium, or

one of several other genera 

=Lepidocychis, possibly

Rhynchotrema 

=Ceraurus or Ceraurinus 

=Calyptaulax 

=Illaemts(we have not identified

any Illaenid trilobite in our

collections from the Hanson

Creek)

Mapping in the Bellevue Peak Quadrangle has disclosed 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) considered 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:

Feet

Limestone, massive, dark-gray, fine-grained 1 __________ 180

Limestone, poorly stratified 2 ________________________ 140

Limestone, slabby and shaly, dark-bluish-gray; weathers 

very light gray; very fossiliferous 2 _______________ 140

Limestone, poorly bedded, noncrinoidal; small black 

chert nodules common 3 __________________________ 45

Limestone, dolomitic, dark-gray; flecked with crinoidal 

columnals 3 ______________________________________ 40

Total _____________________________ 545

1 Probably Combs Canyon Dolomite Member of this report.

- 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, although 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 laminated, thin-bedded graptolite-bearing limestone—on 

the basis of exposures on Martin Ridge, in the Monitor 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 district 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 Formation 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

116°00'

HANSON CREEK FORMATION 

SECTION USGS COLLECTIONS 

8320 CO-8330 CO.D2703 CO, 

D2704 CO.D2713 CO, D2714 CO

PANISH MOUNTAIN

BELLEVUE PEAK 

QUADRANGLE

Dobbin :X:Surnrhii

30 KILOMETERS 

I

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 juxtaposition. 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 southwest 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 intraclastic 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 limestones that normally are conformably above it. At 

some places in the Mountain Boy Range, the brecciated contact between the Eureka Quartzite and 

the lower dolomite member exhibits nearly isoclinal folding.

Above the basal dark dolomite, a sequence of thinand medium-bedded silty limestone and thin-bedded 

light-gray dolomite is found at all the more extensive 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

EXPLANATION

Stream and piedmont gravel

Quartz monzonite

Lone Mountain Dolomite-May include 

some Rabbit Hill Limestone

Combs Canyon Dolomite Member

Limestone and light-gray dolomite-- 

Fossiliferous

Lower dark dolomite member

fQUATERNARY

[• OLIGOCENE

I DEVONIAN!?) 

AND SILURIAN

Hanson Creek 

Formation

SILURIAN AND 

ORDOVICIAN

* ORDOVICIAN

Eureka Quartzite 

Pogonip Group

—————— CONTACT-Dotted where concealed

« ^« •— THRUST FAULT-Dashed where uncertain; dotted where concealed 

Teeth on upper plate. Arrows show direction of movement

D277J CO collection locality, U S. Geological Survey 

A

1000

CONTOUR INTERVAL 200 FT. 

1 FT=0.3048 METERS

2000 METERS

__I

FIGURE 2.—Preliminary geologic map and sections of the Wood Cone Peak area, Bellevue Peak Quadrangle, Eureka 

County, Nevada. Geology by T. B. Nolan, R. J. Ross, Jr., and L. A. Wilson. Base from U.S. Geological Survey 1:62,500 

Bellevue Peak (1956).
THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C5

the limestone. Locally the limestone is abundantly i 

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 intraclastic conglomeratic zones. Fragments as large as 

5-cm across were found in an isolated fossiliferous 

limestone outcrop 1 km southwest of Spanish Mountain at the south end of the Mountain Boy Range 

(fig.l).

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 lightgray 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 Mountain 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 sufficiently well preserved to be recognized as pentamerids. In other places, halysitid corals have been 

preserved. Intraclastic conglomeratic beds and horizons 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 uppermost light-gray dense dolomite of the middle member. 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.

(unsurveyed). Here the member overlies the upper 

very light gray dolomite of the Hanson Creek Formation 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 (Manitoba; Wyoming) (fig. 3; USGS collns. D2707-CO, 

D2708-CO, D2741-CO).

The second area is on the west side of the Mountain 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 m, N. 4,369,500 

m) (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 conodonts. The dark-gray Combs Canyon Dolomite Member (Silurian) is exposed about 150 m southeast of 

the main section (UTM Grid, zone 11: E. 581,880 m, 

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). 

In the lower part of the middle limestone member, 

corals, bryozoans, brachiopods, and a few trilobites 

closely resemble genera found in the Maquoketa 

Shale of Iowa, the Stony Mountain Formation of 

Wyoming (Ross, 1957, p. 446-488, pis. 37-41), and
CORALS BRYOZOANS BRACHIOPODS TRILO

( ———— * ————— } 1 —— —— "" ———— ^ 1 ——————————————— ——————— — i r • 

§ 

l 

^ .C «: 

i If!! i i 

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Combs Canyon Dolomite Member- -Dark-gray 

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Platy limestone intertongue

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Limestone, nodular, fine-grained, 

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dark-gray chert nodules / 

Limestone and limy dolomite, cherty, 

fme-qrained and unevenly laminated

Limestone, nodular, silty, light-gray 

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bedding upward

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dark-gray, massive

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|
THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C7

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C8 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

the Ely Springs Dolomite in the New York Butte 

Quadrangle, Inyo Mountains, Calif. (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 cannot 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 correlation 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 evidence that each of these kinds of fossils provides 

for determining the age of the Hanson Creek Formation 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 concluded that they are of Late Ordovician age. Nelson 

(1963) also found this assemblage in the Hudson 

Bay Lowland; the questionably identified Nyctopora(t) and Saffordophyllum(l) 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 Lobocorallium trilobatum major Nelson; according to 

Nelson, that group is of Richmondian or Gamachian 

age.

Oliver (written commun., Jan. 4, 1978) commented on the distinctive Lobocorallium found in 

USGS colln. D2713-CO:

1. If the Lobocorallium 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 specimens 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 Bighornia, concluding that it was of Richmondian 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), Bighornia 

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 identified by O. L. Karklins (written commun., March 3, 

1976), who stated that preservation was too poor to 

permit determination of species and that no conclusion as to age within the Ordovician was therefore 

possible. On the other hand Sceptropora facula 

Ulrich is readily identified (Ross, 1957, pi. 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 brachiopods listed in figure 3 are 

fragmentary. Several genera are of little use in differentiating rocks of Middle and Late Ordovician 

age. Pionorthis is present in the Stony Mountain 

beds of Manitoba and Wyoming (Ross, 1957; Macomber, 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 Anoptambonites 

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) (Williams, 1962, p. 170-172).

EVIDENCE OF THE TRILOBITES

The ranges of the 18 Hanson Creek trilobite genera are shown in table 1. (See also figs. 4, 5.) The
THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C9

TABLE 1.—Trilobites of the Hanson Creek Formation in the Bellevue 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)

Selected previously reported 

occurrences

Age range 

previously assigned Reference and remarks

Hypodicranotus n. sp . 

D2645a, c, d, e-CO 

D2707-CO 

D2708-CO 

D2771-CO

Robergia sp ____ 

D2645d-C0

Isotelus sp _____ 

D2645d-C0 

D2707-CO 

D2708-CO

Anataphrus sp __ 

D2645a-e-CO 

D2707-CO 

D2708-CO 

D2741-CO 

D2771-CO

Styginid, possibly 

Stygina. 

D2741-CO

Proetid _______.

D2707-CO 

Otarion sp ________

D2645a, c-e-CO

D2707-CO

D2741-CO

D2771-CO

Cryptolithoides sp _. 

D2645b-e-CO 

D2707-CO 

D2741-CO

Calymenid (possibly 

Gravicalymene or 

Calymene). 

D2645d, e-CO 

D2707-CO 

D2708-CO 

Ceraurinus cf. C. 

icarus (Billings). 

D2645e-C0

Ceraurus sp. 1 ____ 

D2645d, e-CO 

D2741-CO

Ceraurus, sp. 2 (radial 

ornament)-n. sp.? 

D2645a-C0 

D2741-CO 

Ceraurus sp. 3 (fine 

ornament).

D2645e-C0

Acanthoparyphinid,

possibly Holia

or Youngia.

D2645e-C0

Genus reported in Trenton Limestone (New York) ; Prosser 

Limestone (Minnesota) ; Viola 

Limestone (Oklahoma) ; uppermost Copenhagen Formation 

(Nevada) ; Silliman's Mount 

(Baffiin Island) ; Whittaker Formation (Dist. of MacKenzie) ; 

Albany and Stinchar beds 

(Ayrshire) ; Tipliteki horizon 

(central Urals).

Road River Formation (Yukon) ; 

Ogygiocaris Shale (Sweden).

Chazy Group (New York) ; Lexington Limestone (Kentucky); 

Ellis Bay Formation (Anticosti 

Island). Numerous other occurrences.

Sillman's Mount (Baffin Island) ; 

Cape Calhoun Formation 

(Greenland) ; Maquokeita Shale 

(Iowa); uppermost Copenhagen 

Formation (Nevada).

Chair of Kildare Limestone

(Eire) ; Upper Ordovician (Poland) ; Boda Limestone (Sweden) ; Killey Bridge (North 

Ireland).

Numerous localities; widespread 

in world.

Many species in United Kingdom, 

Estonia, Sweden, Poland, Quebec, 

MacKenzie, New York, Oklahoma, Iowa, New Jersey, Virginia, Anticosti Island.

Viola Limestone (Oklahoma) ; 

Copenhagen Formation (Nevada) ; Road River Formation 

(Yukon).

Numerous localities in North 

America and Europe.

Richmond Gp. (Indiana and Ohio) 

English Head Formation and 

Vaureal Formation (Anticosti 

Island).

Whittaker Fm. (Dist. 

MacKenzie).

Not known previously (?) _____.

Middle to late 

Caradocian.

Llandelian to

Ashgillian. 

Chazyan to

Richmondian.

Llandeilian to

Ashgillianv

Late Mohawkian 

to Richmondian.

Ashgillian

Middle Ordovician 

to Mississippian.

Llanvirnian to 

Middle Devonian 

(Chazyan through 

Richmondian).

Mid-Caradocian to 

Ashgillian.

Llanvirnian to 

Middle Devonian.

Richmondian 

Maysvillian?

"Black RiverTrenton".

Arenigian to 

Ashgillian.

Whittington, 1952; Tripp, 1965, 

1967; Ross and Shaw, 1972; 

Ancigin, 1973; Ludvigsen, 1975. 

Middle body of hypostome differs markedly from those of H. 

striatulus (Walcott) and H. 

missouriensis (Foerste). Roy, 

1941.

Whittington, 1950; Lenz and

Churkin, 1966. 

Twenhofel, 1927; Ross, 1967;

Shaw, 1968, 1974.

Whittington, 1954; Troedsson, 

1928; Ross and Shaw, 1972.

Dean, 1974, p. 67; Kielan, 1960, p. 

82, pi. 8, figs. 1, 2; Warburg, 

1925, p. 96, pi. 3, fig. 1; Whittington, 1950, p. 547-549.

Owens, 1973.

Lamont, 1935; Opik, 1937; Twenhofel, 1927 (listed as Cyphasphis) ; Begg, 1939; Weller, 1903; 

Tripp, 1954, 1967; Ingham, 

1970; Warburg, 1925; Whittington, 1966; Kielan, 1960; Schuchert and Cooper, 1930; Shaw, 

1968, 1974; Raymond, 1925; Hu, 

1975; Ludvigsen, 1975, 1976a.

Whittington, 1941; Ross and Shaw, 

1972; Lenz and Churkin, 1966.

Shirley, 1936; Twenhofel, 1927; 

Dean, 1962; Ross, 1967.

Ludvigsen, 1977; Twenhofel, 1927.

Ludvigsen, 1975, pi. 5, figs. 8, 9.

For generic range see Lane, 1971, 

text fig. 2.

Not known previously (?)

Kimmswick Ls. (Missouri and 

Illinois); Lincolnshire and Edinburgh Formations (Virginia) ; 

Esbatattaotine and Whittaker 

Formations (Dist. of MacKenzie) .

Chazyan to Edenian. 

Llandeilian to 

Llandoverian.

Schmidt, 1881; Warburg, 1925; 

Bradley, 1930; Whittington and 

Evitt, 1954; Ludvigsen, 1975, 

1976a; Lane, 1971.
CIO SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

TABLE 1.—Trilobites of 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 

D2771-CO

Encrinurus sp. 

(may be 2 sps.) 

D2645b, d-CO 

D2707-CO 

D2771-CO

Cybeloides sp _____ 

D2708-CO

Calyptaulax sp __. 

D2645c, e-CO 

D2707-CO 

D2708-CO 

D2713-CO

Lichid, possibly 

Hemiarges.

D2645a-C0

Odontopleurid, possibly 

Ceratocephala. 

D2645e-C0 

D2708-CO

Uppermost Copenhagen Fm.

(Nevada); Ellis Bay Formation 

(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); 

Stile End beds (United Kingdom) ; Ashgill (Cumbria).

Maquoketa Shale (Iowa); Edinburgh Formation and Effna 

Limestone (Virginia); Lincolnshire Limestone (Tennessee); 

Rhiwlas Limestone (Wales).

Whitehead Formation, Perce 

(Quebec); Tyrone Limestone 

(Kentucky); Galena Dolomite 

and Platteyille Limestone (of 

Mississippi Valley); Middle 

Ordovician of Appalachians 

(Virginia); Copenhagen Formation (Nevada).

Chazy an to 

Gamachian.

(Llandeilian to 

Ashgillian).

Caradocian,? 

Ashgillian to 

Silurian.

Chazyan to 

Richmondian.

Chazyan to 

Richmondian.

Family range 

Ordovician to 

Devonian.

Middle Ordovician 

Late Devonian.

Ross and Shaw, 1972; Shaw, 1968; 

Lane, 1971; Warburg, 1925; 

Ingham, 1974; Twenhofel, 1927. 

Also in middle limestone member of Hanson Creek type 

section.

Twenhofel, 1927; Dean, 1963; 

Ingham, 1974.

Slocum, 1913; Cooper 1953; Shaw, 

1968; Whittington, 1964.

Delo, 1940; Shaw, 1968; Cooper, 

1953; Schuchert and Cooper, 

1930; Ross and Shaw, 1972.

silicified trilobites are almost all immature and therefore 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. missouriensis (Foerste) (fig. 5a) 

differs markedly from the elongate middle body of

the Hanson Creek hypostome (fig. 5/i-fc). The Trenton species H. striatulus (Walcott), as shown by 

Whittington (1952), bears the same middle body as 

H. missouriensis. 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 Richmondian or an older age. If future ciollecting should 

prove that the styginid is really Ashgillian Stygina, 

rather than an older form like Raymondaspis, or 

that the acanthoparyphinid (fig. 4, aa-ac) Youngia, 

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 Bighornia and particularly 

of Lobocorallium major Nelson among the corals, of 

Sceptropora 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 Cll

Ceraurinus 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 formicacid treatment of limestone and dolomite samples, 

respectively. Insoluble residues were washed 

through a 20-mesh sieve nested in a 200-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-/) 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 Richmondian; see Sweet and Bergstrom, 1976, fig. 4). Another indicator of slightly younger Maysvillian 

through Richmondian age, Protopanderodus insculptus (Branson and Mehl) s.f., occurs 20.7 m above 

the base of the section. Other elements of less restrictive but definitely Late Ordovician age that cooccur with or above A. ordovicicus in the section are 

Belodina sp. D of Shatzer (1976) (fig. 6,r, s), 

Oulodus ulrichi (Stone and Furnish) (fig. 6,m-p), 

and Plegagnathus nelsoni Ethington and Furnish s.f. 

(fig. 6,q). These elements alomg 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-CO) also contains elements of Amorphognathus ordovicicus and therefore lies within the 

same age range as collections 8320-CO to 8328-CO 

lof 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 Cone 

Peak.

The uppermost collection (table 2; USGS colln. 

8329-CO) within the measured part of the section

contains conodonts of Late Ordovician (zygognathiform elements) and Silurian aspect (plectospathodiform 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 spathognathodiform and plectospathodiform elements that 

are probably part of the same Ozarkodina apparatus 

(fig. 7, j, k) and Exochognathus keislognathoides 

Pollock, Rexroad, and Nicoll (fig. 7,1) and distomodids (fig. 7,h, i) that probably belong with other 

elements assigned to Aphelognathust n. sp. (fig. 

7a-g). Significantly, A.? n. sp. elements are abundant 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 Mountains, Calif. (A. G. Harris and J. F. McAllister, 

USGS collns., unpub. data) - 1

Conodont data from the measured section demonstrate 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 immediately below the possible fault are of latest Ordovician 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 

oonodonts from central Nevada provide new information 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 section 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 Vaureal 

and overlying Ellis Bay Formations of Anticosti Island, Quebec (Nowlan 

and Barnes, 1979; McCracken and Barnes, 1979). Representatives of what 

we have assigned to Aphelognathust n. sp., distomodids, and Exochognathus 

keislognathoides are among the most abundant elements in the Vaureal 

and Ellis Bay conodont faunas. McCracken and Barnes (1979) placed 

these elements in a new genus and species, Gamachignathus ensifer. This 

species is restricted to the Upper Ordovician Vaureal 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 Ozarkodina 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 immediately 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.
C12 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979
THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C13

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 Belodina sp. A of Sweet, Ethington, and 

Barnes (1971) was recovered from 2 m above the 

base of the Hanson Creek (fig. 6,#). 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.

In the course of our studies, conodonts were also 

collected from, the lower part of the graptolitebearing 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 southwest 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 diagnostic 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 secFIGURE 4.—Trilobites from the Hanson Creek Formation.

[a-j, Q-S, z, aa-ag, from USGS colln. D2645e-CO; fc, t-y, ah-aj, from USGS coDn. D2645d-CO; l-p, ak, from USGS colln. D2645a-CO; ax, from

USGS colln. D2707-CO]

a—k. 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, i, j, cranidium, 

dorsal, antero-dorsal, lateral, and anterior 

views, x 3, USNM 252217. k, hypostome, ventral view, X 8, USNM 252218.

l-p. Ceraurus sp. 2. I, m, n, cranidium, dorsal, lateral, 

and anterior views, x5, USNM 252219; note 

radial ornament around each tubercle, o, partial thorax and pygidium, dorsal view, X 10, 

USNM 252220; lacks radial ornament of 

associated crandium. p, hypostome, fragmentary, ventral view, X 10, USNM 252221; lacks 

radial ornament.

q, r. Ceraurus ? sp. 3. Crandium and pygidium, dorsal 

views, X 7, USNM 252222, 252223.

s, z. Ceraurinus icarus (Billings), s, hypostome, ventral view, x 3, USNM 252224. z, pygidium, 

dorsal view, X 3, USNM 252225.

t-j/. Calymenid, possibly Gravicalymene. t, hypostome, ventral view, X 7, USNM 252226. u, \ 

pygidium, dorsal view, X 6, USNM 252227. v, 

cranidium, fragmentary, dorsal view, X 5, 

USNM 252228. w, x, y, pygidium, posterior, 

lateral, and dorsal views, X 9. USNM 252229. 

aa—ac. Acanthoparyphinid, possibly Youngia or Holia. 

«°» pygidium plus one thoracic segment, 

dorsal view, X 7, USNM 252230. ab, ac, cranidium, dorsal and lateral views, X 4, USNM 

252231.

ad-ag. Odontopleurid, possibly Ceratocephala. ad, ae, 

partial thorax, probably assignable to this 

genus, dorsal and lateral views, X 7, USNM 

252232. af, ag, cranidium, dorsal and anterior 

views, X 7, USNM 252233.

ah-aj, ax. Encrinurus sp. ah, pygidium, fragmentary, dorsal view, X 5, of a small part of the pleura 

and axis, USNM 252234. ai, aj, cranidium, 

fragmentary, dorsal and anterior views, X 9, 

USNM 252235. ax, pygidium, dorsal view, 

X 8, USNM 252236.

ak. Lichid, possibly Hemiarges. Cranidium, dorsal 

view, X 10, USNM 252237.

al-au. Anataphrus sp. and Isotelus sp. Because complete 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. an, ao, ar, as, at, 

au, pygidia, dorsal views, X 10, USNM 2'52240, 

252241, 252244, 252245, 252246, 252247 respectively, ap, pygidium, dorsal view, X 6, 

USNM 252242. aq, av, pygidia, dorsal views, 

X 5, USNM 252243, 252248.

aw. Calyptaulax sp. Cranidium, immature, dorsal 

view, X 7, USNM 252249.

ay. Cybeloides sp. Pygidium, fragmentary, dorsal 

view, X 8, USNM 252250.
C14 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979
THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C15

tion is probably continuous. Ross and Berry (1963) 

concluded that the Hanson Creek graptolites were 

chiefly late Caradocian forms but had some Ashgillian components. This conclusion would be in reasonable 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 limestone 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.

Any apparent uncertainty or discrepancy in dating 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 paleontologically, even in strata deposited at the same 

time.

REFERENCES 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. I., 1973, Trilobita, in Ordovician Stratigraphy and 

fauna of the central Urals: Ministry of Geology SSSR, 

Ural. Terr. Geol. Uprav., Nedra, Moskva, 166 p., 30 pis.

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 Canadian Society of Petroleum Geology Special Volume, p. 

221-240.

Barnes, C. R., and Fahraeus, L. E., 1975, Provinces, communities, and the proposed nektobenthic habit of Ordovician conodontophorids: Lethaia, v. 8, no. 2, p. 133-149.

[References continued on p. C18]

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 missouriensis (Foerste). All X 3. 

a, hypostome, ventral view; note almost circular 

middle body divided almost equally into three 

parts; USNM 252181. 6, c, d, cranidium, lateral, 

anterior, and dorsal views, USNM 252182.

e—p. Hypodicranotus n. sp. e, cranidium, dorsal view, 

X 20, USNM 252183. /, cranidium, smallest 

specimen recovered, X 20, USNM 252184. g, m, 

n, cranidial, dorsal views, respectively, X 10, 

X 20, X 10, 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, p, free 

cheeks, left, dorsal views, respectively X 6, 

X 10, X 10, USNM 252188, 252191, 252192.

q,r. Robergia sp. Fragmentary cranidium and free 

cheek, dorsal views, X 10, USNM 252193, 

252194.

s-u. Isotelus sp. Immature cranidia, deformed tectonically, all dorsal views, X 10, USNM 252195, 

252196, 252197, respectively.

v-ad. Anataphrus sp. v, w, cranidium, dorsal and anterior views, X 5, USNM 252198, x, y, ad, 

cranidia, dorsal views, respectively X 7, X 10, 

X 10, USNM 252199, 252200, 252204. z, 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. 

ai, aj, ak, cranidium, lateral, dorsal, and anterior views, X 6, USNM 252207. al, pygidium, 

dorsal view, X 20, USNM 252208. am, associated hypostome, ventral view, X 20, USNM 

252209.

an-ar. Ceraurus sp. 1. an, ao. cranidium, fragmentary, 

dorsal and anterior views, X 6. USNM 252210. 

ap, aq, ar, cranidium, fragmentary, anterior, 

dorsal, and lateral views, X 5, USNM 252211.

as-aw. Cryptolithoides sp. as, cranidium, dorsal view, 

X 7, of immature specimen; note that eye tubercles 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, fragmentary, ventral view, X 7, USNM 252214.
C16 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979
THE UPPER OEDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C17

FIGURE 6.—Late Ordovician conodonts from the lower dark 

dolomite member (a-;) and the middle limestone and dolomite member (k-ab) of tha 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 specimens illustrated on figures 6 and 7, except 6g, are reposited in the U.S. National Museum of Natural History, 

Washington, D.C. (USNM). Susann Braden assisted in the 

SEM photography.

a-/. Amorphognathus ordovicicus Branson and Mehl, from 

USGS colln. 8320-CO. a-e, x 100; /, 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) ; Maysvillian through Richmondian (Ashgillian).

a, b. Outer lateral views of two holodontiform 

elements, USNM 248444 and USNM 248445.

c. Lateral view of trichonodelliform element, 

USNM 248446.

d. Inner lateral view of eoligonodiniform element, USNM 248448.

e. Upper view of ambalodiform element, USNM 

248447.

/. Upper view of amorphognathiform element,

USNM 248449.

g. Belodina sp. A of Sweet, Ethington, and Barnes 

(1971) (=a new species of Belodina described by 

Craig (1968) from the lower half of the Femvale 

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, Spanish 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 Bergstrom, 

1976; Shatzer, 1976) and the Canadian Arctic 

(Barnes, 1974). It appears to be a component of 

the "Arctic-Red River" fauna of earlier workers. 

h, i. Panderodus gracilis (Branson and Mehl) element, 

from USGS colln. 8321-CO. 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. 

8321-CO. Outer lateral view, USNM 248451, 

XlOO.

k, I. Zygognathiform element, posterior and oblique lateral 

views of two specimens, USNM 248456 and USNM 

248460, x50; USGS colln. 8323-CO. Several different zygognathiform elements occur in our collections without other associated ramiform elements. Accordingly these were not referred to a 

discrete or multielement taxon. 

m-p. Oulodus ulrichi (Stone and Furnish), X 50.

m. Posterior view of oulodiform element, USNM

248452; USGS colln. 8321-CO. 

n. Inner lateral view of prioniodiniform element,

USNM 248454; USGS colln. 8326-CO. 

o. Oblique posterior view of trichonodelliform element, USNM 248455; USGS colln. 8323-CO. 

p. Lateral view of cordylodiform element, USNM

248453; USGS colln. 8323-CO.

q. Plegagnathus nelsoni Ethington and Furnish s.f. Inner lateral view, USNM 248459, X 50; USGS 

colln. 8322-CO. P. nelsoni, P. dartoni Ethington 

and Furnish s.f., and Cordylodus robustus Ethington and Furnish s.f. form one apparatus that is a 

diagnostic component of post-early Edenian Ordovician "Arctic-Red River" conodont faunas. P. dartoni and C. robiistus elements have not been idenfied in our collections.

r, s. Belodina n. sp. D of Shatzer (1976), X 100; USGS 

colln. 8325-CO. Outer and inner lateral views of 

two different belodiniform elements, USNM 248457 

and USNM 248458. This apparatus includes Belodina profunda (Branson and Mehl) of most previously described Late Ordovician conodont faunas. 

This species is a Maysvillian-Richmondian component of the "Arctic-Red River" fauna. 

t. Coelocerodontus trigonius Ethington s.f. Posterior 

view, USNM 248469, X 100, USGS colln. 8323-CO. 

u. Oistodus venustus Stauffer s.f. Outer lateral view,

USNM 248464, X 100; USGS colln. 8322-CO. 

v. Drepanoistodus suberectus (Branson and Mehl). 

Inner lateral view of oistodiform element, USNM 

248463, X 100; USGS colln. 8325-CO. Elements of 

this species comprise nearly 10 percent of the 

conodonts in Ordovician collections from the Hanson Creek Formation in the Mountain Boy Range 

and Wood Cone Peak area.

w, x. Pseudoneotodus mitratus ( Moskalenko). (=0istodtts(?) sp. of Branson and Mehl (1933, pi. 9, fig. 

3). Upper views of two specimens, USNM 248466 

and USNM 248465, X 50; USGS colln. 8322-CO. 

This is a sporadic element in Middle and Late 

Ordovician conodont faunas in the western United 

States.

y. Pseudoneotodus beckmanni (Bischoff and Sannemann) s.f. Upper view, USNM 248467, X 100; 

USGS colln. 8323-CO. P. beckmanni ranges from 

the Late Ordovician through Early Devonian and 

appears to be the only morphotype in its apparatus.

z. Belodina compressa (Branson and Mehl). Inner 

lateral view of belodiniform element, USNM

248461. X 100; USGS colln. 8322-CO. 

aa. Prioniodiform element. Oblique anterior view,

USNM 248470, x 50; USGS colln. 8327-CO. 

ab. Zygognathiform element, posterior view, USNM

248462. x 50; USGS colln. 8327-CO. This distinctive element with a greatly laterally compressed 

cuisp bearing a posterior costa and short but high 

lateral processes appears to be rather common in
CIS SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

FIGURE 6.—Continued.

Late Ordovician conodont faunas of the western 

United States.

ac. 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-minute Quadrangle. The extremely elongate posterior 

process is typical of true North American P. 

insculptus (middle Maysvillian-Richmondian).

This is the most common element in the meager 

conodont fauna recovered from three 5-kg samples 

from the lower part of the graptolite-bearing 

platy limestone facies on Martin Ridge. In contrast, 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 conodont biofacies.

References Cited—Continued

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Bradley, J. H., Jr., 1930, Fauna of the Kiimmswick Limestone 

of Missouri and Illinois: Chicago University, Walker 

Museum of Paleontology Contributions, v. 2, no. 6, p. 

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Branson, E. B., and Mehl, M. G., 1933, Conodonts from the 

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Cooper, B. N., 1953, Trilobites from the lower Champlainian 

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Oraig, W. W., 1968, The stratigraphy and conodont paleontology of Ordovician and Silurian strata, Batesville District, Independence and Izard Counties, Arkansas: Austin, 

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Dean, W. T., 1962, The trilobites of the Caradoc Series in 

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British Museum (Natural History) Bulletin, v. 7, no. 3, 

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London, British Museum (Natural History) Bulletin, v. 

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————1974, The trilobites of the Chair of Kildare limestone

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Delo, D. M., 1940, Phacopid trilobites of North America: Geological Society of America Special Paper 29, 135 p., 13 

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Duncan, Helen, 1956, Ordovician and Silurian coral faunas 

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Hall, James, 1847, Paleontology of New York, containing 

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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, pis. 1-9. [Published as part of v. 124.]

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pis. 10-18. [Published as part of v. 128.]

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Poland and some related forms from Bohemia and Scandinavia: Palaeontologia Polonica, no. 11, i-vi, 198 p., 36 

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Kirk, Edwin, 1933, The Eureka quartzite of the Great Basin 

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Lamont, A., 1935, The Drummuck group, Girvan; a stratigraphical revision, with descriptions of new fossils from 

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Le Fevre, Jean, Barnes, C. R., and Tixier, Michel, 1976, 

Paleoecology of Late Ordovician and Early Silurian 

conodontophorids, Hudson Bay Basin, in Barnes, C. RM 

ed., Conodont paleoecology: Canada Geological Association Special Paper 15, p. 69-89.

Lenz, A. C., and Churkin, Michael, Jr., 1966, Upper Ordovician trilobites from northern Yukon: Palaeontology, v. 

9, pt. 1, p. 39-47, pis. 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 pis.

———1976a, Middle Ordovician trilobite biofacies, southern 

MacKenzie Mountains: Canada Geological Association 

Annual Meeting, P. S. Warren Biostratigraphy Symposium, Program with Abstracts, v. 1, p. 39.

————1976b, New cheirurinid trilobites from the lower Whittaker Formation (Ordovician), southern MacKenzie 

Mountains: Canadian Journal of Earth Sciences, v. 13, no. 

7, p. 947-959, 2 pis. 

-1977, The Ordovician trilobite Ceraurinus Barton in

North America: Journal of Paleontology, v. 51, no. 5, p. 

959-9721, pis. 1, 2.

McCracken, A. D., and Barnes, C. R., 1979, Conodonts and 

the Ordovician-Silurian boundary from the Ellis Bay Formation, Anticosti Island: Canada Geological Survey Bulletin [in press].

[References continued on p. C21J
THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA C19

TABLE 2.—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.

[Collections 8320-8329-CO and 9785-SD from measured section 200 m east of center sec. 31, T. 19 N., R. 53 E., described in text. Collection 8292- 

CO from dark-gray thick-bedded limestone about 200 m west of measured section (UTM Grid, zone 11, E. 581450 m, N. 4369600 m); mapping 

indicates this outcrop may overlie the middle limestone of the measured section. Collections 8332-8334-CO from Wood Cone Peak area. Leaders, _., 

not present]

oe

«o « 

Distance (meters) above base of section ______"=> -5

Collection No. (USGS) ._ __

O 

.......... J

CO 

00

o

O

CQ

CO

oo

o0

CO

oo

o 

o

CO

oo

o

V(N

CO 

00

o 

o

sCO

oo

o 

o

CO 

00

o 

o

00 

05

CO

oo

o0

•"JI 

CO

oo

o0

£CO 

oo

o<?

oo 

o

o0

(N

00

o 

o

CO 

CO

oo

o

VCO 

CO

oo

o0

300

Sample weight ________________kilograms.. 4.9 4.5 3.6 4.3 4.6 5 5.2 3.8 3.5 4.2 4.3 2 3.6 4.2 3.8

Amorphognathus ordovicicus Branson 

and Mehl:

Ambalodiform element ——————— 7 2 —_ __ __ —— __ _— __ __ ___ 2 ___ _— ——

Amorphognathiform element ____ 43 13 _ __ __ __ __ ___ __ __ __ 7 ___ __ __

Eoligonodiniform element _-____ 5 1 __ __ __ __ __ -_ __ __ __ __ __ __ __

Holodontiform element _ — _———— 6 3 __ __ ___ __ __ — __ __ __ __ __ __ -_

Tetraprioniodiform element ___ 1 __ __ __ __ __ __ __ __ __ __ __ __ __ __

Trichonodelliform element ______ 8 2 __ _ __ __ __ __ __ __ __ __ _- _- __

Aphelognathus"! n. sp.:

Eoligonodiniform element _______ _ ___ _ ___ __ __ __ __ __ __ 8 _ ___ __ __

Prioniodiform element _______ __ __ ___ ___ __ __ __ __ _ __ 4 __ __ __ ___

Spathognathodiform element ___- __ __ __ __ __ __ __ __ __ __ 6 __ __ __ ___

Trichonodelliform element _____ __ __ __ __ __ __ __ ___ __ __ 4 ___ __ __ ___

Zygognathiform element _______ __ ___ __ ___ __ __ __ __ __ __ 3 _ __ __ __

Belodina compressa (Branson and 

Mehl) :

Belodiniform element - -

/? /7t QTW711 <?/* ( frl (*Y\ 1 Cii'iAT1 ^ C "F

B. n. sp. D elements of Shatzer (1976) _

Drepanoistodus suberectus (Branson and

Exochognathus keislognathoides Pollock,

Pseudoneotodus beckmanni (Bischoff

Pseudoneotodus mitratus

i *1W r\ clr Q 1 £*Y\ lr rv ^

Oulodus ulrichi (Stone and Furnish) :

Panderodus gracilis (Branson and

Plegagnathus nelsoni Ethington and

Protopanderodus insculptus (Branson

51 T1/1 IWoVll ^ G "F

Trichonodelliform elements with very

1 & t'^f'Sl 1 1 "\7 fWIYITIVAQ QA/1 f*HQTI

Zygognathiform elements with very

1 2

26 12

23 34 32

2

3

138

1

5 4

1

3

1 1

169 405 252

211

1

4 6

1

13

?,

27

3

3

8

6

7

5

6

1

3

288

4

«JQ

11

30

1

1

14

6

3

197

1

3

14 1

2

22 2 _

3 1

631

5 2

461

591

121

150 16 9

1

4

1

7 __

_ _ _ 3 1

1

1

3 2 26 22 4

6

1

2 1

1

1 _ _ _ 31

2 1

1

11 10 12 54 122 26

9

1 1 _ __

16

1

11

1

2

2 _ 2

1 Dark-gray platy limestone. UTM Grid,, zone 11, E. 574750 m, N. 4359550 m (=D2707-CO).

2 Dark-gray flaggy skeletal limestone. UTM Grid zo.ne 11, E. 574800 m, N. 4 359 125 m (=D2708-CO).

3 Dark gray fine-grained limestone. UTM Grid, zone 11, E. 574 630 m, N. 4 359 840 m.

* Cordylodiform and eoligonodiniform elements both appear to occupy the Sc position (apparatus nomenclature of Sweet and Schonlaub, 1975) 

this species of Oulodus. These collections contain no other apparatus in whichan eoligonodiniform or cordylodiform element can be accommodated.
C20 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 /).

a~g. Aphelognathus^ n. sp. The generic assignment is 

queried because Sweet, Thompson, and Satterfield 

(1975) diagnosed the apparatus of Aphelognathus 

as containing a dichognathiform or prioniodiniform 

element in the Oi position. In this species, a prioniodiform element occupies the Oi position. In addition, 

the denticle pattern and profile designated by Sweet, 

Thompson and Satterfield (1975, p. 29) as characteristic 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 elements aside from distomodids, keislognathodids, 

Oulodus sp., and panderodids, Aphelognathus is suph, i.

k.

FIGURE 7.—Continued.

posedly a Middle and Late Ordovician genus. Pollock, Rexroad, and Nicoll (1970), however, designated a p element from Lower Silurian beds in 

northern Michigan and Ontario, Aphelognathus 

siluricus. Moreover, the same form occurs in the B 

and lower C divisions of the Llandovery in the 

Welsh Borderland (Aldridge, 1972). The p element 

of our species closely resembles A. siluricus 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 ancestral to A. siluricus.

a, b. Spathognathodiform (p) 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 prioniodiform 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 available to us from the Ely Springs Dolomite 

of the Funeral Mountains (fig. /), we 

would have assumed this element to be 

dichognathiform.

/. 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, SW^i sec. 33, T. 26 

N., R. 4 E., Ryan 15-minute quadrangle, 

Inyo County, Calif.). Note the denticulate 

anterior process. 

g. Trichonodelliform element, posterior view,

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 characteristic 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 potst-Ordovician 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.
THE UPPER ORDOVICIAN AND SILURIAN HANSON CREEK FORMATION, NEVADA

References Cited—Continued

C21

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, pis. 

77-80.

Merriam, C. W., 1940, Devonian Stratigraphy and Paleontology of the Roberts Mountains Region, Nevada: Geological Society of America Special Paper 25, 114 p.

————1963a, Paleozoic rocks of Antelope Valley, Eureka and 

Nye Counties, Nevada: U.S. Geological Survey Professional Paper 423, 67 p. 

-1963b, Geology of the Cerro Gordo mining district,

Inyo County, California: U.S. Geological Survey Professional 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 Sciences 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. Geological Survey Professional Paper 800-B, p. B21-B24.

Nelson, S. J., 1963, Ordovician paleontology of the northern 

Hudson Bay Lowland: Geological Society of America 

Memoir 90,152 p., 37 pis.

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 Vaureal 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 

pis.

Owens, R. M., 1973, British Ordovician and Silurian Proetidae 

(Trilobita) : Paleontographical Society Monograph, p. 

1-98, pis. 1-15. [Published as part of v. 127.]

Pollock, C. A., Rexroad, C. B., and Nicoll, R. S., 1970, Lower 

Silurian conodonts from northern Michigan and Ontario: 

Journal of Paleontology, v. 44, no. 4, p. 743-764.

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 pis.

Ross, R. J., Jr., 1957, Ordovician Fossils from wells in the 

Williston Basin, eastern Montana: U.S. Geological Survey Bulletin 1021-M, p. 439-510,, pis. 37-43.

————1967, Calymenid and other Ordovician trilobites from 

Kentucky and Ohio: U.S. Geological Survey Professional 

Paper 583-B, p. B1-B18, 5 pis.

————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 

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and Idaho: U.S. Geological Survey Bulletin 1134, 177 p., 

13 pis.

Ross, R. J., Jr., and Dutro, J. T., Jr., 1966, Silicified Ordovician brachiopods from east-central Alaska: Smithsonian Miscellaneous Collections, v. 149, no. 7, 22 p., 3 pis.

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Schuchert, Charles, and Cooper, G. A., 1930, Upper Ordovician

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Shaw, F. C., 1968, Early Middle Ordovician Chazy trilobites

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————1974, Simpson Group (Middle Ordovician) trilobites of

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Department of Natural Resources, Report Investigations 

57. pt. 1, p. 1-59.
C22 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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I-VI.

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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

CONTENTS

Page

Abstract ____________________________________________ Dl

Introduction ___________________________________________ 1

Previous nomenclature ____________________________________ 1

Stratigraphy __————————______________________________———— 1

Extent and thickness _ ___ _______ _________________________ 3

Lithology __________________________________________———— 4

Environments of deposition __________________________________ 5

Depositional history __________________________________________ 6

Summary ____________________________________________ 8

References cited ________________________________________ 8

ILLUSTRATIONS

Page

FIGURE 1. Map showing location of study area and areas of exposure of the Rowland Member of the Drakes Formation and the Bull Fork Formation in north-central Kentucky _____________________ D2

2. Diagrammatic cross section showing relation of the main elements of the Marble Hill Bed to the enclosing

stratigraphic units within the study area ____ ———————— ___________________________ 3

3. Map showing thickness and stratigraphic position of the Marble Hill Bed within the Rowland Member __ 4

4. Cross sections showing position of Marble Hill Bed within Rowland Member _____________________ 5

5. Sketch map of study area showing interpretation of depositional environments ________________ 7

ill

SHORTER CONTRIBUTIONS TO STRATIGRAPHY 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 biosparrudite and biosparite that occur near the northern edge of 

the Rowland Member in north-central Kentucky and southeastern 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 platform 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 limestone in the Rowland Member of the Drakes Formation in north-central Kentucky and southeastern 

Indiana (fig. 1). It consists of several related limestone 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 U.S. 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 herein analyzed on the basis of lithology, field relationships, and fauna! content.

This report is based largely on field data obtained 

during the mapping of six 7 1/2-minute quadrangles 

in north-central Kentucky and reconnaissance mapping 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 buildingstone 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 assigns the pant of the stratigraphic section that includes the Marble Hill to the Waynesville Shale. In 

a summary of stratigraphic nomenclature of Indiana, 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 mudstone. The Marble Hill Bed comprises two clastic 

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 northDl
D2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

86° 85° 84° 83°

39°

38'

37°

Cincinnati

INDIANA OHIO

Area of figure 3

EXPLANATION

Area of exposure of Rowland Member 

of Drakes Formation

Area of exposure of Bull Fork Formation

FIGURE 1.—Map showing location of study area and areas of exposure of the Rowland Member of the Drakes Formation 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 interlayered 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 contents 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; 

echinoderms, 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 D3

N

Saluda Dolomite Member

Bardstown Member

V^ f^arWeHiH Rowland Member 

\^s Bed /-^

Drakes Formation

N 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 calcareous 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 commonly 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 northern 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 detail in following paragraphs.

Overlying the Rowland Member is the Bardstown 

Member of the Drakes, a unit of fossiliferous limestone 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

faunally so much like the Bull Fork Formation that 

it cannot be differentiated from the Bull Fork north 

of the pinch out of the Rowland Member (fig. 2).

The Saluda Dolomite Member is the upper member of the Drakes and consists of 20 m of dolomite 

and dolomitic to limy mudstone. It overlies the 

Bardstown Member and the laterally equivalent part 

of the Bull Fork.

EXTENT AND THICKNESS

The Marble Hill Bed is exposed in a small area of 

north-central Kentucky and southeastern Indiana 

south of Madison, Ind. (fig. 1). It consists of two 

clastic limestone bodies: a narrow, elongate unit at 

the base of the Rowland Member about 11 km south 

of the Rowland pinchout and a wider, 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 diseontinuously 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 section, the body is flat bottomed, convex, and asymmetrical and has the steeper face to the northwest 

(fig. 4, section B-B'). The basal contact appears 

conformable. The thickness averages about 1.2 m; 

the maximum observed thickness is 4 m. The body 

goes into the subsurface at the west end. Its subsurface extent is unknown.

The larger unit of Marble Hill limestone occurs 

along the northern.edge of the Rowland Member as 

an elongate body that parallels the basal unit and 

extends a short distance into Indiana before descending into the subsurface (fig. 3). It is as much as 9 m 

thick, thinning abruptly to the northwest and more 

gradually to the southeast (fig. 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 northwardthinning 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 extends for nearly 10 km 

to the east-southeast. For about half that distance,
D4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

85°22'30" 85°15'

Area of this figure

EXPLANATION

Marble Hill Bed at top of 

Rowland Member

Marble Hill Bed within 

Rowland Member

Marble Hill Bed at base of 

Rowland Member

Northern limit of Rowland

Member (where not coincident 

with edge of Marble Hill Bed

FIGURE 3.—Map showing thickness and stratigraphic position of the Marble Hill Bed within the Rowland Member. Contour 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 IVz -minute geologic quadrangles within which the Marble Hill Bed occurs.

this subunit occurs at the top of the Rowland Member; 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 disconformable. 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 

2m.

LITHOLOGY

The Marble Hill Bed typically consists of gastropod biosparrudite and echinoderm biosparite. Biomicrudite locally is an important component. The
A

30 r

20

o

I 0

c

30 r

Odb

Odr

Ob

Odb

Odr

Ob

THE MARBLE HILL BED IN THE UPPER ORDOVICIAN DRAKES FORMATION, KENTUCKY 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 bryozoan and brachiopod fragments. Grains are fine to medium, and 

sorting is generally good. The cement is mediumto coarse-grained sparry calcite, much of which occurs 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 conspicuous 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 bryozoan 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 consists of a basal unit of crossbedded biosparite overlain by an equal or greater thickness of thick-bedded 

gastropod biosparrudite. At some exposures., the two 

lithologies are interbedded, and infrequently the biosparite 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 

micritic limestone and calcareous mudstone of characteristic 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, wellaerated water of normal salinity. The jumble of 

whole and coarsely broken fossils in a micritic matrix 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

Odb

Odr

Ob

Odb

Odr

m

Ob

EXPLANATION

Bardstown Member 

Rowland Member 

Marble Hill Bed 

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 

brachiopod fragments, common to sparse echinoderm
D6 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

the Bull Fork was probably deposited below normal 

wave base but within the zone affected by storm 

waves. This environment is the shelf fades 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 pre-Rowland shelf is indicated by the southern limit of Bull Fork deposition, 

shown in figure 1.

The dolorniitic to limy mudstone of the type Rowland Member in southern central Kentucky was deposited in very shallow areas along the landward 

(south) side of a marine platform. The great sparsity of the marine fauna there indicates restricted 

circulation, and the doloniitic character of the sedimentary 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 mudstone 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 supported an abundant infauna, as evidenced by numerous burrows in the limestone. There was little wave 

action 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 ito a few tens of meters. The open 

marine platform on which the Rowland was deposited covered central Kentucky and extended 

northward. As. shown, by the superposition of the 

Rowland 001 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 narrow 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 biosparrudlte of the 

Marble Hill indicate deposition in a high-energy environment capable of extensive reworking. The 

abundance of sparry calcite cement in an area 

bounded, on the north and south by micritic limesitone indicates a narrow zone of high energy levels 

needed for thorough winnowing of carbonate sands.

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 platform. 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 generated 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 environment (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 probably affected all of central Kentucky. As a result of 

the shoaling, the shallow, restricted-circulation environment 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 Fork (fig. 1).

Where the wave base impinged on the rising sea 

floor, an offshore bar formed along a northeastsouthwest 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 northwest. 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 

sand. No evidence has been found to indicate that 

the bar followed the margin of the Rowland environment where it swings southeast (fig. 3). No corresponding 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 D7

85°22'30" 85°15'

FIGURE 5.—Sketch map of study area showing interpretation 

of depositional environments. Solid arrows indicate prevailing 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.

era 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 influx of terrigenous clastic 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 accompanying this influx caused the first bar to be abandoned and shifted bar deposition about 10 km seaward. A second bar formed on the basal mudsitone 

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 material carried onto the second bar by incoming waves 

was reworked in the turbulent zone along the bar 

and deposited as biosparite and biosparrudite. Gastropods living on the bar contributed large amounts 

of coarse skeletal material to the biosparrudite. The 

width of the main bar (fig. 3) indicates that deposition 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 buildup of sediments would have caused the Rowland to

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 biomicrudite 

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 finegrained calcite deposited on the platform was probably derived from the disintegration of organisms 

living in the lagoon and from the fine debris produced 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 consist 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 second bar 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, orachiopods, 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 restricted flow from the open shelf to the northwest, 

sio that tidal exchange was concentrated in a channel around the east end of the bar. The southeastward 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 environment, and a tongue of Rowland micrite and 

mudsitone 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 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 intertonguing 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 terrigenous 

clastic 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. Deposition of 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 Bardstown 

Member. The sharp basal contact of the Bardstown 

(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 indicating 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 gastropod 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 underlies the Rowland and is its lateral equivalent to the 

north, and the Bardstown Member of the Drakes 

Formation, which overlies the Rowland, were deposited on an open marine shelf that bordered the 

Rowland platform. The Marble Hill developed in a 

high-energy environment, where the wave base impinged on the sea floor between the open shelf and

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 central Kentucky to migrate northward over the southern edge of the open shelf. Deposition was ended 

by subsidence that brought about a return of the 

shelf environment.

REFERENCES CITED

Kepferle, R. C., 1977, Geologic map of the Ballardsville quadrangle, north-central Kentucky: U.S. Geological Survey 

Geologic Quadrangle Map GQ-1389, 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 1:24,000.

Owen, D. D., 1859, Report of a geological reconnoissance of 

the State of Indiana made in the year 1837, pt. 1: Indianapolis, John C. Walker, 63 p.

Peck, J. H., 1966, Upper Ordovician formations in the Maysville area: U.S. Geological Survey Bulletin 1244-B, 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 Bulletin 43, 229 p.

Swadley, W C, 1977, Geologic map of the Bedford quadrangle, north-central Kentucky: U.S. Geological Survey 

Geologic Quadrangle Map GQ-1409, scale 1:24,000.

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 quadrangle, 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, Galloway Creek Limestone and Ashlock and Drakes Formations (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. Geological Survey Geologic Quadrangle 

Mar> GQ-1334, scale 1:24,000.

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 fades 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 ___________________________——————————————-___ El

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 __________________________________________________ 11

Age __________v __________________________________________________ 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 Peninsula, 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

ill

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 chlorite, vermiculite, smectite, analcime, laumontite, clinoptilolite, 

wairakite, and the rare zeolite yugawaralite. These minerals 

reflect a complex history of alteration involving burial diagenesis, migration of hydrothermal solutions associated with 

intrusion of granodiorite plutons, and local thermal metamorphism 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 interbedded flows and pyroclastic rocks that contain secondary 

minerals (chloritSe, albite, actinolite, muscovite, epidote) characteristic of greenschiist-facies metamorphism. The age of 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 potential 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 metamorphism throughout the arc. However, absence of metamorphism in the Andrew Lake Formation indicates either 

that the Kula ridge was subducted at an earlier time (about 

50 m.y. 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 central Aleutian Islands, consists of late Cenozoic stratovolcanoes 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 intruded 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 tentatively dated as Paleozoic on the basis of leaflike impressions identified as Annularia stellata. Because 

rocks of unequivocal Paleozoic age were not known 

from other Aleutian islands, Scholl, Green, and Marlow (1970) re-examined these strata and found that 

the "Annularia" beds are in fact of middle or late 

Eocene age on the basis of foraminifers, dinoflagellates, and pelecypods. They named the fossiliferous 

strata the Andrew Lake Formation and reported 

that it rests depositionally on the Finger Bay Volcanics, 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 especially 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 basaltic 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).

El
E2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

177°00' 177°30'

52°00' —

51°45' —

BERING SEA

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 environments of the Andrew Lake Formation and the sedimentary and pyroclastic rocks at and near Wedge 

Point. We also assess the potential of these rocks as 

sources of hydrocarbons. We define mineral assemblages formed by hydrothermal processes as distinguished from assemblages developed by regional 

low-grade metamorphism. We speculate on the significance 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 U.S. Naval Station, Adak, for permission to work 

on the Naval Station and Comdr. Elton Himes for

providing transportation and logistic support for 

fieldwork. The U.S. Naval Special Service Corps 

provided logistic support for our works on Yakak 

Peninsula. H. N. Meeks of the National Oceangraphic and Atmospheric Administration Observatory, Adak, provided a vehicle during part of 

our work. Paul T. Fuller was our able field assistant. C. E. Gutmacher provided X-ray diffractograms 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 €orp., and George Claypool, U.S. Geological Survey, provided organic geochemical analyses and interpretation. We benefited from review
PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA

176°38' 176°36'

E3

51°57'-

51°56'-

51°55' —

EXPLANATION

c°

0).»

a> a>

c c

a> a>

(1

f..0) C 

Q. 0)

=>"£

I

( » ,_

L. C V, 0)

0) 0) •> c

Q_ O > Q)

| QTv |

Volconic

rocks

1

| Tol [ '

Andrew

Lake

Formation

f | Tfb |

1 Finger

Bay

a >

>• 1

< 1^

P <

a D

a:

I <

f i—

OL

LU

1—

^Volcanics^

• 803-600 

Isolated outcrop studied

0802-100 

Measured section locality

Strike and dip of beds

Topographic contour, 

in feet

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, Saitama University, Japan; and A. C. Waters, University of California at Santa Cruz.

ANALYTICAL PROCEDURES

Bulk rock samples and mineral separates were 

powdered and examined with a Norelco1 X-ray diffractometer. 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 described the Andrew Lake Formation. They suggested

1 Any 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 section 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 DEPOSITIONS 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 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

176°58' 176°56'

51°42'

51°4T

KILOMETERS

CONTOUR INTERVAL 100 FEET

Area of 

figure 3

Index Map

81 1-300.

Wedge Point

EXPLANATION 

o Measured section locality

FIGURE 3.—Sample locations and topography near Wedge Point, Yakak Pensinsula, southwest Adak Island. Measured sections 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 ashfall tuff attests to coeval volcanism. Numerous dikes 

and sills cut this part of the Andrew Lake Formation. Descriptions of the samples studied are given 

in table 1, in stratigraphic order.

Overlying this relatively coarse grained clastic 

section are the only richly fossiliferous lower Tertiary 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 tuff beds 

interbedded with quartz chert, laminated quartz porcelanite, 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 WEDGE POINT

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 

803-213

802-102

802-101 

802-103 

802-804 

802-803 

802-701 

802-802

802-202 

802-201 

802-204 

802-301 

802-303

802-501 

802-602

802-403 

802-401

^=3L 802-405

802-402

SECTION SOUTH 

OF WEDGE POINT

811-101

811-201a 

811-201b

811-401

811-302

EXPLANATION

^=| SHALE

SILTSTONE 

Jill SANDSTONE

LVo'l CONGLOMERATE, PEBBLY MUDSTONE 

|£\,&j PYROCLASTIC 

£55^ CHERT AND PORCELANITE 

|Vv y| VOLCANIC FLOW OR SILL 

802-301 SAMPLE NUMBER

810-406 

810-403

810-404

810-401

810-402 

810-502

810-301a 

810-301b

810-302

810-202

810-201

810-101

rl5

</> -in

' U

-5

L 0

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. Thickness of covered areas (breaks in sections) is unknown. Top part of Andrew Lake section and upper 7 m of Wedge Point section were measured. 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.—Description 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, actinolite; Al, albite; Am, amphibole; Au, augite; B, biotite; C, calcite; Ca, chalcedony; Ce, celadonite; Ch, chlorite; Cl, clinoptilolite; Cr, chert; 

Cz, clinozoisite; 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, smectite; Sp, sphene; St, stilbite; T, tremolite; V, vermiculite;

Z, undifferentiated zeolites. Tr., trace. He and St

Vg, altered volcanic glass; VRF, volcanic rock fragments; W, wairakite; Y, yugawaralite; 

occur only in sample 803 — 701]

Sample No. Location

802 

802

802 

802

802 

802

802

802 

802 

802

802 

802 

802

802 

802 

802

802

802 

803

803 

803 

803

803 

803

803 

803 

803

803

803 

803

803

- 402 

- 405

- 401 

- 403

- 602 

- 501

- 303

- 301 

- 204 

- 201

- 202 

- 802 

- 701

- 803 

- 804 

- 103

- 101

- 102 

- 209

- 211 

- 213

- 207

- 202 

- 205

- 203 

- 201 

- 204

- 200

- 206 

- 303

- 301

Andrew Lake area.. 

__ do .... _ ......

__ do ________

—.do .... __ . __

_ -do — ———— — .

....do ____ . _ ..

__ do ____ _-..-

__ do _________

__ do ... ____ . ...

....do ... _ . __ ..

__ do __ . __ ---

__ do _________

____ do ----_- —— ___

__ do _________ 

__ do _________

__ do _________

....do ____ ... _

__ do ___ .... __

__ do _________

__ do _...._....__.

. _ do ____ ...... 

-..do ....... _ —

.-.do — ..........

do

__ do .__...____._.

do

....do _______ ..

__ do _________

__ do _________

__ do _________

Rock type

Sandy pebble conglomerate. 

Silty mudstone and finegrained sandstone.

Lithic sandstone and 

silts tone.

Porphyritic dike _ ..-. 

Lithic sandstone and 

sandy and silty mudstone. 

Pebble conglomerate and 

silty mudstone and 

sandstone.

Pumiceous ash-fall tuff?.

Fine-grained sandstone 

and silty mudstone.

Fine-grained lithic 

sandstone.

Volcanic siltstone ______ 

Tuffaceous mudstone ...

Tuff? ____________

__ do __ —— . ____ ....

Chert and porcelanite ._

Chert and porcelanite — 

__ do . __________

__ do _——————__„___

Diatom chert and siliceous shale with chert 

nodules, 

.—do .... __ .........

Major primary 

constituents

VRF, Pr, P, Q, Vg 

Q, P, VRF, Pr, Vg

VRF, P, Vg, Q, I 

VRF, Vg, I, P, Q

P, Pr, Q 

VRF, Vg, Pr, Q, P

Vg, VRF, Q, I, Pr

Vg, P, Pr 

VRF, P, Au, Q, Vg

VRF, P, Q, Pr, Ac, 

E, C, Ph

P, VRF, Au, Ac, E, 

C, Mu 

P, Q, Au, H, B

P, Pr, Q, B

P, Vg, Q 

Q, P, Vg 

Q, P, Vg, Pr, B

P, Vg, Pr, Q

P, Pr, Q 

Q

Vg, P 

Vg, P, Pr 

Q, P, I. F

Q, P, F 

Q, P, F

Q, P, F 

Q, P, F 

T

Q, P, I. F

? 

Q, P, F, Am, VRF

Q, P, F, VRF

Secondary 

minerals

Al, V, Ch, Q, Fe, L 

Al, V, Ch, Q, Fe

Q, V-Ch, Ce, Cl, Ch, 

Fe, Al 

Ch, L, Q, Fe, V-Ch, 

Cl, S

V-Ch, L, Q, Fe, Al 

Q, V-Ch, Ch, L, Al, Fe

L, V-Ch, Q, Fe

L, Q, V, Fe

Q, V, Ch, L, Al, Ph, 

S, Fe

Q, V-Sm, L, Al, Fe

Q, Al, V, Ch, Fe, L 

Ch, V 

Q, V, Ch

V, Sm, Q 

Q, Z, Ch, V, C, Fe

Q, C, Ch 

V-Ch, Fe, Al, Q, Ch

V, Ch, Q 

Q, Tr, Z

Sm.Cl 

Sm-V, Ce, Fe, Q 

Q, Fe, Ch

Q, C 

Q, Py, Mn

Q, Py, Mn 

Q, Py, Mn 

?

Q, C, Py, Ce, V

?

Q, V-Ch, C, Fe

Q, C, V-Ch, Fe

Fossils

....do _„____ —

Rare fragments— 

Not looked for ._

__ do _„__—_.

....do ... __ -_

——do ... — —— 

— .do -------

....do -------

Not looked for ._ 

_-..do ...... — .

Diatoms and 

radiolarians. 

Benthic 

foraminifers.

Not looked for _. 

Dinoflagellate _

—.do ..__——

See table 2 ..— .

Diatoms, 

radiolarians. 

Microfossil molds.

___ do ------

___ do _____ -----

Benthio foraminifers.

See table 2 __ —

Diatoms, 

radiolarians, 

foraminifers. 

Diatoms, 

radiolarians, 

foraminifers, 

and flsh.

Comments

Poorly sorted 

up to 9 mm. 

Burrowed, lenses, 

contorted 

bedding.

Pods of vitric 

ash, lenses, 

contorted 

bedding.

Lenses of organic matter, 

burrowed. 

Grains compactpenetrate. Conglomerate is 

well graded. 

Clasts to 12 

mm. 

Nearly totally replaced by 

laumontite. 

Very poorly 

sorted clasts to 

13 mm. 

Sandy layers are 

poorly sorted 

and crudely 

graded.

K-Ar date 14 m.y.

Pillowlike 

structures.

Pods of tuff.

Completely altered, 14 cm 

below dike.

Much organic 

debris. 

Friable. 

Do. 

Late Eocene, 

upper bathyal.

Burrowed, 

laminated. 

Laminated.

Cretaceous 

through 

Oligocene. 

Laminated, late 

Eocene, upper 

bathyal.

Vermiculite clay 

nodules, shale 

is laminated. 

Laminated shale, 

hematite 

pseudomorphs 

after pyrite.

Lake Formation described by Scholl and others 

(1970) include aquagene tuff, ash-fall tuff and possibly ash-flow tuff, and volcanic dikes (fig. 4). These 

pyroclastic rocks differ texturally and mineralogically (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 underlying rocks and suggest that this uppermost section

may be younger than the underlying, more altered 

part of the section. Therefore, both the lower and 

upper contacts of this formation as originally proposed by Scholl and others (1970) are redefined 

here.

We redefine the base of the Andrew Lake Formation 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.—Description of rock samples from the Andrew Lake Formation, Finger Bay Volcanics, and Yakak Peninsula, Adak

Island, Alaska—Continued

Sample No.

803

803

803

803

803

803

803

810

810

810

810

810

810

810

810

810

810

810

810

812

811

811

811

811

811

811

803

805

QAO

- 302

- 403

- 403

- 401

- 601

- 602

- 604

- 101

- 201

- 202

- 302

- 301a

- 301b

- 502

- 402

- 401

- 404

- 403

- 406

- 104

- 302

- 401

- 201a

- 201b

- 101

- 201

- 101

- 101

- 701

Location

Andrew Lake area ..

....do __ .........

.—do ....... _ ...

....do ___ .......

__ do _____ . __

_ .do _____ . ....

....do ——__- — .

South of Wedge

Point.

__ do ________ .

....do .... .........

....do ______ . ...

.—do .... ..... ... _

__ do ____ . .....

——do — — .— —

__ do .... ______

__ do _________

__ do ____ ... ...

——do .............

— .do .... __ . __

_ .do _ .... .__.__

Wedge Point _ __

__ do — — -.---

__ do ___ — .....

.—do ——— ... ._—

__ do ___ .... __

__ do ____ ______

Southeast of Andrew

Lake.

Gannet Cove __. _ _

Bock type

Volcanic dike _ _ ______

Aquagene tuff _ __ _.

Ash-flow tuff ______._._

Ash-fall tuff _ ........

Clayey siltstone -.. _____

Lithic sandstone — — .

__ do ___.... __________

Welded tuff _.___._.._._

Volcanic dike ____._.___

Lithic sandstone _______

Tuff ____ . ____

Pebbly mudstone .___.__

Volcanic flow ----------

Sandy shale --------- _ .

__ do -. _ . ____ ....

Ash-flow tuff ____ ——___

Pebbly mudstone __- __

Lithic sandstone _ _ _-_

——do ... ... ..__. _-.._.

Silicified ash _ —— _ _

Pebbly mudstone ----- __

Volcanic dike ------ _ -

Pebbly mudstone _____ _

Volcanic dike — -------

Silty sandstone --------

Volcanic wacke ________

Lithic sandstone _-. ____

Quartzite and epidosite—

Major primary 

constituents

P, Pr

Vg, P, VRF, Pr

Vg, VRF, P, B,

Au, Cz

Vg, P, B, Am

Q, P, F

P, Q, VRF

P, VRF, Q

Vg, P, Pr

P, Pr, Vg

VRF, P, Pr

?

Vg, P, Au. VRF

P, Au

?

?

Vg, Au, P

Vg, P, W

VRF, P, Pr

VRF, P, Pr, Q

Vg, P

?

Au, P, Sp, M

Vg, P, Pr, M

P, Au, M

Vg, P, M, Au, Q

VRF, P, Pr, Q

P, Q, B, VFR,

Ho, Cr

Q, P, Cg, F, B

W, Ca, V, Cl, He, St

Secondary 

minerals

Al, Ch, V

Q, W. V, Ce, rare C

Q, W, Sm, Cl, Ch, C. V

Q, Cl, W. V, Sm

Ch, V-Ch, Fe, Q

V-Ch, C, Q, Ch, Al

V-Ch, C, Q, Ch, Al, Fe

Y, Ch, Ce, Fe, I-Mu, L

Vesicles: Q, S-CeCh, L or C; Matrix:

Fe, C, Q

Y, Q, Al, Ch, Fe

?

Y, Q, Fe, Ch, V, Tr,

L,A1

Vesicles: W, S, L;

Matrix: Al, Fe, I

?

?

I, Ch, V, Ce, C, Fe,

Tr, Y, L

An.Q, Ch, Ce, Y(?)

An, Fe, clays

Z, Q, Ch, Fe

Q, Ch, Fe, C

?

Vesicles: Q,

Ch+S+Ce, L, Q,

Tr, Pr, E; Matrix:

Q, L, Al, E, Fe, Ch, I

Q, C, Ch. S, Ce, Fe, A

Vesicles : L,

S+Ch+L, Q;

Matrix: Q, Fe, Ch

Q, Fe, Ce, I-Mu,

Y, L,A

Al, Q, Y, Fe, L, Ch

Ac-T, E, Q, Fe, Ch, V

E, C, Q, Mu, Fe, Ch,

Pi?

Fossils

Not looked for __

__ do ... _ ....

..-do ..........

Barren ._. _ -.-_

Foraminifers,

echinoid.

Barren -------- -

_-__ao -.- . -.--

-..do _ . _ ...

Not looked for ._

Barren __________

— ..do ___„____.

——do __________

Not looked for ._

Barren __.... __

.—do ----- .....

Not looked for __

Barren _-.-. .--_

....do __ ......

——do -------

Spicules? .- _-- _

Barren __________

Not looked for —

Barren __ .....

Not looked for __

Barren . ._.-- —

...-do ------

— ..do ------

Ghosts of microfossils.

T^ni" 1 nnlf f>f{ fnv

Comments

Unaltered minerals and some

glass (?).

Unaltered minerals.

Unaltered minerals and some

glass (?).

Baked by nearby

dike.

Do.

Purple with green

mottling.

Poorly sorted,

compact, grains

penetrate.

Lahar.

Slightly welded.

Pumiceous lahar.

Banded, well

sorted, crudely

graded.

Poorly sorted.

graded.

More than 50 percent altered to

quartz and

laumontite.

Pumiceous.

Poorly sorted,

pumiceous.

Finger Bay Volcanics, compact.

Finger Bay Volcanics, poorly

sorted.

Bay Volcanics.

volcanic rocks are assigned to the Finger Bay Volcanics. 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 containing diatoms or devitrified ash are locally abundant; 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 terrigenous sedimentation. Consequently, mainly biogenic 

material accumulated, although an active infauna 

was not present.

Other sedimentary structures and current-direction indicators are rare. They include ripple laminations and poorly developed graded bedding. Sandstone beds are graded and have flute casts, and sandstone 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 volcanogenic 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 MeDougall, R. Z. Poore, and W. V. Sliter]

Sample No. Benthie forams Other mierofossils Age range Depth of water

802-101

803-200

803-207

803-601 

803-204

Bathysiphon sp. _________ —— __-.

Cyclammina pacifica"! ________.

Cyclammina sp. __________

Dentalina sp. _____________

Eponides sp. ___——— — — — — _

Gyroidina soldanii d'Orbigny ____

Lenticulina sp. ____________.

Valvulineria sp.? _____________

Bathysiphon eocenicus ________

Bathysiphon sp. ___ — __________

Cyclammina pacifica Beck ______.

Dentalina sp. ____________.

Eponides sp. ____________.

Gyroidina soldanii d'Orbigny ____.

Lenticulina sp. _____ — _ — __ — ___.

Melonis iimbilicatulus (Montague) 

Cyclammina sp. _________________

Bathysiphon eocenicus! ________

Rhabdammina eocenica! _______.

Echinoid spines 

Radiolarians __

Indeterminate 

Late Eocene _. Upper bathyal.

Echinoid spines 

Fish debris __-_ 

Radiolarians _

Late Eocene ____ Upper bathyal.

Echinoid spines Intermediate 

Cretaceous 

through 

Oligocene.

laumontite occur as interstitial cement and replace 

volcanic debris along with vermiculite, chlorite, 

smectite2 and vermiculite-chlorite (randomly interlayered). 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, chlorite, and calcite. These minerals, commonly found in 

greenschist-facies metamorphic rocks, occur as detrital minerals in these three samples (table 1). Adjacent to dikes and sills, however, secondary prehnite and epidote have formed (for example, prehnite 

in sample 802-204, table 1).

Because minerologically unstable volcanic debris 

is less abundant, fewer secondary minerals characterize the chert-porcelanite part of the section. 

Quartz and minor calcite and pyrite are the most 

important secondary minerals. Quartz indiscriminately 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 tuff in the Andrew Lake Formation has 

altered to smectite and minor clinoptilolite. In places 

smectite was subsequently converted to vermiculite.

2 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.

Locally vermiculite, chlorite, hematite, and celadonite 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. Vermiculite, 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 common 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 mierofossils and megafossils. 

Our collections of benthic foraminifers (table 2) 

indicate that the Andrew Lake Formation was deposited during the late Eocene, virtually identical to 

the foraminiferal age assigned in Scholl and others
PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA E9

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]

Sample No.

802 - 802

Mineral

Plagioclase

Percent 

K»O

0.079

»Arrad

( moles/g )

1.640X10-12

»Arrad

40Artotal

a 04

Calculated 

age 

(millions 

of years )

14.4±3.5

TABLE 4.—Organic geochemistry of Tertiary sedimentary rocks 

from Adak Island

[Analyses of total organic carbon (TOC) by U. S. Geol. Survey Organic 

Geochemistry Laboratory, Lakewood, Colo. Analyses of chloroform extractable bitumen (EOM), vitrinite reflectance (Ro), and EOM/TOC 

by Mobil Oil Corp., Dallas, Tex.]

*°K decay constants: \e = 0.572 X 10-10/yr, \c' = 8.73 X 10~13/yr, 

\l> = 4.905 X 10-10/yr. Abundance ratio «K/K = 1.167 X 10~* 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±3.5 m.y. (table 3).

We speculate that the pyroclastic rocks that overlie the Andrew Lake Formation as redefined, but included 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 alteration of samples from only slightly lower stratigraphically. 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 granodiorite plutons and related dikes and pyroclastic deposits (Fraser and Snyder, 1959). However, eruption 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 extractable bitumen), and R0 (vitrinite reflectance). 

These quantities, as well as EOM/TOC, 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 R0 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/TOC ratio can be interpreted as resulting 

from the "cracking" of organic material; the breakup of organic compounds probably resulted from the

Sample 

No.

803 - 204 

803 - 206 

810 - 302 

810-402

810 - 502 

811-302

Location

Andrew Lake 

__do _______ 

Wedge Point 

_ do ____

do __ __ 

_ do _____

TOC 

( weight 

percent)

0.41 

.04 

.05 

.09

.07 

.04

EOM

(ppm)

28.6 

28.7 

44.3 

40.2

29.3 

26.4

R0

2.1+ *

*

*

* 

*

EOM/ 

TOC 

(percent)

0.70 

7.10 

8.90 

4.50

4.20 

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/TOC 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 tuff); volcanic dikes cut the section. Sandstone beds are primarily 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 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

monly cemented by laumontite, yugawaralite,3 iron 

oxides, or clays.

Pebbly mudstone beds contain mostly subangular 

grains, although a complete range of grain roundness 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 lapilli?). 

Texturally, it appears massive to weakly flow 

banded. Sample 810-101 (figs. 3 and 4; table 1) is 

mostly glass globules and flattened pumice fragments (greater than 3 mm) replaced by yugawaralite. Sample 810-401 has pumice with an elongation 

ratio of 20:1. Large plagioclase and augite glomerocrysts (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. Vesicles 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, intrusive 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 tuff and boulder lahars 

and the apparent absence of microfossils suggest

3 Yugawaralite, 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 Marshall, 1965) and from Japan (Sakurai and Hayasbi, 19S2; 

Seki and Okumura, 1968; Sameshima, 1969; Seki and others, 1969).

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 tuff 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 deposition; if so, Wedge Point rocks are in part subaqueous 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). Plagioclase 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 dominantly 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 remaining pore spaces (commonly very small, maximum 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 (probably 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 amygdules suggest the following paragenetic sequence: 

quartz (occasionally replaced by laumontite), 

granular material (unidentifiable), analcime, phyl-
PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA Ell

losilicates (stilpnomelane, chloride, celadonite, rarely vermiculite or hematite), laumontite, phyllosilicates 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 hydrothermal 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 determination of vitrinite reflectance was removed from 

any of these samples.

FINGER BAY VOLCANICS

Coats (1956) and Fraser and Snyder (1959) described 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 feldspathic arenite. The main framework grain is plagioclase 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 feldspar 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, epidote, and muscovite are secondary minerals that now

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 generally less than 40° occur (Fraser and Snyder, 

1959).

METAMORPHISM

Secondary mineral assemblages (Fraser and 

Snyder, 1959) suggest that the Finger Bay Volcanics was subjected to regional greenschist-facies 

metamorphism. The diagnostic mineral assemblages 

range from actinolite- or tremolite-epidote-chloritealbite-quartz to epidote-quartz-muscovite-chloritealbite (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 coherent regional pattern of metamorphic facies. Certainly, the actinolite and epidote greenschist assemblages appear to be dominant on Adak Island.

It is not clear whether emplacement of granodiorite plutons contributed significantly to metamorphism (contact metamorphism) of the Finger 

Bay Volcanics or whether metamorphism was dominated by a regional thermal event. Fraser and 

Snyder (1959) described only a thin zone of contact-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 exposed plutonic rocks, these formations show significant differences in metamorphic mineral assemblages. More fieldwork is needed to fully distinguish 

regional patterns from contact metamorphism.

In contrast to the Finger Bay Volcanics, sedimentary and pyroclastic rocks on Yakak Peninsula and 

the Andrew Lake Formation at Andrew Lake have 

not been subjected to greenschist- or even zeolitefacies regional metamorphism. These rocks have 

been moderately altered by low-temperature supergene and hydrothermal processes of the zeolite facies 

and nowhere show greenschist-facies metamorphism. 

Thermal metamorphism associated with emplacement of dikes and sills, together with the supergene
E12 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 apparently devoid of fossils. Fossil findings indicate 

that the Andrew Lake Formation accumulated during 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 Volcanics and associated sedimentary rocks evolved 

through a sequence of deposition, burial, regional 

greenschist-facies metamorphism, uplift, and erosion before the Andrew Lake Formation was deposited. 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 Paleocene (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 Volcanics 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 alteration. Variability of alteration, presence of fresh 

calcic plagioclase, only weakly altered volcanic glass, 

and the occurrence of a variety of temperature-sensitive zeolites argue against regional greenschistfacies 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

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 exposed 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 environment. 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 recognizably younger strata, the Andrew Lake Formation. Probably by middle Eocene time, the growing 

Aleutian ridge had nearly reached sea level, and 

subaerial volcanoes contributed debris to surrounding 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 difficult to identify the origin of the silica in the chert beds. It is not 

clear whether the silica in the chert-porcelanite section of the Andrew Lake Formation is released by 

the dissolution of siliceous biogenic debris, deposition from hydrothermal solutions, or both. Diatoms 

and minor radiolarians from the quartz chert occur 

in all states of preservation, from ghosts to specimens that retain frustule ornamentation. This suggests that at least part of the silica was derived 

from dissolution of siliceous microfossils. The calcite in this section is probably redeposited carbonate
PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA E13

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 temporally equivalent to the Andrew Lake Formation, 

then they possibly represent the subaerial and shallow-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 subaerial flanks of a volcanic complex and in adjacent 

offshore shelf and slope environments. The presence 

of a small enclosed basin, one isolated from turbidity-current deposition, is evident in the laminated chert and porcelanite of the Andrew Lake 

Formation.

ALTERATION OF SEDIMENTARY ROCKS

Burial diagenesis, local thermal metamorphism by 

dikes and sills, and hydrothermal activity contributed to the alteration of Eocene rocks. Burial diagenesis was an important process in the early stages 

of alteration of these rocks, primarily because they 

contain a large fraction of unstable mafic to intermediate 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 wallrock. 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.

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 temperature zeolite, stratigraphically overlies laumontite and clinoptilolite, minerals characteristic 

of a relatively lower temperature metamorphic 

facies (Coombs, 1961; Harada, 1969; Seki and 

others, 1969).

2. Nonequilibrium mineral assemblages are common; 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 laumontite, 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 associated 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 deposits (Bundy and Murray, 1959; Lovering and 

Shepard, 1960; Heystek, 1963; Steiner, 1968), (3) 

secondary mineral assemblages (except in the biogenie 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 Marshall, 1965; Harada and others, 1969).

Hydrothermal solutions associated with emplacement of plutons and with the contemporaneous volcanic activity apparently permeated the Paleocene 

rocks and sealed any available pore space by deposition of secondary minerals. Rocks close to the main 

thoroughfares of circulating fluids were 50-75 percent replaced. Deposition of secondary minerals in
E14 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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). Circulating 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 laumontite, 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, 200-300°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 complex (for example, Atwater, 1970; Grow and Atwater, 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 terminate 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 m.y. 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

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 metamorphosed, subduction of the Kula ridge must have occured before the late Eocene, about 50 m.y. ago. 

Models for North Pacific plate motion allow subduction of the Kula ridge at 50 m.y. or 35 m.y. ago depending on whether relative motions have been discontinuous or continuous, respectively (Cooper and 

others, 1976, fig. 4). Although the plate models are 

approximations, there is increasing evidence of discontinuous 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 m.y. ago (Hayes and 

Pitman, 1970; Hein, 1973; Hamilton, 1973; Scholl 

and others, 1977; also, see Francheteau and others, 

1970; Larson and Pitman, 1972). Therefore, evidence for the timing and the effects of ridge subduction as proposed by DeLong and others (1978) 

can be interpreted variously. More likely, early 

Tertiary (Paleocene or Eocene) metamorphism resulted from the depositional burial and tectonic uplift of more than 4000 m of volcanic and sedimentary 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 equivalent to the upper Eocene, Andrew Lake Formation. 

However, despite the similarity in lithologic composition, degree of alteration, and induration, agediagnostic fossils must be found in rocks in the 

Wedge Point area to justify including them as part 

of the Andrew Lake Formation. These rocks accumulated approximately 40 m.y. 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 significantly reduced porosity and produced authigenic 

clay minerals, iron oxides, and possibly quartz. Additional alteration occurred in conjunction with intrusion of sills, dikes, and especially granodiorite 

plutons. Secondary minerals formed during these 

late-stage thermal and hydrothermal events are 

primarily zeolites (yugawaralite, laumontite, analcime, wairakite), clays, iron oxides, and quartz. 

Most pore spaces that remained after burial diagenesis were filled during this episode, essentially
PALEOGENE SEDIMENTARY AND VOLCANOGENIC ROCKS, ADAK ISLAND, ALASKA E15

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 subsequently "cooked" by the heat of igneous intrusions. 

Consequently, these strata are unlikely sources of 

hydrocarbons.

The Finger Bay Volcanics was regionally metamorphosed to the greenschist facies some time before 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 metamorphism in the Andrew Lake Formation suggests 

other interpretations. If Adak is typical of other 

Aleutian islands, either the Kula ridge was subducted about 50 m.y. ago or greenschist metamorphism need not accompany subduction of a 

spreading center.

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1969, Chemical composition and optical properties of
E16 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

yugawaralite from the type locality: Am. Mineralogist, 

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162-178.

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of Rat Island, Aleutian Islands, Alaska: U.S. Geol. Survey Bull. 1028-Q, p. 555-562.

Lovering, T. S., and Shepard, A. 0., 1960, Hydrothermal 

argillic alteration on the Helen claim, East Tintic District, Utah: Clays and Clay Minerals, v. 8, p. 193-202.

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T. R., 1973, Tectonic history of the central Aleutian Arc: 

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405.

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Amateur Mineralogists), v. 20, p. 70-78.

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151, p. 1-31.

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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

CONTENTS

Page

Abstract __________________________________________ Fl

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

FIGURE 1. Generalized bedrock geologic map of the central part of the Livengood quadrangle, showing the Livengood

	Dome Chert and other major rock units —— ——————————— - ——————————————_____——____ F3

2. Reference section of the Livengood Dome Chert in the Lost Creek borrow pit _______________ 4

3. Map showing the location of the Lost Creek borrow pit and cross section A-A' ________________________ 6

4. Geologic cross section A-A' through the Lost Creek borrow pit, showing the structural interpretation __ 8

5. Photograph of graptolite-bearing shale beds in the Livengood Dome Chert at the Lost Creek borrow pit 9

6. 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, JR.,

and CLAIRE CARTER

ABSTRACT

The discovery of Late Ordovician graptolites in the Livengood 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 formation consisting of chert, limestone, dolomite, shale, and argillite 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 underlain by Cambrian and Preeambrian(?) argillite, slate, quartzite, siltstone, limestone, and chert. Extensive chert units 

similar to the Livengood Dome Chert in lithology and stratigraphic 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 understanding 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 paleontologic knowledge of Paleozoic rocks in central 

Alaska have provided data to allow revision of several 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 Ordovician on the basis of graptolites that were discovered 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 result of geologic mapping and studies in the Livengood 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, Jr., and 

Claire Carter was a major contribution to this study 

(U.S. Geological Survey, 1972, p. A51-A52). Regional 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 (Chapman, Yeend, Brosge, and Reiser, 1975; Chapman, 

Yeend, and Patton, 1975).

We have benefited also from the geologic investigations in the Livengood area and the description of 

several chert thin sections by Robert L. Foster, from 

paleontologic and stratigraphic studies in the Livengood and White Mountains areas by J. Thomas

Fi
F2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

Dutro, Jr., and from geologic mapping by Donald 

Grybeck, who worked in the eastern part of Livengood 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 helpful 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). However, 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 8% miles" 

(Metrie, 1937, p. 105). This belt is within the Livengood quadrangle. Mentie (1937, p. 105-111) identified 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 extensions that occur as isolated beds in the vicinity of 

Sawtooth Mountain and as "metamorphosed equivalents 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 commonly white to cream, crystalline, and in various 

stages of silicification and is less commonly dark 

gray, noncrystalline, thin bedded, and mostly unsilicified. 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."

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 measurement of true thicknessi is possible, but "a considerable thickness of beds, perhaps several thousand 

feet," 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 Athyris sp. This collection was accepted, with 

certain reservations, as of Carboniferous, probably 

Late Mississippian, age, although it lacked the more 

diagnostic Late Mississippian 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 Livengood 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 eastnortheast-trending belt (fig. 1) between the Mud 

Fork and the headwaters of Victoria Creek, a disstance of about 91 km. The width of this belt generally is about 5 km but ranges from 1.6 to 9.6 km. 

Minor amounts, of some other rocks, either unrecognized or too small to be shown at the map scale,
THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT F3 

149°00' 148°30' 148°00'

65°45' ——

65°30'

0 5 10 KILOMETERS 

I 1 I

EXPLANATION

CRETACEOUS

AND 

Graywacke, JURASSIC (?) conglomerate, shale, and quartzite

TRIASSIC

AND 

PERMIAN (?)

Sedimentary and basaltic volcanic rocks of Rampart Group (Permian?),

and intrusive mafic rocks (Triassic)

Graywacke, conglomerate, and shale

Dolomite and limestone, in part silicified, chert, shale, siltstone, and

argillite; also includes some mafic intrusive and volcanic

rocks and serpentinite of uncertain ages

Livengood Dome Chert

Chert, shale, siltstone, claystone, tuff, and minor limestone

QUADRANGLE 

LOCATION

1. Tanana

2. Livengood

3. Circle

4. Charley River

5. Kantishna River

6. Fairbanks

Argillite, slaty shale, quartzite, grit, siltstone, some limestone, and minor chert

CAMBRIAN

AND 

JPRECAMBRIAN(?)

Contact 

Queried where uncertain

Inferred fault Site of Lost Creek borrow pit

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 U.S. Geological Survey quadrangle, scale 1:250,000 (1956).
F4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

500—. METERS

450 —

400—

350-

300 —

A A A A

= -=-=^=( t

Continued next page

North margin of pit 

Maraon shale

A A A

A A 

A A A

Palagonite tuff,dark-olive-green with white carbonate spots,- 

thin section 71 AWr-570-33A from 450 m. Maroon shale, 

ferruginous, slighty silty, thin-bedded

Chert, gray, ond some claystone, olive-drob, 

with splintery fracture

Shale, drob-olive-gray, splintery fracture. Limestone, 

medium-gray, weathers brown, very fine to fine-grained, 

beds <. 30 cm thick, thin white calcite veins. Lithic tuff, 

moderate yellowish brown, fragmentol rock; thin section 

71 AWr-570-31 B from 405 m. Shale and tuff are soft and 

mostly very thin bedded

Y TT 

TY Y

EXPLANATION 

Chert

Claystone, clay shale, 

silt shale, in part tuffaceous

Limestone

Tuff and tuffaceous rocks

Fossiliferous beds 

| Graptolite

<>• Radiolarian 

ts Thin section

—*— Fault zone,- clayey gouge, dark brown

Chert, light-gray to medium-dark-groy, 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 71AWr-570-30 

from 366 m

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, 

Q. and pale-green clay shale partings; white efflorescence on some clayey beds. Very soft, round white clay balls 

o 2.5 cm in diameter, in one medium-gray claystone bed. All beds are 2.5 to 7.5 cm thick and evenly layered

___/ Shale and chert, dark-gray, in alternating beds 2.5 to 5 cm thick; limestone, dark-gray, 

_______sulfide-beoring, rare; thin section 71AWr-570-27 from 280 m___________________

Shale, drab-olive-gray, breaks into tiny splinters, and limestone, medium-dark-gray, weathers brawn; 

beds 5 to 8 cm thick, cut by thin veins of dark-gray calcite

Chert, gray, weathers brawn, and claystone, light-gray, weathers yellow in part; alternate in beds 8 cm thick

FIGURE 2. — Reference section of the Livengood Dome Chert in the Lost Creek borrow pit (see fig. 3 for location). The stratimined. The amount and direction of movement on the fault are unknown,

are included. The exposures are such that a complete section and the exact lower contact could not 

be found, and the upper contact was seen at only

one place. Within this belt the Livengood Dome 

Chert is best exposed in the area between the valleys 

of Lost Creek and the South Fork of Hess Creek,
THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT F5

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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

2_ \ grayish-black, graphitic, greatly sheared /

o • ,,.._/

: 

- Poorly exposed; chert, medium-gray, weathers dull brown except for a 4.6 m yellowish-orange band in 

j 

2 midsection, fairly massive beds 30 to 60 cm thick, some slightly banded

/ Chert, medium-gray and locally

greemsn-gray, weathers gray /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 maroon tints, 

possibly tutfaceous; chert, gray and greenish-gray, and some claystone, dark-gray, with light-gray 

_____ v clayey clasts,- in thin alternating beds. One bed of black shale 8 cm thick

Chert, light-gray, and interbedded claystone, very light gray; one bed of chert, grayish-green, 15 cm thick

Claystone, dusky-yellow to light-olive-gray, and partings of clay shale, pale-olive; form 1 cm chips.

\ One bed of chert, medium-dark-gray / '

Clay shale and interbedded tuffaceous siltstone, various shades of gray and yellow, beds ^.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

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 yellowish-orange, moderately porous, contains 

devitrified glass shards and finely divided iron oxide. Thin section 71 AWr -570-5 from 49 m

Mostly covered; chert, gray, weathers pale yellow to dark yellowish orange, beds possibily ^ 30 cm thick

Chert, light-gray, weathers ochre, black manganese stains, beds 2.5 to 10 cm thick; clay shale.

Chert, greenish-gray, weathers moderate brown, bedded , vitreous, conchoidal fracture, brittle, close jointed

1

South margin of pit

graphic orientation and continuity of this section are uncertain because tops and bottoms of beds rarely could be deterand some beds may be repeated owing to undetected isoclinal folding.

which is herein designated its type area. The forma- prominent topographic feature in the type area. The

tion name is taken from Livengood Dome, which is j thickest section, a reference section (fig. 2), was

underlain largely by this chert and is the most ! uncovered during 1970-71 in a large highway bor-
F6 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

T 9 N 

T 8 N

148"50'

30' R 7 W. R. 6 W

SCALE 163360

5 KILOMETERS

FIGURE 3.—Location of the Lost Creek borrow pit and cross 

section A-A' (fig. 4). Base from U.S. Geological Survey 

Livengood C-4 quadrangle, 1951.

row pit 13.6 km west of the town of Livengiood and 

about 2.4 km west of Lost Creek in the SWi/4 sec. 8, 

T. 8N., R. 6W. (fig. 3).

West of the Mud Fork a belt of chert and metachert, 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, Brosge, 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, Brosge, and Reiser, 1975; Chapman, Yeend, 

and Patton, 1975) and in the northwest corner of the 

Fairbanks quadrangle (Pewe and others, 1966). 

These cherts outside the Livengood quadrangle are 

tentatively correlated with the Livengood Dome 

Chert on the basis of liithologic similarities and stratigraphic and structural position* There is no 

definitive paleontologic evidence that these cherts

are all the same age. However, the recent discovery 

of radiolarians 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 quadrangle, the belt of chert mapped by Mertie (1937, 

pi. 1) between Beaver and Preacher Creeks was examined 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 Tinitina fault, has been reinterpreted 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 consists 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 

tuff, 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 F7

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 natural outcrops and rubble patches, and our observations of them have been made chiefly in manmade 

cuts. A distinctive thin unit of small-pebble conglomerate, consisting of chert pebbles in a siliceous 

or chert cement, is present at several places near 

Livengood and Lost Creek. This chert pebble conglomerate apparently is in the lower part of the 

formation, but owing to complex structure and discontinuous outcrops, its stratigraphic position is 

uncertain.

In thin section the chert is cryptocrystalline but 

includes some micro-crystalline quartz. In part it 

contains altered remnants of radiolarians (71AWr570-30) ; some sponge spicules are also visible in 

hand specimens. Minor amounts of sericite and argillaceous grains are present. Hairline fractures filled 

with quartz, iron oxides, and mica are abundant. One 

section shows evidence of at least three sets of fractures 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 plagioclase, colorless mica, chlorite, a zeolite (probably 

laumontite), and a large amount of what appears tjo 

be altered glass. The veinlets are largely quartz or 

zeolite plus quartz (J. W. Hawkins, written commun., 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 secton shows cryptocrystalline to microcrystalline 

chert, probable devitrified glass shards, and some 

yellow, red, and brown earthy iron oxides and microcrystalline 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 tuffaceous rocks (thin sections 

71AWr-570-31B and -33A) are crystal to vitric 

tuff and vesicular palagonite tuff, which are generally less extensive and derived probably from basaltic 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

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 conglomeratic 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. 60-70° 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 metamorphism 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 borrow pit has an apparent thickness 

of about 459 m, but because this section is cut by a 

fault of unknown displacement and includes dip reversals, it has a smaller and unknown true stratigraphic thickness. An estimated thickness of 300- 

600 m for the Livengood Dome Chert would be compatible with its relatively broad outcrop belt in the 

Livengood quadrangle and with similar size belts of 

probably correlative chert units in adjacent quadrangles.

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° S., 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 position. 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

N A'

METERS 

600-i 

500- 

400- 

300- 

200- 

100-

SEA LEVEL-

**?„

conceded

FEET

-2000

-1500

-1000

-500

100 300 500

Horizontal scale

2000 FEET 

I

EXPLANATION

—— — — Contact - Dashed where inferred

— ?— Fault-Queried where uncertain

FIGURE 4.—Geologic cross 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 darkgray 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 predominantly of very light; gray to light-gray dolomite and 

limestone, is commonly slightly to nearly completely 

silicified, includes a minor amount of medium-darkgray to black chert occurring as thin beds, lenses, 

and nodules and contains some thin beds of slaty 

shale, siliceous siltstone, and argillite. The uniformly 

dark color of this chert contrasts with the various 

shades of gray and other light colors that are characteristic 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 conodonts, corals, and brachiopods made in recent years 

and of collection 18AOF8 (Mertie, 1937, p. 110), 

the maximum age range of the fauna in the apparently 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, Jr., written commun., 

1963). The brachiopods have a possible range of 

Middle Silurian through Early Devonian, but a Middle Silurian age is considered most reasonable by J. 

T. Dutro, Jr., (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 serpentinite 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 unconf ormable and probably 

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. Subsequently this pit has been smoothed by bulldozing; in
THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT F9

1976 most of the bedrock section was not well exposed, and the graptolite beds were covered.

A section having an apparent stratigraphic thickness 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 tuffaceous 

rocks and a few thin beds of limestone. The beds 

strike about N. 70° E., dip 50°-70° N. in the southern part of the section, and dip about 70° S. in the 

north part of the section. Graded beds in one outcrop 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-section specimens, is given in figure 2. A diagrammatic

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 characteristic 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. Therefore, this section is presumably representative of a 

part of the formation and does not include either of 

its contact zones.

PALEONTOLOGY

Graptolites were discovered in and collected (collection 71ACn-391) from thin shale beds interbedded 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 71ACnFIGURE 5.—Graptolite-bearing shale becis, indicated under the hammer, in the Livengood Dome Chert at the Lost Creek

borrow pit.
F10 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

391) was made from this site by others. Both collections of graptolites present problems in identification 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 unfortunately not found here.

Climacograptus aff. C. longispinus supernus Elles 

and Wood (fig. 6A) resembles Climacograptus 

bicornis 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. I. 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. /. supernus lacks. It is 

more than 57 mm long and 0.7-1.8 mm wide. C. I. 

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. QD) 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. miserabilis s.s. in the spacing of its thecae: C. miserabilis 

s.s. has 10-11 thecae per centimeter. C. miserabilis 

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).

Amplexograptusl pacificus pacificus (Ruedemann) (fig. 6fl) is characterized by a short rhabdosome 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 imperfect 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 Kolyma 

region of the northeastern U.S.S.R. A.? p. pacificus 

is found in Upper Ordovician rocks in Idaho (Rue^- 

demann, 1947, p. 429; Ross and Berry, 1963, p. 125 ; 

Riva, 1974b, p. 1457) and in the Climacograptus 

supernus Zone (= Dicellograptus ornatus Zone) of

H

I/'

G

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, Climacograptus sp., USNM241052, collection 71ACn-391. C, Diagram illustrating parts of thecae, as discussed in text: 1, 

apertural excavation, 2, thecal aperture, 3. supragenicular 

wall, 4, geniculum, 5, genicular spine. D, Climacograptus 

cf. C. miserabilis Elles and Wood, USNM241053, collection 

71ACn-391. E, Climacograptus sp., USNM241054, collection 71ACn-391. F, Climacograptus sp., USNM241055, collection 71ACn-391. G, Amplexograptusl sp., USNM241056, 

collection 71ACn-391. H, Amplexograptust pacificus pacificus (Ruedemann), USNM241057, collection 71ACh-287.
THE LIVENGOOD DOME CHERT, ALASKA, AND DISPLACEMENT ON THE TINTINA FAULT Fll

the Kolyma Basin and Kazakhstan, U.S.S.R. (Koren' 

and Soboleyskaya, 1977).

The specimen illustrated in figure 6G is questionably 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 centimeter) and differs in the details of the proximal 

end, but it possesses the outwardly inclined supragenicular 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 consists of numerous large-and medium-size climacograptids (figs. QB, 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 rhabdosome 

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 uniserial 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 correlated the Livengood Dome Chert fauna with 

faunas from the upper part of the Late Ordovician 

Zone of C. supernus (=Dicellograptus 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 Holdsworth 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 Kantishna River quadrangle (D. L. Jones, oral commun., 

1978). Conodonts were looked for but not found in 

the shale interbedded with chert in and near the 

Lost Creek borrow pit.

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 contains 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 limestone and dolomite fragments that have not been 

recognized in the Livengood Dome Chert. The Road 

River in the Charley River area unconf ormably overlies a thick section of mainly limestone (Brabb, 

1967) that in its upper part is rich in Cambrian 

trilobites (Palmer, 1968). The Road River underlies 

the McCann Hill Chert, a thin-bedded chert and 

siliceous shale unit that has a basal limestone unit 

rich in shelly fossils of Early Devonian age (Churkin and Brabb, 1965; Churkin and Brabb, 1967).

The fossiliferous Cambrian limestone and distinctive 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 stratigraphically below the Livengood Dome Chert along 

the north side of its outcrop belt are mainly siliceous. 

However, the presence of a thin fossiliferous limestone immediately overlying the chert, chert wacke, 

and silty shale of the Livengood Dome Chert suggests 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 correlatives in the Charley River area. For example, 

the coarse clastic rocks of Late Devonian age near 

Livengood may be a fades equivalent of the Nation 

River Formation; the volcanic and clastic rocks of 

the Rampart Group of probable Permian age may 

be related to the Circle Volcanics and associated 

clastic 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 graywacke formations exposed in the Livengood and
F12 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 Territory boundary, separates greenschist, quartzite, 

phyllite, and greenstone of Paleozoic age on the 

south from essentially unmetamorphosed Precambrian, 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 apparently splays into the several faults that are mapped 

in the northeastern part of the Livengood quadrangle (Chapman and others, 1971). On the south 

side of these faults near the eastern boundary of the 

Livengood quadrangle, unmetamorphosed rock sequences 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, Roddick (1967) estimated a 352- to 416-km displacement on the Tintina-Rocky Mountain trench system, 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 Tintina 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

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, phyllite, and siltstone that in part contains Oldhamia, a 

fan-shaped trace fossil of either Cambrian or late 

Precambrian age. The structurally complex sequences 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 (U.S. 

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 Ordovician 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: 

U.S. Geological Survey Miscellaneous Geologic Investigations Map 1-573, scale 1: 250,000.

Chapman, R. M., Weber, F. R., and Taber, Bond, 1971, Preliminary geologic map of the Livengood quadrangle, 

Alaska: U.S. Geological Survey open-file report, scale 

1:250,000.

Chapman, R. M., Yeend, W. E., Brosge, W. P., and Reiser, H. 

N., 1975, Preliminary geologic map of the Tanana and 

northeast part of the Kantishna River quadrangles, 

Alaska: U.S. 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: Calgary, 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 pis.

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 F13

Ordovika-Silura (Severe-Vostok S.S.S.R.): Doklady Akademia 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. Geological 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.

Pewe, T. L., Wahrhaftig, Clyde, and Weber, F. R., 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 faunal provinces, in Rickards, R. B., Jackson, D. E., and Hughes, C. P., eds., 

Graptolite studies in honour of O. M. B. Bulman: Paleontology Special Paper 13, p. 107-126, pi. 7.

————1974b, Graptolites with multiple genicular spines from 

the Upper Ordovician of western North America: Canadian 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 

pis.

Ruedemann, Rudolf, 1947, Graptolites of North America: 

Geological Society of America Memoir 19, 652 p., 92 pis.

Tempalman-Kluit, D. J., Gordey, S. P., and Read, B. C. 1976, 

Stratigraphy and structural studies in the Pelly Mountains, 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 Moffat area, south Scotland: 

British Museum (Natural History) Bulletin, Geology, v. 

19, no. 1, p. 1-26, 16 pis.

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 relationship of the

Star Point Sandstone and Blackhawk Formation

and its bearing to coal correlations

CONTENTS

Page

Abstract __________—___—_____________———__————————— Gl

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 of contact and the zone of intertonguing 

between the Star Point Sandstone and the Blackhawk Formation, the location of the panoramic photograph, and the location of the diagrammatic sections __________________________ G2

2. Composite panoramic photograph and diagrams 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., 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

ill

SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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

ABSTRACT

Geologic mapping and detailed measurement and description of closely spaced stratigraphic sections across the contact 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 intertonguing 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 requires 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 Convulsion 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 tihe lower part of the Blackhawk: a lower group 

consisting of the Hiawatha and Upper Hiawatha 

coal beds and aon 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 considered 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 detailed 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 observed at several localities within a 10-km-long and 

1-km-wide belt extending south-southeastward from 

the north wall of Muddy Canyon to a dfownfaulted 

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 photograph 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 recognize 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

Gi
G2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

WASATC H

PLATEAU

EXPLANATION

£§£:££ Zone of intertonguing

Contact between Blackhawk 

Formation (B) and Star Point 

Sandstone (SP)

Fault-Dashed where inferred. 

U, upthrown side; D, downthrown 

side

A——A' Line of diagrammatic section 

^ Coal mine

114° 

42°

40'

38° —

110°

UTAH

STUDY AREA

0 50 100 KILOMETERS 

I I I

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 correlation®, 

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 cut out by a channel sandstone of the 

Blackhawk Formation between sections 464 and 

463. The Muddy No. 2 coal bed of section 464 is correlated with the Hiawatha coal bed of sections 461 

and 463.

Figure 3 and 4 are diagrammatic sections constructed from closely spaced stratigraphic sections 

that were measured along the Wasatch Plateau front 

(A-A't 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 G3

by a tongue of the Star Point Sandstone. The coal 

bed immediately overlying the Star Point in areas 

to the northeast is generally traceable southwestward 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 apparently was unaware of imtertonguing between the 

two rock units within the coal field. Spieker's concept 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 necessary 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), 

as 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 stratigraphic 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 equivalent to the Hiawatha coal bed of the revised correlation. 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 significant paleogeographic and paleoenvironmenital implications that are beyond the scope of this report. 

Some of these implications are presented in preliminary form by Marley, Flores, and Cavaroc 

(1978) and in more detail by Flores (1979) and 

Marley, Flores, and Cavaroc (1979).

REFERENCES CITED

Blanchard, L. F., 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-Sanpete, 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 Tertiary and Cretaceous coal fields in the U.S. Western Interior: International Association of Sedimentology, Organic 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 Report 77-822.

Marley, W. E., Ill, and Flores, R. M., 1977, Descriptions of 

stratigraphic sections, Upper Cretaceous Blackhawk Formation and Star Point Sandstone in the Emery West and 

Flagstaff Peak quadrangles, Utah: U.S. Geological Survey Open-File Report 77-833, 257 p.

Marley, W. E., Ill, Flores, R. M., and Cavaroc, V. V., 1978, 

Lithogenetic variations of the Upper Cretaceous Blackhawk Formation and Star Point Sandstone in the Wasatch Plateau, Utah: Geological Society of America Abstracts 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 OpenFile Report 77-823, scale 1 :24,000.

Spieker, E. M., 1931, The Wasatch Plateau coal field, Utah: 

U.S. Geological Suryey Bulletin 819, 210 p.

Spieker, E. M., and Reeside, J. B., Jr., 1925, Cretaceous and 

Tertiary formations of the Wasatch Plateau, Utah: Geological Society of America Bulletin, v. 36, p. 435-454.
G4 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

464

Muddy No. 2 -T 

coal bed

EXPLANATION

Coal

Si

Bony coal

Shale

^

Siltstone

E23 

Sandstone

Correlation line-Queried 

where extent uncertain

Kb

463

VERTICAL EXAGGERATION X 25 

METERS

461

>-Hiawatha 

coal bed

FIGURE 2.—A, Composite panoramic photograph of the north wall of Muddy Canyon in EV2 sec. 36, T. 20 S., R. 5 E., and 

WV2 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

464 461

Muddy No. 2 / 

coal bed

/-Hiawatha 

I 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 proposed in this paper. (Coal-bed names are those of Spieker.)
G6 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

Kb

Kb

0.1 0.2 0.3 0.4 KILOMETERS

I

EXPLANATION

[^Q&j Sandstone 

5BSiif Carbonaceous shale I I Shale and siltstone

>— Coal--Dashed where 

inferred

FIGURE 3.—Diagrammatic section A-A' (southwest to northeast) along Wasatch Plateau front (see fig. 1) showing relalations 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 13 detailed stratigraphic sections.

Kb

Kb

METERS 1—10

0.1

EXPLANATION

^^ — — Coal-Dashed where inferred 

€>>=E=E-B> Carbonaceous shale

0.2 KILOMETERS

I

Sandstone

Shale and Siltstone

FIGURE 4.—Diagrammatic section B-B' (southwest to northeast) along downfaulted block near Emery (see fig. 1) showing 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

CONTENTS

Page

Abstract ______________________________________________ HI

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

ill

SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

NEW EVIDENCE SUPPORTING NEBRASKAN AGE 

FOR ORIGIN OF OHIO RIVER IN NORTH-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 Madison 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 northern Kentucky and of the glaciated portion of the 

Ohio River has undergone considerable change in 

recent years. Prior to about 1950, the glacial deposits of northern Kentucky were generally considered to be of Illinoian age. In 1956 Rich 

determined that the upland till in southwestern Ohio 

and northern Kentucky was 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 position of the deposits, they identified the lower 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 northeastward 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 glacial 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 U.S. 

and Kentucky Geological Surveys.

Mapping in north-central Kentucky west of Carrollton has revealed three small deposits of deeply 

weathered till and outwash that occur at low 

topographic levels. The low topographic position and 

deep weathering of these deposits, and their distribution 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 Kentucky consisted of three principal streams; the Kentucky, 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 outward-dipping escarpments capped by more resistant 

limestones and dolomites of Silurian age. The preglacial 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 Carrollton 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,

Hi
H2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

85° 84°

39'

Q)^ 

X

"Great Miami, 

River

Cincinnati ^-"'^

/

/

/

/

\ 

/ ^ Area of figure 2

38"

\

KENTUCKY

EXPLANATION 

Preglacial drainage

Modem drainage. Shown only where 

not coincident with preglacial 

drainage

KILOMETERS 50

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 northeastward into central Ohio, where it 

joined the preglacial master stream of the northcentral states, the Teays-Mahomet River.

The preglacial Licking River headed in eastcentral 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 j oin the Kentucky near 

Hamilton, Ohio, about 32 km to the north. The preglacial 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 advance 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 Kentucky 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 RIVER, NORTH-CENTRAL KENTUCKY H3

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 reversing the direction of flow of the preglacial Kentucky. 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 tributary 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 protected localities. The high-level Nebraskan Drift and 

its partial covering of Kansan deposits were subjected to further weathering and erosion that 

destroyed any distinctive topographic expression.

The advance of the Illinoian Glaciation into Kentucky was very limited. Southeast 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 Kentucky 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.

The Sangamon Interglaciation, following the 

Illinoian Glaciation, was another period of erosion 

and weathering. Tributaries of the Ohio were probably deepened and extended headward, but erosion 

along the major stream was limited largely to removal 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 

ponding of tributary streams.

GLACIAL DEPOSITS

The glacial deposits of this study lie along the 

Silurian escarpment at the western edge of the central 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 shales 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 southwestward 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) consist of till and lesser amounts of outwash. Most are 

deeply weathered. A high-level till forms thin, discontinuous 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 produced by 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 TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

85°30' 85°00'

INDIANA

KENTUCKY

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

- 500-foot (152-meter) topographic contour

0 KILOMETERS 10

I I I I I I I I I I I

FIGURE 2.—Map showing location of glacial deposits described in this study

Many igneous rock fragments are thoroughly 

weathered and crumble easily. In all exposures 

studied, the drift was completely leached of carbonate 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 occurrence was reported by Ray (1957). At the first locality (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).

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 elevation of 230 m at the north end to 170 m at the 

south. The drift is poorly exposed but seems to consist largely of weathered sand and gravel that 

contains no calcareous minerals. The drift includes 

chert, quartz, quartzite, and igneous and metamorphic 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 RIVER, NORTH-CENTRAL KENTUCKY H5

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 elevation of about 230 m. The drift consists of 

weathered, noncalcareous silty clay till similar to 

the high-level till but less oxidized. Boulders, including 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, locality 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 outwash. Near the base, an abandoned borrow pit exposes 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 lightbrown sandy clay till with scattered chert, quartz, 

and igneous clasts. The till, which is leached of calcareous minerals and includes igneous boulders that 

crumble easily, rests on ledges of Ordovician limestone.

The third occurrence is about 6.4 km upstream 

from Madison, Ind., at the base of the south wall of 

the Ohio River valley (fig. 2, locality 3). This deposit, 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, Gibbons, 1978). Here drift and loess form a low spur 

about 12 m high that juts northward from the main 

valley wall. Drift is exposed along the nose and east 

side of the spur in natural exposures as well as along 

a drainage ditch that cuts across the north end of 

the spur. A steep face at the end of the sp'iir 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 excavation for a pond spillway exposes 2.1 m of till resting 

on limestone. The till is pale-yellowish-brown noncalcareous 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

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 deposit 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 limonite-stained 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 consists of meduim-gray calcareous clay till that contains 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 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 unavailable 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 report suggests that the major breaching of the Madison divide and the initiation of the Ohio occurred 

after the first glacier had withdrawn from its maximum 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 deposition 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 preglacial drainage, Thus, the elevations would have 

been about 200 meters within the old Kentucky valley 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 highlevel till, after a lengthy period of active stream 

erosion during which the Ohio and its small tributaries 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 upstream along the Ohio valley (Swadley, 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 preIllinoian 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 indications 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 weathering, 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, Kentucky: U.S. Geological Survey Geologic Quadrangle Map 

GQ-1471, scale 1:24,000.

Leigh ton, 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. B126-B131.

Ray, L. L., 1957, Two significant new exposures of Pleistocene 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 H7

Rich, J. L., 1956, Pre-Illinoian age of upland till in southeastern Indiana, southwestern Ohio, and adjacent parts 

of Kentucky [abs.] : Geological Society of America Bulletin, v. 67, no. 12, pt. 2, p. 1756.

Swadley, W C, 1971, The preglacial Kentucky River of northern Kentucky, in Geological Survey research 1971: U.S. 

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: U.S. 

Geological Survey Geologic Quadrangle Map GQ-1123, 

scale 1:24,000.

——1976, Geologic map of part of the Carrollton quadrangle, Carroll and Trimble Counties, Kentucky: U.S. 

Geological Survey Geologic Quadrangle Map GQ-1281, 

scale 1:24,000.

——1977a, Geologic map of the Bedford quadrangle, northcentral Kentucky: U.S. Geological Survey Geologic Quadrangle Map GQ-1409.

——1977b, Geologic map of part of the Bethlehem quadrangle, 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 quadrangle, Trimble County, Kentucky: U.S. Geological Survey 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

CONTENTS

Page

Abstract ________________________________________—____—____ II

Introduction _______________________________————————— 1

Acknowledgments ________________________——————————— 2

Regional setting _____________________________———————— 2

Description of geologic units ______________________________—__——___ 3

Paraopeba Formation _________________________——————— 3

Unit A ________________________________________— 3

Unit B _____________________________________———— 3

Unit C ___________________________________________ 7

Structure __________________________________——————— 8

Folds ______________________________________________ 8

Lineations ____________________________——————— 8

Faults _________________________________________— 8

Mineralization and reserves ___________________________________——___ 10

Summary __________________________________—————— 11

References cited __________________________________________—___ 11

ILLUSTRATIONS

Page

FIGURE 1. Index map of Brazil and the Vazante region _____________________________________ 12

2. Geologic map and cross sections of the Vazante zinc district, Minas Gerais, Brazil _____________ 4

3. 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 _____ 6

4. Photograph of thin-bedded, folded phyllite of unit A ___________________________________ 7

5. Panoramic view of Serra do POC.O Verde, viewed from the south _________________________ 7

6. Diagrammatic examples of ways in which breccia zones may form along step-bedding faults ________ 10 

7u Photograph of resistant ridge of mineralized breccia (primarily hemimorphite, hydrozincite, willemite) on

the north end of Serra do POC.O Verde _________——____——____——_________________ 11

ill

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 *

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 erogenic belt (ca. 600-500 m.y.). Bedding and lowangle 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 thrusting 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 designated 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 weathering 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-producing area in Brazil, is about 130 km. south of Paracatu 

and 120 km northwest of Patos de Minas in the western part of the State of Minas Gerais (fig. 1). Companhia Mineira de Metais at Serra do Pogo Verde 

and Companhia Mercantile Industrial Inga at Serra 

do Ouro Podre (fig. 2), the only companies operating 

in the district, produced approximately 110,000

1 Companhia de Pesquisa de Recursos Minerals, Ministerio das Minas e 

Energia, Rio de Janeiro, Brazil.

metric tons of zinc metal from 1965 to 1974 (White 

and Nagell, 1975; Bohomolete, 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 Projeto Bambui, a reconnaissance geochemical exploration program of the Banir 

bui basin undertaken by the CPRM (Companhia de 

Pesquisa de Recursos Minerals) 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 Vazante-type 

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 

Poco Verde. Subsequently, Carvalho, Dequech, and 

Guimaraes (1962), Branco (1962), Guimaraes 

(1962), and Cassedanne (1968a,b) reported on the 

Vazante deposit. These studies were restricted primarily to the deposit and to the narrow cuesta along 

which mineralization occurs. Amaral (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 photographs (fig. 2). Five days in March and April 1974 

were spent field checking the photogeologic interpretations.

li
12 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

STATE OF 

MINAS GERAIS

• Manaus

• Porto V el ho

BRAZIL

•Cuiaba Brasilia^ 

Goiania*

FIGURE 1.—Index map of Brazil and the Vazante region. Triangulo Mineira area (patterned) south of Vazante, studied 

by Barbosa and others (1970).

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 "mineralized cuesta" to refer to the entire feature.

ACKNOWLEDGMENTS

We wish to acknowledge the many fruitful discussions 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 assistance is gratefully acknowledged. Dr. Percio 

Branco, CPRM made several pertinent, helpful suggestions during the final preparation of the manuscript. 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 Bambui basin, a late Precambrian basin (1,000-600 m.y.) filled with siltstone, marl, conglomerate, sandstone, limestone, and 

dolomite of the Bambui Group. The clastic and pelitic 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 (oa. 600-500 m.y.); the 

orogeny formed a major north-trending deformed 

belt about 1,000 km long, characterized by lowangle thrust faults, bedding-plane faults, large 

folds, and widespread low-grade metamorphism. 

Allochthonous rocks in the belt include the following Precambrian units, in ascending order (Barbosa 

and others, 1970); granite-gneiss complex; mica 

schist, quartzite, and amphibolite of the Araxa 

Group; quartzite and phyllite of the Canastra 

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 easternmost major thrust brings Canastra over the Bambui (fig. 1). Smaller scale thrusting and broad folding 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 13

Previous workers interpreted the structural setting at Vazante as one of normal faulting or tensional 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 structural 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 report follows the work of Braun. We divided the 

Paraopeba Formation into three informal units (A, 

B, and C, in ascending order). Figures 3 A and 3J5 

are diagrammatic stratigraphic sections of the Bambui 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 contact between unit B and C—the Vazante fault—is 

less clear. Along the southern half of the fault's exposure, 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 lime^- 

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 relationships, 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 quartzite 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 bedding. Fracture and crenulation 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 northeast-trending 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 lightgray, aphanitic to fine-grained dolomitic limestone. 

Collenia, an algal stromatolite, occurs in some beds 

(Cassadanne, 1968a), and siliceous layers are locally 

present. A few thin beds of medium-grained, wellsorted, light-colored quartzite were noted near the 

crest of Serra do POC.O Verde. Interbedded with the 

limestone are thin-bedded, in part laminated marl, 

shale, and siltstone that vary in color from red to 

buff to brown to green.

The massive to thick-bedded limestone underlies 

cuestas and isolated hills that alternate with elongate and irregularly shaped valleys underlain by 

marl, shale, silstone, and, in some places, limestone. 

This topography is typical of unit B. A major characteristic 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 containing thin argillaceous partings grades laterally 

into shale and thin-bedded limestone in a distance 

of about 20 m.
14 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

46'J 45'

EXPLANATION 

UPPER PRECAMBRIAN 

PARAOPEBA FORMATION 

Siltstone, limestone, sandstone, 

and marl

Limestone, Siltstone, and shale 

Limestone and shale 

Siltstone, shale, and limestone 

Limestone, Siltstone, and shale 

Limestone 

Marl and shale 

A I Phyllite and quartzite

FIGURE 2.—Geologic map and cross sections of the Vazante district, Minas Gerais, Brazil.
RECONNAISSANCE GEOLOGIC STUDY, VAZANTE ZINC DISTRICT, MINAS GERAIS, BRAZIL 15

NORTHWEST 

D

SOUTHEAST 

D'

-UNITC- -UNITBSANTA CATARINA 

FAULT

METERS

1200 -,

800-

400

SEA 

LEVEL

-400

-400

-UNIT C- -UNITBUNITC'

MINERALIZED CUESTA

SANTA CATARINA 

MINEIRA FAULTx^ FAULT

SEA 

LEVEL

B

METERS * 

1200-,

800

-UNIT C UNITB- -UNIT

B'

MINERALIZED CUESTA SANTA CATARINA 

MINEIRA FAULT FAULT

SANTA CATARINA 

FAULT

-400

3km

NO VERTICAL EXAGGERATION

LITHOLOGIC UNITS ON CROSS SECTION

Sandstone 

Limestone 

Siltstone and shale 

Marl and shale 

Phyllite and quartzite

Thrust fault—Arrows show

direction of relative movement. 

Dashed where inferred

Trace of marker bed from aerial photographs

Trend or lineament from aerial photographs

Contact—Dashed where approximate

Possible facies change—Photogeologic 

interpretation

Thrust fault—Dashed where inferred. 

Teeth on upper plate. Arrows show 

direction of relative movement

MAP AND CROSS SECTION EXPLANATION 

—*-?-*— Possible thrust fault along bedding plane 

^ ~, Tear fault bounded by thrust fault or

bedding fault 

ATTITUDES OF BEDS 

_i_ Strike and dip 

'£~s Strike and dip showing plunge of

lineation 

. i Interpreted from eerial photographs

LARGE-SCALE FOLDS-Showing 

crestline and direction of plunge 

Anticline 

Syncline 

SMALL-SCALE

plunge

~ ~ Mineralized breccia zone 

===== Road 

—— Stream

FIGURE 2.—Continued.
16 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

GROUP

BAMBUI'

FORMATION

TRES 

MARIAS

PARAOPEBA 

(Fig.3fi)

PARANOA

LITHOLOGY

Arkose, micaceous siltstone, graywacke, 

arkosic sandstone, argillaceous limestone 

lenses

LAGUNAR CONGLOMERATE —————— -

SAMBURA CONGLOMERATE —————— - 

Pelitic and carbonate rocks including facies 

named "Sete Lagoas," "Lagoa Jacare," 

"Serra de Santa Helena". Sandstone, 

siltstone, argillite, marl, limestone 

CARRANCAS CONGLOMERATE ———— -

Quartzites with interbedded phyllite and 

metasiltstone

SAG MIGUEL CONGLOMERATE ———— »•

CANASTRA GROUP

--^ESril£n=^

T / / 1

-1 J-Iy-i5

~~L-=:J~~ —

THICKNESS, 

IN METERS

20-400

1000-1400

100-3800

FORMATION THICKNESS, 

IN METERS

INFORMAL

UNITS, THIS

REPORT

LITHOLOGY

500

-VAZANTE 

.FAULT

LITHOLOGIES

DOLOMITIC LIMESTONE

SANDSTONE

IRON FORMATION

> \MINERALIZED 

JZONE 

SMINEIRA 

FAULT

-SANTA 

CATARINA -^.~^r-~] FAULT

B

-•- SI LTSTONE WITH SHALE

:=-:-: PHYLLITE

: ~~ PHYLLITIC SI LTSTONE

FIGURE 3.—Diagrammatic stratigraphic sections. A, Stratigraphic section of the late Precambrian 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 subumits, numbered 

B-l at the base to B-6 at the top (fig. 2), primarily 

on the basis of photointerpretation of geologic and 

geomorphic features. Some of the subunits are 

shown intertonguing with others (fig. 2), a reflection 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 marl 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 17

FIGURE 4.—Thin-bedded, folded phyllite of unit A. These are 

fi 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 characterized by vecry low relief and probably is composed largely of fine-grained clastic and argillaceous 

rocks interlayered with some thin-bedded limestone; 

a few isolated hills are underlain by thin- to mediumbedded limestone. Subunit B-5 is characterized by 

numerous low cuestas and hills of moderately dipping 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 massive limestone beds that are discontinuous; the intervening areas are flat pasture land underlain by 

thin- to medium-bedded limestone, siltstone, shale, 

and marl.

Subunits B-1, B-3, and B-4 were not separated 

in the northern part of the area, because B-3 becomes 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 B-5 and B-6 are not 

distinguishable immediately northeast of Vazante 

because of folding and possibly because of stratigraphic 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 dolomitic limestone. These rocks are lithologically similar to those in unit B, but they have a much higher 

ratio of siltstone and marl to limestone. Limestone 

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 POC.O Verde, viewed from the south. The valley to the right (northeast) is underlain by shale and marl of subunit B-1. The ridge is Serra do POC.O 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.
18 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

STRUCTURE

The Vazante district displays many structures 

typical of a deformed belt. During- the Brazilian 

orogeny (ca. 600-500 m.y.), older Precambrian 

rocks were thrust eastward over the western margin of the Bambui basin along a front approximately 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 orog-eny are widespread at 

Vazante, and at least two phases of tectonism can 

be documented. First, low-angle, bedding, and stepbedding- thrust faults developed, as did some related folding". Low-grade metamorphism and penetrative deformation either preceded or was coeval 

with the thrusting-. Second, broad cross folds deformed 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.

FOLDS

Four sets of folds (fx through ft) are recognized. 

Set fx consists of folds that trend N. 20°-50° W. 

and range in size from microfolds less than 1 mm 

to some having wavelengths of 2-5 cm and amplitudes of 1-3 cm (fig. 4). This fold set is best developed in unit A, is well developed along the contact 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 sandstone beds.

Set f 2 consists of two north-north west-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 wavelengths and amplitudes of about 0.5 km and are at 

least 2 km in length. The sets of fx and f2 folds are 

parallel.

The third fold set, f3 , consists of large-scale structures that trend N. 10° W. to N. 40° E., parallel to 

regional structural trends. All units and fi and f2 

folds are folded about f3 axes. The folds are broad,

open, and slightly asymmetrical. They are overturned 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-northwest-plunging 

folds transverse to the regional trend. All structures 

in the northern part of the area are folded about 

f4 axes.

Set f! folds, the oldest, formed early in the Brazilian 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 fi and f2 axes are more or less normal to 

the direction of Brazilian thrust faulting. Set f;i 

folds are interpreted as being related to the major 

period of Brazilian folding and thrust faulting. 

Transverse folding of all structures by f.t folds is 

the youngest event, postdating all thrust faulting.

LINEATIONS

Three types of mutually parallel lineations were 

noted: streaks of fine-grained sericite, fi 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 fi 

folds were present and are parallel to< fi axial planes. 

Quartz-filled fractures aire 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 carbonate beds. The thinner the argillaceous bed between 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 faulting. 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 pronounced and pervasive penetrative deformation and 

low-grade metamorphism of unit A rocks and the 

generally locally developed and much weaker lowgrade meitamorphism of unit B rocks. Beds on opposite sides of the fault are generally more or less 

concordant; however, minor but important discordances do occur. In the basal member of unit B, 

several marker horizons traceable on aerial photographs 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 stratigraphic 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 involve 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 related 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 responsible must be confined by these marker beds. 

We suspect that multiple bedding faults and stepbedding faults formed, and, differential movement 

of the intervening beds created tear faults that terminated 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 underlying layers.

Two examples of this type of tear-fault relationship 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 alongside the road, a distance of about 10 m from limestone 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 siltstone and shale are present. The change from limestone 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 (Branco, 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 sequence 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 several 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 breccia zone and the regional dip, which is 30° to 45° 

southwest, are folded about f4 fold axes. Removal 

of the regional dip and f4 folding would change the 

breccia zone from a curved and near vertical structure to a, linear, moderately steeply southwest-dipping structure like that of a step-bedding fault (fig.
no SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

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 flf f2, and (or) f3 folding but definitely 

before f4 folding. Some variations in the possible 

forming1 of a sfep-bedding-plane breccia zone are 

shown in figure 6.'Multiple 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 portion of the breccia zone some distance from the 

lower portion, which in turn would be overlain by 

unbrecciated "upper plate" rocks. Other interpretations for the origin of the breccia include: a highangle fault, as proposed by previous workers; a 

fault-zone vein that predated f4 folding but was 

postthrusting, and perhaps was of a rebound tensional nature; a faulted and metamorphosed Mississippi-valley type or similar type of "stratabound" 

deposit; a faulted and metamorphosed massive sulfide 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 (referred to as calamine in Brazilian reports of the 

district), willemite, hydrozincite, cerussite, smithsonite, chalcocite, brochantite, pyromorphite, covellite, zincite, cuprite, native copper, malachite, 

linarite, aurichalcite, acanthite, and native silver 

(Amaral, 1968). The main ore minerals are hemimorphite, 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 111

FIGURE 7.—Resistant ridge of mineralized breccia (primarily 

hemimorphite, hydrozincite, willemite) on the north end of 

Serra do POC.O Verde. Trees do not grow over the mineralized 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 toms 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 reserves seem to indicate that total zinc metal reserves could be as much as 3 million metric tons.

SUMMARY

The Vazante district, which occurs along the leading edge of a deformed belt, has been subjected to 

considerable compression, and the rocks have responded accordingly. Low-grade dynamic metamorphism 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 m.y.).

The breccia zone along the mineralized cuesta is 

only a very small part of the area, but is economically important. The breccia probably formed in response to step-bedding-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

as high-angle faults or collapse-breccia zones. The 

step-bedding-plane faulting resulted in the brecciation of carbonate and argillaceous rocks. The amount 

of stratigraphic separation is unknown but it probably 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 diolomitic limestone of the 

Paraopeba Formation. Cassedanne (1968b) arrived 

at such a conclusion for at least part of the mineralization, pointing out that CoWem'a-bearing dolomite in the district has an abnormally high zinc and 

lead content. Amaral and Kawashita (1967, 1968) 

dated shales at Vazante from the Sete Lagoas Formation (which we refer to as the Paraopeba Formation) as 600 ±22 m.y. old. Amaral (1968) reported 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 concentration of 'the deposit into its present form as an 

epigenetic deposit probably was made possible because Brazilian o'rogenic 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 

oir 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 common in the breccia zone, but they reflect probably 

only the latest history of the zone.

This interpretation suggests that future exploration for Vazante-type deposits should be concentrated along the deformed belt where folded carbonate rocks are interbedded with argillaceous 

rocks. This coincidence of structural and stratigraphic situations is favorable for disharmonic 

folding, which, in turn, leads to bedding and stepbedding faulting. The end result of these events 

may have been the formation of breccia zones that 

were potential depositional sites for ore-bearing 

fluids.

REFERENCES CITED

Amaral, Gilberto, 1968, Contribuigao ao conhecimento dos 

depositos de Zn-Pb-Cu-Ag da Serra do POC.O Verde, Vazante, Estado de Minas Gerais: Sociedade Brasilerra
112 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979

Geologia Congresso, 22d, Belo Horizonte, 1968, Anais,

p. 13-31. 

Amaral, Gilberto, and Kawashita, Koji, 1967, Idade do grupo

Bambui: [abs.]. Parananaense Geociencias Bol. 26, p.

39-40. 

————1968, Idade do grupo Bambui por determinacoes 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, Bol. 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, Principals ocorrencias de zinco

em Minas Gerais: Sociedade Intercambio Cultural e

Estudos Geologia, Ouro Preto, pub. 2, p. 151-184. 

Braun, Oscar P. G., 1968, Contribuicao a estratigrafia do

grupo Bambui: Sociedade Brasileira Geologia Congresso,

22d, Belo Horizonte, 1968, Anais, p. 155-166.

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Brazil Divisao Fomento Produgao Mineral Bol. 110,119 p.

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Brasileira Ciencias Anais, v. 40, no. 2, p. 215-225.

————1968b, Nota sobre o ambiente de sedimentacao das 

rochas encaixando a mineralizagao de Vazante (MG) : 

Sociedade Brasileira Geologia Congresso, 22d, Belo Horizonte, 1968, Anais, p. 33-40.

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Estudos Geologia, Ouro Preto, pub. 2, p. 101-147.

Moore, Samuel L., 1956, Zinc and copper deposits of the 

Vazante area, Minas Gerais, Brazil: U.S. Geological Survey open-file report, 16 p.

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Constraints on the Latest Movements 

On the Melones Fault Zone, Sierra 

Nevada Foothills, California

By J. 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

CONTENTS

Page

Abstract ——_—__—____________—_______—___—_____—__—__—_———- Jl

Introduction ________ ___ __ ___________————————————————— 1

Acknowledgments ________—__—________—_____—___————————- 1

Age of the Mehrten Formation ________________—_————————— 1

Age of dacite intrusions ________________________—————————— 3

Discussion __________________________————————————- 4

References cited ______________________________—_—————__———_ 4

ILLUSTRATIONS

Page

FIGURE 1. Geologic map of the area southwest of Mokelumne Hill, Calaveras County, Calif. ——————————————— J2 

2. Diagram showing the correlation of some Tertiary stratigraphic units in the western Sierra Nevada, showing age control from K-Ar dating of volcanic rocks -____ ——————————————— ———————————— 3

TABLES

Page

TABLE 1. Chemical data on late Tertiary intrusions _____—_—————————————————————————————————— J3 

2. Analytical data for K-Ar age determinations ______— ———————— __——————————————————— 4

III

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 near Mokelumne 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 colluvium 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 Formation, a late Tertiary volcaniclastic unit. Inasmuch 

as the dacite and rhyodacite bodies intrude the 

Mehrten and are locally overlain by it, dating the intrusions would give an age within the Mehrten Formation 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 colluvium 

containing clasts of rhyodacite apparently identical 

with the rhyodacite of the Hamby dome intrusion. 

The minimum offset on the colluvium-Mehrten contact 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 WoodwardClyde 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 clastic 

and pyroclastic material. The formation was originally described by Piper, Gale, Thomas, and Robinson (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 Mobelumne Hill area, the formation consists mainly of 

coarse andesitic conglomerate and minor interbedded 

sandstone (Rose, 1959). The Mehrten unconformably 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) considered 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

Jl
J2 SHORTER CONTRIBUTIONS TO STRATIGRAPHY AND STRUCTURAL GEOLOGY, 1979 

120° 45' 120°42'30"

AMADOR COUNTY CALAVERAS COUNTY

Dacite Rhyodacite

Tm| I L Pliocene and 

Mehrten Formation J Miocene

I Miocene and 

Valley Springs Formation J Qligocene

lone (?) Formation 

Granitic rocks

Slate and Metaother meta- volcanic 

sedimentary rocks 

rocks

Gourley Reservoir

X GV76-30

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 U.S. Geological Survey, Jackson, Mokelumne Hill, San Andreas, and Valley 

Springs 7.5 minute quadrangle.
CONSTRAINTS ON MOVEMENTS, MELONES FAULT ZONE, SIERRA NEVADA FOOTHILLS J3

Mountain, about 15 km southeast of the map area 

(Stirton and Goeriz, 1942), and one of late Hemphillan age above the latite (Stirton and Goeriz, 1942, 

Wagner, 1976) indicating that the Mehrten is Miocene and Pliocene. Correlation of K-Ar dates with 

the North American mammalian chronology by 

Evernden, Savage, Curtis, and James (1964) indicates 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 (Dalrymple, 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

drainages)

MIDDLE 

ALTITUDES

(Stanislaus and

Tuolumne River

drainages)

Andesitic

sedimentary

rocks

Latite of 

Table Mountain

Andesitic

sedimentary

rocks

and

breccia

Upper tuff

r .o 10" o '

Lower tuff

lone Formation 

Basement racks

-Oakdale fauna 

(late Hemphillian)

-Two-Mile Bar fauna 

(early Clarendonian)

_fl 9. 9m.y.-- -------

\21.9m.y.

-22.8m.y.

Andesitic breccia

Latite 

(Dardanelle facies)

Latite of 

Table Mountain

o Andesitic 

| tuff, 

,£ breccia, 

J^ and 

£ sedimentary

•£ rocks 

§

^^^-^y^-^^-w^r^^^^^x^

Upper tuff

Lower tuff

^atiteweldedtuff

AtirTf-erourgravel

Basement rocks

9.0m.y.

9.2m.y.

-9.0m.y.

21.1m.y. 

23.1m.y.

;-33.2m.y.

FIGURE 2.—Correlation of some Tertiary stratigraphic units 

in the western Sierra Nevada, showing age control from 

K-Ar dating of volcanic rocks (from Dalrymple, 1964).

TABLE 1.—Chemical 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

Fe2 0 3

FeO*

MgO

CaO

Na2 0-

K2 Q —

Ti02

P 2 0 5

H2 0+* —— 

H2 0-* —— 

C02 * ——— 

S (total) 

Cr ———— 

Ni ————— 

Ba ———— 

Rb ————

Sr ———— 

V ————— 

Sc ———— 

Y —————

67.92

16.10

1.24

.52

.90

2.24

6.09

2.97

.33

.10

.032

.89

.21

.03

.001

55

26

960

83

1600

27

5

<5

59.97

15.63

5.00

< .1

4.46

5.07

4.24

.78

.81

.27

.076

1.05

.25

.02

.006

155

113

1250

64

1650

116

15

6

intrusive bodies, referred to as Hamby dome and 

McSorley dome, were classified as hornblende rhyodacite. 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 Formation. At McSorley and Hamby domes, the Calaveras 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 surrounded 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

[GV76-30 - Rhyodacite of Hamby dome; GV76-34 - Dacite of Golden Gate Hill. 

Potassium-measurements: J. H. Tillman. Argon measurements and age calculation: E. Sims. ''"K decay constants: X £ = 0.572 x 10~ 10 yr" 1 , X E .' = 

8.78 x 10~ 13 yr" 1 , Xg = 4.962 x 10" 10 yr- 1 . Abundance ratio: [*°K/K = 

l._167X 1Q-1* atom percent.]

Field No.

GV76-30

GV76-34

Mineral

Hornblende

Hornblende

K2 0

1.061

1.057

1.186

rad

(moles/g)

6.283X10" 12

6. 900X10" 12

"°Ar ,

iniAr . 

total

0.20

0.33

Calculated

(10 6 years)

4.12+0.11

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 Formation 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. old 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 described 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. inferred 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 equivalent in age to the upper part of the Mehrten Formation. 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 making 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 containing blocks of rhyodacite from Hamby dome is involved in the faulting, the latest movement must be 

more recent than 4 m.y., but no other time constraints can be placed on the movement at this site.

REFERENCES CITED

Dalrymple, G. B., 1963, Potassium-argon dates of some Cenozoic volcanic rocks of the Sierra Nevada, California: 

Geological Society of America Bulletin, v. 74, p. 379-390.

————1964, Cenozoic chronology of the Sierra Nevada, California; 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 mammalian 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, Calaveras County, California: 

Los Angeles, California University M.A. thesis, 56 p.

Piper, A. M., Gale, H. S., Thomas, H. E., and Robinson, T. W., 

1939, Geology and ground-water hydrology of the Mokelumne area, California: U.S. Geological Survey WaterSupply 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 Volcanology, 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 Ferry, California: California University Publications in Geological 

Science, v. 26, no. 5, p. 447-472.

Wagner, Hugh, 1976, A new species of Pliotaxidea (Mustelidae: Carnivora) from California: Journal of Paleontology, 

v. 50, p. 107-127.

irU.S. GOVERNMENT PRINTING OFFICE: 1980 O— 311-344/24