Iowa State University Capstones, Theses and Retrospective Theses and Dissertations Dissertations

1979 The geology of Sheep Canyon Quadrangle: Wyoming Robert Edward Ladd Iowa State University

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Recommended Citation Ladd, Robert Edward, "The eg ology of Sheep Canyon Quadrangle: Wyoming" (1979). Retrospective Theses and Dissertations. 16864. https://lib.dr.iastate.edu/rtd/16864

This Thesis is brought to you for free and open access by the Iowa State University Capstones, Theses and Dissertations at Iowa State University Digital Repository. It has been accepted for inclusion in Retrospective Theses and Dissertations by an authorized administrator of Iowa State University Digital Repository. For more information, please contact [email protected]. The geology of Sheep Canyon Quadrangle: Wyoming

by

Robert Edward Ladd

A Thesis Submitted to the

Graduate Faculty in Partial Fulfillment of

The Requirements for the Degree of

MASTER OF SCIENCE

Department: Earth Sciences Hajor: Geology

Signatures have been redacted for privacy

Iowa State University Ames, Iowa

1979 ii

TABLE OF CONTENTS

Page

INTRODUCTION 1

Methods of Study 1

Physiography and Location 2

Previous Work 6

MAP CORRELATION BETWEEN SHEEP CANYON AND GREYBULL NORTH QUADRANGLES 10

GEOLOGIC SETTING 15

STRATIGRAPHY 21

Mississippian. 24

Madison Formation 24

Mississippian-Pennsylvanian 26

Amsden Formation 26

Pennsylvanian 30

Tensleep Formation 30

Permian 32

Phosphoria Formation 32

Triassic 35

Dinwoody Formation 35 Red Peak Formation 37

Jurassic 40

Gypsum Springs Formation 40 Sundance Formation 41 Morrison Formation 47 iii

Page

Early Cretaceous 48

Cloverly Formation 48_ Sykes Mountain Formation 53 Thermopolis Formation 55 Muddy Sandstone Formation 57 Shell Creek Shale Formation 58 Mowry Shale Formation 60

Late Cretaceous 62

Frontier Formation 62 Cody Shale Formation 65 Mesaverde Formation 66 Meeteetse Formation 70 Lance Formation 72

Tertiary 75

Paleocene Fort Union Formation 75

Paleocene-Eocene 78

Late Paleocene-Eocene Willwood Formation 78

STRUCTURE 82

Structural Elements of Sheep Canyon Quadrangle 87

Sheep Mountain Anticline 87 Alkali Anticline 88 Synclines 88 Faults 89 Joint fractures 90

Structural History of the Sheep Mountain Area 90

Pre-Laramide tectonic activity 90 Laramide tectonic activity 91 Post Laramide tectonic activity 94

GEOMORPHOLOGY 96

Geomorphic Elements of Sheep Canyon Quadrangle 96

Topographic features 96 Terrace remnants 97 iv

Page

Colluvial deposits 101 Alluvial deposits 101 Sand deposits 102 Landslides 103

SUMMARY 105

ACKNOWLEDGMENTS 108

BIBLIOGRAPHY 109 v

LIST OF FIGURES

Page

Fig. 1 Location Map of Sheep Canyon Quadrangle 4

Fig. 2 Geologic Map of Sheep Canyon Quadrangle, 23 \vyoming

Fig. 3 Nomenclature chart for the Madison Formation 25

Fig. 4 Nomenclature chart for the Amsden and Tensleep Formations 28

Fig. 5 Nomenclature chart for the Phosphoria and Dinwoody Formations 33

Fig. 6 Nomenclature chart for the Red Peak and Gypsum Springs Formations 38

Fig. 7 Nomenclature chart for the Sundance Formation 43

Fig. 8 Photographs of the Sheep Canyon Area 45 a) Sundance crossbedding b) Anomalous dips and tombstone topography in the Fort Union Formation c) Sheep Canyon with Madison Limestone exposed d) Slumps on the margin of Emblem Beach

Fig. 9 Nomenclature chart for Cretaceous and Tertiary Formations 50

Fig. 10 Structural Location Map of Sheep Canyon Quadrangle, (Andrews, Pierce, and Eargle 1947) 84

Fig. 11 Structural Cross-Sections of Sheep Canyon Quadrangle 86

Fig. 12 Terrace Level Correlations 100 vi

LIST OF TABLES

Page

Table 1 List of plants common to vegetation zones in Sheep Canyon Quadrangle 7

Table 2 Boundary problems in Greybull North Quadrangle from north to south 12 1

INTRODUCTION

The Sheep Mountain area north of Greybull Wyoming has long been recognized as an excellent instructional area for field geology.

Barren or sparsely covered rock outcrops and magnificently developed folds and faults are characteristics of the northeastern Bighorn Basin that have made this corner of Wyoming a favorite study area for uni­ versity field camps as well as professional geologists and rock collectors. In 1958, Iowa State University began a geology field camp near Shell, Wyoming. Students mapped areas around Shell and at Sheep

Mountain near Greybull for practice in geologic field methods and for graded problems. These yearly assignments provided new field data but a need for more detailed, comprehensive mapping was felt to gain a better understanding of the geomorphology, stratigraphy, and structure of the area as well as to provide a key to student field problems. The

Geology of Sheep Canyon Quadrangle: Wyoming is the second such detailed investigation by graduate students at Iowa State University. The eventual goal is complete geologic map coverage of Sheep Mountain

Anticline.

Methods of Study

Field observations for this study were begun in June 1976 and were completed by the end of August that same year. Field checks and corre- 2

lations between Sheep Canyon and Greybull North Quadrangle were done during the summer of 1977.

A 7.5 minute U.S.G.S. topographic map of the study area was used as a base on which formational boundries were plotted. Aerial photo­ graphs provided an aid to mapping in areas where formational contacts were obscured by ground cover. Formation thicknesses were determined by stratigraphic sectioning procedures. Most formations in Sheep

Canyon Quadrangle were measured with tape and or Jacob's Staff but thick gently dipping units, were measured with a 100 foot chain gradu­ ated in 10 foot segments. The apparent thickness was then converted mathematically to true thickness.

Pebble counts were routinely performed on Pleistocene and Recent terrace deposits to determine their lithology and to use as an aid in their correlation. One square foot areas were randomly chosen and sampled for clasts one inch or greater in length. These clasts were then sorted as to lithology into piles and relative percents for each kind were calculated.

Physiography and Location

Sheep Canyon Quadrangle is located in the northeastern part of the

Bighorn Basin between 108 37' 30" and 108 15' west longitude and 44 30' and 44 37' 30" north latitude (Fig. 1). The Bighorn Basin itself is an asymmetrical structural depression typical of the basins within the

Northern Rocky Mountain Physiographic Province (Howard and Williams 1972).

The structural axis lies near the western margin of the basin and trends 3

4 5

northwest-southeast. Flanking the basin on three sides are Precambrian

cored anticlinal uplifts - the Pryor and Bighorns Mountains to the

northeast and east, the Owl Creek Mountains to the south, and the now

buried Washakie Range and the Beartooth Mountains to the west and

northwest. Although the basin is topographically open to the north

it is structurally closed by the Nye Bowler Lineament. Northwesterly

and southeasterly plunging anticlines and synclines rim the margins

of the basin. These small folds are developed in Paleozoic and

Mesozoic strata and are usually asymmetrical with their steeper flanks

facing the mountain fronts. Like other Laramide basins in the Rocky

Mountain region, the Bighorn Basin Mesozoic strata and structures were

covered by a thick fill of Tertiary sediments which are now partially

exhumed as a result of Pleistocene dissection.

Sheep Canyon, the major topographic feature in the area, is super-

imposed across Sheep Mountain and is eroded into the }fudison Limestone .

. The Burlington Northern Railroad passes through the canyon from Greybull

enroute to Lovell, Wyoming.

Road access into the area is excellent. Numerous dirt and gravel

roads designed for bentonite trucks traverse the west flank of Sheep

Mountain and prospect trails and power and gas line service roads criss­

cross the more remote portions. In addition, U.S. highway 310 to Lovell

and U.S. highways 14, 16 and 20 to Cody traversef the west and south s1de•

of the quadrangle respectively. 6

The Greybull area, as well as the entire Bighorn Basin, is arid.

Emblem, just west of the study area, had an average annual rainfall of

5.77 inches between 1951 and 1960 while Greybull had an average of 5.79 inches during the same period (U. S. Dept. of Commerce, 1965). The months of highest rainfall are April, May, and June. Temperatures vary considerably both during the day and over the year. The hottest month is July when average temperatures are 71°F. with occasional days of

100°F. January is usually the coldest month of the year with average temperatures of about 18°F. and with occasional "cold snaps" of -40°F.

(Fisser 1964).

Three vegetatio~ zones occur in the Greybull area. The foothills scrub zone occupies the highest elevations of Sheep Mountain especially where Tensleep Sandstone is exposed. The desert and basin zone covers most of the Bighorn Basin while the river bottom zone is restricted to a narrow band of trees and grasses growing along rivers and creeks.

Table 1 lists a few of the most common plants found in these zones

(Porter 1962, Fisser 1964, Hitchcock 1971, Preston 1976).

Previous Work

Oral and written descriptions by trappers, early travelers, and railroad surveys provided nearly all of the topographical and geological information that existed in the American West prior to 1856. In that year F. B. Meek and F. V. Hayden began publishing the results of their explorations along the Missouri and Platte Rivers. From these expe­ ditions came the first geologic maps of the Nebraska Territory (Meek and Table 1. List of plants common to vegetation zones in Sheep Canyon Quadrangle

Desert and Basin Zone River Bottom Zone

Prickly Pear Cactus •••••••••• Opuntia polycantha Lance leafed Rabbitbrush ••••••••. Chrysothamnus viscidiflorus cottonwood Populus augustifolia Shadscale •••••••.•••••••• Atrip'Iex confertifolia Populus sargentii Willow Salix amgdaloides Silverscale saltbush ••••••.•.• Atriplex argentea ••••••••••••••• Box Elder ••.•••••••••••••• Acer negundo Fireweed ..••••••...... ••..•..•.• Kochia scoparia Sandbar willow .•••••••.•.• Salix exigua Foxtail Barley •••.••.•..••.••••• Hordeum jubatum Cocklebur Xanthium chinense Foxtail ••••••••••••••.•••••• Alopecurus aegualis ••.•..•••.•• Sunflower •.•••.•• Helianthus petiolaris Needle grass •••.••••••••••••.•••••• Stipa comata -...J Rabbitbrush •• Chrysothanamnus nauseosus Giant ryegrass •••.••••••••... Stipa occidentalis Sagebrush •••••.•••••••••••• Artemisia tridentata Cat tail .•••...... •.. Typha latifolia

Foothills Scrub Zone Rocky Mountain Juniper .•.....•• Juniperus osteosperma Curlleaf Mountain Mahogany .•.. Cercocarpus ledifolius Prickly Pear Cactus .••...••••••.•• Opuntia polycantha Yucca (Spanish Bayonet) ••••.••••••• Yucca filamentosa 8

Hayden l856a, l856b, Hayden 1857, 1858) as well as the first strati­ graphic columns dealing with the region (Meek and Hayden 1861). The geologic observations of T. B. Comstock (1873) and O. St John (1883) in the Teton-Yellowstone area and Eldrige (1894) in the Bighorn Basin built on Meek and Hayden's earlier work.

In the Bighorn Basin, the earliest geological literature concerned the occurrence of vertebrate fossils, particularly those found in the

Eocene Wasatch Formation. In the 1880's and 1890's, E. D. Cope orga­ nized several paleontological expeditions to the Wasatch badlands of the Wind River and Bighorn Basins. J. L. Wortman, who led Cope's expeditions (Cope 1882), believed the Wasatch to be a lacustrine deposit, a hypothesis later refuted by Fisher (1906) and Loomis (1907) in favor of a fluvial origin. Sinclair and Granger (1911, 1912) and Granger

(1914) continued vertebrate studies in the Bighorn Wasatch. Brown (1935) excavated dinosaur remains from the Jurassic Morrison Formation and

Brainerd and Keyte (1927) and Branson (1939) studied Pennsylvanian invertebrate assemblages at Sheep Mountain.

Darton (1906) made the first comprehensive studies of stratigraphy and structure of the Bighorn area. With the discovery of oil at Greybull

Dome came important papers by Washburne (1908), Hintze (1915), and

Lupton (1916) concerning the stratigraphy of the field. Water resources of the Greybull area were evaluated by Fisher (1906), and coal occur­ rences were noted by Washburne (1909). Hewett and Lupton (1911) located and described anticlinal structures in the Bighorn Basin as an aid to 9

oil exploration. Lupton and Condit (1917) studied the occurrence of bedded gypsum in the Embar and Chugwater Formations near Sheep Mountain and Hewett (1917) attributed economic bentonite deposits in the region to the weathering of volcanic ash.

During the 1930's and 1940's several geologists contributed to development of the stratigraphic nomenclature and an understanding of the erosional history of the Bighorn Basin. J. H. Mackin's hypothesis of downcutting, capture, and stream migration during the Pleistocene and Recent formed the foundation for current geomorphic studies (Mackin

1935, 1936, 1937). Love (1939) divided the Chugwater Formation into several members and later defined the Gypsum Springs Formation. The

Bighorn Wasatch deposits were studied and assigned to the Willwood and

Tatman Formations by Van Houten (1944, 1948).

More recent stratigraphic investigations include those made by

Eicher (1960, 1962) and Paull (1962) on the Thermopolis-Shell Creek interval, Moberly (1960, 1962), regarding the Cloverly and Sykes

MOuntain Formations, and Robinove and Langford's (1963) concerning water quality. A study of the Willwood Formation was also made by Neasham and

Vondra (1972). Published and unpublished reports by students at Iowa

State University which proved useful in this study were written by

DeKoster (1960), Johnson, Garside, and Warner (1965), Stensland (1965)

Stone (1967), Warner (1968), Neasham (1967), and Kozimko (1977). 10

MAP CORRELATION BETWEEN SHEEP CANYON

AND GREYBULL NORTH QUADRANGLES

Along the boundary between Sheep Canyon and Greybull North

Quadrangles several discrepancies in formational or rock body contacts

exist. These problems are listed in table 2 by section number.

Boundaries between Sheep Canyon and Alkali Anticline Quadrangles

have been correlated through the work of Mr. Jim Carlson.

In section 25, T54N. R94W. Mr. Mike Kozimko (1977) was unable to

follow the axial traces of two converging synclines because of Qf and

Qal cover. Mr. Jim Carlson (personal comm. Iowa State Univ. 1978), mapping Spence Dome Quadrangle, has traced the axis of a syncline separating Sheep Mountain from Spence Dome and the Spence Dome axis

until they are buried under Qal in the southern portion of his study

area. By extending the trend of the exposed axial traces and noting

the extent of Sundance Formation (Js) exposed in Spence Dome Quadrangle

the buried formation positions and structural trends were estimated for

Sheep Canyon and Greybull North Quadrangles. Because of the limited

information available to Mr. Kozimko (1977) his interpretation does

not correspond to Mr. Carlson's and the writer's estimations for the

positions of the Sundance (Js), Morrison (Jm), Cloverly (Kcl) , Sykes

Mountain (Ksm) , Shell Creek (Ksc), the axis of Spence Dome, and the

syncline between Spence Dome and Sheep Mountain. 11

In section 1 and section 12, T53N. R94W, Kozimko (1977, page 40) apparently mapped the top of the Amsden Formation at a layer of cherty carbonate rather than at the base of Tensleep Sandstones (Mallory 1967).

Most of the Ranchester Limestone Member of the Amsden was therefore included within the Tens1eep Sandstone. This accounted for most of the discrepancies with Amsden inliers and Tensleep boundaries between the two maps.

In section 12 and 13, T53N. R94W. the Red Peak, Gypsum Springs,

Sundance, Morrison, Sykes Mountain, Thermopolis, and Muddy Formations do not match those in Sheep Canyon Quadrangle. Due to poor outcrop exposure, Kozimko estimated their positions. Good control in Sheep

Canyon Quadrangle later proved these estimates inaccurate. In addition, terrace level Qt5 in section 26, T53N. R94W. was predominantly covered. by colluvium (Qc) in the writers study area and extended slightly into section 25, T53N. R94W. creating an apparent mismatch. Several other possible terrace levels can be recognized in section 26 but were combined into Qt5. Kozimko apparently mislabeled terraces Qt4 in section 2 T52N. R94W. From the map, these terraces should be Qt2.

Correlation problems between Sheep Canyon and Greybull North

Quadrangles were corrected by the writer during the summer of 1977.

It was learned that in any future mapping, rock body boundaries should be extended into adjacent areas and not just up to the edge of the study area. Table 2. Boundary problems in Greybull North Quadrangle from north to south

SW 1/4, Section 25, T54N. R94W.

(1) The synclinal axis found in the SW 1/4, Section 25, T54N. R94W. should probably bend more west northwesterly to connect with a synclinal axis mapped by Mr. Jim Carlson in Spence Dome Quadrangle. The synclinal axis as mapped is roughly in the position estimated for the axis fo Spence Dome.

(2) The Jurassic Morrison and Cloverly formations probably bend around the synclinal axis and bend east-west beneath the Qal in section 25, T54N. R94W.

I-' Section 36, T54N. R94W. N

(1) NW1/4, section 36. The Morrison and Cloverly formations probably trend N.N.W. beneath Qal.

(2) NW1/4, section 36. The upper Dinwoody contact should be located further north.

Section 1, T53N. R94W.

(1) NW1/4, NW 1/4, section 1. The Tensleep inlyer should be enlarged and connect with Tensleep strata exposed on the crest of Sheep Mountain.

(2) NW 1/4, section 1. The trace of the axial plane of Sheep Mountain should be located further north. Table 2. Continued

Section 1, T53N. R94W. (continued)

(3) NW 1/4, section 1. The Tensleep basal contact was chosen too low and therefore occurs too far south on the map.

(4) NW 1/4, section 1. Adjoining Madison inliers should extend into Greybull North Quadrangle.

(5) SW 1/4, SW 1/4, section 1. An adjoining Amsden inlier should extend into Greybull North Quadrangle.

...... w Section 12, T53N. R94W.

(1) NW 1/4, section 12. Qf mapped in Greybull North Quadrangle was interpreted as Qal in Sheep Canyon Quadrangle.

(2) SW 1/4, section 12. The Trrp-Jgs contact is located too far north.

(3) SW 1/4, section 12. The Jgs-Js contact is located too far north.

Section 13, T53N. R94W.

(1) NW1/4, section 13. The Js-Jm contact is located too far north.

(2) SW 1/4, section 13. The Ksm-Kt contact is located slightly too far north.

(3) SW 1/4, section 13. The Kmd map width is too large. The Kmd-Kt contact is too far north. Table 2. Continued

Section 25, T53N. R94W.

(1) SW 1/4, section 25. Qc mapped in Sheep Canyon Quadrangle partially covers Qt5 at this point.

Section 2, T52N. R94W.

(1) W 1/2, SW 1/4, section 2. The Qt4 terrace should be labeled Qt2.

(2) SW 1/4, SW 1/4, section 2. The Qt4 terrace should be Qal or Qs. I-' .j:'-

Section 11 and Section 12, T52N. R94W.

(1) N 1/2, section 12, and N 1/2, section 11. The Qt4 terrace level correlates with Emblem Bench and be Qt • should labeled 2 15

GEOLOGIC SETTING

The Bighorn Basin lies on the Rocky Mountain Foreland or

Cordilleran shelf between the Idaho-Wyoming Thrust Belt on the west and the North American craton to the east. Broad asymmetrical basins containing Tertiary fill, marginal Precambrian cored uplifts trending

NW-SE, and a relatively thin drape (10,000 feet) of Paleozoic and

Mesozoic shelf sediments characterize the area.

The structural style is Laramide and developed as a response to arc-continent collision (Early Triassic) and later ocean continent collision above an eastward subducting oceanic plate (Burchfiel and

Davis 1975). A second subduction zone formed at the leading edge of the North American continent as the continent overrode the oceanic plate (Lowell 1974). As length of underthrusting continued, compression and bouyancy from the subducted oceanic slabs introduced deformation progressively continentward. This deformation eventually reached the foreland area by Late Cretaceous or Early Tertiary time. Bouyancy from the uncoupled oceanic slab may account for late Cenozoic uplift in the Rocky Mountain Foreland (Lowell 1974).

More specifically, the Rocky Mountain region was affected by two major orogenies during Late Jurassic through Eocene time. The Sevier

Orogeny, in areas marginal to the Cordilleran Platform, produced a series of decollement thrust faults from Late Jurassic-Early Cretaceous to Paleocene times (Kent 1972). The Sevier was followed by the Laramide 16

Orogeny in Late Cretaceous (Maestrichtian) to Middle Eocene time. The

Laramide was responsible for major mountain building in the Rocky

Mountains (Kent 1972).

Sediments of the North American craton from Late Precambrian

through the Phanerozoic are separated by six major unconformities that subdivide the rock record into six rock stratigraphic units. These

rock sequences represent major transgressive maxima toward the trans­

continental arch, from either the Cordilleran or Appalachian shelves,

followed by regression. Each major transgressive-regressive cycle is

separated by a hiatus and modified by erosion (vacuity). These

sequences, as proposed by Sloss (1963), are the Sauk (Late Precambrian­

Early Ordovician), Tippecanoe (Middle Ordovician-Early Devonian),

Kaskaskia (Early Devonian-Late Mississippian), Absaroka (Late

Mississippian-Early Jurassic), Zuni (Middle Jurassic-Late Cretaceous),

and Tejas (Late Paleocene-Recent). The deposits of the Rocky Mountain

foreland can therefore be placed in this framework.

A large depositional basin extending from northwestern Montana

and central Idaho, to Utah and Arizona collected sediments which form

the Belt Supergroup (Raun and Kent 1965). From this trough a Lower

Cambrian sea transgressed across an eroded Precambrian surface eastward

toward the Transcontinental Arch in western Iowa (Kent 1972). Lower

Cambrian deposits, predominately clastics derived from Precambrian

structural highs, prograded eastward across the continental shelf

(Kent 1972). Sauk deposition ceased with a final westward regression

in Early Ordovician time (Sloss 1963). 17

During Middle and Late Ordovician time the sea returned to the

Cordilleran shelf and brought on deposition of the Tippecanoe Sequence

(Haun and Kent 1965, Sloss 1963). The most widespread transgression

in the Tippecanoe Sequence occurred in Middle Silurian time in the

Rocky Mountain region (Kent 1972). The Williston Basin was forming in

Mdntana and preserved Silurian sediments which, except for isolated outliers on the Colorado-Wyoming border, have been removed during post

Silurian-Middle Devonian uplift (Haun and Kent 1965).

The Kaskaskia Sequence (Sloss 1963) represents renewed trans­

gression from Middle and Late Devonian through Mississippian time.

Devonian sediments thin eastward on the shelf with basal detrital deposits grading upward into shallow marine dolomites and shales

(Kent 1972). Maximum transgression occurred during Late Devonian

(Fammenian) time (Kent 1972). A smaller Late Devonian or Early

Mississippian transgression produced deposits limited to isolated basins and embayments (Kent 1972). A Kinderhookian through Osagean

transgression re-introduced regional carbonate deposition and formed

the widespread Madison Formation (Sando 1976). Madison deposition continued through the Meramecian as the sea regressed (Sando 1976).

Extensive karst developed on the emerging shelf in some areas during

the Meramecian (Sando 1976), before a Chester age transgressive­ regressive cycle completed Mississippian deposition (Kent 1972). In western and central Wyoming marine transgression, that began in

Chester time, continued into Morrowan producing continuous deposition from Late Mississippian to Early Pennsylvanian (Kent 1972). 18

Pennsylvanian-Permian deposits form the lower part of the Absaroka

Sequence in the Rocky Mountain region (Sloss 1963, Kent 1972). The

Ancestral Rocky Mountains (which attained elevations of 5,000 - 10,000 feet by Des Moines time, Kent 1972), and the Uncompahgre Uplift in

Colorado shed vast amounts of detritus onto the fairly stable Wyoming shelf until Middle Permian time (Kent 1972). A marine regression

(Missouri to Early Virgil time) limited deposition to shallow or re­ stricted marine basins (Kent 1972). Shelf tilting in northern Wyoming produced an irregular Lower Virgil to Early Wolfcampian erosion surface

(Kent 1972). A Middle Permian Cordilleran transgression deposited black shale and phosphate in the miogeocline and tongues of redbeds, carbo­ nates, and evaporites eastward over the shelf (Kent 1972).

The Upper Absaroka Sequence (Kent 1972) in the region represents

Triassic through Early Jurassic deposition in a broad marine trough which stretched from Canada to central Arizona. Fine grained redbeds are the most conspicuous Triassic deposit though westward they grade into carbonate and dark shale (Kent 1972). Triassic shallow marine conditions on the Wyoming shelf continued through Early Jurassic and the

Triassic-Jurassic boundary, where preserved, is probably gradational

(Kent 1972).

The Zuni Sequence (Middle Jurassic-Cretaceous) (Sloss 1963) re­ flects a major shift from west-east to north-south marine transgressions

(Kent 1972). A large depositional basin, extending from the Arctic to the Gulf of Mexico, was bounded on the east by the Canadian Shield and 19

Transcontinental Arch and on the west by the rising western Cordillera

(Kent 1972). The first and second Jurassic marine cycles deposited shales, siltstones, red beds, evaporites, and sandstones on truncated

Paleozoic and Traissic rocks (Kent 1972). The third Jurassic trans­ gression resulted in the Rierdon-Lower Sundance interval while the fourth and largest Jurassic transgression left Swift-Upper Sundance sediments dis conformably overlying the earlier deposits (Kent 1972).

The western Cordillera region became an important source for detritus during the Swift-Upper Sundance transgression and continued to supply sediment for the remainder of the Jurassic and Cretaceous (Kent 1972).

Quiet tectonic conditions on the Cordilleran shelf produced nearly continuous deposition from the Swift-Sundance into the nonmarine

Morrison Formation (Kent 1972).

The Cretaceous deposits of the Rocky Mountain region were deposited in a broad asymmetrically subsiding trough which received sediments from successive mountain uplifts in the western Cordillera as well as the midcontinent region (Kent 1972). The Jurassic-Cretaceous boundary lies at the top of the Morrison Formation in a variegated series of nonmarine shales, mudstones, and sandstones conformable in most cases with the

Cloverly Formation. The first Cretaceous transgression (Middle Albian) extended from the Arctic southward toward the Gulf of Mexico (Eicher

1962) and was followed by a series of transgressive-regressive cycles which lasted until Middle Maestrichtian time (Kent 1972). Basal trans­ gressive or marginal marine sandstones were overlain by thick sequences 20

of dark gray shale in each transgressive phase. By Late Cretaceous

time the seaway shallowed as sand was carried eastward from the

Cordilleran area in increasing amounts by rivers and streams (Kent

1972). By Late Maestrichtian fluvial sedimentation was controlled by

developing intermontane basins (Kent 1972). The Cretaceous-Paleocene

boundary is usually conformable except near rising mountain fronts.

Volcanic activity in central Idaho and western Montana contributed ash

to terrestrial and marine systems during this time (Kent 1972).

The Tejas Sequence (Tertiary-Quaternary)(Sloss 1963), represents

the final period of deposition in the Rocky Mountain region. Large

subsiding asymmetrical basins received sediments from surrounding

rapidly rising mountain ranges until Early Pliocene time. A broad

regional uplift occurred in Middle and Late Eocene as well as volcanic

activity which continued into Oligocene time in the Absaroka and

Yellowstone region (Love 1960).

The Late Pliocene was a time of broad regional uplift and changing

climate. Large amounts of sediments were stripped from the Rocky

Mountain basins and new drainage systems superimposed themselves over

buried Laramide folds as erosion progressed (Kent 1972). Five major

glacial periods separated by interglacials marked the Pleistocene in

the area (Kent 1972). Erosion surfaces on the mountains were dissected,

basin fill continued to be removed, and Yellowstone volcanism became

active at this time. 21

STRATIGRAPHY

Rocks of all systems except the Silurian are present in the study area. Over 12,500 feet of Mississippian through Eocene strata are exposed. In addition to these the Cambrian to Lower Mississippian interval is still buried and the bulk of Eocene deposits outcrop out­ side Sheep Mountain Quadrangle.

Based on measurements by Mr. Gary Bible about 3,000 feet of the

Flathead Sandstone, Gros Ventre, Gallatin, Bighorn, Jefferson-Three

Forks, and Madison formations are exposed near Pyramid Peak in the

Bighorn Mountains (personal comm. Gary Bible, Iowa State Univ. 1978).

Van Houten (1944) measured 2,500 feet of Eocene Willwood and 700 feet of Tatman Formation near the topographic center of the Bighorn Basin.

On the basis of these sources and measured sections taken during the course of this study the Paleozoic is estimated to be 3,450 feet thick, the Mesozoic 8,100 feet thick, and the Cenozoic about 5,950 feet thick in the Bighorn Basin.

In the following discussion formations exposed in the study area are described from notes and measurements taken in the field. Figure 2

is a geologic map and legend giving brief descriptions of the mapping units and showing their distribution in the study area. 22 Fig. 2 Geologic--'Map of Sheep Canyon Quadrangle, Wyoming Missing Page This page was lost from the original text and therefore was unable to be restored. 24

Mississippian

Madison Formation

Peale (1893) named a thick carbonate sequence in the Madison and

Bridger Hountains of Montana, the Madison Formation. N. H. Darton, mapping in the Bighorn Mountains, recognized the same interval of lime­ stone and carried the term Madison Formation into Northern Wyoming.

Since Peal and Darton, the Madison Formation has undergone a number of nomenclatural changes which have been summarized in figure 3.

The four members of the Madison Formation, the Lodgepole, Lower

Hission Canyon, Upper Mission Canyon, and Cave Member, are separated by distinct slope breaks and bands of vegetation on the near vertical walls of Sheep Canyon. On the crest and flanks of Sheep Hountain the

Madison Formation is exposed as a series of inliers surrounded by the

Amsden Formation. The total thickness of the Madison Formation in

Sheep Canyon is about 760 feet. Local variations occur due to an ero­ sional unconformity between the Madison and overlying Amsden Formation.

The Madison Limestone consists of a variety of carbonate lithologies some of which are dolomitized. Interbedded evenly layered, thin to medium bedded, olive-gray to yellowish-brown micrites, bio-micrites, oomicrites, oosparites, and biosparites predominate in the lower three

Madison members. Broken and abraded bryozoan, horned coral, and brachy­ pod remains are also common as well as some small intraformational faults. Carbonate lithologies in the Cave Member are similar to those below though bedding is not as persistent. Collier and Peal Darton Sloss and Denson and Richards Edmisten Sando Sando Kozimko This Report (1893) (1906) Cathcart Hamblin Morrissey (1955) (1961) (1967) (1968) (1977) (1922) (1942) (1952,1954)

Montano Bighorn Mtns. : Montano Montano Montano Bighorn Mtns. Bighorn Mtns. Regional Regional Sheep Mtn. Sheep Mtn. I

Saca- Bull Upper Cave Cave A Cave jawea Ridge Mission Madison Mor. Mor. Mor. MOr. Canyon Fm. Mor. MOr. Fm. Mission Upper Canyon Upper Upper Upper Upper Fm. §- c::t Upper I§- ;:, Mission Mission Mission Mission Mission ~ ~ ~ Mission B .... Canyon Canyon ~ Mbr. ~ Canyon ~ Canyon ~ Canyon ~ Canyon ~ N G G ~ Ii: U1 e:: e:: e:: Mor. e:: Mbr. e:: Mbr, e::: Mbr. Madison Madison e:: e:: Fm. Mbr. e:: Paine ~ ~ ~ ~ ~ ~ ~ ~ ~ \I) \I) .\1) Fm. Fm. \I) Fm. \I) ...... ~ .~ ..... Lower Lower Lower Lower Lower 1:3 1:3 ~ ~ ~ i5 1:3 1:3 ~ Lower ~ ~ Mission Mission ~ Mlsslon Mlsslon ~ ~ Mission C Mlsslon ~ ~ ~ ~ ~ ~ ~ ~ ~ Canyon Mbr. Canyon Canyon Canyon Canyon Canyon Fm. Mbr. Mbr. Mbr. Mbr. MOr. Lodge- pole Wood- Fm. hurst Lodge- Lodge Lodge Lodge Lodge Lodge Fm. pole 0 pole pole pole pole pole Fm. Mbr. Mbr. Mbr. Mbr. Mbr. Albr.

Fig. 3. Nomenclature chart for the Madison Formation 26

The Madison Limestone was originally thought to range from Late

Kinderhookian to Osagian in age (Weller et al. 1948), but Sando (1976) has demonstrated, on the basis of foraminifera that it extends through

Middle Meramecian time. During the Early Kinderhookian, a marine trough adjacent to the Antler Orogenic Belt in central Idaho began receiving land derived sediments from the west. The sea occupying the trough transgressed eastward across previously eroded Cambrian, Ordovician, and

Devonian strata until Late Kinderhookian time. This rapid transgression resulted in deposition of clastic carbonate upon the Cordilleran Shelf and thick carbonate mud upon the continental slope (Sando 1976, Wilson

1969). By Late Kinderhookian, the Madison sea attained its greatest areal extent. The shelf shifted eastward while progradation of shelf deposits over slope deposits began. Osage time saw marine regression and the deposition of dolomite and evaporite in restricted marine environments. Sando (1976) states that the Mississipian regression continued through early Meramecian time. Uplift and draining of the cratonic shelf followed, and the limestone units of the upper Madison were subjected to solution during the remainder of the Meramecian. As a result, an extensive karst topography developed on its upper surface.

Mississippian - Pennsylvanian

Amsden Formation

Darton (1904) originally named the Amsden Formation near Amsden

Creek in the Bighorn Mountains. He described it (Darton 1906) as a basal sandstone, unconformable on the Madison Formation, overlain by a 27

middle sequence of red shales, and followed by an upper series of shale, limestone, and sandstone. Fossils indicated a Mississippian to

Pennsylvanian age. Following Darton, Blackwelder (1918) named the basal Amsden sandstone the Darwin Sandstone, from Darwin Peak, Wyoming.

He included the red shale layer with the interbedded sandstones, shales, and carbonates. The three divisions of the Amsden by Darton and the two divisions by Blackwelder remain the most common subdivisions of the

Amsden Formation. Figure 4 summarizes the development of Amsden termi­ nology and illustrates the problem caused by introduction of the term

Sacajawea Formation by Branson (1937, 1939). Branson attempted to eliminate the term Amsden Formation by including all strata of Penn­ sylvanian age in the Tensleep Formation. He substituted the Sacajewea for the Mississippian portion of the Amsden, including the Darwin

Sandstone and the breccia and cave fillings in the Madison paleo-karst.

Branson's ideas never gained acceptance but resulted in confusion.

Sando et ala (1975) established an Amsden type section on Amsden Creek near Ranchester, Wyoming.

Amsden Formation outcrops in gullies on Sheep Mountain, where it forms lance shaped inliers, and in scarp slopes of Tensleep flatirons.

A thick cherty dolomite within the Amsden often forms resistant ledges and is a distinctive marker unit in the area. The Amsden Formation attains a thickness of 117 feet at Stucco of which lowest 21 feet is the Darwin Sandstone.

The Darwin Sandstone is a light gray, fine grained quartz arenite which is sometimes stained red by the overlying red siltstones. It is Sando Darton fBlackwelder Branson Agatston Gorman Mallory This Report (1904) I (1918) (1937,1939) (1954) (1963) (1967) (1975)

Bighorn Mtns. Regional Owl Creek Mtns. Regional Bighorn Mtns. Bighorn Mtns. Regional Sheep Mtn.

"

Tensleep Fm. Tensleep Fm. Tensleep Fm. Tensleep Fm. Tensleep Fm. Tensleep Fm.

c:: .~ c:: t:) Tensleep Fm. Tensleep Fm...... "" ~ c:: Upper c:: Clastic ~ Sandstone Ranchester Rancheste, Rancll • .,., N Zone OJ Shal. Limestone Limestone limestoll' Limestone Mbr. Mbr. Mbr. Middle Carbonate E E Zone E E IL Amsden Fm. IL IL IL ILE <: <: <: <: <: ., G) G) .. Han.shoe HOI.lshoe G) Horseshoe / '0 Red '0 Lower 'tJ '0 'tJ Shale Sha/. Sha/. / E'" Sho'. E'" C'astic E'" E'" E'" ct ct ct Mbr. ct Mbr. ct Mbr. / Zan eo / Sacajawea Fm. Amsden Fm. c:: .~ Darw;n Darwin Darw;" Darw;n ~ Sandstone .~ Sandstone Sandstone Sandstone (Darw;n) Sandston, Mbr. .~'" Member Mbr. Mb,. ( Sandstone) '" r-- ~ ~ ~ ----- Madison Fm.

Fig. 4. Nomenclature chart for the Amsden and Tensleep Formations 29

large-scale planar and trough crossbedded and is highly variable in thickness. The sandstone rests dis conformably on the Madison Limestone.

Above the Darwin Member are about 60 feet of homogeneous grayish- red siltstones known as the Horshoe Shale Member (Mallory 1967). Thin -----.----- pisolitic hydrated iron oxide layers outcrop within this siltstone in the Sheep Mountain area.

The Horseshoe Shale Member is overlain by a layer of cherty dolo- mite and a series of thin interbedded sandstones, siltstones, shales, and thin carbonates, known as the Ranchester Limestone Member. The cherty dolomite is resistant, about 8 feet thick, and forms a prominent ledge or dips lope with undulating irregular surfaces.

The Amsden Formation represents a transgressive marine sequence developed during Late Mississippian and Early Pennsylvanian time. During the Meremecian and Chesterian karst developed on the Madison Limestone. surface (Sando 1974). Streams transporting large quantities of sediment westward to the Cordilleran sea near the Idaho-Wyoming border deposited silt and sand into cavities and solution channels on/or within the

Madison Formation. Later, as a regional eastward extension of the

Cordilleran sea progressed, fluvial Darwin deposition gave way to re- stricted near shore lagoonal red siltstones of the Horseshoe Shale

Member. Continued eastward transgression during the Morrowan and Early

Atokan produced the deeper water carbonates and shales of the Ranchester

Limestone Member (Sando et al. 1975). 30

Pennsylvanian

Tensleep Formation

Darton (1904) named and described the Tensleep Sandstone in the

Bighorn Mountains near Tensleep Canyon. Darton determined the Tensleep to be Pennsylvanian in age and therefore equivalent to the Ninnelusa

Formation in the Black Hills. Brainerd and Keyte (1927) collected invertebrates from the Tensleep at Sheep Nountain and concluded that it was Pennsylvanian or Early Permian in age. The basal Tensleep con­ tact was lowered by Branson (1939) to include all strata of Pennsyl­ vanian age in the Bighorn, Owl Creek, and Wind River regions (Fig. 4).

Agatston (1954) concurred with Branson by using the present Ranchester

Limestone Nember of the Amsden as the basal unit of the Tensleep For­ mation. Since resolution of the Sacajawea problem (Sando et al. 1975) the basal contact has been redefined in Darton's original sense.

The Tens1eep Sandstone supports thin stands of mountain mahogany and juniper on the flanks and crest of Sheep Nountain. Its thickness varies regionally from 240 feet near the southern margin of the Bighorn

Basin to 95 feet at Sheep Canyon (this report), to 75 feet near the

Montana border (Agatston 1954). This northwesterly thinning as well as paleotopographic irregularities and thickness variations are due to a

Pennsylvanian erosion surface developed on the Tensleep Sandstone

(Agatston 1954, Lawson and Smith 1966, Tenney 1966, and Maughan and

Roberts 1967). 31

The contact of the Tensleep with the underlying Ranchester Lime­ stone Member of the Amsden Formation is transitional. Interbedded thin sandstones, shales, and carbonates give way to progressively thicker beds of quartz arenite. The first thick quartz arenite above the interbedded sequence was chosen as the basal contact of the Tensleep.

In the study area the contact occurs above a distinctive lavender to purple shale described by Kozimko (1977) and Rioux (1958). Above the basal contact the beds of quartz arenites thicken to a massive trough cross bedded sandstone. The uppermost 5 feet of the Tensleep consists of thin reddish to gray silty shale, siltstone, and very thin fossilif­ erous micrite layers.

Fusilinids collected from the Upper part of the Tensleep in Shell

Canyon indicate a Middle Des Moinesian age (Henbest 1954). However,

Mallory (1967), considers the Tensleep to be Des Moines or Late Atokan in age. During this time, uplift of the Ances~ral Rockies in northern

Colorado resulted in a regional shallowing of the Cordillerian epiric sea and deposition of the Tensleep Sandstone. -The source area for the sediments, however, is uncertain. Possible localities range from eastern Idaho and western Montana to northern Colorado and the Canadian

Shield (Agatston 1954, Tenney 1966, and Todd 1964 respectively). As tectonism increased, Tensleep deposits were elevated to sea level where wind and waves reworked the sand into dunes and bars. Interbedded carbonates and mudstone deposits may be records of minor subsidence or quiet water conditions (Mallory 1967). Finally as the Ancestral Rocky 32

Mountain Uplift climaxed in the Late Des Moinesian the Tensleep de­ posits were aerialy exposed. The resulting surface was exposed to erosion through the Permian Wolfcampian and contributed detritus east­ ward to a remnant sea in southeastern Wyoming (~fullory 1967).

Permian

Phosphoria Formation

The Phosphoria Formation was named by Richards and ~nsfield (1912)

from an outcrop of phosphatic shales, cherts, and limestones measured

in Phosphoria Gu1tch, Idaho. At that time, the Phosphoria was equated

to the upper two members of the Park City Formation described earlier

by Boutwell (1907) Embar Formation. A facies change from deepwater

phosphatic shales and limestones in Idaho through thick reef limestone,

transitional limestone, and red shales near Greybull, and shallow water

red shale and gypsum on the flanks of the Bighorn Mountains was recog­

nized as early as 1916 (Condit 1916). These Permian facies changes

have created a complicated terminology in Wyoming. Figure 5 attempts

to summarize the development of Phosphoria nomenclature relevant to the

study area.

Despite the precedent for using Darton's Embar Formation at Sheep

Mountain, it is apparent that the term is outdated. Since the Triassic

Dinwoody Formation occurs and can be easily differentiated in the study

area from Permian Phosphoria carbonates the term Goose Egg Formation

would not be appropriate here. Therefore the suggestion of McKelvey Richards and Condi1 (916) Pierce and Newell and Burk and McKelvey Dorton DeKoster This Report (1906) Mansfield Blackwelder Andrews Kummel 094" Thomas Rioux (958) (1959) (1968) (1912) (1918) (1941) Kummel (1954) (1956)

Big horn Mlns. Idaho Owl Creek Mlns. Bighorn Mlns. Regional S E Wyoming Sheep Min. Reglona I Sheep Min. Sheep Min.

E II- Chugwater Chugwaler Chugwater Chugwater Chugwater Chugwater Red Peak .. Red Peak ~ Fm. Fm. Fm. C Fm. Fm. Fm. Fm. .., Mbr. '-.., • I:) ~'"" u '-~ >- .., C 10loia Zone Dlnwoody Fm . 0 Dlnwoody Dlnwoody Dlnwoody Dinwoody 0 L i "9,,10 ZO"e- fm • Fm. fm. Fm. c• (Clataia Zone) - ~ - '--- :-~ w w

Embar Fm. Phosphoria Park City Embar Fm. Phosphoria Goose Egg Goose Egg Phosphoria Fm. Phosphoria Phosphoria ~ I:) Fm. Fm. Fm, Fm. Fm. wilh Fm. Fm. members '-~ '- ~ I

f-...

Tensleep Fm. Tensleep Fm. Tensleep Fm. Casper Fm. Tensleep Fm. L- ____

Fig. 5. Nomenclature chart for the Phosphoria and Dinwoody Formations 34

et ale (1959) has been followed and the Permian interval in the study area is referred to the Phosphoria Formation.

The Phosphoria Formation is 224 feet thick in the study area and is exposed along the flanks and crest of Sheep Hountain. Regional and local variations in thickness are attributed to deposition upon the irregular erosion surface developed on the Tensleep. Agatston (1954) found that the combined Phosphoria-Tensleep interval in the northeastern

Bighorn Basin produced a fairly even thickness, about 270 feet.

The basal contact of the Phosphoria Formation is locally sharp.

A distinctive red shale, regionally known as the "red shale marker"

(Tenney 1966) lies immediately above the unconformable Tensleep­

Phosphoria contact. This marker bed is about 60 feet thick and con­ sists of evenly bedded reddish-brown siltstone and shale overlain by

25 feet of gypsum. Above the gypsum layer are 25 feet of poorly ex­ posed interbedded gypsum, reddish-brown silty shale and siltstone. The final 155 feet of Phosphoria Formation is dominated by highly fractured, vuggy algal biolithite, micrite, and biosparite which often contains asphaltic residues.

The basal Phosphoria red siltstones are considered Leonardian to

Guada1upian in age (Campbell 1956) and represent the first trans­ gressional marine deposits laid down upon an extensive and irregular

Tensleep surface. By late Des Moines time the Ancestral Rocky Mountain

Uplift had reached its climax. Diastrophism waned in early Permian times, eventually ceasing to be an important controlling factor by the 35

late Wolfcampian (Mallory 1967). A low erosion surface separated a remnant sea in southeastern Wyoming from a new arctic sea located along the Idaho-Wyoming border (Agatston 1954). Though apparently shallow, causing the deposition of calcium phosphate along the Idaho­

Wyoming line (Bromley 1967), the Permian boreal sea spread eastward laying red bed and evaporite deposits by onlap onto the Tensleep surface. Extensive changes in the strand line position were caused by minor sea level fluctuations and may have contributed to the formation of sabkhas. Barriers to normal salt water circulation like algal reefs near wave base or extremely shallow water above a very gentle westward sloping sea floor probably were the cause of the thick gypsum and anhydrite beds seen along the western flank of the Bighorn

Mountains (Campbell 1956).

In the Bighorn-Sheep Mountain region, marine carbonates inter­ finger with shallow water redbed and gypsum deposits. In the study area, the lower red siltstone represents an initial transgression followed by deposition of thick gypsum and red beds in shallow restricted waters. Finally, deeper water algal reef deposits, brecciated by wave and current action, form the upper Phosphoria carbonates at Sheep

MOuntain.

Triassic

Dinwood~ Formation

The term Dinwoody Formation was first used in the Bighorn Basin by

Condit (1916) (Fig. 5) who divided Darton's Embar Formation into the Park 36

City and Dinwoody Formations at the suggestion of Blackwelder.

Blackwelder (1918) inferred a Permian or Early Triassic age for the

Dinwoody Formation and designated "the canyon of Dinwoody Lakes", in the Wind River Mountains as the type locality.

Newell and Kummel (1942) observed evidence of an unconformity between the Phosphoria and Dinwoody Formations. They also redefined the Dinwoody to include silty strata above the Phosphoria carbonates and below the red Chugwater siltstones. Newell and Kummel suggested that the f'ormation could be separated into three distinct zones. They recognized a basal siltstone, a dense fossiliferous limestone (the

Lingula Zone), and an upper series of olive, brown, or buff siltstones

(the Claraia Zone). Kummel (1954) correlated the Dinwoody Formation with other Triassic deposits in the western United States.

The Dinwoody Formation lies at the base of back slopes of flatirons upheld by the Phosphoria. It is well-exposed in cut banks along the

Bighorn River near Stucco and in cuts along the Burlington Northern

Railroad. The Dinwoody is only 54 feet thick near Stucco. Here the basal contact is sharply defined by a lithologic change from limestone to siltstone as well as a color change from light gray to bluish-gray.

Above the basal contact, thin to tabular beds of siltstone are inter­ bedded with thin irregular beds of gypsum. Thin gypsum veinlets criss­ cross the siltstone beds and form a boxwork structure. Siltstone layers become more resistant and thicker upwards in the formation.

In western Wyoming the Lower Dinwoody is Early Triassic in age

(Kummel1954). Newell and Kummel (1942) demonstrated that a regional 37

dis conformity occurs between the Dinwoody and Phosphoria Formations in

Wyoming, and that basal Dinwoody strata were progressively lost eastward by onlap onto the post Phosphoria surface. Eastward transgression from the Idaho-Wyoming border by the shallow Dinwoody sea caused migration of restricted marine environments. The Dinwoody, therefore, is younger to the east as the lowermost siltstone and the Lingula Zone are lost.

Only the Claraia Zone also of Early Triassic age is present in the study area (Newell and Kummel 1942, and Kummel 1954).

Red Peak Formation

Darton (1904) named the Chugwater Formation from exposures in the

~? Laramie Mountains near Chugwater Creek. As originally stated the Chug- ~ water included all strata between the Tensleep Sandstone and the Jurassic

Sundance Formation. Darton (1906) later defined the Embar Formation from Permian strata in the lower portion of the Chugwater. Lee (1927) named the Alcova Limestone. Love (1939) separated the Gypsum Springs

Formation from the upper part of the Chugwater and subdivided the rest of the Chugwater Formation into the Red Peak, Alcov~ Limestone, C~ow

MOuntain Sandstone, and Popo Agie Members from bottom to top. Members were raised to formational status by High and Picard (1967) and each in turn, with the exception of the Alcova Limestone was subdivided into informal units based on lithologic differences (Fig. 6).

The Chugwater Group in the study area is represented by only the

Red Peak Formation. The mudstones and siltstones which constitute the formation are nonresistant and typically form long strike valleys between resistant carbonates of the Phosphoria and Gypsum Springs Formations. Branson and High and Darton Darton Lee Love Love Rioux Kozimko (1977) (904) (1906) (1939) Branson (1957) Picard (958) (1927) (1941) (1967) This Report

Bighorn Mlns. Bighorn Mtns. Regiona I Western Wyo. Central Wyo. Reglona I Regional Sheep Min. Sheep Mtn.

Gypsum Gypsum Gypsum u ..... Springs Fm. Springs Fm. Springs fm . ... Gypsum ... Gypsum Gypsum ~ Springs ~ Springs' Fm. Springs Fm. ::. Mbr. "') ~ ~ ~ - u -~ ..... --Papa Agie ... Popo Agle Papa Agie Popo Agle ... Chugwater Fm. Red Beds Mbr . Fm. ~ E Mbr. Fm. IL. , w ~ "- .. ~ ~ Chugwater Fm. ....'II E -0 ... Crow Mtn. '"c Crow Mtn. ~ ~ Crow Mtn. Crow Mtn. ~ --.I co Sandstone Sandstone :J Sandstone :;; ... - .<:: Sandstone Mbr. Fm. u Fm. ;; ~ Mbr. ~ ., ~ b .<:'" .....U " ----- Limestone Alcova 1.S. Mb,. Alcova Dolomi,. U Alcova l.S. '" Alcovo 1.5. Mbr. ... ~ - ... - .....~ ~ Chugwater Fm. Red Peak Red Peak Red Peak Red Peak Red Peak Red Beds Mbr. Fm. Mbr. Chugwater Fm. ~... Fm. Fm . ~

Fig. 6. Nomenclature chart for the Red Peak and Gypsum Springs Formations 39

Usually, the upper portion of the Red Peak is exposed in steep protected scarp slopes beneath the basal gypsum layer of the Gypsum Springs

Formation.

A total of 570 feet of Red Peak sediments was measured near Stucco.

The basal contact is sharply defined by an olive-gray to reddish-brown color change between the siltstones of the Dinwoody and Red Peak for­ mations. Above the basal contact, Red Peak siltstones become more massive. On fresh fracture, yellow-gray or green mottles often appear

indicating reduction of iron. Induration varies from bed to bed in the

Red Peak Formation which accounts for its ledgy appearance. Thin gypsum veinlets, both parallel and criss-crossing bedding are common. Primary

structures include small scale trough crossbedding and asymmetrical

ripple marks. Generally, however, the finegrained well-sorted aspect

of Red Peak strata preclude easy identification of primary structures.

High and Picard (1967) divided the Red Peak Formation into the

silty claystone member, the lower and upper platy members, the alter­

nating platy member, and the sandy member. The silty claystone member

may have been deposited in a low energy fluvial or coastal plain envi­

ronment which gave way to deep water siltstones, sandstones, and shales

in the Chugwater Group (Picard 1967). A shift to quieter depositional

conditions in the lower and upper platy members may correspond to the

development of a tidal flat complex. The alternating member was

deposited in a shallow marine environment below tide base where fine

grained sediments were reworked and sorted by marine currents (Picard

1967). 40

Jurassic

Gypsum Springs Formation

A series of thick gypsum, red mudstone, and carbonate layers lo­ cated 18 miles southeast of Dubois, Wyoming was defined as the Gypsum

Springs Formation by Love (1939). Love thereby completed the final division of Darton's (1906) Embar Formation and defined a Jurassic formation that is continuous and lithologically similar throughout the

Wind River and Bighorn Basins (Fig. 6).

The Gypsum Springs Formation can be divided into three informal members in the Bighorn Basin (Peterson 1954, Riouoc 1958). The lowest member is a unit of thick massive gypsum and thin red shales which usually outcrop in the scarp faces of large hogbacks. A middle member of interbedded limestone, dolomite, and red shales forms the dip slope of hogbacks in the study area while the third member, a thick red shale, forms a distinct strike valley between hogbacks. The overlying

Sundance Formation is regionally disconformable with the upper red shale member, gradually truncating the Gypsum Springs Formation from north to south along the eastern Bighorn Basin (Imlay 1956).

At Stucco, 219 feet of Gypsum Springs strata were measured. The basal contact is abrupt and unconformable with underlying Red Peak siltstones. Immediately above the contact is a 60 foot bed of white gypsum intercallated with thin wavy red and green siltstone. Forty feet of reddish-brown siltstones complete the lower member. The middle member is composed of 50 feet of thin, vuggy biomicrite containing thin 41

laminations interbedded with layers of red and green siltstones and shales. These deposits are rhythmic in nature and contain at least three cycles of limestone, siltstone, and shale.

The upper member consists of 70 feet of a reddish-brown siltstone and some thin greenish-gray siltstone layers.

The age of the Gypsum Springs Formation was originally thought ~o be Late Triassic (Love 1939). Imlay (1945) however, proposed that its age was Bathonian (Middle Jurassic) based on abundant faunal remains.

During the Middle Jurassic; seas again transgressed eastward across the western continental interior (Downs 1952). The gypsum and red shale of the lower member of the Gypsum Springs Formation were deposited in restricted shallow marine basins extending from the Williston Basin south through much of Wyoming. Peterson (1954) attributes current restriction, necessary for the formation of evaporites, to the Belt

Island Uplift in Montana and isolated high points in Wyoming. Source areas to the west and southwest of Wyoming contributed clastic and limy sediments from exposed Paleozoic formations. As sea level rose in the

Middle Jurassic, evaporite deposition ceased. Cyclic limestone, silt­ stone, and shale deposits in the middle member reflect deeper water deposition and perhaps fluctuating sea levels. Peterson (1954) postu­ lated that a final regression in Gypsum Springs time deposited the upper red shale member.

Sundance Formation

Darton (1899, 1904) named the Sundance Formation for marine Jurassic sediments exposed near Sundance, Wyoming. The Sundance Formation extends 42

from South Dakota through Wyoming and into Montana where it is divided

into the Reirdon and Swift formation. Figure 7 summarizes the develop­ ment of Sundance terminology relevant to the Bighorn Basin. Neely

(1937) first subdivided the Sundance into fossil zones while Peterson

(1954) attempted to extend Montana terminology into the Bighorn Basin.

Imlay (1947, 1956) subdivided the Sundance into lower and upper members, which is the system followed in this report.

The lower and upper members in the Stucco area are 106 and 117 feet

thick respectively. Exposures are generally good with sandstone or limestone layers forming ridges and shales or siltstones forming flat

sparsely vegetated strike valleys. Sandstones and limestones of the

upper Sundance are usually lenticular and form discontinuous ridges.

A sharp color change from red to gray-green shale marks the dis­

conformable Gypsum Springs-Sundance contact (Imlay 1956). The first

resistant beds in the lower Sundance occur about 20 feet above the

basal greenish-gray glauconitic shale unit. This ridge forming unit

consists of 4 feet of crossbedded oolitic biosparite containing

angular shell fragments. This layer abruptly gives way to nearly 70

feet of calcareous glauconitic shale containing interbedded layers

packed with complete shells of Gryphaea nebrascensis. The uppermost

unit of the lower Sundance is a tan, highly lenticular, oolitic cal­

carenite. It is spectacularly crossbedded (Fig. 8) and locally

quarried for flagstone.

The upper part of the Sundance Formation is comprised of two

units. The lower consists of about 80 feet of calcareous gray-green Peterson Imlay Kozimko Darton Darton Crickmay Neely Cobban Imlay Downs (1977) (1954) (1956) (1899,1904) and O'Hara (1936) (1937) (1945) (l94n (1949) This Report (1909)

Black Hills Black Hills Regional Regional Montana Regional NW Wyoming Montana N W Wyoming Sheep Mtn.

Upper Sundance Upper Mbr. Upper Mbr. Mbr. Upper ... Sundance ~ Upper Upper Swift Port Swift Upper Sundance Fm. Lower Fm. Sundance Unit Q. Shale Mbr. <>" .po Middle ~ ~ ~ I Q. ~ ~ Sundance ~ ~ Mbr. ~ W u B~lem";'e ., .. .. ., .. .. <> Fm. I'ochyleu,his <>" <> ., <> <> (Belemnites r::: '-.. .. <> Zone <> r::: <> OenSUI ) ~ r::: r::: r::: I .. <> r::: (Belemnites <> <> \:) r::: r::: <> <> <> "I::J "I::J "I::J <> <> ~ .. "I::J "I::J r::: '- Middle "I::J r::: r::: Upper "I::J r::: r::: r::: r::: ~ II) ~ r::: Mbr. II) II) Red Bed II) Sandstone " I II) II) "-' " " I/)" " I/)" " " Unit Mbr.

I Medial Shale Lower Grypheo (G''fphea) Lower Rierdon I Mbr. Unit Part Zone Fm. Lower Basal (GrypheoJ Sundance Sundance ! Unit Rlerdon Oolitic Lower Lower Fm. Unit Mbr. Sundance '--

Fig. 7. Nomenclature chart for the Sundance Formation 44 Fig. 8 Photographs of the Sheep Canyon Area a) Sundance crossbedding b) Anomalous dips and tombstone topography in the Fort Union Formation c) Sheep Canyon with Madison Limestone exposed d) Slumps on the margin of Emblem Beach 45

c

u 46

shale containing layers of the belernnit~ Pachyteuthis densus near the base, a thin layer of O$t.t'ea engelmanni near the middle, and a thin layer of C$lptonectes bellistri'at'Us near the top. Above the shale a

35 foot series of int~rbedded calcareous gray-green sublitharenites and soft glauconitic shales form low discontinuous ridges. The sandstones are often very foss:iliferous and may grade latterally into arenaceous biocarbanates composed of packed, disarticulated, and broken bivalve shells. Astarte§acotensis and Trigonia sturgisensis have been reported from this interval (Imlay 1956).

The Late Jurassic Sundance Formation is Callovian to Early Oxford­ ian iA age (Imlay 1956) and represent~ a southeastward marine trans­ gression from the Williston Basin and Cordilleran Trough areas (Peterson

1954) Petersnn speculated that as the Sheridan Arch became prominent during the deposition of the lQWer Sundance, shallow water offshore bars (Stone 1967) and barrier bars of oolitic limestone, sandstone, and shale formed adjacent to it. Peterson (1954) postulated that as Early

Sundance seas retreated and deposition closed, a regional disconformity formed· between lower and upper Members of the Sundance. Stone (1967) however, found no field evidence for such an unconformity. The possi­ bil~ty then exists for continuous deposition through the Sundance.

Br:enner and Davies (1973) proposed that deposition of the upper Sundance, especially the upper interbedded sandstone-shale interval, represents concentrations of sand or macerated shell debris in lenticular current channels or wave surge channels generated by periodic storms. Deposits 47

such as these are of shallow marine origin and reflect the final with­ drawal of Jurassic seas (Peterson 1954, Imlay 1956).

Morrison Formation

The Morrison Formation, world famous for dinosaur remains collected by o. C. Marsh near Como Bluff, Wyoming and Morrison, Colorado, was named by G. H. Eldridge in 1896 (Emmons et a1. 1896). The dinosaur remains substantiated evidence for a terrestrial environment of deposi­ tion (Mook 1915) but lack of invertebrate fossils prevented an accurate determination of the age of the Horrison (Berry 1915, Stanton 1915,

Roth 1933, Branson 1935). The Morrison Formation was recognized in the

Bighorn Hountains by Darton (1906) and large sauroped dinosaurs were quarried near the mountains at the Howe Ranch north of Shell, Wyoming

(Brown 1935). Furer (1970), Mirsky (1961), and Moberly (1960) all attempted to clarify the Jurassic Morrison-Cretaceous Cloverly boundary in the Bighorn Basin.

The Morrison Formation usually forms featureless sage and grass covered flats except where erosion causes "badland" areas devoid of vegetation. Northwest of Stucco the Morrison attains a thickness of 296 feet. There the basal contact was sharply defined by a coquinoidal glauconitic limestone which formed an unconformable surface on which non-resistant light olive-gray siltstones of the Morrison Formation were deposited. Though Morrison lithologies vary unpredictably around Sheep

Mountain, a non-resistant clacareous olive-gray siltstone was the pre­ dominate Morrison rock type in the study area. The lower half of the formation is composed of siltstone while the upper portion is composed 48

of nearly 110 feet of variegated gray-green and dark gray silty shale and mudstone. A one foot thick nodular chert layer as well as occa­ sional thin limestone layers and isolated calcareous concretions with fibrous radiating structures are interbedded in the upper one half of the Morrison. Selenite crystals of secondary origin litter the sur­ face.

The Morrison Formation is considered to be Late Jurassic in age with deposition occurring between Oxfordian and Purbeckian time. Un­ certainty occurs from lack of fresh-water invertebrate remains (Moberly

1960). A transitional contact with the overlying Cloverly Formation indicates more or less continuous deposition into the Early Cretaceous.

The Sundance Sea had retreated northward by Late Jurassic time and a series of low energy, prograding river and floodplain sediments were deposited on the stable western interior (Furer 1970). Source areas for Morrison sediments were in Utah and northern Arizona (Furer

1970) while volcanic ash was added in increasing amounts from eruptions in California and Oregon. The Late Jurassic climate was probably warm with intermittent periods of drought aiding the precipitation of evapo­ rites and fresh water limestones in ephemeral lakes and swamps (Moberly

1960).

Early Cretaceous

Cloverly Formation

A series of brightly colored, variegated mudstones, siltstones, and buff sandstones were named the Cloverly Formation by Darton (1899, 49

1904). He considered the Cloverly Formation equivalent to the Lakota

Sandstone, the Fuson Formation, and the Dakota Sandstone of the Black

Hills, but left the stratigraphic description of the Cloverly general and obscure with respect to its lower and upper contacts. Hintze (1915) formally named the intermittently exposed Greybull Sandstone which was a producing horizon at Greybull Dome, near Basin, Wyoming. He specu­ lated that the Greybull Sandstone was the upper sandstone of Darton's

Cloverly Formation. In 1917, Hewett and Lupton raised the Greybull

Sandstone to member status and Hares (1917) named the Pryor Conglomerate.

The Pryor Conglomerate, however, pinches out laterally from the Pryor

Mountains and cannot be used everywhere as the basal Cloverly contact.

Lithologic similarities between mudstones of the upper part of the

Morrison and lower part of the Cloverly as well as similarities between the Greybull Sandstone and the overlying "Rusty Beds" (the Sykes Moun­ tain Formation) caused Rioux (1958) to map the interval as undifferen­ tiated. More recently, Moberly (1960), on the basis of petrographic studies, divided the Cloverly Formation into the Little Sheep Mudstone

Member and the Himes Member with the Pryor Conglomerate and the Greybull

Sandstone Members being occasionally present. Mirsky (1961) studied heavy mineral concentrations within the Morri$on and Cloverly Formations as did Furer (1970). Figure 9 traces the nomenclature of the Cloverly

Formation.

The Cloverly Formation, 152 feet thick northwest of Stucco, is exposed with the Morrison Formation in a narrow band of barren badlands Cobban and Eldridge Darton Fisher Washburne Hewefl Hintze Lupton Hewett and Rioux This 1894 (1906) (1904,1906)1 (1908,1909) (1914) (1915) (1916) 0917) Reeside (1958) Report Lupton (19520)

Bighorn Basin Bighorn Mln •. 1Bi'ghorn Ba.in E. Bighorn B. W. Bighorn B. Ba.in Wyo. BOlin Wyo. E. Bighorn B. Bighorn Ba.in E. Bighorn B. Sheep MIn. Sheep MIn. UJ[ ....

Va., Houten (114« • . ,., ~ Wasalch I Wasatch I Wasatch I Wasatch Wasatch Wasatch Willwood Fm. Willwood Willwood ~ :;. Fm. Fm. Fm. Fm. F m. Fm. Fm. Fm.: ~ ~ ... Fori Union Fort Union , Jepsen (li401 For' Union Fort Union 1';- Undlff. Fort Union Pol'eol B.neh F F ~. Fm. Fm. Fox Hills Fox Hills Ft. Union Fm. Fm. ?m. m.· ~ Fm. Fm. Lorami. Fm 110 Fm. 110 Fm. and Lance Lance Fm. Lance Fm. Lone. Fm. _;

Meeteetse Meeteetse Meeteetse M ~ ~~ etorpow Meeteetse 0: .. tnts, Laramie ., Fm . Shale Fm. Fm. Fm. Fm. ~n and Judith Rivlr ~ c; associated Fm. Undiff. " Mesaverde Montano I "Cl Laramie Laramie formations Fm. Mesaverde Mesaverde Mesaverde Muoverde Ciagoett Gebo Fm. ;> Fm. Fm. Fm. ., Fm. Fm. Fm. Fm. ~IIf ~ " f-r-----tl '" '" ~I I Eo;loFm. I~IEo;I' Fm g LI1 tI .! pierre ~ o ~ ~o-g r')" ~ a.. Shale _ , " " ('"\ - ~ ~I Pierre Pierre Pierre uppor c Cody Cody Cody Cody nCo dv Colorado IE Mbr. ~ Bo.in Shale Shale Shale Shale shale~'2 ... .e Shol, rrl Tn /1 :} ; Shale c z V J-.I-.iJ L.-i... •

~ ~ ~ ~ . . ,,~ 5S B li90$/2 Torchl. lit SS Torch"g~..!....§!Torchl. hi 5S Torel'll' ht S5 T~ c:: ("') ~ Mid~~~. Frontiu fm. Frontier Fm. Frontier Fm. ~rontjerFm. Frontier Fm "b o'!

~ ...... P'QL~~:I~

"Mow,/~ S~1I'"I Mowry Mowry o~ Mowry Mowry MOlllry MOW~ ~. Shal. Shale!. Shal. Shol. ~hQl. Sholl r-- Lower • 1\ II Cr II. SI! I S~,.. _ Mtr. 1:~~i~:t:i~~~3"~ d VC\" Thermopolis Sh. TIl"mopol" Fm i;:i TII,rmopoU' Fm. Thermopolis~ Thermopolis Thtrmopoll' Fm. - "" RII"ya,d," "Rultya,d," ;:; Syk .. MIn. r-m. Shale Fm. ~ ~ ,Gnl~\I\1.,/ '3,.yt\lI\~."'. ~r'1~ . G.(.~I\ ~ ~J~."'f..s f---l~ Ookota ~ ~ ~ Undiff. ~ ~H~Mb,.'~:17//1'~ Fm. Cloverly Cloy.,ly Cloy.,ly Cloyerly Cloverly Cleverly .. >- Lilli, ~ -;: ~ Cloverly .. Fm. Fm. Fm. Fm. Fm. l Fm. '" ~Sh"P; -< ~ S~(lt,Mbr.-< and ~ 10411"on, Mtr:." 0 ." Monilon ".. ~ F'g;aL? u Con~br.? "

Fig. 9. Nomenclature chart for Cretaceous and Tertiary Formations 51

west of Sheep Mountain. Here the transition between the Morrison sediments and the Cloverly Formation is transitional and not easy to pick. Moberly found that the Cloverly differed from the Morrison

Formation by a higher percentage of volcaniclastic sediment and a higher frequency of black chert grains and pebbles in the basal Cloverly.

Furer (1970) also pointed to a coarser percentage of detritus in the basal part of the Cloverly Formation. Mirsky (1961) observed a high zircon, low garnet ratio when compared to the Morrison sediments.

In the study area, bright variegated colors and a rubbly chert lag were characteristics of Cloverly sediments. The basal contact was chosen at a color change where variegated gray-green or gray-reds in

the Morrison gave way to brighter and deeper dusky-reds, grayish- reds, bluish-grays, and grayish-oranges in the Cloverly Formation (Rock

Color Chart Committee 1975). Surprisingly, it was also found that the

Morrison-Cloverly transition was characterized by a change from calcareous to non-calcareous mudstones.

Moberly (1960) recognized three members within the Cloverly For­ mation, the Pryor Conglomerate, which is a basal sandstone or conglom­ erate, the Little Sheep Mudstone, and the Himes Member, in which the

Greybull Sandstone is included. In the study area the Pryor Conglom­ erate is not present so the basal 110 feet of the Cloverly Formation are dusky-red, grayish-red, and dusky red mudstones and Silty fissil shales of the Little Sheep Mudstone Member. Brown carbonaceous fragments and

secondary selenite are common on the surface. A high expandable-clay

content is indicated by the cracked "popcorn" like surface texture of

this unit. 52

The upper 40 feet of the Cloverly in the study area correspond to

Moberly's Himes Member. Abundant dark brown carbonaceous material occurs in conjunction with pale orange to yellowish-brown sublitharenite and thin carbonaceous silty shales. Small oxidized iron concretions color much of this unit a rusty brown to orange. The sublitharenite beds are thick but lenticular and contain small scale crossbeds. The Grey­ bull Sandstone is a local channel sand which crops out just north of the study area, but pinches out laterally into shale and mudstone.

The Early Cretaceous Cloverly Formation is considered Neocomian by Moberly (1960) but Upper Aptian to Lower or Middle Albian by Ostrom

(1970). It represents a period of continued continental aggradation which began with the Morrison Formation in the Late Jurassic. Basal sandstones and the Pryor Conglomerate, present in some areas, are

thought to be deposits related to very early orogenic pulses of the embryonic Idaho-Wyoming thrust belt, but Furer (1970) pointed out that no crystalline rocks were then exposed to contribute feldspars to the

Cloverly sediments. Moberly (1960) hypothesized that Cloverly depo­ sition began with streams transporting an increased amount of wind

blown volcanic material from volcanoes in the Nevadian Orogeny. Even­

tually ash choked the river systems which formed vast swamps and

seasonal lakes where volcanic ash was altered to montmorillonite.

Evaporation of the swamps and lakes during dry seasons may have pro­

duced evaporites. Moberly (1960) also speculated that the highly variegated colors in the Cloverly represent ancient soil profiles. A

regeneration of drainage occurred at the end of Little Sheep Mudstone

, 53

time as increased tectonism to the west produced coarser sediments in the Himes Member Moberly (1960). The resulting sandstones and silty shales are probably derived from stream and floodplain deposits which, like the Greybull Sandstone in the Himes Member, flowed in a northeasterly direction. Cloverly deposition ceased with the first

transgression of the Cretaceous sea (Moberly 1960).

Sykes Mountain Formation

Moberly (1960) named the Sykes Mountain Formation for a distinctive rust colored series of Lower Cretaceous sediments resting on the Cloverly

Formation and covered by the Thermopolis Shale. Sykes Mountain sedi­ ments were previously included within the Benton Formation (Darton 1904),

the Colorado Formation (Darton 1906), and the Thermopolis Formation

(Lupton 1916). Some geologists, however, still prefer to use the in­

formal terms "rusty series" (Darton 1906), "rusty beds" (Washburne 1908), or the "Rusty Series member of the Thermopolis Formation" (Mirsky 1962) when referring to the Sykes Mountain interval.

Moberly's definition of the Sykes Mountain Formation states;

"The Sykes Mountain Formation includes the "Dakota

Silt", "Rusty Beds" and perhaps part of the Greybull

Sandstone and Cloverly of Mills and previous writers."

Moberly continued by describing the thin continuous nature of Sykes

Mountain sandstones as opposed to the lenticular nature of Cloverly

sand bodies. This point confuses the fact that Hintze (1915) named the

Greybull Sandstone for a discontinuous sand body in the Cloverly

Formation. Consistency in Moberly's descriptions would dictate that

the Greybull Sandstone be a Cloverly unit. 54

The Sykes Mountain Formation is 100 feet thick and well-exposed in some parts of the study area. Straight west of Stucco a prominent hogback is upheld by Sykes Mountain sandstones and shales. At this location the basal contact of the Sykes Mountain Formation rests above thin pale red and brown siltstones, shales, and discontinuous sandstones in the Cloverly Formation and lies below continuous reddish-brown calcareous quartz arenites of the lower Sykes Mountain Formation. Work in and around the study area by Moberly (1962) and by the writer has shown that where the Greybull Sandstone is present, the base of the

Sykes Mountain can be defined by the following points.

1. Dark brown to black iron spherulites are present in the

Sykes Mountain Formation and not in Cloverly sandstones.

2. Large scale crossbedding is generally well-developed in

Cloverly sands as compared to small scale crossbedding

in the Sykes Mountain.

3. Cloverly sandstones have lenticular geometries with

flat upper surfaces while Sykes Mountain sandstones are

tabular.

4. Cloverly sandstones pinch out laterally over relatively

short distances while the Sykes Mountain sandstones are

traceable over long distances.

Above the basal contact the Sykes Mountain Formation consists of a cyclical series of brown to reddish-brown, thin fine grained cal­

careous quartz arenites and yellowish-brown to black fissile shales. 55

The sandstones are iron stained, contain brown to black spherulitic concretions, and show both small scale asymmetrical ripple marks and worm trails on many bedding planes. Dark shales become thicker while sandstones diminish upwards in the Formation. The Sykes Mountain beds are transitional with the overlying Thermopolis Shale.

The Middle Albian Sykes Mountain Formation (Ostrom 1970) repre­ sents deposition from the first marine transgression of the Early

Cretaceous sea (Eicher 1962). A number of primary structures such as disturbed laminations, worm burrows, and small scale ripple marks indi­

cate typical tidal flat conditions (Noberly 1960) while thin persistent

sandstones are more indicative of a quiet shallow sublittoral deposit

(Eicher 1962). Eicher postulated that much of the carbonaceous material was reworked from coastal swamps and estuaries and deposited on the sea

floor. Reducing conditions produced by the decaying organic remains

promoted the formation of spherulitic iron concretions common to the

Sykes Mountain. Continued transgression resulted in increased shale

deposition as sand deposition waned.

Thermopolis Formation

Darton (1904, 1906) was the first to describe Thermopolis sediments

in any detail, placing them in a thick unit called the Benton (Darton

1904) or Colorado Formation (Darton 1906). Lupton (1916), however,

subdivided the Colorado Formation into the Thermopolis, Mowry, and

Frontier formations. His definition of the Thermopolis included the

"rusty beds" or Sykes Mountain at the base and the Thermopolis, Muddy,

and Shell Creek formations as defined today. Lupton's terminology was 56

officially adopted by the U.S.G.S. but numerous revisions of his original Thermopolis Formation have since occurred. The term

Thermopolis Formation is now restricted to the dark shales occurring between the Sykes Mountain Formation and the Muddy Sandstone. Figure

10 traces the history of nomenclature applied to the Cretaceous strata in the Sheep Mountain area.

The barren, non-resistent Thermopolis Formation is 194 feet thick in the study area. Its basal contact is transitional with the under­ lying Sykes Mountain Formation, but can be picked at the base of the first dark gray or black shale which contains dahllite concretions.

These calcium phosphate concretions (McConnell 1935) are universally present in the basal Thermopolis sediments in the study area.

Above the basal contact the Thermopolis Formation consists of soft fissile brownish-black shale with interbedded thin fine grained yellowish-brown quartz arenites. The quartz arenites diminish in thickness upward in the formation, gradually being replaced by silty shales or siltstones. A few thin yellowish-green bentonite layers as well as thin beds of light gray sublitharenite occur in a transition zone with the overlying Muddy Sandstone.

The Middle Albian Thermopolis Formation (Eicher 1962, Ostrom 1970) was formed in a reducing deep marine environment associated with the first Cretaceous transgression (Eicher 1962). Foraminiferal collections studied by Eicher indicate a southerly transgression for the Early

Cretaceous sea. Eventually this boreal seaway connected with warm Gulf waters and foraminiferal faunas were mixed. Weak bottom currents, which 57

promoted reducing conditions and unfavorable salinites during the

deposition of the Thermopolis, diminished the foram populations

(Eicher 1962).

Muddy Sandstone Formation

Darton (1906) and Fisher (1906) both noted the occurrence of a

light colored sandstone in the Colorado Formation. Later when oil was

produced from Greybull Dome, drillers named this unit the IIHuddy Sand"

(Hintze 1915). Lupton (1916) included the Muddy Sandstone as an in­

formal member within the Thermopolis Formation. Because of the eco­ nomic importance of the Muddy Sandstone and its distinct lithology.

Eicher (1960, 1962) and Paull (1962) recognized the Muddy Sandstone

as a formation and proposed the N. 1/2, section 26, T53N. R93\.J east

of Sheep Mountain as a type section.

The Muddy Sandstone is 20 feet thick in the study area where it

forms a thin white band between the black Thermopolis and Shell Creek

shales. The basal contact is gradational with the Thermopolis For­ mation and is chosen where light gray bentonitic sublitharenite pre­

dominates over black fissile shale in an interbedded sequence. The main body of Muddy Sandstone is a massive, fine to medium grained

sublitharenite with a light gray bentonitic matrix. Some fine carbo­ naceous material is present within the matrix but crossbedding and other primary structures are not well-developed. Elsewhere around

Sheep Mountain, black chert pebbles, well-developed small scale trough

crossbedding, and vertebrate fossils have been found in the Muddy. 58

The Muddy Sandstone is diachronous. Its equivalent, the Newcastle

Formation, in the Black Hills, is Middle Albian (Skolnick 1958) while the Muddy in central Wyoming is late Albian in age (Paull 1962). This is consistent with the concept that the Muddy Sandstone and its equivalents were deposited along the shoreline of a regressing early

Cretaceous sea. Large deltas in central South Dakota prograded west­ ~~ ward as black shale of the ~ll Creek and Thermopolis Formations were deposited in western South Dakota and eastern Wyoming. Westward tilting ~~ of the early Cretaceous basin resulted in fluvial erosion of the Skull

Creek Shales and a westward migration of the shoreline (Stone 1972).

Volcanic ash, which weathered to bentonite, was supplied from volcanic centers in Idaho while sand for the large spreading deltas was carried by streams from eroded Lower Cretaceous sandstones in the Dakotas and

Nebraska (Stone 1972, Mitchell 1976). Delta deposits, like those which eventually reached the Sheep Mountain area (Paull 1962), were subse- quently reworked by longshore and tidal currents in environments which included fluvial deltaic, lagoon, estuary, swamp, marine bay, and barrier bar (Waring 1976, Mitchell 1976, Berg 1976).

Shell Creek Shale Formation

The Shell Creek Shale Formation was originally proposed by Eicher

(1960) for the black shale interval between the Muddy Sandstone and the

Mowry Formations. Previously the Shell Creek deposits had been assigned to the Colorado Shale (Darton 1908), the Thermopolis Shale (Lupton

1916), the Mowry Shale (Burk 1957), and the Upper Thermopolis Shale 59

(Mills 1956) to mention a few (Fig. 9). Eicher designated the type section of the Shell Creek Shale at N. 1/2, section 14, T53N R94w in

Sheep Canyon Quadrangle.

The Shell Creek Shale is 229 feet thick in the study area. It is soft non-resistant, barren of vegetation, and in combination with the Thermopolis and Muddy Formation, forms broad nearly flat strike valleys between Sykes Mountain and Mowry hogbacks. A sharp color change between light gray bentonitic sublitharenits interbedded with dark gray fissile shale occurs at the transition between the Shell

Creek and Muddy Sandstone Formations. The basal contact of the Shell

Creek Shale is chosen at the base of the first thick dark gray shale above the light to dark color change. Above the basal contact the

Shell Creek Shale becomes a continuous series of soft, dark gray fissile shales. Some thin layers of rusty weathering ferruginous shale occur near the base along with thin layers of greenish-yellow bentonite. The bentonite layers increase in number and thickness toward the top of the

Shell Creek Shale.

The Shell Creek Shale is Albian in age and represents a return ,to deposition in a marine environment similar to the Upper Thermopolis

Shale (Eicher 1960). Volcanic ash from sources in Idaho continued to fall while the Shell Creek sea transgressed rapidly over the Muddy

Sandstone (Paull 1962). Study of foraminifera by Eicher (1960) indi­ cates that the sea was never connected with the Gulf Coast nor was it as deep as it was during the deposition of the Thermopolis. The Shell

Creek foraminifera indicate a cold water environment (Eicher 1960). 60

MOwry Shale Formation

The "Mowrie Beds" were named by Darton (1904) for a series of hard, light gray shales in the Benton Formation. Darton (1905) located a type "Mowrie" at Mowrie Creek near Buffalo, Wyoming and later Lupton

(1916) raised the Mowry Shale to formational status (Fig. 9). With the separation of the Muddy Sandstone from the Thermopolis Formation the basal Mowry contact was lowered to include a soft black shale sequence between the Mowry Shale and Muddy Sandstone (Cobban and Reeside

1952a) (Rioux 1958). With the establishment of the Shell Creek Shale

(Eicher 1960, 1962) the basal Mowry contact, as defined by Darton, was reestablished. Hewett (1917) attributed economic deposits of bentonite in the Mowry to the diagenesis of volcanic ash. ~ubey (1928) proposed that desilicification of volcanic ash caused the extreme induration of the Mowry Formation but Davis (1970) showed that the silica content of the Mowry shales may be due, in part, to a primary incorporation of opaline quartz derived from radiolarian tests.

The MOwry Formation is nearly 400 feet thick in the study area and forms steep hogbacks above the Shell Creek Shale. The basal contact is gradational with the underlying Shell Creek Shale, but a transition from fissile black shale to well-indurated gray platy shales character­ izes the boundary. In some instances cone-in-cone concretions, fossilized fish scales, subtle vegetation bands, and a marked acclivity of slope are good indicators of the Mowry Shale. 61

Above the basal contact the lower 300 feet of Mowry consists of an even bedded uniform series of well-indurated platy gray shales.

Fish spines and scales diminish in concentration upward while thin olive-gray bentonites thicken and become more abundant. Two medium­ bedded dark-gray well-indurated litharenites containing small scale crossbeds and asymmetrical ripple marks on their surface form the crest of Mowry hogbacks in the area. Immediately overlying them, but usually

exposed on the dipslope of the hogback, is an 18 inch thick bentonite

layer. The upper one-fourth of the Mowry consists of dark gray

siliceous shale interbedded with thin light-gray sublitharenite. Shale

induration decreases while the proportion of litharenite layers in­

creases upward in the final 80 feet.

At several locations in the study area, clastic dikes composed of

fine to medium grained sublitharenite intrude the Mowry Formation. They

are commonly slickensided planar bodies which generally taper downward.

Most clastic dikes in the Sheep Mountain area terminate before reaching

the basal Mowry, though some penetrate the Shell Creek Shale. One dike

was observed by Kozimko (1977) to cut the Muddy Sandstone. Smith

(1952) and Lemish (1958) proposed that sand, infilling fractures from

above, could form such dikes. Warner (1968) postulated that a forceable

injection of sand from the Frontier Formation could have filled joints

developed along principal stress directions associated with Sheep

Mountain Anticline.

The Upper Albian Mowry Formation (Reeside and Cobban 1960) was

deposited under deepening water conditions associated with an eastward 62

and southward transgression of the early Cretaceous boreal sea (Eicher

1962). Volcanic ash was contributed from centers in western Idaho while clastic sediments which were brought in from sources northwest of Wyoming, were disseminated by southerly marine currents. The bottom environment was probably reducing over prolonged periods of time, but above the substrata, conditions approached normal marine and supported an abundant fauna of fish and radiolaria (Davis 1970). Eicher (1962) attributed a combination of slow sedimentation with devitrification of volcanic ash as a cause for the highly indurated nature of the Mowry

Shale. Davis (1970) showed that the Mowry Shales could not all have been hardened by the excess silica derived from volcanic ash and that opaline quartz from radiolaria tests, which were common in the Mowry sediments may have contributed to their overall hardness. Interfingering dark shales and thin sandstones in the upper Mowry heralded a regression of the early Cretaceous sea.

Late Cretaceous

Frontier Formation

The Frontier Formation was named by Knight (1902) for a coal bearing series of sandstone, clay, and shale near Frontier Wyoming.

Equivalent strata at Greybull and Torchlight Domes near Basin Wyoming were labeled lIS andstone A" and lIS andstone B" of the Colorado Formation by Washburne (1908) and the Peay and Torchlight Sandstones by Hintze

(1915). Lupton (1916), however, was the first to apply the term Frontier

Formation to strata including the Peay and Torchlight Sandstones in the 63

Bighorn Basin (Fig. 10). Recent work on the boundaries and petrology of the Frontier Formation has been done by Cobban and Reeside (1952b),

Goodell (1962), Stensland (1965), and Siemers (1975).

The Frontier Formation is nearly 200 feet thick in the study area.

The Peay Sandstone forms a prominent hogback innnediately followed by a flat sage covered shale series. The Torchlight Sandstone is a friable white sandstone which forms low barren outcrops subject to wind deflation. In the northern part of the study area small sage anchored dunes are formed from sand derived from the Torchlight Sandstone. A vegetation band of saltsage, shadscale, grass, and other low plants grows at the boundary between barren Mowry Shales and the Frontier

Formation. The basal contact is transitional, with thin light gray litharenite and sublitharenites predominating over interbedded dark gray fissile shale. The sandstones thicken upward becoming massive in the Peay Sandstone. Dark brown calcareous sandstone concretions preferentially erode from the Peay Sandstone and give it distinctive serrated appearance. Large scale trough crossbedding is emphasized by the concretions but individual forests can be traced into the enveloping non-calcareous non-resistant sandstone.

An 8 foot thick bentonite layer known as the Stucco Bentonite overlies the Peay Sandstone and is extensively mined in the study area.

It provides an excellent marker bed and occurs at the boundary between resistant litharenite and sublitharenite of the Peay Sandstone and 220 feet of overlying soft dark gray shale and siltstone. Thin impure bentonite layers and thin friable litharenite occur intermittently 64

throughout the middle unit. The uppermost unit of the Frontier

Formation is informally called the Torchlight Sandstone and consists of 75 feet of fine to medium grained, light gray litharenite. Inter­ bedded sandstone and shale in the lower part of the Torchlight grades upward into massive poorly indurated sandstone containing small scale trough crossbeds at the bottom and large scale trough crossbeds at the top. In addition isolated roller shaped cobbles of andesite and disc

~haped cobbles and pebbles of chert occur interspersed throughout the uppermost Torchlight. A poorly cemented polymictic conglomerate of andesite and black chert pebbles and cobbles caps the Torchlight Sand­ stone. Erosion and deflation of the Torchlight produces a character­ istic lag of andesite and black chert pebbles and cobbles. This lag is a diagnostic outcrop feature of the Torchlight Sandstone.

The Frontier Formation~ like all formations, is diachronous across

Wyoming. The Peay Sandstone may be as old as early Cenomanian while the Torchlight may be middle Cenomainian in eastern \~oming to Turonian or Lower Coniacian in the west (Merewether et al. 1975).

Frontier deposition began in early Late Cretaceous time when the

Mowry Sea shallowed. Coarse detritus from uplifted and eroded sources in the western Cordillera accumulated in marine bar and tidal flat environments and diluted the contribution of volcanic ash from volcanoes in Idaho (Masters 1952, Van Houten 1962, Siemers 1975). Beach and bar deposits formed the Peay Sandstone as the sea regressed but increased sedimentation produced an eastward prograding delta lobe (Van Houten

1962, Siemers 1975). Sediment supply waned and lagoons, embayments, 65

and rapidly fluctuating shoreline beach deposits characterized the final stage of Frontier deposition. Subsidence of the clastic wedge produced a westward transgression and brought about Cody Shale deposition.

Cody Shale Formation

The thick basal Upper Cretaceous shales of the Bighorn Basin were originally assigned to the Pierre Shale by Darton (1904, 1906). Lupton

(1916) renamed them the Cody Shale for exposures near Cody and Basin

Wyoming. Little other work has been published on the Cody Shale except for fau~al studies by Haas (1949), Cobban (1951), and Fox (1954).

The Cody Shale in the study area is barren, non-resistant, and forn~ a broad strike valley between ridges upheld by the Frontier and

Mesaverde formations. Concretionary horizons support isolated colonies of prickly pear cactus or sagebrush. Good exposures of Cody occur along the south nose of Alkali Anticline and along Lovell Draw where approx­ imately 2,400 feet of Cody strata were measured.

The Frontier-Cody contact is sharp. Dark gray easily erodable

Cody shale rests directly upon the Torchlight Sandstone of the Frontier

Formation. Where the Torchlight is non-resistant and unexposed, a lag of black chert pebbles and andesite cobbles marks the contact.

The main body of the Cody Shale can be divided into three gradational units on the basis of erosional and lithologic differences in the study area. The lower 700 feet is composed of soft dark gray to black fissile shale intercallated with some thin impure bentonite layers. Ten to twelve feet above the basal contact is a characteristic zone of calcar- 66

eous septarian concretions. These pale-brown bodies constitute a diagnostic horizon in the basal Cody and have yielded specimens of

Inoceramus and large coiled arnmonities.

The middle unit, about 330 feet thick, is composed of thin hard layers of calcareous mudstone, calcareous concretions, and calcareous light gray shale with interbedded bentonite layers. The increased induration of this unit generally forms a low hill or rise within the non-resistant body of Cody Shale.

The upper unit of the Cody Shale is composed of uniform fissile gray shale and occasional thin bentonite layers. The upper 70 feet of this 1,350 foot unit is interbedded with calcareous shale, thin layers of olive-gray mudstone, and thin light gray sublitharenite. A few isolated concretions, like those previously described, exist below the

Cody-Mesaverde contact.

The Cody Shale is diachronous from east to west in Wyoming. A

Middle Coniacian through Santonian age is assigned to the Cody in the

Bighorn Basin (Merewether et ale 1975). The Cody along with its equivalent, the Pierre Shale of Colorado and South Dakota, represents a uniform sequence of shales deposited during a stable stand of the

Cretaceous Lewis Sea (Love 1960).

Mesaverde Formation

The Mesaverde Formation was named by Holmes (1877) for sandstone and shale exposures near Mesa Verde, Colorado. Washburne (1909) ex­ tended the Eagle, Claggett, Judith River, and Bearpaw Shale Formations 67

from Montana to the east flank (Fig. 10) of the Bighorn Basin on the basis of correlations by Fisher (1908) (Fig. 9). Hewett and Lupton

(1917) were the first to extend the term Mesaverde into the Bighorn region. The U.S.G.S. has retained Mesaverde as a formational name in the area (Pierce and Andrews 1941, and Gill and Cobban 1973) but equiv­ alent deposits across the border in Montana are divided into the Eagle

Sandstone (Wagge 1975 and Randall 1961), Claggett Formation (Hatcher and Stanton 1903), Judith River Formation (Meek and Hayden l856b) and

Bearpaw Shale (Hatcher and Stanton 1903). Washburne (1909), Fisher

(1908), Severn (1961), Miller et al. (1965), and Mackenzie (1975) have attempted to extend the Montana terminology into the Bighorn Basin.

The Mesaverde Formation is 1,350 feet thick in the study area.

Here thick or massive sandstones stand out as sharp low ridges while interbedded sandstone and shale intervals, form low grass and sage covered plains.

Severn (1961) states that the Eagle Sandstone, Claggett,Shale,

Judith River, and Bearpaw Shale Formations can be recognized in good outcrops in the eastern portion of the Bighorn Basin. It was found that in the study area the Bearpaw Shale equivalent was missing and that the Cody Formation-Eagle equivalent and the Eagle-Claggett equivalents interfingered with one another and proved difficult to subdivide

(personal comma Dave Prose, Iowa State Univ. 1978).

The Mesaverde Formation (Eagle equivalent) rests conformably upon the Cody Shale and consists of thin light olive-gray to yellow-brown sublitharenites evenly interbedded with fissile dark gray calcareous 68 shale. The basal contact of the Mesaverde Formation occurs at the base of the first thick sandstone in this interbedded series which overlies the gray shales in the Cody Formation.

In the study area 250 feet of calcareous sandstone and interbedded shale equivalent to the Eagle are exposed forming a low ridge. Sub-

litharenite layers are 8 to 10 inches thick and may contain small rounded chert pebbles. Interbedded calcareous shales increase in

thickness upward as sandstone layers diminish.

The following 650 feet is a largely covered sequence of soft,

fissile calcareous light gray shale and thin discontinuous sublitha-

renite layers. This probably corresponds to the Claggett Shale and

extends upward to the base of massive sandstones belonging to the

Judith River member.

The upper 340 feet of Mesaverde sediments in the study area is , correlative with the Judith River and Teapot Sandstone members (per-

sonal comm. Dave Prose, Iowa State Univ. 1978). Medium brown to light

gray fine grained crossbedded calcareous arenite and sublitharenite, of

variable induration form beds ranging from 3 inches to several feet in

thickness. Olive-gray shales occur interbedded with the middle of the

sandstone and forms erosional depressions separating the sand bodies.

Large calcareous concretions as well as small grayish-red ironstone con-

cretions weather from the upper part of the Judith River sandstone.

The uppermost 50 feet of Mesaverde sediments consists of sandstone

interbedded with gray to olive-gray shale. Primary structures include

small and large scale crossbedding as well as fossilized wood fragments. 69

Since there is no Bearpaw Shale present in the study area the Mesaverde sandstones are immediately overlain by interbedded sandstone, shale, lignite, and bentonite layers of the Meeteetse Formation.

The Campanian Mesaverde Formation (Gill and Cobban 1973) records the deposition of the first coarse clastic debris to enter the Lewis

Sea from the west. Source areas for Mesaverde sandstones were in the

Precambrian Belt Series, the Elkhorn Uplift of Montana, and the Idaho

Batholith, which formed part of the Cordilleran Highland along eastern

Idaho and Utah (Pryor 1961).

Gill and Cobban (1973), in a series of diagrams, hypothesized broad east-west shoreline migrations during 11esaverde time. Strand1ine positions were dependent on the amount of sediment supply and the rates of basin subsidence (Pike 1947). Shoreline transgressions took place more rapidly than shoreline regressions and produced sharp shale­ sandstone boundries compared to interfingered regressional boundries

(Gill and Cobban 1973).

In Montana, the Eagle and Judith River Formations are presumed to represent wide ranging progradational shorelines while the Claggett and

Bearpaw Shale Formations record westward transgressions of the Lewis

Sea. In Wyoming, however, Gill and Cobban (1973) show that the Mesaverde shorelines progressively migrated eastward in a slow regular fashion from Eagle through Claggett time. Strandline positions near the end of

Claggett deposition were in the vicinity of the eastern Bighorn Basin.

An eastward progradation of strandlines during Judith River time produced thick sequences of marine grading to continental sandstones. Mackenzie 70

(1975) postulated a delta plain or prodelta environment for middle and

upper portions of the Eagle and the lower and middle parts of the Judith

River equivalents in the Bighorn Basin. The Claggett Shale equivalent

is probably of marine origin. The variety of planar and trough cross­

beds, ironstone concretions, and thin mudstone and siltstone in the

Judith River are characteristic of non-marine flood plain environments

(Mackenzie 1975). Central Wyoming never experienced a major Bearpaw

Shale transgression. The eastern strandline of the transgressing

Bearpaw Sea never reached the Powder River Basin (Gill and Cobban 1973).

Meeteetse Formation

Hewett (1914) named the Meeteetse Formation for a sequence of sand­

stone, shale, and coal exposed near Meeteetse Wyoming. Pierce and

Andrews (1941) and Keefer (1965b) described the Meeteetse Formation in the western Bighorn and Wind River Basins respectively. Andrews, Pierce and Eargle (1947) and Glass et al. (1975) mapped the combined Meeteetse and Lance Formations as undifferentiated in the study area. Rioux (1958) and Kozimko (1977) separated them based on excellent exposures in Dry and Little Dry Creeks. The Meeteetse Formation in Sheep Mountain Quadrangle is about 850

feet thick and poorly exposed. The basal units of the Meeteetse form

a low strike valley above more resistant sandstones in the Mesaverde

Formation. Sandstones in the lower Meeteetse can be traced along

strike as vegetation bands even though outcrops may not be exposed.

The upper Meeteetse in the study area is not exposed except where pro­

tected by Quaternary terraces or exposed by gully erosion. 71

The Mesaverde-Meeteetse contact appears conformable and gradational in the field. The upper Mesaverde is a thick bedded series of resistant yellowish-brown calcareous sublitharenites containing fossilized wood and small to large trough cross-beds. Mesaverde sandstones contain a distinctive rusty brown to black layer of ironstone concretions below the Mesaverde-Meeteetse contact which may be used as a marker bed in the area. The basal Meeteetse contact lies about 10 feet above the iron­ stone concretions where a transition between resistant sublitharenite and nonresistant carbonaceous shales and sublitharenites occurs.

The lower Meeteetse is a thick valley forming interval of soft carbonaceous shale, poorly indurated non-calcareous sublitharenite, thin bentonite layers, and lignite. Several bentonite layers become fairly thick. One impure bentonite at 190 feet above the basal contact, has a thickness of 11 feet. A low ridge forming unit consisting of 25 feet of calcareous sublitharenite and thin fissile shale divides the

Meeteetse Formation in half. Above this interval the Meeteetse Formation unlike its lower portion exhibits better development of small scale trough and planar crossbeds. Sandstones in this upper unit are slabby and discontinuous occasionally containing small clasts of mudstone in their matrix.

The Meeteetse Formation in the southern Bighorn Basin is Late

Campanian to very early Maestrichtian (Gill and Cobban 1973). The

Meeteetse Formation, unlike its correlatives the marine Bearpaw and

Lewis Shales of Montana and the Powder River Basin, is thought to repre­ sent continental deposition. While the Lewis Sea transgressed westerly 72

in Montana, strandlines in central lvyoming remained nearly stable or prograded slowly eastward (Gill and Cobban 1973). The Meeteetse

Formation in the study area probably reflects deposition in fluvial flood plains, back bar swamps, and lagoons. These conditions began in central Wyoming during the uppermost phase of Mesaverde time (lfuckenzie

1975). Abundant bentonite layers and a clay-ash matrix in Meeteetse sandstones may suggest an increase in volcanic activity in the western

Cordillera or in the Elkhorn Mountains of Montana (Gill and Cobban 1973).

Lance Formation

In 1888 and 1889, J. B. Hatcher working for O. C. Marsh, discovered abundant fossil remains of mammals and dinosaurs along Lance Creek,

Niobrara County, Wyoming, and assigned them to the Laramide Formation

(Hatcher 1893). In 1889 o. C. Marsh, referred to the Lance Creek fossil deposits as the Ceratops Beds. In 1903 Hatcher suggested that the name be changed to the Lance Creek Beds of Converse County. The Lance For­ mation continued to be called the Ceratops Beds, Laramie Formation or

Lance Creek Beds until 1910 when the U.S.G.S. officially recognized the abbreviated term Lance Formation (Stanton 1910). Hewett and Lupton

(1917), by correlating vertebrate fossils, recognized the Lance For­ mation in the Bighorn Basin but later workers including Andrews,

Pierce, and Eargle (1947) and Robinove and Langford (1963) did not differentiate between the Meeteetse and Lance Formations on maps near Greybull. Rioux (1958) and Kozimko (1977), however, mapped the

Meeteetse and Lance Formations separately. 73

Poorly exposed Lance sandstones form long continuous bands of sagebrush while shale units support little or no vegetation. Lance sandstones are very friable and readily weather into loose sand which is blown by the wind into small dunes anchored around sage. Seven-hundred ninety eight feet of Lance Formation were measured but this thickness is considered a close estimate since the basal Lance contact is un­ certain in the study area. Previously the Meeteetse-Lance contact in the Sheep Mountain area had been chosen at the base of a massive con­ cretionary fossil bearing buff sandstone which overlies a massive con­ cretionary white sandstone (Rioux 1958, Kozimko 1977) but in the study area neither the uppermost Meeteetse or basal Lance sandstones are resistant to erosion. Concretions typical of the Lance at Little Dry

Creek are developed in only two small areas, and physical tracing of the basal Lance contact from Little Dry Creek northwestward along strike was not possible due to Quaternary terrace deposits. For the purposes of this study the Lance-Meeteetse contact was chosen at the base of a thick, calcareous sublitharenite which contains large bulbous dark yellowish-brown sandstone concretions. Concretions of this nature have been described from the type locality of the Lance (Clemens 1960) but are common in most Cretaceous and Tertiary sandstones in Wyoming.

Meschter (1958) attributes them to preferential calcification by migrating groundwater.

Lance sandstones are calcareous and contain some beds of quartz arenite. Colors range from light gray near the base to grayish-orange or dark yellowish-orange with rusty spots near the top of the formation. 74

The sandstones predominantly consist of fine to medium subrounded to rounded, moderate to well-sorted quartz and feldspar grains. Bedding thickness varies from three inches to over 10 feet, with most layers being between one to four feet thick. The sandstone beds form isolated slabby outcrops which often are partially covered by wind blown sand.

Asymmetrical ripple marks occur on some bedding planes while small scale trough and planar crossbeds are common throughout the formation.

Fossil bone fragments and the above mentioned round "cannonball" con­ cretions are restricted to the basal one-fourth of the Lance Formation while thin black to dark gray carbonaceous shales, interbedded with sandstone units, become thicker and more common in the upper one-half.

A five-foot layer of black carbonaceous shale and coal underlies a

30 foot light gray to white, fine grained quartz arenite occuring about

10 feet below the unconformable Lance-Fort Union contact.

The Lance Formation is Maestrichtian in age (Gill and Cobban 1973)

and represents a continuation of terrestrial deposition initiated

during the Campanian. In north-central Wyoming, rates of sediment

supply from western sources, e.g. the Elkhorn Mountains and the Western

Cordillera, surpassed the rate of subsidence in the Lewis Sea during

the Early Maestrichtian. A large prograding delta, the Sheridan Delta,

developed in north-central Wyoming and south-eastern Hontana (Gill and

Cobban 1973). Throughout the development of the Sheridan Arch

fluvial deposition occurred in the Sheep Mountain Area. At the

Lance type locality in the Powder River Basin lenticular sandstones

similar to those in the study area were interpreted as point bar 75

deposits of meandering rivers and streams (Dodge 1976). Organic rich mudstones and siltstones were deposited in ponds and marshes or as

silty overbank deposits (Dodge 1976). Vertebrate remains were con­

centrated within channel sandstones (Clemens 1963).

Tertiary

Paloecene Fort Union Formation

The Fort Union Formation was originally described by Meek and

Hayden (1861) near old Fort Union along the Missouri River in western

North Dakota. Since then, the term Fort Union Formation has been applied to strata of Paleocene age throughout much of North Dakota, Montana, and Wyoming. Jepsen (1940) and Jepsen and Van Houten (1947) proposed

the name Polecat Bench Formation to distinguish between Paleocene deposits in the Bighorn Basin and those of the type Fort Union in North

Dakota and Montana. Though the Wyoming Stratigraphic Nomenclature

Committee accepts the term Polecat Bench Formation in the Bighorn Basin,

the United States Geological Survey does not and continues to apply

Fort Union Formation to Paleocene deposits throughout the Bighorn

Basin.

Two thousand seven hundred seventy-five feet of Fort Union strata were described of an estimated 3,750 feet of total Fort Union exposed

in the Sheep Canyon Quadrangle area. The Fort Union Formation is ex­ posed in over one third of the map area but good outcrops occur only under Quaternary terrace remnants, in ephemeral stream beds, or where resistant sandstones may be present. Differential erosion between 76 lenticular sandstones and thin layers of fissile shale in the lower part of the Fort Union in the study area produce a characteristic

"tombstone" topography.

Slabs of sandstone, standing in lines, may tower over 20 feet above the surrounding low relief plain. The intercalation of thin shale with the steeply dipping lenticular crossbedded sandstones results in a variety of anomalous structural attitudes. Strata within distances of

30 feet range from normal to overturned and back again without a change in strike (Fig. 9, photo b). The dip angle progressively shallows basinward in the Fort Union Formation from a maximum of 56 degrees west along the west flank of Sheep Mountain Anticline to 3 degrees west at the Fort Union-Willwood contact. The most rapid change occurs about midway in the Fort Union. Here dip changes from 46 degrees, to 17 de­ grees, to 3 degrees west through about 600 feet of strata. This rapid change in dip appears to correspond to a decrease in sand and an increase in mudstone, carbonaceous shale, and coal. The upper 1,000 feet of Fort

Union mudstones dip from six to three degrees west and form a flat featureless plain. Only sinuous sandstone bodies representing channel deposits of Paleocene streams occasionally stand out in low relief above the plain.

The Lance-Fort Union contact in the Greybull-Sheep Mountain area is mapped at a prominent angular unconformity (Hewett and Lupton 1917,

Kozimko 1977). Rioux (1958), however, placed the contact below this unconformity at the first coal layer encountered in the Lance Formation.

In areas of low relief, angular discordances between the Lance and Fort 77

Union are difficult to ascertain. In this case the contact was placed between lighter colored more massive and continuous Lance sandstones and generally darker brown, lenticular, thinner Fort Union sandstones and interbedded shales.

The lower 457 feet of the Fort Union is composed of yellowish orange to olive-gray sublitharenite or litharenite. Locally, ferruginous concretions and fossil leaf imprints are present. Large scale trough and planar crossbedding is characteristic of the lower sandstones though small scale crossbedding is also common. Sandstone beds range from 2 to 15 feet in thickness and are highly lenticular. Individual sandstone bodies seldom persist over 15 to 20 feet along strike. The lower sand­ stones are intimately interbedded with thin fissile shales and olive­ gray siltstones. Erosional differences between sandstone and softer shales and siltstones account for anomalous dips and the characteristic weathering of the lower part.

Above this lower sandy unit the amount of coarse materials decreases.

Lenticular sublitharenite beds, though common, are thinner and more dis­ persed and small scale trough crossbedding is the characteristic primary structure. Olive-gray siltstones, gray shales, thin black coal, and carbonaceous shales form persistent continuous layers. TIle amount of siltstone, mudstone, and shale increases rapidly above 1,250 feet in the

Fort Union. Thin two-to -four foot thick fine grained calcareous sub­ litharenites form small lenticular ledges within thick light olive-gray mudstones. Above 2,000 feet from the base of the Fort Union, sandstone layers are thin, isolated, and very uncommon. The dominant lithology 78

of the upper 1,700 feet is olive-gray mudstone or variegated gray and olive-gray siltstone. Carbonaceous shale and coal layers are abundant occasionally attaining five feet in thickness. Crossbedding directions

in one sinuous sandstone body indicate a south-easterly current flow.

The Fort Union Formation in the Bighorn Basin represents the entire

Paleocene epoch (Jepsen 1940). Continued terrestrial deposition was

aided by increased Laramide uplift in Wyoming. Deposition occurred in

sluggish st~eam channels and flood plains with bordering wooded swamps

and freshwater marshes (Van Houten 1957). A vast coal swamp existed in

the Powder River Basin indicating that climatic conditions remained warm and moist throughout the Paleocene (Roth 1975). Also, during the

Paleocene, the first expression of the northern Bighorn Mountains

occurred (McGrew 1971) and angular unconformities were developed between

the Lance and Fort Union Formations. A continued subsidence of the

Bighorn Basin increased thicknesses of fine grained deposits derived

from eroding Mesozoic sediments in western Wyoming and southwestern

Montana (Keefer and Rich 1957).

Paleocene-Eocene

Late Paleocene-Eocene Willwood Formation

F. B. Van Houten (1944) named the Eocene deposits in Bighorn Basin

the Willwood Formation. Previously these strata had been assigned to

the Wasatch Group of southwestern Wyoming and Utah by Cope (1882) on

the basis of mammal correlations. J. L. Wortman (1882), who collected

fossils for Cope first mapped the extent of the Wasatch Formation in 79

the Bighorn Basin and hypothesized a lacustrine origin for them.

Fisher (1906) produced a more accurate map of the Wasatch distribution and proposed a fluviatile origin for the Bighorn Basin Wasatch which was later substantiated by Loomis (1907). Sinclair and Granger (1911,

1912) and Granger (1914) subdivided the Wasatch in the Bighorn and Wind

River Basins into the Greybull, Lysite, and Lost Cabin Faunal Zones on

the basis of characteristic mammalian assemblages. Stow (1938) studied heavy mineral suites within the Upper Cretaceous and Lower Tertiary of

the Bighorn Basin. After Van Houten's work (1944, 1948) Neasham (1967)

and Neasham and Vondra (1972) studied the depositional history of the

Willwood sediments.

The Willwood Formation overlies the Fort Union Formation in the

extreme southwest portion of the study area. It is slightly more re­

sistant to erosion than the Fort Union Formation and dips at six to

three degrees west forming distinctive and colorful variegated badlands

in the Bighorn Basin.

Estimated total thickness of the Willwood Formation ranges from

2,500 feet (Van Houten 1944) to 2,300 feet (Neasham and Vondra 1972).

Only 421 feet of Willwood strata exist in the southwestern corner of the

study area. There they are protected from erosion by Quaternary terrace

gravels.

The Fort Union-Willwood contact is conformable in Sheep Canyon

Quadrangle. It was placed at the base of the first red bed in a repe­

titive sequence of variegated red and gray siltstones and mudstones

(personal cornm. Dr. Carl Vondra, Iowa State Univ. 1976). The Willwood 80

sediments are dominantly siltstones and mudstones of a variety of

colors and suggest an ordered sequence of formation. Dark reddish­ brown, greenish-gray, brownish-gray, and light olive-grays are the most common shades. Bedding within the siltstones and mudstones is very regular and easily traceable. Bone chips and fossil teeth are

fairly common in the Willwood of the study area. Willwood siltstones are interbedded with thin lenticular channel

sandstones. The sandstones are calcareous sublitharenites, usually light gray in color, with fine to medium fine-grained well-sorted sub­

rounded grain textures. Small scale trough crossbedding was the only primary structure observed. The Willwood sandstones are ledgy and

lenticular in appearance with a pitted weathered surface. They protect

underlying mudstones from erosion and emphasize the badland weathering

characteristics of the Willwood.

The Willwood Formation, based upon mammal correlations is latest

Paleocene through Early Eocene in age (Neasham and Vondra 1972), and was deposited as a consequence of rapid Laramide uplift in mountain

ranges marginal to the Bighorn Basin. Sediments derived from nearby

Fort Union, Mesozoic, and Paleozoic strata, and quartzites from sources

further west were deposited in alluvial fan conglomerates along the west

flank of the basin. A paleo stream system carried the detritus in a

generally north-northeast direction from the fanglomerates. It pro­

gressively winnowed out finer particles and deposited them in flood

plains marginal to stream channels (Neasham and Vondra 1972). The flora

and fauna of the Willwood Formation indicate growth in a warm temperate 81

lowland or subtropical environment (Van Houten 1944). The developing mountain ranges are not thought to have been barriers for fauna, migration or to have significantly affected the climate of the area

(Van Houten 1944). Variegated bedding within the mudstones of the

Willwood is commonly attributed to soil formation under differing drainage conditions. Red beds are thought to represent well-drained soils where free iron was oxidized, while drab gray and green colors represent poorer drainage conditions and a reducing environment (Neasham and Vondra 1972). 82

STRUCTURE

The structural elements of Sheep Canyon Quadrangle are portions of larger folds contained within a NW-SE trending belt parallel to the Bighorn Mountains (Fig. 10). Both Sheep Mountain and Alkali

Anticline are open, asymmetrical folds with their steeper flanks facing the Bighorn front. A small subsidiary anticline, open and symmetrical, intersects the Sheep Mountain axis at 53°, Field studies and map patterns indicate that formation thicknesses remain constant on the limbs and nose of the folds. Absence of formational thinning supports a flexure-slip origin for these structures (Billings 1972). Cross­ sectional views of the folds in the study area (Fig. 11) suggest that

Sheep Mountain and Alkali Anticline remain open with depth. Defor­ mation may have occurred over a platform of subsiding crystalline rocks. The most competent layers in the study area are 250-900 foot thick carbonates of the Phosphoria, Madison, and Bighorn Formations.

During Latest Cretaceous (Maestrichtian) time the Paleozoic formations were buried under at least 8,400 feet of Mesozoic sediments. The

Paleocene-Eocene Fort Union and Willwood formations add another 6,000 feet of deposit,

Laramide deformation (Late Cretaceous-Paleocene) in the region produced Precambrian cored block uplifts separated by subsiding sedi­ mentary basins. The uplifts rose along vertical faults in the Pre­ cambrian basement, which changed from high angle to low angle reverse within the Paleozoic and Mesozoic cover. Beloussov (1962) proposed 83 Fig. 10 Structural Location Map of Sheep Canyon Quadrang1e,(Andrew8, Pierce, and Eargle 1947) 84 85 Fig. 11 Structural Cross-Sections of Sheep Canyon Quadrangle 86

A

1 8 8

4600 42

d

Symbols used in addition to those on page 23.

Devonian Ordovician Cambrian Precambrian I I 1 I

Duperow Formation Bighorn Dolomite Gallatin Limestone Gros Ventre Formation Flathead Sandstone

Scale I: 2400

Vertical :: Horizontal 87

an origin for basin folds in the Caucasus Mountains which may also apply to the Bighorn Basin. The rising block uplifts loaded the trapped basin sediments vertically by the added weight of sediments stripped into the basin and marginally by the weight of the overthrust mass as well as gravitational sliding along the upturned strata of the mountain front. Stresses such as these could be expected to produce folds, paralleling marginal uplifts in trend, with their steeper 'flanks facing the mountains. Billings (1972) states similar f<;>1ds have been formed experimentally. Folds like those found along the basin margins probably also exist buried beneath the thick Tertiary cover in the center of the

Bighorn Basin.

. Structural Elements of Sheep Canyon Quadrangle

Sheep Mountain Anticline

Sheep Mountain Anticline is the dominant structural feature in the study area. This large Laramide fold involves all of the stratigraphic units through the Paleocene Fort Union Formation and is a striking example of an asymmetrical, doubly plunging anticline. The steeper

0 0 east flank dips nearly 64 eastward whereas the gentle flank dips 17 -

18 0 westward. The axial trace of the fold trends N.43°W, and the axial plane is inclined 55 0 west (Fig. 11). Rioux (1958) and Kozimko

(1977) estimate approximately 400 to 500 feet of closure for the structure.

A small nearly symmetrical anticline is developed on the western margin of Sheep Mountain. The axial trace of the smaller anticline 88

intersects that of the larger one at an angle of 53°. This small anti­ cline trends S.3°W bending sharply to S.22°E. The axial plane (Fig.

11) is nearly vertical at 89° west. The resistant upper carbonate of the Phosphoria Formation forms the flanks and crest of the small fold.

Inliers of red shale or of Tensleep Sandstone are developed where pro­ tecting carbonate layers have been removed. A small landslide is developed on the east flank of the anticline.

Alkali Anticline

The extreme southern nose of Alkali Anticline extends into the northern limit of the study area. As seen developed in the soft Cody

Shales, the axial trace trends S.45°E. The axial plane dips 88° west and indicates a slight asymmetry to the nose of Alkali Anticline.

Strike and dip within the Cody Shale is difficult to assess so the positions of the axial trace of Alkali Anticline and the adjacent syncline could only be approximately placed.

Synclines

In the center of section 34, T.54N. R94W. between Sheep Mountain and the small subsidiary anticline on the west is developed an asymmet­ tical syncline. The axial trace parallels the small anticline trending

S.19°E. The axial plane is incined 10° west. The expression of this syncline is rapidly lost beneath Bighorn River terraces and colluvial deposits in section 3, T.53N. R.(RW.

Another small syncline exists between the small subsidiary anti­ cline and Alkali Anticline. This syncline is barely recognizable due 89

to the soft non-resistant nature of the Cody Formation in which it is developed. The axial trace varies from S.23°E. to S.45°E. and the axial plane dips 69 0 to the east.

A small asymmetrical syncline is inferred to be buried beneath colluvial and alluvial deposits in the north-eastern corner of the study area (Fig. 2 and 12). Field work by Mr. Jim Carlson indicates that the axis of the syncline, exposed north of the study area, approxi­ mately parallels Sheep Mountain's structural axis. This syncline separates Sheep Mountain Anticline from Spence Dome and probably connects with a bifurcating syncline mapped by Kozimko (1977).

Faults

Faulting in Sheep Canyon Quadrangle is chiefly confined to dis­ placements within the Madison Formation. Several small faults were recognized in the walls of Sheep Canyon but as with the largest of these, a reverse fault with 4 to 5 feet of displacement, they could not be traced into overlying formations and were considered related to solution and collapse in the Cave Member of the Madison Formation.

Special attention was paid to the northeast corner of the study area, section 35, T.54N. R.94W. where Sheep Canyon Quadrangle joins

Greybull North Quadrangle. Kozimko (1977) mapped several northwest­ southeast trending faults near the crest of Sheep Hountain Anticline.

These faults were found to be from 5 to 20 feet in vertical displace­ ment and were expressed as offsets of the upper Phosphoria carbonate

unit. Displacement of underlying Tensleep sandstones or the Madison

Formation was not observed and physical tracing of these faults indi- 90

cated that either they did not enter the study area or the faults, if present, had so little displacement that they could not be recognized.

It is thought that these faults are probably intraformational and represent the adjustments of more competent carbonate layers in the

Phosphoria Formation to deformation within underlying red shales.

Joint fractures

Joint fractures are common in almost every competent formation in Sheep Canyon Quadrangle. No measurements of joint trends were made for this study. Johnson et al. (1965) measured joint trends on Sheep

MOuntain in the Tensleep Sandstone and the Cloverly Formation while

Warner (1968) measured joint trends in the Mowry and Frontier Formations.

These authors found that most joints trended northeast to east or north­ west to west and either paralleled, fell perpendicular or at some angle to the structural axes of folds (Warner 1968).

Structural History of the Sheep Mountain Area

Pre-Laramide tectonic activity

A large continental shelf formed in Cambrian time between the developing Cordilleran trough in Idaho and the Transcontinental Arch in eastern South Dakota and northwestern Iowa (Eardley 1962, Thomas 1965).

A continental seaway came to occupy this depression, first depositing sandstones upon Precambrian gneiss and shist and later thick sequences of shale. Subsidence and deposition on this shelf continued until

Silurian times when regional uplift and erosion caused a removal of

Silurian strata (Thomas 1965). Isopachous maps prepared by Thomas (1965) 91

indicate that marine deposition occurred during Devonian and early

Mississippian time but northwesterly tilting and uplift of the shelf area during late Mississippian time produced a regional erosion surface and karst topography in the upper part of the Madison Formation.

Marine deposition and shelf stability were characteristic of most of Paleozoic and Mesozoic time except for four periods of tectonic activity. Regional uplift occurred near the end of Pennsylvanian time.

Associated erosion progressively truncated the Tensleep Formation from south to north. During Triassic time erosion of all but the Red Peak

Formation in the Chugwater Group occurred in the study area. During

Late Jurassic to Early Cretaceous time uplift and erosion with subse­ quent terrestrial sedimentation deposited the Cloverly Formation.

Following gradual subsidence an inland sea occupied the study area un­ til late Cretaceous time when increasing tectonism in the western

Cordillera caused a final marine regression and deposition of the

Judith River equivalent of the Mesaverde Formation in the study area.

Up until this point tectonic movements of the Cordilleran shelf area had been minor compared with the following Laramide structural period

(Fanshawe 1971).

Laramide tectonic activity

Beginning in Late Cretaceous and Early Tertiary time the first expression of regional Laramide deformation could be detected in central

Wyoming. An angular unconformity occurred between the Lance and Fort

Union Formations indicating the beginning development of the Bighorn

Basin. By the end of Paleocene time all of Wyoming's major mountain 92

ranges except the Tetons, Absaroka Range, and the eastern one-half of the Gros Ventre Range had formed (Love 1960). Northeastward compres­ sion by Laramide forces created folds in Paleozoic and Mesozoic strata of the Wyoming basins associated with large scale vertical uplifts of the Precambrian crystalline basement. Within the Bighorn Basin this

Precambrian basement was buried beneath about 10,000 feet of Paleozoic and Mesozoic strata (Chamberlain 1940). Chamberlain hypothesized that

the Paleozoic and Mesozoic sediments were too thin to control structur­ al behavior in the shelf area so crystalline Precambrian rocks, too

competent to fold, deformed by breaking into blocks and either rose or sank to form mountain ranges or asymmetrical basins.

The mountain ranges in Wyoming are bordered by folds, high angle normal and low to high angle reverse faults. Local structural relief

from the top of the Precambrian basement in the Bighorn Basin to the

Precambrian surface exposed on the Bighorn Mountains can exceed 34,000

feet (Prucha et al. 1965). Block faulting of Precambrian crystalline

rocks along high, angle reverse faults are common at Wyoming mountain

fronts. These faults begin vertically at the base of the block but

progressively change to high angle reverse and then to low angle re­

verse faults near the surface (Prucha et al. 1965). The exact shape

of an upthrust fault above a break in the basement rock depends upon

thickness of the sedimentary strata above the basement, the types of

rocks involved in the fault, the attitude of the basal fault plane,

the magnitude of basement displacement and the topographic effects

produced by faulting (Prucha et al. 1965). 93

Sedimentary rocks above moving basement blocks deform in several ways. They may deform by tilting or rotating with the basement block

(Hoppin and Palmquist 1965), by folding due to drag along faults (Hoppin and Palmquist 1965), by forming monoclinic folds (Stearns 1971), or by forming asymmetrical minor folds along basin margins (Chamberlain 1940).

Asymmetrical minor folds were formed parallel to the mountain front and are characteristic of basins like the Bighorn Basin. These folds, like Sheep Mountain and Alkali Anticline, commonly have steep flanks facing the mountain front. Chamberlain (1940) attributed asym­ metrical minor fold formation to northeast-southwest crustial shortening of weak Paleozoic and Mesozoic sediment above more rigid and faulted

Precambrian rocks. Fanshawe (1971) suggested that large minor folds may have formed by folding above a high angle upthrust fault. Beloussov

(1962) attributed sediment compression between vertical mountain up­ lifts as a cause of basin folds, and other authors have proposed folding above small tilted Precambrian blocks.

Work by Hoppin and Palmquist (1965), showed that only a few high angle faults which cut both basement and cover rocks were controlled by

Precambrian foliations or shear zones. Their study concluded that

generally, basement fractures were not propagated upward through weak

Cambrian shales and that northwest to northeast trends common to both

Precambrian and younger rocks were the result of regional stresses during Laramide orogeny. Precambrian shear zones and foliations were found to be sharply discordant with the main trend of the Bighorn 94

Mountains and were not considered responsible for their outline (Hoppin and Palmquist 1965).

Hoppin and Jennings (1971) proposed the following age relationship for Laramide structural features developed in the Bighorn Mountain Area.

1) The development of northwest trending folds.

2) The main uplift of the Pryor, Bighorn, Bridger,

and Owl Creek l1ountains.

3) Development of east-west lineaments crossing

mountains and basins.

4) Development of N.IOoW. to N.lOoE. high angle faults

and associated monoclinic flexures.

5) The formation of flank thrusts along the margins

of the uplifts.

Post Laramide tectonic activity

Tectonic activity progressively declined through Middle and Late

Eocene time in the Bighorn Basin (Love 1960). Volcanism increased along the west flank of the basin in the Absaroka area and large amounts of stream and wind borne ash accumulated in the basins (Love 1960).

Stream drainage formerly northeasterly (Neasham and Vondra 1972) was reversed and water became ponded in Middle Eocene time (Love 1960).

Large lake deposits, like those of the Tatman-Mountain Formation in the Bighorn Basin, were formed.

During Oligocene time eastward stream drainage was restored.

Volcanism in the Province filled the Wyoming basins nearly to the highest mountain tops with volcanic ash (Love 1960). Crustial stability 95

however, existed from Miocene through most of Pliocene time. In

Late Pliocene and Pleistocene times large scale regional uplift resulted in erosion in the Wyoming basins and the Pliocene drainage systems were superimposed on mountain ranges and anticlines. 96

GEOMORPHOLOGY

The Tertiary Period in the Bighorn Basin was principally a period of deposition and aggradation (Love 1960). During the Late Pliocene and Early Pleistocene regional uplift caused river downcutting and excavation of Pliocene, Miocene, Oligocene, and Eocene sediments.

Holocene features, therefore, reflect the responses of these Plio­

Pleistocene stream systems on buried Laramide structures.

Geomorphic Elements of Sheep Canyon Quadrangle

Topographic features

The topographic relief of Sheep Mountain is caused by erosional differences between resistant hard Paleozoic limestones and soft non- resistant Mesozoic sandstone and shale sediments. Inliers of Madison /

Limestone and Amsden Formation occur in gullies incised through over- lying strata near the crest and on the flanks of the anticline. Ero- sional outliers of Tensleep Sandstone occur between parallel gullies cut through the Tensleep and into the Amsden Formation. Flatirons of

Phosphoria carbonates are developed on the steep eastern flank of Sheep

Mountain whereas long dip slopes of Tensleep and Phosphoria Formations form the more gradual western flank.

Sheep Canyon is a classic example of a superposed valley. The canyon, whose walls rise nearly 980 feet above the Bighorn River through exposed Madison, Amsden and Tensleep Formations, was formed when the 97

Bighorn River cut downward through Tertiary basin fill into the folded and buried Paleozoic rocks of Sheep Mountain (Fig. 9, photo C) (Mackin

1937) .

Hogbacks are formed wherever steeply dipping thick resistant strata are interbedded with softer less resistant units. The Phosphoria,

Gypsum Springs, Sykes Mountain and Mowry Formations form the most promi­ nent hogbacks in the study area.

Gravel covered "benches" or terrace remnants occur principally in

the southern one-third of the study area where their relief may be over

200 feet. These benches originally paleo-stream valleys of either the

Greybull or Bighorn Rivers (Mackin 1937), have preserved underlying soft bedrock with their coarse gravel veneer while surrounding areas were eroded.

Terrace remnants

Rock cut, gravel covered, river terraces were formed by stream

erosion at formerly higher levels of the Bighorn and Greybull Rivers.

Mackin (1936, 1937) described the process of downcutting and stream

capture which displaced the Greybull River south from its former course along Emblem Bench as well as criteria to recognize terrace origins

from the characteristic bedload of the parent stream. The Greybull

River carried coarse rounded basaltic and quartzitic debris from the

Absaroka Mountains eastward while the Bighorn River carried a mixed load (

of carbonate, chert, and igneous rock derived from Paleozoic terrains

as well as volcanic cobbles and quartzite contributed by tributary

streams like the Greybull River. 98

For quick terrace identification purposes a one square foot area was arbitrarily chosen on a terrace surface. One hundred clasts larger than one inch in diameter were sorted by lithology into piles where the number of clasts of like lithology were counted. Relative percentages of carbonate, quartzite, andesite, and basalt could therefore be deter­ mined. For correlation purposes terrace remnants were assigned to the

terrace levels used by Palmquist (1978)(Fig. 12).

Gravel and floodplain deposits of present Bighorn River, Dry Creek,

and Little Dry Creek were mapped as Quaternary alluvium Qal. Qtl

terraces (Himes Complex and Red Flat Fig. 12) were formed from Bighorn

gravels (48% volcanic clasts, 30% quartzite, 12% plutonic igneous

clasts, 10% carbonate clasts) and are located about 30 feet above the

present Bighorn River floodplain. Qt2 levels (Sanatarium terrace Fig.

12) consist of both Greybull and Bighorn terraces. Emblem Bench, a

Qt2 level, was the former valley of the Greybull River (Mackin 1937)

and is composed of about 74% mafic volcanic clasts, 26% quartzite

cobbles, and minor amounts of chert. Qt3 terraces (Basin Cemetary

level Fig. 12) are about 130 feet above the present Bighorn River and

are Bighorn in origin. Qt4 terraces (Water Tank level Fig. 12) have

mean heights above the Bighorn River of about 220 feet. They are

Greybull River in origin along Emblem Bench and Bighorn in origin else­

where in the study area. Qt5 terraces (Cottonwood level Fig. 12) lie

about 290 feet above the Bighorn River, Qt5 levels bordering Emblem

Bench are Greybull in origin and all others are Bighorn deposits in the 99

y u ·... v 8 Divide (extrapolated) ...... "v 4t400 '"'l., •••• ~•••• Ten (extrapolated) 2 -...... ~ ...... d ~ Qt6 I 0···· ... . em er -...... I '-' A am ...... (T1

'" " .. ~.. Qt5 Cottonwood CD < . . . '. Qt4 a ..... i-' Emblem Bench ..... t2·::::::l1t~::.",...Water Tank o o o ------...... ~ ..---. Basin Cemetary 4,000 :::l ------Qtl>-- -- __ .. Sanatarium ..-er-. -CD ...... ::,. ----- '~RedFlat CD Greybull -. '-- Himel!..... -...... Terraces River - -.Complex '---J N. or E. side ~ S. or W. side Mod ified from Palmquist (1918) ~----.-----~.-----~------r-----~------~----~3pOO 10 8 6 4 2 o Distance {miles} 101 study area. QtS terraces in section 26, T.S3N. R.94W. are obscured by colluvium and poorly developed. The Qt6 terrace (Alamo level Fig. 12) lies 370 feet above the present Bighorn River and is also Bighorn in origin. Volcanic ash associated with QtS levels indicate 600,000 year old ages (Palmquist 1978). Qt2 terraces have been correlated with

Bull Lake glacial deposits suggesting that Qt2 terraces are no younger than 70,000 years old (Palmquist 1978).

Colluvial deposits

Coarse rock fragments, usually angular carbonate clasts mixed in a silty matrix form thick deposits in sections 25 and 26, T.54N. R.94W.

Where dissected by gullies, the deposit is at least 10 feet thick and is poorly stratified. The poor stratification and imbrication indicates a rudimentary degree of water sorting but the angular nature of the clasts indicates short transport distances. These deposits were mapped as (Qc)

Quaternary colluvium. Ribbon Canyon in sections 33 and 34, T.S4N. R.94W. is also floored with a thick layer of colluvium which obscures the Red

Peak and Dinwoody Formations. This material is fine grained with angular carbonate fragments apparently derived from the Phosphoria Formation.

The silty matrix was probably derived from slope wash from the lower

Phosphoria red beds, the Red Peak, and Gypsum Springs formations. This silty colluvium extends into sections 2 and 3, T.53N. R.94W. where it is well-stratified and overlies a series of Qtl Bighorn River terraces.

Alluvial deposits

Fine grained, well-stratified alluvial deposits occur along Lovell

Draw. Little Dry Creek and Dry Creek. Ten feet of alluvium was arb i- 102 trarily chosen as the minimum thickness necessary to map. Alluvium exceeded 20 feet in section 11, T.S3N. R.94W. near the mouth of Lovell

Draw but decreased rapidly upstream in section 10. Locally thick remnants of alluvium were preserved along stream banks but in general

Lovell Draw was cutting into bedrock and did not have enough Qal to map.

The same situation exists at Little Dry Creek. Mappable Qal ex­ tends to the U.S. 310 highway bridge in section 27, T.S3N. R.94W. where about 12 feet of well-stratified Qal is exposed. Further west, however,

Little Dry Creek is cutting southward into the Fort Union Formation, leaving less than ten feet of Qal on its north bank.

The entire valley of Dry Creek was mapped Qal though alluvium thickness is not known. Dry Creek is well-established and numerous abandoned meander channels scar the valley. Little downcutting by Dry

Creek is apparent so alluvium thicknesses were inferred to be thick enough to map.

The modern gravel covered stream bed as well as the meander belt and floodplain of the modern Bighorn River was also mapped as Qal.

Sand deposits

Sand dune deposits 2 to 3 feet high and from 5 to over 30 feet in length, were found in the SE 1/4 section 29, T.S4N. R.94W. in the

Frontier Formation, in the NE 1/4 section 27, T.S3N. R.94W. and SW

1/4 section 26, T.S3N. R.94W. on the Fort Union Formation, and in the

SE 1/4 section 26, T.53N. R.94W. on the Meeteetse Formation. The loose sand accumulated in low hummocks around sagebrush or grass. The vege­ tation, which seem to prefer the higher permeability of the loose 103

sand, helps to anchor and stabilize the small sand dunes from further wind erosion. The sand dunes in section 29, T.SlN. R.94W. were derived from the erosion of the Torchlight Sandstone in the Frontier Formation.

Small andesite cobbles and rounded chert pebbles were associated with the loose sand cover.

In sections 26 and 27, T.S3N. R.94W. the sand dunes, about 2 to

3 feet thick, were thought to be eroded from the Lance Formation and deposited in depressions, around sagebrush, or in the lee of hills.

Landslides

A landslide occurs on the east flank of the small subsidiary anti­ cline at NW 1/4 section 34, T.S4N. R.94w. It is developed in red

Phosphoria siltstones, shales, and gypsum beds. Large blocks of gypsum are slumping down hill while red beds with anomalous di~s show various degrees of rotation. Stream erosion may have removed the support for the steeply dipping Phosphoria strata and allowed them to slide down dip.

Along the north side of Dry Creek Valley a series of bedrock slumps have been initiated in Fort Union siltstones and overlying terrace gravels of Emblem Bench (Fig. 9, photo. d). These slumps are character­ ized by backward rotation of the slump blocks, relatively fresh fracture planes, rough blocky appearances, and non-eroded toe slopes. Marshes or springs are often associated with the slumps while above, on Emblem

Bench, arcuate cracks are formed in the terrace overlooking Dry Creek.

Fence lines in the area may also be seen dangling in the air indicating recent slump movement. 104

The terrace deposits probably contained little water before 1955 when intensive irrigation was brought to the Greybull-Dry Creek area

(Robinove and Langford 1963). With increased irrigation to support alfalfa and sugar beet production on Emblem Bench, water may have become perched in the gravels overlying the Fort Union Formation or seeped down bedding planes or fractures into the bedrock. When saturated, silt­ stones and thin coal layers of the Fort Union become unstable and fail where unconfined conditions arise. So long as intensive irrigation is practiced on Emblem Bench continued slumping of Emblem Bench into Dry

Creek Valley will probably continue. 105

SUMMARY

Sheep Canyon Quadrangle is located on the west flank of Sheep

Mountain about five miles northwest of Greybull Wyoming. The strata of the study area represents sedimentation upon a stable platform to the east of a miogeocline (continental shelf) situated along the Idaho­

Wyoming border. Early Paleozoic sedimentation began with the depo­ sition of the transgressive marine Flathead Sandstone on an irregular erosion surface developed on Precambrian granites and gneisses. The rest of the Cambrian per~od was characterized by deposition of thick sequences of shale and thin limestone. The Ordovician Bighorn Dolomite,

Devonian Duperow, and Mississippian Madison Limestone added a thick sequence of limestone and dolomite to the platform and produced the most resistant and competent Paleozoic units in the study area.

Silurian strata are not present.

Late Mississippian (Meramecian) uplift produced erosion and a karst topography on the Madison Limestone surface. Periods of regional uplift followed by subsidence, continued through the Paleozoic and

}lesozoic eras to produce unconformities between the Tensleep and

Phosphoria Formations, the Red Peak and Gypsum Springs Formation, the Gypsum Springs and Sundance Formation, and the Sundance and

Morrison Formation.

The Jurassic-Cretaceous boundary occurs between the Morrison and

Cloverly Formations during a period of continuous terrestrial deposition. 106

Silts, muds, and volcanic ash accumulated in tropical swamps and lakes during Morrison and Cloverly times. These deposits were superseded by tidal sediments of the Sykes Mountain and marine shales of the

Thermopolis Formation. The Cretaceous in the Bighorn region was a time of thick shale accumulation. The Thermopolis, Shell Creek, Mowry, and

Cody formations were deposited in a vast boreal seaway. Regressions of this sea are represented by sandstones in the Muddy and Frontier For­ mations. Sandstones in the Eagle Member of the Mesaverde Formation were deposited when Late Cretaceous seas shallowed and increased sediment was supplied from rising uplands to the west. By the Late Campanian terrestrial sedimentation had returned in the study area. Sandstones and shale were deposited in marshes, river channels, and floodplains from Late Campanian through Maestrichtian times, and formed the

Meeteetse and Lance Formations.

East-west crustal compression in the Bighorn Basin created asym­ metrical NW-SE trending folds. The brittle Precambrian granite and gneiss reacted to Laramide forces by rising along high angle reverse faults to form block uplifted mountain ranges or sinking to form asym­ metrical basins. In the study area a significant angular unconformity formed between the Lance and Fort Union formations in response to

Laramide uplift.

The Bighorn Mountains, low or non-existent in Early Paleocene time, were well-developed by Late Paleocene (Love 1960). Sand supply dimin­ ished and streams became sluggish. Paleocene coal swamps were formed in both the Powder River and Bighorn Basins but ceased prior to Eocene 107

deposition. Willwood fanglomerates along the southwest and western margins of the Bighorn Basin developed proximally to rapidly rising source areas. Northeasterly flowing stream systems carried fine grained deposits into the basin center. Volcanoes, active through the Middle Eocene in the Absaroka region, nearly buried the Washakie

Range in volcaniclastic debris. Also during Middle Eocene time Bighorn

Basin streams became impounded and the resulting lake deposits formed the Tatman Formation.

Aggradation continued in the Oligocene, Miocene, and Pliocene as ash was deposited from volcanic centers in the Basin and Range Province.

During the Pliocene, regional uplift rejuvenated stream systems. The basin fill, which buried all but the highest mountain peaks, was excavated. Rivers such as the Shoshoni and Bighorn superimposed them­ selves on Laramide structures and formed a series of rock cut gravel terraces. One such terrace, Emblem Bench, was the lowest elevation attained by the Greybull River before it was captured and displaced south by a small tributary of the Bighorn River. Stream systems in the study area continue to down-cut. Fresh erosional scars in stream alluvium, colluvial deposits, valley walls, and recent slump deposits indicate continued hydraulic adjustments to the environment. 108

ACKNOWLEDGHENTS

The writer wishes to acknowledge the help of Mr. Dave Prose,

Mr. Pete Julstrom, and Mr. Erik Kvale who aided with mapping and sectioning during the course of this field work. Dr. John Lemish and Dr. Thomas Fenton served on the writer's graduate committee, and

------Dr. Carl Vondra offered suggestions on this manuscript and the field work as well as being the writer's major professor. Their help and counsel are deeply appreciated. In addition, the writer would also like to thank his parents Mr. and Mrs. Edward Ladd for their guidance and support. 109

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