Pennsylvanian and Early Permian 

Paleogeography of the Uinta-Piceance 

Basin Region, Northwestern Colorado and 

Northeastern Utah

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Chapter CC

Pennsylvanian and Early Permian 

Paleogeography of the Uinta-Piceance 

Basin Region, Northwestern Colorado and 

Northeastern Utah

By SAMUEL Y. JOHNSON, MARJORIE A. CHAN, and 

EDITH A. KONOPKA

A multidisciplinary approach to research studies of sedimentary 

rocks and their constituents and the evolution of 

sedimentary basins, both ancient and modern

U.S. GEOLOGICAL SURVEY BULLETIN 1787

EVOLUTION OF SEDIMENTARY BASINS UINTA AND PICEANCE BASINS
U.S. DEPARTMENT OF THE INTERIOR 

MANUEL LUJAN, JR., Secretary

U.S. GEOLOGICAL SURVEY 

Dallas L. Peck, Director

Any use of trade, product, or firm names in this publication is for descriptive purposes only and 

does not imply endorsement by the U. S. Government

UNITED STATES GOVERNMENT PRINTING OFFICE: 1992

For sale by

Book and Open-File Report Sales 

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Library of Congress Cataloging-in-Publication Data

Johnson, Samuel Y.

Pennsylvanian and Early Permian paleogeography of the Uinta-Piceance Basin region, northwestern Colorado and northeastern Utah / by Samuel Y. Johnson, Marjorie 

A. Chan, and Edith A. Konopka.

p. cm. (Evolution of sedimentary basins Uinta and Piceance basins; ch. CO 

(U.S. Geological Survey bulletin ; 1787-CC)

Includes bibliographical references.

1. Paleogeography Uinta Basin (Utah and Colo.) 2. Paleogeography Colorado Piceance Creek Basin. 3. Geology, Stratigraphic Pennsylvanian. 

4. Geology, Stratigraphic Permian. 5. Geology, Structural Uinta Basin (Utah and 

Colo.) 6. Geology, Structural Colorado Piceance Creek Basin. 7. Uinta Basin 

Utrah and Colo.) 8. Piceance Creek Basin (Colo.) I. Chan, Marjorie A. 

II. Konopka, Edith. III. Title. IV. Series. V. Series: U.S. Geological Survey bulletin ; 1787-CC. 

QE75.B9 no. 1787-CC 

[QE501.4.P3] 

557.3 s dc20

[551.7'52'09792] 91-43773

CIP
CONTENTS

Abstract CC1

Introduction CC2

Late Paleozoic tectonism and regional geology CC4

Late Paleozoic climate, eustasy, and cyclicity CC6

Paleogeography of the Uinta-Piceance basin region CCS

Sub-Pennsylvanian geology CCS

Early to early Middle Pennsylvanian Morrowan and early Atokan CC10

Middle and late Middle Pennsylvanian Late Atokan and Desmoinesian CC14

Late Pennsylvanian Missourian and Virgilian CC19

Early Early Permian early Wolfcampian CC23 

Discussion CC27 

Conclusions CC30 

References cited CC31 

Appendix Borehole geophysical logs CC35

FIGURES

1, 2. Maps showing:

1. Uinta-Piceance basin region and outcrops and borehole locations in 

study CC2

2. Schematic paleogeography of Uinta-Piceance basin region and surrounding areas CC3

3. Chart showing Pennsylvanian and Permian stratigraphy of sedimentary provinces in Uinta-Piceance basin region CCS

4. Graph showing inferred sea-level curve for the Pennsylvanian and Early 

Permian CC7

5. Drawings showing schematic Pennsylvanian depositional sequences in UintaPiceance basin region CC9

6-9. Maps of Uinta-Piceance basin region showing paleogeography during maximum regression and transgression:

6. Morrowan to early Atokan CC12

7. Late Atokan to Desmoinesian CC16

8. Missourian to Virgilian CC20

9. Early Wolfcampian CC24 

10, 11. Geohistory diagrams, Uinta-Piceance basin region, for:

10. Oquirrh basin, Wyoming shelf, northern Paradox basin, and Eagle 

basin CC28

11. Total subsidence and basement subsidence in Oquirrh basin, Wyoming 

shelf, Paradox basin, and Eagle basin CC29

Contents III

EVOLUTION OF SEDIMENTARY BASINS UINTA AND PICEANCE BASINS

Pennsylvanian and Early Permian Paleogeography of the 

Uinta-Piceance Basin Region, Northwestern Colorado and 

Northeastern Utah

By Samuel Y. Johnson 1 , Marjorie A. Chan2, and Edith A. Konopka3

Abstract

The Uinta-Piceance basin region of northwestern Colorado and northeastern Utah includes parts of four major sedimentary provinces that were active during the late Paleozoic 

ancestral Rocky Mountain orogeny: the Eagle basin, the northern part of the Paradox basin, the southern Wyoming shelf, and 

the southeastern part of the Oquirrh basin. Depositional patterns in these sedimentary provinces were controlled to varying degrees by eustatic and climatic fluctuations (forced by 

expansion and contraction of late Paleozoic continental ice 

sheets), tectonism, and sediment supply. Four sets of paleogeographic maps illustrate the major changes in paleogeography 

and depositional patterns associated with repetitive transgressions and regressions. In general, clastic deposition (mostly 

sandstone in fluvial, deltaic, and eolian systems) dominated 

during regressions, whereas deposition of marine limestone 

and clastic rocks characterized transgressions. The most arid 

climates are correlated with regressions.

Morrowan and lower Atokan strata throughout most of 

the Uinta-Piceance basin region consist of fine-grained clastic 

rocks (regressive deposits) and more abundant limestone 

(transgressive deposits). Significant Morrowan and early Atokan tectonic uplifts include the Front Range and Sawatch 

uplifts.

The late Atokan to Desmoinesian history of the region is 

characterized by unroofing of the Front Range and Sawatch

Manuscript approved for publication October 10, 1991.

'U.S. Geological Survey, MS 939, Box 25046, Denver Federal Center, Denver, Colorado 80225.

Department of Geology and Geophysics, University of Utah, Salt 

Lake City, Utah 84112.

3Department of Geology and Geophysics, University of Wisconsin, 

Madison, Wisconsin 53706. Present address: 541 Fairmount Avenue, Westfield, New Jersey 0790-1358.

uplifts and rapid rise of the Uncompahgre uplift. In the central 

parts of the Eagle and Paradox basins, tectonic activity resulted 

in restricted circulation and low clastic sediment supply, creating conditions suitable for evaporite deposition during regressions. Regressive deposition was also characterized by 

significant progradation of eolian sands across the Wyoming 

shelf en route to the Oquirrh basin. Limestone deposition was 

dominant throughout the region during transgressions.

During the Missourian and Virgilian, progradation of 

clastic sediments into the central parts of the Eagle and northern Paradox basins resulted in termination of evaporite deposition. During regressions, fluvial and eolian deposition 

dominated in the Eagle basin, and sabkha and (or) shallow 

marine deposition dominated in the northern Paradox basin. 

The Wyoming shelf continued to serve as a conveyor belt of 

clastic sediment to the Oquirrh basin, which was characterized 

by a transition from shallow- to deep-water deposition. Deposition of limestone during transgressions was limited to the 

western part of the Uinta-Piceance basin region and a small 

area in the eastern Eagle basin.

In the early Wolfcampian, the Emery uplift became fully 

or mostly submerged, and the Paradox basin ceased to be a discrete geomorphic element. Deep-water clastic deposition in 

the Oquirrh basin continued from the latter part of the 

Missourian-Virgilian. Missourian-Virgilian depositional patterns in the Eagle basin and on the Wyoming shelf also continued into the early Wolfcampian. Deposition of limestone 

during transgressions was limited to the southwestern part of 

the Uinta-Piceance region.

Regionally, rates and magnitude of subsidence were 

greatest in the Oquirrh basin, intermediate and very similar in 

the Eagle and northern Paradox basins, and lowest on the Wyoming shelf. In all four sedimentary provinces, rates of subsidence were lowest in the Early Pennsylvanian, highest in the 

Middle Pennsylvanian, and intermediate in the Late Pennsylvanian and Early Permian. The timing, magnitude, and geometry 

of subsidence in the Uinta-Piceance basin region strongly

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC1
suggest that regional intraplate deformation reflects the 

overlapping influences of interactions along a more distant, 

convergent continental margin to the south and a more proximal, transform-fault (?) continental margin to the west.

INTRODUCTION

The rectangular Uinta-Piceance basin region of northwestern Colorado and northeastern Utah encompasses the 

Laramide (latest Cretaceous to Paleogene) Uinta and 

Piceance basins and surrounding uplifts (fig. 1). The area is 

mostly underlain by Phanerozoic sedimentary rocks that represent at least five different phases of basin evolution 

(Johnson, Tuttle, and others, 1990). Pennsylvanian and 

lower Wolfcampian rocks were deposited during one of

these phases, generally referred to as the ancestral Rocky 

Mountain orogeny. During this orogeny, the Uinta-Piceance 

region included parts of four major sedimentary provinces: 

the Eagle basin, the Paradox basin, the Wyoming shelf, and 

the Oquirrh basin (fig. 2). Episodic deposition also occurred 

on the Emery uplift and the Callville shelf. Pennsylvanian 

and lower Wolfcampian rocks are widespread in the subsurface of the Uinta-Piceance region, but outcrops are restricted 

to a few areas (fig. 1). Late Paleozoic paleogeographic 

reconstructions of all or significant parts of the UintaPiceance region include those by Hallgarth (1967), Mallory 

(1972, 1975), Rascoe and Baars (1972), Chronic (1979), 

Welsh and Bissell (1979), and DeVoto and others (1986).

Recent detailed sedimentologic studies (for example, 

Kite and Buckner, 1981; Driese and Dott, 1984; Johnson 

and others, 1988; Schenk, 1989; Karachewski, 1992)

112 109° WYOMING IOB°

100 KILOMETERS

EXPLANATION

Outcrop area of Pennsylvanian

and Early Permian rocks 

Trace of Wasatch fault

Sevier-age thrust fault east 

of the Wasatch fault zone

Outline of Paleogene sedimentary 

basin

Figure 1. Map showing the Uinta-Piceance basin region in northwestern Colorado and northeastern Utah. Abbreviations for outcrop locations cited in text: CB, Crested Butte, Colorado; CN, Charleston-Nebo thrust sheet; CR, Crystal River, Colorado; CV, Cove 

Fort, Utah; DR, Duchesne River, Utah; EV, Eagle Valley, Colorado; JM, Juniper Mountain, Colorado; M, Minturn, Colorado; MC, 

McCoy area, Colorado; MIC, Miller Creek, Colorado; MO, Moab, Utah; MT, Mount Timpanagos, Utah; OM, Oquirrh Mountains, 

Utah; RC, Ripple Creek, Colorado; RR, Ruedi Reservoir, Colorado; R1S, Route 148 southwest, Utah; SC, Sols Canyon, Utah; SM, 

Split Mountain, Utah; SV, Strawberry Valley, Utah; SY, Sylvan, Colorado. Numbered triangles show locations of boreholes cited 

in text and listed in the appendix.

CC2 Evolution of Sedimentary Basins Uinta and Piceance Basins
UINTA-PICEANCE 

BASIN REGION

SOUTH AMERICA 

(GONDWANA)

u Hypothetical *u *- ,* * transform

TdUlt OT otOHG

and Stevens (1988)

300 KILOMETERS

EXPLANATION

Uplift or neutral area

\

Figure 2. Schematic map showing paleogeography of the Uinta-Piceance basin region, adjacent basins and uplifts of the ancestral Rocky Mountain orogeny, and inferred continental margins to the west and southeast. Map does not restore Mesozoic shortening and Cenozoic extension in the western United States. Abbreviations: AB, Anadarko basin; AOB, Antler overlap basins; CS, 

Callville shelf; DB, Delaware basin; DEB, Denver basin; DMT, Dry Mountain trough; EB, Eagle basin; EU, Emery uplift; FRU, Front 

Range uplift; MB, Midland basin; OB, Oquirrh basin; ORB, Orogrande basin; PB, Paradox basin; PDB, Palo Duro basin; UU, Uncompahgre uplift; WYS, Wyoming shelf. Areas characterized by minimal uplift or subsidence are considered "neutral." Modified 

from McKee and Crosby (1975), Ross (1986), Stone and Stevens (1988), and Smith and Miller (1990).

demonstrate that Pennsylvania!! and lower Wolfcampian 

strata of the Uinta-Piceance region are commonly cyclic, 

characterized by repetitive sequences of lithofacies that 

represent repetitive successions of depositional environments. Comparable cyclic deposition in upper Paleozoic 

strata has been widely recognized globally and is generally attributed to eustatic fluctuations and associated climate changes forced by expansions and contractions of 

Gondwana continental ice sheets (for example, Crowell, 

1978; Heckel, 1986; Ross and Ross, 1987, 1988).

Existing paleogeographic maps and reconstructions for 

the Uinta-Piceance region greatly homogenize geologic history by showing only the dominant lithology deposited during a particular cycle or sequence and not the variations in 

depositional environments and lithofacies represented in the 

sedimentary cycles. The maps presented herein show 

inferred differences between eustatic transgressive and 

regressive deposition and thus differ significantly from earlier paleogeographic reconstructions. Paired maps for four

time intervals (Morrowan and early Atokan, late Atokan and 

Desmoinesian, Virgilian and Missourian, early Wolfcampian) are presented. These maps represent the duration 

of the ancestral Rocky Mountain orogeny and also correspond to the Absaroka I sequence of Sloss (1982). The maps 

are speculative because many relevant rock units have not 

been studied in detail (particularly in the subsurface), and 

they should be regarded as preliminary working models. 

Their purpose is to illustrate the likely large-scale regional 

variations within a time interval and, in so doing, to demonstrate the necessity of this type of approach for reconstructing late Paleozoic paleogeography.

Acknowledgments. This research was supported by the 

U.S. Geological Survey Evolution of Sedimentary Basins 

Program. We thank Robert Hite, Richard Langford, Omer 

Raup, and John Welsh for important discussions of influence 

lithofacies distributions. A. Curtis Huffman, Jr., Christopher 

J. Schenk, and Marjorie E. MacLachlan provided helpful 

reviews.

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC3
LATE PALEOZOIC TECTONISM AND 

REGIONAL GEOLOGY

The Uinta-Piceance basin region (fig. 1) is part of the 

ancestral Rocky Mountains erogenic province of the westcentral United States. This province is somewhat enigmatic 

in that it formed in an intraplate tectonic setting distant from 

active continental margins (fig. 2). Kluth and Coney (1981), 

Kluth (1986), and Budnik (1986) suggested that deformation 

in this province resulted from the collisional suturing of 

North America and South America-Africa (the Laurasia and 

Gondwana plates) along the Marathon-Ouachita-Appalachian orogenic belt. Kluth (1986) showed that the timing of 

deformation along this suture and in the ancestral Rocky 

Mountains was broadly correlative. Many of the major faults 

in the ancestral Rockies, most prominently the Wichita 

megashear (Budnik, 1986), probably are reactivated zones of 

preexisting structural weakness. Stone (1977) and Budnik 

(1986) emphasized sinistral strike-slip offset and deformation on major northwest-trending structures in the ancestral 

Rockies, whereas Stevenson and Baars (1986) postulated 

dextral strike-slip offset on northwest-trending faults and 

sinistral offset on northeast-trending faults. Kluth (1986) 

suggested that ancestral Rocky Mountains deformation was 

in some ways similar to the intraplate Cenozoic deformation 

of Asia associated with its collision with India; however, 

fault geometry and inferred kinematics between the Asia 

analogue and the ancestral Rocky Mountain orogeny are significantly different.

Both Kluth (1986) and Budnik (1986) considered the 

western continental margin of North America as passive during the late Paleozoic and thus incapable of providing a driving force for intraplate deformation. More recently, Stone 

and Stevens (1988) proposed that this continental margin 

was a northwest-trending left-lateral transform fault (fig. 2) 

that had a structural trend similar to that proposed for at least 

parts of the ancestral Rocky Mountains. Smith and Miller 

(1990) similarly presented evidence for significant late 

Paleozoic extension in Nevada that they related partly to a 

western driving force. If the effects of Mesozoic shortening 

and Cenozoic extension are removed from the southwestern 

United States (following Levy and Christie-Blick, 1989), 

then the proposed left-lateral transform fault of Stone and 

Stevens (1988) was about 500-600 km from the northern 

Paradox basin, much closer than the distance (about 

1,050-1,250 km) between the Paradox basin and the 

Marathon-Ouachita-Appalachian orogenic belt. Using the 

Neogene San Andreas fault and Basin and Range extension 

as a very limited analogue, diffuse extensional shearing 

related to transform faulting can occur hundreds of kilometers from continental margins (Atwater, 1970; Wernicke and 

others, 1988) and thus could have affected the ancestral 

Rocky Mountains.

The ancestral Front Range and Uncompahgre uplifts (fig. 

2) are the major late Paleozoic orogenic highlands in the

Uinta-Piceance basin region. Subsiding areas include the 

Eagle, Paradox, and Oquirrh basins, the Wyoming shelf, the 

Emery uplift, and the Callville shelf. Isopach maps showing 

the thickness of Pennsylvanian to Wolfcampian sedimentary 

deposits are in Mallory (1972, 1975), Rascoe and Baars 

(1972), and DeVoto and others (1986). Late Paleozoic 

regional stratigraphy is shown in figure 3. The ancestral 

Front Range and Uncompahgre uplifts, cored by Precambrian crystalline rocks, form the eastern and southwestern 

boundaries of the Eagle basin (fig. 2), respectively. These 

uplifts had considerable structural and topographic relief and 

were bounded by narrow complex fault zones. The ancestral 

Sawatch uplift formed an intrabasinal block in the southern 

part of Eagle basin and was probably characterized by lesser 

positive relief during the late Paleozoic (Johnson and others, 

1988; Johnson, 1989a, b). The Emery uplift of central Utah 

was intermittently emergent during the Pennsylvanian to 

Wolfcampian; at those times it formed the northwestern 

boundary of the Paradox basin. The Emery uplift was never 

deeply eroded (Mississippian strata are preserved below an 

unconformity with Wolfcampian rocks), however, and thus 

probably had minimal topographic relief.

The Gore fault zone forms the boundary between the 

ancestral Front Range uplift and the Eagle basin. Reactivation of this fault at least three times during the Mesozoic and 

Cenozoic (Tweto, 1977) has obscured its late Paleozoic history. The Gore fault is now locally a reverse fault, but 

DeVoto (1980) and DeVoto and others (1986) showed it as 

a normal fault during the late Paleozoic. The normal-fault 

hypothesis suggests a local extensional environment that is 

consistent with the presence of interbedded volcanic rocks in 

the Middle Pennsylvanian Minturn Formation (fig. 3) 

(Koschmann and Wells, 1946) within a kilometer of the 

Gore fault. Alternatively, the presence of a deep late Paleozoic trough adjacent to the ancestral Front Range (Johnson 

and others, 1988, fig. 3) could reflect crustal loading from 

the east and reverse motion on the Gore fault.

The unexposed Garmesa fault zone forms the boundary 

between the ancestral Uncompahgre uplift and the Eagle 

basin (fig. 2). Stone (1977) suggested that the Garmesa fault 

was a late Paleozoic oblique-slip fault; Heyman and others 

(1986, fig. 13) showed it as a reverse fault in the late Paleozoic. Numerous intrabasinal faults (with inferred normal, 

reverse, and oblique-slip displacement) were also active in 

the Eagle basin during the late Paleozoic (for example, 

Stone, 1977, 1986; DeVoto, 1980; Waechter and Johnson, 

1985,1986; DeVoto and others, 1986; Dodge and Bartleson, 

1986; Johnson and others, 1988; Schenk, 1989; Johnson, 

Tuttle, and others, 1990). These structures had significant 

influence on patterns of deposition and subsidence. The 

Eagle basin includes as much as 3,200 m of Pennsylvanian 

and Wolfcampian strata (Johnson and others, 1988). It 

merges to the northwest with the southern part of the Wyoming shelf depositional province (fig. 2). The Wyoming 

shelf contains a comparatively thin section (about 700-800

CC4 Evolution of Sedimentary Basins Uinta and Piceance Basins
NORTHERN

PARADOX

BASIN

Curry 

Peak 

and

Freeman 

Peak 

Fms.

Elephant

Canyon

Formation

Figure 3. Chart showing Pennsylvanian and Early Permian stratigraphy of five sedimentary provinces in Uinta-Piceance region. Units may have different age ranges in areas other than those discussed here. Time scale for age and series boundaries is 

from Haq and Van Eysinga (1987); suggested time scale of Klein (1990) is shown in parentheses. Loope and others (1990, fig. 

2) suggested a different stratigraphic nomenclature for Paradox basin in which Hermosa Group has formation status, lower 

part of upper member of Hermosa Formation is equivalent to Honaker Trail Formation, and upper part of upper member of 

Hermosa Formation and "lower Cutler beds" are equivalent to Elephant Canyon Formation. References: Callville shelf (Brill, 

1963); Oquirrh basin (Tooker and Roberts, 1970; compilation of Hintze, 1988); northern Paradox basin (Baars, 1962; Mallory, 1972; Welsh and Bissell, 1979; compilation of Hintze, 1988; Baars, 1991); Wyoming shelf (Hansen, 1965; Welsh and 

Bissel, 1979; Driese and Dott, 1984; DeVoto and others, 1986; compilation of Hintze, 1988); Eagle basin (Mallory, 1972; DeVoto, 1980; DeVoto and others, 1986). Lower Molas Formation in the northern Paradox basin may be as old as latest Mississippian, following Merrill and Winar (1958) and suggestions of A.C. Huffman, Jr., and S.M. Condon (oral commun., 1991).

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CCS
m) of Pennsylvanian to Wolfcampian strata (Johnson and 

Johnson, 1991) and represents a relatively stable, slowly 

subsiding platform.

The northern part of the Paradox basin, which contains as 

much as about 2,400-2,700 m of Pennsylvanian and lower 

Wolfcampian strata (Mallory, 1972; Rascoe and Baars, 

1972), also is in the Uinta-Piceance region (fig. 2). The 

northern Paradox basin is bounded on the northeast by the 

Uncompahgre fault and the ancestral Uncompahgre uplift. 

Based on borehole, seismic, and structural data, Frahme and 

Vaughan (1983) and White and Jacobson (1983) showed that 

the Uncompahgre fault was a reverse fault in the late Paleozoic. Frahme and Vaughan (1983) estimated at least 9.6 km 

of horizontal displacement and about 6 km of vertical displacement along the northern segment of the Uncompahgre 

fault. The deepest part of the Paradox basin was adjacent to 

the Uncompahgre uplift (Mallory, 1972, 1975), and the 

basin's marked asymmetry is consistent with thrust-induced 

flexural loading along its northeastern margin. The 

Uncompahgre uplift most likely plunges and dies out to the 

northwest in a complex structural zone characterized by 

thrust and reverse faults (Szabo and Wengerd, 1975; Potter 

and others, 1991).

The Emery uplift or platform (fig. 2) formed a low-relief 

shelf on the west flank of the Paradox basin and may (at least 

in part) represent a flexural bulge or "forebulge" coupled to 

the overthrusted Uncompahgre uplift. Szabo and Wengerd 

(1975) suggested that the eastern margin of the Emery shelf 

is fault controlled (the Emery fault), and Potter and others 

(1991) described northwest-trending reverse faults on the 

northern flank of the shelf; however, the vertical displacement on these structures is probably not more than a few 

thousand feet. Permian strata of the northern Paradox basin 

postdate this postulated faulting and onlap the eastern flank 

of the Emery uplift. The northern Paradox basin is similar to 

the Eagle basin in that intrabasinal faults and lineaments had 

significant effects on depositional patterns and local rates of 

subsidence (for example, Szabo and Wengerd, 1975; Stevenson and Baars, 1986).

The Emery uplift is bounded on the west by a structurally 

complex zone characterized by Cretaceous thrust faults 

(Sevier orogeny) and Neogene extensional faults (Basin and 

Range). Pennsylvanian and Wolfcampian strata in this province are considered part of the Callville shelf depositional 

province (fig. 2). Callville shelf deposits probably onlapped 

the western flank of the low-relief Emery uplift, but this relationship is not clear because of the lack of outcrops and boreholes and because upper Paleozoic strata in this area 

apparently have been tectonically displaced 50-100 km eastward from their depositional site (Levy and Christie-Blick, 

1989, fig. 3).

The southeastern part of the Oquirrh basin is in the northwestern part of the Uinta-Piceance basin region (fig. 2) and 

contains as much as 7,100 m of Pennsylvanian and Wolfcampian strata (Tooker and Roberts, 1970). Basinal strata 

have been disrupted by Cretaceous thrust faults and Neogene 

extensional faults. Levy and Christie-Blick (1989, fig. 3) 

suggested net eastward tectonic transport of 30-100 km for 

Oquirrh basin strata in the Uinta-Piceance region. Alternatively, Jordan and Douglas (1980) suggested that, in the area 

west of the Wasatch fault (the easternmost fault in the extensional Basin and Range province), the younger extensional 

deformation may have fortuitously cancelled out older 

Sevier shortening such that late Paleozoic spatial relationships remain relatively intact. Jordan and Douglas (1980) 

suggested that the Oquirrh basin was a northwest-trending 

trough that had its maximum topographic relief in the Late 

Pennsylvanian and Early Permian. The bounding faults of 

this trough are apparently masked by younger sedimentary 

rocks and structural features and have not been positively 

identified. Picha (1986) suggested, however, that the Leamington lineament represents the southern boundary of the 

Oquirrh basin, and Jordan and Douglas (1980) and Bryant 

and Nichols (1986) stated that the "Oquirrh-Uinta arch" 

partly defined the northern margin of the deepest part of the 

basin. The nature of the boundaries between the Oquirrh 

basin and the Paradox basin and the Wyoming shelf to the 

southeast and east are obscured by structural complexities 

and (or) burial under thick Mesozoic and Cenozoic strata.

In summary, the Uinta-Piceance basin region is characterized by a complex mix of compressional and extensional 

tectonic styles from the Early Pennsylvanian through the 

Early Permian. This mix of structural styles is most consistent with regional strike-slip deformation (Christie-Blick 

and Biddle, 1985), probably driven by interactions along distant continental margins. The relative importance of each 

continental margin as a driving force for intraplate deformation is explored from the perspective of subsidence analysis 

later in this paper.

LATE PALEOZOIC CLIMATE, EUSTASY, 

AND CYCLICITY

The late Paleozoic was the time of the last major preQuaternary glaciation (Crowell, 1978). Initiation of this glacial age is generally correlated with the closure of several 

ocean basins and the resultant amalgamation of continental 

landmasses in the Southern Hemisphere to form the western 

half of Pangea (Scotese and McKerrow, 1990). During this 

period, North America was rotated clockwise about 30°-40° 

from its present orientation, and the Uinta-Piceance basin 

region was at a latitude ranging from about 0°-5° N. in the 

Early Pennsylvanian to about 15° N. in the Early Permian 

(Scotese and McKerrow, 1990). Extensive eolianites and 

evaporites in upper Paleozoic deposits of the Uinta-Piceance 

region (see later discussion) indicate an arid climate. PaleoCC6 Evolution of Sedimentary Basins Uinta and Piceance Basins
PROPOSED EUSTASY CURVE 

Fall -^ ^~ Rise

0 100 200 METERS

D) 

>

Figure 4. Inferred sea-level curve of Ross and Ross 

(1987) for the Pennsylvanian and Early Permian.

wind directions were primarily to the west (to the south in 

present-day coordinates) (Parrish andPeterson, 1988).

Using the Quaternary as an analogue, numerous authors 

have suggested that the late Paleozoic ice sheets expanded 

and contracted at fairly regular time intervals, probably 

forced by Milankovitch orbital parameters (for example, 

Heckel, 1986). As in the Quaternary, cyclic eustatic and climatic changes are the inferred effects of these ice-sheet fluctuations. Late Paleozoic sea-level curves (for example, Ross 

and Ross, 1987) (fig. 4) are currently being refined. Sealevel changes are generally interpreted as asymmetric (similar to Quaternary analogues), characterized by rapid transgressions (rapid melting of ice sheets) and slow regressions 

(slow buildup of ice sheets). Ross and Ross (1987) estimated 

that the magnitude of Pennsylvanian cyclic eustatic changes 

ranged from tens of meters to almost 200 m; Kerr (1989) 

suggested an average sea-level fluctuation of 90 m. Coastal 

onlap curves (Ross and Ross, 1988, fig. 9) indicate that 

eustatic fluctuations were superimposed on a significant global sea-level rise from the Morrowan to the end of the Desmoinesian; the magnitude of coastal onlap remained 

approximately the same in the Missourian, Virgilian, and 

Wolfcampian.

Deposition during late Paleozoic transgressions and 

regressions produced sedimentary "cycles" or depositional 

sequences that have been recognized and correlated across 

basins and globally (for example, Wanless and Shepard, 

1936; Heckel, 1986; Ross and Ross, 1987, 1988). Many of 

these depositional sequences are not "cyclic" in the strict 

sense in that preferred ordering of lithofacies may not be 

present (for example, Konopka and Dott, 1982; Driese and 

Dott, 1984). This lack of regular ordering within a transgressive-regressive sequence can reflect the variable effects of 

tectonic activity, sediment supply, basin geometry, and other 

factors. In this report, we use the term "cycle" in the more 

general way to describe a transgressive-regressive depositional sequence.

Heckel (1986) estimated that the periodicity of late 

Paleozoic sedimentary cycles or depositional sequences in 

the North American midcontinent outcrop belt was approximately 235,000-400,000 years, matching the Milankovitch 

eccentricity parameter; however, based on a synthesis of 19 

studies of late Paleozoic cycles, Kerr (1989) estimated a typical periodicity of 1.3 million years, significantly longer than 

inferred late Cenozoic eccentricity, and suggested that a 

longer Milankovitch parameter may have been active in the 

late Paleozoic. Ross and Ross (1988) suggested that late 

Paleozoic transgressive-regressive sequences averaged 

about 2 million years (range of 1.2-4 m.y.) but also noted 

that these sequences could include four or five of Heckel's 

(1986) cyclothems. Klein (1990) however, maintained that 

use of a poorly calibrated time scale by Heckel (1986), and 

by inference, Ross and Ross (1988) and Kerr (1989) resulted 

in an overestimation (by a factor of 1.3-1.8) of cycle or

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC7
sequence periodicity. Klein (1990) challenged the Milankovitch hypothesis and suggested that eustatic changes may 

have been controlled by other processes including tectonism 

and large-scale meltwater discharge. Klein (1990) did not 

discuss a mechanism for repetition of these inferred alternative cycle-controlling mechanisms. Until such a mechanism 

has been described and justified, Milankovitch orbital 

parameters of unknown periodicity provide the most compelling rationale for the repetitive late Paleozoic eustatic and 

climatic changes. Figure 3 shows both the time scale of Haq 

and Van Eysinga (1987) (which is similar to the scale used 

by Heckel, 1986) and that proposed by Klein (1990). The 

Klein (1990) calibrations have been challenged by Langenheim (1991) and Heckel (1991).

PALEOGEOGRAPHY OF THE UINTAPICEANCE BASIN REGION

Late Paleozoic transgressive-regressive cycles or 

sequences, primarily reflecting the eustatic controls discussed above, have been reported from most of the welldescribed stratigraphic units in the Uinta-Piceance basin 

region (Kite and Buckner, 1981; Driese and Dott, 1984; 

Johnson and others, 1988; Schenk, 1989; Karachewski, 

1992) (fig. 5). Because much of the basinal topography in the 

Uinta-Piceance region was at or near sea level in the late 

Paleozoic, inferred eustatic fluctuations significantly 

affected the location of the marine-nonmarine interface and 

other depositional patterns. We attempt to show the changes 

in these depositional patterns on four sets of maps (figs. 6-9) 

representing deposition during maximum transgression and 

regression. In one case (late Atokan-Desmoinesian regressions map; fig 7A), we also show the limits of an evaporite 

facies in the Eagle basin that was deposited during but not at 

the peak of regressions because of the facies1 importance in 

understanding regional depositional patterns. By presenting 

paired maps, our report differs from previous efforts at 

regional paleogeographic reconstruction (for example, Hallgarth 1967; Mallory, 1972, 1975; Rascoe and Baars, 1972; 

Chronic, 1979; Welsh and Bissell, 1979; DeVoto and others, 

1986), which generally show the dominant lithofacies or 

depositional environment in a particular area. This dominant 

lithofacies or environment may represent deposition during 

a transgression, a regression, or an intermediate stage.

The time intervals selected for the maps (Morrowan and 

early Atokan, late Atokan and Desmoinesian, Missourian 

and Virgilian, early Wolfcampian) coincide with inferred 

major changes in regional tectonic and (or) depositional 

patterns. We emphasize that the maps are preliminary models, inhibited by the general lack of outcrop of late Paleozoic 

rocks (fig. 1) and (or) the lack of detailed sedimentologic 

studies. We relied extensively on geophysical or lithologic 

(AmStrat) borehole logs (fig. 1, appendix) for reconstruction

of paleogeography in northwesternmost Colorado and the 

Paradox basin-Emery uplift area; however, these borehole 

logs generally lack well-constrained age data and systematic 

lithologic descriptions. We show deposition on the maps in 

areas where Pennsylvanian rocks are no longer present if 

there is compelling evidence that the rocks were originally 

deposited and later eroded. Essentially no data are available 

from the subsurface of the Paleogene Uinta basin, and the 

inferred facies patterns in this important area where the Paradox basin merges with the Wyoming shelf and the Oquirrh 

basin are particularly speculative. Finally, due to gross 

uncertainties (Jordan and Douglas, 1980; Levy and ChristieBlick, 1989), we did not attempt to palinspastically restore 

strata of the Oquirrh basin and Callville shelf to their depositional sites (pre-Mesozoic shortening and Cenozoic extension). Future studies will undoubtedly refine estimates of the 

amounts of tectonic transport and make accurate palinspastic 

restorations possible.

Pennsylvanian and Early Permian subsidence in four 

depositional provinces in the Uinta-Piceance basin region is 

shown in figures 10 and 11. These diagrams are based on 

considerable imprecise data involving the ages of stratigraphic units and their boundaries, the diagenetic and compaction histories of most units, and calibration of the 

geologic time scale. Additionally, thicknesses vary significantly for different units within the four sedimentary provinces. The diagrams should therefore be regarded as 

approximations that illustrate general trends. The ages of 

units used in the analysis are shown on figure 3. The time 

scale of Haq and Van Eysinga (1985) was used for all analyses; however, figure 10A also shows the data plotted using 

the time scale proposed by Klein (1990). Use of different 

time scales modifies the shapes of the curves but not their 

general trends.

Sub-Pennsylvanian Geology

Morrowan strata unconformably overlie Mississippian 

rocks throughout most of the Uinta-Piceance region (Mallory, 1972, p. 114; Johnson and Johnson, 1991). In the Eagle 

basin and most of northwestern Colorado (fig. 4), Morrowan 

strata generally overlie the Lower Mississippian Leadville or 

Madison Limestone. Pennsylvanian strata (including the latest Mississippian (?) to early Middle Pennsylvanian Molas 

Formation) overlie the Leadville Limestone, the Doughnut

Figure 5 (facing page). Schematic Pennsylvanian depositional sequences in the Uinta-Piceance basin region. Examples 

from the Minturn and Maroon Formations and the Eagle Valley 

Evaporite are modified from Johnson and others (1988), 

Schenk (1989), and Karachewski (1992), from the Morgan Formation modified from Driese and Dott (1984, fig. 20), and from 

the Paradox Formation modified from Hiteand Buckner (1981, 

fig. 7).

CCS Evolution of Sedimentary Basins Uinta and Piceance Basins
Sea level 

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Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC9
Formation, or the Meramecian Redwall Limestone in the 

northern Paradox basin area, and the Redwall Limestone 

underlies Wolfcampian strata on the Emery uplift (Mallory, 

1972; Rascoe and Baars, 1972; Welsh and Bissell, 1979).

The contact between the Upper Mississippian Doughnut 

Formation and the Lower Pennsylvanian Round Valley 

Limestone in the Wyoming shelf area is problematic. Welsh 

and Bissell (1979, p. Y21) suggested that "sedimentation 

was continuous or only slightly interrupted from Chesterian 

(latest Mississippian) to Morrowan time. The Round Valley 

Limestone***represents a marine invasion over the deltaic 

and estuarine environments of the Chesterian Doughnut Formation." Near the Utah-Colorado border, however, some 

post-Doughnut, pre-Morrowan erosion clearly occurred 

because the Morgan Formation (in part laterally equivalent 

to the Round Valley) rests unconformably on the Madison 

Limestone and the Doughnut Formation is missing (for 

example, borehole 39, fig. 1). This unconformity may pass 

westward into a conformity along the Wyoming shelf. To the 

west, in the Oquirrh basin and on the Callville shelf, Morrowan strata are inferred to conformably overlie Chesterian 

rocks (Brill, 1963; Tooker and Roberts, 1970). Isopach maps 

(Poole, 1974; Poole and Sandberg, 1977; Skipp, 1979) indicate that the Oquirrh basin and Wyoming shelf began to subside as discrete tectonic elements in the Late Mississippian, 

significantly before subsidence began in the Eagle and 

Paradox basins.

Early to Early Middle Pennsylvanian 

Morrowan and Early Atokan

Within the Pennsylvanian and Wolfcampian strata of the 

Uinta-Piceance region, depositional cycles or sequences are 

least recognized in the generally thin sections of Morrowan 

and lower Atokan strata (fig. 6). This relative lack of recognition probably reflects several factors, including poor outcrops, little recent detailed sedimentologic analysis, slow 

subsidence rates (figs. 10, 11), and eustatic fluctuations of 

comparatively lesser magnitude. Ross and Ross (1987, 

1988) showed 10 significant eustatic fluctuations in the Morrowan and lower half of the Atokan that had inferred magnitudes of from 20 to about 80 m, about half the size of the 

magnitude inferred for late Atokan through Early Permian 

eustatic fluctuations (fig. 4).

Brill (1963) described cycles from a 265-m-thick section 

of the Callville Limestone on the Callville shelf (Route 148 

southwest section, fig. 1). These cycles consist of a lower 

fine-grained sandstone overlain by an upper thin-bedded 

limestone. Farther south in the Cove Fort area (fig. 1), Morrowan strata (255 m thick) are mainly dolomitic limestone 

but include some sandstone and oolitic limestone. Brill 

(1963) suggested that cyclic sedimentation in this area 

reflects repeated transgression (carbonate platform 

deposition) and regression (shelf sand deposition) over an

unstable border between a shelf to the east and a basin to the 

west.

Brill (1963) noted only a thin Atokan sequence of the 

Callville Limestone in the Callville shelf area, overlain by an 

unconformity. Welsh (1976) argued that Desmoinesian 

strata were deposited across the Callville shelf but subsequently eroded. Brill (1963) described cycles from the Atokan part of the Callville from the Route 148 southwest 

section (fig. 1) that consist of a basal gray to brown shale 

overlain by a middle unit of gray to yellowish-gray sandstone and an upper unit of thin-bedded limestone or 

dolomite. Similar sandstone is crossbedded in the Cove Fort 

section (fig. 1). Brill did not describe the limestone, but it 

probably represents deeper facies (transgressive deposits) 

deposited at the beginnings of cycles, whereas the sandstone 

forms the upper parts of cycles and was deposited during 

regressions.

In the Oquirrh Mountains (fig. 1) of the southeastern 

Oquirrh basin, the Lake Point Limestone and the West Canyon Limestone include about 365-440 m of mainly Morrowan strata. Tooker and Roberts (1970, p. A12) described 

cycles in the Lake Point Limestone that comprise, in ascending order, shaly limestone, clastic calcareous sandstone, calcitic arenaceous limestone, crossbedded arenaceous 

limestone, cherty limestone, and bedded chert. They suggested (p. A24) that the West Canyon Limestone, comprising arenaceous limestone, cherty argillaceous limestone, and 

crossbedded sandstone, is also characterized by cyclic deposition but did not describe a typical cycle. These cycles 

resemble those described for the Callville shelf and have an 

inferred similar origin. Alexander (1978) measured and provided detailed petrographic descriptions of a section 321 m 

thick of similar age and lithology at Cascade Mountain in the 

central Wasatch range. He did not identify cycles but interpreted a variety of shallow-marine depositional environments that changed in response to variations in sea level.

Tooker and Roberts (1970) and Konopka and Dott (1982) 

described a significant change in lithology and depositional 

style in the southeastern Oquirrh basin between Morrowan 

strata and the more sandstone-rich Atokan to Desmoinesian 

Butterfield Peaks and Erda Formations. The timing of this 

change is similar to that noted for the Callville shelf. These 

changes are shown on the late Atokan to early Desmoinesian 

maps (fig. 7) but were probably initiated in the early to middle Atokan.

The Emery uplift, on which Permian rocks unconformably overlie Mississippian strata, bounds the Callville shelf 

to the east. Strata on the Callville shelf are relatively fine 

grained, and it is possible that Morrowan and lower Atokan 

strata were deposited over at least the westernmost part of 

the Emery uplift (as shown in fig. 6) and subsequently 

eroded. The area that now includes the eastern part of the 

Emery uplift, the Paradox basin, and at least part of the 

ancestral Uncompahgre uplift was probably emergent or 

partly emergent. Latest Mississippian (?) through early

CC10 Evolution of Sedimentary Basins Uinta and Piceance Basins
Atokan deposition in the Paradox basin is recorded by the 

Molas Formation, mainly a red regolith that formed by 

weathering of Mississippian limestone. The emergent area is 

here informally named the "Molas platform."

Morrowan and lower Atokan strata of the Wyoming shelf 

depositional province are generally less than about 150 m 

thick (Johnson and Johnson, 1991) and include the Round 

Valley Limestone and the lowest part of the Morgan Formation. The Round Valley Limestone crops out along the south 

and north flanks of the Uinta Mountains and extends east as 

far as Juniper Mountain (fig. 1) (Upton, 1958). It consists 

mainly of thin-bedded to massive limestone, commonly 

dolomitic or cherty (Sadlick, 1955, 1957; Hansen, 1965). 

Siliciclastic deposits, generally gray or purplish-red mudstone, are typically a minor component but locally (for 

example, Sols Canyon, fig. 1) comprise as much as 35 percent of Round Valley sections. Huddle and McCann (1947) 

noted a significant clastic component to the Round Valley 

Limestone (mapped as the lower member of the Morgan Formation) in the Duchesne River area (fig. 1).

The proportion of clastic material in Morrowan and 

lower Atokan strata (or strata that overlie Mississippian 

rocks and are of inferred Morrowan and early Atokan age) 

increases in the subsurface south and southeast of the Uinta 

Mountains outcrop belt (for example, boreholes 34-38, fig. 

1), in strata assigned to the Morgan Formation. This increase 

suggests that the "Molas platform" contributed clastic detritus to the Wyoming shelf. The more abundant limestone in 

this province was probably deposited during transgressions, 

when clastic detritus was trapped closer to its source. 

Regressions were probably characterized by lowering of 

base level and more widespread clastic deposition.

The Morrowan and lower Atokan Belden Formation (as 

thick as 200 m or more) of the Eagle basin represents deltaic 

and prodeltaic deposition off the ancestral Front Range. The 

best exposures of cyclic deltaic deposits are in the Sweetwater Creek area about 30 km northeast of Glenwood Springs 

(fig. 1). Typical cycles coarsen upward from (1) prodeltaic 

deposits of gray to black shale, dolomite, and limestone to 

(2) delta-front deposits of wave-rippled, fine- to coarsegrained sandstone and interbedded dark mudstone. Some 

cycles are capped by fluvial channel deposits of trough 

crossbedded, coarse-grained to conglomeratic sandstone. 

Paleocurrent data (n=60, X=237°, a=41°) indicate that sediment transport was to the southwest. The texture and sedimentology of these deltaic rocks indicate contemporaneous 

uplift in the ancestral Front Range. Houck (1991) noted a 

large delta system in the overlying Minturn Formation in the 

McCoy area (fig. 1), in the same location as the large Belden 

delta shown in figure 6A. She inferred that the large Minturn 

delta was fed by a major drainage system channeled through 

a west-trending transfer zone in the Gore fault, the interpretation shown on figure 6.

Throughout most of Eagle basin, the Belden Formation 

consists of interbedded gray limestone and dark-gray,

commonly calcareous, mudstone. These mudstones are the 

inferred regressive deposits and represent maximum progradation of deltas and associated prodelta clastic sediments. 

During transgressions, clastic sediments were trapped closer 

to uplifted sources and limestones were deposited over 

broader areas. Beds of anhydrite and (or) gypsum in the 

Belden Formation have been found in the subsurface on the 

north flank of the White River uplift (for example, boreholes 

41-43, fig. 1), and in outcrops on the west side of the 

Sawatch uplift (Freeman, 1972; Bryant, 1979). Interbedded 

facies (shale and limestone) at each location suggest a basinal setting. We infer that these evaporites precipitated from 

basinal brines in fault-bounded sub-basins (following 

DeVoto and others, 1986; Dodge and Bartleson, 1986) during eustatic lowstands.

DeVoto and others (1986) presented evidence from an 

area southeast of the Uinta-Piceance region that major 

coarse-grained delta systems in the Belden Formation had a 

source in the Sawatch uplift. The lack of comparable deltaic 

deposits on the west side of the Sawatch uplift (Bryant, 

1979) in the Uinta-Piceance region suggests either that the 

uplift had markedly asymmetric relief or that it was significantly narrower than the modern Sawatch Range.

Based on the dominance of limestone and dark mudstone 

in the Belden Formation in the Crested Butte area (fig. 1), 

Langenheim (1978) concluded that the ancestral Uncompahgre uplift had not developed in the Morrowan and early Atokan. The "Molas platform" apparently did not shed 

significant sediment eastward into the Eagle basin. The eastern boundary of this low-relief source terrane was probably 

west of the Middle and Late Pennsylvanian margin of the 

Uncompahgre uplift.

Due to a lack of subsurface data, the location of the 

Morrowan-early Atokan structural boundary between the 

Front Range and the Eagle basin is not well constrained. 

North and northwest of the McCoy area (fig. 1), coarsegrained deltaic deposits have not been recognized adjacent to 

the basin margin, either in outcrops (for example, Sharps, 

1962) or in the subsurface (for example, Thomas and others, 

1945; Hallgarth, 1959) (boreholes 44, 45, 47, 51, fig. 1). 

Morrowan and lower Atokan strata are probably not present 

in borehole 53 (fig. 1), where a thin (120 m) Pennsylvanian 

section rests on Mississippian rocks. It may be that Morrowan and lower Atokan rocks were never deposited east of 

borehole 53, as suggested by DeVoto and others (1986, fig. 

4). Alternatively, Morrowan and lower Atokan rocks could 

have been deposited east of borehole 53 and then eroded 

later in the Pennsylvanian.

Summarizing, Morrowan and lower Atokan strata 

across most of the Uinta-Piceance basin area consist of 

alternating fine-grained clastic rocks (inferred regressive 

deposits) and more abundant limestone (transgressive 

deposits). Significant Morrowan and early Atokan tectonic 

uplifts include the ancestral Front Range and Sawatch 

uplifts.

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC11
CC12 Evolution of Sedimentary Basins Uinta and Piceance Basins
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Middle and Late Middle Pennsylvanian 

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Explanation for figures 6 through 9.

The late Atokan and Desmoinesian paleogeographic 

maps (fig. 7) highlight three major changes that occurred 

during the early and (or) middle part of the Atokan in the 

Uinta-Piceance region: (1) uplift of the ancestral Uncompahgre uplift and the Emery uplift; (2) subsidence of the Paradox 

basin; and (3) significant progradation of clastic sediments 

across the Wyoming shelf. Ross and Ross (1987, 1988) 

recognized 24 major eustatic fluctuations during the last half 

of the Atokan (2) and Desmoinesian (22) having inferred 

magnitudes of as much as 200 m (fig. 4). The sea-level curve 

of Heckel (1986) is quite similar.

As described earlier, upper Atokan and Desmoinesian 

strata probably have been eroded from the Callville shelf 

(Welsh, 1976). Accordingly, figure 7 shows the early Atokan depositional pattern (described above) as continuing into 

the late Atokan and Desmoinesian. Brill (1963) suggested 

that the increased clastic detritus on the Callville shelf was 

derived from both the Uncompahgre and Emery uplifts. The 

Emery uplift was emergent, but preservation of Mississippian rocks below a Permian unconformity indicates that this 

uplift had minimal relief, probably was not a major sediment 

source, and may have been onlapped significantly during 

maximum transgressions.

The Butterfield Peaks and Erda Formations represent 

Atokan and Desmoinesian deposition in the Oquirrh basin 

(Tooker and Roberts, 1970). From detailed sedimentologic 

studies in the Mount Timpanagos area (fig. 1) on the upper 

plate of the Charleston-Nebo thrust and in the southeastern 

Oquirrh basin, Konopka and Dott (1982) and Konopka 

(1985) described lithofacies deposited in environments ranging from normal-marine carbonate shelf to eolian dunes, 

inferred to represent deposition during maximum transgressions and regressions, respectively. The statistical analysis 

presented in these studies suggests that these rocks are characterized by several preferred vertical facies transitions but 

were not cyclic in the strict sense. Suggested factors that 

might have obscured better development of cycles include 

(1) rapid, possibly episodic subsidence, (2) rapid deposition 

of windblown sand and silt, and (3) rapid deposition of carbonate rocks in environments not overwhelmed by sand 

(Konopka and Dott, 1982). Eolian deposits are much less 

common in the Oquirrh Group west of the Wasatch fault than 

they are at Mount Timpanagos. Shoal and emergent carbonate deposits are present at several locations west of the 

Wasatch fault. These deposits, along with too numerous 

alternations of sandy and carbonate facies, show that there is 

not a one-to-one correspondence between lithology and 

transgressive-regressive depositional patterns.

Tooker and Roberts (1970), Morris and others (1977), 

and Konopka (1985) described similar lithofacies in the 

Oquirrh Group in the Oquirrh Mountains and at several other

CC14 Evolution of Sedimentary Basins Uinta and Piceance Basins
locations west of the Wasatch fault. A number of authors (for 

example, Nygreen, 1958; Bissell, 1962; Roberts and others, 

1965; Tooker and Roberts, 1970) suggested that these strata 

are cyclic, mainly on the basis of the obvious alternation of 

siliciclastic and carbonate sediments. Stevens and Armin 

(1983) described similar Middle Pennsylvanian rocks immediately west of the Uinta-Piceance basin region. They 

described an idealized transgressive-regressive facies 

sequence within the carbonate strata, bounded by siliciclastic units that represent the most regressive facies.

Based on extensive paleocurrent measurements, 

Konopka and Dott (1982) and Konopka (1985) reported that 

sand transport in the Oquirrh basin during the Atokan and 

Desmoinesian was to the south. In addition, sandstone 

deposition was slightly more dominant in the northern and 

eastern parts of the basin than in the southern and western 

parts. The sands probably were derived from the craton to 

the north and northeast and transported to the Oquirrh basin 

primarily via the Wyoming shelf, and significant sand input 

from either the Emery or Uncompahgre uplift seems 

unlikely. Mean composition of siliciclastic grains (2-3 phi) 

is within the quartz-rich end of the subarkose field, a mature 

composition for this grain size (Odom and others, 1976). Jordan (1979) and Jordan and Douglas (1980), however, 

pointed out that, because the Oquirrh basin subarkosic sands 

(Late Pennsylvanian and Early Permian) and Pennsylvanian 

subarkosic sands derived from crystalline rocks in the Front 

Range uplift are compositionally similar, the source of sands 

in the Oquirrh basin may have been the crystalline rocks of 

the nearby Uncompahgre uplift. Welsh and Bissell (1979, 

fig. 8) inferred that a delta system derived from the 

Uncompahgre uplift fed the Oquirrh basin during the Middle 

and Late Pennsylvanian, but their hypothesis is difficult to 

assess because neither outcrops nor subsurface data are 

available in the critical area between the Uncompahgre uplift 

and the Oquirrh basin and fluvial-deltaic strata have not been 

recognized in the Oquirrh basin. Potter and others (1991) 

documented the presence of a complex structural zone on the 

northwest flank of the Uncompahgre uplift that might have 

served to channel Uncompahgre detritus to the northwest. If 

Uncompahgre-derived sand did enter the southern part of the 

Oquirrh basin (as shown in fig. 1A), then it probably was 

reworked and transported south onto the Callville shelf.

Driese and Dott (1984) described sedimentary cycles in 

the upper Atokan and Desmoinesian upper member of the 

Morgan Formation of the Wyoming shelf depositional province (fig. 5). These cycles typically pass upward from (1) 

carbonate rocks deposited in normal-marine shelf environments (maximum transgression deposits), to (2) carbonate 

rocks deposited in restricted-shelf and nearshore environments, to (3) sandstones deposited in eolian dunes (maximum regression deposits). Sands were derived from the 

north, across the Wyoming shelf.

Transgressive-regressive depositional patterns on the 

Wyoming shelf and in the Oquirrh basin (and by inference,

on the Callville shelf) were thus fairly similar. AtokanDesmoinesian strata on the Wyoming shelf (Morgan Formation and lower part of Weber Sandstone) have a maximum 

thickness of about 460 m, whereas correlative strata in the 

Oquirrh basin (the Butterfield Peaks Formation) are as thick 

as 2,765 m (Tooker and Roberts, 1970). Because sediment 

supply and subsidence were in balance in the Oquirrh basin, 

the basin was not characterized by significant negative topographic relief at this time. If these two areas experienced the 

same number of transgressive-regressive events during this 

period, then the Oquirrh basin depositional cycles should be 

much thicker than those that formed on the Wyoming shelf.

The Paradox basin formed as a discrete tectonic element 

in the Atokan by fragmentation of the "Molas platform." The 

low-relief Emery uplift (discussed above) and the 

Uncompahgre uplift (of inferred high relief) formed the 

western and east-northeastern margins of the Paradox basin, 

respectively. Erosion of Paleozoic cover rocks from the 

Uncompahgre uplift exposed Precambrian crystalline rocks, 

considered to be the main source of clastic detritus in the 

Paradox basin. As described earlier, seismic data (Frahme 

and Vaughan, 1983) show that the late Paleozoic boundary 

between the northern Paradox basin and the Uncompahgre 

uplift was a reverse fault, and basin asymmetry is inferred to 

reflect flexural loading along the northeastern margin of the 

basin.

Mallory (1972, figs. 6, 7) showed a maximum thickness 

of about 1,700 m for the Atokan and Desmoinesian in the 

northern Paradox basin, including strata of the Pinkerton 

Trail and Paradox Formations of the Hermosa Group and 

undivided rocks of the Hermosa Formation. The Atokan to 

lower Desmoinesian Pinkerton Trail Formation is the lowest 

unit and consists of mixed carbonate and fine-grained clastic 

rocks. The Pinkerton Trail is thickest on the western margin 

of the basin and thins toward the center of the basin, where 

it intertongues with and is overlain by the Paradox Formation, which consists of evaporites, black shale, and limestone 

(Hite and Liming, 1972). The undivided rocks of the Hermosa Formation crop out on the northeastern margin of the 

basin and are coarse-grained clastic rocks derived from the 

Uncompahgre uplift.

Atokan and Desmoinesian sedimentary cycles in the Paradox basin are well described (fig. 5) (Peterson and Hite, 

1969; Hite, 1970; Hite and Buckner, 1981). Hite and Liming 

(1972, p. 134-135) identified 29 evaporite cycles on a Paradox basin cross section that includes boreholes 7 and 13 (fig. 

1). These basinal evaporite cycles ideally consist of highly 

calcareous, organic-rich black shale (inferred maximum 

transgression deposits), anhydrite, dolomite and silty dolomite (intermediate sea-level stand), and halite and potash salt 

deposits (maximum regression). Some halite beds exhibit 

evidence of subaerial exposure and erosion. Hite and Cater 

(1972) showed the limits of evaporite occurrence. On the 

western margin of the basin (for example, boreholes 7 and 

13, fig. 1), Hite and Buckner (1981) described cycles that

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC15
I 

i

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o. 

 o 

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112 109° 108°

MINTUR

FORMATI

MORGAN 

FORMATION

WEBER 

SANDSTONE

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A

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IHERMOSA 

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UNDIFFERENTI

PARADOX 

FORMATIO

50 100 KILOMETERS
109° 108° 107°

no_o

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o

,-v , __ ., FRONT RANGE

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FORMATION SANDSTONE! T ..j , __T.. .,f

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FORMATION EAGLE VALLEY

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ERMOSA 

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PARADOX 

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Figure 7 (above and facing page). Maps showing the late Atokan to Desmoinesian paleogeography of the Uinta-Piceance basin region during maximum regression (A) and 

transgression (E). See explanation on page CC14.
ideally consist of calcareous black shale (inferred maximum 

transgression deposits), dolomite and siltstone (intermediate 

sea level), and algal and fossiliferous limestone (maximum 

regression). Strata north of the evaporite limit in the northernmost Paradox basin (for example, boreholes 16, 18, 26, 

fig. 1) consist mainly of interbedded fossiliferous limestone 

(inferred transgressive deposits), and variegated finegrained clastic deposits of inferred sabkha or playa origin 

(inferred regression deposits).

It should be noted that the basinal and western-basinmargin cycles of the Paradox described above are anomalous 

in comparison to other regionally correlative cycles in that 

the most clastic rich intervals have been assigned to transgressive phases (that is, raised base level), whereas clastic 

sediments elsewhere are inferred to have been trapped in 

basin-margin settings. Kite and Buckner (1981) suggested 

that the clastic component in the Paradox basin reflects 

reworking of basin-margin clastic sediments by transgressing seas. Perhaps a more suitable explanation involves sediment accumulation rates. The thick halite intervals in the 

cycles may have deposited at rates of 5 cm per year or more 

(O.B. Raup, U.S. Geological Survey, oral commun., 1991), 

whereas the highly calcareous, organic-rich, black shales 

probably accumulated at rates of a few millimeters per year 

(Kite and Buckner, 1981) or much less. Given these rates, it 

may be that input of clastic sediment to the basin actually 

increased during regressions, but that this increase is masked 

by the great volume of the rapidly accumulating evaporites.

Kite and Buckner (1981) postulated that the evaporites 

were deposited subaqueously in a deep basin through precipitation from a concentrated brine, an interpretation that 

requires a basin with restricted circulation. The Emery uplift 

was emergent and provided a partial barrier to circulation. 

Szabo and Wengerd (1975, fig. 11) showed several possible 

inlets to the Paradox basin, including (1) the "Oquirrh sag," 

a presumed connection with the Oquirrh basin in the structurally complex zone between the northwestern end of the 

Uncompahgre uplift and the northern Emery uplift (Potter 

and others, 1991), and (2) the "Fremont sag," a possible eastwest connection between the Paradox basin and the Callville 

shelf, to the south of the Uinta-Piceance region. Given the 

nonmarine to marginal-marine regressive facies described 

above for the Oquirrh basin and Callville shelf, it is unlikely 

that either inlet was active during maximum regressions. The 

presence of fossiliferous limestone in the northernmost Paradox basin, however, requires normal or near-normal marine 

waters in the basin during transgressions, and this water must 

have entered the basin without mixing with the evaporitic 

brines. This nonmixing requirement is more easily met for 

the Oquirrh sag, and we therefore show it as an inlet into the 

Paradox basin during maximum transgressions. Boreholes in 

the area of the Oquirrh sag (for example, boreholes 24, 25, 

27, 28, fig. 1) contain anomalously thin sections (less than 

220 m) of inferred but undated Pennsylvanian strata and do 

not provide sufficient data to test the hypothesis.

Tectonic activity was a major control on evaporite deposition in the Paradox basin. Based on apparent rates of subsidence in the Paradox basin (figs. 10, 11), the Uncompahgre 

uplift probably experienced its highest rates of overthrusting, and extended farthest to the northwest, during the Desmoinesian. Decreased activity on the Uncompahgre would 

have resulted in decreased sedimentation rates (and accommodation space for evaporites), greater progradation of clastic sediments from basin margins, and widening of the 

Oquirrh sag, all effects detrimental to evaporite deposition 

and preservation.

The boundary between the Eagle basin and Wyoming 

shelf depositional provinces is gradational north of the 

Uncompahgre uplift near the Colorado-Utah State line. In 

general, varicolored clastic rocks become more fine grained 

and are interbedded with greater proportions of limestone 

with increasing distance from the Uncompahgre uplift (boreholes 33-39, fig. 1). Anomalous limestone-poor strata in 

borehole 35 indicate lateral lithologic variability and the possible presence of a north-flowing fluvial-deltaic system. 

During regressions, alluvial clastic sediments derived from 

the Uncompahgre probably merged northward with southprograding eolian deposits. The greater width of the belt of 

Uncompahgre-derived clastic sediments north of the uplift in 

Utah than in the Paradox and Eagle basins probably reflects 

the lower subsidence rates characteristic of the Wyoming 

shelf depositional province. In more rapidly subsiding areas, 

Uncompahgre-derived clastic sediments are inferred to have 

been trapped closer to basin margins.

During regressions, alluvial clastic sediments derived 

from the northern Front Range uplift probably interfingered 

with south-prograding dunes in the northeastern Eagle basin, 

but there is little subsurface data to document this relationship. The southern limit of the Wyoming shelf dune field 

probably was in the Juniper Mountain area (fig. 1). Thick 

beds of white and red sandstone, presumably eolian, are 

present in the upper part of the Morgan Formation in boreholes 48 and 51 (fig. 1) and at Juniper Mountain (Upton, 

1958). Thick beds of well-sorted sandstone are almost absent 

farther southeast in boreholes 47 and 49.

Upper Atokan and Desmoinesian strata of the Eagle 

basin include the Minturn and Gothic Formations on the 

eastern and western margins of the basin, respectively, and 

the Eagle Valley Evaporite in the central part of the basin. 

Karachewski (1992) described highly variable shoalingupward sequences (fig. 5) in the Minturn Formation in the 

Minturn area (fig. 1). Maximum regression deposits are 

coarse alluvial and fan-delta clastic rocks; maximum transgression deposits are offshore marine mudstone and biostromal limestone. During transgressions, limestone was 

deposited within a few kilometers of the Gore fault 

(Houck, 1991; Karachewski, 1992) (borehole 46, fig. 1). 

Karachewski (1992) inferred that relief was significant on 

the Front Range uplift, the source for the Minturn Formation, and minor on the Sawatch uplift, which he considered

CC18 Evolution of Sedimentary Basins Uinta and Piceance Basins
an influence on depositional patterns but not a significant 

sediment source. Houck (1991) inferred that a large deltaic 

system in the McCoy area (fig. L) was channelled through 

a west-trending transfer zone in the Gore fault.

Leighton (1986) and Levorson (1986) reported shoalingupward sequences similar to those in the Minturn Formation 

from the Gothic Formation in the Crested Butte and Crystal 

River areas (fig. 1), respectively. Each concluded that both 

the Uncompahgre and Sawatch uplifts were significant sediment source terranes and that relief was greater on the 

Sawatch. Leighton inferred that the main input of marine 

waters into the "Gothic trough" was from the south.

In the central part of Eagle basin (Eagle Valley location 

of fig. 1), Schenk (1989) described shoaling-upward cycles 

in the Eagle Valley Evaporite (fig. 5). Micritic carbonates 

are inferred maximum transgression deposits. Because Desmoinesian subsidence rates were greater in the Eagle basin 

than on the Wyoming shelf, sea water was apparently 

trapped in the Eagle basin during regressions and gypsum 

was precipitated from an evaporative brine for much of the 

duration of the regressive phase. Although gypsum is widespread over the central part of Eagle basin, halite is present 

in only a few discrete areas (for example, DeVoto and others, 

1986, p. 43; Dodge and Bartleson, 1986, p. 117). These areas 

of thick halite deposition have been interpreted as topographically low areas produced by syndepositional faulting 

(DeVoto and others, 1986; Dodge and Bartleson, 1986). 

During maximum regressions, shallow marine to nonmarine 

clastic sediments prograded over evaporitic deposits, and 

access to normal marine water was completely cut off to the 

north by Morgan eolianites. Consistent with the inferred 

paleogeography, the lack of a marine fauna in clastic rocks 

of the Eagle Valley Evaporite suggests that lowstands 

(regressive phases) were characterized by lacustrinelike circulation (Schenk, 1989). Schenk (1989) noted that the geometry and facies of clastic wedges at the tops of cycles in the 

Eagle Valley Evaporite indicate progradation from both the 

Uncompahgre and Front Range uplifts.

The location of the northern part of the Gore fault zone 

may have moved eastward from its Morrowan to early Atokan position. The very thin Pennsylvanian section in boreholes 52 and 53 (fig. 1), closest to the Front Range uplift, is 

consistent with a complex history of basin-margin faulting.

Summarizing, the late Atokan to Desmoinesian history of 

the Uinta-Piceance region is characterized by increased tectonic activity reflected in higher rates of basin subsidence 

and in continued or initial uplift and unroofing of the Front 

Range, Sawatch, and Uncompahgre uplifts. The combined 

effects of tectonism and eustasy led to restricted circulation 

in the Eagle and Paradox basins in which evaporite deposition was significant. Deposition during regressions was also 

characterized by significant progradation of eolian sands 

across the Wyoming shelf. Deposition of limestone dominated during transgressions.

Late Pennsylvanian Missourian and Virgilian

The major uplifts and basins of the Uinta-Piceance region 

continued to be active during the Missourian and Virgilian 

(fig. 8). Evaporite sedimentation essentially ended in the 

Paradox and Eagle basins, probably in response (1) a general 

rise in sea level (Ross and Ross, 1987,1988), leading to less 

restricted circulation, (2) decreasing rates of subsidence in 

evaporite basins (figs. 10, 11), leading to increased progradation of clastic sediments derived from basin-margin 

uplifts, and (3) changes in tectonic geomorphology, leading 

to less topographic restriction of basins. Ross and Ross 

(1987) suggested there were 17 major eustatic fluctuations 

during the Missourian and Virgilian, with magnitudes ranging from about 60 to 180 m (fig. 4). Alternatively, Heckel 

(1986) recognized evidence for 36 major eustatic fluctuations in the North American midcontinent during the Missourian and the lower part of the Virgilian.

Missourian-Virgilian rocks are not preserved on the 

Emery uplift, and Brill (1963) noted no Missourian and only 

a thin (<100 m) section of Virgilian limestone on the northern Callville shelf. Brill (1963) suggested that these areas 

were characterized by minor uplift during most of this interval, whereas Welsh and Bissell (1979) suggested that Missourian and Virgilian strata may have been deposited and 

then subsequently removed during a latest Pennsylvanian 

erosional event. Neither hypothesis suggests more than 

minor subaerial or subaqueous relief in these areas, perhaps 

close to the inferred magnitude of eustatic fluctuations. 

Accordingly, our maps show emergence during maximum 

regressions and partial submergence and carbonate deposition during maximum transgressions. Periodic emergence 

and erosion of the Callville shelf is consistent with the presence of clasts in correlative rocks in the adjacent Oquirrh 

basin that include fossils slightly older than their host strata 

(Jordan and Douglas, 1980).

Relatively rapid subsidence continued in the southeastern 

Oquirrh basin, where Missourian and Virgilian strata as 

thick as 2,200 m include the Bingham Mine Formation and 

the lower part of the Kessler Canyon Formation (fig. 3) 

(Tooker and Roberts, 1970). In contrast, strata of inferred 

Missourian and Virgilian age on the east flank of the Oquirrh 

basin on the Wyoming shelf (the middle part of the Weber 

Sandstone) are at most a few hundred meters thick (Bissell 

and Childs, 1958; Welsh and Bissell, 1979).

Both the Bingham Mine and Kessler Canyon Formations 

consist of sandstone and lesser amounts of interbedded limestone. Sandstone is typically massive or crossbedded and 

commonly bioturbated (Tooker and Roberts, 1970; Jordan 

and Douglas, 1980; Jordan, 1981). The Cruziana trace-fossil 

community in sandstones in the Oquirrh Mountain area (Jordan and Douglas, 1980, fig. 8; Jordan, 1981) (fig. 1) suggests 

a shallow-marine environment subject to wave energy. In the 

same area, ripple-laminated and crossbedded sandstones 

were interpreted as shallow-marine tractive-current deposits

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC19
CC20 Evolution of Sedimentary Basins Uinta and Piceance Basins
112 109° WYOMING 108°

no_o

i.o

Figure 8 (above and facing page). Maps showing the Missourian to Virgilian paleogeography of the Uinta-Piceance basin region during maximum regression (A) and transgression (8). See explanation on page CC14.

S
and massive sandstones as the deposits of density-modified 

grain flows in deeper water (Jordan, 1981). On our maps 

(fig. 8), we show shallow-water sandstones as regressive 

deposits and limestones as transgressive deposits. The 

transition from deposition of limestone to deposition of 

deep-water clastic rocks during transgressions (see fig. 9) 

occurred some time during the Missourian and Virgilian.

The inferred Missourian and Virgilian part of the Weber 

Sandstone on the Wyoming shelf crops out along the south 

flank of the Uinta Mountains. Western exposures (west of 

the Duchesne River area, fig. 1) consist of thick-bedded, 

rarely crossbedded, calcareous sandstone and minor limestone and dolomitic limestone (Huddle and McCann, 1947; 

Bissell and Childs, 1958). The sandstones are interpreted as 

mainly shallow marine in origin but may also in part reflect 

redeposition of eolian dune sediments deposited during the 

largest regressions. Accordingly, we show the contact 

between eolian and shallow-marine facies slightly south of 

the Duchesne River area. Limestones are the inferred 

maximum transgression deposits in the western Uinta 

Mountains area.

East of the Duchesne River area, the inferred Missourian 

and Virgilian part of the Weber Sandstone is dominated by 

thick units of massive and crossbedded sandstone, and limestone is uncommon (Huddle and others, 1951; Kinney, 1955; 

Bissell and Childs, 1958; Bissell, 1964; Fryberger, 1979). 

The thick sets of crossbedded strata are eolian in origin (Fryberger, 1979) and represent regressive deposition. We infer 

that the Weber eolian dune field was transgressed during 

maximum highstands and that eolian sands were redeposited 

by mass flows to form massive and (or) deformed beds (see 

Doe and Dott, 1980; Eschner and Kocurek, 1986). This 

hypothesis is strengthened by the presence of marine fossils 

in similar sandy facies in the upper (Wolfcampian) part of 

the Weber Sandstone (Bissell and Childs, 1958; Bissell, 

1964). A few transgressions may have resulted in deposition 

of limestone; however, Driese (1985) argued that Weber 

limestones in the eastern Uinta Mountains area are nonmarine and formed in interdune ponds. To the south, 'Weber 

Sandstone strata interfmger with fluvial (transgressive) and 

eolian sand-sheet (regressive) deposits of the Maroon Formation that were derived from the northern flank of the 

Uncompahgre uplift. This facies change is well documented 

in the Rangely area of northwestern Colorado (for example, 

Fryberger, 1979).

The Missourian and Virgilian of the northern Paradox 

basin consists mainly of mixed clastic and carbonate rocks of 

the Honaker Trail Formation (the upper formation of the 

Hermosa Group) of Wengerd and Matheny (1958), and the 

lower part of the "Elephant Canyon Formation" of Baars 

(1962, 1987, 1991). These strata were considered by Loope 

and others (1990) as the lower part of the upper member of 

the Hermosa Formation. Missourian and Virgilian strata in 

this area may be as thick as about 600 m. Coarse alluvial 

clastic rocks of the upper part of the Hermosa Formation

crop out on the northeastern margin of the basin. Limestone 

is the most common lithology in the Honaker Trail Formation (for example, Melton, 1972) and is inferred to represent 

transgressive deposition. Honaker Trail clastic rocks in the 

northern Paradox basin consist of varicolored siltstone and 

shale and uncommon sandstone (for example, boreholes 10, 

12, 17, fig. 1), which we infer were deposited during regressions in sabkha to shallow-marine environments. Loope 

(1984) described eolian sandstones from outcrops of the 

Honaker Trail Formation (he assigned these sandstones to 

the upper part of the Hermosa Formation) in the Moab area 

(fig. 1) in the southeastern part of the northern Paradox 

basin. Thick sandstone beds of possible eolian origin are 

common in nearby boreholes but pinch out to the north and 

west. These eolianites are also inferred regressive deposits 

and may be more abundant in the main part of the Paradox 

basin to the south.

Relief in the structural zone on the northwestern flank of 

the Uncompahgre uplift (Potter and others, 1991) is inferred 

to have been less than in the Desmoinesian, and Missourian 

and Virgilian strata are inferred to have onlapped the northwest-plunging uplift (for example, borehole 29, fig. 1). This 

decreased relief probably led to improved circulation 

through the area of the Oquirrh sag (the eastern Uinta basin) 

(Szabo and Wengerd, 1975). Several depositional facies 

converge in and north of the Oquirrh sag, an area for which 

there is essentially no late Paleozoic outcrop or subsurface 

control.

In the Eagle basin, Missourian and Virgilian strata 

include part of the nonmarine Maroon Formation and, in the 

center of the basin, the uppermost, nonevaporitic part of the 

Eagle Valley Evaporite (Johnson, 1987a; Johnson and others, 1988). A Missourian and Virgilian age for the upper 

Eagle Valley Evaporite (not previously inferred) is based on 

the following. Johnson (1987a) measured a 987-m-thick 

stratigraphic section of the upper Eagle Valley Evaporite in 

the Eagle Valley (fig. 1) in the central part of Eagle basin. 

This section begins at the top of the highest thick (60 m) gypsum bed in the Eagle Valley Evaporite, lacks unconformities, and is a continuation of a stratigraphic section measured 

by Schenk (1989). Schenk traced the upper Desmoinesian 

Jacque Mountain Limestone Member, which forms the 

uppermost bed in the Minturn Formation (fig. 3), from the 

Minturn area (fig. 1) on the east flank of the Eagle basin into 

the middle of the Eagle basin, where it immediately underlies the highest thick gypsum. In the Minturn area, Tweto 

and Lovering (1977) estimated a Maroon Formation thickness of 1,147 m, using the top of the Jacque Mountain Limestone Member as its base. This Maroon Formation thickness 

is similar to that of the post-Jacque Mountain section of the 

upper Eagle Valley Evaporite and the Maroon Formation 

along the Eagle River (1,047 m), and the two sections are 

considered correlative. The stratigraphic cross section of 

Johnson and others (1988, fig. 3) shows this relationship and 

illustrates how the upper Eagle Valley Evaporite in the

CC22 Evolution of Sedimentary Basins Uinta and Piceance Basins
central part of Eagle basin interfingers with the lower part of 

the Maroon Formation on both basin margins. Although 

lacking diagnostic fossils, the Maroon Formation is considered late Desmoinesian to mid-Wolfcampian in age based on 

regional considerations. If relatively constant sediment accumulation rates are assumed for the Maroon throughout this 

period, then the upper Eagle Valley Evaporite (which also 

lacks diagnostic fossils and, as discussed above, interfingers 

with the lower Maroon) would have a late Desmoinesian to 

Virgilian age.

The Eagle basin became progressively more nonmarine 

with time throughout the late Paleozoic, a pattern that probably reflects decreasing rates of subsidence (figs. 10,1 1) and 

associated increased progradation of clastic rocks from the 

basin margins. Facies and paleocurrent patterns indicate that 

the axis of Eagle basin was strongly skewed to the northeast 

during Maroon time (Johnson and others, 1988, fig. 7), constituting a reversal in basin polarity from the Morrowan and 

Atokan. The Uncompahgre uplift was the main source of 

clastic detritus in the Maroon (Johnson, 1987b). The 

Sawatch uplift influenced depositional patterns (Johnson, 

1987a, 1989a; Johnson and others, 1988) but probably had 

low relief.

The Maroon Formation has a maximum thickness of 

about 1,000 m in the main part of Eagle basin (Johnson and 

others, 1988) and consists of interbedded fluvial and eolian 

deposits. Near the basin margins, Maroon fluvial deposits 

(Johnson, 1987b) consist of channel architectural elements 

(Miall, 1985) of conglomeratic sandstone. These proximal 

channel deposits grade down the paleoslope into laminated 

sand elements of very fine to fine grained sandstone. The 

facies changes reflect decreases in flow strength, depth, and 

discharge, characteristics of "terminal fan" fluvial systems 

in arid basins (Friend, 1978).

Maroon Formation eolianites are mainly sand-sheet 

deposits (Johnson, 1987a; Johnson and others, 1988). These 

sand-sheet deposits typically comprise about 30 percent of 

Maroon sections and appear to cyclically alternate with fluvial intervals (fig. 5). Some fluvial deposits (shown schematically in fig. 8) were also deposited lateral to the eolian sand 

sheets, in wadi streams. Silt eroded from the sand-sheet 

deposits in the Maroon Formation was transported to the 

downwind basin margin, the northwest flank of the Sawatch 

uplift (Ruedi Reservoir area, fig. 1), where it was redeposited 

as a thick section of loess (Johnson, 1989a).

Johnson (1987a) and Johnson and others (1988) suggested that the inferred nonmarine cyclicity in the Maroon 

Formation was controlled by cyclic climatic changes synchronous with global glacio-eustatic changes. In this interpretation, fluvial deposition dominated in the Maroon during 

wetter climatic periods (correlating with eustatic highstands), and eolian sand-sheet deposition dominated during 

drier climatic periods (eustatic lowstands).

The upper, nonevapori tic part of the Eagle Valley 

Evaporite, which crops out along the Missourian-Virgilian

axis of the Eagle basin, consists of several gradational facies 

(Johnson, 1987a; Johnson and others, 1988): (1) plane-bedded or low-angle-bedded, varicolored, fine- to coarsegrained sandstone; (2) ripple-laminated siltstone to finegrained sandstone; (3) massive, extensively bioturbated 

sandstone and siltstone; (4) grayish-black, thinly laminated, 

calcareous siltstone and mudstone; and (5) grayish-black 

micritic limestone and lime mudstone. These facies represent a continuum from sabkha and shallow-water (regressive) to offshore marine (transgressive) deposition. As in the 

late Atokan and Desmoinesian, the Eagle basin was probably 

completely cut off from normal marine waters during regressions, when the central part of the basin may have resembled 

a playa lake. During transgressions, marine waters entered 

the basin from the north. Marine limestone of inferred Missourian and Virgilian age (based on stratigraphic position) is 

present at Ripple Creek, Miller Creek, and Sylvan (Johnson, 

1987a) and in boreholes 40, 41, 45, 48, and 49 (fig. 1). The 

contact between these limestones and sandy marine environments (the transgressed Weber sand sea) in northwestern 

Colorado is between boreholes 40 and 50.

In summary, the Missourian and Virgilian in the UintaPiceance region was characterized by decreasing tectonic 

activity, reflected in lower rates of subsidence (figs. 10, 11). 

In the Eagle basin, fluvial and eolian deposition dominated 

while evaporite deposition ended. Evaporite deposition similarly ended in the Paradox basin and was replaced by deposition of sabkha and (or) shallow-marine deposits during 

regressions. Progradation of sands across the Wyoming shelf 

continued during regressions. Transgressive deposition of 

limestone was limited to the western part of the UintaPiceance region and a small area in the eastern Eagle basin.

Early Early Permian Early Wolfcampian

The major basins of the Uinta-Piceance region (fig. 9) 

continued to subside during the early Wolfcampian (figs. 10, 

11). As the Emery uplift subsided, however, the Paradox 

basin became connected to the Callville shelf and its history 

as a narrow trough ended. Subsidence exceeded sediment 

supply in the southeastern Oquirrh basin, so that it was both 

a depositional and topographic basin. Depositional patterns 

on the Wyoming shelf and in the Eagle basin remained similar. Ross and Ross (1987, 1988) recognize 11-15 major 

eustatic changes in the Wolfcampian, with inferred magnitudes of as much as 170 m (fig. 4).

The Pakoon Limestone (50-70 m thick) represents early 

Wolfcampian deposition on the Callville shelf; the Queantoweap Sandstone (50-90 m thick) represents late Wolfcampian deposition (Brill, 1963; Welsh, 1976). Few lithologic 

descriptions are available for the Pakoon in the UintaPiceance region. It apparently consists of dolomitic limestone and minor sandstone (borehole 1 and Cove Fort area 

outcrops, fig. 1), inferred to represent transgressive and

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC23
CC24 Evolution of Sedimentary Basins Uinta and Piceance Basins
109° 108C 107 e

6)Q.s

era

£oi

o

^i-2i?:?-;-.V:iV..o'»-o 4C^-^^%°-Jf RYINGPAjSl'7 : >,_TT 4-iiO »..«--Q-- '..:.-,. -)*x^^-o; .» * ' ;'. , ._ !_' ' -± nl . & ., > -V>-v-MEMBE .J !,...r±."rr.,'mrrr ±r.,

B 50 100 KILOMETERS

Figure 9 (above and facing page). Maps showing the early Wolfcampian paleogeography of the Uinta-Piceance basin region during maximum regression (A) and transgression (B). See explanation on page CO 4.
regressive deposition, respectively. J.E. Welsh (oral commun. to M.A. Chan, 1989) suggested that features of the Pakoon carbonate rocks indicate a somewhat restricted, shallowmarine environment. George (1985) reported that Pakoon 

sandstones are crossbedded and that they resemble sandstones in the unconformably underlying Callville Limestone. 

These sandstones could be shallow marine (as shown in fig. 

9) or possibly eolian in origin. Because Wolfcampian subsidence rates on the Callville shelf were apparently very low 

(suggested by the small thickness of preserved section), it is 

also possible that the Callville shelf was an emergent, erosional area during larger regressions. There are apparently 

no transitional facies between Pakoon and Queantoweap 

strata in the Cove Fort area and the thick section of Oquirrh 

basin Wolfcampian strata to the north in the Route 148 

southwest area (Brill, 1963; Morris, 1977) (fig. 1).

The Pakoon Limestone apparently grades eastward into 

the upper part of the Elephant Canyon Formation of Baars 

(1962, 1987, 1991) in the northern Paradox basin; these latter strata were considered the Rico Formation by Loope 

(1984) and the upper part of the upper member of the Hermosa Formation and "lower Cutler beds" by Loope and others (1990). Although debate concerning the stratigraphic 

nomenclature continues (for example, Baars, 1991), the 

name Elephant Canyon Formation is used here for convenience and because of its prevalence in the literature. In the 

main part of the northern Paradox basin, the Elephant Canyon Formation is as thick as about 460 m and consists of 

interbedded carbonate rocks, fine-grained clastic rocks, and 

minor anhydrite (for example, boreholes 8, 14, 16, 17, 24, 

26, fig. 1) (Baars, 1962, 1987; Terrell, 1972; Loope, 1984). 

Limestones are commonly fossiliferous and are interpreted 

as shelf sediments deposited during transgressions. Interbedded clastic rocks are interpreted as sabkha, playa, and shallow-marine regressive deposits.

Sandstones of inferred eolian origin are also present at a 

few locations in the upper part of the Elephant Canyon Formation in the southern part of the northern Paradox basin (for 

example, boreholes 4 and 6, fig. 1) and in a narrow belt near 

the contact with the Cutler Formation on the east flank of the 

basin (Terrell, 1972; Loope, 1984; Campbell, 1987). Chan 

and Langford (1987) described cycles (about 6-40 m thick) 

consisting of marine and erg-margin deposits in the upper 

part of the Elephant Canyon Formation in the central to 

northern Paradox basin. Transgressive deposits consist of 

normal-marine fossiliferous shelf limestone and tidal sandstone and are as thick as a few tens of meters. Regressive 

deposits consist of fluvial and eolian sandstone. The relatively sharp lower contacts of limestone beds indicate rapid 

eustatic rises and inundation of erg margins. The marine 

incursions are thought to have entered the Paradox basin 

from the west (Baars, 1979; Campbell, 1987). Closely 

spaced (every 4-15 m) supersurfaces in the eolian and fluvial units may have formed during wet climatic intervals by 

river flooding (Langford and Chan, 1988, 1990). Based on

the abundance of the supersurfaces, Chan and Langford 

(1987) suggested that climatic fluctuations of shorter periodicity than late Paleozoic eustatic fluctuations may have influenced sedimentation in this area.

The Cutler Formation comprises coarse alluvial conglomerate and arkosic sandstone derived from the 

Uncompahgre uplift. As in the Missourian and Virgilian, 

narrowing of the width of the band of coarse elastics derived 

from this highland to the northwest suggests diminished 

relief on the Uncompahgre in that direction. Onlap of the 

northwestern flank of the Uncompahgre uplift apparently 

continued in the Wolfcampian.

In the area of the Emery uplift, transitional strata between 

the Pakoon Limestone to the west and the upper part of the 

Elephant Canyon Formation to the east consist of dolomite 

and limestone, thick to thin-bedded sandstone, varicolored 

siltstone and shale, and minor anhydrite (boreholes 2, 3, 5, 

19-23, 25, fig. 1). The proportion of sandstone is variable 

but probably is significantly greater than in the upper part of 

the Elephant Canyon Formation. Most likely the sands were 

not transported westward across the Paradox basin from the 

Uncompahgre uplift but instead had northern sources in 

sand-rich Weber environments of the Wyoming shelf or 

southern sources in the eolian sands of the Cedar Mesa Sandstone of Baars (1962). The sedimentology of these Elephant 

Canyon sands is unknown, but we infer that they were 

deposited by eolian processes during larger regressions and 

graded westward into shallow-marine environments. The 

Cedar Mesa Sandstone interfingers with the upper part of the 

Elephant Canyon Formation to the south in the main part of 

the Paradox basin and overlies the upper Elephant Canyon 

Formation in the northern Paradox basin.

Lower Wolfcampian strata in the southeastern Oquirrh 

basin include the Curry Peak and Freeman Peak Formations 

and undifferentiated strata of the Oquirrh Group (Welsh and 

James, 1961; Jordan and Douglas, 1980). These Wolfcampian strata are as thick as 1,500 m thick and are unconformably overlain (in sequence) by the Diamond Creek 

Sandstone of Wolfcampian age and the Park City or Phosphoria Formations of late Early (Leonardian to Wordian) 

Permian age. Lower Wolfcampian rocks are mainly massive, 

crossbedded, graded, and ripple-laminated sandstone and 

minor interbeds of conglomerate and limestone (Larson, 

1979; Larson and Clark, 1979; Jordan and Douglas, 1980; 

Jordan, 1981). Conglomerate crops out in a northwesttrending band along the southwest margin of the basin and 

contains limestone clasts older than host strata. Chamberlain 

and Clark (1973), Larson (1979), Jordan (1979, 1981), and 

Jordan and Douglas (1980) interpreted the sandstones and 

conglomerates as the deposits of sediment gravity flows, 

including turbidity currents, density-modified grain flows, 

and debris flows. Trace-fossil communities in the sandstones 

suggest that deposition occurred below wave base at bathyal 

depths (Chamberlain and Clark, 1973; Larson, 1979; Larson 

and Clark, 1979; Jordan and Douglas, 1980; Jordan, 1981).

CC26 Evolution of Sedimentary Basins Uinta and Piceance Basins
Jordan and Douglas (1980) suggested a slope or basin-floor 

depositional environment at depths of a few hundred meters 

or more. These depths exceed the magnitude of inferred sealevel changes (Ross and Ross, 1987) and thus provide a 

rationale for similar depositional processes and environments during both eustatic highstands and lowstands. Lowstand deposits are probably more conglomeratic, containing 

clasts derived from periodically exposed platforms and carried basinward through erosional channels (see Larson and 

Clark, 1979). On figure 9, the location of the transition from 

shallow- to deep-water environments on the northeastern 

basin flank is based on the maps of Jordan and Douglas 

(1980) west of the Wasatch fault and on the facies descriptions of Larson (1979) and Larson and Clark (1979) in the 

area of the Charleston-Nebo thrust sheet. On the southwestern flank of the basin, this transition must lie northeast of the 

Lower Permian limestones of the Oquirrh Group mapped by 

Morris (1977) in the Route 148 southwest area (fig. 1).

The upper part of the Weber Sandstone consists almost 

entirely of sandstone and comprises the early Wolfcampian 

of the Wyoming shelf depositional province (Bissell and 

Childs, 1958). It is unconformably overlain by the Park City 

Formation. Sandstones of the Weber are predominantly 

crossbedded in the eastern Uinta Mountains (east of the 

Duchesne River area, fig. 1) and predominantly massive in 

the western Uinta Mountains. These strata were probably 

deposited in eolian (maximum regressions) and shallowmarine (maximum transgressions) environments, and the 

relative proportions of the two bedding types reflect the 

dominant depositional process. Shallow-marine conditions 

probably dominated in the western Uinta Mountains; eolian 

dunes probably formed only during a few of the larger 

regressions and were rapidly reworked in subsequent transgressions. Similarly, the large proportion of crossbedded 

strata in the eastern Uinta Mountains suggests that eolian 

conditions dominated and that only the largest transgressions 

submerged the Weber dune field. The presence of marine 

fossils (Bissell and Childs, 1958), massive beds, contorted 

stratification, and extensive burrowing (Fryberger, 1979) all 

support intermittent marine conditions for the Weber in the 

eastern Uintas. The Weber Sandstone interfingers to the 

south and southeast with fluvial and eolian sand-sheet deposits of the Maroon Formation (boreholes 31-35, fig. 1), which 

were derived from the northern flank of the ancestral 

Uncompahgre uplift.

The upper part of the Maroon Formation comprises the 

early Wolfcampian of the Eagle basin. These strata are 

unconformably overlain by the Upper Permian to Lower Triassic State Bridge Formation. Similar to the Missourian and 

Virgilian lower part of the Maroon, these strata consist of 

mixed fluvial and eolian sand-sheet deposits in most of the 

basin. The basin continued to be asymmetric, with its depositional axis skewed strongly to the east. During transgressions, fluvial systems draining the Uncompahgre and Front 

Range uplifts flowed across the basin, converged along the

depositional axis, and flowed north to exit the basin 

(Johnson, 1987a, b; Johnson and others, 1988). Johnson 

(1987a) described rare marine beds from the upper Maroon 

in the Sylvan, Miller Creek, and Ripple Creek areas (fig. 1), 

and Hallgarth (1959) showed dolomitic and gypsiferous 

mudstones from the uppermost Maroon in boreholes 41,45, 

and 47 (fig. 1). These strata were probably deposited during 

the largest transgressions in a sabkha or marginal-marine 

setting.

As in the Missourian and Virgilian, eolian sand-sheet 

deposits were dominant in the Eagle basin during regressions 

and merged to the north with the Weber dune deposits. Loess 

continued to be deposited in the Ruedi Reservoir area (fig. 1) 

on the downwind basin margin adjacent to the low-relief 

Sawatch uplift. Near the end of Maroon deposition, however, deposition of eolian dune sediments of the Fryingpan 

Member of the Maroon Formation over the loessite formed 

a basin-margin dune field (Johnson, 1989b). The termination 

of loess deposition and the initiation of dune deposition were 

probably forced by a cessation or slowing of subsidence in 

the Eagle basin, the source area for the basin-margin eolianites. Rather than being rapidly buried beneath the Maroon 

alluvial-eolian plain, sands in the source area were exposed 

at the surface and made susceptible to eolian erosion and 

transport for longer time intervals. Consistent with this interpretation, thick beds of coarse to granular sand interpreted as 

deflation lags are present at the top of the Maroon Formation 

in many places (Johnson and others, 1988).

In summary, depositional patterns initiated in the Missourian and Virgilian in the Eagle basin and on the Wyoming 

shelf continued into the early Wolfcampian. The Emery 

uplift became fully or mostly submerged and the Paradox 

basin ceased to be a discrete geomorphic element. Deepwater clastic deposition began in the Oquirrh basin (or continued from the latter part of the Missourian and Virgilian). 

Transgressive limestone deposition was limited to the southeastern part of the Uinta-Piceance basin region.

DISCUSSION

The geohistory diagrams (figs. 10, 11) show that the 

Oquirrh basin experienced the greatest rates and magnitude 

of subsidence, the Eagle and northern Paradox basins experienced intermediate and similar rates, and the Wyoming 

shelf experienced the lowest rates and magnitude of subsidence. All four sedimentary provinces experienced their lowest rates of subsidence in the Early Pennsylvanian, their 

highest rates in the Middle Pennsylvanian, and intermediate 

rates in the Late Pennsylvanian and Early Permian. The 

decrease in subsidence rates from the Middle to Late Pennsylvanian in the Oquirrh basin is minor in contrast to 

decreases in the other three depositional provinces.

Subsidence in the northern Paradox basin (fig. 10C) is 

inferred to have been mainly controlled by flexural loading

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC27
325

2-

4H

6-

AGE, IN MILLIONS OF YEARS

300 

i i I i i

275

Basement subsidence

\\ Basement subsidence (Klein)

Total subsidence 

(Klein)

Early Middle

PENNSYLVANIAN

Late Early 

PERMIAN

AGE, IN MILLIONS OF YEARS 

300 275

Early 

PERMIAN

AGE, IN MILLIONS OF YEARS

300

I

275

Early Middle

PENNSYLVANIAN

Late Early 

PERMIAN

325

AGE, IN MILLIONS OF YEARS 

300 275

D

Early Middle

PENNSYLVANIAN

Late Early 

PERMIAN

driven by westward overthrusting of the Uncompahgre 

uplift. As stated earlier, the basin axis was skewed strongly 

to the east adjacent to the Uncompahgre uplift (Mallory,

Figure 10 (facing column). Geohistory diagrams for four sedimentary provinces in the Uinta-Piceance basin region. A, 

Oquirrh basin; B, Wyoming shelf; C, northern Paradox basin; 

D, Eagle basin. The origin of each curve at the beginning of 

the Pennsylvanian is assumed to be at sea level because the 

Late Mississippian in each province was characterized either 

by shallow-marine carbonate deposition (A, B) or by minimal 

emergence (C, D). Solid lines represent total subsidence. Dotted lines represent subsidence of the Late Mississippian "basement" corrected for the incremental load induced by the 

weight of sediment through time and, thus, the subsidence due 

to tectonics. Solid circles represent the time-thickness data 

points used in the analysis. Time scale used is from Haq and 

Van Eysinga (1987). Diagram A also shows the same data plotted using the time scale proposed by Klein (1990). Corrections 

for compaction are based on lithology and follow the exponential porosity function presented by Sclater and Christie (1980). 

No corrections were made for bathymetry or eustatic fluctuations. Because of uncertainties involving the ages of units and 

their compaction and diagenetic histories, the plots should be 

regarded as approximations. Sources of data: A, south-central 

Oquirrh Mountains (Tooker and Roberts, 1970); B, Split Mountain anticline (Kinney, 1955; Thomas and others, 1945); C, 

northern Paradox basin (jay Roberts no. 1 Whitecloud, sec. 34, 

T. 22 S., R. 17 E.); D, White River uplift (Johnson and others, 

1988).

1972, 1975), consistent with this inference. Subsidence in 

the eastern Eagle basin (fig. 10D) closely resembles that in 

the northern Paradox basin, supporting the contention that 

the Gore fault (for which the sense of late Paleozoic offset is 

uncertain) on the eastern margin of Eagle basin was also a 

reverse fault. Alternatively, two different yet synchronous 

tectonic processes (flexural loading in the northern Paradox 

basin, oblique or normal rifting in the Eagle basin) would 

need to be invoked to produce their almost identical subsidence histories.

Evaporite deposition in both the northern Paradox basin 

and the Eagle basin probably was controlled by subsidence 

rate, drainage restriction, arid climate, and fluctuating 

eustasy. Evaporite deposition in each basin occurred only 

during the Middle Pennsylvanian, the time of most rapid 

subsidence. Rapid subsidence commonly results in trapping 

of clastic sediment adjacent to basin margins (Blair and Bilodeau, 1988; Heller and others, 1988), causing basin interiors 

to become isolated from supplies of clastic sediment (or sediment starved), a partial requirement for evaporite deposition. Drainage restriction took different forms in the two 

basins, with an inferred tectonic barrier at the northern end 

of the Paradox basin and a combined tectonic (slowly subsiding) and sedimentary (the positive relief on the Morgan 

Formation dune field) barrier at the northern end of the Eagle 

basin (fig. 7A). Regular eustatic fluctuations led to submergence of these barriers and replenishment of the water column. Re-emergence of barriers during eustatic regressions 

led to isolation of water bodies and climate-induced evaporite deposition. All of the above conditions were required for 

evaporite deposition; evaporite deposition ceased in the Late 

Pennsylvanian when subsidence rates diminished (figs. 10C,

CC28 Evolution of Sedimentary Basins Uinta and Piceance Basins
D, 11), and clastic sediments began to prograde into the 

basin interiors.

The slowly subsiding Wyoming shelf (fig. 105) served as 

a bypass zone across which sediment was transported southwestward to the rapidly subsiding Oquirrh basin. The much 

larger amount of total subsidence in the Oquirrh basin (6-7 

times greater than on the Wyoming shelf) (fig. 11) suggests 

that significant structures (faults or ramps) were present 

between the two depositional provinces. This inferred zone 

of structural weakness is not presently recognizable; it was 

presumably reactivated during the Cretaceous Sevier and

325

2-

4-

6-

AGE, IN MILLIONS OF YEARS

300 

j___I___I___i___I___

275

EXPLANATION

 Oquirrh basin

 - Wyoming shelf

 . Paradox basin

 Eagle basin

Early Middle Late 

PENNSYLVANIAN

Early 

PERMIAN

325

1-

2-

3-

AGE, IN MILLIONS OF YEARS 

300 275

_i____i____i_ i___i____i

EXPLANATION

- - Oquirrh basin 

-- Wyoming shelf 

Paradox basin 

Eagle basin

Early Middle Late 

PENNSYLVANIAN

Early 

PERMIAN

Figure 11. Ceohistory diagrams comparing total subsidence 

(A) and basement subsidence (B) in four sedimentary provinces 

in the Uinta-Piceance basin region. See figure 10 for details 

concerning construction of the diagrams.

early Tertiary Laramide orogeny or has been buried by 

younger sediments. Despite the contrasts in subsidence history, shallow-marine deposition in both the Wyoming shelf 

and the Oquirrh basin through most of the Pennsylvania!! 

indicates that sediment supply kept up with subsidence in the 

Oquirrh basin. Deepening of the Oquirrh basin in the latest 

Pennsylvania and the Early Permian (fig. 9) was apparently 

not accompanied by a significant increase in subsidence 

rates (fig. 10/4). Diminishing sediment supply may therefore 

have led to the formation of deep-water depositional environments.

The Oquirrh basin is distinguished from the three other 

depositional provinces plotted on figure 11 on the basis of its 

higher rates and magnitude of subsidence and its continuing 

high subsidence rates through the Late Pennsylvanian and 

the Early Permian. All of the Uinta-Piceance basin region 

was located east of the ISr=0.706 isopleth for Mesozoic and 

Cenozoic igneous rocks (Kistler and Peterman, 1978); thus 

its basement rocks are similar. The Oquirrh basin, however, 

does overlie the hinge zone of the Late Proterozoic to midPaleozoic miogeocline (Hill, 1976), and weakened crust in 

this zone might in part be responsible for the anomalously 

high subsidence in the Oquirrh basin. More importantly, the 

tectonic controls on subsidence in the Oquirrh basin were 

probably different from those in the Eagle and Paradox 

basins.

Both the timing (figs. 10,11) and geometry of subsidence 

suggest that late Paleozoic deformation in the UintaPiceance region of the ancestral Rocky Mountains resulted 

from interactions along both the western and southeastern 

continental margins (fig. 2). As stated earlier, initial subsidence of the Oquirrh basin and the Wyoming shelf as discrete basinal elements began in the Late Mississippian. This 

subsidence predates the onset of significant deformation in 

the more proximal parts of the foreland province associated 

with the Marathon-Ouachita erogenic belt (the collisional 

mountain belt of the southeastern convergent margin) by 

10-20 million years and predates initial subsidence in the 

Eagle and Paradox basins by 25-30 million years (Kluth, 

1986). Assuming that the effects of the continent-continent 

collision to the southeast propagated inland with time as the 

collision broadened or even that deformation was initiated 

synchronously across the entire foreland, the early phases of 

subsidence in the Oquirrh basin and Wyoming shelf are too 

old to be the result of the collision and demand an independent driving force. The inferred extensional or transtensional 

origin of the Oquirrh basin is consistent with the structural 

style noted to the west by Ketner (1977) and Smith and 

Miller (1990), and a western driving force is highly likely. In 

contrast, the inferred contractional or transpressional style of 

the Paradox basin and possibly the Eagle basin may be more 

consistent with the foreland deformation associated with a 

convergent margin.

Can the effects of the western and southeastern driving 

forces be isolated? Because significant subsidence in the 

Eagle and Paradox basins did not begin until the Atokan,

Pennsylvanian and Early Permian Paleogeography, Colorado and Utah CC29
these basins may never have been significantly affected by 

the western driving force. The Desmoinesian acceleration of 

subsidence in the Eagle, Paradox, and Oquirrh basins is 

roughly correlative and therefore may in each case reflect the 

growing influence of the southeastern force (the dominant 

influence in the Eagle and Paradox basins and a supplemental influence in the Oquirrh basin). The Late Pennsylvanian 

decrease in subsidence in the Eagle and Paradox basins indicates that the influence of the southeastern force diminished; 

continued rapid Late Pennsylvanian and Early Permian subsidence in the Oquirrh basin supports the continued importance of west-induced extensional deformation in the 

western Uinta-Piceance region during this interval.

While considering only interactions on the southern continental margin as a driving force for deformation in the 

ancestral Rocky Mountains, Kluth (1986) noted the high 

angle between the uplifts and basins of the ancestral Rocky 

Mountain orogeny and the Marathon-Ouachita orogenic 

belt. Although Kluth presented a rationale for these anomalous trends in the context of northward-imposed paleostress, 

the geometry of structures in at least the western ancestral 

Rocky Mountains is more consistent with paleostress 

imposed by transtensional deformation along a western continental margin. Regional transtension in the ancestral 

Rocky Mountain province is supported by the mix of compressional and extensional structural styles. Future work on 

Pennsylvanian and Permian deformation and subsidence in 

the western United States (including locating and describing 

the structures bounding the Oquirrh basin), along with better 

documentation of modern analogues, will further clarify the 

nature of intraplate deformation in the ancestral Rocky 

Mountains.

CONCLUSIONS

Expansion and contraction of late Paleozoic Gondwana 

ice sheets resulted in repetitive global eustatic and climatic 

fluctuations. These eustatic and climatic fluctuations, along 

with tectonic activity and variations in sediment supply, 

were major controls on the development of transgressiveregressive depositional sequences in the Uinta-Piceance 

basin region of northwestern Colorado and northeastern 

Utah. Major changes in depositional patterns in the UintaPiceance region associated with the repeated transgressions 

and regressions are shown on four sets of paleogeographic 

maps representing key time periods in the late Paleozoic. 

The contrasts between maps representing maximum transgressions and regressions are substantial and indicate the 

necessity of this approach for adequately representing late 

Paleozoic paleogeography.

The Uinta-Piceance basin region includes four major 

sedimentary provinces: the Eagle basin, the northern Paradox basin, the southern Wyoming shelf, and the southeastern 

Oquirrh basin. The Oquirrh basin experienced the greatest 

rates and magnitude of subsidence, the Eagle and northern

Paradox basins experienced intermediate and very similar 

rates, and the Wyoming shelf experienced the lowest rates 

and magnitude of subsidence. All four sedimentary provinces experienced their lowest rates of subsidence in the 

Early Pennsylvanian, their highest rates in the Middle Pennsylvanian, and intermediate rates in the Late Pennsylvanian 

and Early Permian. The timing, magnitude, and geometry of 

subsidence in the Uinta-Piceance region suggests that 

regional transtension was driven by the overlapping influences of a more distant, convergent continental margin to the 

southeast and a more proximal, transform-fault (?) continental margin to the west. The effects of this complex intraplate 

deformation are reflected in the considerable variations 

within transgressive-regressive depositional sequences.

Morrowan and lower Atokan strata throughout most of 

the Uinta-Piceance basin region consist of alternating finegrained clastic rocks (regressive deposits) and more abundant limestone (transgressive deposits). Significant Morrowan and early Atokan tectonic uplifts include the ancestral 

Front Range and Sawatch uplifts. Intrabasinal faulting characterized the Eagle basin (Waechter and Johnson, 1986) and 

locally influenced deltaic and prodeltaic deposition.

The late Atokan to Desmoinesian history of the UintaPiceance region was characterized by increased tectonic 

activity, which is reflected in higher rates of basin subsidence, continued unroofing of the Front Range and 

Sawatch uplifts, and initial uplift of the Uncompahgre uplift. 

The combined effects of tectonism and eustasy led to 

restricted circulation and evaporite deposition in the Eagle 

and Paradox basins. Regressive deposition was also characterized by significant progradation of eolian sands across the 

Wyoming shelf en route to the Oquirrh basin. Limestone 

deposition dominated during transgressions.

The Missourian to Virgilian history of the UintaPiceance basin region was characterized by decreasing tectonic activity and rates of subsidence. Progradation of clastic 

sediments into the central parts of the Eagle and northern 

Paradox basins resulted in cessation of evaporite deposition. 

During regressions, fluvial and eolian deposition dominated 

in the Eagle basin, whereas sabkha and (or) shallow-marine 

deposition dominated in the northern Paradox basin. The 

Wyoming shelf continued to serve as a conveyor belt of clastic sediment to the Oquirrh basin, which was characterized 

by a transition from shallow- to deep-water deposition. 

Transgressive deposition of limestone was limited to the 

western part of the Uinta-Piceance basin region and a small 

area in the eastern Eagle basin.

Depositional patterns initiated in the Missourian and Virgilian in the Eagle basin and on the Wyoming shelf continued into the early Wolfcampian. The Emery uplift was fully 

or mostly submerged, and the Paradox basin ceased to be a 

discrete geomorphic element. Deep-water clastic deposition 

continued in the Oquirrh basin. Transgressive limestone 

deposition was limited to the southwestern part of the UintaPiceance region.

CC30 Evolution of Sedimentary Basins Uinta and Piceance Basins
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Atwater, T., 1970, Implications of plate tectonics for the Cenozoic 

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CC34 Evolution of Sedimentary Basins Uinta and Piceance Basins
Appendix Borehole Geophysical Logs

[Locations shown by number on fig. 1 ]

1. Shell Oil Company Sunset Canyon (W) No. 1, sec. 21, T. 22 S., R. 4 W., Millard County, Utah.

2 Phillips Petroleum Company United States "E" No. 1, sec. 27, T. 19 S., R. 3 E., Sanpete County, Utah.

3. Skelly Oil Company Emery Unit No. 1, sec. 34, T. 22 S., R. 5 E., Sevier County, Utah (Irwin, 1976).

4. Mountain Fuel Supply South Last Chance Unit No. 1A, sec. 18, T. 26 S., R. 7 E., Emery County, Utah.

5. Pan American Oil Company No. 3 Perron Unit, sec. 21, T. 20 S., R. 7 E., Emery County, Utah.

6. Pan American Petroleum Corporation No. 9 Nequoia Arch, sec. 25, T. 26 S., R. 13 E., Emery County, Utah.

7. Delhi Oil Company No. 1 Russell, sec. 34, T. 25 S., R. 12 E., Emery County, Utah.

8. Superior Oil Company North Springs Wash No. 31-15, sec. 15, T. 25 S., R. 15 E., Emery County, Utah.

9. Jay Roberts No. 1 Whitecloud, sec. 14, T. 26 S., R. 20 E., Grand County, Utah (Hite and Liming, 1972).

10. Billiard Oil and Gas Company Klondike No. 1, sec. 24, T. 23 S., R. 18 E., Grand County, Utah.

11. Texaco Oil Company No. 1 Government MacKinnon, sec. 15, T. 23 S., R. 19 E., Grand County, Utah.

12. Superior Oil Company No. 22-34, sec. 34, T. 22 S., R. 17 E., Grand County, Utah (Hite and Liming, 1972).

13. Amax Petroleum Company No. 24-1 Green River Desert, sec. 24, T. 22 S., R. 13 E., Emery County, Utah.

14. Superior Oil Company No. 14-24, sec. 24, T. 21 S., R. 15 E., Emery County, Utah.

15. Gulf Oil Company No. 1 Norris, sec. 8, T. 18 S., R. 16 E., Emery County, Utah.

16. Humble Oil and Refining Company No. 2 Woodside Unit, sec. 30, T. 18 S., R. 14 E., Emery County, Utah.

17. Placid Oil Company No. 1 Marsh Flat, sec. 29, T. 17 S., R. 14 E., Emery County, Utah.

18. El Paso Natural Gas Company No. 1 Pack Saddle, sec. 12, T. 18 S., R. 12 E., Emery County, Utah.

19. Jake Hamon No. 8-1 USA Federal, sec. 8, T. 19 S., R. 9 E., Emery County, Utah.

20. Tiger Oil Company No. 4 1-15 Jack Curtis, sec. 15, T. 18 S., R. 7 E., Emery County, Utah.

21. Phillips Petroleum Company Huntington Anticline No. 1, sec. 15, T. 17 S., R. 8 E., Emery County, Utah.

22. Atlantic Richfield No. 1 Hiawatha, sec. 13, T. 15 S, R. 7 E., Emery County, Utah.

23. Tenneco Oil Company Clear Creek No. 1, sec. 17, T. 14 S., R. 7 E., Carbon County, Utah.

24. Shell Oil Company No. 1 North Springs Unit, sec. 27, T. 15 S., R. 9 E., Carbon County, Utah.

25. Pure Oil Company No. 1-A Washboard, sec. 12, T. 16 S., R. 9 E., Emery County, Utah.

26. Pure Oil Company No. 1 Desert Lake, sec. 1, T. 17 S., R. 10 E., Emery County, Utah.

27. Shell Oil Company No. 1 Miller Creek, sec. 26, T. 15 S., R. 10 E., Carbon County, Utah.

28. Pan American Oil Company Federal Mounds No. 1, sec. 11, T. 16 S., R. 11 E., Emery County, Utah.

29. Shell Oil Company No. 1-A Farnham Dome, sec. 8, T. 15 S., R. 12 E., Carbon County, Utah.

30. Mountain Fuel Supply No. 1 Sunnyside, sec. 17, T. 15 S., R. 13 E., Carbon County, Utah.

31. Chevron Oil Company No. 1 Stone Cabin, sec. 29, T. 12 S., R. 15 E., Carbon County, Utah.

32. Gulf Energy and Minerals Graynolls Federal No. 1, sec. 22, T. 12 S., R. 21 E., Uintah County, Utah.

33. Phillips Petroleum Company Two Waters No. 1, sec. 22., T. 14 S., R. 25 E., Uintah County, Utah.

34. Continental Oil Company No. 22-1 Federal, sec. 22, T. 9 S., R. 20 E., Uintah County, Utah.

35. Phillips Petroleum Company Watson "B" No. 1, sec. 34, T. 9 S., R. 25 E., Uintah County, Utah.

36. Phillips Petroleum Company No. 1 Hells Canyon, sec. 12, T. 2 S., R. 104 W., Rio Blanco County, Colorado.

37. Superior Oil Company No. 1 Hellman-Government, sec. 5, T. 3 S., R. 101 W., Rio Blanco County, Colorado (Hallgarth, 1959).

38. Phillips Petroleum Company No. 15 Douglas Creek Arch, sec. 14, T. 1 S., R. 102 W., Rio Blanco County, Colorado.

39. The Texas Company No. 70-32 UPRR, sec. 32, T. 2 N., R. 102 W., Rio Blanco County Colorado (Hallgarth, 1959).

40. The Texas Company and the California Company No. 20 Government, sec. 34, T. 3 N., R. 94 W., Rio Blanco County, Colorado 

	(Hallgarth, 1959).

41. Stanolind Oil and Gas Company Scott No. 1, sec. 20, T. 1 N., R. 93 W., Rio Blanco County, Colorado.

42. Buford Oil Company Government No. 1, sec. 16, T. 1 N., R. 91 W., Rio Blanco County, Colorado.

43. Phillips Petroleum Company No. 1 Linville Government, sec. 1, T. 1 N., R. 92 W., Rio Blanco County, Colorado.

44. Benedum Trees Oil Company Daugherty Government No. 1, sec. 7, T. 1 N., R. 88 W., Garfield County, Colorado.

45. Phillips Petroleum Company Poose Creek No. 1, sec. 10, T. 2 N., R. 88 W., Rio Blanco County, Colorado.

46. Sam Gary No. 11-15 Aspegren, sec. 2, T. 1 S., R. 85 W., Routt County, Colorado.

47. Miami Oil Producers No. 1 O'Brien, sec. 14, T. 4 N., R, 90 W., Moffat County, Colorado.

48. Texaco Oil Company No. 7 Government Treleaven, sec. 31, T. 5 N., R. 90 W., Moffat County, Colorado.

49. Stanolind Oil and Gas Company No. 24 Parkinson Government, sec. 22, T. 4 N., R. 92 W., Moffat County, Colorado (Hallgarth, 1959).

50. Stanolind Oil and Gas Company Blue No. 1, sec. 35, T. 6 N., R. 96 W., Moffat County, Colorado (Hallgarth, 1959).

51. Humble Oil Company No. 1 Lay Creek, sec. 13, T. 8 N., R. 93 W., Moffat County, Colorado.

52. Gulf Energy and Minerals Company No. 1 Slater Creek, sec. 7, T. 11 N., R. 89 W., Moffat County, Colorado.

53. Phillips Petroleum Company No. 8 Baggs Unit, sec. 10, T. 12 N., R. 92 W., Carbon County, Wyoming.

Pennsylvania!) and Early Permian Paleogeography, Colorado and Utah CC35

rtU.S. GOVERNMENT PRINTING OFFICE: 1992-673-049/66012

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