Geochemical, Biogeochemical, and 

Sedimentological Studies of the 

Green River Formation, 

Wyoming, Utah, and Colorado

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ERRATA

U.S. Geological Survey Bulletin 1973-A-G, Geochemical, biogeochemical, and sedimentological 

studies of the Green River Formation, Wyoming, Utah, and Colorado.

The final paragraph on page E7 of chapter E should read as follows.

The large solid in the upper left quadrant in figure 44 is iron rich but is not the 

original bladed goethite. Its massive morphology is common for iron-bearing solids 

early in most experiments. Three SEM-EDS analyses of this product gave a mean of 

753 weight percent Fe and 24.6 weight percent S. The mole ratio of Fe:S of 1.75 is 

much greater than the Fe:S ratio for known crystalline iron sulfides (table 2). 

Therefore, the exact composition of this substance is unknown, but it is probably a 

mixture of Sg, poorly crystalline or oxidized iron monosulfide(s), and massive or 

indistinguishable crystallites of a-FeOOH. Sulfur forms other than Sg were not 

distinguished on the basis of SEM-EDS analysis. X-ray diffraction of the 6N HC1- 

leached solids from this sample detected only goethite and sulfur.

Geochemical, Biogeochemical, and 

Sedimentological Studies of the 

Green River Formation, 

Wyoming, Utah, and Colorado

Edited by MICHELE L TUTTLE

Chapters A-G are issued as a single volume and are not 

available separately. Individual chapter titles are listed in 

volume contents

U.S. GEOLOGICAL SURVEY BULLETIN 1973-A-G
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: 1991

For sale by the

Books and Open-File Reports Section

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Library of Congress Cataloging-in-Publication Data

Geochemical, biogeochemical, and sedimentological studies of the Green River

formation, Wyoming, Utah, and Colorado / edited by Michele L. Turtle, 

p. cm. (U.S. Geological Survey bulletin ; 1973-A-G)

Supt. of Docs, no.: I 19.3:1973-A-G

1. Geology, Stratigraphic Paleogene. 2. Geology Wyoming.

3. Geology Utah. 4. Geology Colorado. 5. Diagenesis. 6. Green River 

Formation. I. Turtle, M.L. II. Series. 

QE75.B9 no. 1973-A-G 

[QE691.5]

557.3 s dc20 91-10904 

1557.87] CIP
CONTENTS

[Letters designate the chapters]

(A) Introduction, by Michele L. Turtle.

(B) Sulfur geochemistry and isotopy of the Green River Formation, Wyoming, Utah, 

and Colorado, by Michele L. Tuttle and Martin B. Goldhaber.

(C) A preliminary study of the carbon and nitrogen isotopic biogeochemistry of

lacustrine sedimentary rocks from the Green River Formation, Wyoming, Utah, 

and Colorado, by James W. Collister and J.M. Hayes.

(D) Trace elements in pyrites of the Green River Formation oil shales, Wyoming,

Utah, and Colorado, by Wendy J. Harrison, D.R. Pevear, and P.C. Lindahl. 

(E) An experimental study of goethite sulfidization Relationships to the diagenesis

of iron and sulfur, by Mark R. Stanton and Martin B. Goldhaber. 

(F) Effects of source, depositional environment, and diagenesis on characteristics of

organic matter in oil shale from the Green River Formation, Wyoming, Utah,

and Colorado, by Walter E. Dean and Donald E. Anders. 

(G) Petrography of iron sulfide minerals in the Green River Formation of Wyoming,

Utah, and Colorado, by Michele L. Tuttle.

Volume Contents III

Chapter A

Introduction

By MICHELE L TUTTLE

U.S. GEOLOGICAL SURVEY BULLETIN 1973

GEOCHEMICAL, BIOGEOCHEMICAL, AND SEDIMENTOLOGICAL STUDIES OF THE 

GREEN RIVER FORMATION, WYOMING, UTAH, AND COLORADO
CONTENTS

Abstract Al

Introduction Al

The Green River Formation: Background Al

Regional geology A2

Genetic models A4 

References cited A5 

Appendix Sample collection, preparation, descriptions, and analyses A7

FIGURES

1. Map showing Laramide uplifts and basins in parts of Wyoming, Utah, Colorado, 

and Idaho A3

2. Chart showing stratigraphy of cores used in study A4

TABLE

1. Resources in the Green River Formation A2

Contents
Introduction

By Michele L. Turtle

Abstract

During the Paleogene, the Green River Formation was 

deposited in two large lakes that occupied the greater Green 

River basin of Wyoming, Piceance basin of Colorado, and 

Uinta basin of Utah. New geochemical, biogeochemical, and 

sedimentological data for samples from these three 

depositional basins provide a basis for a much needed 

interdisciplinary approach with which to better understand the 

deposition and diagenesis of the Green River Formation.

INTRODUCTION

In 1922, Wilmot H. Bradley launched a long and

illustrious career committed to describing and

understanding all aspects of deposition and diagenesis of the

Paleogene (Eocene and locally Paleocene) Green River

Formation. He, better than any other geologist, realized that

the origin of the Green River rocks must be viewed as an

interdisciplinary problem. In 1970, Bradley wrote (p. 985):

A much fuller understanding of the Green River oil shale

and its organic chemistry will emerge when the

geologists, paleontologists, organic chemists, biologists,

paleolimnologists, and biogeochemists, who are now

working on it, integrate their findings with those of the

others.

Twenty years later, Bradley's recommendation is still valid. 

Many questions about the origin of the Green River 

Formation remain, and little progress has been made in 

testing conflicting hypotheses regarding its formation.

The geochemical and biogeochemical studies 

described in chapters A-G of this volume represent the 

broadest interdisciplinary collection of comparable data on 

the Green River Formation known to the author. They span 

a wide range of disciplines: inorganic and organic 

chemistry; stable isotopy of sulfide minerals (34S/32S) and

Manuscript approved for publication January 28, 1991.

organic matter ( 13C/12C, 15N/14N, 34S/32S); petrography of 

sulfide minerals; evaluation of a modern lake as a geochemical analogue for the ancient Green River lakes, Uinta 

and Gosiute; and an experimental study on iron diagenesis 

in saline, high pH environments such as the Green River 

lakes. Additional chapters are planned on detailed major and 

trace chemistry, stable isotopy of carbonate minerals 

( 13C/12C, 18O/16O), and sedimentology of core material.

A broad goal of these studies is to provide results that 

can be integrated into an interdisciplinary understanding of 

depositional and diagenetic processes, as suggested by 

Bradley. Because many of the studies used splits from the 

same sample suite (appendix), results can be directly 

compared without having to assume similar depositional or 

diagenetic conditions.

This introductory chapter describes the regional 

geology of the Green River Formation and summarizes the 

two most popular genetic models explaining its origin. For 

the most part, writers of other chapters have drawn their 

conclusions independent of results from the other studies; 

no attempt has been made at this time to integrate individual 

conclusions.

Acknowledgments. This chapter has greatly benefitted from the constructive reviews of Ronald Johnson, 

George Breit, and Karen Franczyk. Larry Trudell (Laramie 

Research Institute, Laramie, Wyoming) arranged access to 

the Blacks Fork core. The work on which this chapter is 

based was done as part of the U.S. Geological Survey 

Evolution of Sedimentary Basins Program.

THE GREEN RIVER FORMATION: 

BACKGROUND

The geology and geochemistry of the Paleogene 

Green River Formation have received considerable attention 

since the mid-1920's (for a detailed bibliography on the 

Green River Formation, see Smith, 1986). Most studies

Introduction A1
Table 1 . Resources in the Green River Formation

Shale oil.............................................................................

Sodium carbonate minerals

NaHCOs (Piceance basin)..........................................

Na3(CO3)(HCO3)*2H2O (greater Green River basin). 

Oil (Altamont-Blucbell field, Uinta basin).........................

. 1.5x10*3 barrels of oil-equivalent*

. 2.9x1010 metric tons2 

.7.5xl010 metric tons3 

. 1.9X1010 barrels4

'National Petroleum Council (1973). 

2Dyni (1974).

3Bradley and Eugster (1969). 

4Anders and Gerrild (1984).

have focused on aspects of the formation's economic 

resources (table 1), which include shale oil and sodium 

carbonate minerals. The formation is also an important 

source rock for oil-producing fields in Utah. In addition to 

its economic importance, the formation has interested 

researchers because it represents the largest and one of the 

most complex saline-lake deposits in the geologic record 

(Eugster and Hardie, 1978).

Regional Geology

The Green River Formation was deposited in lakes 

occupying three Laramide fault-bounded basins (fig. 1): the 

Uinta basin of Utah and the Piceance basin of Colorado 

were occupied by ancient Lake Uinta, and the greater Green 

River basin of Wyoming was occupied by ancient Lake 

Gosiute. These basins existed for 45 m.y. in the latest 

Cretaceous and early Tertiary. The Green River Formation 

was deposited during 10-23 m.y. of this time interval 

(Franczyk and others, in press). During transgressive stages, 

the lakes covered more than 62,150 km2 (the present-day 

areal extent of the formation) ( Bradley, 1929). Because 

most of these lacustrine rocks are bounded by an erosional 

edge, the lakes undoubtedly covered a much larger area, 

particularly in Utah. Figure 2 shows the stratigraphy of 

Green River rocks in cores from the depocenters of these 

three depositional basins.

Lacustrine sedimentation (North Horn Formation) 

began as early as the Late Cretaceous (Maastrichtian) in the 

southwestern part of the Uinta basin and in the Wasatch 

Plateau (Fouch and others, 1983). By the late Paleocene, 

much of the Uinta basin and the Wasatch Plateau was 

covered by interspersed shallow lakes and wetlands. The 

Flagstaff Limestone in the Wasatch Plateau and the 

Flagstaff Member of the Green River Formation in the 

Uinta basin formed in these environments. These units 

consist of fossiliferous limestone, nonfossiliferous dolomicrite, mudstone, sandstone, and bedded and nodular 

gypsum (Stanley and Collinson, 1979). Similar depositional

environments were present in the Piceance basin at this time 

(Johnson, 1985). An influx of clastic sediments to both 

basins occurred in the latest Paleocene through earliest 

Eocene.

Subsequently, but still during the early Eocene, large, 

relatively freshwater lakes were established in both the 

Uinta and Piceance basins (Johnson, 1985); the upper part 

of the Flagstaff Member (Bradley's Basal Tongue of the 

Green River Formation) (Johnson, 1985) was deposited in 

the Uinta basin, and the Cow Ridge Member of the Green 

River Formation was deposited in the Piceance basin 

(Johnson, 1985). During this time, the initial phase of 

lacustrine sedimentation began in the greater Green River 

basin, with deposition of the freshwater Luman Tongue of 

the Green River Formation. These and older lake deposits 

were not penetrated by the cores shown in figure 2 and are 

not included in the studies presented in these chapters.

During the middle of the early Eocene, the lakes in all 

three basins expanded. The lakes in the Uinta and Piceance 

basins coalesced across the Douglas Creek arch to form one 

large lake. Sandstone, siltstone, shale, and limestone formed 

in the marginal lacustrine areas (Uinta basin and lake 

margins in Piceance and greater Green River basins), and 

illitic-rich oil shale formed in offshore lacustrine areas 

(center of Piceance and greater Green River basins). During 

this time, the water of both Lake Uinta and Lake Gosiute 

was saline enough to kill most of the freshwater mollusc 

populations that existed during the early freshwater stages, 

but not saline enough to precipitate saline minerals. The 

offshore oil shale deposited in the Piceance basin part of 

Lake Uinta during this time is included in the Garden Gulch 

Member of the Green River Formation. Marginal lacustrine 

rocks in the Uinta and Piceance basins are generally 

included in the Douglas Creek or Anvil Points Members of 

the Green River Formation. Both offshore and marginal 

lacustrine rocks deposited in Lake Gosiute comprise the 

Tipton Shale Member of the Green River Formation. This 

early Eocene interval (beginning with the middle early 

Eocene transgression) lasted between 0.5 and 1 m.y. and 

will be referred to in these chapters as lake stage 1 (fig. 2).

A2 Studies of the Green River Formation, Wyoming, Utah, and Colorado
ABSAROKA VOLCANIC FIELD

iotLORAno.-lM? \ N

50 100 KILOMETERS

EXPLANATION

Present-day limit of Tertiary rocks Q

Upper Cretaceous outcrop

Present-day outcrop of Lower 

Cretaceous through Cambrian rocks

Present-day outcrop of Precambrian rocks

Figure 1. Laramide uplifts and basins in parts of Wyoming, Utah, Colorado, and Idaho. Core 

holes (solid circles): 1, Blacks Fork no.1 core, greater Green River basin, Wyoming; 2, 01A core, 

Piceance basin, Colorado; and 3, Coyote Wash core, Uinta basin, Utah. Modified from Johnson 

(1985).

(Lake stages are used in these chapters to define sections of 

the Green River Formation that were deposited in similar 

depositional environments. In some cases, these lake stages 

correspond to lithostratigraphic members.)

During the latter part of the early Eocene, an increase 

in the salinity of the lakes coincided with a shift to a drier 

climate (Sullivan, 1985). As a result, Lake Gosiute shrank 

significantly (Bradley, 1964, 1973; Eugster and Hardie, 

1975). Lake Uinta probably did not decrease significantly in

size, but broad shallow carbonate shelves formed around its 

margins. The initial phase of oil-shale formation occurred in 

Lake Uinta in the depocenter of the Uinta basin, and a 

compositional change in oil shale from illitic rich to 

dolomitic rich occurred in Lake Uinta in the depocenter of 

the Piceance basin. Halite (NaCl) and nahcolite (NaHCO3) 

are present in oil shale in the Piceance basin but not in the 

Uinta basin. Marginal lacustrine rocks from this interval are 

generally included in the Douglas Creek or Anvil Points

Introduction A3
460

Coyote Wash

669- 

681-

£710-

950

1,055

a E 

i"!

220

01A Blacks Fork

273 - 

291 ~

329 7

 706

774

(830)

C3 C3

460

Lake 

stage 3

Lake 

stage 2

50 100 METERS

Figure 2. Stratigraphy of cores from depocenters of greater 

Green River (Blacks Fork), Piceance (01 A), and Uinta (Coyote 

Wash) basins. Mz, Mahogany zone; [ ] Mahogany bed. 

Location of core holes shown in figure 1.

Members and consist of carbonate-rich mudstone, 

sandstone, and stromatolitic and ostracodal limestone. The 

Eocene Wilkins Peak Member of the Green River 

Formation was deposited in the greater Green River basin 

about this time and consists of greenish-gray dolomitic 

marlstone and interbedded layers of mudstone, sandstone, 

oil shale, and volcanic ash. The Wilkins Peak Member is 

rich in the saline minerals shortite [Na2Ca2(CO3)3], trona 

[Na3(CO3)(HCO3)-2H2O], and halite. This saline deposition 

lasted 4 5 m.y. and is hereafter referred to as lake stage 2 

(fig. 2).

Soon after the beginning of the middle Eocene, the 

lakes expanded. The Mahogany zone of the Parachute Creek 

Member in the Piceance and Uinta basins was deposited 

during this transgression. In the Piceance basin, the zone is 

made up of carbonate-rich, exceptionally organic carbon 

rich, laminated oil shale (for a description of sedimentary 

structures see Bradley, 1929, 1931). This unit contains the 

richest oil shale in the Green River Formation (shale oil 

yields average more than 125 L/tonne and some values are 

as high as 300 L/tonne; Trudell and others, 1970). Saline 

minerals are not present in this interval although cavities 

and brecciation formed by dissolution of salt-minerals have

been observed in the lower part of the interval (Bradley, 

1931; Trudell and others, 1970; Dyni, 1981). Oil shale in the 

Uinta basin contains less organic matter (oil yields 

averaging 80 L/tonne; Cashion, 1967) than that in the 

Piceance basin and is interlayered with mudstone, siltstone, 

and sandstone. Johnson (1985) attributed the lower organic 

contents of the Uinta basin oil shale to dilution by clastic 

sediments. In the greater Green River basin, the Eocene 

Laney Member of the Green River Formation overlies the 

Wilkins Peak Member. The Laney Member was deposited 

in freshwater or near freshwater conditions; no saline 

minerals are reported in these rocks. Laterally extensive 

beds of moderately rich oil shale (averaging 40 L/tonne; 

Smith and others, 1968) are interbedded with buff to gray 

limestone, dolomitic limestone, siltstone, mudstone, and 

volcaniclastic sedimentary rocks. Deposition of the 

Mahogany zone in the Uinta and Piceance basins is 

tentatively correlated with deposition of the Laney Member 

in the greater Green River basin. The stage, hereafter 

referred to as lake stage 3 (fig. 2), lasted about 0.6 m.y. in 

the greater Green River basin and a little less than 2 m.y. in 

the Piceance basin.

Volcaniclastic sedimentation increased in the greater 

Green River basin toward the end of lake stage 3; the basin 

was filled, and Lake Gosiute was extinguished about 44 

Ma (Mauger, 1977; Surdam and Stanley, 1979). After the 

end of Lake Gosiute, volcaniclastic wedges that prograded 

southward in the Piceance basin and intertongue with oil 

shale eventually filled this basin. Volcaniclastic 

sedimentation spread into the eastern part of the Uinta basin, 

but its depocenter was not filled (Johnson, 1985). Most of 

the part of Lake Uinta in the Uinta basin was filled about 

41 Ma (Mauger, 1977); however, in the western part of the 

basin Lake Uinta persisted into the late Eocene when 

lacustrine shale, siltstone, sandstone and zones of saline 

minerals formed.

Genetic Models

Certain features of Lakes Uinta and Gosiute are 

common to all hypotheses regarding the genesis of the 

Green River Formation, including the extremely high pH of 

lake waters, sodium-carbonate-dominated brine chemistry, 

and high salinity, especially during lake stage 2. In addition, 

all models assume that the lakes occupied essentially closed 

basins (no outflow) during most of their history. These 

basins would have coalesced into one large closed basin 

when the lakes became connected from time to time 

(Bradley, 1929).

Actively debated genetic models for the Green River 

Formation include the playa-lake and stratified-lake models. 

The playa-lake model was originally proposed by Eugster 

and Surdam (1973) for the Eocene Tipton Shale and 

Wilkins Peak Members of the Green River Formation in the 

greater Green River basin in Wyoming. The depositional

A4 Studies of the Green River Formation, Wyoming, Utah, and Colorado
model they proposed (p. 1120) was "***that of a vast 

alkaline earth playa fringing an alkaline lake...This puts the 

Green River rocks in direct genetic relationship with the 

marine tidal flats deposits of the sabkha type***" They 

stated that the best modern analog is a vastly expanded 

Deep Springs playa but without high sulfate content. Deep 

Springs playa, California, is a small playa (5 km2) that has 

a highly saline shallow lake in its center and precipitation of 

sodium chloride, sulfate, and carbonate salts around its 

perimeter.

The stratified-lake model was originally proposed by 

Bradley (1929, 1931) based on data from the Green River 

Formation in all three basins. Bradley proposed that 

minerals, including sodium carbonate minerals, were 

precipitated from the saline bottom waters of a "deep" 

stratified lake. He also suggested that during periods when 

the climate became drier, the waters in Lake Gosiute may 

have been quite shallow. Sedimentological and geochemical 

arguments supporting the stratified-lake model in the 

Piceance basin are presented in Smith (1974), Cole and 

Picard (1978), Desborough (1978), and Johnson (1981).

REFERENCES CITED

Anders, D.E., and Gerrild, P.M., 1984, Hydrocarbon generation in 

lacustrine rocks of Tertiary age, Uinta Basin, Utah Organic carbon, pyrolysis yield, and light hydrocarbons, in 

Woodward, J., Meissner, F.F., and Clayton, J.L., eds., 

Hydrocarbon source rocks of the greater Rocky Mountain 

Region: Denver, Rocky Mountain Association of 

Geologists, p. 513-529.

Bradley, W.H., 1929, The varves and climate of the Green River 

epoch: U.S. Geological Survey Professional Paper 158-E, 

p. 87-110.

___ 1931, Origin and microfossils of the oil shale of the Green 

River Formation of Colorado and Utah: U.S. Geological 

Survey Professional Paper 168, 58 p.

____ 1964, Geology of the Green River Formation and 

associated Eocene rocks in southwestern Wyoming and 

adjacent parts of Colorado and Utah: U.S. Geological 

Survey Professional Paper 496-A, 86 p.

___ 1970, Green River oil shale Concept of origin extended: 

Geological Society of America Bulletin, v. 81, p. 985-1000.

___ 1973, Oil shale formed in desert environment Green 

River Formation, Wyoming: Geological Society of America 

Bulletin, v. 84, p. 1121-1124.

Bradley, W.H., and Eugster, H.P., 1969, Geochemistry and paleolimnology of the trona deposits and associated authigenic 

minerals of the Green River Formation of Wyoming: U.S. 

Geological Survey Professional Paper 496-B, 71 p.

Cashion, W.B., 1967, Geology and fuel resources of the Green 

River Formation, southeastern Uinta basin, Utah and 

Colorado: U.S. Geological Survey Professional Paper 548, 

48 p.

Cole, R.D., and Picard, M.D., 1978, Comparative mineralogy of 

nearshore and offshore lacustrine lithofacies, Parachute 

Creek Member of the Green River Formation, Piceance

Creek Basin, Colorado, and eastern Uinta Basin, Utah: 

Geological Society of America Bulletin, v. 89, 

p. I44i_i454.

Curry, H.D., 1964, Oil-content correlations of Green River oil 

shales, Uinta and Piceance Creek basins: Intermountain 

Association of Petroleum Geologists, 13th Annual Field 

Conference, Guidebook, p. 169-171.

Desborough, G.A., 1978, A biogenic-chemical stratified lake 

model for the origin of oil shale of the Green River 

Formation An alternative to the playa-lake model: 

Geological Society of America Bulletin, v. 89, p. 961-971.

Dyni, J.R., 1974, Stratigraphy and nahcolite resources of the saline 

facies of the Green River Formation in northwest Colorado, 

i/i Murray, O.K., ed., Rocky Mountain Association of 

Geologists guide to the energy resources of the Piceance 

Creek basin, Colorado: Denver, Rocky Mountain 

Association of Geologists, p. 111-121.

___ 1981, Geology of the nahcolite deposits and associated oil 

shales of the Green River Formation in the Piceance Creek 

basin, Colorado: Boulder, University of Colorado, Ph.D. 

thesis, 144 p.

___ 1983, Distribution and origin of sulfur in Colorado oil 

shale, i/i Sixteenth Oil Shale Symposium proceedings: 

Golden, Colorado School of Mines, p. 144-159.

Eugster, H.P., and Hardie, L.A., 1975, Sedimentation in an ancient 

playa-lake complex The Wilkins Peak Member of the 

Green River Formation of Wyoming: Geological Society of 

America Bulletin, v. 86, p. 319-334.

____ 1978, Saline lakes, m Lerman, A., ed., Lakes Chemistry, 

geology, physics: New York, Springer-Verlag, p. 237-294.

Eugster, H.P., and Surdam, R.C., 1973, Depositional environment 

of the Green River Formation of Wyoming A preliminary 

report: Geological Society of America Bulletin, v. 84, p. 

1115-1120.

Fouch, T.D., Lawton, T.F., Nichols, D.J., Cashion, W.B., and 

Cobban, W.A., 1983, Patterns and timing of synorogenic 

sedimentation in upper Cretaceous rocks of central and 

northeast Utah, in Reynolds, N.W., and Dolly, E.D., eds., 

Mesozoic paleogeography of west-central United States: 

Denver, Society of Economic Paleontologists and 

Mineralogists, p. 305-336.

Franczyk, K.J., Fouch, T.D., Johnson, R.C., and Molenaar, C.M., 

in press, Cretaceous and Tertiary paleogeographic 

reconstructions for the Uinta-Piceance basin study area: 

U.S. Geological Survey Bulletin 1787-Q.

Johnson, R.C., 1981, Stratigraphic evidence for a deep Eocene 

Lake Uinta, Piceance Creek basin, Colorado: Geology, v. 9, 

p. 55-62.

___ 1985, Early Cenozoic history of the Uinta and Piceance 

Creek basins, Utah and Colorado, with special reference to 

the development of Eocene Lake Uinta, in Flores, R.M., and 

Kaplan, S.S., eds., Cenozoic paleogeography of western 

United States: Denver, Society of Economic Paleontologists 

and Mineralogists, p. 247-276.

King, P.B., 1977, The evolution of North America: New Jersey, 

Princeton University Press, p. 101-132.

Mauger, R.L., 1977, K-Ar ages of biotites from tuffs in Eocene 

rocks of the Green River Formation, Piceance Creek basin, 

Colorado: Geology, v. 3, p. 493-497.

Introduction A5
Miesch, A.T., 1976, Sampling designs for geochemical surveys 

Syllabus for a short course: U.S. Geological Survey OpenFile Report 76-772, 128 p.

National Petroleum Council, 1973, U.S. energy outlook Oil shale 

availability: Oil Shale Task Group of the Other Energy 

Resources Subcommittee of the National Petroleum 

Council's Committee on U.S. Energy Outlook, A.E. Kelley, 

chairman., 87 p.

Scott, R.W., Jr., and Pantea, M.P., 1982, Results of USGS oil-shale 

core drilling in eastern Uinta basin, Utah Coyote Wash-1 

drill hole: U.S. Geological Survey Open-File Report 

82-966, 58 p.

Smith, J.W., 1974, Geochemistry of oil shale genesis in Colorado's 

Piceance Creek basin, in Guidebook to the energy resources 

of the Piceance Creek basin, Colorado: Rocky Mountain 

Association of Geologists, 25th Annual Field Conference, 

p. 71-79.

Smith, J.W., Trudell, L.G., and Stanfield, K.E., 1968, 

Characteristics of Green River Formation oil shales at 

Bureau of Mines Wyoming Corehole No. 1: Bureau of 

Mines Report of Investigations 7172, 92 p.

Smith, M.C., 1986, Bibliography of the geology of the Green River 

Formation, Colorado, Utah, and Wyoming, to July 1, 1986: 

U.S. Geological Survey Open-File Report 86-466, 81 p.

Snyder, R.P., and Terry, S.S., 1977, Subsurface geology and 

structure in the vicinity of three drill holes in Piceance 

Creek basin, Rio Blanco County, Colorado: U.S. Geological 

Survey Open-File Report 77-271, 78 p.

Stanley, K.O., and Collinson, J.W., 1979, Depositional history of 

Paleocene-lower Eocene Flagstaff limestone and coeval 

rocks, central Utah: American Association of Petroleum 

Geologists Bulletin, v. 63, p. 311-323.

Sullivan, Raymond, 1985, Origin of lacustrine rocks of Wilkins 

Peak Member, Wyoming: American Association of 

Petroleum Geologists Bulletin, v. 69, p. 913-922.

Surdam, R.C., and Stanley, K.O., 1979, Lacustrine sedimentation 

during the culminating phase of Eocene Lake Gosiute, 

Wyoming (Green River Formation): Geological Society of 

America Bulletin, v. 90, p. 93-110.

Trudell, L.G., Beard, T. N., and Smith, J.W., 1970, Green River 

Formation lithology and oil shale correlations in the 

Piceance Creek Basin, Colorado: Bureau of Mines Report of 

Investigations 7357, lip.

Tuttle, M.L., 1988, Geochemical evolution and depositional 

history of sediment in modem and ancient saline lakes Evidence from sulfur geochemistry: Golden, Colorado School 

of Mines, Ph.D. thesis, 312 p.

A6 Studies of the Green River Formation, Wyoming, Utah, and Colorado
APPENDIX SAMPLE COLLECTION, 

PREPARATION, DESCRIPTIONS, AND 

ANALYSES

Samples for most of the studies described in the 

following chapters were collected from core from three 

wells, one in each of the three depositional basins of ancient 

Lakes Uinta and Gosiute (fig. 1). This sample distribution 

allows for comparison of depositional conditions within and 

among basins. The cores are from the centers of the basins 

and penetrate the three major depositional units referred to 

as lake stages 1 through 3. They comprise rocks from within 

each basin that are little affected by basin-margin processes. 

The lateral persistence of lithologically and geochemically 

distinct, laminated beds over many kilometers within each 

basin, and even between the Uinta and Piceance basins 

(Curry, 1964; Dyni, 1974, 1983; Johnson, 1985), indicates 

that samples from the three cores most likely are 

representative of rocks from each basin depocenter.

Sample Collection

The cores from which most of the samples used in 

these studies reported in this volume are the Energy 

Research and Development Administration Blacks Fork 

core no. 1 (BF), Green River basin, Wyoming; U.S. Bureau 

of Mines core 01A (01A), Piceance basin, Colorado; and 

U.S. Geological Survey Coyote Wash no. 1 core (CW), 

Uinta basin, Utah (locations shown in fig. 1). All of the 

cores are stored at the U.S. Geological Survey Core 

Research Center, Lakewood, Colorado. The Blacks Fork 

core was described by L.G. Trudell of the Laramie Energy 

Research Center (written commun., 1983), the OlA core by 

Snyder and Terry (1977), and the Coyote Wash core by 

Scott and Pantea (1982). Geologic sections of the cores are 

shown in figure 2. The three cores represent sediment 

deposited during 14 of the 21 m.y. of the Green River lake 

system (Tuttle, 1988).

Hierarchical, unbalanced analysis of variance 

sampling designs (Miesch, 1976) allowed for objective 

sampling and optimal coverage of the cores given the 

limitations on the number of samples that could be

analyzed. Samples from the Piceance basin OlA core were 

collected for the major- and trace-element chemistry study 

using a design that divided the core into 30.5-m sections; 

within every other 3.5-m section in the 30.5-m section, two 

sample intervals were chosen; 20 percent of the ground 

samples were split for analytical duplication. A total of 190 

samples were analyzed in this study. For most of the other 

studies the cores were divided into sections 18m long. One 

sample was collected approximately 15 m from the top of 

each cored section; a second sample was collected 6 m from 

the top of about 40 percent of the sections; 10 percent of the 

samples were split after grinding for analytical duplication. 

Thirty-five samples were collected from the Blacks Fork 

core, 41 from the OlA core (a subset of those collected for 

the major- and trace-element chemistry study), and 39 from 

the Coyote Wash core. Each sample was about 3 cm long 

and lithologically homogeneous. Because some Green 

River depositional rates were very slow (0.04 mm/yr for 

organic-carbon-rich oil shale to 0.17 mm/yr for organiccarbon-poor marlstone; Bradley, 1929), each sample 

represents a depositional time span on the order of hundreds 

of years.

Sample Descriptions, Preparation, and Analyses

The presence of visible sulfide minerals, organic 

material, tuffaceous material, and laminations and the 

sample color and grain size were noted in the OlA subset, 

Coyote Wash, and Blacks Fork samples prior to any further 

treatment. Descriptions are given in table Al. The samples 

were described as being calcareous if they reacted 

vigorously with 10 volume percent HC1 and dolomitic if 

they reacted slowly with HC1. Before grinding, a plug was 

removed from the sample for thin-section preparation. 

Samples were then ground with ceramic plates to pass a 115- 

mesh screen (130-|0,m mean diameter) and splits for each 

study removed.

The whole-rock mineral composition for each sample 

was determined on a random powder mount using an X-ray 

diffractometer with nickel-filtered CuKa radiation. The 

mineral analyses are qualitative because standards were not 

run with the samples. X-ray diffraction results are presented 

in table Al.

Introduction A7
Table A1 . Macroscopic descriptions and X-ray diffraction mineralogy for 35 samples from the Blacks Fork core no. 1, 

Wyoming, 41 samples from the 01A core, Colorado, and 39 samples from the Coyote Wash no.1 core, Utah

[Abbreviations used in the table are explained below. X-ray mineralogy is listed in order of frequency, not abundance. Core holes: BF, Blacks Fork 

no. 1; 01,01A; CW, Coyote Wash]

Abbreviations

anal

aptt

arag

b

brec

brn

c

cal

calc

cly

cr-gr

d

daw

dc

dk

dol

dolo

ds

analcime

apatite

aragonite

bladed

brecciated

brown

continuous laminae

calcareous

calcite

clay (other than illite)

coarse grained

disseminated

dawsonite

discontinuous laminae

dark

dolomitic

dolomite

distinct

f

f-gr

fos

grn

gy

hlit

horn

ill

inter

Kfld

1

lam

It

m

mdst

muse

Nafld

nahc

faint

fine grained

fossiliferous

green

gray

halite

hornblende

illite

interbedded

potassium feldspar

lens

laminated

light

medium

mudstone

muscovite

sodium feldspar

nahcolite

org organic matter rich

ost ostracods

po pyrrhotite

py pyrite

qtz quartz

rd red

sh shortite

sid siderite

sltst siltstone

slty silty

ss sandstone

sul sulfides

tr trona

tuf tuffaceous

v vuggy

X trait present

Depth Sample description

Sample (meters) sul org cal dol tuf lam Color, grain size Mineralogy

BLACKS FORK CORE NO. 1, GREATER GREEN RIVER BASIN, WYOMING

Lake stage 3

BF-1

BF-2

BF-3

BF-4

BF-5

BF-6

BF-7

63 d

70 d X X

91 d,l X X

98 X

107 X

125 d

135 d,l X X

X X

X

X X

X

X

It gy sltst to f-gr ss

ds,c dk brn mdst

f,c m brn mdst

ds.dc m grn to m brn mdst

f,c It brn mdst

ds.dc It-m gy sltst

ds,c m-dk gy-brn mdst

qtz,calc,dolo,Kfld,Nafld,anal,py,

ill

qtz,calc,dolo,Kfld,anal,py,ill

qtz,calc,dolo,Kfld,anal,py,ill

qtz,calc,dolo,Kfld,anal,ill,cly

qtz,calc,dolo,Kfld

qtz,calc,Kfld,anal,py,cly

qtz,calc,dolo,Nafld, anal, py

Lake stage 2

BF-8

BF-9

BF-10

BF-11

BF-12

BF-13

BF-14

BF-15

BF-16

BF-17

BF-18

BF-19

BF-20

BF-21

144 d,b X

162 d X

180 d X

189 X

199 d X

217 d X

235 X X

244 XX

254 d,l

272 X

290 X

299 XX

309

333 d X

X

X

X

X

X

dk grn mdst to sltst

m-dk brn mdst

dk brn mdst

m grn to gy mdst

m grn to gy mdst

m grn mdst

m to dk brn mdst

d-c It to dk brn mdst

qtz,dolo,Kfld,sh,po

qtz,dolo,Nafld,sh,py

qtz,dolo,sh,py

qtz,dolo,Kfld,Nafld,sh,ill

qtz,calc,dolo,Kfld,Nafld,sh,py,po,

ill

qtz,dolo>Nafld,sh,py,ill,cly

qtz,calc,dolo,Nafld,anal,sh,ill,cly

qtz,dolo,Nafld,sh

d-dc m grn mdst to m br to grn qtz,calc,dolo,Kfld,Nafld,anal,ill,

sltst

m grn mdst

dk grn mdst

f-c m brn to grn mdst

grn slty mdst

d-c m brn to grn mdst

cly

qtz,calc,dolo,Kfld,Nafld,anal,ill,

cly

qtz,calc,dolo,Kfld,Nafld,anal,ill,

cly

qtz,calc,dolo,Nafld,ill

qtz,calc,Nafld,anal,ill,cly

qtz,dolo, Kfld,Nafld,sh,py ,ill

A8 Studies of the Green River Formation, Wyoming, Utah, and Colorado
Table A1. Continued

Sample

Depth

(meters)

Sample description

sul org cal dol tuf lam Color, grain size Mineralogy

Lake stage 2 Continued

BF-22

BF-23

BF-24

BF-25

BF-26

BF-27

BF-28

BF-29

BF-30

344

353

363

381

400

409

418

437

454

d

d,l

d,b

X

X

X

X

X

X

X

X

X

X

X

f-c

d-c

X

X

m grn to gy mdst

m grn to brn mdst

dk brn mdst

m grn mdst to m gy slty

mdst

It grn mdst

m grn to gy slty mdst

m grn to gy slty mdst

m grn mdst

It grn to gy slty mdst

qtz,dolo,sh,tr

qtz,dolo,Nafld,anal,cly

qtz,dolo,Nafld,sh,ill

qtz,dolo,Kfld,Nafld,sh,tr,py

qtz,dolo,Nafld,sh,cly

qtz,dolo,Kfld,Nafld,sh,ill,cly

qtz,dolo,Kfld,Nafld,sh,ill,cly

qtz,dolo,Kfld,Nafld,sh,py,ill

qtz,calc,dolo,Kfdl,Nafld,anal,py,

po.ill.cly

Lake stage 1

BF-31

BF-32

BF-33

BF-34

BF-35

464

473

491

500

509

d,l

d,l

d

d,l

d

X

X

X

X

X

X

X

X

X f-dc

X X d-c

d-c

X X d-c

It brn mdst

dk brn mdst

dk gy to brn mdst

dk brn ost mdst

dk brn fos mdst

qtz,calc,dolo,Kfld,Nafld,anal,py,

ill

qtz,calc,dolo,Kfld,anal,py,ill

qtz,calc,dolo,Nafld,sid,py,ill,cly

qtz,calc,dolo,Kfld,anal,py,ill, cly

qtz,calc,arag,Nafld,py,cly

OIA CORE, PICEANCE BASIN, COLORADO

Lake stage 3

01-6

01-8

01-12

01-14

01-18

01-24

01-26

268

272

285

290

304

321

328

d

d

X

X

X

X

X

X

X

X

X

X

X ds,c

X ds,c

X ds.c

X X ds,c

ds,c

f.dc

f.dc

dk brn mdst

It brn mdst

m brn mdst

dk brn mdst

dk brn mdst

m brn mdst, v

It to m brn mdst, v

qtz,calc,arag,dolo,Nafld,anal,cly

qtz.dolo, Kfld.daw.cly

qtz,dolo, Kfld,Nafld,daw,py ,cly

qtz.calc.dolo, Kfld,Nafld,daw,py ,

cly

qtz,calc,dolo,Kfld,Nafld,anal,py,

cly

qtz,calc,dolo,Kfld,Nafld,anal,ill,

cly

qtz,calc,dolo,Kfld,Nafld,cly

Lake stage 2

01-30

01-36

01-42

01-48

01-54

01-59

01-62

01-65

01-72

01-74

01-78

01-84

340

358

378

395

415

429

437

448

468

474

486

504

d

d

d

d

d,b

d,l

d

1

d

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X X ds,c

X ds.c

X ds,c

X

X f,c

ds.cd

ds.dc

X ds.dc

X ds,dc

ds,dc

m brn mdst

dk brn mdst

m brn mdst

brec It brn mdst

dk brn mdst

inter m brn mdst and nahc

inter dk brn mdst

It brn mdst

m bra mdst

m brn mdst

dk brn mdst

dk bra mdst, bedded hlit,

nahc

qtz,calc,dolo,Kfld,Nafld,cly

qtz,calc,dolo,Kfld,Nafld,ill

qtz,dolo,Kfld,Nfld,anal,ill,cly

qtz,dolo,Kfld,Nafld,py,cly

qtz,dolo,Kfld,Nafld,anal,py,ill

qtz,dolo,Kfld,Nafld,anal,daw,

nahc, py.ill.cly

qtz,calc,dolo,Nafld,nahc,hlit

qtz,dolo,Kfld,Nafld,daw,nahc,

py.cly

qtz,calc,dolo,Kfld,Nafld,daw,

nahc.py

qtz.calc.dolo, Kfld,Nafld,daw,

nahc.py.cly

qtz,dolo, Kfld,Nafld,daw,py,cly

qtz,calc,dolo,Nafld,daw,nahc,

hlit,cly

Introduction A9
Table A1. Continued

Sample

Depth

(meters)

Sample description

sul org cal dol tuf lam Color, grain size Mineralogy

Lake stage 2 Continued

01-86

01-89

01-96

01-98

01-102

01-108

01-110

01-113

01-120

01-123

01-126

01-132

01-134

01-138

01-144

510

521

541

547

559

578

583

595

614

625

632

651

657

669

687

d

d

1

d,l

d

1

d

d

d

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

ds,c

ds.c

ds,dc

ds.dc

ds,c

ds,c

dk brn mdst

f-gr nahc, It brn mdst

dk brn mdst

dk brn mdst

dk brn mdst

It to m brn mdst

dk brn mdst and cr-gr nahc

dk brn mdst

cr-gr nahc, dk brn mdst

dk brn mdst

cr-gr nahc

inter nahc, m brn mdst

dk brn mdst

It brn mdst

m brn mdst and cr-gr nahc

qtz,dolo, Kfld,Nafld,daw,nahc,

hlit,py,cly

qtz,dolo,Nafld,daw,nahc,hlit, ill,

cly

qtz,dolo,Kfld,Nafld,daw,py,cly

qtz,dolo,Kfld,Nafld,daw,py,cly

qtz,dolo,Kfld,daw,py,cly

qtz,calc,dolo,Kfld,daw,py

qtz,dolo,nahc,ill,cly

qtz,dolo,Kfld,Nafld,daw,py,cly

qtz,dolo, Kfld,Nafld,daw,nahc,ill

qtz,dolo, Kfld,Nafld,daw,py

qtz,daw,nahc,ill,cly

qtz,dolo,Kfld,Nafld,daw,nahc,cly

qtz,calc,dolo, Kfld,Nafld,daw,

nahc,py,cly

qtz.dolo, Kfld,Nafld,daw,py,cly

qtz.dolo, Kfld,Nafld,nahc,py,cly

Lake stage 1

01-150

01-154

01-155

01-162

01-165

01-168

01-172

705

717

728

743

755

760

773

1

d

X

X

X

X

X

X

X

X

X

X

X

X

X

X

ds,c

ds.c

ds,c

f,c

ds,c

f,c

dk brn to gy mdst

dk brn mdst

dk brn to gy mdst

dk gy to brn mdst

dk brn mdst

m gy to brn ost mdst

m gy to brn mdst

qtz,calc,dolo,Nafld,py,ill

qtz,dolo,Nafld,py,ill

qtz,dolo,Nafld,py,ill,

qtz,dolo,Kfld,Nafld,py,ill,cly,

aptt

qtz,dolo,Nafld,py,ill,cly

qtz,calc,dolo,py,ill

qtz,dolo,Nafld,py,ill

COYOTE WASH NO. 1 CORE, UINTA BASIN, UTAH

Lake stage 3

CW-38

CW-37

CW-36

CW-35

CW-34

CW-33

CW-32

CW-31

CW-30

CW-29

571

587

597

606

624

642

652

661

670

681

d

d

d

d

d

d

d

1

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

ds,c

ds.c

ds,c

X ds,c

ds,c

X ds,c

X f.dc

f.dc

ds,dc

X ds,dc

It to m brn mdst

m brn mdst

m to dk brn mdst

m to dk brn mdst

dk brn mdst

m rd, m brn mdst

m to dk brn mdst

m to dk brn mdst

m brn mdst

dk brn mdst

(Mahogany bed)

qtz,calc,dolo,Nafld,sid

qtz,calc,dolo,Nafld,ill

qtz,calc,dolo, Kfld,Nafld,sid,ill

qtz,calc,dolo,Nafld

qtz,calc,dolo,Nafld,daw,sid,py, ill

qtz,calc,dolo,Kfld,Nafld,daw,py

qtz,calc,dolo,Nafld,daw,py,cly

qtz,calc,dolo,Kfld,Nafld,daw,py

qtz,calc,dolo,Nafld,daw,sid

qtz,calc,dolo,Nafld,daw,py

Lake stage 2

CW-28

CW-27

CW-26

CW-25

CW-24

CW-23

697

706

716

726

734

752

d

d

d

d

X

X

X

X

X

X

X

X

X

X

X

X

ds,c

ds,c

ds.dc

ds,c

ds,c

It brn mdst

m to dk brn mdst

It to m brn mdst

It brn mdst

It to m brn mdst

m brn mdst

qtz,calc,dolo,Kfld,Nafld,anal,ill

qtz,calc,dolo,Kfld,Nafld,anal,ill

qtz,calc,dolo,Nafld

qtz,calc,dolo,Kfld,Nafld,anal,

daw.ill

qtz,calc,dolo,Kfld,Nafld,daw

qtz,calc,dolo,Kfld,Nafld,anal,

sid.py

A10 Studies of the Green River Formation, Wyoming, Utah, and Colorado
Table A1. Continued

Sample

Depth

(meters) sul

Sample description

org cal dol tuf lam Color, grain size Mineralogy

Lake stage 2 Continued

CW-22

CW-21

CW-20

CW-19

CW-18

CW-17

CW-16

CW-15

CW-14

CW-13

CW-12

CW-11

CW-10

CW-9

CW-8

762

771 d,b

789 d

807 d

815 d,b

825

843 d

862

871 d

880 d,l

898

908

917 d

926

935 d

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

X

ds,c

f,c

ds,c

ds,c

ds.dc

ds.c

ds.c

f.dc

ds.dc

ds,c

ds,dc

It brn mdst

m brn mdst

m brn mdst

It gy mdst

It to m brn mdst

It to m brn mdst

m to dk brn mdst

It grn, gy mdst

It to m brn mdst

dk brn mdst

It to m brn ost mdst

m gr to brn mdst

m gr to m brn mdst

m gr mdst

dk grn gy slty mdst

qtz,calc,dolo,Kfld,Nafld,daw

qtz,calc,dolo,Kfld,Nafld,anal,

daw.sid.ill

qtz,dol,Kfld,Nafld,anal,ill

qtz,calc,dolo,Kfld,Nafld,daw,

musc.ill

qtz,calc,dolo,Kfld,Nafld,anal,ill

qtz,calc,dolo,Kfld,anal,ill,cly

qtz,calc,dolo,Kfld,Nafld,anal,py,

m,cly

qtz,calc,dolo,Kfld,Nafld,anal,

horn.ill.cly

qtz,calc,dolo,Kfld,Nafld,anal,ill,

cly

qtz,dolo, Kfld.anal.py.ill

qtz,calc,dolo,Kfld,Nafld,anal,ill

qtz,calc,dolo,Kfld,Nafld,anal,ill

qtz,calc,dolo,Kfld,Nafld,anal,

py,ill,cly

qtz,calc,dolo,Kfld,Nafld,anal,

m,cly

qtz,calc,dolo,Kfld,Nafld,anal,

py.ill.cly

Lake stage 1

CW-7

CW-6

CW-5a

CW-5

CW-4

CW-3

CW-2

CW-1

953

971 d

981

990

1,008

1,026

1,035

1,043 d

X

X

X

X

X

X

X

X

X

X

-

X

ds,c

ds,dc

ds,c

X

f.dc

m brn to gy mdst

m to It gy slty mdst

m gy slty mdst

m brn mdst

It gy sltst

m brn mdst

m gy brn mdst to sltst

dk to m brn mdst

qtz,calc,dolo,Kfld,Nafld,anal,ill,

cly

qtz,calc,dolo,Kfld,Nafld,anal,py,

qtz,calc,dolo,Kfld)Nafld,anal,sid,

qtz,calc,dolo,Kfld,Nafld,anal,ill,

cly

qtz,calc,dolo,Kfld,Nafld,anal,ill

qtz,calc,dolo,arag,Nafld,ill,cly

qtz,calc,dolo,Kfld,Nafld,anal,sid,

ill.cly 

qtz,dolo,Kfld,Nafld,anal,py,cly

Introduction A11

Chapter B

Sulfur Geochemistry and Isotopy of the 

Green River Formation, 

Wyoming, Utah, and Colorado

By MICHELE L TUTTLE and MARTIN B. GOLDHABER

U.S. GEOLOGICAL SURVEY BULLETIN 1973

GEOCHEMICAL, BIOGEOCHEMICAL, AND SEDIMENTOLOGICAL STUDIES OF THE 

GREEN RIVER FORMATION, WYOMING, UTAH, AND COLORADO
CONTENTS

Abstract Bl

Introduction Bl

Sulfur in the Green River Formation Bl

Methods B3

Sampling methods B3

Analytical methods B3 

Results B4 

Discussion B5

Sulfate and its reduction in Lakes Uinta and Gosiute B5

Sulfidization products BIO

Sulfur geochemistry of Lakes Uinta and Gosiute Bll 

Applications to geologic problems B12 

References cited B13 

Appendix 1 Analytical data B15 

Appendix 2 Calculation of gypsum thickness B20

FIGURES

1. Map showing extent of rocks of Green River Formation and location of cores 

used in study B2

2. Diagram showing method used to collect and analyze forms of sulfur B4

3. Depth profiles for total sulfur and 534S values of disulfides in depocenter 

cores B5

4. Rayleigh fractionation plot and frequency histogram for S^S values of disulfides in depocenter cores B7 

5-6. Graphs showing:

5. Regression lines for sulfur in sulfide minerals versus organic carbon for 

marine, freshwater, and euxinic sediments and for rocks of Green River 

Formation B9

6. 834S values of organosulfur versus disulfide sulfur BIO 

7-8. Charts summarizing sulfur geochemistry of the Green River Formation during:

7. Euxinic deposition Bll

8. Shallow deposition Bll

9. Diagram showing mean 834S values in Green River Formation during each lake 

stage B12

TABLES

1. Summary of lake stages represented by samples in this study B3

2. Average chemical and isotopic parameters for samples from the three lake stages 

in the Green River Formation B6

3. Dissolved water-column sulfate, sediment sulfur concentrations, and 534Sdisulfide 

values in samples from the Green River Formation and other lacustrine 

environments B8

Contents
Sulfur Geochemistry and Isotopy of the Green River 

Formation, Wyoming, Utah, and Colorado

By Michele L. Tuttle and Martin B. Goldhaber

Abstract

Sulfur geochemistry was used to gain insight into 

conditions during deposition and diagenesis of the lacustrine 

Green River Formation (Paleogene). The abundance and 

isotopic composition of major forms of sulfur (sulfate, monosulfides, disulfides, and organosulfur) in samples from three 

cores indicate that during the time represented by the sampled 

core interval (about 8 m.y.) the two ancient Green River lakes 

(Lakes Uinta and Gosiute) intermittently had similar 

depositional conditions.

Both lakes received a continuous supply of sulfur, most 

of which was supplied by inflow. An intermittent source of 

sulfur was associated with volcanic ash. This volcanic sulfur 

was depleted in 34S relative to the sulfur reservoir in the lakes, 

and its isotopic signature is particularly evident in the samples 

deposited during the low stands of Lake Gosiute (shallow 

deposition).

The enrichment of 34S in sulfide minerals and organic 

matter in the Green River Formation supports the concept of 

reduction of sulfate to hydrogen sulfide (H2S) in anoxic lake 

bottom water (monimolimnion) or in anoxic pore water. 

During periods of stratification, the H 2S:sulfate ratio of the 

monimolimnion was high. Throughout the saline stage of Lake 

Uinta, long-term stratification resulted in the progressive 

enrichment of 34S in the H2S reservoir. During the same period 

in Lake Gosiute, progressive enrichment did not occur and 

534S values in rocks deposited during this time are highly 

variable. In both lakes, the high pH of the pore water likely 

promoted sulfidization of organic matter and kinetically 

inhibited sulfidization of iron.

These results support the hypotheses that a connection 

existed between the two Green River lakes coincident with 

progressive infilling of the basins; that pH affects the sulfidization of organic matter and iron; that sulfidization of organic 

matter affects the sulfur content of hydrocarbons generated 

from lacustrine source rocks; and that existing sulfur models 

and paleosalinity indicators may not be applicable for ancient 

lacustrine rocks.

INTRODUCTION

Rocks of the Paleogene (Eocene and locally Paleocene) Green River Formation accumulated in two large 

lakes (fig. 1): Lake Gosiute in the greater Green River basin 

(southern Wyoming) and Lake Uinta in the Piceance 

(northwestern Colorado) and Uinta basins (northeastern 

Utah). Sulfur geochemistry has been used extensively to 

interpret deposition and diagenesis in marine and 

freshwater-lacustrine rocks because of the sensitivity of 

sulfur abundance, residence, and isotopic composition to 

chemical and biological changes. An investigation of sulfur 

geochemistry of the Green River Formation therefore 

should increase our understanding of the overall geochemical and biogeochemical conditions controlling the 

accumulation and preservation of Green River sediments.

Acknowledgments. This chapter has benefited 

greatly from the thorough reviews of Linda Bonnell and 

Paul Lillis. We wish to thank Robert Rye, U.S. Geological 

Survey, for access to his isotope laboratory. This work was 

conducted as part of the U. S. Geological Survey Evolution 

of Sedimentary Basins Program.

SULFUR IN THE GREEN 

RIVER FORMATION

Most published sulfur data on the Green River 

Formation are from rocks in the Uinta and Piceance basins. 

Reported concentrations of total sulfur range from 0 weight 

percent in algal carbonate rocks in the Uinta basin, Utah 

(Boyer and Cole, 1983), to 5 weight percent in organiccarbon-rich shale from the saline zone in the Piceance basin, 

Colorado (Dyni, 1983). The sulfur is known to be present as 

sulfate, monosulfide, and disulfide minerals, as well as 

organosulfur. Primary sulfate minerals have not been 

identified in the formation (Milton and Eugster, 1959); the

Sulfur Geochemistry and Isotopy B1
113 106°

ABSAROKA VOLCANIC FIELD

REATER: GREEN

.UINTA-.---X-BAS

EXPLANATION

Present-day limit of Tertiary rocks

Upper Cretaceous outcrop

//// Present-day outcrop of Lower

Cretaceous through Cambrian rocks

;::V. Present-day outcrop of Precambrian rocks

50 100 KILOMETERS

Figure 1. Map showing extent of rocks of Green River Formation (Tertiary rocks) and location of 

cores used in study. Core holes (solid circles): 1, Blacks Fork core; 2, 01 A core; and 3, Coyote 

Wash core. Modified from Johnson (1985).

abundance of sulfate accounts for less than 4 percent of the 

total sulfur (Stanfield and others, 1951; Smith and others, 

1964). Monosulfide minerals (pyrrhotite and wurtzite) and 

disulfide minerals (pyrite and marcasite) have been 

described by Milton and Eugster (1959), Pabst (1970), Cole 

and others (1978), Cole and Picard (1981), and Melchior 

and others (1982); monosulfides are reported to be minor

sulfur phases. According to Stanfield and others (1951) and 

Smith and Young (1983), 50-90 percent of the sulfur in 

most Green River oil shale is in pyrite and the remaining 

10-50 percent is as organosulfur. Sulfur-containing organic 

compounds in the Green River oil shale include thiophenes, 

benzothiophenes, and polycyclic thiols (Ingram and others, 

1983).

B2 Green River Formation, Wyoming, Utah, and Colorado
Tabl« 1 . Summary of lake stages 1,2 (2B in the Green River basin), and 3 represented by samples in this study 

[Samples are from the Blacks Fork core no. 1 (BF), greater Green River basin, Wyoming; 01A core (01A), Piceance basin, 

Colorado; Coyote Wash no. 1 core (CW), Uinta basin, Utah. Number of samples in parentheses is number of samples for which 

sulfur forms and their isotopic compositions were analyzed]

Lake 

stage Green River basin

Greater 

Piceance basin Uinta basin

Fresh water 

Transgression 

High productivity 

Organics preserved 

Open-water fades

High salinity 

Regression 

Low productivity 

Organics mineralized 

Shallow-water fades

Fresh water 

Transgression 

High productivity 

Organics preserved 

Open-water fades

Moderate salinity 

Transgression 

Low productivity 

Organics preserved 

Open-water fades

High salinity 

Regression 

High productivity 

Organics preserved 

Open-water fades

Moderate-high salinity 

Transgression 

High productivity 

Organics preserved 

Open-water fades

Fresh water 

Transgression 

Low productivity 

Lake-margin fades

Number of 

samples

5 

(5)

7 

(4)

8 

(4)

High salinity 

Regression

Moderate productivity 

Organics preserved 

Open-water fades

Number of 

samples

23 

(23)

27 

(11)

21 

(15)

Low-moderate salinity 

Transgression 

High productivity 

Organics preserved 

Open-water fades

Number of 

samples

7 

(7)

7 

(5)

10 

(6)

METHODS

Sampling Methods

The sampling design, cores, and samples used in this 

study are described in the appendix in Turtle (this volume). 

A suite of 115 samples from the three Green River 

depositional basins was analyzed 35 samples from the 

Blacks Fork core, greater Green River basin, Wyoming; 41 

from the 01A core, Piceance basin, Colorado; and 39 from 

the Coyote Wash core, Uinta basin, Utah (fig. 1).

The sample suite represents continuous lacustrine 

deposition during three major evolutionary stages of the 

lakes during the Eocene. These stages and their associated 

rock types are discussed in detail in Turtle (this volume) and 

summarized in table 1. The first stage (lake stage 1) 

represents freshwater to brackish-water deposition during 

early expansion of the lakes. Lake stage 2 was a very saline 

stage during which the areal extent of Lake Uinta changed 

little (Johnson, 1985) but that of Lake Gosiute drastically 

decreased (Bradley, 1973; Eugster and Hardie, 1975). 

Salinity increases were probably initiated by a regional 

climatic shift to more arid conditions (Sullivan, 1985). 

During lake stage 3, the lakes expanded again. Lake Gosiute 

became a freshwater lake and Lake Uinta became less

saline, probably brackish. During this stage, the richest 

Green River oil shales were deposited in Lake Uinta.

Analytical Methods

All 115 samples were analyzed for total sulfur, 

organic and carbonate carbon, and total iron. Seventy-nine 

of the samples were analyzed for HCl-soluble iron (used to 

calculate reactive iron concentrations). These last four 

variables were determined because of the importance of 

organic carbon on sulfate reduction, carbonate carbon on 

dilution of iron and sulfur in rocks, and iron species on 

sulfur processes in sediment. Sulfate, monosulfide sulfur 

(acid-volatile sulfur), disulfide sulfur, organosulfur, and their 

isotopic compositions were analyzed in 79 of the samples. 

The analytical techniques are detailed in Turtle and others 

(1986) and Turtle (1988).

Samples ground to pass a 115-mesh screen (130 ^im) 

were used in all analytical procedures. Concentrations of 

total sulfur, carbonate carbon, and organic carbon were 

measured using a commercially available apparatus 

consisting of an induction furnace coupled to an infrared 

detection system (sulfur) and thermal conductivity cell 

(carbon). Total iron concentrations were determined by 

fusing the sample with lithium metaborate (Ingamells,

Sulfur Geochemistry and Isotopy B3
Filtrant

'4 (aq)

+2Ag+(aq)

+Ba2+

(aq) +2Ag- (aq) (aq)

Ag2s-j sj BaSO «s,w4 jsj Ag2S (S) 

volatile ^sulfate ^disulfide

Figure 2. Method used to collect and analyze forms of sulfur. Modified from Tuttle and others (1986).

BaSO4

(S)

Organic

1970) and dissolving the fusion product in 10 volume 

percent HC1. Both total iron and HCl-soluble iron (see 

procedure below) were measured using flame atomic 

absorption spectrometry. Sulfur species were sequentially 

collected and gravimetrically analyzed by the separation 

scheme described in Tuttle and others (1986) and 

summarized in figure 2. The scheme removes monosulfide 

sulfur as H2S gas with hot 6W HC1 while solubilizing the 

sulfate sulfur and reactive iron not bound in disulfides; it 

reduces and removes disulfide sulfur as H2S gas with a hot 

Cr^-HCl solution; and it collects all residual sulfur (defined 

as organosulfur) using an Eschka-fusion procedure. Sulfur 

products from the separation are converted to SO2, which is 

then isotopically analyzed using mass spectrometry. All 

isotope results are reported relative to the troilite phase of 

the Canon Diablo meteorite (CDT standard).

RESULTS

Chemical (dry-weight basis) and sulfur isotopic data 

are tabulated in appendix 1. Because key geochemical 

relationships can be obscured by the dilution effect that

authigenic carbonate minerals may have on other element 

abundances, carbonate-free chemical data are used 

hereafter. Averages of chemical (carbonate-free basis) and 

isotopic data for each lake stage in each basin are shown in 

table 2. In the greater Green River basin core, three samples 

in lake stage 2 strongly resemble (both geochemically and 

physically) those in stages 1 and 3. These samples are 

interpreted to represent fresher water conditions 

(transgressive stages, lake stage 2A) than those inferred for 

other stage 2 samples from the core (shallow stages, lake 

stage 2B).

Depth profiles for total sulfur (weight percent) and 

834S (%o, per mil) of disulfides are presented in figure 3. 

The abundance of sulfur in Piceance basin samples exhibits 

no apparent trend with time, and values cluster around 1 

weight percent. Similarly, the abundance of sulfur in Uinta 

basin samples shows no time-dependent trends, and values 

of about 0.5 weight percent are about half those of Piceance 

basin samples. In contrast, the abundance of sulfur in the 

greater Green River basin shows a distinct trend with time: 

abundances are significantly lower in lake stage 2B relative 

to stages 1 and 3. In the lower part of lake stage 2, one 

sample from each basin has a very high sulfur content (fig.

B4 Green River Formation, Wyoming, Utah, and Colorado
3). Although the significance of these anomalies is not 

known, the samples most likely represent an event that 

affected the entire Green River lake system.

The predominant sulfur species in most of the Green 

River samples is iron disulfide (pyrite and less frequently 

marcasite) (Turtle, this volume). All disulfide minerals are 

enriched in 34S relative to the CDT standard (fig. 3, table 2); 

in many samples, this enrichment is very large. The depth 

profile of 834S values in Lake Uinta stage 2 disulfides shows 

that the sediments incorporated progressively more 34S with 

time. The 834S values for stage 2 of Lake Gosiute are 

likewise enriched in 34S but show no systematic enrichment 

with decreasing depth. Instead, 34S-depleted values occur 

near ash layers. 34S-depleted values are present in the late 

portion of stage 2 in the greater Green River basin and in the 

early portion of stage 3 in the Piceance and Uinta basin. The 

magnitude of this depletion decreases from the greater 

Green River basin to the Piceance basin and from the 

Piceance basin to the Uinta basin.

DISCUSSION

The hundreds of meters of lithologically and mineralogically similar lacustrine rocks indicate that in their 

depositional centers Lakes Uinta and Gosiute were 

continuous bodies of water through time. Similar stages of 

geochemical evolution occurred within each lake: a saline 

stage 2 bracketed by two expansive, fresher water stages 1 

and 3. Stage 1 and much of stage 2 are time correlative 

between the two lakes; however, the time relationship 

between stage 3 of Lakes Uinta and Gosiute is less certain 

(R.C. Johnson, written commun., 1989). Sulfur chemical 

and isotopic data from our study indicate that (1) although 

the lakes were indeed similar in many aspects, key 

differences are observed, and (2) the geochemistry of the 

lakes slowly evolved during the 5.1-8 m.y. intervals of the 

three lake stages.

Sulfate and Its Reduction in 

Lakes Uinta and Gosiute

Sulfate-bearing saline minerals are rarely found in 

Green River rocks, although nonsulfate saline minerals are 

common at certain stage 2 horizons. The paucity of sulfatesaline minerals has been attributed to low sulfate in lake 

water in which brine evolution results in a sulfate-poor, 

sodium-bicarbonate-rich water (Eugster and Hardie, 1975, 

1978). This hypothesis, however, is not supported by sulfur 

abundances in Green River sedimentary rocks; the rocks 

deposited in basin depocenters contain more sulfur than 

sediment deposited in sulfate-rich environments such as the 

oceans and sulfate-rich saline lakes (table 3).

0123 4 5 16

Wt% Total S (cf)

10 20 30 40 50

CO

oc

UJi

600

700

800

Lake stage

- 900

UJ 

O

1,000

1,100

0 1 2 601020304050

Wt% Total S (cf) g34s

Figure 3. Depth profiles for total sulfur (carbonate-free basis) 

and 534S values of disulfides in depocenter cores of Green 

River Formation. A, Greater Green River basin; B, Piceance 

basin; C, Uinta basin. Lake stages are designated on plots.

Another approach to evaluating lake sulfate 

concentrations is to estimate the amount of bedded gypsum 

(CaSO4-2H2O) that could have preciptated during the saline

Sulfur Geochemistry and Isotopy B5
Table 2. Average chemical and isotopic parameters for samples from the three lake stages in the Green River Formation of 

the greater Green River basin, Wyoming, Piceance basin, Colorado, and Uinta basin, Utah

[All chemical values (except carbonate carbon) are on a carbonate-free, dry-weight basis. All isotopic data are reported relative to Canon Diablo 

troilite standard, x, arithmetic mean; s, standard deviation; n, number of analyses; nd, not determined; leaders (--), too many samples below the limit of

Greater Green River basin

Stage 1 Stage 2 Stages Stage 2A

Piceance basin

Stage 2B Stage 1 Stage 2 StageS Stage 1

Uinta basin

Stage 2 StageS

Percent sulf orto|a|

X

s

n

2.4 

1.2

5

0.71 

0.57 

22

1.4 

0.51 

7

0.96 

0.33 

3

0.67 1.9 1.7 

0.60 1.0 0.23 

19 7 25

1.0 

0.65 

7

0.68 

0.67 

8

0.78 

1.3 

21

0.76 

0.69 

10

Percent sulfur^,^

X

s 

n

1.8 

0.71 

5

0.44 

0.46 

22

0.98 

0.52 

7

0.39 

0.24 

3

0.45 1.5 0.86 

0.49 0.95 0.54 

19 4 11

0.78 

0.52 

5

0.96 

0.60

4

0.77 

1.2 

15

0.43 

. 0.32 

6

Percent sulfuradd_volatile

X

s 

n

 ~ ** ~~ 0.02 

0 

3

__ -- ~~ " "

Percent sulfursulfide

X

s 

n

1.9

0.72 

5

0.52 

0.46 

22

0.99 

0.51 

7

0.41 

0.24 

3

0.54 1.6 0.89 

0.48 0.94 0.52 

19 4 11

0.78 

0.52 

5

0.96 

0.60 

4

0.90 

1.2 

15

0.45 

0.31 

6

Percent sulfursu]f ate

X

s 

n

" ~~ 0.15 

0.11 

7

0.02 

0 

3

0.23 

0.23 

11

0.04 

0.03 

5

 " " "

Percent sulfurorgan|c

X

s 

n

0.24 

0.09 

5

0.05 

0.07 

23

0.21 

0.12 

7

0.17 

0.14 

3

0.04 0.19 0.39 

0.04 0.08 0.50 

20 4 11

0.25 

0.28 

5

0.06 

0.05 

4

0.06 

0.05 

15

0.17 

0.26 

6

Percent Ironreact|ve

X

s 

n

2.7 

0.41 

5

2.2 

1.0 

22

2.0 

0.76 

7

0.88 

0.32 

3

2.3 2.8 2.4 

1.0 1.0 1.1 

19 4 11

2.5 

0.78 

5

2.8 

1.2

4

3.2 

1.1 

15

1.9 

0.62 

6

Percent ironj^i

X

s 

n

3.2 

0.63 

5

2.6 

1.3 

21

2.3 

0.84 

7

1.9 

0.78 

3

2.7 3.3 2.8 

1.4 0.88 0.99 

18 7 25

11 

0.87 

7

3.4 

1.0 

8

3.7 

1.2 

15

2.6 

0.47 

10

Percent carbonorganlc

X

s 

n

9.4 

4.1 

5

3.8

4.9 

22

7.6 

4.1 

7

15 

1.0 

3

2.0 13 15 

1.7 8.1 7.5 

20 7 27

16 

12 

7

2.2 

3.2 

8

4.9 

5.1 

21

14 

11 

10

Percent c.rboncarboiiale

X

s

n

4.8 

2.3 

5

6.3 

2.6

23

4.6 

2.7 

7

7.3 

0.53 

3

6.1 3.6 6.2 

2.7 2.3 2.5 

20 7 27

5.7 

1.7 

7

3.5 

2.1 

8

4.8 

2.3 

21

5.4 

1.0 

10

phase (stage 2) of Lakes Uinta and Gosiute. The amount of 

sulfur contained in our stage 2 samples, on a gypsum 

equivalence (calculations in appendix 2), indicate that 29 m 

of pure gypsum could have been deposited at the

depositional center of Lake Uinta in the Piceance basin and 

14 m at the center of Lake Gosiute in the greater Green 

River basin. These calculated thicknesses of sulfate salt are 

large compared to the aggregate thickness of bedded

B6 Green River Formation, Wyoming, Utah, and Colorado
Table 2. Average chemical and isotopic parameters for samples from the three lake stages in the Green River Formation of 

the greater Green River basin, Wyoming, Piceance basin, Colorado, and Uinta basin, Utah Continued

Greater Green River basin Piceance basin Uinta basin

Stage 1 Stage 2 Stage 3 Stage 2A Stage 2B Stage 1 Stage 2 Stage 3 Stage 1 Stage 2 Stage 3

DOS,'reactive

X

s

n

0.60

0.19

5

0.27

0.23

22

0.46

0.24

7

0.33

0.19

3

0.26

0.24

19

0.49

0.20

4

0.39

0.24

11

0.35

0.33

5

0.32

0.21

4

0.25

0.25

15

0.33

0.22

6

disulfide

X

s

n

14.1

8.7

5

25.6

8.1

22

17.6

10.9

7

31.4

7.3

3

24.5

8.0

16

16.8

8.9

4

28.5

5.0

11

34.7

12.1

5

12.7

10.4

4

26.4

9.2

15

27.6

9.5

6

'sulfate

X

s

n

11.0

-

2

23.2

9.2

3

22.5

16.6

7

nd 23.2

9.2

3

nd 27.9

3.3

9

29.4 nd

-

2

nd nd

'organic

X

s

n

17.9

6.4

5

26.9

7.0

8

18.9

9.7

7

29.9

.

2

26.0

8.0

6

17.1

8.0

4

27.9

4.0

6

37.0

11.2

4

15.4

4.9

3

24.6

9.9

11

29.9

7.3

4

+80

-60

100

+80 i

+60 -

+40

DCUJ +20

Q.

CO

v 

co CO

-20 -

-40 -

-60

10 20 30 40

PERCENT REDUCTION FREQUENCY

Figure 4. Rayleigh fractionation plot and frequency histogram for 834S values of disulfides in depocenter cores of Green River 

Formation from greater Green River, Piceance, and Uinta basins. Curves were calculated assuming 834S of inflow was +20%o.

Sulfur Geochemistry and Isotopy B7
sodium carbonate minerals and halite in the cored interval 

from the Piceance basin (approximately 33 m); saline 

minerals in the greater Green River basin core are 

disseminated, not bedded.

This discrepancy between low sulfate concentrations 

(absence of sulfate salts in otherwise evaporite-type rocks) 

in the Green River lakes and high sulfur in the rocks could 

be reconciled if the sulfate in the two ancient lakes had been 

extensively reduced by sulfate-reducing bacteria. This process would have lowered water-column/porewater sulfate 

concentrations below sulfate-mineral saturation, yet 

allowed a buildup of aqueous sulfide and therefore reduced 

forms of sulfur in the sediment. Our isotope data (discussed 

later) clearly support such extensive reduction.

Sulfur isotopes are fractionated during the bacterial 

reduction of sulfate to sulfide; the accumulated reservoir of 

sulfide becomes progressively enriched in 34S as a function 

of the percent sulfate reduced. This isotopic evolution is 

quantified by the Rayleigh fractionation equations (Nakai 

and Jensen, 1964) and is plotted in figure 4. If only a small 

net portion of the sulfate in lake water is reduced, the 

accumulated sulfide is depleted in 34S relative to sulfate as 

shown on the left side of the plot in figure 4. This situation 

is similar to that in marine and some lake environments 

where H2S has formed from large, essentially infinite 

sulfate reservoirs and thus the sulfide is depleted in 34S 

relative to the initial sulfate. If the sulfate is quantitatively 

reduced (right side of Rayleigh plot in fig. 4), then the 

sulfide has the same isotopic composition as the initial 

sulfate and is enriched in 34S (chosen as +20%o, the most 

34S-enriched isotopic composition conceivable for inflow 

waters). A frequency histogram for 834S values of the 

disulfide in the Green River Formation is shown in figure 4. 

The 34S-enriched values in the histogram indicate that, 

unless there was an input of sulfate extremely enriched in 

34S (that is, 534S between +20%o and +50%o), the sulfur 

isotopic composition of the disulfide cannot be reconciled 

by simply invoking complete reduction of porewater 

sulfate.

Plots of sulfur in sulfide-mineral phases versus 

organic carbon (organic-C) are commonly used to 

distinguish between unlimited and limited sulfate reservoirs 

in marine and freshwater-lake systems (fig. 5A). A linear 

relation is observed in "normal" marine sediments, where 

the sulfate reservoir is unlimited and sulfide minerals form 

diagenetically (Sweeney, 1972). The amount of H2S that is 

trapped as sulfide minerals is dependent on the amount of 

organic carbon in the sediments available for the bacteria 

during sulfate reduction (extent of reduction is organiccarbon limited). The slope of this line decreases in 

freshwater sediments because sulfate reduction in the 

sediment is sulfate limited (Berner and Raiswell, 1984). 

Under euxinic conditions (H2S-bearing water column), a 

linear relation with a positive intercept is observed; the 

sulfate reservoir is unlimited and sulfide minerals formed in

Table 3. Dissolved water-column sulfate, sediment sulfur 

concentrations (on a carbonate-free basis), and 

d^Sdjsulfide values in samples from the Green River 

Formation and other lacustrine environments 

[NA. indicates not available]

SO42" water

concentration Total S ^Sdjsuffide 

______________(mAf) (weight percent) (per mil) 

Green River Formation

Piceance basin................NA. 1.6 +27.7

Uintabasin.....................NA. .69 +24.6

Greater Green 

River basin .................NA. 1.1 +223

Walker Lake1*2 ...................... 20 0.94 -14.7

Great Salt Lake1'3

Stage I (saline)................. 180 0.50 -21.4

Stage II (saline) .............NA. 0.44 -19.2

Stage H (fresh)...............NA. 0.90 -14.5

Soap Lake1

Noneuxinic .........................30 0.08 -29.4

Euxinic..............................360 0.20 -20.1

Lake Ontario4 ..........................0.10 0.13 +18.6

Marine environment5 ...........^_____1_____-25___

Buttle (1988).

2Benson and Spencer (1983).

3Sturm (1980).

4Calculated from Nriagu and Coker (1976).

5Calculated from data in Goldhaber and Kaplan (1974).

anoxic bottom water are added to those forming in the 

sediments during diagenesis (Leventhal, 1983). Because 

most of the mineral-sulfide and organic-carbon data for the 

Green River Formation (figs. 5B-D) show none of these 

correlations or trends, models for marine, freshwater, and 

euxinic sediments shown in figure 5A are limited in their 

ability to determine the size of the sulfate reservoir or 

explain controls on sulfate reduction and sulfide-mineral 

formation in the Green River lake system.

The similar isotopic compositions of organosulfur 

and coexisting sulfide minerals in the Green River samples 

suggest that a high proportion of the sulfate reservoir in the 

Green River lakes was reduced prior to any formation of 

sulfidization products. Once reduced, sulfur is very reactive 

and the dissolved sulfide (H2S) generally reacts readily with 

available iron to form iron sulfide minerals (Pyzik and 

Sommer, 1981). If H2S is in excess relative to reactive iron, 

organic matter in the sediment can be sulfidized (Nissenbaum and Kaplan, 1972). In marine and freshwater 

sediments, the sulfidization of organic matter is slow 

relative to the sulfidization of iron. The H2S formed at the 

start of sulfate reduction is depleted in 34S relative to the 

H2S formed during later stages of the reduction process (fig.

B8 Green River Formation, Wyoming, Utah, and Colorado
,7i4

 Lake stage 1

 Lake stage 2

D Lake stage 3

D

02 4 6 8 10 12 14 16 18 20 

Wt. % Organic C

n-»-

29

02 4 6 8 10 12 14 16 18 20 

Wt. % Organic C

B .r

 Lake stage 1 

X Lake stage 2A

 Lake stage 2B 

D Lake stage 3

D

B

 Lake stage 1

 Lake stage 2 

D Lake stage 3

D

10 12 14

Wt. % Organic C

16 18 20 6 8 10 12 14 16 

Wt. % Organic C

18 20

Figure 5. Regression lines for sulfur in sulfide minerals versus organic carbon. A, Marine sediments (Sweeney, 1972), freshwater 

sediments (Berner and Raiswell, 1984), and sediments deposited in euxinic water columns (Leventhal, 1983). Rocks of Green River 

Formation from B, greater Green River, C, Piceance, and D, Uinta basins. All data for Green River samples are weight percent 

calculated on a carbonate-free basis.

4). Therefore, sulfur in sulfide minerals is expected to be 

depleted in 34S relative to that of the more slowly forming 

organosulfur, and such is the case in marine sediments (see 

band of marine data in fig. 6). In contrast, the isotopic 

compositions of organosulfur in samples from all three 

Green River cores are quite similar to those of coexisting 

disulfide sulfur (individual points in fig. 6). Smith and 

Young (1983) argued that the similar isotopic compositions 

indicate sulfide-mineral formation from diagenetically 

released, assimilatory (inherited from living organisms) 

organosulfur. Organosulfunorganic-carbon ratios (OS/C) 

for the Green River samples are greater than in living 

analogs. Thus, the organic matter must actually have been 

enriched, not depleted, in sulfur.

As argued above, a reservoir of isotopically 

homogeneous H2S must have been available during both 

sulfidization of iron and organic matter. The existence of

such a reservoir is inconsistent with the formation of 

sulfidization products during progressive reduction of 

sulfate in the sediment column (as in marine sediments). An 

isotopically homogeneous H2S reservoir requires almost 

complete reduction of the sulfate, such as would occur in 

either stagnant, anoxic bottom waters of a stratified lake or 

anoxic pore waters within sediments prior to sulfidization of 

either iron or organic matter (controls on sulfidization 

kinetics are discussed in the following section). Almost 

complete reduction of sulfate in either of these 

environments would allow for accumulation in the 

sediments of abundant 34S-enriched sulfide minerals and 

organosulfur but not sulfate saline minerals.

The presence of H2S in an anoxic water column 

(euxinic conditions) would inhibit the aerobic destruction of 

organic matter and allow for accumulation and preservation 

of organic matter in the lake sediments. Sediments

Sulfur Geochemistry and Isotopy B9
cc

LU 

CL

z

3 

(0

o

c

CO

>

CO

20

10

§34S Disulfide, IN PER MIL

Figure 6. 834S values of organosulfur versus 834S values of disulfide sulfur. Shaded band shows marine data (Orr, 1985); 

individual points are data for Green River Formation from this study.

deposited in Lake Uinta during all stages and in Lake 

Gosiute during stages 1, 2A, and 3 contained an amount of 

well-preserved organic matter sufficient to form oil shale 

(as much as 39 weight percent organic carbon; fig. 5). These 

sediments probably were deposited under euxinic 

conditions. Although some organic-carbon-rich layers are in 

stage 2 rocks from the greater Green River basin (stage 2A 

samples), thick sequences of rocks in stage 2B contain 

relatively little or no organic matter (average 2 weight 

percent organic carbon; table 2). Low sulfur abundances 

(average 0.67 weight percent; table 2) are characteristic of 

these relatively organic-carbon-poor rocks. Organic matter 

and sulfide minerals were either not produced or not 

preserved within the depositional center of the greater 

Green River basin during stage 2B intervals. The few 

organic-carbon-rich samples of stage 2A suggest that 

euxinic conditions were intermittent during this otherwise 

shallow-lake stage. The cause of the decrease in the size of 

Lake Gosiute and the organic-matter and sulfur abundances 

in its sediment is unknown. As mentioned earlier, climatic 

changes may have been a key factor. Another possibility 

consistent with our data is a change in drainage patterns 

within the hydrologic basin of Lake Gosiute that would

have resulted in a decrease in water inflow, incoming sulfur, 

and possibly organic matter production. The absence of 

stratification and euxinic conditions during this shallowing 

period resulted in the poor preservation of deposited organic 

matter and in the formation of very lean oil shale.

The lower amounts of sulfur in Lake Gosiute 

sediment during this time caused the system to be isotopically sensitive to the influx of 34S-depleted sulfur associated 

with volcanic ash; Rye and others (1984) showed that 

volcanic ash has 834S values of from 0%o to 6%o. This 

additional supply of sulfur explains the spikes to lower 834S 

values in stage 2 samples near ash layers in the greater 

Green River basin.

Sulfidization Products

After the Green River lake sulfate was reduced to H2S 

in the water column (euxinic conditions) or sediments 

(shallow conditions), it would have diffused into and 

through the sediments along concentration gradients 

controlled by sulfidization of iron minerals and organic 

matter. Sulfidization of iron in Green River lake sediments

B10 Green River Formation, Wyoming, Utah, and Colorado
Euxinic deposition

Inflow Volcanic ash

Water column

Chemocline

FeS (po)

FeS2 (py), FeS (po) FeS2 (py) FeS2 (framboid) 

FeS2 (mrcs) + Fe2*

Organosulfur

Figure 7. Summary of sulfur geochemistry during euxinic 

deposition of Green River Formation in Lake Uinta (all stages) 

and Lake Gosiute (stages 1, 3, and 2A). py, pyrite; mrcs, 

marcasite; po, pyrrhotite. No depth scale inferred.

was probably initially hindered by high pH, such as 

observed in the sediments of modern Soap Lake, which has 

a pH of 9.8 and water chemistry similar to that proposed for 

the Green River lakes (Turtle and others, this volume), and 

in experimental laboratory studies on the sulfidization of 

goethite (Stanton and Goldhaber, this volume). Eventually, 

stepwise iron sulfidization occurred in Green River 

sediments and is recorded in the mineralogy and 

morphology of the sulfide minerals (Turtle, this volume).

Pyrite (FeS2) is the predominant sulfide mineral in the 

Green River Formation and is present as early diagenetic 

framboids and later diagenetic anhedral and euhedral grains. 

Pyrrhotite (FeS) is locally abundant, and its morphology 

indicates that it formed after compaction. Pyrrhotite is 

stabilized under very reducing conditions where elemental 

sulfur is not available to transform precursor FeS phases to 

FeS2 (Rickard, 1975). Marcasite (FeS2; mrcs) is commonly 

associated with pyrrhotite (po) in samples of Green River 

Formation and probably replaces pyrrhotite according to the 

equation (De Waal, 1984)

2FeS(po) FeS2(mrcs)+Fe2+.

Contemporaneous with formation of the earliest 

sulfide minerals (precursor FeS and framboidal pyrite), 

organic matter in the Green River sediments was also 

sulfidized. As previously mentioned, OS/C ratios indicate 

enrichment of the organic matter in sulfur, and 834S values

Shallow deposition

Inflow Volcanic Ash 

\______/_______

Water column

FeS2 (py), 

FeS2 (mrcs) + Fe2+

FeS (po) FeS2 (py) Organosulfur

Figure 8. Summary of sulfur geochemistry during shallow 

deposition of Green River Formation in Lake Gosiute (stage 

2B). py, pyrite; mrcs, marcasite; po, pyrrhotite. No depth scale 

inferred.

of the organosulfur reflect the isotopic composition of the 

sulfidizing H2S. Sulfidization of organic matter was likely 

promoted by the high pH of the pore water. Organic-matter 

sulfidization increases as pH increases, in part due to the 

stabilization of polysulfides at higher pH (Aizenshtat and 

others, 1981; Francois, 1987; LaLonde and others, 1987). 

Such early sulfidization of organic matter in high-pH 

environments is well documented in modern Soap Lake 

sediments (Turtle, 1988).

SULFUR GEOCHEMISTRY OF 

LAKES UINTA AND GOSIUTE

Collectively, the results of our study show that the 

water columns at the depositional centers of Lake Uinta 

during all stages and Lake Gosiute during stages 1,2A, and 

3 were stratified and probably euxinic. The sulfur geochemistry under these euxinic conditions is summarized in 

figure 7. Sulfate entered the lake dissolved in inflow waters 

or adsorbed onto abundant volcanic ash. The H2S was 

produced in the water column or pore water by bacterial

Sulfur Geochemistry and Isotopy B11
42

34

QC 

111 

Q_

Z

26 -

CO

18 -

10

Lake Stage

Figure 9. Mean 834S values in Green River Formation from 

the greater Green River (circles), Piceance (squares), and Uinta 

(diamonds) basins during each lake stage. Only values from 

within Mahogany zone were used to calculate stage 3 means 

for Piceance and Uinta basins.

reduction of sulfate. The H2S:sulfate ratio in these waters 

was high, and the accumulation of H2S in the reducing 

sediments inhibited aerobic destruction of organic matter. 

High porewater pH prevented rapid sulfidization of iron and 

promoted sulfidization of organic matter. The stable sulfur 

phases preserved in rocks deposited during these stages are 

pyrite, pyrrhotite, marcasite, and organosulfur.

During stage 2 in Lake Gosiute, shallow conditions 

prevailed and the water column was only intermittently 

stratified and euxinic. The sulfur geochemistry of ancient 

Lake Gosiute during periods of shallow deposition (stage 

2B) is summarized in figure 8. Again, sulfate entered the 

lake dissolved in inflow waters or adsorbed to volcanic ash. 

Its reduction took place in the sediments. The H2S slowly 

reacted with iron and, if present, organic matter. The stable 

sulfur phases in these rocks are the same as in rocks 

deposited under euxinic conditions in the lake, although 

they are generally much less abundant.

APPLICATIONS TO GEOLOGIC PROBLEMS

Data from this study not only aid in understanding the 

depositional and diagenetic conditions during the formation

of rocks of the Green River Formation but also are useful in 

developing or testing working hypotheses regarding other 

aspects of the Green River lake system and ancient lake 

systems in general.

1. Work done by Bradley (1929) and Surdam and 

Stanley (1979) suggests that Lakes Uinta and Gosiute may 

have been hydrologically connected at the onset of basin 

infilling, but the key geologic evidence has been eroded. 

Geologic studies of Surdam and Stanley (1980) and Johnson 

(1985) indicate that infilling of the basins with volcaniclastic rocks occurred from north (greater Green River 

basin) to south (Piceance basin) and then to the west (Uinta 

basin). Correlation between the degree of change in lakestage isotopic values from stage 2 to stage 3 (fig. 9) and the 

sequence of infilling supports such a connection during lake 

stage 3. Infilling caused a shift in the chemistry of Lake 

Gosiute as evidenced from the marked decrease in 834S 

values. To a lesser degree, this shift in chemistry influenced 

Lake Uinta in the Piceance basin (by damping the overall 

evolution of the sulfur isotopes to heavier values) as the 

water flowed into Lake Uinta in the northeastern corner of 

the basin. The isotope evolution in Lake Uinta in the Uinta 

basin was only slightly influenced (essentially no change in 

the evolution of the isotopes) because the infilling only 

proceeded to the eastern half of the basin.

2. The hypothesis proposed in this study regarding 

the effect of pH on sulfidization of organic matter may have 

broad implications for the origin of sulfur in oils generated 

from lacustrine source rocks. Sulfur-rich petroleum is 

generated from ancient, iron-deficient marine source rocks 

(Vairavamurthy and Mopper, 1987). The shale oil retorted 

from the Green River rocks has moderate sulfur contents, 

not because the rocks are iron deficient but because the high 

pH of the pore water in Lakes Uinta and Gosiute sediments 

inhibited iron sulfidization and promoted organic-matter 

sulfidization. The effect of porewater pH during 

sedimentation on the sulfur content of generated 

hydrocarbons has not previously been investigated but may 

be of considerable importance.

3. Our results show that working models developed 

to study sulfur geochemistry in marine and freshwater 

systems (Sweeney, 1972; Leventhal, 1983; Berner and 

Raiswell, 1984) have limited use in interpreting most of the 

sulfur geochemistry of the Green River Formation. In 

addition, paleosalinity indicators developed for these two 

sedimentary environments (such as FeS2 :FeS ratios; Berner 

and others, 1979) are not applicable to these rocks. High 

FeS2:FeS ratios are reported to occur in high-salinity 

depositional environments; the opposite relation is observed 

in our results, and, in fact, saline stage 2 samples have the 

lowest ratios.

The failure of these established models and indicators 

to interpret sulfur geochemistry in these saline lacustrine 

rocks reinforces the tenet that models developed for one 

system may not be applicable for others, even when the 

systems share similar characteristics.

B12 Green River Formation, Wyoming, Utah, and Colorado
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geochemical sulphur enrichment of recent organic matter by

polysulfides in the Solar Lake, in Bjoroy, M., ed., Advances

in organic geochemistry 1981: Chichester, England, John

Wiley and Sons Limited, p. 279-288. 

Benson, L.V., and Spencer, R.J., 1983, A hydrochemical

reconnaissance study of the Walker River basin, California

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Berner, R.A., 1970, Sedimentary pyrite formation: American

Journal of Science, v. 268, p. 1-23. 

____ 1984, Sedimentary pyrite formation An update:

Geochimica et Cosmochimica Acta, v. 48, p. 605-615. 

Berner, R.A., Balwin, T., and Holdren, G.R., 1979, Authigenic iron

sulfides as paleosalinity indicators: Journal of Sedimentary

Petrology, v. 49, p. 1345-1350. 

Berner, R.A., and Raiswell, R., 1984, C/S method for

distinguishing freshwater from marine sedimentary rocks:

Geology, v. 12, p. 365-368. 

Boyer, D.L., and Cole, R.D., 1983, Variations in sulfur

mineralization in the Parachute Creek Member of the Green

River Formation, Colorado and Utah, in Sixteenth Oil Shale

Symposium proceedings: Golden, Colorado School of

Mines, p. 160-175. 

Bradley, W.H., 1929, The varves and climate of the Green River

epoch: U.S. Geological Survey Professional Paper 158-E,

p. 87-110. 

____ 1973, Oil shale formed in desert environment; Green River

Formation, Wyoming: Geological Society of America

Bulletin, v. 84, p. 1121-1124. 

Cole, R.D., Liu, J.H., Smith, G.V., Hinckley, C.C., and Saporoschenko, M., 1978, Iron partitioning in oil shale of the

Green River Formation, Colorado A preliminary Mossbauer study: Fuel, v. 57, p. 514-520. 

Cole, R.D., and Picard, M.D., 1981, Sulfur-isotope variations in

marginal-lacustrine rocks of the Green River Formation,

Colorado and Utah: Society of Economic Paleontologists

and Mineralogists Special Publication 31, p. 261-275. 

De Waal, S.A., 1984, Experimental simulation of the supergene

reaction pyrrhotite to marcasite + Fe2+ + electrons:

Transactions of the Geological Society of South Africa, v.

87, p. 273-279. 

Dyni, J.R., 1983, Distribution and origin of sulfur in Colorado oil

shale, in Sixteenth Oil Shale Symposium proceedings:

Golden, Colorado School of Mines, p. 144-159. 

Eugster H.P., and Hardie, L.A., 1975, Sedimentation in an ancient

playa-lake complex The Wilkins Peak Member of the

Green River Formation of Wyoming: Geological Society of

America Bulletin, v. 86, p. 319-334. 

____ 1978, Saline lakes, in Lerman, Abraham, ed., Lakes 

Chemistry, geology, physics: New York, Springer-Verlag,

p. 237-294. 

Francois, R., 1987, A study of sulphur enrichment in the humic

fraction of marine sediments during early diagenesis:

Geochimica et Cosmochimica Acta, v. 51, p. 17-27. 

Goldhaber, M.B., and Kaplan, I.R., 1974, The sulfur cycle, in

Goldberg, E.D., ed., The sea, v. 5: New York, John Wiley

and Sons, p. 569-655.

Ingamells, C.O., 1970, Lithium metaborate flux in silicate 

analysis: Analytica Chimica Acta, v. 52, p. 323-334.

Ingram, L.L., Ellis, J., Crisp, P.T., and Cook, A.C., 1983, 

Comparative study of oil shales and shale oils from the 

Mahogany zone, Green River Formation (U.S.A.) and 

Kerosene Creek seam, Rundle Formation (Australia): 

Chemical Geology, v. 38, p. 185-212.

Johnson, R.C., 1985, Early Cenozoic history of the Uinta and 

Piceance Creek Basins, Utah and Colorado, with special 

reference to the development of Eocene Lake Uinta, in 

Flores, R.M., and Kaplan, S.S., eds., Cenozoic paleogeography of the west-central United States: Denver, 

Society of Economic Paleontologists and Mineralogists, 

p. 247-276.

LaLonde, R.T., Ferrara, L.M., and Hayes, M.P., 1987, Lowtemperature, polysulfide reactions of conjugated ene carbonyls A reaction model for the geologic origin of Sheterocycles: Organic Geochemistry, v. 11, p. 563-571.

Leventhal, J.S., 1983, An interpretation of carbon and sulfur 

relationships in Black Sea sediments as indicators of 

environments of deposition: Geochimica et Cosmochimica 

Acta, v. 47, p. 133-137.

Melchior, D.C., Wildeman, T.R., and Williamson, D.L., 1982, 

Mossbauer investigation of the transformations of the iron 

minerals in oil shale during retorting: Fuel, v. 61, 

p. 516-522.

Milton, C., and Eugster, H.P., 1959, Mineral assemblages of the 

Green River Formation, in Abelson, P.H., ed., Researches in 

geochemistry: New York, John Wiley, p. 118-150.

Nakai, N., and Jensen, M.L., 1964, The kinetic isotope effect in the 

bacterial reduction and oxidation of sulfur: Geochimica et 

Cosmochimica Acta, v. 28, p. 1893-1912.

Nissenbaum, A., and Kaplan, I.R., 1972, Chemical and isotopic 

evidence for the in-situ origin of marine humic substances: 

Limnology and Oceanography, v. 17, p. 570-582.

Nriagu, J.O., and Coker, R.D., 1976, Emission of sulfur from Lake 

Ontario sediments: Limnology and Oceanography, v. 21, 

p. 485-489.

Orr, W.L., 1985, Kerogen/asphaltene/sulfur relationships in sulfurrich Monterey oils: Advances in Organic Geochemistry, v. 

10, p. 499-516.

Pabst, A., 1970, Pyrite of unusual habit simulating twinning from 

the Green River Formation of Wyoming: The American 

Mineralogist, v. 56, p. 133-145.

Pyzik, A.J., and Sommer, S.E., 1981, Sedimentary iron monosulfides Kinetics and mechanism of formation: Geochimica 

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Smith, J.W., and Young, N.B., 1983, Stratigraphic variation of 

sulfur isotopes in Colorado Corehole Number 1, in 

Sixteenth Oil Shale Symposium proceedings: Golden, 

Colorado School of Mines, p. 176-188.

Sulfur Geochemistry and Isotopy B13
Smith, J.W., Young, N.B., and Lawlor, D.L., 1964, Direct 

determination of sulfur forms in Green River oil shale: 

Analytical Chemistry, v. 36, p. 618-622.

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1951, Properties of Colorado oil shale: U.S. Bureau of 

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J.W., ed., Great Salt Lake A scientific historical, and 

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Bulletin 116, p. 147-162.

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Member, Wyoming: American Association of Petroleum 

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during the culminating phase of Eocene Lake Gosiute, 

Wyoming (Green River Formation): Geological Society of 

America Bulletin, v. 90, p. 93-110.

____ 1980, Effects of changes in drainage-basin boundaries on 

sedimentation in Eocene lakes Gosiute and Uinta of 

Wyoming, Utah, and Colorado: Geology, v. 8, p. 135-139.

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sediments: Los Angeles, University of California, Ph.D. 

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history of sediment in modern and ancient saline lakes Evidence from sulfur geochemistry: Golden, Colorado School 

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analytical scheme for determining forms of sulphur in oil 

shales and associated rocks: Talanta, v. 33, p. 953-961.

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sedimentary organic matter: Nature, v. 329, p. 623-625.

B14 Green River Formation, Wyoming, Utah, and Colorado
APPENDIX 1 ANALYTICAL DATA

The analytical data for 35 samples from the Blacks 

Fork core, greater Green River basin, Wyoming; 41 samples 

from the 01A core, Piceance basin, Colorado; and 39 

samples from the Coyote Wash core, Uinta basin, Utah, are 

tabulated in table Al. Chemical data are reported on a 

dry-weight percent basis; isotopic values are reported as 

834S relative to the troilite phase in the Canon Diablo 

meteorite (CDT standard). The limit of determination on the 

sulfur species data is 0.01; < signifies that concentrations 

are below that level.

Reactive iron (Fer) and the degree of sulfidization 

(DOSr), both calculated parameters, are also given in table 

Al. Fer is operationally defined as iron that is initially 

available for sulfidization by H2S (Berner, 1984). It 

includes the measured HCl-soluble iron and iron residing in 

disulfide minerals and is calculated as

Fer=FeHC1+Fedi,

where FeHC1 is HCl-extractable iron (iron in carbonate, 

oxide, acid-volatile sulfide, and possibly clay- mineral 

phases), Fedi is iron in disulfide minerals, and Feav is the 

iron in acid-volatile sulfide minerals. Fedi and Feav are

calculated from the sulfur data, assuming stoichiometric 

FeS and FeS2. DOSr is a measure of the fraction of the 

reactive iron sulfidized and is calculated as

DOSr = [Fedi+Feav] / Fer

DOSr is similar to Berner's (1970) degree of pyritization 

(DOP). The former, however, also includes acid-volatile 

sulfides.

Abbreviations used in this appendix are as follows.

Meters Depth interval (in meters)

Di-S Disulfide sulfur

SO4-S Sulfate sulfur

R-Fe Reactive iron

Org-C Organic carbon

DOSr Degree of sulfidization

T-S Sulfur

Av-S Acid-volatile sulfur

Org-S Organosulfur

T-Fe Total iron

CO3-C Carbonate carbon

del 834S

Sulfur Geochemistry and Isotopy B15
w

Q.

9

o

Table A1. Analytical data

[BF, Blacks Fork core, greater Green River basin, Wyoming; unlettered numbers, OIA core, Piceance basin, Colorado; cw

Boxhead abbreviations explained in appendix 1, proceeding]

Sample Meters T-S% Di-S% Av-S% SO4-S% Org-S% R-Fe% T-Fe% Org-C% CO3-C% DOS

BLACKS FORK CORE

Lake stage 3

BF-1 63

BF-2 70

BF-3 91

BF-4 98

BF-5 107

BF-6 125

BF-7 135

1.8

1.2

1.5

0.56

0.79

1.2

1.1

1.4

0.65

1.1

036

0.27

0.65

0.87

0.02 0.15

< 0.05

< 0.12

< 0.03

0.01 0.16

0.01 034

< 0.04

0.02

0.26

0.15

0.20

0.26

0.14

0.12

2.9

0.94

2.0

1.9

1.2

1.2

0.93

3.0

1.1

2.4

2.2

1.5

1.4

1.2

039

6.5

4.9

83

8.8

4.1

7.0

2.1

5.6

5.1

4.4

7.2

0.24

7.7

0.44

0.62

0.49

0.18

0.17

0.49

0.84

Lake stage 2

BF-8 144

BF-9 162

BF-10 180

BF-11 189

BF-12 199

BF-13 217

BF-14 235

BF-15 244

BF-16 254

BF-17 272

BF-18 290

BF-19 299

BF-20 308

BF-21 333

BF-22 344

BF-23 353

BF-24 363

BF-25 381

BF-26 400

BF-27 409

BF-28 418

BF-29 437

0.61

0.68

0.75

0.60

0.75

1.2

0.02

0.02

0.32

0.18

0.02

0.25

0.02

0.97

0.51

032

0.02

0.97

0.26

0.13

0.86

8.5

0.17

038

0.09

0.51

0.61

1.0

0.01

0.01

0.18

0.17

0.01

0.19

0.01

0.83

039

0.22

0.01

0.77

0.21

0.10

0.56

7.4

034 0.01

< <

< <

< <

0.01 <

< <

< <

< <

0.07 0.01

< 0.01

< <

< <

< <

< <

< <

0.01 <

< 0.02

< 0.08

< 0.02

< <

< 0.08

0.01 0.06

0.03

0.18

0.02

0.04

0.03

0.06

0.02

0.01

0.01

0.01

<

0.04

0.01

0.04

0.01

0.10

0.01

0.01

<

0.03

0.12

<

13

0.67

0.43

0.90

13

1.5

0.56

039

3.2

1.7

1.7

0.87

3.2

1.7

0.77

1.2

1.4

2.5

1.2

1.5

2.5

8.4

1.4

0.87

0.89

1.1

1.2

1.5

0.71

0.53

3.7

1.9

1.7

1.1

3.8

2.0

1.0

1.5

1.5

2.8

1.4

2.0

23

9.0

1.5

8.2

9.7

2.2

1.5

2.8

0.99

0.70

0.19

0.15

0.14

4.1

0.14

33

0.58

8.3

1.2

0.12

0.49

2.9

1.2

1.2

6.2

7.9

6.9

7.2

73

7.2

9.9

11

1.2

5.4

5.7

7.7

1.1

7.6

9.1

7.1

8.1

23

7.6

4.7

6.0

2.7

0.57

0.55

0.25

0.52

0.41

0.60

0.02

0.04

0.09

0.09

0.01

0.17

0.01

0.43

0.47

0.20

0.01

0.27

0.07

0.06

0.20

0.77
Table A1 . Analytical data Continued

2-

n

Sample Meters T-S% Di-S% Av-S% SO4-S%Org-S% R-Fe% T-Fe% Org-C% CO3-C% DOSr 

Lake stage 2 Continued

BF-30 454 0.97 0.01 0.80 0.02 0.02 3.9 4.4 0.47 4.5 0.35 

Lake stage 1

BF-31

BF-32

BF-33

BF-34

BF-35

464

473

491

500

509

1.1

1.7

2.2

1.5

2.8

0.87

1.4

1.9

1.1

1.7

0.02

<

<

<

<

<

0.01

0.06

0.01

0.07

0.16

0.13

0.30

0.23

0.15

2.7 2.6

2.2 2.9

2.3 2.9

1.3 1.8

1.9 2.3

12

7.3

5.8

11

2.5

23

3.2

3.9

7.0

7.4

0.30

0.57

0.75

0.61

0.77

01ACORE

Lake stage 3

6

8

12

14

18

24

26

268

272

285

290

304

321

328

0.71

0.25

0.38

2.0

0.83

0.55

0.59

0.54

0.18

0.33

1.4

0.52

0.01

0.01

0.01

0.01

0.01

0.04

0.03

0.01

0.08

<

0.12

0.03

0.03

0.64

0.08

1.6 1.7

2.7 2.8

2.4 2.5

1.4 1.8

1.5

1.2

13 1.5

9.6

13

5.4

34

13

10

7.5

6.0

43

5.6

2.6

7.2

6.4

7.5

0.30

0.06

0.13

0.89

0.35

Lake stage 2

30

36

42

48

54

59

62

65

72

74

78

84

86

89

96

98

340

358

378

395

415

429

437

448

468

474

486

504

510

521

541

547

1.5

13

0.61

2.5

5.0

1.1

0.24

0.9

1.3

0.93

1.4

0.78

2.4

0.42

1.2

1.6

0.32

0.85

0.69

1.3

0.24

0.97

0.42

0.20

0.01

0.01

0.01

0.01

0.01

0.03

0.03

0.04

0.13

0.08

0.07

0.06

0.45

0.06

0.22

0.07

0.01

0.10

0.12

0.28

2.9

1.8

3.1 3.0

2.3

4.1

1.6 1.7

0.85

1.5 1.8

2.3

2.1

2.4 2.4

1.0

3.4

0.9 0.96

2.4 2.4

1.8 2.3

14

23

10

7.0

5.6

12

18

10

13

11

11

7.4

14

1.2

6.8

12

53

53

3.8

4.5

2.2

5.4

7.4

63

5.3

5.2

4.9

4.7

4.7

7.4

6.7

4.5

0.20

0.47

0.41

0.49

0.23

0.37

0.23
Table A1. Analytical data Continued

39

8.

n

^

o

i

o

Sample Meters T-S% Di-S% Av-S% SO4-S% Org-S% R-Fe% T-Fe% Org-C% C03-C% DOSr 

Lake stage 2 Continued

102

108

110

113

120

123

126

132

134

138

144

559

578

583

595

614

625

632

651

657

669

687

1.4

0.94

038

1.8

1.2

1.4

0.08

0.47

1.2

0.52

11

0.93

0.06

032

4.7

0.08 0.42 037

0.01 0.01 <

0.02 039 0.47

0.59 4.6 13

23

0.1

1.7

8.8

1.9

1.4

0.41

2.5

1.2

3.5

0.10

1.0

2.0

13

93

15

17

0.94

14

11

8.9

0.55

8.5

13

17

6.2

5.8

5.2

13

5.6

8.9

3.8

14

8.6

4.9

73

5.7

0.41

0.79

0.19

0.59

Lake stage 1

150

154

155

162

165

168

172

705

717

728

743

755

760

773

1.1

1.5

1.7

3.7

1.9

0.83

0.73

1.1

0.85

2.5

0.59

0.01 0.06 0.18

0.02 0.01 0.15

0.01 0.01 0.23

0.01 < 0.71

1.5

1.9

3.5

2.1

1.5

2.4

3.4

4.2

3.4

2.0

23

22

12

12

14

8.6

3.2

3.7

5.0

1.5

13

1.8

2.9

6.2

6.5

0.68

0.42

0.62

0.25

COYOTE WASH CORE

Lake stage 3

cw-38

cw-37

cw-36

cw-35

cw-34

cw-33

cw-32

cw-31

cw-30

cw-29

571

587

597

606

624

642

652

661

670

681

0.95

0.46

0.08

0.07

0.25

0.63

0.79

036

030

1.8

0.66

0.01

0.16

0.70

030

0.76

< < 0.03

< < 0.05

0.05 < 0.08

< < 0.08

0.03 < 0.08

< 0.23 Q.69

2.4

1.6

1.6

1.5

0.90

23

2.6

23

1.8

1.5

1.8

1.5

1.7

2.0

1.6

2.0

3.4

3.5

6.6

6.2

11

7.2

8.2

5.6

26

27

4.7

5.2

73

63

5.2

3.8

63

6.0

4.7

4.9

0.24

0.01

0.14

0.40

034

0.28

Lake stage 2

cw-28 697 0.04 < < < 0.04 1.6 1.6 2.5 8.4 0.00 
Table A1 . Analytical data Continued

I

CO

 L

vO

Sample Meters T-S% Di-S% Av-S% SO4-S% Org-S% R-Fe% T-Fe% Org-C% CO3-C% DOS

Lake stage 2 Continued

cw-27

cw-26

cw-25

cw-24

cw-23

cw-22

cw-21

cw-20

cw-19

cw-18

cw-17

cw-16

cw-15

cw-14

cw-13

cw-12

cw-11

cw-10

cw-9

cw-8

706

716

725

734

752

762

111

789

807

815

825

843

862

871

880

898

907

917

926

935

0.10

0.02

0.50

0.17

1.0

0.02

0.25

1.8

0.89

1.1

0.13

0.03

0.40

0.12

4.0

0.04

0.04

1.0

0.13

0.52

0.01

0.39

0.31

0.93

0.17

0.64

0.72

0.29

0.01

0.38

3.3

0.03

0.84

0.46

< < 0.03

< 0.06 0.03

< < 0.04

< < 0.08

0.05 < 0.04

1.0 < 0.10

< 0.04 0.02

0.67 0.02 0.03

< < 0.09

< < 0.02

< < 0.12

< < 0.02

< < 0.01

< < 0.02

3.5

1.8

1.4

2.0

2.8

3.6

1.4

3.0

1.0

3.6

3.4

1.7

3.1

3.3

0.99

3.5

2.4

1.3

2.1

1.0

3.7

4.1

1.6

3.3

1.3

2.0

3.9

1.8

3.7

2.6

2.3

3.7

3.4

4.0

4.9

2.4

0.90

3.2

4.9

2.4

2.8

7.6

0.70

1.2

8.6

19

1.1

4.8

5.8

1.3

1.8

0.60

0.50

0.40

2.9

4.4

4.2

8.0

6.4

8.6

2.7

2.2

6.2

4.3

5.3

3.1

4.3

6.9

6.0

7.5

3.6

4.1

1.3

1.2

0.00

0.19

0.20

0.42

0.09

0.63

0.43

0.47

0.01

0.09

0.83

0.02

0.25

0.12

Lake stage 1

cw-7

cw-6

cw-5a

cw-5

cw-4

cw-3

cw-2

cw-1

953

971

981

990

1008

1026

1035

1043

0.06

1.5

0.47

0.37

0.11

0.80

0.04

0.86

1.3

0.16

0.59

0.78

< < 0.01

< 0.07 0.02

< 0.02 0.08

< < 0.05

2.9

2.2

0.90

2.6

2.7

3.7

3.8

3.0

2.1

1.2

2.9

2.9

0.90

0.60

0.50

0.60

0.30

6.2

0.30

3.3

2.6

3.4

1.9

2.1

2.1

6.1

2.4

7.5

0.40

0.07

0.56

0.25
APPENDIX 2 CALCULATION OF 2) was converted to moles of sulfur per gram rock. A 

GYPSUM THICKNESS density of 2.3 for Green River oil shale (Smith, 1969) was

used to convert the amount of grams to a volume. The moles

The equivalent bedded gypsum thickness in the stage equivalent gypsum and the density of gypsum were used to 

2 (saline) interval of the 01A and Blacks Fork cores was convert the volume into a thickness (meters) of bedded 

calculated as follows. The average amount of sulfur (table gypsum equivalence.

B20 Green River Formation, Wyoming, Utah, and Colorado
Chapter C

A Preliminary Study of the Carbon and 

Nitrogen Isotopic Biogeochemistry of 

Lacustrine Sedimentary Rocks from the 

Green River Formation, Wyoming, Utah, 

and Colorado

By JAMES W. COLLISTER and J.M. HAYES

U.S. GEOLOGICAL SURVEY BULLETIN 1973

CEOCHEMICAL, BIOGEOCHEMICAL, AND SEDIMENTOLOGICAL STUDIES OF THE 

GREEN RIVER FORMATION, WYOMING, UTAH, AND COLORADO
CONTENTS

Abstract Cl 

Introduction Cl 

Methods C2

Class fractionation of total extract C2 

Inorganic nitrogen analysis C3 

Combustion procedure and cryogenic distillation C3 

Results C3 

Discussion C5

Depositional models for the Green River Formation C6 

Isotopically enriched organic nitrogen C8 

Models of nitrogen cycling in the Green River lakes C9 

Saline (playa-lake) stages CIO 

Fresh (meromictic lake) stages Cll

Comparison of Piceance, Uinta, and greater Green River basins C12 

Isotopically depleted total organic carbon C12 

Sources of isotopically depleted organic carbon C13 

Anaerobic photosynthetic bacteria C13 

Methanogens and methylotrophs C13 

Summary C14 

References cited C14

FIGURES

1-3. Depth profiles of total organic carbon content and 813C and 815N compositions 

of kerogen in Green River Formation for:

1. 01A core, Piceance basin, Colorado C6

2. Blacks Fork core, greater Green River basin, Wyoming C7

3. Coyote Wash no. 1 core, Uinta basin, Utah C7 

4-6. Diagrams showing:

4. Distribution of 15N in freshwater environments C9

5. Proposed nitrogen cycle for saline stages of Green River lakes CIO

6. Microbial nitrogen cycle for meromictic lakes Cll

TABLES

1. Observed fractionation factors for the major pathways in the nitrogen

cycle C2 

2-4. Data bases for the:

2. 01A core, Piceance basin, Colorado C4

3. Blacks Fork core no. 1, greater Green River basin, Wyoming C5

4. Coyote Wash no. 1 core, Uinta basin, Utah C5

5. Isotopic composition of organic and inorganic nitrogen of evaporites of the 

Green River Formation from the 01A core, Piceance basin, Colorado C8

Contents
A Preliminary Study of the Carbon and Nitrogen 

Isotopic Biogeochemistry of Lacustrine Sedimentary 

Rocks from the Green River Formation, Wyoming, 

Utah, and Colorado

By James W. Collister anc/J.M. Hayes 1

Abstract

The Paleogene Green River Formation in the tri-State 

region of Wyoming, Utah, and Colorado contains the world's 

largest known reserve of oil shale. The unusually high organic 

content of these lacustrine rocks is a striking example of the 

productivity and preservation potential of hydrocarbons 

deposited in modern saline lacustrine environments. In order 

to reconstruct the biogeochemistry of this lacustrine setting, 

carbon and nitrogen isotopic compositions of Green River 

organic material were studied. The results of this study indicate 

that a complex nitrogen cycle existed in this lake system. 

Isotopic compositions of organic nitrogen in the Green River 

sedimentary rocks are highly enriched and significantly 

variable (+10.8 to +20.7%o). Because isotopic fractionation 

accompanying nitrogen fixation is negligible, organisms 

utilizing this pathway to meet their nitrogen requirements 

should reflect the isotopic composition of atmospheric N2 

(0%o). The existence of 15 N-enriched organic matter in 

these rocks requires a complex nitrogen cycle in the lake 

system and indicates that nitrogen was not a limiting nutrient 

for primary growth. Cyanobacterial assimilation of ammonia 

and nitrate enriched in 15 N from the microbial processes of 

denitrification and nitrification probably accounts for the 

isotopic enrichment.

The carbon isotopic composition of Green River kerogens is from -41.4 to -28.8%o. Geoporphyrins can serve as 

isotopic biomarkers for primary photosynthate and can 

therefore provide an estimated range of 13C compositions of 

primary producers in the Green River lakes. Nickel geopor-

'Biogeochemical Laboratories, Departments of Chemistry and of 

Geology, Geology Building, Indiana University, Bloomington, Indiana, 

47405-5101.

phyrins from two samples have 813C values of -23.6 and 

-28.6°/oo. A large isotopic depletion (-17%o) in at least one 

kerogen relative to the associated geoporphyrins indicates 

large contributions of isotopically depleted organic material to 

the sediments. Possible sources of isotopically depleted carbon 

include photosynthetic sulfur bacteria at the chemocline of a 

stratified lake and methane cycling by methylotrophs.

INTRODUCTION

Organic matter comprises an unusually large volume 

fraction (1-40 percent; Robinson, 1976) of the total rock in 

the lacustrine Paleogene (Eocene, locally Paleocene) Green 

River Formation of Wyoming, Utah, and Colorado. Modem 

hypersaline environments are some of the most productive 

on earth (Warren, 1986), and the accumulation of organic 

carbon in the Green River Formation suggests that the saline 

Green River lake system supported a highly productive 

community and provided an environment that favored 

preservation of organic material. Production of biomass 

requires phosphorus, sulfur, and nitrogen, as well as the 

carbon, hydrogen, and oxygen that are plentiful in lacustrine 

environments. Although concentrations of phosphorus 

commonly are biolimiting, particularly in unpolluted lacustrine environments, the presence of large quantities of 

organic material in the Green River rocks indicates that 

fluvial inputs of phosphate must have been substantial. (We 

can speculate that the remarkable organic richness of the 

Green River sediments is due to fertilization of the lake 

system by erosion of the Permian Phosphoria Formation.) 

Given adequate supplies of phosphorus, lacustrine 

communities can obtain nitrogen by fixation of N2. In order

Carbon and Nitrogen Isotopic Biogeochemistry C1
Table 1 . Observed f ractionation factors for the major pathways in the nitrogen cycle

Isotope effect

Observed fractionation factors 

OC(react/prod) Reference

Kinetic isotope effects

Nitrogen fixation

Atmospheric N2 »Fixed nitrogen

Nitrification: NH4+-»NC>2 

Denitrification: NO3

Inorganic nitrogen assimilation

Nitrate assimilation: NO3+-»Cell-N

Ammonium assimilation: NH4+ >Cell-N 

Decomposition: Sediment organic-N-»NH4+

1.000

1.004

0.996-1.009

1.0022

1.0024

1.02

1.02

1.04

1.011-1.013

1.013-1.023

1.013

0.993-1.010

1.001

Hoering and Ford (1960). 

Delwiche and Stein (1970). 

Wada and others (1975). 

Takigiku (1987). 

Macko and others (1987).

Miyake and Wada (1971).

Delwiche and Stein (1970), Miyake and 

Wada (1971), Wellman and others (1968). 

Cline and Kaplan (1975).

Macko and others (1987). 

Wada and Hattori (1978). 

Macko and others (1987). 

Wada and others (1975). 

Miyake and Wada (1975).

Equilibrium Isotope effects

Ammonia volatilization: NH4+(aq)-> NH3(gas), 25 *C 1.034 

Solution of nitrogen gas: N2(Solution)->N2(gas). ° ° c 1.00085

Kirschenbaum and others (1947). 

Klots and Benson (1963).

to study the production of organic material in the Green 

River lake system, we chose to focus on the nutrition of the 

producer community and, in particular, on the cycling of 

nitrogen within these lakes. Nitrogen isotopic analyses 

provide a new view of this phenomenon.

Unlike phosphorus, nitrogen has an active redox 

cycle and passes through many mobile forms. The biogeochemical nitrogen cycle is comprised of processes in 

which nitrogen is present in a combination of oxidation 

states (-3 to +5). A restricted group of microorganisms 

mediates most of the reactions in this cycle. Organic 

material is decomposed to produce ammonia and CO2 in a 

process termed mineralization. Ammonia produced by 

mineralization has several possible fates. In anaerobic 

environments, ammonium is directly assimilated by plants 

and bacteria. In aerobic environments, ammonium is 

oxidized to nitrate by nitrifying bacteria. Evaporation of 

surface waters and concomitant volatilization of ammonia 

removes ammonia from the lacustrine system. Nitrate is 

either assimilated directly by plants, or, in anaerobic 

environments, is reduced to N2 by the process of denitrification, in which denitrifying bacteria utilize NO^ as a 

terminal electron acceptor. The loss of N2 from sediments to 

the atmosphere by denitrification is balanced by nitrogenfixing bacteria that convert N2 to organic nitrogen (herein

referred to as organic-N). The observed isotopic 

fractionation factors for the major pathways in the nitrogen 

cycle are summarized in table 1.

METHODS

For isotopic analysis of total organic carbon, 5-10 mg 

of powdered, decalcified sample was combusted in a sealed 

quartz combustion tube at 850 °C (Knoll and others, 1986). 

The percentage of total organic carbon in the sample was 

determined from the yield of CO2. Kerogen was isolated 

using the method described by Wedeking and others (1983), 

and carbon, hydrogen, and nitrogen elemental contents were 

obtained by use of a Carlo Erba model 1106 elemental 

analyzer. Total extractable organic matter was obtained by 

Soxhlet extraction with 50/50 (v/v) dichloromethanemethanol for about 72 hours. Excess solvent was removed 

by rotary evaporation at 20-30 °C.

Class Fractionation of Total Extract

Subtractions of the total extracts enriched in geoporphyrins were prepared following the method outlined by 

Takigiku (1987). A modification of this method was

C2 Green River Formation, Wyoming, Utah, and Colorado
employed to isolate large quantities of geoporphyrins from 

gilsonite. A 2.5 cm by 40 cm column was packed with 

activated silica gel. About 200 mg of gilsonite (a pyrobitumen from the Uinta basin present as dike and fracture 

fillings; Hunt and others, 1954) was taken up in a minimal 

amount of hexane and transferred to the top of the column. 

The saturate fraction was eluted using 250 mL of n-hexane, 

followed by 250 mL of a 2 percent solution of dichloromethane in petroleum ether to elute saturated 

hydrocarbons and some aromatic compounds. The fraction 

enriched in geoporphyrins was then eluted using 250 mL of 

a 70/50/5 (v/v/v) mixture of hexane, toluene, and ethyl 

acetate. This subfraction was further purified by thin-layer 

chromotography (TLC) as outlined by Takigiku (1987). 

Purities of the porphyrin subtractions were checked using 

ultraviolet-visible spectrophotometry. Identification of 

nickel geoporphyrins was based on analysis of visible 

spectra. Final purification of the porphyrin fractions was 

carried out using high-pressure liquid chromotography 

(HPLC) (Altech 605RPD CIS reverse-phase column, 4.6 

mm by 250 mm, 5-(J,m particle size, eluted isocratically 

using 80/20 v/v acetonitrile/methanol at 1 mL/min).

Inorganic Nitrogen Analysis

Three varieties of evaporites from the Green River 

Formation were examined: bedded nahcolite, nodular 

nahcolite, and bedded halite. Weighed samples of powdered 

evaporite mineral were dissolved in deionized water and 

filtered through a 0.2-2 (im cellulose filter to remove 

particulate material. The filtrate was analyzed using ion 

exchange chromatography to detect the presence of nitrate 

and nitrite.

Separation of the organic fractions from the evaporite 

mineral phase was carried out using a separatory funnel 

extraction. Several grams of powdered evaporite mineral 

was placed in a 500-mL separatory funnel together with 100 

mL of dichloromethane and 250 mL of deionized water. 

Upon dissolution of the mineral phase, the water layer was 

recovered and filtered. The remaining material was then 

washed twice with approximately 200 mL of deionized 

water. The water layer from each washing was filtered and 

added to the original aqueous phase. Purified saline 

minerals were recovered by evaporation of the water. Total 

extractable organic material was recovered by draining and 

filtering the dichloromethane layer through a 2-(im Teflon 

filter. Excess solvent was removed by rotary evaporation at 

20-30 °C. The remaining insoluble material was transferred 

to a Nalgene bottle and centrifuged for 30 minutes. Kerogen 

was isolated from insoluble material using the procedure of 

Wedeking and others (1983).

Combustion Procedure and 

Cryogenic Distillation

Throughout this study, a modified Dumas combustion 

(Macko, 1981) was utilized for preparation of samples of N2 

from kerogens and extracts. An organic sample of 

approximately 15-25 mg was placed in a Vycor combustion 

tube (9 mm outside diameter by 150 mm long). 

Approximately 2 g of CuO wire and 1 g copper shavings 

were placed in the tube together with a 0.5-cm square of 

silver foil. The combustion tubes were evacuated, sealed, 

and placed in a muffle furnace. Combustion of the samples 

was carried out at 850 °C for at least 12 hours, followed by 

at least 2 hours at 550 °C to allow for complete reduction of 

nitrogen oxides.

After combustion, an O-ring compression fitting was 

used to adopt tubes to the cracker described by Baker 

(1987). Operation of the cracker released gaseous contents 

of each tube to a vacuum system. Nitrogen gas was 

collected on 2-3 pellets (5 mm by 3 mm diameter) of 

5-angstrom molecular sieve (-196 °C, 10 minutes; Kobayakawa and others, 1962) placed in a Pyrex collection tube (6 

mm by 120 mm). Prior to collection of N2, the sieves were 

degassed on line by heating at 200 °C for approximately 15 

minutes. Because of the large amounts of CO2 and H2O 

produced during combustion of large samples required for 

nitrogen analysis, the end of the combustion tube was 

immersed in liquid nitrogen for several minutes prior to 

admitting the sample into the vacuum line. Samples of CO2 

were subsequently cryogenically distilled and collected for 

mass spectrometric analysis.

Samples of N2 were released from the molecular 

sieve by heating at 90 °C for 2 minutes. The sample was 

then introduced into the mass spectrometer, and the 

molecular sieves were kept at 90 °C to ensure that none of 

the nitrogen remained on the sieves.

All isotopic analyses were carried out on a Finnigan 

Delta-E mass spectrometer. Overall analytical uncertainties 

are +0.2%o for 813C and +0.6%o for 815N (95 percent 

confidence limits).

RESULTS

Results of the geochemical analyses are shown in 

tables 2-4 (sample descriptions are given in Tuttle, this 

volume). Two distinct features are readily apparent: (1) the 

organic matter is extremely enriched in 15N compared to 

other reported values, and (2) the same material is depleted 

in 13C compared to most modern lacustrine organic 

material. Organic material from lake sediments shows a 

broad range of isotopic compositions (-8 to -38%o), but 

organic matter from most lacustrine sediments has 813C 

values in the range of-18 to -28%o ( Stuiver, 1964, 1975; 

Deines, 1980). Figures 1-3 show total organic carbon

Carbon and Nitrogen Isotopic Biogeochemistry C3
Table 2. Data base for 01A core, Piceance basin, Colorado

[Location of core hole shown in Turtle (this volume, fig. 1). TOC is total organic carbon; KER is kerogen; DOL is dolomite. Leaders (--)

indicate not determined]

Sample

01-6

01-8

01-12

01-14

01-24

01-26

01-36

01-59

01-62

01-65

01-72

01-78

01-84

NaCl

01-89

01-96

01-98

01-102

01-113

01-126

NaHCO3

01-132

01-134

b NaHC03

01-144

01-150

01-154

01-165

01-172

Depth 

(meters)

268

272

285

290

321

328

358

429

437

448

468

486

504

512

521

541

547

559

595

632

646

651

657

680

687

705

717

755

773

Weight percent 

TOC

9.60

1.30

5.40

34.00

10.00

7.50

23.00

12.00

18.00

10.00

13.00

11.00

7.40

1.20

6.80

12.00

15.00

14.00

0.55

8.50

13.00

6.20

22.00

12.00

12.00

14.00

813CTOC

-29.9

-28.7

-30.9

-29.8

-31.4

-29.2

-33.5

-31.4

-30.6

-30.9

-30.1

-31.7

-31.1

-31.0

-32.3

-32.0

-32.9

-30.0

-30.9

-29.9

813CKER

-31.2

-41.4

-32.0

-31.4

-30.4

-32.1

-29.4

-33.7

-32.3

-33.0

-31.1

-31.6

-31.4

-36.5

-31.5

-32.5

-29.9

-31.7

-31.5

-34.9

-33.6

-32.7

-31.0

-35.5

-31.8

-33.4

-30.5

-31.3

-30.3

S15NKER

20.7

15.6

19.1

18.4

19.4

17.0

14.7

15.2

15.8

16.3

16.4

16.5

16.6

11.5

17.3

16.1

17.8

17.6

16.4

10.8

12.9

13.7

16.8

12.0

18.0

13.4

14.3

16.0

14.5

C:N 

ratio

33

91

45

45

38

31

48

37

29

28

36

26

42

~

31

31

34

34

45

26

36

43

38

36

38

53

43

32

H:C 

ratio

1.67

1.02

1.57

1.71

1.56

1.43

1.61

1.66

1.61

1.54

1.56

1.60

1.60

1.60

1.49

1.62

1.70

1.68

1.04

1.50

1.66

1.64

1.94

1.56

1.44

1.65

1.55

S^CDOL1

9.1

6.0

8.6

6.3

7.5

6.0

10.4

-5.7

1.5

3.3

5.7

7.2

4.1

«

2.9

4.5

6.8

6.4

7.2

6.5

--

6.1

7.8

4.5

7.6

3.1

5.7

11.4

From M.L. Ttittle (U.S. Geological Survey, written commun., 1987).

(TOC, in percent) content and carbon and nitrogen isotopic 

compositions of samples of kerogen from the three cores 

plotted versus depth.

Kerogen isolates from the Green River Formation 

show a broad range of nitrogen isotopic compositions; 815N 

values are from +10.8 to +20.7%o. The S15N values for the 

associated total extracts are depleted in 15N by 0.7-4.8%o 

relative to the kerogens. Nitrogen isotopic compositions of 

organic material, total inorganic nitrogen (inorganic-N), and 

nitrate from Green River Formation evaporites are 

compared in table 5. S15N values for total inorganic-N in 

these samples are from +8.0 to +9.5%o. These values are 

depleted in 15N relative to kerogen and total extractable 

organic matter isolated from the same samples (table 5).

Another striking observation is the depletion of 13C in 

the organic matter in the Green River Formation; 513C 

values for kerogens are from -41.4 to -28.8%o. Total

extracts are enriched in 13C relative to associated kerogens 

by l-4%o (tables 2-4), and dolomite in the core samples is 

enriched in 13C. 813C values for dolomite in the Blacks Fork 

and 01-A cores are shown in tables 2 and 3 (core locations 

shown in Tuttle, this volume, fig. 1). These values, from 

-5.7 to +11.4%o versus PDB, are enriched in 13C as 

compared to many lacustrine carbonate rocks. Stuiver 

(1970) measured S13C values for carbonate minerals from 

four lakes in North America of from -7.5 to +4%o. Other 

studies of lacustrine carbonate minerals indicate 813C values 

from -7.7 to +6.3%o (Broecker and Olson, 1959; Oana and 

Deevey, 1960; Stuiver, 1975; Lee and others, 1987). The 

most 13C enriched value was for tufa along the lakeshore of 

Pyramid Lake (Broecker and Olson, 1959). Stuiver (1964) 

measured S13C values for carbonate minerals from Searles 

Lake (a saline lake in California) of from +1 to +4.6%o and 

one value of +10.1%o.

C4 Green River Formation, Wyoming, Utah, and Colorado
Table 3. Data base for Blacks Fork core no. 1, Green River basin, Wyoming

[Location of core hole shown in Tuttle (this volume, fig. 1). TOC is total organic carbon; KER is kerogen; DOL is dolomite. Leaders (-)

indicate not determined]

Sample

BF-4

BF-6

BF-7

BF-8

BF-13

BF-19

BF-22

BF-23

BF-31

BF-34

Depth 

(meters)

98

125

134

144

217

299

344

353

464

500

Weight percent 

TOC

830

4.10

7.00

1.50

2.80

4.10

0.58

8.30

12.00

11.00

513C TOC

-32.4

-30.4

-29.8

-293

-31.1

-31.2

-25.7

-32.1

-33.9

-31.8

513C KER

-32.9

-31.1

-30.0

-30.8

-32.2

-323

-32.8

-34.5

-32.2

515N KER

14.9

10.1

12.1

16.7

18.0

17.7

11.4

183

18.9

13.3

C:N 

ratio

50

30

-

H:C 

ratio

1.88

1.68

-

-

-

513C DOL1

10.1

-1.5

5.7

4.0

2.5

2.1

4.3

3.1

4.0

9.1

1From M.L. Tuttle (U.S. Geological Survey, written commun., 1987)

Table 4. Data base for Coyote Wash no. 1 core, Uinta basin, Utah

[Location of core hole shown in Tuttle (this volume, fig. 1). TOC is total organic carbon; KER is kerogen; DOL is dolomite.

Leaders (-) indicate not determined]

Sample

CW-38

CW-31

CW-29

CW-28

CW-26

CW-20

CW-18

CW-13

CW-11

CW-1

Depth 

(meters)

571

661

681

697

716

789

815

880

908

1,043

Weight percent 

TOC

3.40

5.60

27.00

2.50

2.40

7.60

1.20

5.80

1.80

330

513C TOC

-29.6

-29.8

-28.5

-30.8

-283

-303

-30.8

-30.8

-30.5

513C KER

-29.5

-29.5

-28.8

-32.4

-32.1

-28.8

-31.1

-31.0

-31.8

-31.4

515N KER

18.6

18.2

17.9

20.5

16.2

16.9

10.4

16.6

15.0

143

C:N 

ratio

26

31

34

32

34

24

29

36

33

H:C 

ratio

1.47

1.48

1.72

1.62

1.74

1.83

1.71

1.77

1.98

Nickel geoporphyrins were isolated from the extracts 

of two samples from the 01A core (Colorado). Geoporphyrins from samples 01-8 and 01-36 have 513C values of 

-23.6%o, and -28.6%o, respectively. Nickel geoporphyrins 

isolated from gilsonite from the Uinta basin have a 513C 

value of -28.5%o.

DISCUSSION

The variability of the nitrogen isotopic compositions 

of the organic matter in this sequence of rocks must carry 

information regarding (1) the nutrition of the organisms in 

these lakes and (2) the cycling of nitrogenous compounds in 

this ecosystem; however, more significant than the 

variability is the enrichment in 15N (see Letolle, 1980). Such 

extreme enrichment in 15N for total organic-N is known 

from only a few ancient and modern sediments (fig. 4). In a 

study of nitrogen isotopes in the Scheldt estuary, Mariotti 

and others (1984) measured 515N values for NH4+ and

suspended organic matter in the water column. These 

authors observed extreme 15N enrichments, as well as 

dramatic seasonal variability, in the 515N values for both 

fractions. 515N values of NH4+ were from +10 to greater 

than >+30%o, and those for suspended organic matter were 

from +2 to +18%o. Mariotti and others (1984) speculated 

that algae were assimilating a pool of 15N-enriched 

ammonium produced by nitrification. Cifuentes and others 

(1988) measured 515N values of paniculate organic matter 

(POM) in the Delaware estuary and found large seasonal 

effects on the nitrogen isotopic composition of this material. 

In the winter, 515N values for POM were between +5.5 and 

+12.2%o, and a maximum 515N value of +18.7%o was 

observed in April. These authors suggested that algae were 

assimilating a pool of ammonium enriched in 15N (>30%o) 

generated by fractionations accompanying previous 

assimilation. Kerogen isolated from the Eocene Messel 

shale (West Germany) was chosen for comparison of a 

stratified lake that existed during the same epoch as the

Carbon and Nitrogen Isotopic Biogeochemistry C5
LU 

Cfl

LU

LU

LU

Ofi

u

LU 

H

^

X

u

Qi

<l

Q.

GARDEN GULCH MEMBER

iln

04

LU

o

H

V3

LU

<

_J

LAKE STAGE 

1

-200r-

-300 -

-400 -

-500 -

-600 -

-700 -

-800

-200

-300

-400

-500

-600

-700

0 10 20 30 40

TOTAL ORGANIC 

CARBON, IN PERCENT

J -800

-200 r

-300 -

-400

-500

-600

-700

-45-40 -35 -30 -25 

8 13C,INPERMIL

-800

10 14 18 22 

8 15 N,INPERMIL

Figure 1. Depth (in meters) profiles of total organic carbon content and 813C and 815 N 

compositions of kerogen in Green River Formation for 01A core, Piceance basin, Colorado. Open 

symbols represent samples whose mineralogy indicates extreme evaporitic conditions. Location of 

core shown in Tuttle (this volume, fig. 1).

Green River lakes. This material has a 815N value of +6.8%o 

(Collister, 1989).

If, as is likely, isotopic compositions of inorganic-N 

sources were roughly constant, the observation of wide 

variations in the abundance of 15N must indicate either 

significant variations in the biological cycling of nitrogen. 

Further, the presence of 15N-enriched organic matter 

indicates that nitrogen inputs either were, on average, highly 

enriched in 15N or were extensively reworked and 

fractionated. As shown by the fractionation factors 

summarized in table 1, processes capable of causing 

significant isotopic shifts involve assimilation, oxidation, 

reduction, and volatilization of nitrogen. Of course, these 

processes occur to one degree or another in almost all 

environments, yet they do not cause 15N-enrichments 

approaching those recorded by Green River rocks. Processes occurring in the Green River lake system must have 

been unusual in that they allowed the escape of considerable 

quantities of 15N-depleted material. Moreover, the escape of 

15N-depleted material occurred while large quantities of 

organic-N were being sequestered in accumulating 

sediments. The significant isotopic fractionation is, 

therefore, a signal that nitrogen was not the limiting nutrient 

in the Green River lake system. Because the 01A core is the

most densely sampled of the three cores analyzed, most of 

our interpretation of the biogeochemistry of the Green River 

Formation is based on the geochemical data for this core.

Depositional Models for the 

Green River Formation

Two models of deposition have been proposed for the 

lacustrine origin of the Green River Formation. The first, 

involving a meromictic or permanently stratified lake, has 

been discussed in detail by Bradley (1948, 1964), Bradley 

and Eugster (1969), Smith (1974), and Desborough (1978). 

The second, involving a playa-lake complex, was 

introduced by Eugster and Surdam (1973) and has been 

elaborated in more recent articles (Eugster and Hardie, 

1975; Lundell and Surdam, 1975; Surdam and Stanley, 

1979).

In the meromictic-lake model, the lakes are thought to 

have been deeper than 30 m (100 ft) and have chemically 

and thermally stratified water columns (Sullivan, 1980). At 

their margins, the lakes gradually shallowed to shoals and 

other shallow-water environments (Sullivan, 1980). 

Shallow-water sedimentary rocks include calcareous and

C6 Green River Formation, Wyoming, Utah, and Colorado
UJ QQ

<I

UJ

Cfl

UJ «J

^

<;

UJ Q.

Z

r>

-j

TIPTON SHALE MEMBER

UJ M

^r

cs

UJ O

H

V3

UJ *:

_j

LAKE STAGE 

1

-100

-200

-300

-400

-500

-600

-100

-200

-300

-400

-500

0 4 8 12

TOTAL ORGANIC

CARBON, IN PERCENT

-600

-100

-200

-300

-400

-500

J -600

-35 -33 -31 -29 

8 13C,INPERMIL

10 12 14 16 18 

8 15 N,INPERMIL

20

Figure 2. Depth (in meters) profiles of total organic carbon content and 813C and 815 N compositions of kerogen 

in Green River Formation for Blacks Fork core, greater Green River basin, Wyoming. Location of core shown in 

Tuttle (this volume, fig. 1).

Pi

UJ

«

1

s

UJ 

UJ

u

UJ 

H

ac 

u

3

a.

DOUGLAS 

CREEK MEMBER

O

£L

UJ

^1-J

cs

UJO

H

UJ

_J

LAKE STAGE 

1

-500

-600

-700

-800

-900

-1000

-1100

-500

-600

-700

-800

-900

-1000

-500

-600

-700

-800

-900

-1000

10 20

TOTAL ORGANIC 

CARBON, IN PERCENT

-I -1100 

30 -33 -32 -31 -30 -29 -28

8 13 C,INPERMIL

I -1100 I__I

10 12 14 16 18 20 22 

8 15 N,INPERMIL

Figure 3. Depth (in meters) profiles of total organic carbon content and S13C and 815N compositions of kerogen in Green River 

Formation for Coyote Wash no.l core, Uinta basin, Utah. Location of core shown in Tuttle (this volume, fig. 1).

Carbon and Nitrogen Isotopic Biogeochemistry C7
Table 5. Isotopic composition of organic and inorganic 

nitrogen of evaporites of the Green River Formation from 

the 01A core, Piceance basin, Colorado

[Location of core shown in Tuttle (this volume, fig. 1). Leaders (--) 

indicate not determined. 515N in per mil]

Sample

Bedded halite. .........

Nodular nahcolite....

Bedded nahcolite ....

Depth NO3~ 

(meters) detected

512 +

.......648 +

....... 680

515N 515N 

(kerogen) (NO3~)

11.5 9.5

10.8 8.0

12.0

dolomitic mudstone, fine clastic sandstone, marlstone, and 

limestone of the carbonate facies, and deeper water rocks 

include laminated carbonate (oil shale) rocks (Sullivan, 

1980). Organic material preserved in the oil shale facies was 

derived from abundant plant and animal life in the upper 

euphotic zones (Sullivan, 1980). The lack of bioturbation in 

the Green River Formation indicates that monimolimnetic 

waters in the Green River lakes were a hostile environment 

for metazoan grazers, probably because they were anoxic 

(Sullivan, 1980). As a result, the bottom waters of these 

lakes had a high preservation potential for organic matter 

that originated in the surface waters and accumulated on the 

lake bottoms (Sullivan, 1980). Bradley (1929) interpreted 

the varves of the oil shale facies as representing a yearly 

cycle in these lakes. The organic-rich accumulation on the 

bottom of the centers of the basins was modified in the 

summer months by a large influx of carbonate due to 

increased photosynthetic activity of blue-green algae in the 

warmer surface waters. This process is thought to account 

for the characteristic varves of the Green River Formation, 

which consist of organic-rich laminae alternating with 

carbonate laminae (Sullivan, 1980).

Although the meromictic-lake model accounts for 

many aspects of the Green River Formation, Sullivan 

(1980) listed several observations that the model does not 

adequately explain: (1) the absence of a modern analogue; 

(2) sedimentary structures indicative of shallow water; (3) 

the presence of Magadi-type chert similar to that found in 

alkaline lakes of East Africa; (4) the close association of oil 

shale and trona, which suggests related depositional 

environments for these rocks; and (5) the predominance of 

dolomite as compared to calcite in these rocks, which is 

consistent with the playa-lake model.

The playa-lake model was proposed by Eugster and 

Surdam (1973) to explain inconsistencies of the deep 

stratified model. This model is based mostly on 

observations made at the margins of the basins. In this 

model, broad alluvial fans flanked the mountains 

surrounding the lakes and graded into broad playa flats 

(Sullivan, 1980). Shallow perennial lakes maintained by 

ground water and springs were located in the basin centers. 

The playa flats were areas of high evaporation, and

evaporation in the capillary zone above the water table on 

the playa produced a crust of alkaline salts (Sullivan, 1980) 

and resulted in deposition of dolomitic and calcareous 

mudstone of the carbonate facies (Eugster and Surdam, 

1973).

In the playa-lake model, only evaporites and 

laminated carbonate rocks (oil shale) are of truly lacustrine 

origin. Evaporite minerals were formed in the lakes because 

of the high evaporation rates, and the oil shales probably 

formed in the shallow lakes under less arid conditions 

(Sullivan, 1980). Such an environment would support 

abundant microscopic aquatic plants due to the high nutrient 

concentration of the waters, and the accumulation of this 

algal ooze would lead to the formation of oil shale (Sullivan, 

1980).

Boyer (1982) suggested that these two models of 

deposition are not necessarily mutually exclusive but may 

be complementary. He proposed that periods of meromixis 

in the Green River lakes may have resulted directly from 

prior episodes of brine generation in playa-lake 

environments.

Isotopically Enriched Organic Nitrogen

If the primary producers, mainly cyanobacteria 

(Bradley, 1931, 1970; Robinson, 1979), in the Green River 

lakes had met their nitrogen requirements solely by means 

of nitrogen fixation, then the nitrogen isotopic composition 

of the organic matter should be similar to that of 

atmospheric nitrogen (0%o). Alternatively, the isotopic 

composition of cyanobacteria assimilating combined 

inorganic-N would closely reflect the isotopic composition 

of the inorganic-N source (Mariotti and others, 1984).

Mineralogic evidence strongly suggests that the 

depositional environment of the Green River lakes was 

highly evaporitic and alkaline and had a pH of from 8 to 9.5 

(Smith and Robb, 1973). In such an environment, ammonia 

volatilization is likely to be an important physical process. 

The fractionation factor associated with this process is the 

largest for any in the nitrogen cycle [oc(NH4/NH3)= 1.034]. 

This fractionation, which enriches the residual ammonium 

in 15N, can be described as an open-system fractionation. 

The concentrations of NH3 and NH^ in solution are pH 

dependent and are described by the Henderson-Hasselbalch 

equation,

PH=pKa+log [base] 

[acid]'

The Ka for ammonium is 5.6xlO~10, and pKa is 9.25. 

For the equilibrium of ammonia in an aqueous solution,

NH3(aq)+H20(1) ^ 4 (aq)+OH (aq)'

C8 Green River Formation, Wyoming, Utah, and Colorado
Lakes

Total inorganic nitrogen 

Phytoplankton 

Algae 

Cyanobacteria 

Kerogen 

Total organic nitrogen

Rivers

Total organic matter 

Kerogen

Rain

Nitrate 

Ammonia

-1

 

 

 

* Green River Formation

~Ti 1

. . Messel Shale 

Green River Formation

 

 

 

 

1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 

2 -10 -8 -6 -4 -2 0 2 4 6 8 10 12 14 16 18 20 22

8 15 N

Figure 4. Distribution of 15N in freshwater environments. Data from Hoering (1957), Pang and Nriagu (1976, 1977), Peters and 

others (1978), Sweeney and others (1978), Wada and others (1975), Estep and Vigg (1985), and Collister (1989, Messel shale).

the concentrations of ammonia and ammonium at differing 

pH can be calculated by

pH=9.25+log

The rate of ammonia volatilization,

NH3(aq)-> NH3(g),

depends on the concentration of NH3 in solution, which is 

controlled by pH. As pH increases, a larger fraction of the 

total ammonia is in the un-ionized form and volatilization 

proceeds more rapidly. At pH 8, 5 percent of the total 

ammonium is in the form of NH3; at pH 9.5, the upper limit 

suggested by Smith and Robb (1973), 65 percent of the total 

ammonium is in the form of NH3. Accordingly, in such a 

system losses of NH3 will be extensive. Because NH3 is 

always depleted in 15N by 34%o relative to NH4+, residual 

NH4+ will be increasingly enriched in 15N. Thus, changes in 

the pH of the Green River lakes would have significantly 

affected rates of volatilization of NH3 and, therefore, the 

isotopic composition of the ammonium available for 

assimilation.

Assimilation of isotopically heavy residual 

ammonium would yield algal material enriched in 15N. But 

if such assimilation was the only important step in the 

fractionation of nitrogen isotopes of the Green River lake 

system, algae growing during the most saline and alkaline 

stages ought to be the most enriched in 15N. In fact, isotopic 

data from the 01A core (table 2) show the opposite

relationship. Samples of kerogens isolated from shale 

representing "fresh water" conditions are the most enriched 

in 15N; those from the most evaporitic horizons are depleted 

in 15N. Thus, although ammonia volatilization probably was 

pronounced and must have produced an isotopically heavy 

pool of aqueous ammonium during saline stages, other 

biologically imposed fractionations of the nitrogen isotopes 

must also have occurred and played an important role.

Models of Nitrogen Cycling in the 

Green River Lakes

We would expect the biogeochemical nitrogen cycle 

in a playa lake to be different from that in a stratified lake. 

Thus, models explaining the isotopic composition of 

organic matter in the Green River Formation are based upon 

these end members. Five samples (shown separately in fig. 

1) possessing mineralogy indicative of extreme evaporitic 

conditions were chosen to represent the playa-lake end 

member: three evaporites (bedded halite, bedded nahcolite, 

and nodular nahcolite) and samples 01-126 and 01-132 (for 

mineralogic compositions see Tuttle, this volume). We 

chose the Mahogany zone of the Parachute Creek Member 

of the Green River Formation (lake stage 3) in the Piceance 

basin as being representative of the meromictic-lake end 

member. The following discussion presents models of 

nitrogen cycling for these two environments based upon 

nitrogen isotopic compositions of preserved organic 

material. It must be kept in mind, however, that these two

Carbon and Nitrogen Isotopic Biogeochemistry C9
sets of samples were chosen to represent end members of 

depositional conditions and that many horizons in the Green 

River Formation may represent transitional stages between 

these two environments.

Saline (playa-lake) Stages

A proposed nitrogen cycle for the shallow, saline 

stages (well-mixed water column) of the Green River lakes 

is shown in figure 5. During these stages, reduced 

freshwater influx decreased nutrient supply and limited 

primary production. Rates of ammonification would be 

expected to decrease because of the decrease in detrital 

organic material being supplied to the sediments. Levels of 

salinity under these conditions were apparently almost 22 

percent (Bradley and Eugster, 1969); nitrification is 

inhibited at salinities above 15 percent (Rubenchik, 1929; 

Post and Stube, 1988). During the most saline stages, 

therefore, fluvial input would be the principal source of 

NO3~. Any nitrate entering the lakes in streams probably 

was quickly assimilated by the algal population before it 

could diffuse into the anoxic zone of the sediments. At the

resulting low levels of nitrate, denitrification would also be 

severely limited.

According to Stanier and others (1986), heterocystous 

cyanobacteria, when grown in the presence of nonlimiting 

concentrations of combined inorganic-N (NH4+ or NO3~), 

do not produce heterocysts (the specialized cells containing 

nitrogenase, the enzyme responsible for nitrogen fixation). 

Further, in the presence of combined inorganic-N, nitrogenase synthesis is inhibited in nonheterocystous cyanobacteria (Stanier and others, 1986). Under similar nonlimiting 

conditions, both phytoplankton and bacteria preferentially 

assimilate ammonium in the presence of both ammonium 

and nitrate (Mariotti and others, 1984). If the concentration 

of ammonium becomes small (as would be the case in the 

alkaline stages due to loss through volatilization), a larger 

fraction of the nitrogen used by the biota will be derived 

from nitrate or N2 (by nitrogen fixation). Control of the 

nitrogen isotopic composition of the organic material would 

pass from ammonium to nitrate (from fluvial input and 

reduced levels of nitrification) and N2 (from nitrogen 

fixation). The nitrogen isotopic composition of organic 

material in the Green River rocks indicates that nitrogen 

fixation was not the major source of nitrogen. Nitrate from

WATER COLUMN

SEDIMENT

Figure 5. Proposed nitrogen cycle for saline stages of Green River lakes. 

CIO Green River Formation, Wyoming, Utah, and Colorado
evaporites has an isotopic composition of +8 to +10%o 

(table 5). Use of this inorganic-N source would produce 

algal material having nitrogen isotopic compositions in the 

range of+6 to +10%o (Mariotti and others, 1984), and, even 

if rates of evaporation were high, the 815N of organic-N 

would decline. Isotopic compositions of nitrogen in kerogens from the most evaporitic facies are from +10 to about 

+12%o. These materials are marginally enriched in 15N 

relative to organic-N derived entirely from nitrate at +8 to 

+10%o, but this enrichment can be attributed to assimilation 

of a mixed inorganic-N source comprised of a small pool of 

ammonium (enriched in 15N) and a much larger pool of 

nitrate and dissolved N2 (depleted in 15N). Even if 

concentrations of nitrate were so low as to be nutritionally 

insignificant, assimilation of a mixture of isotopically heavy 

ammonium and atmospheric N2 could account for the 

relative depletion of 15N in organic matter from the saline 

facies of the Green River Formation.

Fresh (meromictic lake) Stages

During fresher stages, flooding of the lake basins 

from increased precipitation produced a stratified meromictic lake having a saline monimolimnion (Desborough, 1978; 

Sullivan, 1980; Boyer, 1982). A proposed nitrogen cycle for 

such a lake is shown in figure 6. Increased nutrient supply 

(from increased precipitation delivering higher loads of 

dissolved nutrients) would support increased primary

productivity. A larger input of detrital organic matter to the 

sediments would support greater rates of ammonification, 

and greater amounts of ammonium would be released into 

the overlying water column. "Fresh-water" conditions in the 

mixolimnion (the salinity was still approximately that of 

seawater; Bradley and Eugster, 1969) would be compatible 

with the salinity tolerance of nitrifying bacteria, and 

ammonium from the monimolimnion would support high 

rates of nitrification. Nitrate produced during nitrification 

would become available to denitrifiers upon diffusion into 

the monimolimnion. The increased supply of detrital 

organic material and the generation of nitrate by 

nitrification would support a large population of denitrifying bacteria. A similar coupling of nitrification and denitrification rates has been shown to occur in modern estuarine 

sediments (Jenkins and Kemp, 1984) and in modern lake 

sediments (Knowles, 1979).

Several potential sources of isotopically enriched 

inorganic-N species are indicated in figure 6. The 

ammonium remaining after volatilization and nitrification 

would be enriched in 15N. In fresh stages, loss of ammonia 

to the atmosphere would be decreased, and consequently a 

larger pool of ammonium would be available for 

assimilation. Assimilation of 15N-enriched ammonium 

would, in turn, produce isotopically enriched algal cellular 

nitrogen. Another source of isotopically enriched inorganicN during the fresher stages of these lakes is the residual pool 

of nitrate from denitrification, assimilation of which would

NH 4 assimilation 

a =0.993-1.013

MIXOLIMNION

(AEROBIC)

MONIMOLIMNION

(ANAEROBIC)

SEDIMENT

Figure 6. Microbial nitrogen cycle for meromictic lakes.

Carbon and Nitrogen Isotopic Biogeochemistry C11
also cause enrichment of total cell nitrogen. Finally, because 

of the elevated levels of combined inorganic-N species 

expected during these lake stages, nitrogen fixation would 

be expected to be less important to primary producers.

Therefore, extreme 15N enrichments during the meromictic stages of these lakes probably was due to 

assimilation of 15N-enriched ammonium and nitrate. The 

15N-enriched nitrate was produced by the process of denitrification, and the 15N-enriched ammonium was produced 

by the processes of volatilization and nitrification (fig. 6).

Comparison of Piceance, Uinta, and 

Green River Basins

Although our discussion up to this point has focused 

primarily on the 01A core (Piceance basin), some 

generalizations can be made regarding the other basins. 

Organic material from all three cores shows both extreme 

15N enrichments and a broad range of 815N values. Kerogens from both the 01A and the Coyote Wash (Lake Uinta) 

cores show an overall 15N-enrichment through time, 

although large deviations from this trend exist for individual 

samples. In contrast, kerogens from the Blacks Fork (Lake 

Gosiute) core show the most enriched values in rocks from 

the saline lake stage 2, and kerogens from a few samples 

from lake stages 1 and 3 suggest depletion of 15N in those 

stages. The overall pattern of 15N enrichment suggests that 

nitrogen was not generally a limiting nutrient in either lake; 

however, differences in the nitrogen isotopic signal through 

time indicate that details of the nitrogen cycles in these 

lakes were different.

Lake Gosiute (Wyoming) was considerably shallower 

than Lake Uinta and during lake stage 2 was significantly 

reduced in surface area (Turtle, this volume). Sample BF-22 

(Lake Gosiute), which contains the saline minerals shortite 

and trona (Turtle, this volume, appendix), is the most saline 

sample analyzed from the Blacks Fork core and may 

indicate a playa-lake environment. Kerogen isolated from 

this sample has a 515N value of +11.4%o, consistent with 

our model of nitrogen cycling in playa-lake environments. 

The remainder of the samples from lake stage 2, however, 

are more enriched in 15N (+16.7 to +18.3%o, table 3) and 

perhaps the community was utilizing one or more sources of 

15N-enriched inorganic-N. In both lakes, volatilization 

produced residual ammonium enriched in 15N. The extent to 

which this enrichment also occurred in organic-N would 

depend upon the importance of NH4+ as a nitrogen source. 

We can, therefore, speculate that during lake stage 2 NH4+ 

supplied a larger fraction of total nitrogen requirements in 

Lake Gosiute than in Lake Uinta.

Isotopically Depleted Total Organic Carbon

Isotopic variations in dissolved inorganic carbon 

(DIG) can be pronounced in lakes, and one consequence is

that phytoplankton can have very negative 513C values (-35 

to -45%o; Oana and Deevey, 1960; Deevey and others, 

1963; Deevey and Stuiver, 1964; Rau, 1978,1980). Organic 

material in the Green River Formation has a wide range of 

513C compositions. Interestingly, the range of isotopic 

compositions for the kerogens (-29.4 to -41.4%o) is much 

broader than that for total organic carbon (-28.7 to

-33.5%o).

Kerogens and total organic carbon, commonly 

analyzed in sedimentary isotopic studies, derive not only 

from primary producers but also from many other 

organisms including photosynthetic bacteria, a wide variety 

of heterotrophs, chemolithotrophs, methanogens, and methylotrophs. Consequently, the isotopic composition of the 

total organic material reflects not only that of the primary 

photosynthate but also isotope effects associated with a 

variety of secondary processes. In an attempt to resolve 

primary and secondary processes, sedimentary geoporphyrins were analyzed to more accurately determine the 

isotopic composition of the primary producers in these 

lakes.

Geoporphyrins are derived from the heteroaromatic 

tetrapyrrole nucleus of chlorophyll and can be considered to 

be "isotopic biomarkers" for primary producers. The tetrapyrrole nucleus is enriched in 13C relative to total biomass 

by about 0.5%o (Hayes and others, 1987; Takigiku, 1987); 

thus the isotopic composition of photosynthate of ancient 

primary producers can be estimated by the equation

5^-0.5,

where 5p is the isotopic composition of the primary photosynthate and 5por is the isotopic composition of the related 

porphyrins. Because of their low abundances in Green River 

Formation extracts, geoporphyrins from only three samples 

were analyzed. Consequently, isotopic variations in porphyrins from the Green River Formation may be considerably 

greater than that observed. The heaviest porphyrin (sample 

01-8) has a 513C value of -23.6%o, whereas that of the 

lightest (sample 01-36) is -28.6%o. Nickel geoporphyrins 

isolated from gilsonite (Uinta basin) have a 513C value of

-28.5%o. Corresponding values for total primary organic 

material are -24.1 and -29.1, respectively. Assuming the 

fractionations between aqueous CO2 and total cell carbon 

observed in oceanic plankton (18-19%o; Deuser and others, 

1986) and in neutral carbonate springs (19%o; Estep, 1984) 

are similar to those of algae in the Green River lakes, the 

isotopic composition of the CO2 dissolved in the photic 

zone of the Green River lakes must have varied at least 

between -4.6 and -9.6%o.

Because structures of the nickel geoporphyrins in the 

Green River Formation are not known, their precise origins 

are uncertain. Assuming the predominant organism to be 

blue-green algae (Bradley, 1970), it is reasonable to suggest 

that the majority of the porphyrins derive from algal 

chlorophylls (Bradley, 1963). However, significant amounts

C12 Green River Formation, Wyoming, Utah, and Colorado
of porphyrins may have been produced by photosynthetic 

bacteria. Because DIG in the anaerobic part of the photic 

zone of meromictic lakes is commonly depleted in 13C 

(Oana and Deevey, 1960; Deevey and others, 1963), 

photosynthetic bacteria fixing this CO2 should synthesize 

13C-depleted tetrapyrroles.

An alternate explanation for the isotopic range of the 

nickel geoporphyrins is that they reflect different sources of 

organic material having distinctive isotopic compositions. 

The geoporphyrins from algal material would be expected 

to be isotopically enriched relative to those from photosynthetic bacteria. Isotopic variations of porphyrin fractions 

from the Green River Formation may reflect changes in the 

abundance of algal versus bacterial primary production in 

response to changing lake conditions.

Sources of Isotopically Depleted 

Organic Carbon

Anaerobic Photosynthetic Bacteria

One source of isotopically depleted carbon material in 

meromictic lakes is photosynthetic sulfur bacteria (Fry, 

1986). In such lakes, the isotopic composition of the DIG is 

increasingly depleted with depth (Deevey and others, 1963; 

Takahashi and others, 1968). In the aerobic photic zone of 

Green Lake, Fayetteville, New York, the isotopic 

composition of the dissolved bicarbonate is between -6 and 

-12%o (Deevey and others, 1963; Takahashi and others, 

1968); however, at depths where photosynthetic sulfur 

bacteria occur (17-20 m; Fry, 1986), the isotopic 

composition of the DIG is between -17 and -21%o (Deevey 

and others, 1963; Torgersen and others, 1981). Thus, the 

photosynthetic sulfur bacteria (Chromatium and Chlorobium), which occur below the aerobic photoheterotrophs, 

fix 13C-depleted CO2 and consequently their cell material 

can be extremely depleted in 13C (513C, -40%o; Fry, 1986).

Sample 01-36, in which the 13C-depleted primary 

material is found, is from the saline facies of the Parachute 

Creek Member in the Piceance basin. This member reflects 

a lake stage in which evaporation rates greater than inflow 

resulted in increased salinity and decreased surface area of 

the lake (Tuttle, 1988). Sample 01-8, however, in which the 

isotopically heaviest primary organic material was 

observed, is from the Mahogany zone and represents a 

period in which inflow rates increased and the lake 

expanded and freshened. In the Piceance basin, with the 

exception of sample 01-8, the most 13C-depleted kerogens 

are from the Parachute Creek Member below the Mahogany 

zone (table 1). Within this part of the member, the 

mineralogy of the samples containing the most depleted 

kerogens include the saline minerals dawsonite, nahcolite, 

and halite (table 2).

The 13C-depletion of kerogens from the saline facies 

of the Parachute Creek Member in the Piceance basin may

reflect nutrient-limited primary production in the aerobic 

zone. Algal contribution to the overall organic matter would 

be greatly reduced, and the contribution from 

photosynthetic bacteria would become more important. 

These factors would both serve to produce 13C-depleted 

total organic material in the sediments representing these 

saline lake stages.

Methanogens and Methylotrophs

Enrichment of 13C in dolomite from the Green River 

Formation relative to that estimated for dissolved CO2 in the 

photic zone of these lakes indicates that aqueous CO2 in the 

sediment pore waters was not in isotopic equilibrium with 

dissolved CO2 in the overlying water column. This disequilibrium is particularly striking if the CO2 in the monimolimnia of these lakes had isotopic compositions similar to those 

of dissolved CO2 in the monimolimnia of modern-day 

meromictic lakes (-21 to -27%o; Oana and Deevey, 1960; 

Takahashi and others, 1968), and it requires a mechanism 

for enrichment of 13C in dissolved inorganic carbon in 

sedimentary pore waters.

The presence of 13C-enriched dolomites in anoxic 

pore waters from marine and lacustrine systems has been 

attributed to activities of methanogenic bacteria in the 

sediment (Irwin and others, 1977). Extreme 13C enrichment 

in nahcolite from the Green River Formation of the 

Piceance basin (813C values from +10 to +20%o) was 

suggested by Reitsema (1980) to indicate that bacterial 

fermentation by methanogenic bacteria was the source of 

this 13C-enriched CO2 . Reitsema pointed out that nahcolite 

precipitation depends on the presence of an excess of CO2 

and that amounts of CO2 below the concentration in 

equilibrium with the atmosphere lead to the formation of 

trona. Reitsema concluded that methanogenesis produced 

locally high concentrations of 13C-enriched CO2 that led to 

precipitation of 13C-enriched nahcolite. The presence of 

dolomite having 513C values as high as +ll%o also 

indicates that methanogenesis was an important process in 

these lakes.

Kerogen from one sample (01-8) is depleted in 13C 

relative to the associated geoporphyrin by more than 17%o. 

Takigiku (1987) suggested that the isotopic difference 

between geoporphyrin precursors and related lipids in fresh 

biomass is approximately -4%o. Therefore, in newly 

deposited primary organic material, the lipid fraction should 

be depleted relative to the porphyrins by 4%o. The 17%o 

depletion of kerogen in sample 01-8 relative to the geoporphyrin in this sample suggests that the isotopic composition 

of this kerogen (and probably other kerogen in the cores) 

reflects a secondary input of 13C-depleted organic matter. 

Assimilation of 13C-depleted methane from methanogenesis 

by micro-aerophilic methylotrophs living at the aerobicanaerobic boundary has been suggested to explain similar 

isotopic "mismatches" observed in the Messel shale (Hayes 

and others, 1987; Takigiku, 1987). Incorporation of this

Carbon and Nitrogen Isotopic Biogeochemistry C13
isotopically light methane would in turn result in 13Cdepleted biomass. The same process in the Green River 

lakes may account for the observed depletion of 13C in 

kerogen relative to geoporphyrins.

SUMMARY

The Paleogene Green River Formation was deposited 

in a highly productive system of alkaline, saline lakes in 

which algae and cyanobacteria were probably the main 

primary producers. The accumulation of organic material in 

this formation indicates that conditions favored preservation 

of organic material. Very large quantities of nutrients must 

have been supplied to the lake basins, and fluvial inputs of 

phosphorus from erosion of the Phosphoria Formation 

probably were substantial. Extreme enrichments in 15N in 

preserved organic matter from the Green River Formation 

suggest that nitrogen was not limiting in this lake system.

The broad variability of 815N values for organic 

material in the Green River Formation provides information 

regarding (1) the nutrition of the organisms in these lakes 

and (2) the cycling of nitrogenous compounds in this 

ecosystem. Although volatilization of ammonia in this 

alkaline environment probably produced 15N-enriched 

ammonium in the surface waters of these lakes, other 

biological fractionations must have been superimposed on 

this effect to produce the observed trends in nitrogen 

isotopic composition of the organic matter. Kerogens from 

the fresher stages of this lake system are the most enriched 

in 15N and those from the most saline and evaporitic stages 

are relatively depleted. Evaporation of NH3 ought to have 

been maximized during the latter stages. It is, therefore, 

surprising that evaporitic stages do not display the highest 

levels of 15N enrichment. To explain this paradox, it is 

suggested that isotope effects associated with nitrification 

and denitrification must also have played an important role 

in the control of overall nitrogen isotopic compositions. The 

isotope effect associated with nitrification would tend to 

increase 15N in ammonium and that effect associated with 

denitrification would increase 15N in nitrate. Taken 

together, these processes would enhance 15N enrichment 

associated with volatilization of ammonia. During salin'e 

stages, high salinity limited nitrification. This in turn limited 

the process of denitrification because these two microbial 

processes are coupled through NO3~. This simpler nitrogen 

cycle resulted in less reworking of the organic material 

during the saline stages. Consequently, 15N was less 

strongly enriched during the saline stages in spite of higher 

levels of evaporation.

The carbon isotopic composition of kerogens from 

the Green River Formation is from -41.4 to -28.8%o, 

whereas that of nickel geoporphyrins from two samples is 

-23.6 and -28.6%o. Geoporphyrins can be regarded as 

isotopic biomarkers for primary producers and therefore 

may provide an estimated range of 13C compositions of 

primary producers in the Green River lakes. In at least one

sample, a large isotopic depletion (17%o) in a kerogen 

relative to the associated geoporphyrin indicates significant 

contributions of 13C-depleted organic material into the 

sediments. Possible sources of isotopically depleted carbon 

include photosynthetic bacteria at the chemocline of a 

stratified lake and capture of isotopically light 

methanogenic-methane by methylotrophs.

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of southwestern Wyoming: Geological Survey of Wyoming

Report of Investigations 20, 50 p. 

Surdam, R.C., and Stanley, K.O., 1979, Lacustrine sedimentation

during the culminating phase of Eocene Lake Gosiute,

Wyoming (Green River Formation): Geological Society of 

America Bulletin, v. 90, p. 93-110.

Sweeney, R.E., Liu, K.K., and Kaplan, I.R., 1978, Oceanic 

nitrogen isotopes and their uses in determining the source of 

sedimentary nitrogen, in Robinson, R.W., ed., Stable 

isotopes in the earth sciences: New Zealand Department of 

Scientific and Industrial Resource, p. 9-26.

Takahashi, T., Broecker, W., Li, Y.H., and Thurber, D., 1968, 

Chemical and isotopic balance for a meromictic lake: 

Limnology and Oceanography, v. 13, p. 272-292.

Takigiku, R., 1987, Isotopic and molecular indicators of origins of 

organic compounds in sediments: Bloomington, Indiana 

University, Ph.D. thesis, 219 p.

Torgersen, T., Hammond, D.E., Clarke, W.B., and Peng, T.H., 

1981, Fayetteville, Green Lake, New York 3H-3He water 

mass ages and secondary chemical structure: Limnology 

and Oceanography, v. 26, p. 110-122.

Tuttle, M.L., 1988, Geochemical evolution and depositional 

history of sediment in modern and ancient saline lakes 

Evidence from sulfur geochemistry: Golden, Colorado 

School of Mines, Ph.D. thesis, 312 p.

Wada, E., and Hattori, A., 1978, Nitrogen isotope effects in the 

assimilation of inorganic nitrogenous compounds by marine 

diatoms: Geomicrobiology Journal, v. 1, p. 85-10.

Wada, E., Kadonaga, T., and Matuso, S., 1975, 15N abundance in 

nitrogen of naturally occurring substances and global 

assessment of denitrification from isotopic viewpoint: Geochemical Journal, v. 9, p. 139-148.

Warren, J.K., 1986, Shallow-water evaporitic environments and 

their source rock potential: Journal of Sedimentary 

Petrology, v. 56, p. 442-454.

Wedeking, K.W., Hayes, J.M., and Matzigkeit, U., 1983, 

Procedure of organic geochemical analysis, in Schopf, J.W., 

ed., Earth's earliest biosphere Its origin and evolution: 

Princeton, Princeton University Press, p. 428-441.

Wellman, R.P., Cook, F.D., and Krouse, H.R., 1968, Nitrogen-15 

microbial alteration of abundance: Science, v. 161, 

p. 269-270.

C16 Green River Formation, Wyoming, Utah, and Colorado
Chapter D

Trace Elements in Pyrites of the 

Green River Formation Oil Shales, 

Wyoming, Utah, and Colorado

By WENDYJ. HARRISON, D.R. PEVEAR, and P.C. LINDAHL

U.S. GEOLOGICAL SURVEY BULLETIN 1973

CEOCHEMICAL, BIOCEOCHEMICAL, AND SEDIMENTOLOGICAL STUDIES OF THE 

GREEN RIVER FORMATION, WYOMING, UTAH, AND COLORADO
CONTENTS

Abstract Dl

Introduction Dl

Sample descriptions D3

Pyrite separation D5

Trace-element analytical methods D12

Results and discussion D12

Electron microprobe analyses D14 

Geochemical interpretation D15 

Conclusions D17 

References cited D17

Appendix 1 Petrographic description of shale samples in West Willow no. 3 

core D19 

Appendix 2 Analytical methods D19

FIGURES

1. Map showing location of West Willow no. 3 core hole, Colorado D2

2. Chart showing location of pyrite-rich samples in West Willow no. 3 core D3

3. Graph showing mineralogy of samples of Green River Formation from West 

Willow no. 3 core D4

4. Drawings showing classification of common macroscopic and microscopic ironsulfide morphotypes in Green River Formation D5

5. Photomicrographs of morphotypes of pyrite in Green River Formation D6

6. Scanning electron micrographs of pyrites in Green River Formation D8 

7-9. Element correlation diagrams for :

7. Copper and cobalt in pyrites extracted from Green River oil shale D13

8. Cobalt, molybdenum, nickel, and copper in whole-rock samples of Green 

River oil shale D14

9. Copper in pyrite extracts and in whole-rock samples of Green River oil 

shale D14

10. Elemental X-ray intensity images of pyrite-rich oil shale D15

11. Graph showing fields of stability of solids as a function of pH and Eh in system 

Mn-S-CO2-H2O D16

TABLES

1. Semiquantitative X-ray diffraction mineralogy of shale D3

2. X-ray mineralogy of pyrite extracts D3

3. Pyrite morphotypes observed in petrographic section D8

4. Trace-element contents of pyrites D12

5. Measured, calculated, and published trace-element, H2O, and ash contents of 

whole-rock samples D13

Contents
Trace Elements in Pyrites of the Green River Formation 

Oil Shales, Wyoming, Utah, and Colorado

. Harrison, 1 D.R. Pevear,2 and P.C. Lindahl3

Abstract

Pyrites extracted from oil shale of the Paleogene Green 

River Formation (West Willow prospect, Colorado) were 

analyzed for As, Cd, Co, Cu, Mn, Mo, Ni, Pb, Se, and Zn by 

atomic absorption spectrophotometry. The purpose of this 

study was to confirm a mode of occurrence for these elements 

for use in chemical equilibrium modeling of leachates from oil 

shale.

Concentrations in the pyrite separates (>99 percent pure) 

are (in parts per million) As, 120-1,700; Cd, 0.16-9.1; Co, 

38-220; Cu, 243-1,130; Mn, 12-1,900; Mo, 25-185; Ni, 

69-223; Pb, 110-710; Se, <20; and Zn, 17-1,280. Electron 

microprobe analysis was used, where possible, to confirm that 

the trace metals were substituted into the pyrite structure and 

are not present as discrete metal sulfide phases in the pyrite 

extracts. Results of electron microprobe analysis also show that 

pyrites forming at different times during diagenesis have 

different trace-element suites.

The trace elements studied may be subdivided into two 

groups. Cadmium, cobalt, manganese, and zinc show a 

change in mode of occurrence from sulfide hosted in shale 

from depths below 610 m to carbonate hosted in samples from 

above 610 m. The remaining elements (As, Cu, Mo, Ni, Pb) 

show no stratigraphic variations. Mass-balance calculations 

show that substitution into pyrite is the principal mode of 

occurrence for As, Cu, Ni, and Pb in all samples and for Cd, 

Co, Mn, and Zn in the deeper samples. Molybdenum has 

another as yet undetermined mode of occurrence in addition 

to pyrite, and the possibility that cadmium and zinc may also

'Department of Geology and Geological Engineering, Colorado 

School of Mines, Golden, Colorado 80401.

2Exxon Production Research Company, P.O. Box 2189, Houston, 

Texas 77001.

3Argonne National Laboratory, 9700 South Cass Avenue, Argonne, 

Illinois 60439-4815.

be present in wurtzite (ZnS) cannot be precluded because this 

phase was probably lost during pyrite separation.

INTRODUCTION

The possibility that trace elements associated with oil 

shale can cause significant problems in processing and 

environmental pollution is widely recognized (Margheim, 

1975; Stollenwerk and Runnels, 1977; Fox, 1979; 

McWhorter, 1980; Klieve and others, 1981; Jackson and 

Jackson, 1982; Peterson and Wagner, 1982). Numerous 

studies of trace-element abundances in oil shale from the 

Paleogene (Eocene, locally Paleocene) Green River 

Formation of Colorado have been made (for example, 

Cooke, 1973; Meddaugh and Salotti, 1983), but few studies 

have addressed the more significant and difficult task of 

determining modes of occurrences for trace elements in the 

shales. Notable contributions to this problem are those of 

Desborough and others (1976) and Saether and others 

(1981).

Mode of occurrence is an essential item of 

information in assessing the processing and environmental 

behavior of trace elements. The knowledge that a given 

trace element is associated with a sulfide mineral, for 

example, makes possible prediction of the behavior of that 

element both in processing (such as volatilization or 

catalytic poisoning) and during subsequent storage and (or) 

disposal of the shale.

The geochemical association of the group of elements 

known as heavy metals (Pb, Zn, Ni, Co, Cu, Cd, Mo, As, 

Se) and sulfide minerals (pyrite, FeS2; pyrrhotite Fe^S) is 

well known (Fleischer, 1955; Hawley and Nichol, 1961; 

Springer and others, 1964; Roberts, 1981). Several of these 

elements can cause health problems, and their emission 

levels are regulated by the Environmental Protection 

Agency. Arsenic is known to be a catalytic poison during

Trace Elements in Pyrites D1
44 ABSAROKA VOLCANIC FIELD

ROCK . 

PRINGS

 UPLIFT 

0 50 100 KILOMETERS

EXPLANATION

Present-day limit of Tertiary rocks

Upper Cretaceous outcrop

Present-day outcrop of Lower 

Cretaceous through Cambrian rocks

Present-day outcrop of Precambrian rocks

Figure 1. Location of West Willow no. 3 core hole (sec. 7, T. 3 S., R. 97 W., Rio Blanco County, 

Colorado). Core is from Exxon Coal Resources West Willow Lease Block.

shale retorting (M. Siskin, Exxon Research and Engineering 

Company, written commun., March 1981).

The purpose of our study was to determine the 

concentrations of As, Cd, Co, Cu, Mn, Mo, Ni, Pb, Se, and 

Zn in pyrite extracted from Green River oil shale. Pyrite is 

the principal sulfide mineral in the shale (Milton, 1977), 

and, because its stability temperature is exceeded in many 

shale processing schemes, the potential for heavy-metal 

release is enhanced. The oxidation of pyrite when exposed

to groundwater provides a further opportunity for tracemetal contamination. Knowledge of trace-metal contents in 

pyrite is thus needed before quantitative prediction of the 

contamination potential of oil shale can be made by 

chemical equilibrium modeling (Harrison, 1990).

Acknowledgments. Discussions with HJ. Gluskoter 

(U.S. Geological Survey), GJ. Grabowski, Jr. (Exxon 

Production Research Company), and K.O. Stanley (Exxon 

Company, International) were helpful in constructing this

D2 Green River Formation, Wyoming, Utah, and Colorado
DEPTH IN METERS

BARIOD 

OIL YIELD SAMPLE

10 377 

26 415.7 

L4

8 612.5 L3 

14 682.1 

L1

Tgp u

Mz

R6

L5

R5

R4

R3

R2

R1

Tgg,

UINTA FORMATION

' Io, 

MEMBER PARACHUTE CREEK 

-700

GARDEN 

GULCH 

MEMBER

-800

DOUGLAS 

CREEK 

MEMBER

FORMATION RIVER GREEN

Figure 2. Location of pyrite-rich samples in West Willow no. 

3 core of Green River Formation. Oil shale zones: mz, 

Mahogany; R, oil shale rich, L, oil shale lean. Bariod oil yield 

in gallons per ton; depth in meters. Location of core hole 

shown in figure 1.

study. This study was entirely completed between 1981 and 

1983, while all three authors were employed in the Mining 

and Synthetic Fuels Division at Exxon Production Research 

Company. Walter J. Holzwarth operated the electron microprobe and gave valuable assistance in obtaining those 

analyses. G.B. Dreher, A.P. Sullivan, and J.R. McGee 

provided trace-element analyses of whole-rock samples by 

inductively coupled plasma-atomic emission spectrometry. 

Reviews by J. Crock, G.A. Desborough, and M.L. Turtle 

improved the quality of this report.

SAMPLE DESCRIPTIONS

Ten samples of shale from the Parachute Creek 

Member of the Green River Formation were selected from 

a core (West Willow no. 3) taken in the central Piceance 

basin (fig. 1). Samples were selected from a number of 

different zones (fig. 2), partly on the basis of observable 

high pyrite content, in attempt to identify variations in 

trace-element contents that might relate to stratigraphic 

position in the formation.

Sample mineralogy was determined using a 

petrographic microscope and X-ray powder diffraction 

analysis (XRD) of whole-rock samples. Petrographic

Table 2. X-ray mineralogy of pyrite extracts

[Location of core hole West Willow no. 3 shown in figure 1; location of 

samples shown in figure 2]

Sample

377.3

397.3

470.1*

608.6

612.5

682.1

682.4

Pyrite

Yes

Yes

Yes

Yes

Yes

Marcasite Pyrrhotite

Yes

Yes Yes

Yes

Yes

Yes

Other

Yes

Yes

No material available for XRD analysis.

Table 1. Semiquantitative X-ray diffraction mineralogy of shale

[Locations of core hole West Willow no. 3 shown in figure 1. Location of samples shown in figure 2. Values in weight percent]

Sample

377.3

397.3

415.7

470.1

557.7

608.6

612.5

682.1

682.4

Dawsonite

1

0

0

0

0

10

19

13

20

Clay1

10

18

0

0

0

5

5

0

6

Aragonite

0

0

3

0

0

0

0

0

0

Quartz

14

4

6

0

0

23

24

23

28

Potassium 

feldspar

17

19

0

14

11

5

7

7

7

Albite

6

7

23

32

28

11

2

2

2

Calcite

0

4

42

2

0

1

2

2

3

Dolomite

33

44

22

38

44

32

11

37

10

Siderite

19

0

0

1

2

8

19

15

17

Pyrite

0

4

4

14

15

6

10

2

7

Total phyllosilicates.

Trace Elements in Pyrites D3
230-

300

400-

a 500-

UJQ

UJ

cc 

O 

O

600 

700-

800-

z

ag %< 5u <O O PYR SJD DOL CAL ALB KSP QT2 ARA ANA CLA DAW

 A GROOVE

MAHOGONY 

ZONE

'-B GRO( VE

KEY

PYR-PYRITE 

SID-SIDERITE 

DOL- DOLOMITE

10 PERCENT =D

CAL-CALCITE 

ALB-ALBITE

KSP-POTASSIUM 

FELDSPAR 

QTZ-QUARTZ

ARA-ARAGONITE

ANA- ANALCIME

CLA-TOTAL PHYLLOSILICATES

DAW-DAWSONITE

Figure 3. Mineralogy of samples of Green River Formation from West Willow no. 3 core. Baroid oil yield in gallons 

per ton.

descriptions are presented in appendix 1. Percentages of 

the minerals dawsonite, aragonite, quartz, potassium feldspar, albite, calcite, dolomite, siderite, pyrite, and total 

phyllosilicates are shown in table 1 and are probably 

accurate to ±20 percent at best, perhaps ±50 percent for 

some minerals such as pyrite.

Our data are a subset of a larger data set of mineral 

abundances for the whole core. The larger data set also 

shows a change in mineralogy at a depth of about 607 m, 

below which dawsonite, quartz, and siderite are significant 

rock-forming phases (fig. 3). Calcite, clay, and aragonite are 

more abundant in higher parts of the formation.

D4 Green River Formation, Wyoming, Utah, and Colorado
COMMON IRON-SULFIDE MORPHOTYPES IN THE GREEN RIVER FORMATION

MACROSCOPIC (1 mm-10 m) MICROSCOPIC (less than 1 mm)

PODS AND LENSES BLADES (spicular)

CONCRETIONS COMPOUND BLADE CLUSTERS AND BANDS

BOUDINAGES SPECKS (less than 4 microns)

STRINGERS

GRANULAR BLEBS

MEGAFRAMBOIDS

IRREGULAR MASSIVE BLEBS

CUBES (few encountered)

Figure 4. Classification of common macroscopic and microscopic iron-sulfide morphotypes in Green 

River Formation. Modified from Boyer and Cole (1983).

PYRITE SEPARATION

Pyrite was removed, together with some other sulfide 

phases, from powdered shale samples by a series of acid 

dissolutions and density separations described in appendix 

2. Pyrite extracts were examined by scanning electron 

microscopy (SEM) and XRD to assess purity. Results of

XRD analysis of the pyrite extracts are shown in table 2; all 

samples contain pyrite and small amounts of marcasite (see 

Milton, 1977). In addition, pyrrhotite was identified in one 

sample, and two samples had several very low intensity 

peaks that could not be assigned to any mineral. Pyrrhotite 

is soluble during the acid decomposition treatment and may 

have been removed from many of our samples.

Trace Elements in Pyrites D5
Figure 5 (above and facing page). Morphotypes of pyrites in Green River Formation. A, 

Bladed morphology; pyrite replacing salt minerals or pyrrhotite, 1 cm=46 jam. B, Pyrite 

replacement of dolomite rhomb, 1 cm=46 jam. C, Diagenetic pyrite blebs, 1 cm==28 jam. D, 

Finely disseminated, syndepositional pyrite, 1 cm=103 (im.

D6 Green River Formation, Wyoming, Utah, and Colorado
Trace Elements in Pyrites D7
Table 3. Pyrite morphotypes observed in petrographic section

[Location of core hole West Willow no. 3 shown in figure 1. Location of samples shown in figure 2]

Sample

3773

3973

415.7

470.1

557.7

608.6

612.5

682.1

682.4

Disseminated

yes

yes

yes

Irregular 

blebs

yes

yes

yes

yes

Cubes

yes

yes

yes

Fracture 

Blades filling

yes

yes

A very few fragments of organic matter and undissolved minerals (tentatively identified as rutile TiO2 and 

feldspar KAlSi3O8) were found, and the samples are 

estimated to be more than 99 percent pure. Wurtzite (ZnS), 

reported to be widespread (Milton, 1977), was not observed 

in SEM or XRD investigations, but it is soluble in most 

dilute acids and, if present, may have been dissolved during 

pyrite extraction.

A number of morphotypes of pyrite have been 

identified from the Green River Formation (Pabst, 1971; 

Boyer and Cole, 1983). Figure 4 shows the classification 

scheme devised by Boyer and Cole (1983). In our samples 

we observed abundant blades (a replacement texture after 

some saline minerals and (or) pyrrhotite blades), specks 

(finely disseminated pyrite), bands, and blebs (irregular 

pyrite masses). Examples are shown in figure 5. A distinct 

stratigraphic distribution of these morphotypes was noted in 

the West Willow no. 3 core (table 3). Disseminated pyrite 

can be distinguished from tiny pyrite cubes, which are 

scattered throughout samples and concentrated into lenses 

and blebs, by lack of crystal morphology (either not present 

or too fine grained to see).

An SEM investigation of the pyrite extracts yielded 

information about pyrite morphology on a different scale: 

(1) spongy pyrite masses, (2) small euhedral crystals, (3) 

spheres, (4) fragments of large crystals and pseudomorphs, 

(5) replacement of organic material, and (6) pyrrhotite 

hexagonal plates (fig. 6). (Our identification of pyrrhotite is 

based on the hexagonal crystal morphology and XRD data. 

Sweeney and Kaplan (1973) also identified the hexagonal 

plate morphology as characteristic for pyrrhotite.) The 

textural features of extracted pyrite observed using SEM are 

difficult to relate to those observed at lower magnification 

in petrographic thin sections. It is of interest that the fine 

textures illustrated in figures 6C and D have been preserved 

despite the physical and chemical dissolution procedures 

used to extract the sulfide minerals. We do not believe that 

the spheres represent abraded crystals, rather they are a 

morphology typical of a pyrite precursor, greigite (Fe3S4), 

now preserved as pyrite (Sweeney and Kaplan, 1973).

Spheres and crystals have a comparable range of sizes; 

abrasion likely would have reduced the range observed in 

the sphere population relative to that of the crystal 

population. The spheres could represent disaggregated 

framboids; however, the general absence of framboidal 

pyrite in this particular core and the absence of any partially 

disaggregated fragments argue against this hypothesis. 

(Framboidal pyrite was not observed in petrographic thin 

sections of these rocks; however, it commonly is associated 

with organic matter, and any small quantities of framboidal 

pyrite present might have been lost during the separation 

procedure.) Blades (fig. 5A) observed in thin section 

comprise many tiny pyrite crystals as do some of the blebs 

and larger irregular masses.

We suggest that the spheres and tiny euhedral pyrite 

crystals formed at the time of deposition of the organic-rich 

sediments, either as pyrite or an iron-sulfide precursor (Lord 

and Church, 1983). All other pyrite morphologic types 

formed later, during and after diagenesis, and all show 

replacement features of various kinds. Our conclusion is in 

agreement with Dyni's (1983) identification of three stages 

of sulfide mineralization in the Green River Formation:

(1) early diagenetic phase (our finely disseminated pyrite),

(2) late diagenetic phase (blades and the replacement of 

other minerals formed during sediment dewatering), and

(3) posflithification phase (pyrite precipitated in fractures 

and cracks).

Figure 6 (following pages). Scanning electron micrographs of 

pyrites in Green River Formation A, Small euhedral crystals, 

spheres, and spongy pyrite masses. B, Small pyrite spheres. C, 

Pyrite replacement of crystal fragments and spongy pyrite 

masses. D, Shell fragment replaced by pyrite. E, Hexagonal 

pyrrhotite plates. F, Energy-dispersive X-ray intensities from 

sample shown in C

D8 Green River Formation, Wyoming, Utah, and Colorado
Trace Elements in Pyrites D9
D10 Green River Formation, Wyoming, Utah, and Colorado
Trace Elements in Pyrites D11
Table 4. Trace-element contents of pyrites

[Location of core hole West Willow no. 3 shown in figure 1; location of samples shown in figure 2. Concentrations in parts per million; oil

yield in gallons per tonne]

Sample

3773

3973

415.7

470.1

557.7

608.6

612.5

682.1

682.4

Zone

Mahogany

L6

R6

L5

R4

R4

L3

R3

R3

Baroid 

oil 

yield

9.57

5.41

26.37

7.55

1234

23.0

8.42

13.74

4.16

As

220

360

130

120

1,770

200

220

340

200

Cd

0.26

0.16

034

0.18

0.52

0.20

1.4

9.1

2.9

Co

99

93

75

113

50

38

150

161

220

Cu

882

552

516

243

328

292

878

771

1,130

Mn

41

25

17

14

34

12

1,420

1,900

1,800

Mo

41

25

185

124

47

179

151

111

181

Ni

223

155

96

103

219

110

127

69

158

Pb

710

230

200

240

110

230

150

110

270

Zn

72

76

39

22

46

17

158

1,280

441

TRACE-ELEMENT ANALYTICAL METHODS

Pyrite mineral separates obtained from the samples of 

oil shale were dissolved in nitric acid, and solutions were 

analyzed using atomic absorption (AA) spectrophotometry. 

Conventional flame AA was used to determine Fe and the 

trace elements Co, Cu, Mn, Ni, Pb, and Zn. Flameless AA 

using a heated graphite atomizer (HGA) was employed to 

determine the trace elements As, Cd, Mo, and Se because 

concentrations of these elements are below the flame AA 

detection limits. Unfortunately, the selenium content proved 

to be below that of the flameless AA detection limit due to 

interference caused by the high iron content in the samples. 

Further details of instrumental parameters are given in 

appendix 2.

RESULTS AND DISCUSSION

Trace-metal concentrations in pyrites extracted from 

the oil shale are given in table 4. Only selenium is present 

in amounts below the detection limit of the primary 

analytical techniques (< 10-20 ppm). Two groups of trace 

elements can be identified: (1) those for which there is no 

systematic stratigraphic variation (Cu, Pb, Ni, As, Mo) and 

(2) those for which a distinct increase in concentration is 

observed below 608.6 m depth (Mn, Zn, Cd, Co). The 

magnitude of the increase, from a factor of 100 for 

manganese, 10 for zinc and cadmium, to 5 for cobalt, argues 

against analytical or sample preparation error. Also of note 

is that a considerable amount of the increase is in a short 

vertical stratigraphic interval; for example, compare data for 

samples 682.1 and 682.4 (table 4) (all sample numbers 

represent depth, in meters), separated by only a third of a 

meter. There is no correlation between element content in 

pyrite and whole-rock oil yield in these samples and no 

correlation between individual elements other than that

between zinc and cadmium, which is very common, and 

copper and cobalt (fig. 7).

Trace-element concentrations in whole-rock samples 

were determined for mass-balance considerations. These 

data are shown in table 5 and analytical details are described 

in appendix 2. Although the number of samples is limited, 

correlations between molybdenum and cobalt (fig. 8A), 

nickel and cobalt (fig. 8#), cobalt and copper (fig. 8Q, and 

nickel and copper (fig. 8D) most likely are significant. 

Comparison of pyrite extract contents with whole-rock 

contents show only one significant correlation, that for the 

element copper (fig. 9). The lack of correlation between 

whole-rock and pyrite trace-element contents indicates 

complexities in the modes of occurrences for these 

elements.

Our pyrite extraction technique is not quantitative 

because very fine grained pyrite remains with the kerogen 

during the sink/float separations. Nevertheless, massbalance calculations were attempted with the intent of 

estimating the percentage of each trace element that is 

sulfide associated. In these calculations (table 5), wholerock trace-element contents were determined by assuming 

that all of each trace element is in the pyrite; the percentage 

of pyrite in the rock was determined from XRD analysis. 

The theoretical whole-rock contents were then compared 

(table 5) with the analytical data and a compilation of 

literature data from Cook (1973), Desborough and others 

(1976), Giaugue and others (1981), and Meddaugh and 

Salotti (1983).

Although detailed comparison of the calculated and 

analytically determined whole-rock content was not 

possible because of the errors inherent in the XRD 

determination of pyrite abundance, a high percentage of the 

Cu, Pb, Co, Ni, and As in our samples most likely is sulfide 

associated, held in some way in the pyrite structure. Below 

608.6 m in the West Willow no. 3 core, Mn, Zn, and Cd are 

also, predominantly sulfide associated; above this level,

D12 Green River Formation, Wyoming, Utah, and Colorado
z 

uiQ 1,000

Ed

£j| 800

ZU D.

Uj</> 600

£5 

8| 400

200

Figure 7.

extracted

20 40 60 80 100 120 140 160 180 200 220 240 

COBALT IN PYRITE, IN PARTS PER MILLION

Correlation between copper and cobalt in pyrites 

from Green River oil shale.

some other phase, probably carbonate, is the principal 

host of these three elements. Part of the molybdenum is also 

sulfide associated, but our calculated values are generally 

lower than those reported in the literature (table 5) and 

molybdenum may have a second mode of occurrence (see 

also Giaugue and others, 1981).

Our working hypothesis is that all the heavy metals 

are substituted directly into the pyrite structure; however, a 

very small amount of a sulfide mineral such as ZnS, MnS, 

or PbS in the pyrite extract could easily account for the 

measured trace-element contents. X-ray diffraction data 

show that the pyrite extracts are composed primarily of 

pyrite and some marcasite (a polymorph of FeS2) and 

pyrrhotite (Fei.xS), but the XRD technique is not sensitive 

to small quantities of minerals. No separate sulfide phases

Table 5. Measured, calculated, and published trace-element, H>O, and ash contents of whole-rock samples

[Location of core hole West Willow no. 3 shown in figure 1; location of samples shown in figure 2. Trace-element contents except Se in parts per 

million; Se, H2O, and ash contents in weight percent; nd, not determined; blank indicates no data available]

Whole-rock analyses

Sample As

377.3 8.2

397.3 16

415.7 9.0

470.1 9.4

557.7 52

608.6 99

612.5 23

682.1 15

682.4 26

Cd

<0.10

<0.10

<0.10

<0.10

<0.10

<0.10

<0.10

<0.10

<0.10

Co Cu

9.2 78

6.9 41

5.4 30

7.2 49

6.2 42

6.8 41

19 100

9.1 43

15 98

Mn

330

325

160

250

295

240

175

170

150

Mo

16

3.1

18

14

17

7.5

73

42

93

Ni

31

25

18

31

33

36

100

29

41

Pb

24

12

4.1

14

9.6

7.7

17

8.3

4.8

Se

<0.00

<0.00

<0.00

<0.00

<0.00

<0.00

<0.00

<0.00

<0.00

Zn

95

88

48

99

120

89

125

73

115

H20

0.28

0.37

0.30

0.61

0.52

0.47

0.23

0.35

0.26

Ash

76

75

58

81

77

71

75

79

76

Calculated trace-element contents

Percent pyrite

Sample X-ray

377.3 0

397.3 4

415.7 4

470.1 14

557.7 15

608.6 6

612.5 10

682.1 2

682.4 7

Assumed

5

5

5

15

15

10

10

5

10

As

44

72

26

8

120

20

22

68

20

Cd Co

0.05 20

0.03 19

0.07 15

0.01 8

0.03 3

0.02 4

0.14 15

2.0 32

0.3 22

Cu

175

110

100

16

22

29

88

155

115

Mn

8

3

3

41

2.3

1.2

140

380

180

Mo

8

5

37

8

3

18

15

22

18

Ni

45

31

19

7

15

11

13

14

16

Pb

140

46

40

16

7

23

15

22

27

Zn

14

15

8

1.5

3

2

16

255

44

Published values

Percent assumed

Reference pyrite

Desborough and others (1976) Low

Cook (1973)

High

Meddaugh and Salotti (1983) Average

Giauque and others (1981)

High

Low

High

As

25

75

7.2

10.8

71

5

133

Cd

<1

1.2

nd

nd

nd

Co

5

20

39

4

14

Cu

33

114

15

11

63

Mn

181

296

34

190

460

Mo

7

40

nd

47.5

83

4

53

Ni

20

40

11

11

31

Pb

19

58

10

5

41

Zn

49

112

13

38

153

Trace Elements in Pyrites D13
A 

100

80

40

20 V '

7 9 11 13 15 17 

COBALT, IN PARTS PER MILLION

O 21

d 19

S

OC 17

Ul

0- 15

to

9

. * *

30 40 50 60 70 80 90 

COPPER, IN PARTS PER MILLION

100

50

30

9 11 13 15 17 

COBALT, IN PARTS PER MILLION

19

70

50

 30 

Ul X.

O 10

SO 60 70 80 90 

COPPER, IN PARTS PER MILLION

100

Figure 8. Correlation between cobalt, molybdenum, nickel, and copper in whole-rock samples of Green River oil shale.

80100 

ccj 

iu=! -«Z 80

 ^LUiu 5o.

2<2 

trtr

O^ 

O 20

..

240 360 480 600 720 840 960 

COPPER IN PYRITE, IN PARTS PER MILLION

1,080

Figure 9. Correlation between copper in pyrite extracts and 

in whole-rock samples of Green River oil shale indicates that 

pyrite may be primary mode of occurrence for this element.

were found when the samples were examined by SEM. The 

most conclusive evidence of mode of occurrence comes 

from a preliminary study of pyrite using the electron 

microprobe.

Electron Microprobe Analyses

To obtain more conclusive evidence for our 

hypothesis that trace metals are substituted into pyrite and

are not present as trace amounts of metal-sulfide phases, we 

analyzed some of the pyrites for a suite of trace elements. 

Preliminary data indicate that some of the elements are 

substituted in the pyrite structure.

We analyzed for Mn, Zn, As, Cu, and Pb in samples 

453.1, 612.5, and 682.1 and in two samples not from the 

interval originally studied, 344.8 and 810. Operating 

conditions for the Cameca electron microprobe were an 

accelerating voltage of 30 kV and a beam current of 100 

nanoamps. Counting times were 30 seconds for all elements 

except arsenic, for which 10-second counting times were 

employed because significant volatilization was observed. 

Detection limits are about 100 ppm for As, Mn, and Zn, 200 

ppm for Cu, and 500 ppm for Pb. Pyrites having known 

contents of the elements analyzed were used as standards. 

Lead and copper concentrations commonly were not 

sufficiently above the detection limits for quantitative data 

to be obtained, and no concentrations can be quoted for 

these two elements until we obtain more appropriate standards. The presence or absence of lead and copper, however, 

can be reliably noted. Elemental X-ray maps of a polished 

pyrite-rich section of oil shale from 810m depth in the core 

are shown in figure 10. Although the resolution is poor for 

the arsenic distribution (concentration about 650 ppm), its 

association with pyrite, in a rather uniform pattern, is clear.

D14 Green River Formation, Wyoming, Utah, and Colorado
Figure 10. Elemental X-ray intensity images (maps) of polished pyrite-rich oil shale (sample WW-810; sample location shown 

in fig. 2). Antipathetical relationship between Fe Ka distribution, exposed for 5 min, and Si Ka distribution, exposed for 5 min, 

illustrates distribution of pyrite and silicate material in section. Correlation of S Ka and As La distribution with Fe Ka distribution 

shows the quite homogeneous arsenic substitution in pyrite.

Different morphologic forms of pyrite contain 

distinctive suites of trace elements. Fine-grained, euhedral, 

disseminated pyrite, typically only in samples below 608.6 

m in the West Willow core (see appendix 1), contains large 

amounts of manganese, from 2,000 to 3,500 ppm, zinc of as 

much as about 160 ppm, and detectable copper and lead. 

Arsenic was very rarely detected. Larger pyrite grains in 

replacement or diagenetic textures, such as fracture filling 

or replacement of pre-existing crystals, contain no 

manganese (or rarely 100-200 ppm), but many contain 

arsenic, copper, and lead. Arsenic contents are from 200 to 

1,000 ppm. We could not distinguish different trace-element 

suites among these later formed pyrites.

We tentatively conclude that there are stratigraphic 

and morphologic controls on trace-element contents in the 

pyrites, both of which we can account for in the geochemical model described below.

GEOCHEMICAL INTERPRETATION

Our assignment of the metals As, Co, Cu, Mo, Ni, and 

Pb to a sulfide mode of occurrence, Mn, Zn, and Cd to both 

carbonate and sulfide, and Mo to another, yet undetermined 

(in this study) mode of occurrence agree with published data 

on the Green River Formation (Desborough and others, 

1976; Saether and others, 1980, quoted in Giaugue and 

others, 1981; Giaugue and others, 1981). The incorporation 

of manganese in pyrite is surprising. More that 50 percent of 

the 900 pyrite analyses reported by Fleischer (1955) show 

less than 10 ppm Mn, although contents as high as 10,000 

ppm are recorded. Pyrrhotite, on the other hand, readily 

accepts manganese and zinc into its lattice (Wedephol, 

1969). The geochemical coherence of manganese and cobalt 

is well documented (Burns, 1976), and the elements Mn,

Trace Elements in Pyrites D15
o o.oo

-0.20

-0.60

-aeo

pH

Figure 11. Fields of stability of solids as a function of pH and 

Eh at 25 °C and 1 atm in system Mn-S-CO2-H2O. Total CO2 

species activity is 2,000 mg/L as HCO3~, total sulfur species 

activity is 2,000 mg/L as SO42~(Hem, 1972).

Zn, Co, and Cd are all enriched in the pyrites from the lower 

part of the core.

Sweeney and Kaplan (1973) showed experimentally 

that one mechanism for the formation of pyrite from 

aqueous solution involves addition of sulfur to the precursor 

mackinawite (FeS0 9) to produce greigite (Fe3S4) or pyrrhotite (Fe^S). Further addition of sulfur results in pyrite 

formation. The morphologic types typical of greigite (spherules 1-5 Jim) and of pyrrhotite (hexagonal plates), which we 

also observed in the West Willow no. 3 core, were observed 

by Sweeney and Kaplan (1973) to be retained by the pyrite 

except when it formed at high temperature 

(>170 °C) when the spherules formed euhedral crystals. We 

propose that Mn, Zn, Cd, and Co were incorporated into the 

greigite and pyrrhotite precursors and retained during 

subsequent pyrite formation.

Pyrite that formed later, during diagenesis, has a 

different trace-element chemistry. This pyrite, precipitated 

directly from pore water, contains As, Pb, and Cu that may 

have been remobilized from organic material and clays.

The incorporation of manganese in pyrite may have 

implications for the water chemistry at the time of sediment 

accumulation. Manganese sulfide (MnS-alabandite) forms 

only under very restricted conditions. In order to precipitate 

MnS, the aqueous sulfide to carbonate activity ratio must be 

> 10~~5 (compare ratios of 1(T15 for precipitation of PbS, 

ZnS, FeS2, and CuS; Boden, 1957). The presence of 

abundant carbonate minerals in the Green River Formation 

implies that the carbonate activity ratio was high; thus, the 

formation of MnS would require exceptionally high sulfide 

activities. Because we cannot estimate the activity of either 

aqueous sulfide or carbonate with any accuracy, it is not 

easy to improve on the Eh-pH diagram shown in figure 11, 

reproduced from Hem (1972). The stability field of MnS 

moves to lower values of pH and higher values of Eh as 

sulfur activity increases (Carrels and Christ, 1965).

Two alternative hypotheses can be put forward for the 

change, or apparent change, in mode of occurrence of Mn, 

Co, Cd, and Zn with depth in the core. First, a reduction in 

aqueous sulfide activity, pH, or alkalinity would favor the 

stability of these four elements in carbonate phases relative 

to sulfide phases. Alternatively, small pyrite grains may 

have been removed by leaching from higher parts of the 

section that is, they were not preserved and the change 

in mode of occurrence is only apparent. In support of the 

former is the change in major mineralogy of the deeper 

samples, notably the presence of the mineral dawsonite 

[NaAl(CO3)(OH)2], the formation of which presumably 

requires high alumina solubilities that are present under 

alkaline conditions. In support of the latter is the presence of 

the well-documented "leached zone" (406-720 m depth, fig. 

3). It is difficult, however, to envisage the selective removal 

of tiny pyrite crystals (<l-5 Jim) from such a low-porosity 

formation, and presumably any leaching of the pyrite must 

have occurred during sediment dewatering and compaction. 

We believe that a change in lake-water and possibly 

porewater chemistry is the best interpretation for our data: 

small pyrite crystals formed in much smaller quantities in 

samples above 607 m, and Mn, Co, Cd, and Zn have a 

different mode of occurrence in this part of the formation.

An important generalization can be made from our 

analysis of trace-element data: the Green River Formation is 

not a homogeneous, uniform mineral assemblage. Rather, 

changes in trace-element mode of occurrence indicate that 

chemically significant variations exist on a large scale 

within the formation. Furthermore, the observation that, 

depending on its paragenesis, one mineral (pyrite) may have 

various trace-element contents is indicative of a grainto-grain-scale heterogeneity that reflects a complex diagenetic history for the oil shale. The homogeneity noted by 

many workers is in part related to averaging whole-rock 

analyses over significant intervals, a process that masks 

information that can be obtained by more detailed study.

D16 Green River Formation, Wyoming, Utah, and Colorado
CONCLUSIONS

Pyrites extracted from oil shales of the Green River 

Formation were analyzed by atomic absorption spectrophotometry for the elements As, Cd, Co, Cu, Mn, Mo, Ni, Pb, 

and Zn. Two distinct types of pyrite were identified on the 

basis of electron microprobe analysis. Manganese-rich, 

finely disseminated pyrite formed during sediment 

deposition (stage 1 of Dyni's (1983) classification), and 

manganese-free pyrite containing As, Pb, and Cu and 

showing a variety of replacement textures formed during 

and after diagenesis (stages 2 and 3 of Dyni's (1983) 

classification).

A stratigraphic variation in the mode of occurrence of 

Mn, Zn, and Cd was observed. Below about 610 m in the 

West Willow no. 3 core, these three elements are sulfide 

hosted. Above this depth, carbonate minerals are probably 

the host mineral for the trace elements. This change is 

related to evolution of the alkalinity and salinity of the 

sedimentary environment that favored the stability of 

manganese, zinc, and cadmium carbonate minerals relative 

to sulfide minerals.

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composition of the ore forming fluids: Economic Geology, 

p. 333-353.

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mineralization in the Parachute Creek Member of Green 

River Formation, Colorado and Utah, m Sixteenth Oil Shale 

Symposium proceedings: Golden, Colorado School of 

Mines, p. 160-175.

Burns, R.G., 1976, The uptake of cobalt into ferromanganese 

nodules, soils and synthetic manganese (IV) oxides: 

Geochimica et Cosmochimica Acta, v. 40, p. 95-102.

Chao, T.T., and Sanzolone, R.F., 1977, Chemical dissolution of 

sulfide minerals: U.S. Geological Survey Journal of 

Research, p. 409-412.

Cole, R.D., 1975, Sedimentology and sulfur isotope geochemistry 

of Green River Formation (Eocene), Uinta Basin, Utah, 

Piceance Creek Basin, Colorado: Salt Lake City, University 

of Utah, Ph.D. thesis, 276 p.

Cook, E.W., 1973, Elemental abundances in Green River oil shale: 

Chemical Geology, v. 11, p. 321-324.

Czamanske, G.K., and Ingamells, C.O., 1970, Selective chemical 

dissolution of sulfide minerals A method of mineral 

separation: American Mineralogist, v. 55, p. 2131-2134.

Desborough, G.A., Pitman, J.K., and Huffman, C, Jr., 1976, 

Concentration and mineralogical residence of elements in 

rich oil shales of the Green River Formation, Piceance 

Creek Basin, Colorado, and the Uinta Basin, Utah A 

preliminary report: Chemical Geology, v. 17, p. 13-26.

Dolezal, J., Povondra, P., and Sulcek, Z., 1968, Decomposition 

techniques in inorganic analysis: London, Iliffe, p. 224.

Dyni, J.R., 1983, Distribution and origin of sulfur in Colorado oil 

shale, in Sixteenth Oil Shale Symposium proceedings: 

Golden, Colorado School of Mines, p. 144 159.

Fleischer, M., 1955, Minor elements in some sulfide minerals: 

Economic Geology 50th Anniversary Volume, 

p. 970-1024.

Fox, J.P., 1979, Water quality effects of leachates from an in-situ 

oil shale industry: Lawrence Berkeley Laboratory, Report 

LBL-9987, Berkeley, California.

Garrels, R.M., and Christ, C. L., 1965, Solutions, minerals, and 

equilibria: San Francisco, Freeman, Cooper & Company, 

450 p.

Giauque, R.D., Fox, J.D., and Smith, J.W., 1981, Characterization 

of two core holes from the Naval Oil Shale Reserve Number 

1: U.S. Department Energy Report PB81-167736, 176 p.

Harrison, W.J., 1990, Modelling fluid-rock interactions in 

sedimentary basins, m Cross, T.A., ed., Quantitative 

dynamic stratigraphy: New Jersey, Prentice-Hall, 

p. 195-231.

Hawley J.E., and Nichol, I., 1961, Trace elements in pyrite, 

pyrrhotite and chalcopyrite of different ores: Economic 

Geology, v. 56, pp. 467-487.

Hem, J.D., 1972, Graphical methods for representing form and 

stability of aqueous metal ions: Chemical Geology, v. 9, 

p. 119-132.

Hubert, A.E., and Lakin, H.B., 1973, Atomic absorption 

determination of thallium and indium in geological 

materials, m Geochemical Exploration 1972: International 

Geochemistry Symposium, 4th, London, 1972, 

Proceedings, p. 383-387.

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shale and process waters Effect on shale leachate 

composition, m Fifteenth Oil Shale Symposium 

proceedings: Golden, Colorado School of Mines, 

p. 505-515.

Klieve, J.R., Rawlings, G.D., and Hoeflein, J.R., 1981, Assessment 

of oil shale retort waste water treatment and control 

technology: Monsanto Research Corporation, National 

Technical Information Service, PB 81-187288.

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filter used in the deuterium arc background corrector: 

Atomic Spectroscopy, v. 2, p. 134.

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marshes Sedimentary ion diffusion, sulfate reduction, and 

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p. 1381-1393.

Lynch, J.J., 1971, The determination of copper, nickel, and cobalt 

in rocks by atomic absorption spectrometry using cold 

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Metallurgy Special Volume, v. 11, p. 313-314.

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Golden, Colorado School of Mines, p. 113-123.

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Trace Elements in Pyrites D17
Pabst, A., 1971, Pyrite of unusual habit simulating twinning from 

the Green River Formation of Wyoming: American 

Mineralogist, v. 56, 133-145.

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p. 516-528.

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p. 636-652.

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p. 187-194.

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of the Mahogany Zone of the Green River Formation, 

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D18 Green River Formation, Wyoming, Utah, and Colorado
APPENDIX 1 PETROGRAPHIC DESCRIPTION OF 

SHALE SAMPLES IN WEST WILLIOW NO. 3 CORE

[Sample number indicates depth in core in meters. Core location 

shown in figure 1]

WW 377.3 Fine-grained (10-15 um) dolomite; 

pronounced organic-rich/organic-poor layering; organic 

material golden and amorphous; scattered larger clasts of 

organic-poor dolomite; abundant disseminated euhedral 

crystals of pyrite.

WW 367.8 Very fine grained (<5 um) dolomite. 

Scattered detrital quartz (<50 um) and untwinned feldspar 

grains; brown biotite and a colorless micaceous mineral; 

fragments of coaly material and pale-golden amorphous 

organic material. Pyrite blebs, made up of small euhedral 

crystals clumped together, are distributed irregularly 

through sample.

WW 415.7 Tuffaceous fragments (quartz and albitic 

feldspar) in fine-grained (<5 um) dolomite. Abundant 

grains of quartz and feldspar (<50 um); golden amorphous 

organic material concentrated in some layers; thin brown 

"stringers"; small disseminated euhedral pyrite crystals; 

abundant larger pyrite (as large as 100 um) associated with 

tuffaceous fragments.

WW 470.1 Fine-grained (<5 um) dolomite containing 

detrital quartz and feldspar grains (<50 um) and very little 

organic material. Pale-yellow coloration caused by cryptocrystalline phosphate (apatite). Bladed salt crystals pseudomorphed by dolomite, calcite, and pyrite; pyrite in irregular 

blebs.

WW 554.8 Fine-grained dawsonite-bearing oil shale; 

dawsonite partly replaced by calcite and dolomite, but 

numerous leached areas remain. Micaceous minerals 

abundant golden amorphous organic material distributed in 

distinct layers; pyrite as small blebs and scattered crystals 

and as fracture filling.

WW 557.7 No thin section available.

WW 608.6 Fine-grained (<5 um) dolomitic oil shale 

containing abundant partly leached dawsonite crystals 

(<100 um). Quartz, feldspar, and a micaceous mineral; 

uniform yellow amorphous organic material and rare 

stringers and brown blobs; small scattered crystals of pyrite 

(<10 um), no larger blebs or clumps of crystals.

WW 612.5 Fine-grained dolomitic oil shale 

containing abundant large (<50 um) dawsonite crystals and 

quartz grains. Golden amorphous organic material 

prominently layered; abundant finely disseminated pyrite 

grains and rare larger blebs.

WW 682.1 Fine-grained dolomitic oil shale (<5 um) 

containing large feldspar grains (<50 um). Small dawsonite 

crystals concentrated in layers several millimeters thick; 

distinct layering of golden amorphous material; pyrite finely 

disseminated throughout, locally higher concentrations in 

more organic-rich layers; disseminated pyrite crystals (<10 

um diameter).

WW 682.4 No thin section available.

APPENDIX 2 ANALYTICAL METHODS

Separation of Pyrite Samples from Shale

Pyrite was extracted from powdered samples by a 

series of acid dissolutions and density separations. To 

dissolve the carbonate fraction, a 30-50-g sample was 

heated in HCI until no reaction occurred on the addition of 

more acid (3-4 hrs). The residue was separated from the 

HCI by centrifugation. Next, the silicate minerals were 

dissolved by treating the residue with HF, and the sample 

was left overnight in HF. It was water washed and dried at 

50 °C. The hardened residue (kerogen) was removed by two 

successive sink/float separations in 75 percent bromoform 

and a third in 100 percent bromoform. The sink fraction, 

primarily pyrite, was washed in acetone and then water and 

finally fired at 50 °C. A few tenths of a gram of sulfides was 

generally obtained. Because a significant amount of pyrite 

encapsulated in kerogen was lost in the float fraction, the 

resulting pyrite separates do not represent the total pyrite 

content of the sample.

Preparation of Pyrite Samples for Analysis

Several chemical reagents have been used for 

dissolution of sulfide minerals, including aqua regia (Stanton, 1966), nitric and tartaric acids (Rubeska, 1968), 

hydrogen peroxide-ascorbic acid (Smirnova and others, 

1968; Lynch, 1971), water-bromine (Czamanske and 

Ingamells, 1970), a solution of bromine in carbon 

tetrachloride and nitric acid (Hubert and Lakin, 1973), and 

potassium chlorate-hydrochloric acid (Dolezal and others, 

1968). Recently, Chao and Sanzolone (1977) evaluated the 

effectiveness of these chemical treatments. They reported 

that dissolution of pyrite using either 4N nitric acid or 

potassium chlorate-hydrochloric acid followed by 4N nitric 

acid was very effective (86-100 percent dissolution by 

sulfur determination).

Because the pure pyrite-mineral separates are 

extremely precious, one sample dissolution procedure had 

to suffice for the 11 elements to be determined. The nitric 

acid dissolution procedure was selected because it provides 

a relatively uncomplicated sample solution matrix that can 

be used in both flame and flameless AA analysis.

The pyrite mineral separates, about 0.1 g, two 

additional pyrite samples (a South Dakota pyrite and RICO 

pyrite (Ward Scientific)), and an ASK Reference Material 

(sulfide ore, ASK-3) were prepared for spectrochemical 

analysis by treatment with 10 mL of 4N nitric acid in a 

50-mL Teflon beaker. This mixture was heated to boiling 

for 15 min on a hot plate. After cooling to ambient, the 

sample was diluted to volume, 50 mL, in linear 

polyethylene volumetric flasks. The sample solution was 

stored in a linear polyethylene bottle.

Trace Elements in Pyrites D19
Table A1. Atomic absorption spectrophotometry instrument parameters for pyrite analyses

Element

Arsenic

Cadmium

Cobalt

Copper

Iron

Manganese

Molybdenum

Nickel

Lead

Selenium

Zinc

Wavelength 

(nm)

193.7

228.8

240.7

324.8

248.8

279.5

313.3

231.1

283.3

196.0

213.9

Slit 

(nm)

0.7

0.7

0.2

0.7

0.2

0.2

0.7

0.2

0.7

0.7

0.7

Source

EDL/11W

HCL/8ma

HCL/30ma

HCL/15ma

HCL/30ma

HCL/25ma

HCL/29ma

HCL/25ma

EDL/10W

EDL/10W

HCL/15ma

Atomizer'

HGA

HGA

Flame/Air-C2H2

Flame/Air-C2H2

Flame/Air-C2H2

Flame/Air-C2H2

HGA

Flame/Air-C2H2

Flame/Air-C2H2

HGA

Flame/Air-C2H2

Background 

correction

Yes

Yes

Yes

No

No

No

Yes

Yes

Yes

Yes

Yes

Ipor HGA parameters see table A2.

Table A2. Atomic absorption spectrophotometry HGA parameters for pyrite analyses

Element

Arsenic 

Cadmium 

Molybdenum 

Selenium

Temperature (°C) 

Dry Char Atomize

100 

100 

100 

100

1,000 

250 

1,000 

800

2,700 

2,100 

2,800 

2,700

Sample 

volume (uL)

15 

10 

20 

50

Gas1 flow 

(mL/min)

40 

40 

40 

40

Graphite 

tube

Standard 

Standard 

Pyro 

Standard

'Argon.

Analytical Methods for Pyrite Analyses

Atomic absorption spectrophotometry (AA) is a 

commonly used instrumental method for the quantitative 

determination of inorganic elements in geological samples. 

In addition, the capabilities of AA such as high sensitivity, 

low detection limits, and flame and flameless capacity make 

it an ideal instrumental technique for trace-element 

determinations. Rubeska (1968) used AA to determine 

selected trace elements in various sulfide samples including 

pyrites. Teseneer and Brown (1979) used AA to determine 

Co, Au, Hg, and Ni in pyrites from the North Carolina Slate 

belt.

Flame AA measurements were performed using a 

Perkin-Elmer Model 5000 atomic absorption spectrophotometer equipped with a data station printer. Flameless AA 

measurements were performed using a Perkin-Elmer Model 

703 atomic absorption spectrophotometer equipped with a 

Model 400 HGA with AS-40 autosampler, PRS-10 printer, 

056 recorder, and D2 arc background corrector. The D2 arc 

background corrector was modified as described by Lindahl 

(1981). Instrument parameters for the AA analyses of these 

pyrite samples are given in table Al. Supplemental 

information for those elements determined by flameless AA 

(HGA) is given in table A2. No matrix modification was 

needed.

Results for the determinations of iron and the trace 

elements in the ASK Reference Material sulfide ore, 

ASK-3, are reported in table A3 along with results for 

duplicate samples of a South Dakota cluster pyrite and 

RICO pyrite. Instrument precision for both flame and 

flameless measurements was better than 2 percent relative, 

and method precision for duplicate samples, including 

sampling, sample homogeneity, sample preparation, and 

instrumental imprecisions, is from better than 1 percent 

relative to 18 percent. Typically, the agreement between 

duplicate sample analyses for these trace elements is 5-13 

percent. The method precision for each element, based on 

duplicate analyses of the three samples, is Fe (0-2 percent), 

As (3-5 percent), Cd (0-13 percent), Co (0-6 percent), Cu 

(0-12 percent), Mn (0-18 percent), Mo (5 percent), Ni 

(0-11 percent), Pb (5 percent), and Zn (0-7 percent). The 

accuracy of the data was determined by comparing our 

results with the recommended results for the ASK 

Reference Material ASK-3. The bias and accuracy, 

expressed at percent relative error, for the elements 

determined is Fe (-10 percent), As (+5 percent), Cd (-3 

percent), Co (-2 percent), Cu (-16 percent), Mn (-7 

percent), Mo (-3 percent), Ni (-32 percent), Pb (-22 

percent), and Zn (-8 percent). Values for bias and accuracy 

could not be determined for selenium. Values determined

D20 Green River Formation, Wyoming, Utah, and Colorado
CD>i^o * * i-i « " a-hl £3 ^&fl> CD

s ?l 8 11 s i I S g 1 - §. & 3

!?ftHthHi»!i?

« i« $ >5 sli cr- cr * »" tj " CoO>-''~a

h[?iillUl*ifg£ ;gi£ ? . S s M 8* s 2. ^ a Q »£8 s.g-S-^? 3

o 

n 

p

>

i

-<

Table A3. Results for determinations of iron and trace elements in duplicate samples of ASK Reference Material (sulfide or

South Dakota cluster pyrite (SD)

[EPR, concentration measured in this study; Lit., published concentration. Analytical reference materials are described in text. All values excep

Blank in column indicates no data available]

Fe As

Sample EPR Lit EPR

ASK-3 32.2 35.6 650

ASK-3 32.1 35.6 670

RICO 42.3 46.63 <10

RICO 423 46.6 <10

SD 353 3,140

SD 35.8 3310

1Not detected < 20 ppm.

Not determined.

Lit

630

630

Cd

EPR Lit

290 310

290 310

0.16

0.14

6.3

63

Co

EPR Lit

84 86

84 86

10

10

16

15

Cu

EPR

1,250

1,110

14

14

961

948

Lit

1,400

1,400

125

111

Mn Mo

EPR

242

241

18

15

64

61

Lit EPR Lit

260 40 40

260 38 40

ND

ND

ND

ND

Ni 

EPR

25

25

17

19

14

14

Lit 

37 

37 

^Theoretical percent Fe based on FeS^.
Table A4. Inductively coupled plasma emission 

spectrometer operating parameters

Table A5. Wavelengths, detection limits, and'maximum 

concentrations of linear ranges for inductively coupled 

______________________________ plasma emission spectrometry analyses 

Instrument...........................Jarrell-Ash Model 1155V vacuum __________________________________

spectrometer Concentration 

Nebulizer.............................MAK(Herritt-Gordon Mines Element Wavelength (A) Limit (mg/L) (mg/L)

Ltd., Ft. Saskatchewan, Alberta, - 

Canada T8L2P2) Cobalt 228<U56 °-05 ^ 

Sample uptake rate............... 1.9 mL/min Co^r 324^° .05 500

Argon flow rates Manganese 2576.104 .005 500 

Sample ...........................52 L/min Molybdenum 2020.300 .01 100

Coolant.......................... 14.5 IVmin Nickel 2316'037 '°5 10°

Plasma ..........................0.2 L/min a"c 2138'560 -Q1_______50°

Forward power..................... 1.05 kW

Reflected power...................<5 W

Plasma viewing height......... 16 mm above top of load coil

Exposure time...................... 10 sec_______________

Table A6. Atomic absorption spectrophotometric analysis operating parameters for determination of arsenic

Spectrophotometer...............Perkin-Elmer Model 4000 (PRS-10 printer, 056 recorder, and D2 arc background corrector

Autosampler........................ Perkin-Ebner AS-40

Atomization source..............Perkin-Ebner heated graphite atomizer with Model 400 programmer

Wavelength.......................... 193.7 nm

Detection limit.................... 1 ng/L

Light source.........................EDL, 8 W, background corrected

Graphite tube....................... Standard

Sample aliquot.....................20 ^L

Purge gas............................. Argon_______________________________________________

_______________________________Programmer keyboard entries________________________________________________________

Temp 

Step number 

CC) 125 

1234

1,100 2,700 100

Ramp(s) 13 25 1 5

Hold(s) 30 30 5 0

Read

Record X X

Gas flow (mL/min)_______*____________*____________50____________*_________________

*Purge gas flow rate controlled by internal pressure regulator.

D22 Green River Formation, Wyoming, Utah, and Colorado
Table A7. Atomic absorption spectrophotometric analysis operating parameters for determination of cadmium

Spectrophotometer ........

Autosampler .................

Wavelength...................

Detection limit.......... . .

Light source..................

Graphite tube................

Sample aliquot..............

Purge gas ......................

....... Perkin-Elmer Model 4000 (PRS-10 printer, 05

....... Perkin -Elmer Model AS 40

....... Perkin-Elmer heated graphite atomizer with M

....... 228.8 nm

. ...O.lug/L

....... HCL, 8 ma, background corrected

.......Standard

.......20 nL

.......Argon

6 recorder, and D2 arc background corrector 

odel 400 programmer

Programmer keyboard entries

Step number 

Temp CO 

Ramp (s) 

Hold (s) 

Read 

Record 

Gas flow (mL/min)

1 2 

125 350 

13 10 

30 60

* *

3 

1,900 

1 

5

X 

X

50

4 

2,500

1 

5

*

* Purge gas flow rate controlled by internal pressure regulator.

Table A8. Atomic absorption spectrophotometric analysis operating parameters for determination of lead

Spectrophotometer ........

Autosampler .................

Wavelength...................

Detection limit..............

Light source..................

Graphite tube ................

Sample aliquot..............

Purge gas ......................

....... Perkin-Elme

....... Perkin-Elme

....... Perkin-Elme

.......283.3

.......lug/L

.......EDL, 10W,

.......Standard

....... 20 uL

.......Argon

;r Model 4000 (PRS-10 printer, 056 recorder, and D2 arc background corrector 

T Model AS-40 

T heated graphite atomizer with Model 400 Programmer

background corrected

Programmer keyboard entries

Step number 

Temp CC) 

Ramp (s) 

Hold (s) 

Read 

Record 

Gas flow (mL/min)

1 

125 

13 

30

*

2 

500 

10 

60

*

3

2,700 

1 

5

X 

X

50

4 

100 

5 

0

*

*Purge gas flow rate controlled by internal pressure regulator.

Table A9. Atomic absorption spectrophotometric analysis operating parameters for determination of selenium

Spectrophotometer ...............Perkin-Elmer Model 5000 (PRS-10 printer, 056 recorder, and D2 arc background corrector

Hydride generator ................Perkin-Elmer Model MHS-10

Recorder.............................. Perkin-Elmer Model MHS-10

Wavelength.......................... 196.0 nm

Detection limit..................... 1 ug/L

Light source.........................EDL, 6 W, not background corrected

Sample aliquot..................... 10 uL

Flame...................................Air/acetylene (lean)

Reductant.............................3 percent NaBrLj. in 1 percent NaOH

Trace Elements in Pyrites D23

Chapter E

An Experimental Study of Goethite 

Sulfidization Relationships to the 

Diagenesis of Iron and Sulfur

By MARK R. STANTON and MARTIN B. GOLDHABER

U.S. GEOLOGICAL SURVEY BULLETIN 1973

CEOCHEMICAL, BIOCEOCHEMICAL, AND SEDIMENTOLOGICAL STUDIES OF THE 

GREEN RIVER FORMATION, WYOMING, UTAH, AND COLORADO
CONTENTS

Abstract El 

Introduction El 

Methods E2 

Results and discussion E5

Low pH (<7.0) results E5

Goethite E5

X-ray diffraction E5

Saturation state E6

Pyrite formation rates E6

Pyrite morphology and texture E6

Development of pyrite morphology E7

Spheres produced at low H2S concentrations E10

Sphere diameters E10

High pH results Ell 

Relationships to iron and sulfur diagenesis E12

Rates of pyrite formation E14

Pyrite texture E14

Other reaction products E15

Absence of other sulfide reaction products E15 

Conclusions E18 

Summary E18 

References cited E19

FIGURES

1. Diagram showing analytical method for determination of iron and sulfur

products from reaction of goethite and H2S E3 

2-3. Graphs showing:

2. Time versus concentration for iron and sulfur species E5

3. FeS saturation curve E6 

4 5. Scanning electron microscope photomicrographs illustrating:

4. Solids formed at low pH E7

5. Pyrites formed at low H2S(g) concentration E12

6. Histogram of modal sphere diameters E13

7. Graph showing modal sphere diameter versus reaction time at two ionic 

strengths E13

8. Scanning electron microscope photomicrographs illustrating decrease in pyrite 

sphere size with increasing ionic strength E15

9. Graph showing modal sphere diameter versus ionic strength E18

TABLES

1. Initial reaction conditions, rates of pyrite formation, reaction products, and pyrite 

morphology E3

2. Fe2+:S2~ mole ratios for early formed samples showing composition of FeS 

solids E4

3. Modal sphere diameter and size range for pyrite spheres formed in experiments 7 

and 12 E14

Contents
An Experimental Study of Goethite Sulfidization 

Relationships to the Diagenesis of Iron and Sulfur

By Mark R. Stanton and Martin B. Goldhaber

Abstract

The geochemical characteristics of iron sulfide minerals 

have been used to examine iron and sulfur diagenesis of the 

Eocene Green River Formation. To better understand the origin 

of iron sulfide minerals, we experimentally studied formation 

of these minerals under conditions that may have existed 

during deposition of the Green River Formation. Our study 

determined the types of products generated during goethite 

sulfidization and the rates of pyrite formation when goethite 

was converted to pyrite.

Goethite (oc-FeOOH) was reacted with aqueous 

hydrogen sulfide (H2S) over a range of pH, ionic strength, and 

total iron and total sulfide concentrations. Pyrite (FeS2), 

elemental sulfur (S°), and iron monosulfide (FeS) were the 

dominant solid products formed at pH less than or equal to 

7.0. Rough-textured spheres are the dominant morphology of 

pyrite; framboids and euhedral crystals of pyrite were not 

produced. At pH greater than 7.0, elemental sulfur and FeS 

formed, but pyrite was not detected; unreacted goethite 

remained as a solid phase. Other sulfide minerals present in 

some Green River sedimentary rocks, such as pyrrhotite and 

marcasite, were not detected.

Authigenic pyrite formed rapidly under conditions of low 

pH (<5.0). At neutral pH, the rate of pyrite formation was slow 

compared to that of acidic pH. Pyrite was not detected in 

experiments at pH values greater than 9.0. Formation of pyrite 

in alkaline solution requires either extended time of reaction or 

reaction pathways that operate only in complex natural 

settings. The formation of euhedral and framboidal pyrite in 

unconsolidated sediments likely requires processes or 

conditions that were not observed in our experiments.

INTRODUCTION

Authigenic iron sulfide minerals are present in 

virtually all subaqueous sediments in which organic matter 

and sulfate are present (Goldhaber and Kaplan, 1974). They

form at low temperature (<50 °C) and pressure (1 atm), 

mainly through the action of aqueous hydrogen sulfide 

(H2S(aq)) produced by bacterial sulfate reduction. This 

H2S(aq) reacts with iron-bearing minerals to produce iron 

"monosulfides" such as mackinawite (FeS0 9) or greigite 

(Fe3S4) (Berner, 1970). These iron monosulfide minerals 

(FeS) are converted to pyrite (FeS2) by reaction with 

zerovalent sulfur (S°): FeS+ S°-»FeS2. Pyrite may also 

directly precipitate from a ferrous polysulfide solution 

without the formation of an intermediate FeS precursor 

(Howarth, 1978); however, direct precipitation may be 

inhibited below 100 °C (Schoonen and Bames, 1988). The 

kinetics and mechanisms of pyrite formation have been 

experimentally studied by many researchers; for example, 

Berner (1964a, 1969, 1970), Rickard (1969a, b), Roberts 

and others (1969), Hallberg (1972), Pyzik and Sommer 

(1981), Murowchick and Bames (1987), and Schoonen and 

Bames (1988). Nevertheless, the mechanistic details of 

FeS2 formation are not clearly understood, and the origin of 

iron sulfide minerals probably varies from one geochemical 

environment to another. The "reactivity" of various iron 

sources toward sulfide has, until recently, received little 

attention (Berner, 1964a; Canfield and Berner, 1987; Canfield, 1989). The transitional role of iron monosulfides 

during their conversion to iron disulfides is also poorly 

understood (Berner, 1972; Sweeney and Kaplan, 1973; 

Rickard, 1975). In particular, insufficient experimental 

information is available concerning the overall sequence of 

reactions during iron sulfidization.

An aspect of our ongoing experimental studies of 

pyrite formation, applicable to the Eocene Green River 

Formation, is to identify the nature of iron sulfidization in 

environments such as alkaline saline lakes. Sediments in 

these lakes contain sulfate, H2S(aq), and iron, yet only minor 

amounts of authigenic FeS and FeS2 are produced (Turtle, 

1988). The paucity of sulfide minerals may reflect a lack of

Experimental Study of Goethite Sulfidization El
reactive iron sources (Canfield, 1989) or kinetic hindrance 

due to high pH (Berner, 1964a, 1969). Because the factors 

governing the formation of iron sulfide minerals in 

environments of this type are not well documented, a 

companion purpose of this study was to examine goethite 

sulfidization and iron sulfide formation at high pH (7-10).

Textural and morphological variations of pyrite may 

be indicative of distinct pathways of pyrite formation in 

unconsolidated sediments, but the geochemical controls on 

these differences are not well known (Raiswell, 1982). The 

two most commonly observed pyrite textures in sediments 

are well-developed single euhedral crystals (mainly cubes) 

and framboids or polyframboids (Sweeney and Kaplan, 

1973; Kribek, 1975; Raiswell, 1982). Framboids are 

aggregates of euhedral crystallites and are from 1 to more 

than 100 (am in diameter (Kribek, 1975); aggregates of 

several framboids result in the "polyframboidal" texture 

described by Love and Amstutz (1966). Euhedral pyrite 

may result from direct precipitation of pyrite or conversion 

of an FeS precursor phase. Framboidal pyrite textures, on 

the other hand, are believed to be produced only from the 

conversion of an FeS precursor (Raiswell, 1982). This 

report examines pathways and products qualitatively; a 

detailed chemical and kinetic treatment of all experimental 

parameters is in progress.

Acknowledgments. This work was part of the U.S. 

Geological Survey Evolution of Sedimentary Basins 

Program. We extend our appreciation to N. Fishman, J. 

Fitzpatrick, and G. Breit for critical review of this 

manuscript.

METHODS

Synthetic a-FeOOH was prepared by the method of 

Atkinson and others (1968), which employs the alkaline 

hydrolysis of Fe(NO3)3 H2O in KOH to form the iron 

oxyhydroxide precipitate. After aging at 90-100 °C for 10 

hours, the precipitate was rinsed with distilled water until 

dissolved nitrate was less than 1 ppm, and then it was stored 

under N2 until needed. Flame atomic absorption spectrophotometry analysis (AA) for dissolved iron from a 0.05- 

|4,m filtration of this rinsed suspension was below the 

detection limit of 3.5|4,M/L Fe. X-ray powder diffraction 

(XRD) of the iron oxyhydroxide detected only a-FeOOH. 

Scanning electron microscope (SEM) observations of the 

(X-FeOOH solids revealed that the solids are acicular blades 

approximately 1 ^lm long and 0.2 jam wide.

Two types of experiments were performed during this 

study. The first type consisted of "free-drift" inorganic 

reactions in which the pH was allowed to change as the 

experiment proceeded. In the free-drift reactions, the initial 

pH (pH(init)) was less than 5.0 (table 1). The pH change did 

not exceed ±1.3 and generally remained within ±0.8 units of 

the initial pH. The second type of experiment consisted of 

inorganic reactions that were maintained at constant pH

(±0.2) by means of a pH-stat. In this second group, the pH 

ranged from 5.6 to 10.0 (table 1).

The experimental conditions are representative of 

natural environments such as estuaries or saline lakes that 

have ionic strengths between 0.01 M (freshwater) and 0.7 M 

(seawater) (Garrels and Christ, 1965). The ionic strength (I) 

was 0.1 M in all but five experiments; in those, ionic 

strengths were from 0.3 to 2.0 M (table 1). The amount of 

total added iron (1.3xlO~3-5.7xlO~3 M/L) was two to three 

orders of magnitude greater than total paniculate iron in 

oxidizing waters (Kester and others, 1975) or dissolved iron 

in porewater of reducing sediments (Morse and others, 

1987). Values for total dissolved sulfide in sedimentary 

environments vary but are generally less than 5.0xlO~3 M/L 

(Pyzik, 1976). Calculated total dissolved sulfide used in 

these experiments, the sum of H2S(aq), HS~(aq), and S2~, was 

from 8.3xlO~5 to 8.3xlO~3 M/L.

All reagents and solutions were prepared from 

reagent-grade chemicals and CO2-free, O2-free distilled 

water. Experiments were carried out in a gas-tight, 1.5-L 

glass reaction vessel. After deoxygenation of the reaction 

water with ultra high purity N2, KC1 was added to adjust the 

ionic strength. N2 was then flowed through for an additional 

8-16 hours. A mixture of H2S and N2 gas was then bubbled 

through the mixture for the duration of the experiment. The 

pressure and composition of the H2S-N2 gas mixture were 

precisely maintained through the use of a mass flow 

controller system (MKS Instruments). The goethite 

suspension was rapidly stirred and deoxygenated with N2 

for a minimum of 24 hours, then added to the reaction vessel 

after stabilization of the H2S flow and solution pH. The 

addition of goethite to the solution was taken as the starting 

point for the reaction (time=0). For the constant-pH 

experiments, pH was maintained at a set value by the 

addition of HC1 or KOH (depending upon the initial 

experimental conditions) using a pH-stat. Stirring rates were 

kept high in all experiments to minimize surface-controlled 

reactions.

Analyses for the various forms of iron and sulfur were 

performed as shown in figure 1 and described next. This 

series of chemical extractions and gravimetric techniques 

was based on the fact that iron monosulfides are soluble in 

6N HC1, whereas pyrite is insoluble in this acid (Berner, 

1964a). The pH of each sample was measured using a 

double junction reference/AgCl glass electrode pair. 

Generally, one sample per day was taken from the bulk 

mixture over a 1-2 week period. Samples were withdrawn 

into a N2-purged 50-mL syringe that formed a gas-tight seal 

with the sampling port. The gas flow and stirring were not 

interrupted for the sampling process. Sample treatment was 

carried out in a glove box that was continuously purged with 

nitrogen.

The concentration of AVS, acid-volatile sulfides 

(FeS), was determined by filtering a raw sample through a 

0.05-|4,m membrane filter and washing excess sulfide from

E2 Green River Formation, Wyoming, Utah, and Colorado
Table 1 . Initial reaction conditions, rates of pyrite formation, reaction products, and pyrite morphology

Expt. 

no. 1

5(F)

6(F)

7(F)

8(F)

9(F)

10<F)

11(F)

12(F)

13(F)

15(S)

16(S)

18(S)

19(S)

28(F)

29(F)

30(F)

31(S)

32(S)

33(S)

pH(init)

4.62

4.11

4.62

4.75

4.91

4.45

4.99

4.69

4.42

6.01

7.10

9.50

7.00

4.92

4.61

4.59

5.69

9.00

10.0

H2S(g) 

(vol. percent)

10

10

10

10

10

10

10

10

10

0.1

0.1

10

1.0

1.0

0.1

0.1

0.1

10

1.0

Ionic 

strength (M)

0.1

0.1

0.1

0.1

0.1

03

0.7

2.0

0.3

0.1

0.1

0.3

0.1

0.1

0.1

0.1

0.1

0.1

0.1

Fe(tot) Rate FeS2 Initial 

(mM) formation (M/min)2 solids3

4.75

4.75

4.75

2.55

132

2.55

2.55

2.55

2.55

2.55

2.55

4.33

3.27

3.27

3.27

1.70

5.70

1.70

1.70

1.8 xlO'7

5.0xlO-7

6.0xlO-7

1.2 xlO"8

2.2 xlO'7

6.6 xlO"8

6.0 xlO"8

1.0 xlO"8

6.1 xlO-8

4.4X10-8

dnf

dnf

2.6 xlO"9

23 xlO"8

3.2 xlO"8

1.3 xlO-8

dnf

dnf

dnf

M,S

M,S

M,S

M,S

M,S

M,S

M,S

M,S

M,S

M,G

M,G,S

M,G,GR

M,G,S

M,G,S

M,G,S

M,G,S

M,G,S

M,G,S,GR

M,G,S,GR

Final FeS2 

solids3 morphology4

P,S

P,S

P,S

P,S

P,S

P,S

P,S

P,S

P,S

P,G,M

G,M

S,G,GR,M

P,S,M

P,M

P,S,M

P,S,G,M

G,S,M

S,M

S,GR,M

Sph

Sph

Sph

Sph

Sph

Sph

Sph

Sph

Sph

Sph

na

na

Sph

Sph

Sph

Sph

na

na

na

Sphere 

sizes5

nd

4.7

4.4

3.2

nd

1.9

0.7

0.5

2.1

1.5

na

na

nd

1.4

13

1.8

na

na

na

1Free-drift (pH unconstrained);S, constant pH (±0.2 units).

2Overall rate of amount formed from when pyrite was first detected to end of experiment. Rate values are subject to experimental error of 

±20 percent, dnf, did not form.

3M, mackinawite or FeS(am); S, elemental sulfur; G, goethite; P, pyrite; OR, green rust. Initial products are those present in first sample; 

final products are those present in last sample. No quantities are given because amount and type of product depend on time of sampling of 

reaction mixture, which differed for each experiment. Generally, initial products were present within first 24 hours.

^ph, spherical; na, not applicable. 

Sphere size given is modal diameter (um) of final sample, na, not applicable; nd, not determined.

Raw sample

Filter

Solids + 6N HCI

Iron analysis = Aqueous iron (Fe2+(aq)) 

Ag2S = Sulfur in FeS (S2')

^Solution 

I

Iron analysis = Iron in FeS (Fe2*

Filter 

1

Solids 

4.

Acetone extraction 

1 I 

Solids Solution 

J, 4,

HCI/HNO3 Elemental Sulfur (S°) 

I 1 

Solution Sulfate precipitated as BaSO4

J, I 

Total iron = Pyrite iron + BaSC>4 = Sulfur in pyrite

Goethite iron Calculated equivalent Fe = Pyrite iron 

j, 4, 

Total iron - Pyrite iron = Goethite iron

Figure 1. Analytical method for determination of iron and 

sulfur products from reaction of goethite and H2S.

the solids using deoxygenated distilled water (fig. 1). Next, 

H2S(g) evolved from a 25 °C, 6N HCI leach of the solids was 

precipitated as Ag2S. This Ag2S was filtered, dried, and 

weighed to obtain the total sulfide present in the iron 

monosulfide. The iron content of this 6N HCI filtrate was 

determined using AA and represents iron in the monosulfide. The concentration of aqueous iron in the original 

filtrate, Fe2+(aq), also was determined using AA and is iron 

released into solution during the Fe(III) reduction-goethite 

dissolution steps (Pyzik and Sommer, 1981). The accuracy 

of the AA method for all iron analyses is ±10 percent. The 

crystalline goethite was only slowly soluble in 6N HCI and 

was not dissolved in the short time necessary for this sample 

treatment step (<10 minutes). (Fe2+,Fe3+)-oxyhydroxides 

formed at high pH were also removed in this step.

Total elemental (zerovalent) sulfur (S°) was 

determined by cyanolysis in acetone (Bartlett and Skoog, 

1954). An aliquot of the dried, 6N HCl-treated solids was

Experimental Study of Goethite Sulfidization E3
refluxed in acetone to extract the solid S°. The solution was 

separated by filtration, then NaCN and FeCl3 were added to 

form the (Fe3+SCN)2+ complex. The concentration of 

(Fe3+SCN)2+ was determined spectrophotometrically at 465 

nm. Based on iron sulfide mineral mixtures (FeS+FeS2) 

spiked with known amounts of elemental sulfur, this 

method was effective in extracting 95 percent of the solid S° 

from the products. Sulfur was not extracted from the iron 

sulfide phases.

The sulfur extracted in this step represents total 

zerovalent sulfur of unspecified chain length or structure. 

This sulfur formed by the redox reaction of Fe(in) and 

sulfide (2Fe3++H2S(aq) (or HS~(aq))^S°+2Fe2++2H+). 

Sulfur species such as S8 the stable form of elemental sulfur 

at room temperature, may subsequently form from this 

zerovalent sulfur. Solid dipyramidal crystals of sulfur seen 

in microscopic examination are S8 and contain all zerovalent sulfur atoms. The distribution of S°-bearing species in 

the reactant solution was not determined.

The amounts of pyrite and goethite were estimated by 

oxidation in aqua regia of the solids that remained after 

extraction of S°. Previous extractions had removed all but 

pyrite and goethite. FeS and (Fe2+,Fe3+)-oxyhydroxides 

present in high pH samples were removed in the 6N HC1 

step.

Aqua regia oxidized the sulfide sulfur in pyrite to 

sulfate (SO42~), which was then precipitated as BaSO4. The 

BaSO4 solids were separated by filtration, dried, and 

weighed. Iron originally in pyrite was calculated as the 

difference between sulfur in barite and iron in the filtrate. 

Residual iron is attributed to goethite (fig. 1).

X-ray diffraction was used after each chemical 

treatment for qualitative mineral identification and to 

determine the effectiveness of each successive extraction. A 

Philips XRG-3000 diffractometer equipped with a graphite 

focusing monochromator using Ni-filtered Cu Ka radiation 

at 30 kV and 20 mA was used for all diffraction work. 

Samples were scanned at 2°/26 per minute from 10° to 70° 

26. Samples not leached with 6N HC1 (generally duplicate 

raw samples) were scanned after mixing the filtered sample 

with high-vacuum lubricant in an attempt to inhibit 

oxidation; unfortunately, this technique did not always 

prevent oxidation of FeS or produce useful X-ray results. 

X-ray diffraction was able to detect mineral species at 

concentrations of 5 volume percent or greater.

Spatial and textural relationships of the solid 

products, based on SEM observations, werfe used to infer 

some possible pathways for the transformation of goethite 

to pyrite. Polished and unpolished mounts of the sample 

were prepared from runs in which pyrite was detected by 

XRD. Unpolished gold-coated mounts were used to view 

the solids to determine morphology, size, and texture. 

Carbon-coated polished grain mounts were used to obtain a 

two-dimensional surface for iron and sulfur analysis 

Goldstein and others, 1981). Because FeS and

(Fe2+,Fe3+)x(OH)x solids are highly susceptible to oxidation 

in air, few of these solids were examined. Diameters of the 

product FeS2 spheres, measured to 0.1 um, were determined 

visually from SEM photomicrographs. SEM energydispersive spectroscopy (EDS) analyses of the polished 

spheres are within ±5 weight percent of the iron and sulfur 

values obtained from a pyrite standard. Polished mounts 

were also examined under reflected light to look for 

mineralogical features not detected by SEM.

Due to poor diffraction characteristics of the 

amorphous, early formed FeS, Fe2+:S2~ mole ratios obtained 

from AA and Ag2S gravimetry in the AVS analysis were 

used to characterize the mineralogy of the early FeS 

fraction. The Fe:S ratios for crystalline iron sulfide minerals 

are given in table 2.

The saturation state of the solution and distributions 

of aqueous and mineral species were modeled using the 

computer program PHREEQE (Parkhurst and others, 1980). 

This equilibrium model uses Debye-Huckel activity 

coefficients to calculate activities for aqueous iron and 

sulfur species for all ionic strengths. The log of the ratio of 

the ion activity product (LAP) to the solubility product (Ksp) 

is used to determine the saturation state of the solution with

Table 2. Fe2+ :S2~ mole ratios for early formed samples 

(<24 hours) showing composition of FeS solids 

[Fe2+ is from 6N HC1 leach filtrate. S2' is sulfide precipitated as 

Ag,S from 6N HC1 leach, nd, not determined. Fe:S compositions of

crystalline iron sulfide minerals are shown below. FeS composition 

based on these ratios is primarily that of mackinawite (Fe:S = l.l)]

Experiment no.

5

6

7

8

9

10

11

12

15

29

30

Time (hours)

2

1

2

22

21

5

2

2

3

2

2

Fe2+ :S2' ratio

1.10

1.11

1.09

1.00

1.34

1.15

1.24

1.15

0.95

0.%

0.89

Mean value is 1.09 (s±0.12)

Mineral name

Mackinawite1

Pyrite

Marcasite

Greigite

Pyrrhotite2

Formula

FeSl-x

FeS2

FeS2

Fe3S4

Fel-xS

Fe:S mole ratio

1.1

0.5

0.5

0.75

0.9

1Weight percentages and Fe:S ratios based on FeSQ 9. 

2Weight percentages and Fe:S ratios based on Fe0 9S.

E4 Green River Formation, Wyoming, Utah, and Colorado
respect to mineral phases in the system. Negative values of 

this ratio (log(IAP/Ksp)) indicate solution undersaturation 

and probable mineral dissolution; positive values point to 

solution saturation and thus mineral precipitation. Values of 

zero indicate equilibrium or metastability of the solution 

and solid phases.

RESULTS AND DISCUSSION

Low pH (<7.0) Results

Experimental conditions and results are summarized 

in table 1. Because pyrite was not detected at high pH 

(>7.0), most of the discussion on pyrite formation focuses 

on the experiments from pH 4 to 7. Unless specified, low 

pH refers to pH values less than or equal to 7.0, whereas 

high pH values are those greater than 7.0. This separation 

was chosen because pyrite was not detected at pH values 

greater than 7.0. Similarly, low H2S refers to experiments 

run at 1.0 or 0.1 H2S(g) volume percent, whereas high H2S 

refers only to those experiments run at 10 percent H2S(g). 

The high pH results are discussed by comparing the reaction 

products to the pyrite-yielding sequence of low pH.

All the low pH experiments displayed the same 

reaction sequence. First, rapid reduction of Fe3+ and 

complete dissolution of oc-FeOOH occurred within 2-8 

hours at pH(init) of 4.11^.99 to form FeS, S°, and Fe2+(aq). 

Dissolution was followed by a period of slow reaction of 

FeS, S°, and Fe2+(aq) to form FeS2. Finally, pyrite formed 

rapidly once undersaturation with respect to FeS was 

attained. The rate of this overall process varied with 

conditions but was most rapid at pH of 4-5 and 10 percent 

H2S(g) (table 1). Figure 2 illustrates the reaction sequence 

and changes in the concentrations of iron and sulfur species 

with time.

oc-FeOOH dissolved rapidly (0-2 hours) at pH of less 

than 5.0 and 10 percent H2S(g). The resulting solution was 

calculated to be saturated with mackinawite, amorphous 

FeS, greigite, marcasite, and pyrite. The major iron species 

after goethite dissolved was Fe(H) in the aqueous and FeS 

phases. More than 90 percent of the iron added as goethite 

was incorporated into the aqueous and FeS phases under 

these conditions. The initial, rapid redox reaction of goethite 

and H2S is consistent with the results of Pyzik and Sommer 

(1981), who reported that Fe(IQ) was reduced by bisulfide 

(HS~(aq)) prior to goethite dissolution and release of Fe(H). 

The reduced iron subsequently reacted with H2S(aq) or 

HS"(aq) to form FeS.

Goethite

oc-FeOOH was detected as a residual component of 

the solids under almost all experimental conditions. In

0.005

¥ 0.004

tr hz

HI 

O

z 

o 

o

0.003 -

0.002 -

0.001 -I

1 00 

TIME, IN HOURS

200

Figure 2. Time versus concentration for iron and sulfur 

species, experiment 7. Results are typical of a reaction 

sequence at low pH (<7.0). Goethite is rapidly consumed by 

reaction with H2S, and FeS, S°, and fe+2iaq) are formed. The 

part of pyrite curve labeled "a" represents slow formation of 

pyrite while solution is saturated with FeS. The part labeled "b" 

represents rapid formation of pyrite once solution becomes 

undersaturated with FeS.

solids from the earliest stages of low pH reactions, ccFeOOH was detected visually after treatment with 6N HC1. 

In low H2S runs, it was detected by XRD in unleached 

samples or was visible as an unreacted solid (table 1). And, 

in most high pH reactions, goethite was a major, unreacted 

constituent. X-ray diffraction studies suggest that this goethite is not a poorly crystalline oxidation or alteration 

product of reaction or sample treatment but rather the 

material added at the start of the experiment. Thus, 

oc-FeOOH was consumed rapidly at low pH but remained as 

a major phase for the duration of the run at high pH.

X-Ray Diffraction

In the first 8 hours of low pH experiments, XRD 

patterns of unleached samples were characteristic of an 

amorphous substance (wide, poorly defined peaks). As the 

reactions progressed beyond approximately the 8-hour 

point, mackinawite peaks became evident. Detection of FeS 

in unleached samples by XRD at pH greater than 9.0 was 

difficult because of the presence of ferrous-ferric oxyhydroxides. Goethite was present in some early samples from 

low pH experiments (0-8 hours) and in all high pH samples. 

Elemental sulfur was detected as a product of every 

experiment in amounts that varied depending on the 

conditions and extent of reaction. Similarly, pyrite was first 

detected at different elapsed times depending on conditions, 

but once detected it became more abundant and more

Experimental Study of Goethite Sulfidization E5
crystalline (relative to the earlier samples). Diffraction 

patterns from some early samples contain pyrite peaks 

having greater half-widths (generally an indicator of poor 

crystallinity) than pyrite peaks from samples near the end of 

a run.

Mole ratios of Fe2+:S2~ are not given for the 

experiments run at pH 9.0 or higher (experiments 18, 32, 

and 33) because the FeS and iron oxyhydroxides ("green 

rust") that formed were both dissolved by 6N HC1. Thus, the 

filtrate contained total iron from two solids that could not be 

assigned to individual minerals. Additionally, FeS formed 

so slowly in these experiments that it was not useful to 

sample the reaction mixture until after 24 hours.

Fe2+:S2~ mole ratios from the earliest stages of the 

low pH experiments gave ratios of AVS iron to sulfide of 

between 0.8 and 1.3 (table 2). The mean of these ratios is 

1.09 (s±0.12). The ratio of Fe:S in mackinawite (FeS09) is 

1.1; in pyrite and marcasite (FeS2), it is 0.5. If only greigite 

were present, this ratio should have been 0.75 (3Fe:4S). 

Based on the Fe2+:S2~ ratios obtained by analysis and XRD 

data, mackinawite and (or) amorphous FeS are the dominant 

sulfide phase(s) produced early in these experiments. Ratios 

of Fe:S of less than 1.1 might represent several possible 

mixtures of mackinawite, pyrite, and greigite (Hallberg, 

1972), but the existence and proportions of such mineral 

groupings could not be resolved using the available data. 

High Fe:S mole ratios (>1.1) may be due to variations in the 

stoichiometry of the FeS or FeS(am); amorphous FeS may 

contain as little as 0.87 S per 1.0 Fe (Morse and others, 

1987), an Fe:S ratio of 1.15. Adsorption of sulfide or iron 

also niay have altered the Fe:S ratios (Rickard, 1969a; 

Pyzik, 1976).

Saturation State

Figure 3 depicts the change over time in the saturation 

state of the solution with respect to FeS for experiment 7. 

The sample at 68 hours is the last solution that is oversaturated with FeS; after this time, the rate of FeS2 formation 

increases. Samples at 93 and 115 hours indicate a 

transitional period of FeS metastability, and beyond 115 

hours the solutions are undersaturated with respect to FeS. 

Because the concentration of H2S(aq) and other parameters 

were constant, FeS undersaturation resulted from the 

decrease in iron concentration by its incorporation into 

pyrite. The solution remained saturated with pyrite, the only 

other iron mineral detected. In figure 2, the first slope 

increase of the pyrite formation line ("a") represents the 

continuous slow rate of production of FeS2 during the time 

when FeS and S° are dominant in the system. The part of the 

line labeled "b" indicates the rapid formation of pyrite 

following the undersaturation of FeS.

<D 

U_

0-*

-1 -

CL

< -2

O)

o -3-

1 00 

TIME, IN HOURS

200

Figure 3. FeS saturation curve, experiment 7. Values greater 

than zero indicate FeS oversaturation; values less than zero 

indicate FeS undersaturation. A change toward FeS undersaturation begins at approximately 68 hours. Note rapid FeS 

dissolution coincident with rapid FeS2 formation shown in 

figure 2.

Pyrite Formation Rates

In experiments run at or below pH 7.0, the overall rate 

of pyrite formation that is, the slope of FeS2 versus 

time varies from 6.0x10~7 M/min at pH(init) of 4.11 to 

2.6xlO~9 M/min at pH(init) of 7.0 (table 1). Pyrite formed 

most rapidly in experiments 5, 6, and 7, which had 10 

percent H2S(g) at low initial pH. The slowest rates of pyrite 

formation were in experiments run at pH(init) of almost 7.0, 

or at H2S(g) of 1.0 or 0.1 volume percent. The rate of pyrite 

formation is undetectable at pH values greater than 7.0. The 

slower rates or lack of pyrite formation (for example, 

experiment 19) at high pH are probably a result of 

retardation of the goethite dissolution/FeS formation process that proceeded rapidly at low pH values. The overall 

reaction represented as FeS+S° »FeS2 does occur at 

alkaline pH but is very slow at low temperatures (Berner, 

1969; Rickard, 1975).

Pyrite Morphology and Texture

Pyrite produced in these experiments is mainly 

spherical, from less than 1 to slightly more than 5 microns 

in diameter. According to semiquantitative SEM-EDS 

analysis, pyrite spheres consist only of iron and sulfur in the 

proportions of one mole of iron to two moles of sulfur. 

Departures from this ratio were observed in nonspherical 

solids, in samples that were incompletely reacted (as in the

E6 Green River Formation, Wyoming, Utah, and Colorado
earliest stages of an experiment), or in samples that were not 

6N HC1 and acetone leached. Solids possessing this lFe:2S 

ratio produce XRD patterns characteristic of pyrite.

Development of Pyrite Morphology

Contrasting experimental conditions and extent of 

reaction resulted in systematic differences in pyrite 

morphology. Figure 4 is a sequence of SEM photomicrographs from experiment 7 depicting the formation and 

change of pyrite morphology and texture. This sequence is 

typical of all low pH experiments. Euhedral pyrite (that is, 

well-developed crystal faces) was not seen as a product 

under any conditions.

Except for the sample shown in figure 4A, the solids 

were leached with 6N HC1 and acetone before they were

examined. Based on SEM-EDS analysis, chemical 

extraction, and XRD, the solids shown in figures 4B-E are 

pyrite. The dipyramids in figure 4A (2 hours) are crystalline 

elemental sulfur; the smallest sulfur crystal is approximately 

2 (jrn in diameter and the largest is almost 22 |im (measured 

along the minor axis of the dipyramid).

The large solid in the upper left quadrant in figure 4A 

is iron rich but is not the original bladed goethite. Its 

massive morphology is common for iron-bearing solids 

early in most experiments. Three SEM-EDS analyses of 

this product gave a mean of 75.3 weight percent Fe and 24.6 

weight percent S. The mole ratio of Fe:S of 1.75 is much 

greater than the Fe:S ratio for known crystalline iron 

sulfides (table 2). Therefore, the exact composition of this 

substance is unknown, but it is probably a mixture of S8, 

poorly crystalline or oxidized iron monosulfide(s), and

Figure 4 (above and following pages). Scanning electron microscope photomicrographs of solids 

formed in experiment 7. Sequence illustrates development of pyrite texture and morphology with 

increasing time of reaction at low pH (4-5). A, Elemental sulfur dipyramids and massive 

"intermediate" iron-rich sulfur-bearing phase formed during goethite dissolution/FeS formation. Bar, 

10 urn; time, 2 hours. B, Spherical pyrite morphology typical of FeS2 formed early in reaction. Bar, 1 

urn; time, 22 hours. C, Pyrite microspheres showing platy surface texture of spheres. This spherical 

shape with rough-surfaced texture was the dominant form of pyrite. Bar,1 jim; time, 43 hours. D, 

Range of pyrite sphere sizes approximately halfway through the experiment. Bar, 1 jim; time, 68 

hours. E, Large pyrite spheres at the end of experiment 7. The rough surface and spherical shape are 

typical of spheres formed in experiments run at low pH values and high H2S (10 percent) 

concentration. Bar, 1 jim; time, 139 hours.

Experimental Study of Goethite Sulfidization E7
E8 Green River Formation, Wyoming, Utah, and Colorado
Experimental Study of Goethite Sulfidization E9
Sulfur forms other than S8 were not distinguished on the 

basis of SEM-EDS analysis. X-ray diffraction of the 6N 

massive or indistinguishable crystallites of oc-FeOOH. HC1- 

leached solids from this sample detected only goethite and 

sulfur.

Pyrite agglomerates that have incipient spherical 

morphology had formed by 22 hours (fig. 4B). At this point, 

pyrite was a minor component (<lxlO~5 M Fe as pyrite) of 

the total iron in the solids. The sphere size is about 1 ^im, 

smaller than the majority of the sulfur dipyramids shown in 

figure 4A. These spheres did not form directly from either 

the elemental sulfur dipyramids or the large, iron-rich 

"intermediate phase" shown in figure 4A, and a dissolution- 

/precipitation mechanism must have been operating at this 

time if the solids shown in figure 4A were being consumed.

X-ray diffraction and ultraviolet-visible spectrophotometry of each sample obtained during experiment 7 show 

that elemental sulfur was present while pyrite was being 

formed (fig. 2). Later samples were leached with acetone 

prior to SEM examination, so sulfur dipyramids were 

removed. Although solid sulfur was not seen in subsequent 

micrographs in this series, other methods demonstrate that 

acetone-extractable elemental sulfur was present.

Uniform pyrite spheres formed after 43 hours (fig. 

4C). These spheres are atypically large for this sample but 

exhibit in detail the platy rough surface texture 

characteristic of most spheres that formed by 48 hours. 

These rough-surfaced spheres are the major type of pyrite in 

all experiments in which pyrite formed. This spherical 

texture may have been inherited by continued growth on the 

smaller, spherical grains shown in figure 46.

Figure 4D shows the surface texture of the spheres 

after 68 hours. The spheres became rougher and more 

nodular as the experiment progressed. Figure 4E shows the 

texture typical of the pyrite by 139 hours. This series of 

photomicrographs shows that the spherical morphology of 

pyrite developed early in the reaction. Although an 

increasingly rough surface texture developed over time, the 

spherical shape persisted throughout the experiment. 

Additionally, the diameters of the spheres increased 

(discussed later).

Spheres Produced at Low H2S Concentrations

The concentration of H2S(aq) may affect sphere size 

and texture. Pyrite spheres formed in experiments run at low 

H2S(g) (for example, experiments 29 and 30, table 1) are 

texturally and morphologically similar to spheres formed at 

high I^S^ concentrations. However, development of the 

rough sphere texture was slower and modal sphere 

diameters (MD) smaller. Experiment 30 was run at low 

H2S(s) at an initial pH of 4.59 (table 1). The pyrite spheres 

at 126 hours (fig. 5A) have a smoother, less platy surface

than spheres formed in experiment 7 (compare with fig. 4E). 

Thus, for a similar amount of reaction time, the rough 

texture seen in experiment 7 had not yet developed.

Experiments 30 and 7 were similar except for the 

H2S(g) concentration but produced spheres of different 

diameters. The modal diameter of the spheres from 

experiment 30 (1.8 ^im, table 1) is smaller than in 

experiment 7 (4.4 ^im, table 1) for comparable reaction 

times.

Spheres formed in experiments at low H2S(g) acquire 

the rough texture later than in experiments at high H2S(g). 

Figure 5B shows a sample taken at 188 hours from 

experiment 29, which was also run at low H2S(g). The 

sample was treated with 6N HC1 but not acetone leached, 

and XRD indicated that only sulfur and pyrite were present. 

SEM-EDS analysis detected only sulfur in the large 

dipyramid-shaped solid, and analysis of one sphere in the 

field of view is 44.1 weight percent Fe and 55.9 weight 

percent S, similar to the expected values for pyrite (46.5 

weight percent Fe and 53.4 weight percent S). The spheres 

shown in figure 5B also exhibit the rough, nodular surface 

features that resemble the physical arrangement of crystallites in framboids. Internal nucleation of pyrite crystals has 

been invoked as a possible mechanism for the formation of 

pyrite framboids from spherical greigite precursors 

(Sweeney and Kaplan, 1973); however, the reactions and 

processes that lead to framboidal pyrite formation remain 

unknown.

Sphere Diameters

The modal sphere diameter of FeS2 increases with 

time within individual experiments. Figure 6 shows the 

particle size distribution from a sample at 128 hours from 

experiment 6; the modal diameter is 4.7 ^m, and the range 

of sizes is from 0.6 to 5.3 (xm. The reasons for the range in 

sphere sizes are not known, but the range may represent a 

dependence of growth rate on the size of the initial 

precipitate nuclei or changing saturation state of the solution 

(Skoog and West, 1982). Spheres larger than approximately 

5 |im may result from growth at the expense of the smaller, 

early formed spheres (Shaw, 1980; Stumm and Morgan, 

1981). Spheres having diameters of 3-4 (im are larger than 

those that precipitated at the start of FeS2 formation 

(0.5-1.7 |im at 20 hours).

Comparison of size distributions from similar 

experiments at 10 percent H2S(g) shows that the modal 

sphere diameter increases with time of reaction. The modal 

sphere diameter in experiment 7 markedy increased after 68 

hours, at which time the solution was no longer saturated 

with FeS. Spheres from experiment 6 at 128 hours (fig. 6) 

have a modal diameter of 4.7 Jim. Spheres from experiment 

7 at 139 hours have an modal sphere diameter of 4.4 pn 

(fig. 7, table 3), and the minimum diameter increased from 

0.7 nm at 68 hours to 2.4 (im at 93 hours. The maximum

E10 Green River Formation, Wyoming, Utah, and Colorado
diameter increased only by small amounts after 93 hours; 

apparently, the maximum sphere diameter (approximately 

5.0-5.5 jam) was attained shortly after FeS undersaturation 

was reached. The rate of size increase coincides with the 

increase in the rate of FeS2 formation and perhaps reflects 

growth due to the more rapid formation of FeS2.

The size of the pyrite spheres decreases in diameter 

with increasing ionic strength. Data presented in figure 7 

and table 3 contrast the sphere sizes of experiment 7 (1=0.1 

M) and experiment 12 (1=2.0 M).

The modal diameter of the spheres in experiment 7 

(1=0.1) is 4.4 jam, (range 2.5-5.7 jam). In experiment 12 

(1=2.0), the modal size is from 0.5 to 0.6 jam, (range 0.4-1.1 

jam) (see table 3). The modal sizes in the other two 

experiments of intermediate ionic strength (10 and 11) are 

intermediate between these end members, 1.9-2.0 jam and 

0.7-0.8 jam for experiments 10 (1=0.3) and 11 (1=0.7), 

respectively.

The contrasting size relationships for experiments 7, 

10, 11, and 12 are also illustrated by the series of SEM 

micrographs in figure 8. Particle size analysis, based on 

visual measurements of almost 100 spheres from these four 

experiments, shows that the modal diameter is consistent 

within individual runs and decreases with increasing ionic 

strength (figs. 8,9). Differing size ranges were obtained for 

samples taken at almost the same time during each 

experiment ( 140-150 hours). The rate of FeS2 formation 

was more rapid in experiment 7 but similar for experiments 

10, 11, and 12.

Either nucleation or precipitate growth (but not both) 

is the dominant mechanism for controlling the sizes of 

solids. Generally, increases in the saturation state of a 

solution lead to increased nucleation rates, smaller particles, 

and slower precipitate growth (Skoog and West, 1982). 

More nucleation sites may develop as a result of higher 

ionic strength influences on FeS, S°, or other possible 

nuclei. Alternatively, the same number of nuclei may form, 

but slower growth adds less pyrite to these nuclei (Adamson, 1976). In addition, specific cations or anions (such as 

Cl~, NO3~, K+) may selectively influence growth or 

nucleation processes and, thus, particle sizes (van Straten 

and others, 1985). It is unknown which of these processes is 

dominant in our experimental system.

High pH Results

At high pH (>7.0), iron was slowly incorporated into 

FeS and (Fe2+,Fe3+)-oxyhydroxide phases or remained as 

unreacted a-FeOOH. Solid reaction products are FeS, S°, 

and ferrous-ferric solids known as green rust. FeS and S° 

formed at all pH values, but in experiments at pH greater 

than 7.0 FeS formed in smaller amounts compared to low 

pH experiments and was only a minor component of the 

solids as the pH exceeded 9.0. Elemental sulfur formed 

slowly during all high pH runs and was most abundant and

becoming a dominant product (relative to iron monosulfide) 

only toward the end of most experiments. By contrast, at 

low pH, maximum sulfur values were reached about the 

same time (2-8 hours) as maximum FeS values (fig. 2). 

Thus, formation of sulfur by sulfide oxidation by Fe(III) 

occurred more slowly at high pH as compared to low pH. 

Goethite remained as a major constituent of the solids 

during most high pH experiments.

At the neutral pH values of experiments 16 (pH 7.1) 

and 19 (pH 7.0), a black solid slowly formed on the goethite 

surface and persisted for most of the experiment. X-ray 

diffraction of an unleached, late-stage (94 hours) sample 

from experiment 19 indicated that the black coating was 

FeS. This layer of FeS was removed by treatment with 6N 

HC1 and orange-yellow goethite was beneath it. X-ray 

diffraction of the 6N HCl-leached solids from experiment 

19 indicated that sulfur and goethite were the major solids 

from early in the experiment (3 hours) until 119 hours later. 

This is in marked contrast to the low pH experiments in 

which goethite completely dissolved and FeS formed within 

the first 2-8 hours. After 119 hours in experiment 19, 

goethite was not present in the 6N HCl-leached solids. 

Beyond 119 hours, S° was the dominant solid product in 

addition to FeS. It should be pointed out that experiment 19 

had the slowest rate of pyrite formation of any run in which 

pyrite formed. This slow rate is probably a reflection of the 

slow initial sulfide-Fe(IH) reaction at high pH and low 

H2S(g). Solid green rust was not seen in experiments 16 and 

19, but the light-green color of the raw sample filtrates until 

late in the reaction (>100 hours) indicates the presence of 

ferrous iron. The dominant form of aqueous iron is 

predicted to be Fe(OH)+ in the pH range 7.0-8.5 (Pyzik and 

Sommer, 1981).

At the highest pH values (pH>9.0) (experiments 18, 

32, and 33, table 1), mixed (Fe2+,Fe3+) -oxyhydroxide 

solids known as green rust formed (Rickard, 1969a; Newman, 1987). Reduced iron (Fe2+) in these minerals imparted 

a green color to the solid, although Fe3+ was also present. 

The color of these solids and the reaction solution 

intensified from light green to a dark blue green with higher 

solution pH. Oxidation of the Fe2"1" in these minerals to Fe3+ 

after a few hours of exposure to air is indicated by a change 

in color from green to brown. These solids were not well 

characterized due to their instability, variable composition, 

and poor crystallinity.

Small black particles coexisted with the green rust. 

These particles were likely FeS, but this could not be 

verified because XRD results were inconclusive. The 

solubility of the black coating in 6N HC1 and the high pH of 

formation is consistent with FeS.

The green rust is a thin layer coating the a-FeOOH. 

X-ray diffraction of an untreated green rust sample from an 

early sample (24 hours) in experiment 33 detected only 

goethite and a few peaks characteristic of sulfur. The 

presence of sulfur was verified by ultraviolet analysis of the

Experimental Study of Goethite Sulfidization E11
Figure 5 (above and facing page). Scanning electron microscope photomicrographs of pyrites 

formed in experiments 29 and 30. A, Pyrite spheres formed at low H2S(g) concentration (0.1 

percent), experiment 30. Spheres have smoother surfaces than spheres from experiment 7 (fig. 4£) 

for similar reaction times. Bar, 1 um; time, 126 hours. B, Pyrite spheres from an unleached sample 

at 188 hours from experiment 29 run at low H2S(g) concentration (0.1 percent). With increasing 

reaction time, spheres formed at low H2S(g) became similar to rough-textured spheres formed at 

high H2S(g) concentration. Large dipyramid in center is S°. Nodular surface of FeS2 spheres is 

similar to framboidal texture of naturally formed sedimentary pyrite. Bar, 2 um.

acetone-extracted solids. Treatment of this sample with 6N 

HC1 resulted in immediate removal of the green rust and 

revealed orange-yellow goethite underneath. X-ray 

diffraction of the 6N HCl-treated sample again revealed 

only goethite and sulfur; the green rust apparently was 

present in amounts less than 5 volume percent in the raw 

sample. As reaction time increased, more of the goethite 

was converted to green rust, but goethite remained as a 

major component of the 6N HCl-leached solids. Elemental 

sulfur was also present; thus, Fe(III) was reacting with 

aqueous sulfide to produce S° but at a rate slower than 

similar experiments at low pH. The layer of green rust 

seemingly inhibited Fe3+ reduction and goethite dissolution. 

Thus, over the pH range 7-10, the rate of a-FeOOH 

reduction-dissolution decreases, S° and FeS form more 

slowly, and FeS2 formation is inhibited. Both the green rust 

and FeS inhibit goethite dissolution (and presumably 

subsequent pyrite formation) at high pH values. Pankow 

and Morgan (1979,1980) noted that the rate of mackinawite 

dissolution is dependent on [H+] if the pH is less than 5.3. 

Adsorption of Fe2+ onto goethite at acidic pH values is

enhanced as the pH increases (LaKind and Stone, 1989) and 

may act as an additional barrier to goethite dissolution. 

Similarly, Tewari and Campbell (1976) have shown that the 

rate of anoxic dissolution of FeS (troilite) is slower as the 

pH increases to 6.5. This inhibition of reduction and 

dissolution of goethite is consistent with increased FeS 

saturation and therefore diminished tendency of FeS to 

dissolve at pH greater than 7.0. The slowness of pyrite 

formation as the solution becomes more alkaline may be 

due to changes in the solubility and rate of dissolution of 

both green rust and FeS.

RELATIONSHIPS TO IRON AND 

SULFUR DIAGENESIS

Various iron minerals react differently during sulfidization due to variations in crystallinity, solubility, or 

surface charge. For example, at neutral pH values 

crystalline Fe(OH)3 is more readily sulfidized to FeS than is 

crystalline a-FeOOH (Pyzik, 1976). The reactivity of iron 

minerals depends upon several factors, some of which are

E12 Green River Formation, Wyoming, Utah, and Colorado
cc

LU

m

SIZE INTERVAL, IN MICRONS

Figure 6. Histogram of modal sphere diameters from 

experiment 6 (128 hours). Modal sphere diameter is centered 

in the 4-5 fj.m interval. Increase in modal sphere diameters as 

compared to those of earlier samples from same experiment 

indicates an increase in sphere size with reaction time.

4-

03 z. 

o 

cc 

o

CC 

LU

H 

LU

Q 2H

LU 

CC 

LU 

X 

D. 

03

1 -

< 0

§ °

50 100 

TIME, IN HOURS

1 50

Figure 7. Modal sphere diameter versus reaction time at 

two ionic strengths : experiment: 7 (1=0.1 M) (squares) and 

experiment 12 (1=2.0 M) (circles). At 1=0.1 M sphere 

diameter increases with reaction time, whereas at 1=2.0 M 

sphere diameter remains constant with reaction time. 

Increasing ionic strength promotes formation of smaller 

sized spheres for otherwise similar experimental 

conditions.

Experimental Study of Goethite Sulfidization E13
Table 3. Modal sphere diameter and size range for pyrKe 

spheres formed in experiments 7 and 12

[na, not applicable; nd, not determined. Ionic strength experiment 7,0.1 

M; experiment 12,2.0 M]

Sample no. Time (hours) Modal diameter (|jm) Size range (\m\)

Experiment 7

1

2

3

4

5

6

7

8F

2

22

43

68

93

115

118

139

na

1.3

2.6

2.7

4.0

nd

4.2

4.4

na

0.6-1.4

0.7-3.5

0.7-3.4

2.4-5.2

nd

2.2-5.1

2.5-5.7

Experiment 12

4

5

6

7

9

46

72

97

118

144

0.4

0.5

0.5

0.6

0.6

0.3-0.5

0.3-0.8

0.4-0.8

0.3-1.1

0.4-1.1

detailed by Canfield (1988, 1989). Iron oxyhydroxides such 

as goethite or lepidocrocite are common sources of iron 

available for sulfidization and are generally delivered to a 

depositional site as coatings on clay minerals (Carroll, 

1958). The sum of adsorbed and detrital goethite may be as 

much as 50 percent of the total iron load carried by rivers 

(Gibbs, 1973). Our experimental results show that the 

reactive nature of goethite in a wide range of conditions 

results in formation of a variety of dissolved and solid 

reaction products. Goethite is completely converted to FeS 

and FeS2 at low pH or partly converted to these sulfides as 

pH increases or H2S decreases. As pH exceeds 7.0, goethite 

remains as an unreacted solid; at pH greater than 9.0, 

modeling indicates that goethite does not dissolve and may 

even form.

Rates of Pyrite Formation

FeS, FeS2 and S° concentrations in the upper layers 

of marsh sediments are subject to temporal variation as a 

result of seasonal oxidative recycling (such as 

FeS »Fe(OH)3 >FeS), whereas the concentration of FeS2 

remains generally constant below approximately the top 

10-15 cm of sediments (Luther and others, 1982; Cutter and 

Velinsky, 1988). Depending on the predominance of 

sulfide-mineral-forming reactions as seen in this study, 

rapidly formed FeS and S° are either sinks or sources for 

iron and sulfur over short time periods (less than one year), 

and short-term shifts in the iron and sulfur budgets of these 

sediments may result. The results here reinforce the fact that 

pyrite formation from iron monosulfide and zerovalent 

sulfur is very slow at neutral to slightly alkaline pH and

possibly inhibited at pH of 9.0 or more. As diagenesis 

proceeds, however, there is more time for the slow 

conversion of FeS to pyrite at high pH (Berner, 1969). 

Pyrite formation and subsequent burial below the oxidizing 

zone would remove some of the available iron and sulfur 

from further diagenetic cycling because pyrite is the 

thermodynamically stable iron-sulfur mineral in reducing 

sediments (Berner, 1964a).

The buffering of porewaters maintains most natural 

environments within the pH range of approximately 6.9-8.3 

(Ben-Yaakov, 1973). Thus, goethite sulfidization and pyritization should occur more slowly under natural conditions 

than in the very low pH (<5) experiments. The major effect 

of high pH is to slow down the initial Fe(III)-sulfide redox 

reaction, which is rapid at acidic pH. Acidic conditions can 

be generated, however, by bacterial processes in microenvironments or in organic-rich swamps and marshes 

(ZoBell, 1954; Birnbaum and Wireman, 1984). Thus, the 

very low pH of some experiments could be encountered in 

these natural environments, and the formation of FeS and 

FeS2 may be rapid; however, in most terrestrial 

environments pyrite formation is limited by sulfate 

availability and not by iron or pH controls (Goldhaber and 

Kaplan, 1975).

The rates of iron monosulfide and pyrite formation in 

our study are more rapid than would occur in natural 

settings because diffusion of reactants in a heterogeneous 

sediment-water system is slow (Rickard, 1969a, b; Berner, 

1970; Canfield, 1989). Because our system was well mixed, 

unlike natural sediments, diffusive control of the reactions 

was minimized.

Pyrite Texture

The morphology of the solids produced in our 

experiments implicates chemical dissolution of 

a-FeOOH, S°, and FeS in the sequence cc-FeOOH+H2S-» 

FeS+S° »FeS2. The discrete blade morphology of the 

a-FeOOH was not retained by the FeS or S8 that formed (as 

shown in fig. 4A). In addition, pseudomorphic replacement 

of FeS or S8 by FeS2 was not observed in the early stages of 

experiments. Instead, pyrite spheres were formed that did 

not resemble FeS or S8 textures or sizes. The rapid 

formation of pyrite as the result of undersaturation of FeS 

indicates that dissolution of FeS is important in the growth 

of pyrite. Low-temperature FeS2 formation likely involves 

dissolution of both FeS and S°, but as temperature increases 

the solid-solid reaction of FeS and S° may become more 

important (Rickard, 1975). The spatial association of FeS2 

and S° shown in figure 5B suggests that solid surfaces may 

be involved in pyrite formation, but our results do not allow 

for conclusive identification of the mechanisms involved in 

the conversion of FeS to FeS2 under these conditions.

The only apparent difference in pyrite specifically 

attributable to the slow or fast FeS2 precipitation sequences

E14 Green River Formation, Wyoming, Utah, and Colorado
is an increase in sphere size with time (figs. 6, 7). This 

increase .suggests that iron and sulfur continue to be added 

to the early formed pyrite spheres. The spherical texture 

may represent the result of one of the major, early FeS2- 

forming pathways in natural environments. Pyritic microspheres have been observed in different sedimentary 

environments, including near-shore sediments of Long 

Island Sound (Canfield and Berner, 1987) and the Gulf of 

California (Berner, 1964b, 1969) and organic-rich swamps 

(Postma, 1982). Our ionic strength results indicate that the 

size of a pyrite sphere may depend, in part, on salinity and 

thus reflect the environment of formation. Euhedral pyrite 

may result from pyrite precipitation without formation 

of an intermediate FeS solid (for example, Fe(|q)+ 

Sn"(aq)^FeS2+Snl2(aq)) (Howarth, 1978; Raiswell, 

1982) rather than from conversion of monosulfides to pyrite 

(Goldhaber and Kaplan, 1974). The lack of euhedral pyrite 

under these experimental conditions that is, in the 

presence of an FeS precursor supports this observation.

Other Reaction Products

The inorganic formation of iron mono- and disulfides from iron oxyhydroxide produces.and consumes 

elemental sulfur, a recognized mineral phase in marine 

sediments (Berner, 1964b). Pyzik (1976) reported that 

elemental sulfur comprises as much as 90 percent of the 

sulfide oxidation products of (X-FeOOH sulfidization at pH 

7.0-8.5; however, the quantitative significance of this 

inorganic sulfur formation mechanism in natural systems is 

not known because many reactions involving sulfur are 

biologically mediated. The distribution of sulfur among the 

pyrite, monosulfide, sulfate, and elemental sulfur phases 

represents the difference between zerovalent sulfur 

production and consumption, including oxidation and 

reduction of sulfur species by microorganisms (Troelsen 

and Jorgensen, 1982). Because elemental sulfur is not 

abundant in most marine sediments (2-4 weight percent), it 

may become limiting in the conversion of monosulfides to 

pyrite (Berner, 1964b).

Pyrite did not form in any low pH experiment until 

after elemental sulfur was produced. Berner and others 

(1979) noted that the conversion of mackinawite and 

greigite to FeS2 in marine sediments requires excess H2S. 

Thus, excess zerovalent sulfur or H2S(aq), or both, may be 

critical to the subsequent formation of pyrite under these 

conditions (Berner, 1970; Sweeney and Kaplan, 1973). In 

the absence of excess sulfide, mackinawite and greigite are 

metastable phases in marine sediments.

The reduction of iron(III) by sulfide species at low pH 

and high H2S produces aqueous iron(II). Measured values 

of dissolved iron and H2S in marsh sediments indicate that 

greigite solubility controls aqueous iron concentrations in 

the top 10-15 cm of sediment and FeS solubility controls

aqueous iron in the 50-55 cm interval (Boulegue and others, 

1982). The concentration of aqueous iron(II) in our 

experiments (lO^-lO^6 M) indicates that FeS, not FeS2, is 

the controlling phase for dissolved iron in most sediments. 

Aqueous iron can react with different sulfide species in 

solution to form a variety of iron sulfides (Berner, 1964a; 

Rickard, 1969a, b). Once Fe3+ is reduced, the more mobile, 

soluble Fe2+(aq) may form sulfides at sites removed from the 

point of dissolution of the iron source (goethite).

Green rust phases are formed primarily through 

reaction of Fe(II)-hydroxy and Fe(III)-hydroxy species 

(Newman, 1987) and commonly are intermediate-stage 

products resulting from the incomplete oxidation or 

reduction of iron-bearing substances. In natural 

environments, anaerobic conditions provide soluble 

iron(II)-hydroxy species necessary for formation of these 

minerals. The mottling of some gleyed soil horizons and the 

gray-green color in the uppermost layers of some lake and 

marine sediments have been attributed to the presence of 

mixed reduced and oxidized iron minerals, which may 

include green rust phases (Buurman, 1980; Newman, 1987). 

Soils and lake sediments containing available iron and 

having seasonal fluctuations of photosynthetic and 

respiratory processes and thus unstable redox conditions are 

likely environments in which green rusts may form. 

However, green rusts probably are transient phases that are 

unstable with respect to goethite and lepidocrocite (Newman, 1987). Our experimental results point to variations in 

pH and redox conditions as key factors in green rust 

formation.

Absence of Other Sulfide Reaction Products

Marcasite was not detected by X-ray diffraction or 

reflected-light petrography. Marcasite, which is commonly 

found in association with pyrite, requires a low pH to form 

(Goldhaber and Stanton, 1987), but it did not form in our 

experiments despite the low pH(init) (4-5) in the free-drift 

runs. Evidently, under the conditions studied, marcasite 

formation is inhibited by some unknown mechanism if 

a-FeOOH is the iron source.

Figure 8 (following pages). Scanning electron microscope 

photomicrographs illustrating decrease in pyrite sphere size 

with increasing ionic strength. Micrographs are arranged in 

order from lowest to highest ionic strength and from highest to 

lowest modal sphere diameter. Bar scale on all micrographs is 

2 urn. A, Pyrite spheres from experiment 7. Ionic strength, 0.1 

M; modal diameter, 4.4 urn; time, 139 hours. B, Pyrite spheres 

from experiment 10. Ionic strength, 0.3 M; modal diameter, 

2.7 um; time, 143 hours. Small needles are probably melanterite, an oxidation product of pyrite. C, Pyrite spheres from 

experiment 11. Ionic strength, 0.7 M; modal diameter, 0.7 um; 

time, 144 hours. D, Pyrite spheres from experiment 12. Ionic 

strength, 2.0 M; modal diameter, 0.6 mm; time, 144 hours.

Experimental Study of Goethite Sulfidization E15
E16 Green River Formation, Wyoming, Utah, and Colorado
12 015

Experimental Study of Goethite Sulfidization El 7
C/) 

O 5

cc

O

2 4H

DC 

LU HLU

5 2-

LU

cc

LLJ 

X 

D.

O 

O

1 - 1 2

0.0 1 .0 2.0 3.0

MOLAR IONIC STRENGTH

Figure 9. Modal sphere diameter (MD) versus ionic strength 

for experiments 7, 10, 11, and 12. Sphere diameters are from 

samples taken at approximately the same time from each 

experiment.

X-ray diffraction data suggest that greigite either did 

not form during the experiment, formed in low quantities 

(<5 volume percent), or formed and was converted rapidly 

to another phase. Rickard (1969b) noted that greigite forms 

in the pH range of 2.5-6.5, well within the range employed 

in this study. The formation of greigite may require, in a 

manner similar to marcasite, a different mechanism at low 

pH if goethite is the iron source. It has been postulated that 

greigite contains ferric iron and therefore requires small 

amounts of oxygen for formation (Goldhaber and Kaplan, 

1974). The rigorous exclusion of air and the strongly 

reducing conditions employed at low pH and 10 percent 

H2S(g) should rapidly reduce ferric iron.

The stable form of pyrrhotite (Fe^S) under 

sedimentary conditions, hexagonal pyrrhotite (Arnold, 

1969), has been synthesized at low temperatures, but there 

is little evidence for its rapid formation in modern sediments 

(Pyzik, 1976). In our study, pyrrhotite was not detected. 

Similar to greigite, it could not be determined if it was not 

formed or if it was rapidly produced and converted to some 

other sulfide phase. The ratio of iron to sulfur in pyrrhotite 

is approximately 0.9, based on the mineral formula Fe^S, 

where x is less than 0.125 (Ward, 1970). Pyrrhotite contains 

less iron than mackinawite and iron:sulfur ratios of our 

experimental products indicate that the initial FeS phases 

were iron rich (Fe:S >1.0) rather than iron deficient 

(table 2).

CONCLUSIONS

A variety of solid and aqueous iron-bearing products 

are initially formed from the reaction of goethite and 

aqueous hydrogen sulfide. The transformation of goethite to 

pyrite takes place slowly at the near-neutral pH of most 

natural sediments and perhaps not at all at the high pH 

(>9.0) of some alkaline lake systems. Thus, the early 

diagenesis of iron and sulfur may follow any of several 

pathways that are controlled primarily by pH; however, 

other factors operating in natural depositional settings (for 

example, rate of sulfate reduction) that were not taken into 

account in this inorganic system may influence these 

pathways.

It has been postulated that the Green River Formation 

was deposited in an alkaline environment that had a pH of 

more than 8.0 (Turtle, 1988). High pH results of our study 

indicate that iron probably was not immediately or directly 

incorporated into pyrite. High pH may also explain the lack 

of pyrite in modern analogs of the Green River lacustrine 

system, such as Walker Lake and Soap Lake (Tuttle, 1988). 

A more complex, long-term sequence of diagenetic events 

must have produced the present-day suite of iron sulfide 

minerals in the Green River Formation. The absence of 

pyrrhotite, greigite, and marcasite indicates that these 

minerals may form only after extended time periods or 

under different conditions than examined in our study.

Pyrite produced during the experiments has a 

spherical morphology that superficially resembles early 

stages of framboidal development. Euhedral crystals were 

not formed in our study. Thus, formation of euhedral or 

framboidal forms of pyrite may also require extended 

reaction time or other mechanistic pathways. Increasing 

ionic strength produces smaller spheres for otherwise 

similar conditions. Pyrite formed under controlled 

conditions, however, may differ from that in natural unconsolidated sediments; other factors not considered here may 

significantly affect pyrite formation rates and textures 

(Goldhaber and Kaplan, 1974; Raiswell, 1982; Berner and 

Westrich, 1984; Canfield and Berner, 1987).

SUMMARY

The pathways of formation of iron monosulfide and 

pyrite from goethite, with qualitative indicators of the 

reaction rates, are summarized below.

pH(init)<5.0

a-FeOOH+H2S(fast)->FeS+S°+Fe2+(aq)(slow)-»FeS2 

pH(init)5.0-7.0

a-FeOOH+H2S(slow)-»FeS+S°+Fe2+(aq)(slow)-»FeS2 

pH>7.0<9.0

a-FeOOH+H2S(slow)^FeS+S°+Fe(OH)++cc-FeOOH

El 8 Green River Formation, Wyoming, Utah, and Colorado
pH>9.0

oc-FeOOH+H S(slow)-»FeS+S° +Fe(OH)++a-FeOOH+

(Fe2+,Fe3+)x(OH)x

Differing amounts of solid a-FeOOH were reacted 

with H2S over a range of chemical conditions at 30 °C. 

Variations in the pH, ionic strength, and concentrations of 

iron and sulfur species led to a variety of reaction products 

that formed at differing rates. The solid products of the 

goethite-hydrogen sulfide reaction at acidic pH (4-<7) 

were zerovalent sulfur, iron monosulfide, and, eventually, 

pyrite. Pyrite was not a product of experiments conducted at 

pH greater than 7.0. The overall rate of pyrite formation for 

varied total dissolved sulfide and iron concentrations was 

from 10~7 to 10~9 M/min for pH 4-7.

The major solids formed during experiments 

conducted at higher pH (>7-10) were iron monosulfide, 

zerovalent sulfur, and ferrous-ferric oxyhydroxides. Unreacted goethite was a common constituent of the solids over 

the time span of high pH experiments.

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Colloid and Interface Science, v. 106, p. 175-185.

Ward, J.C., 1970, The properties of some iron sulphides: Pure and 

Applied Chemistry, v. 20, p. 175-206.

ZoBell, C.E., 1954, Changes produced by microoorganisms in 

sediments after deposition: Journal of Sedimentary 

Petrology, v. 12, no. 3, p. 127-136.

E20 Green River Formation, Wyoming, Utah, and Colorado
Chapter F

Effects of Source, Depositional Environment, 

and Diagenesis on Characteristics of Organic 

Matter in Oil Shale from the Green River 

Formation, Wyoming, Utah, and Colorado

By WALTER E. DEAN and DONALD E. ANDERS

U.S. GEOLOGICAL SURVEY BULLETIN 1973

GEOCHEMICAL, BIOGEOCHEMICAL, AND SEDIMENTOLOGICAL STUDIES OF THE 

GREEN RIVER FORMATION, WYOMING, UTAH, AND COLORADO
CONTENTS

Abstract Fl 

Introduction Fl 

Analytical methods F3 

Results and discussion F3

Pyrolysis and gas chromatography data F3

Carbon isotope data Fll 

Conclusions F14 

References cited F15

FIGURES

1. Map showing Laramide basins and uplifts, southwestern Wyoming, northwestern 

Colorado, and northeastern Utah F2

2. Chart showing stratigraphy of cores of Green River Formation F3 

3-6. Plots of:

3. Depth versus total organic carbon, hydrogen index, oxygen index, and 

513C F7

4. Oxygen index versus hydrogen index F8

5. Pyrolysis data F9

6. Pyrolysis Sj+ S2 yield for samples of oil shale of different geologic ages

and from different geologic environments F9 

7-9. Gas chromatograms of saturated-hydrocarbon fraction of:

7. Samples from three depths in Blacks Fork core F10

8. A sample from Mahogany zone in Coyote Wash core Fll

9. Samples from three depths in WOSCO EX-1 core F12

10. Chart showing range of 513C values in lake sediments, peat, lake plankton, 

plants, and samples of Green River Formation F12

11. Plots of hydrogen index versus 513C in samples of Green River 

Formation F13

TABLES

1. Fischer assay oil yield, total organic carbon content, Rock-Eva! pyrolysis results, 

and values of 513C of organic carbon for samples of oil shale from the Coyote 

Wash, Blacks Fork, and 01A cores of the Green River Formation F4

2. Fischer assay oil yield, total organic carbpn content, and Rock-Eva! pyrolysis 

results for samples of oil shale from the WOSCO EX-1 core of the Green River 

Formation F6

Contents
Effects of Source, Depositional Environment, and 

Diagenesis on Characteristics of Organic Matter in 

Oil Shale from the Green River Formation, Wyoming, 

Utah, and Colorado

By Walter E. Dean and Donald E. Anders

Abstract

Results of Rock-Eval pyrolysis and stable-carbon isotope 

analysis were used to assess the richness, source, and degree of 

degradation of oil shale in the Green River Formation sampled 

in basin-center cores from the Green River basin of Wyoming, 

Piceance basin of Colorado, and Uinta basin of Utah. By far 

the richest oil-shale samples are from the Piceance basin, 

particularly from the Mahogany zone near the top of the 

Parachute Creek Member of the Green River Formation. Some 

samples of oil shale from the other two basins are as rich, but, 

in general, greater oxidation of organic matter and dilution of 

organic matter by terrigenous clastic material has taken place 

in these shallower lake basins.

Hydrogen-rich organic matter accumulated in deeper, 

less saline phases of Eocene Lake Gosiute, represented by the 

Tipton Shale and Laney Members of the Green River 

Formation in the Green River basin, but much of the organic 

matter that accumulated during deposition of the shallower, 

more saline intervening Wilkins Peak Member was from both 

algal and terrestrial sources. Extremely high algal productivity 

and rapid burial of 13C-depleted organic carbon in Eocene 

Lake Uinta, culminating in the highly organic-carbon-rich 

strata of the Mahogany zone in the Piceance and Uinta basins, 

produced enrichment of 13C in surface waters of the lake that 

eventually affected the isotopic composition of organic carbon 

fixed by phytoplankton.

INTRODUCTION

The Green River Formation (Eocene) of the Green 

River, Piceance, and Uinta basins of Wyoming, Colorado, 

and Utah, respectively, is one of the most remarkable 

accumulations of organic matter in the world. The "oil

shales" of this formation commonly contain more than 10 

weight percent organic carbon in a suite of sedimentary 

rocks that is as varied as any imaginable, from dolomitic 

marlstone to mudstone, siltstone, sandstone, and volcanic 

ash and containing a multitude of exotic sulfide, evaporite, 

and authigenic silicate minerals (Desborough and Pitman, 

1974). With all of this lithologic variability, it is surprising 

that few of the rocks in the Green River Formation called 

"oil shale" can be classified as shales. The dominant 

lithology of the richest oil shales in the Piceance and Uinta 

basins is dolomitic marlstone, but even that classification 

belies the complexity of the rock, which is a mixture of 

almost every possible combination of carbonate minerals in 

the system Ca-Mg-Fe-CO3 (Desborough and Pitman, 1974; 

Desborough, 1978). The complex mineralogy and geochemistry of the Green River make it even more remarkable 

because these minerals, most of which are autochthonous or 

authigenic, act as dilutants to the organic matter; without 

these diluting minerals, the Green River Formation would 

consist almost entirely of solid kerogen. The purposes of 

this paper are to characterize kerogen in the Green River 

Formation using results of Rock-Eval pyrolysis, gas 

chromatography, and stable-carbon isotope analysis and to 

interpret these results in terms of source, depositional 

environment, and diagenesis of organic matter.

Splits of selected samples of the Green River 

Formation from the Blacks Fork core in the Green River 

basin, Wyoming, the 01A core in the Piceance basin, 

Colorado, and the Coyote Wash core in the Uinta basin, 

Utah (figs. 1, 2), collected for inorganic geochemical 

analysis (P.H. Briggs and M.L. Tuttle, written commun., 

1985) were used. Descriptions of these samples are given in

Characteristics of Organic Matter in Oil Shale F1
Turtle (this volume). Samples were selected from the larger 

suite of samples analyzed for inorganic geochemistry based 

on organic-carbon content and stratigraphic position in 

order to provide results over a wide range of organic-carbon 

concentrations and the complete stratigraphic sequence 

sampled in each core.

The cores are from the centers of the three 

depositional basins of Lakes Uinta and Gosiute (Turtle, this 

volume), and the geochemical characteristics of the rocks in

113°

these cores should be representative of the Green River 

Formation deposited in the deepest parts of each of the 

basins. We will also use some geochemical characteristics 

of thermally immature oil-shale samples from other basincenter cores from the Piceance and Uinta basins to 

demonstrate the geochemical similarities of Green River oil 

shale in each of these two basins and to provide additional 

data for interpretation. From the Uinta basin, we will use 

additional data from the Mahogany zone in the WOSCO

106°

ABSAROKA VOLCANIC FIELD

EXPLANATION ___ _. 

Present-day limit of Tertiary rocks 0 50 100 KILOMETERS

Upper Cretaceous outcrop

Present-day outcrop of Lower 

Cretaceous through Cambrian rocks

Present-day outcrop of Precambrian rocks

Figure 1. Laramide basins and uplifts in southwestern Wyoming, northwestern Colorado, and 

northeastern Utah. Core holes: 1, Blacks Fork no. 1; 2, 01 A; 3, Coyote Wash; 4, WOSCO EX-1; 5, 

Colorado no. 1. Modified from Johnson (1985).

F2 Green River Formation, Wyoming, Utah, and Colorado
460

Coyote Wash

669- 

681 ~

710-

950

1,055

o fc 

J§E

O) <?

220

01A Blacks Fork

273 - 

291 -

329 -7

 706

774

'ESI 

ro 3 <» C5C5-

55

460 ^

Lake 

stage 3

Lake 

stage 2

Lake 

stage 1

50 100 METERS

Figure 2. Stratigraphy of cores of Green River Formation 

analyzed in this study. Location of core holes shown in figure 1. 

Mz, Mahogany zone.

EX-1 core (fig. 1) (Anders and Gerrild, 1984), and from the 

Piceance basin we will use additional data from the entire 

Green River Formation in the Colorado no. 1 core (fig. 1) 

(Smith and others, 1968; Anders and Robinson, 1973; 

Robinson and Cook, 1975).

ANALYTICAL METHODS

The Rock-Eva! pyrolysis method provides a rapid 

estimate of the type and hydrocarbon-generating potential 

of sedimentary organic matter (Espitalie" and others, 1977). 

Programmed heating of a sample in a helium atmosphere 

results in the generation of hydrocarbons (HC) and carbon 

dioxide. Free or adsorbed HC yield is recorded as the area 

under the first peak on a program (Sj) (mg HC/g sample) 

during heating of the sample to 250 °C for 5 min. The 

second pyrolysis peak (S2) represents hydrocarbons 

generated by thermal breakdown of kerogen during heating 

of the sample from 250 to 550 °C. CO2 is also generated by 

kerogen degradation and is retained during the heating 

interval from 250 to 390 °C and analyzed as the integrated 

area under the third peak on the pyrogram (S3) (mg CO^g 

sample). The Rock-Eval n instrument used for these

analyses also measures total organic carbon (TOC) as the 

sum of carbon in the S x and S2 hydrocarbons plus that 

determined by oxidizing the residual carbon at 600 °C. In 

addition to the S lt S2, and S3 integrated peaks, the results 

are expressed in terms of a hydrogen index (HI), which is 

the S2 peak normalized to the total organic carbon (TOC) 

content of the sample (mg HC/g TOC), and an oxygen index 

(OI), which is the S3 peak normalized to TOC content of the 

sample (mg CO2/g TOC).

The soluble-organic-matter (SOM) fraction of wholerock samples was obtained by Soxhlet extraction with 

chloroform for 24 hours. The extract was concentrated 

under nitrogen to remove the chloroform, then diluted with 

iso-octane. A concentrate of the SOM in iso-octane was 

fractionated by column chromatography on silica gel/ 

alumina by eluting successively with iso-octane, benzene, 

and benzene-methanol to collect the saturated hydrocarbon, 

aromatic hydrocarbon, and resin fractions, respectively. The 

saturated hydrocarbon fractions were analyzed by capillarycolumn gas chromatography using a Hewlett-Packard 5880 

instrument with a flame-ionization detector. Peaks on chromatograms were identified by co-elution using standards or 

by fragmentation patterns of gas chromatography/mass 

spectrometry (GC/MS) electron impact spectra. Relative 

abundances of hydrocarbons are indicated by the relative 

intensities of detector responses.

Stable-carbon isotope ratios of organic carbon were 

determined by standard techniques (for example, Pratt and 

Threlkeld, 1984). Powdered samples were oven dried at 40 

°C and reacted with Q.5N HC1 for 24 hours to dissolve 

carbonate minerals. If all of the carbonate was not removed 

by the first acid leach, the sample was reacted with acid for 

an additional 24-hour period. The residue was centrifuged, 

decanted, washed three times with deionized water, dried 

under a stream of flowing nitrogen at 50 °C, and combusted 

at 1,000 °C under oxygen pressure in a LECO induction 

furnace. The resulting CO2 was dehydrated and purified in 

a high-vacuum gas-transfer system, and the isotope ratios 

determined using a Finnigan MAT 251 6 in., 90° sector, 

isotope-ratio mass spectrometer. Results are reported in the 

usual delta notation as the per mil difference relative to the 

Pee Dee belemnite marine-carbonate standard (PDB):

513C (o/oo) = [(Rsample/RPDB) - I]xl03 , 

where R is the ratio 13C : 12C.

RESULTS AND DISCUSSION

Pyrolysis and Gas Chromatography Data

Results of Rock-Eval pyrolysis of samples from the 

Coyote Wash, Blacks Fork, and 01A cores are given in table 

1 along with TOC values determined by LECO (Tuttle and

Characteristics of Organic Matter in Oil Shale F3
Table 1. Fischer assay oil yield, total organic carbon content, Rock-Eval pyrolysis results, and values 

of 513C of organic carbon for samples of oil shale from the Coyote Wash, Black Forks.and 01A cores 

of the Green River Formation

(TOC, total organic carbon; CW, Coyote Wash; BF, Blacks Fork; LECO indicates using a LECO induction furnace]

Sample 

number

CW-38

CW-31

CW-29

CW-28

CW-26

CW-20

CW-18

CW-13

CW-11

CW-01

BF-04

BF-06

BF-07

BF-08

BF-13

BF-19

BF-22

BF-23

BF-31

BF-34

01A-006

01A-014

01A-024

01A-036

01A-054

01A-062

01A-072

01A-084

01A-089

01A-102

01A-113

01A-132

01A-150

01A-154

01A-165

01A-172

Depth 

(meters)

571

661

681

697

716

789

815

880

907

1,043

98

125

134

144

217

299

344

353

464

500

268

290

321

358

415

437

468

504

521

559

595

651

705

717

755

773

Oil 

(9Pt)

8

16

65

16

3

18

2

13

3

14

17

12

17

2

5

5

1

23

28

3

15

70

21

15

16

17

34

4

4

30

37

17

41

23

12

12

Oil 

(L/mt)

33

67

271

66

12

75

8

54

13

58

69

52

73

9

21

19

5

98

117

10

63

292

88

63

67

71

142

17

17

125

154

71

171

%

50

50

Percent TOC 

LECO Rock-Eval

3.4

5.6

26.5

2.5

2.4

7.6

1.2

5.8

1.8

33

83

4.1

7.0

1.5

2.8

4.1

0.58

83

12.0

10.6

9.6

33.7

10.2

23.0

5.6

17.9

133

7.4

1.2

15.0

14.1

8.5

21.6

11.8

8.6

3.7

3.17

5.08

28.9

2.59

2.05

7.25

1.48

6.41

1.71

431

6.82

3.64

6.82

2.18

2.75

4.65

0.56

8.69

12.0

12.0

10.8

37.5

10.7

273

5.42

20.10

14.8

7.51

1.23

18.0

16.9

9.06

24.1

12.5

8.52

3.67

Goldhaber, this volume). Results of Rock-Eval pyrolysis 

and TOC analyses of samples from the WOSCO EX-1 core 

are given in table 2. The results show that, in general, the 

agreement is good between TOC determined by the two 

methods. Values of TOC, HI, and OI are plotted versus 

depth for each of the three cores in figure 3. Values of HI 

and OI are plotted as cross plots for the three primary cores 

and the WOSCO core in figure 4. These cross plots are

equivalent to van Krevelen-type cross plots of H/C versus 

O/C and can be used to identify types of organic matter 

(types I, H, and HI of Espitah'6 and others, 1977).

The richness of oil shale can be measured using our 

results by the organic-carbon content (TOC), the pyrolytic 

hydrocarbon yield (Sj+ S2), or the hydrogen index (HI). The 

most commonly used measure of richness of Green River 

oil shale, however, is the amount of oil produced by

F4 Green River Formation, Wyoming, Utah, and Colorado
Table 1 . Fischer assay oil yield, total organic carbon content, Rock-Eval pyrolysis results, and values of 513C of 

organic carbon for samples of oil shale from the Coyote Wash, Black Forks, and 01A cores of the Green River 

Formation Continued

Sample 

number

CW-38

CW-31

CW-29

CW-28

CW-26

CW-20

CW-18

CW-13

CW-11

CW-01

BF-04

BF-06

BF-07

BF-08

BF-13

BF-19

BF-22

BF-23

BF-31

BF-34

01A-006

01A-014

01A-024

01A-036

01A-054

01A-062

01A-072

01A-084

01A-089

01A-102

01A-113

01A-132

01A-150

01A-154

01A-165

01A-172

S1 

(mg/g)

3.88

5.71

10.0

3.05

3.63

1.21

0.51

0.78

0.78

1.80

1.27

1.07

1.66

035

0.94

0.87

0.29

1.63

0.92

1.95

1.51

5.63

132

5.20

0.83

436

5.75

2.96

0.91

6.82

7.40

3.55

8.28

531

3.44

1.58

S2 

(mg/g)

20.1

35.4

287

15.0

12.2

61.6

5.0

42.8

103

24.6

553

33.5

51.8

73

14.9

26.0

2.7

603

97.0

113

79.5

365

76.7

246

44.0

169

121

60.7

73

146

138.6

773

240

119

63.9

27.6

S3 

(mg/g)

4.23

5.23

7.59

1.84

1.29

2.02

0.68

2.69

0.91

1.68

5.10

2.28

538

1.%

3.77

3.50

9.95

6.88

6.11

3.15

4.72

163

5.71

7.10

3.05

713

30.0

333

75.7

27.7

22.8

73.2

4.14

2.55

1.81

139

H-index 

(mg/g)

635

698

993

579

595

849

339

667

603

570

810

921

816

333

541

559

480

694

809

938

734

973

715

901

811

839

819

808

595

815

812

852

997

955

749

751

O-index 

(mg/g)

133

102

26

71

63

27

45

41

53

38

74

62

84

89

137

75

1,800

79

50

26

43

43

53

25

56

354

203

443

6,200

153

134

808

17

20

21

37

a13Corg 

(per mil PDB)

-29.6

-29.8

-28.5

-30.8

-283

-303

-303

-30.8

-30.5

-32.4

-30.4

-29.8

-293

-31.1

-31.2

-25.7

-32.1

-33.9

-31.8

-29.9

-30.9

-29.8

-29.2

-30.1

-31.4

-30.6

-30.9

-30.1

-31.1

-31.0

-32.3

-32.9

-30.0

-30.9

-29.9

retorting by Fischer assay analysis, expressed as percent oil 

yield, gallons per ton (gpt), or liters per metric ton (L/mt) of 

rock. For comparison, Fischer assay results (tables 1 and 2), 

(S!+S2) yield, and HI for Green River oil shale are plotted in 

figure 5, all versus TOC. Results of Fischer assay analysis 

most closely approximate those of pyrolytic hydrocarbon 

yield (SjH^) from Rock-Eval pyrolysis. Dyni and others 

(1990) presented Fischer assay and (Sj+Sj) hydrocarbonyield data for 74 samples of oil shale from different ages

and geologic settings from around the world. These data are 

plotted in figure 6 along with a plot of data for Green River 

oil shale. Using the plots in figures 5 and 6 it should be 

possible to convert values of any one richness parameter to 

values of any other richness parameter.

By far the richest oil-shale samples in the three 

primary cores studied are from the 01A core from the 

Piceance basin. This is in line with resource estimates that 

show the richest oil shales of the Green River Formation are

Characteristics of Organic Matter in Oil Shale F5
Table 2. Fischer assay oil yield, total organic carbon content, and Rock-Eval pyrolysis results for samples of oil shale from 

the WOSCO EX-1 core, Utah, of the Green River Formation

[TOC, total organic carbon; LECO indicates using a LECO induction furnace]

Depth, 

(feet)

2,181.5

2,200.4

2,240.5

2,270.5

2,280.5

2,290.5

2,300.6

2,320.5

2,341.0

2,359.5

2,377.2

2,408.5

2,422.6

2,444.5

2,459.5

2,480.5

2,500.5

2,521.5

2,536.5

2,560.5

2,573.5

2,607.5

2,618.5

2,640.5

2,660.5

2,680.5

2,706.5

2,723.5

2,746.5

2,760.5

2,780.5

2,800.5

2,820.5

2,844.5

2,860.5

2,880.5

2,901.5

2,923.5

2,940.5

2,962.5

Depth 

(meters)

664.92

670.68

682.90

692.05

695.10

698.14

701.22

707.29

713.54

719.18

724.57

734.11

738.41

745.08

749.66

756.06

762.15

768.55

773.13

780.44

784.40

794.77

798.12

804.82

810.92

817.02

824.94

830.12

837.13

841.40

847.50

853.59

859.69

867.00

871.88

877.98

884.38

891.08

896.26

902.97

Oil 

(gpt)

9.1

9.6

11

34.2

22.6

9.4

37

15.8

30.3

5.7

_

1.4

0.9

5.9

8.4

20.8

27.2

15.2

8.5

2

2.6

6.6

5.5

14.8

5.3

6.7

6.7

10.9

19.1

19.1

6.9

8.5

5.3

--

_

-

Oil

(L/mt)

37.95

40.03

45.87

142.61

94.24

39.20

154.29

65.89

126.35

23.77

5.84

3.75

24.60

35.03

86.74

113.42

63.38

35.45

8.34

10.84

27.52

22.94

61.72

22.10

27.94

27.94

45.45

79.65

79.65

28.77

35.45

22.10

-

_

--

Percent TOC 

LECO Rock-Eval

4.3

4.4

5.2

14.3

9.9

3.8

15.8

6.9

13.6

2.9

3

0.8

0.9

3.1

3.7

9.3

12.1

7.7

4

1.3

1.4

3

2.2

5.9

2.6

3.7

3.2

5.3

8.5

8.4

2.9

3.8

2.3

0.4

0.2

0.4

2.5

5

8.7

7.2

3.7

4

4.4

11.5

8.3

3.7

12.8

5.7

11.2

2.8

2.6

0.7

0.9

2.7

4

7.6

10

5.8

2.9

1.9

1.3

2.2

1.8

4.1

2.2

3

3.4

4.7

6.5

6.6

2.4

2.8

2.2

0.4

0.2

0.4

1.3

4.2

7.4

6.7

S1 

(mg/g)

1.73

3.99

3.25

1.4

2.15

0.63

1.91

2

1.33

1.14

0.27

1.79

2.09

0.54

1.97

1.37

2.94

1.08

0.3

0.06

0.17

0.24

0.34

0.23

1.08

0.76

0.73

0.57

0.63

2.62

0.27

0.2

0.43

0.03

0.03

0.08

0.09

0.85

0.39

1.03

S2 

(mg/g)

26.34

30.53

34.05

105.96

78.62

32.57

118.29

48.38

99.3

18.97

18.88

2.3

3.67

15.76

20.61

66.5

85.12

47.52

21.63

4.9

7.9

12.83

9.91

32.32

13.41

18.11

20.05

34.57

55.69

50.5

15.23

20.4

15.75

0.72

0.09

0.4

5.46

35.72

65.62

18.29

S3

(mg/g)

3.65

1.95

3.39

10.52

5.17

2.89

14.04

5.52

10

2.05

2.51

1.48

1.77

4.24

7.53

9.25

8.71

5.84

3.46

1.34

1.84

3.08

2.52

4.88

4.08

3.76

3.34

4.57

6.07

9.67

3.33

3.2

2.99

0.93

1.13

1.13

1.21

3.63

5.46

4.21

H-index 

(mg/g)

717

772

766

921

946

873

924

851

885

677

723

353

412

594

516

877

851

826

748

526

589

580

538

786

615

607

597

741

863

761

634

725

574

167

50

108

426

773

883

740

O-index 

(mg/g)

99

49

75

91

62

77

109

97

89

73

96

227

198

163

188

122

87

101

119

144

137

139

136

118

187

126

99

98

94

145

138

113

134

216

627

305

94

85

73

170

in the Piceance basin (Donnell, 1980). The only rich sample 

in the Coyote Wash core from the Uinta basin is a sample 

from the Mahogany zone at the top of the Parachute Creek 

Member (fig. 3). Samples from other cores in the Uinta

basin, however, generally are more organic carbon rich; 

rocks from the Mahogany zone typically contain about 10 

percent TOC (Anders and Gerrild, 1984). The richest 

samples from the Blacks Fork core are from the Tipton

F6 Green River Formation, Wyoming, Utah, and Colorado
TOC H-index O-index 13-Corg

0 10 20 30 200 600 1000 0 500 1000 -35 -30 -25

100-

200-

Z 300-

400-

500-

600

BF

100-

200-

300-

400-

500-

600

\

\

100

200-

300-

400-

500

600

100

200-

300-

400-

500-

600

BF

\

TOC H-index O-index 13-Corg

0 10 20 30 200 600 1000 0 500 1000 -35 -30 -25

200

300-

400-

500-

600-

700-

800

200

300-

400-

500-

600-

700-

800

200

300-

400-

500-

600

700-

800-

200

300

400

500

600

700-

800

TOC H-index O-index 13-Corg

0 10 20 30 200 600 1000 0 500 1000 -35 -30 -25

500 500

600-

700-

800-

900-

1000-

1100 1100- 1100

CW

500

600-

700-

800

900

1000

1100

CW

L

Figure 3. Plots of depth versus percent total organic carbon (TOC), hydrogen index 

(H-index, in milligrams of hydrocarbons per gram organic carbon), oxygen index 

(O-index, in milligrams of carbon dioxide per gram organic carbon), and 813C organic 

carbon (13-Corg, in per mil difference relative to PDB marine carbonate standard) in 

samples of Green River Formation from Blacks Fork core (BF), Green River basin, 

Wyoming; 01A core, Piceance basin, Colorado; and Coyote Wash core (CW), Uinta 

basin, Utah. Numbers (1, 2, and 3) to right of each panel refer to three lake stages of 

Turtle (this volume).

Characteristics of Organic Matter in Oil Shale F7
1,000

800

600

400

200

50 100 150 200 

0-INDEX

,uuu

800

600

400

200

0 - 

I- ( L; | I I I I [ I I I I | I I I I

. y

_ / 0 00 _

°/ °0 

/ °

-

-

-

^- "

-

-

I I I I I I I I I I I I I I I I

0 50 100 150 20

0-INDEX

1,000

800

600

400

200

50 100 150 200 

O-INDEX

1,000 i i i i | i i i i | i i i i | i i i r

800

600

X

LLJ 

Q

z

400

200

50 100 

0-INDEX

150 200

Figure 4. Plots of oxygen index (in milligrams of carbon dioxide per gram organic carbon) versus hydrogen index 

(in milligrams of hydrocarbons per gram organic carbon) in samples of Green River Formation. Curves marked I, II, 

and 111 represent thermal maturation pathways of type I, type II, and type III kerogen, respectively (Espitali£ and 

others, 1977). A, Blacks Fork core. B, 01A core. C, Coyote Wash core. D, WOSCO core. Arrows to right of the plots 

for Blacks Fork, 01 A, and WOSCO cores represent points having oxygen-index values of more than 200 

(tables 1, 2).

Shale and Laney Members, both of which were deposited in 

fresher deeper waters than the intervening Wilkins Peak 

Member, which was deposited during a shallower, much 

more areally restricted phase of Eocene Lake Gosiute 

(Eugster and Hardie, 1975; see summary by Turtle, this 

volume).

Values of percent TOC in the 01A core tend to 

increase above 400 m. Some of the lower TOC values 

between about 420 and 680 m are due to dilution of oil shale 

with nahcolite and halite. High values of HI throughout the 

saline zone (lake stage 2) indicate that the diluted organic 

matter is still hydrogen-rich, type n kerogen. The nahcolite

F8 Green River Formation, Wyoming, Utah, and Colorado
3UU

250

200

150

100

50

n

J I

^j-

£_

_

^-

Z

^~

Z

^-

#

1 1 | 1 1 M 1 1 M | 1 M 1 | M 1 1 1 1 II | II 1 1 | 1 P L

 Z

x ~Z

K ° -

_Q& 

£ £v ° X ~

A^ GX x ^

i cd$& ° o -

x^x ° A~

.yv x 

H M i \°\ \ if if 1 i i i i 1 1 1 1 i 1 1 i 1 i ii 1 1 i 1 1 r

0 5 10 15 20 25 30 35 40 

TOC, IN PERCENT

w

4UU

350

300

250

200

150

100

50

0

4 I II MM MM I I M | M

EEEE o

r o°

E- oj* +^

E- ,0^

i&Vri I 1 1 ii MM i M 1 1 i i

II I I I I MM II !- o -

1

O ° "=

-E

-E

-E

B-E

ii i i i i i i i i ii F

10 15 20 25 30 35 

TOC, IN PERCENT

40

1,000

800

X 600

01

Q

400

200

1 1 II 1 1 M 1 1 Ul 1 1 1 M 1 1

A ' * +' + '

11 H i ii M i itpi i rrr

11

0 5 10 15 20 25 30 35 40 

TOC, IN PERCENT

O 01A

A Blacks Fork

+ WOSCO

 Coyote Wash

x Colorado 1

Figure 5. Plots of pyrolysis data for samples of Green River 

Formation. Data are from tables 1 and 2. A, Total organic 

carbon (TOC) versus Fischer assay oil yield. B, TOC versus 

(ST+SJ) pyrolytic hydrocarbon yield from Rock-Eval pyrolysis 

(milligrams of hydrocarbons per gram of sample). C, TOC 

versus hydrogen index from Rock-Eval pyrolysis (milligrams of 

hydrocarbons per gram organic carbon).

300-

200-

100 200

S1+S2

300

300-

100-

01A

BLACKS FORK 

WOSCO 

COYOTE WASH

1 00 200

S1+S2

300 400

Figure 6. Plots of pyrolysis (S.,+S2) yield (milligrams of 

hydrocarbons per gram of sample) versus Fischer assay oil 

yield (liters per metric ton). A, 69 samples of oil shale of 

different ages and from different geologic environments 

(modified from Dyni, 1988). B, Green River Formation (data 

from tables 1 and 2).

also has had a large effect on the oxygen index (figs. 3,45) 

because of the evolution of CO2 from this sodium carbonate 

mineral during pyrolysis. Most carbonate minerals evolve 

CO2 at temperatures much higher than that reached during 

Rock-Eva! pyrolysis (550 °C). For example, the Green 

River Formation, particularly in the Piceance basin, 

contains abundant calcium, magnesium, and iron carbonate 

minerals (Desborough, 1978), but evolution of CO2 from 

these minerals apparently does not affect the S3 peak as 

judged by the low OI values outside the saline zone (fig. 3); 

however, nahcolite in samples from lake stage 2 in the 01A 

core must be contributing significant amounts of CO2 in 

order to explain the unusually high OI values. One sample 

from the Blacks Fork core also has an unusually high OI 

value (1,800; table 1, fig. 3), which is probably due to 

evolution of CO2 from sodium carbonate minerals (trona

Characteristics of Organic Matter in Oil Shale F9
and shortite) in this sample (Tuttle, this volume). 

Apparently no sodium carbonate or other saline minerals 

were deposited in the Uinta basin during lake stage 2 of 

Eocene Lake Uinta (Johnson, 1985). Very high hydrogen 

indices and very low oxygen indices for all samples from 

the 01A core outside lake stage 2 (figs. 3, 45) indicate that 

the organic matter in these samples is very well preserved 

algal (type I) organic matter with little oxidation.

Some samples from the other two basins have 

hydrogen indices that are as high as those from the Piceance 

basin, but their much wider range of values of both HI and 

OI (fig. 4) suggests that they may contain varying mixtures 

of algal (type I) and terrestrial (type IE) organic matter. 

Alternatively, oxidation of organic matter also can produce 

a pyrolysis response equivalent to organic matter of type HI 

composition (Tissot and Welte, 1984).

Additional evidence for organic-matter type is 

provided by gas chromatograms of the soluble organic 

fraction (figs. 7-9). Gas chromatograms of saturatedhydrocarbon fractions of immature Green River oil shale 

typically are characterized by a strong odd-carbon-number 

predominance and a bimodal distribution, with rc-C17 

dominating the rc-alkanes made up of less than 20 carbon 

atoms, and higher molecular weight hydrocarbons (>20 

carbon atoms) dominated by C27, C28, and C29 steranes, C30 

hopane, and J3-carotene (figs. 1A, 8). The chain isoprenoids 

in the intermediate-weight hydrocarbons generally contain 

more phytane than pristane (figs. 1'A, 8), characteristic of a 

highly reducing environment. The predominance of n-Cl7 in 

the rc-alkanes is particularly characteristic of Green River oil 

shale (Tissot and others, 1978), and many aquatic algae 

produce hydrocarbons in which n-Cl7 is the dominant 

hydrocarbon (see, for example, Clark and Blumer, 1967). 

Anders and Robinson (1973) concluded that most of the 

principal isoprenoid compounds (pristane, phytane, steranes, and pentacyclic triterpanes) and |3-carotene in the 

Colorado no. 1 core do not come from thermal cracking of 

kerogen but are the chemical, catalytic, and microbial 

hydrolysis products of lipids generated during the beginning 

phase of diagenesis. Normal alkanes of high molecular 

weight (C25-C33 range) and with odd-carbon predominance 

that are characteristic of higher plant waxes of terrestrial 

origin (Simoneit, 1978; Tissot and Welte, 1984) have been 

identified in Green River oil shale (Robinson and others, 

1965; Tissot and others, 1978), but their peaks would be 

masked by those of the steranes and pentacyclic triterpanes 

(mostly hopane derivatives) that are common in Green 

River oil shale (Anderson and others, 1965; Tissot and 

Welte, 1984).

The overwhelming predominance of C17 normal 

alkanes shows that most organic matter in Green River oil 

shale is of algal origin (see review by Tissot and Welte, 

1984). The lower HI values in samples from cores from the

Ph

BF-4

Depth 98m 

TOC 6.8 

HI 810 

Lake stage 3

Pr

23

JLlJjllUiL-~-

Retention Time

Ph

0) 

CO

c 

o

Q.

0) 

CC

BF-23 

Depth 353m 

TOC 8.7 

HI 694 

Lake stage 2

17

Pr 19

21 23

25 27

JLilL-i 

Retention Time

L

D 

5

5

i 

n

D

C

5

j

D_»

D 

3

BF-31

Depth 464m

TOC 12.0

HI 809 Ph

Lake stage 1

Pr

V, . .. .i. LLJLL (ML

21

19

JM. * .J .^

23

ML .i

25

1

27

C

4-jJ _____

Retention Time

Figure 7. Gas chromatograms of saturated-hydrocarbon 

fraction of samples from three depths in Blacks Fork (BF) core. 

Location of core hole shown in figure 1. Selected peaks are 

labeled by carbon number. Ph, phytane; Pr, pristane.

F10 Green River Formation, Wyoming, Utah, and Colorado
Steranes and 

triterpanes

CD 

W

c 

o

Q. 

(/> 

CD

rr

l_

o

t>

CD *- 

CD

Q

Ph

CW-26 

Depth 681m 

TOC 29 percent 

HI 993 

Lake stage 3

'18

Pr

17

Retention Time

Figure 8. Gas chromatogram of saturated-hydrocarbon fraction of a sample from the Mahogany zone in Coyote Wash (CW) core. 

Location of core hole shown in figure 1. Selected peaks are labeled by carbon number. Ph, phytane; Pr, pristane; P, beta carotene.

Uinta and Green River basins thus probably are the result of 

greater oxidation of type I (algal) organic matter rather than 

dilution by type IQ (terrestrial) organic matter.

Another characteristic of gas chromatograms of 

Green River oil shale is the change with depth from a 

bimodal distribution of hydrocarbons having an odd-carbon 

predominance (fig. 1A) to a bell-shaped distribution 

centered on C21-C23 and a loss of odd-carbon predominance 

(fig. 1C), Anders and Robinson (1973) concluded that 

development of this bell-shaped distribution of normal 

alkanes in the Colorado no. 1 core was due to increasing 

thermal maturity with depth; however, subsequent thermal 

maturity measurements in samples from this core (RockEval, Tmax, and vitrinite reflectance) indicate that there is 

no significant difference hi level of maturity between

shallow samples having a bimodal n-alkane distribution and 

deeper samples having a bell-shaped n-alkane distribution. 

Perhaps the change in n-alkane distribution is the result of 

increasing microbial degradation with increasing age, as 

suggested by Tissot and others (1978). In addition, there is 

always the possibility of contamination with core-box wax, 

which usually contains rc-alkanes in the range of C20-C24; 

however, contamination would not explain the progressive 

changes with depth.

Carbon Isotope Data

The sources of organic matter in lake sediments may 

include planktonic algal debris, rooted aquatic macrophytes,

Characteristics of Organic Matter in Oil Shale F11
Depth 545m

C17

"XjJ

Depth 659m

Depth 805m

Retention Time Retention Time Retention Time

Figure 9. Gas chromatograms of saturated-hydrocarbon fraction of samples from three depths in WOSCO EX-1 core. Location 

of core hole shown in figure 1. Selected peaks are labeled by carbon number.

and terrestrial vegetation washed in from the drainage basin. 

Each of these sources has a carbon-isotope signature that is 

not necessarily unique (fig. 10). Probably the most unique 

isotopic signature of source organic matter is the highly 

13C-depleted character of lacustrine plankton, which has 

values of 813C generally in the low negative 30's (fig. 10). 

The sedimentary signature of 13C-depleted algal debris is 

illustrated by the algal-rich end of the distribution of 13C 

values in surface sediments of Minnesota lakes (fig. 10) 

Dean, unpublished data). Within this framework of 813C 

values (fig. 10) in modern organic matter, the carbonisotope signature of organic matter in Green River oil shales 

clearly is dominated by algal debris (type I and n organic 

matter).

Changes in the carbon-isotope composition of 

sedimentary organic matter with time in a lake system (fig. 

3) can be related to changes in source of the organic matter, 

changes in climate and associated environmental conditions 

in the lake, particularly temperature, and (or) changes in 

organic productivity. For example, the steady increase in 

values of 813C through the Parachute Creek Member in the 

01A core and through the Wilkins Peak Member in the 

Blacks Fork core may be due to an influx of more 

13C-enriched terrestrial organic matter, higher temperature 

and a lower pCO2 in the lake water, or an increase in 

productivity. Each of these possibilities needs to be 

explored.

A plot of 813C versus some other indicator of organicmatter source, such as HI, can provide a means of separating 

source and environmental effects in carbon-isotope data (for 

example, Dean and others, 1986). On a plot of 813C versus 

HI, organic matter that is a mixture of terrestrial and algal 

origin shows as a negative slope reflecting 13C depletion

BF 

01A

 CW

Algal Oxidized

GREEN RIVER 

FORMATION

Surface sediments,

Minnesota lakes

C-4 plants

C-3 plants

Lake plankton

Peat

Lake sediments

I \\ I I "I I

-40 -30 -20 

C(%c,PDB)

| I 

-10

I I I

Figure 10. Range of 513C values in lake sediments, peat, lake 

plankton, C-3 plants, C-4 plants, Minnesota lake sediments, 

and samples of the Green River Formation. Data for lake 

sediments, peat, lake plankton, and C-3 and C-4 plants were 

compiled by Deines (1980); data for Minnesota sediments are 

from Dean (unpublished); data for Green River Formation are 

from table 1.

and hydrogen richness of algal organic matter and 13C 

enrichment and low hydrogen content of terrestrial organic 

matter. If all of the organic matter is of algal origin but has 

undergone varying degrees of oxidation or decomposition,

F12 Green River Formation, Wyoming, Utah, and Colorado
-24

m -26-

Q

Q_

J8 -28 -

E> 

o

O

CO

-32-

-34

200

B

400 600 800 

H-INDEX

1,000

-24

Q 

Q_

-26-

-28-

O ° -*f "JL

CO

to -32- 

-34

0

200

C

400 600 800 1,000 

H-INDEX

-24

CD Q 

o_

-26-

-28-

E>

° -i O J

CO

to -32-

-34

200

I

400 600 

H-INDEX

800 1,000

Figure 11. Plots of hydrogen index (in milligrams of 

hydrocarbons per gram organic carbon) versus 513C of organic 

matter in samples of Green River Formation. A, Blacks Fork 

core. B, 01A core. C, Coyote Wash core. Location of core 

holes shown in figure 1.

then an ffl-613C plot shows a large variation in values of HI 

and low and uniform values of 813C. If all values of HI are 

high, and values of 813C vary somewhat, then all of the 

organic matter was algal, but some environmental changes 

probably have altered the isotopic composition of that 

organic matter. The isotopic fractionation between algal 

cells and dissolved CO2 is highly dependent on pCO2 in that 

less fractionation occurs at lower pCO2 (Deuser and others, 

1968; Calder and Parker, 1973). Variations in pCO2 caused 

by changes in productivity or temperature could induce 13C 

enrichment or depletion in organic carbon fixed by 

photosynthesis (Herczeg and Fairbanks, 1987).

The three data sets from this investigation show three 

distinct distributions on an ffl-613C plot (fig. 11). Data from 

the Blacks Fork core show the greatest variability in both 

513C and HI (fig. 1 IA), Values of HI decrease and values of 

813C increase throughout the shallow-water Wilkins Peak 

Member so that the most 13C-enriched, hydrogen-poor 

organic matter is in the upper part of the Wilkins Peak (fig. 

3). We interpret this distribution as representing dilution of 

algal organic matter with more terrestrial organic matter 

(negative slope on HI-813C plot, fig. IIA). The 813C value 

of-25.7 (table 1, fig. 3) is for a sample containing shortite 

and trona and, therefore, may reflect extreme environmental 

conditions (such as high alkalinity or temperature), although 

such conditions probably have not affected the isotopic 

composition of organic matter in the saline zone in the 

Piceance basin (fig. 3). More likely this 13C-enriched 

sample represents a pulse of influx of more terrestrial 

organic matter during a particularly low lake level. The 

Tipton Shale and Laney Members in the Blacks Fork core 

contain the highest values of HI and most 13C-depleted 

values of 813C, and the organic matter in these two members 

was mostly algal rich.

The plot of 813C versus HI for samples from the 

Colorado 01A core (fig. 1 IB) shows that most of the organic 

matter in this core is very hydrogen rich and very 13C 

depleted. There is a slight trend toward less 13C-depleted 

organic matter in the upper part of the Parachute Creek 

Member, similar to the trend observed for the Wilkins Peak 

Member in the Blacks Fork isotope data. The samples 

having lower HI values are scattered throughout the core 

(fig. 3), are not part of a systematic stratigraphic trend, and 

do not correspond to samples containing 13C-enriched 

organic matter. These observations suggest that the lower 

values of HI are due to greater oxidation of organic matter, 

perhaps during lower lake levels. If we were to apply the 

interpretation of the data for Wilkins Peak Member in the 

Blacks Fork core to the data for the Parachute Creek 

Member in the 01A core (fig. 3), the stratigraphic trend 

toward increasing values of 813C should indicate a 

progressive increase in the amount of terrestrial organic 

matter during deposition of this unit. However, this 

interpretation goes against the organic-carbon and HI data, 

which suggest that production of hydrogen-rich algal

Characteristics of Organic Matter in Oil Shale F13
organic matter was high throughout the Parachute Creek 

Member and, if anything, increased through the unit to 

culminate in the accumulation of very rich algal organic 

matter in the Mahogany zone (fig. 3). If the enrichment of 

13C in the Parachute Creek were due to some environmental 

factor such as increased alkalinity or temperature, then we 

would expect the maximum values to be within the saline 

zone where presumably the alkalinity and temperature of 

the water were greatest.

Increase in the richness of oil shale upward within the 

Parachute Creek Member suggests that organic productivity 

in the lake increased, which, in turn, caused an increase in 

the rate of burial of 13C-depleted organic carbon. This 

increase in burial rate culminated during deposition of the 

extremely rich Mahogany zone, which represents maximum 

expansion of Lake Uinta. Rapid photosynthetic removal of 

13C-depleted organic carbon during such a burial event 

would leave the surface-water carbon reservoir enriched in 

13C, as was demonstrated for increased organic carbon 

burial as a result of eutrophication of Lake Greifen, 

Switzerland (McKenzie, 1985). Such an event certainly 

would have affected the isotopic composition of any 

carbonate minerals precipitated from the lake water and 

may have affected the isotopic composition of organic 

carbon. Unfortunately because most of the carbonate 

minerals in the Green River Formation are diagenetic 

(Desborough, 1978), the carbon isotopic composition of 

carbonate minerals is dominated by a diagenetic signal (J.K. 

Pitman, U.S. Geological Survey, written commun., 1990). 

Values of 813C of carbonate minerals become heavier with 

depth in the 01A core rather than lighter, as do values of 

813C of the organic matter (fig. 3), presumably due to a 

greater influence of methanogenesis and fermentation with 

increasing burial depth (J.K. Pitman, written commun., 

1990).

Scholle and Arthur (1980) and Arthur and others 

(1985) showed that periods of accumulation of 13C-depleted 

marine carbonate rocks throughout the Phanerozoic can be 

correlated with periods of rapid burial of marine organic 

matter (so-called oceanic anoxic events). During periods of 

extremely high burial rate of organic carbon, such as 

occurred at the Cretaceous Cenomanian/Turonian (C/T) 

boundary, the isotopic composition of organic carbon fixed 

by photosynthesis also can be enriched in 13C by as much as 

4-6%o on a global scale (Pratt and Threlkeld, 1984; Arthur 

and others, 1988). If an organic-carbon burial event, such as 

the C/T boundary event, can alter the isotopic composition 

of the surface-water carbon reservoir of the world ocean, 

surely a burial event such as that recorded in the Mahogany 

zone of the Green River Formation would have altered the 

isotopic composition of the much smaller surface-water 

reservoir of a lake, even a lake as large as Lake Uinta. We 

conclude that the increase in 813C values at the top of the 

Parachute Creek Member in the 01A core, and perhaps

some of the increase in 813C values at the top of the Wilkins 

Peak Member in the Blacks Fork core, is due to extreme 13C 

enrichment of the entire surface-water carbon reservoir.

Most of the data in the ffl-8l3C plot for the Coyote 

Wash core (fig. 11C) show little variation in values of 813C 

but considerable variation in values of HI. The small 

variation of 813C values suggests that there was little 

variation in the source of organic matter, and therefore the 

large variation in values of HI probably is due to different 

degrees of decomposition of algal organic matter; however, 

the two samples having the highest values of HI also are the 

most 13C enriched. One of these samples is from the 

Mahogany zone and the other is from an organic-carbonrich unit below the Mahogany zone (fig. 3). These two 

anomalously high 813C values, relative to other values in the 

core, suggest that organic carbon fixed by photosynthesis in 

the Uinta basin, as in the Piceance basin, was at times 

affected by extreme 13C-enrichment of the surface-water 

reservoir in Lake Uinta.

CONCLUSIONS

The richest oil shale samples from the three cores of 

the Green River Formation studied are from the Piceance 

basin in Colorado. These samples are from a core in the 

depocenter of the saline fades in the Piceance basin and 

would be even richer if the kerogen was not diluted by 

evaporite minerals (nahcolite and halite). The organic 

matter in the Piceance basin is very well preserved and little 

oxidized. The organic matter in samples of oil shale that 

accumulated in shallower water in the Uinta and Green 

River basins is lower in amount and is more highly 

oxidized.

Variations in source of organic matter, degree of 

preservation of that organic matter, and environmental 

conditions during formation and accumulation of organic 

matter commonly can be detected by consideration of type 

(such as by using pyrolysis hydrogen index, HI) and carbon 

isotopic composition (813C) of the organic matter. On a plot 

of HI versus 813C, a mixture of terrestrial and algal organic 

matter shows as a negative slope reflecting 13C-depletion 

and hydrogen-richness of algal organic matter and 13Cdepletion and low hydrogen content of terrestrial organic 

matter. An example of such a mixture is the organic matter 

in the shallow-water Wilkins Peak Member of the Green 

River Formation in Wyoming.

If all of the organic matter is of algal origin but has 

undergone varying degrees of oxidation or decomposition, 

then an HI-513C plot will show a large variation in HI but 

low and uniform values of 813C. Such a distribution is 

exemplified by most samples of oil shale from the Coyote 

Wash core from the Uinta basin. Algal organic matter that 

is well preserved will be very hydrogen rich and very 13C 

depleted. Much of the organic matter in the Piceance basin 

core has these characteristics; however, there is an unusual

F14 Green River Formation, Wyoming, Utah, and Colorado
enrichment in 13C toward the top of the Parachute Creek 

Member of the Green River Formation in the Piceance 

basin, culminating in the extremely algal rich organic matter 

of the Mahogany zone. This 13C enrichment is interpreted as 

due to rapid burial of 13C-depleted organic matter to the 

point that the entire surface-water carbon reservoir became 

13C enriched.

REFERENCES CITED

Anders, D.E., and Gerrild, P.M., 1984, Hydrocarbon generation in 

lacustrine rocks of Tertiary age, Uinta Basin, Utah Organic carbon, pyrolysis yield, and light hydrocarbons, in 

Woodward, J., Meissner, F.F., and Clayton, J.L., eds., 

Hydrocarbon source rocks of the greater Rocky Mountain 

region: Rocky Mountain Association of Geologists, Denver, 

p. 513-529.

Anders, D.E., and Robinson, W.E., 1973, Geochemical aspects of 

the saturated hydrocarbon constituents of Green River oil 

shale Colorado No. 1 core: U.S. Bureau of Mines, Report 

of Investigations 7737, 23 p.

Anderson, P.C., Gardner, P.M., Whitehead, E.V., Anders, D.E., 

and Robinson, W.E., 1969, The isolation of steranes from 

Green River oil shale: Geochimica et Cosmochimica Acta, 

v. 33, p. 1304-1307.

Arthur, M.A., Dean, W.E., and Pratt, L.M., 1988, Geochemical 

and climatic effects of increased marine organic carbon 

burial at the Cenomanian/Turonian boundary: Nature, v. 

335, p. 714-717.

Arthur, M.A., Dean, W.E., and Schlanger, S.O., 1985, Variations 

in the global carbon cycle during the Cretaceous related to 

climate, volcanism, and changes in atmospheric CO2, in 

Sundquist, E.T., and Broecker, W.S., eds., The carbon cycle 

and atmospheric CO2 Natural variations Archean to 

Present: American Geophysical Union Monograph 32, 

p. 504-529.

Calder, J.A., and Parker, P.L., 1973, Geochemical implications of 

induced changes in 13C fractionation by blue-green algae: 

Geochimica et Cosmochimica Acta, v. 37, p. 133-140.

Clark, R., and Blumer, M., 1967, Distribution of n-paraffins in 

marine organisms and sediment: Limnology and 

Oceanography, v. 12, p. 79-87.

Dean, W.E., Arthur, M.A., and Claypool, G.E., 1986, Depletion of 

13C in Cretaceous marine organic matter Source, diagenetic, or environmental signal?: Marine Geology, v. 70, 

p. 119-157.

Deines, P., 1980, The isotopic composition of reduced carbon, in 

Fritz, P., and Fontes, J.C., eds., Handbook of environmental 

isotope geochemistry, v. 1, The terrestrial environment: 

Amsterdam, Elsevier, p. 329-406.

Desborough, G.E., 1978, A biogenic-chemical stratified lake 

model for the origin of oil shale of the Green River 

Formation An alternative to the playa-lake model: 

Geological Society of America Bulletin, v. 89, p. 961-971.

Desborough, G.E., and Pitman, J.R., 1974, Significance of applied 

mineralogy to oil shale in the upper part of the Parachute 

Creek Member of the Green River Formation, Piceance 

Creek basin, Colorado, in Murray, D.K., ed., Energy

resources of the Piceance Creek basin, Colorado: Rocky 

Mountain Association of Geologists, Field Conference, 

25th, Denver, p. 81-89.

Deuser, W.G., Degens, E.T., and Guillard, R.R.L., 1968, Carbon 

isotope relationships between plankton and sea water: 

Geochimica et Cosmochimica Acta, v. 32, p. 657-660.

Donnell, J.R., 1980, Western United States oil-shale resources and 

geology, m Synthetic fuels from oil shale: Chicago, Institute 

for Gas Technology, p. 17-19.

Dyni, J.R., 1988, Review of the geology and oil-shale resources of 

the tripolitic oil-shale deposits of Sicily, Italy: U.S. 

Geological Survey Open-File Report 88-270, 23 p.

Dyni, J.R., Anders, D.E., Rex, R.C., Jr., 1990, Comparison of 

hydroretorting, Fischer assay, and Rock-Eval analyses of 

some World oil shales, in Proceedings of the 1989 Eastern 

Oil Shale Symposium: University of Kentucky, Institute for 

Mining and Minerals Research, p. 270-286.

Espitali6, J., LaPorte, J.L., Madec, M., Marquis, F., Leplat, P., 

Paulet, J., and Boutefeu, A., 1977, Methode rapide de 

caracterisation des roches m6res de leur potential petrolier 

et de leur degr6 d'evolution: Revu Institut Francais du 

Petrole, v. 32, p. 23-42.

Eugster, H.P., and Hardie, L.A., 1975, Sedimentation in an ancient 

playa-lake complex The Wilkins Peak Member of the 

Green River Formation of Wyoming: Geological Society of 

America Bulletin, v. 86, p. 319-334.

Herczeg, A.L., and Fairbanks, R.G., 1987, Anomalous carbon 

isotope fractionation between atmospheric CO2 and 

dissolved inorganic carbon induced by intense 

photosynthesis: Geochimica et Cosmochimica Acta, v. 51, 

p. 895-899.

Johnson, R.C., 1985, Early Cenozoic history of the Uinta and 

Piceance Creek basins, Utah and Colorado, with special 

reference to the development of Eocene Lake Uinta, in 

Flores, R.M., and Kaplan, S.S., eds., Cenozoic paleogeography of the western United States: Denver, Society of 

Economic Paleontologists and Mineralogists, Rocky 

Mountain Symposium 3, p. 247-276.

McKenzie, J.A., 1985, Carbon isotopes and productivity in the 

lacustrine and marine environments, in Stumm, W., ed., 

Chemical processes in lakes: New York, WileyInterscience, p. 99-118.

Pratt, L.M., and Threlkeld, C.N., 1984, Stratigraphic significance 

of 13C/12C ratios in mid-Cretaceous rocks of the western 

interior, U.S.A., in Scott, D.F., and Glass, D.J., eds., The 

Mesozoic of middle North America: Canadian Society of 

Petroleum Geology Memoir 9, p. 305-312.

Robinson, W.E., and Cook, G.L., 1975, Compositional variations 

of organic material from Green River oil shale WOSCO 

EX-1 core (Utah): U.S. Bureau of Mines, Report of 

Investigations 8017, 40 p.

Robinson, W.E., Cummins, J.J., and Dinneen, G.U., 1965, 

Changes in Green River oil-shale paraffins with depth: 

Geochimica et Cosmochimica Acta, v. 29, p. 249-258.

Scholle, P.A., and Arthur, M.A., 1980, Carbon isotope fluctuations 

in Cretaceous pelagic limestones Potential Stratigraphic 

and petroleum exploration tool: American Association of 

Petroleum Geologists Bulletin, v. 64, p. 67-87.

Characteristics of Organic Matter in Oil Shale F15
Simoneit, B.R.T., 1978, The organic chemistry of marine Tissot, B., Deroo, G., and Hood, A., 1978, Geochemical study of

sediments, in Riley, J.P., and Chester, R., eds., Chemical the Uinta Basin Formation of petroleum from the Green

oceanography (v. 7): New York, Academic Press, River Formation: Geochimica et Cosmochimica Acta, v. 42,

p. 233-311. p. 1469-1485.

Smith, J.W., Trudell, L.G., and Dana, G.F., 1968, Oil yields of

Green River oil shale from Colorado corehole No. 1: U.S. Tissot, B.G., and Welte, D.H., 1984, Petroleum formation and

Bureau of Mines, Report of Investigations 7071, 28 p. occurrence (2nd ed.): New York, Springer-Verlag, 699 p.

F16 Green River Formation, Wyoming, Utah, and Colorado
Chapter G

Petrography of Iron Sulfide Minerals in 

the Green River Formation of Wyoming, 

Utah, and Colorado

ByMICHELELTUTTLE

U.S. GEOLOGICAL SURVEY BULLETIN 1973

GEOCHEMICAL, BIOGEOCHEMICAL, AND SEDIMENTOLOGICAL STUDIES OF THE 

GREEN RIVER FORMATION, WYOMING, UTAH, AND COLORADO
CONTENTS

Abstract Gl

Introduction Gl

Methods G2

Results G3

Blacks Fork core G3 

01A core G3 

Coyote Wash core G4

Discussion G5

Pyrite formation G8 

Pyrrhotite formation G10 

Marcasite formation Gil

Conclusions Gil

References cited G12

FIGURES

1. Map showing Laramide uplifts and basins in parts of Wyoming, Utah, Colorado, 

and Idaho G2

2. Chart showing stratigraphy of cores used in study G3 

3-5. Diagrams showing iron sulfide mineralogy and morphology for samples of 

Green River Formation from:

3. Blacks Fork core G5

4. 01A core G6

5. Coyote Wash core G7

6. Photomicrographs showing typical sulfide-mineral morphology G8

7. Chart showing paragenetic sequence of sulfide mineral formation G10

8. Eh-pS2~ stability field diagrams for pyrite, pyrrhotite, siderite, and 

hematite Gil

TABLE

1. Summary of lake stages represented by the samples in this study G4

Contents
Petrography of Iron Sulfide Minerals in the Green River 

Formation of Wyoming, Utah, and Colorado

By Michele L. Tuttle

Abstract

The Paleogene Green River Formation contains three 

iron sulfide minerals that are easily examined using reflectedlight microscopy: pyrite, pyrrhotite, and marcasite. At least 

four generations of pyrite were identified. Framboidal pyrite 

formed in organic stringers by the sulfidization of iron oxide 

minerals that coat clay minerals and were very reactive to H2S 

(episode 1). Euhedral and anhedral pyrite grains and pyrite 

infilling framboids formed when less reactive iron oxide 

minerals were dissolved, possibly during iron reduction by 

bacteria (episode 2).

Pyrrhotite then formed from unstable iron sulfide phases 

that were never converted to pyrite because of the extremely 

reducing conditions in the sediments (episode 3). Later, iron 

loss from pyrrhotite and partial oxidation of sulfur led to 

pseudomorphic replacement of pyrrhotite by marcasite 

(episode 4). Subsequent alteration of some marcasite to the 

more stable disulfide phase pyrite produced the unusual 

bladed pyrite.

INTRODUCTION

The Paleogene Green River Formation, deposited in 

ancient Lakes Uinta and Gosiute in Utah, Colorado, and 

Wyoming, contains iron sulfide minerals that have highly 

unusual mineralogy and morphology. Iron disulfide 

minerals are the most common sulfide minerals; of these, 

pyrite (cubic FeS2) (described in Milton and Eugster, 1959; 

Pabst, 1970; Cole and others, 1978; Cole and Picard, 1981) 

is more abundant than marcasite (orthorhombic FeS2) (Milton and Eugster, 1959; Cole and others, 1978; Cole and 

Picard, 1981). Some pyrite has an unusual bladed 

morphology (Cole, 1975; Boyer and Cole, 1983; Dyni,

1983). Pyrrhotite (monoclinic Fe^S) is locally abundant 

(Milton and Eugster, 1959; Cole and others, 1978; Cole and 

Picard, 1981; Melchior and others, 1982), and wurtzite 

(hexagonal ZnS) is rare (Milton and Eugster, 1959).

The presence of marcasite in unmetamorphosed 

lacustrine rocks is unknown outside the Green River 

Formation. Marcasite commonly forms in sedimentary 

environments where the pH is less than 4 (Goldhaber and 

Kaplan, 1974; Murowchick and Barnes, 1986), 

environments much more acidic than those postulated for 

pore waters in Green River sediments (pH~10; Bradley 

and Eugster, 1969). Authigenic pyrrhotite is rare in 

sedimentary environments and, similar to marcasite, has not 

been reported in unmetamorphosed lacustrine rocks other 

than those in the Green River Formation.

In this report, I discuss the petrography (mineralogy 

and morphology) of iron sulfide minerals in core material of 

the Green River Formation from the centers of the three 

depositional basins occupied by the Green River lakes (fig. 

1). The petrographic characteristics indicate a complex 

paragenetic sequence of iron sulfide formation. 

Identification of four distinct episodes allows a separation, 

in time, of processes affecting sulfur in the Green River 

Formation. Determining the paragenetic sequence is crucial 

for guiding interpretation of sulfur chemistry and isotopy 

(Tuttle and Goldhaber, this volume), which in turn is used to 

reconstruct the depositional and diagenetic history of the 

Green River Formation, the primary goal of the studies 

presented in this volume.

Acknowledgments, This report has benefited greatly 

from the thorough reviews of Richard Reynolds, Neil 

Fishman, and Robert Zielinski. The work on which this 

chapter is based was funded by the U.S. Geological Survey 

Evolution of Sedimentary Basins Program.

Petrography of Iron Sulfide Minerals G1
113 106°

ABSAROKA VOLCANIC FIELD

50 100 KILOMETERS

EXPLANATION

L Present-day limit of Tertiary rocks Q

Upper Cretaceous outcrop

Present-day outcrop of Lower 

Cretaceous through Cambrian rocks

Present-day outcrop of Precambrian rocks

Figure 1. Map showing Laramide uplifts and basins in parts of Wyoming, Utah, Colorado, and 

Idaho. Core holes (solid circles): 1, Blacks Fork no.1 core, greater Green River basin, Wyoming; 2, 

01A core, Piceance basin, Colorado; 3, Coyote Wash core, Uinta basin, Utah. Modified from 

Johnson (1985).

METHODS

A hierarchical, unbalanced analysis of variance 

sampling design was used to assure optimal coverage of 

cores given the limitations on the number of samples 

analyzed. Polished thin sections were prepared from plugs 

of 114 samples: 41 (39 petrographically examined) samples 

from Piceance basin, Colorado (01A core), 35 samples from

the greater Green River basin, Wyoming (Blacks Fork 

core), and 38 samples from the Uinta basin, Utah (Coyote 

Wash core). Each original sample was about 3 cm long and 

lithologically homogeneous. Locations of the three depocenter cores are shown in figure 1. The stratigraphy of rocks 

in each core and the division of each core into the three 

major lake stages are shown in figure 2. The cores, samples, 

and lake stages are described in detail in Turtle (this

G2 Green River Formation, Wyoming, Utah, and Colorado
460

Coyote Wash

669- 

681 ~

710-

950 

1,055

2 I'*-' 

O

a E

I1 !

220

273

291

329 -}

01A Blacks Fork

-706

774

55

460

Lake 

stage 3

Lake 

stage 2

stage 1

0 50 100 METERS

Figure 2. Stratigraphy of cores from depocenters of greater 

Green River (Blacks Fork), Piceance (01 A), and Uinta (Coyote 

Wash) basins. Mz, Mahogany zone; -Mahogany bed. Location 

of core holes shown in figure 1.

volume), and lake stages (and number of samples within 

each stage examined in this study) are summarized in 

table 1.

The polished thin sections were commercially 

prepared by cutting and grinding in oil because many 

samples contain water-soluble minerals. Pyrrhotite, marcasite, and pyrite were optically identified using color, bireflectance, and anisotropy under reflected light. The relative 

sulfide abundance was visually estimated under the 

microscope. In most samples, the carbonate and silicate 

minerals are too fine grained to be identified using 

transmitted-light techniques.

1.6 weight percent. Representative examples of iron sulfide 

morphology in the three cores are shown in figure 6 and 

referred to in the discussion of each core. Wurtzite (ZnS) 

was not observed in any of the thin sections, consistent with 

its reported "microscopic" occurrence (Milton and Eugster, 

1959).

Blacks Fork Core

Pyrite is the dominant sulfide mineral in sediments of 

lake stage 1 in the Blacks Fork core (fig. 3). All stage 1 

samples contain either pyrite framboids (as much as 30 |im 

in diameter; as represented in fig. 6A) or anhedral pyrite 

grains. Fine-grained pyrite fills shell fragments in the 

deepest sample (509 m). One sample (464 m) also contains 

minor amounts of marcasite and pyrrhotite. The marcasite is 

within pyrrhotite blades as long as 400 |im.

Pyrite is the dominant sulfide in sediments of lake 

stage 2. It is commonly anhedral, although some samples 

contain euhedral pyrite cubes as much as 75 Jim on edge 

(fig. 65) and bladed pyrite as long as 500 [4,m. Bladed 

pyrrhotite was observed in six samples; the blades 

commonly are in splays (as long as 1 Jim) that cut across 

bedding (fig. 6C). Pyrrhotite in one sample (144 m depth; 

fig. 6£>) borders shortite (Na2Ca2(CO3)3) grains that, in turn, 

enclose pyrite cubes similar to that shown in figure 6B. Both 

anhedral and euhedral marcasite are associated with pyrite 

in three samples and with bladed pyrrhotite in two other 

samples (fig. 6E). No pyrite framboids were identified in 

samples of stage 2 sediments.

Sulfide minerals in lake stage 3 are predominantly 

anhedral pyrite; several samples contain euhedral and (or) 

framboidal pyrite. Euhedral marcasite was identified in 

three samples containing pyrite; one of these samples also 

contains sparse anhedral pyrrhotite. Several samples contain 

what probably is large (as much as 100 |im in diameter) 

iron-titanium-oxide aggregates or concretions that have 

been replaced by sulfide mineral(s) (fig. 6F); the sulfide 

mineral(s) in these samples is too fine grained to identify 

petrographically.

01A Core

RESULTS

Depth profiles for sulfide-mineral abundances based 

on visual inspection of thin sections are similar to profiles 

based on measured sulfide chemical abundance data (figs. 

3-5). In figures 3-5 an estimated abundance of one dot 

corresponds to less than 0.03 weight percent sulfur as 

sulfide minerals (dry-weight basis), and the highest 

estimated abundance (five dots) corresponds to greater than

Most iron sulfide minerals in lake stage 1 samples of 

the 01A core are framboids and euhedral or anhedral grains 

of pyrite (fig. 6). The framboids (fig. 6A) are as much as 30 

Jim in diameter; the euhedral and anhedral grains are 

commonly as much as 100 \im in diameter. Two samples 

contain euhedral and anhedral marcasite. Bladed textures 

were not observed.

Sediments from lake stage 2 contain framboidal, 

euhedral, and anhedral pyrite similar to that in stage 1 

sediments; framboidal pyrite is rare in the upper 100 m of

Petrography of Iron Sulfide Minerals G3
Table 1. Summary of lake stages 1,2 (2B in the Green River basin), and 3 represented by samples in this study

[Samples are from the Blacks Fork core no. 1 (BF), greater Green River basin, Wyoming; 01A core (01A), Piceance basin, 

Colorado; Coyote Wash no. 1 core (CW), Uinta basin, Utah. Number of samples in parentheses is number of samples for which 

sulfur forms and their isotopic compositions were analyzed]

Lake 

stage Green River basin

Greater 

Piceance basin Uinta basin

Fresh water 

Transgression 

High productivity 

Organics preserved 

Open-water fades

High salinity 

Regression 

Low productivity 

Organics mineralized 

Shallow-water fades

Fresh water 

Transgression 

High productivity 

Organics preserved 

Open-water fades

Moderate salinity 

Transgression 

Low productivity 

Organics preserved 

Open-water fades

High salinity 

Regression 

High productivity 

Organics preserved 

Open-water fades

Moderate-high salinity 

Transgression 

High productivity 

Organics preserved 

Open-water fades

Fresh water 

Transgression 

Low productivity 

Lake-margin fades

Number of 

samples

5 

(5)

7 

(4)

8 

(4)

High salinity 

Regression

Moderate productivity 

Organics preserved 

Open-water fades

Number of 

samples

23 

(23)

27 

(11)

21 

(15)

Low-moderate salinity 

Transgression 

High productivity 

Organics preserved 

Open-water fades

Number of 

samples

7 

(7)

7 

(5)

10 

(6)

stage 2 (the leached zone1 ). Fine-grained pyrite has replaced 

some unidentified grains similar to those in figure 6F. 

Euhedral or anhedral grains of marcasite, commonly within 

pyrrhotite blades (fig. 6£), are in the upper part of stage 2 

beds. In the 01A core pyrrhotite is exclusive to stage 2 and 

is concentrated near the top of the sequence (leached zone). 

It is present as blades in four samples, two of which also 

contain anhedral or euhedral pyrrhotite. The pyrrhotite 

blades commonly are present together in splays that cut 

bedding and are as long as 1 Jim (fig. 6Q.

Lake stage 3 samples contain pyrite and marcasite. 

Pyrite framboids and anhedral pyrite or marcasite are 

common; fine-grained replacement sulfide minerals (fig. 

6F) and pyrite blades were observed in samples from near 

the base of stage 3. The bladed texture is similar to that 

observed for pyrrhotite; pyrite and marcasite probably are 

pseudomorphs after pyrrhotite.

'The leached zone is in the upper part of the saline stage of the 

Green River Formation in the Piceance basin of Colorado and represents a 

zone leached of its water-soluble salts as evidenced by vugs, crystal 

cavities, and marlstone solution breccias (Dyni, 1981). There is no 

evidence of sulfide-mineral leaching such as iron oxide staining.

Coyote Wash Core

Pyrite is the predominant sulfide mineral in lake stage 

1 of the Coyote Wash core. It is present as anhedral and 

euhedral grains, framboids, and infilled framboids (fig. 6). 

Infilled framboids (Love and Amstutz, 1966) formed by the 

addition of pyrite to the interstices of the framboidal 

structure. Framboid infilling is illustrated in figure 6A. 

Minor amounts of marcasite are present as anhedral and 

euhedral grains.

Lake stage 2 samples contain pyrite, marcasite, and 

pyrrhotite. Two samples in the lower part of stage 2 contain 

only pyrite as anhedral and euhedral grains. Bladed pyrrhotite (fig. 6C) is the predominant sulfide mineral in two 

samples in the middle part of the stage; these samples also 

contain minor euhedral and anhedral marcasite and (or) 

pyrite. The other lake stage 2 samples contain marcasite and 

pyrite; bladed pyrrhotite may be present. Pyrite is present as 

anhedral or euhedral grains, framboids, or infilled framboids, and marcasite is present as (pither anhedral or euhedral grains. Pyrite and marcasite are typically present as 

replacement sulfide minerals within pyrrhotite blades (fig. 

6G).

In stage 3 samples, anhedral pyrite is the dominant 

sulfide mineral; framboids were not observed. Marcasite is

G4 Green River Formation, Wyoming, Utah, and Colorado
Sulfide Visual 

weight percent sulfide abundance Mineralogy and morphology

100

C/)

cc

UJ 

LLJ

Q_ 

UJ 

Q

200

300 I

400 

500 

o none

 sparse

 abundant

***

CD - *i_ 

>, 

Q_

1-A

1--F.A

1-F.E.A

1--A

1--E.S

1--A.S

1--A

2--E

1--A.S

1--SAE

1--E

1--E

1--A

2-A

1-A

1--A

1-A

1-A.E

1-R

1-A

1-A.E

1-B

1--E.A

1-E.A

1-A.L

1-EAB, L

1-S.A

1-A.F.E

1-A.F

1-A

arcasite

2-E

2-E

2-E

1-B.E

2--A

2-E

2-E.B,

0) 

O

3-A

1-B

2-B

1-B

2-B

3--B

CD 

O)

03

03

O) 

CO

03

O)

03

03

Figure 3. Iron sulfide mineralogy and morphology for samples of Green River Formation from Blacks Fork core, 

greater Green River basin, Wyoming. A, anhedral; B, blades; E, euhedral; F, framboids; L, lens; S, iron-sulfide minerals 

replacing (iron-titanium oxide?) grains. Numbers refer to relative abundance with 1 being most abundant. Location 

of core hole shown in figure 1.

common in lower beds but absent at depths shallower than 

about 620 m. The only stage 3 sample devoid of pyrite (642 

m) contains euhedral and bladed marcasite and bladed 

pyrrhotite.

DISCUSSION

Iron sulfide minerals in the Green River Formation 

formed in response to changing diagenetic conditions and

Petrography of Iron Sulfide Minerals G5
Sulfide Visual 

weight percent sulfide abundance Mineralogy and Morphology

300 

400

500

tr LUhLU

I

Q_ 

LU 

Q

600

700

0123

o None

 Sparse

i

T

 Abundant

**

CD *- 

>, 

Q_

1--A. F?

1-A. -

' --A. -

--,

1--A, E

1--A

3--A

1--A. F

1--A

1--A.E.S

1--E.A.B

1-A

1-E.S

1-S.F

1--A.L

1-F

.1=L

1--S.F

1 

1-F.A

2-A

1--A.E

1--A

1--FAE

I-A.F.S.E.L

1-E

1-S.A

1-S.A.F.F

1-A

1-E.A

1-A

1-E.F

1-A

1-F

1-F

1--F

1-A.F

arcasite

9-A

2-E

2-E. A

2-E

1-E

2-E

1-A, E

2~E,A,B,S

2-S

2-E.A

2-A.E

2-A

2-E

CD -H- ' 

" * 

o

1-A. E. B

2-E. B

3-B

1-B

o>

^ CO 

CO

QJ

03

<D 

O)

 <N 

CO

0) ^

03

O)

ro

03

Figure 4. Iron sulfide mineralogy and morphology for samples of Green River Formation from 01A core, Piceance 

basin, Colorado. A, anhedral; B, blades; E, euhedral; F, framboids; L, lens; S, iron-sulfide minerals replacing 

(iron-titanium oxide?) grains. Numbers refer to relative abundances with 1 being most abundant. Location of core hole 

shown in figure 1.

thus have different mineralogies and morphologies. The 

history of iron sulfide formation in a typical lake bed of 

ancient Lakes Uinta and Gosiute was divided into four

episodes (fig. 7) that span the periods of sediment 

deposition (episode 1) through subsequent compaction and 

possibly deeper burial (episodes 3 and 4). The paragenetic

G6 Green River Formation, Wyoming, Utah, and Colorado
Sulfide Visual 

weight percent sulfide abundance Mineralogy and Morphology

550

650 

CO

cr

UJ 

LJJ

^ 750 \

i

o_

LLI 

Q

850 

950 

1050

oNone

 Sparse

T

 Abundant

**

«

CD f *

>*

CL

1-E.L

1--A

1-A

1--A.L

P-A

1--A

1-A

1--A

1--A, F

1--E

1--E, F 

P-A

1-A

1--E.F

2-B.A

1-A F R

1-E, F

1-F

2-A. E.L

1-E^A^

--A. E.F

-A. F?

1-E. A. F

1-E, F

1-E.F

1-E, F

1-F

1-A

arcasite

2-A

1-E

-A

2-E,

2-A

2-E

2--E 

1-A

2-A

2-E

2--A

3-B,A

P-A

2-E

1-E

2-A

2-A

2-A

2-A

P-F

2-E

2-A

0)

" 4 '

0

2-E

3-A

3-B

1-B

1-B

03

O)

H co 

to 

o> 

cc

CD

O)

cc

CD 

CC

Figure 5. Iron sulfide mineralogy and morphology for samples of Green River Formation from Coyote Wash core, Uinta 

basin, Utah. A, anhedral; B, blades; E, euhedral; F, framboids; L, lens; S, iron-sulfide minerals replacing (iron-titanium 

oxide?) grains. Numbers refer to relative abundances with 1 being most abundant. Location of core hole shown in 

figure 1.

sequence was constructed using petrographic observations 

and inferences from sulfide-mineral genesis in modern 

marine sediments. The formation of iron sulfide minerals 

probably began with fine-grained, unstable iron monosulfides (FeS), which are thought to be the precursors of all 

sedimentary iron sulfide minerals (Roberts and others, 

1969; Rickard, 1974; Pyzik and Sommer, 1981). Based on 

analogy with modern marine sediments, these ephemeral

Petrography of Iron Sulfide Minerals G7
/j m _

Figure 6 (above and facing page). Photomicrographs showing typical sulfide-mineral morphology in the Green River Formation. 

Location of core hole shown in figure 1. A, B, E, F, G, reflected-light photomicrographs under oil; C, D, under air. Py, pyrite; M, 

marcasite; Po, pyrrhotite; Sh, shortite. A, Pyrite framboids and framboids in which pyrite has filled spaces between crystals. Coyote 

Wash core, 971 m. B, Euhedral pyrite cube. Blacks Fork core, 144 m. C, Splays of pyrrhotite blades. Coyote Wash core, 815 m. 

D, Pyrrhotite blades rimming shortite grain. Blacks Fork core, 144 m. E, Euhedral marcasite that may have replaced pyrrhotite. 

Blacks Fork core, 464 m. F, Sulfidized (iron-titanium oxide?) mineral grains having pyrite centers surrounded by fine-grained 

sulfide minerals. 01A core, 437 m. C, Pyrrhotite blades partly replaced by pyrite; crossed nicols. Coyote Wash core, 815 m.

FeS phases probably formed very near the sediment-water 

interface. They are not present in any of our samples nor are 

they shown in figure 7.

Pyrite Formation

Pyrite is the predominant sulfide mineral in the Green 

River Formation. In order of decreasing abundance, it is 

present as euhedral and anhedral grains, framboids, infilled 

framboids, and lenses (accumulations of pyrite cement). 

The various forms of pyrite of the Green River Formation

probably formed in a sequence similar to that observed in 

marine sediments even though the overall geochemical 

controls on pyrite formation in the two environments may 

differ.

One of the earliest crystalline iron sulfides to form in 

marine sediments is microscopic pyrite (0.5-100 [4,m) (Love 

and Amstutz, 1966) that commonly aggregates to form 

framboids (Raiswell, 1982). Raiswell associated framboidal 

morphology with sulfidization of reactive "in situ" iron in 

the sediment. Likely sources of what Raiswell called in situ 

iron are iron oxyhydroxides that coat clay-sized minerals 

and very likely contain ferrihydrite (approximately 

Fe(OH)3) and lepidocrocite (y-FeOOH) the two iron

G8 Green River Formation, Wyoming, Utah, and Colorado
250

SO

phases most reactive to reduction by H2S (Pyzik and 

Sommer, 1981; Canfield, 1989). Framboidal pyrite almost 

always is found in organic stringers (filaments) in Green 

River Formation samples. The intimate association of framboids and organic stringers is not surprising because claysize particles, and their iron oxyhydroxide coatings, and 

organic matter have similar slow settling rates through the 

water column (Anderson and others, 1985). No framboids 

were observed in organic-rich samples from stage 3 in the 

Uinta basin core (Coyote Wash). The relative rarity of clay 

minerals throughout this stage suggests that associated iron 

oxide minerals were insufficient to form framboids. In the 

model proposed here, framboids formed near the sedimentwater interface by sulfidization of iron oxide coatings on 

clay particles associated with organic stringers. This period 

of sulfide mineral generation is referred to as episode 1.

The large size of euhedral and anhedral pyrite grains 

in most Green River Formation samples (fig. 6fi) suggests 

slow growth. Raiswell (1982) hypothesized that euhedral 

pyrite in marine sediments grows after the most reactive 

local source of iron is exhausted and depends on diffusive 

input of additional iron. A similar process could have 

occurred in the Green River sediment. After the most 

reactive iron oxyhydroxide coatings were sulfidized to form 

framboids, other iron minerals in the sediments provided a 

reduced flux of dissolved iron for further sulfidization. 

Mineralogical sources of iron, listed in decreasing order of 

reactivity toward H2S, include goethite (oc-FeOOH), hematite (Fe2O3), magnetite (Fe3O4), iron-bearing clays, and iron 

silicates. The relatively large size of the pyrite grains 

indicates that the sediment never became supersaturated 

with respect to pyrite, possibly because iron was very 

slowly released during diagenetic dissolution of the above 

mineral phases. I suggest that both anhedral and euhedral 

pyrite in the Green River Formation are probably cogenetic 

and formed from iron that was slowly released from 

minerals such as magnetite (episode 2). Bacterial reduction 

of ferric iron, as discussed by Canfield (1989), may have 

been the dominant dissolution mechanism because ironreducing bacteria are known to exist in relatively saline 

sedimentary environments (11 per mil salinity) (Lovley and 

Phillips, 1986). Sediment matrix probably controlled the 

morphology of these grains. Where pyrite growth was 

unconstrained by competitive crystallization or pore-space 

volume, euhedral crystals formed; constrained growth, 

however, resulted in irregularly shaped crystals (Carstens, 

1986).

Many organic stringers contain framboids that are 

variably infilled with pyrite (fig. 6A) (framboidal infilling is 

described in Love and Amstutz, 1966). Complete infilling 

of some framboids has produced spherical grains of pyrite. 

The controls on infilling processes are difficult to assess but 

may be linked to amalgamation of distinct pyrite crystals 

within the framboids by addition of new pyrite (Love and 

Amstutz, 1966). The process of pyrite infilling framboids

Petrography of Iron Sulfide Minerals G9
Early Late

Deposition

Increased 

compaction

Burial 

Episode 1

Episode 2

Episode 3

Episode 4

Pyrite f ramboids

Anhedral / euhedral pyrite

Pyrite infill of framboids

Pyrrhotite

Marcasite replacing

Pyrrhotite 

Pyrite replacing

pyrrhotite and marcasite

Figure 7. Paragenetic sequence of sulfide-mineral formation in Green River sediments deposited in ancient Lakes Uinta and 

Gosiute.

in Green River rocks is assigned to episode 2 because the 

iron source for the additional pyrite is probably the same as 

for the euhedral and anhedral pyrite; that is, the dissolution 

of relatively less reactive iron minerals.

The massive sulfide lenses in Green River samples 

probably are pyritic cement in concretionlike shapes. The 

fine-grained nature of these cements suggests that pyrite 

was supersaturated locally, possibly due to rapid release of 

iron. The time of lens formation is unconstrained by the data 

in this study and therefore is not shown on figure 7.

Pyrite that replaces marcasite and pyrrhotite 

represents the most recent generation of iron sulfide 

mineral, episode 4 (fig. 7). Such pyrite is generally anhedral 

or euhedral, and it is commonly bladed where it pseudomorphically has replaced bladed pyrrhotite.

Sulfide minerals that fill fractures in rocks of the 

Green River Formation probably formed during a late 

postlithification episode (Dyni, 1983). This episode of 

pyrite formation is not shown in the paragenetic sequence 

(fig. 7) because fracture fillings were not observed in the 

core intervals of this study.

Pyrrhotite Formation

Pyrrhotite has not been reported in unmetamorphosed 

lacustrine rocks outside the Green River Formation. Its local 

abundance in the Green River rocks indicates unusual

chemical conditions at the time of formation. Eh-pS2 

stability fields of pyrite, pyrrhotite, siderite, and hematite 

for two CO2 partial pressures (Pcoz) are plotted in figure 8 

(the equations used to construct these diagrams are in 

Berner, 1964). The pyrrhotite stability field is constrained 

by very reducing conditions (Eh <-0.5 volts) and by 

relatively high H2S concentrations (pS2~<9) within the 

sediment (fig. 8). At an assumed upper limit of PCO2=1 atm, 

the pyrite and siderite fields are large relative to that of 

pyrrhotite. At a value of PCO2=10~2-5 atm, the estimated 

value for shallow-marine porewater (Berner, 1964), the 

pyrrhotite field expands relative to the siderite field, but the 

slope of the pyrite-pyrrhotite boundary remains unchanged. 

Pyrrhotite is generally restricted to saline lake stage 2; 

thus, high salinity is added to the above requirements of 

very reducing conditions, low PCO2» and relatively high H2S 

concentrations. Commonly, these conditions are found in 

lakes containing euxinic hypolimnions (H2S-bearing bottom 

waters) in which the oxidative conversion of precursor FeS 

phases (S2~) to disulfides (S22~) is inhibited (see discussion 

on Soap Lake in Turtle, 1988). Failing to undergo any 

sulfur-oxidation reaction, unstable FeS phases will 

eventually recrystallize to pyrrhotite, the most 

thermodynamically stable FeS phase. Figure 8 indicates that 

pyrrhotite becomes more stable than siderite at low PCO2 a 

condition expected in sediments subsequent to 

decomposition of most organic matter and to formation of

G10 Green River Formation, Wyoming, Utah, and Colorado
0.0 -i

-0.2 -

-0.4 -

-0.6

pH=10

Pco2=10'2-5 atm

Pyrite

Hematite

Siderite H2O 

«-H2

4.0 6.0 8.0 10.0 12.0 

pS=

14.0 160

0.0 -|

-0.2 -

-0.4 -

-0.6

pH=10 

PC02= 1 atm

Pyrite

Siderite

Pyrrhotite H2o

4.0 6.0 8.0 10.0 12.0 14.0 16.0 

pS=

Figure 8. Eh-pS2~ stability field diagrams for pyrite, pyrrhotite, siderite, and hematite at pH 10. A, PC02 of 1 atm. B, PC02 

of 10~2-5 atm. Equations used in calculations are from Berner 

(1964).

carbonate minerals. Some pyrrhotite in the greater Green 

River basin surrounds shortite (Na2Ca2(CO3)3) (fig. 6D). 

Based on examination of sediment flow lines around shortite crystals, Fahey (1962) proposed that shortite formed in 

"partially compressed" Lake Gosiute mud. Bradley and 

Eugster (1969) used experimental data to argue that the 

shortite formed at temperatures of about 90 °C 

(corresponding burial depth of 1,525 m). The Green River 

Formation, however, never reached temperatures as high as 

90 °C according to vitrinite reflectance data (B. E. Law, 

U.S. Geological Survey, written commun., 1990). The 

timing of pyrrhotite formation, therefore, is not well 

constrained by its relation with shortite, but pyrrhotite likely 

formed in very reducing sediment during or after 

compaction (episode 3). Episode 3 could overlap with 

episode 2 (euhedral and anhedral pyrite formation) if the 

pyrrhotite and pyrite were forming in different microenvironments. These episodes are not shown as overlapping in 

figure 7 because the spatial distribution of the two minerals 

does not suggest that such microenvironments existed.

Marcasite Formation

Marcasite is present in about 40 percent of the Green 

River Formation samples (figs. 3-5). The occurrence of

marcasite in lacustrine rocks (or lake sediments) is 

unprecedented. Marcasite generally forms under acid 

conditions (Goldhaber and Kaplan, 1974; Murowchick and 

Barnes, 1986). Its formation in sediments deposited under 

high pH conditions such as those inferred for Lakes Uinta 

and Gosiute (pH estimated to be about 10) is inconsistent 

with experimental data, and its formation suggests that it did 

not form as a primary mineral during disulfide formation in 

episodes 1 or 2. Marcasite replacement of pyrrhotite has 

been reported to occur by the loss of iron from pyrrhotite 

and oxidation of the sulfide to disulfide (Kelley and 

Turneaure, 1970; Fleet, 1978; De Waal, 1984).

Euhedral and anhedral marcasite is present in two 

types of samples those containing pyrite and large 

amounts of organic matter and those containing pyrrhotite. 

The latter occurrence suggests that marcasite is a late 

alteration product of the pyrrhotite that formed during 

episode 3. The pyrrhotite is optically continuous along the 

blades, and the textures depicted in figure 6G suggest that 

marcasite and pyrite preferentially are in etched pits within 

the pyrrhotite blades. Oxidants made available during 

pyrrhotite alteration but not during its earlier formation 

include oxidants released during later stages of organicmatter mineralization or oxidants introduced by 

groundwater. Whatever the source, the amount was 

insufficient in most cases to completely oxidize the pyrrhotite. Possible exceptions are samples high in organic matter 

that contain only marcasite and do not exhibit relict pyrrhotite or pyrrhotite morphology. If pyirhotite was originally 

present in these samples, either the replacement was not 

pseudomorphic or the original pyrrhotite did not occur as 

blades. Perhaps high concentrations of organic matter 

promote formation of pyrrhotite having a different 

morphology that readily alters to marcasite, but this 

hypothesis remains to be investigated.

CONCLUSIONS

Petrographic characteristics of samples from three 

cores of the Green River Formation indicate that iron 

geochemistry and redox conditions in the sediment 

controlled the mineralogy and morphology of the iron 

sulfide minerals. Four episodes of sulfide-mineral formation 

can be defined. Although the exact timing of these episodes 

is impossible to determine, the relative timing has been 

inferred by combining petrographic observations and 

observations of sulfide-mineral formation in marine 

sediments. Near the sediment-water interface, iron oxide 

coatings on clays were sulfidized and produced framboidal 

pyrite (episode 1). Euhedral and anhedral pyrite grains and 

pyrite infilling of framboids then formed when less reactive 

iron-mineral phases were dissolved, possibly by ironreducing bacteria (episode 2). Pyrrhotite likely formed late 

in the sequence (episode 3) from unstable FeS phases that 

were not converted to disulfide because of extremely

Petrography of Iron Sulfide Minerals G11
reducing conditions in the sediments. Pyrrhotite was 

eventually and variably replaced by marcasite and (or) 

pyrite (episode 4).

REFERENCES CITED

Anderson, R.Y., Dean, W.E., Bradbury, J.P., and Love, D., 1985, 

Meromictic lakes and varved lake sediments in North 

America: U.S. Geological Survey Bulletin 1607, 19 p.

Berner, R.A., 1964, Stability fields of iron minerals in anaerobic 

marine sediments: Journal of Geology, v. 72, no. 6, 

p. 826-834.

Boyer, D.L., and Cole, R.D., 1983, Variations in sulfur 

mineralization in the Parachute Creek Member of the Green 

River Formation, Colorado and Utah, in Sixteenth Oil Shale 

Symposium proceedings: Golden, Colorado School of 

Mines, p. 160-175.

Bradley, W.H., and Eugster, H.P., 1969, Geochemistry and paleolimnology of the trona deposits and associated authigenic 

minerals of the Green River Formation of Wyoming: U.S. 

Geological Survey Professional Paper 496-B, 71 p.

Canfield, D.E., 1989, Reactive iron in marine sediments: 

Geochimica et Cosmochimica Acta, v. 53, p. 619-632.

Carstens, H., 1986, Displacive growth of authigenic pyrite: Journal 

of Sedimentary Petrology, v. 56, p. 252-257.

Cole, R.D., 1975, Sedimentology and sulfur isotope geochemistry 

of the Green River Formation (Eocene), Uinta basin, Utah, 

Piceance Creek basin, Colorado: Salt Lake City, University 

of Utah, Ph.D. thesis, 274 p.

Cole, R.D., Liu, J.H., Smith, G.V., Hinckley, C.C., and Saporoschenko, M., 1978, Iron partitioning in oil shale of the 

Green River Formation, Colorado A preliminary Mossbauer study: Fuel, v. 57, p. 514-520.

Cole, R.D., and Picard, M.D., 1981, Sulfur-isotope variations in 

marginal-lacustrine rocks of the Green River Formation, 

Colorado and Utah: Society of Economic Paleontologists 

and Mineralogists Special Publication 31, p. 261-275.

De Waal, S.A., 1984, Experimental simulation of the supergene 

reaction, pyrrhotite to marcasite+Fe2++electrons: 

Transactions of the Geological Society of South Africa, v. 

87, p. 273-279.

Dyni, J.R., 1981, Geology of the nahcolite deposits and associated 

oil shales of the Green River Formation in the Piceance

Creek Basin, Colorado: Boulder, University of Colorado, 

Ph.D. thesis, 144 p. 

_ 1983, Distribution and origin of sulfur in Colorado oil

shale, in Sixteenth Oil Shale Symposium proceedings:

Golden, Colorado School of Mines, p. 144-159. 

Fleet, M.E., 1978, The pyrrhotite-marcasite transformation:

Canadian Mineralogist, v. 16, p. 31-35. 

Goldhaber, M.B., and Kaplan, J.R., 1974, The sulfur cycle, in

Goldberg, E.D., ed., The sea, v. 5: New York, John Wiley

and Sons, p. 569-655. 

Kelly, W.C., and Turneaure, F.S., 1970, Mineralogy, paragenesis

and geothermometry of tin and tungsten deposits of Eastern

Andes, Bolivia: Economic Geology, v. 65, p. 609-680. 

Love, L.G., and Amstutz, G.C., 1966, Review of microscopic

pyrite: Fortschritte der Mineralogie, v. 43, p. 273-309. 

Lovley, D.R., and Phillips, E.J.P., 1986, Organic matter

mineralization with reduction of ferric iron in anaerobic

sediments: Applied and Environmental Microbiology, v. 51,

p. 683-689. 

Melchior, D.C., Wildeman, T.R., and Williamson, D.L., 1982,

Mossbauer investigation of the transformations of the iron

minerals in oil shale during retorting: Fuel, v. 61,

p. 516-522. 

Milton, C., and Eugster, H.P., 1959, Mineral assemblages of the

Green River Formation, in Abelson, P.H., ed., Researches in

geochemistry: New York, John Wiley and Sons,

p. 118-150. 

Murowchick, J.B., and Barnes, H.L., 1986, Marcasite precipitation

from hydrothermal solutions: Geochimica et Cosmochimica

Acta, v. 50, p. 2615-2629. 

Pabst, A., 1970, Pyrite of unusual habit simulating twinning from

the Green River Formation of Wyoming: The American

Mineralogist, v. 56, p. 133-145. 

Pyzik, A.J., and Sommer, S.E., 1981, Sedimentary iron monosulfides Kinetics and mechanism of formation: Geochimica

et Cosmochimica Acta, v. 45, p. 687-698. 

Raiswell, R., 1982, Pyrite texture, isotopic composition and the

availability of iron: American Journal of Science, v. 282,

p. 1244-1263. 

Rickard, D.T., 1974, Kinetics and mechanism of sulfidization of

goethite: American Journal of Science, v. 274, p. 941-952. 

Roberts, W.M.B., Walker, A.L., and Buchanan, A.S., 1969, The

chemistry of pyrite formation in aqueous solution and its

relation to the depositional environment: Mineralium

Deposita, v. 4, p. 18-29.

G12 Green River Formation, Wyoming, Utah, and Colorado

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