U.S. GEOLOGICAL SURVEY BULLETIN 1578 
Cover shows scanning electron micrograph of late-stage smectite (lower left) coating 

earlier chabazite crystals formed on early-stage smectite, in altered Pliocene vitric 

tuff, Durkee, Oregon. See article by A.j . Gude 3rd, and R.A. Sheppard, (this volume, 

p. 301-333). 
Studies in Diagenesis 

FREDERICK A. MUMPTON, Editor 

MARC W. BODINE, JR., 

Workshop Convenor and Organizing Committee Chairman 

Papers presented at a U.S. Geological Survey workshop on 

diagenesis in sedimentary rocks, Golden, Colorado, March 1982 

U.S. GEOLOGICAL SURVEY BULLETIN 1578 
DEPARTMENT OF THE INTERIOR 

DONALD PAUL HODEL, Secretary 

U.S. GEOLOGICAL SURVEY 

Dallas L. Peck, Director 

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON: 1986 

For sale by the 

Books and Open-File Reports Section 

U.S. Geological Survey 

Federal Center 

Box 25425 

Denver, CO 80225 

Library of Congress Cataloging in Publication Data 

Studies in diagenesis. 

(U.S. Geological Survey bulletin ; 1578) 

Bibliography: p. 

Supt. of Docs. no.: 119.3:1578 

1. Diagenesis-Congresses. I. Mumpton, Frederick Albert, 1932- . II. 

Bodine, Marc W. Ill. Series. 

QE75.B9 no. 1578 557.3 s 87-600128 

[QE571] [551.3] 
PREFACE 

The collection of research papers in this volume evolved from a workshop on 

diagenesis in sedimentary rocks sponsored by the U.S. Geological Survey. The 

workshop was held March 16--18, 1982, in Golden, Colorado, and was attended 

by 248 participants from the Survey's Geologic and Water Resources Divisions. 

The idea for the workshop came to four of us-Walt Dean, Phoebe Hauff, 

Harry Tourtelot, and me-during a series of informal discussions in late 1980. 

The complexity and interdisciplinary nature of the broad range of diagenetic 

processes were clearly recognized, as was their importance in the genesis, 

transport, emplacement, and mode of occurrence of a wide variety of organic 

fuels and mineral resources. The advantages of bringing together scientists who 

were actively engaged in diagenesis studies from several branches in two 

divisions of the Survey were readily apparent, and so a proposal for such a 

multidisciplinary and interdivisional workshop was formulated. The proposal was 

accepted and approved by the Survey in early 1981 . 

The workshop comprised two days of technical sessions consisting of oral 

papers and concurrent poster displays. A third day was set aside so that 

participants could schedule one-on-one conferences with any of the invited guest 

speakers or arrange small group discussions on topics of mutual interest. Each 

of the four half-day technical sessions represented a different aspect of 

diagenesis---<=hemical, carbonate, fine clastic, and coarse clastic-and began with 

an invited review paper on the theme of that session. The program opened with 

an overview "keynote" paper by Robert M. Garrels on the role of diagenesis in 

the geochemical cycle. In the first session, Robert A. Berner reviewed advances 

in several aspects of sediment-fluid interaction accompanying early diagenesis, 

and Lynton Land surveyed diagenesis in carbonate rocks to lead off the second 

session. In the third session John Hower reviewed the advancements in clay 

mineral diagensis, and in the final session Raymond Siever discussed thermal 

and burial relationships during the history of sedimentary basins. 

During these sessions 50 Survey authors presented 22 talks and 9 poster 

displays on topics ranging from diagenesis in limestones to uranium deposits to 

zeolitic tuffs. The authors reported findings of field, laboratory, and theoretical 

investigations dealing with such disparate approaches as organic geochemistry of 

hydrocarbon maturation, stable isotope fractionation, petrographic analysis, 

mineral equilibria and dissolution kinetics, mapping and field relations, and trace 

element geochemistry. The concurrent poster sessions presented results of more 

intensely focused studies. 

The workshop's flavor and organization have been retained in this volume. In 

addition to overview papers by four of the participants, the volume contains 

most of the Survey research papers that were presented at the workshop. They 

range in length from expanded versions of extended abstracts to condensations 

of lengthier treatises. The remaining papers have either been published elsewhere 

or incorporated into works of broader scope. Only a minimum effort has been 

made to formalize the stimulating presentation style of the authors in order to 

preserve and emphasize their unique approaches to this complex and difficult 

subject. 

Preface Ill 
The orgamzmg committee is grateful to the distinguished scientists whose 

informative review papers lead off each of the sections of this volume. The 

untimely death of John Hower precludes inclusion of his review paper in this 

volume. All who participated in the workshop keenly feel his passing. The 

success of the workshop was in large part due to the contributions of time and 

talent of the other members of the organizing committee: Bruce Bohor, Lorna 

M. Carter, Walter E. Dean, Jr., Donald L. Gautier, Martin B. Goldhaber, 

Phoebe L. Hauff, Zell E. Peterman, Orner B. Raup, Betty A. Skipp, and Harry 

A. Tourtelot. My appreciation is also extended to Frederick A. Mumpton, who 

accepted the task of editing this volume, and to Diane C. Schnabel, who 

guided it through its publication hurdles. My thanks also go to Artis M. 

Schack, Jane A. Edwards, and their coworkers for timely and accurate 

manuscript preparation. Arthur L. Isom designed the attractive cover. 

IV Preface 

Marc W. Bodine, Jr., Workshop Convenor 

and Organizing Committee Chairman 
CONTENTS 

PREFACE, by Marc W. Bodine, Jr. Ill 

PART I. INTRODUCTION 

Sediment cycling and diagenesis, by Robert M. Garrels 1 

PART II. CHEMICAL DIAGENESIS 

A kinetic approach to chemical diagenesis, by Robert A. Berner 13 

Seasonal diagenetic changes in salts of Owens Lake, California, 1970-77, by George 

I. Smith and Irving Friedman 21 

Influence of iron and manganese on the chemical partitioning of copper, zinc, and 

chromium during early diagenesis in outer continental-shelf sediments from the Gulf 

of Mexico, by Lorraine H. Filipek 31 

Mobility and distribution of heavy metals during the formation of first-cycle red beds, 

by Robert A. Zielinski, Salman Bloch, and Theodore R. Walker 51 

Stoichiometry of organic decomposition in sediments of the Potomac River estuary, 

by David L. Parkhurst and Steven D. Goodwin 61 

Relation of hydrocarbon type to maturity of organic matter in Upper Cretaceous chalks, 

eastern Denver basin, by Dudley D. Rice 75 " 

Relationship between vitrinite reflectance, metamorphic grade, and temperature in the 

Cerro Prieto, Salton Sea, and East Mesa geothermal systems, Salton Trough, United 

States and Mexico, by Charles E. Barker, Bonnie L. Crysdale, and Mark J. 

Paw lewicz 83 

Origin and diagenesis of Cretaceous organic-carbon-rich lithofacies in the Atlantic 

Ocean, by Walter E. Dean and Michael A. Arthur 97 

PART III. CARBONATE DIAGENESIS 

Limestone diagenesis-Some geochemical considerations, by Lynton S. Land 129 

A comparative study of the dissolution and crystal growth kinetics of calcite and 

aragonite, by Eurybiades Busenberg and L. Niel Plummer 139 

Effect of magnesium ions on calcium carbonate nucleation and crystal growth in dilute 

aqueous solutions at 25°C, by Michael M. Reddy 169 

Authigenic potassium feldspar in ribbon rocks of the Cambrian Conococheague Limestone, western Maryland, by Paul P. Hearn, Jr., Roy C. Lindholm, and John F. 

Sutter 183 

Diagenesis of the Todilto Limestone Member of the Wanakah Formation, Chama basin, 

New Mexico, by Jennie Ridgely 197 

Mineralogy and stable isotope geochemistry of carbonate and sulfate minerals in 

diagenetically altered Tertiary and Cretaceous sandstones, Uinta basin, Utah, by Janet 

K. Pitman, Martin B. Goldhaber, and Thomas D. Fouch 207 

Diagenetic relationships in a hydrocarbon-productive chalk-The Cretaceous Niobrara 

Formation, by Richard M. Pollastro and Peter A. Scholle 219 

Contents V 
PART IV. DIAGENESIS IN CLASTIC ROCKS 

Burial diagenesis of sandstones, by Raymond Siever 237· 

Diagenesis of the Morrison Formation, Smith Lake uranium district, McKinley County, 

New Mexico, by NeilS. Fishman and Richard L. Reynolds 249 

Diagenesis and uranium mineralization in the Westwater Canyon Member of the 

Morrison Formation, Grants uranium region, northwestern New Mexico, by Paula L. 

Hansley 265 

Diagenesis and the sources of the ore-forming fluids, Tony M. vanadium-uranium 

deposit, Henry structural basin, Utah, by H. Roy Northrop, Martin B. Goldhaber, 

C. Gene Whitney, Fred Peterson, Richard L. Reynolds, and John A. Campbell 281 

Clay mineral diagenesis in lacustrine sediments, by Blair F. Jones 291 

Zeolitic diagenesis of tuffs in an Upper Miocene lacustrine deposit near Durkee, Baker 

County, Oregon, by Arthur J. Gude 3rd, and Richard A. Sheppard 301 

Magadi-type chert-A distinctive diagenetic variety from lacustrine deposits, by 

Richard A. Sheppard and Arthur J. Gude 3rd 335 

Salt-induced diagenesis of argillaceous sediments, by Charles W. Holmes 347 

Sodium-potassium ion exchange during smectite diagenesis-A theoretical discussion, 

by Dennis D. Eberl 363 

VI Contents 
Sediment Cycling and Diagenesis 

By ROBERT M. GARRELS, Department of Marine Science, University of South 

Florida, St. Petersburg, Florida 33701 

CONTENTS 

Abstract 2 

Introduction 2 

Acknowledgments 3 

Diagenetic changes in shales 3 

Isotopes as guide to diagenesis 6 

o180 problems 6 

Carbon and sulfur isotopes 6 

Carbonate problems 8 

Paleozoic carbonates 10 

Selective loss or preservation 10 

Mercury in Early Proterozoic shales 10 

The calcium problem 10 

Other similar problems 10 

Summary 10 

References cited 10 

FIGURES 

1. Schematic diagram showing that original minerals of a shale could react to yield 

a simple three-phase system-quartz, illite, chlorite 4 

2. Compositions of fine-grained rocks of various ages plotted as a function of 

N~O/ Al20 3 ratio versus K20/ Al20 3 ratio 5 

3. Variation in o180 in carbonates and cherts as function of geologic age 7 

4. Plot of calcium and magnesium in carbonate rocks as function of their age 9 

TABLE 

1. Comparison of the major elements of Archean, Early Proterozoic, and Phanerozoic shales 6 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 1 
Abstract 

Global compositional and mineralogic differences in rocks 

as a function of their ages can seldom be unequivocally interpreted as primary or diagenetic. Most mineralogic differences 

between Mesozoic-Cenozoic and older rocks are diagenetic 

and represent a tendency for the original aggregates of many 

phases to change to equilibrium assemblages. For shales, this 

assemblage is generally chlorite-quartz-illite-feldspar; for carbonate rocks, it is calcite-dolomite. Average differences in bulk 

chemical compositions of shales with age are small; carbonate 

rocks show an irregular continuous increase in the Mg/Ca ratio 

with age. 

The 5345 isotope composition of calcium sulfate deposits 

and the o13C isotope values for carbonate rocks, if carefully 

chosen, reveal the isotopic sulfur and carbon values of the 

oceans from which they came, and are currently the most trustworthy and useful tools of the geologist studying sedimentary 

cycling. Global oxidative transfer of organic carbon to carbonate 

is counterbalanced by reductive transfer of sulfate to sulfide. 

INTRODUCTION 

Any reasonable estimate of the rates of erosion of 

the land areas of the world through geologic time yields 

a total mass of eroded sediments several to many times 

the mass of sediments in existence today. Also, 80 percent or more of the material currently being transported 

to the oceans by streams is derived from sedimentary 

rocks. The inevitable conclusion is that sedimentary rocks 

do not exist forever; instead, they are removed from the 

Earth's surface environment by subduction of oceanic 

plates, they are metamorphosed by classical processes 

involving folding and intrusion, or they are uplifted, 

eroded, and reborn as new sediments. 

Although data on the masses of sedimentary rocks 

remaining today, as a function of their ages, are not nearly 

as complete as one could wish, there is no doubt that, 

to a first approximation, the mass of rock remaining of 

a given age diminishes exponentially with increasing age. 

This relation is consistent with an erosion-depositiondestruction system, in which the amount of sediment 

deposited per year is nearly constant (averaged over tens 

of millions of years), and the amount of sediment 

destroyed, by whatever processes, has been proportional 

to the amount in existence at the time of destruction. 

The simplest system that can be imagined to explain 

the present-day distribution of the mass of sedimentary 

rocks as a function of their ages is an entirely cannibalistic 

system in which a sedimentary mass was created early in 

the Earth's history, about 4 billion years ago, by interaction of volatiles exuded by a differentiating Earth with 

the primitive crystalline crust. This sedimentary mass 

would then have been eroded from the land areas into the 

seas and eventually returned to the land areas for reerosion. A model of such an unrealistic system, in which 

depositional rates were constant and in which the amount 

2 Studies in Diagenesis 

of eroded material required to achieve the constant depositional rate was derived from the remnants of older 

sediments in proportion to their remaining masses, produces a first-order replica of what exists now in terms of 

sediment mass distribution as a function of sediment age. 

The model, although it reproduces sediment mass-age 

relations qualitatively, is inherently incomplete because 

sedimentary rocks are derived today not only from preexisting sedimentary rocks, but from the 20 percent or so 

of the land area that is underlain by crystalline rocks, 

mostly of igneous rather than sedimentary composition. 

The production of new sedimentary rocks from these 

crystalline rocks requires reaction with acid volatiles, such 

as C02, HCl, HF, and H2S. Thus, the answer to this 

question must be that the acidic v;olatile compounds 

necessary to convert igneous rocks intb sedimentary rocks 

(for example, CaSi03 to CaC03

) are continously resupplied to the surface environment by the removal of such 

volatiles from sedimentary rocks that have been heated 

and squeezed during subduction or classical metamorphism. This process recreates the original igneous minerals 

and returns the volatiles to the surface environment via 

volcanoes, hot springs, and similar venting systems. 

The gross aspects of the Earth-sediment system can 

be outlined with some confidence. If the deposits of a 

given period, such as the Cambrian, are considered from 

their time of deposition to the present and compared to 

the properties of the Cambrian rocks surviving today, the 

following relations must hold: (1) the sample of originally 

deposited Cambrian rocks available for inspection today 

is only a fraction, and probably a small fraction, of the 

rocks originally deposited; (2) the proportions of sedimentary rock types in the remaining Cambrian may or may 

not be a fair sample of the original proportions; and (3) 

the remaining Cambrian rocks have existed for 500 million 

years or more and have been subjected to various degrees 

of metamorphism, ground-water circulation, and other 

processes that may have changed their bulk chemical and 

isotopic compositions, as well as their mineralogies. 

A dilemma exists in trying to discuss long-term 

diagenetic or secondary changes in sedimentary rocks. If 

differences are found between Cambrian shales and 

Jurassic shales, what parts of those differences are between the Cambrian shales as originally deposited and the 

Jurassic shales as originally deposited? Furthermore, what 

part can be attributed to chemical and mineralogic 

changes that have taken place in these shales since they 

were deposited, and what part can be attributed to selective preservation of various types of shales because of the 

vagaries of selective erosion or other destruction? 

When faced with problems like this, there is a tendency for scientists to become polarized and to embrace 

one view or the other without much compromise, as was 

true a few generations ago when the problem of continental drift or permanence became an issue. In the United 

States in the 1930's and 1940's, anyone who proposed 
continental drift was forced to disprove continental permanence. To some extent, the same situation exists today 

with respect to the chemical compositions and mineralogies of ancient sediments. One extreme view is that the 

remaining sediments of a given age are a fair sample of 

the total mass of sediments of that age that once were 

deposited, and that any secular trends of sediment composition deduced from this conclusion are valid if one 

wants to compare what was deposited in the Cambrian 

Period with that in the Jurassic Period. The other school, 

of course, can be termed the Uniformitarianism School 

and claims that all sediments are identical when deposited; 

differences now observed between Cambrian and Jurassic 

sediments are the products of selective preservation of 

those sediments plus selective chemical changes engendered by greater exposure to post-depositional agents 

(for example, ground water) that could leach some components, increase others, and in general tend to produce a false impression of the relations of the original 

deposits. 

Garrels and Mackenzie (1971), in their book 

"Evolution of Sedimentary Rocks," emphasized postdepositional change as the reason for apparent secular 

changes in sedimentary rocks. They suggested that 

chemical analyses of sedimentary rocks had been used 

uncritically as a means of revealing true primary changes 

in chemical compositions of sediments through time. 

They therefore emphasized the data on diagenetic changes 

and may have produced a document that, read by itself, 

advocates a viewpoint overly strong in favor of diagenetic 

change. 

Garrels and MacKenzie (1971) tried to make clear 

that current data are not adequate to permit discrediting 

the hypothesis of the constant original sediment composition and mineralogy. They were able to show that the differences between Mesozoic and Paleozoic shales could 

be attributed to post-depositional changes, or perhaps 

that post-depositional changes tend to obliterate real 

initial compositional and, especially, mineralogic differences. They posed a challenge-what are the true 

secular changes in sedimentary compositions (and they 

must exist), and what are the changes superimposed by 

aging processes and selective erosion? 

In the interpretation of the history of the Earth, 

it is of critical importance that the nature of the original 

sedimentary deposits of a given age be known. One comes 

to the question of which elements, which isotopes, which 

data of any kind can be trusted to record and preserve 

their original values to a given time later in Earth history. 

The readers of any scientific paper cannot ever be 

really up to date on such currently controversial issues. 

This problem of the interpretation of the Earth's history 

is one of the most basic in geology and has been and is 

now being worked on by many able researchers. The 

origin of dolomite is just one aspect of the problem, and 

it is rumored that the mass of scientific literature on 

dolomite exceeds the existing mass of dolomite. In the 

present paper, the number of references has been kept 

small, but those selected are representative sources that 

will give the interested reader extensive bibliographies. 

True secular changes in sedimentary rocks, versus changes 

resulting from post-depositional processes, are actively 

being sorted out, and important progress is anticipated. 

This paper focuses on some of these continuing problems, 

wherein sediment cycling is related to diagenetic change. 

Acknowledgments 

Ideas presented in this paper have a long history. 

Through the years they have evolved from discussions 

with many persons, especially R. A. Berner, J. I. Drever, 

B. Gregor, W. Hoefs, W. Holser, H. D. Holland, the 

late John Hower, C. Junge, I. R. Kaplan, J. Lawrence, 

A. Lerman, F. T. Mackenzie, J. McKenzie, the late E. A. 

Perry, Jr., E. C. Perry, Jr., A. B. Ronov, the late T. J. 

Schopf, M. Schidlowski, and J. Veizer. Support for the 

preparation of this material came from National Science 

Foundation Grant #EAR-8107736. 

DIAGENETIC CHANGES IN SHALES 

The record of post-depositional changes in rocks 

is well documented in shaly rocks (see Perry and Hower, 

1970; Hower and others, 1976). As initially deposited, 

they contain a wealth of mineral phases, most of which 

are clearly out of equilibrium with most of the others in 

any final equilibrium sense. The original muddy or shaly 

sediment that is deposited contains a variety of minerals, 

plus a most important set of organic materials. The 

chemistry of the changes that take place shortly after 

deposition, especially as it is related to the complexities 

of the decomposition of organic materials, has been 

documented in detail elsewhere (see Berner, 1981). 

Muds that eventually become shales are mixtures 

of many mineral phases. These phases are brought to the 

depositional environment physically, as solid phases in 

the suspended load of streams. They then coagulate as 

a result of encountering the high salinity of sea water, 

and deposit as a remarkable conglomeration of tiny 

quartz grains; a variety of clay minerals, including montmorillonites, mixed-layer minerals, vermiculites, degraded 

micas, and chlorites; detrital feldspars; and odds and ends 

of other mineral species. The most important aspect of 

a freshly deposited shale is that a variety of mineral compounds is involved. All these phases cannot simultaneously be in equilibrium with the aqueous phase. The pore 

waters respond to the most reactive and, in general, the 

most soluble. 

What happens in detail in the pore waters of shales 

is not known (it is almost impossible to sample such 

Sediment Cycling and Diagenesis 3 
waters), but chemical analyses of the waters are consistent with the achievement of equilibrium between the pore 

waters and the more reactive (and perhaps metastable) 

constituents of the shale. For example, fresh shales (or 

muds) have dissolved silica values much higher than those 

in equilibrium with quartz. 

Fresh muds or shales violate the phase rule-many 

more phases are present than components, and these phases 

cannot all be in equilibrium. One can think of diagenesischemical and mineralogical changes prior to severe metamorphism-as progress toward true equilibrium among 

mineral phases at low temperature and pressure. The equilibrium assemblage for the multimineralic original shales 

is a chlorite-illite-quartz rock with either feldspar or, more 

rarely, kaolinite. It takes time to produce this simple 

assemblage, but the assemblage does indeed describe the 

clay mineralogy of Paleozoic and many Proterozoic shales. 

The clay mineralogy of these old shales must be secondary. Montmorillonite is a common weathering product in 

modem soils. It changes to illite after burial by substitution 

RADIOLARIA 

DIATOMS 

of K + in the interlayer positions, coupled with an addition of AIH and a loss of Si4+ from the tetrahedral positions. All these changes are well documented as being postdepositional-the mixed-layer illite/montmorillonite 

minerals of Paleozoic shales must once have been montmorillonites, like those found in Mesozoic shales. 

The above relations show that the illite-chloritequartz monotonous mineral assemblage in Paleozoic shales 

is expected as an equilibrium end product for the multitude 

of minerals originally deposited in shales. Figure 1 is a 

schematic diagram adapted from Garrels and Mackenzie 

(1974) to show how the variety of phases present in a newly 

deposited mud could react through time to produce an 

equilibrium assemblage of illite, chlorite, and quartz. The 

bulk compositions of Paleozoic, Mesozoic-Cenozoic, and 

Proterozoic shaly rocks are plotted to show that the final 

product falls into the illite-chlorite-quartz triangle. 

Figure 2 has been adapted from a figure by Garrels 

and Mackenzie (1971) chiefly by adding Cameron and Garrels' (1980) data for Early Proterozoic and Archean shales. 

QUARTZ 

OPAL 

• PALYGORSKITE 

Pz CzMz \ ~ MONTMORILLONITE • K-FELDSPAR • SEPIOLITE 

I ~. \. GLAUCONITE 

PHILLIPSITE NONTRONITE 

ILLITE 

.MICA 

CHLORITE 

GIBBSITE (Na, Fe, Mg, K) 

Figure 1. Schematic diagram showing that a mixture of original minerals in a shale with bulk compositions lying 

in the circle could react to yield a simple three-phase assemblage-quartz + illite + chlorite. E?, Proterozoic; Pz, 

Paleozoic; CzMz, Cenozoic-Mesozoic. Modified from Garrels and Mackenzie (1974). 

4 Studies in Diagenesis 
0.3 IGNEOUS ROCKS 

0.1 

AVERAGE 

ARCHEAN * 

AVERAGE 

IGNEOUS 

ROCK 

CLARKE AVERAGE 

[!) MESOZOIC-CENOZOIC 

LATE PROTEROZOIC 

~ U.S.S.R. 

RONOV AVERAGE 

MESOZOIC-CENOZOIC 

rJ 8 EARLY PROTEROZOIC 

U.S.S.R. 

PRESENT 

SUSPENDED 

STREAM LOAD 0° 8 f:l. LATE PROTEROZOIC BELT 

CLARKE AVERAGE 

PALEOZOIC 

L.;J U.S.A. 

& RONOV AVERAGE 

PALEOZOIC PLATFORM 

0.1 0.2 0.3 0.4 

WEIGHT RATIO K20/AI 20 3 

Figure 2. Compositions of fine-grained igneous and sedimentary rocks of various ages plotted as function of Na20/AI20 3 ratio 

versus K20/AI2

0 3 ratio. Modified from Garrels and Mackenzie (1971, fig. 9.2, p. 229). Clarke average from Clarke (1924); Ronov 

average from Ronov and Migdisov (1971 ). 

The similarity of composition of Paleozoic shales to those 

of both the Early and Late Proterozoic suggests similar 

source areas and the attainment of similar stable 

assemblages. On the other hand, Archean shales reveal 

their kinship to the basaltic crust from which many of 

them were derived. Interestingly, they are high in N~O, 

but nearly as low in CaO as Proterozoic shales, perhaps 

pointing to a breakdown during weathering of highcalcium plagioclase, with recrystallization of albitic 

material (see Land and Milliken, 1981). One common 

interpretation is that the Mesozoic-Cenozoic shales 

(Ronov values) still show their compositional inheritance 

from the same types of materials being carried by rivers 

today. Paleozoic and Proterozoic shales exhibit postdepositional gain in ~0, related perhaps in part to slow 

addition of K+ to highly charged clay minerals. The 

reduction of FeH in the octahedral sites of phyllosilicates may increase their negative charge and cause them 

to fix K + and to lose interlayer water. 

Alternatively, a shift in provenance has taken place 

with Archean and Mesozoic-Cenozoic shales representing 

times of more mafic crystalline sources. Veizer and Jansen 

(1979), in an important paper on basement and sedimentary recycling, and Engel and others (1974) presented in 

detail the case for more mafic sources for Archean and 

Mesozoic sediments, as compared to Paleozoic and Proterozoic ones. At the other extreme are Garrels and 

Mackenzie (1972), who constructed a cycling model (not 

including Archean sediments) that produced all sedimentary rocks by deposition of sediments like those of today. 

Diagenetic changes were held accountable for differences 

in composition as a function of age. 

Sediment Cycling and Diagenesis 5 
The general chemical characteristics of shales as a 

function of age are shown in table 1 (Cameron and Garrels, 1980). The basic similarities are obvious; indeed, if 

one compares the three sets of analyses, oxide by oxide, 

no overall trends are apparent. For some oxides the 

Archean shales are more like Phanerozoic than Early Proterozoic shales; for other oxides the reverse is true. The 

differences among the analyses easily could be the result 

of sampling procedures and selection of shale types. 

Perhaps the two most interesting features are the lack of 

an age trend in the iron content or in the organic carbon 

content. 

ISOTOPES AS GUIDE TO DIAGENESIS 

o180 problems 

As shown in figure 3, o180 values of carbonates 

diminish irregularly with increasing age. The curve for 

silicates is roughly parallel to the o180 carbonate curve, 

averaging about 5 per mil higher. The modern value for 

carbonates in equilibrium with sea water is about + 30 

per mil o180, that for siliceous organisms about + 35 per 

mil. The mean value for all carbonates is probably about 

Table 1. Comparison of Archean, Early Proterozoic (Aphebian), and Phanerozoic shales 

[Modified from Cameron and Garrels (1980), table III] 

Archean Aphebian Phanerozoic 

(carbonate 

free) 

Si02 60.52 65.00 63.11 

Ti02 0.76 0.76 0.89 

A1 203 17.63 16.22 18.80 

Fe2o3 7.91 6.37 7.79 

MgO 3.47 2.51 3.11 

CaO 2.16 0.52 0.07 

Na 2o 2.62 1.44 1.27 

K20 2.84 4.83 3.69 

c 0.77 1.68 0.87 

s 0.66 0.65 0.38 

Totals ---- 99.98 99.98 99.98 

o34s (0 /oo) +1.31 +11. 79 -12 

(per mi 1) :t0.32 :t4.37 

6 Studies in Diagenesis 

+ 20 per mil, that for silicate minerals, about + 17 per 

mil. 

At least four explanations for these isotope-time 

trends exist. One is exchange with ground waters, which 

are typically much lighter isotopically than sea water; a 

second is metamorphism and consequent isotope 

exchange among minerals; a third is formation in an 

ocean that had a o180 value smaller than today's; and a 

fourth is precipitation of the minerals from a much 

warmer ocean. 

Li (1972) performed a clever isotope balance that 

suggests that the ocean has been about the same 

isotopically, or heavier, through time. He showed that 

reaction of an original ocean, with an expected o180 

value of + 6 to + 9 per mil (values representative of water 

emitted from the crystallization of igneous rocks), to give 

an ocean with the zero o180 value of today would indeed 

produce a mass of sediment about equal to that existing 

today with the mean o180 value of that sedimentary 

mass. An important implication of Li's calculation is that 

no major net loss of silicate-containing sediments has 

taken place from the exogenic cycle by subduction or 

related processes. 

It is surprising that there has been so little followup on the work of Knauth and Epstein (1976), who used 

deuterium-oxygen isotope plots of Si02 minerals of 

various ages to reconstruct a generalized temperature 

history of the ocean. If their methods and results are 

valid, the temperature history of the earth must have 

been a major control of the course of organic evolution. 

If values of 60 OC or higher during the Archean and 

Early Proterozoic are correct, they provide a simple 

explanation for the approximately 2 billion years that 

seem to have elapsed between the invention of photosynthesis by procaryotes and the appearance of eucaryotic 

organisms. Despite the warnings of many of his bio-

\ogically oriented colleagues, this author is still impressed with the possibility that the only kinds of 

organisms that could have survived above about 65 °C 

would be classified as procaryotes, and that physicochemical temperature barriers to the development of 

eucaryotic cells exist that defy even the flexibility of evolutionary change. 

Carbon and sulfur isotopes 

Values for carbon and sulfur isotope ratios for carbonates, organic carbon, sulfates, and sulfides that are 

indeed representative of the conditions of their formation are becoming available in increasing numbers. They 

are the most important and trustworthy information 

about ancient environments (except for fossils) available 

to us today. Holser and Kaplan's (1966) work on ancient 

evaporites demonstrated firmly that the o34S values for 

calcium sulfates were indeed close to the o34S values of 
32 0 

K 

0 

J 

e 

T 

z C'.tt, 

<( 

w 0 ~-9)'6' 

e u 

0 c 

z K 

<( 

w 0 0 :iE 26 p D 

0 

a: e e 0 <( 

(;if?. 0 0 

z J ~6'~ s <( 0 1- 011. ~ ~6' 

p 24 

~ 

e 0 

c 

c e e e s 

22 D 0 

CARBONATE FRESHWATER MEAN ---------------

li 

e 

PRESENT 100 20L---------~--------~----------L---------~--------~----c~--~ 200 300 400 500 600 

TIME, IN MILLIONS OF YEARS 

Figure 3. Variations in o ' 180 values of carbonates and cherts as a function of geologic age. Data from 

Degens and Epstein (1962). Q, Quaternary; T, Tertiary; K, Cretaceous; J, Jurassic; Tr, Triassic; P, Per-· 

mian; C, Carboniferous; D, Devonian; S, Silurian; 0, Ordovician, f:, Cambrian. 

contemporary oceanic sulfate. Their paper must therefore 

be regarded as one of the most important geological 

publications of the 20th century. It is apparently the only 

current unequivocal source of information concerning 

ancient ocean compositions. A recent update of that 

paper (Claypool and others, 1980) adds valuable information but does not change the basic patterns established 

in 1966. 

It has slowly become apparent also that the o13C 

values of ancient carbonate minerals and organic materials are trustworthy in the sense that they have generally 

not changed significantly from the time they were buried 

a meter or so in sediment. 

Variations in o34S values with time are much 

greater than those of o13C, chiefly because percentage 

variations in carbon fluxes are much smaller than those 

for sulfur fluxes. Recently, Veizer and others (1980) 

accumulated all available o13C values for carbonate minerals and showed that variations definitely exist within 

Phanerozoic time. They also showed that these variations 

are consistent with the predictions of Garrels and Perry 

(1974), who suggested that the carbon and sulfur cycles 

must be so closely linked that changes in sulfur isotopes 

of sea water (and thus in evaporites) should be reflected 

in changes of carbon isotopes. It is remarkable that the 

o13C values of marine carbonate minerals have fluctuated a few per mil around zero, and that the mean 

o13C values of organic materials, although much more 

variable than those of carbonates, have averaged about 

25 per mil lower than those of contemporaneous carbonates for more than 3 billion years (Schidlowski and 

others, 1984, p. 158). The conclusion from these data is 

that the ratio of carbonate carbon deposited to organic 

carbon deposited has not changed much from about 4: 1 

Sediment Cycling and Diagenesis 7 
through all of recorded Earth history, and that the effective isotope fractionation between inorganic and organic 

carbon has averaged about 25 per mil. Because the 

percentage of organic carbon in sedimentary rocks also 

has been nearly constant with time (Cameron and Garrels, 1980), there is a further, but less certain, implication that the rates of deposition of inorganic and organic 

carbon have also changed little during more than 3 billion 

years. 

Several articles have appeared in the last 10 years 

that use the age relations of o13C and o34S to interpret 

the global oxygen budget (see Schidlowski and others, 

1977). Garrels and Lerman (1981) came to the following 

conclusions: (1) The total oxygen produced, as measured 

by the total amount of buried organic carbon in sedimentary rocks, is greater than that stored in sulfate in 

sediments and in the ocean plus that represented by ferric 

iron in various minerals. Several estimates of the oxygen 

budget find this excess of oxygen produced over oxygen 

stored to be perhaps as much as 400 X 1018 moles 

(roughly equivalent to the total oxygen in calcium sulfate 

deposits of all ages). (2) Oxygen stored in iron minerals 

may have been produced chiefly at about the time of formation of the banded iron formations, some 2 billion 

years ago. One kind of evidence is the disappearance at 

that time of thick strata of ferrous carbonates and 

silicates. (3) Since the beginning of the Phanerozoic Era, 

the Earth's surface has been dominated by the carbonate 

carbon-organic carbon-sulfate sulfur-sulfide reciprocal 

oxidation-reduction system (organic carbon + sulfate = 

sulfide + carbonate). (4) The amount of oxygen stored 

in sulfate has fluctuated by amounts several times larger 

than the amount now in the atmosphere, but there seems 

to be no consistent trend over the last billion years. 

Oxygen storage was minimal at the beginning of Cambrian time, irregularly rose to a maximum in the Permian, 

and has since declined somewhat. 

A current major mystery is the nature of the sulfur 

cycle during the Archean and Early Proterozoic. Apparently, the sulfate-reducing bacteria did not evolve importantly until about 2.3 billion years ago (Cameron, 1983). 

This conclusion is based on consistent mean o34S values 

of sulfides of about zero for all Archean sedimentary 

sulfides. Many deposits have a range of values, but the 

mean values are always close to that of the primordial 

meteorite sulfur standard. No values have been published 

for the o34S of bedded sedimentary anhydrite or gypsum 

deposits older than about 2.3 billion years (Cameron, 

1983). 

Cameron and Garrels (1980) found an amazing 

average value of + 12 per mil for the o34S of the sulfide 

of 400 Early Proterozoic shales. Early Proterozoic shales 

have consistently positive values for the o34S of their 

pyrite sulfur. Where has the light sulfur necessary to 

produce these high positive values gone? Is it possible 

that bacterial oxidation of sulfides during the Early 

8 Studies in Diagenesis 

Proterozoic yielded sulfates with o34S values lower than 

their parent sulfides, and that the isotopically light sulfate 

was removed by reduction at high temperature by circulation through seafloor basalts? Some sulfide oxidation 

pathways, although apparently not important today, 

selectively utilize light sulfur (Schidlowski and others, 

1984, p. 167). 

CARBONATE PROBLEMS 

A major secular trend of sedimentary rocks that 

indeed may be primary is the change over the past several 

billion years from the Early Proterozoic deposition of 

iron carbonates and silicates, to early Phanerozoic deposition of dolomites and magnesium-rich clays, to the 

Mesozoic-Cenozoic calcite limestones and clays. If these 

changes are considered as a chemical problem of the relations of three sets of carbonates and silicates, clearly no 

given value of atmospheric carbon dioxide can have more 

than one of the three pairs in equilibrium. The thermodynamic values required to decide exactly in terms of 

parts per million the C02 pressure in equilibrium with 

the three systems are not well enough known, but these 

systems certainly are in the order, from highest C02 

required to lowest, iron carbonate-silicate, dolomitesilicate, calcite-silicate (see Holland, 1968). The problem 

is complicated by the oxidation state of the iron in the 

iron carbonates and silicates. The loss of ferrous carbonate from sedimentary systems was probably caused 

by an increase of atmospheric oxygen, rather than a 

decrease in atmospheric COr Nevertheless, the basins of 

deposition of the iron carbonates and silicates of the 

banded iron formations required a higher C02 pressure 

for codeposition of both phases than the Early 

Phanerozoic basins producing dolomites and Mg clays; 

the C02 pressure for modern basins producing calcite 

and clays is still lower. 

This author's interpretation considers that Early 

Proterozoic time had low oxygen pressures and relatively 

high C02 pressures (perhaps 10-2 atmospheres). Iron 

behaved like calcium and magnesium and produced 

extensive deposits of ferrous carbonates and aluminumfree silicates. As the sulfide reservoir diminished as an 

oxygen sink, because of the irreversible oxidation to 

sulfate, more oxygen became available for oxidation of 

ferrous compounds. The oxidation of ferrous carbonate 

produced C02, and the global balance of oxygen and 

C02 changed significantly. The change to an oxygenated 

atmospheric system during the Proterozoic must have had 

far-reaching consequences, strangely enough not well 

documented by studies of sediments of the 2-billion- to 

1-billion-year range. We are far from understanding these 

secular changes, but it is not inconceivable that the lowering of C02 in the atmosphere, because of a smaller contribution from depth as the thermal gradient diminished, 
was important in shifting from an anaerobic world in the 

Archean to an aerobic one in the Late Proterozoic. 

Whatever the reasons, iron changed from a dominantly 

2 + valent element to a 3 + valent element at about 2 

billion years ago. 

In figure 4, the Ca/Mg ratio of carbonate rocks is 

plotted as a function of time. If the data relate to primary 

deposition, the percentage of dolomite in carbonate rocks 

has been declining for 600 million years or more. As 

pointed out by Holland (1968), the occurrence of 

dolomite, whatever its origin, is promoted by high C02 

pressure in the atmosphere; low C02 pressure favors 

clay minerals rather than dolomite. Little dolomite has 

PCm Cm 

32 

28 

formed over the last 100 million years. All these relations 

make it important to wonder if the C02 content of the 

atmosphere has not been declining slowly but steadily 

over vast time intervals. If the C02 content of the 

atmosphere is conceived as a steady state balance between 

use of C02 by weathering and production by deep processes, such as volcanism, metamorphism, and subduction, 

it is reasonable that the aging Earth, despite its obviously 

important tectonic activity, is in fact senescent and not 

performing today in terms of its transfer of energy and 

materials from depth to surface as well as it did billions 

of years ago. If so, the carbonate changes suggested may 

be real long-term trends. 

02 Ct 

s2 o1 o3 c2 c3 Pt P2 T J Cr TrQ 

• ~ ,, 

ft. ,'• 1\ 

I\ I\ • /\ 

I \ I \ 1\ I 

..... 

z 

w 

u 

I\ f '. f\ I / \/ \1 \ I 

a: 24 ~-._ II \I '/. \ I w 

0.. 

z 

ui 

u 

z 20 

£3 

z 

::::> 

CXI 

~ 

~ 16 ~ 

a: 

w 

~ 

12 

8 

4 

""" ._ • \ I / ......... """ \ I 

,""" \ I 

•""" \ I 

CALCIUM 

\I 

• 

MAGNESIUM 

o~--------------------------5~1~0--------------4~3-o-----------------3~1-o-----------~2~25~------~------~----~ 

ABSOLUTE TIME, IN MILLIONS OF YEARS 

Figure 4. Calcium and magnesium in carbonate rocks as a function of their age; from Vinogradov and others (1957). 

PCm, Precambrian; Cm, Cambrian; 51

, Silurian 1; 52, Silurian 2; D1, Devonian 1; D2, Devonian 2; D3, Devonian 3; 

C1, Carboniferous 1; C2, Carboniferous 2; C31 Carboniferous 3; P1

, Permian 1; P2, Permian 2; T, Triassic; J, jurassic; 

Cr, Cretaceous; Tr, Tertiary; Q, Quaternary. 

Sediment Cycling and Diagenesis 9 
Paleozoic carbonates 

An important facet of the general relations among 

carbonate minerals in terms of what they once were and 

what they are now is the possibility that many species of 

organisms that today secrete aragonite or high-magnesium 

calcite may have secreted nearly pure calcite during 

Paleozoic time (see Sandberg, 1975). This possibility is 

particularly important to the general theme of diagenetic 

change. If, as generally assumed, present-day aragonitesecreting organisms produced aragonite 400 million years 

ago, the fact that the remains of these organisms today 

are calcite means that diagenetic changes have managed 

to cause recrystallization. This problem was discussed 

above, but it is worth reinforcing the point here. Until the 

kinetics of precipitation of the organisms are understood, 

the interpretation of the fossil record cannot be unequivocal. Chave (1960) was among the first to state clearly 

the importance of secondary changes of fossil remains 

in the interpretation of the organism relations of the 

original assemblage, especially the selective loss of more 

soluble varieties of skeletons. 

SELECTIVE LOSS OR PRESERVATION 

In the interpretation of differences in the compositions of sedimentary rocks as a function of their ages, 

the effects of diagenetic changes have been stressed as 

causing changes in mineralogy as well as in composition. 

Another factor that is potentially just as important can 

be called selective loss or preservation. Geologists are just 

beginning to come to grips with this problem as information emerges on subduction of sediments and on 

removal of constituents from sea water at ocean ridges. 

Most of the sediment models that have been made assume 

either a closed cycling system or one in which the 

chemistry of new sediments, created by acid volatiles recirculated into the exogenic cycle, is the same as that of the 

lost materials. A few of the relations that are puzzling in 

light of current knowledge are discussed below. 

Mercury in Early Proterozoic shales 

Cameron and Jonasson (1972) showed that mercury 

in Early Proterozoic shales is enriched by a factor of about 

10 over younger shales, and that enrichment is quite 

general. The age range of the Early Proterozoic shales 

sampled is several hundred million years, so that at the 

end of Early Proterozoic time a mass of mercury-enriched 

shales presumably existed comparable to the current 

Mesozoic-Cenozoic shale mass. Of course, most of the 

original Early Proterozoic shales have disappeared, but 

in a cannibalistic sediment system ~heir constituent 

elements should still be cycling. Somehow a large mass 

of mercury has been removed from the exogenic system. 

10 Studies in Diagenesis 

(One might guess subduction of the shales or removal by 

seawater circulation at ridges.) 

The calcium problem 

Sedimentary rocks have a much lower percentage 

of Na 0 than the igneous rocks from which they came; 

but th~ missing sodium seems to be balanced by sodium 

in the ocean and in evaporites. On the other hand, sedimentary rocks, by most measures, have about twice as 

much calcium as their igneous forebears. Calcium from 

basalts is being exchanged for magnesium during the circulation of sea water through ridges-one way to account 

for the calcium excess-but there seems to be no corresponding magnesium depletion in sedimentary rocks as 

compared to the average igneous rock. Sibley and Vogel 

(1976) suggested that the calcium excess may be the result 

of subduction of pre-Mesozoic, calcium-free deep-sea 

sediments. If their sole deficiency was carbonates, calcium 

would indeed be concentrated in the surviving shallowwater sediments. 

Other similar problems 

Selective loss into the mantle of sulfate sulfur and 

oxygen would affect the mean isotopic composition of the 

sulfur remaining in the exogenic cycle, as well as cause 

imbalances in the exogenic oxygen and carbon reservoirs. 

Li's (1972) mass and isotope balance for sediments and 

sea water described above has implications that no large 

fractional loss of sediments to the mantle has taken place. 

Many more relations of this kind must exist that could 

be investigated to examine the degree to which the 

exogenic cycle is a closed system with respect to the mass 

of the circulating elements. 

SUMMARY 

To read the environments of the geologic past and 

to reconstruct the flows of energy and materials, we must 

be able to deduce the original compositions, structures, 

textures, and mineralogies of sedimentary rocks. Diagenetic changes have played a major role in changing 

mineralogies and bulk compositions as well, and an 

attempt has been made here to show that we are in the 

midst of sorting out primary and secondary features. Also, 

our reading of the record can be warped by selective losses 

or gains of materials from the exogenic cycle. 

REFERENCES CITED 

Berner, R. A., 1981, Authigenic mineral formation resulting from 

organic matter decomposition in modern sediments: Fortschritte der Mineralogie, v. 59, p. 117-135. 
Cameron, E. M., 1983, Evidence from early Proterozoic anhydrite for sulfur isotopic partitioning in Precambrian oceans: 

Nature, v. 304, p. 54-56. 

Cameron, E. M., and Garrels, R. M., 1980, Geochemical compositions of some Precambrian shales from the Canadian 

Shield: Chemical Geology, v. 28, p. 181-197. 

Cameron, E. M., and Jonasson, I. R., 1972, Mercury in Precambrian shales of the Canadian Shield: Geochimica et 

Cosmochimica Acta, v. 36, p. 985-1005. 

Chave, K. E., 1960, Carbonate skeletons to limestonesProblems: Transactions of the New York Academy of 

Sciences, Series II, v. 23, p. 14-24. 

Clarke, F. W., 1924, Data of geochemistry: U.S. Geological 

Survey Bulletin 770, 841 p. 

Claypool, G. E., Holser, W. T., Kaplan, I. R., Sakai, H., and 

Zak, 1., 1980, The age curves of sulfur and oxygen isotopes 

in marine sulfate and their mutual interpretation: Chemical 

Geology, v. 28, p. 199-260. 

Degens, E. T., and Epstein, S., 1962, Relationship between 

0 18/016 ratios in coexisting carbonates, cherts, and 

diatomites: American Association of Petroleum Geologists 

Bulletin, v. 46, p. 534-542. 

Engel, A. E. J., Itson, S. P., Engel, C. G., and Stickney, D. M., 1974, 

Crustal evolution and global tectonics; A petrogenic view: 

Geological Society of America Bulletin, v. 85, p. 843-858. 

Garrels, R. M., and Lerman, A., 1981, Phanerozoic cycles of 

sedimentary carbon and sulfur: Proceedings of the National 

Academy of Sciences, U.S.A., v. 78, p. 4652-4656. 

Garrels, R. M., and Mackenzie, F. T., 1971, Evolution of sedimentary rocks: New York, W. W. Norton, p. 229. 

___ 1972, A quantitative model for the sedimentary rock 

cycle: Marine Chemistry, v. 1, p. 27-41. 

___ 1974, Chemical history of the oceans deduced from postdepositional changes in sedimentary rocks, in Hay, W. W., 

ed., Studies in paleooceanography: Society of Economic 

Paleontologists and Mineralogists Special Publication 20, 

p. 193-204. 

Garrels, R. M., and Perry, E. A., Jr., 1974, Cycling of carbon, 

sulfur, and oxygen through geologic time, in Goldberg, E. 

D., ed., The sea, v. 5: New York, John Wiley & Sons, 

p. 303-336. 

Holland, H. D., 1968, The abundance of C02 in the Earth's 

atmosphere through geologic time, in Ahrens, L. H., ed., 

Origin and distribution of the elements: Oxford and New 

York, Pergamon Press, p. 949-954. 

Holser, W. T., and Kaplan, I. R., 1966, Isotope geochemistry 

of sedimentary sulfates: Chemical Geology, v. 1, p. 93-135. 

Hower, J., Eslinger, E. V., Hower, M. E., and Perry, E. A., 

1976, Mechanism of burial metamorphism of argillaceous sediment; 1. Mineralogical and chemical evidence: Geological Society of America Bulletin, v. 87, 

p. 725-737. 

Knauth, L. P., and Epstein, S., 1976, Hydrogen and oxygen 

isotope ratios in nodular and bedded cherts: Geochimica 

et Cosmochimica Acta, v. 40, p. 1095-1108. 

Land, L. S., and Milliken, K. L., 1981, Feldspar diagenesis in 

the Frio Formation, Brazoria County, Texas Gulf Coast: 

Geology, v. 9, p. 314-318. 

Li, Y. H., 1972, Geochemical mass balance among lithosphere, 

hydrosphere, and atmosphere: American Journal of 

Science, v. 272, p. 119-137. 

Perry, E., and Hower, J., 1970, Burial diagenesis in Gulf Coast 

pelitic sediments: Clays and Clay Minerals, v. 18, 

p. 165-177. 

Ronov, A. B., and Migdisov, A. A., 1971, Geochemical history 

of the crystalline basement and the sedimentary cover of 

the Russian and North American platform: Sedimentology, 

v. 16, p. 137-185. 

Sandberg, P. A., 1975, New interpretations of Great Salt Lake 

ooids and of ancient nonskeletal carbonate mineralogy: 

Sedimentology, v. 22, p. 497-537. 

Schidlowski, M., Junge, C. E., and Pietrek, H., 1977, Sulfur 

isotope variations in marine sulfate evaporites and the 

Phanerozoic oxygen budget: Journal of Geophysical 

Research, v. 82, p. 2557-2565. 

Schidlowski, M., Hays, J. M., and Kaplan, I. R., 1984, Isotopic 

inferences of ancient biochemistries-Carbon, sulfur, 

hydrogen, and nitrogen, in Schopf, J. W., ed., Earth's 

earliest biosphere: Princeton, New Jersey, Princeton 

University Press, p. 149-183. 

Sibley, D. G., and Vogel, T. A., 1976, Chemical mass balance of the Earth's crust-The calcium dilemma{?) 

and the role of pelagic sediment: Science, v. 192, 

p. 551-553. 

Veizer, J., Holser, W. T., and Wilgus, C. K., 1980, Correlation of 13C/1

2C and 34S/32S secular variations: Geochimica 

et Cosmochimica Acta, v. 44, p. 579-587. 

Veizer, J., and Jansen, S. L., 1979, Basement and sedimentary 

recycling and continental evolution: Journal of Geology, 

v. 87, p. 341-370. 

Vinogradov, A. P., Ronov, A. B., and Ratynskii, V. M., 1957, 

Variations in the chemical composition of carbonate rocks 

of the Russian platform: Geochimica et Cosmochimica 

Acta, v. 12, p. 273-276. 

Sediment Cycling and Diagenesis 11 

A Kinetic Approach to Chemical Diagenesis 

By Robert A. Berner, Department of Geology and Geophysics, Yale University, 

New Haven, Connecticut 06511 

CC>NTENTS 

Abstract 14 

Introduction 14 

Diagenetic equations and steady state 14 

Kinetics of diagenetic organic matter decomposition 15 

Conclusion 20 

Ac:lcnowledgment 20 

References cited 20 

FIGURES 

1. Diagramatic representation of steady state diagenesis 15 

2. Distribution of dissolved nitrate, manganese, and iron versus depth in a pelagic 

sediment from the equatorial Atlantic Ocean 16 

3. Distribution of dissolved sulfate and methane vs. depth for anoxic sediments from 

the FOAM s:lte of Long Island Sound 18 

4. Plot of rate of organic matter decomposition Ri versus concentration of added 

organic carbon A.G for anoxic sediments from Long Island Sound 19 

5. Log-log plot of the rate constant k for organic matter decomposition via bacterial 

sulfate reduction versus sedimentation rate w for a variety of marine sediments 20 

TABLE 

1. Sequence of organic matter decomposition reactions and corresponding standard 

state free energy changes 16 

F. A. Mumpton, eel., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 13 
Abstract 

The study of chemical diagenesis can be approached from 

three viewpoints: observational, thermodynamic, and kinetic. 

In this paper the kinetic approach is briefly presented in terms 

of one-dimensional concentration-vs-depth diagenetic equations, which include expressions for diffusion, particle bioturbation, sediment burial, pore water flow, and chemical reaction. A simplifying concept for deducing rates of diagenetic 

chemical reactions is that of steady state diagenesis. Application of steady state diagenetic equations to the problem of 

microbiological organic matter decomposition reveals that 

organic matter can be divided into a number of fractions, each 

of which undergoes decomposition via first order kinetics and 

is characterized by a reactivity rate constant, k, whose value 

can range worldwide over a factor of 107• 

INTRODUCTION 

The study of chemical diagenesis (diagenesis is 

defined here as all chemical changes taking place in a 

subaqueous sediment both during and after deposition) 

can be approached in several different ways. First, and 

foremost, one must discern what changes actually take 

place during burial over time and this requires detailed 

mineralogical, petrographic, and chemical observations. 

One might call this the observational approach; it forms 

the foundation for all other approaches. Diagenetic 

changes are studied by examination of ancient rocks which 

have undergone replacement, cementation, compaction, 

etc. or by tracing changes occurring with depth (time) in 

modern, unlithified sediments. 

Once sufficient data have been gathered, one normally encounters difficulties in discriminating between 

diagenetic changes and historical changes, in other words, 

those occurring at the time of deposition. At this point, 

chemical thermodynamics becomes a very useful tool. 

Through calculations of free energy changes and 

equilibrium constants, one can determine where true 

equilibrium lies, and this points to the direction in which 

diagenesis must proceed. For example, if a stratigraphic 

succession shows younger calcitic fossils overlying older 

aragonitic fossils, then the difference cannot be due to 

progressive diagenesis because thermodynamic calculations indicate that calcite is more stable than aragonite 

under diagenetic conditions and therefore cannot convert 

to aragonite with time. This kind of approach can be 

called the thermodyamic approach to diagenesis, and it 

directly complements the observational approach. In fact, 

observed changes over long periods sometimes can be used 

to check thermodynamic predictions. 

Once sedimentologic observations and thermodynamic calculations have been made, the question still 

remains as to how fast and by what chemical mechanisms 

diagenesis take place. This leads to a theoretical or quantitative kinetic approach to diagenesis, a topic which will 

be pursued in the present paper. I will present some rather 

14 Studies in Diagenesis 

simple models which illustrate how the fundamental principles of chemical kinetics can be applied to diagenetic 

problems. Equations are derived which are then applied 

to a common problem, the bacterial decomposition of 

organic matter, to illustrate the power of the kinetic 

approach. For further discussion of diagenetic modeling, 

the reader should consult books on early diagenesis such 

as Berner (1980) and Lerman (1979) which apply kinetics 

to a number of low-temperature geologic processes. 

DIAGENETIC EQUATIONS AND STEADY STATE 

Although there are numerous exceptions, one can 

safely say that, in general, diagenetic changes are mainly 

vertical in nature; that is, diagenesis takes place during 

burial and affects wide areas at the same time. This is 

especially true of fine-grained marine sediments (muds) 

during early diagenesis. Here, vertical changes occur on a 

scale of a few centimeters to a few tens of meters as opposed to lateral changes on the scale of many kilometers. 

If diagenetic changes are predominantly vertical in 

nature, one can describe them simply in terms of depth 

and time and this leads to the concept of the diagenetic 

equation (Berner, 1980). All processes affecting a given 

layer of sediment during burial can be described in terms 

of three processes: diffusion, advection, and chemical 

(including radiochemical) reaction. If changes occur only 

with depth and time, this can be expressed as follows: 

For solid components within the layer: 

~l = D a2cs +ER. 

\ at )L B ax2 s 

(1) 

For dissolved components within the layer: 

( 

a c) = D a

2

c _ v a c + ER, <2> 

at L s ax2 g ax 

where: cs = concentration of a solid component (in 

terms of mass per unit mass of total 

solids) 

c = concentration of a dissolved component 

(in terms of mass per unit volume of 

pore water) 

t = time 

x = depth (measured positively downward) 

v = rate of vertical velocity of the pore water g 

relative to the layer 

Rs' R = rate of each chemical reaction affecting 

cs and c, respectively 

Ds = molecular diffusion coefficient for the 

whole sediment (pore water and 

particles) 

DB = biodiffusion coefficient used to describe 

bioturbational mixing of particles on 

the assumption that it can be treated as 

a random walk process. 
Here, the derivatives (a c; a t)L and (a cl a t)L refer to all 

diagenetic changes occurring in the sediment layer (L) as 

it is buried. Also, to simplify matters it is assumed that 

compaction, and diffusion or flow of pore water due to 

biological activity (irrigation), are all negligible. 

The above equations imply a reference frame for 

depth x fixed to a given sediment layer. When studying 

early diagenesis, a more useful reference is the sedimentwater interface and this is what will be adopted here. 

Diagenesis, by definition, begins at the sediment-water 

interface and continues as sediment layers are buried 

below it. In this case the layer can be considered as being 

buried at a velocity which relates depth to time: 

w = dx/dt. (3) 

With this definition equations 1 and 2 can be transformed 

to: 

-w- + I.J san 

( 

a cs ) -- D a 2cs a cs ~R d 

at X B ax2 ax 

(4) 

-- = D --- (w + v )-+ ER. 

( 

ac ) a

2

c ac 

at X S ax2 g ax 

(5) 

Here the derivatives (a C/ a f) X and (a cl a t)x record all 

changes, both diagenetic and historical, occurring at a 

fixed depth below the sediment-water interface in contrast to equations 1 and 2 which describe only diagenetic 

changes that occur within a fixed layer. 

Adoption of the sediment-water interface as 

reference frame (equations 4 and 5) leads to a most useful 

generalization, that of steady state diagenesis. In steady 

state diagenesis all diagenetic changes are so adjusted that 

no net change in concentration occurs at a fixed depth 

x. This means that for all depths: 

~) = 0. 

at x 

(6) 

Under this situation all changes are diagenetic in origin 

and none are historical. (Strictly speaking constant properties can also be due to the maintenance of chemical 

equilibrium, but this will not concern us here since we 

are interested in the diagenetic approach to equilibrium.) 

This can be seen most clearly for x = 0, the sedimentwater interface. If for all values of t, dc/dt = 0 at x 

= 0, there can be no changes in cs, after burial, due to 

historical fluctuations at the time of deposition. This 

means that a layer presently at a depth x, at the time of 

deposition, originally exhibited the properties displayed 

by the layer presently undergoing deposition. This is 

illustrated in figure 1. The usefulness of the steady-state 

assumption is that diagenetic changes, which occur over 

time, can be witnessed as changes with depth and related 

to time via the parameter w (equation 3). By doing this 

Property p ---+ 

Property p --+ 

Figure 1. Diagrammatic representation of steady state diagenesis. Note that sediment property p within a layer (A) changes 

as the layer is buried, but that the value of p at a fixed depth 

x;, remains constant with time (modified from Berner, 1980). 

we can study chemical kinetics "in the field" and thereby 

extract data on the rates and mechanisms of reactions 

from the sediments themselves. 

KINETICS OF DIAGENETIC ORGANIC MATTER 

DECOMPOSITION 

As a way of illustrating the use of the diagenetic 

equation, a subject is chosen here which is of great interest 

to those studying early diagenesis in modern sediments. 

This is the microbiological decomposition of organic matter. The discussion will focus on the complete transformation of complex, polymeric sedimentary organic matter to simple dissolved molecules and ions and not be 

concerned with the various intermediate steps along the 

way which may involve many different compounds and 

many different microorganisms. In other words, only the 

overall process of organic matter decomposition will be 

considered without involving subjects which more properly fall under the disciplines of organic geochemistry 

and geomicrobiology. 

Organic matter undergoes decomposition via biologically mediated reactions with a variety of oxidizing 

agents, the actual decomposition process depending on 

the chemistry of the pore water. Under oxic conditions, 

the reaction can be simply described as: 

A Kinetic Approach to Chemical Diagenesis 15 
(7) 

where CH20 represents complex, polymeric (solid) 

sedimentary organic matter. This reaction produces abundant energy for metabolic processes and is the preferred 

one, providing 0 2 is present. However, in most marine 

sediments 0 2 becomes completely consumed in the 

uppermost few em to few tens of em, and, continued 

organic matter decomposition proceeds under anoxic conditions (e.g. Berner, 1980). 

Anoxic decomposition processes succeed 0 2 reduction and one another in the order of free energy release 

(e.g. Claypool and Kaplan, 1974; Froelich and others, 

1979). In other words, of a number of available oxidizing agents present, the one yielding the most energy is 

used first until it is depleted. This is followed by utilization of the next most energy-producing agent until it is 

depleted and so on. The processes in order of decreasing 

energy yield are: nitrate reduction (denitrification), 

manganese and iron oxide reduction, sulfate reduction, 

and fermentation. These processes result in the liberation 

to solution of N2 , Mn2+, Fe2+, H2S, and CH4 , as well 

as C02 and HC03 -. Characteristic reactions along 

with the standard state free energy yields are shown in 

table 1. 

In most marine sediments nitrate reduction is unimportant because of the low concentration of dissolved 

nitrate which is rapidly consumed upon burial. Also, iron 

and manganese oxide reduction is very slow, because solid 

compounds are involved. This results in continual liberation of Mn2+ and Fe2+ during the ensuing processes of 

sulfate reduction and fermentative methane formation 

because of the persistence of the iron and manganese 

oxides. Nitrate reduction and iron and manganese oxide 

reduction are most notable in sediments deposited slowly 

in the deep sea where organic matter decomposition is 

slow and the processes of sulfate reduction and methane 

formation occur only at great depth. Some examples, 

taken from the work of Froelich and others (1979), are 

shown in figure 2. 

Table 1. Sequence of organic matter decomposition reactions and corresponding standard 

state free energy changes 

[Organic matter represented by CH2

0 (sucrose)]. [Modified from Berner, 1980] 

Reaction 

CH20 + 0 2 - C02 + H20 

5CH20 + 4N03 - - 2N2 + 4HC03 - + C02 + 3H20 

CH20 + 3C02 + H20 + 2Mn02 - 2Mn+2 + 4HC03 -

CH20 + 7C02 + 4Fe(OH)3 - 4Fe+ + + 8HC03 - + 3H20 

2CH20 + SO/- - H2S + 2HC03 -

2CH20- CH4 + C02 

No; 

CONCENTRATION, IN MICROMOLARITY 

Mn 

0 20 0 r----r-r----.-=-'1 10 20 

CJ) 

a: 

~ 20 

• 20 20 w 

~ 

i= 

z • w 

u 

~ 

::i 

1-

c.. 

~ 40 40 40 

60 60 60 

AG 0 , in kilojoules 

per mole of CH20 

Fe 

2 

• 

-475 

-448 

-349 

-114 

-77 

-58 

4 

Figure 2. Distribution of dissolved nitrate, manganese, and iron versus depth 

in a pelagic sediment from the equatorial Atlantic Ocean (modified from 

Froelich and others, 1979). 

16 Studies in Diagenesis 
The characteristic process of anoxic organic matter decomposition in marine sediments, especially those 

deposited near the continents, is bacterial sulfate reduction (see, for example, Goldhaber and Kaplan, 1974; 

Goldhaber and others, 1977). Most shallowly buried · 

marine muds deposited on the continental shelves exhibit 

evidence for sulfate reduction in terms of decreases with 

depth in dissolved sulfate and the presence of H2

S and 

(or) iron sulfides. The overall reaction for bacterial sulfate 

reduction results in H2S formation: 

and this H2S reacts with detrital iron minerals to form 

various iron sulfides, with the end product being pyrite, 

FeSr 

In marine sediments where unusually high concentrations of readily metabolized organic matter are deposited or in fresh water sediments where initial concentrations 

of sulfate are normally low, sulfate can be completely 

removed, and high concentrations of dissolved methane, 

formed by fermentation, may then appear. An example 

of sulfate reduction and methane formation, as evidenced 

by interstitial water chemistry, is shown in figure 3. 

The microbiological decomposition of organic matter, regardless of process, can be treated quantitatively. 

One approach (see, for example, Jorgensen, 1978; Berner, 

1974, 1980; Westrich and Berner, 1984)_ is to assume that 

a given group of compounds of similar reactivity are 

destroyed at a rate directly proportional to the amount 

present, in other words, decomposition is assumed to 

follow first order kinetics. Mathematically this can be 

expressed as: 

(9) 

where: G; = concentration, expressed in terms of mass 

of carbon per unit volume of pore water, 

of a group of compounds i of similar 

reactivity which can undergo bacterial 

decomposition. 

k; = first order rate constant for the group 

R; = rate of decomposition of group i. 

Here, G; can be thought of as representing the amount 

of organic matter and k; its reactivity. The total organic 

matter in the sediment can be divided into a number of 

different groups of different reactivities (k;) such that: 

(10) 

and 

(11) 

where G T = total organic carbon 

R T = rate of total organic carbon destruction. 

Assuming that the organic matter is present only 

as a solid (that is, it is non-diffusable), which is reasonable 

for the complex polymeric material represented as 

CH20 in table 1, one obtains from the diagenetic equation for solids (equation 4) and equation 9: 

(12) 

Assuming steady state ( (a G/ a t>x = 0) and using the 

boundary conditions, X = 0, Gi = Go-' X - 00 ' G; - 0, 

solution of equation 12 yields: ' 

(13) 

Below the top few tens of centimeters, or for sediments 

overlain by anoxic waters, bioturbation can be ignored. 

Under these conditions, DB = 0, and (equation 13) 

reduces to: 

G1 = G., exp - ( : ) x . (14) 

In terms of rate of decomposition, R;, equations 13 and 

14 can be recast, using equation 9, as: 

(15) 

and: 

R. = R exp - (!!__) x . I 0; W (16) 

By applying equations 13-16 to data for organic matter 

content or decomposition rate in sediments undergoing 

steady state diagenesis, and knowing the values of w and 

DB, one can obtain values fork;, the intrinsic reactivity 

of the organic matter. 

In most studies sediments at a given locality have 

been treated as though they contained only one group of 

organic compounds of one reactivity k. This has led to 

the use of so-called "one-G" models (Berner, 1974; 

Lasaga and Holland, 1976; Toth and Lerman, 1977; Murray and others, 1978; Muller and Mangini, 1980). The 

normal procedure has not been to fit diagenetic equations 

such as 13 and 14 to organic carbon data, but to work 

with pore water data for the concentration of an oxidizing agent, mainly sulfate. For example, the value of k; 

can be obtained, if bioturbation is ignored, by substituting equation 14 into a diagenetic equation for dissolved 

sulfate. This can be done because the rate of sulfate 

Kinetic Approach to Chemical Diagenesis 17 
METHANE CONCENTRATION, IN MILLIMOLARITY 

0 0.1 0.2 0.3 0.4 0.5 

SULFATE CONCENTRATION, IN MILLIMOLARITY 

o~O-----------T5 ____________ 1T0 ____________ 1,5 __________ _,2,0 ____ ~----~25 

(/) 

0::: 

w 

1-

w 

~ 

i= 

20 

40 

60 

ffi 80 u 

z 

:X: 

ll: 

w 

c 

100 

120 

140 

160~----------~----------~------------~----------~----------~ 

Figure 3. Distribution of dissolved sulfate and methane vs. depth for anoxic sediments from the 

FOAM site of Long Island Sound. Vertical bars represent sampled depth ranges (data from Goldhaber 

and others, 1977; Martens and Berner, 1977). 

reduction is directly proportional to the rate of organic 

matter decomposition by this process. From reaction 8: 

(18) 

(17) 

Inserting equation 18 into a diagenetic equation for 

sulfate: 

Where Rs = rate of sulfate reduction, so that, from 

equations 9, 14, and 17: 

18 Studies in Diagenesis 

8c w- - 1/2 k;G0 . exp-(k/w)x, ax I 
where c refers to the concentration of dissolved sulfate 

in interstitial waters. This equation can be solved for 

steady state, and the resulting expression fitted to dissolved sulfate-vs-depth data to obtained values of k;. 

Detailed sampling at some sediment locations has 

revealed that the entire profiles of rate of organic matter 

decomposition-vs-depth cannot be fitted to a simple 

exponential as represented by equations 15 and 16, (e.g. 

see Jorgensen, 1978; Berner, 1980; Westrich, 1983). In 

this case the general or multi-G situation, as represented 

by equations 10 and 11, provides a better description of 

organic matter decomposition and a better fit to the data. 

In other words, there is a successive disappearance with 

depth of progressively less reactive (lower k) compounds. 

Although organic matter reactivity normally varies from 

sediment site to sediment site, it can also vary with depth 

(time) at a single site. 

The necessity for utilizing multi-G models has been 

brought out most forcefully by the work of Westrich 

(1983) (see also Westrich and Berner, 1984). Like most 

previous workers, Westrich also studied organic matter 

decomposition by examining depth changes in the concentration of oxidizing agent, in this case sulfate, rather 

than changes in organic matter itself. Unlike most previous workers Westrich actually measured rates of sulfate 

reduction using a radiotracer technique. In this way he 

could obtain directly, via the stoichiometric relation 8, 

the rate of organic matter decomposition R;. Westrich 

then showed that over successive depth portions of a sediment from Long Island Sound (the FOAM site) the values 

of R; followed exponential relations predicted by equations 15 and 16, but with a different k value for each 

depth portion. The k values obtained by him by curve 

fitting to equations 15 and 16 also agreed with independent rate experiments performed in the laboratory using 

the same muds. 

Westrich and Berner (1984) reported the results of 

additional laboratory experiments which test the validity of equation 9. By subjecting marine plankton to oxic 

degradation in seawater for various periods .of time it was 

possible to prepare freeze-dried samples of marine organic 

matter of varying reactivity. Addition of successively 

greater amounts of each dried sample to anoxic Long 

Island Sound mud resulted in corresponding increases in 

rates of organic matter decomposition via bacterial sulfate 

reduction. A linear correlation was found for each of the 

dried samples which verifies the assumption of first order 

kinetics embodied in equation 9. An example is shown 

in figure 4. Not only was a linear correlation found, but 

the more ( oxically) degraded plankton was found to 

decompose more slowly and, hence, have a lower k value 

as predicted by the multi-G model. 

By piecing together rate constants based on laboratory experiments, measured in situ rates of 0 2 and 

sulfate reduction, and the use of diagenetic equations for 

a wide variety of sediments, Westrich (1983) was able 

"F 

Ill 

Q) 

> 

8 

0 

Q) 

a. 

u 

CONCENTRATION OF CADDED, ll.G 

(grams of C per 1 000 grams) 

Figure 4. Plot of rate of organic matter decomposition R; 

(equal to twice the rate of sulfate reduction) versus the concentration of added organic carbon ~G for anoxic sediments 

from Long Island Sound. Temperature = 22°C. Added organic 

carbon represents dried Long Island Sound plankton which had 

previously undergone oxic bacterial degradation in the 

laboratory for 72 days. Note the linear correlation which verifies 

equation 9. Vertical bars represent total range of measurement 

reproducibility (data from Westrich and Berner, 1984). 

to deduce that sedimentary organic matter, on a 

worldwide basis, can be divided into approximately eight 

groups whose reactivity ranges over seven orders of 

magnitude in k;- However, further measurements may 

show that these are only minimum numbers. The major 

point is that reactivity can vary greatly from place to 

place. 

Regardless of the actual number of groups, it can 

be shown that the rate constant for organic matter decomposition, for marine sediments where only one group 

dominates, is closely correlated with sedimentation rate. 

This is shown in figure 5. The basic reason for this is that 

as depositional burial slows, there is more time for the 

oxic destruction of reactive compounds at the sedimentwater interface so that only the less reactive, or more 

refractory, compounds become buried to undergo further 

decomposition via anoxic processes within the sediment. 

By contrast, with fast sedimentation highly reactive 

substances are buried and react quickly with interstitial 

dissolved sulfate. It is in these rapidly deposited sediments 

A Kinetic Approach to Chemical Diagenesis 19 
where sulfate can become quickly exhausted and methane 

can consequently appear at shallow depths in the sediment. 

CONCLUSION 

From the application of diagenetic equations to the 

problem of microbiological organic matter decomposition 

in sediments, rate constants have been obtained which 

characterize the reactivity of various organic fractions 

towards decomposition. These constants are not only 

useful in themselves for classifying sedimentary organic 

matter, but they can be related to other factors such as 

rates of deposition. This approach should be applicable 

to other diagenetic problems such as authigenic mineral 

formation, dissolution, and replacement. From such 

modeling combined with laboratory rate studies, it should 

be possible to derive much more information on the rates 

and mechanisms of sedimentary chemical reactions. In 

this way we may ultimately gain a better idea of the time 

frame over which diagenesis occurs. 

ACKNOWLEDGMENT 

This research was supported by National Science 

Foundation Grant OCE 81-09454. I am indebted to 

Joseph T. Westrich for helpful discussions and for making available to me some of his unpublished data on 

bacterial sulfate reduction. 

REFERENCES CITED 

Berner, R. A., 1974, Kinetic models for the early diagenesis of 

nitrogen, sulfur, phosphorus and silicon in anoxic marine 

sediments, in Goldberg, E. D., ed., The sea, v. 5: New York, 

Wiley, p. 425-450. 

Berner, R. A., 1978, Sulfate reduction and the rate of deposition 

of marine sediments: Earth and Planetary Science Letters, 

v. 37, p. 492-498. 

Berner, R. A., 1980, Early diagenesis; A theoretical approach: 

Princeton, New Jersey, Princeton University Press, 241 p. 

Claypool, G., and Kaplan, I. R., 1974, The origin and distribution of methane in marine sediments, in Kaplan, I. R., ed., 

Natural gases in marine sediments: New York, Plenum, 

p. 99-139. 

Froelich, P. N., Klinkhammer, G. P., Bender, M. L., Luedtke, 

N. A., Hartman, B., and Maynard, V., 1979, Early oxidation 

of organic matter in pelagic sediments of the eastern 

equatorial Atlantic; Suboxic diagenesis: Geochimica et 

Cosmochimica Acta, v. 43, p. 1075-1090. 

Goldhaber, M. B., Aller, R. C., Cochran, J. K., Rosenfeld, J. 

K., Martens, C. S., and Berner, R. A., 1977, Sulfate reduction, diffusion, and bioturbation in Long Island Sound 

sediments; Report of the FOAM group: American Journal 

of Science, v. 277, p. 193-237. 

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

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

p. 569-655. 

Jorgensen, B. B., 1978, Comparison of methods for the quantification of bacterial sulfate reduction in coastal marine 

20 Studies in Diagenesis 

5 

en 

~ 4 w 

>-

z 

0 

~ 3 

~ 

..J 

<( 

g 2 a: 

a.. 

u 

w 

a: 

z 

~ 

(.!) 

9 0 

-1 

-2L---~----~----~--_. ____ _. ____ ~ 

-2 -1 0 1 2 3 4 

LOG w, IN CENTIMETERS PER 1000 YEARS 

Figure 5. Log-log plot of the rate constant k for organic matter decomposition via bacterial sulfate reduction versus 

sedimentation rate w for a variety of marine sediments ranging 

from harbor muds to deep sea clays. (Modified from Toth and 

Lerman, 1977; Berner, 1978; Berner, 1980.) 

sediments II. Calculation from mathematical models: Geomicrobiology Journal, v. 1, p. 20-47. 

Lasaga, A. C., and Holland, H. D., 1976, Mathematical aspects 

of nonsteady state diagenesis: Geochimica et Cosmochimica 

Acta, v. 40, p. 257-266. 

Lerman, A., 1979, Geochemical processes: Water and sediment 

environments: New York, Wiley, 481 p. 

Martens, C. S., and Berner, R. A., 1977, Interstitial water 

chemistry of anoxic Long Island Sound sediments. 1. 

Dissolved gases: Limnology and Oceanography, v. 22, 

p. 10-25. 

Muller, P. J., and Mangini, A., 1980, Organic carbon decomposition rates in sediments of the Pacific manganese nodule 

belt dated by 23<Yfh/231Pa: Earth and Planetary Science 

Letters, v. 51, p. 94-114. 

Murray, J. W., Grundmanis, V., and Smethie, W. M., 1978, Interstitial water chemistry in the sediments of Saanich inlet: 

Geochimica et Cosmochimica Acta, v. 42, p. 1011- 1026. 

lbth, D. J., and Lerman, A., 1977, Organic matter reactivity and 

sedimentation rates in the ocean: American Journal of 

Science, v. 277, p. 265-285. 

Westrich, J. T., 1983, The consequences and controls of bacterial 

sulfate reduction in near shore marine sediments: 

unpublished Ph.D. dissertation, New Haven, Connecticut, 

Yale University, 530 p. 

Westrich, J. T., and Berner, R. A., 1984, The role of sedimentary organic matter in bacterial sulfate reduction; the G 

model tested: Limnology and Oceanography, v. 29, 

p. 236-249. 
Seasonal Diagenetic Changes in Salts of Owens Lake, 

California, 1970-77 

By GEORGE I. SMITH, U.S. Geological Survey, 345 Middlefield Road, MS 902, 

Menlo Park, California 94025, and 

IRVING FRIEDMAN, U.S. Geological Survey, Box 25046, Denver Federal 

Center, MS 963, Denver, Colorado 80225 

CONTENTS 

Abstract 22 

Introduction 22 

Study methods 22 

Salt crystallization 22 

Mineralogic and isotopic evidence of diagenesis 24 

Conclusions 28 

References cited 28 

FIGURES 

'!I 

1. Map showing location of Owens Lake and positions of lake shorelines February 1913, August 1969, and 

April 1971 23 

2. Plot of changes observed at Owens Lake between September 1970 and August 1971 26 

3. Plot showing ranges of oD values of salts and interstitial brines, 1970-77 28 

TABLE 

1. Data for Owens Lake brines and salts, 1970-77 24 

F. A. Mumpton, eel., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 21 
Abstract 

After the 1968-69 flooding and partial solution of salts 

on the floor of normally dry Owens Lake, salts were redeposited 

over a 12-month period starting September 1970. Over the next 

several years, seasonal changes in the mineralogy and 

deuterium/hydrogen (D/H) ratios of the new hydrated salts and 

their interstitial brines indicated continuing postdepositional 

reactions, or diagenesis. Winter cooling promoted crystallization of natron and mirabilite; summer warming promoted 

recrystallization ofthese minerals to trona and burkeite. Natron 

and mirabilite crystallized late each autumn, incorporating Na, 

C03' and 504 from trona and burkeite, and 10 moles of H20 

each from the isotopically heavy lake brine produced by 

summer-long evaporation. Trona and burkeite crystallized each 

spring, with trona deriving Na and C03 from natron, and H2

0 

mostly from the decomposing natron and mirabilite rather than 

from the overlying lake, which was isotopically lighter at the 

end of the wet season. During diagenesis, the initial wide range 

in the D/H ratios in trona narrowed, and the ratios in both trona 

and interstitial brines approached those of the underlying 

undissolved salts that crystallized-and presumably followed 

the same diagenetic path-earlier this century. 

INTRODUCTION 

Owens Lake, in southeast California (fig. 1), desiccated in 1921 after its water supply was diverted to Los 

Angeles, and a 1.5- to 2.0-m-thick layer of salts was 

deposited. At the end of the abnormally wet 1968-69 

winter, Owens River flood waters partially dissolved that 

surface salt layer, and those saline components 

recrystallized during 1970 and 1971 as the lake redesiccated. Samples of the lake brines and crystallizing 

salts were collected during 1970-71 (Friedman and others, 

1976; Smith and others, in press), and samples and cores 

of the older undissolved salts and the recrystallizing new 

salts were taken periodically between '1970 and 1977. 

Mineral compositions of the salts were monitored, as were 

the deuterium concentrations in the hydrogen-bearing 

saline minerals and interstitial brines. Seasonal and 

longer-term changes in both mineral compositions and 

isotope ratios were observed. The deuterium concentrations in the 1970-71 salts and brines initially differed 

greatly from those of the older undissolved salt layers, 

but over several years, their values slowly converged as 

a result of diagenesis. The diagenetic process was driven 

by seasonal temperature changes in the saline layers, and 

our observations suggest that it was approaching completion only after 7 years of repeated cycles. What follows 

is a description of those observations. 

This process may characterize the early histories of 

most saline deposits that contain sodium sulfate and 

sodium carbonate mineral suites which include phases 

whose solubilities are markedly temperature dependent. 

Care must be exercised in extracting paleoclimate information from the mineralogic and isotopic character of 

22 Studies in Diagenesis 

deposits of these types, because minerals containing these 

components are unlikely to be the crystals that originally 

precipitated from lake waters. Evidence for diagenesis of 

this type, however, documents minimum limits on 

seasonal temperature changes at the time of original 

deposition, and the suite that is eventually buried to a 

depth at which annual temperature variations cease 

should reflect mean annual temperatures. 

STUDY METHODS 

Saline minerals settling from the lake surface were 

collected by underwater traps, and sediments accumulating on the lake floor during and after desiccation were 

sampled as cores. All samples were collected at a station 

near the lowest part of the lake floor (fig. 1). Coring, 

done with several types of portable equipment, was difficult because of the varied physical properties of these 

salts. Layers dominated by halite and burkeite were hard, 

tended to plug the core barrel, and prevented the recovery 

of salts from deeper layers. Mud and layers composed 

of trona, nahcolite, natron, or mirabilite tended to be soft 

and were commonly pushed aside by the core barrel or 

fell out as the core barrel was raised. 

The interstitial brines were extracted from freshly 

cored salts for isotope analysis by squeezing them in a 

reinforced syringe through a screen-supported layer of 

filter paper into ~bottle which was promptly capped. The 

dewatered salts were left in the sealed syringe until they 

reached the laboratory. Although those samples proved 

to be a little more stable than samples collected and 

transported wet, most samples composed of natron or 

mirabilite were partially or totally altered to other 

minerals by the time they reached the laboratory. For 

deuterium analysis, this alteration posed no problem 

because all water of hydration that was expelled or recrystallized during spontaneous phase changes remained 

in the syringe and was reflected in the isotope analysis 

of the bulk sample. Determination of the original lakefloor mineralogy was more difficult. Fortunately, most 

samples were also studied in the field immediately after 

collection using a petrographic microscope, so the lakefloor mineralogy could be reconstructed from field descriptions of the salt-crystal habits and the physical properties of the original samples, plus laboratory determinations of the percentage of crystallization water and the 

X-ray diffraction mineralogy of the secondary suites. 

SALT CRYSTALLIZATION 

Calculations based on data of Gale (1914, p. 258, 

fig. 58),Friedman and others (1976, table 1), and Smith 

and others (in press) show that about 20 percent of the 

1.5-m- to 2.0-m-thick bed of salts that crystallized from 
118° 00' 

5 MILES I I I I 

I I I I I 

5 KILOMETERS 

Figure 1. Location of Owens Lake and positions of lake shorelines February 1913 (prior to its first 

desiccation), August 1969 (maximum area during period discussed here), and April1971 (4 months 

prior to reaching near-dryness). Sample site (black circle) is near lowest point in lake floor. 

the lake earlier this century was dissolved in 1969-70. The 

first new salts precipitated from the lake in mid-September 

1970 when brine concentrations (table 1) reached about 

300 g/L (grams per liter). In most places, the new salts 

were deposited on an older layer of hard or muddy salts 

that apparently inhibited deeper solution and later served 

as a stratigraphic marker in the cores. 

The species of crystallizing salts and the deuterium 

concentrations of the hydrated species are also listed in 

table 1, and the sequential changes in the compositions 

of solids and brines during this crystallization period are 

shown diagrammatically in figure 2. In general, Nacarbonates crystallized first, Na-sulfates second, and NaCl 

last, the classical crystallization sequence for systems 

having these compositions. The species of crystallizing 

salts, however, and the time of day that crystallization was 

most rapid, depended on the season. During winter, 

radiative overnight cooling of the lake surface water 

accelerated the predawn crystallization of two 10-hydrate 

minerals on the surface, natron (Na2C03 ·10H20) and 

mirabilite (Na2SO 4 ·10H2

0). Their crystals coalesced into 

large floating rafts, some of which dissolved as the 

surface warmed the next morning although others 

foundered earlier and sank. Crystals were not forming on 

the lake surface at the time of the early morning January 

1971 collection, but several centimeters of rain had fallen 

on the lake in late November and December, and an 

abnormally warm day preceded the collection date. Hence, 

the lake surface water did not cool sufficiently overnight 

to reach its new saturation temperature of about 5 OC. By 

April1971, however, the lake brine and the accumulating 

salt layer had warmed, and all the mirabilite on the lake 

floor had returned to solution and most of the natron had 

dissolved or reacted diagenetically to form nahcolite 

(NaHC03

) or trona (Na2C03 • NaHC03 • 2H20), thus 

reducing the thickness of the new salt layer and increasing both the total amount of salt in solution and the brine 

concentrations (fig. 2B). By July, most of the dissolved 

sulfate had crystallized from the brine as burkeite 

(Na2

C03 • N~SO J which is stable only above 14 °C; all 

other Na-carbonates had crystallized to trona, and 

evaporation that reached its daily maximum in midafternoon was producing large rafts of crystals dominated by 

halite (NaCl). By August, the lake was nearly dry. 

Seasonal Changes in Salts, Owens Lake, Calif. 23 
Table 1. Data for Owens Lake brines and salts, 1970-77 

[g/L, grams per liter; g/mL, grams per milliliter; OC, degrees Celsius; leaders (---) indicate no measurements] 

Brine1 Salts 

Examples of diagenetic 

Collection date Total dissolved solids Density Temperature 1502 Initial salts salts observed later 

~antity 1 in same horizon {1970-71} Concentration (g/ml) (OC) {1970-71} 

(g/L) (tonnes x 106) Mineral (s) 602 Minerals 602 

(per mil) (per mil) 

Sept. 16, 1970 296.6 26.7 1.20 15-16 -21 Nahco 1 ite, -106 Trona------ -89 to -126 

trona. 

Nov. 18, 1970 6-11 -18 Nat ron. +5 Trona, -85 to -95 

Mi rabil i te. -20 burkeite. 

Jan. 19, 1971 239.8 14.6 1.17 5-8 -44 Natron, -2 to -15 Trona, -85 to -95 

mi rabi 1 ite. burkeite. 

March 10, 1971 11-12 -36 Trona, -26± Trona------ -85 to -95 

nahcolite 

nat ron. 

April 13, 1971 387.5 17.7 1.26 13-20 -18 Nahcolite, -56± Trona------ -85 to -95 

trona. 

July 7, 1971 450.0 12.5 1.30 27-28 -2 Halite, Halite, 3-55? 

burkeite, burkeite, 

trona. trona. 

Aug. 20, 1971 470.3 8.4 1.31 +3 Ha 1 ite, Halite, 3-55? 

burkeite, burkeite, 

trona. trona. 

1oata from Friedman and others (1976, tables 1, 2, and 4, and fig. 3); "Quantity" of total dissolved solids calculated by assuming quantities of dissolved Li remained constant. 

2values of 60 are per mil SMOW; values listed for salts are on mixtures of ~drated minerals except for November "Initial 

salts" samples which were determined on separated samples of listed species. 

3Percentage of trona is about 15 percent; 60 values probably also reflect brine inclusions and film adhering to crystals of 

slightly moist samples. 

MINERALOGIC AND ISOTOPIC EVIDENCE OF 

DIAGENESIS 

Isotopic characterization of the deuterium/ 

hydrogen (D/H) ratios in the hydrated salts and brines 

collected during the desiccation and diagenetic stages of 

Owens Lake is expressed as oD per mil, relative to Standard Mean Ocean Water (SMOW), where 

oD = D/H (sample)- D/H (standard) X 1000. 

D/H (standard) 

Laboratory studies of the D/H fractionation factors for 

the minerals trona, nahcolite (Matsuo and others, 1971), 

and mirabilite (Stewart, 1974) show that relative to the 

crystallizing solution, nahcolite and trona are depleted 

in deuterium and mirabilite is enriched; these data allow 

us to make this preliminary isotopic evaluation of the 

processes operating during crystallization and diagenesis 

of those minerals. Isotope fractionation factors for 

natron, a mineral discovered to be important in this process, have not yet been determined satisfactorily in the 

laboratory, but the following field evidence suggests that 

at the nighttime temperatures frequently reached by 

24 Studies in Diagenesis 

Owens Lake surface water during the autumn and winter 

seasons, natron was enriched in deuterium by 20-25 per 

mil relative to the brine. The lake surface brines during 

October and November had oD values between -16 and 

-20 per mil, and natron, recovered in a core obtained in 

mid-November that represented crystallization after midSeptember, had a oD value of + 5 per mil. Between late 

November 1970 and January 1971, about 8 em of rain 

fell, altering the oD composition of the lake to -45 per 

mil, reducing the concentration of solids in its brines, and 

slowing or suspending salt crystallization (fig. 2). By midJanuary, mixtures of natron and mirabilite, crystallized 

at unknown times during the brine's isotopic transition, 

had oD values ranging from -2 to -17 per mil, further 

suggesting that both minerals concentrate deuterium. 

Salts crystallizing at the surface and collected in 

underwater sediment traps in September 1970 were composed mostly of nahcolite (fig. 2), which had a deuterium 

concentration (oD) value of -106 per mil. In November, 

a layer representing the same salts collected in the 

September trap was recovered in a core. By then, in the 

new environment produced by burial in an accumulating 

salt layer, nahcolite had reacted diagenetically to form 

trona, which also had a oD value of -106 per mil SMOW. 
The reaction of nacolite requires the addition of new 

hydrogen from water: 

3NaHC03 + H20 

= Na2C03 ·NaHC03 ·2H20 + C02 • (1) 

Calculations based on the known isotope fractionation 

factor of trona (Matsuo and others, 1971) show that 

recrystallization took place under equilibrium conditions 

at lake-floor temperatures averaging 120C (observed range, 

8°-16°C), and that the required hydrogen was derived 

almost entirely from lake-bottom water which served as 

an open-system reservoir of nearly constant composition 

(oD = -16 to -20 per mil) during this period. In a closed 

system, with hydrogen derived solely from the saline layer 

itself (a mix of approximately three moles of nahcolite and 

one mole of interstitial brine having oD = -21 per mil), 

calculations based on equation 1 show that the secondary trona would have had a oD value near -78 per mil. 

By April1971, as the lake water and floor warmed, 

more than half the mirabilite had dissolved; about 10 percent of the natron had also dissolved, some was preserved, 

and the rest reacted in place to become trona and 

nahcolite. The mixtures of Na-carbonate minerals had oD 

values ranging from -26 to -101 per mil. Those reactions 

release hydrogen in the form of water and proceed as 

follows: 

3Na2C03 ·10H20+C02 

= 2Na2C03 • NaHC03 • 2H20+25H20 (2) 

and 

The large amounts of water expelled by reactions 2 and 

3, plus that liberated by the solution of mirabilite, were 

isotopically heavier (oD = +5 to -17 per mil) than the 

lake brine (oD = -44 to -18 per mil) during the period 

in which the trona and nahcolite crystallized. The isotope 

composition of those crystals, while growing 

diagenetically in the accumulating salt bed, had to be 

influenced by the brines released by reactions 2 and 3, and 

we conclude that they were in fact almost unaffected by 

the isotope chemistry of the overlying lake water. At 

5°-15 °C, the estimated temperature range in the upper 

part of the salt bed during this season, trona crystallizing under equilibrium conditions would have oD values 

that were depleted 95-85 per mil relative to the brine from 

which they crystallized (Matsuo and others, 1971). 

Diagenetic trona collected from these zones in July 1971 

had oD values ranging from -95 to -85 per mil, implying 

equilibrium crystallization from brines having oD values 

ranging from + 10 to -10 per mil. These values are close 

to those of the mixtures of 10-hydrate minerals that 

crystallized the previous winter, but they do not resemble 

those of the overlying lake brine during the period of trona 

crystallization. 

Mineralogic documentation of seasonal diagenetic· 

processes taking place between 1971 and 1977 was more 

difficult. The existence in cores of 10-hydrate phases during winter was never confirmed by X-ray diffraction 

because of their instability after collection. Repeated 

seasonal change in the salt mineralogy and brine composition {)f Ow,ens Lake has been reported previously 

(Dub, 1947, p. 2), and for the following three additional 

reasons we conclude that both 10-hydrate phases formed 

during most winters of the 1971-77 period. The first 

reason is that in the Na2C03-NaHC03-NaCl and 

Na2C03-Na2SOrNaCl systems, the interstitial brine 

(regardless of whether it resembled its September 1970 

composition when crystallization started or its August 

1971 composition when it ceased) was in the natron 

stability field, or on the boundary between it and the 

mirabilite field, at all temperatures lower than about 200C 

(Freeth, 1923; Teeple, 1929); temperatures of the salt layer 

at depths of about 1 min January of both 1971 and 1972 

were about 13 °C. 

The second reason is that in cores collected in 

January 1972 (samples and data we have not yet interpreted or otherwise introduced here), a layer composed 

entirely of trona in April 1971 was composed of salts that 

were observed dehydrating in the air shortly after being 

brought to the surface. Natron and mirabilite are the only 

Na-carbonate and Na-sulfate minerals having this property. By the time the wet cores reached the X-ray diffraction laboratory, trona and burkeite were the only minerals 

in these suite that were detected. For isotope analysis, 

however, interstitial brines were also extracted in the field 

from a duplicate set of samples, leaving a nearly dry and 

less reactive mix in the sealed syringe. The percentages 

of hydrogen in the dewatered salts, determined during 

isotope analysis and expressed as H2

0, ranged from 2 to 

18 percent in all but one sample which had a value of 27 

percent. In pure samples, these percentages are nahcolite 

11, trona 20, mirabilite 56, and natron 60. It appears 

therefore, that a small amount of natron or mirabilite was 

still preserved in that one sample, and that larger amounts 

were probably present in the original core. No cores were 

collected during the midwinter season in later years. 

The third reason for concluding that repeated 

annual cycles involving phase changes were responsible 

for these observations is our lack of an alternative 

explanation for some of them. In late 1970, a range of 

more than 100 per mil in oD values existed in various mixtures of hydrated minerals on the lake floor (fig. 3). By 

197 4, this wide variation in oD values had decreased 

markedly, and between 1974 and 1977, the variation continued to decrease and approached the oD values observed 

throughout the study in the undissolved pre-1970 salt bed 

(fig. 3). The oD values of the 1970-77 interstitial brines 

also approached those in the older deposits. Thus, some 

Seasonal Changes in Salts, Owens Lake, Calif. 25 
BRINE 

NAHCOLITE 

TRONA 

C/) 

m NATRON 

<( 

::I: 

Cl. 

MIRABILITE 

BURKEITE 

HALITE 

A 

I 

~lr 

' ' 

I 

SEPT 

I 

I I I I 

'......._ 

, , ,lr 

I 

1r 1r/ 

I I I I 

OCT NOV DEC JAN 

I I I I 

1970 

i I j t I I I I 

I ~r I I 

I i I~'"' 

I I 1• I ! ' .........._.......,.. 1r ~lr 

I I ~ I l 

I I ,

1 , I i 

? I 

I 

• ,lr 

I I I I I ,j, I ,r 

FEB MAR APR MAY JUNE JULY AUG 

I I I I I I 

1971 

Figure 2 (above and facing page). Changes observed in Owens Lake between September 1970 (when salts first crystallized) 

and August 1971 (when lake had nearly dried). A, Phase changes: crystallization of listed mineral phases from brine phase 

(solid lines) and their diagenesis or resolution (dashed lines). 8, Brine changes: variations in concentrations and quantities 

of solids remaining in solution, in oD values, and in lake-floor temperatures. SMOW, Standard Mean Ocean Water. 

form of seasonal evolution that included the destruction 

and recrystallization of hydrated salt species had to be 

responsible for the continuing shifts in oD values toward 

the older ones. 

Together, these observations suggest a scenario by 

which the original range of observed oD values would 

converge on the theoretical values that should result from 

equilibrium crystallization at the observed temperatures, 

and the smaller shifts in the oD values of the interstitial 

brines can be reasonably explained. The seasonal evolution of mirabilite and burkeite is not reconstructed 

because they account for only about 15 percent of the 

total assemblage and the warm-season sulfate contains 

no water of hydration; this scenario traces only the evolution of the natron and trona: 

1. Early winter diagenetic recrystallization of trona 

to natron is represented by equation 2 reacting toward 

the left. In addition to 2 moles of trona, it requires 25 

moles (67 volume percent) of water which probably 

exhausted the interstitial brines and extracted water from 

the overlying lake. In midsummer to late autumn, prior 

to significant rainfall on the lake, the oD values of lake 

and interstitial brines collected in 1973, 1974, and 1977 

averaged -33, -34, and -41 per mil (range, -30 to -50 

per mil). Lake waters became isotopi~clly lighter after the 

rains began. Assuming a fractionation factor of -25 per 

mil, crystallization of natron from a -35 to -50 per mil 

26 Studies in Diagenesis 

brine would produce crystals with oD values of -10 to 

-25 per mil. 

2. Regeneration of trona the following spring, at 

5° -15 °C, in a bed with its interstitial brines dominated 

by the expelled water of crystallization from natron would 

have produced trona with oD values between -85 and 

-95 per mil lighter-values at the isotopically light end of 

all ranges observed during the study and very close to the 

final range of values observed in September 1977 (fig. 3). 

The shift toward isotopically lighter interstitial brines after 

July 1971 (fig. 3) suggests that by the end of the 6-year 

period, about half the 25 moles of water required for each 

mole of natron had come from the overlying lake after 

it had received some rain (average oD values about -75 

per mil; Friedman and others, 1976, p. 509) or a smaller 

amount of inflowing surface water (which is isotopically 

even lighter). The heavy end of the range of isotopic values 

in the salts also shifted greatly toward the light end in response to the introduction of this water. The isotopic compositions of the salts and interstitial brines in November 

197 4 represent an intermediate point in such change. 

3. We finally hypothesize that prior to each cool 

season, the deuterium content of lake water again 

increased as a result of summer and autumn evaporation, 

bringing the lake and shallow interstitial brines back to 

oD values of about -35 per mil and allowing the same 

mineralogic and isotopic cycle to be repeated. 
a: 500 w 

..... 

::::; 

a: 

w 

a.. 

1/'J 

:::E 

<( 

a: 

(!) 

~ 

z -(Start of rain season 0 

~ o----- a: 1 ..... 

z 1 w ' u 

z ' 0 

u 200 

cg 

0 30 

x 

1/'J " w 

z ' z ' 0 ' ..... 

~ ' ' z ' 0 ' i= 

...I 

::::> '--- 0 

1/'J 

~ 

1/'J 

~ <( 

1/'J 

0 

...I 

~ 

a: 10 

w 

a.. 

~ 0 

~ 0 

:::E -10 1/'J 

0 

GO 

z 

~ 

-20 ~1 w 1 ..... -30 z \ 0 

u 1 

:::E -40 1 ::::> ' a: 

w \... ........ ..... -so ::::> 

w 

0 

30 

SEP OCT NOV DEC JAN FEB MAR APR MAY JUNE JULY AUG 

B 

Seasonal Changes in Salts, Owens Lake, Calif. 27 
NOV. 1970-

JAN. 1971 

APR. 1971-

JULY1971 

NOV. 1974 

SEPT. 1977 

-100 -so 

8D(SMOW) (PER MIL) 

0 +25 

Figure 3. Ranges of oD values measured in samples of salts (solid lines) and 

interstitial brines (dashed lines, 1970-77). Values measured from salts and brines 

representing undissolved (pre-1970) salt beds, collected on indicated dates, 

fall within ranges indicated by stippled areas; those measured from samples 

representing post-1970 salt beds, collected on the same dates, fall within ranges 

bounded by hachured lines. 

With time, the variable distribution of isotopes 

created during original crystallization from the lake was 

replaced by the less variable distribution created at the times 

of diagenesis. The original hydrogen-bearing phases represented crystallization produced by short-term, mostly 

diurnal, extremes in atmospheric temperature that needed 

to affect only the surface millimeter of lake water, whereas 

the diagenetic phases represented crystallization in response 

to seasonal temperature changes that lasted long enough 

to warm or cool the entire brine body and salt bed. With 

further burial, seasonal temperature variations will be 

reduced, and the "last" diagenetic changes will occur under 

conditions that approximate mean annual values. 

CONCLUSIONS 

The mineral species observed crystallizing from the 

brines of Owens Lake in 1970-71, and the subsequent mineralogic and isotopic changes in the recrystallizing layer of 

salts, were results of seasonal variations in the temperatures 

of the lake brines and near-surface deposits. Winter cooling produced the 10-hydrate minerals, natron and mirabilite, 

and summer warming caused these minerals to dissolve and 

28 Studies in Diagenesis 

their componenets to recrystallize as trona and burkeite. 

Once crystallized, very little halite dissolved seasonally. 

With time, all original carbonate and sulfate minerals 

underwent one or more cycles of diagenetic reactions, and 

the isotopic compositions of the brines and Na-carbonate 

salts became more homogeneous; oD values for trona 

evolved toward -100 per mil, for natron toward -10 per mil, 

and for late summer lake and interstitial brine toward -35 

per mil. They also evolved toward the isotopic composition of the brines and undissolved salts in the pre-1970 

saline layers, which probably attained their isotopic makeup 

by the same seasonal diagenetic process after they crystallized earlier this century. Clearly, however, the seasonal 

mineralogic changes and the accompanying partitioning 

of deuterium and hydrogen between brines and salts is a 

complex process, and additional data are needed to refine 

these conclusions further. 

REFERENCES CITED 

Dub, G. D., 1947, Owens Lake-Source of sodium minerals: 

American Institute of Mining and Metallurgical Engineers 

Thchnical Publication 2235, 13 p. 
Freeth, F. A., 1923, II. The system N~O-C0 -NaCl-H 0 considered as two four-component systems: Philosophicallhmsactions of the Royal Society of London, Series A, v. 223, 

p. 35-87. 

Friedman, Irving, Smith, G. 1., and Hardcastle, K. G., 1976, 

Studies of Quaternary saline lakes-11. Isotopic and compositional changes during desiccation of the brines in Owens 

Lake, California, 1969-1971: Geochimica et Cosmochimica 

Acta, v. 40, p. 501-511. 

Gale, H. S., 1914, Salines in the Owens, Searles, and Panamint 

basins, southeastern California: U.S. Geological Survey 

Bulletin 580-L, p. 251-323. 

Matsuo, Sadao, Friedman, Irving, and Smith, G. 1., 1971, Studies 

of Quaternary saline lakes-1. Hydrogen isotope 

fractionation in saline minerals: Geochimica et Cosmochimica Acta, v. 36, p. 427-435. 

Smith, G. I., Friedman, Irving, and McLaughlin, R. J., 1987, 

Studies of Quaternary saline lakes-III. Mineral, chemical, 

and isotopic evidence of salt solution and crystallization 

processes in Owens Lake, California, 1969-1971: Geochimica et Cosmochimica Acta v. 51, no. 4. 

Stewart, M. K., 1974, Hydrogen and oxygen isotope fractionation during crystallization of mirabilite and ice: Geochimica 

et Cosmochimica Acta, v. 38, p. 167-172. 

Teeple, J. E., 1929, The industrial development of Searles Lake 

brines with equilibrium data: American Chemical Society 

Monographic Series 49, 182 p. 

Seasonal Changes in Salts, Owens Lake, Calif. 29 

Influence of Iron and Manganese on the Chemical 

Partitioning of Copper, Zinc, and Chromium during 

Early Diagenesis in Outer Continental Shelf 

Sediments from the Gulf of Mexico 

By LORRAINE H. FILIPEK, U.S. Geological Survey, 913 National Center, 

Reston, Virginia 22092 

CONTENTS 

Abstract 32 

Introduction 32 

Acknowledgments 32 

Methods 32 

Sampling area 32 

Sampling procedure 34 

Analytical methods 34 

Results and Discussion 34 

Iron 34 

Manganese 38 

Copper 40 

Zinc 41 

Chromium 43 

Conclusions 43 

References cited 43 

FIGURES 

1. Map of south Texas continental shelf 33 

2. Flow chart of extraction technique 38 

3-6. Graphs showing distribution of: 

3. Nonlithogenous iron and dissolved iron 39 

4. Nonlithogenous manganese and dissolved manganese 40 

5. BOAc-extractable manganese and calcium 41 

6. Nonlithogenous zinc, copper, and chromium 42 

TABLE 

1. Correlation coefficients of nonlithogenous metals in cores 35 

APPENDIXES 

1. Concentrations of sulfur, carbon, calcium, sulfate, iron, and manganese in 

cores 45 

2. Concentration of iron in extracts from cores 46 

3. Concentration of manganese in extracts from cores 47 

4. Concentration of copper in extracts from cores 48 

5. Concentration of zinc in extracts from cores 49 

6. Concentration of chromium in extracts from cores SO 

F. A. Mumpton, eel., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 31 
Abstract 

To study how iron and manganese influence the chemistry 

of copper, zinc, and chromium in continental-shelf sediments 

during early diagenesis, a sequence of six extractions was used 

to isolate various forms of these five metals from the uppermost 1.5-2 m of sediment from the outer continental shelf off 

the coast of Texas. Three cores were collected from lowcarbonate sediments in which sulfate reduction was minor. 

Concentrations in the extracts suggest that the geochemistry 

of all five metals changed during diagenesis; however, only 

manganese showed signs of vertical remobilization and only 

at the shallowest site (water depth = 32m), which possessed 

the strongest redox (oxidation-reduction) gradient. In all cores 

the redox conditions appeared to be too low for the formation of extensive manganese oxides; hence, most of the 

nonlithogenous manganese was present as a carbonate or in 

an exchangeable form soluble in 1 molar acetic acid (HOAc). 

Hydrous iron oxides and organically bound iron were the 

major nonlithogenous iron components in the more oxidized 

parts of the cores. Iron in pyrite increased with depth 

throughout the core lengths. Coprecipitation and dissolution 

of trace metals with hydrous iron oxides was in the order 

Zn>Cu>Cr. In reducing parts of the core, copper sulfides were 

also a dominant form of non lithogenous copper, especially at 

the shallowest site. Both organic materials and sulfides appeared 

to be important in the diagenesis of zinc, whereas organic 

materials appeared to be the most important sink for chromium. 

Exchangeable (HOAc-soluble) Zn and Cr correlated strongly 

with HOAc-soluble Fe, which appeared to be a reduced form 

of iron; however, HOAc-soluble Cu, where measurable, was 

independent of other metals. Of the total chromium content, 

10-20 percent was in non lithogenous components; of iron, 

40-50 percent; of zinc, 50-60 percent; of copper, 30-65 percent; and of manganese, 75-90 percent. Only copper, however, showed a significant trend with grain size of the sediment 

and distance from shore; the percentage of the total copper 

content present in non lithogenous components increased from 

the nearest-shore sandy silts to the shelf-edge muds. 

Outer continental-shelf sediments provide a geographically large environmental transition between organic-poor, 

suboxic pelagic sediments and organic-rich, anoxic nearshore 

sediments. Based on the present results, the chemistry of this 

shelf environment appears to be controlled mainly by the 

precipitation and dissolution of hydrous iron oxides, accompanied by minor sulfate reduction. Even though reduced iron 

is in excess over sulfide, copper and zinc appear to compete 

successfully for the sulfide so that the shelf sediments provide 

a sink for these trace metals and act as an important control 

in their geochemical cycling. 

INTRODUCTION 

Many surficial enrichments of metals in sediments 

have been attributed to pollution (Nriagu and others, 

1979; Hamilton-Taylor, 1979); however, natural 

diagenetic processes may also redistribute metals to the 

surface of, or within, a sediment column. For example, 

the diagenetic remobilization of manganese is a well32 Studies in Diagenesis 

documented phenomenon (Robbins and Callender, 1975; 

Holdren and others, 1975). Similarly, iron can also 

undergo extensive chemical redistribution during 

diagenesis due to its sensitivity to changes in redox 

(oxidation-reduction) potential. Because the oxides of 

both iron and manganese are highly effective scavangers 

of trace metals, the reduction of these oxides would probably result in the mobilization of associated trace metals. 

The purpose of this study was to investigate the 

influence of iron and manganese diagenesis on copper, 

zinc, and chromium in low-carbonate, outer continentalshelf sediments in which sulfate reduction is minor. The 

outer continental-shelf environment was chosen for study 

because it is the transition zone between nearshore 

sediments, which become anoxic within a few centimeters 

of the sediment surface (see, for example, Goldhaber and 

others, 1977), and pelagic muds, which remain oxic to 

suboxic over depths of several meters (see, for example, 

Froelich and others, 1979). This transition zone may be 

an important sink for metals released from the continents, 

either naturally or through pollution. 

Cores of the uppermost 1.5-2 m of sediments were 

collected from the outer Texas continental shelf of the 

Gulf of Mexico. A sequential extraction scheme was used 

that removes successively more tightly bound forms of 

the metals. Statistical correlations were then performed 

to determine relationships between metals and anionic 

species and among metals in the various chemically 

extracted fractions. 

Acknowledgments 

I thank R. M. Owen for his guidance and advice 

and the U.S. Geological Survey in Corpus Christi, Texas, 

and the crew of the R. V. Decca Pro filer for assistance 

in collecting cores and grab samples used for this study. 

Special thanks go to C. W. Holmes for his guidance during the cruise and to Cyndi Rice for her able assistance 

in the field laboratory. 

This work was based on part of the author's Ph. D. 

dissertation and was completed while she was a National 

Research Council Postdoctoral Research Associate with 

the U.S. Geological Survey. 

METHODS 

Sampling area 

The sampling area has been described previously 

(Filipek and Owen, 1980; Holmes and Martin, 1978), so 

only a brief review is given here. Two-meter gravity cores 

and surface grab samples were collected at each of the 

three sites shown in figure 1. Core 19 (lat 28 00' N., long 

96~4' W.), from a water depth of 32m, consists of 
Figure 1. Map of south Texas continental shelf of Gulf of Mexico showing numbered sampling localities (solid dots). Contours 

are water depths in meters. 

Influence of Fe and Mn on Chemical Partitioning of Cu, Zn, and Cr in Shelf Sediments ~l3 
sandy silt with sand lenses. The sedimentation rate in the 

vicinity of this station is estimated to be 1.44-4 mm/yr 

(Berryhill, 1976). Core 32 (lat 26 °54 1 N., long 96 °52 1 

W.), from a depth of 61 m, is a silty mud. A core from 

this location showed no excess 210Pb, suggesting no 

sedimentation during the last 150 years, but estimates 

based on organic decomposition rates suggest an average 

sedimentation rate of 0.3-1 mm/yr throughout the present core (Filipek and Owen, 1980). Core 9 (lat 26°53 1 

N., long 960J3 1 W.) is a muddy sediment from a water 

depth of 112m. The sedimentation rate at this location 

is 0.54 mm/yr (Berryhill, 1976; Holmes and Martin, 

1978). 

Sampling procedure 

The cores were sectioned on board ship. Interstitial 

water was expressed using N2, and the water samples 

were immediately acidified and frozen. The sectioned 

sediment was placed in plastic bags and refrigerated. 

In all three cores, a thin, oxidized surface layer and 

a subsurface oxidation maximum at about 10 em were 

recognized on the basis of interstitial sulfate concentrations, porosity, and other sediment parameters. This maximum is believed to be due to some activity related to benthic organisms, such as irrigation or the existence of a 

preferred feeding depth (Filipek and Owen, 1980). Living worms were found concentrated at a depth of 35 em 

in core 19, which is the core having the highest pyrite concentrations and the highest rate of sulfate reduction 

(Filipek and Owen, 1980). 

Analytical methods 

The sequential extraction scheme used here has been 

published previously (Filipek and Owen, 1978, 1981) and 

is outlined in figure 2. The extracts, together with the 

labels used to simplify discussion, are as follows: (1) 

dissolved (Diss); (2) acetic acid (HOAc)-dissolves 

exchangeable metals, acid-soluble sulfides, carbonates; 

(3) cold hydroxylamine-HCl (CHxl)-dissolves easily 

reducible species (mainly hydrous manganese oxides; 

from Chao, 1972); (4) hydrogen peroxide (H2

0 2)-

dissolves metals associated with organic matter and 

pyrite; (5) hot hydroxylamine-HCl (HHxl)-dissolves 

moderately reducible species (mainly hydrous iron oxides; 

see Chester and Hughes, 1967); and (6) hydrofluoric, perchloric, and nitric acids (Lith)-dissolves metals in 

residual minerals or incorporated in the structure of 

detrital silicate minerals (lithogeneous metals). 

Each of the extracts was analyzed for Fe, Mn, Zn, 

Cu, and Cr, and the HOAc-extract for Ca, by flame 

atomic absorption spectrophotometry (AA), using 

appropriate blanks and standards. Associated ligands 

34 Studies in Diagenesis 

(phosphate-P, sulfate-S, total sediment-S (or solid-S), 

organic-C, and carbonate-C) were measured using standard techniques discussed by Filipek and Owen (1981). 

Solid-S was labeled "pyrite-S" because analyses of 

distilled water-rinsed sediment and HOAc-treated sediment resulted in values that generally agreed to within 

10 percent, except for the upper 15 em of core 9, in which 

about 40 percent of the sulfur was acid-soluble. Based 

on the solid-S determinations, the H2

0 2-soluble Fe component was further subdivided into pyrite-Fe and OrgFe (organic Fe); that is, pyrite-Fe was estimated as onehalf of the molar value of solid-S, and Org-Fe was 

estimated by subtracting the amount of Fe in pyrite from 

H2

0 2-soluble Fe. Errors in Org-Fe due to inclusion of 

monosulfides are probably quite small because acidsoluble Fe-sulfides were only present to an appreciable 

extent in near-surface sediments where total sulfur was 

much less than H20 2-soluble Fe. 

RESULTS AND DISCUSSION 

Iron 

Between 50 and 60 percent of the total iron at all 

locations was found to be in the lithogenous fraction; that 

is, in crystal structures of, or in mineral inclusions in, 

detrital mineral grains (appendix 2). As shown in figure 3, the most dominant nonlithogenous form was 

HHxl-soluble Fe (hydrous iron oxides). In all three cores, 

the dominant diagenetic processes appear to have been 

the formation of hydrous iron oxides at the surface and 

subsurface oxidation maxima (at about 10 em) and the 

reduction of the oxides below these zones (appendix 2). 

Concentrations of dissolved iron increased just below 

zones of maximum HHxl-soluble Fe concentration, 

indicating a release of Fe2+ during reduction of the 

oxides. In core 9 (fig. 3, table 1), reduction of HHxlsoluble Fe apparently resulted in the formation of the 

HOAc-soluble Fe, based on the inverse relationship between the two profiles. Concentrations of HOAc-soluble 

S were too low for acid-soluble iron sulfides to be a 

significant fraction of the HOAc-soluble Fe. In addition, 

solubility estimates suggest that the interstitial waters were 

generally undersaturated with respect to siderite and 

vivianite, the most common nonsulfide ferrous iron 

phases. On the other hand, the association of BOAcsoluble Fe with HOAc-soluble Mg, K, and P-in a 

Mg:K:Fe:P atomic ratio of about 16:6:6:1 in all three 

cores-suggests that these elements were present in highly 

reactive phases adsorbed on clay minerals similar to the 

6: 1 iron hydroxide-phosphorus complex reported in certain reducing lacustrine sediments (Williams and others, 

1971) and discussed in detail elsewhere (Filipek and Owen, 

1981). The reduced or mixed (F&+, FeH) component is 

presumed to have formed because sulfate reduction was 
:I 

c -ttl 

:I 

~ 

a 

;:r 

11.1 

:I 

Cl. 

~ :I 

0 

:I 

t""' 

:r 

ttl 

3 

;::;· 

e!.. 

;: 

::l 

... 

s· 

:I 

;· 

QQ 

a 

t""' 

~ 

N 

E' 

11.1 

:I 

Cl. 

t""' 

... 

;· 

Ill 

:r 

ttl 

::;: 

[ 

~r tl) 

:I 

;-

w 

VI 

Table 1. Correlation coefficients for nonlithogenous metals and other selected sediment variables for cores 9, 32, and 19 

[P • 0.01, R ~ .64; P • 0.05, R ~.51] 

Core 9 

Depth 

Org-C 

Carb-C 

Solid-S 

Diss-Fe 

Diss-Mn 

HOAc-soluble Fe 

CHxl-soluble Fe 

H2o2-soluble Fe 

Org-Fe 

HHxl-soluble Fe 

HOAc-soluble Mn 

CHxl-soluble Mn 

H2o2-.(loluble Mn 

HHxl-soluble Mn 

HOAc-soluble Zn 

CHxl-soluble Zn 

H2o 2-soluble Zn 

HHxl-soluble Zn 

H2o 2-soluble Cu 

HHxl-soluble Cu 

HOAc-soluble Cr 

H2o 2-soluble Cr 

HHxl-soluble Cr 

1.00 

-.26 1.00 

-.58 -.28 1.00 

.92 -.28 -.60 

-.30 -.41 .48 

-.16 .12 -.38 

.48 .06 -.64 

-.45 .04 .so 

-.39 -.01 .36 

-.79 .17 .57 

-.04 -.40 .63 

.62 -.06 -.46 

.53 -.23 -.15 

-.19 -.18 .25 

-.09 -.24 .61 

.70 -.01 -.67 

.35 -.25 .28 

.17 -.37 .12 

.12 -.43 .56 

.55 -.30 -.03 

.47 -.41 .22 

-.18 .09 .24 

-.05 -.06 .09 

-.06 -.48 .60 

1.00 

-.33 1.00 

-.07 -.14 1.00 

.55 -.08 .30 1.00 

-.60 .18 -.03 -.81 

-.48 .41 -.15 -.25 

-.88 .42 -.04 -.48 

-.18 .21 -.32 -.66 

.70 -.41 .20 .49 

.53 -.11 .13 .11 

-.30 .68 .12 -.06 

-.24 .10 -.34 -.74 

.62 -.14 .13 .80 

.17 .23 -.26 -.09 

.02 .35 .32 .16 

-.04 .20 -.34 -.57 

.33 .12 -.04 .23 

.37 .26 -.27 .02 

-.26 .25 -.08 .39 

-.22 .54 -.28 .12 

-.18 .22 -.56 -.66 

Depth Org- Carb- Solid- Diss- Diss- HOAcC C S Fe Mn soluble 

Fe 

1.00 

.34 

.56 

.57 

-.40 

.14 

.42 

.60 

-.59 

.21 

.26 

.51 

-.08 

.02 

-.19 

.08 

.43 

CHxlsoluble 

Fe 

1.00 

.83 

-.01 

-.45 

-.20 

.67 

.17 

-.37 

-.08 

.28 

.20 

-.24 

.02 

.39 

.57 

.20 

1.00 

.11 

-.68 

-.44 

.54 

.24 

-.59 

-.15 

.13 

.13 

-.33 

-.22 

.37 

.44 

.22 

1.00 

-.19 

.39 

.18 

.94 

-.37 

.43 

.30 

.93 

.44 

.so 

-.10 

-.22 

.85 

HzOz- Org- HHxlsoluble Fe soluble 

Fe Fe 

1.00 

.59 

-.22 

-.27 

.46 

-.05 

.17 

-.12 

.27 

.36 

-.23 

-.33 

-.40 

1.00 

.29 

.29 

.28 

.25 

.55 

.so 

.43 

.56 

-.14 

-.31 

.13 

HOAc- CHxlsoluble soluble 

Mn Mn 

1.00 

.15 

-.01 

.14 

.69 

.33 

.11 

.29 

.32 

.52 

.09 

s!l~s~e Mn 

1.00 

-.so 

.36 

.22 

.93 

.37 

.47 

-.09 

-.19 

.84 

HHxlsoluble 

Mn 

1.00 

.09 

.22 

-.29 

.47 

.23 

.22 

.24 

.42 

HOAcsoluble 

Zn 

1.00 

.45 

.49 

.59 

.45 

.13 

.07 

22 

CHxlsoluble 

Zn 

1.00 

.44 

.61 

.49 

.40 

.17 

.05 

s!l~s~e Zn 

1.00 

.49 

.64 

-.03 

-.11 

.81 

HHxlsoluble 

Zn 

1.00 

.72 

.23 

.07 

.21 

s!l~~~e Cu 

1.00 

.03 

.20 

.29 

HHxlsoluble 

Cu 

1.00 

.36 

-.02 

HOAc 

soluble 

Cr 

1.00 

-.03 

s!l~G~e Cr 

1.00 

HHxlsoluble 

Cr 
w 

a'\ 

~ c 

a. 

~-

:r 

0 

~-

:I 

rJ 

;;· 

Table 1. Correlation coefficients for nonlithogenous metals and other selected sediment variables for cores 9, 32, and 19 (c)--continued 

[P • 0.01, R ~.64; P s 0.05, R ~.51] 

Core 32 

Depth 

Org-C 

Car b-e 

Solid-S 

Diss-Fe 

Diss-Mn 

HOAc-soluble Fe 

CHxl-soluble Fe 

H2o2-soluble Fe 

Org-Fe 

HHxl-soluble Fe 

HOAc-soluble Mn 

CHxl-soluble Mn 

H2o2-soluble Mn 

HHxl-sol'u'ble Mn 

HOAc-soluble Zn 

CHxl-soluble Zn 

H2o2-soluble Zn 

HHxl-soluble Zn 

H2o2-soluble Cu 

HHxl-soluble Cu 

H2o2-soluble Cr 

HHxl-soluble CT 

1.00 

.18 1.00 

-.46 -.10 1.00 

.91 .28 -.61 1.00 

-.32 -.44 .41 -.42 1.00 

-.32 -.62 

.21 .37 

-.69 .11 

.62 .04 

-.57 -.29 

.02 .14 

.75 .40 

.83 .47 

.76 .47 

-.29 .10 

-.15 -.03 

-.34 .06 

.83 .39 

-.32 -.50 

.74 .43 

.35 .52 

.11 .33 

-.21 .17 

.58 -.42 

-. 75 .47 

.34 -.67 

-.49 .54 

.32 -.73 

.35 -.10 

-.49 • 77 

-.58 .87 

-.31 .62 

.26 -.28 

-.43 .04 

-.17 -.13 

-.18 .66 

.19 -.28 

-.44 .73 

.02 .39 

-.66 .26 

.34 -.29 

.85 1.00 

-.43 -.49 1.00 

.16 .21 .06 

-.07 -.26 .22 

.43 .28 -.37 

-.04 .04 -.41 

-.29 -.49 .27 

-.43 -.55 .40 

-.29 -.51 .02 

.10 .08 -.14 

.48 .18 .44 

.27 .18 .36 

-.08 -.17 .06 

-.08 .18 -.20 

-.40 -.47 .38 

-.31 -.29 .16 

-.44 -.64 .60 

.11 .23 -.08 

1.00 

-.51 

.37 

-.17 

-.48 

-.44 

-.47 

-.05 

.13 

.27 

-.36 

-.07 

-.40 

-.32 

-.08 

.27 

Depth Org- Carb- Solid- Diss- Diss- HOAc- CHxlC S Fe Mn Fe soluble soluble 

Fe Fe 

1.00 

.18 1.00 

-.20 -.05 

.31 -.66 

.44 -.65 

.49 -.34 

-.08 .27 

.14 .07 

-.29 -.8 

.53 -.34 

-.31 .08 

.34 -.58 

-.15 -.57 

.38 .oo 

-.32 .08 

1.00 

-.09 

-.11 

.02 

.75 

-.40 

.07 

.06 

.43 

-.13 

.42 

-.30 

.47 

s=~~~~e ~!- s=~~~e Fe Fe 

1.00 

.89 

.so 

-.36 

.01 

-.26 

.67 

-.50 

.70 

.37 

.19 

-.39 

HOAcsoluble 

Mn 

1.00 

.84 

-.36 

-.02 

-.22 

.73 

-.47 

.so 

.46 

.32 

-.21 

CHxlsoluble 

Mn 

1.00 

-.32 

-.20 

-.44 

.so 

-.55 

.77 

.42 

.28 

-.21 

s=~~~~e Mn 

1.00 

-.09 

.25 

-.28 

.51 

-.43 

.28 

-.19 

.52 

HHxlsoluble 

Mn 

1.00 

.73 

-.09 

-.27 

-.01 

-.26 

.24 

-.13 

HOAcsoluble 

Zn 

1.00 

-.28 

.14 

-.10 

.02 

.02 

.31 

CHxlsoluble 

Zn 

1.00 

-.65 

.73 

.26 

.04 

-.08 

s=~~~~e Zn 

1.00 

-.47 

.08 

-.27 

.24 

HHxlsoluble 

Zn 

1.00 

.48 

.46 

-.27 

s=~~~~e Cu 

1.00 

-.oo 

.43 

HHxlsoluble 

Cu 

1.00 

-.45 

s=~~~~e Cr 

1.00 

HHxlsoluble 

Cr 
:; 

c 

-I'D 

:I 

~ 

s. 

;r 

I» 

:I 

Q. 

~ :I 

Q 

:I 

n 

::r 

I'D 

3 

[ 

-: 

~ cr 

:I 

:;-

OQ 

s. 

n 

~ 

N 

~= 

I» 

:I 

Q. 

0 

:;-

VI 

::r 

! 

[ 

a· 

I'D 

a 

Ill 

{N 

...... 

Table 1. Correlation coefficients for nonlithogenous metals and other selected sediment variables for cores 9, 32, and 19--<:ontinued 

[P • 0.01, R L .66; and P • 0.05, R L .53] 

Core 19 

Depth 

Org-C 

Carb-<: 

Solid-S 

Diss-Fe 

Diss-Mn 

HOAc-soluble Fe 

CHxl-soluble Fe 

H2o2-soluble Fe 

Org-Fe 

HHxl-soluble Fe 

HOAc-soluble Mn 

CHxl-soluble Mn 

H2o 2-solulile Mn 

HHxl-soluble Mn 

HOAc-soluble Zn 

CHxl-soluble Zn 

H2o 2-soluble Zn 

HHxl-soluble Zn 

H2o2-soluble Cu 

HHxl-soluble Cu 

HOAc-soluble Cr 

H2o2-soluble Cr 

HHxl-soluble Cr 

1.00 

-.15 1.00 

-.39 -.22 1.00 

.61 -.04 -.47 1.00 

.31 .25 -.35 .32 

.33 -.17 .44 .35 

-.04 -.24 -.14 .65 

-.19 .81 -.07 -.22 

.48 -.14 -.18 .14 

1.00 

.07 1.00 

-.14 .13 

-.09 -.07 

.14 -.25 

1.00 

-.41 

-.16 

-.11 -.os .21 -.67 -.12 -.49 -.61 

-.19 .74 -.30 .oo .24 -.43 .01 

-.18 .66 -.13 -.28 .07 -.24 -.45 

-.03 .70 

.04 .62 

-.34 .54 

-.55 .14 

.18 .60 

.45 .09 

-.16 .66 

.45 -.03 

.os .69 

-.21 -.09 

.67 .os 

-.83 .31 

-.26 -.10 

-.17 -.36 

-.21 -.20 

-.10 .19 

-.45 .56 

-.26 .09 

-.37 .06 

-.40 .81 

-.41 .25 

.26 .09 

-.16 .23 

.24 -.35 

.11 -.17 

.12 -.32 

.38 -.58 

-.23 -.31 

-.06 .14 

.32 -.10 

.25 -.48 

.40 .27 

.26 -.16 

-.19 -.01 

.18 .11 

-.03 -.41 

-.38 

-.67 

-.14 

• 73 

.33 

-.20 

-.03 

.61 

-.01 

.59 

-.19 

.19 

1.00 

-.26 

-.01 

.56 

.86 

.88 

.73 

.32 

-.07 

.53 

-.10 

.53 

-.33 

.71 

-.37 

-.06 

.06 

Depth Org- Carb- Solid- Diss- Diss- HOAc- CHxlC C S Fe Mn soluble soluble 

Fe Fe 

1.00 

.63 1.00 

-.06 -.03 1.00 

-.19 .07 .74 

-.16 -.04 

.28 .47 

.04 .24 

-.20 -.24 

-.17 -.57 

.72 .43 

.18 .12 

.29 -.40 

-.03 -.21 

-.14 -.17 

.76 .31 

-.17 .20 

.73 

.61 

.so 

.32 

.36 

.04 

.91 

-.06 

.85 

.22 

.oo 

.42 

s=~~~~e ~; s!~!~e Fe Fe 

1.00 

.97 

.84 

.58 

-.12 

.27 

-.09 

.71 

-.44 

.81 

-.29 

-.os 

-.10 

HOAcsoluble 

Mn 

1.00 

.82 

.52 

-.13 

.42 

-.OS 

.72 

-.28 

.88 

-.32 

.01 

-.03 

CHxlsoluble 

Mn 

1.00 

.51 

-.34 

.13 

.38 

.65 

-.33 

.65 

-.39 

.39 

-.oo 

s=~~~~e Mn 

1.00 

.29 

-.oo 

-.04 

.81 

-.20 

.56 

.14 

-.15 

.59 

HHxlsoluble 

Mn 

1.00 

.36 

-.31 

.34 

.22 

.13 

.47 

-.37 

• 71 

HOAcsoluble 

Zn 

1.00 

-.07 

.36 

.37 

.ss 

-.12 

.11 

-.05 

CHxlsoluble 

Zn 

1.00 

.09 

.48 

.OS 

-.17 

.84 

-.19 

s=~~~~e Zn 

1.00 

-.01 

.87 

.04 

.07 

.42 

HHxlso~ble 

Zn 

1.00 

.10 

.14 

.37 

-.14 

s=~~~~e Cu 

1.00 

-.10 

.08 

.07 

HHxlsoluble 

Cu 

1.00 

-.21 

.41 

HOAcsoluble 

Cr 

1.00 

-.38 

s=~~~~e Cr 

1.00 

HHxlsoluble 

Cr 
STARTING SEDIMENT 

(Refrigerated) Digest with HF, HN03, HCI04 > Total elements 

I 

Leach with 

~' Leach with 

0.1 M NH20HHCI 

~~o, 

Leach with 

30 percent H202 

~ Leach with 0.25M 

NH20H·HCI + 25 

percent HOAc at sooc 

$ 

Digest with HF, 

HN03,HCI04 

HOAc: 

Exchangeable metals, 

carbonates, iron 

monosulfides 

CHxl: 

Hydrous manganese 

oxides 

H2Q2: 

Organic materials 

Pyrite 

HHxl: 

Hydrous iron 

oxides 

Lith: 

Lithogenous metals 

Figure 2. Flow chart of extraction treatments. See Filipek and Owen (1978, 1981) for details. HOAc, extractable in acetic acid; 

CHxl, extractable in 0.1 M NH2

0H · HCI; H20 2

, extractable in 30 percent H2

0 2; HHxl, extractable in 0.25 M NH20H · HCI + 25 

percent acetic acid; Lith, extractable after digestion in HF, HN03

, HCI04 • 

not intense enough to supply the sulfide necessary to react 

with all the reduced iron in these sediments. The BOAcsoluble Fe and HHxl-soluble Fe do not show the inverse 

relationship in core 19 that they show in core 9 (table 1), 

probably because of the more intense sulfate reduction 

and concomitant pyrite formation in core 19. 

In all three cores, pyrite-Fe increased with depth 

because of organic decomposition (Filipek and Owen, 

1980); however, pyrite-Fe accounted for less than 30 percent of the nonlithogenous iron in all cores, even in core 

19, which had the highest concentrations of solid-S. Concentrations of Org-Fe were erratic but generally decreased 

with depth in the sediment column. 

Manganese 

In terms of relative amounts in each extract, 

38 Studies in Diagenesis 

manganese behaved quite differently from iron. The 

absolute amount of manganese was much less than iron, 

and most (75-90 percent) of the manganese was in 

nonlithogenous components. In all three cores, BOAcsoluble Mn was by far the dominant fraction of the element and generally accounted for more than 60 percent 

of the total metal (fig. 4, appendix 3). All the nonlithogenous manganese components showed slight maxima at 

the surface and subsurface oxidation maxima, at about 

the same depths as the minima in dissolved manganese 

concentrations (fig. 4, appendixes 1 and 3). At the 

shallowest site (site 19), nonlithogenous manganese was 

more than twice as concentrated in the uppermost centimeter of sediment than in the remainder of the core, 

indicating significant diagenetic remobilization of 

manganese here (appendix 3). 

The pore-water profiles of all three cores (appendix I) and the nonlithogenous manganese profiles of core 
en 

a: 

w 

1-

w 

:E 

~ 

z 

50 

~ 10() 

~ 

:i 

b: 

w 

c 

SOLID IRON, IN 

0.4 

I 

\ 

\ 

DISSOLVED IRON, 

PERCENT IN MICROGRAMS PER MILLILITER 

0 2 

o I 

50 

100 

X 

\ \ X 

1 \ 150 150 

>c: I \ X 

I 

200 HOAc ORG HHxl PYRITE 200 

Figure 3. Distribution of nonlithogenous iron and d;1ssolved iron from core 9. CHxl, CHxl-extractable iron; ORG, iron in 

organic: materials; HHxl, HHxl-extractable iron; Pyritt~, iron in pyrite. (See fig. 2 for extraction treatments.) 

19 (appendix 3) suggest that manganese oxides should 

have been significant in the more oxidizing portions of 

the cores. The CHxl (manganese oxide) extract, however, 

contaim~d only a small percentage of manganese, whereas 

the HOAc extract was the richest in this element, even 

at the surface. Based on solubility considerations, 

manganese soluble in cold weak acid (such as HO.Ac) is 

more likely to be manganese carbonate (rhodochrosite) 

or a mixed (Ca,Mn)-carbonate than manganese oxid,e. At 

equilibrium, the carbonate phases are usually the most 

stable forms of manganese in reducing sediments (Stumm 

and Morgan, 1970; Holdren and others, 1975). (Ca,Mn)-

carbonates of varying Mn:Ca ratio have been reported 

in anoxllc marine sediment by Suess ( 1979) and Pedersen 

and Price (1982). Therefore, the presence of a carbonate 

phase could explain the dominance of BOAc-soluble Mn 

in redudng portions of the present cores, but not in the 

oxidizing zones. 

Two explanations for high concentrations of 

BOAc-soluble Mn in oxidizing sediments are possible: 

First, the formation of rhodochrosite or a mixed 

(Ca,Mn)-carbonate from Mn2+ may have been kinetically favored in the presence of carbonates, even in oxidizing sediments. Binstock (1978) reported X-ray diffraction evidence for such materials in oxygenated, 

carbonate-rich clay suspensions. Indirect evidence in the 

form of strong correlations between manganese and carbonate was reported by Luoma and Bryan (1981) for 

oxidizing estuarine sediments low in either organic materials or total iron. These authors suggested that manganese carbonate was common wherever other substrates 

that bind manganese were in low concentration. Filipek 

and Owen (1979) found that BOAc-soluble Mn was the 

dominant manganese form in highly oxidized, carbonaterich, surficial lacustrine sediments not containing visible 

manganese nodules; in contrast, in sediments from 

Influence of Fe and Mn on Chemical Partitioning of Cu, Zn, and Cr in Shelf Sediments 39 
SOLID MANGANESE, IN MICROGRAMS PER GRAM 

DISSOLVED MANGANESE, 

MICROGRAMS PER MILLILITER 

CJ) 

a: 

w 

~ 

w 

:::?: 

i= 

z 

~ 100 

~ 

:i 

li: 

w 

0 

150 

200 

100 200 300 2 

100 

150 

200 

Figure 4. Distribution of nonlithogenous manganese and dissolved manganese in core 9. CHxl, CHxl-extractable manganese; 

HHxl, HHxl-extractable manganese; H20 2, H20 2-extractable manganese; HOAc, HOAc-extractable manganese. (See fig. 2 for 

extraction treatments.) 

neighboring sites containing visible (about 1 mm) nodules, 

the HOAc extraction dissolved very little, and the CHxl 

extraction removed most of the manganese. 

A second possible explanation has been suggested 

by G. R. Holdren (University of Rochester, Rochester, 

New York, oral commun., 1980). Holdren found that 

finely ground manganese nodules were normally insoluble in dilute acids; however, the nodules became soluble 

if an aliquot of oxic marine sediment was added to the 

dilute acid suspension with the nodules. Holdren inferred from these results that the lowered pH of the extraction aided in the reduction of manganese oxides by 

organic materials present in the sediment. 

In the present sediments, two distinct manganese 

phases may have been dissolved by HOAc. In the surficial oxidizing sediments, the HOAc-soluble Mn could 

have originated mainly from noncrystalline hydrous 

manganese oxides, whereas in the lower reducing portions, the HOAc-soluble Mn was probably from a 

(Ca,Mn)-carbonate. The inclusion of manganese oxides 

40 Studies in Diagenesis 

in the HOAc extract from oxidizing layers, but not from 

reducing layers, explains the variable relationship with 

depth between HOAc-soluble Ca and HOAc-soluble Mn 

in the cores (fig. 5). Based on high correlations with 

carbonate-e and an approximately 1:1 organic-C:HOAcsoluble Ca molar ratio, HOAc-soluble Ca represents 

mainly calcium carbonate in the sediments. Below about 

35 em in all three cores, both HOAc-soluble Ca and 

HOAc-soluble Mn had similar patterns, suggesting 

response to pH changes, or, more likely, the formation 

of a (Ca,Mn)-carbonate with a Ca:Mn molar ratio of 

about 80:1 (fig. 5, appendix 1). 

Copper 

Copper was the only element studied that varied in 

chemical partitioning from one core to the next. The 

percentage of the total copper content in the 

nonlithogenous components (appendix 4) varied from a 
CALCIUM, IN 100 MICROMOLES PER GRAM 

MANGANESE, IN MICROMOLES PER GRAM 

2 4 6 8 0 2 4 6 8 0 2 4 6 8 10 ~~~--~~~--.--.-.--.-~ 

9 

50 -

(J) 

a: 

w 

1-

w 

::!: 

~ z 

w 100 - u 

~ 

:i -_Sa~ 1- - Q.. 

w 

c 

150 

200 

Figure 5. Distribution of HOAc-extractable manganese and calcium in cores 9, 32, and 19 (in micromoles and 100 micromoles 

per gram of dry sediment, respectively). (See fig. 2 for extraction treatment.) In core 19, sand lens is indicated by horizontal 

dashed lines. 

low of about 30 percent of the total copper throughout 

the core taken from the shallowest, sandy-silt site (core 

19, water depth = 32 m) to about 50 percent at the 

intermediate site (core 32, water depth = 61 m) to about 

65 perc~~nt at the deepest, muddy site (core 9, water 

depth == 112 m). In all three cores, most of the 

nonlithogenous copper was in the H2

0 2

- and HHxlsoluble forms. In the subsurface, reduced portions of the 

shallow<!st site (site 19), H20 2-soluble Cu was clearly the 

dominant nonlithogenous form; at site -32, the amount 

of H20 2-soluble Cu was equal to or slightly greater than 

the amount of HHxl-soluble Cu; whereas at the de1!pest 

site (site 9), the amount of HHxl-soluble Cu was slightly 

greater than that of H20 2-soluble Cu (appendix 4, 

fig. 6). In cores 19 and 32, H2

0 2-soluble Cu correlated 

significantly with solid-S (r = .81 and . 73, respectively; 

table 1). These correlations strongly suggest the incorporation of copper in pyrite or the formation of copper 

sulfides (covellite or chalcocite?), which are geochem[cally 

plausibl~~ based on the similar redox stability fields of copper sulfides and pyrite (Sato and Mooney, :l960; 

Bolviken, 1977). 

In cores 9 and 19, HOAc-soluble Cu and CHxlsoluble Cu were below the flame AA detection limit 

(about 0.01 p.mol Cu/g dry sediment). In the top 10 em 

of core 32, HOAc-soluble Cu was a significant portion 

of the nonlithogenous copper but decreased below this 

depth (appendix 5) and showed no relation to other 

sediment components. This trend suggests that the 

HOAc-soluble Cu in core 32 was in an easily exchangeable 

form not related to the HOAC-soluble Fe phase discussed 

previously. 

Statistical analyses (table 1) indicate generally high 

positive correlations between HHxl-soluble Cu and HHxlsoluble Fe, HHxl-soluble Mn, and CHxl-soluble Mn; 

however, the overall correlation between copper and iron 

was not as strong as between zinc and iron, suggesting 

that iron is less influential in copper diagenesis. 

Moreover, the relatively large number of high correlations between copper and manganese in various extracts 

suggests a rather strong relation between sediment 

manganese and copper, even though manganese oxide 

was not an important phase in these sediments. 

Zinc 

In all three cores, the dominant nonlithogenous zinc 

components were H2

0 2-soluble Zn and HHxl-soluble Zn 

(fig. 6, appendix 5). Statistical analyses of the Zn components and other sediment variables (table 1) indicate 

that HHxl-soluble Zn is highly correlated with both 

HHxl-soluble Fe and HHxl-soluble Mn in all three cores, 

although less significantly so in core 32, where HHxlsoluble Zn and H20 2-soluble Zn show a strong inverse 

relation. Also, HOAc-soluble Zn has a high positive correlation with HOAc-soluble Fe, especially in low-sulfide 

cores 9 and 32, suggesting that HOAc-soluble Zn was 

Influence of Fe and Mn on Chemical Partitioning of Cu, Zn, and Cr in Shelf Sediments 41 
CONCENTRATIONS, IN MICROGRAMS PER GRAM 

0 10 20 30 0 2 3 4 

0 

5 0 2 3 4 5 6 7 8 

'x ~ ~ 

~X ) J l 

\ \ 

50 ~ l X/ 

\ ~ j en 

a: 

w 

\ I \ 1-

w 

:i!: 

i= 

z 

w 100 

j u 

~ 

z 

1-

j \ 0.. 

w 

0 

150 

Zn Cu Cr 

200 \ \ I 

CHxl HOAc HHxl H202 HOAc HHxl H202 

Figure 6. Distribution of nonlithogenous Zn, Cu, and Cr in core 9 (in micrograms per gram of dry sediment). CHxl, CHxlextractable; HHxl, HHxl-extractable; H2

0 2-extractable; HOAc, HOAc-extractable; H20 2, H20 2-extractable. (See fig. 2 for extraction treatments.) 

associated with the reduced iron component discussed 

previously rather than existing as a zinc carbonate. Here 

also, the relation between zinc and iron indicates the large 

control that iron had on zinc diagenesis, due to coprecipitation and( or) adsorption processes. The strong relationship between zinc and iron has been reported 

previously for pelagic sediments by Forstner and Stoffers (1981). 

Iron was not the only controlling factor in zinc partitioning, however, because a large part (about 30-50 percent) of the nonlithogenous zinc was in the H2

0 2-soluble 

extract of all three cores. The zinc in this extract could 

have come from either an organic complex or a zinc 

sulfide. Significant correlations between H2

0 2-soluble 

Zn and solid-S in core 32 (table 1) suggest the presence 

of zinc sulfide (or Zn substitution in pyrite) in this core. 

Sphalerite was probably not in equilibrium in these 

sediments because its redox stability field is much lower 

than that expected for these sediments (Bolviken, 1977). 

On the other hand, wurtzite has been reported in marine 

sediments (R. Koski, U.S. Geological Survey, Menlo 

Park, California, oral commun., 1982), and extrapolation of the work of Scott and Barnes (1972) to low 

temperatures suggests a stability field for wurtzite at 

higher redox potentials than for sphalerite. For cores 9 

and 19, H20 2-soluble Zn shows no correlation with 

42 Studies in Diagenesis 

solid-S (table 1), nor is it significantly correlated with 

organic-C. Presley and others (1972) also found no correlation between zinc and organic-C in sediments where 

they expected that zinc should be the metal most closely 

associated with organic matter, due to the high concentration of zinc in plankton. 

In the present study, no element in any one extract 

was found to correlate consistently with organic-C content. This lack of correlation could be due to the fact that 

none of the extractions used effectively removed all metals 

complexed with organic matter, because the type and 

strength of the metal-organic bond varied both with the 

metal and the organic compound. Another possible 

explanation is that the method used for measuring 

organic-C, in which sediments were left overnight in 1 

M HCl to dissolve carbonates, gave inaccurate results for 

organic-C (Froelich, 1980). Alternatively, the lack of correlation between organic-C and metals in the H20 2 

extract (or any extract) could have been due to differential breakdown of the various types of bonds during 

diagenesis. In this connection, Price (1976) reported that 

the metal/organic-C ratios for Zn, Cu, and Pb decreased 

sharply with depth in rapidly accumulating marine 

sediments of uniform lithology from Loch Fyne, 

Scotland. Because the Loch Fyne sediments were oxic, 

yet contained little manganese or iron hydroxides, Price 
suggested that this decrease was due to breakdown and 

release of metal-organic substances from the sediment 

during buiial, with rates of release decreasing in the order 

Zn>Cu>Pb. If such differential breakdown is indeed 

common during diagenesis, organic-C and organically 

bound metals should not be linearly correlated throughout the sediment column. 

Chromium 

Nonlithogenous chromium accounted for only 10-

20 percent of the total metal content, making chromium 

the least chemically available of the metals studied. In 

all three t:=ores, H20 2-soluble Cr was the dominant nonlithogenous chromium component, followed by HHxlsoluble Cr (fig. 6, appendix 6). The amount of chromium 

extracted by HOAc was small in cores 19 and 9 and below 

the flame AA detection limit (about 0.01 ~'mol Cr/g dry 

sediment) in core 32, as was the CHxl-soluble Cr in all 

three cores. The H20 2-soluble Cr showed generally high 

correlations with organic-Fe, suggesting that much of the 

chromium was associated with organic material. In certain parts of all three cores, there was an indication of 

opposing behavior of H20 2-soluble Cr and HHxl-soluble 

Cr (fig. 6, appendix 6), although there was no significant 

negative correlation throughout the cores. 

Chromium is sensitive to redox conditions and can 

exist in natural environments in both the CrH and Cr6+ 

states. The higher oxidation state exists mainly as the 

chromate: (CrO/-) or dichromate (Cr20/-) species in 

oxidizing aqueous environments (Canter and Gloyna, 

1968); however, upon contact with sediments, these 

species may be reduced to the much more reactive Cr3 + 

state, which complexes readily with organics and sorbs 

in hydrous form to other sediment particles (Cotton and 

Wilkinson, 1972). The transfer of chromium between the 

H20 2-soluble and HHxl-soluble components with 

changes in redox environment in the present cores could 

have been due to competition between organic materials 

and hydrous iron oxides for the reactive CrH. 

CONCLUSIONS 

The outer continental-shelf sediments of the Gulf 

of Mexico represent an environment transitional between 

(1) organllc-poor pelagic sediments, in which organic matter decomposition involved reduction of oxygen, hydrous 

manganese oxides, and hydrous iron oxides over a large 

depth range in the sediment column, and (2) organic-rich 

nearshon~ sediments, in which sulfate is the major oxidant and pyrite production is essentially complete within 

the uppe:rmost few centimeters. In these outer-shelf 

sediments, hydrous iron oxides were the major oxidant 

of organic matter throughout the uppermost two meters, 

and sulfate reduction was relatively minor. Reduction of 

manganese oxides appears to have been confined to the 

sediment surface; thus most manganese in the sediment 

was HOAc-soluble and probably represents a (Ca,Mn)-

carbonate phase. Therefore, in the outershelf sediments, 

iron diagenesis was able to exert a significant control over 

the chemical partitioning of copper, zinc, and chromium 

due to coprecipitation, dissolution, and adsorption processes, whereas the role of manganese was minor, except 

possibly at the sediment surface. The extent of control 

by iron, however, varied with both the metal and, to a 

lesser extent, location on the shelf. 

The major control on the mobility of copper and 

zinc in the subsurface anoxic sediments appears to have 

been the formation of sulfide phases. These trace metals 

either were incorporated into pyrite or formed separate 

sulfide minerals, and they did so even though sulfate 

reduction was minor and reduced Fe was in excess over 

sulfide. Such an affinity of sulfide for copper and zinc 

suggests that anoxic sediments could provide an efficient 

sink for these metals in other relatively low-sulfur 

environments, such as estuaries or swamps. 

RE FE RE NCES CITED 

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Texas Outer Continental Shelf, 1975, Geology: National 

Technical Information Service Report PB251-341, 269 p. 

Binstock, J. H., 1978, Experimental investigation of silica and 

manganese sorption onto clay, carbonate, and amorphous 

silica: 1978 Annual Meeting, Geological Society of 

America, Abstracts with Programs 10, p. 367. 

Bolviken, Bjorn, 1977, The redox potential field of the earth, 

in Ahrens, L. H., ed., Origin and distribution of the 

elements, Proceedings of 2d Symposium, Paris, May 1977: 

New York, Pergamon Press, p. 649-665. 

Canter, L. W., and Gloyna, E. F., 1968, Transport of 

chromium-51 in an organically polluted environment: 

Engineering Bulletin, Purdue University, Engineering 

Extension Service, no. 132, pt. 1, p. 374-387. 

Chao, T. T., 1972, Selective dissolution of manganese oxides 

from soils and sediments with acidified hydroxylamine 

hydrochloride: Soil Science Society of America Proceedings, v. 36, p. 764-768. 

Chester, R., and Hughes, M. J., 1967, A chemical technique 

for the separation of ferromanganese minerals, carbonate 

minerals, and adsorbed trace elements for pelagic 

sediments: Chemical Geology, v. 2, p. 249-262. 

Cotton, F. A., and Wilksinson, G., 1972, Advanced inorganic 

chemistry (3d ed.): New York, Wiley-Interscience, 1145 p. 

Filipek, L. H., and Owen, R. M., 1978, Analysis of heavy-metal 

distributions among different mineralogical states in 

sediments: Canadian Journal of Spectroscopy, v. 23, p. 

31-34. 

___ 1979, Geochemical associations and grain-size partitioning of heavy metals in lacustrine sediments: Chemical 

Geology, v. 26, p. 105-117. 

Influence of Fe and Mn on Chemical Partitioning of Cu, Zn, and Cr in Shelf Sediments 43 
---- 1980, Early diagenesis of organic carbon and sulfur 

in outer-shelf sediments from the Gulf of Mexico: 

American Journal of Science, v. 280, p. 1097-1112. 

____ 1981, Diagenetic controls of phosphorus in outer 

continental-shelf sediments from the Gulf of Mexico: 

Chemical Geology, v. 33, p. 181-204. 

Forstner, Ulrich, and Stoffers, Peter, 1981, Chemical fractionation of transition elements in Pacific pelagic sediments: 

Geochimica et Cosmochimica Acta, v. 45, p. 1141-1146. 

Froelich, P. N., 1980, Analysis of organic carbon in marine 

sediments: Limnology and Oceanography. v. 25, p. 

564-572. 

Froelich, P. N., Klinkhammer, G. P., Bender, M. L., Luedtke, N. A., Heath, G. R., Cullen, D., Dauphin, P., Hammond, D., Hartman, B., and Maynard, V., 1979, Early 

oxidation of organic matter in pelagic sediments of the 

eastern equatorial Atlantic-Suboxic diagenesis: Geochimica et Cosmochimica Acta, v. 43, p. 1075-1090. 

Goldhaber, M. B., Aller, R. C., Cochran, J. K., Rosenfeld, 

J. K., Martens, C. S., and Berner, R. A., 1977, Sulfate 

reduction, diffusion, and bioturbation in Long Island 

Sound sediments-Report of the FOAM group: American 

Journal of Science, v. 277, p. 193-237. 

Hamilton-Taylor, John, 1979, Enrichments of zinc, lead, and 

copper in recent sediments of Windermere, England: 

Environmental Science and Technology, v. 13, p. 693-697. 

Holdren, G. R., Jr., Bricker, 0. P., 3d, and Matisoff, G., 1975, 

A model for the control of dissolved Mn in the interstitial 

waters of Chesapeake Bay, Washington, D. C., in Church, 

T., ed., Marine chemistry in the coastal environment: 

Washington, D.C., American Chemical Society, 

p. 364-381. 

Holmes, C. W., and Martin, E. A., 1978, Migration of 

anthropogenically induced trace metals (barium and lead) 

in a continental shelf environment, in Proceedings of the 

4th joint conference on sensing environmental pollutants, 

New Orleans, November 1977: Washington, D.C., 

American Chemical Society, p. 672-676. 

44 Studies in Diagenesis 

Luoma, S. N., and Bryan, G. W., 1981, A statistical assessment of the form of trace metals in oxidized estuarine 

sediments employing chemical extractants: Science of Total 

Environment, v. 17, p. 165-196. 

Nriagu, J. 0., Kemp, A. L. W., Wong, H. K. T., and Harper, 

N., 1979, Sedimentary record of heavy-metal pollution in 

Lake Erie: Geochimica et Cosmochimica Acta, v. 43, p. 

247-258. 

Pedersen, T. F., and Price, N. B., 1982, The geochemistry of 

manganese carbonate in Panama Basin sediments: 

Geochimica et Cosmochimica Acta, v. 46, p. 59-68. 

Presley, B. J., Kolodny, Y., Nissenbaum, A., and Kaplan, I. 

R., 1972, Early diagenesis in a reducing fjord, Saanich 

Inlet, British Columbia-H. Trace element distribution in 

interstitial water and sediment: Geochimica et 

Cosmochimica Acta, v. 36, p. 1073-1090. 

Price, N. B., 1976, Chemical diagenesis in sediments, in Riley, 

J.P., and Chester, R., eds., Chemical oceanography, v. 

6: New York, Academic Press, p. 158. 

Robbins, J. A., and Callender, E., 1975, Diagenesis of 

manganese in Lake Michigan sediment: American Journal of Science, v. 275, p. 512-522. 

Sato, Motoaki, and Mooney, H. M., 1960, The electrochemical 

mechanism of sulfide self-potential: Geophysics, v. 25, 

p. 226-249. 

Scott, S. D., and Barnes, H. L., 1972, Sphalerite-wurtzite 

equilibria and stoichiometry: Geochimica et Cosmochimica 

Acta, v. 36, p. 1275-1295. 

Stumm, W., and Morgan, J. J., 1970, Aquatic chemistry: New 

York, Wiley Interscience, 583 p. 

Suess, Erwin, 1979, Mineral phases formed in anoxic sediments 

by microbial decomposition of organic matter: Geochimica 

et Cosmochimica Acta, v. 43, p. 339-352. 

Williams, J.D. H., Syers, J. K., Shukla, S. S., Harris, R. F., 

and Armstrong, D. E., 1971, Levels of inorganic and total 

phosphorus in lake sediments as related to other sediment 

parameters: Environmental Science and Technology, 

p. 1113-1120. 
Appendix 1. Concentrations of organic-C, carbonate-e, HOAc-soluble Ca, and solid-S; and 

dissolved sulfate, Fe, and Mn 

[mM, millimolar; ~M, micromolar] 

So 1 ids, in micromoles per gram of dry sediment Dissolved 

Depth Org-C Carb-C HOAc- Solid-S so2- Fe Mn 

(em) soluble (m~) ( ll-1) ( ll-1) 

Ca 

Core 9/ 0-1 675 642 561 6.2 29.0 2.7 14.2 

1-2 525 742 554 10.3 24.7 13.4 6.4 

2-4 617 617 536 13.2 24.4 4.5 1.1 

4-7 508 992 511 14.1 27.2 2.5 2.9 

7-11 417 975 477 14.1 26.8 26.0 5.1 

11-16 617 567 501 18.8 29.9 8.1 20.7 

20-25 417 682 634 21.6 27.5 6.6 15.3 

30-35 533 550 492 25.1 27.7 2.5 20.0 

45-50 483 592· 482 34.4 28.2 12.4 16.0 

70-75 550 492 414 49.9 27.6 3.8 14.2 

95-100 542 608 534 67.3 27.4 2.1 8.2 

120-125 558 492 467 55.0 25.2 3.8 11.8 

145-150 425 433 387 79.5 22.9 6.3 8.7 

170-175 442 633 467 68.8 24.0 3.2 9.1 

195-200 525 467 441 59.5 21.4 5.4 3.6 

Core 32/ 0-1 533 525 454 12.5 29.6 2.5 0.7 

1-2 308 553 447 14.3 10.0 68.4 41.9 

2-4 167 567 397 17.5 22.5 14.2 27.8 

4-7 450 492 442 16.5 23.5 5.4 10.9 

7-11 375 450 384 23.7 27.9 7.2 18.2 

11-16 433 550 442 28.1 27.4 3.6 18.2 

20-25 375 175 359 53.0 26.8 2.3 2.7 

30-35 525 358 319 52.4 25.1 2.5 8.2 

45-50 392 158 302 49.3 26.4 3.2 4.6 

70-75 517 333 342 71.8 24.9 2.3 12.2 

95-100 483 333 342 59.5 22.6 3.1 2.9 

120-125 358 150 297 65.5 20.0 2.2 8.4 

145-150 392 233 282 87.3 20.5 4.0 10.2 

170-175 467 483 492 93.6 18.9 2.9 10.3 

195-200 417 292 299 84.2 16.6 2.3 9.1 

Core 19/ 0-1 450 350 274 9.4 29.5 4.8 .0.9 

1-2 442 317 309 30.2 27.0 6.1 2.0 

2-4 258 783 297 36.8 24.0 2.9 3.8 

4-7 183 925 242 35.6 27.6 1.8 10.0 

7-11 183 292 245 44.3 27.4 1.8 1.1 

11-16 267 333 220 57.5 27.1 2. 7· 1.1 

20-25 242 283 302 56.1 26.1 2.7 1.1 

30-35 242 367 232 77.4 26.4 1.8 1.1 

45-50 233 117 210 90.4 26.3 17.9 4.4 

70-75 375 267 310 102.7 26.3 2.1 6.2 

95-100 325 275 205 115.6 25.1 2.0 6.6 

118-123 150 217 140 57.5 24.2 1.8 4.0 

123-128 358 233 210 93.8 24.8 27.8 4.6 

145-150 217 400 232 62~2 24.2 3.2 4.4 

Influence of Fe and Mn on Chemical Partitioning of Cu, Zn, and Cr in Shelf Sediments 45 
Appendix 2. Concentrations of iron in extracts from cores 9, 32, and 19 

[in micromoles Fe per gram dry sediment; n.a., not analyzed] 

Depth (em) HOAc CHxl H202 Org HHxl Lith Sum1 Total 1 

Core 9/ 0-1 28.7 9.2 93.8 90.7 122 406 660 574 

1-2 31.3 8.5 114. 109 95.9 378 627 541 

2-4 48.9 5.6 101. 94.8 94.4 326 576 530 

4-7 28.2 8.3 94.7 87.6 132. 407 669 577 

7-11 39.3 7.9 108. 101. 138. 399 692 557 

11-16 59.0 6.0 96.7 87.4 90.7 368 621 548 

20-25 51.7 6.1 102. 91.3 98.7 321 580 577 

30-35 37.9 8.0 97.3 84.8 117. 401 661 578 

45-50 63.5 6.2 93.3 76.1 92.7 363 619 585 

70-75 60.8 6.8 111. 86.1 88.3 403 670 593 

95-100 52.3 5.5 90.9 57.3 99.4 449 697 598 

120-125 60.2 6.5 74.3 46.8 100.5 382 624 582 

145-150 50.1 5.7 81.1 41.4 104.4 369 610 529 

170-175 50.1 6.1 95.5 61.1 121. 342 616 549 

195-200 54.1 6.9 103. 73.0 110. 417 690 592 

Core 32/ 0-1 16.1 8.7 87.1 80.8 93.5 299 504 533 

1-2 19.9 8.3 91.6 84.4 93.3 288 501 386 

2-4 17.0 8.2 83.5 74.7 82.7 n.a. n.a. n.a. 

4-7 34.6 10.3 84.6 76.3 118. n.a. n.a. 529 

7-11 30.9 8.1 73.1 61.2 121. 316 549 432 

11-16 39.7 8.1 89.7 75.7 97.1 n.a. n.a. 509 

20-25 51.6 6.0 95.5 69.0 90.6 290 534 511 

30-35 59.4 8.4 89.3 63.1 93.8 n.a. n.a. 541 

45-50 66.5 7.7 101. 76.5 88.7 n.a. n.a. 466 

70-75 66.7 9.9 84.0 47.8 75.6 318 554 554 

95-100 49.1 7.3 93.6 63.8 88.6 576 815 588 

120-125 55.2 7.2 89.1 55.8 81.9 345 579 538 

145-150 35.6 6.0 101. 56.6 98.2 388 628 515 

170-175 29.0 4.2 90.1 43.3 117. 394 634 588 

195-200 40.4 6.0 116.9 74.5 97.7 316 577 530 

Core 19/ 0-1 10.4 13.6 84.0 79.2 122. 218 448 413 

1-2 17.7 9.0 102. 87.4 75.3 246 451 381 

2-4 22.5 6.2 89.8 71.4 61.5 193 373 323 

4-7 27.6 5.0 79.5 61.7 49.5 162 324 315 

7-11 32.1 4.8 105. 83.1 67.5 201 410 331 

11-16 35.6 4.8 88.3 59.6 80.4 194 403 383 

20-25 34.1 3.9 89.2 61.2 83.8 223 434 408 

30-35 34.7 5.4 119. 80.6 90.4 227 477 397 

45-50 30.0 5.0 84.3 39.0 70.9 213 403 400 

70-75 40.2 8.5 71.7 20.4 95.9 287 503 419 

95-100 35.0 9.0 102. 44.9 69.3 308 524 417 

118-123 17.8 5.2 95.5 66.8 34.7 156 310 293 

123-128 24.3 5.4 113. 66.3 94.2 272 509 476 

145-150 20.4 4.7 129. 85.7 67.9 183 406 417 

1sum = sum of values of all extractions; Total =value of a separate total 

extraction. 

46 Studies in Diagenesis 
Appendix 3. Concentrations of manganese in extracts from cores 9, 32, and 19 

[in micromoles Mn per gram dry sediment; n.a., not analyzed] 

Depth (em) HOAc CHxl H202 HHxl Lith Sum1 Totall 

Core 9/ 0-1 5.11 0.71 0.85 0.48 1.41 8.56 7.54 

1-2 4.75 .53 .87 .38 1.27 7.80 7.18 

2-4 4.62 .47 .75 .38 1.19 7.41 6.76 

4-7 4.73 .67 • 78 .50 1.36 8.04 7.41 

7-11 4.80 .74 .96 .52 1.34 8.37 7.40 

11-16 4.86 .58 .84 .35 1.23 7.85 7.49 

20-25 5.50 .56 .82 .36 .97 8.21 8.25 

30-35 5.50 .76 .84 .45 1.29 8.83 8.54 

45-50 5.50 .69 .87 .28 1.15 8.49 8.17 

70-75 5.55 .77 .89 .35 1.26 8.83 8.04 

95-100 6.72 .77 .80 .36 1.34 9.98 9.37 

120-125 6.32 .71 .75 .35 1.24 9.36 8.25 

145-150 5.13 .67 .76 .37 1.21 8.14 7.43 

170-175 5.90 .74 .78 .49 1.14 9.05 7.56 

195-200 5.64 .77 .90 .41 1.23 8.95 8.31 

Core 32/ 0-1 5.73 0.78 0.93 0.43 1.38 9.26 8.70 

1-2 5.13 .60 .73 .49 1.69 8.65 4.55 

2-4 4.53 .54 .64 .40 n.a. n.a. n.a. 

4-7 5.51 .73 .75 .56 n.a. n.a. 7.60 

7-11 4.64 .63 .71 .51 1.54 8.03 5.22 

11-16 4.55 .64 .71 .52 n.a. n.a. 5.83 

20-25 5.75 .72 .71 .49 1.26 8.93 6.36 

30-35 5.26 .85 .75 .52 n.a. n.a. 7.63 

45-50 5.33 .83 .76 .50 n.a. n.a. 7.39 

70-75 6.28 .97 .78 .34 1.27 9.65 8.17 

95-100 6.57 1.02 1.00 .42 1.90 10.9 9.18 

120-125 6.72 1.01 .87 .39 1.47 10.5 8.31 

145-150 6.64 1.20 .98 .43 1.49 10.8 7.18 

170-175 7.03 1.03 .93 .52 1.42 10.9 7.77 

195-200 5.84 .96 .95 .44 1.11 9.30 8.07 

Core 19/ 0-1 8.61 1.59 1.38 0.40 1.12 13.1 11.5 

1-2 3.24 .59 .99 .26 1.18 6.26 5.28 

2-4 2.97 .44 .76 .22 1.00 5.39 4.58 

4-7 2.51 .40 .69 .17 .97 4.75 4.24 

7-11 2.71 .47 .84 .21 1.29 5.54 4.32 

11-16 2.69 .46 .71 .30 1.20 5.36 4.22 

20-25 2.89 .49 .76 .33 1.09 5.56 4.85 

30-35 3.24 .61 .84 .30 1.20 6.20 5.09 

45-50 3.51 .64 .75 .2A 1.06 6.21 5.41 

70-75 3.73 .80 .76 .17 1.35 6.82 5.52 

95-100 3.84 .86 .88 .22 1.36 7.16 5.11 

118-123 2.31 .50 • 76 .12 1.04 4.73 4.21 

123-128 3.42 .72 .95 .35 1.23 6.67 5.96 

145-150 3.15 .55 1.04 .18 .88 5.81 5.43 

1sum = sum of values of all extractions; Total = value of a separate total 

extraction. 

Influence of Fe and Mn on Chemical Partitioning of Cu, Zn, and Cr in Shelf Sediments 47 
Appendix 4. Concentrations of Cu in in extracts from cores 9, 32, and 19 

[in micromoles Cu per gram dry sediment; HOAc-soluble and CHxl-soluble 

Cu generally below detection limit; n.a., not analyzed] 

Depth (em) HOAc H202 HHxl Lith Sum1 Total 1 

Core 9/ 0-1 <.012 0.055 0.070 0.070 0.190 0.205 

1-2 .045 .065 .050 .160 .190 

2-4 .050 .055 .060 .165 .200 

4-7 .055 .065 .065 .180 .200 

7-11 .065 .080 .055 .200 .190 

11-16 .055 .065 .045 .165 .200 

20-25 .055 .065 .040 .160 .185 

30-35 .050 .055 .060 .170 .185 

45-50 .050 .065 .050 .165 .180 

70-75 .050 .065 .060 .180 .205 

95-100 .050 .070 .055 .180 .205 

120-125 .065 .070 .055 .185 .195 

145-150 .055 .065 .045 .160 .180 

170-175 .070 .080 .055 .205 .185 

195-200 .065 .070 .060 .195 .205 

Core 32/ 0-1 0.025 0.055 0.040 0.095 0.190 0.220 

1-2 .030 .045 .030 .090 .160 .215 

2-4 .030 .045 .020 n.a. n.a. n.a. 

4-7 .035 .045 .040 .155 .240 .240 

7-11 .040 .060 .060 .090 .210 .235 

11-16 .020 .050 .050 n.a. n.a. .245 

20-25 .010 .055 .020 .110 .185 .240 

30-35 .010 .065 .065 n.a. n.a. .230 

45-50 <.01 .050 .050 n.a. n.a. .200 

70-75 <.01 .075 .040 .095 .210 .285 

95-100 <.01 .100 .070 .115 .285 .300 

120-125 <.01 .070 .045 .085 .200 .225 

145-150 .010 .075 .050 .105 .235 .240 

170-175 .015 .070 .085 .115 .270 .240 

195-200 .015 .080 .030 .095 .210 .220 

Core 19/ 0-1 <0.012 0.015 0.050 0.085 0.150 0.140 

1-2 .040 .015 .100 .155 .185 

2-4 .030 .010 .085 .120 .155 

4-7 .035 .010 .085 .120 .155 

7-11 .050 .015 .100 .160 .140 

11-16 .035 .010 .090 .135 .150 

20-25 .040 .015 .055 .105 .150 

30-35 .040 .030 .085 .155 .165 

45-50 .045 .025 .070 .140 .185 

70-75 .050 .035 .090 .170 .215 

95-100 .055 .030 .090 .175 .220 

118-123 .030 <.01 .p75 .105 .090 

123-128 .055 .030 :o95 .180 .130 

145-150 .040 .015 .085 .140 .145 

1sum = sum of values of all extractions; Total = value of a separate 

total extraction. 

2H0Ac-soluble Cu is less than 0.01 micromoles per gram for all samples 

in cores 9 and 19. 

48 Studies in Diagenesis 
Appendix 5. Concentrations of Zn in extraction cores 9, 32, and 19 

[in micromoles Zn per gram dry sediment; n.a., not analyzed] 

Depth (em) HOAc CHxl H202 HHxl Lith Sum1 Total 1 

Core 9/ 0-1 0.095 0.050 0.375 0.320 0.675 1.52 1.22 

1-2 .075 .055 .360 .280 .650 1.42 1.35 

2-4 .100 .045 .305 .260 .570 1.28 1.23 

4-7 .080 .065 .375 .345 .690 1.55 1.36 

7-11 .090 .060 .435 .365 .670 1.61 1.30 

11-16 .110 .060 .390 .265 .645 1.47 1.26 

20-25 .100 .050 .405 .280 .580 1.41 1.29 

30-35 .085 .050 .390 .315 .765 1.61 1.39 

45-50 .125 .050 .390 .260 .650 1.47 1.29 

70-75 .105 .050 .415 .280 .735 1.58 1.40 

95-100 .100 .050 .350 .270 .725 1.51 1.29 

120-125 .115 .060 .385 .265 .685 1.51 1.27 

145-150 .110 .055 .340 .290 .600 1.39 1.14 

170-175 .110 .060 .395 .345 .620 1.53 1.26 

195-200 .135 .060 .420 .325 .605 1.61 1.24 

Core 32/ 0-1 0.100 0.045 0.235 0.250 0.640 1.27 1.35 

1-2 .180 .070 .255 .270 .610 1.39 1.35 

2-4 .095 .050 .145 .415 n.a. n.a. n.a. 

4-7 .105 .060 .240 .350 n.a. 0.75 1.35 

7-11 .120 .080 .195 .385 .615 1.40 1.14 

11-16 .090 .045 .225 .305 n.a. n.a. 1.13 

20-25 .160 .070 .170 .330 .550 1.28 1.22 

30-35 .155 .085 .235 .325 n.a. n.a. 1.49 

45-50 .145 .060 .170 .370 n.a. n.a. 1.42 

70-75 .150 .065 .280 .235 .665 1.40 1.33 

95-100 .125 .055 .315 .270 .835 1.60 1.35 

120-125 .130 .055 .295 .260 .715 1.45 1.32 

145-150 .125 .050 .315 .300 .770 1.56 1.35 

170-175 .090 .050 .310 .350 .740 1.54 1.38 

195-200 .115 .050 .400 .265 .765 1.60 1.34 

Core 19/ 0-1 0.080 0.050 0.235 0.280 0.545 1.19 1.07 

1-2 .095 .055 .295 .180 .715 1.34 1.08 

2-4 .085 .040 .225 .120 .615 1.08 0.90 

4-7 .085 .040 .220 .100 .610 1.05 0.88 

7-11 .095 .040 .325 .150 .450 1.06 0.91 

11-16 .115 .050 .230 .160 .450 1.01 0.96 

20-25 .100 .045 .230 .180 .435 0.99 1.10 

30-35 .110 .050 .235 .250 .550 1.19 1.13 

45-50 .090 .050 .235 .175 .500 1.06 1.10 

70-75 .100 .065 .240 .175 .665 1.24 1.25 

95-100 .095 .065 .260 .180 .730 1.34 1.35 

118-123 .060 .045 .230 .075 .355 0.77 0.81 

123-128 .075 .045 .325 .210 .650 1.30 1.22 

145-150 .070 .045 .350 .155 .470 1.09 1.07 

1sum = sum of values of all extractions; Total = value of separate total 

extractions. 

Influence of Fe and Mn on Chemical Partitioning of Cu, Zn, and Cr in Shelf Sediments 49 
Appendix 6. Concentrations of Cr in extracts from cores 9, 32, and 19 

[in micromoles Cr per gram dry sediment; CHxl-soluble Cr below detection 

1 imit; n .a., not analyzed] 

Depth (em) HOAc H202 HHxl Lith Sum1 Total 1 

Core 9/ 0-1 0.015 0.125 0.080 0.975 1.20 1.09 

1-2 .015 .140 .070 0.960 1.19 1.16 

2-4 .025 .130 .085 0.905 1.14 1.12 

4-7 .025 .110 .100 0.970 1.22 1.14 

7-11 .025 .130 .100 1.05 1.30 1.12 

11-16 .025 .120 .050 0.910 1.10 1.10 

20-25 .025 .125 .070 0.850 1.07 1.13 

30-35 .015 .100 .085 1.02 1.22 1.22 

45-50 .025 .135 .065 0.935 1.16 1.15 

70-75 .025 .125 .065 0.975 1.19 1.22 

95-100 .015 .110 .060 0.935 1.12 1.18 

120-125 .020 .115 .060 0.980 1.17 1.19 

145-150 .015 .113 .080 0.890 1.10 1.05 

170-175 .020 .125 .090 0.940 1.18 1.14 

195-200 .025 .135 .080 1.02 1.26 1.17 

Core 32/ 0-1 2<0.01 0.085 0.035 0.94 1.05 1.17 

1-2 .065 .035 0.82 0.92 1.59 

2-4 .060 .025 n.a. n.a. n.a. 

4-7 .075 .035 n.a. n.a. 1.45 

7-11 .080 .045 0.81 0.94 1.37 

11-16 .085 .040 n.a. n.a. 1.39 

20-25 .105 .015 0.88 1.00 1.55 

30-35 .075 .050 1.54 1.66 1.63 

45-50 .100 .035 n.a. n.a. 1.29 

70-75 .090 .025 1.15 1.26 1.44 

95-100 .110 .025 1.32 1.45 1.50 

120-125 .080 .035 1.09 1.20 1.35 

145-150 .090 .030 1.19 1.31 1.55 

170-175 .060 .035 1.18 1.27 1.54 

195-200 .090 .030 1.08 1.20 1.32 

Core 19/ 0-1 0.010 0.110 0.030 1.05 1.20 0.79 

1-2 .010 .125 .040 1.07 1.24 1.21 

2-4 .010 .105 .035 0.770 0.92 1.00 

4-7 .010 .110 .030 0.700 0.85 0.94 

7-11 .010 .120 .030 0.860 1.02 0.92 

11-16 .010 .115 .035 1.02 1.19 1.08 

20-25 .015 .100 .040 0.96 1.12 1.16 

30-35 .010 .115 .035 1.06 1.22 1.20 

45-50 .010 .110 .030 0.99 1.14 1.27 

70-75 .015 .115 .025 1.20 1.34 1.44 

95-100 .010 .130 .020 1.27 1.43 1.46 

118-123 .010 .110 .010 J).74 0.87 0.59 

123-128 .010 .135 .025 1.36 1.53 0.90 

145-150 .010 .170 .015 1.02 1.21 0.79 

1sum = sum of values of all extractions; Total = value of separate 

total extraction. 

2H0Ac-soluble is less than 0.01 micromoles per gram for all samples 

in core 32. 

50 Studies in Diagenesis 
Mo(,ility and Distribution of Heavy Metals 

During the Formation of First Cycle Red 

Beds, 

By RC>BERT A. ZIELINSKI, U.S. Geological Survey, 

MS 424, Denver Federal Center Denver, Colorado 80225, 

SALMAN BLOCH, Oklahoma Geological Survey, 830 

VanVleet Oval, Norman, Oklahoma 73019, and 

THEODORE R. WALKER, University of Colorado, 

Department of Geology, Boulder, Colorado 80307 

CONTENTS 

Abstract 52 

Introduction 52 

Experimental 52 

Results :and discussion 54 

Conclusions 58 

References Cited 59 

FIGURES 

1. Generalized geographic map showing sampled outcrop locations 53 

2. Plot showing average composition of the three types of samples studied 55 

3. Graph showing whole-rock percentage of heavy metals dissolved by each of the 

six l<!ach solutions (see table 1) 56 

4. Plot showing total leached whole-rock percentage of heavy metals of the three 

types of samples studied 57 

5. Plot showing comparison of percentage of whole-rock iron in leachates 5, 6, and 

7 versus type of sample studied 58 

TABLE 

1. Summary of leaching conditions 54 

1

Summarized from Zielinski, R. A., Bloch, Salman, and Walker, T. R., 1983, The mobility 

and distribution of heavy metals during the formation of first cycle red beds: Economic Geology, 

v. 78, p. 1574-1589. 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 51 
Abstract 

A geologically young (Holocene-Pliocene) red-bed 

sequence displaying successive stages of development of the 

red pigmentation was sampled in northern Baja California for 

the purpose of studying the mobility and distribution of heavy 

metals during aging of iron oxyhydroxides from the noncrystalline to the crystalline state. Twenty-seven whole-rock 

samples and a series of successive leachates chosen to dissolve 

specific forms of iron oxides (noncrystalline, poorly crystalline, 

crystalline) were analyzed by inductively coupled plasmaoptical emission spectrometry for V, AI, Cr, Mn, Fe, Co, Ni, 

Cu, and Zn, and by a delayed neutron technique for U. These 

geochemical data were supported by petrographic, X-ray diffraction, and scanning electron microscopic observations of 

mineralogical transformations, and by fission-track radiography, 

and uranium decay-series measurements. The results indicate 

that the metal content of the studied samples was inherited 

predominantly from detrital minerals and that reddening of the 

rock did not promote major open-system migration of heavy 

metals. Rather, contained metals redistributed on an intergranular scale, moving from detrital mineral hosts to secondary iron 

oxides. The amount of secondary iron oxides and the fraction 

of whole-rock metals associated with the oxides increased during the reddening of these beds. Thirteen fine-grained red beds 

of Pliocene age sampled in this study contain 10-63 percent 

of their whole-rock metal abundance in leachable sites, that 

is, secondary iron oxides, minor manganese oxides, and minor 

exchangeable sites on clays. The average leachable fractions 

(in percent) increase in the order Cr (10), V (15), Pb (?20), Zn 

(22), U (27), Fe (30), Ni (41 ), Cu (44), Co (47), Mn (63). These 

results suggest that red beds that are well flushed by suitable 

pore fluids (oxidizing or reducing) may be sources of significant quantities of heavy metals. 

INTRODUCTION 

Research conducted within the last two decades has 

shown that the pigmentation of red beds is produced by 

diagenetic alteration of constituent iron-bearing minerals 

(Walker, 1967, 1974; Walker and others, 1978; Larsen 

and Walker, 1975). In-place precipitation and aging of 

secondary iron oxides (noncrystalline iron oxides, 

goethite, hematite) during diagenesis of the host sediments has important implications for the geochemical 

mobility of contained heavy metals. Secondary iron 

oxides, particularly hydrous forms, are well-documented 

scavengers of heavy metals in soils and sediments; 

however, some changes in solution chemistry, such as 

lowered pH or the introduction of complexing agents, can 

promote the release of previously adsorbed metals. In 

addition, solutions of low Eh may cause the release of 

metals by the dissolution of secondary iron oxides 

(bleaching). The concept of mobilization of metals 

previously adsorbed by red beds suggests that red beds 

may be a source as well as a concentrator of economically 

important metals. Observations in support of this 

52 Studies in Diagenesis 

hypothesis are the close association of red beds with 

sediment-hosted stratiform deposits of copper and, to a 

lesser extent, similar deposits of lead and zinc (Gustafson and Williams, 1981). Other heavy metals such as Co, 

Ag, U, and V are common as minor or major components 

of metal deposits in and near red beds. · 

The geochemical mobility of economically important metals during early diagenetic reddening of host 

sediments is addressed in this study. Whole-rock and 

leachate analyses for Fe, Mn, AI, Cu, Pb, Zn, Co, Ni, 

Cr, V, Pb, and U are combined with petrographic, 

radiographic, and isotopic data to determine the distribution and mobility of these metals as a function of 

sedimentary facies, sample age, and degree of reddening. 

In addition, the leachate results are used to speculate on 

the relative effectiveness of subsequent leaching of red 

beds for liberating metals. 

EXPERIMENTAL 

The 27 samples chosen for study are increasingly 

reddened (with age) alluvial fan and tidal-flat sediments 

of Holocene to Pliocene age from northeastern Baja 

California (fig. 1). The area is the site of previous red 

bed studies (Walker, 1967; Walker and others, 1978; 

Larsen and Walker, 1975) and represents a depositional 

environment that is analogous to that of many ancient 

red beds, such as those of Permian and Triassic age in 

Europe and North Africa and those of Pennsylvanian to 

Triassic age in the western interior of North America. 

Whole-rock powders (less than 100 mesh), leach solutions, and residues were analyzed by inductively coupled 

plasma-optical emission spectrometry and by a delayed 

neutron procedure for uranium. 

To complement the whole-rock data, each sample 

was crushed, sieved, and subjected to a series of six successive leaching procedures, each chosen to be moderately 

specific for a particular metal site (table 1). Most leach 

solutions were acidic and (or) contained complexing 

agents to stabilize dissolved species. Eh ranged from 

strongly oxidizing (leach solution 3) to strongly reducing 

(leach solution 6). In general, leach solutions 1-3 removed 

loosely bound metals and metals associated with minor 

organic matter and manganese oxides, whereas leach solutions 4-6 were stronger and promoted iron oxide dissolution and attendant bleaching of samples. Additional 

details of the experimental procedure were described by 

Zielinski and others (1983). 

The use of site-specific leaching for determining the 

distribution of metals is admittedly imprecise, and the 

yield of metals from a given leaching may not be complete. Nevertheless, such results are useful for comparing the relative leachability of metals from similarly 

treated samples, under a variety of conditions. 
116° 114° 330r---------------------------~,_=-------------~~--~~-----r------~ .~ 

~;.,_ . ~..;,;.... ~ .,~;;;\ 

El Centro- -:-:-.....:.;... 

Calexico UN\I~P §I~1_E.§.-: ~ -------~- MEXICO ·· -- ~ Mexicali 

OUTCROP AREA OF 

DEPOSITS OF LATE 

PLIOCENE AGE 

EXPLANATION 

I ~ I Motmtains 

j ;_.;;;;_. : I Sand dtmes 

t==_:==_:=j Playa and mudflat 

t ; : ; :I Saltpan 

0 10 20 30 KILOMETERS I I I I I I I 

0 10 20 MILES 

Figure ·a. Generalized geographic map showing sampled outcrop locations. Modified from Larsen and Walker (1975). 

Mobility and Distribution of Heavy Metals During the Formation of First Cycle Red Beds 53 
Table 1. Summary of leaching conditions 

[h, hours; M, molar; "'• approximate value] 

Leach Solution Contact 

solution composition pH time (h) 

No. 

1 1 _!!MgC1 2 5 2 

2 0.1 _!! NH 2

0H•HC1 4 2 

3 ""15% H202 --4 5 

4 0.2!! ammonium oxalate, 3 5 

0.2 M oxalic acid. 

5 1 _!! NH20H·HC1, 25% ""2 5 

(V/V) acetic acid. 

6 Na-dithionite (solid), 8.3 0.5 

""0 .3 _!! NaHC03, 

""0.4 _!! Na-citrate. 

RESULTS AND DISCUSSION 

For the purpose of this study it was most informative to divide the samples into three groups-alluvial 

fan sediments (7) of Holocene and Pleistocene age, tidalflat sediments (7) of the same age range, and tidal-flat 

sediments (13) of Pliocene age (fig. 2). Alluvial fan sediments of Holocene to Pleistocene age (column 1, fig. 2) 

were consistently found to contain more aluminum but 

less iron and heavy trace metals than contemporaneous 

tidal-flat sediments (column 2). The large difference in 

abundance of aluminum, a major element that is 

relatively insoluble, indicates that the difference in bulk 

composition of the sediments was probably due to differences in source composition or to mineral fractionation during transport. For example, the alluvium is rich 

in locally derived granitic detritus, whereas the tidal-flat 

sediments are rich in detritus transported by the Colorado 

River from more distant and compositionally varied 

sources. Alternatively, mineral fractionation during 

transport increased the content of slow-settling phases, 

such as clay minerals and biotite, in fine-grained 

sediments. 

In contrast to these results, the average composition of tidal-flat sediments of Holocene-Pleistocene age 

(column 2) was similar to that of Pliocene equivalents 

(column 3). This similarity in composition supports the 

assumption of constant depositional setting for these 

sediments and suggests that chronologically progressive 

reddening was not accompanied by major fluxes of heavy 

metals, at least on a hand-specimen scale. 

54 Studies in Diagenesis 

Site or phase attacked Comments 

Exchangeable sites on clay; 

carbonates. 

Manganese oxides------------- Dissolves some 

noncrystalline 

oxides. 

Organic matter--------------- At 50°C; highly 

oxidizing. 

Noncrystalline iron oxides--- Treat in darkened 

room; dissolves 

some magnetite. 

Poorly crystallized secondary iron oxides. 

Well-crystallized May attack ironsecondary iron oxides. bearing clays; 

7 5° c; hi gh 1 y 

reducing. 

The amount of each element dissolved by each leach 

solution is expressed as a percentage of its original wholerock abundance. Results for all27 samples are summarized in figure 3. The total percentages of lead and 

uranium leached from whole-rock samples are minimum 

values because of the relatively high detection limits for 

these elements (Pb = 5-17 percent and U = 2-5 percent 

of their whole-rock abundances). Leached percentages of 

aluminum provide an upper-limit estimate of the extent 

of aluminosilicate dissolution. Elements detected in 

leachate 1 included Mn, Co, V, Ni, and Cu. These 

elements, plus Zn, were also detected in the leachates 

chosen to attack manganese oxides and organic material. 

Although these latter two materials have high adsorptive 

capacity, they are apparently too sparse to exert major 

influence on the heavy-metal budget of the studied 

samples. In contrast to the first three leach solutions, the 

leach solutions chosen to attack secondary iron oxides 

dissolved major percentages of whole-rock metals. In 

addition to iron, the metals with the greatest relative 

affinities for the secondary iron oxides appear to be Mn, 

Cu, Ni, Co, and, to a lesser extent, U, Zn, and Pb; AI, 

Cr, and V consistently showed the smallest percentages 

of whole-rock abundance in secondary iron oxides. 

Additional information was derived by subdividing 

total leached whole-rock percentages into the three sample 

groups of figure 2, as shown in figure 4. In spite of lower 

metal abundances in alluvium (fig. 2) and a coarser average grain size in leached powders of alluvium, most of 

the measured heavy metals were more effectively leached 

from the alluvium. This difference in leachability was 
90 

80 

70 

z 

0 

:::::i 

..J 

~ 

a: 60 w 

a.. 

en 

1-

a: 

<{ 

50 a.. 

~ 

en' 

..J 

<{ 

1-

w 40 ~ 

w 

u 

<( 

a: 

1-

30 

20 

AI 

Zn 

v 

Mn 

10 

Cr 

Pb 

Fe 

Cu 

Ni 

9 

8 

7 

1- 6 z 

w 

u 

a: 

w 

a.. 

1-

5 :I: 

(.!) 

jjj 

~ 

~ 

4 ~ 

<f 

3 

2 

10 

cor-------------------~------------------~ Co 

5 

u•e------------------~e.------------------.eu 

AVERAGE AVERAGE 

ALLUVIUM (7) 

HOLOCENE- PLEISTOCENE 

TIDAL FLAT (7) 

AVERAGE 

TIDAL FLAT ( 13) 

HOLOCENE- PLEISTOCENE PLIOCENE 

Figure 2. Average composition of samples, plotted as a function of their depositional setting and age. The number of 

samples in each subgroup is shown in parentheses. Representative error bars encompass ± 1 standard error of the mean. 

observed in most leach steps and probably indicates fundamental differences in the distribution of metals and the 

accessibillity of leach solutions to metal sites. In the finer 

grained tidal-flat materials, more of the metals were in 

unleachable structural sites of the relatively abundant clay 

minerals and biotite, and access of leach solutions to surfaces of iron-bearing grains was hindered by clay coatings 

that were not removed during grinding. This effect was 

partly an experimental problem, but is probably also 

operative in nature. 

The leachability of Pb, U, and, particularly, Mn was 

not as sensitive to these proposed differences in intergranular permeability, perhaps because more of these 

elements were dispersed in minerals that are coarser 

grained or more insoluble. Possible minerals include 

feldspar for lead; zircon, sphene, and apatite for uranium; 

Mobility and Distribution of Heavy Metals During the Formation of First Cycle Red Beds 55 
60 

50 

0 

w 

I 

u 

<( 

w 40 .....1 

w 

(!) 

<( 

1-

z 

w 30 u 

a: 

w 

a.. 

w 

(!) 

<( 

a: 20 w 

> <( 

10 

5 

0 

AI Cr v Zn Pb Fe Ni u Cu Co 

ELEMENT 

Mn 

EXPIANA TION 

Well crystallized 

iron oxides 

Poorly crystallized 

iron oxides 

Noncrystalline 

iron oxides 

Organic matter 

Manganese oxides 

Figure 3. The average whole-rock percentage of heavy metal dissolved by each of six leach solutions. Leach solutions 

are coded according to the particular metal site under attack. Vertical arrows identify elements with relative high detection 

limits and therefore underestimated leachability. 

and feldspars and titanium phases for manganese. The 

exceptional leachability of manganese from finer grained 

samples suggests an additional association with clay 

surfaces. This hypothesis is supported by the results 

of leaching that show a relatively high percentage of 

readily exchangable manganese in the tidal-flat sediments. These results and the results shown in figures 2 

and 4 indicate that studies of subtle compositional 

changes related to progressive diagenesis must be 

restricted to sediments that are not only products of 

similar depositional settings but that are also of similar 

grain size. 

With the possible exception of lead, all measured 

metals were leached in greater percentage from the Pliocene tidal-flat sediments than from the younger equivalents (fig. 4). Increases in leachability with age averaged 

about 6 percent of whole-rock abundances; iron showed 

the greatest increase (about 10 whole-rock percentage 

units). This increased leachability is attributed to timedependent, intergranular-scale redistribution of metals 

from detrital mineral hosts to leachable sites, primarily 

secondary iron oxides. As noted, this redistribution of 

metals did not produce major compositional changes on 

a hand-specimen scale (fig. 2). 

56 Studies in Diagenesis 

The amount of iron dissolved by leach solutions 4-6 

indicates a marked increase in the abundance of wellcrystallized secondary iron oxides (for example, hematite), in the older samples (fig. 5). In another recasting 

of the data, heavy-metal/iron ratios calculated for each 

of the final three leachates indicate generally lower ratios 

in the more abundant, well-crystallized secondary iron 

oxides. Thus, the apparent closed system behavior of the 

heavy metals during the diagenetic reddening of the tidalflat sediments was in part the result of distribution of 

metals between a coexisting population of metal-rich noncrystalline iron oxides and relatively metal-poor, but more 

abundant, crystalline iron oxides, such as goethite and 

hematite. 

Fission-track radiographs (not presented here) 

confirmed the association of uranium with uraniferous 

accessory minerals (zircon, sphene), as well as with 

secondary iron oxides. Uranium decay-series measurements indicated a dominantly closed-system behavior for 

uranium in the sedimentary environment; however, 

samples with a relatively large abundance of uranium in 

noncrystalline iron oxides (leach solution 4) retained an 

isotopic signature indicative of minor adsorptive uptake 

of uranium. 
w 

(!) 50 ~ Co z 

w 

u 

w Cu 

a: Ni a.. 40 Ni 

0 

w 

I 

u 

<( 

w 

....J Mn ....J 30 Fe 

<( Fe 

1- v u 0 

1- Zn 

u Zn 

20 Cr Pb 

Pb v 

10 Cr 

5 

AVERAGE ALLUVIUM (7) 

HOLOCENE -PLEISTOCENE 

AVERAGE TIDAL FLAT (7) 

HOLOCENE- PLEISTOCENE 

AVERAGE TIDAL FLAT (13) 

PLIOCENE 

Figure 4. Total leached whole-rock percentage of heavy metals, plotted as a function of depositional setting and 

age of samples. The number of samples in each subgroup is shown in parentheses. Representative error bars encompass ± 1 standard error of the mean. 

The results of this study indicate a dominant closedsystem behavior of iron and associated metals during 

the early diagenetic reddening of the tidal-flat sediments. Short-range, perhaps intergranular-scale, movement of metals was favored because of the intimate 

association of metal-bearing detrital grains and adsorptive grain coatings of secondary iron oxides, the favorable chemical environment (basic-oxidizing) for metal 

adsorption, and the limited permeability of host sediments. The leachate data, however, also suggest that 

subsequent changes in the Eh, pH, or accessibility of 

pore fluids during more prolonged burial and diagenesis might promote renewed movement of previously 

fixed metals. Given such changes, young, coarsegrained red beds should react more rapidly than old, 

fine-grained red beds because of their high intergranular permeability and their greater content of more reactive, noncrystalline and poorly crystalline secondary iron 

oxides. Old red beds, however, are potentially a greater 

source of metals than young red beds because secondary 

iron oxides are more abundant and contain higher percentages of whole-rock heavy metals. Reducing solutions that promoted dissolution of secondary iron 

oxides (bleaching) were the most effective for liberating 

heavy metals, but, in general, the geochemical mobility 

of released metals may be hindered by the formation of 

insoluble oxides, carbonates, or sulfides. Moderate 

amounts (1-20 percent) of heavy metals were released 

during leach stages 1 through 4, prior to obvious bleaching of samples. As in most of the leaching steps, Mn, 

Co, and Cu were liberated in the greatest percentages. 

Anomalously low concentrations of these elements in 

some red beds compared with regional or worldwideaveraged values for red beds may be an indication of 

incipient open-system leaching (Zielinski and others, 

1983). 

Mobility and Distribution of Heavy Metals During the Formation of First Cycle Red Beds 57 
30 

w 

I-

<( 

I 

u 

<( 20 w 

..J 

~ 

z 

0 

Q: 

~ 

u 

0 

a: 

I 

w 

..J 

0 

I 

~ 

LL 

0 

I- 10 z 

w 

u 

a: 

w 

a.. 8 

6 

........ Poorly ---... crystallized · --....._, " 0 4 --.____ n oxid es 

--........................ 2 

• . Noncrystalline iron oxides 

------,-------· 

AVERAGE HOLOCENE + 

PLEISTOCENE TIDAL FLAT (7) 

AVERAGE PLIOCENE 

TIDAL FLAT (13) 

Figure 5. Comparison of percentage of whole-rock iron in leachates 5, 6, and 7 with type 

of sample studied. 

CONCLUSIONS 

1. Large and systematic chemical differences between alluvial fan and tidal-flat sediments of similar age 

and depositional setting were primarily controlled by differences in provenance and (or) mineral fractionation during deposition. The two sediment types also displayed 

large differences in their laboratory leaching characteristics that were related to fundamental differences in 

metal-site distribution and intergranular permeability. 

These findings emphasize the need to consider a single 

58 Studies in Diagenesis 

sediment type when evaluating compositional differences 

related to progressive diagenesis. 

2. The tidal-flat sediments of Holocene-Pliocene 

age chosen for further study showed little analytically 

detectable change in heavy-metal content as a function 

of age, indicating that they had not experienced major 

open-system fluxes of heavy metals during progressive 

diagenetic reddening. Laboratory leaching of the same 

samples, however, indicated major within-sample redistribution of heavy metals that could be related to reddening. For example, as the abundance of secondary iron 
oxides increased, so did the fraction of whole-rock metals 

associated with them. Well-crystallized forms of secondary 

iron oxides increased relative to noncrystalline and poorly 

crystalline forms in the older, redder samples. Heavy-metal/ 

iron ratios in secondary iron oxides decreased as a function of crystallinity, but the possible loss of heavy metals 

related to the crystallization of iron oxides was apparently 

compensated by the consistently greater abundance of wellcrystallized secondary iron oxides in samples of all ages. 

3. Hydrous manganese oxides or organic matter were 

not sufficiently abundant in the studied samples to have 

a significant effect on the distribution and mobility of 

metals. 

4. Laboratory leaching of all the red-bed sediments 

(alluvium and tidal-flat) by low-temperature reducing solutions was particularly effective for liberating heavy metals 

from secondary iron oxides, but the long-range geochemical 

mobility of heavy metals released in nature is probably 

hindered by their precipitation as relatively insoluble oxides, 

sulfides, or carbonates. In contrast, the experimental lowtemperature leaching of red beds by oxidizing solutions was 

less effective for liberating metals, but less effective leach 

solutions may be compensated for in nature by increased 

duration of leaching and by the availability of large 

volumes of solutions. 

REFERENCES CITED 

Gustafson, L. B., and Williams, Neil, 1981, Sediment-hosted 

stratiform deposits of copper, lead, and zinc, in Skinner, 

B. J., ed., Economic geology, Seventy-fifth anniversary 

vqlume, 1905-1980: Economic Geology Publishing Co., 

p. 139-178. 

Larsen, E. E., and Walker, T. R., 1975, Development of 

chemical remanent magnetization during early stages of 

red-bed formation in late Cenozoic sediments, Baja California: Geological Society of America Bulletin, v. 86, 

p. 639-650. 

Walker, T. R., 1967, Formation of red beds in modern and 

ancient deserts: Geological Society of America Bulletin, 

v. 78, p. 353-368. 

__ 1974, Formation of red beds in moist tropical climates; 

a hypothesis: Geological Society of America Bulletin, v. 85, 

p. 633-638. 

Walker, T. R., Waugh, B., and Crone, A. J., 1978, Diagenesis 

in first-cycle desert alluvium of Cenozoic age, southwestern 

United States and northwestern Mexico: Geological Society 

of America Bulletin, v. 89, p. 19-32. 

Zielinski, R. A., Bloch, S., and Walker, T. R., 1983, The mobility and distribution of heavy metals during the formation 

of first-cycle red beds: Economic Geology, v. 78, p. 1574-

1589. 

Mobility and Distribution of Heavy Metals During the Formation of First Cycle Red Beds 59 

Stoic:hiometry of Organic Decomposition in 

Sediments of the Potomac River Estuary 

By DAVID L. PARKHURST, U.S. Geological Survey, Room 621, 215 Dean A. 

McGee St., Oklahoma City, Oklahoma 73102, and 

STEVEN D. GOODWIN, U.S. Geological Survey, MS 432, National Center, 

Reston, Virginia 22092 

CONTENTS 

Abstract 62 

Introduction 62 

Setting 62 

Sampling and analytical methods 63 

Stoichiometry determined from pore-water concentrations 64 

Mass balance calculations 65 

Stoichiometry determined from concentrations in the solid phases 69 

Summary 71 

References cited 72 

FIGURES 

1. Map showing location of sampling stations V3 and V6 in the Potomac River 

estuary 63 

2. Pore·-water data used to determine the C:N ratio of organic decomposition 66 

3. Pore·-water data used to determine the C:S ratio of organic decomposition 67 

4. Concentration profiles of sulfate, alkalinity, and ammonium in bottom 

sediments at station V3 for May 1978 and September 1978 68 

5. Concentration profiles of carbon and nitrogen in the solid phases from 

stations V3 and V 6 70 

6. Concentration of carbon versus nitrogen in the solid phases from stations 

V3 and V6 71 

TABLES 

1. Mass accounting for the period May to September 1978 for station V3 69 

2. Rates of processes calculated for the period May to September 1978 for 

station V3 69 

3. Flux rates calculated from analyses of pore water and solid phases 70 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 61 
Abstract 

The decomposition of organic matter in sediment was 

investigated at two stations near the mouth of the Potomac River 

estuary. Using the assumption of steady state, equations were 

derived to determine the C:N ratio (carbon to nitrogen ratio) 

of the reactive organic matter from (1) the pore-water concentrations of alkalinity and ammonium, and (2) the concentration of carbon and nitrogen in the solid phases of the sediment. 

From the pore-water data, the C:N ratio was 7.7:1 for decomposition of organic matter by the pathway of sulfate reduction 

and 7.2:1 by the pathway of methanogenesis. From the solidphase data, the C:N ratio was 5.9:1; no distinction could be 

made betweem the pathways of sulfate reduction and 

methanogenesis from the solid-phase data. These estimates of 

the C: N ratio of the reactive part of the organic matter are lower 

than the average C:N ratio, 9.6:1, of the organic matter atthese 

sites, suggesting that organic matter with a low C:N ratio is 

decomposed preferentially to organic matter with a high C:N 

ratio. 

Pore-water concentrations of alkalinity and sulfate were 

used to determine the C:S ratio (ratio of carbon released to 

sulfate consumed) in the decomposition of organic matter by 

sulfate reduction. The estimated C:S ratio, -2.1:1, was close 

to the theoretical value of -2.0:1. 

A mass balance accounting was made for carbon, 

ammonium, and sulfate at one station for May 1978 to 

September 1978. Between 10 and 20 percent of the carbon 

and ammonium, produced by the decomposition of organic 

matter, was retained in the sediment pore water. More than 

80 percent of the carbon and ammonium, produced by decomposition, was released from the sediment pore water to the 

overlying water. The decrease in sulfate concentration in the 

pore water during this period accounted for 19 percent of the 

sulfate consumed in organic decomposition reactions. About 

80 percent of the sulfate consumed in organic decomposition 

reactions was supplied by the overlying water. 

The accumulation of carbon, ammonium, and sulfate in 

the pore water of the sediment due to burial by the accumulating sediment particles was insignificant. Burial accounted 

for less than 1 percent of the carbon, ammonium, and sulfate 

involved in decomposition reactions. Only a small fraction 

of the total ammonium released by the decomposition of 

organic matter was adsorbed or exchanged onto the solid 

phases of the sediment. The mass balance calculations 

demonstrated that nonsteady-state effects accounted for less 

than 20 percent of the total rate of organic decomposition at 

one site. These calculations support the steady-state assumption made in the derivation of the equations used to determine 

the stoichiometry of the reactive part of the organic matter. 

INTRODUCTION 

One of the most important chemical reactions in 

estuarine sedimentary environments is the decomposition 

of organic matter. The decomposition process controls 

the rate of release of nutrients from the sediment to the 

overlying water of the estuary and determines the distribution of chemical species within the solid phases and pore 

62 Studies in Diagenesis 

water of the sediment. The goal of this study was to 

characterize the decomposition of organic matter in the 

sediment of the Potomac River estuary by determining 

the stoichiometry of the decomposition reactions, that 

is, the ratios in which carbon, ammonium, and sulfate 

are produced or consumed in the decomposition process. 

The Potomac River estuary sediment contains 

organic matter which is derived from both marine and 

terrigenous sources. The compositions of these materials 

differ. Redfield (1958) determined an average C:N ratio 

(carbon to nitrogen ratio) in a sample of marine plankton 

to be 106:16 (6.6:1). Pockington (1976) found that terrigenous organic matter had significantly higher C:N 

ratios. The overall C:N ratio of the depositing organic 

matter will affect the stoichiometry of the decomposition 

reaction, but workers have found that the C:N ratio of 

the reactive part of the organic matter may not be the 

same as the overall C:N ratio. Several studies (HolmHansen and others, 1966; Kamatani, 1969; Gordon, 1971) 

have suggested that nitrogen may be preferentially released in the decomposition of organic matter near the 

sediment-water interface. Thus, the stoichiometry of the 

carbon, ammonium, and sulfate in the decomposition 

reaction depends on the stoichiometry of the reactive part 

of the organic matter. 

The concentrations of carbon, ammonium, and 

sulfate in the pore water and the concentrations of carbon and nitrogen in the solid phases contain a record of 

the decomposition reactions in the sediment. The present 

investigation used both pore-water and solid-phase data 

to determine the stoichiometry of decomposition reactions. First, the reaction stoichiometry was determined 

by a steady-state analysis of pore-water concentrations. 

Two reaction pathways, sulfate reduction and methanogenesis, were considered. Next, a mass balance was calculated for carbon, ammonium, and sulfate at one station. The calculations estimated the importance of 

nonsteady-state conditions. The importance of 

ammonium adsorption also was estimated. Finally, the 

reaction stoichiometry of organic decomposition was 

determined by a second method which used the concentrations of carbon and nitrogen in the solid phases of the 

sediment. 

SETTING 

The data were obtained from stations V3 and V6 

in the lower part of the Potomac River estuary (fig. 1). 

Samples were collected at both of the stations in May 1978 

and September 1978 and from station V3 in March and 

August 1979. Detailed data for these collections are 

available in Goodwin and others (1984). 

Station V3 is 9 km from the mouth of the estuary 

and is situated in 10m of water. The salinity of the overlying water ranges from 10 to 15 parts per thousand 
0 10 15MILE3 

~----~~~----~L-------~ 

0 10 15 KILOMETERS 

Figure ·1. Location of sampling stations V3 and V6 in the Potomac River estuary. 

(Blanchard and Hahl, 1981), and the overlying water was 

oxygenatc~d during all the samplings, even in mid-summer. 

A brown· flocculent layer was present on the surface of 

the sediment at each sampling. Station V 6 is 29 km from 

the mouth and is in 25 m of water. The salinity of the 

overlying water is slightly higher than at station V3 but 

has nearly the same range. During the summer samplings, 

H2S was present in the overlying water and no brown 

layer was observed on the surface of the sediment. 

Organic decomposition is the dominant reaction in 

the sediment at these stations. Sulfate concentrations 

decrease with depth in the sediment, and alkalinity and 

ammonium increase with depth, indicating decomposition 

of organic matter by the pathway of sulfate reduction. 

After complete depletion of sulfate, bubble formation and 

high ammonium ion concentrations suggest that decomposition 1;;ontinues at depth at station V 6 by the pathway 

of methanogenesis. 

The concentration profiles at stations V3 and V6 

vary seasonally. Steep gradients, indicating rapid reaction, 

occur in the summer, and less steep gradients occur in the 

winter and early spring. At station V3, during summer 

sampling (but not in spring), a minimum in sulfate and 

a maximum in alkalinity and ammonium occur within 10 

em of the sediment-water interface. These concentration 

changes with time are obvious nonsteady-state characteristics, and must be considered in an analysis of the stoichiometry of the reactions. 

SAMPLING AND ANALYTICAL METHODS 

Sediment cores for this study were obtained with 

a Benthos• gravity coring device. The sediment samples 

1

Use of brand names is for descriptive purposes only and does 

not constitute endorsement by the U.S. Geological Survey. 

Organic Decomposition in Sediments, Potomac River Estuary 63 
were extruded from the core liner under a nitrogen 

atmosphere in a glove box. All further processing took 

place under a nitrogen atmosphere. Either 2- or 2.5-cm 

sections of sediment were placed into nylon squeezers 

(Reeburgh, 1967), and the pore water was expressed by 

the pressure of nitrogen gas against a sheet of dental dam 

rubber. The pore water was collected in 30-mL syringes. 

The remaining sediment was freeze-dried for later analysis. The dissolved constituents in the pore water were 

analyzed within 12 h after collecting the samples. 

Ammonium was determined colorimetrically using 

the phenolhypochlorite method of Berthelot as described 

by Solorzano (1969). The ion-exchange-titrimetric method 

of DoHman (1968) was used to determine sulfate. The 

alkalinity was determined by a single addition of acid and 

pH measurement, followed by a correction to account 

for the excess acid added (Almgreen and Fonselius, 1976). 

The carbon of the solid phases was analyzed using 

a Model DC-12 Leco Carbon Determinator. Preliminary 

acidification of the samples to remove inorganic carbon 

indicated that inorganic carbon was not a significant fraction of the total carbon present in the samples. Therefore, 

the total carbon determination was taken as a measure 

of the organic carbon of the samples. Nitrogen concentrations in the solid phases were determined by Kjeldahl 

digestions (Bremner, 1965). The dissolved and adsorbed 

ammonia concentrations were found to be small, and the 

total value was used as an estimate of the organic nitrogen 

concentration. 

STOICHIOMETRY DETERMINED FROM 

PORE-WATER CONCENTRATIONS 

Berner (1975) gave the following diagenetic equation for sediments: 

Dr a2ci aci ---·---w-- + --- R;' 

1 + K; ax2 ax 1 + K; 

(1) 

where C; is the concentration in moles per cubic centimeter of pore water for species i (mollcm3); DJ is the 

sediment diffusion coefficient in sediment pore water for 

species i, which includes the effects of tortuosity but not 

of adsorption, in centimeters squared per second 

(cm2

/s); R; is the net rate of production (or consumption) of species i, in moles per cubic centimeters per 

second (mol· cm-3s-1

); x is depth from the sedimentwater interface, measured positively downward, in centimeters (em); tis time, in seconds (s); w is the velocity 

of burial of particles relative to the sediment-water interface, in centimeters per second (cm/s); and K is a dimensionless coefficient accounting for adsorption. This equation ignores the effects of compaction and assumes 

constant diffusion coefficients. Further, Berner (1975) 

64 Studies in Diagenesis 

stated that if the pore-water concentrations are at steady 

state and the sedimentation term is small, equation 1 

reduces to 

a2C; DJ -- = R;. (2) 

ax2 

Note that the factor accounting for adsorption, 

1/(1 + K), cancels out under these assumptions. For the 

reaction terms Rc (carbon species) and RN (nitrogen 

species), there is no evidence of any appreciable chemical 

reaction other than organic decomposition (the adsorption of ammonium ion is discussed below). Thus, Rc 

and RN are assumed to depend only on th~ rate of 

decomposition, and the rate of carbon release Is related 

to the rate of nitrogen release by the stoichiometric ratio, 

Lc:N' as follows: 

(3) 

During sulfate reduction the predominant carbon 

species released to the pore water are H2C03 and HC03, 

and the predominant nitrogen species is NH~. For the 

pH range of these pore waters, co;- and NH3 concentrations are negligibly small. Assuming a simple carbohydrate as the source of carbon with nitrogen present 

in a C:N ratio of 6:1, the reaction can be written as 

follows: 

(CH20)6NH3 + 3SO!- - H2C03 

+ 5HC03 + 2HS- + H2S + NH~. (4) 

For sulfate reduction, equation 3 can be represented as 

(5) 

The chemical equation for methanogenesis can be written as follows: 

(CH20)6NH3 + 3H20 - 3CH4 

+ 2H2C03 + HC03 + NH~ . (6) 

The molar amount of CH4 produced equals the sum of 

H2C03 and HC03. Thus, for methanogenesis, 

Lc·NRNH+ = RH co + RHco- + RcH . 4 2 3 3 4 

= 2{RH co + RHco-)· 2 3 3 

Combining equation 5 with equation 2 gives: 

a2C a2CHCO- D! H2C03 + D! - 3 

H2co3 a r HC03 a X2 

= Lc:AH~ a2 CNHt 

ax2 

(7) 

(8) 
For methanogenesis, the left-hand side of the equation 

has the additional factor of 2 from equation 7. The diffusion coefficients of aqueous species in the pore water 

of sedim(~nt generally are assumed to be related to the 

tracer diffusion coefficients by a common factor, which 

accounts for tortuosity and viscosity (Li and Gregory, 

1974): 

k!YJ = DJ. (9) 

Substitutilng for the sediment diffusion coefficients, .a., • • I 

m equation 8, the factor k cancels throughout leaving the 

tracer diffusion coefficients in place of the sediment diffusion coefficient. Equation 8 is a function only of depth, 

and the partial derivatives can be replaced with ordinary 

derivativ(~S. Assuming the diffusion coefficients are constant, equation 8 can be integrated giving: 

where A is a constant of integration. The value of A must 

be zero because the gradients approach zero at infinite 

depth. Integrating a second time gives the following relationship for concentrations: 

(.IJll.. co )( CH co ) + <.IJll..co-)( CHco-) 2 3 2 3 3 3 (11) 

where B is also a constant. 

Berner (1977) presented a method for determining 

the stoichiometry of sulfate-reduction organic decomposition. His method assumed steady-state concentration 

profiles and included the effects of sedimentation; 

however, it required that the data fit an exponential curve. 

Equation 11 is derived assuming steady state, ignoring 

the effects of sedimentation, but does not assume any 

particular mathematical form for the concentration 

profiles. 

Equation 11 implies that, given the preceding 

assumptions, the C:N ratio (Lc:N) can be determined 

from the pore-water concentrations. A plot of the quantity .IJll.. co CH co + <.IJll..co-)(CHco-) versus (l~m+ )(CNH+) 23 23 3 3 4 4 

should produce a linear relationship with a slope of 

Lc:N" By the same reasoning, a plot of .IJll.. co CH co 2 3 2 3 

+ (.IJll..c0 -)(CHco-) versus (Dg0 2- ) (C802-) can be used to .3 3 4 4 

determine the carbon to sulfate ratio, Lc:s· 

Figure 2 shows the carbon-nitrogen plot for cores 

from stations V3 and V6. The distinct change in slope 

corresponds to the depletion of sulfate in these cores. All 

points on the greater sloping line have measureable 

amounts of sulfate ion. All points on the lesser sloping 

line are at or below the limit of detection for sulfate. The 

sulfate ion is assumed to be the electron acceptor (sulfate 

reduction) along the former line, and organic carbon is 

assumed to be the electron acceptor (methanogenesis) on 

the latter line. 

For sulfate reduction the slope of the line is the C:N 

ratio of the reacting organic matter. For methanogenesis, 

equation 7 indicates that the slope of the plot must be 

multiplied by 2 to obtain the C:N ratio. The resultant C:N 

ratio calculated for sulfate reduction is 7. 7: 1; that 

calculated for methanogenesis is 7.2: 1. These values are 

close to one another and indicate that the net stoichiometry of the organic matter used during decomposition 

via both sulfate reduction and methanogenesis is similar. 

These values for the reactive part of the organic matter 

are lower than the average C:N ratio of 9.4:1 for the 

organic matter at these sites, suggesting a selective utilization of low-C:N organic matter. 

Figure 3 is a plot of the data from stations V3 and 

V6 for sulfate ion and carbon. The sloping linear relation indicates a C:S ratio of -2.1: 1, a value that is consistent with the theoretical value of -2.0:1 for the 

stoichiometry of sulfate reduction. Note that carbon continues to enter solution at low values of sulfate due to 

continued methanogenic decomposition following the 

depletion of sulfate from the pore water. 

MASS BALANCE CALCULATIONS 

The preceding analysis was based on the assumption of steady-state profiles for the dissolved constituents. 

For nonsteady-state profiles, the alkalinity and amount 

of ammonium ion produced during organic decomposition can (1) be transported out of the sediments (flux, 

F), (2) be buried below the zone of organic decomposition (burial, B), or (3) remain in the pore water resulting 

in a change in concentration (storage, S). The importance 

of the steady-state assumption can be investigated by 

comparing the masses involved in each of these three 

processes. For a given time period, mass balance 

stipulates: 

R = -F + S + B, (12) 

where R is the mass of a species released to solution during the period; F is the mass movement into or out of 

the sediment due to diffusion during the period (a 

negative value for the quantity F represents diffusion out 

of the sediment); S is the mass of a species accumulating 

in the pore water due to concentration changes during 

the period (this term excludes the effects of sedimentation); and B is the mass of a species accumulating in the 

pore water due to sedimentation during the period. All 

Organic Decomposition in Sediments, Potomac River Estuary 65 
LO 

0 

X 

.---., 

~ 

0 

u 75 01 

I 

'-"' 

u 

X 

r---. ('I) 

8 

01 

6 50 

0 

+ 

,.-..... 

('I) 

0 

u 

I 

'-"' 

u 

X 25 ,.-.... 

('I) 

0 

u 

I 

'-"' 

0 

1...--.J 

0 7809 V6A 

* 7809 V6B 

v 7809 V6C 

+ 7809 V3A 

A 7809 V3B 

~ 7809V3C 

o~~~--~~~--~~~--~~~--~~----~------------

0 5 10 15 20 

D(NH4) X C(NH4) X 105 

Figure 2. Pore-water data used to determine C:N ratio of organic decomposition. For sulfate reduction, C:N ratio is equal 

to slope of line. For methanogenesis, slope must be multiplied by 2 to account for production of reduced carbon (methane) 

in addition to oxidized carbon. Symbols refer to numbered core samples from stations V3 and V6 in the Potomac River 

estuary (fig. 1). 

terms are in moles per square centimeter (mol/cm2). 

Equation 12 can also be applied to the reduction of sulfate 

ion during decomposition; a negative value for R 

represents the amount of sulfate ion reduced. The concentration gradients presented in figure 4 indicate that 

ammonium ions and alkalinity should be diffusing out 

of the sediment and that sulfate ions should be diffusing 

into the sediments. 

Remember that alkalinity does not represent a 

single species. At these two stations, bisulfide, as 

estimated from total sulfide measurements, accounts for 

less than 1 meq/L of the alkalinity. Other species contributing to the alkalinity, CO/~ HP04

2

-, PO~, also 

are present only in low concentrations. Bicarbonate ion 

is by far the principal species contributing to the alkalinity; its concentration was assumed to be equal to the 

alkalinity. 

66 Studies in Diagenesis 

Of the terms (equation 12) flux, storage, and burial, 

storage is the most difficult to estimate. Its estimation 

requires concentration profiles at different times of year 

in the same location taken from cores that have identical 

reaction histories except for the period from one sampling to the next. We have assumed that only the top 20-

30 em are affected by seasonal variation in reaction rates 

and that the concentrations at depth are an indicator of 

past reactions. Thus, the same sequence of reaction was 

assumed in cores with equal concentrations at depths of 

40-60 em. Cores from station V3 have concentrations at 

depth which are nearly equal, whereas concentrations at 

depth at station V6 vary widely. The differences at station V6 have not been attributed to any substantial difference in processes, only to spatial variation in sedimentation, microbial abundance, and extent of reaction. 

The storage term, and consequently the importance of 
Lt) 

0 

X 

;---, 

'(Y) 

8 75 

0. 

I 

"-' 

u 

X 

,....... ("t) 

0 

u 

0. 

I 50 

"-' 

0 

+ 

,....... ("t) 

0 

u 

I 

(/) 

·v; 

~ 

c 8 ~ 

01 II 0 (./') c 

1'0 

£. u ...., 

~ 

D 

v 

* v D 

CJ 7809 V6A 

.. 7809 V6B 

v 7809 V6C 

+ 7809 V3A 

A 7809 V3B 

<> 7903 V3B 

"-' 

u 

X 

,....... 

("t) 

25 

C:sSulfat --2.1 

e reduct· 

I On 0 

u 

I 

"-' 

0 

L...-...J 

CJ v 

* 

o~~~~~~~~~~~~~-L~~~~~~~~~~~--~~ 

0.0 2.5 5.0 7.5 10.0 12.5 

D(S04) x C(S04)x 105 

Figure 3. Pore-water data used to determine C:S ratio of organic decomposition via sulfate reduction and methanogenesis. 

Symbols refer to numbered core samples from stations V3 and V6 in the Potomac River estuary (fig. 1). 

nonsteady state, could be calculated at station V3 and 

not at station V6. 

The reaction, flux, storage, and burial terms were 

estimated for station V3 for the period May to September 

1978. The steepest gradients occurred during this period 

of the year indicating that organic decomposition was 

most rapid at this time. This reasoning is consistent with 

the obs(~rvation of Nedwell and Floodgate (1972a, 1972b) 

that sulfate reduction is temperature dependent. 

The flux terms were estimated to be the average of 

fluxes calculated from the May and September pore-water 

data. The following equation (Berner, 1980, p. 38) was 

used to estimate the flux of a constituent, J;, at the sediment water interface: 

J. = c/>VS. _I, (

a c.) I I ax (13) 

where ct) is the porosity, Ci is the concentration of species 

i, and D; the diffusion coefficient of species i in sediment pore water. 

The following relation (Andrews and Bennett, 1981) 

was used to estimate D;: 

D~ D; = _ 1 , (14) 

c/>J 

where fY! is the tracer diffusion coefficient and f is the 

form factor. Andrews and Bennett (1981) found that 

f = +-3 was a good approximation of the form factor in 

fine-grained muds. The porosity in the upper 2 em of 

cores from station V3 was determined to be 0.91. The 

values of .D'fwere obtained from Li and Gregory (1974) 

at 25 OC, and the gradient was calculated linearly, using 

the first two intervals from core 7805 V3 and core 7809 

V3B (fig. 4). The linear approximation probably 

underestimates the gradient at the sediment-water interface. The flux term, F, was calculated by multiplying the 

flux rate, J, by the time interval betweem samplings. 

Organic Decomposition in Sediments, Potomac River Estuary 67 
10 

~ 20 

w 

tl:i 

~ 30 

z 

w 

u 

~ 40 

I' 

b: 

~ 50 

10 10 

20 20 

30 30 

40 40 

50 50 

60 60 

70o.o 2 .5 5.0 7.5 10.0 70o 5 10 15 20 70o.o 0 .5 1 . 0 1.5 

Ml LLIMOLES PER LITER MILLIEQUIVALENTS PER LITER Ml LLIMOLES PER LITER 

Figure 4. Concentration profiles of sulfate, alkalinity, and ammonium in bottom sediments at station V3 for May 1978 

(plus sign) and September 1978 (open triangle). Shaded areas represent storage terms. 

Few benthic organisms were observed at station V3 

during the present study. The scarcity of benthic macrofauna and the steep concentrations gradients near the surface of the sediment suggest that bioturbation is not an 

important process here. This conclusion is supported by 

the work of Callender and Hammond (1982) who showed 

that the diffusive flux calculations for NH: at station 

V3 are, at most, a factor of two lower than in-place 

measurements. These results indicate that if the estimate 

of the flux term is in error, it is probably an underestimate. The actual flux is likely to be even more important relative to storage and burial (see tables 1 and 2). 

Storage was estimated from the differences in the 

concentration profiles from May and September over the 

top 30 em. The differences were calculated by assuming 

that concentration changes betweem measured points 

were linear. On figure 4, the shaded areas represent the 

storage terms. 

Burial for the same period was estimated as the difference betweem the mass of a species in the pore water 

that is associated with the sediment accumulating at the 

surface of the sediment column and the mass of the 

species that is in the pore water that is buried below 30 

em, as follows: 

where csurf and c30cm are the concentrations of a species 

in the overlying water at the surface of the sediments and 

in the pore water at 30 em, in millimoles per cubic 

68 Studies in Diagenesis 

centimeter (mmoll cm3); w is the sedimentation rate, in 

centimeters per year (cm/yr); cf>surr and c/>30cm are the 

porosities at the surface of the sediment and at 30 em; 

and !::t.t is the time interval, in years (yr). The sedimentation rate in the vicinity of this station is 0.16 cm/yr 

(Knebel and others, 1981). 

The reaction terms (net consumption or production 

of a species) were obtained as the sum of the other terms 

as given by equation 12. Thble 1 presents the reaction term 

as the total mass reacting in the 30-cm sediment column 

over the entire period. The rate of reaction per day is presented in table 2. The ratio of alkalinity produced to sulfate 

reduced is -2.1:1, very close to the -2:1 ratio commonly 

assumed for sulfate reduction (Berner, 1979). The ratio 

of alkalinity produced to ammonium produced is 7 .5:1. 

Burial constitutes an insignificant fraction of the 

total reaction. The total storage terms range between 10 

and 19 percent of the total reaction terms. More than 80 

percent of the alkalinity and ammonium released to the 

pore water during organic decomposition is then released 

to the overlying water column by diffusion. Similarly, 

more than 80 percent of the sulfate reduced is supplied 

by the overlying water. 

The possibility of adsorption of ammonium ion 

onto the solid phases has been excluded from these 

calculations. Measurements of adsorbed ammonium 

(Nancy Simon, U.S. Geological Survey, National Center, 

MS 432, Reston, VA 22092, written commun., 1982) show 

a maximum concentration of 6 ~-tmollg (micromoles per 

gram) in sediment from station V3. At a sediment 
Table 1. Mass accounting for the period May to September 1978 for station V3 

[Units: microequivalents per square centimeter (p.eq/cm2) and micromoles per square centimeter (p.mollcm2)] 

Constituent Flux Storage Burial Reaction 

( -£) (_~) (~_) (~) 

Alkalinity, in ~eq/cm2 •••••••••••••• 320 53 0.21 373 

Ammonium, in ~mol/cm2 ••••••••••••••• 44 5 0.004 49 

Sulfate, in ~mol/cm 2 •••••••••••••••• -146 -33 0.15 -179 

Table 2. Rates of processes calculated from the period May to September 1978 for station V3 

[Units: microequivalents per square centimenter per day (p.eq/cm2-day) and micromoles per square centimenter 

per day (p.mol/cm2 -day)] 

Constituent 

Alkalinity, in ~eq/cm2-day ••••••••••••• 

Ammonium, in ~mol/cm2-day •••••••••••••• 

Sulfate, in ~mol/cm ~day ••••••••••••••• 

accumulation rate of 0.15 g/cm2-year (grams per square 

centimeter per year) (Ann Martin, U.S. Geological Survey, 

Box 25046, DFC, Denver, CO 80225, written commun., 

1981), the rate of adsorption is 6(0.15)/365 = 2.5 X 1Q-3 

J,tmollcm2-day. Comparison with table 2 shows that ammonium adsorption accounts for less than 1 percent of the 

total amount of ammonium released by decomposition. 

These calculations provide some insight into the 

importance of several sedimentary processes. The concentration profiles are obviously not at a steady state; however, 

they imply that storage could be an important process. The 

preceding calculations demonstrate that storage accounts 

for 20 percent or less of the organic decomposition reaction. In these sediments the rates of burial and adsorption are insignificant relative to the rate of organic decomposition. An adequate accounting of the rates of reaction 

can be obtained if the fluxes across the sediment-water 

interface are known through a yearly cycle. These calculations also imply that the assumptions made in deriving the 

equation to calculate the stoichiometry of organic decomposition (that is, the assumptions leading to equation 2) 

are reasonable approximations even with these nonsteadystate conditions. 

STOICHIOMETRY DETERMINED FROM 

CONCENTRATIONS IN THE SOLID PHASES 

The concentrations of organic carbon and nitrogen 

in the solid phases at stations V3 an4 V6 systematically 

Flux Storage Burial Reaction 

( -£) (~ (~ (~ 

2.41 0.40 0.002 2.81 

0.33 0.04 3x1o-5 0.37 

-1.10 -0.25 0.001 -1.35 

decrease with depth (fig. 5). If the sedimentation rate and 

the sediment composition have been constant, these 

decreases reflect the amount of organic decomposition. 

Profiles of 210Pb (Ann Martin, U.S. Geological Survey, 

Box 25046, DFC, Denver, CO 80225, written commun., 

1981) can be used to corroborate the assumption of steadystate sediment deposition. For the area near station V3, 

the 210Pb profiles have a smooth exponential decrease as 

expected from steady-state sedimentation, and a consistent 

sedimentation rate of 0.16 cm/yr was found. In the area 

of station V 6, the sedimentation rate determined from 210Pb was less well defined, but a trend indicating a rate 

of 2 cm/yr was found (Knebel and others, 1981). Carbon 

and nitrogen in the solid phases of cores collected at station V6 are more erratic than at station V3, but a general 

decrease in concentration with depth was observed. The 210Pb data suggest that the approximation of steady-state 

sedimentation can be used at stations V3 and V 6. 

The flux of carbon and nitrogen out of the sediment 

can be calculated from the concentrations in the solid 

phases. The change in concentration of carbon (or 

nitrogen) multiplied by the sediment accumulation rate 

(g/cm2-yr) is an estimate of the flux rate. The fluxes 

calculated from the solids in this manner are compared in 

table 3 with fluxes calculated from pore-water concentration gradients. 

The fluxes calculated from the solid-phase data are 

lower than those from the pore-water data. The calculations with the solids are expected to represent long-term 

Organic Decomposition in Sediments, Potomac River Estuary 69 
10 10 

20 20 

(/) 

C! 

UJ 

tu 

~ 30 

1-- V3 30 

z 

UJ 

u + Carbon 

z A Nitrogen x 1 0 -.. 40 40 I 

b: V6 

UJ 

0 +Carbon 

50 50 A Nitrogen x 10 

60 60 

70--~~~~~--~~----~~ 7 0 ~-L--1---L--L-L~...__._--L..-JL--1--'---4-'-"'--'---"--' 

0 1 2 3 4 0 1 2 3 4 

MILLIMOLES PER GRAM MILLIMOLES PER GRAM 

Figure 5. Concentration profiles of carbon and nitrogen in the solid phase for stations V3 (May 1978) and V6 (May 1978). 

Table 3. Flux rates calculated from analyses of pore water and solid phases 

[Units: micromoles per square centimeter per day (J.Lmol/cm2-day)] 

Flux calculated from Flux calculated 

Station Element 

pore-water data 

( llmo 1 I cm2-day) 

from solid phase 

( llffiO 1 I cm2-day) 

May 

V3 Carbon •••••••••• 1.31 

Nitrogen •••• ~ ••• 0.25 

V6 Carbon •••• ~ ••••• 4.52 

Nitrogen •••••••• 0.55 

(tens of years) average fluxes. These fluxes should be lower 

than those calculated from the pore water, which reflect 

reactions during the late spring and summer, the period 

with the most rapid rates of decomposition. The large 

fluxes calculated from the pore-water data indicate that 

the rate of decomposition is fast enough to produce the 

observed concentration changes in the solids. 

70 Studies in Diagenesis 

August 

3.38 0.63 

0.42 0.09 

6.38 1.37 

1. 37 0.23 

An alternative hypothesis is that the high carbon 

and nitrogen concentrations near the sediment-water interface are due to a historical increase in the organic fraction of the sediment. If this were true, station V6 should 

to have high carbon and nitrogen concentrations to a 

greater depth than station V3 because the sedimentation 

rate is nearly 10 times greater at station V6 than at station 
V3. In fact, the carbon and nitrogen at station V6 decrease 

over a depth range similar to that at station V3 as shown 

on figun~ 5. These data suggest that organic decomposition is responsible for the concentration decreases in the 

solid phases with depth, not historical changes in the composition of the sediment. 

A plot of organic carbon versus nitrogen in the solid 

phases should demonstrate the loss of carbon and nitrogen 

from th(: sediment solids, if the steady-state assumptions 

apply. The slope of the plot should be the ratio in which 

carbon is lost relative to nitrogen, that is, the C:N ratio 

of the reaction. Figure 6 is a plot of such data from stations V3 and V6. There is some scatter in the points, but 

the data points for each individual core follow a trend from 

the upper right of the diagram (for samples near the 

sediment-water interface) to the lower left on the diagram 

(for samples at depth). The best-fit line for all the data 

points has a slope of 5.9. 

The C:N ratio of 5.9:1 is in reasonable agreement 

with the values derived from the pore water concentrations, 

7.2:1 and 7.7:1. The intercept of the line on the plot of solid 

carbon and nitrogen concentrations (fig. 6) is not zero, 

demonstrating that the stoichiometry of the organic decomposition reactions is not the same as the composition of the 

bulk organic matter. The average C:N ratio of the solids, 

9.6:1, is higher than the estimates of the C:N ratios of the 

reactive part of the organic matter. The data presented 

indicate that part of the organic matter, with a C:N ratio 

in the range of 5.5:1 to 8:1, is the reactive fraction of the 

organic matter in the Potomac River estuary. The more 

refractory organic matter must have a greater C:N ratio. 

Spiker (1981) and Hunt (1968) concluded on the basis 

of o13C measurements that marine plankton dominate the 

organic matter in the lower portion of the Potomac River 

estuary, although Spiker noted a winter input of terrigenous 

material. Lyons and Gaudette (1979) suggested that marine 

organic matter in the estuarine sediments of Great Bay, 

New Hampshire, is more readily metabolized than the terrigenous organic matter. The data presented in the present 

study are consistent with the hypothesis that low-C:N 

marine plankton are the reactive fraction of the organic 

matter in the sediment near the mouth of the Potomac 

River estuary. Further work is necessary to determine the 

nature of the compounds actually decomposed. 

SUMMARY 

The decomposition of organic matter in sediment 

was investigated at two stations near the mouth of the . 

Potomac River estuary. Concentration data for pore water 

4.-r-r-r-.-.-.-.-.-~,-,-.-,-,-~-r-r~-r~~-T-T-T~ 

~ 

<( 3 

a: ~~ + 

(!) ~ .. a: 

w c,··~ <) c.. 

en 

w + 

....J 

0 

~ 2 <) 

+ :::J 

....J .. ~ 

z .. z 

0 .. + 7805 V3 al 

a: 1 <( .. 7908 V3 

u <) 7809 V3B 

A 7805 V6 

NITROGEN, IN MILLIMOLES PER GRAM 

Figure 6. Concentration of carbon versus concentration of nitrogen in solid phases for four cores from stations V3 and V6 in the Potomac River estuary (fig. 1). Slope of line through points is an estimate of C:N 

ratio of the reacting organic matter. Symbols refer to numbered core samples. 

Organic Decomposition in Sediments, Potomac River Estuary 71 
and solid phases of the sediment were used to determine 

the stoichiometry of decomposition reactions, and mass 

balance calculations were made to estimate the importance 

of storage, burial, and ammonium adsorption. 

The pore-water concentrations of alkalinity and ammonium were used to determine the stoichiometry of decomposition reactions. Assuming steady-state conditions, 

equations were derived to determine the C:N ratio of the 

reactive part of the organic matter. The C:N ratio was 7.7:1 

for decomposition of organic matter by sulfate reduction 

and 7.2:1 for decomposition by methanogenesis. By a 

similar analysis, the C:S ratio was -2.1:1 for sulfate reduction. This ratio is close to the theoretical value of -2.0:1. 

A mass-balance accounting was made for carbon, 

ammonium, and sulfate at one station for May 1978 to 

September 1978. Between 10 and 20 percent of the carbon and ammonium produced by decomposition of 

organic matter was retained in the sediment pore water. 

More than 80 percent of the carbon and ammonium produced by decomposition was released from the sediment 

pore water to the overlying water. The decrease in sulfate 

concentration in the pore water during this period 

accounted for 19 percent of the sulfate consumed in 

organic-decomposition reactions. About 80 percent of the 

sulfate consumed in organic-decomposition reactions was 

supplied by the overlying water. 

:rhe rate that carbon, ammonium, and sulfate 

accumulated in the pore water due to burial was insignificant compared with the rate of decomposition of organic 

matter. The rate that ammonium was adsorbed onto solid 

phases of the sediment was less than 1 percent of the rate 

that ammonium was released by the decomposition of 

organic matter. 

Calculations indicated that storage, a nonsteadystate effect, burial, and ammonium adsorption accounted 

for less than 20 percent of the total rate of organic decomposition. The calculations support the assumptions made 

in the derivation of the equations used to determine the 

stoichiometry of decomposition from pore-water data. 

The stoichiometry of organic decomposition reactions was estimated using the concentrations of carbon 

and nitrogen in the solid phases of the sediment. The 

estimate of the C:N ratio of the reactive organic matter 

was 5.9:1. The organic-decomposition pathways of sulfate 

reduction and methanogenesis were not distinguishable 

using these data. The data presented indicate that part 

of the organic matter, having a C:N ratio in the range 5.5:1 

to 8:1, is the reactive fraction of the organic matter in the 

Potomac River estuary. The total C:N ratio of the organic 

matter is 9.6:1. The more refractory organic matter must 

have a greater C:N ratio. 

REFERENCES CITED 

Almgreen, T., and Fonselius S. H., 1976, Determination of 

alkalinity and total carbonate, in Grasshoff, Klaus, ed., 

72 Studies in Diagenesis 

Methods of seawater analysis: New York, Verlag Chemie, 

p. 97-115. 

Andrews, D., and Bennett, A., 1981, Measurements of diffusivity 

near the sediment water interface with a fine scale resistivity 

probe: Geochimica et Cosmochimica Acta, v. 45, 

p. 2169-2175. 

Berner, R. A., 1975, Diagenetic models of dissolved species in 

the interstitial waters of compacting sediments: American 

Journal of Science, v. 275, p. 88-96. 

___ 1977, Stoichiometric models for nutrient regeneration in 

anoxic sediments: Limnology and Oceanography, v. 22, 

p. 781-786. 

___ 1979, A new look at biogenous material in deep sea 

sediments: Ambio Special Report, v. 6, p. 5-10. 

___ 1980, Early diagenesis-A theoretical approach: 

Princeton, N. J., Princeton University Press, 241 p. 

Blanchard, S. F., and Hahl, D. C., 1981, Water quality of the 

tidal Potomac River and estuary hydrologic data report, 

1979 water year: U.S. Geological Survey Open-File Report 

81-1074, 149 p. 

Bremner, J. M., 1965, Methods of soil analysis-Part 2: 

Madison, Wisconsin, American Society of Agronomy, 

p. 1149-1176. 

Callender, Edward, and Hammond, D. E., 1982, Nutrient 

exchange across the sediment-water interface in the Potomac 

River estuary: Estuarine, Coastal and Shelf Science, v. 15, 

p. 395-413. 

DoHman, G. W., 1968, Determination of sulfate and phosphate 

in water by ion exchange-titrimetric method: Environmental Science and Technology, v. 2, p. 1027-1029. 

Goodwin, S.D., Schultz, B. 1., Parkhurst, D. L., Simon, N. S., 

and Callender, Edward, 1984, Methods for the collection 

of geochemical data from the sediments of the tidal 

Potomac River and estuary, and data for 1978-1980: U.S. 

Geological Survey Open-File Report 84-074, 129 p. 

Gordon, D. C., 1971, Distribution of particulate organic carbon and nitrogen at an ocean station in the Central Pacific: 

Deep-Sea Research, v.18, p. 1127-1134. 

Holm-Hansen, 0., Strickland, J. D., and Williams, P. M., 1966, 

A detailed analysis of biologically important substances in 

a profile off Southern California: Limnology and 

Oceanography, v: 11, p. 548-561. 

Hunt, J. M., 1968, The significance of carbon isotope variations 

in marine sediments, in Hobson, G. D., ed., Carbon in 

biochemistry: London, Pergamon Press, p. 27-35. 

Kamatani, A., 1969, Regeneration of inorganic nutrients from 

diatom decomposition: Journal of the Oceanographic 

Society of Japan, v. 25, p. 63-74. 

Knebel, H. J., Martin, E. A., Glen, J. L., and Needel, S. W., 

1981, Sedimentary framework of the Potomac River estuary, 

Maryland: Geological Society of America Bulletin, pt. I, 

v. 92, p. 578-589. 

Li, Y.-H., and Gregory, S., 1974, Diffusion of ions in sea water 

and in deep-sea sediments: Geochimica et Cosmochimica 

Acta, v. 38, p. 703-714. 

Lyons, W. B., and Gaudette, H. E., 1979, Sulfate reduction and 

the nature of organic matter in estuarine sediments: Organic 

Geochemistry, v. 1, p. 151-155. 

Nedwell, D. B., and Floodgate, G. D., 1972a, Thmperatureinduced changes in formation of sulphide in a marine sediment: Marine Biology, v. 14, p. 18-24. 
___ 1972b, The effects of microbial activity upon the 

sedimentary sulphur cycle: Marine Biology, v. 16, p. 

192--200. 

Pockington, R., 1976, Terrigenous organic matter in surface 

sediments from the Gulf of St. Lawrence: Journal of 

Fisheries Research Board of Canada, v. 33, p. 93-97. 

Redfield, A. C., 1958, The biological control of chemical factors in the environment: American Scientist, p. 205-221. 

Reeburgh, W. S., 1967, An improved interstitial water sampler: 

Limnology and Oceanography, v. 12, p. 163-165. 

Solorzano, L., 1969, Determination of ammonia in ·natural 

waters by phenol-hypochlorite method: Limnology and 

Oceanography, v. 14, p. 799-801. 

Spiker, E. C., 1981, Carbon isotopes as indicators of the source 

and fate of carbon in rivers and estuaries, in Carbon dioxide 

effects research and assessment program, CONF-8009140, 

Report of a workshop, Woods Hole, Massachusetts, September 21-25, 1980, Flux of organic carbon by rivers to 

oceans: Washington, D.C., U.S. Department of Energy, 

p. 73-108. 

Organic Decomposition in Sediments, Potomac River Estuary 73 

Relation of Hydrocarbon Type to Maturity of Organic 

Matter in Upper Cretaceous Chalks, Eastern Denver Basin 

By DUDLEY D. RICE, U.S. Geological Survey, Box 25046, Denver Federal Center, 

MS 934, Denver, Colorado 80225 

CONTENTS 

Abstract 76 

Introduction 76 

Hydrocarbon reservoirs 76 

Gas geochemistry and origin 76 

Source-rock geochemistry 77 

Thermal maturity 80 

Summary and conclusions 81 

References cited 81 

FIGURES 

1. Index map of Denver basin showing contours on Precambrian 77 

2. Graph showing changes of o13C1 with depth of natural gases from Niobrara 

Formation in Colorado and Kansas 79 

3. Modified van Krevelen diagram showing type and maturity of kerogen from three 

wells 80 

TABLES 

1. Chemical and isotope composition of natural gas from selected wells 78 

2. Details of wells from which core samples were analyzed 79 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 75 
Abstract 

Oil and wet gas-condensate production has been established from chalk beds of the Upper Cretaceous Niobrara Formation and underlying Codell Sandstone Member of the Carlile 

Shale in the deeper part of the Denver basin. Dry gas has been 

produced from chalk reservoirs in the Niobrara on the shallow 

east flank of the basin. The shallow east-flank gases are chemically dry, isotopically light with respect to o13C1 values, and 

are probably of local biogenic origin. The Niobrara contains 

organic-rich laminations characterized by type II kerogen. This 

type of organic matter is capable of generating oil and wet gascondensate when subjected to thermal degradation and cracking processes. The thermal maturity pattern present in the 

Niobrara Formation appears to explain the occurrence of 

thermogenic hydrocarbons in the deeper part of the basin and 

biogenic gases on the shallow east flank. 

INTRODUCTION 

Oil and wet gas-condensate are produced from 

chalk beds of the Upper Cretaceous Niobrara Formation 

and underlying Codell Sandstone Member of the Carlile 

Shale in the deep Wattenberg area of the Denver basin 

in north-central Colorado (fig. 1). Dry natural gas has 

also been produced from chalk reservoirs of the Niobrara 

Formation in several shallow fields on the eastern flank 

of the basin in northeastern Colorado and northwestern 

Kansas. The east-flank gas fields trend northwest-southeast; present-day depths of burial increase from about 

250 to 850 m toward the northwest. The following data 

and discussion are summarized from a more complete 

report (Rice, 1984) and concern primarily the nature and 

origin of gas in this trend with respect to source rocks, 

kerogen types, and maturity patterns present in the 

Niobrara. The results of this work, when combined with 

existing data on hydrocarbon type and thermal maturity 

from the deeper Wattenberg area of the basin (Clayton 

and Swetland, 1980; Tainter, 1982; Rice and others, 1984) 

are of value in predicting regional patterns of thermal 

maturity and of hydrocarbon distribution related to both 

thermogenic and biogenic processes of generation. 

HYDROCARBON RESERVOIRS 

The gas-bearing reservoirs producing in shallow 

fields on the east flank of the Denver basin are finegrained chalks consisting predominantly of calcareous 

nanofossils and microfossils. The chalks are characterized 

by high porosity (30-45 percent) and low permeability 

(average 1 millidarcy) at these shallow depths (Lockridge 

and Scholle, 1978). Core samples of wells from the productive east-flank area (fig. 1) average 30 percent insoluble residue which is concentrated with organic matter in 

alternating laminations. 

Gas-bearing Niobrara chalk reservoirs on the shallow east flank of the basin have apparently lost most of 

76 Studies in Diagenesis 

their original high matrix porosity because of depthrelated diagenetic processes (Lockridge and Scholle, 

1978). As a result, most productive oil and wet gascondensate reservoirs found in the deeper part of the 

basin are characterized by fracture-type porosity and 

permeability (Mallory, 1977; Pollastro and Scholle, 1984). 

The deep Wattenberg area of the basin in Colorado has 

areally extensive production from the Codell Sandstone 

Member of the Carlile Shale, a unit which immediately 

underlies the Niobrara. Productive Codell reservoirs 

generally consist of fine-grained shaley and calcareous 

sandstone and are characterized by low matrix porosity 

and permeability. Poor matrix reservoir characteristics 

have been enhanced locally by the formation of natural 

fracture systems (Petroleum Information, 1983). 

GAS GEOCHEMISTRY AND ORIGIN 

Gas samples were collected from seven wells along 

a southeast-northwest belt that parallels the productive 

trend of shallow gas fields as shown in figure 1. Presentday depths of burial to the top of the Niobrara chalk reservoirs increase to the northwest from 328 to 842 m. All 

the gases are uniformly dry; methane generally accounts 

for more than 98 percent of the hydrocarbon fraction 

(table 1). In addition, the methane component of the gases 

is enriched in the light isotope 12C, and o13C1 values are 

lighter than -55 per mil. As shown in figure 2, the most 

significant change in the gases from the Niobrara is that 

they become isotopically heavier with increasing depth of 

burial; o13C1 values range from -65 per mil at 328 m to 

-55 per mil at 842 m. 

The chemical and isotope composition of these shallow natural gases suggest that they are dominantly of 

microbiological or biogenic origin (Rice and Claypool, 

1981). Biogenic gas was generated during the immature 

stage of hydrocarbon generation prior to onset of major 

oil and gas generation by thermal cracking. The biogenic 

gas was generated by the breakdown of organic matter 

at low temperatures by microorganisms in an anaerobic, 

presumably sulfate-free, environment at low temperatures. 

In-place generation of shallow biogenic gas is suggested by the facts that the chalks (1) have low permeability that inhibits long-range migration, (2) are overlain by a 

thick section of shale (Pierre Shale) containing many bentonites that served as a seal for the gas, and (3) contain 

organic-rich laminae that provided ample source for gas. 

The isotopically light, dry gases characteristic of the 

shallow east-flank fields are in marked contrast to those 

encountered in oil and condensate production from reservoirs of Niobrara Formation and underlying Codell Sandstone Member in the Wattenberg area of the Denver basin 

in Colorado (Rice and others, 1984) (fig. 1). Gases from 

the Codell in this part of the basin become isotopically 

heavier (o13C1 values range from -47.8 to -43.5 per mil) 

as they become chemically drier (C/C1

_5 range from 0.76 
-500 

Wyoming Nebraska 

---c~l~d~l 

0 50 

KILOMETERS 

Figure 1. Index map of Denver basin showing structure contours on Precambrian. Contour interval is 500 

m. Stipple patterns indicate areas of hydrocarbon production from Niobrara chalks. Location of wells from 

which gas samples were analyzed are labeled by numbers (table 1) and from which core samples were 

analyzed are labeled by letters (table 2). 

to 0.8) at depths ranging from 2,082 to 2,265 m. These 

gases are probably of thermogenic origin and were formed 

during the mature stage of hydrocarbon generation. 

SOURCE-ROCK GEOCHEMISTRY 

The~ amount, type, and level of maturation of 

organic matter within the Niobrara chalks of the eastern 

Denver basin were evaluated by analyses of core samples 

from five wells in northeastern Colorado. Location of the 

wells is shown on figure 1; details of the wells and cored 

intervals are given in table 2. Organic carbon values in 

chalk beds of the Niobrara Formation range from 0.43 

to 5.8 percent and average 3.2 percent. The organic 

Hydrocarbon Type Versus Organic-Matter Maturity in Chalks, Denver Basin 77 
~ 

~ c 

Q. 

~· 

:r 

0 

~· 

:I 

II 

;;;· Table 1. Chemical and isotopic composition of natural gas from selected wells shown in figure 1 

[tr., trace; isotope composition in per mil] 

Map Pay zone N2 and 

No. Well name Location (meters) (or) air CH4 

Kansas: 

1. Mountain Petroleum 

Topliff No. 1-10-- NW lf4 SE 1f4 sec. 10, T. 9 S., R. 40 W. 322-328 5.18 93.91 

Colorado: 

2. Kansas-Nebraska 

-Natural Gas 

Engel No. 1-10------ sw114sw1/4sec. 10, T. 3 s •• R. 42 w. 478-482 6.83 91.18 

3. Mountain Petroleum 

Ekberg No. 2-34----- NW 1f4 NW 1f4 sec. 34, T. 2 S., R. 43 W. 481-491 5.08 92.9 

4. Mountain Petroleum 

State No. 1-16------ NE 1f4 SW 1f4 sec. 16, T. 2 S., R. 43 W. 513-518 5.46 92.53 

5. Kansas-Nebraska 

Natural Gas 

Allison No. 1-29---- SE 1f4 SE lf4 sec. 29, T. 1 S., R. 44 W. 642-647 5.72 92.1 

6. Kansas-Nebraska 

Natural Gas 

Fonte No. 1-27---- SW1f4NE1/4sec. 27, T. 1 N., R. 45 W. 688-695 6.6 91.4 

7. Kansas-Nebraska 

Natural Gas 

State No. 1--------- NE 1f4 NE 1f4 sec. 16, T. 3 N., R. 47 W. 838-842 5.12 92.85 

C02 C2H6 C3Hs i-c4H1o n-c4H1o t-csn12 n-CsH12 C1/C1-5 a13c1 

0.25 0.67 tr. tr. tr. tr. tr. 0.993 -65.4 

.13 1.24 0.39 0.06 0.06 0.02 0.01 .981 -62.5 

.29 1.16 .41 .07 .06 tr. tr. .982 -60.8 

.23 1.06 .48 .09 .08 .03 .02 .981 -60.1 

.24 1.07 .55 .13 .11 .04 .03 .979 -58.8 

.2 .91 .54 .12 .12 .05 .04 .981 -59.7 

.43 1.05 .34 .09 .07 .03 .02 .976 -54.7 
(/) 

L.. 

Q) 

+"' 

Q) 

E 

c 

::t 

1-

a.. 

w 

0 

-66 

300 1 

• 

400 

500 

600 

700 

' 

' ' ' 

-62 

' 

' ' ' ' 

2 

• 

' 

' ' 

' 

' 

-60 

' 3., 

4 .. 

' ' 

' ' ' 

' 6 

'~ 

' 

' 

-58 -56 

' ' 5 ' 

. ' ' ' ' ' ' ' ' ' ' ' ' ' ' ' ' 

' ' 

-54 

' ' 7 ' • 

-52 

Figure 2. Changes of o13C1 with depth of natural gases from Niobrara Formation in Colorado and 

Kansas. Dots and numbers correspond to wells shown in figure 1 and described in table 1. 

Table 2. Details of wells from which core samples were analyzed 

Maximum 

Top recorded 

of Pay Cored bottomhole 

Map Status Niobrara interval interval temperature 

letter Well name Location of well (meters) (meters) (meters) (degrees Celsius) 

A. Kansas-Nebraska NE 1/4 NE% sec. 32, Gas well--- 454 457-466 447-463 36 

Natural Gas T. 2 s.' R. 43 w. 

Whomble No. 1-32. 

B. Mountain Petroleum SW 1/4 NE 1;4 sec. 29, 

State No. 1-29. T. 2 s.' R. 43 w. Gas well--- 443 451-454 445-452 32 

c. Kansas-Nebraska 

Natural Gas 

Po·well No. 1-33. 

NE 1;4 SW 1;4 sec. 33, 

T. 3 N., R. 47 w. 

Dry and 

abandoned. 

853 853-871 47 

D. J-W Operating , SE lf4 SE 1/4 sec. 29, Gas well--- 836 839-842 842-863 46 

Brophy No. I". T. 4 N., R. 46 w. 

E. Excelsior Oil NW 1;4 SE lf4 sec. 14, Oil well 

Alice G. Nay No. 1. T. 9 N., R. 58 w. 

matter is concentrated with larger amounts of insoluble 

residue in the darker colored laminae. 

The hydrocarbon generation potential of organic 

carbon found in the Niobrara can be evaluated by 

pyrolysis using a ''Rock-Eval'' instrument (Espatalie and 

others, 1977). Both the type and level of maturation of 

1,646 1,681-1,687 1,655-1,667 74 

the organic matter can be determined by use of a modified 

van Krevelen diagram. Samples from three wells (A, C, 

and E, fig. 1) are plotted on a modified van Krevelen 

diagram shown in figure 3. These samples are typical of 

those from all five wells. All samples shown in figure 3 cluster along the type II path. Type II kerogen is 

Hydrocarbon Type Versus Organic-Matter Maturity in Chalks, Denver Basin 79 
900 

l 750 

rn 

c 

0 

..0 

I... 

cu 

() 

0 

I... c 

"'0 0 600 ~..0 

~ I... 

rn cu 

E u 

cu () I... ·- C)c cu C) 

E I... 

0 450 

c E 

cu 

X I... 

Q) C) 

"'0 I... 

c Q) 

a. 

c 

Q) 300 C) 

0 

I... 

"'0 

~ 

I 

150 

0 

0 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I 

I TYPE I 

I 

Kansas- Nebraska 

Powell No. 1-33 well 

854-871 m 

' ,, ,~ 

I 

/ ~ )-TYPE II ""'~ Kansas- Nebraska 

~whomble No. 1-32 well 

J 

/ 

/.. 

/ 

' 

/ 

/ 

,.... 

r..... 1'\ ' ...... ........ _LY 

""""---= ~v 

~Excelsior 

Alice G. Nay No. 1 well 

1655-1668 m 

.,.,...-------- ........ 

447-461 m 

---TYPE III I / 

I 

50 100 150 200 

Oxygen index, in milligrams C02 __ ._ 

per gram organic carbon 

Figure 3. Modified van Krevelen diagram showing type and maturity of kerogen from three selected 

wells. Pyrolysis data for all five wells are presented in Rice (1984). 

hydrogen-rich and is made up predominantly of 

sapropelic material that is typical of an open-marine 

environment. Data from visual kerogen analyses and 

elemental analyses confirm this interpretation of the type 

of kerogen (Rice, 1984). Type II kerogens are generally 

considered to generate oil and associated gas during the 

intermediate stages of thermal history. As temperatures 

increase, lighter hydrocarbons (wet gas-condensate and 

80 Studies in Diagenesis 

finally dry gas) are formed more rapidly because of thermal cracking processes affecting residual kerogen and 

previously formed heavier hydrocarbons. 

THERMAL MATURITY 

In general, the level of maturation in the Niobrara 

on the shallow east flank of the basin increases to the 
northwest. This increase is consistent with the following 

changes: (1) present-day depth of burial increases to the 

northwest, (2) porosity and permeability of the chalks 

decrease 1to the northwest (Lockridge and Scholle, 1978), 

(3) carbon isotope compositions of biogenic gases become 

heavier within the producing trend to the northwest 

(fig. 2), and (4) oil has been produced from Niobrara Formation in well E (fig. 1) northwest of the gas-productive 

area. 

Pyrolysis data were used to determine thermal 

maturity by the position of samples on the evolutionary 

path of a particular type of kerogen using a modified van 

Krevelen diagram (fig. 3). As discussed earlier, the kerogen 

from all five wells examined was classified as type II. 

Samples from wells A and C, which are considered typical 

for wells B and D also, plot on the immature part of the 

evolutionary path. Kerogen from samples of well E, 100 

km northwest of the gas-productive area, has a substantially lower "hydrocarbon index" and is probably in the 

mature stage. The same levels of maturity were indicated 

by clay-mineral transformations (Pollastro and Scholle, 

in press) and thermal alteration index, vitrinite reflectance, 

elemental analysis, and solvent extraction and elution 

chromatographic data (Rice, 1984). Present-day bottomhole temperatures from wells penetrating the Niobrara 

Formation suggest that the geothermal gradient was 

higher and (or) that more overburden was present to 

account for the generation of liquid hydrocarbons in the 

vicinity of well E. The pattern of maturity indicated by 

these data is consistent with that presented for the 

Niobrara by Tainter (1982) and for the slightly deeper 

Carlile-Graneros-Mowry interval by Clayton and Swetland (1980). The pattern provides evidence that oil and 

wet gas-condensate produced from Niobrara and adjacent 

Codell reservoirs in the deep Wattenberg area of the basin 

were probably formed by thermal degradation and cracking that affected type II kerogen concentrated in the 

Niobrara.. In contrast, the dry gases found on the shallow 

eastern nank have been derived from low-temperature 

microbiological processes. 

SUMMARY AND CONCLUSIONS 

Data and interpretations presented in this paper suggest the following conclusions: 

1. Organic-rich laminations containing an average 

of 3.2 percent organic carbon representing type II kerogen 

are prestmt in the Niobrara Formation. 

2. The Niobrara Formation and underlying Codell 

Sandstone Member of the Carlile Shale produce oil and 

wet gas-condensate in the deeper part of the Denver basin 

and dry gas on the shallow eastern flank. 

3. The shallow eastern-flank gases are chemically 

dry and isotopically light (o13C1 values range from -65 

to -55 per mil) and are probably of biogenic origin. Oil 

and wet gas-condensate produced from deeper in the basin 

were probably formed by thermal cracking and degradation processes that affected kerogen concentrated in the 

Niobrara. This interpretation is consistent with the occurrence of isotopically heavier methane gases (o13C1 values 

range from -47.8 to -43.5 per mil) found in this area. 

4. The pattern of source-rock maturity in parts of 

the Niobrara Formation containing shallow east-flank gas 

fields, when combined with available data from the deeper 

part of the basin, is consistent with the geochemistry and 

types of hydrocarbons present throughout the basin. 

REFERENCES CITED 

Clayton, J. L., and Swetland, P. J., 1980, Petroleum generation 

and migration in Denver basin: American Association of 

Petroleum Geologists Bulletin, v. 64, p. 1613-1633. 

Espitalie, J., Madec, M., Tissot, B., Mennig, J. J., and Leplat, 

P., 1977, Source rock characterization method for petroleum exploration: Ninth Annual Offshore Technology Conference, Proceedings, v. 3, p. 439-444. 

Lockridge, J. P., and Scholle, P. A., 1978, Niobrara gas in eastern 

Colorado and northwestern Kansas, in Pruitt, J. D., and 

Coffin, P. E., eds., Energy resources of the Denver basin: 

Denver, Colorado, Rocky Mountain Association of 

Geologists, p. 35-49. 

Mallory, W. W., 1977, Oil and gas from fractured shale reservoirs in Colorado and northwest New Mexico: Rocky 

Mountain Association of Geologists Special Publication 1, 

38 p. 

Petroleum Information, 1983, Special report-Codell production; practices and problems: August 25, 1983, p. 7-12. 

Pollastro, R. M., and Scholle, P. A., 1984, Hydrocarbons exploration, development from low-permeability chalks-Upper 

Cretaceous Niobrara Formation, Rocky Mountains region: 

Oil and Gas Journal, April 12, p. 140-145. 

Pollastro, R. M., and Scholle, P. A., in press, Diagenetic relationships in a hydrocarbon-productive chalk; The Cretaceous Niobrara Formation, in Spencer, C. W., and Mast, 

R. F., eds., Geology of tight gas reservoirs: American 

Association of Petroleum Geologists Memoir. 

Rice, D. D., 1984, Occurrence of indigenous biogenic gas in 

organic-rich, immature chalks of Late Cretaceous age, 

eastern Denver basin, in J. G. Palacas, ed., Geochemistry 

and source-rock potential of carbonate rocks: American 

Association of Petroleum Geologists Studies in Geology 

18, p. 135-150. 

Rice, D. D., and Claypool, G. E., 1981, Generation, accumulation, and resource potential of biogenic gas: American 

Association of Petroleum Geologists Bulletin, v. 65, p. 5-15. 

Rice, D. D., Threlkeld, C. N., and Vuletich, A., 1984, Nature 

and origin of "vent gases" in the La Salle area, northeastern Colorado: U.S. Geological Survey Open-File Report 

85-220, 9 p. 

Tainter, P. A., 1982, Investigation of stratigraphic and paleostructural controls on hydrocarbon migration and entrapment in Cretaceous D and J sandstones of the Denver basin: 

Boulder, Colorado, University of Colorado M.S. thesis, 

235 p. 

Hydrocarbon Type Versus Organic-Matter Maturity in Chalks, Denver Basin 81 

The Relationship Between Vitrinite Reflectance, 

Metamorphic Grade, and Temperature in the Cerro Prieto, 

Salton Sea, and East Mesa Geothermal Systems, 

Salt<Jn Trough, United States and Mexico 

By CI--IARLES E. BARKER, BONNIE L. CRYSDALE, and MARK j. PAWLEWICZ, 

U.S. Geological Survey, Box 25046, MS 921, 

Denver, Colorado 80225 

CONTENTS 

Abstract 84 

Introduction 84 

Acknowledgments 84 

Methods of investigation 84 

Comparison of metamorphic grade, temperature, and rank 86 

Kaolinite 86 

Smectite and illite/smectite 87 

Illite 88 

Chlorite 89 

Dolomite 89 

Wairakite 90 

Prehnite 90 

Epidote 91 

Biotite 91 

Discussion 91 

Conclusions 93 

References cited 93 

FIGURES 

1. rvlap showing geothermal systems, Salton trough, Mexico and United 

States 85 

2-5. Graphs showing: 

2.. Borehole temperature and temperature computed from vitrinite reflectance measured at changes in metamorphic grade within the Cerro 

Prieto geothermal system 87 

3.. Borehole temperature and temperature computed from vitrinite reflectance measured at changes in metamorphic grade within the Salton Sea 

and East Mesa geothermal systems 88 

4.. Borehole temperature range of each metamorphic grade given in figure 

2, Cerro Prieto geothermal system 89 

5. Temperature range of selected metamorphic grades in geothermal 

systems except for the Cerro Prieto system 90 

6. Plot comparing of borehole temperature data and other geothermometers 

in borehole M -94, Cerro Prieto system 92 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 83 
Abstract 

The vitrinite reflectance of sedimentary organic matter in 

the Cerro Prieto, Salton Sea, and East Mesa geothermal systems 

was found to be strongly related to measured borehole 

temperature. Vitrinite reflectance was used to estimate the maximum temperature (T max) to which rocks in these systems had 

been exposed because other methods of estimating T max and 

thermal history reconstruction showed that these systems are 

now at maximum temperature. Vitrinite reflectance has an advantage over some inorganic-based geothermometers because 

the temperature of the borehole can usually be estimated at 

any depth, giving a continuous thermal record throughout the 

system. By comparison, an estimation ofT max from changes in 

mineral assemblages (metamorphic grade) was found to be 

relatively insensitive. Metamorphic grade is defined as the 

existence of a particular assemblage of index minerals; however, 

these assemblages have been found to react or persist over wide 

temperature ranges; thus, metamorphic grade is only a crude 

geothermometer. 

The degree of inaccuracy possible in estimating T max 

using metamorphic grade is illustrated by a short-term thermal 

event at a depth of 1,300 m in a borehole in the Cerro Prieto 

field, Mexico. This thermal event induced a reequilibration of 

oxygen isotopes and formed high-temperature fluid inclusions 

in veins and host rock. Vitrinite reflectance increased in the 

rock, and fission tracks in detrital apatite were completely 

annealed, indicating a T max of 250°-300°C. Metamorphic 

grade, as estimated by an assemblage change of smectite + kaolinite to illite + chlorite, suggests a T max of only 

150°-180°C. 

The wide temperature range over which certain mineral 

assemblages can form and (or) persist indicates that metamorphic grade is less accurate than vitrinite reflectance in estimating 

Tmax· 

INTRODUCTION 

Thermal history of a rock can be inferred by comparing geothermometers that are set by different events 

in a geothermal system's history. The present study compares the accuracy of geothermometers based on vitrinite 

reflectance and metamorphic grade with measured temperatures in three Salton Trough geothermal systems: 

Cerro Prieto (Mexico) and Salton Sea and East Mesa 

(United States). Geothermal systems are model environments for comparing various geothermometers because 

of their strong thermal imprint on the rocks and their 

relatively short, well-known thermal histories. 

Thmperature in geothermal systems apparently can 

be estimated by changes in both organic matter (Barker, 

1983) and mineral phases (see, for example, Elders and 

others, 1980; Schiffman and others, 1984). The ability of 

such geothermometers to estimate the maximum temperature to which a system has been exposed (T max) is important because a comparison of past temperatures, as 

estimated by changes in the rock, and contemporary 

temperatures indicates whether or not the system has 

84 Studies in Diagenesis 

cooled. This is a crucial observation in thermal history 

reconstruction. 

Temperature a,ppears to be the dominant factor in 

the metamorphism of sedimentary organic matter (OM) 

(Tissot and Welte, 1984; Durand, 1980). Barker (1983) 

found that the rank of OM thermal metamorphism in six 

geothermal systems (including those in the Salton Trough) 

was solely dependent on T max after a heating duration of 

10,000 years. The stability of minerals in geothermal 

systems is also dominantly temperature dependent; pressure seems to have only a minor effect on the degree of 

organic or inorganic metamorphism in these systems 

(Browne, 1978; Tissot and Welte, 1984). Several other factors, however, such as rock and fluid composition (that 

is, cation supply) and permeability, affect the metamorphism of minerals, but apparently do not affect organic 

metamorphism (Stach and others, 1982). The general 

dependence of metamorphic grade on temperature is 

illustrated by the consistent occurrence of certain mineral 

assemblages at similar temperatures in different geothermal systems that have been heated for widely different 

lengths of time. 

Numerous studies indicate that mineral assemblages 

in geothermal systems result from equilibrium reactions 

at particular temperatures (Zen and Thompson, 1974; 

Browne, 1978). Changes in metamorphic grade are therefore indicated by the appearance or disappearance of certain index minerals (as used by Hoffmann and Hower, 

1979) and can be correlated with T max in geothermal 

systems (see, for example, Elders and others, 1980; Schiffman, 1984). This relationship has therefore been used to 

infer maximum temperatures in systems where direct 

temperature measurements have not been made. Metamorphic grade, however, is a discontinuous geothermometer, at best, because temperatures can be estimated only 

at specific depths in the borehole where changes of 

mineral assemblages are identified. 

Acknowledgments 

Borehole samples and temperature logs were provided by Comision Federal de Electricidad (Cerro Prieto, 

Mexico) through W. A. Elders (University of California, 

Riverside) and Marcello Lippman (Lawrence Berkeley 

Laboratory, Livermore, California). 

METHODS OF INVESTIGATION 

Comparison of organic and inorganic metamorphism was limited to three geothermal systems in the 

Salton Trough-Salton Sea, East Mesa, and Cerro Prieto 

(fig. 1), because these were the only systems for which adequate Rm, T max' and mineral-assemblage data were 

available. Mineralogy and temperature data for the Salton 
32° 

CJ 

CJ 

El CentroO 

EXPLANATION 

ALLUVIUM 

CLASTIC SEDIMENTARY 

ROCKS 

IGNEOUS AND 

METAMORHIC ROCKS 

10 20 

L:;HEBER 

• 

30 KILOMETERS 

GEOTHERMAL 

CENTER 

VOLCANIC 

CENTER 

FAULT 

GULF OF 

CALIFORNIA 

Figure 1. Geothermal systems, Salton Trough, Mexico and United States. The Cerro Prieto, Salton Sea, and East Mesa systems 

were selected for study. 

Cerro Prieto, Salton Sea, and East Mesa Geothermal Systems 85 
Sea system were reported by Muffler and White (1969), 

Helgeson (1968), Clayton and others (1968), and 

McDowell and Elders (1979, 1980). Hoagland (1976) and 

Miller (1980) reported similar data for the East Mesa 

system. Vitrinite reflectance data for the Salton Sea and 

East Mesa systems were taken from Barker (1979, 1983). 

For the Cerro Prieto system, mineralogical data are from 

Elders and others (1977, 1978, 1979), temperature data are 

from unpublished borehole measurements by the Comision Federal de Electricidad (Mexico), and vitrinite reflectance data are from Barker (1979), Barker and Elders 

(1981), Barker and others (1981), and unpublished data. 

Mineral assemblage data in the Cerro Prieto system were 

compiled from sandstone samples only, except as otherwise noted. 

Borehole temperatures (T b) were measured using 

recording thermometers after the well had recovered from 

cooling effects of the drilling fluid. Thermal history 

reconstruction (Barker, 1983) indicates that present-day 

borehole temperatures represent the maximum temperatures reached in these systems, and they are used here as 

measured T max values. The T b at the depth where a 

change in metamorphic grade occurred was interpolated 

from thermometer-recorded temperature logs having the 

longest borehole re-equilibration time. This depth was 

used as a reference point to compare the borehole temperature and geothermometer estimates. Rm values of OM 

collected near this same depth were estimated from a 

manually smoothed Rm -depth curve. The Rm data were 

converted to a temperature estimate (T vr) using the 

temperature-Rm curve for western North America geothermal systems (Barker, 1983). The Tb and Tvr estimates 

were used as x-y coordinates to plot the change in 

metamorphic grade in the Cerro Prieto field (fig. 2) and 

the other Salton Trough fields (fig. 3). A linear equation 

T vr = T b was plotted on these figures to compare perfect 

correlation with the data. 

Published data on authigenic minerals from 

geothermal fields were variable in quality. In particular, 

little information was given on the procedure used to identify the clay minerals, which was highly dependent on the 

method of sample preparation. Most of the clay mineral 

identifications in geothermal systems outside the Salton 

Trough were made by comparing the X-ray diffraction 

(XRD) patterns of glycolated and unglycolated samples. 

Diagnostic tests for distinguishing montmorillonite from 

the rest of the smectite group were apparently not performed, and the general term smectite was used for all 

these minerals. Illite was variously called K-mica or 

sericite, and discrete illite and smectite were commonly 

grouped with mixed-layer illite/smectite (1/S). Likewise, 

kaolinite was used in many of the papers to indicate only 

members of the kaolin group and not a specific mineral. 

Some metamorphic-grade index minerals also occur as 

detrital grains in the Cerro Prieto sediments, making it 

difficult to determine the depth of the first authigenic 

86 Studies in Diagenesis 

occurrence of a particular mineral. These problems caused 

us some difficulty in determining the temperature range 

over which certain index minerals persist. 

A change in metamorphic grade is defined as the 

appearance or disappearance of an index mineral or 

assemblage of index minerals with increasing depth 

(temperature) in a borehole. The appearance of an 

authigenic index mineral was recognized by the first 

significant increase in abundance above background level, 

as indicated by XRD, or by its identification based on textural criteria in thin-section studies. The disappearence 

of an index mineral was recognized by its decrease in 

abundance to below XRD detection limits. These methods 

were described in detail by Elders and others (1978). 

COMPARISON OF METAMORPHIC GRADE, 

TEMPERATURE, AND RANK 

A relationship between metamorphic grade, T vr' 

and T b is indicated by the data cluster about the T vr 

= Tb line in all three geothermal systems (figs. 2 and 3). 

Linear regression of these data suggests that correlation between T vr and T b is moderate (r2 = 0.5 at Cerro 

Prieto) to strong (r2 = 0.8 at Salton Sea/East Mesa). 

The difference in correlation between T vr and T b in the 

Cerro Prieto and the other Salton 'Ifough fields is attributed to (1) the greater variability in Tb data from Cerro 

Prieto, which is probably due to widely different borehole 

recovery times before measurement (Barker and others, 

1981), and (2) the fact that some Cerro Prieto data were 

taken from the margins of the system where short-term 

heating and (or) cooling events have taken place (Barker 

and Elders, 1981; Sanford, 1981). 

The temperature ranges of metamorphic grades in 

the Cerro Prieto system (fig. 4) versus those found in other 

geothermal systems (fig. 5) are discussed in the following 

sections. This discussion involves comparing the maximum and minimum temperature for the same metamorphic grade in all the boreholes from each geothermal 

system and the maximum and minimum values of T b 

and T vr· The T b range is listed first, followed by the T vr 

range (in parentheses) for each metamorphic grade. The 

temperature ranges for changes in metamorphic grade 

during burial diagenesis from Hoffman and Hower (1979) 

are also listed, if available. Metamorphic· grades in the 

Salton Sea and East Mesa systems are not discussed in 

detail here because these metamorphic grades are not well 

defined by the sparse data. 

Kaolinite 

The Cerro Prieto geothermal reservoir consists of 

sediments deposited in the Colorado River Delta (Lyons 

and Van de Kamp, 1979). Kaolinite is a common detrital 
300 

250 D. 51 

(/) 

::J 

en 

...J 

&51 w 

u 

(/) 

w 

w 200 a: 

(.!) 

w 

c 

~ 

.....: 

~ 

w 

a: 150 ::J 

1-

<( 

a: 

w 

0.. 

~ 

w 

1-

100 

50 

0 25 

0 10 

8_84 

53o&25 

<fu53 

510 0105 

~;.53 

256. 19A0 

106. g0 

[!]51 

EXPLANATION 

D Kaolinite disappearance 

8. Smectite disappearance 

1;. Illite/smectite disappearance 

0 Illite disappearance 

• Chlorite increase 

D. Dolomite disappearance 

[!] Wairakite appearance 

• Epidote appearance 

0 Biotite appearance 

+ Prehnite appearance 

BOREHOLE TEMPERATURE (Tb), IN DEGREES CELSIUS 

Figure 2:. Borehole temperature (T b) and temperature computed from vitrinite reflectance (Tvr) measured at changes in 

metamorphic grade within the Cerro Prieto geothermal system. Mineral data from Elders and others (1978). Numbers indicate 

the borehole in which the metamorphic grades were measured. The line corresponds to Tvr = Tb. 

component of unmetamorphosed sediments in the Cerro 

Prieto system (Muffler and Doe, 1968; Elders and others, 

1978). The depth at which kaolinite disappears in the 

borehole (fig. 4) was used as a reference point to compute Tb and to compare with T vr (fig. 2). Kaolinite disappearance persists over a Tb range of 80°-250°C (compared with a Tvr range of 125 ° -260°C} in the Cerro 

Prieto system, a slightly wider range than the 55°-200°C 

T b range found for other geothermal systems (fig. 5). 

Weaver (1979) concluded that kaolinite persists to 375 °C 

in low-grade metamorphic environments, and that its 

disappearance is in part dependent on the available reaction time, the chemistry of associated clay minerals, and 

the organic content of the rock. These considerations and 

the wide temperature range over which the mineral 

apparently disappears in geothermal systems suggest 

that the presence or absence of kaolinite is a poor 

geothermometer. 

Smectite and illite/smectite 

Smectite is also a common detrital constituent in 

the Cerro Prieto system (Elders and others, 1978). Smectite (fig. 4) disappears over a Tb range of 80° -240°C 

(125°-215°C Tvr) in the Cerro Prieto system. This Tb 

range is not significantly different from the 90 °-265 OC 

(fig. 5) range for smectite disappearence found in other 

systems, but the T vr range (125 °-215 °C) over which 

smectite disappears is narrower. Weaver (1979) noted that 

Cerro Prieto, Salton Sea, and East Mesa Geothermal Systems 87 
300 

(/) 

~ 250 ~ 

w 

u 

(/) 

w 

w 

a: 

~ 200 Q 

z 

t? 

w 150 

a: 

W-1 (

3787) 0 MM-2 (1461) 

0 •,eRR-1 (2481) 

.MM-2 (615) '\MM-2 (1 000) 

W-1 0 (3333) Q 

.RR-1 (1077) 

EXPLANATION 

::::l 

~ 

M- .6-2 (606) 

SP-1 0,6(304) b A ,&M-.6-2 (841) 

(304) """' 

.6 Kaolinite disappearance 

& Smectite disappearance 

0 Illite/smectite disappearance 

0 Illite appearance 

a: 

w 

c.. 

~ 100 

1-

50 

W-1 (1272)& ~ M-.6-2 (673) 

W-1 (1741) 

.& Chlorite appearance 

D Chlorite disappearance 

0 Dolomite disappearance 

• Epidote appearance 

• Biotite appearance 

(1654) = depth in meters to this 

metamorphic grade 

S-4 =well 

50 100 150 200 250 300 350 400 

BOREHOLE TEMPERATURE (T b). IN DEGREES CELSIUS 

Figure 3. Borehole temperature (T b) and temperature computed from vitrinite reflectance (T) 

measured at changes in metamorphic grade within the Salton Sea and East Mesa geothermal systems. 

The line corresponds to Tvr = T b' Sportsman-1 (annotated SP-1) and Wilson-1 (W-1) mineral data 

from Muffler and White (1969). River Ranch-1 (RR-1 ), Magmamax-2 (MM-2), and Sinclair-4 (S-4) 

mineral data from McDowell and Elders (1979, 1980). East Mesa 6-2 (M6-2) mineral data from 

Hoagland (1976) and Miller (1980). 

some ambiguity exists in the method of determining whether smectite is present or absent in a sample, which may 

explain the difference in the apparent wide temperature 

range of smectite disappearence. He suggested that the 

absence of the 17 -A peak in the XRD pattern after 

glycolation is generally accepted as evidence for the absence of smectite; however, mixed-layer clays may contain 

as much as 60 percent 10-A layers (40 percent 17-A layers) 

and still produce a 17-A peak. Thus, discrete smectite 

could have disappeared at a lower temperature and at a 

much shallower depth than that found in these geothermal systems, a contention that is supported by the upper 

limit of Tvr being about 35°-50°C lower than the upper 

limit of the T b range. Published data on geothermal 

systems are vague, but the broad temperature range over 

which smectite seems to disappear is probably due in part 

to this identification problem. The wide temperature range 

for smectite disappearance and the ambiguity in its identification make this mineral a poor geothermometer. 

In some studies I/S has been lumped with smectite, 

obscuring the distribution of both minerals; the inability 

to recognize liS hinders the determination of smectite 

88 Studies in Diagenesis 

disappearance. The disappearance of I/S is defined in 

some papers to be equivalent to the loss of expandable 

layers upon glycolation. Because of ambiguity in the identification of I/S, no attempt was made here to determine 

the first appearance of I/S, which should be at about 

50°-600C (Weaver, 1979; Hoffman and Hower 1979). In 

the Cerro Prieto system, I/S disappearance was taken to 

be the absence of I/S having more than 10 percent expandable layers (Elders and others, 1978). Accordingly, I/S 

disappeared over a Tb range of 130°-255°C {Tvr = 

160° -270°C} in this field, a range comparable to the 

120 °-265 °C found in other geothermal systems. I/S disappearance is a poor geothermometer in geothermal systems because I/S disappears over a wide temperature 

range and can apparently persist far above the lowest temperature at which it disappears in other systems (130 °C). 

Illite 

Illite is also a common detrital component of the 

Colorado River delta (Muffler and Doe, 1968) and of the 

unmetamorphosed sediments in the Cerro Prieto system 
Kaolinite disappearance 

Smectite disappearance 

lllite/!;mectite disappearance 

Illite disappearance 

Chlorite increase 

Dolomite disappearance 

Wairakite appearance 

Prehnite appearance 

Epidote appearance 

Biotite appearance 

75 100 150 200 250 300 350 

BOREHOLE TEMPERATURE (Tb), IN DEGREES CELSIUS 

Figure 4. Borehole temperature range of each metamorphic grade, Cerro Prieto geothermal system. Long dashes, range 

of each metamorphic grade using temperature from vitrinite reflectance (T) given in figure 2. Short dashes, temperature 

range of each metamorphic grade measured using temperature estimated from oxygen isotpe analyses (Tiso) of calcite 

in sandstone (Schiffman and others, 1984). 

(Elders and others, 1978). Consequently, only the last 

occurrence (disappearance) of illite is important in determining metamorphic grade. Illite disappears over a T b 

range of 180° -310°C (Tvr = 230° -300°C) at Cerro 

Prieto (fig. 4). This range is narrower than the 90° -290°C 

T b range;~ observed in other geothermal systems (fig. 5). 

Hoffman and Hower (1979) indicated that illite disappears 

at about 300 °C in lower Tertiary and Upper Mesozoic 

shales. Illite disappearance is a poor geothermometer in 

geothermal systems because the mineral seems to disappear over a wide temperature range and can apparently 

persist in some systems far above the lowest temperature 

(100°C) at which it disappears in others. These data 

indicate, however, a consistent upper limit of 300° ± IOOC 

for illite disappearance; thus, illite disappearance may 

accurate:ly indicate a minimum temperature that the 

system reached. 

Chlorite 

D<!trital chlorite is also a common constituent in 

Cerro Prieto rocks. Consequently, chlorite grade is defined 

as the shallowest depth at which a significant increase in 

chlorite content occurs relative to background levels. 

Chlorite: appears over a Tb range of 80°-280°C (Tvr = 

130° -270°C) in the Cerro Prieto system and over aT b 

range of 70° -300°C in other systems (figs. 4 and 5). 

Outside of geothermal systems, the appearance of chlorite 

has been observed at depths corresponding to temperatures as low as 80 °C and as high as temperatures 

equivalent to greenschist metamorphism (greater than 

300°C) (Weaver, 1979; Hoffman and Hower, 1979). The 

appearance of chlorite is a poor geothermometer in 

geothermal systems because it appears over a wide temperature range, and can apparently persist in some systems 

far above the lowest temperature (80 OC) at which it disappears in others. 

Dolomite 

Dolomite is a detrital component of Salton Trough 

geothermal systems sediment. Dolomite disappearance, 

defined here as the last occurrence of detrital dolomite, 

takes place over a Tb range of 130°-3100C (Tvr = 150°-

2600C) in the Cerro Prieto system (fig. 4) and a Tb 

range of 50° -280°C in other Salton Trough systems 

(fig. 5). In the Salton Sea system, dolomite reacts with 

kaolinite, ankerite, and ferrous iron to produce chlorite, 

calcite, and C02 at temperatures as low as 1800C and 

possibly as low as 125 OC (Muffler and White, 1969). This 

mineral has not been found in other geothermal areas 

(Browne, 1978), so no comparison of this metamorphic 

Cerro Prieto, Salton Sea, and East Mesa Geothermal Systems 89 
Kaolinite disappearance 

Smectite disappearance 

Illite/smectite disappearance 

Illite disappearance 

Chlorite increase 

Dolomite disappearance 

Wairakite appearance 

Epidote appearance 

Prehnite appearance 

Biotite appearance 

50 100 150 200 250 300 350 

BOREHOLE TEMPERATURE (Tb), IN DEGREES CELSIUS 

Figure 5. Temperature range of selected metamorphic grades in geothermal systems, except for the Cerro Prieto system 

(given in figure 4). Data from Bird and Elders (1975); Browne (1978); Cavarretta and others (1980); Funicello and others 

(1979) Hayashi (1973); Kristmannsdottir (1975); McDowell and Elders (1980); Muffler and White (1969); Naboko and 

Piip (1961); Sheridan and Maisano (1975); Shimazu and Yajima (1973); Sigvaldason (1962); Steiner (1977); Sumi (1968); 

White and Sigvaldson (1962). 

grade with other systems is possible. Dolomite disappearance is a poor geothermometer in geothermal systems 

because the mineral disappears over a wide temperature 

range, and can apparently persist in some systems far 

above the lowest reaction temperature (about 125 OC) at 

which it disappears in others. 

Wairakite 

Detrital wairakite has not been found in the Cerro 

Prieto system (Muffler and Doe, 1968). Authigenic 

wairakite in other geothermal systems forms at temperatures as low as 140°C (and perhaps as low as 75 °C 

(Browne, 1978)), or as high as 260°C (fig. 5). At Cerro 

Prieto (fig. 4), wairakite appears over a Tb range of 

210° -290°C {Tvr = 170°-295 °C). Schiffman and others 

(1984) reported a wairakite stability range, based on 

temperatures computed from oxygen isotope (Tiso) 

analyses of calcite cement, of about 200° -300°C in the 

Cerro Prieto system. Browne (1978) suggested that 

wairakite usually appears at a temperature of about 

220 °C, but cites examples from the Katayama, Japan, 

geothermal system in which wairakite apparently formed 

as low as 7 5° -17 5 °C. He suggested that these lower 

90 Studies in Diagenesis 

temperatures resulted from formation at relatively low 

pressure. In any case, the apparent wide temperature range 

for wairakite appearance is 75 ° -300°C, suggesting it is 

a poor geothermometer in geothermal systems. A better 

indicator of the minimum temperature a system has 

reached is the disappearance of wairakite, which occurs 

at about 300 °C in geothermal systems. 

Prehnite 

Prehnite is not a detrital phase in the Colorado 

River delta sediments (Muffler and Doe, 1968) and has 

not been reported in unmetamorphosed sediments at 

Cerro Prieto. The first appearence of prehnite at Cerro 

Prieto (fig. 4) is over a Tb range of 290° -370°C {Tvr = 

290°-325 °C). Schiffman and others (1984) found that 

prehnite, based on temperatures computed from the 

oxygen isotope calcite-water geothermometer {Tiso), 

appears at a temperature of 275 °C and persists to at least 

370°C, the maximum Tb in the Cerro Prieto system. 

Prehnite appears at about 250 °C in other geothermal 

systems (fig. 5). The wide temperature range for the 

appearence of prehnite indicates that it is a poor geothermometer in geothermal systems. 
Epidote 

Epidote-group minerals occur as rare detrital grains 

in the Cerro Prieto system. Authigenic epidote can be 

distinguished petrographically by its occurrence as 

euhedral, pore-filling crystals, distinct from rounded 

detrital grains. Epidote appearance is defined as the first 

occurrenc:e with depth of authigenic epidote in the borehole. Epidote appears over a Tb range of 185°-325°C 

(Tvr = 170°-315°C) in the Cerro Prieto system (fig. 4) 

and over a T b range of 135 °-305 OC in other geothermal 

systems (fig. 5). Schiffman and others (1984) reported 

that epidote appears at a Tiso of 230 °C and persists to 

at least 3700C, the maximum observed temperature in the 

Cerro Prieto system. The lower temperature estimates for 

the other systems appears to be the result of epidote forming at higher temperature in systems that have now 

cooled (Browne, 1978). The first appearence of epidote 

in low-presssure geothermal systems is at about 230 °C 

(Seki, 1972). Assuming this temperature to be the lower 

limit of epidote appearance, this reaction apparently 

occurs over a temperature range of 230°-325 °C. The 

wide temperature range for epidote appearance indicates 

that this mineral is a poor geothermometer in geothermal syste:ms. 

Biotite 

Biotite is a common detrital component in the 

Cerro Prieto geothermal reservoir rock. The T b range of 

the first appearance of authigenic biotite (as determined 

petrographically) is 320°-340°C (315°-335°C Tvr> at 

Cerro Prieto (fig. 4) and 305 ° -340°C in other geothermal systems (fig. 5). This narrow temperature range for 

the first appearance of biotite suggests that it is a good 

geothermometer; however, it usefulness is limited because 

authigenic biotite is a relatively rare mineral in geothermal systems (Browne, 1978) and occurs near the bottom 

of only the hottest boreholes. 

DISCUSSION 

The~ wide temperature range for the occurrence of 

each metamorphic grade noted here could be due to cooling in th€~ system. Evidence for this cooling is the lower 

Tb compared with Tvr because, with cooling, Tvr retains 

the maximum temperature the system has reached, 

whereas T b decreases. The depths at which these metamorphic grades change, however, are typically marked 

by similar Tb and Tvr ranges. Thus, the wide temperature range for each metamorphic grade change appears 

not to be caused by problems in temperature 

measurements, such as cooling in the system. Changes 

of metamorphic grade appear to take place over a 

considerable temperature range. Consequently, metamorphic grade is a less accurate geothermometer than 

temperature computed from Rm because the OM rank 

appears to be a direct function of temperature. 

The general lack of response of metamorphic grade 

to increasing temperature may be due to other factors that 

significantly affect mineral stability, including (1) 

pressure, (2) bulk-rock composition, (3) fluid composition, (4) permeability, and (5) reaction duration (Zen, 

1974; Browne, 1978; Hower, 1981). The role of pressure 

is minor because geothermal systems generally conform 

to a hydrostatic gradient and pressure seldom exceeds 200 

bars (Browne, 1978). Furthermore, Zen (1974) showed 

that pressure exerts little direct effect on mineral alteration in geothermal systems. 

Bulk-rock composition significantly controls metamorphic mineral formation in geothermal systems. It 

directly affects the resultant metamorphic mineral asemblages and the apparent degree of diagenesis, even in 

rocks that have been subjected to similar temperatures 

(Rettke, 1976; Hower, 1981). At Cerro Prieto, most of 

the reservoir rock was derived from a single source (Muffler and Doe, 1968}, but the depositional environment has 

sorted this material into sand-rich and mud-rich rock 

(Lyons and van de Kamp, 1979). These two rock types 

show moderate differences in metamorphic response, 

mainly in that with increasing temperature, the index 

minerals in the permeable sandstone react faster than 

those in the low permeability mudstone (Elders and 

others, 1978). 

Fluid composition and permeability also significantly affect mineral formation in geothermal systems 

because these reactions are rarely isochemical and mass 

transfer is necessary for hydration-dehydration and (or) 

cation diffusion into and out of a mineral (Browne, 1978; 

Weaver, 1979). Mass transfer is accomplished through 

meteoric waters that react with the country rock and 

hydrothermally altered reservoir rock (Ellis, 1979). Movement of these fluids and their access to the reservoir rock 

are controlled by permeability. Consequently, in relatively 

impermeable rocks, equilibrium between the solid and 

fluid phase is seldom achieved, and authigenic or detrital 

minerals in the reservoir rock can persist to temperatures 

past their normal equilibrium stability range. Elders and 

others (1980) demonstrated that permeability, fluid flow 

patterns, and temperature are controlling factors in the 

metamorphism at Cerro Prieto. 

The effect of reaction duration on inorganic metamorphism in geothermal systems is important. Weaver 

(1979) indicated that the formation of illite-rich shales 

requires temperatures of 200° -3000C for millions of 

years. Hoffman and Hower (1979) indicated that reaction duration is significant for increasing metamorphic 

grade at low temperatures and that these reactions do not 

approach completion in 107 years. Laboratory experiments on smectite, one of the most reactive clay minerals 

Cerro Prieto, Salton Sea, and East Mesa Geothermal Systems 91 
APPARENT AGE, IN MILLIONS OF YEARS 

0 100 200 300 

0 

• 

EXPLANATION 

• 

• Tb Log T-11 

• .. Th 

• D. Tvr 

0 Tiso 

0.5 • 1----+--1 F-T age 

• 

D. 

6 • 

6 • 

D. 

6 ~ 

• 

6 

1.0 6 • 

D. I •• I I (/) 

a: D. • w D. 0 1-

w 

~ D. • 0 

...J D. 1292m K+S :;2 D. eo ~ 

::i • D. 1-

a.. 

w D. 0 

1.5 0 • D. 

6 

• D. .. D. 

• D. 

D. 

• 0 ~ 

D. 

• D. 

D. 

2.0 • D. 

• D. 

0 • 

• 

\ D. 

0 • D. 

.. D. 

06 

D. 

2.5 0 D. 

0 100 200 300 

TEMPERATURE, IN DEGREES CELSIUS 

Figure 6. Comparison of borehole temperature (T b) data and other geothermometers in borehole M-94, Cerro Prieto system. Fission track apatite age (F-T), 

fluid inclusion (T h), and oxygen isotope (Ti) data are from Sanford (1981). Mineral 

data are from Elders and others (1978). Vitrinite reflectance data are from Barker 

and Elders (1981). F-T dates are plotted with ± one standard deviation error bars. 

92 Studies in Diagenesis 
(Weaver, 1979), show that it can probably persist for at 

least 108 years at earth-surface conditions (Eberl and 

Hower, 1976). Geothermal systems exist for as long as 

1()6 years (Barker, 1983), and short-term heating events 

along the margins of a system may last for less than 100 

years (Sanford, 1981). The life span of geothermal 

systems is at least an order of magnitude smaller than 

that required for some mineral reactions to equilibrate. 

Thus, mineral reactions in geothermal systems where 

temperature can change rapidly should be significantly 

time dependent. This time dependency is illustrated by 

a short-term heating event in borehole M-94 at 1,300 m 

(Cerro Prieto) that produced temperatures of 250°-

300 OC, based on mineral fluid inclusions, mineral oxygen 

isotopes, and vitrinite reflectance (fig. 6). A study of 

fission-track annealing in apatite indicated that this 

heating event lasted only about 10-100 years (Sanford, 

1981). Metamorphic grade changes in borehole M-94, 

indicated by the disappearence of smectite + kaolinite 

and the appearence of illite + chlorite, suggested a temperature of only 150°-l80°C (Elders and others, 1980, 

their fig. 6). 

The time dependency of these mineral reactions 

suggests that any situation that increases reaction duration in a geothermal system will change the relationship 

of metamorphic grade to temperature. Thus, reactions 

should have come closer to completion in the deeper parts 

of a geothermal system than in the shallower parts, 

because the deeper parts have been heated longer and to 

a higher temperature (which promotes reaction). Increasing time also changes the apparent metamorphic grade. 

Hoffman and Hower (1979) observed that deeply buried 

Pliocene-Pleistocene shale has a much lower grade mineral assemblage than shale buried since the early Tertiary, 

although these Pliocene-Pleistocene rocks have attained 

the same temperature. 

The strong correlation between Rm and T b indicates that temperature is the most important variable in 

organic metamorphism (Barker, 1983), a conclusion that 

is also suggested by the fair to good agreement between 

Tvr and Tb at the endpoints of the range for each metamorphic: grade. Vitrinite reflectance geothermometry, 

however, is also flawed because of subjective selection 

of vitrinite for analysis (Barker and Pawlewicz, 1986), 

suppression of Rm in OM-rich rocks (Price and Barker, 

1984), and increasing bireflectance with rank that makes 

the determination of Rm progressively less precise 

(Stach and others, 1982). Further, vitrinite is only found 

in sedimentary systems, which rigidly restricts its 

application. 

CONCLUSIONS 

G1enerally, metamorphic grade is a less accurate 

geothermometer than temperature computed from Rm. 

Exceptions are the appearance of biotite and the disappearance of wairakite and prehnite. Low-temperature 

phase changes primarily in the clay minerals occur over 

wide temperature ranges and are unsuitable for accurate 

geothermometry. At best, metamorphic grades in geothermal systems are discontinuous geothermometers. 

They can be interpreted precisely only at points where 

reactions are changing the mineral assemblage. 

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79/10, 127 p. 

Barker, C. E., 1983, Influence of time on metamorphism of 

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519 p. 

Cerro Prieto, Salton Sea, and East Mesa Geothermal Systems 93 
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Elders, W. A., Olson, E. R., Hoagland, J. R., Barker, C. E., 

Johnson, P. D., and Collier, P., 1977, A comprehensive study 

of samples from geothermal reservoirs-Petrology and light 

stable isotope geochemistry of wells M48, M84, M90, M92, 

and Prian No. 1, in the Cerro Prieto geothermal field, Baja 

California, Mexico: University of California, Riverside, 

Institute of Geophysics and Planetary Physics, UCR/IGPP 

Report 77/37, 274 p. 

Elders, W. A., Hoagland, J. R., Olson, E. R., McDowell, S. D., 

and Collier, P., 1978, A comprehensive study of samples 

from geothermal reservoirs-Petrology and light stable 

isotope geochemistry of 23 wells in the Cerro Prieto geothermal field, Baja California, Mexico: University of California, Riverside, Institute of Geophysics and Planetary Physics, 

UCR/IGPP Report 78/26, 274 p. 

Elders, W. A., Hoagland, J. R., Valette-Silver, J. N., Williams, 

A. E., Barker, C. E., and Collier, P., 1979, A comprehensive 

study of samples from geothermal reservoirs-Petrology 

and geochemistry of surface and subsurface samples from 

the Cerro Prieto geothermal field, Baja California, Mexico: University of California, Riverside, Institute of 

Geophysics and Planetary Physics, UCR/IGPP Report 

79/27, 203 p. 

Elders, W. A., Hoagland, J. R., and Williams, A. E., 1980, 

Hydrothermal alteration as an indicator of temperature and 

flow regime in the Cerro Prieto geothermal field of Baja 

California: Geothermal Resources Council, Transactions, 

v. 4, p. 121-125. 

Ellis, A. J., 1979, Explored geothermal systems, in Barnes, H. 

L., ed., Geochemistry of hydrothermal ore deposits, 2d ed.: 

New York, Wiley and Sons, p. 632-678. 

Funiciello, R., Mariotti, G., Parotto, M., Preite-Martinez, M., 

Tecce, F., Toneatti, R., and Thri, B., 1979, Geology, mineralogy, and stable isotope geochemistry of the Cesano 

geothermal field: Geothermics, v. 8, p. 55-73. 

Hayashi, M., 1973, Hydrothermal alteration in the Otake 

geothermal area, Kyushu, Japan: Japan Geothermal Energy 

Association, v. 10, p. 9-46. 

Helgeson, H. C., 1968, Geologic and thermodynamic characteristics of the Salton Sea geothermal system: American 

Journal of Science, v. 266, p. 129-166. 

Hoagland, J. R., 1976, Petrology and geochemistry of hydrothermal alteration in borehole Mesa 6-2, East Mesa geothermal area, Imperial Valley, California: University of California, Riverside, Institute of Geophysics and Planetary 

Physics, UCR/IGPP Report 76/12, 90 p. 

Hoffman, J., and Hower, J., 1979, Clay mineral assemblages as 

low grade metamorphic geothermometers-Application to 

the thrust faulted disturbed belt of Montana, U.S.A, in 

Scholle, P. A., and Schluger, P. R., eds., Aspects of 

diagenesis: Society of Economic Paleontologists and 

Mineralogists Special Publication 26, p. 55-80. 

Hower, J., 1981, Shale diagenesis, in l..ongstaffe, F. J., ed., Short 

course in clays and the resource geologist: Mineralogical 

Association of Canada, p. 60-80. 

Kristmannsdottir, H., 1975, Hydrothermal alteration of basaltic 

rocks in Icelandic geothermal areas: Proceedings, second 

United Nations symposium on development and use of 

geothermal resources, v. 1, p. 444-446. 

94 Studies in Diagenesis 

Lyons, D. J., and van de Kamp, P. C., 1979, Subsurface geological 

and geophysical study of the Cerro Prieto geothermal field: 

Proceedings, second symposium on the Cerro Prieto 

geothermal field, Baj? California, Mexico, Sponsored by 

Comision Federal de Electricidad, Mexicalli, Mexico, 

October, 1979, v. 1, p. 197-186. 

McDowell S. D., and Elders, W. A., 1979, Geothermal metamorphism of sandstone in the Salton Sea Geothermal system, 

in Elders, W. A., ed., Geology and geothermics of the Salton 

Trough: Geological Society of America, Guidebook 7, San 

Diego, California, p. 70-76. 

McDowell, S. D., and Elders, W. A., 1980, Authigenic layer 

silicate minerals in borehole Elmore 1, Salton Sea geotherl 

mal field, California, USA: Contributions to Mineralogy 

and Petrology, v. 74, p. 293-310. 

Miller, K. R., 1980, Petrology, hydrothermal mineralogy, stable 

isotope geochemistry and fluid inclusion geothermometry 

of borehole Mesa 31-1, East Mesa geothermal field, 

Imperial Valley, California: University of California, Riverside, Institute of Geophysics and Planetary Physics, 

UCR/IGPP Report 80/1, 113 p. 

Muffler, L. J. P., and Doe, B. R., 1968, Composition and mean 

age of detritus of the Colorado River delta in the Salton 

Trough, southeastern California: Journal of Sedimentary 

Petrology, v. 38, p. 384-399. 

Muffler, L. J. P., and White, D. E., 1969, Active metamorphism of Upper Cenozoic sediments in the Salton Sea 

geothermal field and the Salton nough, southern California: Geological Society of America Bulletin, v. 80, 

p. 157-82. 

Nakobo, S. 1., and Piip, B. 1., 1961, Recent metamorphism of 

volcanic rocks in the region of the Pauzhetsk Hot Springs 

(Kamchatka): Akademiya Nauk S.S.S.R. lfudy Lab 

Vulkanologiya, v. 19, p. 99-114. [In Russian]. 

Price, L. C., and Barker, C. E., 1984, Suppression of vitrinite 

reflectance in amorphous-rich kerogen-A major unrecognized problem: Journal of Petroleum Geology, v. 8, 

p. 59-84. 

Rettke, R. C., 1976, Clay mineralogy and clay mineral distribution patterns in Dakota Group sediments, northern Denver 

basin, eastern Colorado and western Nebraska: Cleveland, 

Ohio, Case Western Reserve University, Ph. D. dissertation, 

147 p. 

Sanford, S. J., 1981, Dating thermal events by fission track 

annealing, Cerro Prieto geothermal field, Baja California, 

Mexico: University of California, Riverside, Institute of 

Geophysics and Planetary Physics, UCR/IGPP Report 

8114, 105 p. 

Schiffman, P., Elders, W. A., Williams, A. E., McDowell, S. D., 

and Bird, D. K., 1984, Active metasomatism in the Cerro 

Prieto geothermal system, Baja California, Mexico-A 

telescoped low pressure, low temperature metamorphic 

facies series: Geology, v. 12, p. 12-15, and p. 255. 

Seki, Y., 1972, Lower grade stability limit of epidote in the light 

of natural occurrences: Journal of the Geological Society 

of Japan, v. 78 p. 405-413. 

Sheridan, M., and Maisano, M., 1975, Zeolite and sheet silicate 

zonation in a late Tertiary geothermal basin, near 

Hassayampa, central Arizona: Proceedings, second United 

Nations symposium on development and use of geothermal resources, v. 1, p. 597-609. 
Shimazu, M., and Yajima, J., 1973, Epidote and wairakite in 

drill cores at the Hachimanti geothermal area, northeastern Japan: Journal of Japan Association of Mineralogy, 

Petrology, and Economic Geology, v. 68, p. 363-371. 

Sigvaldason, G. E., 1962, Epidote and related minerals in two 

deep geothermal drill holes, Reyjavik and Hveragerdi, 

Iceland: U.S. Geological Survey Professional Paper 450-

E, p .. 77-79. 

Stach, E., Mackowsky, M.-Th., Teichmuller, M., Taylor, 

G. H., Chandra, D., and Teichmuller, R., 1982, Stach's 

textbook of coal petrology, 3d ed.: Berlin, Gebruder Borntraeger, 535 p. 

Steiner, A., 1977, The Wairakei geothermal field, North Island, 

New Zealand: New Zealand Geological Survey Bulletin 90, 

36 p .. 

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geothermal area, northeast Japan: Geological Survey of 

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Tissot, B. P., and Welte, D. H., 1984, Petroleum formation 

and occurrence, 2d ed.: New York, Springer-Verlag, 

699 p. 

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and shales; diagenesis; technical report: Office of Nuclear 

Waste Isolation Technical Report 21, 176 p. [Available from 

National Technical Information Service, Springfield, VA 

22161.] 

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springs systems and depth of formation of propylitic 

epidote in epithermal ore deposits: U.S. Geological Survey 

Professional Paper 450E, p. 80-84. 

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Annual Review of Earth and Planetary Science, v. 2, 

p. 179-212. 

Cerro Prieto, Salton Sea, and East Mesa Geothermal Systems 95 

Origin and Diagenesis of Cretaceous 

Deep-Sea, Organic-Carbon-Rich 

Lithofacies in the Atlantic Ocean 

By WALTER E. DEAN, U.S. Geological Survey, Box 25046, Denver Federal 

Center, MS 940, Denver, Colorado 80225, and 

MI~:HAEL A. ARTHUR, Graduate School of Oceanography, University of 

Rhode Island, Narragansett, Rhode Island 02882 

CONTENTS 

Abstract 99 

Introduction 99 

Acknowledgments 100 

Oxidation-reduction cycles in the Cretaceous Atlantic Ocean 100 

Description of the cycles 100 

Cycle periods 103 

Diagenetic processes and products 103 

Sulfur diagenesis 103 

Carbonate diagenesis 104 

Phosphate diagenesis 108 

Trace eh!ments in organic-carbon-rich strata 108 

Models of deposition 118 

Bottom-water oxygenation 118 

Supply of organic carbon to sediments 118 

Conclusions 122 

References cited 123 

FIGURES 

1. Plots of relative concentrations of ions and gases in pore water versus depth 

within anoxic sediment 100 

2. Continental reconstruction map showing locations of Deep Sea Drilling Project sites where Cretaceous organic-rich strata have been recovered 101 

3. Summary stratigraphic column for strata recovered from Deep Sea Drilling 

Project sites in North Atlantic Ocean 102 

4. Plot of concentrations of organic carbon and total sulfur in Cretaceous strata 

rich in organic carbon from Atlantic Deep Sea Drilling Project sites 105 

5. Plots showing relationship between o13C and total dissolved C02 in pore 

waters of anoxic sediments during sulfate reduction and methane 

generation 106 

6. Diagram showing ranges and means of oxides and trace element concentrations in claystone and black shale from Deep Sea Drilling Project site 530 110 

7. Diagram showing ranges and means of oxides and trace element concentrations of interbedded black shale and green claystone from Deep Sea Drilling 

Project site 367 112, 113 

8. Plots of concentrations versus depth for V, Cu, Ni, and Ba in black shale 

from cores, Deep Sea Drilling Project site 530 114 

9. Illustrations showing how influx of terrigenous organic matter, surface-water 

organic productivity, upwelling, and turbidity currents can contribute to 

accumulation and redeposition of sediments containing varying amounts of 

organic carbon 120, 121 

F. A. Mumpton, eel., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 97 
TABLES 

1. Generalized reactions involving organic matter in pore water and their 

products 99 

2. Average or representative chemical analyses for modern organic-carbon-rich 

sediments, oxic deep-sea sediments, and Cretaceous black shales and 

interbedded strata 109 

3. Mass accumulation rates of selected components in red, green, and black 

beds at Deep Sea Drilling Project site 530 116, 117 

4. Concentrations of major-element oxides and trace elements and mass 

accumulation rates for modern and ancient Black Sea sediments 119 

98 Studies in Diagenesis 
Abstract 

Deep-water sedimentary sequences of Cretaceous age, 

rich in organic carbon (OC), have been recovered at many Deep 

Sea Drilling Project (DSDP) sites in the Atlantic Ocean. The OCrich strata occur within cyclic interbeds of shale, marlstone, 

and (or) limestone which have cycle periods that range from 

about 20,000 to 100,000 years and average 40,000 to 50,000 

years. These cycles may be due to the effects of climate change 

on the influx of terrigenous organic matter. The variations in 

supply and (or) preservation of organic matter had pronounced 

effects on the potential of the sediment for subsequent diagenetic changes and geochemical partitioning in adjacent beds. 

Many beds rich in OC contain authigenic iron sulfides, 

barite, gypsum, dolomite, siderite, rhodochrosite, and (or) other 

Mg-, Fe-, and Mn-rich carbonate minerals. Authigenic gypsum 

and barite, precipitated near the lower boundaries of some OCrich layers, may indicate later oxidation of dissolved sulfide. 

The geochemistry and stable isotope composition of the 

authigenic carbonate minerals suggestthat they were formed 

under anoxic conditions during and after sulfate reduction, and 

that some formed from dissolved C02 in equilibrium with 

methane. Dolomite is the most common carbonate mineral in 

OC-rich strata. Siderite is common in sequences that have 

relatively low concentrations of OC or high sedimentation rates. 

Manganese carbonates are rare, probably because manganese 

is highly mobile under reducing conditions and diffuses out of 

the anoxic sediment. Although phosphate minerals have not 

been reported from Cretaceous OC-rich strata, some strata contain as much as 2 percent P20 5, probably in the form of 

phosphatic fish debris. 

Many trace elements, such as Cr, Ni, V, Cu, Zn, and Mo, 

are concentrated in OC-rich lithologies relative to interbedded 

OC-poor lithologies. The association of high trace-element concentrations with organic matter may be the result of a concentration of these elements by organisms and (or) by chemical 

sorption and precipitation processes. Detailed trace-element 

profiles in a thick sequence of OC-rich strata from the South 

Atlantic Ocean suggest that the source of the organic matter 

may have been the main controlling factor on the concentration of trace elements; strata containing predominantly marine 

organic matter tend to have higher trace-element concentrations than those containing predominantly terrestrial organic 

matter. 

Most depositional models that have been used to explain 

the accumulation of the OC-rich strata imply that reducing conditions in the sediments (and therefore an increased degree of 

preservation of organic matter) were the result of anoxic or nearanoxic conditions in oceanic bottom waters or in a midwater 

oxygen-minimum zone. Evidence from several DSDP sites in 

the Atlantic, however, indicates that anoxia or near-anoxia probably occurred within midwater oxygen-minimum zones and 

possibly, under extreme conditions, throughout much of the 

bottom-water mass. Although bottom-water anoxia may have 

been reached during periods of accumulation of OC-rich strata, 

it was not necessarily a cause for accumulation of these strata. 

The main reason for the accumulation of OC-rich strata was 

an increase in the relative amount of organic debris being 

deposited. This organic debris was derived from continental 

margins and transported to slope and basinal sites by turbidity 

currents. 

INTRODUCTION 

Deep-water sedimentary sequences of Cretaceous 

age that contain beds rich in organic carbon (OC) (more 

than 1 percent) have been recovered at many Deep Sea 

Drilling Project (DSDP) sites, particularly in the Atlantic Ocean (Tissot and others, 1980; Summerhayes, 1981; 

Arthur, Dean, and Stow, 1984; Authur and Dean, 1986). 

Deposition of OC-rich strata in the Atlantic began in the 

Early Cretaceous and continued intermittently to the present. Most of these sequences display a marked cyclicity 

in amount of OC that is manifested as interbedded lighter 

and darker shale, madstone, and (or) limestone. Lower 

Cretaceous (Neocomian) OC-rich strata occur within carbonate sequences, but most of the other Cretaceous strata 

in the Atlantic are carbonate poor and clay rich, resulting 

in cyclic interbedding of red, green, and black claystone 

or shale. 

Organic matter is preserved in relatively high concentrations in deep-sea, OC-rich strata because of a high 

rate of supply of organic matter, an enhanced preservation under anoxic conditions, an enhanced preservation 

due to high sedimentation rate, or some combination of 

these factors. Once the organic matter reaches the sediment, some part of it usually is decomposed. Various 

oxidizing agents are reduced in a series of oxidationreduction reactions that are largely mediated by bacteria. 

Some of the possible oxidation-reduction reactions, highly 

simplified, are shown in table 1, and their general sequence 

and distribution with depth below the sediment-water 

interface is illustrated in figure 1. The byproducts of these· 

reactions and of the decomposition of organic matter are 

Table 1. Selected important generalized reactions involving 

organic matter in pore water and their products 

[Generalized equations from Berner (1980) and Curtis (1980) listed in 

order of occurrence in anoxic pore waters. CH2

0 is a generalized formula for organic matter] 

CH2

0 + 0

2 ~ C02 + H

2

0 (or H+ + HCO;) 

5CH2

0 + ijNO; ~ 2N2 + ijHCO; C02 +3H2

0 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 99 
MILLIMETERS 

TO 

CENTIMETERS 

CENTIMETERS 

TO 

METERS 

:I: 

t 

w 

Cl 

METERS 

TO 

TENS 

OF 

METERS 

LOW 

"NUTRIENT', 

RELATIVE 

CONCENTRATION 

/:..._Mn2+ 

I , i/ .r/ 

,/I 

/ I I I 

/--·Fe2+ 

I 

I 

I 

SOl-, HzS, CH4, 

RELATIVE CONCENTRATION 

HIGH LOW -------------7 HIGH 

I 

I 

I 

I 

I 

I 

I 

I 

/ 

I 

/ 

/ / / 

I 

I 

I 

I 

r-sor 

I 

I 

I 

OXYGEN 

DRAW DOWN 

+ 

NITRATE REDUCTION 

l 

SULFATE 

REDUCTION 

METHANOGENESIS 

Figure 1. Profiles of relative concentrations in pore water versus depth within anoxic sediment for 0 2, N03, Mn2+, Fe2+, PO~-, 

alkalinity + total C02 (TC02), SO~-, H25, and CH4 • Zones of oxygen drawdown and nitrate reduction, sulfate reduction, and 

methanogenesis are indicated. 

various reduced metal species, ammonia, phosphate, 

sulfide, and large amounts of carbon dioxide and bicarbonate that contribute to increases in the alkalinity of the 

interstitial waters (Berner, 1980). 

Based upon published data collected by many 

workers, the present paper attempts to: (1) outline the 

typical chemical and mineralogical compositions of OCrich facies of Cretaceous age recovered at DSDP sites in 

the North and South Atlantic Oceans and explain how 

diagenesis has affected these compositions; (2) discuss the 

extent to which environment of deposition determines the 

supply of OC to deep-water sediments and therefore predetermines the extent to which diagenetic reactions and 

enrichment of certain trace elements can occur; and (3) 

contrast the chemistry of OC-rich facies from different depositional settings and outline briefly the potential effects 

of organic carbon accumulation on paleocean chemistry. 

100 Studies in Diagenesis 

Acknowledgments 

We thank Joseph Hatch, Douglas Kirkland, Philip 

Meyers, Seymour Schlanger, Harry Tourtelot, and 

Douglas Waples for helpful reviews of the manuscript. 

OXIDATION-REDUCTION CYCLES IN THE 

CRETACEOUS ATLANTIC OCEAN 

Descriptions of cycles 

The oldest OC-rich strata in the North Atlantic 

Ocean, recovered at more than one site, are interbeds of 

laminated dark-olive to black marlstone within bioturbated white to light-gray limestone of Tithonian to Neocomian age recovered from DSDP sites 101, 105, 367, 387, 

391, 417, 418, 534, and 603 (figs. 2 and 3; Jansa and 
w 

c 

::::> 

1-

~ 15° 

0 

w 

...J 

<( 

a.. 

603.&.&105 

390.& ... 392 .&387 

391 .. 534 .&386 

01 ... 417/418 140 

... ... 4/5 137... 138 

100 .&368 

384 

... ... 135 

136 ~.&120 

370/415 

.&369 

ANGOLA/SRAZIL 

BASIN 

.&328 

511.327/330 

\ FALKLAND 

PLATEAU 

Figure 2. Locations of Deep Sea Drilling Project sites (solid triangles) where Cretaceous organic-carbon-rich strata have 

been recovered plotted on a continental reconstruction at 100 ± 10 million years before present (modified from Sclater 

and others, 1977). 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 101 
(/) 

a: 

w 

..... 

w 

~ 

~ 

:i 

b: 

w 

c 

~ 

0 

~ 0 

a:l 

a:l 

:::> 

(/) 

LITHOLOGY AGE 400-r---------,---------------~--~ --;--__: • .:.._,----=._ 

~ :._--=-____:_ :-_.:.__ 

c-:- --:-~-=--=---:--=-- - ~--=-_:_ :-__:_--: _:_ ---·-·- . -·-·-· 

. -·----

- ·-·-· ·-- ·-· 

·--. -· -. . - -·- . - . -

600 - --:-___:__----=-~ ---=-

Multicolored 

silty clay 

~~~_:__-=..:__-=+---(gradational)--- ~ 

- - f---~-

ltl 

E 

0 

c 

Q) 

.. .. .. u 700 -Jiiiii Biack shale 1------

- ··...:..= Aibian 

1------

early Albian 

800 ~=-~-§~~~l----- -~~ptia~ 

- -_ -= early 

--- -- -----~-

-=I-=rT-=-

-

- +- + 

Variegated claystone AptianBarremian 

-------r----

c 

<0 ·e 

~=L:o::::r::::::c:::q Light gray nannog-~~2-~-~-a limestone interbedded with olive-gray marlstone 

0 

u 

0 

Q) 

z 

1 000 1-------'----'--....L ..L. --

1-1-1-1-1-

11 00 ---,.....,=:.,.-=::;--::::::r=--FI ----

-1-1-1-1-

Reddish brown nannobearing argillaceous 

limestone 

Late 

Jurassic 

Figure 3. Summary stratigraphic column for strata recovered 

from Deep Sea Drilling Project sites in North Atlantic Ocean 

(modified from Jansa and others, 1979). 

others, 1978, 1979; Dean and Gardner, 1982; Arthur and 

Dean, 1986). Upper Jurassic (Callovian-Oxfordian) strata 

containing as much as 3 percent OC were recently 

recovered at DSDP site 534 in the Blake-Bahama basin 

(Sheridan and others, 1982). The Neocomian laminated 

dark marlstone beds generally contain more than 5 per.; 

cent OC, and one bed at site 367 in the Cape Verde basin 

off northwest Africa contains 33 percent OC (Dean and 

others, 1978; Dean and Gardner, 1982). 

The end of deposition of the Neocomian carbonates 

apparently was caused by a sudden rise in the carbonate 

compensation depth (CCD) (Thierstein, 1979; Arthur and 

102 Studies in Diagenesis 

Dean, 1986), and the remainder of the section in the North 

Atlantic through the Eocene is dominated by multicolored 

clay-rich lithologies. At most DSDP sites in the North 

Atlantic, the Neocomian carbonates are overlain by the 

so-called middle Cretaceous (Aptian-Albian to Turonian) 

black shale facies, but at some sites the two are separated 

by a thin (less than 10 m) unit of interbedded red and 

green claystones of late Aptian to early Albian age (fig. 3; 

Jansa and others, 1978; Dean and Gardner, 1982). 

The middle Cretaceous so-called ''black shale'' 

facies represents the main period of accumulation of 

organic matter in the Atlantic, and at most DSDP sites 

it consists of interbedded, green and black clay-rich 

lithologies (see reviews by Arthur, 1979a; Tucholke and 

Vogt, 1979; Thierstein, 1979; Arthur and Natland, 1979; 

Tissot and others, 1979, 1980; Dean and Gardner, 1982; 

Weissert, 1981; Arthur, Dean, and Stow, 1984; Arthur 

and Dean, 1986). This facies has been recovered at DSDP 

sites 4, 5, 101, 105, 135, 137, 138, 144, 367, 368, 369, 

370,386,387,391,398,400,402,415,416,417,418,534, 

and 603 in the North Atlantic Ocean, and at sites 327, 

328, 330, 356, 357, 361, 363, 364, 511, and 530 in the 

South Atlantic Ocean (fig. 2). The green and black 

claystone sequences are overlain by Upper Cretaceous 

green and red claystone. Concentrations of OC in the 

black lithologies in the North Atlantic range from less 

than 1 percent to 37 percent and usually are more than 

c percent (Lancelot and others, 1978; Tissot and others, 

1979, 1980; Summerhayes, 1981; Dean and Gardner, 

1982; Arthur, Dean, and Stow, 1984a; Arthur and Dean, 

1986). Concentrations of OC generally are lower in black 

lithologies from western North Atlantic sites than in those 

from eastern North Atlantic sites. True black shale 

usually amounts to less than half of the so-called ''black 

shale" facies. 

The organic detritus in OC-rich layers may be from 

autochthonous marine planktonic production or from 

terrestrial organic debris. According to Tissot and others 

(1979, 1980), and Summerhayes (1981), autochthonous 

marine organic matter (that is, organic matter containing type II kerogen of Tissot and others, 1974) in continuous sections of Lower Cretaceous rocks in the North 

Atlantic is found only in the Cape Verde basin (site 367), 

on the Cape Verde rise (site 368), and on the Demerara 

rise (site 144). Everywhere else in the North Atlantic, terrestrial organic matter with admixtures of highly 

degraded, residual, marine organic matter dominates the 

Lower Cretaceous OC-rich strata (Tissot and others, 

1979, 1980; Summerhayes, 1981). The Cretaceous OCrich strata in the southeastern North Atlantic, therefore, 

appear to have more in common with similar strata in 

the South Atlantic, where most of the organic matter is 

derived from marine sources, than with equivalent strata 

in the rest of the North Atlantic (Tissot and others, 1979, 

1980; Summerhayes, 1981; Gilbert, 1984). 
In the South Atlantic sequences, OC concentrations average about 3-4 percent (values as high as 21 

percent have been reported for the Aptian-lower Albian) 

at site 361. in the Cape basin (Arthur and Natland, 1979; 

Gilbert, 1.984). They average about 4-5 percent (values 

as high as 24 percent have been reported for the lower 

Albian) at site 364 in the eastern Angola basin. Organic 

carbon concentrations average about 4 percent in the 

Hauterivian-lower Aptian strata at site 511 in the Argentine basin, and are less than one percent in the AptianAlbian (Dick and others, 1983; Gilbert, 1984). Black 

shale beds of Albian to Coniacian-Santonian age at 

site 530 contain as much as 19 percent OC and contain 

overall an average of 5.4 percent OC, as compared with 

less than 0.5 percent OC in intercalated beds of red and 

green claystone (Deroo and others, 1984; Meyers and 

others, 1984). Most of the organic matter in black shales 

at site 530 is of marine origin (Deroo and others, 1984; 

Gilbert, 1984). Much of the organic matter at site 361 

in the Cape basin apparently is redeposited terrestrial 

organic matter, whereas the organic matter in black shales 

at sites 364 and 511 is mostly autochthonous marine 

organic matter (Gilbert, 1984). In spite of these differences in sources of organic matter, the OC-rich strata 

in all Atlantic basins are interbedded with bioturbated, 

OC-poor strata. 

Cycle periods 

Calculations of periodicities of diagenetic 

oxidation-reduction cycles, ranging in age from Early 

Cretaceous to Tertiary, are subject to several assumptions, particularly regarding sedimentation rate. These 

calculations, however, show that the oxidation-reduction 

cycles generally have periods ranging from 20,000 to 

100,000 years and averaging about 40,000-50,000 years 

(Dean and others, 1978; Arthur and Fischer, 1977; 

Arthur, l979a; McCave, 1979; Arthur, Dean, Bottjer, 

and Scholle, 1984). The cyclicity of black shale deposition at site 530 in the southern Angola basin (see discussion in section on Models of Deposition) is much 

more variable, ranging from 9,600 to 393,000 years, 

with most values between 40,000 and 220,000 years. The 

more-regular periodicities at other sites are similar to 

periodicities of carbonate dissolution, dilution, seasurface temperature fluctuations, and other parameters 

that are <:orrelated with climatic changes in the Neogene 

and Quaternary (Broecker and Van Donk, 1970; 

Ruddiman, 1971; Matthews, 1974; Gardner, 1975; Hays 

and others, 1976; Dean and others, 1981; Gardner, 1982; 

Gardner and others, 1984). The cyclic patterns of 

sedimentation that are so distinctive of the entire 

stratigraphic sequence in the North and South Atlantic 

Oceans also may be a response to climatic forcing similar 

to Quaternary analogues that are largely driven by 

climate-induced changes in ice volume and bottom-water 

production and flow. Cycles of climate-induced carbonate 

dissolution in the North Atlantic apparently extend 

through the Neogene and possibly the entire Tertiary 

(Dean and others, 1978, 1981). Matthews and Poore 

( 1980) presented isotope evidence suggesting that changes 

in sea level related to ice volumes were important 

throughout the Tertiary. 

Regardless of the exact mechanism involved in the 

formation of the sedimentary cycles, the pronounced 

cyclicity of the sedimentary section in the Atlantic, particularly in the middle Cretaceous, suggests that there 

probably were cyclic variations in the influx of terrigenous 

organic matter and detrital material, the sea-surface production of organic matter and CaC03, the concentration 

of dissolved oxygen in deeper water masses, and (or) the 

preservation of CaC03• 

DIAGENETIC PROCESSES AND PRODUCTS 

Sulfur diagenesis 

Sulfur is incorporated into anoxic, OC-rich sediments mainly through bacterial sulfate reduction. This 

process is relatively well known (see for example, Berner, 

1971, 1974, 1978, 1980, 1982, 1984; Goldhaber and 

Kaplan, 1974, 1975, 1981; Goldhaber and others, 1977; 

Jorgensen, 1977; Oremland and Taylor, 1978) and 

involves the utilization of pore-water sulfate as an electron donor for the oxidation of sedimentary organic matter by bacteria in the absence of dissolved molecular 

oxygen. Sulfate reduction usually proceeds after all other 

oxidants have been consumed. For example, Froelich and 

others (1979) found that oxidants in pelagic sediments in 

the eastern equatorial Atlantic were consumed in the 

following order of decreasing energy produced per mole 

of organic carbon oxidized: 0 2 > Mn-oxides > N02 

> Fe-oxides >SO~-. Most OC (at least 90 percent) in the 

pelagic rain is oxidized by 0 2 before reaching the 

sediment-water interface, but the secondary oxidants are 

of great importance in regulating the burial and preservation of OC within the sediments (Bender and Heggie, 

1984). In most anoxic sediments, sulfate accounts for 

most of the oxygen consumption and decomposition of 

organic matter (Berner, 1984). Jorgensen (1977) found 

that sulfate reduction accounted for 53 percent of the 

mineralization of organic matter in anoxic sediments of 

a fjord in Denmark. He also found that only 10 percent 

of the sulfide produced by sulfate reduction was 

precipitated as metal sulfides and that the rest diffused 

to the surface where it was reoxidized. Berner (1978) 

suggested than an average of 0.2-0.4 weight percent of 

the OC is lost from sediments during sulfate reduction, 

but this amount may vary depending upon the rate of 

sedimentation and the amount of sulfate that is available 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 103 
by diffusion into the sediment, as well as upon the type 

and reactivity of organic matter present (Toth and Lerman, 1977; Berner, 1980; Goldhaber and Kaplan, 1980). 

The reduced sulfur is free to diffuse upward to the 

sediment-water interface, where it can be oxidized by 

bacteria, or to diffuse to sites of sulfide precipitation, a 

major sink for sulfur in anoxic sediments. Modern anoxic 

sediments contain as much as several weight percent total 

sulfur. The weight ratio of sulfur to OC in anoxic 

sediments averages 0.4, but varies significantly from this 

value (Berner, 1982; Leventhal, 1983; Berner and 

Raiswell, 1983). Based on data for sulfur:carbon (S:C) 

ratios from the Black Sea and Devonian black shales of 

eastern North America, Levanthal (1983) suggested that 

S:C ratios greater than 0.4 in black shales, with a slope 

intersect at a positive value of sulfur at zero carbon, 

signify deposition under an anoxic water column, rather 

than having only anoxic sediment pore waters. 

A plot of OC versus S for those samples of Cretaceous strata for which both OC and S analyses are 

available (fig. 4) suggests that a S:C ratio of about 0.4 

is a reasonable average, similar to the S:C ratio in most 

modern anoxic sediments (Berner and Raiswell, 1983). 

Kendrick (1979) measured concentrations of total sulfur 

as much as 4.3 weight percent in OC-rich strata from 

DSDP sites 386 and 387. As much as 40 percent of this 

sulfur apparently is in the organic fraction; the rest is in 

pyrite. Kendrick interpreted this sulfur enrichment as 

evidence for the deposition and preservation of OC under 

an anoxic water column. Low concentrations of S and 

OC were interpreted as an indication of oxic bottomwater conditions. 

Metal sulfides of varying morphologies are present 

in nearly all the Cretaceous OC-rich strata recovered from 

DSDP sites in the Atlantic. Pyrite is the most commonly 

reported sulfide mineral and occurs as framboids, testfillings of foraminifera, micronodules, replacements or 

molds of radiolarian tests, and large discrete nodules and 

lenses. Localized concentrations of pyrite and (or) marcasite are common at contacts between black shale and 

underlying green claystone beds (Dean and Schreiber, 

1978; Arthur, Dean, Bottjer, and Scholle, 1984a; Dean 

and others, 1984). In fact, sulfide sulfur appears to be 

highly mobile in sequences of interbedded OC-rich and 

OC-poor strata (Berner, 1969, 1970), and green claystone 

beds commonly contain as much total sulfur as adjacent 

beds of black shale. 

Authigenic sulfate minerals such as gypsum and 

barite also occur within or near the boundaries of some 

black shale layers. Small (usually less than 10 mm) 

euhedral crystals of barite are common in modern oxic 

deep-sea sediments, particularly in the equatorial Pacific 

Ocean (Cronan, 1974; Dean and Schreiber, 1978; Church, 

1979), but larger concentrations of barite as nodules, 

104 Studies in Diagenesis 

lenses, rosettes, and laminae are not common in modern 

deep-sea sediments. Such larger aggregates have been 

described from Cretaceous OC-rich strata from several 

DSDP sites along the continental margin of northwest 

Africa (Dean and Schreiber, 1978) and on the Vigo Seamount off Portugal (Arthur, 1979a). Dean and Schreiber 

(1978) observed that the barite was always associated with 

organic matter, usually just below an OC-rich layer. They 

also noted that the barite occurred just below major 

unconformities and in sections in which the salinity of 

the interstitial water increased downward, probably 

related to suspected underlying Jurassic evaporites. Dean 

and Schreiber (1978) concluded that the main source of 

barium probably was organic matter, but that the sulfate 

may have come from the diffusion of seawater-sulfate 

into the sediments, the oxidation of preexisting pyrite, 

or the upward diffusion of sulfate from the dissolution 

of evaporites at depth. 

Gypsum has been reported from DSDP cores in a 

variety of lithologies (Nesteroff, 1973; Briskin and 

Schreiber, 1978), particularly from anoxic sediments. 

Gypsum also is intimately associated with OC-rich continental slope sediments from off the coast of Southwest 

Africa (Seisser and Rogers, 1976). 

Carbonate diagenesis 

Many Cretaceous deep-sea black shale sequences 

contain dolomite, siderite, rhodochrosite, and other Mg-, 

Fe-, and Mn-rich carbonates. These carbonate minerals 

also are found in Holocene marine sediments where they 

are thought to be the products of anaerobic decomposition of organic matter and even to be associated with 

methane generation (Suess, 1979; Hein and others, 1979; 

Pedersen and Price, 1982). In fact, the abundance of 

siderite in some mid-Cretaceous black shales can be taken 

as evidence for complete reduction of sulfate and 

mineralization of sulfide (Berner, 1971 ). Methane generation is common in anaerobic, OC-rich sediments after 

sulfate reduction has ceased. Methanogenesis, like sulfate 

reduction, is a bacteria-mediated process that reduces OC, 

acetate, and C02 in pore waters and converts them to 

methane (see, for example, Claypool and Kaplan, 1974; 

Martens and Berner, 1974; Oremland and Taylor, 1978). 

During the production of methane, a substantial fractionation of carbon isotopes (50-90 per mil, o13C) takes 

place between methane and C02, leaving the residual 

C02 enriched in 13C (Presley and others, 1972; Rice and 

Claypool, 1981; Fuchs and others, 1979). During methane 

production, values of o13C in total C02 in pore waters 

may progressively shift from highly negative to highly 

positive (fig. 5). Carbonate minerals precipitated from 

pore waters during methanogenesis therefore may have 
1--

z 

w 3 u 

a: 

w 

0.. 

z 

a: 

::::> 

u.. 

...J 

::::> 2 

(/) 

...J 

<( 

1--

0 

1--

0 

• 

0 

• 

• 

0 

• • / .1 

• 

2 3 4 5 6 7 

ORGANIC CARBON, IN PERCENT 

DEEP SEA DRILLING PROJECT SITES-ATLANTIC OCEAN 

Lower Cretaceous: Groups 1, 2, and 3 

of Brumsack ( 1 980) 

Upper Cretaceous (1) and Cenozoic (2) 

of Scharf ( 1 980) 

Cretaceous black shales (sites 386 and 

387; Kendrick, 1979) 

Upper Cretaceous black shales (site 

36 7, core 1 7; Dean and Gardner, 1982) 

Upper Cretaceous green claystones (site 

367, core 17; Dean and Gardner, 1982) 

Upper Cretaceous green and red claystones 

(site 36 7, cores 1 5 and 16; Dean and Gardner, 1982) 

Lower Cretaceous olive marlstone (site 36 7, 

cores 25 and 26; Dean and Gardner, 1982) 

Lower Cretaceous white limestones (site 36 7, 

cores 25 and 26; Dean and Gardner, 1982 

• 

• 

8 9 

Figure 4. Concentrations of organic carbon and total sulfur in Cretaceous strata rich in organic carbon from 

Atlantic Deep Sea Drilling Project sites. 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 105 
TOTAL DISSOLVED CARBON (TDC), 

IN MILLMOLES PER LITER 

030 35 40 45 

(/) TDC 

a: 813cco2 w 

1-

w 

~ 100 1-

z 

w 

u 

~ 

ui 

a: 

0 

u 

~ 200 

:I: 

1-

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w 

0 

A 

30010 -2 

813Cco2. in per mil (PDB) 

20 

•• • 

• • 

iii • . ,. c 

~ 0 ®Seawater • .E 

iii • a. 

·= • 

0 

u 

u -20 Eo 

B -40L-----~·~--~--------~--------~-L----~ 0 20 40 60 

TOTAL DISSOLVED CARBON, IN MILLIMOLES PER LITER 

Figure 5. Relationships between o13C and total dissolved C02 

(TDC) in pore waters of anoxic sediments during sulfate reduction and methane generation. A, Core 4 from anoxic Saanich inlet 

(data of Nissenbaum and others, 1972) showing rapid change to 

heavier values of o13C in TDC in pore waters with increasing TDC 

during methane generation. 8, Plot of o13C in TDC in pore water 

versus concentration of TDC for all data from Saanich inlet 

reported by Nissenbaum and others (1972) showing isotopically 

light values of o13C at low concentrations of TDC due to decomposition of organic matter during sulfate reduction and a gradual 

increase in o13C as methane generation proceeds. C, Example of 

deeper methane generation at Deep Sea Drilling Project site 147 

in the Cariaco trench (Presley and others, 1973) where values of 

o13C in pore-water TDC increases over tens of meters ratherthan 

over several meters, as in A or B. 

106 Studies in Diagenesis 

0 

20 

40-

60 

(/) 

a: 

w 

1-

w 

~ 

~ 

ui 

a: 

0 80 u 

~ 

:I: 

1-

0.. 

w 

0 

100 

120 

140 

c 160L-------------~------------~------------~ 

-20 -10 0 10 

813C, in per mil (PDB) 
a large range of o13C values. Diffusion and oxidation of 

methane within more oxic sediments may complicate carbon isotope values and may lead to very light (highly 

negative) values of one in total COr 

During periods when OC is being fixed in the 

sediment at unusually high rates (for example, during 

the middle Cretaceous), the oceanic surface-water and 

total-dissolved-carbon reservoirs may become isotopically heavier because of the net transfer of isotopically 

light OC to the sediment. Consequently, any biogenic 

carbonate formed in surface waters during these times 

would be: isotopically heavy (more positive onC) (Scholle 

and Arthur, 1980). For most of the Cretaceous, the one 

of carbonate carbon apparently was between + 1.0 and 

+4.5 per mil (relative to the PDB standard). In anoxic 

sediments that are relatively depleted in carbonate, 

diagenetic reactions control the carbon isotope composition of carbonate. 

As the decomposition of organic matter proceeds 

from aerobic degradation through sulfate reduction, the 

total alkaJinity of the pore waters increases (Berner, 1980). 

Ammonia and phosphate liberated from organic matter 

and sulfide species resulting from sulfate reduction contribute to the increase in alkalinity. Concentrations of total 

dissolved carbon (TDC) increase in proportion to the 

amount of oxidation of OC. The dissolved carbon added 

to intersltitial water has a carbon isotopic composition 

close to that of the organic matter (one of about -23 per 

mil, on average). The initial TDC content in bottom waters 

in the open ocean today ranges from 2.2 to 2.5 mmol/L, 

and ranges from about -0.5 to + 1.5 per mil in one 

(PDB). The addition of dissolved isotopically light carbon during sulfate reduction causes pore waters to have 

more negative values of one, which can approach -20 

per mil as the pore-water TDC values increase to as much 

as 35 to 40 mmol/L. 

Concentrations of dissolved Fe, Mg, and Mn also 

increase in anoxic pore waters (see, for example, Manheim, 

1976) which, with increased alkalinities, can lead to saturation with respect to a variety of carbonate minerals 

(Presley and others, 1972). For example, carbon in 

diagenetic carbonate minerals in highly anoxic sediment 

from th(;: Baltic Sea has highly negative values of one 

(about -13 per mil; Suess, 1979). Hein and others (1979) 

measured values of one ranging from about -7 to about 

-21 per mil (average of about -17 per mil) in diagenetic 

carbonates (high-Mg calcite, l)rotodolomite, siderite, 

rhodochrosite, low-Mg calcite, and manganosiderite) in 

diatom-rich sediment in DSDP cores from the Bering Sea. 

Because of the highly negative values of one, Hein and 

others (1979) concluded that the carbon in the diagenetic 

carbonates was derived from the oxidation of organic matter under anaerobic conditions with the production of 

methane. 

Dolomite is the most common carbonate mineral 

in OC-rich sediments in general; it typically occurs as 

isolated rhombs less than 62 p.m in diameter, but also may 

occur as discrete laminae~ beds, and nodules. Sources of 

magnesium for dolomite formation may be from 

magnesium dissolved in interstitial waters, magnesium 

desorbed from clay minerals or biogenic silica (see, for 

example, Donnally and Merrill, 1977), and (or) upward 

diffusion from underlying evaporite beds. The latter may 

have been a source of magnesium for the dolomite found 

associated with barite in OC-rich strata at DSDP sites 364, 

369, and 370 (Dean and Schreiber, 1978; Bolli and others, 

1978). The chlorinity of interstitial waters at these three 

sites increase" downward in response to underlying 

evaporites. 

Friedman and Murata (1969) demonstrated that carbon in authigenic dolomite in the OC-rich Monterey Formation (Miocene) of California has a range of one 

values reflecting formation in anoxic waters. Pisciotta and 

Mahoney (1981) and Kelts and McKenzie (1982) analyzed 

dolomite samples from Holocene OC-rich sediments from 

DSDP cores from the California borderland and Gulf of 

California, respectively. Both workers found large ranges 

in one values of carbonate carbon, apparently reflecting 

dolomite precipitation both during and after sulfate reduction in pore waters. Pisciotta and Mahoney (1981) found 

a strong correlation between sedimentation rate and o13C 

values of dolomite; sediment deposited rapidly apparently 

began methane generation earlier, had heavier one 

values of dissolved carbon in pore waters, and, therefore, 

had heavier one values in precipitated dolomite. 

The exact mechanism of dolomite precipitation in 

anoxic sediments is unknown, but Baker and Kastner 

(1980) suggested that dolomite precipitation is favored by 

a lack of dissolved sulfate in pore waters. Presumably, 

most of the dolomite precipitates after reduction of sulfate 

in pore waters is complete. 

Siderite is common in many Cretaceous OC-rich 

sequences at Atlantic DSDP sites, particularly sites 105, 

386, 387, and 398. Siderite at sites 105, 386, and 387 occurs 

mainly as hard, yellow laminae or lenses. The siderite at 

site 398 occurs in a variety of morphologies including 

laminae, lenses, and replacements of biogenic calcite 

(Arthur, 1979a). Stable-isotope values of the siderite 

reflect an early diagenetic formation (one values as 

negative as -5.0 per mil) and (or) replacement of biogenic 

calcite. 

The type of diagenetic carbonate formed in 

Cretaceous OC-rich sequences probably depends on the 

chemistry of the initial detrital sediment, the sedimentation rate, and the extent of OC enrichment and sulfate 

reduction. Siderite appears to be the most common 

diagenetic carbonate mineral in sequences that have 

relatively low concentrations of OC (for example, the 

western North Atlantic basin) or high sedimentation rates 

(for example, sites 370 and 398). This distribution may 

be due to the fact that sulfate is reduced quickly without 

much incorporation of sulfur into metal sulfides leaving 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 107 
a large supply of unmineralized reduced iron. Rhodochrosite and other Mn-rich carbonates are not common 

in Cretaceous OC-rich strata, although manganese substitution in authigenic carbonates is common. Manganese is 

highly mobile in anoxic pore waters, and, in sediments 

deposited more slowly, it probably diffuses out of the sediment too rapidly to be precipitated in the form of carbonate 

minerals. Impure Mn-rich carbonate phases, usually containing varying amounts of Ca, Mn, and (or) Fe, occur in 

reduced marine sediments from many different localities 

(see, for example, Zen, 1959; Manheim, 1961; Lynn and 

Bonatti, 1965; Calvert and Price, 1970; Froelich and others, 

1979; Rein and others, 1979; Suess, 1979; Pederson and 

Price, 1982). Most of the manganese in these carbonates 

may be derived from the dissolution of manganese oxides 

after burial. The carbonate may be derived from oxidation of organic matter under reducing conditions as long 

as this reduction is not accompanied by a major decrease 

in pH. Given a sufficient supply of manganese and 

dissolved carbonate and a sufficiently high pH, a Mn-rich 

carbonate layer may form under reducing conditions. 

The Cretaceous OC-rich strata in the Atlantic vary 

considerably in amounts and types of diagenetic carbonate minerals. These minerals deserve more systematic 

study than has been devoted to them so far; in particular, 

the chemistry and mineralogy of diagenetic carbonates 

related to initial sediment properties deserves further study. 

Phosphate diagenesis 

We know of no reported carbonate-fluorapatite 

pellets, nodules, or other phosphate aggregates from Atlantic Cretaceous OC-rich strata, although these strata usually 

have a large range of phosphorous concentrations (0.005-

1.8 percent P20 5). Most phosphorous is supplied to 

sediments in productive areas of the oceans in the form 

of organic matter and fish debris (Froelich and others, 

1982). Phosphate is released from organic matter by decomposition both in the water column during settling of organic 

matter and fecal pellets to the bottom and in pore waters 

by aerobic and anaerobic decomposition. The carbon:phosphorus (C:P) ratio in the organic matter increases 

systematically as the organic matter settles through the 

water column and during early diagenesis (Muller and 

Suess, 1980; Krom and Berner, 1981; Froelich and others, 

1982). Phosphate produced by decomposing organic matter in the sediment can diffuse out of the sediment into 

overlying seawater, precipitate in the sediment as a 

phosphatic mineral, or become adsorbed on ironoxyhydroxides at an oxic-anoxic interface (Berner, 1974). 

In general, dissolved phosphate in pore waters of OCrich sediment increases with depth to a maximum 

that is likely controlled by the precipitation of carbonatefluorapatite. This increase is due to the decomposition of 

organic matter and possibly the dissolution of fish debris 

108 Studies in Diagenesis 

(Baturin and others, 1970; Berner, 1980; Suess, 1981). Fish 

debris usually is well preserved in OC-rich strata, suggesting 

that its dissolution is buffered by the increase in phosphate 

concentration due to the decomposition of organic matter. Baturin and others (1970) found a maximum phosphate 

concentration of 9.7 1-'mol/L in the pore waters of Walvis 

Bay muds, more than three times the highest concentration in open oceanic deep waters. Phosphate concentrations several orders of magnitude greater than those in 

overlying seawater have been reported from sediments in 

nearshore environments (Baturin, 1972, Nissenbaum and 

others, 1972; Sholkovitz, 1973; Berner, 1974; Suess, 1976; 

Goldhaber and others, 1977; Krom and Berner, 1981). A 

limit to the amount of phosphate that can be buried in 

sediment is the initial C:P ratio in the sedimenting organic 

matter. Deep-water OC-rich sediments may not be a major 

sink for phosphorous because, although much organic matter may be preserved in anoxic strata, much of the organic 

phosphorous may have been released by decay of organic 

matter that settles to the bottom. Thus, the C:P ratio of 

sedimented organic matter may already be elevated, and 

the total phosphorous content of deeper-water OC-rich 

sediment will be low (Glenn and Arthur, 1982) relative to 

that of OC-rich sediment in shallow water. The average 

P20 5 content of modern OC-rich sediment from the 

Black Sea is 0.17 percent, which is similar to or lower than 

P20 5 concentrations in oxic sediments where phosphate is 

adsorbed onto ferric oxyhydroxides (Berner, 1974). 

The average concentration of P 20 5 in black shale 

from site 530 in the southern Angola basin is 0.19 percent 

(data of Dean and Parduhn, 1984). The average concentration of OC in all samples of black shale reported by 

Meyers and others (1984) is 5.4 percent. These two values 

yield an average C:P atomic ratio of 168 which is high 

relative to the average C:P ratio of 106 found by Redfield 

(1958) for freshly fixed planktonic organic matter, and confirmed by more recent studies on plankton and particulate 

organic matter (for example, Krom and Berner, 1981; 

Froelich and others, 1982). The C:P ratio of 168 for organic 

matter in site 530 black shales, however, is lower than the 

values reported for modern OC-rich sediments (see, for 

example Sholkovitz, 1973; Krom and Berner, 1981; Froelich 

and others, 1982). Phosphorus accumulated at rates between 0.013 and 0.58 g/cm2

/m.y. (average of about 0.2) in 

the black shales at site 530-not exceptionally high values. 

Phosphate probably occurs mainly in the form of preserved 

phosphatic fish debris in most deep-sea OC-rich strata. 

TRACE ELEMENTS IN 

ORGANIC-CARBON-RICH STRATA 

The Cretaceous OC-rich strata in the Atlantic provide 

insight into the general problem of the incorporation of 

trace metals in OC-rich sediments. Such marine sediments 

and rocks are known to be significantly enriched in a 
variety of oxidation-reduction-sensitive trace elements (see, 

for example, Vine and Tourtelot, 1970) relative to average 

marine sediment and shale, but the degree of enrichment 

varies greatly. Few systematic studies have been made of 

major and trace element concentrations in Cretaceous 

deep-sea OC-rich lithologies, but such studies are important because these extensive lithologic units are potential 

sinks for OC, sulfur, and certain trace elements. During 

the deposition of OC-rich sediments, trace elements, OC, 

and sulfur may be extracted from the oceans at greater than 

steady-state rates, thereby perturbing the chemical and 

isotope balances of the ocean. The possible effects of Cretaceous OC accumulation on the global carbon cycle was 

discussed by Scholle and Arthur (1980), Arthur (1982), and 

Arthur and others (1985) who found that average o13C 

values of oceanic TDC increased during mid-Cretaceous 

black shale deposition, as did the rates of accumulation 

of OC in the deep sea. Average deep-sea accumulation rates 

of OC were as much as seven times those of today during 

the peak of Aptian-Albian black shale deposition. Anderson and others (1982) suggested that increased rates of 

burial of sulfur as metal sulfides in Cretaceous black shales 

caused significant changes in the oceanic sulfur cycle. 

Brumsack (1980) also considered the possible effects of 

Cretaceous black shale deposition on trace metal and carbon cycling, but his estimates of the fluxes of those components to the black shales, probably were too high because 

he considered the entire thickness of a so-called "black 

shale'' unit to have the composition of his "average'' black 

shale (see table 2). As pointed out earlier, black-shalebearing units recovered from Atlantic DSDP sites commonly contain less than 10 to 20 percent of beds with an 

OC concentration of more than one percent, but some 

intervals have more. The geochemistry of interbedded 

Table 2. Average or representative chemical analyses for modern organic-carbon-rich sediment, oxic deep-sea sediments, and 

Cretaceous black shales and interbedded strata 

[Values in parentheses are recalculated on a carbonate-free basis. Leaders (---) indicate not analyzed or not reported] 

South West Deep-sea Cretaceous Cretaceous 

Element Bunnefjord, Africa shelf pelagic Atlantic black shales 

or oxide Black Sea1 Norway2 diatom clay3 Baltic Sea4 clay5 black shales6 DSDP site 5307 

Percent 

Sio2 32.5 (51. 8) 34.9 (40.4) 51.2 (57.6) 39.0 (42. 0) 60.4 56.7 

Al2o3 8.97 (14. 3) 8. 77 (10.1) 2.92 (3. 28) 14.2 (15.3) 19.1 11.8 

Tio2 .29 (.46) .43 (.SO) .23 (.26) .72 (. 77) .92 1.11 

Fe 3o3 4.82 (7. 69) 10.4 (12.1) 1.62 (1.82) 7.7 (8. 3) 8.93 5.06 8.5 

Cao 20.9 7.3 7.26 2.2 1. 01 

MgO 2. 21 (3.53) 1. 83 (2. 10) 1.33 (1. 49) 2. 7 3.73 2.8 

Na2o 1. 61 ( 2. 57) .93 (1. 08) .3 (. 3) 1.67 1.39 

K2o 1. 72 (2. 75) 2.27 (2.63) .71 (.80) 4.3 (4.6) 3.56 3.17 

MnO .066 (.11) 4.78 (5.54) .014 5.2 (5.6) .52 .12 .07 

P2os .27 (. 43) .61 (. 71) 1. 59 (1. 79) .23 (.25) .16 .18 

s 2.90 (4.61) 4.26 (4.94) 1.59 (1. 79) 2.8 (3. 0) 2.19 2.36 

Org-C 2.25 (3. 58) 5.44 (6.30) 9.35 (10. 5) 4.6 (5.0) 6.18 4. 77 

Parts per million 

As 71 (82) 16 13 32 

Ba 265 (423) 552 (640) 198 (222) 750 (806) 2,300 530 660 

Cd 35 0.21 0.92 8 

Co 60 (70) 4.6 20 (24) 74 27 59 

Cr 107 (171) 107 (124) 91 90 (97) 90 202 180 

Cu 45 (72) 135 (157) 68 (7 6) 78 (84) 250 156 150 

Mo 26 (42) 26 (30) 53 (60) 35 (38) 27 59 26 

Ni 108 (172) 45 (52) 108 (121) 43 (46) 225 186 130 

Pb 10 (16) 167 (194) 12 (13) 25 (27) 80 16 27 

v 276 (320) 130 (140) 120 822 460 

Zn 76 (121) 944 (1 ,090) 68 (76) 110 (118) 165 828 320 

1

Average of modern Black Sea deep-water sediments; data of Hirst (1974). 2

Average of anoxic sediments from Bunnefjord, Norway (Calvert, 1976). 3 Average of diatomaceous clays, South West Africa shelf (Calvert, 1976). Values for MoO, As, Cd, Co, and Cr are averages of samples from 

the tops of 10 cores of green diatomaceous muds from South West Africa shelf; data of Brongersma-Sanders and others, 1980. 4

Average of analyses of sediment from Gotland basin, Baltic Sea; data of Manheim (1961). 

s Average oxidized, deep-sea pelagic clay on a carbonate-, water-, and organic-free basis (Chester and Aston, 1976). 6

Average of 38 samples of Atlantic Cretaceous black shales; data of Brumsack, 1980. 7 Average of 28 samples of black shale from Deep Sea Drilling Project (DSDP) site 530, southern Angola basin; data of Dean and Parduhn 

(1984) recalculated on a carbonate-free basis. 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 109 
1 

Mo 

CONCENTRATION, IN PARTS PER MILLION 

10 100 1,000 10,000 100,000 

• 

MnO 

Fe203 ~ 

MgO~ 

I --cs--- ~ Na20 --o-

-6-

I 

K20 :::t- 1 ~ 

Ti02~ --1::7-

Ba ---o~-----------

Cu ---oN i --o-------- ~ 

B ----oL. • 

I ~ 

---6--

Co --o---

0.001 0.01 

PERCENT 

0.1 

1 PERCENT 

Site530A 

Cores 89-105 

SYMBOL FOR BED COLOR 

~black 

~green 

--b- red 

Figure 6. Observed range (length of bar) and geometric mean (symbol) of concentrations of major-element oxides 

and trace elements in samples of red (open triangle) and green (open square) claystone and black shale (solid circle) 

from Deep Sea Drilling Project site 530, cores 87-105. 

110 Studies in Diagenesis 
black, green, and red Cretaceous strata are contrasted 

below, and their geochemistry and metal accumulation 

rates are compared with those of modern anoxic sediments. 

The enrichment of certain trace elements-especially 

Cu, Zn, Mo, V, Ni, and Cr-in sediments and rocks enriched in organic matter has been reported by many 

investigators (for example, Tourtelot, 1964; Wedepohl, 1964; 

Brongersma-Sanders, 1965; Calvert and Price, 1970; Vine 

and Tourtelot, 1969, 1970; Volkov and Fomina, 1974; lange 

and otht~rs, 1978; Chester and others, 1978; Leinen and 

Stakes, 1979; Brongersma-Sanders and others, 1980; Dean 

and Gardner, 1982; Dean and Parduhn, 1984). The association of high trace-element concentrations with organic matter may reflect the concentration of these elements by 

organisms (bioconcentration), by chemical sorption and 

precipitation with organic detritus, clays, and so forth falling through the water column, or formation of sulfides 

under anoxic conditions created by a flux of excess OC. 

Marine plankton concentrate trace elements, especially Ba, 

Pb, Ni, Cu, Zn, Mn, and Fe (Vinogradov, 1953; Goldberg, 

1957; Boyle and Lynch, 1968; Knauer and Martin, 1973; 

Martin and Knauer, 1973; Vinogradova and Kovalskiy, 

1962; Bostrom and others, 1974; Chester and others, 1978; 

Moore and Bostrom, 1978; Leinen and Stakes, 1979; Collier and Edmond, 1984), suggesting that bioconcentration 

is a potentially important mechanism for incorporation of 

certain trace elements in OC-rich marine sediments. The 

great effectiveness of the adsorption of trace metals by clay 

minerals and organic matter, however, and coprecipitation 

of trace metals, particularly as sulfide minerals, suggests 

that these processes also may play an important role in the 

removal of trace metals from seawater (Tourtelot, 1964, 

1979; Volkov and Fomina, 1974; Holland, 1979). 

From studies of the minor-element compositions of 

contemporaneous Upper Cretaceous marine and nonmarine shales in the western interior of the United States, 

Tourtelot (1964) found that element concentrations were 

consistently higher in the marine shales. He concluded that 

sorption and syngenetic mineral formation in the marine 

environment were important, if not the most important, 

processes affecting the minor element content of shales. 

The final minor element composition is thus the result 

of cycles of sorption-desorption and (or) precipitationdissolution as the sediments are subjected to varying chemical conditions in depositional and diagenetic environments. 

The greatest control is probably exerted by oxidationreduction conditions during diagenesis. Vine and Tourtelot 

(1970) also concluded that sorption processes are important in metal enrichment in black shales, but suggested that 

some prior concentration of metals from seawater (for 

example,. by bioconcentration) was required. They stated 

that the availability of metals in the various solutions with 

which the organic matter comes in contact throughout its 

history is probably the most significant factor in determining metal enrichment, and that interstitial pore fluids are 

probably the most important of these solutions. 

If thin beds of black shale are interstratified with 

other lithologies, they may become sinks for elements 

mobilized and transported in pore waters during compaction (Berner 1969, 1975). The importance of interbedded 

more-reduced and less-reduced strata on the migration and 

accumulation of trace elements and the relative mobility 

of different elements under reducing conditions was 

pointed out by Dean and Parduhn (1984) for Cretaceous 

black shale interbedded with red and green claystone in 

the Angola basin at DSDP site 530. This 160-m-thick 

interbedded sequence is Albian through ConiacianSantonian in age and represents the upper part of midCretaceous black-shale deposition in the Atlantic. The 

sequence contains 260 beds of black shale which average 

5.4 weight percent OC, compared with less than 0.5 percent for red and green claystones in the sequence (Meyers 

and others, 1984). Geometric-mean concentrations of trace 

elements from each lithology show that concentrations 

of As, Cd, Co, Cr, Cu, Mo, Ni, Pb, V, and Zn are all 

higher in black shales than in red and green claystones 

(fig. 6). Concentrations of Ba and Fe generally are higher 

in red claystone than in green claystone and black shale. 

Concentrations of manganese generally are higher in green 

claystone than red claystone, and much lower in black shale. 

Dean and Gardner (1982) found the same general relationships in cycles of black shale and green and red claystone 

from DSDP site 367 off northwest Africa; Cr, Mo, Ni, Zn, 

and V were enriched in black shale, whereas Co and Cu 

were of about equal concentrations in all lithologies (fig. 7). 

The maximum concentrations of elements in the site 

530 black shales, however, are not all at the same depth 

(fig. 8). Concentrations of V and Cu (as well as Cd, Zn, 

and Mo, not shown in fig. 8) are highest in the part of 

the section where black shale beds are most abundant. 

Concentrations of several other trace elements, especially 

Ni, Ba, and Co (not shown in fig. 8) are higher in black 

shales in the upper part of the section. These differences 

may be related to differences in amount and type of organic 

matter in the black shale beds. Data of Deroo and others 

(1984; fig. 8) show that the black shale beds in the upper 

part of the section contain more organic matter (higher 

percentage of OC) and that this organic matter contains 

more lipid-rich marine organic matter (higher H-index). 

Brumsack (1980) analyzed 38 samples of 

Cretaceous black shale from the Atlantic that had more 

than 1 percent OC (average of 6.12 percent; table 2) and 

compared the element concentrations in these black shales 

to those from 11 samples of associated lithologies that 

contained 0.5-1.0 percent organic carbon and 38 samples 

of associated lithologies that contained less than 0.5 percent OC. His results showed that Cu, Cr, Mo, Ni, and 

Zn were enriched in OC-rich beds relative to associated 

OC-poor beds. Iron and manganese concentrations were 

highest in black and red beds. Table 2 compares Brumsack's (1980) data with those from site 530 (Dean and 

Parduhn, 1984). The mean values for samples from site 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 111 
PARTS PER MIL LION 

1 10 

Site367 

Core 17 

100 1,000 10,000 

totai-C ---o--

total-5 

Mg 

Na .(::,. 

A 

K 

Ti 6 

• 

100,000 

• 

Fe 

6. ... 

6 .. 

Mn • -t::r 

Ba,...! 

v -a- ........ 

1 PERCENT 

Cu-r 

Cr ~_._ 

Ni · --b-.._ 

~Be 

CoT 

Pb =::-

0.001 

--6--

Zn • 

Zr ~ 

s: 

0.01 

PERCENT 

530 given by Dean and Parduhn and plotted in figure 6 

are geometric means which are more realistic measures 

of central tendency because trace element concentrations 

usually have frequency distributions that more closely 

approximate log-normal than normal distributions. In 

table 2, however, the site 530 data are expressed as 

arithmetic means for purposes of comparison. The 

average concentrations of Cr, Cu, Ni, and Pb in both 

sets of black shale samples are comparable, whereas 

112 Studies in Diagenesis 

0.1 

SYMBOL FOR BED COLOR 

--{:r- Green 

--6- Black 

A 

Brumsack's (1980) average concentrations of Mo, V, and 

Zn are about twice those of the site 530 black shales; concentrations of Ba, Cd, and Co are higher in site 530 black 

shales. Some of these differences may reflect analytical 

uncertanties, but, in general, Cretaceous black shales 

from the Atlantic are similarly enriched in trace elements. 

Compared with modern anoxic sediments from a 

variety of environments, the site 530 black shales are at 

least as enriched or more enriched in Ba, Cd, Co, Cr, 
PARTS PER MILLION 

10 

Site 367 

Cores 15-16 

100 1,000 

• 

10,000 

I 

totai-C 

total-S 

100,000 

Fe .o. • 

Mg~ 

No~ 

K~ 

T1. . -oMn----o--

Ba~ 1 PERCENT 

I_.,__ ·~ 

B 

--o-M 

~·--41·--- 0 

0.000'1 

Co--o-- ....... 

0.001 

PERCENT 

0.1 PERCENT 

SYMBOL FOR BED COLOR 

--o-- Green 

Zr --..-Red 

B 

0.01 

Figure 7 (above and facing page). A, Observed range (length of bar) and geometric mean (symbol) of concentrations of 

major-element oxides and trace elements on a carbonate-free basis for samples of interbedded black shale (solid triangle), 

and green claystone (open triangle), from Deep Sea Drilling Project site 367, core 17. 8, Observed range (length of bar) 

and geometric mean (symbol) of concentrations of major-element oxides and trace elements on a carbonate-free basis for 

samples of interbedded green (open circle) and red (solid circle) claystone from Deep Sea Drilling Project site 367, cores 

15 and 16. 

Cu, Ni,. Pb, V, and Zn (on a carbonate-free basis; 

table 2). In general, concentrations of trace elements in 

site 530 deep-water black shales, as well as most of those 

reported by Brumsack ( 1980) and Dean and Gardner 

(1982), are significantly higher than those in sediments 

from modern anoxic environments from the Black Sea 

and Saanich Inlet central basin (table 2). These two 

modern anoxic environments have anoxic water columns 

above the sediment rather than only anoxic sedimentinterstitial waters. The ranges and mean concentrations 

of most trace elements in site 530 black shales also are 

larger than those in very OC-rich green muds deposited 

within an intense oxygen-minimum zone in Walvis Bay, 

South West Africa (table 2). 

Enrichments of Cd, Mo, Ni, Co, and Zn can be 

related to enhanced supply and preservation of organic 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 113 
.... 

.... 

""' 

1;11 

c 

c. 

~· 

"' :;-

0 

~· (JCI 

~ 

::I 

~ 

"' (ii' 

C/) 

[C 

w 

1-

w 

::?! 

~ 

::i 

li: 

w 

0 

::?! 

0 

1: 

0 

CD 

CD 

::> 

C/) 

I I 

940 

960 

-

- 980 -

-

- 1-000 

1020 ---

1040 -

- 1060 ---

-

~ 

.... 1080 

1100 f'(':\1

: I I I I I I 1 . I I 

0 25 50 0 5 10 

BLACK SHALE, 

IN PERCENT 

ORGANIC CARBON, 

IN PERCENT 

15 0 200 400 600 0 1500 3000 0 250 500 0 300 600 0 2000 4000 

H-INDEX. VANADIUM. COPPER, NICKEL. BARIUM. 

IN MILLIGRAMS IN PARTS PER IN PARTS PER IN PARTS PER IN PARTS PER 

OF HYDROCARBONS MILLION MILLION MILLION MILLION 

PER GRAM OF 

ORGANIC CARBON 

Figure 8. Plots of carbonate-free (cf) concentrations (in parts per million) versus depth for V, Cu, Ni, and Ba in samples of black shale from Deep Sea Drilling 

Project site 530, southern Angola basin (from Dean and Parduhn, 1984). Percentage of black shale beds in each 9.5-meter core is shown. Values of percent 

organic carbon and hydrogen index (H-index, in milligrams of hydrocarbons per gram of organic carbon) are from Deroo and others (1984). 
matter (see, for example, Bruland, 1980), with subsequent 

liberation during decomposition of the organic matter and 

scavenging by sulfides during early diagenesis in anoxic 

sediment. The high concentrations of zinc and cobalt also 

may havt;:~ been supplied directly to the sediment as metal 

sulfide particles produced in an anoxic, sulfide-rich water 

column as in the modern anoxic Black Sea (see for example, Spencer and others, 1972; Brewer and Spencer, 1974), 

but other evidence suggests that the black shales at site 

530, and probably in much of the Cretaceous Atlantic 

Ocean, were not deposited under an entirely anoxic water 

column. 

A possible example of trace element enrichment by 

pore waters in oxidized sediment overlying reduced sediment was recovered at DSDP site 105 in the North 

American basin of the western North Atlantic Ocean. 

Here, th(~ middle Cretaceous black shale unit is overlain 

by Upper Cretaceous-lower Tertiary red and green clays 

(Lancelot and others, 1972). This unit has been recovered 

at eight other sites in the North American basin, and is 

interpreted as having been deposited in an oxygenated, 

pelagic deep-sea environment similar to that presently 

accumulating pelagic clays in the central North Pacific 

(Jansa and others, 1979). Lancelot and others (1972) considered that the multicolored clays at site 105 were 

enriched in certain heavy metals, especially Mn, Zn, Cu, 

Pb, Cr, Ni, and V, and that this enrichment was a result 

of hydrothermal volcanic exhalations. If the analyses of 

the site 105 multicolored clays by Lancelot and others 

(1972) are compared with analyses of modern and ancient 

deep-sea sediments and rocks from a wide variety of 

environments (table 2, fig. 7), however, the site 105 clays 

are clearly not enriched in any elements, even when compared with modern pelagic red clay and Creteceous red 

claystones from other Atlantic sites. On the other hand, 

bands enriched in goethite, manganese oxides, and 

sphalerite in the site 105 multicolored clays do require 

explanation. Because iron, manganese, and certain trace 

elements, particularly the transition elements, are easily 

mobilized under reducing conditions, Arthur (1979b) proposed that much of the excess iron and manganese in 

some bands within the site 105 multicolored clays was 

derived from upward diffusion of dissolved ions and 

advection of pore waters from compaction of black shales 

underlying the multicolored clays. When the chemically 

reduced ions derived from the underlying black shales 

reached a sufficiently oxidizing interface within the 

multicolored clays or at the sediment/water interface, 

iron and (or) manganese oxides precipitated in response 

to lower solubilities under increased Eh conditions. 

Authur (1979b) suggested that the sphalerite, which is 

mainly at the base of the multicolored clay sequence, also 

originated by upward diffusion of zinc from the underlying black shales. 

For each 9.5-m cored interval in the black-shalebearing sequence at site 530, mass accumulation rates 

(MAR; g/cm2/m.y.) were calculated for bulk sediment, 

for each different colored lithology, for OC, for Zn, Cr, 

and Ni in each different colored lithology, and for the 

sum of each of the three trace elements (table 3). Mass 

accumulation rates were calculated (table 3) using 

estimates of the percentages of black, red, and green beds 

in the recovered part of each 9.5-m cored interval, 

biostratigraphically determined accumulation rates and 

porosity (Hay and others, 1984), and an average grain 

density of 2.7 g/cm3• The MAR for OC varies from 0 

to 28 g/cm2/m.y. As expected, maximum values for OC 

and trace element MARs occur in cores 97 and 98 where 

black shales are much more abundant than in any other 

part of the section (fig. 8). Maximum MARs for Zn, Cr, 

and Ni in black shale beds are 1.5, 0.22, and 0.11 

g/cm2/m.y., respectively, in cores 97 and 98. MARs for 

these elements are all about 0.01-0.02 g/cm2/m.y. in 

black shale beds in the lower part of the black shale 

sequence, and 0.01-0.4 g/cm2/m.y. in the upper part of 

the sequence (table 3). The changes in zinc MAR are the 

most extreme. Maximum values of total MARs (all 

lithologies) for Zn, Cr, and Ni are 1.57, 0.67, and 0.41 

g/cm2/m.y., respectively, and only the maximum for 

zinc occurs in the zone of maximum black shale concentration in cores 97 and 98. Total MARs for all three 

elements are uniformly low in the lower part of the black 

shale section and relatively high in the upper part of the 

section (cores 87 to 96). 

In comparison, Leinen and Stakes (1979) calculated 

average MARs for Zn and Ni of 0.17 g/cm2/m.y. and 

0.08 g/cm2/m.y., respectively, for sediment on the crest 

of the East Pacific rise, and 0.06 g/cm2/m.y. and 0.03 

g/cm2/m.y., respectively, for pelagic sediment in the 

eastern equatorial Pacific Ocean. The average MARs of 

Leinen and Stakes (1979) are somewhat higher than the 

average MARs of site 530 black shales, but lower than 

total MARs in all cores except those in the lower part 

of the section. As expected, the percentages of total MAR 

for each element that are contributed by black shale are 

highest in cores 97 and 98. Black shales, however, also 

contribute significant proportions of the total MARs for 

Zn and Ni in cores above 96 (as much as 62 percent of 

the zinc MAR in core 87), although black shales comprise less than 8 percent of this part of the section. 

Site 530 probably is an example of the lower limit 

of trace-element accumulation in Cretaceous black shales. 

The OC accumulation rates are low (maximum of 28 

g/cm2/m.y.) compared with that of other Atlantic Cretaceous black shale sequences. For example, Arthur and 

Natland (1979) calculated that OC in the Aptian-lower 

Albian black shale sequence at site 361 in the Cape basin 

accumulated at an average rate of 258 g/cm2/m.y., 

and that the OC in lower to middle Albian black, 

dolomitic sapropels and marly limestones accumulated 

at rates of about 500 g/cm2/m.y. We have calculated 

that OC in Albian-Turonian black shales at site 367 in 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 115 
... 

aVI 

.. 

c 

Q. 

;:;-

(II 

:r 

0 

;· Table 3. Site 530, mass accumulation rates of selected components in red, green, and black lithologies ~ :I 

I'D 

[OC, organic carbon; leaders (---) indicate no analyses for that core] (II 

w;· 

Core Depth interval (m) Recovery Colored beds (percent) Percentage in black beds 2 Percentage in red beds 2 Percentage in green beds 2 

No. From To (em) Black Red Green oc 1 Zn Cr Ni Zn Cr Ni Zn Cr Ni 

87 940 949 703 8 72 20 6.8 0.180 0.029 0.036 0.010 0.011 0.004 0.008 0.015 0.010 

88 949 958 448 6 74 20 8.5 .140 .045 .039 .007 .011 .004 .013 .019 .005 

89 958 967 922 3 87 10 10.0 .140 .020 .040 .025 .025 .OIS .013 .017 .006 

90 967 976 446 5 85 IO 9.0 .110 • OI8 .049 .020 .020 .010 .013 .OIS .008 

9I 976 985 547 0 95 5 --- --- --- --- .011 .010 .006 .OIS .OIS .034 

92 985 990 0 

93 990 999 9I8 4 90 6 I.O .OI4 .025 .030 .007 .008 .006 .OI6 .OI2 .007 

94 999 I ,008 273 6 79 IS 7.I .042 .025 .OIS .007 .008 .006 .OI6 .OI2 .007 

95 I,008 1,017 656 5 80 15 9.6 .009 .092 .065 .009 .009 .007 .031 .012 .005 

96 1,017 I,026 795 9 45 46 3.0 .023 .033 .008 .009 .009 .009 .OI4 .008 .006 

97 I ,026 1,035 607 52 0 48 5.1 .275 .040 .006 --- --- --- .013 .010 .005 

98 1,035 I,044 SOI 36 0 64 5.0 .300 .016 .030 --- --- --- .024 .007 .006 

99 I,044 1,053 735 9 20 71 I. 7 .019 .034 .009 .009 .009 .008 .007 .007 .006 

100 1,053 1,062 672 12 25 66 I. 7 .012 .020 .010 .009 .009 .008 .010 .008 .006 

101 1,062 1,071 995 5 35 60 3.0 .013 .042 .OI2 .009 .009 .009 .010 .010 .006 

102 1,071 1,080 800 3 70 25 3.9 .013 .027 .020 .006 .006 .004 .010 .010 .005 

103 1,080 1,085 600 2 60 37 3.5 .014 .008 .011 .006 .006 .004 .010 .006 .005 

104 1,085 I,094 790 8 0 94 2.9 .012 .010 .010 --- --- --- .010 .009 .008 

105 1,094 1,102 800 10 20 72 4.0 .014 .010 .005 .009 .006 .008 .008 .008 .005 
Mass accumulation rate (in grams per square centimeter per million years) 

Core Claystone beds Element in colored bed4 

0 No. Black Red Green Total ... Bulk ciQ" 3 Black Red Green oc Zn Cr Ni Zn Cr Ni Zn Cr Ni Zn Cr Ni :r 

I» 

= 87 2,786 223 2,006 557 15.2 0.401 0.065 0.080 0.200 0.221 0.080 0.045 0.084 0.056 0.646 0.369 0.216 Q. 88 2, 786 167 2,062 557 14.2 • 234 .075 .065 .144 .227 .082 .072 .106 .028 .451 .408 .176 0 

Ai" 89 2,786 84 2,424 279 8.4 .117 .017 .033 .606 .606 .364 .036 .047 .017 .759 .670 .414 

~ 90 2, 786 139 2,368 279 12.5 .153 .025 .068 .474 .474 .237 .036 .042 .022 .663 .540 • 327 

I'D = 91 2,786 0 2,647 139 --- --- --- --- .291 .265 .159 .021 .021 .047 .312 .286 .206 

;;;· "' 

0 92 --- 0 0 0 0 

n - 93 2,786 111 2,507 167 1. 1 .016 .028 .033 .176 .201 .150 .027 .020 .012 .218 .249 .196 

i 94 2, 786 167 2,201 418 11.9 .070 .04~ .025 .154 .176 .132 .067 .050 • 029 .291 • 268 .186 I» 

n 95 1,049 52 839 157 5.0 .005 .048 .034 .076 .076 .059 .049 .019 .008 .129 .143 .101 

~ 96 1,049 94 472 483 2.8 .022 .031 .008 .042 .042 .042 .068 • 039 .029 .132 .112 .079 c 

"' 0 97 1,049 545 0 504 27.8 1.500 .218 .033 --- --- --- .065 .050 .025 1.565 .268 .058 ~ 98 1,049 378 0 671 18.9 1.133 .060 .113 --- --- --- .161 .047 .040 1.294 .107 .153 i' 

~ 99 1,049 94 210 745 1. 6 .018 .032 .008 .019 .019 .017 .052 .052 .045 .089 .103 .070 

~ 100 1,049 126 262 692 2.1 .015 .025 .013 .024 .024 .021 .069 .055 .042 .108 .104 .on 

0 101 1,049 52 367 629 1. 6 .007 .022 .006 .033 .033 .033 .063 .063 .038 .103 .118 .076 ... 

~ 

= 102 1,049 31 734 262 1. 2 .004 .008 .006 .044 .044 .029 .026 .026 .013 .074 .079 .049 

~ 103 1,049 21 629 388 • 7 .003 .002 .002 .038 .038 .025 .039 .023 .019 .080 .063 .047 

I» 104 1,049 84 0 986 2.4 .010 .008 .008 --- --- --- .099 .089 .079 .109 .097 .087 

a- 105 1,049 105 210 755 4.2 .015 .010 .005 .019 .013 .017 .060 .060 .038 .094 .084 .060 0 

= 

~ ;:;-

:::r 1

Average percentage of organic carbon (OC) in black shale beds in each core is based on data of Deroo and others (1984) and Meyers and others (1984). 

s; 2

Data of Dean and Parduhn (1984). :::r 3

Mass accumulation rate (MAR) for bulk sediment = (sedimentation rate in meters per million years) X 100 (1 - porosity/100) x grain density (Thiede and Rea, 1981) s. 

I» where sedimentation rate = 15.4 meters per million years for cores 87-94 and 5.8 for cores 95-105. Porosity = 33 percent for all cores, and grain density is assumed to be n 

;· 2.7 g/cm3• Sedimentation rates and porosity from Hay and others (1984). MAR for OC = MAR of black beds X (percent OC/100) for each core. ~ 4

MAR for trace elements = MAR of each colored bed x (percentage of trace element in each colored bed/100). ~ 

iii 

a 

;::;· 

0 

~ I» 

= 

...... 

...... 

...... 
the southeastern North Atlantic off northwest Africa 

accumulated at rates of about 460 g/cm2/m.y. using 

measured OC, porosity, and sedimentation-rate data of 

Lancelot and others (1978). For site 367, Zn and Ni 

MARs of 1.16 and 0.53 g/cm2/m.y., respectively, were 

calculated using average Zn and Ni concentrations of 390 

and 180 ppm, respectively (data of Lange and others, 

1978). Maximum concentrations of OC and trace 

elements in the black shale section at site 367 occur in 

cores 18 and 19. Our unpublished analyses of 16 samples 

from these two cores show average zinc and nickel concentrations of 1280 and 246 ppm, respectively, and MARs 

of 3.8 and 0.73 g/cm2/m.y., respectively. These values 

clearly are at least one order of magnitude higher than 

those in pelagic clays studied by Leinen and Stakes (1979). 

The trace element MARs for site 367 also are higher than 

those estimated for modern anoxic deep-basin sediments 

from the Black Sea (table 4), although the MARs for OC 

are about the same as those for site 367. 

Black shales obviously were an important sink for 

trace metals in the Cretaceous Atlantic Ocean. The concepts and observations mentioned here deserve further 

research and may ultimately explain much of the 

variability in metal enrichment in black shales in general. 

Studies of major and trace element geochemistry of black 

shale sequences, in conjunction with studies of amount 

and type of sulfur and OC, and changes in sedimentation rate, also are needed. 

MODELS OF DEPOSITION 

Bottom-water oxygenation 

Inter bedding of sediments alternately rich and poor in 

organic matter can be the result of either fluctuating 

preservation of organic matter or fluctuating rate of 

supply of organic matter. Differential preservation can 

result if bottom waters at the site of accumulation are 

alternately oxic and anoxic (or near-anoxic). Therefore, 

alternating periods of oxygen-deficient and relatively well 

oxygenated bottom waters may be required if the site of 

accumulation is on a basin floor. Cyclic variation in 

amount of organic matter deposited at a continental 

margin site can be explained by cyclic fluctuations in the 

thickness and intensity of a midwater oxygen-minimum 

layer. Both depositional models (basin deoxygenation and 

expansion and intensification of an oxygen-minimum 

layer) have been proposed to explain the accumulation 

of OC-rich strata (see, for example, Schlanger and 

Jenkyns, 1976; Ryan and Cita, 1977; Fischer and Arthur, 

1977; Thiede and van Andel, 1977; Arthur and Schlanger, 

1979). Both models require almost no dissolved oxygen 

in part of the water column as a result of reduced advection of oxygenated water and (or) increased supply of 

organic matter, and both models imply that reducing 

118 Studies in Diagenesis 

conditions in the sediments (and therefore the increased 

degree of preservation of organic matter) are the result 

of anoxic or near-anoxic conditions in the overlying 

waters. 

Black, laminated, OC-rich Cretaceous strata are so 

widespread in the Atlantic that the closed-basin model 

has been proposed as the most likely analog for the Early 

to middle Cretaceous South and North Atlantic basins. 

The implication is that a stable salinity stratification 

existed, either because of lower salinity surface water with 

bottom water of more normal salinity as in the modern 

Black Sea, or because of normal-salinity surface water 

with more-saline bottom water, which led to increased 

residence time of deep-water masses and gradual oxygen 

depletion with or without increased surface productivity. 

Higher-salinity bottom water would have a lower oxygen 

solubility which, together with an increased residence 

time, would result in greater oxygen depletion. The 

salinity-stratification model is made even more plausible 

by the fact that isolated Atlantic basins were accumulating 

great thicknesses of evaporites in the Late Jurassic and 

again in the Aptian, and therefore probably were more 

saline as well as more stagnant in the Early Cretaceous. 

In contrast, Brass and others (1982) suggested that the 

sinking of warm, saline bottom water produced over 

warm, arid, extensive shelf areas could generate more 

vigouous bottom-water turnover rates that may have been 

as high as those of modern thermohaline deep circulation. Model studies by Southam and others (1982) suggest that a totally anoxic water column in the open ocean 

is unlikely because of the delicate balance between oxygen 

demand and supply to deeper waters and nutrient upwelling to support surface productivity. The Black Sea sediment record, however, is proof that rapid variations in 

oxygenation and anoxia can occur in large, isolated basins 

(Deuser, 1974; Degens and Stoffers, 1976). 

Supply of organic carbon to sediments 

All the interbedding of more-reduced and lessreduced strata by fluctuating deoxygenation of bottom 

waters is difficult to explain, particularly in deep-basin 

settings. Most of the Jurassic and Cretaceous strata in 

the Atlantic contain abundant evidence that they were 

deposited in well-oxygenated bottom waters with at least 

some benthic infauna; only the interbedded, darkcolored, OC-rich lithologies, which are a minor part of 

the total middle Cretaceous stratigraphic record, suggest 

anaerobic depositional conditions because they commonly 

are laminated and apparently unbioturbated. Ekdale 

(1980) showed that trace-fossil diversity varies within 

Cretaceous deep-water oxidation-reduction cycles, 

decreasing in OC-rich beds. Chondrites, a shallowburrowing, branching ichnofossil, is typically the last 

trace fossil found at the transition between a green and 
Table 4. Concentrations of major-element oxides (in percent) and trace elements (in parts 

per million) and mass accumulation rate for modern and ancient Black Sea sediments 

[Leaders (---) indicate no analyses. Mass accumulation rates (MAR)3 are in grams per square centimeter 

per million years (g/cm2

/m.y.)] 

Modern Black Sea sediments 1 Ancient Black Sea sapropels2 

Element Mean Mean Mean 

or bulk bulk carbonate-free 

oxide sediment 

Mean 

carbonatefree sediment sediment sediment MAR3 

Si02 

A1 203 

Fe 2o3 

MgO 

CaO 

4caco3 

Na2o 

K2o 

Ti02 

MnO 

s 

c 

P205 

Li 5 

Mo 

Ni 

Pb 

Rb 

v5 

y 

Zn 

Zr 

28.6 

7.66 

4.46 

2.15 

24.7 

44.1 

1.05 

1.53 

.23 

.06 

2.8 

2.6 

.31 

70 

280 

27 

115 

43 

13 

90 

17 

120 

8 

64 

188 

6 

55 

53 

Major elements and oxides 

50.9 

11.3 

7.87 

2. 93 

1.60 

2.72 

.57 

.10 

6.0 

4.8 

.70 

118 

537 

49 

145 

88 

24 

90 

36 

200 

14 

111 

316 

12 

121 

118 

2,860 

766 

446 

215 

2,470 

4,410 

105 

153 

23 

6.1 

280 

260 

31 

Trace elements 

0.7 

2.8 

.27 

1.15 

.43 

.13 

.90 

.17 

1.20 

.08 

.64 

1.88 

.05 

.55 

.53 

43.4 

12.6 

6.07 

2.15 

7.20 

12.9 

2.37 

2.54 

.57 

.09 

4.1 

6.2 

.17 

215 

569 

37 

168 

108 

28 

92 

84 

151 

12 

127 

378 

17 

120 

112 

49.9 

14.6 

6.97 

2.87 

2.72 

2.92 

.65 

.10 

4.7 

7.1 

.20 

242 

639 

42 

192 

123 

31 

103 

95 

172 

14 

145 

425 

20 

138 

129 

1 Data for top 20 em in Black Sea cores 1432 and 1462 (Hirst, 1974). 2 Data for old Black Sea sapropels (3,000 to 7,000 yr BP) (Hirst, 1974). 

6,510 

1,890 

910 

377 

1,080 

1,935 

356 

381 

86 

14 

615 

930 

24 

3.2 

8.5 

.56 

2.52 

1.62 

.42 

1.38 

1.26 

2.27 

.18 

1.91 

5.67 

.26 

1.80 

1.68 

3 Mass accumulation rates (MAR) (grams per square centimeter per million years) are based on the 

mean bulk-sediment concentrations of elements or oxides, an average sediment accumulation rate of 9 

em per 1,000 yr for modern Black Sea sediments, an average sediment accumulation rate of 15 em per 

1,000 yr for old Black Sea sapropels (data of Ross and Degens, 1974), and a dry bulk density of 1.1 gram 

per cubic centimeter (Keller, 1974). 4 Percent CaCO 3 was calculated from percent CaO. 5 Data for B, Co, Ga, Li, and V are for core 1432 only. 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 119 
a black bed, and the first to colonize the bottom following the deposition of laminated black clays. Changes in 

trace-fossil diversity suggest that at least some of the 

oxidation-reduction cycles may be due to a change in 

bottom-water oxygen concentration at or above the 

sediment-water interface. Interbedding of oxic-anoxic 

strata in the North Atlantic began in the Jurassic and continued at least into the Eocene; however, only the Lower 

to middle Cretaceous and Eocene strata contain more 

than 1 percent OC. It is not difficult to visualize abrupt 

changes in bottom-water circulation in a restricted basin 

such as the Early Cretaceous Atlantic Ocean, but it is difficult to imagine total stagnation of the Atlantic during 

the Late Cretaceous and Tertiary. 

We therefore favor a third depositional model to 

explain many examples of interbedded more-reduced and 

less-reduced strata in which varying diagenetic oxidationreduction conditions in the sediments are caused by varying oxygen demand within the sediments, largely in 

response to varying rates of accumulations of marine and 

(or) terrestrial organic detritus. This model does not 

necessitate major changes in the oxygen content of the 

bottom water, only that the supply of organic matter 

increases, either by periodic increases in surface productivity or by periodic redeposition of organic matter in the 

deep basin from sites on the outer shelf or slope where 

more organic matter is preserved in sediments, for example, under an oxygen-minimum zone. Similar environmental interpretations of cyclic facies of the Jurassic of 

Yorkshire, United Kingdom, were made by Morris (1979) 

and Demaison and Moore (1980). 

If the rate of supply of organic matter is low, the 

bottom waters will be more or less well oxygenated, the 

zero-oxygen isopleth will be at some distance below the 

sediment-water interface, and red clay will accumulate. 

If the rate of supply of organic matter is higher, the bottom waters still will be more or less well oxygenated, but 

the increased oxygen demand in the sediments, mainly 

from the increased decomposition of the organic matter, 

will result in a zero-oxygen isopleth at or near the 

sediment-water interface, and green, reduced clay will 

accumulate. If the rate of accumulation of organic matter is so great that the sediments are completely anoxic 

and reduced, the zero-oxygen isopleth may be some 

distance above the sediment-water interface, and H2

S 

accumulates locally in the bottom waters. The extent of 

deoxygenation of the bottom waters would be controlled 

by the rate of oxygen consumption, the rate of H2

S production, the rate of oxygen supply to the bottom waters, 

the strength of bottom currents (advection of oxygen), 

and the rate of sedimentation. The sluggish circulation 

of the bottom waters and (or) the extreme drawdown of 

oxygen by oxidation of OC in the water column would 

aid in the accumulation of H2S, but the primary cause 

of anoxic conditions is the high rate of supply of organic 

matter and concomitant high biological oxygen demand 

120 Studies in Diagenesis 

within the sediments. The black color is due mainly to 

organic matter and FeS. The FeS would eventually be 

converted to FeS2 (pyrite or marcasite) which would 

tend to lighten the color of the sediments. Maturation 

of the organic matter, however, would tend to darken 

the sediments such that the black muds probably become 

black or at least dark-gray shales through increasing thermal alteration and diagenesis. If the initial deposition 

resulted in an interbedding of layers more and less 

enriched in organic matter but otherwise homogeneous, 

migration of so~- into layers enriched in organic matter to replace SO~- consumed by sulfate reduction would 

produce local concentrations of FeS and an intensification of color differences (Berner, 1969). With time, these 

local sulfide concentrations would produce concentrations of pyrite at the boundary between interbeds of 

more- and less-OC-rich mud rocks. 

One problem with this model is how the OC supply 

to deep-water sediment can increase without creating an 

excess of production over consumption of OC in the 

water column. In the modern ocean, less than 0.1 percent of the OC produced survives to a depth of 4,000 m 

(Muller and Suess, 1979, 1980) because of high rates of 

consumption and oxidation in the water column. Changes 

in surface productivity alone probably could not account 

for OC enrichment in the black shales without concomitant depletion of dissolved oxygen in the bottom waters. 

Enough OC, however, could be supplied periodically to 

cause interbedding of green and -red sediment because 

such changes in oxidation state are found in Holocene 

oceanic deep-water sediments (Lynn and Bonatti, 1965; 

McGeary and Damuth, 1973; Gardner and others, 1982). 

Figure 9 (facing page). Illustrations showing how influx of terrigenous organic matter, surface-water organic productivity, 

upwelling, and turbidity currents may contribute to accumulation and redeposition of sediments containing varying amounts 

of organic carbon and, therefore, ranging in color from red to 

black. 

A, Example of a period of moderate supply of organic detritus 

from autochthonous-marine and terrigenous sources and 

moderate upwelling resulting in accumulation of green sediment in shelf areas, sediment with high concentrations of 

organic carbon, perhaps black in color, within the 0 2-

minimum zone, and dominantly red clays in the deep basin. 

Such conditions might be representative of the Late Cretaceous (for example, Santonian-Coniacian) in the Atlantic 

Ocean. 

B, Example of a period of rapid rate of supply of organic detritus 

and stronger upwelling resulting in an expanded and more 

intense 0 2

-minimum zone that may include much of the 

water mass and allow organic-carbon (OC)-sediment rich in carbon to accumulate over a much greater area and provide a 

greater source area for supply of OC-rich sediment to the deep 

basin by downslope transport. Such conditions might be 

representative of periods of accumulation of OC-rich strata in 

the Atlantic Ocean (Aptian-Albian and Cenomanian-Turonian). 
A 

B 

Black sediments 

(organic carbon > 2 percent) 

Late Albian to Early 

Cenomanian and Coniacian 

Dark green to 

black sediments 

(organic carbon > 2 percent) 

Late Cenomanian 

to Early Turonian 

Terrigenous influx 

\ ~ High productivity @I 0 2 >>0 ~ ~· "' 8> &) 

- -----------~-------------------------

~--~~~~~~-----------~-~~------------

Red sediments 

(organic 

carbon < 0.5 percent) 

More terrigenous influx 

Expanded and intensified 

Green or 

black sediment 

(depending on amount 

of organic carbon) 

Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich Lithofacies, Atlantic Ocean 121 
Presumably the periodicity in the initial OC-preservation 

would be aided by periodic increases in sedimentation rate 

(Heath and others, 1977; Muller and Suess, 1979) which 

could also explain the red and green cycles. 

From the preceding discussion, it is apparent that 

the origin of interbedded more- and less-reduced lithologies with variable amounts of organic matter and 

variable amounts of pelagic, hemipelagic, and terrestrial 

sediment is complex and probably is not due to any one 

simple process. Figure 9 is a diagram of many of the 

variables during a period of time of accumulation of less 

organic matter (fig. 9A; for example, AlbianCenomanian), and during a period of accumulation of 

more organic matter (fig. 9B; Aptian-Albian or 

Cenomanian-Turonian). The main variables are the rate 

of influx of terrigenous sediment and organic matter and 

surface-water productivity. The amounts of organic matter from these two sources (terrigenous and marine) in 

turn determine the thickness and intensity of a mid-water 

oxygen-minimum zone. Surface-water productivity is 

determined, at least in part, by the intensity of upwelling 

of nutrient-rich water from the oxygen-minimum zone. 

Another variable is the frequency and magnitude of 

turbidity currents. Holding this variable constant and 

assuming that turbidity currents were more or less continuously redepositing shallow-water or slope-derived, 

continental-margin sediment to deeper sites, accumulation at these deeper water sites was a combination of 

pelagic, hemipelagic, and turbiditic processes. 

In the situation shown diagrammatically in figure 

9A, surface-water productivity is fairly high so that a midwater oxygen minimum exists. Shelf sediments consist 

largely of terrigenous clastic material and organic matter 

with some contributions from authochthonous marine 

organic matter and biogenic calcareous and siliceous 

debris. Sediment that accumulated where the continental margin intersected the oxygen-minimum zone are OCrich, black, and laminated where the dissolved oxygen 

concentration was zero throughout a sufficient thickness 

of the water column. The composition of redeposited sediment depends on local sources. Slope sediments along arid 

continental margins that are characterized by coastal 

upwelling are primarily fine grained and rich in carbonate 

or biogenic silica (for example, sites 398, 400, 402, and 

370). Sediment accumulating in well-oxygenated bottomwater environments are mostly red clay. ·Locally, green clay 

interbeds exist that represent periods of accumulation of 

slightly higher concentrations of organic matter, high 

enough to consume all the dissolved oxygen within the 

sediment and impart a green color (mainly from reduced 

iron) to the sediment, but not enough to have a marked 

effect on the benthic fauna. The result is interbedded red 

and green clay units that are both intensely bioturbated. 

In the situation shown diagrammatically in figure 

9B, there is higher surface-water productivity in response 

to a greater supply of nutrients from more intense 

122 Studies in Diagenesis 

upwelling. In addition, an increase in the influx of terrigenous organic matter takes place as a result of a rise 

in sea level and (or) an increase in amount of terrestrial 

vegetation. This increased supply of organic matter from 

marine and (or) terrestrial sources increases the oxygen 

demand in the water column and in the sediments which 

causes an expansion and intensification of the oxygen 

minimum zone. As a result, sediment accumulating along 

the continental margin is more reduced and contains more 

organic matter. This sediment covers a much greater area. 

Because of the presence of more sediment that contains 

more organic matter available for down-slope transport 

into deeper-water environments, the basin sediment also 

is more reduced and contains more organic matter. In 

extreme cases, when bottom waters are delicately poised 

at a low oxygen level, oxygen-deficient conditions may 

extend into deeper waters, but most of the time the 

bottom-water mass is sufficiently oxic to support at least 

a benthic epifauna. The resulting sediment facies consist 

of dark-green to black, OC-rich sediment deposited in 

continental-margin settings and interbedded green and 

black, OC-rich basinal sediment. Sediment deposited in 

either the continental margin or basinal environments at 

times of extreme oxygen deficiency may be laminated, 

indicating complete elimination of the benthic fauna. 

CONCLUSIONS 

1. Most of the organic matter in OC-rich Cretaceous 

strata in the western and northern North Atlantic Ocean 

is of terrestrial origin, whereas most of the organic matter in equivalent strata in the southeastern North Atlantic and South Atlantic is of marine origin. Mixtures of 

terrestrial and marine organic matter, however, are common in both the North and South Atlantic. 

2. Periodicities of the oxidation-reduction cycles 

usually range from 20,000 to 100,000 years per cycle and 

average about 40,000-50,000 years per cycle. These cycles 

may be due to climatic forcing, similar to Cenozoic 

climate-induced depositional cycles, but they may also be 

due, in part, to periodic sediment redeposition. 

3. Local concentrations of sulfide, sulfate, and carbonate minerals usually are produced by diagenetic 

oxidation-reduction reactions. These reactions are generally related to anaerobic decomposition of organic matter. Phosphate is not always enriched in deep-water black 

shales, and no phosphate minerals have been reported 

from these strata. 

4. Certain trace elements, especially Cd, Cu, Zn, 

Mo, V, Ni, and Cr, are enriched in OC-rich strata. This 

enrichment may be due to the uptake of these elements 

by clays and organic matter and (or) to their coprecipitation with sulfide minerals under reducing conditions. 

The enrichments may reflect deposition of sediment 

under anoxic conditions in the deep sea, or they may be 
inherited from the composition of organic matter originally deposited upslope under low-oxygen conditions and 

subsequently redeposited in deeper water. 

5. Compactional dewatering of some black shale 

sequences resulted in upward diffusion of some reduced 

metal ions after deposition of high concentrations of 

organic matter, resulting in anoxic conditions within the 

sediment. Slowly deposited, oxidized claystones overlying some black shale sequences are enriched in metals, 

particularly Fe, Mn, Zn, and Cu, relative to normal 

pelagic days; this enrichment may be due to upward 

migration of metals from the underlying black shale. 

6. The origin of interbedded more- and less-reduced 

lithologies containing varying amounts of organic matter and varying amounts of pelagic, hemipelagic, and terrigenous sediment is complex and probably is not due to 

any one process. The main variables are the rate of influx 

of terrigenous sediment and organic matter and surfacewater productivity. Variations in rate of supply of these 

two sources of organic matter from shallow-water areas 

to deeper, basinal sites of accumulation produced cyclic 

interbeds of more-reduced and less-reduced lithologies. 

During several periods of time (for example, AptianAlbian and Cenomanian-Thronian), a coincidence of 

several dimatic and oceanographic factors apparently 

resulted in maximum production and preservation of 

organic matter in deep-sea environments that otherwise 

were fairly well oxygenated and were preserving lesser 

amounts of organic matter. Bottom-water anoxia may 

have occurred during these periods of maximum accumulation of OC-rich strata, but the commonly observed cyclic 

interbeds of green and black strata are mainly the result 

of diagenetic fluctuations in oxidation-reduction potential in the sediment in response to fluctuation in amount 

of organic debris being deposited, regardless of source. 

Much of the organic debris may have been transported 

to basinal sites from continental-margin settings within 

oxygen-minimum zones by turbidity currents. 

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Arthur, M.A., 1979a, North Atlantic Cretaceous black shales; 

the record at site 398 and a brief comparison with other 

occurrences, in Sibuet, J-C., Ryan, W. B. F., and others, 

Initial reports of the Deep Sea Drilling Project, v. 47, part 

2: W'ashington, D.C., U.S. Government Printing Office, 

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___ 1979b, Origin of Upper Cretaceous multicolored 

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___ 1982, The carbon cycle-Controls on atmospheric 

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P. R., and Tucholke, B. E., eds., The geology of North 

America, Volume M-The western North Atlantic region: 

Boulder, Colo., Geological Society of America, p. 617-630. 

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Milankovitch-like cycles, in Berger, A., Imbrie, J., Hays, 

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Archean to Present: American Geophysical Union 

Monograph 32, p. 504-529. 

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D. J., eds., Fine-grained sediments, processes and products: 

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sedimentation rates in the ocean: American Journal of 

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of Late Cretaceous age in the western interior of the 

United States: Geochimica et Cosmochimica Acta, v. 28, 

p. 1579-1604. 

____ 1979, Black shale-Its deposition and diagenesis: Clays 

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sedimentary evolution and aspects of tectonic history, in 

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Origin and Diagenesis of Cretaceous Deep-Sea, Organic-Carbon-Rich lithofacies, Atlantic Ocean 127 
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128 Studies in Diagenesis 

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Limestone Diagenesis-Some Geochemical Considerations 

By LYNTON S. LAND, Department of Geological Sciences, 

University of Texas at Austin, Austin, Texas 78713 

CONTENTS 

Abstract 130 

Introduction 130 

Limestone diagenesis 130 

Mc!teoric stabilization of shallow-water sediments 130 

Burial diagenesis of chalk 132 

lnltermediate situations 135 

Conclusions 135 

References cited 135 

FIGURES 

1. Scatter plot of o13C and o180 for Holocene sediments and Pleistocene 

limestones 133 

2. Strontium and oxygen isotopic changes during diagenesis of Miocene chalk, 

northern Jamaica 134 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 129 
Abstract 

Carbonate sediments are deposited as metastable assemblages of aragonite, Mg-calcite, and (or) Ca-rich dolomite. The 

transformation of such assemblages of metastable minerals 

toward rocks composed of stable assemblages of pure calcite 

and ideal dolomite (stabilization) takes place continuously, but 

episodically, until the rocks are destroyed by weathering or 

metamorphism. Stabilization in surficial environments, especially those associated with meteoric recharge, dominates 

limestone diagenesis. Geochemical techniques have not yet 

evolved to the stage where environments of diagenesis can be 

uniquely determined. 

INTRODUCTION 

Experimental work in the 1950's established the 

basic phase equilibria for the system CaC03

-MgC03• 

Calcite was demonstrated to be thermodynamically stable 

with respect to aragonite over all normal (not tectonically 

overpressured) burial conditions (Jamieson, 1953). Dolomite was shown to be stable with respect to the assemblage 

calcite + magnesite (Graf and Goldsmith, 1955; Goldsmith and Heard, 1961). The finding that modern shallowwater carbonate sediments are composed of the metastable minerals aragonite and magnesian calcite 

(Mg-calcite) was first met with skepticism. It has become 

clear, however, that these metastable minerals formed the 

precursor sediment for many ancient limestones and that 

massive stabilization reactions are necessary (accompanying or in addition to compaction and cementation) to 

form carbonate rocks from metastable precursor 

sediments. 

The concept of progressive, stepwise stabilization 

(that is, the transformation of metastable mineral assemblages toward stable assemblages), first proposed by 

Ostwald (1911, p. 444), dominates all aspects of diagenesis, 

not only the diagenesis of carbonate sediments. Most 

sediments are deposited as metastable assemblages of 

minerals which transform by stepwise stabilization as 

temperature and fluid compositions change during 

diagenesis. The question is: can the temperature and fluid 

compositions which induce the diagenetic changes be 

documented? 

Although the mineralogy of lower Paleozoic 

shallow-water sediment (especially carbonate mud) is 

unknown, we do know that metastable phases existed. 

Early molluscs were composed of aragonite just as most 

are today, and early marine aragonite cements and ooids 

did indeed form (Sandberg, 1983). There is growing 

evidence that either calcite or Mg-calcite may have been 

much more common components of lower Paleozoic sediment than was aragonite (Wilkinson, 1982). PostPaleozoic carbonate sediments probably had a mineralogy 

not too dissimilar from today's sediment-that is, about 

two-thirds aragonite and one-third Mg-calcite. Skeletal 

130 Studies in Diagenesis 

pelagic sediment, which first appears abundantly in the 

Mesozoic, has always been composed of calcite, yet it too 

is subjected to stabilization reactions. 

LIMESTONE DIAGENESIS 

Recognizing that overlap exists, carbonate sediments 

can conveniently be subdivided into shallow-water sediment, initially composed of Mg-calcite and aragonite, and 

pelagic sediment composed primarily of calcite. The 

stabilization of sediment dominated by aragonite and Mgcalcite to limestone composed of calcite is commonly, but 

not necessarily universally, accomplished by early influx 

of meteoric water. Pelagic sediment, in contrast, is compacted or transformed by convection-dominated burial 

diagenesis in sea water. These two diagenetic realms, (1) 

nearly open-system meteoric dissolution-precipitation, 

and (2) more nearly closed-system stabilization dominated 

by compaction, are the "end member" processes that form 

the basis for the present discussion. 

Meteoric stabilization of shallow-water sediments 

The stabilization of aragonite and Mg-calcite as the 

result of the influx of meteoric water has been intensively 

studied in Pleistocene limestones. Unfortunately, too 

many authors have used these examples as actualistic 

models for ancient limestone diagenesis without realizing the sample bias which has been introduced by the complexity of Pleistocene eustatic sea-level changes. Most 

early studies concentrated on Pleistocene rocks subareally 

exposed on the coasts of tropical and subtropical land 

masses (for example, Bermuda, Gross, 1964; Barbados, 

Matthews, 1968; the Red Sea, Gavish and Friedman, 1969; 

Jamaica, Land and Epstein, 1970). These sediments were 

deposited during interglacial highstands of sea level, 

especially at + 5 m, 120,000 years ago (see, for example, 

Harmon and others, 1981). All these deposits have been 

in the vadose zone (the unsaturated zone above the water 

table) for most of their history. Many early studies were 

therefore biased toward vadose meteoric, rather than 

phreatic meteoric diagenesis. Diagenesis within phreatic 

ground-water lenses certainly dominates most limestone 

diagenesis, yet few studies of Pleistocene diagenesis deal 

with it, except where extensive uplift has occurred, because 

drilling is commonly required to obtain samples. Moreover, the rapid oscillations of sea level during the 

Pleistocene did not deposit extensive tabular beds, nor did 

they allow large-scale, stable, steady-state, meteoricphreatic systems to be established. The overprinting of 

early Pleistocene phreatic events by later sea-level oscillations, which crested near the position of modern sea level, 

further complicates the identification and interpretation 

of phreatic events. 
Studies of Holocene phreatic-meteoric diagenesis 

are rare (see, for example, Halley and Harris, 1979; Budd, 

1985) because the modern rise of sea level has been so 

rapid that only small quantities of sediment have been 

generated. Most modern carbonate sediment has been 

deposited below present-day sea level. Insufficient time 

has been available for extensive coastal recharge to 

permeat(: most modern sediments. Land (1973a) and 

Steinen and Matthews (1973), in addressing phreatic as 

opposed to vadose diagenesis, concluded that diagenesis 

in the phreatic zone is much more rapid than in the vadose 

zone. Phreatic-meteoric diagenesis produces both replacement and cementation textures more similar to ancient 

limestones than does vadose diagenesis. Extrapolating 

Holocene stabilization and cementation rates observed in 

the Schooner Cays, Budd (1985) estimated that phreaticmeteoric diagenesis could accomplish destruction of all 

metastable aragonite and generate a limestone composed 

of calcite in less than 50,000 years. 

Petrographic and geochemical data typically show 

large changes in whole-rock elemental and isotopic 

chemistry and in the chemistry of individual components 

during alternation to calcite. Depletion in strontium, 

magnesium, 13C, and 180 are typically observed. The 

stabilization of aragonite and Mg-calcite takes place in 

relatively open aqueous systems by dissolutionreprecipitation. Although various words have been used 

to describe the processes (for example, inversion (incorrectly!), neomorphism, recrystallization, incongruent solution), it is convenient to categorize reactions as proceeding 

either by replacement or by wholesale dissolution. In 

reality, re:placement is simply dissolution, with subsequent 

precipitation soon afterward, commonly by "templating" 

the phase which is being dissolved in submicroscopic 

voids. Thus, it is clear why replacement need not occur 

in truly open chemical systems. The degree of fluid communication between local solutions causing dissolution, 

those causing precipitation, and the larger reservoir of 

water from which they are all derived may vary considerably depending on such factors as local rock 

permeability, fluid flow rates, diffusion gradients, and 

crystal nucleation rates. Precipitated (replacement) crystals 

may form far from equilibrium with the bulk-water reservoir and be affected strongly by the chemistry of the phase 

which is being dissolved (Brand and Veizer, 1980, 1981). 

A good example of this problem is the stabalization of high-strontium marine aragonite found in corals, 

modern ooids, and some marine cements (but not 

molluscs). This kind of aragonite typically has strontium 

contents of 7,000-10,000 ppm (parts per million), and 

o180 values of about -2 to -4 per mil (PBD). An "equilibrium" marine aragonite, assuming 25 OC and o180 

seawater = + 1 per mil, should contain about 8,200 ppm 

strontium and have a o180 value of -2.5 per mil (PDB). 

Natural departures from "equilibrium" are caused by local 

variations in temperature, o180 of the water, kinetic 

factors, and "vital effects" in the case of skeletal precipitates. Pleistocene aragonite replaced by calcite (that is, 

displaying relict textures such that wholesale dissolution 

can be ruled out) varies from 140 to 2,000 ppm strontium, 

and from -0.5 to -3.7 per mil o180 (Land, 1973b; Land 

and Epstein, 1970; Martin and others, 1986). Calcite in 

equilibrium with coastal meteoric water (25 OC, o180 

water = -3.5 per mil) should have a o180 value of about 

-5 per mil and a strontium content of less than 200 ppm. 

Clearly, most of the calcite which diagenetically replaced 

Pleistocene aragonite is not in equilibrium with an open 

meteoric system. Almost all replacement calcite seems to 

form in partly open systems, the "openness" of which was 

locally highly variable (Pingitore, 1976), and locally related 

to the permeability of the allochem being replaced (Martin and others, 1986). Calcite that formed from aragonite 

in a rigorously closed system has not been identified. At 

the present state of the art, modeling (see, for example, 

Morrow and Mayers, 1978) has not yet proceeded to the 

point that this process can be quantitatively duplicated 

mathematically. Modeling of diagenesis (including all 

other phases), using both carbon and oxygen isotopes (and 

strontium isotopes for some samples) and several trace 

elements (Sr, Mg, Mn, Fe) simultaneously holds considerable promise for the future. 

In an example from Permian rocks, Given and 

Lohmann (1985, p. 430) proposed that nonluminescent 

calcite replaced aragonite and Mg-calcite marine cements 

in a "low water/rock, 'closed' system" such "that the 

isotopic signature of the nonluminescent calcite also 

represents the original isotopic composition of the marine 

cement~' They did not demonstrate, however, that the 

replacement took place in a system strictly closed to 

cations such as strontium and magnesium. A comparison 

of the reservoir sizes between calcite and water in a rock 

of reasonable porosity demonstrates that whereas the 

reservoirs of carbon and cations in the rock overwhelms 

the carbon (HC03

-) and cation reservoirs in the pore 

water, the reservior of oxygen is more nearly of the same 

size. Therefore, the oxygen isotope composition of calcite 

will be most sensitive to alteration during diagenesis. If 

the oxygen isotope composition is strictly conserved during diagenesis, no other component can be modified whatsoever. It is impossible to convert strontium-rich aragonite 

or Mg-calcite to strontium- and magnesium-depleted 

calcite and rigorously conserve the oxygen isotope composition of the precursor phase. 

It is possible, however, to alter a primary aragonite 

or Mg-calcite component to calcite, either by advective 

or diffusive diagenesis, and cause no change in the o180. 

If diagenesis takes place by slow stabilization in seawater 

(Saller, 1984), the oxygen isotope composition of the 

calcite replacement product might be identical to the 

precursor aragonite or Mg-calcite, whereas the cation content might be quite different. In this situation, however, 

the o180 of the replacement calcite is in no sense 

Limestone Diagenesis-Some Geochemical Considerations 131 
"original:' It is the same as the original phase only by happenstance in that diagenesis (stabilization) occurred in the 

same solution (seawater) from which the metastable phase 

initially precipitated. Given and Lohmann (1985) did not 

document a composition for primary Permian marine 

cement; rather, they documented the conditions of 

diagenesis, which may have been induced by 180-

depleted Permian seawater (Given and Lohmann, 1985) 

or by meteoric water (Given and Lohmann, 1986). 

Many authors have emphasized the depletion of 

Pleistocene limestones in 13C with increased diagenesis 

(Allen and Matthews, 1977, 1982) and the correlation between o180 and o13C (fig. 1). Although these observations 

are correct, they are almost all based on sampling rocks 

which have been subaereally exposed (that is, immediately 

adjacent to soils) for varying lengths of time. Soil carbon dioxide depleted in 13C because of oxidation of 

organic matter is produced during vadose diagenesis in 

semitropical (humid) climates. A similar large-scale input 

of organic-derived carbon dioxide is not to be expected 

in an arid climate (few plants), or during phreatic 

diagenesis in which the carbon reservoir in the rocks is 

dominant. Additionally, the incorporation of significant 

amounts of organic (13C-depleted) carbon in carbonate 

cement is not limited to oxidation associated with soils. 

In fact, the isotopic signature of vadose diagenesis ( depletion in 13C and 180) can be nearly identical to that of 

burial cementation of some sequences, such as the Texas 

Gulf Coast Teritary (Land, 1984; Lundegard and others, 

1984). 13C-depleted cements are neither unique to or 

necessarily characteristic of vadose diagenesis. 

Many Pleistocene and older limestones composed 

completely of calcite are still inhomogeneous in their 

distribution of trace elements and stable isotopes (see, for 

example, Benson and Matthews, 1971, their fig. 5; Prezbindowski, 1984, his figs. 7 and 8). Such inhomogeneities 

can provide valuable clues to diagenetic processes; but the 

inhomogeneous nature of the crystals indicates that they 

have not yet fully stabilized. The capacity for continued 

stabilization continues to exist until thermal metamorphism homogenizes (or destroys) the limestone. 

Burial diagenesis of chalk 

Although many authors have assumed calcite to be 

a stable constituent of limestones, it clearly need not be. 

Small differences in magnesium or strontium concentration, for example, significantly change the solubility of 

calcite. Given geologic times for reaction, the more soluble phases, even though they are only slightly more soluble, do react. Schlanger and Douglas (1974) usefully 

expressed this idea in the form of a "diagenetic potential!' Small, trace-element-rich crystals have more 

"diagenetic potential" (they are more soluble) than large 

pure crystals, all other things being equal. 

132 Studies in Diagenesis 

Pelagic chalk consists of a mixture of skeletal components dominated by coccoliths and pelagic foraminifera. Each taxon exhibits slightly different skeletal microarchitecture and skeletal chemistry (see, for example, Savin 

and Stehli, 1974), and pelagic sediment is highly inhomogeneous even though an X-ray diffraction pattern 

shows it to be pure calcite (aragonitic pteropods are 

ignored in this discussion because their extreme "diagenetic potential" is obvious; Berner and Honjo, 1981). 

The Deep Sea Drilling Program has demonstrated that 

carbonate ooze transforms into chalk and that chalk 

transforms into limestone with increasing burial. The 

process occurs by "solution-compaction" (Scholle, 1977) 

as more soluble components dissolve (aided by increased 

overburden pressure) and reprecipitate "locallY,' Suprisingly, massive chemical changes in both the trace element 

composition and isotopic composition of the chalk occur 

as burial diagenesis proceeds. In some cases, the burial 

diagenesis of chalk is indistinguishable on the basis of 

oxygen isotope and strontium content from meteoric 

diagenesis (fig. 2; Land, 1979). 

During burial diagenesis of chalk, trace elements 

such as strontium clearly must be removed from the 

compacting system by diffusion and (or) fluid flow. The 

depletion in 180 toward the bottom of some chalk sections (fig. 2) cannot be due solely to replacement at 

elevated temperatures, but must involve introduction of 180-depleted (and 87Sr-depleted, Elderfield and others, 

1982) water from beneath (Land, 1980, his fig. 3), probably driven by thermal convection. Even though the fluids 

participating in burial diagenesis of deep-sea chalk 

sequences are introduced at the bottom of the section (and 

at high temperatures), the chemistry of limestones that 

result strikingly resemble chalk sequences altered by 

meteoric water. Cretaceous chalk exposed by submarine 

erosion is also replaced by a Sr-depleted (but not 180-

depleted, because of the cold temperatures) calcitereaction rind in essentially unmodified seawater (Schlager 

and James, 1978). 

What is important here is that although the 

"diagenetic potential" of calcite chalk is much smaller 

than that of sediment composed of aragonite and Mgcalcite, stabilization still occurs, albeit over a much longer 

time. By analogy, the burial diagenesis of platform carbonate sequences (where solution compaction is accomplished by megascopic stylolitization) can almost certainly 

continue long after the original mineralogy of the sediments has been converted to calcite. Using currently 

available data, it may be impossible to distinguish early 

surficial meteoric diagenesis from burial compaction. 

Scholle and Halley (1985) advocated burial 

diagenesis of platform carbonate sequences as being of 

considerably greater significance than has been previously 

recognized. They pointed out (p. 312) that some studies 

of Pleistocene diagenesis concluded "that there has been 

little or no net addition of carbonate material to the 
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Figure 1. Scatter plot of o13C and o180 for unlithified bulk Holocene sediment (open circle), lithified bulk Holocene sediment 

(plus sign), and Pleistocene limestones altered by meteoric water (closed circle). Data summarized from numerous sources, cited 

in Land, (in press). The large depletion in 13C is due to strong bias toward samples dominated by vadose diageneis. Samples with 

o180 more enriched than -5 per mil are from temperate climates, samples with o180 more depleted than -5 per mil are from 

tropical climates. 

Limestone Diagenesis-Some Geochemical Considerations 133 
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Figure 2. Strontium and oxygen isotopic changes during diagenesis of Miocene chalk, northern Jamaica. Open circles 

are from marine chalk which has never been subareally exposed. Solid circles are from subareally exposed samples. 

The line is generalized from Matter and others (1975) for Shatsky rise limestones recovered by JOIDES, which have 

never been subareally exposed. The similarity between subarea! meteoric and marine burial diagenesis of chalk is 

striking. 

134 Studies in Diagenesis 
section despite as much as 120,000 years of near-surface 

alteration~' Here again, as in interpreting 13C, some commonly studied Pleistocene rocks are not good actualistic 

models for ancient limestones. Scholle and Halley 

emphasized that the 120,000-year-old Miami Oolite has 

lost substantially no porosity compared with modern 

ooids. This fact is not surprising because the rock has been 

exposed for at least 110,000 years in the vadose zone where 

gravitational removal of dissolved CaC03 is the dominant process. Scholle and Halley (1985) ignored data 

which clearly demonstrate that significant porosity reduction can take place in meteoric phreatic lenses. In addition to the examples from Barbados and Bonaire cited by 

~cholle and Halley, the Pleistocene Hope Gate Formation of Jamaica not only consists of 37 percent dolomite, 

but has an average porosity of 5 percent (Land, 1973c, 

his tables 3 and 1 ). Low-porosity limestones can indeed 

form without burial compaction. 

Intermediate situations 

If shallow-water carbonate sediments are buried 

quickly or isolated from recharge zones, they enter the 

subsurface in a highly metastable state. Their diagenesis 

thus differs substantially from sediments that were converted to calcite by recharging meteoric water. 

Rec~ckmann (1981) documented such a sequence in 

which updip grainstones were massively altered by meteoric influx while down-dip wackestones were relatively 

unaltered. If the sequence she studied were to be buried, 

continued stabilization of the down-dip wackestones would 

be inevitable. Burial-driven stabilization of the down-dip 

rocks could result in a very different type of diagenesis 

than the meteoric stabilization that is exhibited by the updip rocks. Burial stabilization has been invoked by Moore 

and Druckman (1981) for the Jurassic Smackover of the 

Gulf Coast and convincingly documented by Saller (1984) 

for the subsurface Tertiary limestones of Enewetak atoll. 

The opposing view, that stabilization of the Jurassic 

Smackover limestones to calcite was accomplished in contemporaneous meteoric systems, was espoused by Wagner 

and Matthews (1982a), Wagner and Matthews (1982b), 

Wagner and Matthews (1983), and Humphrey and others 

(1986). Clearly, the stabilization of aragonite-Mg-calcite 

assemblages in other than surficial meteoric hydrogeochemical environments needs additional attention, 

both from a theoretical (modeling) perspective, and from 

additional well-studied examples. 

CONCLUSIONS 

Current knowledge of diagenesis is highly biased 

toward those environments which have been intensively 

studied, especially the vadose alteration of Pleistocene 

grainstones. Burial diagenesis is only currently receiving 

the attention it deserves, but it is hardly "out of mind" 

(Scholle and Halley, 1985). If studied critically (Prezbindowski, 1984), burial diagenesis rarely approaches the 

importance of early marine and meteoric diagenesis in 

molding limestones petrographically or chemically. Most 

authors have accepted the fact that thick platform 

sequences accumulate at net rates of about 50 m/Ma, 

more than an order of magnitude more slowly than 

analogous modern carbonate sediments accumulate 

(Fisher, 1969). Therefore, meteoric infiltration is inevitable 

under most circumstances as shallow-water carbonate 

sediments "shoal up" to sea level, and it is not surprising 

that meteoric diagenesis dominates the diagenesis of carbonates. In contrast, the diagenesis of shales and many 

sandstones is dominated by thermally driven processes 

during burial. There is no question that subsurface compaction and stylolitization on a range of scales can 

accomplish considerable mass transfer during the 

diagenesis of limestones; however, just as the degree of 

openness of the chemical system which prevails during 

aragonite or Mg-calcite stabilization is argued, the degree 

of openness of compacting limestone (or dolomite, Land, 

1985) sequences is unproven. Geochemistry has the potential to unravel these problems and to document the fluid 

compositions and temperatures that prevail during 

diagenesis. But quantitative studies which integrate a 

variety of different techniques are needed, coupled with 

rigorous mathematical models, to quantify diagenetic 

environments in a convincing manner. 

REFERENCES CITED 

Allen, J. R., and Matthews, R. K., 1977, Carbon and oxygen 

isotopes as diagenetic and stratigraphic tools-Surface and 

subsurface data, Barbados, West Indies: Geology, v. 5, 

p. 16-20. 

__ 1982, Isotope signatures associated with early meteoric 

diagenesis: Sedimentology, v. 29, p. 797-817. 

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Limestone Diagenesis-Some Geochemical Considerations 137 

A Comparative Study of the Dissolution and Crystat Growth 

Kinetics of Calcite and Aragonite 

By EURYBIADES BUSENBERG, U.S. Geological Survey, MS 432, National Center, 

Reston, Virginia 22092, and 

L. NIEL PLUMMER, U.S. Geological Survey, MS 432, National Center, 

Reston, Virginia 22092 

CONTENTS 

Abstract 141 

Introduction 141 

Acknowledgments 141 

Plummer, Wigley, and Parkhurst model for the dissolution of calcite 141 

Experimental 143 

Procedures 143 

Analytical methods 144 

Gas, gas mixtures, and Pco 144 

Calcites and aragonites us~d 144 

Summary of the aqueous model 144 

Calculations 14S 

Experimental results for calcite and aragonite dissolution and crystal 

growth 146 

Forward reactions 146 

Bac:kward reactions 147 

Bac:kward reactions in KOH solutions ISO 

Bac:kward reactions in KHC03 solutions ISO 

Bac:kward reactions in CaC12 solutions lSl 

Bac:kward reactions in KHC03-KOH solutions 1S2 

Bac:kward reactions in Ca(HC03) 2 solutions ISS 

The dissolution region 1S6 

The precipitation region 160 

The backward rate constant k 4 160 

Effects of stirring on rate 161 

Comparison of the kinetics of dissolution and precipitation of calcite and 

aragonite 163 

Comparison of the dissolution kinetics of calcite and aragonite with 

dolomite 163 

Summary and conclusions 166 

References cited 168 

F. A. Mumpton, eel., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 139 
FIGURES 

1-11. Graphs showing: 

1. Rate of dissolution of calcite and aragonite as function of pH 147 

2. Dissolution rates of aragonite and calcite as function of pH and 

Pco 148 2 

3. Backward rate effects of hydroxyl ion activity 151 

4. Rates of calcite dissolution in the absence and presence of bicarbonate 

ion activities 153 

5. Effect of HC03

- activites on calcite and aragonite dissolution 156 

6. Effect of CaC12 concentrations on the rate of dissolution of 

calcite 158 

7. Effect of C03 

2- activity on rates of dissolution of calcite and 

aragonite 159 

8. Rate of dissolution and growth of calcite minus k 1

aH+ and activity 

product of Ca2+ and HC03

- 161 

9. Dissolution of calcite and aragonite as function of -log(l - 0) 165 

10. Dissolution and precipitation rates of calcite and aragonite 166 

11. Comparison of dissolution and precipitation rates of calcite and 

aragonite 167 

TABLES 

1. Description of the calcium carbonate used, locality, and source of 

sample 145 

2. Composition and surface areas of the samples used 146 

3. Rate constants that can be used to calculate the rate of dissolution and 

crystal growth of calcite and aragonite 149 

4. Procedures used to evaluate the backward reactions 150 

5. Observed and calculated minus log rates of dissolution of aragonite as a 

function of pH and CO2 152 

6. Observed and calculated -log rates of dissolution of calcite as a function 

of pH and C02 154, 155 

7. Observed and calculated rates of dissolution of calcite and aragonite in 

CaC12 solutions 157 

8. Observed and calculated log rates of dissolution of calcite and aragonite in 

KHC03-KOH solutions 160 

9. Observed and calculated log rates of dissolution and crystal growth of 

aragonite in Ca(HC03

) 2 solutions 162, 163 

10. Observed and calculated log rates of dissolution and crystal growth of 

calcite in Ca(HC03

) 2 solutions 164 

140 Studies in Diagenesis 
Abstract 

The dissolution and crystal-growth kinetics of calcite and 

aragonite were studied in pure water and in the presence of 

HCI, CaCI2, KHC03' and KOH at various C02 partial pressures 

and at 25°C. In C02-H20 solutions, the forward rate of reaction (R0) of the CaC03 minerals is described by 

where k1, k2, and k3 are rate constants. In the presence of 

Ca2+, HCO -, CO/-, or OH- species far from equilibrium, 

the forward rate decreases due to backward reactions. The 

backward rate contribution of these four ions is described 

by a Langmuir isotherm and is equal to Rb<il = Rcf1; /fA; 

+ B;a), where Rb<il is the backward rate contribution of the 

ith species, A and 8 are rate constants, and a; is the activity 

of the ith species. Combined, R = R0 - Rb describes the net 

rate of dissolution (R) of calcite and aragonite in a wide range 

of solution compositions. The dissolution equation can be used 

in the 0 (0 = IAP/K) range of 0.0 to about 0.6 for calcite and 

aragonite, where lAP is the ion activity product and K is the 

equilbrium constant. 

The rate of crystal growth of calcite and aragonite was 

investigated in Ca(HC03)

2 solutions at C02 partial pressures 

of 1 o-3

·

5 to 1 atmosphere and is described by the equation 

R = k4 (cE - aca2+aHC03-), 

where k4 and care constants and E is the equilibrium activity 

product of Ca2+ and HC03- of the CaC03 polymorph at the 

bulk C02 partial pressure. This crystal growth equation 

accurately describes the rate of precipitation of calcite in the 

0 range of 1.5 to more than 100 and aragonite in the 0 range 

of 2.0 to more than 100. A Freundlich isotherm describes the 

undersaturation region near equilibrium in the 0 range of 0.6 

to 1.0 for calcite and aragonite. A similar Freundlich equation 

may describe the supersaturated region near equilibrium in the 

0 range of 1.0 to 1.45 for calcite and 1.0 to 2.0 for aragonite. 

The rates of dissolution of aragonite are slightly greater 

than for calcite in solutions of the same composition; however, 

the rate of crystal growth of aragonite is significantly slower 

than that of calcite in supersaturated Ca(HC03

)

2 solutions. 

These kinetic factors may explain the near absence of aragonite 

in freshwater environments. The same forward rate equation 

can be used to describe the dissolution of dolomite except that 

the rates are half order with respect to H+, H2

C03*, and H20. 

The rates of dissolution of dolomite are 100 times slower than 

those of calcite and aragonite, and relatively small concentrations of HC03- can almost completely stop dolomite dissolution far from "equilibrium." 

INTRODUCTION 

Understanding the time-space relationships of the 

diagenesis of carbonate sediments and rocks requires fundamental knowledge of the kinetics and mechanisms of 

reactions at the surfaces of calcite, aragonite, and 

dolomite. An extensive literature on the solution kinetics 

of calcite was reviewed by Plummer and others (1979) 

and compared with results and predictions based on a 

model of calcite dissolution (Plummer and others, 1978). 

The dissolution model (reviewed later) was derived from 

experiments in the C02-HC1-H20 system. Most predicted rates were found to agree within a factor of 20 of 

the observed literature values and many within a factor 

of two or better over widely varying chemical conditions 

beyond those investigated by Plummer and others (1978). 

Near-equilibrium predicted dissolution rates were found 

to be systematically faster than the observed. The effects 

of small additions of as much as a few millimoles of salts 

such as CaC12 and KHC03 were largely accounted for 

via thermodynamic corrections calculated by the aqueous 

model used to define the saturation state of the bulk solution and the calcite surface equilibrium pH (both essential factors in the dissolution model). No data exist, 

however, on the effects of higher concentrations of added 

salts. Reddy and others (1981) showed that measured 

rates of crystal growth of calcite in calcium bicarbonate 

solutions were closely predicted by the dissolution model 

at C02 partial pressures of 0.03 to 0.30 atm (atmosphere), but again larger departures were found near 

equilibrium. No studies of the dissolution and crystal 

growth kinetics of aragonite in C02-H20 solutions are 

known, although the dissolution kinetics of dolomite 

was recently investigated by Busenberg and Plummer 

(1982). 

The present study examines more closely the effects 

of larger concentrations (as much as 2 M) of CaC12, 

KHC03, K2C03 and KOH on the dissolution kinetics of 

calcite and aragonite far from equilibrium in C02-H20 

solutions at 25 OC. The work of Reddy and others (1981) 

has been extended to a wider range of P co and supersaturation in an investigation of the crystal2

growth rate 

of calcite and aragonite in Ca(HC03) 2 solutions. Also, 

the dissolution of calcite and aragonite has been studied 

in detail near equilibrium in an attempt to interpret departures from the dissolution model. Finally, the dissolution 

mechanisms of calcite and aragonite are compared with 

Busenberg and Plummer's (1982) recent study of the 

mechanisms of dissolution of dolomite. 

Acknowledgments 

The authors are grateful to Eric T. Sundquist and 

Michael M. Reddy for their careful reviews of the 

manuscript and their many useful comments. 

PLUMMER, WIGLEY, AND PARKHURST MODEL 

FOR THE DISSOLUTION OF CALCITE 

Plummer and others (1978) studied the kinetics of 

dissolution of Iceland spar calcite in dilute solutions far 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 141 
from equilibrium as a function of pH, P co and temper- 2 

ature using the pH-stat method. The dissolution behavior 

of calcite ''near'' equilibrium was evaluated using the 

free-drift method. The experimental procedures of these 

methods were described in detail by these authors. The 

results of their investigation suggested that three independent (but simultaneous) forward reactions and one 

backward reaction give rise to the observed overall rate. 

At low pH and in the absence of C02 the dissolution rate was found to be proportional to the activity of 

hydrogen ion in the bulk solution. At high pH in the near 

absence of hydrogen ions and C02, the forward rate 

was constant. In the intermediate pH range, the rate 

increased in proportion to the C02 partial pressure. The 

forward rate for the dissolution of calcite (RJ was given 

by the expression: 

where k1, k 2 and k 3 are rate constants dependent on 

temperature, and H2CO; is equal by convention to 

H2C03° + C02(aq)' The units of rate are in mmol·cm·2·s·1 

of calcite dissolved. The temperature dependences of the 

forward rate constants at 1,800 revolutions per minute 

of Plummer and others (1978) are given by the following 

expressions: 

log k 1 = 0.198 - 444/T 

log k2 = 2.84- 2,177/T, and 

log k 3 = 1.10- 1,737/T, 

(2) 

(3) 

(4) 

where Tis temperature in Kelvins. The Arrhenius activation energies were found to be about 2, 10, and 7.9 

kcal·mol-1 for k., k2, and k3, respectively. 

Using the free-drift method, Plummer and others 

(1978) found a direct relationship between the activity 

product of Ca2+ and HCO; and the decrease in the net 

rate relative to the forward rate given by equation 1. The 

backward rate (Rb) was given by: 

(5) 

where k4 is a rate constant. Equations 1 and 5 were 

combined to yield the overall rate (R) for the dissolution 

of calcite, 

A significant achievement of the mechanistic model 

of calcite dissolution of Plummer and others (1978) was 

the theoretical derivation of the rate constant k using • 4 

a reaction mechanism model. This dissolution model is 

an extension of the adsorption heterogeneous reaction 

model of Mullin (1972). It assumes that a thin, relatively 

142 Studies in Diagenesis 

immobile surface layer exists between the surface of the 

crystal and the shear layer (the hydrodynamic boundary 

layer), and that the reactions occur on a small number 

of sites on the mineral surface. The model assumes that 

the reaction of surface CaC03 with H+ is much faster 

than reaction with H2CO; and H20. The model then 

assumes that the flux of bulk solution H20 and C02 

across the boundary layer is faster than the heterogeneous 

reaction at the mineral surface. Therefore, the activities 

of H20 and H2CO; at the mineral surface are about the 

same as their bulk solution activities. Because the H+ 

flux across the boundary layer is slower than the 

heterogeneous surface reaction, the activity of H + the 

surface layer is assumed to be the equilibrium (saturation) value at the bulk solution C02 partial pressure. 

For the backward reactions to occur, interaction 

of the boundary layer reactants and products with the 

surface layer is necessary. Several steps may be involved 

in the interaction of the boundary layer ions with the solid 

su~face. The initial step may involve a collision of 

CaHCOj with the surface layer anions CO~-, HCO;, 

and OH- and the formation of an uncharged growth 

unit on the surface. A second step may involve dehydration and incorporation of the growth unit into the crystal 

lattice. The reactions can be written to show the relationship between the boundary layer (0) and the surface layer 

(s) ions for the forward reactions: 

For the backward reaction between CaHC03 + <O> and the 

surface layer anion co~-(s)' 

This reaction can be simplified to read: 

CaHC03+<o>-+CaC03 + H+<sr (11) 

The other backward reactions involved are: 

CaHC03 + <o> + OH<;> -+ CaC03 + H20
' and (12) 

CaHC03 + <o> + HC03- 
 -+ CaC03 + H2C03 °
' (13) 

The rates of reaction from equations 7-9 can be written: 

RO(t) = k; aH+ - k~ acaHco + ' (14) (0) 3 (0) 

R0(2) = k2 aH co 0 - k4 acaHco + aHco- ' and (15) 2 3 (0) 3 (0) 3 (s) 

R = k a - k a + a (16) 0(3) 3 H20(0) 4 CaHC03 (O) OH-(s)' 
The total rate of dissolution is equal to 

(17) 

Because the surface layer is assumed to be in equilibrium 

with the calcite surface, the net rate of reaction of the 

surface species is zero. From the assumption that calcite 

reacts fast with H+ and more slowly with H2

co; and 

H20, the surface layer activities of H2CO; and H20 are 

about the same as their bulk fluid activities, but 

significantly different for H+. Therefore, k~ aH co* and 2 3(0) 

k; aH 0 can be replaced by k2 aH co • and k3 aH 0 , 2 (0) 2 3(b) 2 (b) 

where the subscript (b) designates bulk solution. At low 

pH, the relationship 

(18) 

is a good approximation. At high pH the k; aH+ term is 

(0) 

small in comparison to the other two forward terms and 

can be neglected. From equations 14-16, 

Because the backward and forward terms must be equal 

at equilibrium as required by the principle of microscopic 

reversibility, equations 14-16 yield: 

(20) 

where K2, K W' and Kc are equilibrium constants for the 

reactions: HC03

- = H+ + CO~-; H2

0 = H+ +OH-; 

and CaC03 (calcite) = Ca2+ + CO~-. Substituting 

equation 20 into equation 19 gives 

(21) 

The ba<:kward rate constant k4 can be calculated from 

equation 21 by replacing k; with the experimentally 

determined value k1 • 

The model assumes that k; = k2 and k; = k3 , and 

that the bulk activities of H2co; and H20 are the same 

as the surface values. With these approximations, the 

theoretical temperature and P co dependence of k 4 can 2 

be calculated. The Plummer, Wigley, and Parkhurst 

(1978) model for calcite dissolution has been used to 

explain the results of many published dissolution studies 

(Plummer and others, 1979) and to calculate the rates of 

precipitation of calcite from supersaturated calcium bicarbonate solutions (Reddy and others, 1981). The model 

has also been used to quantify other experimental and 

natural phenomena involving calcite dissolution and 

precipitation (Dreybrodt, 1980, 198la, 198lb, 1981c; 

House, 1981a, 1981b). 

EXPERIMENTAL 

Because the work of Plummer and others (1978) 

was based only on dissolution measurements of calcite 

in the presence of C02

-H20-HC1, a new study was made 

to determine the rates of dissolution and growth of the 

polymorphs, calcite and aragonite, at 25 °C and a stirring rate of 260 revolutions per minute in the presence 

of C02-H20-HCl-KHC03-KOH-CaC12 • The single 

crystals used were suspended with platinum wires in the 

reaction vessels containing solutions of known composition. Every effort was made to maintain a constant composition of the solution by limiting the change in weight 

to 0.1-0.2 mg, by using large volumes of solution (800 

milliliters) and by maintaining the pH constant (to ± 0.03 

pH units). All reference to the rate constants ki' k2, and 

k 3 listed here pertain to new revised values determined 

in this experimental study. 

Procedures 

All experiments were performed in jacketed oneliter Pyrex reaction vessels maintained at 25 OC ± 0.05 OC 

with Haake (Model FE and FK2) constant temperature 

circulating pumps1• In experiments performed at less 

than pH 4 or greater than pH 9, changes in pH caused 

by the dissolution of CaC03 crystals were less than 0.03 

pH units. These low pH changes were achieved by using 

large volumes of solution (800 millileters) and dissolving 

small amounts of CaC03 (0.1-0.2 mg). In the pH range 

4-9, the pH was maintained to within 0.03 pH units by 

adding dilute HCl into the reaction vessels with pH-stats 

(Radiometer pH meter PHM 64, titrator TTT 60, and 

autoburette ABU 13). 

Experiments were performed using (1) single 

crystals, and (2) particles or powders. The first type of 

material consisted of single crystals, cleavage rhombohedrons (calcite), or aragonite crystals cut into welldefined geometrical shapes. The crystals were reacted with 

dilute HCl and washed several times with distilled water 

before initial use. These crystals were suspended with 

platinum wires into the reaction vessels, and the rates of 

dissolution or crystallization were determined by weight 

loss or gain of the samples. The crystals were washed 

carefully with deionized water immediately after each 

experiment, dried between 10 and 60 min at 110 OC and 

1

Use of trade and company names is for descriptive purposes only 

and does not imply endorsement by the U.S. Geological Survey. 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 143 
allowed to cool overnight before weighing. Differences in 

weight used to calculate rates were determined with a 

Mettler electronic balance (Model ME22-BA25-BE22) 

with an accuracy of better than 2 p.g. Weight losses were 

checked by chemical analysis of the Ca released into the 

known volume of solution. ]jxcellent agreement was found 

between these two methods. A more detailed account of 

the procedures used is given by Busenberg and Plummer 

(1982). The surface areas of the crystals were calculated 

from the dimensions and the geometrical shape of the 

crystals and ranged from about 2 to 10 cm2. 

In the second group of experiments, rates were 

measured in stirred suspensions. These experiments were 

used mainly to study dissolution near equilibrium. The 

material used in some experiments were size-sorted fragments of calcite and aragonite. The surface areas were 

calculated from the sizes and geometry of the fragments 

by the method described by Plummer and others (1978). 

Size-sorted oolites of aragonite were used in some experiments. The surface area was calculated from the mean 

diameter of the spheres. Two oolite samples, C1 and C2, 

(table 1) had calculated surface areas of 85.6 and 56.4 

cm2·g-1, respectively. About 20 g of oolite sample was 

used per experiment. Other materials used in batch experiments were reagent-grade calcite powder (J. T. Baker 

Chemical Co., No. 5-1294), and a pure synthetic aragonite 

powder prepared and aged in artificial seawater at 7 5 °C 

(Plummer and Busenberg, 1982). The BET surface areas 

of the two powders were 0.19 m2·g-1 for the calcite and 

1.2 m2·g-1 for the aragonite. Reaction progress in the 

batch experiments was calculated from changes in pH or 

Ca of the solution as a function of time. 

Analytical methods 

The Ca released in the free-drift experiments was 

determined from aliquots filtered through 0.45-p.m membrane filters, and obtained at regular intervals throughout 

the experiments. Calcium was determined by ethylenediamine tetraacetic acid (EDTA) using calcein as the endpoint indicator by the procedure of Diehl (1964). The 

EDTA was standardized with a primary CaC03 standard 

(J. T. Baker Chemical Co., Ultrex CaC03 No. 4-4918), 

and the EDTA concentration was determined with an 

accuracy of better than 0.1 percent. The accuracy of Ca 

analyses was better than 0.5 percent. The Ca released from 

single crystals was determined by standard atomic absorption procedures using a Perkin-Elmer Corporation Model 

503 instrument. Strontium and magnesium were also 

determined by standard atomic absorption procedures. 

The pH was measured with Radiometer Model 

PHM 64, Orion Research, Inc. Model 801 and Fluke 

Model 8800A instruments. Radiometer Cat. No. OK 

2104C and Orion Research Model 81-0200 combination 

pH electrodes were used. The electrodes were standardized 

144 Studies in Diagenesis 

with 4.01 ± 0.01, 6.86 ± 0.01, and, when necessary, 

9.18 ± 0.01 Fisher Scientific Co. buffers. 

Gas, gas mixtures, and Pc02 

CO , N2, and C02

-N2 gas mixtures prepared by Air 

Products,

2 

Inc. were used in all the experiments. Analysis 

of some of the mixtures by gas chromatography indicated 

that the actual compositions were well within one percent 

of the stated compositions. Before bubbling the gas into 

the reaction vessel, all gases were bubbled through 

deionized water at the same temperature as the reaction 

vessel to saturate the gas with water vapor. The C02 partial pressure was calculated from the measured room 

barometric pressure (B) using the equation: 

p _ B- V + 0.4 hiD X (22) co2 - • C02' 

760 

where B is the measured barometric pressure in millimeters 

Hg, corrected for temperature and latitude; V is the 

saturation vapor pressure of water, in millimeters of Hg, 

at the experimental temperature; h is the depth in 

millimeters of the gas bubbler in the reaction vessel; D 

is the specific gravity of mercury (13.6); and Xc02 is 

the volume fraction of C02 in the gas (Plummer and 

Busenberg, 1982). 

Calcites and aragonites used 

A description of the calcites and aragonites used in 

this study is given in table 1. Table 2 is a compilation of 

Sr and Mg concentrations, surface areas, and weights of 

the samples used per 800 milliliters of solution. 

Summary of the aqueous model 

Plummer and Busenberg (1982) evaluated the 

equilibrium constants for the system CaC03

-C02-H20 

between 0 ° and 90 OC and determined the solubility constants of calcite and aragonite over a wide range of C02 

partial pressures in this temperature range. The same 

aqueous model and data base are used in this paper. The 

necessary reactions and the logarithms of the equilibrium 

constants for this system at 25 OC (Plummer and 

Busenberg, 1982) are as follows: 

= H+ + OH- log Kw = -13.992 (23) 

log KH = -1.468 (24) 
Table 1. Description of the calcium carbonate used and locality and source of samples 

Sample 

A 

B 

M, M1 

N 

0 

s 

C1, C2 

AA 

CM, CM1 

CA 

HC03

-

Mineral 

Aragonite-----

Aragonite-----

Aragonite-----

Aragonite-----

Aragonite-----

Aragonite-----

Aragonite-----

Aragonite-----

Calcite-------

Calcite--.. ----

= H++c~-

Ca2++ HC03

- = CaHCOj 

Ca2++ COi- = CaC03 ° 

Ca2++ OH- = CaOH+ 

Description 

Pseudohexagonal twin crystal, 1 by 4 em, clear. 

Pseudohexagonal twin crystals, 

2.5 by 3 em, translucent, · gray. 

Pseudohexagonal twin crystals, 

3 by 3 em, translucent, 

light gray. 

Pseudohexagonal twin crystals, 

0.8 by 1 em, translucent. 

Pelecypod genus Area---------

Massive, fine grained, 

opaque, white. 

Oolites---------------------

Synthetic, needles 

prepared in artificial 

seawater at 75°C. 

Large, clear cleavage fragments. 

Synthetic, rhombohedrons, 

analytical reagent 

chemical. 

Locality 

Aragon, Spain---------

Bastennes, France-----

Molina de Aragon, 

Spain. 

Artesia, N. Mex.-----: 

South Carolina--------

Compton Ma 11 et, 

Somerset, United 

Kingdom. 

Stabl i zed oolite 

bar--North of 

Andros Island, · Bahamas. 

Chihuahua, 

Mexico. 

Source 

Peter C. Avery, SUNY at Buffalo, 

Buffalo, N.Y. 

Ward •s Natural 

Science Establishment, Inc., 

Rochester, N.Y. 

Ward •s Natural 

Science Establishment, Inc., 

Rochester, N.Y. 

Walter F. White, 

u.s. Geo 1 ogi ca 1 

Survey, 

Towson, Md. 

Present paper. 

Peter G. Avery, SUNY at Buffalo, 

Buffalo, N.Y. 

Randolph Steinen, 

University of 

Connecticut, 

Storrs, Conn. 

Present paper. 

Ward • s Natural 

Science Establishment 

Inc., Rochester, N.Y. 

J. T. Baker 

Chemical Co., 

Phillipsburg, N.J. 

log K2 = -10.329 (26) were calculated from the expression of Reardon and 

Langmuir (1976). A detailed evaluation of the aqueous 

log KCaHco+ 1.11 (27) model used and equilibrium constants can be found in 

3 Plummer and Busenberg (1982.) 

log Keaco o 3.22 (28) 3 

log KeaoH+ 1.3 (29) 

Calculations 

For single crystals suspended into the various soluCaC03 (calcite) = Ca2+ + c~- log Kc -8.480 (30) tions, rates were calculated from the changes in weight 

CaC03 (aJcagonite) = Ca2++ c~-Iog KA -8.336 (31) 

Individual ion-activity coefficients were calculated 

using the modified version of the extended Debye-Hiickel 

equation given by Truesdell and Jones (1974). The activity 

coefficient of CaHC03 +was assumed to be equal to that 

of HC03-. Individual ion activities of uncharged species 

using the equation 

R = (W/G)/(t X S) (32) 

where R is the rate in mmol·cm-2·s-1, W is the weight 

gained or lost in milligrams, G is the gram formula weight 

of calcite or aragonite in mg·mmol-1, t is the duration of 

the experiment in seconds, and S is the surface area in 

square centimeters. Equation 32 was also used to obtain 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 145 
Table 2. Compositions and surface areas of the samples used 

[Leaders ( ---) indicate not determined; ppm, parts per million; cm2, square centimeters; 

and mL, milliliters] 

Stronti urn Magnesium 

Sample (ppm) (ppm) 

A 700 9.0 

B 720 5.2 

M 660 12 

N 720 23 

0 160 21 

s 64 1 

C2 930 87 

C1 940 92 

AA 15 53 

M1 660 12 

CM 

CA 10 9 

rate from the analytical data; here, W represents the total 

calcium released in the solution in milligrams, and G the 

weight of calcium per mole of calcite or aragonite in 

mg·mmol-1• 

Two procedures were used to calculate rates from 

the free-drift runs. In runs in which pH was measured 

as a function of time, rates were calculated by the stepwise central-difference method similar to that of Plummer and Wigley (1976) and Plummer and others (1978). 

For the free-drift runs in which Ca was measured as a 

function of time, the experimental Ca data were fitted 

to the equation: 

Ca = A + Bt + Ct 0·

5 + D log t + Et 0·

33 , (33) 

where A, B, C, D, and E are fitted parameters for a given 

run and tis time. Rates were obtained by differentiation 

of equation 33 at regular time intervals. 

EXPERIMENTAL RESULTS FOR CALCITE 

AND ARAGONITE DISSOLUTION AND 

CRYSTAL GROWTH 

Forward reactions 

Figure 1 shows the effect of bulk solution H+ 

activity on the dissolution rate of single aragonite and 

calcite crystals in the near absence of C02 (N2 

atmosphere) and calcium between pH 2 (HCI-H20) and 

11 (KOH-H20). The slope of log rate as a function of 

pH is -1.0 between pH 2 and 4.5. In this pH range, the 

hydrogen ion dependence of the forward rate is then 

RH+ = k1 OH+. (34) 

146 Studies in Diagenesis 

Total Weight 

surface Surface of 

area area (g/800 ml 

(cm2 ) (cm 2/g) solution) 

3.0-3.7 

2.0-5.5 

2.3-6.7 

4.0-4.4 

2.7-4.8 

5.3-6.4 

56.4 20-35 

85.6 5-20 

12,000 2.0 

190 7.0 

3-10 

1,900 5.0 

Figure 1 shows a region between pH 5 and 9 where the 

rate is independent of pH in the absence of COr In this 

region both H+ and OH- are nearly absent (their 

activities do not exceed I0-5 mol·L-1

), and the forward 

rate is constant according to: 

(35) 

The region of decreasing rate at pH values greater than 

9. 5 has been interpreted as backward reaction between 

the mineral surface and OH- and will be treated in a 

later section of this paper. 

Figure 2A shows the effect of increasing C02 partial pressure on the forward rate for aragonite. This effect 

was evaluated for both aragonite and calcite using: 

which is obtained by rearranging equation 1. The forward 

rate dependence of RH co. was found to be directly proportional to the co: pitrtial pressure for aragonite, 

similar to that found earlier for calcite (Plummer and 

others, 1978): 

(37) 

Figure 2B is a comparison of the results for calcite and 

aragonite. The calculated values for k 1 and k 2 for calcite 

and aragonite are identical within experimental error 

(table 3). The apparent larger value for the total forward 

rate (Rc) on figure 2B results from the larger k 3 constant 

for aragonite. 

The total forward rate far from equilibrium observed 

for both calcite and aragonite in C02-H20 solutions 

obeys the equation given by Plummer and others (1978), 
-4~--~----~----~--~----~----~----~--~----~--~ 

w 

a: 

<( 

::J 

0 

fJ) 

ffi Cl -5 

a..Z oo8 

W UJ 5(1) 

::Effi 

::ic.. 

- -6 :EO: 

w 

ZJ- -w 

u.i:E 

..... - <(Ja:: z 

UJ 

~ (.) -7 

::E 

::I: 

1-

a:: 

<( 

C) 

(!) - Aragonite 

x -Calcite 

g ~~--~----~----~----~--~----~----~----~----~--~ 

2 4 6 8 10 12 

pH 

Figure 1. Dissolution rates of calcite and aragonite single crystals as a function of pH in near absence of C02 (N2 

atmosphere) far from equilibrium. 

where the rate constants k 1, k 2, and k 3 are given in table 

3 for the present experimental conditions. 

Backward reactions 

In the model of Plummer and others (1978), the 

backward rate is interpreted as a reaction between bulk 

solution CaHCO/ and negative surface sites (eqs. 11-

13); however, positive sites in addition to negative sites 

exist on the calcite and aragonite surfaces. Foxall and 

others (1979) showed that the surface positive sites on 

calcite in aqueous solutions are mainly Ca2+ and that 

the Ca2+-anion pair rather than the H+-OH- pair are the 

surface-potential determining ions for calcite. Therefore, 

backward reactions between the positive surface sites and 

bulk solution anions should also occur. All the other 

possible backward reactions between the surface and ions 

not included in the reactions given by equations 10-13 

in the system Ca2+-H20-C02 are evaluated in the next 

few sections of this paper. 

To isolate and determine the backward rate effects 

of individual species, the solution parameters (pH, P co , 

2 

and composition) were adjusted so as to eliminate or 

minimize the effects of all other interfering ions. If this 

was not possible, the combined backward rate effects of 

more than one species were determined. The backward 

rate effect of the species of interest was calculated by subtracting all the effects of the other interfering ions. The 

assumption that backward rates are additive is implied 

by the procedures used. Our results suggest that this 

additive assumption is valid in dilute solutions, but may 

be less reliable in complex concentrated solutions and 

when the saturation 0 for calcite and aragonite exceeds 

0.6 in pure Ca(HC03

) 2 solutions. The failure may result 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 147 
-4.-----.-----~--~----~----~----~----~----~-----.-----, 

0 

z 

0 

(.) 

w 

C/) 

a: 

w 

a.. 

a: 

w 

tu 

~ 

1-

z 

w 

(.) 

w 

a: 

<( 

::::> 

g 

a: 

w 

a.. 

C/) 

w 

...J 

0 

:E 

~ 

...J 

:E 

z 

w 

~ a: 

LL 

0 

:E 

J: 

1-

a: 

<( 

-5 Aragonite 

-6 

A 

§ -7~--~----~----~----~----~----~----~----~--~~--~ 

2 3 4 5 6 7 

pH 

Figure 2. Dissolution rates of single crystals of A, aragonite as a function of pH and gas composition (numerical 

values represent mole percent C02 in the C02

-N2 gas mixtures) far from equilibrium, and 8, calcite and aragonite 

as a function of pH, in near absence of C02 (N2 atmosphere) and 100 percent C02 far from equilibrium. 

from the inability of the model to account for ion-ion 

and ion-surface interactions. 

The effects of OH-, HC03

-, CaH, and COi- procedures and experimental conditions used are outlined in 

table 4. Presumably, the first step in all backward reactions involves the adsorption of a reactive ion into the 

148 Studies in Diagenesis 

surface from the bulk solution. If the rate of backward 

reaction is directly proportional to the activity of the 

adsorbate and the adsorption is described by some 

isotherm, a direct relationship will exist between the bulk 

solution concentration of the ion and the isotherm that 

will be defineo by the changes in the theoretical forward 
Table 3. Rate constants used to calculate the rate of dissolution (equations 53 and 54) and crystal growth (equation 56) 

of calcite and aragonite 

[Unit of rate calculated from, constants is mmol·cm-2·s-1; m, molar] 

Rate constants 

1 Log k1 

Log k2 

Log k3 

k4 2 

c 

3 Log k5 

n 

Cal cite 

i At Bi 

Calcite 

-2.063 

-4.32 

-6.63 

Equation 21 

1.45 

-6.198 

0.8962 

At 

Aragonite 

-2.063 

-4.32 

-6.46 

(2) 

2.0 

-6.436 

1.161 

Aragonite 

Bi 

Concentration 

range 

4 ca2+ 0.0025 0.09[10 + 2( 9-pH)J 0.00075[2( 9-pH)] 0.10[10 + 2(9-pH)J 0-0.3 M 

OH- 0.00014 2.0 0.00014 2.0 <pH 11.2 

HC03 0.01 1. 65 0.001 1. 65 0-0.1 M 

HC0 3 0.05 1.0 0.046 0.96 0.1-2.0 M 

co2-

3 0.000001 4.5 0.000001 17.2 (5) 

1 At a stirring rate of 260 revolutions per minute. 2 k4 is a function of P co as given theoretically by equation 21, the experimental value of k4 for aragonite is the theoretical value 

of callcite divided by 2.14. 2 

3 Formula can be used between pH 4 and 10 and Ca2+ > 0.1 molar. 4 For 0.96 atmospheres P co . 

2 5 Concentration ranges are 0 to 0.2 M and 0 to 0.05 M for calcite and aragonite, respectively. 

rate given by equation 1. To obey an adsorption isotherm, 

the backward reaction rate will have to be proportional 

to the fraction of the surface sites covered by the adsorbate. The rates of reaction in the presence and absence 

of the adsorbate are given by R and R0, respectively. 

The effect of an adsorbate on the rate can be expressed 

by the ratio (R0 - R)l R0 which is proportional to the 

fraction of adsorption sites occupied by the adsorbate. 

The fraction of adsorption sites free from the adsorbate 

is then proportional to R/ R0• 

The Langmuir isotherm equation (Langmuir, 1918) 

states that 

8 = bX/(1 + bx), (39) 

where 8 is the fraction of adsorption sites covered by the 

adsorbalte, b is a constant, and X is the concentration of 

the adsorbate in solution. Because (1 - Rl R0

) is proportional to 8, 

(1 - R/Ro) = a8 = abX/(1 + bX), (40) 

where a is a constant. Dividing by X and taking the 

reciprocal, 

(41) 

where the subscript i represents species i, A is equal to 

the constant 1/ab, and B is a second constant equal to 

1 I a. Equation 41 describes a straight line and is a common representation of a Langmuir isotherm. Inasmuch 

as the chemical reaction on the surface depends on activity 

rather than concentration, the problem involves chemisorption and X; should be replaced by the activity of the 

ith adsorbed species in equation 41. The rate of the initial 

adsorption reaction can be obtained by rearranging equation 41 as follows: 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 149 
Table 4. Procedures used to evaluate the backward reactions 

[Concentrations are in moles (mol) per liter (L); C02 partial pressures are in atmospheres] 

Ion 

i 

Solution 

preparation pH 

Equation used to 

Pco evaluate backward rate 

2 

OH KOH added. 8-11.3 Pco about O; R0H- = R0 - R 

2 

bubbled N2 • 

KHC03 added 

0-2 mol/L. 

Determined by 

composition 

and Pco 

Variable, determined RHCO - = R0 - R - RoHby composition of 3 

the gas bubbled in 

2 the solution. 

4-7.4 0.96-0.03. 

CaC1 2 added. 4-10 

co 2- 3 KHC0 3-KOH 

adaed. 

Adjusted to 9.8 

by adding KOH. 

Closed system. 

1 R = experimental rate of dissolution of calcite or aragonite. R0 = theoretical forward rate calculated from equation 1. 

Ri = backward rate of the ith species. 

(42) 

The backward rate contribution (R,) of adsorbed species 

is equal to 

(43) 

where Rb = R0 - R. Langmuir isotherm equations 

similar to equation 41 were used by Sjoberg (1978) and 

Reddy (1977) and Reddy and Wang (1980) to describe 

backward rate effects and inhibition by various ions in 

the dissolution or precipitation of calcite, respectively. 

The logical sequence that must be used to evaluate the 

backward rate constants in the system Ca2+ -H2

0-C02 

(table 4) is followed in the next sections of this paper. 

Backward reactions in KOH solutions 

At pH values greater than about 9.5, as shown in 

figure 1, the observed forward rate decreased below the 

rate predicted by equation 1. The decrease in rate can be 

explained as a backward reaction between the surface and 

adsorbed OH-. This reaction may be written: 

where the subscripts (ads) and (s) denote adsorbed and 

surface species, respectively. If the backward reaction rate 

is proportional to the activity of the adsorbate and the 

adsorption follows the Langmuir equation, a plot of 

X/(1 - R/ R0) as a function of X will yield a straight 

line with a slope equal to B and an intercept equal to 

A (eq. 41). Such a plot, strongly suggesting a direct 

150 Studies in Diagenesis 

relationship between the amount of OH- absorbed and 

the backward-reaction rate, is shown on figure 3. Also, 

a single Langmuir relation for OH- applies equally to 

both calcite and aragonite dissolution. 

The experimental results, including the observed 

and calculated rates as a function of pH and P co , 

2 

are given in tables 5 and 6 for aragonite and calcite, 

respectively. 

Backward reactions in KHC03 solutions 

A significant backward reaction between the surface of aragonite and calcite and HC03-<0> was observed 

above pH 5 at P co of 0.1 to 1.0. Figure 4 shows the 2 

extent of this backward reaction on calcite dissolution at 

0.96 atm COr The solid line represents the forward 

rates calculated from equation 38 and the constants of 

table 3, which exclude the HC03- effect. For aHco-3 

values greater than 1 Q-3, the calcite dissolution rate is 

below that calculated by equation 38 (fig. 4). Similar 

results were obtained for aragonite. The high HC03

-

activities were obtained by adding KHC03 salt. 

The observed decrease in the forward rate due to 

added KHC03 may have resulted from the following 

reactions: 

(45) 

CaHCOt 
 + HC03- (ads>- CaC03 + H2C03
' and (46) 

(47) 
~ 3.0 ,.... 

X 

-0 

~ 2.5 

,.... 

' 

~ 2.0 

> 

~ z 

0 

....J 

g 

~ >- :I: 

1.5 

1.0 

0.5 

(!)-Calcite 

x- Aragonite 

(!) 

x/~ (!) X 

yx 

(!)(!) 

X (!) 

ok_-----L------~----~------~----~------~------~----~ 

0 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 

HYDROXYL ION ACTIVITY X 1 ()3 

Figure 3. Langmuir isotherm showing backward rate effects of hydroxyl-ion activity (a0 w) for calcite and aragonite 

single crystals. 

Reaction 45 was probably the most important because 

Ca2+ is the most common cation on calcite and aragonite surfaces in aqueous solutions (Foxall and others, 

1979); reactions 46 and 47 probably occurred to a lesser 

extent. Two Langmuir isotherms are required to describe 

the backward rate effects of HC03

-; the first describes 

the backward rate for HC03

- in the activity range 0 to 

0.1, and the second, at activities greater than 0.1 (fig. 5). 

The isotherms were independent of P co as shown by 2 

the data given in tables 4 and 5 for aragonite and calcite, 

respectively. The rate constants are presented in table 3. 

Backward reactions in CaCI2 solutions 

The decrease in forward rate caused by the presence 

of Ca2+ in the bulk solution may have resulted from the 

following two reactions: 

and 

+ Ca2+ (ads) - CaC03 • (49) 

Both of these reactions probably occurred simultaneously; however, it is reasonable to expect that under differing conditions of surface pH (that is, surface P co ), 

2 

one reaction may have been significantly more important 

than the other. At high and intermediate activities of 

H+ in the stability fields of H2C03* and HC03-
 it is 

reasonable to assume that HC03

-
 is significantly higher 

than co~-(s) and that the reaction represented by equation 48 dominates the backward rate contribution due 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 151 
Table 5. Observed (Obs.) and calculated (Calc.) -log rates of dissolution of aragonite as a function of pH and C02 partial 

pressure 

[Unit of rate is mmol·cm-2·sec-1: pH was adjusted by addition of HCI, KOH, or KHC03• Omega value of the solutions was near zero] 

Cbs. Calc. Cbs. calc. Cbs. calc. 

Pco2 -log -log PC0:2 -log -log Pea -log -log Run sample pH rate rate Run sample pH rate rate Run Sample pH 2 rate rate 

3 o-2 2.00 0.0000 4.17 4.06 111 A-3 4.16 0.2843 5.85 5.85 53 8-l 4.24 0.9588 5.62 5.63 

265 A-1 2.00 0.0000 4.18 4.06 126 A-1 4.40 0.2848 5.95 5.94 59 M-2 4.25 0.9572 5.62 5.63 

9 M-1 2.23 0.0000 4.35 4.29 115 A-2 4.60 0.2843 5.94 6.00 81 G-2 4.25 0.9625 5.61 5.63 

6 M-2 2.50 0.0000 4.62 4.56 130 A-3 4.72 0.2860 6.02 6.02 48 N-1 ·4.50 0.9585 5.63 5.68 

8 N-1 2.77 0.0000 4.88 4.82 129 A-1 4.75 0.2860 6.02 6.03 58 B-1 4.75 0.9572 5.71 5.72 

1 8-2 3.00 0.0000 5.10 5.05 112 A-1 4.96 0.2843 6.15 6.06 83 8-1 4.78 0.9609 5.75 5.72 

266 8-l 3.00 0.0000 5.10 5.05 123 M-1 5.25 0.2848 6.17 6.10 54 M-1 s.oo 0.9588 5.75 5.76 

14 B-1 3.20 0.0000 5.27 5.24 116 M-2 s.so 0.2843 6.17 6.13 69 B-1 5.00 0.9585 5.74 5.76 

5 G-1 3.40 0.0000 5.46 5.42 125 M-2 5.76 0.2848 6.23 6.17 88 N-1 5.00 0.9600 5.77 5.75 

267 8-2 3.50 0.0000 5.53 5.51 131 M-2 5.92 0.2860 6.23 6.20 90 A-2 5.01 0.9600 5.73 5.76 

7 B-1 3.75 0.0000 5.78 5.72 127 A-3 6.00 0.2848 6.24 6.22 75 M-2 5.26 0.9585 5.79 5.80 

16 o-1 3.80 0.0000 5.75 5.77 114 N-1 6.03 0.2843 6.30 6.23 68 A-2 5.31 0.9592 5.85 5.81 

2 8-1 4.00 0.0000 5.91 5.92 132 N-1 6.14 0.2860 6.28 6.25 57 A-2 5.50 0.9572 5.90 5.85 

13 N-1 4.25 0.0000 6.13 6.08 119 s-1 6.26 0.2848 6.29 6.28 67 A-1 5.51 0.9592 5.88 5.85 

12 M-1 4.50 0.0000 6.14 6.21 128 A-2 6.26 0.2848 6.27 6.28 82 o-1 5.53 0.9625 5.83 5.86 

17 M-2 4.50 0.0000 6.10 6.21 135 M-1 6.35 0.2860 6.23 6.31 66 A-3 5.75 0.9592 5.92 5.91 

18 N-1 4.75 0.0000 6.32 6.30 113 M-1 6.52 0.2843 6.27 6.35 74 A-3 5.85 0.9585 5.85 5.94 

19 M-1 5.50 0.0000 6.32 6.43 117 B-1 7.00 0.2843 6.39 6.46 86 B-1 5.95 0.9609 5.95 5.97 

11 8-2 6.00 0.0000 6.39 6.45 97 M-1 3.50 0.4607 5.40 5.42 so G-2 6.02 0.9587 5.90 5.99 

20 N-1 6.50 0.0000 6.52 6.46 93 A-1 4.06 0.4799 5.67 5.73 84 A-3 6.11 0.9609 5.96 6.01 

25 8-l 7.00 0.0000 6.45 6.46 102 B-1 4.24 0.4600 5.78 5.80 65 M-1 6.25 0.9592 5.97 6.05 

10 8-1 7.03 0.0000 6.44 6.46 94 A-2 4.53 0.4791 5.86 5.87 51 s-2 6.30 0.9587 6.02 6.06 

34 N-1 7.50 0.0000 6.44 6.46 104 A-1 4.60 0.4765 5.80 5.88 80 N-1 6.35 0.9625 6.02 6.07 

35 8-l 7.50 0.0000 6.43 6.46 100 s-1 4.81 0.4791 5.98 5.92 60 A-1 6.51 0.9593 6.04 6.10 

37 N-1 8.00 0.0000 6.46 6.46 96 M-2 5.04 0.4799 5.91 5.95 56 A-1 6.74 0.9572 6.16 6.13 

32 B-1 8.52 0.0000 6.49 6.46 106 A-2 5.24 0.4765 6.03 5.98 61 A-2 6.74 0.9593 6.09 6.13 

33 M-2 8.52 0.0000 6.43 6.46 98 A-3 5.53 0.4607 6.11 6.03 72 8-2 6.88 0.9585 6.24 6.23 

21 G-1 9.00 0.0000 6.44 6.46 109 M-1 5.54 0.4574 6.14 6.04 79 M-2 6.99 0.9625 6.26 6.30 

22 M-2 9.00 0.0000 6.47 6.46 101 o-1 5.76 0.4791 6.07 6.07 78 M-1 7.07 0.9625 6.22 6.37 

444 M-1 9.40 o.oooo 6.49 6.46 108 N-1 5.82 0.4574 6.12 6.09 77 A-2 7.20 0.9625 6.37 6.47 

453 I-1-2 9.58 0.0000 6.50 6.46 99 N-1 6.00 0.4607 6.13 6.14 76 A-1 7.26 0.9625 6.48 6.52 

447 M-1 9.59 0.0000 6.49 6.46 107 A-3 6.27 0.4765 6.13 6.20 87 A-3 7.62 0.9600 6.63 6.85 

450 M-2 9.79 0.0000 6.50 6.46 95 s-2 6.50 0.4799 6.17 6.26 89 A-1 7.81 0.9600 7.02 7.06 

446 M-2 10.04 0.0000 6.59 6.46 lOS M-2 6.50 0.4765 6.16 6.26 

454 A-2 10.20 0.0000 6.65 6.46 91 M-1 2.25 0.9600 4.35 4.30 

24 B-1 10.20 0.0000 6.58 6.46 43 A-1 2.49 0.9585 4.59 4.52 

449 A-2 10.22 0.0000 6.63 6.47 92 M-2 2.75 0.9600 4.77 4.76 

443 A-1 10.46 0.0000 6.72 6.47 44 A-2 2.99 0.9585 4.99 4.97 

448 A-1 10.61 0.0000 6.70 6.47 62 N-1 3.24 0.9593 5.25 5.16 

451 A-1 10.61 0.0000 6.73 6.47 73 M-1 3.35 0.9585 5.20 5.24 

445 A-2 10.84 0.0000 6.80 6.48 46 B-1 3.50 0.9585 5.28 5.34 

124 B-1 3.24 0.2848 5.23 5.24 47 M-2 3.68 0.9585 5.42 5.43 

118 8-2 3.75 0.2848 5.55 5.63 85 s-2 3.75 0.9609 5.40 5.47 

122 N-1 3.75 0.2848 5.63 5.63 45 M-1 3.92 0.9585 5.48 5.54 

133 A-2 4.07 0.2860 5.81 5.81 52 o-1 3.93 0.9588 5.47 5.54 

to Ca <!c~s>· Similar reasoning can be used to define there- reflect these changes in surface speciation of calcite. The 

gion of low activity of H+ in the stability field of CO~- experimental data and the calculated effects of Ca2

+ 

(dashed lines, fig. 6) are in good agreement. The solid.line 

as the region where reaction 49 predominates. Our data represents the forward rate in the absence of backward reacare insufficient to evaluate the surface C02 partial pres- tions as given by equation 38 and the constants of table 

sure and to calculate the exact HC03-
 to CO~-
 ratio. 3. Similar results were obtained for aragonite, and the conAll results for calcite and aragonite are presented in table 7. stants of equation 41 describing the Ca2+ effect for both 

Figure 6 shows the effect of bulk solution pH on the calcite and aragonite are given in table 3. 

backward reaction between Ca2+ and calcite under nitrogen atmosphere. As the bulk solution pH increased from 4 Backward reactions in KHC03-KOH solutions 

to 9, the surface p co likely decreased and c~-(s) became 2 . 

came The stability field of c~- ion begins at pH of 10.3 increasingly more important in the surface speciation. The changes in backward rate shown on figure 5 and is beyond the pH range of most natural environments. 

152 Studies in Diagenesis 
~ -4.0 

0 

(.) 

w 

(/) 

a: 

w -4.5 

a.. 

a: 

w 

1-

w 

~ 

1- -5 . .0 

z 

w 

(.) 

w 

a: 

~ -5.5 

0 

(/) 

a: 

w 

a.. 

(/) -6.0 

w 

_J 

0 

~ 

~ _J 

~ -6.5 

z 

w 

~ a: -7.0 lL 

0 

~ 

::r: 

1-

a: <( -7 5 (!) • 2 

g 

-6.0 

3 4 

-5.0 -4.0 

5 

pH 

-3.0 -2.0 -1.0 

LOGARITHM OF BICARBONATE ION ACTIVITY 

(!) 

(!) 

(!) 

(!) 

8 

0.0 

Figure 4. Dissolution rates of calcite single crystals in absence (solid line) and presence of different bicarbonate ion activities 

(aHco -) and 0.96 atmosphere C02 partial pressure far from equilibrium. For experimental conditions, pH.= 7.84 + 3 

log aHCO -. 

3 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 153 
Table 6. Observed (Obs.) and calculated (Calc.) -log rates of dissolution of calcite as a function of pH and C02 partial pressure 

[Unit of rate is mmol·cm-2·sec-1; pH was adjusted by addition of HCl, KOH, or KHC03• Omega value of the solutions was about zero] 

OJs. Calc. (l)s. Calc. (l)s. Calc. 

Pco -log -log Pco2 -log -log Pco -log -log Run Sample pH 2 rate rate Run Sample pH rate rate Run Sample pH 2 rate rate 

369 16 2.33 0.0000 4.44 4.39 236 10 5.00 0.0000 6.60 6.49 474 15 4.53 0.0991 6.15 6.21 

3 1 2.45 0.0000 4.50 4.51 286 13 5.10 0.0000 6.61 6.52 488 13 4.72 0.0979 6.28 6.28 

25 2 2.50 0.0000 4.54 4.56 275 11 5.20 0.0000 6.68 6.54 473 14 4.78 0.0991 6.27 6.30 

30 2 2.59 0.0000 4.59 4.65 287 14 5.30 0.0000 6.70 6.56 489 14 4.87 0.0979 6.33 6.34 

368 17 2.65 0.0000 4.77 4.71 308 19 5.35 0.0000 6.67 6.56 472 13 4.98 0.0991 6.37 6.37 

4 2 2.65 0.0000 4.61 4. 71 293 9 5.40 0.0000 6.68 6.57 490 15 4.99 0.0979 6.35 6.37 

26 1 2.70 0.0000 4.82 4.76 290 22 5.50 0.0000 6.73 6.58 215 9 5.00 0.0989 6.46 6.37 

1 1 2.80 0.0000 4.85 4.86 28 2 5.50 0.0000 6.60 6.58 216 10 5.00 0.0989 6.48 6.37 

31 1 2.84 0.0000 5.00 4.89 288 15 5.55 0.0000 6.50 6.59 217 11 5.00 0.0989 6.41 6.37 

32 2 2.84 0.0000 4.98 4.89 302 13 5.60 0.0000 6.62 6.59 218 12 5.00 0.0989 6.38 6.37 

363 11 2.90 0.0000 4.92 4.95 289 16 5. 75 0.0000 6.68 6.60 491 16 5.20 0.0979 6.37 6.43 

24 1 3.00 0.0000 5.05 5.05 303 14 6.00 0.0000 6.57 6.61 471 12 5.28 0.0991 6.33 6.45 

2 2 3.08 0.0000 5.08 5.13 291 23 6.20 o.oooo 6.75 6.62 492 17 5.35 0.0979 6.50 6.47 

358 22 3.10 0.0000 5.18 5.15 301 12 6.40 0.0000 6.63 6.62 493 18 5.51 0.0979 6.55 6.51 

367 15 3.15 0.0000 5.25 5.20 277 15 6.70 0.0000 6.58 6.63 470 11 5.57 0.0991 6.47 6.53 

362 10 3.25 0.0000 5.25 5.29 304 15 7.00 0.0000 6.62 6.63 494 19 5.62 0.0979 6.50 6.54 

11 1 3.25 0.0000 5.24 5.29 292 24 7.20 0.0000 6.66 6.63 469 10 5.78 0.0991 6.52 6.58 

357 21 3.30 0.0000 5.39 5.34 280 9 7.50 0.0000 6.68 6.63 495 20 5.78 0.0979 6.57 6.58 

14 2 3.40 0.0000 5.47 5.43 305 16 7.75 0.0000 6.56 6.63 231 9 6.00 0.0985 6.60 6.64 

15 1 3.40 0.0000 5.43 5.43 628 13 7.87 0.0000 6.65 6.63 232 10 6.00 0.0985 6.61 6.64 

89 1 3.50 0.0000 5.55 5.53 629 14 7.96 0.0000 6.60 6.63 233 11 6.00 0.0985 6.57 6.64 

90 2 3.50 0.0000 5.53 5.53 19 2 8.00 0.0000 6.60 6.63 234 12 6.00 0.0985 6.56 6.64 

95 7 3.50 0.0000 5.57 5.53 276 12 8.20 0.0000 6.57 6.63 431 11 6.02 0.0983 6.61 6.64 

111 7 3.50 0.0000 5.52 5.53 306 17 8.40 0.0000 6.66 6.64 266 10 6.07 0.0988 6.61 6.66 

112 8 3.50 0.0000 5.50 5.53 630 15 8.59 0.0000 6.67 6.64 468 9 6.27 0.0991 6.59 6.70 

114 2 3.50 0.0000 5.54 5.53 307 18 8.60 0.0000 6.73 6.64 432 12 6.54 0.0983 6.53 6.74 

119 7 3.50 0.0000 5.51 5.53 294 17 9.00 0.0000 6.69 6.66 187 9 6.66 0.0989 6.65 6.75 

120 8 3.50 0.0000 5.50 5.53 631 16 9.00 0.0000 6.64 6.66 188 10 6.66 0.0989 6.66 6.75 

5 1 3.51 0.0000 5.59 5.54 274 10 9.38 0.0000 6.63 6.69 189 11 6.66 0.0989 6.64 6.75 

366 14 3.60 0.0000 5.58 5.62 624 9 9.41 0.0000 6.68 6.69 190 12 6.66 0.0989 6.62 6.75 

361 9 3.75 0.0000 5.64 5.75 295 18 9.45 0.0000 6.78 6.70 191 9 6.66 0.0987 6.60 6.75 

22 1 3.80 o.oooo 5.75 5.79 625 10 9.60 0.0000 6.69 6.72 192 10 6.66 0.0987 6.61 6.75 

83 2 3.83 0.0000 5.80 5.82 627 12 9.61 0.0000 6.69 6.72 193 11 6.66 0.0987 6.61 6.75 

356 20 3.85 0.0000 5.72 5.84 626 11 9.67 0.0000 6.67 6.73 194 12 6.66 0.0987 6.53 6.75 

365 13 3.97 o.oooo 5.80 5.94 278 13 9.85 0.0000 6.75 6.76 438 18 6.84 0.0983 6.64 6.77 

12 2 4.00 0.0000 5.91 5.96 296 19 10.01 0.0000 6.83 6.78 433 13 7.05 0.0983 6.56 6.79 

360 24 4.05 0.0000 5.84 6.00 279 14 10.33 0.0000 6.81 6.84 268 12 7.21 0.0988 6.60 6.79 

355 19 4.15 0.0000 6.13 6.07 297 20 10.55 0.0000 6.79 6.86 439 19 7.22 0.0983 6.64 6.79 

21 2 4.19 0.0000 6.18 6.10 298 21 10.80 0.0000 6.97 6.89 440 20 7.50 0.0983 6.66 6.80 

364 12 4.27 0.0000 6.09 6.16 281 16 11.03 0.0000 6.88 6.91 455 19 7.68 0.0985 6.91 6.81 

283 10 4.30 0.0000 6.16 6.18 300 11 11.04 0.0000 6.94 6.91 270 10 7.80 0.0986 7.15 6.81 

18 2 4.30 0.0000 6.15 6.18 299 10 11.18 0.0000 6.85 6.91 448 12 7.81 0.0985 6.93 6.81 

359 23 4.40 0.0000 6.17 6.24 487 12 3.16 0.0981 5.04 5.20 456 20 7.92 0.0985 7.14 7.01 

354 18 4.50 0.0000 6.22 6.29 479 20 3.27 0.0990 5.10 5.30 449 13 8.04 0.0985 7.25 7.14 

273 9 4.50 0.0000 6.34 6.29 486 11 3.33 0.0981 5.27 5.35 272 12 8.12 0.0986 7.48 7.22 

17 1 4.50 0.0000 6.28 6.29 485 10 3.50 0.0981 5.46 5.50 457 21 8.16 0.0985 7.35 7.26 

282 9 4.55 0.0000 6.42 6.32 478 19 3.52 0.0990 5.36 5.52 450 14 8.28 0.0985 7.60 7.42 

23 2 4.65 0.0000 6.48 6.37 484 9 3.68 0.0981 5.65 5.66 265 9 5.60 0.2964 6.36 6.39 

284 11 4.73 0.0000 6.54 6.40 477 18 3.77 0.0990 5.52 5. 73 267 11 6.90 0.2962 6.42 6.54 

36 2 4.75 0.0000 6.53 6.41 483 24 3.82 0.0981 5.60 5. 78 269 9 7.30 0.2959 6.62 6.55 

285 12 4.95 0.0000 6.58 6.48 476 17 4.00 0.0990 5.81 5.91 271 11 7.63 0.2959 6.99 6.93 

235 9 5.00 0.0000 6.64 6.49 482 23 4.13 0.0981 5.80 5.99 403 16 . 4.50 0.4711 5.92 5.97 

236 10 5.00 0.0000 6.60 6.49 475 16 4.23 0.0991 5.96 6.06 405 18 5.14 0.4711 6.12 6.18 

241 10 5.00 0.0000 6.62 6.49 481 22 4.37 0.0981 6.01 6.13 400 20. 5.83 0.4693 6.26 6.33 

235 9 5.00 0.0000 6.64 6.49 480 21 4.47 0.0981 6.13 6.18 374 22 6.53 0.4725 6.37 6.39 

Possible reactions between calcite and aragonite surface CaHC03t> + co~-(ads) -+ CaC03 + HC03-
" (52) 

sites and co~-(ads) are: 

Of these, reaction 50 is probably the most important. The 

Ca2+(s) + co~-(ads) -+ CaC03, (50) contribution of the co~-(ads) to the backward reaction 

was evaluated at pH 9.8 in solutions ranging from lQ-4 

to 0.3 M total carbon. The rates for calcite as a function 

CaOH+
 + co~-(ads) -+ CaC03 + OH-
' and (51) of the total dissolved carbon in solution is shown in 

154 Studies in Diagenesis 
Table 6. Observed (Obs.) and calculated (Calc.) -log rates of dissolution of calcite as a function of pH and C02 partial 

pressure--Continued 

O:>s. Calc. O:>s. Calc. O:>s. Calc. 

Pco2 -log -log p~ -log -log Pco2 -log -log 

Run Sample pH rate rate Run Sample pH rate rate Run Sample pH rate rate 

383 15 6.80 0.4737 6.30 6.39 322 22 3.19 0.9955 5.17 5.14 339 19 6.01 0.9799 6.12 6.12 

388 20 6.81 0.4700 6.37 6.40 53 2 3.30 0.9791 5.25 5.22 417 13 6.04 0.9892 6.06 6.11 

410 10 6.98 0.4684 6.32 6.40 323 16 3.40 1.0025 5.27 5.29 340 20 6.10 0.9799 6.07 6.12 

375 23 7.05 0.4725 6.48 6.40 324 20 3.50 0.9959 5.32 5.36 261 9 6.11 0.9885 6.08 6.12 

411 11 7.06 0.4684 6.31 6.40 50 2 3.50 0.9772 5.40 5.36 341 21 6.21 0.9799 6.08 6.12 

412 12 7.17 0.4684 6.47 6.55 325 15 3.60 0.9959 5.43 5.42 342 22 6.31 0.9799 6.11 6.13 

381 13 7.30 0.4737 6.60 6.65 245 9 3.70 0.9859 5.66 5.48 424 20 6.35 0.9876 6.14 6.13 

389 21 7.40 0.4700 6.72 6.75 326 21 3.70 0.9955 5.48 5.48 262 10 6.39 0.9885 6.14 6.13 

390 22 7.44 0.4700 6.80 6. 79 247 11 3.70 0.9857 5.36 5.48 343 23 6.41 0.9799 6.15 6.13 

376 24 7.52 0.4725 7.14 6.88 246 10 3.70 0.9859 5.40 5.48 465 17 6.48 0.9849 7.26 6.13 

382 14 7.75 0.4737 7.30 7.18 248 12 3.70 0.9857 5.41 5.48 460 24 6.49 0.9867 6.13 6.13 

399 15 2.47 0.4884 4.55 4.52 39 2 3.75 0.9965 5.57 5.50 459 23 6.49 0.9867 6.26 6.13 

395 11 2.70 0.4885 4.68 4.74 327 19 3.80 0.9959 5.53 5.53 466 18 6.49 0.9849 8.12 6.13 

398 14 2.82 0.4884 4.87 4.85 309 20 3.85 0.9892 5.53 5.55 461 9 6.49 0.9867 6.03 6.13 

394 10 3.03 0.4885 5.02 5.05 328 14 3.90 0.9959 5.56 5.58 344 24 6.53 0.9879 6.21 6.13 

393 9 3.49 0.4885 5.39 5.42 43 2 4.00 0.9950 5.63 5.62 418 14 6.53 0.9892 6.14 6.13 

397 13 3.91 0.4884 5.62 5.71 310 21 4.10 0.9892 5.60 5.66 345 9 6.64 0.9879 6.13 6.13 

396 12 4.31 0.4885 5.78 5.89 66 2 4.10 0.9847 5.60 5.66 263 11 6.66 0.9884 6.23 6.13 

404 17 4.63 0.4907 5.95 6.01 65 4 4.13 0.9847 5.64 5.68 419 15 6.74 0.9892 6.25 6.13 

392 24 4.89 0.4885 6.02 6.10 311 22 4.25 0.9890 5.66 5.72 346 10 6.74 0.9879 6.19 6.13 

427 23 5.25 0.4917 6.16 6.20 312 23 4.48 0.9890 5.72 5.80 347 11 6.84 0.9879 6.28 6.28 

391 23 5.38 0.4884 6.15 6.23 44 1 4.50 0.9950 5.79 5.81 264 12 6.84 0.9884 6.22 6.28 

428 24 5.54 0.4917 6.21 6.27 313 24 4.51 0.9890 5.70 5.81 377 9 6.89 0.9868 6.31 6.32 

401 21 5.67 0.4889 6.25 6.29 253 9 4.52 0.9834 5.79 5.82 348 12 6.94 0.9879 6.27 6.36 

429 9 5.73 0.4917 6.23 6.30 315 10 4.58 0.9892 5.79 5.83 378 10 6.94 0.9868 6.28 6.36 

430 10 5.85 0.4917 6.25 6.32 52 1 4.70 0.9791 5.86 5.87 425 21 6.98 0.9876 6.37 6.40 

413 9 6.04 0.4946 6.26 6.34 314 9 4.73 0.9890 5.81 5.88 384 16 7.00 0.9792 6.35 6.41 

402 19 6.09 0.4907 6.30 6.35 316 11 4.84 0.9892 5.84 5.91 370 18 7.00 0.9843 6.47 6.41 

414 10 6.28 0.4946 6.25 6.36 255 11 4.84 0.9834 5.86 5.92 349 13 7.03 0.9868 6.39 6.44 

406 22 6.50 0.4879 6.37 6.37 317 12 4.93 0.9896 5.82 5.94 379 11 7.10 0.9868 6.52 6.51 

415 11 6.55 0.4946 6.27 6.37 329 9 4.99 0.9818 5.90 5.96 350 14 7.13 0.9868 6.44 6.54 

407 23 6.64 0.4879 6.37 6.38 454 18 5.00 0.9853 5.83 5.96 371 19 7.16 0.9843 6.57 6.57 

408 24 6.75 0.4879 6.39 6.38 452 16 5.00 0.9853 5.86 5.96 380 12 7.20 0.9868 6.57 6.61 

416 12 6.76 0.4946 6.23 6.38 453 17 5.00 0.9853 5.81 5.96 351 15 7.22 0.9868 6.64 6.63 

409 9 6.86 0.4879 6.31 6.38 464 16 5.00 0.9849 6.15 5.96 426 22 7.23 0.9876 6.76 6.64 

420 16 6.95 0.4938 6.39 6.38 462 10 5.00 0.9849 5.92 5.96 385 17 7.25 0.9792 6.58 6.67 

434 14 7.07 0.4913 6.46 6.39 463 11 5.00 0.9849 6.03 5.96 372 20 7.26 0.9843 6.83 6.69 

421 17 7.15 0.4938 6.44 6.53 45 2 5.00 0.9950 5.92 5.96 352 16 7.32 0.9868 6.77 6.75 

435 15 7.16 0.4913 6.54 6.54 67 3 5.00 0.9724 5t92 5.96 386 18 7.35 0.9792 6.91 6.79 

436 16 7.29 0.4913 6.61 6.65 256 12 5.10 0.9834 5.89 5.99 373 21 7.35 0.9843 7.07 6.79 

422 18 7.38 0.4938 6.88 6.74 318 13 5.13 0.9896 5.91 5.99 353 17 7.42 0.9868 6.83 6.89 

437 17 7.40 0.4913 6.76 6.76 331 11 5.14 0.9818 5.92 6.00 387 19 7.44 0.9792 6.99 6.93 

441 21 7.48 0.4874 7.06 6.84 330 10 5.22 0.9818 5.95 6.01 445 9 7.55 0.9749 7.12 7.11 

423 19 7.58 0.4938 7.16 6.95 258 10 5.26 0.9885 5.98 6.02 446 10 7.63 0.9749 7.17 7.29 

442 22 7.69 0.4874 7.16 7.10 332 12 5.36 0.9818 5.94 6.04 447 11 7.72 0.9749 7.23 7.55 

443 23 7.79 0.4874 6.93 7.25 257 9 5.45 0.9885 6.01 6.05 

444 24 7.87 0.4874 7.46 7.41 333 13 5.47 0.9818 6.00 6.06 

319 18 2.50 0.9959 4.54 4.54 336 16 5.50 0.9801 6.00 6.06 

41 1 2.54 0.9965 4.62 4.57 49 1 5.50 0.9763 6.01 6.07 

55 2 2.69 0.9801 4.69 4.71 259 11 5.60 0.9885 6.03 6.08 

320 23 2.74 0.9955 4.83 4.76 335 15 5.65 0.9801 6.02 6.08 

40 1 2.74 0.9965 4.86 4.76 334 14 5.74 0.9801 6.02 6.09 

54 1 2.90 0.9801 5.05 4.90 260 12 5.80 0.9885 6.00 6.10 

321 17 3.00 0.9959 4.99 4.98 337 17 5.82 0.9801 6.07 6.10 

42 2 3.15 0.9965 5.14 5.11 338 18 5.95 0.9801 6.10 6.11 

figure 7. Line 1 in figure 7 represents the forward rate are given in table 8 as a function of the total molarity 

(RJ for the dissolution of calcite in the absence of back- of co~-. 

ward reactions. The calculated rates as R0 H-, RHco -, 

3 

and Rc0 z- effects are subtracted successively from R0 3 Backward reactions in Ca(HC03) 2 solutions 

and represented by lines 2, 3, and 4, respectively, as a 

function of log activity CO~- The calculated rate (line 4) No single model describes the rates of dissolution 

agrees well with the experimental points for calcite. The of calcite and aragonite in Ca(HC03

) 2 solutions in the 

experimental and calculated rates for calcite and aragonite saturation region of 0.0 to 1.0 0. The Plummer, Wigley, 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 155 
-0 

cr: 

' a: 

or-

' ~ 

> 

..... 

~ 

z 

0 

w 

~ z 

0 

co 

a: 

<( 

(.) 

co 

0.9 

0.8 

0.7 

0.6 

0.5 

0.4 

0.3 

0.2 

0.1 

• - CALCITE 1 0% C02 

(!)- CALCITE 30% C02 

• - CALCITE 50% C02 

m - CALCITE 1 00% C02 

• - ARAGONITE 1 00% C02 

o~~~~~~--~~--~~--~~--~~--~~--~~--~~--~-

0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 

BICARBONATE ION ACTIVITY 

Figure 5. Langmuir isotherm showing backward rate effects of bicarbonate ion activities (aHco -)for calcite and aragonite 3 

single crystals at different gas compositions (numerical values represent mole percent C02 in the C02-N2 gas mixtures) 

far from equilibrium. 

and Parkhurst model describes accurately the rate of dissolution of calcite and aragonite in the saturation region 

of 0.005 to 0.90 0. The model fails to predict accurately 

rates very far from equilibrium and particularly very near 

equilibrium. The Plummer, Wigley, and Parkhurst model 

can predict reasonably accurate crystal growth rates above 

0 of 1.5 but not near equilibrium (Reddy and others, 

1981). The hybrid model incorporating many features of 

the model of Plummer, Wigley, and Parkhurst used in 

the present paper can describe accurately rates in the 

saturation region of 0.0-0.6 n for both calcite and 

156 Studies in Diagenesis 

aragonite. This model fails in the saturation region of 

0.6-1.0 0. An additional deficiency of the models its 

failure to predict rates of crystal growth from supersaturated solution. An asset of the model is its ability to 

predict rate far from equilibrium in relatively complex 

(mixed salt) solutions. 

The dissolution region 

Figure 8 shows the rate of dissolution and precipitation of single crystals of calcite as a function of the 
Table 7. Observed (Obs.) and calculated (Calc.) log rates of dissolution of calcite and aragonite 

in CaCI 2 solutions as a function of pH 

[Concentrations are in mmol·L-1; rate is in mmol·cm-2·sec-1

] 

Calcite Aragonite 

Cbs. Calc. Cbs. Calc. 

log log log log 

Run Sample CaC12 pH rate rate Run Sample CaC12 pH rate rate 

15 24 0.0005 8.60 -6.77 -6.78 39 M-6 0.0010 6.00 -6.39 -6.55 

16 23 0.0005 8.60 -6.81 -6.78 73 M-3 0.0010 6.80 -6.62 -6.57 

5 19 0.0010 6.00 -6.76 -6.66 74 B-9 0.0010 6.80 -6.60 -6.57 

6 20 0.0010 6.00 -6.62 -6.66 41 A-1 0.0010 7.00 -6.58 -6.57 

1 9 0.0010 7.00 -6.74 -6.72 76 B-8 0.0010 7.00 -6.59 -6.57 

2 10 0.0010 7.00 -6.77 -6.72 28 B-4 0.0010 7.37 -6.61 -6.57 

11 9 0.0010 7.40 -6.74 -6.75 77 M-5 0.0010 8.00 -6.58 -6.58 

12 10 0.0010 7.50 -6.75 -6.76 78 B-7 0.0010 8.00 -6.66 -6.58 

17 22 0.0010 8.85 -6.93 -6.90 33 M-3 0.0010 8.00 -6.73 -6.58 

18 21 0.0010 8.85 -6.82 -6.90 34 B-7 0.0010 8.00 -6.66 -6.58 

34 9 0.0100 5 .. 00 -6.67 -6.59 79 M-6 0.0010 9.00 -6.78 -6.58 

35 10 0.0100 5.00 -6.68 -6.59 25 N-1 0.0010 9.00 -6.58 -6.58 

7 21 0.0100 6.00 -6.84 -6.80 26 B-3 0.0010 9.00 -6.64 -6.58 

8 22 0.0100 6.00 -6.90 -6.80 62 B-1 0.0100 6.00 -6.65 -6.73 

31 22 0.0100 7.00 -7.00 -6.94 65 A-2 0.0100 6.00 -6.63 -6.73 

3 11 0.0100 7.00 -7.03 -6.94 45 N-1 0.0100 6.85 -6.82 -6.82 

30 21 0.0100 7.00 -6.92 -6.94 46 B-8 0.0100 6.85 -6.84 -6.82 

13 11 0.0100 7.50 -7.00 -7.02 44 B-9 0.0100 7.30 -6.86 -6.86 

14 12 0.0100 7.50 -7.15 -7.02 50 B-6 0.0100 8.00 -6.88 -6.92 

22 15 0.0100 8.00 -7.08 -7.12 35 M-4 0.0100 8.00 -6.70 -6.92 

25 19 0.0100 8.00 -7.15 -7.12 36 B-8 0.0100 8.00 -6.75 -6.92 

23 16 0.0100 8.00 -7.01 -7.12 47 N-2 0.0100 8.00 -6.89 -6.92 

24 18 0.0100 8.00 -7.07 -7.12 23 A-2 0.0100 9.00 -6.95 -6.97 

26 20 0.0100 8.00 -7.14 -7.12 59 M-1 0.1000 5.00 -6.67 -6.59 

20 18 0.0100 8.82 -7.11 -7.30 60 B-1 0.1000 5.00 -6.58 -6.59 

21 17 0.0100 8.82 -7.17 -7.30 67 M-2 0.1000 5.10 -6.61 -6.63 

40 15 0.1000 5.00 -6.73 -6.67 72 B-2 0.1000 6.10 -6.86 -6.85 

9 23 0.1000 6.00 -7.09 -6.91 31 M-2 0.1000 7.00 -6.98 -7.05 

10 24 0.1000 6.00 -7.06 -6.91 32 B-6 0.1000 7.00 -6.88 -7.05 

37 18 0.1000 7.00 -7.27 -7.10 37 M-5 0.1000 8.00 -7.21 -7.31 

4 13 0.1000 7.00 -7.23 -7.10 38 B-9 0.1000 8.00 -7.18 -7.31 

28 11 0.1000 7.00 -7.20 -7.10 21 A-1 0.1000 9.00 -7.33 -7.62 

29 12 0.1000 7.00 -7.27 -7.10 22 B-1 0.1000 9.00 -7.45 -7.62 

33 23 0.1000 7.50 -7.35 -7.21 

38 14 0.1000 8.00 -7.45 -7.33 

32 15 0.1000 8.50 -7.71 -7.45 

19 19 0.1000 8.86 -7.72 -7.55 

39 15 0.1000 9.00 -7.85 -7.58 

27 9 0.1000 9.00 -7.85 -7.58 

36 21 0.3000 5.00 -6.81 -6.69 

activity product of Ca2+ and HC03

- at P co = 0.96 atm. contributions of Ca2+, HC03

-, COi-, and OH- (eq. 43). 2 

The The contributions of OH- and COi- at the solution con- effect of pH on rate was eliminated by subtracting 

k1aH+ from the observed rate. Figure 8 is divided into ditions are negligible. The equation describing rate is then 

dissolution (undersaturated part) and crystal growth 

(oversaturated part) segments. The dissolution segment R = R, {I -EA. :' B a ). (53) 

can be subdivided into two dissolution regions. I I I 

In dissolution region I, the observed rates are equal 

to the forward rates (eq. 1) minus the backward rate where R0 is the forward rate in the absence of all 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 157 
c -4 z 

0 

0 

w 

(f) 

a: 

w 

a. 

a: 

w 

tu -5 ~ 

tz 

w 

0 

w 

a: 

<( 

::> 

0 

(f) -6 

a: 

w 

a. 

(f) 

w 

...J 

0 

~ 

:::J 

...J 

~ -7 

~ .. w 

~ a: 

u. 

0 

~ 

::c -8 to: 

<( 

(!) 

g 

2 3 4 5 

Calcite 

CaCI2 

x - 0.0005 M (molar) 

(!)- 0.001 M 

6 -0.01 M 

rn- 0.1 M 

+- 0.3 M 

6 7 

pH 

(!) (!) 

8 9 

Figure 6. Effect of different CaCI2 concentrations on rate of dissolution of calcite single crystals far from equilibrium. 

Solid line is calculated rate in absence of CaCI 2 (eq. 38). Dashed lines are rates calculated from equation 42 and constants of table 3. 

1 58 Studies in Diagenesis 
a: 

w 

tu 1 ~ 

1--

z 

w 

0 

.................•..................................................................................................... 

-6.5 

w 2 a: 

<C 

::> 

0 

C/) 

---------------------------------------

a:c 

~ ~ -7.0 

C/)(J 

ww 

...JC/) 

Oa: 

~w 

:::Ja.. 

d 

~ 

z 

u.i 

~ a: 

u. 

0 

~ 

:I: 

1--

a: 

<C 

-7.5 

-8.0 

[!) 

[!) 

Calcite 

3 

4 

§ -Ei 4 ~ ~ -1 

LOGARITHM OF CARBONATE ION ACTIVITY 

Figure 7. Dissolution rates of single crystals of calcite and aragonite as function of activity of carbonate ion (aco 2-). Lines 3 

1, 2, 3, and 4 represent log R0, log (R0 - R0 w), log (R0 - R0w- RHco -), and log (R0 - R0 H-- RHco-- Rc0 2-), respectively. 3 3 3 

backward reactions (eq. 1), A and Bare Langmuir isotherm 

constants (table 3), and a; represents the activity of the ith 

ions in the solutio:1. The experimental data, saturation, and 

experimental and calculated rates are presented in tables 

9 and 10 for aragonite and calcite, respectively. 

In dissolution region II (1 > 0 > 0.6) the observed 

rates in Ca(HC03) 2 solution follow the empirical 

equation: 

(54) 

where k5 and n are constants and 0 is the ratio IAPIK, 

where lAP is the ion activity product of CaC03 in solution and K is the equilibrium constant. Equation 54 was 

first used by Morse (1978) and Morse and others (1979) 

to describe the dissolution of calcite and aragonite near 

equilibrium in sea water. Sundquist (1979) proposed that 

CO~- can be substituted for 0 in equation 54. This form 

of the equation does not fit our data as well and is not 

used in this paper. 

Thking the logarithm of equation 54, 

log R = log k 5 + n log (1 - 0), (55) 

which is the equation of a straight line. Figure 9 is a plot 

of aragonite and calcite data near equilibrium; the last 

point in this illustration has an 0 value of 0.992. 

Equation 54 resembles a Freundlich isotherm 

(Adamson, 1976). The relationship (1 - 0) could con~ 

ceivably describe the decrease in the number of sites with 

high enough free energy to initiate backward reactions 

near equilibrium. This equation is valid for the system 

Ca2+ ~H 0-C02 from 0 of 0.6 to at least 0.99 for both 

calcite and aragonite. The rate constants needed to 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 159 
TableS. Observed (Obs.) and calculated (Calc.) log 1.161 in calcium bicarbonate solutions (table 3) near 

rates of dissolution of calcite (C) and aragonite (A) in equilibrium. 

KHC03-KOH solutions as a function of total carbon 

in solution 

The precipitation region 

[Rates were calculated using equation 51 and constants of 

table 3. Units of concentration and rate are mmol·L -I and The kinetics of precipitation of calcite and aragonite mmol·cm-2·sec-1, respectively] on single crystals at various P co was investigated. The 2 

supersaturated solutions were prepared as described by 

Cl>s. Calc. Reddy and other ( 1981) with the exception of dissolving 

Total log log vaterite rather than calcite to obtain the supersaturated Run Sample carbon pH rate rate Phase solutions. Single crystals of aragonite and calcite were 

22 22 0.0001 9.80 -6.89 -6.83 c suspended in the stirred metastable solutions and the rate 

23 24 0.0001 9.80 -6.81 -6.83 c of crystallizations determined from the weight gained. 18 17 0.0003 9.80 -6.88 -6.90 c The phase forming on aragonite was carefully checked 20 19 0.0003 9.80 -6.94 -6.90 c 

21 by staining (Friedman, 1959), optical microscopy, and 20 0.0003 9.80 -6.85 -6.90 c 

X-ray diffraction to make sure it was aragonite. No calcite 

16 13 0.0006 9.80 -7.13 -6.98 c as detected forming on aragonite single crystals. The 17 14 0.0006 9.80 -7.07 -6.98 c results are shown in figures 8 and 10. 11 23 0.0010 9.80 -7.26 -7.05 c 

12 24 0.0010 9.80 -7.17 -7.05 c Two precipitation regions were recognized and are 

10 22 0.0030 9.80 -7.41 -7.25 c shown in figure 8. The rate of growth of calcite and 

14 11 0.0030 9.80 -7.33 -7.25 c aragonite in precipitation region I can presumably be 

15 12 0.0030 9.80 -7.47 -7.25 c described by equation 54; however, we were unable to 

9 19 0.0060 9.80 -7.45 -7.37 c evaluate these constants because of the extremely slow 

13 10 0.0060 9.80 -7.49 -7.37 c rates of precipitation and the metastable nature of the 

7 17 0.0100 9.80 -7.45 -7.45 c supersaturated solutions. 

8 18 0.0100 9.80 -7.38 -7.45 c In precipitation region II, the rate of CaC03 

6 16 0.0300 9.80 -7.51 -7.55 c growth is proportional to the activity product of Ca2+ 

4 12 0.0600 9.80 -7.46 -7.58 c and HC03

-. The rate expression that describes the 5 14 0.0600 9.80 -7.59 -7.58 c 

3 11 0.1000 9.80 -7.54 -7.60 c precipitation of both calcite and aragonite is: 

1 9 0.3000 9.80 -7.54 -8.30 c R = kicE - Oc8 2+ aHCO -), (56) 2 10 0.3000 9.80 -7.61 -8.30 c 3 

1 M-6 0.0001 9.80 -6.62 -6.73 A where k 4 2 A-1 and care constants, and E is the equilibrium 0.0003 9.80 -6.84 -6.74 A 

3 A-2 0.0006 9.80 -6.86 -6.76 A product of Ca2+ and HC03

- for the appropriate CaC03 

4 M-1 0.0010 9.80 -6.99 -6.78 A polymorph at the bulk solution C02 partial pressure. 

5 B-4 0.0010 9.80 -6.97 -6.78 A Equation 56 was found to be valid for C02 partial 

10 N-2 0.0030 9.80 -7.11 -6.87 A pressures from 10-3·

5 to 0.96 atm and for calcite in the 

11 B-7 0.0030 9.80 -7.04 -6.87 A 0 range of 1.5 to more than 100 and aragonite in the 0 

8 N-1 0.0060 9.80 -7.02 -6.98 A range of 2.0 to more than 100. In terms of 0, an alter9 B-6 0.0060 9.80 -7.04 -6.98 A native form of equation 56 is: 

6 M-2 0.0100 9.80 -7.31 -7.09 A 

7 B-5 0.0100 9.80 -7.13 -7.09 A 

13 M-5 k4 K ( OH+ ) 0.0300 9.80 -7.21 -7.47 A R = -- OH+(eq)c-0--, (57) 14 M-7 0.0300 9.80 -7.09 -7.47 A K2 °H+(eq) 

12 M-3 0.1000 9.80 -7.21 -8.60 A 

where K is the equilibrium constant for calcite or 

aragonite, the subscript (eq) refers to the equilibrium 

evaluate equation 54 at 0.96 atm C02 partial pressure value, and K2 is given by equation 26. The experimenare given on table 3. The effects of P co (if any) on the tal results for both calcite and aragonite are shown in 2 figure 10. rate constants have not been evaluated. 

A problem with using an empirical equation such 

as equation 54 is that the values of k 5 and n vary with The backward rate constant k4 the composition of the aqueous solution. For example, 

in seawater Morse and other (1979) found a value of n Plummer and others(1978) experimentally deterfor aragonite of 2.93 which compares with our values of mined the rate constant k 4 of calcite for a large number 

160 Studies in Diagenesis 
X 2.0 ,.......----.---"""T"r-'"-.---r------.--.-----,----.-------,----.,r---.-------,-----, 

c 

z 

0 

(.) 

w 

en 

a: 

w 

Q.. 

a: 

w 

~ ~ 

1.0 

Oil :;solution 

reqion I j_l Dissolution 

region li 

I 

I 

I 

j.-- Equilibrium line 

1-

z 

w 

(.) 

0 ----- (!D~ 

I 

w 

a: 

< 

::::> 

0 

~ -1.0 

w 

Q.. 

I 

I 

Precipitation~ 

region I I 

I 

I 

Calcite 25°C 

0.96 Pc~ 

~ I 

I 

Precipitation 

region li. 

~ :::i -2.0 

...J 

I 

I 

I ~ I 

~ l 

I + I ,f 

~~.OL---~--~--~~--~--~--~----~--~--~----~--~--~ 

~ 0 5 10 15 20 25 30 

ACTIVITY PRODUCT OF CALCIUM AND BICARBONATE IONS X 1Q6 

Figure 8. Obse ·ved rate of dissolution and growth of calcite single crystals minus k1aH+ as function of activity product 

of Ca2+ and HC03- at 0.96 P co . See text for details. 2 

of C02 partial pn~ssures. They also evaluated this constant theoretically using equation 21 and found close agree-

·ment with the experiemental results. The theoretical value 

of k4 for calcite as a function of C02 partial pressure can 

be calculated from equation 21 and is given in Plummer 

and others (1978). The value for k4 calculated for 

aragonite from figure 10 is about half the theoretical value 

of calcite and about three times smaller than the theoretical 

value for aragonite calulated from equation 21 at 0.96 atm 

C02 partial pressure. These values suggest that kinetic 

factors favor the growth of calcite at the expense of 

aragonite in super:iaturated Ca(HC03

)

2 solutions. 

Effects of sti rri 111g on rates 

The effects of stirring on the rate constant k 1 are 

large and have been well documented by Plummer and 

others (1979). In the present study, experiments were 

performed to determine the stirring effects on the rate 

constants k 2 and k 3 in pure C02-H20 solutions before 

the large backward reaction effects of Ca2+ an HC03

-

were recognized; the results are, therefore, ambiguous. 

The values of k 1 and k2 obtained for both calcite and 

aragonite via the weight loss method (table 3) are the 

same, which are expected if the H+ and C02 reactions 

were transport controlled. A new study of stirring 

dependence in a different part of dissolution region I is 

required to resolve this question. The rate constant k3 of 

aragonite is 1.4 times greater than that of calcite under 

the same hydrodynamic conditions, strongly suggesting 

that the water reaction is not predominantly transport 

controlled for the two CaC03 polymorphs. 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 161 
Table 9. Observed (Obs.) and calculated (Calc.) log rates of dissolution and crystal growth of aragonite in Ca(HC03) 2 solutions 

[Units of concentration and rate are mmol· kg-1 and mmol· em -2 • sec-1, respectively. Rates of dissolution and crystal growth were calculated using 

equations 51 and 54, respectively, and the constants in table 3. Crystal growth rates are in square brackets; leaders(---), indicate values not calculated] 

Log Obs. Calc. Log Obs. Calc. 

Pco Calc. log log Pco Calc. log log 

Run sample Ca 2 pH Onega rate rate Run Sample Ca 2 pH Onega rate rate 

22 N-2 0.582 -0.016 4.88 0.0004 -5.77 -5.80 32 ~2 5.000 -0.540 6.29 0.5008 -7.19 -6.75 

23 ~5 0.582 -0.016 4.88 0.0004 -5.88 -5.80 40 r+-2 5.560 -0.540 6.34 0.6602 -7.65 

24 ~4 0.582 -0.016 4.88 0.0004 -5.80 -5.80 41 A-1 5.560 -0.540 6.34 0.6602 -6.86 

20 A-2 1.120 -0.016 5.16 0.0026 -5.84 -5.89 46 ~1 5.940 -0.539 6.36 0.7813 -7.14 

21 ~2 1.120 -0.016 5.16 0.0026 -5.88 -5.89 47 ~5 5.940 -0.539 6.36 0.7813 -8.23 

12 ~1 2.620 -0.016 5.51 0.0270 -6.04 -6.07 66 ~1 9.930 -0.539 6.56 2.8848 [-6.70] [-8.17] 

13 ~5 2.620 -0.016 5.51 0.0270 -6.04 -6.07 67 t-1-5 9.930 -0.539 6.56 2.8848 [-7.43] [-8.17] 

11 A-1 3.720 -0.016 5.65 0.0692 -5.95 -6.18 60 N-1 11.390 -0.539 6.62 4.0584 [-6.62] [-6.94] 

14 ~2 3.720 -0.016 5.65 0.0692 -6.12 -6.18 61 ~3 11.390 -0.539 6.62 4.0584 [-6.85] [-6.94] 

7 N-1 3.790 -0.016 5.66 0.0726 -5.98 -6.19 42 ~1 12.380 -0.540 6.65 5.0012 [-6.45] [-6.70] 

8 ~4 3.790 -0.016 5.66 0.0726 -6.01 -6.19 43 ~6 12.380 -0.540 6.65 5.0012 [-6.68] [-6.70] 

3 A-2 4.630 -0.016 5.74 0.1231 -6.08 -6.27 37 ~4 13.370 -0.540 6.68 6.0409 [-6.54] [-6.53] 

4 ~1 4.630 -0.016 5.74 0.1231 -6.23 -6.27 50 N-1 15.000 -0.539 6.73 7.9752 [-6.16] [-6.35] 

6 ~2 5.780 -0.016 5.83 0.2195 -6.37 -6.37 51 ~6 15.000 -0.539 6.73 7.9752 [-6.15] [-6.35] 

9 ~7 5.780 -0.016 5.83 0.2197 -6.34 -6.37 29 A-1 15.650 -0.540 6.74 8.8633 [-6.29] [-6.29] 

2 A-1 6.780 -0.016 5.89 0.3313 -6.39 -6.46 30 t+-1 15.650 -0.540 6.74 8.8633 [-6.27] [-6.29] 

5 N-2 6.780 -0.016 5.89 0.3313 -6.30 -6.46 44 A-1 17.700 -0.539 6.79 11.8914 [-6.11] [-6.14] 

15 A-2 7.390 -0.016 5.93 0.4135 -6.59 -6.52 45 N-2 17.700 -0.539 6.79 11.8914 [-6.17] [-6.14] 

10 ~6 8.480 -0.016 5.98 0.5867 -7.10 -6.62 84 ~2 17.940 -0.545 6.80 12.4640 [-5.90] [-6.13] 

1 ~3 9.510 -0.016 6.03 0.7824 -7.37 83 B-2 18.820 -0.545 6.82 13.9765 [-6.07] [-6.07] 

17 ~3 14.260 -0.015 6.18 2.1349 [-7 .98] [-7.56] 96 ~1 19.090 -0.541 6.82 14.3264 [-5.76] [-6.06] 

19 N-1 14.260 -0.016 6.18 2.1367 [-8.37] [-7.56] 97 ~6 19.090 -0.541 6.82 14.3264 [-5.92] [-6.06] 

16 A-1 15.630 -0.015 6.22 2.6706 [-7.36] [-7.02] 81 ~6 19.680 -0.545 6.84 15.5472 [-5.87] [-6.03] 

18 ~1 15.630 -0.016 6.22 2.6728 [-7 .30] [-7.02] 82 ~7 20.650 -0.545 6.85 17.4292 [-6.02] [-5.98] 

64 ~6 16.920 -0.016 6.25 3.2402 [-6.55] [-7.08] 77 M-5 20.720 -0.545 6.86 17.5695 [-6.12] [-5.98] 

65 ~7 16.920 -0.016 6.25 3.2402 [-6.92] [-7.08] 78 B-1 20.720 -0.545 6.86 17.5695 [-5.95] [-5.98] 

68 ~2 18.290 -0.016 6.28 3.9101 [-6.35] [-6.55] 92 M-5 21.140 -0.541 6.86 18.2552 [-6.02] [-5.96] 

69 B-1 18.290 -0.016 6.28 3.9101 [-6.56] [-6.55] 93 B-2 21.140 -0.541 6.86 18.2552 [-5.91] [-5.96] 

58 A-2 18.360 -0.016 6.28 3.9462 [-6.31] [-6.54] 79 ~4 21.520 -0.545 6.87 19.2175 [-5.95] [-5.94] 

59 M-2 18.360 -0.016 6.28 3.9462 [-6.49] [-6.54] 80 N-2 21.520 -0.545 6.87 19.2175 [-6.06] [-5.94] 

85 M-1 19.900 -0.023 6.32 4.8618 [-5.95] [-6.28] 88 N-1 23.180 ·0.541 6.89 22.6847 [-5.72] [-5.87] 

62 N-2 20.150 -0.016 6.32 4.9330 [-6.47] [-6.25] 89 M-4 23.1.80 -0.541 6.89 22.6847 [-5.83] [-5.87] 

63 ~4 20.150 -0.016 6.32 4.9330 [-6.33] [-6.25] 161 N-2 1.015 -1.012 6.11 0.0195 -6.70 -6.48 

73 M-2 20.680 -0.023 6.33 5.3299 [-6.08] [-6.19] 159 B-1 2.000 -1.012 6.39 0.1277 -6.61 -6.63 

76 N-1 20.820 -0.023 6.34 5.4164 [-6.22] [-6.18] 160 ~7 2.000 -1.012 6.39 0.1277 -6.77 -6.63 

74 A-2 21.030 -0.023 6.34 5.5476 [-6.33] [-6.16] 157 ~5 3.060 -1.012 6.57 0.4037 -6.95 -6.77 

75 l+-3 21.100 -0.023 6.34 5.5918 [-6.38] [-6.15] 158 N-1 3.060 -1.012 6.57 0.4037 -6.75 -6.77 

72 A-1 21.610 -0.023 6.35 5.9191 [-6.04] [-6.11] 155 ~2 4.010 -1.012 6.68 0.8284 -7.02 

71 ~1 21.770 -0.023 6.35 6.0240 [-5.97] [-6.09] 156 B-2 4.010 -1.012 6.68 0.8284 -7.04 

94 ~2 23.100 -0.018 6.37 6.8689 [-5.92] [-5.99] 146 B -1 5.060 -1.011 6.77 1.5207 [-8.46] 

95 A-2 23.100 -0.018 6.37 6.8689 [-6.27] [-5.99] 147 l1 -7 5.060 -1.011 6.77 1.5207 [-8.05] 

90 B-1 23.140 -0.018 6.37 6.8971 [-6.00] [-5.99] 144 A-2 6.080 -1.011 6.84 2.4453 [-7.83] [-7.97] 

91 l-1-7 23.140 -0.018 6.37 6.8971 [-6.17] [-5.99] 145 ~4 6.080 -1.011 6.84 2.4453 [-7.88] [-7.97] 

86 A-1 23.170 -0.018 6.37 6.9183 [-5.90] [-5.99] 148 N-2 6.700 -1.011 6.88 3.1362 [-7.82] [-7 .83] 

87 l-1-3 23.170 -0.018 6.37 6.9183 [-6.16] [-5.99] 142 A-1 8.060 -1.011 6.95 5.0129 [-6.82] [-6.95] 

56 A-1 23.530 -0.016 6.38 7.1343 [-6.33] [-5.96] 143 ~3 8.060 -1.011 6.95 5.0129 [-6.73] [-6.95] 

57 M-1 23.530 -0.016 6.38 7.1343 [-6.00] [-5.96] 140 M-5 10.010 -1.011 7.04 8.6181 [-6.55] [-6.57] 

38 M-7 1.060 -0.540 5.66 0.0074 -6.40 -6.27 141 N-1 10.010 -1.011 7.04 8.6181 [-6.54] [-6.57] 

39 A-2 1.060 -0.540 5.66 0.0074 -6.31 -6.27 153 A-2 10.980 -1.012 7.08 10.8519 [-6.37] [-6.45] 

35 N-2 2.000 -0.540 5.92 0.0432 -6.35 -6.40 154 t-1-4 10.980 -1.012 7.08 10.8519 [-6.40] [-6.45] 

33 N-1 3.000 -0.540 6.09 0.1295 -6.33 -6.52 138 ~2 11.900 -1.011 7.11 13.1957 [-6.37] [-6.36] 

34 l+-3 3.000 -0.540 6.09 0.1295 -6.39 -6.52 139 B-2 11.900 -1.011 7.11 13.1957 [-6.40] [-6.36] 

54 M-4 3.250 -0.539 6.12 0.1600 -6.49 -6.55 151 A-1 11.980 -1.012 7.11 13.4392 [-6.26] [-6.35] 

55 ~2 3.250 -0.539 6.12 0.1600 -6.41 -6.55 152 M-3 11.980 -1.012 7.11 13.4392 [-6.24] [-6.35] 

52 N-2 3.510 -0.539 6.15 0.1965 -6.67 -6.58 149 l-1-1 13.110 -1.012 7.14 16.7382 [-6.08] [-6.26] 

53 M-7 3.510 -0.539 6.15 0.1965 -6.70 -6.58 150 ~6 13.110 -1.012 7.14 16.7382 [-6.13] [-6.26] 

36 r+-5 3.980 -0.540 6.20 0.2750 -6.86 -6.63 136 ~1 14.010 -1.011 7.17 19.6176 [-6.04] [-6.19] 

48 A-2 4.010 -0.539 6.20 0.2797 -6.77 -6.63 137 M-6 14.010 -1.011 7.17 19.6176 [-6.06] [-6.19] 

49 ~3 4.010 -0.539 6.20 0.2797 -6.80 -6.63 110 N-2 4.060 -1.545 7.21 2.8761 [-8.06] [-8.97] 

31 A-2 5.000 -0.540 6.29 0.5008 -7.15 -6.75 108 B-1 5.040 -1.545 7.30 5.0430 [-7.31] [-7 .10] 

162 Studies in Diagenesis 
Table 9. Observed (Obs.) and calculated (Calc.) log rates of 

dissolution and crystal growth of aragonite in Ca(HC03

)

2 

solutions-Continued 

Run Sample Ca 

Log 

Pco 

2 

109 fvl-7 

106 f!.r-2 

107 l-'1-4 

104 f!.r-1 

105 M-3 

102 l-'1-5 

103 1\f-1 

100 fv:l-2 

101 B-2 

98 fv:i-1 

99 fv:i-6 

123 1\:...2 

121 Er-1 

122 H-7 

119 A~2 

120 M:-4 

134 A-2 

135 M:-4 

132 B-1 

133 M-7 

130 A~l 

131 M-3 

128 M-1 

129 M-5 

115 M-5 

116 N-1 

126 M-6 

127 N-1 

113 M-2 

114 B-2 

124 t-1-2 

125 B-2 

166 A-2 

167 M-4 

180 M-7 

178 N-1 

179 B-8 

164 A-1 

165 M-3 

176 M-2 

177 B-1 

174 B-2 

175 M-5 

172 A:-1 

173 M--3 

162 M--1 

163 M--6 

170 A:-2 

171 M--4 

168 H:·-1 

169 H·-6 

5.040 -1.545 

6.030 -1.545 

6.030 -1.545 

7.000 -1.545 

7.000 -1.545 

7. 700 -1.545 

7. 700 -1.545 

8.940 -1.545 

8.940 -1.545 

9.980 -1.545 

9.980 -1.545 

2.000 -2.545 

3.010 -2.545 

3.010 -2.545 

3.950 -2.545 

3.950 -2.545 

3.990 -2.543 

3.990 -2.543 

4.500 -2.543 

4.500 -2.543 

5.050 -2.543 

5.050 -2.543 

5.610 -2.543 

5.610 -2.543 

5.960 -2.545 

5.960 -2.545 

5.970 -2.543 

5.970 -2.543 

6.950 -2.545 

6.950 -2.545 

7.400 -2.543 

7.400 -2.543 

1.000 -3.519 

1.000 -3.519 

1.500 -3.519 

1.750 -3.519 

1.750 -3.519 

2.000 -3.519 

2.000 -3.519 

2.250 -3.519 

2.250 -3.519 

2.500 -3.519 

2.500 -3.519 

2.750 -3.519 

2.750 -3.519 

3.000 -3.519 

3.000 -3.519 

3.250 -3.519 

3.250 -3.519 

3.500 -3.519 

3.500 -3.519 

Obs. 

Calc. log 

pH Onega rate 

7.30 5.0430 [-7.01] 

7.37 7.9791 [-6.66] 

7.37 7.9791 [-6.87] 

7.43 11.6258 [-6.57] 

7.43 11.6258 [-6.57] 

7.46 14.7480 [-6.43] 

7.46 14.7480 [-6.32] 

7.52 21.3151 [-6.18] 

7.52 21.3151 [-6.51] 

7.56 27.8624 [-6.07] 

7.56 27.8624 [-6.14] 

7.91 4.0933 [-7.30] 

8.07 11.8039 [-6.50] 

8.07 11.8039 [-6.56] 

8.18 23.1848 [-6.26] 

8.18 23.1848 [-6.27] 

8.18 23.6539 [-6.34] 

8.18 23.6539 [-6.41] 

8.22 31.6017 [-5.99] 

8.22 31.6017 [-6.04] 

8.26 41.4876 [-5.94] 

8.26 41.4876 [-6.05] 

8.30 52.9141 [-5.82] 

8.30 52.9141 [-5.99] 

8.32 60.9786 [-5.98] 

8.32 60.9786 [-5.82] 

8.32 60.9620 [-5.77] 

8.32 60.9620 [-5.74] 

8.38 85.8562 [-5.74] 

8.38 85.8562 [-5.93] 

8.40 98.0406 [-5.60] 

8.40 98.0406 [-5.81] 

8.57 4.9308 [-7 .11] 

8.57 4.9308 [-6.93] 

8.72 13.4562 [-6.46] 

8.78 19.3214 [-6.36] 

8.78 19.3214 [-6.37] 

8.82 26.1761 [-6.19] 

8.82 26.1761 [-6.29] 

8.86 33.9502 [-6.27] 

8.86 33.9502 [-6.54] 

8.89 42.5737 [-6.53] 

8.89 42.5737 [-6.53] 

8.92 51.9797 [-6.02] 

8.92 51.9797 [-6.02] 

8.95 62.1056 [-5.79] 

8.95 62.1056 [-5.92] 

8.97 72.8949 [-6.01] 

8.97 72.8949 [-6.28] 

8.99 84.2951 [-5.58] 

8.99 84.2951 [-5.71] 

Calc. 

log 

rate 

[-7.10] 

[-6.77] 

[-6.77] 

[-6.58] 

[-6.58] 

[-6.47] 

[-6.47] 

[-6.32] 

[-6.32] 

[-6.22] 

[-6.22] 

[-7.43] 

[-6.66] 

[-6.66] 

[-6.38] 

[-6.38] 

[-6.37] 

[-6.37] 

[-6.26] 

[-6.26) 

[-6.17] 

[-6.17] 

[-6.08] 

[-6.08] 

[-6.04] 

[-6.04] 

[-6.04] 

[-6.04] 

[-5.93] 

[-5.93] 

[-5.89] 

[-5.89] 

[-7.17] 

[-7.17] 

[-6.58] 

[-6.43] 

[-6.43] 

[-6.32] 

[-6.32] 

[-6.22] 

[-6.22] 

[-6.15] 

[-6.15] 

[-6.08] 

[-6.08] 

[-6.02] 

[-6.02] 

[-5.97] 

[-5.97] 

[-5.92] 

[-5.921 

Comparison of the kinetics of dissolution and 

precipitation of calcite and aragonite 

Figure 11 is a comparison of the kinetics of dissolution and precipitation of calcite and aragonite at 0.96 atm 

C02 partial pressure in Ca(HC03

) 2 solutions. As shown, 

the rates of dissolution are comparable (those of aragonite 

slightly faster), but the rates of precipitation are significantly different (aragonite precipitates significantly slower 

than calcite). Three factors contribute to this large difference in the rates of precipitation of these two CaC03 

minerals. First, the solubility of aragonite is greater than 

that of calcite; thus, the equilibrium points are not the 

same (fig. 11). Second, k4 of aragonite is smaller than 

that of calcite. Third, the intercept of the backward rate 

line with the zero rate line (fig. 11) differs for the two 

minerals, occurring at 0 values of 1.45 and 2.0 for calcite 

and aragonite, respectively, over the range of P co inves2 

tigated. In terms of our rate equations, these three factors favor the precipitation of calcite at the expense of 

aragonite in Ca(HC03

) 2 solutions. Supersaturation (0) 

of at least 1.45 for calcite and 2.0 for aragonite are 

necessary before the rate becomes proportional to the 

activity product of Ca2+ and HC03

-. Below these supersaturation values, backward reactions probably occur 

only at higher energy sites. 

Even though both calcite and aragonite can 

precipitate from dilute supersturated solutions, the rate 

of formation of calcite is significantly faster than that 

of aragonite under comparable conditions. The rate of 

dissolution of aragonite from undersaturated solutions 

is slightly faster than that of calcite. These two kinetic 

processes may explain the near absence of aragonite in 

natural fresh water environments at low temperatures. 

Comparison of the dissolution kinetics of calcite 

and aragonite with dolomite 

The model describing the dissolution of dolomite 

(Busenberg and Plummer, 1982) is similar to that for 

calcite and aragonite. The main difference between the 

CaC03 minerals and dolomite is the presence of two different components in dolomite that behave as separate 

entities on the surface. The rate of reaction is rapid for 

the CaC03 component of dolomite, but significantly 

slower for the MgC03 component; thus, the surface 

becomes enriched with the MgC03 component. The rate 

determining step in dolomite dissolution is the removal 

of the MgC03 component from the surface. 

The forward reaction mechanisms for dolomite 

dissolution have been interpreted as two consecutive 

steps in three parallel reactions. For the CaC03 

minerals, there are only three parallel reactions. The 

generalized forward equation for the dissolution of all 

three minerals is: 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 163 
Table 10. Observed (Obs.) and calculated (Calc.) rates of dissolution and crystal growth of calcite in Ca(HC03

)

2 solutions 

[Units of concentration and rate are mmol·kg-1 and mmol·cm-2·sec-1, respectively. Rates of dissolution and crystal growth were calculated using 

equations 51 and 54, respectively, and the constants in table 3. Crystal growth rates are in square brackets; leaders(---), indicate values not calculated] 

Log Chs. Calc. Log Chs. Calc. 

Pco2 calc. log log Pco2 Calc. log log 

Run sample Ca pH Qnega rate rate Run sample ca pH Qnega rate rate 

821 24 0.130 0.9628 4.32 0.0000 -5.75 -5.75 785 20 5.957 0.4819 6.14 0.6595 -7.39 

817 20 0.205 0.9632 4.47 0.0000 -5.75 -5.81 839 10 6.570 0.4812 6.18 0.8502 -7.65 

813 16 0.222 0.9639 4.50 0.0000 -5.81 -5.82 784 19 7.939 0.4819 6.26 1.3769 [-9.52] 

809 12 0.312 0.9639 4.63 0.0001 -5.81 -5.87 827 14 8.650 0.4822 6.29 1.7101 [-7.35] [-7.67] 

758 9 0.915 0.9562 5.08 0.0021 -5.93 -6.03 795 14 9.261 0.4832 6.32 2.0269 [-7.23] [-7.05] 

759 10 1.777 0.9562 5.35 0.0132 -6.04 -6.15 807 10 9.550 0.4804 6.33 2.2023 [-7.44] [-6.85] 

762 13 2.640 0.9553 5.51 0.0388 -6.08 -6.24 791 10 10.023 0.4839 6.35 2.4681 [-6.87] [-6.64] 

763 14 3.547 0.9553 5.64 0.0856 -6.19 -6.32 803 22 10.850 0.4806 6.38 3.0286 [-6.42] [-6.41] 

770 21 3.985 0.9643 5.68 0.1155 -6.20 -6.35 794 13 10.860 0.4832 6.38 3.0193 [-6.51] [-6.41] 

766 17 4.397 0.9641 5.72 0.1497 -6.21 -6.38 799 18 11.538 0.4806 6.40 3.5283 [-6.53] [-6.28] 

771 22 4.971 0.9643 5.77 0.2065 -6.28 -6.43 823 10 11.600 0.4814 6.41 3.5696 [-6.69] [-6.27] 

767 18 5.350 0.9641 5.80 0.2502 -6.34 -6.46 790 9 11.852 0.4839 6.41 3. 7454 [-6.38] [-6.23] 

778 13 5.952 0.9626 5.84 0.3304 -6.38 -6.51 819 22 12.490 0.4816 6.43 4.2834 [-6.19] [-6.14] 

774 9 6.394 0.9629 5.87 0.3974 -6.46 -6.54 815 18 13.000 0.4819 6.45 4.7241 [-6.23] [-6.08] 

779 14 7.179 0.9626 5.92 0.5351 -6.64 -6.60 811 14 13.980 0.4819 6.48 5.6459 [-6.16] [-5.98] 

775 10 7.407 0.9629 5.93 0.5795 -6.85 -6.62 806 9 14.266 0.4804 6.49 5.9509 [ -6.03] [-5.95] 

783 18 7.761 0.9639 5.95 0.6520 -7.02 822 9 14.970 0.4814 6.50 6.6790 [-5.98] [-5.89] 

838 9 8.270 0.9624 5.97 0.7674 -7.15 802 22 15.230 0.4806 6.51 6.9763 [-5.96] [-5.87] 

842 19 8.758 0.9618 5.99 0.8878 -9.38 798 17 15.347 0.4806 6.51 7.1073 [-5.94] [-5.87] 

843 20 8.758 0.9618 5.99 0.8878 -9.48 18 21 15.600 0.4816 6.52 7.3801 [-5.67] [-5.85] 

840 17 8.886 0.9618 6.00 0.9210 -9.58 814 17 16.740 0.4819 6.55 8.7496 [-5.82] [-5. 77] 

841 18 8.886 0.9618 6.00 0.9210 -9.76 810 13 17.420 0.4819 6.56 9.6318 [-5.81] [-5.73] 

844 21 9.355 0.9618 6.02 1.0484 [-9.87] 837 24 18.240 0.4815 6.58 10.7676 [-5.75] [-5.68] 

845 22 9.355 0.9618 6.02 1.0484 [-9.63] 835 22 18.990 0.4817 6.60 11.8553 [-5.71] [-5.65] 

825 12 9.710 0.9644 6.03 1.1482 [-9.18] 833 20 19.740 0.4819 6.61 13.0006 [-5.71] [-5.61] 

782 17 10.283 0.9639 6.06 1.3262 [-8.83] 831 18 20.810 0.4820 6.63 14.7408 [-5.63] [-5.56] 

793 12 10.759 0.9665 6.07 1.4809 [-8.60] [-7.70] 829 16 22.040 0.4822 6.65 16.8872 [-5.61] [-5.51] 

789 24 11.779 0.9678 6.11 1.8525 [-7.30] [-7.26] 826 13 23.320 0.4822 6.67 19.2990 [-5.53] [-5.47] 

805 24 11.779 0.9608 6.11 1.8659 [-7.44] [-7.26] 857 20 0.209 0.0958 5.45 0.0003 -6.49 -6.51 

792 11 12.658 0.9665 6.14 2.2163 [-6.54] [-6.55] 860 23 0.274 0.0958 5.56 0.0007 -6.56 -6.54 

801 20 12.809 0.9611 6.14 2.2948 [-6.84] [-6.49] 858 21 0.415 0.0958 5.74 0.0022 -6.67 -6.60 

797 16 13.281 0.9612 6.16 2.5084 [-6.57] [-6.35] 859 22 0.892 0.0958 6.06 0.0192 -6.83 -6.74 

788 23 14.069 0.9678 6.18 2.8700 [-6.50] [-6.18] 855 18 1.570 0.0958 6.30 0.0929 -6.92 -6.88 

804 23 14.190 0.9608 6.18 2.9521 [-6.27] [-6.16] 856 19 2.100 0.0958 6.42 0.2063 -7.15 -6.98 

800 19 14.720 0.9611 6.20 3.2280 [-6.22] [-6.07] 876 24 2.530 0.0951 6.50 0.3445 -7.09 -7.07 

796 15 15.220 0.9612 6.21 3.5018 [-6.16] [-6.00] 853 16 3.130 0.0959 6.58 0.6057 -7.46 

820 23 15.340 0.9628 6.21 3.5635 [-6.13] [-5.99] 865 13 3.460 0.0952 6.63 0.7972 -7.73 

786 21 15.470 0.9641 6.22 3.6326 [-6.06] [-5.97] 864 12 4.740 0.0952 6.75 1.8280 [-8.95] [-8.35] 

787 22 15.470 0.9641 6.22 3.6326 [-6.05] [-5.97] 848 9 5.290 0.0959 6.79 2.4140 [-7.36] [-7.14] 

816 19 16.060 0.9632 6.23 3.9824 [-6.02] [-5.91] 874 22 5.860 0.0951 6.84 3.1715 [-7.09] [-6.83] 

812 15 16.790 0.9639 6.25 4.4325 [-5.95] [-5.84] 852 15 6.520 0.0959 6.88 4.1383 [-6.75] [-6.62] 

808 11 17.770 0.9639 6.27 5.0837 [-5.88] [-5.76] 854 17 6.520 0.0958 6.88 4.1417 [-6.73] [-6.62] 

836 23 18.560 0.9629 6.29 5.6510 [-5.81] [-5.70] 851 14 7.960 0.0959 6.96 6.8714 [-6.53] [-6.34] 

834 21 18.930 0.9635 6.29 5.9219 [-5.75] [-5.67] 871 19 8.500 0.0952 6.98 8.1670 [-6.41] [-6.26] 

832 19 19.870 0.9639 6.31 6.6489 [-5. 75] [-5.62] 868 16 9.660 0.0952 7.03 11.2430 [-6.30] [-6.12] 

830 17 21.300 0.9640 6.34 7.8486 [-5.62] [-5.54] 850 13 9.990 0.0959 7.04 12.1263 [-6.27] [-6.09] 

828 15 21.920 0.9644 6.35 8.4000 [-5.63] [-5.51] 872 20 10.860 0.0952 7.08 15.0257 [-6.12] [-6.01] 

824 11 23.040 0.9644 6.37 9.4550 [-5.55] [-5.46] 873 21 10.860 0.0952 7.08 15.0257 [-6.12] [-6.01] 

760 11 0.696 0.4781 5.26 0.0019 -6.19 -6.25 863 11 11.230 0.0952 7.09 16.3100 [-8.30] [-5.98] 

761 12 1.521 0.4781 5.59 0.0171 -6.24 -6.39 846 10 11.690 0.0959 7.10 17.8625 [-6.25] [-5.95] 

764 15 2.095 0.4777 5.72 0.0413 -6.37 -6.46 869 17 12.930 0.0952 7 .·15 23.0185 [-5.96] [-5.86] 

765 16 2.709 0.4777 5.83 0.0829 -6.39 -6.53 870 18 12.930 0.0952 7.15 23.0185 [-6.01] [-5.86] 

772 23 3.171 0.4821 5.89 0.1254 -6.48 -6.58 866 14 13.650 0.0952 7.17 26.2476 [-5.95] [-5.82] 

768 19 3.539 0.4821 5.93 0.1681 -6.49 -6.62 867 15 13.650 0.0952 7.17 26.2476 [-5.97] [-5.82] 

773 24 3.821 0.4821 5.96 0.2061 -6.55 -6.65 861 9 14.460 0.0952 7.19 30.1548 [-5.91] [-5.77] 

769 20 4.168 0.4821 6.00 0.2593 -6.59 -6.68 862 10 14.460 0.0952 7.19 30.1548 [-6.03] [-5.77] 

780 15 4.648 0.4813 6.04 0.3460 -6.67 -6.74 847 11 16.840 0.0959 7.24 43.0696 [-5.71] [-5.65] 

776 11 5.026 0.4815 6.07 0.4244 -6.76 -6.77 849 12 22.860 0.0959 7.36 87.9868 [-5.39] [-5.43] 

781 16 5.267 0.4813 6.09 0.4797 -6.86 -6.80 

777 12 5.605 0.4815 6.12 0.5637 -6.99 -6.84 

164 Studies in Diagenesis 
a: 

w 

c.. 

en 

w 

...J 0 Q 11.!) ~z -o ::Jo 

-w 

~CJ) 

Za:: 

.. w 11.0 we.. 

~a: a::W 

u..l--

0~ 

~1-- :::c z 10.!5 

~---~ a: 

<CW 

(!)0:: 

o<C 

...J::J 

w 0 10.0 

>en 

~ C) 

w 

z 

Aragonite 

9 1- .,) ~----~----~----~------~-----L----~------~----~----~----~ 

0.0 0.5 1.0 1.5 2.0 2.5 

NEGATIVE LOGARITHM OF (1-Q) 

Figure 9. Dissolution rates of calcite and aragonite as function of -log(l - 0) r:~ear equilibrium. Rate was determined 

from measurement of pH in free-drift experiments between 0 of 0.6 and 0.99. 

where n is 1 for calcite and aragonite and 0.5 for 

dolomite. The rate-limiting step in dolomite, the removal 

of the MgC03 component, give rise to the observed 

half-order kinetic!i at 25 °C. 

The forward rate constants k1

, k2, and k3 for 

dolomite~ are highly variable among different dolomite 

samples. Under identical experimental conditions, the rate 

constants vary by as much as one order of magnitude between different samples. Clear hydrothermal dolomites 

dissolve much more slowly than do imperfect cloudy 

crystals, and the rates probably vary as a function of 

crystal defect density. No differences in forward rate constants were observ,~d between synthetic and seven natural 

aragonites from different locations and origins. Similarly, 

no differences in rate were observed between a natural 

and two different synthetic calcites. 

No reversible backward reactions were observed for 

dolomites. The most important backward reaction in dolomites results from the reaction of adsorbed HC03

- with 

positive sites at the surface. This reaction effectively stops 

dissolution far from "equilibrium" and can be described 

by a Langmuir isotherm. Our results show that the rate 

of dissolution of dolomite in extremely dilute solutions is 

about 100 times slower than that of calcite and aragonite. 

Relatively low concentrations of HC03

- can almost completely stop the dissolution of dolomite far from "equilibrium~' Our combined results on the dissolution of dolomite 

suggest that reversible dolomite-solution equilibrium 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 165 
U) 

0 

,.... 

+~ 

~ 

~ 

2r---~----~--~----~--~--~ 

i 

0.0317 percent 

100 

30 

10 

3 

0.298 

0.0317 

Aragonite 

8.55 

3.96 

1.95 

0.893 

0.198 

0.0457 I -3 ....._____---L_ __ _._._ _ __J __ ___J__ __ ...J.__ _ ___J 

Calcite 

EOUILBRIUM POINTS 

Percent C02 aca2t aHCOi X 1 0 5 

100 6.80 

50 4.36 

10 1.55 

Q:: 0 10 20 30 0 10 20 30 

ACTIVITY PRODUCT OF CALCIUM AND BICARBONATE IONS X 105 

Figure 10. Dissolution and precipitation rates of calcite and aragonite as function of R - k 1 aH+, gas composition 

(numerical values represent mole percent C02 in the C02-N2 gas mixtures), and activity product of Ca2+ and HC03 -. 

Each point represents an independent measurement of weight gain or weight loss of single crystal of aragonite or calcite. 

cannot be achieved in the laboratory at less than 45 <>c. 

Backward reactions have been demonstrated that can form 

calcite and aragonite in calcium bicarbonate solutions. 

SUMMARY AND CONCLUSIONS 

This paper presents experimental results, at 25 °C, 

of (1) dissolution of calcite and aragontie in C02-H20-

HCl-KHC03-KOH-CaC12-Ca(HC03)2 solutions, and (2) 

crystal growth of calcite and aragonite in Ca(HC03

)2 

solutions. In C02-H2

0 solutions far from equilibrium, 

the rate of dissolution of CaC03 is described by 

where the rate constants k 1, k 2 , and k 3 are given in table 

3 for the experimental conditions. With the addition of 

166 Studies in Diagenesis 

single salts of KOH, KHC03, K2C03, or CaC12 far from 

equilibrium, the rate of dissolution decreases, and the 

change in rate of dissolution is accounted for by Langmuir 

isotherms for OH-, HC03

-, CO~-, and Ca2+ on the 

calcite or aragonite surface. The decrease in rate is attributed to backward reactions between adsorbed species and 

surface species and is described by terms of the form 

Rb = R0 a; /(A; + B; a) for the ith species adsorbing. 

Values of A and B for OH-, HC03

-, CO~-, and Ca2+ are 

given in table 3. The backward adsorption isotherms 

appear to be additive during dissolution in the range 0 

to 0.6 0 corresponding to the rate equation 

R = Ro(1- EA. :;B.a.)· I I I 

For example, the backward reaction term, k 4 aca2+ aHco -

3 
2.0~---~--~--~----~--~--~--~----~--~--~--~--~ 

a: 

w 

tu 

~ 

!z 

W«> 

oo ,.... 

1.0 \ 

(!)~ 

'~~ (!)'~x Aragonite 

!\.~ equilibrium 

~X 

<Co 

=>z 00 

cno 

0.0 '~~ - f -=..-~---><--x--...x Aragonite 

Cal 't (!)(!) (!) (!) X~~ X 

c.1e . >-<~~ 'L 

a:w 

equilibrium ·~.,_ x ~ 

~(/.) 

en a: wW 

...J Q.. 

0 

~ 

~ 

...J 

~ 

z 

+~ 

~ " 

tt 

-1.0 ~(!) X 

"-...(!) 

~- -2.0 

-3.0 L...._ __ ....~-__--L...... _ ___J,_ _ ___._ __ l...___....I-__--J._ _ ___._ _ ___,_ __ ..__ _ _,___ (!)""' ...... 

0 5 10 15 20 25 30 

ACTIVITY PRODUCT OF CALCIUM AND BICARBONATE IONS 

Figun! 11. Comparison of dissolution and precipitation rates of calcite and aragonite single crystals as function of activity 

product of Ca2+ and HC03- at 0.96 P co . Note that rates of precipitation of calcite are significantly faster than those of 2 

aragonite, whereas aragonite dissolves only slightly faster than calcite. 

of Plummer and ·:>thers (1978) is completely accounted 

for by summing two backward reaction isotherm expressions (for Ca2+ and HC03

- were determined separately 

in CaC12 and KHC03 solutions) in describing dissolution 

rate in Ca(HC0 2 solutions to 0 of 0.6. Conceivably, 

every bulk solution species that can react with surface sites 

to produce calcite or aragonite does so and contributes 

to the overall observed backward rate. Bulk solution 

anions and cations precipitate in this process. The inital 

step in all back reactions may always involve the adsorption of the reactiv~~ species into the Helmholtz layer. The 

next step may be the reaction between the adsorbate and 

the surface reactive sites. This reaction was found to be 

proportional to the activity of the adsorbate. 

In Ca(HC01

) 2 solutions at undersaturations higher 

than 0 of 0.6, the isotherm relations are no longer simply 

additive, and we have found that an empirical relation 

R = k 5(l - O)n adequately describes dissolution rate to 

equilibrium. The constants k 5 and n are given in table 3 

for calcite and aragonite in Ca(HC03

}2 solutions, but in 

general these constants are expected to depend on composition of the solution. A similar empirical relation may 

describe growth of calcite and aragonite near equilibrium. 

In supersaturated Ca(HC03) 2 solutions of 0 = 1.5 to 

more than 100 for calcite and 2.0 to more than 100 for 

aragonite, the rate of crystal growth is given by 

over the range of P co investigated (10-3·

5 to 1.0 atm). 2 

Our results show that kinetic factors account for 

the absence of aragonite in fresh-water environments. In 

subsaturated water aragonite dissolves faster than calcite, 

and in supersaturated waters we have found that calcite 

grows significantly faster than aragonite. 

Dissolution and Crystal Growth Kinetics of Calcite and Aragonite 167 
REFERENCES CITED 

Adamson, A. W., 1976, Physical chemistry of surfaces [3d ed.]: 

New York, Wiley, 698 p. 

Busenberg, E., and Plummer, L. N., 1982, The kinetics of 

dissolution of dolomite in C02-H20 systems at 1.5 ° to 

65 "C and 0 to 1 atm P co : American Journal of Science, 

v. 28~ p. 45-78. 2 

Diehl, H., 1964, Calcein, calmagite, and o, o' -dihydroyazobenzene. Titrimetric, colorimetric, and fluorometric reagents for calcium and magnesium: Columbus, Ohio, The 

G. Frederick Smith Chemical Co., 24 p. 

Dreybrodt, W., 1980, Deposition of calcite from thin films of 

natural calcareous solutions and the growth of speleothems: 

Chemical Geology, v. 29, p. 89-105. 

____ 1981a, Mixing corrosion in CaC03-C02-H20 systems and its role in the karstification of limestone areas: 

Chemical Geology, v. 32, p. 221-236. 

____ 1981 b, The kinetics of calcite precipitation from thin 

films of calcareous solutions and the growth of speleothems; revisited: Chemical Geology, v. 32, p. 237-245. 

____ 1981c, Kinetics of the dissolution of calcite and its 

applications to karstification: Chemical Geology, v. 31, 

p. 245-269. 

Foxall, T., Peterson, G. C., Rendall, H. M., and Smith, A. 

L., 1979, Charge determination at calcium salt/aqueous 

solution interface: Journal of the Chemical Society, Faraday Transactions I, v. 75, p. 1034-1039. 

Friedman, G. M., 1959, Identification of carbonate minerals 

by staining methods: Journal of Sedminentary Petrology, 

v. 29, p. 87-97. 

House, W. A., 1981a, Kinetics of crystallization of calcite from 

calcium bicarbonate solutions: Journal of the Chemical 

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Surfaces, v. 2, p. 119-131. 

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of glass, mica, and platinum: Journal of the American 

Chemical Society, v. 40, p. 1361-1403. 

Morse, J. W., 1978, Dissolution kinetics of calcum carbonate 

in sea water. V. Effects of natural inhibitors and the position of the chemical lysocline: American Journal of Science, 

v. 274, p. 638-647. 

Morse, J. W., Kanel, J.D., and Harris, K., 1979, Dissolution 

kinetics of calcium carbonate in seawater. VII. The dissolution kinetics of aragonite and pteropod tests: American 

Journal of Science, v. 9, p. 488-502. 

168 Studies in Diagenesis 

Mullin, J.W., 1972, Crystallization [2d ed.]: Cleveland, Ohio, 

CRC Press, 415 p. 

Plummer, L. N., and Busenberg, E., 1982, The solubilities of 

calcite, aragonite, and vaterite in C02-H20 solutions between 0 ° and 90 "C, and an evaluation of the aqueous 

model fore the system CaC03-C02-H20: Geochimica et 

Cosmochimica Acta, v. 46, p. 1011-1040. 

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of calcite in C02-saturated solutions at 25 "C and 1 

atmosphere total pressure: Geochimica et Cosmochimica 

Acta, v. 40, p. 191-202. 

Plummer, L. N., Wigley, T. M. L., and Parkhurst, D. L., 1978, 

The kinetics of calcite dissolution in C02-water systems at 

5° to 60° and 0.0 to 1.0 atm C02: American Journal of 

Science, v. 278, p. 179-216. 

____ 1979, Critical review of the kinetics of calcite dissolution and precipitation, in Jenne, E. A., ed., Chemical 

modeling in aqueous systems: American Chemical Society 

Symposium Series 93, p. 537-573. 

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of MgC03 ° and CaSO 4 ° ion pairs as a function of ionic 

strength: Geochimica et Cosmochimica Acta, v. 40, 

p. 549-554. 

Reddy, M. M., 1977, Crystallization of calcium carbonate in 

the presence of trace concentrations of phosphoruscontaining anions. I. Inhibition by phosphate and glycerophosphate at pH 8.8 and 25 "C: Journal of Crystal Growth, 

v. 41, p. 287-295. 

Reddy, M. M., Plummer, L. N., and Busenberg, E., 1981, 

Crystal growth of calcite from calcium bicarbonate solutions at constant P co and 25 "C; A test of a calcite disso- 2 

lution model: Geochimica et Cosmochimica Acta, v. 45, 

p. 1281-1289. 

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calcium carbonate in the presence of metal ions. I. Inhibition by magnesium ion at pH 8.8 and 25 "C: Journal of 

Crystal Growth, v. 50, p. 470-480. 

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waters: U.S. Geological Survey Journal of Research, v. 2, 

p. 233-248. 
Effect of Magnesium Ions on 

Calcium Carbonate Nucleation and Crystal Growth 

Dilute A~~ueous Solutions at 25°C 

. 1n 

By MICHAEL M. REDDY, U.S. Geological Survey, Box 25046, Denver Federal Center, 

MS 407, Denver, Colorado 80225 

CONTENTS 

Abstract 170 

Introduction 17'1) 

Materials and methods 170 

Results and discussion 171 

Nucleation in labile solutions 172 

Nucleation in metastable solutions 174 

Calcite-seeded crystallization 17 5 

Approach to calcite equilibrium 176 

Conclusions 17~• 

References cited 182 

FIGURES 

1. Plot of caldte supersaturation versus time for calcium carbonate precipitates 

from stirred and quiescent labile solutions 171 

2. Scanning electron micrographs of calcium carbonate precipitates from stirred 

and unstirred solutions 172, 173 

3-8. Plots: 

3. Number of nucleation particles and induction interval versus calcite supersaturation for precipitation without and with magnesium ions 176, 177 

4. Calcium carbonate nucleation rate versus calcite supersaturation in solution without and with magnesium ions 178 

5. Calcite 1;rystallization reduced-rate constant as a function of concentration 

of magnesium ion in solution 179 

6. Langmuir isotherm plot of k/(k0 -k) versus reciprocal of magnesium ion 

concentration 179 

7. Solution calcite supersaturation versus time for calcite-seeded crystallization wit bout and with magnesium ions l80 

8. Steady-:;tate supersaturation versus initial total magnesium ion concentration and initial seed-crystal concentration 181 

TABLES 

1. Initial solution concentration and calcite supersaturation values for calcium carbonate nucleation from labile-supersaturated solutions at pH 10.3 172 

2. Calcium carbonate nucleation in presence and absence of magnesium ion at 

25°C 174 

3. Experimental1;onditions, seed-crystal concentrations, and rate constants for calcite 

growth in the presence and absence of magnesium ion 178 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 169 
Abstract 

Calcium carbonate formation has been examined within 

a range of solution supersaturations and magnesium-to-calcium 

ratios. Potentiometry, turbidimetry, and chemical analysis were 

used to monitor changes in solution composition; solid phases 

were characterized by optical and scanning electron microscopy, specific surface area measurements, and X-ray diffractometry. Nucleation rates, particle numbers, precipitate composition, and crystal morphology were found to be a function 

of supersaturation level, solution hydrodynamics, and magnesium ion concentration. The crystallization of calcite from 

metastable supersaturated solutions was studied using a seededgrowth technique. Under all conditions, calcite crystallization 

rates were proportional to the square of relative supersaturation and were controlled by a surface reaction. Kinetic analysis 

demonstrated that the observed reduction in the rate of calcite 

crystallization in the presence of magnesium ion resulted from 

a decrease in the reaction rate constant. This observed reduction was consistent with a Langmuir adsorption model, implying that the inhibition mechanism involved reversible adsorption of magnesium ions at growth sites on the calcite crystal 

surface. The formation of an epitaxial magnesium-rich surface 

phase was not significant. Crystallization experiments generally 

did not reach calcite equilibrium. Supersaturation at the termination of these experiments depended on the initial magnesium ion concentration and seed crystal-to-solution ratio, 

demonstrating a kinetically controlled reaction. 

INTRODUCTION 

Rates and mechanisms of carbonate mineral reactions in surface and subsurface waters are of considerable 

interest in low-temperature aqueous geochemistry and are 

important aspects of diagenesis in sedimentary rocks. 

Heterogeneous reactions of calcium carbonate have been 

studied extensively (Reddy, 1978, and references therein). 

Several investigators have demonstrated the importance 

of kinetic factors in the approach of carbonate minerals 

to equilibrium and the influence of kinetic inhibitors on 

the formation and dissolution of carbonate minerals. 

Magnesium ion has commonly been cited as a mediator 

of calcium carbonate mineral kinetics. For example, 

Thorstenson and Plummer (1977, 1978) and Berner (1978) 

examined the influence of magnesium ion on the formation of calcium carbonate over a range of natural and 

experimental conditions; however, relatively little is known 

about the reaction rates or mechanistic details of this 

inhibitory process. 

Calcite-seeded growth in moderately supersaturated 

solutions follows a parabolic rate equation in both the 

presence and absence of magnesium ion (Reddy, 1978). 

In these experiments, magnesium ion reduced the rate of 

calcite crystallization by a mechanism involving the blockage of growth sites on the crystal surface. No evidence 

was obtained to support the hypothesis that a magnesiumrich surface phase formed during crystallization. Reddy 

170 Studies in Diagenesis 

and Nancollas (1976) reported that high concentrations 

of dissolved magnesium ion markedly influenced the 

formation of calcium carbonate by spontaneous 

precipitation. 

In the present study, the formation of calcium carbonate has been characterized by chemical analysis, potentiometry, X-ray diffraction, and optical and electron 

microscopic methods over a wide range of supersaturation and solid/solution ratios in the presence and absence 

of magnesium ion. Supersaturation, Oc, is herein defined 

as the calcium carbonate ion activity product (lAP) 

divided by the calcite thermodynamic solubility product, 

(Ksp), that is, Oc = /API Ksp· 

Supersaturation values are commonly divided into 

two classes: labile solutions and metastable solutions. This 

differentiation is based on the time required to form a 

precipitate in the solution. Labile solutions have high 

supersaturations; nucleation occurs within a brief (typically a few seconds) interval after solution preparation. Metastable solutions, on the other hand, have 

intermediate supersaturations; nucleation follows a specific well-defined interval, t1 (referred to as the induction 

interval), typically from a minimum of several minutes 

to several hours. Designation of supersaturation values 

for each class is dependent on the mineral being studied 

and the experimental conditions. For calcium carbonate, 

labile solutions have supersaturations greater than 100, 

and metastable solutions have supersaturations less than 

100. Metastable solutions can be prepared with low supersaturation values (5-7), which do not nucleate within 24 

hours. These metastable solutions are used for seededgrowth experiments. 

MATERIALS AND METHODS 

Grade-A glassware, reagent-grade chemicals, and 

deionized, distilled water filtered through Millipore1 

filters were used in all experiments. Prior to each experiment, reagent solutions were filtered through prewashed 

0.22-~-tm Millipore filters to minimize the presence of 

heteronuclei. 

Details of the nucleation experiments in labilesupersaturated solutions have been presented elsewhere 

(Reddy and Nancollas, 1976). These solutions were 

prepared by mixing thermostated sodium carbonate and 

calcium chloride solutions in a large reaction vessel as 

rapidly as possible. The progress of the precipitation 

reaction was followed by the analyses of filtered samples withdrawn from the reaction cell. The calcium concentration in the filtered aliquots was determined by 

ethylene-diamine tetraacetic acid (EDTA) titration with 

a micrometer buret in the presence of a Calcine indicator. 

1

Use of trade and company names is for descriptive purposes only 

and does not imply endorsement by the U. S. Geological Survey. 
Magnesium ion analyses were performed in the same 

manner (by difference) with the appropriate indicator 

(Eriochrome Black T). Total carbonate concentrations in 

solution were calcuJated from dilution of stock solution 

and separately from titration of the filtrate with 0.01 N 

sulfuric acid. 

pH was measured at 25.00 ± 0.02°C with a Corning l\tlodel12 pH meter, a Sargent Combination Electrode, 

and a Sargent .M:odel SP recorder. Electrodes were 

calibrated prior to and following each experiment with 

National Bureau of Standards standard buffer solutions 

at pH 6.865 (0.025 M potassium dihydrogen phosphate, 

0.025 M disodium hydrogen phosphate) and pH 9.18 (0.01 

M sodium borate) at 25 °C. 

Tht! preparation of metastable supersaturated solutions has been described previously (Reddy, 1978). 

Calcium, magnesium, and total carbonate concentrations 

determined by analysis of the metastable solution shortly 

after preparation agreed with the concentration values 

calculated from the dilution of stock solution. Solution 

pH was monitored continuously, and turbidity was determined by a light-seattering technique on aliquots withdrawn from time to time during the nucleation experiments. Estimates of the length of the induction period 

from pH and turbidity measurements agreed within 5 percent for the same experiment. 

Precipitates for chemical and X-ray diffraction 

analysis and for optical and electron microscopy were 

obtained by filtering 10 to 500 milliliters of the growth 

suspension through a 0.22-p.m Millipore filter, followed 

by air-drying. Precipitates on the filter surface were coated 

with carbon and gold and examined by scanning electron 

microscopy (SEM) using an ETEC Autoscan instrument. 

Dried precipitates for optical microscopy were slurried in 

a small amount of ethylene glycol and examined with a 

Zeiss RA microscope. Specific surface areas of calcium 

carbonate precipitates were determined with a Quantachrome Monosorb surface-area analyzer. X-ray diffraction (XRD) analyses were done using Ni-filtered CuKa 

radiation with a Siemens diffractometer or with a DebyeScherrer powder camera with a Norelco X-ray generator. 

Seeded-growth experiments began by inoculating a 

stirred supersatura~ed solution with a known weight of 

calcite seed at seed concentrations of 0.05 to 0.5 g/L for 

magnesium-free solutions. Inasmuch as the rate of formation of calcite iL solutions containing magnesium ion 

can be considerably slower than in magnesium-free solutions, two modifications of this laboratory scheme were 

adopted for selected experiments. Several runs used seed 

crystal concentrations greater than 0.5 g/L to enhance 

calcite formation and the approach to equilibrium. Other 

runs were: reinoculated with calcite seeds after a specific 

growth interval, and the subsequent reaction was followed 

to completion. 

Several measurements of the crystallization rate in 

seeding solutions employed a constant pH of 9 using the 

Radiometer pH stat system: TTT2 titrator, titrograph, 

ABU 12 auto buret, a closed reaction cell with stirrer, and 

separate pH and reference electrodes. 

RESULTS AND DISCUSSION 

Nucleation in labile solutions 

Ionic concentrations and supersaturations were 

calculated from the measured pH and from the total 

calcium and carbonate concentrations by using massaction and mass-balance equations (Reddy, 1978) (table 

1). The solubility product of calcite and the stability constants of CaC03 and CaHC03 +were taken from Jacobson and Langmuir (1974). Supersaturation dropped 

rapidly in continuously stirred labile solutions; visible 

onset of precipitation occurred within 2-10 seconds of 

mixing (fig. 1). Initial pHs of 10.4 decreased to 9.5 

500~--------------------------------~ 

z 

0 

~ a: 

::::> 

1-

<( 

(/) 

a: 

w 

a.. 

::::> 

(/) 

w 

1-

u 

...J 

<( 

u 

0 STIRRED SOLUTION 

• UNSTIRRED SOLUTION 

• UNSTIRRED SOLUTION 

WITH, 10-3M ADDED 

MAGNESIUM ION 

CONCENTRATION 

0 

TIME, IN MINUTES 

Figure 1. Calcite supersaturation (Oc) versus time for calcium 

carbonate nucleation from stirred and quiescent labile solutions. 

Effect of Mg Ions on CaC03 Nucleation and Crystal Growth in Dilute Aqueous Solutions at 25°C 171 
Table 1. Initial solution concentrations and calcite supersaturation values (Oc) for calcium carbonate nucleation from labilesupersaturated solutions at pH 10.3 

[Total carbonate concentration 5 x w-3 Mat 250C. Leaders (---) 

indicate no added magnesium ion; Oc, supersaturation] 

Initial concentration (moles per 1 iter X 104} 

Experi- Total 

ment ca2+ Mg2+ ca2+ co 2- 3 Caco3o Qc 

13 50 26.83 17.16 21.82 462 

15 54 28.83 16.81 22.47 476 

7 54 29.57 16.16 22.25 471 

8 55 10 31.16 16.42 22.13 468 

9 53 30.25 16.05 22.45 475 

11 53 29.65 16.12 22.30 472 

12 46 10 24.63 20.30 17.54 371 

13 58 33.72 15.63 23.97 508 

1 Experiments 3 and 5 were stirred continuously during precipitation; all other experiments were quiescent after intial mixing. 

during the first 10 minutes of reaction and to 8.6 during 

the first 100 minutes. Long-term solution pH's approached 

8.2. The total calcium concentration dropped from 5 X 

10-3 M to a steady-state value of less than 1 x 10-3 M 

within 5 minutes of mixing and solutions rapidly approached calcite equilibrium. In unstirred experiments, 

the calcium concentration dropped more slowly with time; 

the experiments approached a steady-state calcium carbonate lAP significantly greater than that for calcite 

equilibrium (fig. 1). lAP's during this phase of the 

precipitation reaction may correspond to the steady-state 

values of the initially formed precipitate. 

The initial precipitate had characteristics of a noncrystalline or microcrystalline material. X-ray diffraction 

patterns had only a few broad peaks, and examination 

by optical microscopy showed submicrometer, isotropic 

particles. In many experiments, however, vaterite and 

calcite were identified in the XRD pattern, and under the 

microscope the samples consisted of a mixture of birefringent crystals in a submicrometer-size, optically 

isotropic matrix. By SEM, most of these initial precipitates were seen to consist of uniform, sub micrometer-size, 

spherical particles that displayed no euhedral character 

(fig. 2). Specific surface areas calculated for these initial 

precipitates agreed well with the measured value of 

15 m2/g. 

The formation of calcium carbonate polymorphs 

and their subsequent transformation and incorporation 

of magnesium ion differed between the stirred and unstirred solutions, suggesting that the processes occurring 

during or shortly after calcium carbonate nucleation from 

highly supersaturated solutions were influenced by solution hydrodynamics. For example, both the change in total 

calcium concentration and the nucleation particles 

172 Studies in Diagenesis 

number (the number of particles formed in the supersaturated solution) differed between the stirred and unstirred solutions. Particle number differences were most 

pronounced for calcite rhombs and vaterite spheres. The 

precipitates that formed in stirred solutions did not appear 

to have a noncrystalline component (figs. 2B and 2D), 

presumably because of the rapid transformation of any 

such material that did form. Scanning electron 

micrographs of precipitates collected on filters showed 

that in stirred solutions, containing low concentrations 

of magnesium ion, calcite rhombs and vaterite spheres 

formed in nearly equal numbers. The recrystallization of 

what may have initially been noncryst~line material and 

unstable polymorphs of CaC03 yielded calcite with a 

uniform, twinned-rhombohedral morphology and a narrow size distribution (figs. 2C and 2E). XRD, optical 

microscopy, and SEM demonstrated that for concentrations of dissolved magnesium ion at or less than 1Q-4 M, 

the final crystalline product was calcite. Sufficiently high 

concentrations of magnesium ion (1Q-3 M) caused the 

percentage of vaterite spheres to increase relative to calcite 

rhombs. 

The degree of incorporation of magnesium ion in 

the precipitate varied with the presence or absence of stirring and with the initial magnesium concentration. Initial 

total concentrations of magnesium ion were selected to 

eliminate the formation of magnesium carbonate phases. 

Magnesium-to-calcium ratios thus were low, ranging from 

0.2 to 0.02 in the initial solution. The calcium ion concentration decreased markedly during precipitation, 

whereas the magnesium ion concentration changed only 

slightly. Thus, as the precipitation experiment approached 

equilibrium the magnesium-to-calcium ratio in the solution increased. For an initial unstirred solution containing 1 x 10-3 M total concentration of magnesium ion, 

the precipitate contained about 1 percent magnesium. 

Experiment 12 (table 1) yielded a precipitate containing 

1 mole percent magnesium ion 3 minutes after the onset 

of nucleation. Four subsequent analyses of the precipitate 

at intervals over the next 100 minutes showed that the mole 

percent magnesium decreased to 0. 7 percent. Precipitation from a similar solution that had an initial concentration of magnesium ion of 5 X 10-4 M showed 0.7 

mole percent magnesium in the precipitate 3 minutes after 

nucleation, whereas after 10 minutes the magnesium 

Figure 2 (facing page). Scanning electron micrographs of 

calcium carbonate precipitates from stirred and unstirred solutions. A, unstirred solution, 3 minutes after nucleation. T Mg = 

5 x w-4 M. 8, stirred solution, 5 minutes after nucleation, no 

added magnesium ion. C, as in B but 1 day after nucleation. 

0, stirred solution, 5 minutes after nucleation. T Mg = 1 x w-4 

M. E, as in 0 but 1 day following nucleation. Bar marker is one 

micrometer long. 
Effect of Mg Ions on CaC03 Nucleation and Crystal Growth in Dilute Aqueous Solutions at 25°C 173 
content of the precipitate had decreased slightly to 0.6 

mole percent. 

Continuous stirring following nucleation yielded 

precipitates with somewhat higher magnesium ion contents. Experiment 2, which was equivalent to experiment 

12 except that it was unstirred, produced an initial 

precipitate that contained 2.3 mole percent magnesium 

5 minutes after nucleation. The magnesium content of 

the precipitate decreased to 1.6 mole percent within the 

first 100 minutes of reaction. 

In each nucleation experiment (stirred and unstirred) that was sampled successively during the reaction, 

the magnesium content of the initial precipitate was at 

a maximum. As the precipitate aged, its magnesium content decreased, typically several tenths of 1 percent over 

the course of the experiment. In contrast, the precipitates 

that formed in the sealed-bottle experiments showed 

higher magnesium contents at the termination of the reaction. Inasmuch as copius amounts of calcium carbonate 

precipitated in the first few minutes of the reaction, 

changes in the magnesium ion content of the precipitate 

were probably due in large part to transformation and 

recrystallization processes. 

The variation in the magnesium content of the 

precipitate during the formation of CaC03 may be due 

to a number of factors. The reaction rate is important 

because at high rates of precipitation (at the start of reaction) more magnesium was incorporated in the precipitate than later in the reaction when the reaction rates were 

significantly lower. This observation is in agreement with 

that of Berner (1978), who reported that experiments at 

high supersaturation and high precipitation rates produced 

calcites richer in magnesium. The incorporation of 

magnesium during the precipitation of calcium carbonate 

may have involved polymorphs other than calcite which 

have dissimilar magnesium ion distribution coefficients, 

thus leading to a variation in the distribution coefficient 

with changing polymorph composition. 

Long-term experiments involved the recrystallization 

of the calcium carbonate precipitate in a solution having 

a magnesium-to-calcium ratio significantly higher than 

that at the initiation of the precipitation. Thus, because 

of the higher ratio, magnesium ion may have been incorporated into the precipitate either via a solid state 

exchange of surface magnesium or by the formation of 

a magnesium-rich surface phase. 

Nucleation in metastable solutions 

The nucleation and subsequent crystallization of 

calcium carbonate in stirred, metastable-supersaturated 

solutions (table 2) formed less complex precipitates than 

those that formed from labile-supersaturated solutions. 

Lower initial supersaturations in metastable solutions 

reduced the amount of initial noncrystalline precipitate 

174 Studies in Diagenesis 

Table 2. Calcium carbonate nucleation without and with 

magnesium ion at 25°( 

[Leaders(---) indicate no added magnesium ion; Oc, supersaturation; 

t, it:J.duction interval, in seconds] 

Total initial concentration 

(moles per liter) X 104 

Experi- ca2+ Mg2+ co32- pH nc !.r 

ment 

3 26.64 41.39 8.565 27.8 480 

4 13.32 41.39 8.567 15.5 2,100 

6 33.30 41.39 8.502 29.2 420 

7 13.32 41.39 8. 654 18.6 1,500 

8 26.64 41.39 8.505 24.5 660 

9 33.30 41.38 8.545 32.0 300 

12 46.62 41.38 8.448 33.9 300 

13 6.66 41.38 8.832 14.2 4,500 

16 19.98 41.38 8.647 26.1 720 

M2 40.05 3.214 41.64 8. 585 39.9 300 

M3 26.70 2.142 41.64 8.665 34.2 450 

M4 13.35 1.071 41.64 8.830 25.6 960 

M12 43.74 15.01 41.55 8.419 28.7 1,050 

M14 26.05 25.91 41.64 8.523 22.4 1,740 

M17 52.09 51.82 41.64 8.302 22.6 1,680 

M22 6.34 12.75 41.69 8. 725 13.9 8,640 

M25 50.00 252.09 41.71 8.135 10.7 15,120 

M27 72.67 385.34 41.71 8.025 10.8 15,600 

and led to smaller amounts of unstable polymorphs. 

Direct nucleation (that is, crystal formation from an 

initially supersaturated solution without intervention of 

an antecedent) appeared to be important in the development of the ultimate crystal assemblage. 

The predominate crystalline components of the 

precipitate, over a range of magnesium ion concentrations 

and supersaturations, were vaterite and aragonite. The formation of calcite from the metastable solutions was 

retarded. Some aged solutions, containing low concentrations of magnesium ion, developed calcite as a final 

precipitate. In these experiments, initially formed vaterite 

transformed to calcite during the course of the experiment. 

Precipitates from solutions containing high concentrations 

of magnesium ion were aragonite. 

An abrupt change in solution pH and turbidity was 

used to identify calcium carbonate nucleation. The time 

required to initiate this change after the component solutions were mixed was considered to be the induction interval (t1

). At low supersaturations, changes in the solution 

pH and turbidity during nucleation were less pronounced. 

Induction intervals and the number of particles formed 

during nucleation (particles per milliliter) exhibited little 

systematic variation with changing magnesium ion concentration (fig. 3). 

Both the induction interval and the number of particles per milliliter followed a log-log linear relationship 

with calcite supersaturation; the best-fit line for solutions 

containing magnesium ion lay slightly above that for the 

magnesium-free solution (fig. 3). The presence of magnesium ion in the supersaturated solutions lengthened 

the induction interval and led to a larger number of 
nucleation particles. Nucleation rates, defined as Nlt1

, 

were thus used to identify the metastable region of calcium 

carbonate supersaturation and to quantify the influence 

of magnesium ion on this region (fig. 4). 

As discussed previously, metastable solutions can 

be divided into two groups: (1) those having a very small 

nucleation rate, and (2) those having a measurable nucleation rate, with subsequent crystalline phase formation in 

a specific time interval t r A rate of 50 particles per cubic 

centimeter per second has been chosen as the boundary 

between these two regions. For the experiments shown in 

figure 4, the calcium carbonate metastable region extends 

from Oc = 1 to Oc = 10. Inasmuch as the presence of 

magnesium ion increased both the induction interval and 

the nucleation particle number in roughly the same proportion, magnesium ion did not appear to have altered 

the calcium carbonate nucleation rate appreciably. 

Calcite-seeded crystallization 

The initial conditions for calcite-seeded growth 

experiments are given in table 3. Typical plots of the concentration of calcium in solution and pH against time for 

calcite growth in the presence and absence of magnesium 

ion illustrate that seed crystals induced immediate crystal 

growth. Total calcium concentrations and solution pH's 

demonstrate that magnesium ions did not prevent crystallization, but that they reduced the calcite crystallization 

rate markedly. The role of magnesium ions is most clearly 

shown with the second addition of seed crystal to solutions with the high magnesium-ion concentration. Initially, the concentration of calcium in solution decreased 

only slightly, but if more seeds were added the growth rate 

nearly equaled that of the magnesium-free control reaction. The formation of calcite from supersaturated solutions with a high concentration of magnesium ions was 

thus kinetically inhibited and proceeded only when sufficient reactive surfaces were available. Magnesium exhibited no marked changes in solution concentration during 

the course of the reaction. The crystallization rates 

followed an expression found to hold in earlier work 

(Reddy, 1978): 

dN/dt = ksN2 , (1) 

where N is the amount of calcium carbonate in moles per 

liter at time t (from the start of the reaction by addition 

of seed crystal) in minutes to be precipitated to reach 

(calcite) equilibrium; dN/dt is the calcite crystallization 

reaction rate; k in [(moles per liter) x minutes X milligrams of seed crystal per liter]-1 is the crystal growth-rate 

constant; and s in milligrams of seed per liter is the added 

seed concentration and is proportional to the surface area 

available for growth. 

Analysis of the data for the seeded crystallization 

experiments was facilitated by use of an integrated form 

of equation 1: 

(2) 

where N0 (the initial theoretical crystal yield) is calcite 

(moles per liter) to be precipitated from the stable supersaturated solution at the time of seed inoculation. Linear 

plots of N-1 - N0 - 1 versus time demonstrate the validity 

of equation 2 (Reddy, 1978) for interpreting the 

experimental results. Rate constants for calcite crystallization in solutions with and without magnesium ion are 

listed in table 3. The parabolic dependence of the crystal 

growth rate (the N2 term in equation 1) is consistent with 

a calcite crystallization mechanism involving a ratelimiting surface reaction step at the crystal face (Reddy, 

1978). 

Equation 2 enables a quantitative description to be 

made of the influence of magnesium ion on calcite crystallization kinetics. Plots of N-1 - N0-

1 versus time showed 

a decrease in slope (ks from equation 2) with increasing 

magnesium ion concentration. The influence of magnesium ion on the calcite crystallization rate was compared 

for different experiments by plotting the reduced reactionrate constant, k/ k0 (where k is for solutions containing 

magnesium ion and k0 is for magnesium-free control 

experiments), as a function of the magnesium ion concentration (fig. 5). This plot presents the data shown in 

table 3 as well as two sets of published data. The good 

agreement of the four independent sets of experimental 

results is evidence of the reliability of the seeded-growth 

technique. Rate-constant reduction occurred at magnesium ion concentrations of lQ-5 M; pronounced decrease 

in the rate constant was evident above 2 X 10-4 M. Previous investigations have shown that calcite growth-rate 

reduction follows a Langmuir-type adsorption model 

(Reddy, 1978). If this treatment is appropriate, the 

crystallization rate constant in the presence and absence 

of magnesium ion is 

where k1 and k2 are the adsorption and desorption rate 

constants, respectively, for magnesium ion at the growth 

sites on the crystal surface, and 2+ is the total concentration of dissolved magnesium ion. A linear relationship between kof(k0 - k) and the reciprocal of the 

total magnesium ion concentration (fig. 6) demonstrates 

that the Langmuir adsorption isotherm model described 

the inhibition satisfactorily. This agreement over a wide 

range of solid to solution ratios implies that inhibition 

of calcite growth was due to the adsorption of magnesium 

ion at growth sites on the crystal surface. 

Effect of Mg Ions on CaC03 Nucleation and Crystal Growth in Dilute Aqueous Solutions at 25°C 175 
WITHOUT MAGNESIUM ION 

10 ~------~----~--~~~~~~ 

z 

0 

f= 

:::> 

....J 

0 

Cl) 

u.. 

0 

a: 

w 

1-

:::J 

:::J 

00 ....J 106 100 ~ 

a: 

w 

a.. 

~ Cl) 

w 

1-

z :::> 

0 z f= ~ <( 

....J 

w ~ 

u 

:::> .:=: z 0 w (!) ~ 

~ f= a: 

:::> z c 0 

c f= w u 

~ :::> 

c a: 

0 ~ u.. 0 Cl) 

w 

....J 105 u 10 

f= 

a: 

<( 

a.. 

u.. 

0 

a: 

w 

Cll 

~ 

:::> 

z 

104 L--4~L-~----L---L-~_L_L~~ 

10 100 100 

CALCITE SUPERSATURATION (!lc) 

Figure 3 (above and facing page). Number of nucleation particles and induction interval versus calcite supersaturation for precipitation without and with magensium ions. 

Approach to calcite equilibrium 

Whereas an adsorption mechanism is compatible 

with calcite growth-rate inhibition by magnesium ion during most of the reaction, equation 1 does not describe the 

crystallization process near equilibrium. In the absence 

of magnesium ion, and at an intermediate seed-crystal 

176 Studies in Diagenesis 

concentration, calcite equilibrium was attained within 100 

minutes of reaction. In the presence of magnesium ion, 

even with significantly higher seed concentration, the 

crystallization did not reach calcite equilibrium (fig. 7). 

In several experiments, the solution composition after 

extended reaction time changed only slightly and appeared to approach steady-state calcite supersaturations 
WITH MAGNESIUM ION 

0 

0 

z 0 0 

i= 

::> 

...J 

0 

(/) 

u.. 

0 

a: 

w 

I-

:::i 

:::i 106 0 100 ...J 

~ 

a: 0 w (/) 

a... w 

I-

~ ::> 

z 

0 

z ~ 

i= ~ <( 

w 2 ...J 

u 

::> w 

z ~ 

(!) i= z z a: 0 ::> i= Cl u Cl ::> 

w 0 

~ ~ a: 

0 

u.. 

(/) 

w 105 10 ...J 

u 

i= 

a: 

<( 

a... 

u.. 

0 

a: 

w 

IXl 

~ 

::> 

z 

10 ~------~----~~--~~~~~ 1L_ ______ _L ____ L_ __ L_~~~~~ 

10 100 10 100 

CALCITE SUPERSATURATION (!"lc) 

significantly greater than equilibrium. This steady-state 

calcite supersaturation value was related to both the initial 

magnesium ion concentration and the seed-crystal concentration (fig. 8). Changes in the concentration of 

magnesium ion during calcite growth were not statistically 

significant over the initial phase of the reaction. Longterm equilibration and successive seed additions, however, 

led to the loss of magnesium ion from solution. 

Steady-state supersaturation, the saturation value 

approached in solutions at long reaction times, was clearly 

related to magnesium-ion concentration (fig. 7). At higher 

magnesium-ion concentrations, steady-state supersaturation was higher. With a second addition of seeds, new 

steady-state supersaturations were reached which were 

closer to calcite equilibrium. Steady-state supersaturation 

values for several replicate experiments at intermediate 

seed concentrations were in good agreement. Crystallization experiments, at similar seed-crystal concentrations, 

exhibited steady-state supersaturation, which was a function of the log magnesium ion concentration (fig. 8). This 

Effect of Mg Ions on CaC03 Nucleation and Crystal Growth in Dilute Aqueous Solutions at 25°C 177 
Table 3. Experimental conditions, seed crystal concentrations (s), and rate constants 

(k) for calcite growth in the presence and absence of magnesium ion 

[sin milligrams per liter; rate constant (k) in (moles per liter)· (minutes)· (milligrams of seed per liter)] 

Initial concentration of ions in 

solution 

Experiment ca2+ 

SC-1 2.258 

SC-2 2.238 

SM-1 2.315 

SM-2 2.172 

SM-3 2.078 

SM-4 2.243 

SM-5 2.197 

36 1.862 

38 2.297 

40 2.165 

0 WITHOUT MAGNESIUM 

• WITH MAGNESIUM 

• 

~ 

• • 

• 

0 

(moles eer 1 iter} 

Total 

Mg2+ co 2- 3 

0 40.97 

0 40.97 

10.979 40.97 

10.979 40.97 

5.227 38.62 

2.233 40.97 

5.723 40.97 

0 41.86 

13.08 42.30 

6.066 39.86 

0. 

• 

4 8 1 2 16 20 24 28 32 36 40 

CALCITE SUPERSATURATION (OcJ 

Figure 4. Calcium carbonate nucleation rate versus calcite 

supersaturation in solutions without and with magnesium 

ions. 

178 Studies in Diagenesis 

X 104 

ca 2+ co 2- 3 pH ~ k 

1.874 2.307 9.01 248 1.40 

1.848 2.357 9.02 24.8 1.42 

1.924 2.393 9.00 1247 0.0124 

1.837 2.394 9.00 1247 0.0084 

1. 749 2.279 9.02 2327 0.020 

1.875 2.288 9.00 249 0.162 

1.828 2.382 9.01 248 0.037 

1.679 1.035 8.638 249 1.49 

2.031 1.385 8. 736 1370 0.0012 

1.938 1.262 8.736 2410 0.025 

situation may have arisen because of irreversible adsorption of magnesium ion on the calcite surface forming 

localized areas of a magnesium-rich phase. This explanation is consistent with the suggestion of Fajans (1931) that 

an ionic species is strongly adsorbed on a crystalline solid 

if that species forms an insoluble compound with opposite 

charged ions of the crystal structure. A quantitative 

description of this process is difficult to develop on a 

theoretical basis, because the incorporation of magnesium 

ion during growth led to changing surface properties with 

time. The incorporation of magnesium ion (or subsequent 

exclusion from a growth phase) may have been a function of growth rate. During rapid crystallization, ions may 

be incorporated in a crystal by a process that does not 

correspond to simple adsorption (Adamson, 1960). Such 

adsorption would be accompanied by surface exchange. 

In the present experiments, magnesium ions may have 

been continually taken up or released as each growth layer 

came into exchange equilibrium with the substrate crystal 

and the solution. The final steady-state supersaturation 

value decreased linearly (fig. 8) with increasing log of the 

seed concentration. 

Other authors have emphasized the kinetic restrictions associated with the heterogenous reactions of 

aqueous calcium carbonate in the presence of magnesium 

ion. For example, Thorstenson and Plummer (1977, 1978) 

termed the steady-state solution composition in such 

systems ''stochiometric saturation.'' The solid phase in 

contact with this solution remains invariant, even though 

it may be part of a continuous compositional series. They 

proposed that the magnesium content of solid calcium 

carbonate should be governed primarily by the activity 

product of calcium and carbonate ions in solution. From 

the thermodynamic requirements of stoichiometric 

saturation, the solution magnesium-to-calcium ratio 

should not be a major regulating factor in determining 
a: 0 08 0~ . u..-~o: 

~--~ 

zz 

<{0 

t;i= z <{ 0.6 

o!:::! u:::l 

w<t 

t-1- <{(1) 

a:>- 0.4 

Cla: 

wu 

~~ ClU 

~ ;£ 0.2 u 

2 7 8 9 10 11 12 13 

MAGNESIUM ION CONCENTRATION, IN MOLES PER LITER x1 04 

Figure 5. Calcite crystallization reduced-rate constant (k/k0 ) as a function of the concentration of magnesium ion in solution. 

1.2.---------------~----------------~--------------~--------------~--------------~ 

0.1 0.2 0.3 0.4 0.5 

MAGNESIUM ION CONCENTRATION, IN MOLES PER LITER x1 04 

Figure 6. Langmuir isotherm plot of k0/(k0 - k) versus the reciprocal of the magnesium ion concentration. 

the composition of the solid phase. Results obtained in 

the seeded-crystallization experiments are consistent with 

the hypothesis of Thorstenson and Plummer, because 

calcite containing little or no magnesium ion formed over 

the range of experimental conditions used. 

Berner (1978), in discussing the magnesium content 

of calcite precipitated from seawater, emphasized that the 

composition of the solid phase is controlled by kinetic 

factors, and that it does not represent exchange equilibrium of calcium and magnesium ions between seawater 

and calcite. Morse and others (1979) observed a magnesium calcite growth onto a calcite crystal suspended in 

seawater. This incorporation of magnesium in the growth 

phase in the seawater experiments differs from the dilute 

solution results presented in this paper. The differences 

may be due in part to the higher ionic strength of the 

seawater medium or to naturally occurring calcite-growth 

inhibitors present in seawater, which could have facilitated 

the formation of magnesium calcite. Additional work on 

the comparison of laboratory and field calcium carbonate 

formation kinetics seems needed. 

CONCLUSIONS 

Diagenetic transformations of sedimentary carbonate minerals arise from numerous processes including 

calcium carbonate nucleation and crystal growth. Results 

obtained in this study demonstrate that calcium carbonate 

mineral formation was characterized reliably and reproducibly over a range of magnesium ion concentrations. 

Both magnesium ion distribution and calcium carbonate 

Effect of Mg Ions on CaC03 Nucleation and Crystal Growth in Dilute Aqueous Solutions at 25°C 179 
? 3 z 

0 

~ a: 

::> 

I-

<( 

(/) 

a: 

w 

a.. 

::> 

(/) 

w 

1-

g 2 <( 

u 

0 

ADDED MAGNESIUM 

ION CONCENTRATION 

---------------~CI-- 2nd seed 

addition 

13X1 0-4M 

1L-----~-------L------~------~L-------~-------L------~--------~------~------~ 

0 10 20 30 40 50 

TIME,IN MINUTES 

60 70 80 90 100 

Figure 7. Calcite supersaturation versus time for calcite-seeded crystallization without and with magnesium ions. 

formation exhibited systematic variations over a wide 

range of calcium carbonate formation rates and 

supersaturations. 

Solution supersaturation, hydrodynamics, magnesium ion concentration, and reaction path were involved 

in regulating the course of calcium carbonate formation. 

Supersaturation level was of paramount importance in the 

experiments examined here. At high supersaturations 

nucleation proceeded rapidly, yielding an unstable and 

reactive initial precipitate, which rapidly transformed to 

more stable carbonate minerals. Solution hydrodynamics 

and the concentration of magnesium ion had a secondary influence on the nucleation reaction at high 

supersaturation. 

At intermediate supersaturations (10-40) the 

influence of supersaturation on the precipitation reaction 

was also important: nucleation particle number and 

induction interval followed a log-log linear relationship 

with supersaturation. As observed at high supersaturation, magnesium-ion concentration level had a secondary 

influence on the rate and course of reaction. High 

magnesium-ion concentrations regulated the polymorphic 

composition of the final CaC03 precipitate. 

For calcite-seeded growth from metastable solutions (supersaturations of less than 10), nucleation processes were unimportant. The growth rate followed a 

second-order rate equation, consistent with a surface 

180 Studies in Diagenesis 

reaction rate-determining step. Magnesium ion reduced 

calcite formation under these conditions by blocking 

growth sites on the seed crystal surface. Variation of the 

reaction rate constant with magnesium ion concentration is consistent with a Langmuir adsorption isotherm 

model. Because many sediment reactions occur in metastable solutions, the quantitative description of the 

influence of magnesium ion on the calcite crystallization rate may be useful in interpreting field observations. Additional support for a kinetic mechanism for 

the retardation of the growth rate of calcite by magnesium ion is that the steady-state supersaturation, 

reached at the apparent end of the crystallization experiment, was a function of the experimental solution-tosolid ratio. The formation of a magnesium-containing 

surface phase did not appear to be associated with the 

mechanism of magnesium ion inhibition of calcite crystal 

growth. 

Bathurst (1983) pointed out several puzzling aspects 

of carbonate cement mineralogy which may be clarified 

by a fuller understanding of the chemical processes 

involved in calcite crystal growth. Cement formation and 

dissolution reactions are regulated in large measure by 

the solution concentration of magnesium ion; however, 

supersaturation levels and the presence of a reactive 

growth surface can be more influential than magnesium 

ion concentration in some situations. 
2.5X104 

B 

ADDED MAGNESIUM ION 

2X104 CONCENTRATION. IN 

MOLES PER LITER X1 0 4 

15 Q12 

A 

05 

10 Seed concentration = 104 

9 250 mg/1 \ .62 

8 

b 7 

a: x 6 w 

1- a: 5 • ::J w 

1- a: ::J w 4 a.. a: (/) w 

a.. ~ 

(/) 3 <{ 

w a: ...J (.!) 

0 ::J ~ ::::! 

~ 2 ~ 

z ~ 

0 z i= 0 <{ i= a: <{ 1-

z a: w 1-

u 1.0 z z 0.9 w 

u 0 0.8 z u 0 z 0.7 u Q ...J 

0.6 <{ 

~ 1-

::J (/) 

0.5 >- 103 

Ci5 a: w u z 0.4 I (.!) 0 <{ w 

~ w 

(/) 

0 0.3 w 

0 .6 0 0 

<{ 

0.2 

0.1L---~----~----~----~----~----~ 

1 2 3 4 

CALCITE SUPERSATURATION (nc) 

Figure 8. Steady-state calcite supersaturation versus (A) initial total magnesium-ion concentration and (8) initial seed-crystal 

concentration. 

As an example of the application of laboratory 

results to a field situation, consider the formation of lowmagnesian calcite from marine sediments that have been 

uplifted and bathed in fresh water. Mackenzie and others 

(1983) summarized the results of numerous investigations 

which showed magnesium loss and fabric retention during freshwater diagenesis of high-magnesian calcites. This 

reaction depends strongly on solution supersaturation 

level and the presence of crystallization inhibitors. 

Although the high-magnesian calcites are less stable than 

low-magnesian calcite, the transformation from the 

unstable phase to the stable phase was not observed until 

the sediment pore water changed from a marine to a 

freshwater composition. 

It is unclear whether the rate-limiting step in the 

transformation is the dissolution or the crystallization 

process. If the crystallization reaction is rate-limiting and 

calcite supersaturation values are fixed by the solubility 

of the high-magnesian calcite, calcite will not form until 

the solution concentration of magnesium ion decreases to 

a value consistent with an appreciable calcite growth rate. 

If a minimum necessary rate constant is assumed to be 

1 percent of the value in the absence of magnesium ion, 

from the data in table 3, the magnesium concentration 

associated with this calcite growth-rate constant is about 

1-2 millimoles of magnesium ion per liter. To reduce a 

seawater magnesium-ion concentration to this value 

requires a 50-fold dilution, a value consistent with circulation of a low-magnesium fresh water in the uplifted sediment. Conversion rates would also be regulated by 

hydrodynamic factors, including the pore-water flushing 

rate. 

The observed microstructure retention during the 

transformation implies low-solution supersaturations and 

Effect of Mg Ions on CaC03 Nucleation and Crystal Growth in Dilute Aqueous Solutions at 25°C 181 
low growth rates. Release of magnesium ion to the pore 

water during recrystallization will limit supersaturation 

and growth rates at the microstructure level. 

The important point concerning the application of 

laboratory studies to carbonate-sediment diagenesis is that 

information about the mechanism of magnesium ion 

regulation of carbonate-sediment diagenetic reactions is 

available. Laboratory studies, such as those presented in 

this paper, give reliable, quantitative estimates of the 

influence of magnesium ion on mineralogical changes that 

occur in sediments. Additional work is needed to extend 

these results to actual sedimentary environments. 

REFERENCES CITED 

Adamson, A. W., 1960, Physical chemistry of surfaces: New 

York, Wiley-Interscience Publishers, 629 p. 

Bathurst, R. G. C., 1983, Early diagenesis of carbonate 

sediments, in Parker, A., and Sellwood, B. W., eds., Sediment diagenesis: Boston, Mass., D. Reidel Publishing Company, p. 349-377. 

Berner, R. A., 1978, Discussion-Equilibrium, kinetics, and the 

precipitation of magnesium calcite from seawater: American 

Journal of Science, v. 278, p. 1475-1477. 

Fajans, Kasimir, 1931, Radio elements and isotopes; chemical 

forces and optical properties of substances: New York, 

McGraw-Hill, 125 p. 

182 Studies in Diagenesis 

Jacobson, R. L., and Langmuir, D., 1974, Dissociation constants 

of calcite and CaHCOt from 0 to 50°C: Geochimica et 

Cosmochimica Acta, v. 38, p. 301-318. 

Mackenzie, F. T., Bischoff, W. D., Bishop, F. C., Loijens, M., 

Schoonmaker, J., and Wollast, R., 1983, Magnesian 

calcites-Low-temperature occurrence, solubility, and solidsolution behavior, in Reeder, R. J., ed., CarbonatesMineralogy and chemistry: Washington, D.C., Mineralogical 

Society of America Reviews in Mineralogy, v. 11, p. 97-144. 

Morse, J. W., Mucci, A., Walter, L. M., and Kaminsky, M. S., 

1979, Magnesium interaction with the surface of calcite in 

seawater: Science, v. 205, p. 904-905. 

Reddy, M. M., 1978, Kinetic inhibition of calcium carbonate 

formation by wastewater constituents, in Rubin, A. J., ed., 

The chemistry of wastewater technology: Ann Arbor, 

Michigan, Ann Arbor Science Publishers, Inc., p. 31-58. 

Reddy, M. M., and Nancollas, G. H., 1976, The crystallization 

of calcium carbonate-IV. The effect of magnesium, strontium and sulfate ions: Journal of Crystal Growth, v. 35, 

p. 33-38. 

Thorstenson, D. C., and Plummer, L. N., 1977, Equilibrium 

criteria for two-component solids reacting with fixed 

composition in an aqueous phase-Example, the 

magnesium calcites: American Journal of Science, v. 277, 

p. 1203-1223. 

___ 1978, Discussion:-Equilibrium criteria for twocomponent solids reacting with fixed composition in an 

aqueous phase-Example, the magnesium calcites: 

American Journal of Science, v. 278, p. 1478-1488. 
Authigenic Potassium Feldspar in Ribbon Rocks of the 

Cambrian-Conococheague Limestone, Western 

Maryland 

By PAUL P. HEARN, Jr., U.S. Geological Survey, MS 957, National Center, 

Reston, Virginia 22092, 

ROY C. LINDHOLM, George Washington University, 2021 G Street N.W., 

Washington, D.C. 20002, and 

JOHN F. SUTTER, U.S. Geological Survey, MS 981, National Center, 

Reston, Virginia 22092 

CONTENTS 

Abstract 185 

Introduction 185 

Methods 186 

Results and discussion 186 

Petrography of ribbon rocks 186 

Accessory minerals 188 

Microprobe analyses 189 

Chemistry of insoluble residues 190 

X-ray diffraction mineralogy 190 

Evidence of interstitial brines 190 

Thermochemical and mass-balance considerations 192 

Timing of feldspathization 193 

Conclusions 195 

References cited 195 

FIGURES 

1. Map of study area showing locations of outcrops where samples were 

collected 185 

2-3. Photographs: 

2. Ribbon rock in outcrop, showing dolostone and limestone 

interbeds 186 

3. Slabbed sample of ribbon rock, showing detail of individual 

dolostone and limestone interbeds 186 

4. Photomicrograph of interface between individual fine-grained 

dolostone and coarser-grained limestone interbeds 187 

5-8. Scanning electron micrographs: 

5. Adjacent dolostone and limestone interbeds, showing differences 

in grain size and microporosity of dolostone 187 

6. Adjacent dolostone and limestone interbeds taken with backscattered electron detector 188 

7. Adjacent dolostone and limestone interbeds showing electronexcited cathodoluminescence of detrital phases 188 

8. Detrital quartz and potassium feldspar, showing relative luminescent intensity 189 

9. Diagram showing analyses of ethylenediamine tetraacetic acid-insoluble 

residues of individual limestone and dolostone interbeds 191 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 183 
10. X-ray diffractograms of ethylenediamine tetraacetic acid-insoluble 

residue, showing characteristic reflections of fine-grained potassium 

feldspar in dolostone interbeds, sand-size potassium feldspar with 

overgrowths in limestone interbeds, and detrital cores of potassium 

feldspar grains from limestone interbeds following removal of 

overgrowths by dissolution in dilite hydrofluoric acid 192 

11. Scanning electron micrograph of halite daughter crystal in fluid inclusion cavity in calcite breached during polishing 193 

12. Plot of 40Ar/39Ar age spectra for detrital potassium feldspar with 

overgrowths and detrital grains whose overgrowths had been removed 

by dissolution in 5 percent hydrofluoric acid, generalized age spectra 

for each, and synthetic age spectra for overgrowths alone 194 

TABLES 

1. Microprobe analyses of potassium feldspar 189~ 

2. Representative chlorine and bromine analyses of dolostone and 

limestone interbeds 192 

184 Studies in Diagenesis 
Abstract 

Thin alternating beds of coarse-grained limestone and finegrained dolostone, commonly known as ribbon rock, are a 

characteristic lithology in the Cambrian Conococheague 

Limestone of western Maryland as well as in other lower 

Paleozoic carbonate rocks of the central Appalachian basin. 

The depositional environment of these rocks has been interpreted as a mixed sand-mud tidal flat. Abundant potassium 

feldspar has been identified in both the limestone and dolostone 

interbeds; on the basis of cathodoluminescence properties, high 

potassium/rubidium and potassium/barium ratios, and chemical 

purity, much of this feldspar is believed to be authigenic. 

Authigenic potassium feldspar in the limestones occurs 

as overgrowths on sand-size grains of detrital microcline and 

orthoclase; the total potassium feldspar content of this lithology 

averages 9 weight percent. In contrast, potassium feldspar in 

the dolostones occurs as a matrix of intergrown clay and siltsize grains, and comprises as much as 30 weight percent of these 

interbeds. This fine-grained feldspar appears to be the product 

of the metasomatism of preexisting clay minerals. 

Simple mass-balance calculations indicate that the authigenic potassium feldspar could not have formed isochemically, 

but had to involve the movement of multiple pore volumes of 

potassium-bearing fluids through these rocks. Bulk-rock chlorine/bromine ratios and the presence of halite daughter crystals 

in some fluid inclusions suggest that these fluids were brines 

with salinities considerably greater than seawater. 40Arf39Ar age 

spectrum analyses of overgrowths on detrital grains indicate that 

the authigenic potassium feldspar is Late Pennsylvanian-Early 

Permian in age. These results strongly suggest that the authigenic 

potassium feldspar in the Conococheage Limestone resulted 

from the mobilization and migration of basinal brines during 

the Alleghanian orogeny. 

INTRODUCTION 

Thin, interbedded dololutites and pelleted limestones 

(ribbon rocks) are a common subfacies in lower Paleozoic 

carbonate rocks throughout the central Appalachian basin. 

Their presence, together with stromatolites, mudcracks, 

flat-pebble conglomerates, and other diagnostic sedimentary structures, indicates deposition in a shallow-platform 

lagoon-tidal flat complex, which is believed to have 

paralleled the Middle Cambrian to Middle Ordovician 

mainland for several thousand miles (Reinhardt and Hardie, 1976). Ribbon rocks were apparently formed by sedimentation in a zone transitional between supratidal mudflats and subtidal to intertidal sand flats. Demicco and 

Mitchell (1982) described the sedimentary processes 

recorded in these rocks as a wave and current "turn-on, 

turn-off' system, where alternating high- and low-energy 

conditions produced layers of rippled sand draped by 

laminae of fine silt and mud. The extensive occurrence of 

this lithology is attributable to both the low slope and the 

rapid subsidence of the broad carbonate shelf, which 

allowed short-term changes in current and wave energy to 

be recorded and preserved in sediments over a large area. 

Numerous occurrences of authigenic potassium 

feldspar (K-feldspar) in lower Paleozoic rocks of the 

Appalachian basin have been described during the past 

50 years (see, for example, Daly, 1912; Tester and Atwater, 

1934; Berg, 1954; Woodard, 1972). Much of this work was 

summarized by Buyce and Friedman (1975), who 

recognized authigenic K-feldspar as a characteristic component of the Cambrian-Ordovician rocks that formed 

on the east-facing shelf of the proto-Atlantic Ocean. The 

common association of authigenic K-feldspar with 

dolomitized sequences of tidal-flat carbonates led some 

workers to suggest that the feldspar was the product of 

early diagenesis (Mazzullo, 1976). The results of our work, 

however, suggest that the authigenic K-feldspar in the 

Conococheague Limestone formed some 200 million years 

after deposition due to the migration of tectonically 

mobilized brines. 

This paper describes the results of petrographic and 

geochemical analyses of Conococheague ribbon rocks and 

suggests a scenario for the formation of authigenic Kfeldspar. Although limited in scale, the study should serve 

as a basis for reconstructing the diagenetic history of 

similar carbonate units elsewhere in the Appalachian 

basin. 

The study area is located about 15 km west of 

Hagerstown, Maryland, south of the town of Clear Spring 

(fig. 1). Samples were collected from outcrops (fig. 2) 

exposed along the tracks of the Western Maryland 

Railroad. 

Area of 

~ MARYlAND 

2 KILOMETERS 

2 MILES 

Figure 1. Study area, showing location of outcrop (X) where 

samples were collected. 

Authigenic Potassium Feldspar in Ribbon Rocks, Conococheague Limestone, Maryland 185 
Figure 2. Ribbon rock in outcrop, showing dolostone (D) and 

limestone (L) interbeds. Note dime in upper right-hand corner 

for scale. 

<ISS 

Figure 3. Slabbed sample of ribbon rock, showing detail of 

individual dolostone (0) and limestone (L) interbeds. 

186 Studies in Diagenesis 

METHODS 

A single oriented sample (fig. 3) about 20 em invertical thickness was chiseled from the outcrop shown in 

figure 2. Uncovered, oil-polished thin sections oriented 

normal to the bedding plane were prepared covering the 

entire vertical sequence in the sample. Subsamples of 

individual limestone and dolostone interbeds were separated with a rock saw. The noncarbonate fraction of 

samples was separated by boiling crushed samples in a 

solution of trisodium ethylenediamine tetraacetic acid 

(EDTA), as described by Glover (1961); the weight percent insoluble residue was determined gravimetrically. 

X-ray diffraction analysis of bulk samples and 

K-feldspar separates was performed on a Diana XRD-5 

X-ray diffractometer 1

, using copper radiation and a 

graphite monochrometer. The potassium content of the 

insoluble residues was determined by wavelength-dispersive X-ray fluorescence (XRF) on a Diana model 8600 

automated XRF spectrometer; rubidium and barium were 

determined by energy-dispersive XRF on a Kevex model 

0700 automated XRF spectrometer. Polished thin sections 

were examined with an ETEC Autoscan scanning electron microscope (SEM) equipped with an EDAX 9100 

X-ray analyzer; quantitative analyses of major elements 

of feldspars were made using an ARL SEMQ electron 

microprobe. 

Age-estimates for authigenic overgrowths in pure 

mineral separates of K-feldspar were _made using the 40 Ar/39 Ar incremental-heating technique (Merrihue and 

Turner, 1966; Dalrymple and Lanphere, 1971; Lanphere 

and Dalrymple, 1974). This technique produces a series 

of dates, or age spectra, by releasing argon gas from a 

sample in steps of increasing temperatures. In contrast to 

the conventional potassium-argon method, the incremental-heating technique can yield information about 

argon loss due to thermal events or argon gained by the 

diagenetic addition of potassium during the history of the 

sample. 

RESULTS AND DISCUSSION 

Petrography of ribbon rocks 

The contrasting textures of the dolostone and limestone 

interbeds are shown in the photomicrograph in figure 4. 

The transition upward from limestone to dolostone is 

usually gradual, reflecting a slow decrease in current and 

wave energy. In contrast, the transition upward from 

dolostone to limestone is sharp, indicating an abrupt 

increase in energy and the partial erosion of the underlying mud. 

1

Use of trade and company names is for descriptive purposes only 

and does not imply endorsement by the U.S. Geological Survey. 
Dolostone layers range in thickness from 1 mm to 

3 em. A typical interbed contains about 50 percent 

dolomite and 25 percent K-feldspar, the remainder being 

quartz, calcite, and minor pyrite. Oxidation of pyrite 

causes the dolostone to weather to a grayish orange, but 

Figure 4. Photomicrograph of interface between individual 

fine-grained dolostone (0) and coarser pelleted limestone (L) 

interbeds. 

freshly broken surfaces are medium-dark gray. The 

dolomite is finely crystalline; rhomb sizes range from 20 

to 50 J-tm. Quartz grains are about the same size. Kfeldspar appears in thin section as dark-brown, clay- and 

silt-size grains which form the matrix of the dolostones. 

Limestone layers range in thickness from 2 to 5 mm 

and are finely laminated. Both weathered and fresh surfaces are dark gray. In the thicker layers, the limestone 

is a medium-grained calcarenite cemented by sparry 

calcite. The allochems are mostly intraclasts, but pellets, 

ooids, and biogenic fragments are also common. In the 

thinner layers the limestone is a micaceous, medium- to 

coarse-grained calcilutite cemented by sparry calcite; the 

allochems are dominantly pellets. K-feldspar occurs as 

discrete sand-size grains, most of which have overgrowths. Subequal amounts of sand-size quartz are also 

present. 

The SEM micrograph shown in figure 5, taken with 

a secondary electron detector, reveals several textural 

features. The difference in grain size between K-feldspars 

in the limestone and dolostone layers is clear. Differences 

in polishing hardness have caused feldspar, quartz, and 

accessory mineral grains in the limestone to stand out in 

Figure 5. Scanning electron micrograph (secondary electron excitation) of adjacent 

dolostone and limestone interbeds showing difference in grain size and microporosity 

of dolostone. 

Authigenic Potassium Feldspar in Ribbon Rocks, Conococheague Limestone, Maryland 187 
Figure 6. Scanning electron micrograph of adjacent dolostone 

and limestone interbeds taken with backscattered electron 

detector. Grain brightness is proportional to average atomic 

weight. 

positive relief. One notable feature of the dolostone layers 

is the well-developed microporosity, which is not readily 

apparent with a standard petrographic microscope. Pore 

size ranges from less than 1 I-'m to 40 I-'m; most pores are 

less than 10 I-'m across. 

The SEM micrograph shown in figure 6 was taken 

with a backscattered electron detector, which reveals differences in the average atomic weight of the sample. 

Zones of higher average atomic weight are indicated by 

lighter shades of gray. In combination with point analyses 

by energy-dispersive X-ray fluorescence (EDX), backscattered electron imagery reveals the distribution of separate 

mineral phases which are indistiguishable with a secondary electron detector. K-feldspar, quartz, dolomite, 

calcite, and pyrite are easily identified, their gray level 

increasing in the following order: quartz < dolomite 

< K-feldspar <calcite< pyrite. In contrast to the larger 

K-feldspar grains in the limestone, the fine-grained Kfeldspar in the dolostone clearly forms a porous matrix 

which surrounds larger grains of dolomite, quartz, and 

pyrite. 

Additional information was obtained from SEM 

analysis using electron-excited cathodoluminescence 

(CL). Detrital or igneous K-feldspar strongly luminesces 

when exposed to ionizing radiation, whereas authigenic 

K-feldspar shows no appreciable luminescence. Kastner 

(1971) suggested that certain trace elements or structural 

188 Studies in Diagenesis 

Figure 7. Scanning electron micrograph of adjacent dolostone 

and limestone interbeds showing electron-excited 

cathodoluminescence of detrital phases. 

defects in igneous K-feldspars act as activators, causing 

them to luminesce under an electron beam. 

Luminescing detrital cores and nonluminescing 

authigenic overgrowths are readily apparent in the 

limestone interbeds in CL images (fig. 7). Differences in 

luminescent intensity and the apparent lack of a detrital 

core in some grains is due to the relatively shallow ( rv 1 

J.tm) depth from which cathodoluminescence can be 

detected. K-feldspar in the dolostone is also a mixture 

of luminescing and nonluminescing grains. In contrast 

to the limestone interbeds, however, the fine-grained, 

matrix-forming K-feldspar shows no evidence of having 

precipitated around preexisting detrital grains; rather it 

appears to be an alteration product of clay minerals. 

Nonluminescing overgrowths on quartz grains were 

observed in both the limestone and dolostone interbeds 

(fig. 8), but they were less well developed and present on 

fewer grains than were overgrowths of K-feldspar. 

Accessory minerals 

Accessory minerals were identified on the basis of 

chemical composition and morphology by SEM/EDX. 

The accessory suite includes muscovite, zircon, monazite, 

titanium oxide, apatite, pyrite, and siderite. Zircon, 
Figure 8. Scanning electron micrograph (cathodoluminescence detector) showing relative luminescent intensity of 

detrital quartz (Q) and potassium feldspar (KF). 

titanium oxide, monazite, and apatite were found only 

in limestone interbeds; pyrite and siderite were found in 

both lithologies but were more abundant in the 

dolostones. Pyrite and siderite are almost certainly 

secondary and indicate reducing conditions. These phases 

commonly coexist with organic matter and may reflect 

the former presence of stromatolitic algae. Apatite and 

titanium oxide could be either detrital or secondary. 

Zircon and monazite, however, are most probably 

detrital, and, together with detrital K-feldspar, quartz, 

and muscovite, they suggest that the siliciclastic material in these rocks was derived from the craton to the 

west. 

Microprobe analyses 

K-feldspars in both limestone and dolostone 

interbeds were analyzed in polished and carbon-coated 

thin sections. Igneous K-feldspar was distinguished from 

authigenic K-feldspar on the basis of luminescence or 

nonluminescence under the electron beam. The average 

compositions of authigenic and igneous K-feldspar from 

each lithology are given on table 1. 

Table 1. Microprobe analyses of potassium feldspars (averages based on analyses of 10 grains) 

[Leaders ( ---) indicate absent or undetermined] 

Oxide (in percent) 

Type of Si02 Al 2o3 K20 Na 2o CaO FeO BaO Total potassium feldspar 

Det rita 1 cores, in 

limestones------------ 63.93 18.55 15.64 0.61 0.65 0.01 0.39 99.78 

Authigenic overgrowths, on 

potassium feldspars 

in limestones--------- 64.15 18.82 15.89 .14 .81 .02 .10 99.74 

Detrita 1, in 

dolostones------------ 65.20 18.62 15.04 .48 .57 .44 100.35 

Authigenic, i n 

dolostones------------ 64.15 18.14 14.96 .19 .55 --- .21 98.20 

Average composition of 5 authigenic 

microclines 

reported by 

Kastner ( 1971) -------- 64.84 18.67 16.96 .08 .01 --- 100.56 

Theoretical 

composition----------- 64.76 18.32 16.92 --- 100 

Authigenic Potassium Feldspar in Ribbon Rocks, Conococheague Limestone, Maryland 189 
Although the chemical composition of authigenic 

K-feldspar in the Conococheague Limestone is closer to 

the theoretical end-member composition than most 

igneous K-feldspars (Deer and others, 1967), authigenic 

K-feldspar in these rocks still contains significant amounts 

of impurities relative to the pure specimens analyzed by 

Kastner (1971). The low potassium values are probably 

due to the presence of small amounts of albite and 

anorthite and the substitution of barium for potassium. 

This explanation is consistent with the findings of Ali and 

Turner (1982), who reported an inverse relationship between potassium and barium-sodium contents in 

microprobe traverses across K-feldspar overgrowths. The 

low totals obtained for authigenic K-feldspar in the 

dolostones may possibly be due to the presence of organic 

matter or structural water. The presence of remnant 

organic matter in the dolostones is quite possible, if not 

likely, based on the algal stromatolites present in nearby 

strata and the reducing conditions suggested by the occurrence of pyrite and siderite. The presence of water as both 

finely disseminated fluid inclusions and as structurally 

bound H30+ is plausible in light of recent infrared 

studies of igneous feldspars (Solomon and Rossman, 

1982), where values were reported for fluid-inclusion and 

structurally bound water as high as 1.0 and 0.80 percent, 

respectively. 

Microprobe results reveal no significant differences 

between authigenic K-feldspar in limestone and dolostone 

interbeds. Authigenic K-feldspar in both lithologies are 

noticeably depleted in sodium and barium relative to 

igneous K-feldspar. This depletion would be expected in 

the typically purer authigenic phases. Calcium values, 

however, show no appreciable difference between 

authigenic and detrital phases. Calcium in the K-feldspar 

in the limestone interbeds could not have been contributed 

by carbonate inclusions, as these zones were easily identified and avoided during analysis. The contribution of 

carbonate to calcium values in analyses of feldspars in 

dolostones is uncertain. 

Chemistry of insoluble residues 

The total amount of K-feldspar (authigenic and 

detrital) in each sample was estimated from the potassium 

content of the insoluble residue and the average potassium 

content of K-feldspars determined by the electron probe. 

The assumption that all the potassium was located inKfeldspar is justified on the basis of petrographic and XRD 

analyses. 

Both rubidium and barium commonly substitute 

for potassium in feldspars. This substitution is less pronounced at low temperatures; hence authigenic K-feldspar 

is depleted in rubidium and barium relative to hightemperature polymorphs. For this reason, potassium/ 

rubidium (K/Rb) and potassium/barium (K/Ba) ratios 

190 Studies in Diagenesis 

can be used to distinguish samples containing appreciable 

authigenic K-feldspar. 

The results of the insoluble residue analyses are 

shown for samples of three limestone and three dolostone 

interbeds in figure 9. Dolostone interbeds contain two to 

three times more insoluble residue than limestone 

interbeds. Potassium/rubidium and K/Ba ratios are 

significantly greater than the range of values reported for 

igneous rocks and igneous K-feldspars. Potassium/ 

rubidium ratios indicate no difference between insoluble 

residues of dolostone and limestone interbeds, whereas 

the K/Ba ratios are noticeably lower in limestone residues, 

due to an enrichment in barium. The reason for this 

enrichment is not clear. Excess barium could be present 

as finely disseminated barite. Although barite was not 

detected in these samples, it has been identified in the 

Conococheague Limestone at another location (unpublished data from this laboratory). 

X-Ray diffraction analysis 

X-ray diffraction (XRD) analyses of EDT A-soluble 

residues show the noncarbonate fraction of the Conococheague ribbon rocks to be dominated by K-feldspar, 

with lesser amounts of quartz. Minor amounts of 

muscovite were identified in some samples. Monoclinic 

and triclinic K-feldspars were distinguished by the XRD 

criteria of Goldsmith and Laves (1954); the presence of 

a single 131 reflection is indicative of monoclinic Kfeldspar, whereas a 131-131 doublet suggests the presence 

of triclinic K-feldspar. XRD patterns of insoluble residues 

of dolostone interbeds show only a single 131 monoclinic 

reflection, whereas patterns of limestone interbeds show 

a 131 monoclinic reflection flanked by less intense triclinic 

131 and 131 reflections (fig. 10). XRD patterns of detrital 

grains whose overgrowths have been removed by dissolution in dilute hydrofluoric acid also show both monoclinic 

and triclinic reflections, but the intensity of the monoclinic reflection is substantially reduced. These results suggest that the authigenic K-feldspar in both the dolostone 

and limestone interbeds is monoclinic, and that the 

detrital cores of grains with overgrowths are a mixture 

of triclinic and monoclinic K-feldspar. 

Evidence of interstitial brines 

Authigenic K-feldspar has commonly been reported in 

conjunction with rock features indicating hypersaline conditions (Fuchtbauer, 1972; Buyce and Friedman, 1975). 

The Conococheague Limestone is characterized by several 

features, most notably stromatolites and mudcracks, 

which are suggestive of a sabkha-like environment 
n '"[- I 

- I 

700 1 I 

2 50 EXPLANATION 

30 f-- 12 f- 35 ~ 160 1 600 800 • Limestone interbeds 

• • • • Dolostone interbeds 

,.. t ~ • 11 f-- ,J: 

• • Range of K/ Rb ratios • • reported for igneous 700 • rocks and meteorites by • • Taylor and others (1956) 25 f-- 10 f-- • 500 • 200 

• • + Range of K/Ba ratios 

• reported for igneous • • microclines by Deer and • 120 • 600 others (1967) • 

> 25 

s. 20 8 

100 ~ 400 ::r 5oo r- 150 1 QQ. 

tD • 

= 7 r;· 20 I "'1:::1 

0 

S' 15 • 6 80 300 400 

II> • II> c:· 5 15 3 100 

"TT 60 300 tD • i:i 10 4 200 II> 

"0 

AI 10 • ... 

:r 3 40 200 

;111:1 I 

·oo u Uf s: 

g- 5 2 100 

5 

= 20 ;111:1 

1 ~ 'Ill:" 

~ 

n 0 0 0 0 0 0 0 

= WEIGHT WEIGHT WEIGHT Rb(ppm) Ba (ppm) K/Rb K/Ba 

0 PERCENT PERCENT PERCENT 

a INSOLUBLE POTASSIUM POTASSIUMn RESIDUE FELDSPAR ::r (WHOLE-ROCK) tD 

AI 

OQ 

c Figure 9. Analyses of insoluble residues after ethylenediamine tetraaretic acid treatment of individual limestone and dolostone tD 

,.... interbeds. 3" 

tD 

II> 

5 

:I 

!D 

~ AI 

~ 

&'" 

Q. = 

... 

~ 
Detrital 

cores of 

Potassium feldspar 

in 

dolostone 

interbed 

in 

limestone 

interbed 

potassium feldspar 

in 

limestone 

interbed 

29.0 29.5 

m1crochne 1 Maximum 

(131) 

Monoclinic 

( 131) potassium feldspar 

DEGREES 28 

CuKa 

30.0 30.5 

Figure 10. Partial X-ray diffractograms of ethylenediamine 

tetraacetic acid-insoluble residues, showing characteristic reflections of fine-grained potassium feldspar in dolostone interbeds, 

sand-size potassium feldspar with overgrowths in limestone 

interbeds, and detrital cores of potassium feldspar grains from 

limestone interbeds following removal of overgrowths by 

dissolution in dilute hydrofluoric acid. Positions of characteristic 

reflections taken from Wright (1968). (CuKa radiation). 

192 Studies in Diagenesis 

exposed to periodic dessication. Other evidence indicates 

the former presence of interstitial fluids with salinities 

appreciably greater than normal seawater. The ratio of 

chloride to bromine was used by Carpenter (1978) to 

evaluate the history of oilfield brines. Because bromine 

does not precipitate during the evaporation of seawater 

and is progressively concentrated in the residual brine, 

the chlorine/bromine (Cl/Br) ratio is an indicator of the 

degree of evaporative concentration which a fluid has 

undergone. Experiments have shown that brines produced 

from the evaporation of seawater should have Cl/Br 

ratios ranging from 182 to 70 during the initial stages of 

halite precipitation (Zherebtsova and Volkova, 1966). 

Cl/Br ratios for bulk samples of limestone and dolostone 

interbeds, which are assumed to reflect the composition 

of fluid inclusions, range from 111 to 140 (table 2). These 

data suggest that these rocks contained a brine which was 

concentrated by evaporation well past the point of saturation with respect to halite. This conclusion is supported 

by the presence of daughter crystals of halite in a few 

fluid inclusion cavities in calcite exposed during polishing 

(fig. 11). The crystals are too large to have been precipitated by evaporation after the inclusions were breached, 

indicating that they were in equilibrium with the trapped 

fluids. 

Table 2. Representative chlorine and bromine analyses of 

dolostone and limestone interbeds. 

[Analysts: J. Gillison and C. Palmer Analyses in parts per million] 

Interbed Cl Br Cl/Br 

Do 1 ostone: 200 1.7 118 

260 2.3 113 

240 2.0 120 

Limestone: 190 1.6 119 

210 1.5 140 

200 1.8 111 

Thermochemical and mass-balance considerations 

The data of Helgeson (1974) indicate that a brine 

in equilibrium with both K-feldspar and albite should 

have a potassium/sodium (K/Na) activity ratio of 0.18 

at 300°C and 0.06 at 200°C. Extrapolation of Helgeson's 

data to 100°C yields an equilibrium ratio of 0.02. Solutions having K/Na activity ratios higher than the value 

for a given temperature should be supersaturated with 
Figure 11. Scanning electron micrograph of halite daughter crystal in fluid inclusion cavity in calcite breached during polishing. Smaller crystals probably represent 

residue from evaporation of fluid. 

respect to K-feldspar. According to the data of Zherebtsova and Volkova (1966), the evaporation of seawater 

should produce a brine with a KINa activity ratio of 0.02 

at the first point of gypsum precipitation, and ratios of 

0.04 to 0.17 during the initial stages of halite precipitation. Conodont alteration indices (Epstein and others, 

1977) indicate that the maximum temperatures 

experienced by the Conococheague Limestone were in the 

range of 200° to 250°C. Thus, the precipitation of Kfeldspar at these temperatures in a brine supersaturated 

with respect to halite is feasible from a thermochemical 

standpoint. 

Simple mass-balance calculations indicate that the 

observed mass of authigenic K-feldspar could not have 

formed isochemically, but had to have involved a flux 

of potassium-bearing fluids through these rocks. A typical 

dolostone interbed in the Conococheague Limestone contains in excess of 25 weight percent K-feldspar (see fig. 9). 

Estimates from cathodoluminescence images indicate that 

at least 75 percent of this K-feldspar is authigenic. 

Accordingly, 1 cubic centimeter of dolostone should contain about 80 mg of diagenetic potassium. If the initial 

sediment had a porosity of 0.70 and the pore fluids were 

supersaturated with respect to halite, less than 12 mg of 

dissolved potassium would have been available to participate in diagenetic reactions. Even if the initial 

siliciclastic mud in these interbeds was pure illite, there 

would not have been sufficient potassium for the Kfeldspar to have formed isochemically. Rather, the 

feldspathization had to involve the migration of multiple pore volumes of fluid through the rocks. 

Timing of feldspathization 

Initially, a 40 Ar / 39 Ar age spectrum was generated 

from a sample of K-feldspar grains having authigenic 

overgrowths. The overgrowths were then removed from 

Authigenic Potassium Feldspar in Ribbon Rocks, Conococheague Limestone, Maryland 193 
(J) 550 a: 

<I: 

w 

>-

u.. 

0 

(J) 

z 

0 

:::i 

_J 

~ 450 

~ 

u1 

(!) 

<I: 

fz 

w 

a: 

<I: 

Cl. 

Cl. 

<I: 350 

20 

----r- -------t-------

40 60 80 

I 

I 

I 

I 

100 

PERCENTAGE OF 39ArK RELEASED 

Figure 12. 40ArJ39Ar age spectra for detrital potassium feldspar grains with overgrowths and detrital 

grains whose overgrowths had been removed by dissolution in 5 percent hydrofluoric acid (step-like 

patterns), generalized age spectra for each (smooth dashed curves), and synthetic age spectrum for 

overgrowths alone (solid curve) obtained by subtracting one generalized age spectrum from the other. 

a second split of this sample by dissolution in a 5 percent 

solution of hydrofluoric acid and a second 40 Ar / 39 Ar 

age-spectrum analysis performed. The results of these two 

experiments are shown in figure 12. To obtain an approximate age spectrum for the authigenic overgrowths alone, 

a smooth curve was fitted to both of the experimentally 

produced age spectra and the "cores + overgrowths" 

curve subtracted from the "cores only" curve. Because 

estimates from cathodoluminescence images indicated 

that 45 to 50 percent of these grains were authigenic, no 

weighting factors were applied. The resulting synthetic 

curve is an approximation of the age spectrum that would 

result from the analysis of overgrowths alone if such a 

sample could. be analyzed. 

The synthetic age spectrum for overgrowths alone 

shows an age plateau at about 300 million years, 

indicating a Late Pennsylvanian-Early Permian age for 

the authigenic overgrowths. In addition, the age spectra 

of the "cores only" and "cores + overgrowths" indicate 

that the detrital K-feldspars still retain some portion of 

194 Studies in Diagenesis 

their predepositional history. Inasmuch as the core Kfeldspar fragments are partly monoclinic, this age spectrum indicates that the sample was not heated above 

250 °C following deposition, because above this temperature monoclinic K-feldspar loses all of its accumulated 

radiogenic 40Ar (Harrison and Mc-Dougall, 1982). This 

conclusion is consistent with the estimates of thermal 

history based on conodont alteration indices. 

The interpretation of the last 30-35 percent of the 

argon released from the samples is more difficult, as 

argon begins to exsolve in the argon-extraction apparatus 

at the temperatures represented by these steps on the age 

spectrum (Harrison and Mc-Dougall, 1982). The subtraction of the upper portion of these age spectra is 

probably not justifiable; probably only the first 65-70 

percent of the argon released represents true volume diffusion loss. 

The conclusion that the authigenic K-feldspar is 

Late Pennsylvanian-Early Permian in age is consistent 

with reports of low-temperature secondary magnetization 
with late Carboniferous pole positions in early to middle 

Paleozoic rocks of the central Appalachians (McCabe and 

others, 1983). These authors showed that this magnetization is not due to the thermal resetting of a detrital component, but rather is the result of the precipitation of 

secondary magnetite that was contemporaneous with 

folding, that is, Carboniferous in age. 

CONCLUSIONS 

We believe that the unusually large quantities of 

authigenic K-feldspar in ribbon rocks of the Conococheague Limestone reflect a chemical event associated 

with the mobilization and migration of basinal brines 

during the Alleghanian orogeny. This conclusion is supported by 40 Ar / 39 Ar age determinations on authigenic Kfeldspar overgrowths, petrographic and chemical evidence 

for the presence of hypersaline brines, and mass-balance 

and thermochemical calculations. The extensive nature 

of stratigraphically equivalent and lithologically similar 

carbonate units throughout the Appalachians and 

numerous reports of authigenic K-feldspar in these rocks 

suggests that orogenic activity may have been responsible for brine migration throughout the entire basin. 

Accordingly, the results of this study may have important implications for research on oil and gas migration 

pathways and the evolution of brine-derived mineral 

deposits. 

REFERENCES CITED 

Ali, A. D., and Turner, P., 1982, Authigenic K-feldspar in the 

Bromsgrove Sandstone Formation (Triassic) of central 

England: Journal of Sedimentary Petrology, v. 52, 

p. 187-197. 

Berg, R., 1954, Franconia Formation of Minnesota and Wisconsin: Geological Society of America Bulletin, v. 65, 

p. 857-882. 

Buyce, M. R., and Friedman, G. M., 1975, Significance of 

authigenic K-feldspar in Cambrian-Ordovician carbonate 

rocks of the Proto-Atlantic shelf in North America: Journal of Sedimentary Petrology, v. 45, p. 808-821. 

Carpenter, A. B., 1978, Origin and chemical evolution of brines 

in sedimentary basins: Oklahoma Geological Survey Circular 79, p. 60-77. 

Daly, R. A., 1912, Stratigraphy and structure of the Clarke 

Range, in American Cordillera-Forty-ninth parallel: 

Geological Society of Canada Memoir 38, p. 47-57. 

Dalrymple, G. B., and Lanphere, M. A., 1971, 40Ar/39Ar 

technique of K/ Ar dating-A comparison with the conventional technique: Earth and Planetary Science Letters, 

v. 12, p. 300-308. 

____ 1974, 40Ar/39Ar age spectra of some undisturbed 

terrestrial samples: Geochimica et Cosmochimica Acta, 

v. 38, p. 715-738. 

Deer, W. A., Howie, R. A., and Zussman, J., 1967, An introduction to the rock-forming minerals: New York, John Wiley 

and Sons, 528 p. 

Demicco, R. V., and Mitchell, R. W., 1982, Facies of the great 

American bank in the central Appalachians, in P. L. Lyttle, ed., Central Appalachian geology: NE-SE Geological 

Society of America 1982 Field Trip Guidebooks, p. 171-266. 

Epstein, A. G., Epstein, J. B., and Harr~s, L. D., 1977, Conodont color alteration-An index to organic metamorphism: 

U.S. Geological Survey Professional Paper 995, 27 p. 

Fuchtbauer, Hans, 1972, Influence of salinity on carbonate rocks 

in the Zechstein Formation, northwestern Germany: 

UNESCO, 1972, Geology of saline deposits, Proceedings 

of the Hanover Symposium, 1968, p. 23-31. 

Glover, E. D., 1961, Method of solution of calcareous materials 

using the complexing agent, EDTA: Journal of Sedimentary Petrology, v. 31, p. 622-626. 

Goldsmith, J. R., and Laves, F., 1954, The microcline-sanidine 

stability relations: Geochirnica et Cosmochimica Acta, v. 5, 

p. 1-19. 

Harrison, T. M., and Mc-Dougall, 1., 1982, The thermal 

significance of potassium feldspar K-Ar ages inferred from 40 Ar/39 Ar age spectrum results: Geochimica et Cosmochimica Acta, v. 46, p. 1811-1820. 

Helgeson, H. C., 1974, Chemical interaction of feldspars and 

aqueous solutions, in W. S. Mackenzie and J. Zussman, eds., 

The feldspars: Proceedings of a NATO Advanced Study 

Institute, p. 184-217. 

Kastner, M., 1971, Authigenic feldspars in carbonate rocks: 

American Mineralogist, v. 56, p. 1403-1442. 

Lanphere, M. A., and Dalrymple, G. B., 1971, A test of the 40 Ar/39 Ar age spectrum technique on some terrestrial 

materials: Earth and Planetary Science Letters, v. 12, 

p. 359-372. 

Mc-Cabe, C., Van der Voo, R., Peacor, D. R., Scotese, C. R., 

and Freeman, R., 1983, Diagenetic magnetite carries ancient 

yet secondary remanence in some Paleozoic sedimentary 

carbonates: Geology, v. 11, p. 221-223. 

Mazzullo, S. J., 1976, Significance of authigenic K-feldspar in 

Cambrian-Ordovician carbonate rocks of the Proto-Atlantic 

Shelf in North America-A discussion: Journal of Sedimentary Petrology, v. 46, p. 1035-1040. 

Merrihue, C., and Thrner, G., 1966, Potassium-argon dating 

by activation with fast neutrons: Journal of Geophysical 

Research, v. 71, p. 2852-2857. 

Reinhardt, J., and Hardie, L. A., 1976, Selected examples of carbonate sedimentation, Lower Paleozoic of Maryland: 

Guidebook for the 1976 Northeast-Southeast Joint Section 

Meeting of the Geological Society of America, Geological 

Society of America Guidebook 5, 53 p. 

Solomon, G. C., and Rossman, G. R., 1982, Water in feldspars: 

Geological Society of America Abstracts with Program, 

p. 622. 

Authigenic Potassium Feldspar in Ribbon Rocks, Conococheague limestone, Maryland 195 
Taylor, S. R., Emilius, C. H., and Exley, C. S., 1956, Some 

anomalous K/Rb ratios in igneous rocks and their 

petrological significance: Geochimica et Cosmochimica 

Acta, v. 10, p. 224-229. 

Tester, A. C., and Atwater, G. 1., 1934, The occurrence of 

authigenic K-feldspars in sediments: Journal of Sedimentary Petrology, v. 4, p. 23-31. 

Woodard, H. H., 1972, Syngenetic sanidine beds from the Middle Ordovician St. Peter Sandstone: Wisconsin Journal of 

Geology, v. 80, p. 323-332. 

196 Studies in Diagenesis 

Wright, T. L., 1968, X-ray and optical study of alkali 

feldspar-H. An X-ray method for determining the composition and structural state from measurement of 20 values 

for three reflections: American Mineralogist, v. 53, 

p. 88-104. 

Zherebtsova, I. K., and Volkova, N. N., 1966, Experimental 

study of the behavior of trace elements in the process 

of natural solar evaporation of Black Sea water and 

Sasyk-Sivash brine: Geochemistry International, v. 3, 

p. 656-670. 
Diagenesis of the Todilto Limestone Member of the 

Wanakah Formation, Chama Basin, New Mexico 

By JENNIE L. RIDGLEY, U.S. Geological Survey, Box 25046, Denver Federal 

Center, MS 916, Denver, Colorado 80225 

CONTENTS 

Abstract 198 

Introduction 198 

Methods of study 198 

Depositional textures 198 

Diagenetic alterations 199 

Micritization 199 

Types of cementation 199 

Authigenic evaporites 199 

Calcite replacement of evaporites 199 

Dolomite and dedolomitization 201 

Stylolitization 202 

Conclusions 205 

References cited 205 

FIGURES 

1. Map showing sample localities and structural features, southern part of 

Chama basin, New Mexico 199 

2. Photomicrographs showing micritization, early cementation, authigenic 

evaporites 200 

3. Photomicrographs showing calcite replacement of evaporites, dolomite, 

dedolomite, stylolites 202, 203 

4. Cross plot of bulk density versus compensated neutron porosity 204 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 197 
Abstract 

The Middle Jurassic Todilto Limestone Member of the 

Wanakah Formation in the Chama basin, north-central New 

Mexico, is dominantly a carbonate-gypsum sequence. 

Preliminary petrographic studies indicate that most of the carbonates in the southern part of the basin were originally 

deposited as carbonate mud that contained little bioclastic and 

detrital material. Diagenetic alteration of these sediments has 

resulted in a variety of textures. Early stages of diagenesis are 

marked by incomplete cementation of the sediments and 

micritization of biotic debris. Fine-grained mosaics of calcite 

crystals were precipitated as cement, partially or totally filling 

the pore spaces between detrital, larger carbonate, and 

bioclastic grains. Later stages of diagenesis are marked by the 

formation of moldic porosity followed by various stages of 

cementation by sparry calcite. Closure of vugs and voids by 

later stages of calcite cement is also apparent. Much of the later 

calcite is ferroan, in contrast to early calcite cements which 

are nonferroan. Authigenic gypsum or anhydrite crystals fill 

pore spaces within the sediment or replace the micrite. Locally 

some authigenic gypsum has been replaced by calcite. Some 

of the originally bedded gypsum has also been partially to totally 

replaced by sparry calcite. Minor dolomitization of both the 

micrite and sparry calcite occurred during late-stage diagenesis. 

Much of this dolomite has been subsequently dedolomitized. 

Clay minerals, quartz, organic material, and iron-rich minerals 

are concentrated along the boundaries of horizontal, vertical, 

and network stylolites. 

INTRODUCTION 

In the Chama basin of north-central New Mexico, 

the Middle Jurassic Todilto Limestone Member of the 

Wanakah Formation (Condon and Huffman, in press) 

overlies the Entrada Sandstone and is overlain by an 

upper unnamed clastic member of the Wanakah Formation (Ridgley, 1977). The Todilto consists of a carbonateevaporite sequence containing local interbedded sandstone and mudstone. The carbonate and clastic rocks 

form the lower part of the member and are continuous 

throughout the basin. The evaporites form the upper part 

of the member and locally may be absent from large areas 

of the basin. The carbonates consist of multistage calcite 

and minor remnant dolomite; the evaporites consist of 

gypsum at the surface and anhydrite in the subsurface. 

The carbonate facies may be further subdivided into 

a basal platy, laminated limestone, a middle crinkly 

limestone, and an upper microfragmental brecciated 

limestone. Where the limestone section is overlain by gypsum, the crinkly limestone is a more rhythmically bedded 

sequence of limestone and gypsum. The brecciated 

limestone is generally present only where there is no 

overlying gypsum; however, locally it has been found 

within the more massive gypsum facies. 

Preliminary petrographic studies of 36 samples 

from the carbonate facies from the southern part of the 

198 Studies in Diagenesis 

Chama basin (fig. 1) were made to ascertain the original 

mode of deposition and the subsequent diagenetic alterations of these facies. These samples were originally collected at random for use in other studies, and thus, the 

position of the samples are not tied to measured sections. 

The results of this study represent the first detailed 

petrographic depositional and diagenetic data to be 

published on the Todilto Limestone Member. For the 

purposes of this paper diagenetic alterations include all 

those changes that occur between the time of deposition 

and the time that conditions of metamorphism occur 

(Murray and Pray, 1965). 

METHODS OF STUDY 

Samples used in this study were originally collected 

for isotope analysis from widely spaced outcrop localities 

in the southern part of the Chama basin (fig. 1). These 

samples are from representative sections of the platy, 

crinkly, and brecciated carbonate facies and the gypsum 

and rhythmically bedded limestone-gypsum evaporite 

facies. Thin sections of each sample were examined with 

the petrographic microscope; rock chips from selected 

samples were examined with a scanning electron microscope (SEM) in order to identify the relationships of 

diagenetic minerals at the greater resolution capability of 

the SEM. 

To check for the presence of dolomite, thin sections 

_ were stained with a combined solution of alizarin red-S 

and potassium ferricyanide (Dickson, 1965). This technique permits the distinction of calcite and dolomite and 

indicates the distribution of ferrous iron between these 

two minerals. It does not, however, indicate the amount 

of iron present in ferroan calcite or ferroan dolomite. 

DEPOSITIONAL TEXTURES 

Petrographic and SEM examination indicate that 

in most samples much of the carbonate was deposited 

originally as micrite or microcrystalline carbonate mud 

(1-4/lm) with a low content of allochem and terrigenous 

components (less than 10 percent). The allochem components are pellets, superficial ooliths, bioclastic fragments, and glauconite grains. The terrigenous components are quartz and detrital clays. The allochem 

components occur singly or in clusters; the terrigenous 

components occur singly, but more commonly have been 

concentrated along later stylolite boundaries. In contrast, 

a few samples contain a higher proportion of terrigenous 

material (between 10 and 50 percent) and may be 

classified as sandy or silty micrites and calcilutites (Folk, 

1959). In these samples, the terrigenous components, 

generally quartz and detrital clays, are concentrated along 

laminae less than 1 mm thick. The silty micrites and 
107° 00" 

San Juan basin 

Mesa Alta 

OA 

Chama basin 

Jemez volcanic field 

10 KILOMETERS 

I 

106° 30" 

•o 

Ghost 

Ranch 

Figure 1. Index map of the southern part of the Chama basin, 

New Mexico. Triangles represent sample localities. Heavy lines 

separate adjacent structural features. 

calcilutites commonly occur with thin interbedded sandstones or mudstones that are locally present in the basal 

platy and middle crinkly facies. Petrographic observations 

indicate that for most of the samples the original depositional fabric has been altered or obscured by subsequent 

diagenesis. 

DIAGENETIC ALTERATIONS 

Several types of diagenetic events have been 

recognized and are discussed separately as follows. 

Micritization 

One of the earliest diagenetic events was the micritization of bioclastic shell material (fig. 2A). Micritization took place by the repeated boring of shell material 

by algae and the subsequent infilling of empty bores with 

micritic carbonate (Bathurst, 1966). The result of this 

process was the formation of a micrite envelope that surrounds the remaining shell material. 

Types of cementation 

Cements in the carbonate and clastic rocks of the 

Todilto are of various ages. The earliest type of cement 

is rim cement around grains, especially where allochem 

and terrigenous grains are moderately to densely packed. 

The rim cements occur as fibrous fringing cements and 

are especially abundant around pellets and superficial 

ooliths (fig. 2B). The rim cements also occur as mosaics 

of small crystals around quartz grains and calcite grains 

greater than 4 ~tm. 

Other types of cements precipitated at various stages 

of lithification of the carbonates. Some cements fill pore 

spaces that formed by the dissolution of bioclasts 

(fig. 2A) and evaporite nodules. Other cements fill 

intergranular pore space, late fractures, and voids formed 

when evaporite minerals dissolved. Late calcite cements 

are commonly ferroan, in contrast to early calcite cements 

which are nonferroan. 

Authigenic evaporites 

Examination of thin sections and rock chips and 

field relations of the carbonate-evaporite rocks indicate 

the presence of primary bedded and later authigenic 

evaporites. To date, gypsum and anhydrite are the only 

evaporite minerals that have been identified. Authigenic 

gypsum and anhydrite replace micrite (figs. 2C, W, and 

2£); replacement began prior to complete lithification. 

Gypsum also occurs as an authigenic precipitate in open 

pore space (fig. 2F). 

Two forms of authigenic gypsum and anhydrite 

replace partially lithified micrite. In one, anhydrite grew 

as euhedral crystals (0.08-0.4 mm in length), either singly 

or in clusters. Presently, in the rhythmically bedded crinkly 

limestone facies these euhedral crystal volumes are 

occupied by masses of tiny gypsum crystals (3-50 ~tm) 

(fig. 2C). This textural relationship suggests that the 

anhydrite either dissolved and created moldic porosity that 

was subsequently filled by numerous small gypsum 

crystals or that during hydration the single anhydrite 

crystals were replaced by many small gypsum crystals. 

In the other form, gypsum grew as lath-shaped 

crystals of various sizes (0.2-0.08 mm in length) from 

numerous point sources within the partially lithified 

micrite. These lath-shaped crystals grew at a variety of 

angles to one another and eventually coalesced to form 

nodular gypsum or masses of gypsum enclosing remnant 

micrite patches (figs. 2D and 2£). The gypsum in these 

masses is both displacive and replacive. It is displacive in 

that it pushed aside carbonate laminae and stylolites as 

the nodules grew. It is replacive in that it incorporated 

some of the previously deposited carbonate as it grew (figs. 

Wand 2£). This particular type of diagenetic alteration 

is characteristic of the rhythmically bedded limestonegypsum facies that are overlain by more massive gypsum. 

Calcite replacement of evaporites 

Gypsum replaced by calcite is present in all facies 

but is most abundant where overlying gypsum is absent. 

Diagenesis of the Todiito Limestone Member of the Wanakah Formation, New Mexico 199 
L_Jlmm C ' ~ 

Figure 2. Photomicrographs of micritization, early cementation, and authigenic evaporites. A, Micritized bioclast. Probable 

ostracode shell replaced with micritic envelope. Sometime after formation of the micrite envelope the remaining fossil was leached. 

Later sparry calcite cement filled the interior. Crossed nicols. Field of view 1.0 mm. 8, Early radial rim cement of calcite around 

pellets or superficial ooliths. Light areas around ooliths or pellets are the radial rim cement which has undergone some neomorphism. Field of view 1.0 mm. C, Gypsum replacement of anhydrite. Cluster of euhedral grains set around large euhedral area 

occurs in micrite. Euhedral grains, once anhydrite, are now filled with mosaics of small gypsum crystals. Field of view 1.0 mm. 

0, Gypsum replacement of micrite. Lath-shaped gypsum (lighter areas) replacing micrite (center dark mass and dark wispy areas 

above). Dark wavy bands are stylolites. Fractures oriented perpendicular to stylolites are filled with gypsum. Field of view 5.1 

mm. E, Gypsum replacement of micrite. Lath-shaped gypsum (lighter areas) replaces micrite (dense gray areas at left and right). 

Dark bands are microstylolites with quartz grains (clear grains) along seams. Field of view 5.1 mm. F, Scanning electron micrograph 

of authigenic gypsum. Gypsum (large light-gray grain in center) probably precipitated in pore. Later quartz and calcite overgrowth 

can be seen near bottom of crystal. Field of view 12 11 m. 

200 Studies in Diagenesis 
Many of the resulting textures mimic textures produced 

by the growth of authigenic anhydrite or gypsum 

(fig. 3A). For example, in the brecciated facies large areas 

of sparry calcite contain patches of original micrite in a 

pattern similar to that seen where gypsum has replaced 

portions of the original micrite (compare figs. Wand 2E 

to fig. 3A). These patches of micrite do not appear to 

be supported by framework or matrix, but rather appear 

to float within the sparry calcite areas. Thus, the micrite 

patches are not true breccia clasts and the sparry calcite 

that surrounds them is not totally cement. Most of the 

sparry calcite crystals are replacements of gypsum by 

calcite. Occasional relict gypsum within the large sparry 

calcite crystals is evidence of replacement rather than 

dissolution followed by precipitation of cement in the 

resulting void. Invariably large, single calcite crystals have 

patchy extinction and contain inclusions that may be gypsum along crystallographic boundaries. Also, the manner in which many of the sparry calcite crystals interlock 

resembles the random interlocking pattern seen in larger 

gypsum laths. 

One principal difference between the calcite and 

gypsum replacement textures is the presence of more 

quartz in the calcite. Some of this quartz undoubtedly was 

part of the original sediment; the remainder appears to 

be authigenic. The authigenic quartz commonly occurs 

as crystals with well-defined terminations and hexagonal 

outlines or as fibrous chalcedony. The presence of 

authigenic quartz in these rocks and the textural relations 

between the quartz and calcite suggest the presence of 

former evaporite minerals (Friedman, 1980). 

Calcite also occurs as mosaics of crystals (about 1-3 

,urn in width) filling well-defined euhedral volumes 

(fig. 3B; see also, Stapor, 1968) that were once occupied 

by gypsum or anhydrite. This diagenetic texture is more 

abundant in the brecciated facies. As was noted previously, 

similar euhedral volumes in the bedded, crinkly limestone 

facies contain mosaics of gypsum crystals. Whether the 

calcite actually replaced the gypsum or whether the gypsum was selectively leached and the resulting moldic 

porosity cemented with mosaics of calcite crystals has not 

been determined. Similar textures in the Permian Tansill 

Formation, Texas and New Mexico, were described by 

Lucia (1961) as representing either direct calcite replacement of gypsum or anhydrite or precipitation of calcite in 

voids created by the leaching of these evaporite minerals. 

Direct calcite replacement of primary bedded anhydrite or gypsum was also observed (fig. 3C). In figure 3C, 

the large sparry calcite is cloudy, and remnants of gypsum and crystal shape reminiscent of gypsum are apparent. 

Dolomite and dedolomitization 

Several authors (for example, Rawson, 1980; Stapor, 

1968) who have studied the Todilto have commented on 

the lack of dolomite; however, Gabelman (1970) reported 

finding a minor amount of dolomite in the Todilto at the 

Flat Top uranium mine near Grants, New Mexico. His 

identification of dolomite was based on the lack of reaction of portions of the limestone to weak solutions of 

hydrochloric acid. In the present study, petrographic and 

SEM examination of thin sections and rock chips found 

only minor amounts of dolomite. The presence of 

dolomite in the Todilto is also suggested by the results 

obtained by plotting bulk density against compensated 

neutron porosity. The values were obtained by interpretation of density and compensated neutron-porosity data 

formation shown on compensated neutron-density logs 

of holes that penetrate the Todilto Limestone Member in 

the east-central part of the San Juan basin. Several of 

these cross plots are shown in figure 4. Successive changes 

in composition, from the base to the top of the member, 

are shown for each hole. These compositional changes 

suggest the presence of anhydrite and mixtures of 

dolomite and limestone. 

In the Todilto the dolomite generally occurs as remnant patches in micrite and microspar matrix (fig. 3D). 

A minor amount of dolomite was also observed in one 

thin section in the form of overgrowths on gypsum crystals 

present within micrite patches. The rest of the dolomite 

occurs as remnants within large sparry calcite crystals, 

some of which formed as replacement of evaporites. 

The age relations of the dolomite are not totally 

clear. Groundmass dolomite of rhombic and anhedral 

shape replaces micrite and microspar and may have 

formed early, perhaps contemporaneous with primary 

evaporite precipitation. The remnant dolomite in the large 

sparry calcite crystals may have formed somewhat later 

under different geochemical conditions. The presence of 

zoned dolomite, later dedolomitized, indicates fluctuations in the geochemical environment. The remnant 

dolomite may have formed as partial replacement of gypsum and was subsequently incorporated in the calcite 

when it replaced the gypsum; however, no detectable 

dolomite was found in the coarse gypsum patches in the 

two thin sections of highly gypsiferous limestone 

examined. An alternative explanation is that early in the 

diagenetic history of the limestone calcite began to replace 

the gypsum as a result of fluctuation in calcite-gypsum 

solubilities (Wigley, 1973). This calcite was then subsequently partially dolomitized. 

The minor amount of dolomite present is attributed 

to extensive dedolomitization-that is, replacement of 

dolomite by calcite. Dedolomitization is a diagenetic process that occurs at low temperature and low partial pressure of carbon dioxide and is apparently promoted by the 

dissolution of sulfates, in particular gypsum (de Groot, 

1967; Yanat'eve, 1955). It is therefore a near-surface 

process related to present exposure or past unconformities. 

Dedolomitization of the Todilto Limestone Member may 

be a recent phenomenon related to present exposure; 

Diagenesis of the Todilto Limestone Member of the Wanakah Formation, New Mexico 201 
however, some dedolomitization may have occurred 

shortly after early burial when the ground-water 

chemistry changed. Pipiringos and O'Sullivan (1978) and 

Lupe (1981) postulated the presence of an unconformity 

at the top of the Summerville Formation-Wanakah Formation interval-the immediately overlying formationthroughout this and adjacent regions. This unconformity, 

however, is not readily apparent at the outcrop. Examination of core of the Todilto from the central part of the 

depositional basin might help resolve the age of the 

dedolomitization. 

Evidence for dedolomitization is seen in a variety 

of textures. In most thin sections some matrix calcite 

occurs as rhombic or incomplete rhombohedral crystals. 

These crystals may be primary or they may represent 

replacement of dolomite. Calcite replacement of zoned 

dolomite is also apparent (fig. 3£). In several thin sections clotted or pseudomicropelletal texture is present 

(fig. 3F) which is similar to the grumeleuse texture 

reported by Cayeux (1935). Evamy (1967) suggested that 

some clotted textures might be due to the dedolomitization of cloudy-center dolomite described by Murray 

(1964). 

Another type of dedolomitization texture observed 

was the progressive replacement of dolomite crystals by 

calcite. This process led to the formation of irregular remnants of dolomite within the original crystal. Complete 

replacement of the dolomite crystals may have resulted 

in a texture that is difficult to distinguish from that of 

the original limestone. Similar textures of dedolomitized 

dolomitic limestones were noted by Munn and Jackson 

(1980, fig. 2, plate 1) in their examination of the Lower 

Carboniferous Hoptonwood and Matlock Formations, 

Wirksworth, United Kingdom. They reported the progressive replacement of dolomite by calcite that began 

with selective replacement at the margins or along planes 

of crystallographic weakness, such as cleavages, and 

culminated in partial or total replacement. This type of 

dedolomitization texture was most noticeable in the upper 

2 m of the limestones and was considered to be related 

to present-day weathering. 

Stylolitization 

The formation of microstylolites is, perhaps, the 

most striking visible form of diagenetic alteration in the 

Todilto Limestone Member. Horizontal, vertical, and inclined stylolites are all present, although horizontal 

stylolites predominate. The microstylolite seams are 

generally less than 0.1 mm thick; suture amplitudes are 

also usually less than 0.1 mm. Residual material concentrated along the stylolite seam consists of clay minerals, 

quartz, organic material, and iron-bearing minerals. The 

morphology of the stylolites is varied. Simple or primitive 

wavelike types, sutured types, and sharp-peaked types, 

Figure 3. Photomicrographs of calcite replacement of evaporites, dolomite, dedolomite, and stylolites. A, calcite replacement 

of gypsum. Darker patches and wisps are original micrite. Sparry calcite (light area) is for the most part replacement of gypsum, 

although some may represent cement. Most of the calcite shows patchy extinction and remnants of gypsum. Sparry calcite fills 

fractures in micrite near center of photograph. Circular areas in center are bubbles; clear patches are voids. Field of view 1.64 

mm. 8, Calcite filling euhedral volumes once occupied by anhydrite crystals. Large lath-shaped area near the center of the photograph 

and clusters of smaller euhedral volumes are now filled with mosaics of small calcite crystals or large sparry calcite crystals. This 

assemblage may represent calcite replacement of anhydrite or gypsum or calcite cement filling voids that resulted from leaching 

of anhydrite. Crossed nicols. Field of view 2.5 mm. C, Calcite replacement of bedded evaporites. Micrite layers (dark areas) 

202 Studies in Diagenesis 
interbedded with sparry calcite (light areas) that has replaced primary evaporites. Field of view 2.5 mm. 0, Ground mass dolomite 

in micrite. Light areas are dolomite, quartz, and late kaolinite overgrowths. Dark areas are micrite. Dolomite replaced part of 

the micrite and has subsequently been partly dedolomitized. Field of view 0.26 mm. E, Calcite replacement of zoned dolomite. 

Crossed nicols. Field of view 0.41 mm. F, Clotted or grumeleuse texture due to dedolomitization of cloudy-center dolomite. Thin, 

dark, irregular band between grains may be due to pressure solution. Field of view 0.41 mm. G, Network microstylolites. Residual 

material of clay, organic material, iron-rich minerals, and quartz are concentrated along stylolite seams. Field of view 1.64 mm. 

H, Tension gashes or extension fractures. Coarse sparry calcite fills wedge-shaped fractures that begin at stylolite seams (dark 

wavy bands). Field of view 2.5 mm. 

Diagenesis of the Todilto Limestone Member of the Wanakah Formation, New Mexico 203 
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Figure 4 (above and facing page). Cross plot of bulk density versus compensated neutron porosity for several holes that penetrated 

the Todilto Limestone Member of the Wanakah Formation. Bulk density and compensated neutron porosity values shown were 

obtained from examination of density and compensated neutron porosity and formation density porosity logs. Successive vertical 

changes in composition from base to top of formation for each hole are shown as occurring from top to bottom of the figure. 

Base diagram is from Schlumberger (1972) and is for freshwater, liquid-filled holes. 

204 Studies in Diagenesis 
LOCATION OF DRILL HOLES IN THE TODILTO LIMESTONE MEMBER OF THE WANAKAH FORMATION, 

CHAMA BASIN, NEW MEXICO 

Company Well name Location Symbol 

Damson Oil Corporation Fernandez No. 1-3 Sec. 3, T. 16 N • , R. 8 w. • 

Hanson Oil Corporation Candy Butte 3 Sec. 8, T. 17 N • , R. 2 w. • 

Rij an Oil Company, Inc. Federal No. 1-26 Sec. 26, T. 19 N., R. 3 w. + 

Gulf Energy and Minerals Panos Investment Sec. 21, T. 20 N • , R. 5 w. • Company--u.s. Company "D" Federal 

Dome Petroleum Santa Fe 33-20-10 

as defined by Park and Schot (1968, fig. 1), are present. 

Subparallel horizontal and vertical stylolites were noted 

in some thin sections; in others, interconnecting networks 

of stylolites are more common (fig. 3G). These network 

stylolites are similar to those discussed by Park and Schot 

(1968; fig. 5), who suggested that they might be of tectonic origin. 

Associated with many of the stylolites, but particularly with the network type, are fractures that are 

nearly perpendicular to the seam. These fractures are 

filled with sparry calcite and originate at a peak or valley 

of the stylolite seam. Many of these fractures are wedge 

shaped; they are wider where they abut against the seam 

(fig. 31/). Fractures extending from a single stylolite seam 

in opposite directions are not uncommon. Choukroune 

(1969) called these types of fractures tension gashes and 

stated that they were related to the same stress conditions 

that formed the stylolites. 

More recently, Nelson (1979) redefined "tension 

gashes'' as extension fractures and concurred with 

Choukroune's theory of origin. Nelson (1981) further 

suggested that the pattern of stylolites with extension fractures are similar to stress patterns produced in laboratory 

triaxial compression and extension tests and are therefore 

of probable tectonic origin. He envisioned these types of 

stylolite patterns as having been produced by a slow tectonic process operating on soluble rocks being compacted 

in one direction and extended in another direction at 90°. 

Complete evaluation of the stylolite patterns and 

time of formation will require more detailed study; 

however, certain observations as to timing of stylolite formation suggest that it began during early burial. In the 

bedded crinkly limestone, remnant patches of micrite 

enclosed within the growing gypsum contain stylolites of 

various forms (figs. 2D and 2E}. Some of these stylolites 

contain gypsum-filled fractures oriented perpendicular to 

the seam (fig. W). Locally, the stylolites extend out from 

the micrite into the gypsum. They may completely cross 

the gypsum, where the gypsum mass is small, or terminate 

within it. This phenomenon suggests that the growing 

gypsum was unable to incorporate all the seam material 

as it grew and not that late stylolites cut across the gypsum. Thus, the formation of stylolites appears to have 

Ill Sec. 33, T. 20 N • , R. 10 w. 0 

preceded the formation of the gypsum that replaced the 

micrite. 

CONCLUSIONS 

These preliminary studies indicate a varied and 

complex diagenetic history for the Todilto Limestone 

Member of the Wanakah Formation. The most important diagenetic alterations are a complex cementation 

history, dolomitization followed by dedolomitization, an 

evaporite replacement of carbonates and subsequent carbonate replacement of these evaporites, and an early 

formation of stylolites. This preliminary study lays the 

groundwork for the future investigations needed to 

explain more fully the diagenetic history of the Todilto 

Limestone Member. 

REFERENCES CITED 

Bathurst, R. G. C., 1976, Carbonate sediments and their 

diagenesis: New York, Elsevier, 658 p. 

Cayeux, L., 1935, Les roches sedimentaires de France; les roches 

carbonatees: Paris, Masson, 463 p. 

Choukroune, P., 1969, An example of mesoscopic analyses of 

concentric folding in limestone series: Tectonophysics, v. 7, 

p. 57-70. 

Condon, S. M., and Huffman, A. C., in press, Revisions in 

nomenclature of the Middle Jurassic Wanakah Formation, 

northwest New Mexico and northeast Arizona: U.S. 

Geological Survey Bulletin 1633-A. 

de Groot, K., 1967, Experimental dedolomitization: Journal 

of Sedimentary Petrology, v. 37, p. 1216-1220. 

Dickson, J. A. D., 1965, A modified staining technique for carbonates in thin section: Nature, v. 205, p. 587. 

Evamy, B. D., 1967, Dedolomitization and the development 

of rhombohedral pores in limestone: Journal of Sedimentary Petrology, v. 37, p. 1204-1215. 

Folk, R. L., 1959, Practical petrographic classification of 

limestones: American Association of Petroleum Geologists 

Bulletin, v. 43, p. 1-38. 

Friedman, G. M., 1980, Dolomite is an evaporite mineral; 

Evidence from the rock record and from sea marginal 

ponds of the Red Sea: Society of Economic Paleontologists 

and Mineralogists Special Publication 28, p. 69-80. 

Diagenesis of the Todilto Limestone Member of the Wanakah Formation, New Mexico 205 
Gabelman, J., 1970, The Flat Top uranium mine, Grants, New 

Mexico: U.S. Atomic Energy Commission, RME-4112, 81 p. 

Lucia, F. J., 1961, Dedolomitization in the Tansill (Permian) Formation: Geological Society of America Bulletin, v. 72, 

p. 1107-1110. 

Lupe, R. D., 1981, Stratigraphic sections of subsurface Jurassic 

rocks in the San Juan basin, New Mexico, Colorado, Utah, 

and Arizona: U.S. Geological Survey Oil and Gas Chart 118. 

Munn, D., and Jackson, D. E., 1980, Dedolomitization of Lower 

Carboniferous dolostone in the Wirksworth area, Derbyshire, England: Geological Magazine, v. 117, p. 607-612. 

Murray, R. C., 1964, Preservation of primary structures and 

fabrics in dolomite, in Imbrie, J., and Newell, N.D., eds., 

Approaches to paleoecology, 1st ed.: New York, Wiley and 

Sons, 432 p. 

Murray, R. C., and Pray, L. C., 1965, Dolomitization and 

limestone diagenesis-An introduction: Society of 

Economic Paleontologists and Mineralogists Special 

Publication 13, p. 1-2. 

Nelson, R. A., 1979, Natural fracture systems, description and 

classification: American Association of Petroleum 

Geologists Bulletin, v. 63, p. 2214-2221. 

___ 1981, Significance of fracture sets associated with stylolite 

zones: American Association of Petroleum Geologists 

Bulletin, v. 65, p. 2417-2425. 

Park, W. C., and Schot, E. H., 1968, Stylolites; their nature and 

origin: Journal of Sedimentary Petrology, v. 38, p. 175-191. 

206 Studies in Diagenesis 

Pipiringos, G. N., and O'Sullivan, R. B., 1978, Principal unconformities in Triassic and Jurassic rocks, western interior 

United States-A preliminary survey: U.S. Geological Professional Paper 1035-A, 29 p. 

Rawson, R. R., 1980, Uranium in Todilto Limestone (Jurassic) 

of New Mexico-Example of a sabkha-like deposit, 

in Rautman, C., compiler, Geology and mineral technology of the Grants uranium region 1979: New Mexico 

Bureau of Mines and Mineral Resources Memoir 38, 

p. 304-312. 

Ridgley, J. L., 1977, Stratigraphy and depositional environments 

of Jurassic-Cretaceous rocks in the southwest part of the 

Chama Basin, New Mexico: New Mexico Geological Society 

Guidebook, 28th Field Conference, San Juan Basin III, 

p. 153-158. 

Schlumberger, 1972, Log interpretation charts: Houston, Texas, 

Schlumberger Limited, 92 p. 

Stapor, F. W., 1968, Stratigraphy of the Todilto Formation in 

the Ghost Ranch area, New Mexico: Madison, Wisconsin, 

University of Wisconsin, M.S. thesis, 75 p. 

Wigley, T. M. L., 1973, Chemical evolution of the system calcitegypsum-water: Canadian Journal of Earth Science, v. 10, 

p. 306-315. 

Yanat'eva, 0. K., 1955, Effect of aqueous solutions of gypsum 

on dolomite in the presence of carbon dioxide: Doklady 

Akadimiya Nauk S.S.S.R., v. 101, p. 911-912. 
Mineralogy and Stable Isotope Geochemistry of 

Carbonate and Sulfate Minerals in Diagenetically 

Altered Tertiary and Cretaceous Sandstones, 

Uinta Basin, Utah 

By JANET K. PITMAN, U.S. Geological Survey, Box 25046, DFC, MS 916, 

Denver, Colorado 80225, 

MARTIN B. GOLDHABER, U. S. Geological Survey, Box 25046, DFC, MS 

916, Denver, Colorado 80225, and 

THOMAS D. FOUCH, U.S. Geological Survey, Box 25046, DFC, MS 916, 

Denver, Colorado 80225 

CONTENTS 

Abstract 208 

Introduction 208 

Acknowledgments 208 

Methods 208 

Geology 209 

Whole-rock mineralogy 209 

Carbonate mineralogy and diagenesis 212 

Porosity development 212 

Stable isotopes 214 

Carbon isotopes 214 

Oxygen isotopes 216 

Dolomite 217 

Summary 217 

References cited 217 

FIGURES 

1. Index map of Uinta basin, Utah, showing location of Natural Buttes 

field 208 

2. Chart showing correlation of major Tertiary and Cretaceous rock units in 

Uinta basin, Utah 210 

3. Ternary diagram showing mineralogic composition of sandstones in cored 

part of Wasatch and Tuscher Formations 211 

4. Photomicrographs of authigenic anhydrite, iron-free and iron-bearing carbonate, and associated dissolution features in the Wasatch and Tuscher 

Formations 213 

5. Cross plot showing o13CPDB versus o180PDB for carbonate grains and cement in 

sandstones from the Tuscher and Wasatch Formations 216 

TABLE 

1. Stable isotope data for calcite, dolomite, and anhydrite in sandstones from 

the Wasatch and Tuscher Formations 215 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 207 
Abstract 

Nonmarine Tertiary and Cretaceous sandstones in the 

Uinta Basin, Utah, have been extensively modified by the formation of a carbonate mineral assemblage composed of ironfree calcite, ferroan calcite, dolomite, and ankerite. Iron-free 

calcite was the earliest carbonate to form followed by ferroan 

calcite which is now the most abundant carbonate phase. Most 

of the calcite appears to have formed during early burial before 

significant compaction had taken place. 

The o13C compositions of calcite in the Tuscher Formation vary between -3.1 and -8.1 per mil. The source of heavy 

carbon (near -3 per mil) was dissolved marine carbonate that 

was derived locally from depositional waters. In the Wasatch 

Formation, o13C values are distinctly light, ranging from -5.2 

to -10.6 per mil. The isotopically light carbon may have been 

derived from atmospheric C02 or it may have originated from 

organic matter in kerogenous shales, as o34S values ( + 30 per 

mil average) for associated anhydrite are similar to those for 

organic matter and pyrite in the Eocene Green River Formation. 

o180 compositions of calcite are similar in the Tuscher 

(-9.7 to -14.2 per mil) and Wasatch Formations (-9.4 to -13.1 

per mil). These data suggest that the variation in pore-water 

composition was a more important factor than the effects of 

temperature with increasing burial depth in determining o180 

systematics. 

INTRODUCTION 

Thick sequences of nonmarine Tertiary and 

Cretaceous rocks in the eastern part of the Uinta Basin, 

Utah, are unconventional reservoirs for large amounts 

of natural gas (fig. 1). Individual sandstones have 

undergone extensive postdepositional mineral alteration 

that resulted in the formation of authigenic carbonate and 

minor anhydrite. This paper is a study of the carbonate 

and sulfate mineral assemblage that characterizes the 

lower Tertiary Wasatch and Upper Cretaceous Tuscher 

Formations. The recognition of individual carbonate 

phases and the diagenetic environment in which they 

formed is important in reconstructing the burial history 

of the sandstones. The purpose of the paper is to (1) identify and describe the discrete carbonate mineral phases 

that are present in Tertiary and Cretaceous sandstones, 

(2) present data for carbon and oxygen isotope compositions of carbonate grains and cement, and (3) discuss the 

origin of carbonate and its relation to the occurrence of 

anhydrite. 

Acknowledgments 

We thank Lisa Pratt for her helpful suggestions and 

critical review of the manuscript. Isotope analyses were 

performed by Global Geochemistry Corp. The U.S. 

Geological Survey research was supported by a grant 

from the U.S. Department of Energy. 

208 Studies in Diagenesis 

0 50 100 KILOMETERS 

EXPLANATION 

Tertiary volcanic and Intrusive 

rocks 

Tertiary sedimentary rocks 

Pr~'rsbrian sedimentary 

Precambrian metamorphic 

rocks 

Precambrian Intrusive rocks 

--Contact 

...u...u..uJJ. Fault 

~ Thrust fault 

e Gas well that provided 

core samples: 

Figure 1. Index map of Uinta basin, Utah, showing location 

of Natural Buttes field. 

METHODS 

About 80 sandstone samples from the CIGE 

Natural Buttes 21 well were analyzed for their mineralogic 

composition and diagenetic character using standard 

X-ray diffraction and petrographic techniques. Cored 

intervals that were sampled included the Wasatch Formation between 1,342 and 1,373 m (4,389 and 4,508 ft) 

and the Tuscher Formation between 1,586 and 2,227 m 

(6,350 and 6,509 ft). In thin section, mineral abundances 

were determined by counting 300 grains per section. Each 

section was impregnated with blue epoxy and stained with 

alizarin red-S and potassium ferricyanide to differentiate 

between individual carbonate phases. Scanning electron 

microscopy was used to identify authigenic minerals on 

the basis of their morphology. 

Carbon and oxygen isotope compositions of carbonate minerals were determined for 27 samples by 

Global Geochemistry Corporation.• All samples were 

1 Use of company names is for descriptive purposes only and does 

not imply endorsement by the U.S. Geological Survey. 
first treated with Clorox to mtmmtze contamination 

from reactive organic compounds. For samples that contained both calcite and dolomite-ankerite, an attempt was 

made to collect C02 from each of the two carbonate 

fractions based upon the relative reactivity of calcite and 

dolomite in 100-percent H3

P04 (Walters and others, 

1972). Carbon dioxide was collected after 6, 30, and 102 

hours of reaction time at 25 °C. The 30-hour aliquot 

representing mixed calcite and dolomite was discarded. 

The 6- and 1 02-hour aliquots were assumed to represent 

the calcite and dolomite-ankerite fractions, respectively. 

Dolomite data were not corrected for the difference in 

fractionation factor between 100-percent H3

P04 and 

dolomite compared with calcite (Land, 1980). Carbon 

( 13C/12C) and oxygen (1 80/160) stable isotope ratios are 

reported as o13C and o180 values with units in parts per 

thousand (per mil) relative to the PDB standard (Cretaceous belemnite from the Peedee Formation). The 

samples chosen for isotope analyses were studied 

petrographically in more detail than the rest of the 

samples. 

Four samples of anhydrite were analyzed for o34S. 

Anhydrite was dissolved in 6 N HCl; sulfate was precipitated as BaSO 4 which was decomposed directly to S02 

for analysis of the sulfur isotope ratio. The barite was 

purified by reaction with hot, alkaline sodium carbonate. 

The resulting sodium sulfate solution was reprecipitated 

as BaS04 and decomposed to S02 to determine o34S. 

GEOLOGY 

Nonmarine rocks of the lower Tertiary Wasatch 

Formation and Upper Cretaceous Tuscher Formation in 

the Uinta Basin, Utah, were deposited over a time span 

of 25-30 million years in subaqueous depositional settings 

that record the evolution from a Late Cretaceous marine 

foreland trough to an internally drained Tertiary 

hydrographic basin (Fouch and others, 1983). This transition occurred in response to Laramide tectonism during 

late Maestrichtian and early Paleocene time. Accompanying Laramide deformation was widespread erosion, 

which is reflected in a regional unconformity that 

separates Cretaceous from Tertiary strata. Locally, the 

unconformity represents a hiatus of about 15 million 

years (Fouch and others, 1983). The stratigraphic and age 

relations between the Wasatch and Tuscher Formations 

and adjacent rock units are shown in figure 2. 

In the eastern part of the Uinta Basin, the upper 

part of the Wasatch Formation is composed of finegrained lenticular sandstones that contain small-scale 

cross beds. Commonly the sandstones contain red coloration and are interbedded with variegated mudstones 

and shales. Some sandstone beds locally are conglomeratic. The presence of numerous red sandstone beds 

as well as their stratification and grain size indicate that 

these rocks formed from sediment deposited in deltaic 

plain environments peripheral to ancient Lake Uinta. 

Fouch (1975) and Ryder and others (1976) showed that 

the upper part of the Wasatch Formation in the eastern 

part of the basin is gradational with lacustrine rocks of 

the Eocene Green River Formation. Much of the Wasatch 

Formation also is temporally equivalent to rocks between 

the middle and carbonate markers of the Eocene Green 

River Formation that have yielded abundant gas and 

some liquid hydrocarbons in the north-central and 

southeastern parts of the basin (Fouch, 1981; Pitman and 

others, 1982). 

The Tuscher Formation of the Mesa Verde Group 

contains fine-grained gray and brown sandstone (some 

beds are calcareous}, gray-green sandstone, minor fossil 

mollusk debris, and carbonaceous and coaly shale beds. 

The sandstone units are laterally discontinuous and 

formed from sediments deposited in sinuous streams on 

the lower part of an alluvial braidplain (Fouch and others, 

1983). The Tuscher Formation contains large amounts 

of thermogenic gas in the western and central parts of 

the basin where the rocks are deeply buried. In the eastern 

part of the basin, large amounts of early thermal gas are 

stratigraphically trapped in diagenetically altered lenticular sandstones (Pitman and others, 1987). The origin 

and distribution of rocks in the Wasatch and Tuscher Formations at Natural Buttes have been discussed in more 

detail by Pitman and others (1987). 

WHOLE-ROCK MINERALOGY 

On the basis of petrographic and X-ray diffraction 

examination, fine-grained, moderately sorted to wellsorted sandstones in the Wasatch Formation are feldsparrich, ranging from lithic arkoses to feldspathic litharenites 

(fig. 3). Sandstones in the Tuscher Formation generally 

are finer grained than those in the Wasatch and contain 

more quartz. These sandstones range between feldspathic 

litharenites and litharenites (fig. 3). Variations in 

mineralogic composition between the units reflect admixtures of sediment eroded from Laramide highlands that 

lay south and east of the Uinta Basin. Tertiary and 

Cretaceous rocks of interest in this study are similar in 

mineralogic composition to fluvial sandstones studied by 

Keighin and Fouch (1981}, Pitman and others (1982), 

Stanley and Collinson (1979), Zawiskie and others (1982), 

and Lawton (1983). 

Sandstones in the Wasatch and Tuscher Formations 

contain different amounts of framework grains, matrix 

material, pore-filling and replacement cements, and clay 

minerals. Monomineralic quartz, feldspar, and a variety 

of sedimentary lithic fragments and minor igneous 

fragments constitute the detrital fraction of the sandstones. Quartz grains commonly are angular to 

subrounded and contain euhedral syntaxial overgrowths 

Mineralogy and Stable Isotope Geochemistry of Diagenetically Altered Sandstones, Utah 209 
AGE 

IN 

Ma 

45 

50 

u f w 

0 ~ 55 N <( C!1 

0 

~ § ~ u 1- :. w 

~ u 60 § 

:. 

65 

70 

75 

IJl u 80 ::> 

N 

0 ~ w 0 u 1-

IJl 

~ ~ 

~ 85 u 

90 

95 

~ ~ 

AGE 

MIDDLE 

(part) 

EARLY 

LATE 

EARLY 

MAESTRICHTIAN 

CAMPANIAN 

SANTONIAN 

CONIACIAN 

TURONIAN 

CENOMANIAN 

ALBIAN 

(part) 

1 PRICE 

CANYON 

NORTH HORN 

FORMATION 

SUNNYSIDE GREEN RIVER 

4 SEGO 

CANYON 

CEDAR MOUNTAIN FORMATION 

(part) 

5 WESTWATER 

CANYON 

6 SAN AR~OYO 

CANYON 

Figure 2. Correlation of major Tertiary and Cretaceous rock units in Uinta basin, Utah, modified from Fouch and 

others (1983). See figure 1 for location of sections shown in columns 1-8. Drill hole represented in column 5 has 

been projected southward to coeval rocks exposed near Green River. Dark sections of hole represent cored intervals. 

210 Studies in Diagenesis 
Quartz 

Quartzarenite 

• 

Arkose Lithic arkose 

• 

Feldspathic 

litharenite Litharenite 

Feldspar 3:1 1:1 

EXPLANATION 

1:3 Rock fragments 

• Wasatch Formation 

0 Tuscher Formation 

Figure 3. Ternary diagram (modified from Folk, 1974) showing mineralogic composition of sandstones in cored part 

of Wasatch and Tuscher Formations. 

that are embayed due to dissolution associated with 

replacement carbonate and clay minerals. Feldspar is 

dominantly plagioclase, although smaller yet significant 

amounts of potassium feldspar are also present. Feldspar 

grains, particularly the potassium varieties, are weathered 

or show the effects of chemical alteration due to carbonate replacement and dissolution. Sedimentary rock 

fragments are of diverse origin and include sand-size, 

partly abraded carbonate clasts (mostly dolomite), and 

rhombic, detrital(?) dolomite grains which commonly 

contain an anhedral rim of ankerite. 

Authigenic mineral cement, common to most sandstones, includes secondary silica as syntaxial quartz 

overgrowths, a mineralogically complex carbonate 

assemblage, and local barite and anhydrite. Authigenic 

carbonate minerals fill primary and secondary pores and 

replace framework grains, principally feldspar. Pervasive 

dissolution features and relicts of undissolved carbonate 

are evidence that authigenic carbonate formerly occupied 

much of the intergranular and intragranular porosity. 

Authigenic clay minerals in sandstones consist of 

different proportions of kaolinite, illite, mixed-layer 

illite/smectite, and chlorite. Small amounts of corrensite 

are present in sandstones of the Wasatch Formation. 

Kaolinite typically occurs as well-crystallized booklets in 

secondary dissolution pores that are lined with fibrous 

illite or mixed-layer illite/smectite. Chlorite, usually 

associated with illite, forms pseudohexagonal platelets 

Mineralogy and Stable Isotope Geochemistry of Diagenetically Altered Sandstones, Utah 211 
that are oriented perpendicular to grain surfaces. Corrensite may be associated with volcanic lithic fragments. 

CARBONATE MINERALOGY AND DIAGENESIS 

Sandstones in the Wasatch and Tuscher Formations 

in the eastern part of the Uinta basin have had a complex diagenetic history. The formation of several generations of carbonate during burial altered the sandstone 

fabric more than other diagenetic events. X-ray diffraction analysis and thin-section staining for iron content 

revealed a carbonate mineral assemblage composed of 

iron-free calcite, ferroan calcite, dolomite, and minor to 

trace amounts of ankerite. In most sandstones, ferroan 

calcite is the dominant carbonate cement and makes up 

between 0 and 35 percent of the mineral fraction. The 

relative abundance of iron-free calcite was not estimated 

because this mineral occurs only as isolated relicts in 

dissolution pores. Dolomite and small amounts of 

ankerite vary in abundance from 0 to 15 percent. Because 

detrital and authigenic dolomite are complexly 

intergrown, these percentages are only approximate. Textural relations between the individual carbonate phases 

and other mineral constituents in the sandstones are 

shown in figure 4. 

Iron-free calcite was the earliest carbonate to form 

in Tertiary and Cretaceous sandstones and was followed 

by ferroan calcite which is now the most abundant carbonate phase. Iron-free calcite occurs in intergranular 

pores as isolated relicts of grains and cement which commonly are surrounded by ferroan calcite. The iron-free 

calcite commonly displays polysynthetic twinning, 

indicating that the mineral formed soon after deposition. 

During early burial, leaching of iron-free calcite locally 

may have created secondary porosity that was later 

recemented by ferroan calcite. 

Ferroan calcite partly to completely replaces 

framework grains, particularly feldspar and rock 

fragments, and fills intergranular pores. Occurrences of 

ferroan calcite commonly are syntaxial with relict ironfree calcite grains and cement. Where ferroan calcite is 

widespread, it forms a fine- to coarse-crystalline, 

poikilotopic cement that encloses detrital framework 

grains which were typically not subjected to mechanical 

compaction or chemical alteration prior to its emplacement. Within cemented areas, detrital quartz rarely 

exhibits quartz overgrowths and feldspar grains commonly are angular and fresh in appearance, implying that 

they were protected from alteration by chemically active 

pore-fluids. 

The distribution of ferroan calcite is highly variable; 

it is most widespread in sandstones that contain clastic 

carbonate grains or where there is evidence for preexisting 

iron-free calcite cement. Remnants of iron-free calcite 

surrounded by optically continuous ferroan calcite 

212 Studies in Diagenesis 

suggest that ferroan calcite either grew by nucleating on 

detrital carbonate grains and remnants of iron-free calcite 

cement or by incompletely replacing early-formed ironfree calcite. The relative abundance of ferroan calcite in 

Tertiary and Cretaceous sandstones may reflect the 

existence of these nucleation sites for carbonate growth. 

Patchy ferroan calcite, common in many sandstones, suggests that dissolution removed most of the replacment 

cement. 

Authigenic dolomite in the sandstones is relatively 

rare and is distributed as minor anhedral pore-filling and 

replacement cement. Ankerite occurs in relatively small 

amounts as anhedral syntaxial rims surrounding dolomite 

grains. The relations between grains and cement suggest 

that the formation of ankerite required preexisting detrital 

dolomite grains for its growth; most ankerite probably 

formed along with the dolomite clasts from the source 

area. 

Coarse-crystalline anhydrite and barite occur locally 

in sandstones of the Wasatch Formation as replacement 

minerals. Anhydrite typically forms large poikilotopic 

crystals that surround detrital grains and replace carbonate cement and secondary quartz overgrowths. The 

textural relations between the sulfate minerals and the 

other mineral cements suggest that anhydrite and barite 

formed during a later stage of diagenesis, following the 

precipitation of carbonate and silica cement. 

POROSITY DEVELOPMENT 

The relative abundance of ferroan calcite in sandstones appears to reflect initial sandstone porosity and 

permeability. Coarser grained, cleaner sandstones tended 

to be relatively permeable and permitted the circulation 

of large volumes of reactive fluids and, hence, more complete mineral reaction. In contrast, finer-grained sandstones and siltstones typically were less permeable. 

Mineral reactions were slower in these rocks due to 

retarded fluid circulation rates. 

Many sandstones in the Wasatch and Tuscher Formations show moderate to high porosity which resulted 

in part from the dissolution of calcite that formerly 

occupied many intergranular and intragranular pores. 

Partly dissolved feldspar grains containing relicts of ferroan calcite, corroded quartz overgrowth and grain 

boundaries, and isolated remnants of iron-free and ferroan calcite cement in pores are common in many of the 

sandstones and are evidence for dissolution of preexisting 

carbonate. 

A possible explanation for carbonate dissolution 

may be changes in pore-water chemistry that occurred 

when waters of low pH came in contact with the sandstones. Physiochemical conditions that promoted dissolution probably existed during periods of subaerial erosion 

associated with the formation of the Cretaceous-Tertiary 
Figure 4. Photomicrographs of authigenic anhydrite, iron-free and iron-bearing carbonate, and associated dissolution features in the Wasatch and Tuscher Formations. 

A, Early iron-free calcite (c) surrounded by late-formed ferroan-calcite (f). 

8, Early diagenetic, poikilotopic ferroan calcite (c) enclosing detrital unaltered quartz grains. 

C, Detrital abraded dolomite fragments (d) surrounded by optically continuous ankerite (a). 

0, Poikilotopic anhydrite (a) replacing framework grains in mineral cements. 

E, Embayed quartz grain (q) and overgrowth (o) boundaries caused by dissolution of replacement carbonate. 

F, Partly dissolved feldspar grains (f) and lithic fragment (I) that resulted from leaching of replacement carbonate. 

unconformity and during hydrocarbon generation. During periods of regional tectonism, exposed sandstone beds 

that were being eroded were exposed to ground waters 

of low pH which were capable of dissolving carbonate. 

Meteoric waters moved into the subsurface via these 

porous sandstones and possibly incipient fractures where 

they intermingled with migrating formation fluids. Mixing of waters of different chemistry may have created an 

environment at depth favorable for the dissolution of 

carbonate. During diagenesis, the production of C02 

in response to the degradation of organic matter and 

the expulsion of interlayer shale waters during burial 

Mineralogy and Stable Isotope Geochemistry of Diagenetically Altered Sandstones, Utah 213 
compaction also may have contributed to the dissolution 

of carbonate. 

In some sandstones there is no evidence of extensive carbonatization. Either the observed porosity is of 

primary origin, which seems unlikely, or it is secondary 

and was caused by the direct dissolution of framework 

grains, such as feldspar and unstable lithic fragments. 

STABLE ISOTOPES 

Carbon and oxygen isotope compositions were 

determined separately for selected samples of ferroan 

calcite and mixtures of dolomite and subordinate ankerite 

from the Wasatch and Tuscher Formations (table 1). 

o13C values for calcite from the Wasatch range from 

- 5.2 to -10.6 per mil, and o180 values range from -9.4 

to -13.1 per mil. One calcite sample from the Wasatch 

Formation, U80TF249, is unusually light in carbon and 

heavy in oxygen compared with the other samples. 

In the older Tuscher Formation, o13C values for 

calcite range from -3.1 to -8.1 per mil and o180 ranges 

from -9.7 to -14.2 per mil. Two samples, Sandia 8 and 

Sandia 13, have unusually light isotope values relative to 

the other samples, with o13C = -7.6 and -8.1 per mil 

and o180 = -12.9 and -14.2 per mil, respectively. These 

isotope compositions are similar to those of calcite in the 

Wasatch Formation. 

The isotope compositions of dolomite are fairly 

constant for both the Wasatch and Tuscher Formations. 

Values range from -0.5 to -4.2 per mil for o13C and 

from -5.5 to -8.8 per mil for o180. The dolomite fractions for the Sandia 8 and Sandia 13 samples from the 

Tuscher are enriched in the light isotopes of carbon and 

oxygen compared with the other samples. 

A crossplot of o13C and o180 values for the calcite 

and dolomite fractions is shown in figure 5. These data 

reveal that the isotope compositions of calcite in the 

Wasatch (with the exception of sample U80TF249) fall 

along a linear trend that is distinct from the trend along 

which all but the two anomalous calcite values for the 

Tuscher cluster. Isotope values for the dolomite-ankerite 

fraction plot along a separate trend. Distinct isotopic 

trends for calcite -in Tertiary and Cretaceous sandstones 

indicate that the fluids which diagenetically altered these 

sediments had quite different isotope compositions. The 

initial porosity and permeability of individual sandstone 

beds controlled fluid migration pathways and water-rock 

interaction and, thus, the degree of isotope exchange between carbonate and water. 

Carbon isotopes 

Carbon isotope compositions for ferroan calcite in 

the Wasatch Formation are consistently more negative 

214 Studies in Diagenesis 

than those for calcite in the older Tuscher Formation. 

o13C values, from -5 to -10 per mil in the Wasatch, may 

reflect a component of atmospheric carbon, or more 

likely, an influx of organically derived carbon that was 

produced during bacterial degradation of organic matter. Petrographic analysis indicates that most of the carbonate in the Wasatch Formation formed early in the 

burial history; thus, some dissolved carbonate for carbonate precipitation may have been generated from the 

bacterial oxidation of organic matter during the early 

stages of diagenesis. It is also possible that at some later 

diagenetic stage, a small fraction of dissolved HC03

-

formed during nonbacterial (catagenic) methane production; however, this is uncertain inasmuch as sequences 

of lacustrine shale interbedded with Wasatch strata 

generally did not reach the early stages of thermal 

maturity with respect to hydrocarbon generation until late 

in the burial history. 

One calcite sample from the Wasatch Formation 

(U80TF249) contains unusually light o13C (-10.6 per mil) 

compared with the o13C of the other samples and is 

displaced from the o13C versus o180 trend line shown in 

figure 5. This isotope composition may reflect a sandstone characterized by a high water:rock ratio and isotope 

exchange due to an influx of water (possibly from a different source) that contained a large proportion of 

isotopically light carbon, or it may represent a local effect 

of a larger than average abundance of organic matter 

indigenous to that sample. 

Isotopically light o13C compositions for calcite in 

the Wasatch Formation may possibly be related to the 

sulfate mineral anhydrite that occurs locally in some sandstones. Calcite in rocks that contain anhydrite typically 

is displaced toward lighter o13C (and o180) ratios compared with calcite in rocks devoid of anhydrite (fig. 5). 

Markedly light o13C compositions for calcite in sandstones containing anhydrite suggest that the precipitation 

or recrystallization of calcite may have been associated 

with the formation of anhydrite. 

o34S compositions for anhydrite in sandstones with 

associated isotopically light o13C calcite are consistently 

heavy, averaging + 30.3 per mil (table 1). These values 

are close to the o34S compositions reported by Harrison 

and Thode (1958), Mauger (1972), Cole and Picard 

(1981), Smith and Young (1983), and D. E. Anders and 

P. Gerrild (U.S. Geological Survey, Denver, Colorado, 

written commun., 1980) for both organic and pyrite 

sulfur in kerogen-rich rocks of the lacustrine Green River 

Formation. o34S values of anhydrite, however, are 

dissimilar to typical sulfates present in weathering 

environments. The apparent relation between the isotope 

composition of sulfate sulfur in Wasatch beds and 

organic and pyrite sulfur in lacustrine shales of the Green 

River Formation suggests that fluids derived from 

kerogenous shale beds in the Green River may have 

migrated during burial into nearby porous and permeable 
Table 1. Stable isotope data for calcite, dolomite, and anhydrite in sandstones from the Wasatch 

and Tuscher Formations 

[Asterisks (*), anomalous samples; isotope values, per mil; leaders (---), not determined] 

Sample Depth Anhydrite 

No. (feet) o34s 

Wasatch 

U80TF240 4402.7 

U80TF248 4415.7 

*U80TF249 4416.4 

U80TF253 4443.0 

U80TF256 4458.9 

U80TF261 4463.0 +30.8 

U80TF263 4465.0 +28.7 

U80TF258 4470.0 

Sandia 6 4486.8 

U80TF270 4487.0 +31.5 

Tuscher 

U80TF272 5406.0 

U80TF275 6423.7 

U80TF277 6425.0 

*Sandia 8 6469.0 

U80TF279 6472.8 

U80TF284 6487.0 

*Sandia13 6507.0 

U80TF294 6510.0 

strata of the Wasatch Formation. Smith and others 

(1964) showed that sulfur in the Green River occurs 

only in a reduced state, as pyrite and organic sulfur. 

Hence, to derive sulfur for anhydrite, the oxidation 

of reduced sulfur was required to form sulfate. This 

process suggests that oxygenated meteoric waters may 

have played a role in the formation of anhydrite in 

Wasatch sandstones. 

Another possibility is that the oxidation of sulfur 

involved the oxidation of organic matter to soluble 

carbonate species. If this were true, isotopically light 

organic carbon was remobilized into adjacent sandstone beds of the Wasatch Formation and precipitated 

as carbonate. Such a mechanism could explain the 

isotopically light carbonate in sandstone beds containing anhydrite and can also account for the formation of 

a smaller component of isotopically light calcite in sandstones devoid of anhydrite. 

Most o13C compositions for calcite in the Tuscher 

Formation are isotopically heavy (near -3 per mil) compared with those for calcite in the Wasatch Formation, 

indicating that dissolved biogenic carbon was only a 

Calcite Dolomite 

13 

o cPDB o18oPDB 13 o cPDB 18 o oPDB 

Formation 

-6.4 -12.4 -2.0 -7.0 

-7.0 -12.5 -4.2 -7.9 

-10.6 -9.4 -1.7 -6.3 

-6.5 -11.8 -2.4 -7.9 

-5.3 -10.7 -1.5 -6.9 

-7.6 -12.6 -3.7 -8.8 

-8.5 -13.1 -3.1 -8.0 

-5.2 -10.8 -1.3 -6.8 

-5.3 -10.9 -2.3 -7.9 

-7.9 -12.1 -0.8 -6.1 

Formation 

-3.7 -10.9 -1.3 -6.6 

-3.4 -9.8 -0.5 -5.5 

-3.9 -11.6 -0.8 -6.0 

-7.6 -12.9 -3.8 -12.4 

-3.1 -9.9 -1.3 -7.5 

-3.1 -9.7 -1.1 -7.1 

-8.1 -14.2 -8.2 -18.0 

-3.1 -10.5 -0.7 -6.8 

minor carbon source for these carbonates. Similar o13C 

values throughout the Tuscher suggest that the calcite 

precipitated from a large reservoir of dissolved carbonate 

of nearly constant isotope composition. The isotopically 

heavy carbon can be attributed largely to dissolved carbonate species of marine origin that were derived locally 

from depositional waters. Calcite with light carbonisotope signatures typically is absent in the Tuscher 

despite the presence of interbedded carbonaceous and 

coaly shales. 

o13C and o180 compositions for the two calcite 

samples in the Tuscher Formation (Sandia 8 and Sandia 

13) that are isotopically light compared with those of the 

other samples, closely correspond to the carbon and 

oxygen values that characterize the average isotope composition of calcite in the Wasatch Formation. 

Petrographic analysis revealed that isotopically light 

calcite in the Tuscher formed during early diagenesis in 

sandstones that had high initial porosity and permeability. 

Thus, this early calcite probably formed locally from 

reactive fluids that were similar in isotope composition 

to those that deposited calcite in the Wasatch. 

Mineralogy and Stable Isotope Geochemistry of Diagenetically Altered Sandstones, Utah 215 
....I 

~ 

0::: 

w 

a.. 

u 

CW) 

-1 

GO -6 

-7 

-11 

-15 -14 -13 -12 -11 -10 -9 -8 -7 -6 -5 -3 

81So , PER MIL 

EXPLANATION 

Wasatch Formation Tuscher Formation 

0 Calcite • Calcite 

(]) Calcite and associated • Anomalous calcite 

anhydrite • Dolomite/ ankerite 

!:;,. Anomalous calcite 

0 Dolomite/ ankerite 

Figure 5 Crossplot showing o13C versus o180 for carbonate grains and cement in sandstones from the Tuscher and • PDB PDB 

Wasatch Formations. 

Oxygen isotopes 

The o180 compositions of calcite in sandstones of 

the Wasatch and Tuscher Formations are similar to those 

of diagenetic carbonate cement in sandstones such as have 

been reported by Hudson (1977) and Pitman and others 

(1982). The range, between -9.4 and -14.2 per mil, can 

be attributed either to the influence of isotopically light 

meteoric waters during diagenesis (Rothe and others, 

1974) or to variations in temperature with increasing 

depth of burial which caused a decrease in the fractionation between the oxygen of carbonate ions and water. 

The o180 compositions for calcite in the Wasatch 

are, on the average, nearly identical to those for calcite 

in the older Tuscher Formation. The expected general 

trend of decreasing o180 compositions with increasing 

burial depth and age of rock is not apparent, suggesting 

that temperature was not the major influence governing 

216 Studies in Diagenesis 

the isotope composition of oxygen during carbonate 

diagenesis. If temperature had been an important factor, 

o180 values for calcite should be systematically lighter 

with increasing depth of burial. An inverse relation between o180 composition and burial depth suggests that 

a flux of isotopically light pore waters through permeable 

horizons was a more important factor than depth of 

burial and temperature in determining 180 systematics. 

The two calcite samples from the Tuscher with 

unusually light o180 compositions (-12.9 and -14.2 

per mil) and the three samples from the Wasatch 

containing anhydrite probably reflect enhanced interaction with meteoric water related to high initial sandstone 

permeability coupled with reactions involving an input 

of organically derived aqueous COr Much of the 

Wasatch apparently was influenced by meteoric water, 

inasmuch as the isotope compositions of calcite in sandstones with associated anhydrite plot on the same linear 

trend as those for calcite in sandstones with no anhydrite. 
Dolomite 

The trend of o13C and o180 ratios for dolomite in 

Tertiary and Cretaceous sandstones extrapolates close to 

the approximate isotope composition of average marine 

dolomite (near zero per mil; fig. 5). Petrographic analysis 

indicated that the dolomite fraction in both formations 

contains abundant detrital dolomite and minor authigenic 

dolomite and ankerite. The similar isotope compositions 

of dolomite-ankerite in the two formations are attributed 

principally to a detrital component of these minerals common to both. 

The two dolomite samples in the Tuscher Formation (Sandia 8 and Sandia 13) that are isotopically light 

relative to the other samples may reflect the presence of 

early authigenic dolomite and possibly ankerite, inasmuch 

as each contains an unusually high amount of these 

minerals. 

SUMMARY 

The formation of a carbonate mineral assemblage 

composed of iron-free calcite, ferroan calcite, dolomite, 

and ankerite significantly modified Tertiary and 

Cretaceous sandstones in the Uinta basin, Utah. Iron-free 

calcite occurs as relicts and was the earliest carbonate to 

form. Ferroan calcite followed the precipitation of ironfree calcite and now is the dominant cement. Commonly, 

ferroan calcite forms a poikilotopic cement that occludes 

pores and replaces framework grains. Locally dissolved 

ferroan calcite reflects either a late stage of diagenesis due 

to the expulsion of interlayer shale waters, the production of C02 during degradation of organic matter, or 

the influx of meteoric water during a period of erosion. 

Dolomite and ankerite occur principally as detrital grains 

and less commonly as authigenic pore-filling cement. 

Stable isotope analysis of carbonate grains and 

cement indicate that calcite in the Tertiary Wasatch Formation formed from fluids that were of lighter isotope 

composition than fluids that precipitated calcite in the 

older Cretaceous Tuscher Formation. Isotopically light 

carbon in calcite from both formations is attributed 

primarily to an influx of organic carbon derived from 

nearby beds of lacustrine shale in the Green River Formation. o34S compositions of anhydrite from sandstones 

in the Wasatch that contain calcite enriched in 13C support this interpretation, inasmuch as they are similar to 

compositions reported for organic and pyrite sulfur in 

the Green River Formation. Isotopically heavy o13C 

values approach zero per mil in the Tuscher, reflecting 

a marine origin except locally where there was an influx 

Of isotopically light meteoriC Water. o18Q ValueS for 

calcite in the Wasatch are similar to those for calcite in 

the Tuscher, suggesting that variations in pore-water composition rather than the effects of temperature with 

increasing burial depth controlled oxygen composition. 

Carbon and oxygen isotope ratios of dolomite in the 

Wasatch and Tuscher Formations approximate the trend 

of marine dolomites. This relationship suggests that a 

component of detrital dolomite is common to both 

formations. 

REFERENCES CITED 

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. 

Folk, R. L., 1974, Petrology in sedimentary rocks: Austin, 

Texas, Hemphill's Publishing Company, 182 p. 

Fouch, T. D., 1975, Lithofacies and related hydrocarbon 

accumulations in Tertiary strata of the western and central Uinta basin, Utah, in Bolyard, D. W., ed., Symposium 

on deep drilling frontiers in the central Rocky Mountains: 

Rocky Mountain Association of Geologists Bulletin, 

p. 163-173. 

1981, Chart showing distribution of rock types, 

lithologic groups and depositional environments for some 

lower Tertiary and Upper Cretaceous rocks from outcrops 

at Willow Creek-Indian Canyon through the subsurface of 

the Duchesne and Altamont oil fields, southwest to northcentral parts of the Uinta basin, Utah: U.S. Geological 

Survey Oil and Gas Investigations Chart OC-81, 2 sheets. 

Fouch, T. D., Lawton, T. F., Nichols, D. 1., 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, M., and Dolly, E., eds., 

Mesozoic paleogeography of west-central United States: 

Society of Economic Paleontologists and Mineralogists, 

Rocky Mountain Section Symposium 2, p. 305-334. 

Harrison, A. G., and Thode, H. G., 1958, Sulphur isotope 

abundances in hydrocarbons and source rocks of Uinta 

Basin, Utah: American Association of Petroleum 

Geologists Bulletin, v. 42, p. 2642-2649. 

Hudson, J. D., 1977, Stable isotopes and limestone lithification: Geological Society of London Journal, v. 133, 

p. 637-660. 

Keighin, C. W., and Fouch, T. D., 1981, Depositional 

environments and diagenesis of some nonmarine Upper 

Cretaceous reservoir rocks, Uinta basin, Utah, in Ethridge, 

F. G., and Flores, R. M., eds., Recent and ancient nonmarine depositional environments; Models for exploration: 

Society of Economic Paleontologists and Mineralogists 

Special Publication 31, p. 109-125. 

Land, L. S., 1980, The isotopic and trace element geochemistry 

of dolomite; The state of the art: Society of Economic 

Paleontologists and Mineralogists Special Publication 28, 

p. 87-110. 

Lawton, T. F., 1983, Late Cretaceous fluvial systems and the 

age of foreland uplifts in central Utah, in Lowell, J.D., 

and Gries, R., eds., Rocky Mountain Foreland basin and 

uplifts: Denver, Colorado, Rocky Mountain Association 

of Geologists, p. 181-199. 

Mineralogy and Stable Isotope Geochemistry of Diagenetically Altered Sandstones, Utah 217 
Mauger, R. L., 1972, A sulphur isotope study of bituminous 

sands from the Uinta basin, Utah: 24th International 

Geological Congress Proceedings, Comptes Rendus, sec. 5., 

p. 19-27. 

Pitman, J. K., Fouch, T. D., and Goldhaber, M. B., 1982, Depositional setting and diagenetic evolution of some Tertiary 

unconventional reservoir rocks, Uinta basin, Utah: 

American Association of Petroleum Geologists Bulletin, v. 

66, p. 1581-1596. 

Pitman, J. K., Anders D. A., Fouch T. D., and Nichols D. J., 

1987, Depositional environments, diagenesis, and hydrocarbon potential of nonmarine Upper Cretaceous and lower 

Thrtiary rocks, eastern Uinta basin, Utah, in Spencer, C. W., 

and Mast, R. F., eds., Geology of tight gas reservoirs: 

American Association of Petroleum Geologists Studies in 

Geology 24, p. 235-252. 

Rothe, P., Hoefs, J., and Sonne, V., 1974, The isotopic composition of Thrtiary carbonates from the Mainz basin, an example of isotopic fractionations in "closed basins": Sedimentology, v. 21, p. 373-395. 

Ryder, R. T., Fouch, T. D., and Elison, J. H., 1976, Early Ter218 Studies in Diagenesis 

tiary sedimentation in the western Uinta Basin, Utah: 

Geological Society of America Bulletin, v. 87, p. 496-512. 

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. 

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

of sulfur isotopes in Colorado corehole number 1, in Gary, 

J. H., ed., Sixteenth oil shale symposium proceedings: 

Golden, Colorado, Colorado School of Mines, p. 176-188. 

Stanley, K. 0., and Collinson, J. W., 1979, Depositional history 

of the Paleocene-lower Eocene Flagstaff Limestone and 

coeval rocks, central Utah: American Association of 

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

Walters, L. J., Jr., Claypool, G. E., and Choquette, P. W., 1972, 

Reaction rates and o180 variation for the carbonatephosphoric acid preparation method: Geochimica et 

Cosmochimica Acta, v. 36, p. 129-140. 

Zawiskie, J., Chapman, D., and Alley, R., 1982, Depositional 

history of the Paleocene-Eocene Colton Formation, northcentral Utah: Utah Geological Association Bulletin, v. 10, 

p. 273-284. 
Diagenetic Relationships in a 

Hydrocarbon-Productive Chalk-the 

Cretaceous Niobrara Formation 

By RICHARD M. POLLASTRO, U.S. Geological Survey, Box 25046, 

Denver Federal Center, MS 960, Denver, Colorado 80225, and 

PETER A. SCHOLLE, Southern Methodist University, Dallas, Texas 75275 

CONTENTS 

Abstract 221 

Introduction 221 

Acknowledgments 222 

Stratigraphy and hydrocarbon production 222 

Methods of study 222 

Samples 222 

Analytical methods 223 

Results and discussion 224 

Porosity, chalk fabric, and oxygen isotopes 224 

Mineralogy 226 

Detrital-diagenetic relations 226 

Nature and origin of regularly interstratified chlorite/smectite 230 

Geologic and economic implications 230 

Summary and conclusions 233 

References cited 233 

FIGURES 

1. Index map and cross section of wells studied, Denver basin, eastern 

Colorado 223 

2. Plots showing weight percentage of insoluble residue as a function of 

frequency in chalk samples 224 

3. Plots showing porosity versus depth in cores studied 225 

4. Scanning electron micrographs showing progressive cementation during 

burial 226 

5-7. Plots of: 

5. Whole-rock oxygen isotopes versus burial depth 227 

6. Composition of mixed-layer illite/smectite in Pierre Shale and 

insoluble residue from chalk of Niobrara Formation versus burial 

depth 228 

7. Composition of mixed-layer illite/smectite versus depth in residues 

and bentonites 229 

8-10. X-ray diffraction traces of: 

8. Mixture of random and ordered illite/smectite in Niobrara insoluble 

residue 230 

9. Kaolinitic bentonites showing presence of regularly interstratified 

mixed-layer illite/smectite and chlorite/smectite in Niobrara strata 231 

10. <2-J.tm fraction of aluminous bentonite showing evidence for the 

presence of interlayer aluminum complexes 232 

F. A. Mumpton, eel., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 219 
11. Regional petrophysical facies map of Niobrara Formation 233 

12. Paleotherm map of Niobrara Formation in the Denver basin based on 

degree of ordering of mixed-layer illite/smectite in bentonites 234 

TABLE 

1. Names, locations, status, core intervals, and bottom-hole temperatures of 

wells studied 222 

220 Studies in Diagenesis 
Abstract 

An understanding of diagenesis is critical in predicting 

reservoir or source-rock potential of the Niobrara Formation 

because the primary properties of the chalks and calcareous 

shales have been greatly modified by burial. Porosity and 

isotope composition of the chalks are consistently related to 

burial depth (or paleoburial depth) and thermal history. Initial 

porosities of 60 percent or more have been reduced by progressive cementation to about 25 percent with 1 km of burial 

and to less than 10 percent with burial depth of about 2 km 

or more. The oxygen isotope composition of the whole rock 

reflects these changes with shifts to progressively more negative 

o180 values with greater burial. 

The clay mineral composition and degree of ordering of 

mixed-layer illite/smectite (1/5) clay in bentonite beds and 

insoluble residues from the chalks can also be related to 

porosity loss due to increasing temperatures with burial depth. 

Although the starting composition of 1/S clay originally deposited in chalk residues was highly variable, the expandability and range of expandability of the 1/S became increasingly 

more limited with greater burial. A systematic increase in the 

percentage of illite-like layers in the 1/S of bentonites as well 

as a progressive increase in the degree of ordering of 1/S was 

noted in more deeply buried samples. Regularly interstratified 

chlorite/smectite (C/5) was also formed in aluminous bentonites 

at elevated temperatures and pressures during deep burial 

(>100°C). The composition and degree of ordering of 1/S and 

the occurrence of CIS in bentonites are in good agreement with 

the level of maturation of kerogen and hydrocarbons within 

the Niobrara Formation, Denver basin. The clay minerals can, 

therefore, be used as geothermometers for predicting indigenous biogenic gas and thermogenic oil from the Niobrara. 

These depth-dependent relationships also suggest that parts of 

the Denver basin were more deeply buried in the past than 

formerly thought and (or) had paleogeothermal gradients higher 

than those of today. 

INTRODUCTION 

Natural gas and oil are produced from the organicrich chalk beds of the Upper Cretaceous Niobrara Formation in eastern Colorado and northwestern Kansas. 

Because of the increased production of oil and gas from 

these pelagic sediments during the past 15 years (Scholle, 

1977a, 1977b; Pollastro and Scholle, 1984), the diagenetic 

alteration of chalk has received considerable attention. 

Although most scientific studies have focused on the 

diagenesis of the chalk itself, few have looked at the 

diagenesis of clay minerals or other inorganic components 

within insoluble residues or bentonites from the chalk. 

Chalks are fine-grained limestones (micrites) composed largely of calcareous nannofossils, called coccoliths, and lesser amounts of calcareous microfossils, 

such as Foraminfera and calcispheres. Chalk forms only 

where organisms that form these calcareous tests can live 

in the overlying water and where the rate of input is low 

from clastic terrigenous material or other carbonate 

grains which might otherwise overwhelm and mask the 

nannofossils and microfossils (Scholle, 1977a). 

Chalks formed from coccoliths and tests of planktonic Foraminifera are composed of chemically stable 

low-magnesium calcite and have inherently low 

permeability due to their very small average grain size. 

These two factors combine to make chalk fairly stable 

under near-surface conditions. Chalk oozes can be 

deposited with as much as 60-80 percent porosity. 

Dewatering and initial mechanical compaction (grain 

reorientation and breakage) of a chalk ooze typically 

reduce this initially high porosity to about 50 percent 

under relatively shallow burial (Schlanger and Douglas, 

1974). Scholle (1977a) showed that further loss in porosity 

in chalk is due to pressure solution of calcium carbonate 

at grain contacts and along solution seams as the chalk 

is subjected to an increase in lithostatic pressure due to 

depth of burial. This calcium carbonate is then 

reprecipitated locally at sites of lower differential stress 

(such as open pores). Permeability in chalks also decreases 

linearly with porosity (Scholle, 1977a). These relationships have been demonstrated in the Niobrara Formation 

by Lockridge and Scholle (1978). 

The chalks of the Niobrara Formation contain a 

relatively large amount of insoluble residue. The insoluble 

residue consists chiefly of both detrital and authigenic clay 

minerals, quartz, pyrite, plagioclase feldspars, and organic 

matter (Pollastro, 1981a; Pollastro and Martinez, 1985). 

Numerous bentonite beds are also common throughout 

most of the Niobrara strata (Hattin, 1976, 1981; Kauffman, 

1977; Pollastro, 1981a, 1981b). Studies of clay-rich rocks 

have demonstrated specific depth-dependent mineralogical 

changes (see reviews by Weaver, 1979; Hower, 1981). In particular, clay-mineral assemblages and their extent of diagenetic reaction have been used as geothermometers for 

determining maximum paleotemperature in sedimentary 

rocks (Hoffman and Hower, 1979; Pollastro and Bader, 

1983; Pollastro and Scholle, 1984). The most significant 

reaction within the clay-mineral assemblages of deeply 

buried clay-rich rocks is the conversion of smectite to illite 

through a mixed-layer illite/smectite (liS) series. Pollastro 

(1985) discussed the nature of the I/S reaction within the 

Niobrara of the Denver basin. 

The present paper focuses on post-depositional 

diagenetic relations within a single chalk lithofacies of the 

Niobrara Formation, Denver basin, Colorado. In the cores 

studied, the petrophysical properties, fabric, and oxygen 

isotope composition of the chalk are consistently related 

to burial depth. In addition, the composition and ordering of I/S clay show a similar progression of diagenesis 

with burial and temperature (or paleoburial and paleotemperature). These relationships, when combined with studies 

of organic geochemistry and maturation, demonstrate how 

chalk and clay-mineral diagenesis can be used as a predictive tool in regional paleogeographic and economic 

reconstructions. 

Diagenetic Relationships in Hydrocarbon-Productive Chalk, Niobrara Formation 221 
Acknowledgments 

The manuscript benefited greatly from critical 

reviews by Leonard G. Schultz and Donald L. Gautier. 

We also thank Dudley Rice and the late John Hower for 

their helpful discussions. 

STRATIGRAPHY AND HYDROCARBON 

PRODUCTION 

Chalk and chalky shale of the Niobrara Formation 

(Coniacian-Campanian) were deposited over a wide area 

of the Western Interior of the United States in a relatively 

shallow epicontinental seaway during an overall transgressive episode. The Niobrara Formation of eastern Colorado and adjacent areas consists of two members: the 

lower, Fort Hays Limestone Member consists of 12-26 m 

of highly bioturbated chalk and chalky shale; the upper 

and major part of the Niobrara Formation, represented 

by the Smoky Hill Chalk Member, consists of about 150-

200 m of gray to dark-gray chalky shale and locally 

massive chalk beds apparently deposited under poorly 

oxygenated conditions (Arthur and others, 1985). 

Gas production comes from relatively pure chalk 

beds (that is, containing a small acid-insoluble fraction) 

at the top of the Smoky Hill Member. The productive 

interval is informally referred to as the "Beecher Island 

zone" (Lockridge, 1977). This uppermost, relatively 

"clean" chalk unit, ranging from about 6 to 15 m in 

thickness, was deposited over a large region. Using subsurface geophysical logs, Hann (1981) correlated the 

Beecher Island zone from outcrops along the Front Range 

near Fort Collins, Colorado, across the Denver basin to 

the outcrops in west-central Kansas described by Hattin 

and Cobban (1977). Biostratigraphic and radiometric 

studies of such widespread lithofacies of the Niobrara 

cyclothem are believed to represent time-equivalent 

sediments (Kauffman, 1977). These uppermost chalk beds 

are overlain by a thick section of bentonitic Pierre Shale 

that serves as the seal for the reservoir. Most of the 

hydrocarbons produced from the Niobrara are believed 

to have been generated from type II organic matter within 

the Niobrara itself (Rice and Claypool, 1981; Rice, 1984). 

METHODS OF STUDY 

Samples 

This paper contains data on Niobrara samples collected from numerous well cores and several outcrops. 

Most of the data, however, were obtained from core 

samples of five wells recovered from three specific depths 

along the eastern flank of the Denver basin, eastern 

Colorado. The origin, geochemistry, and maturity of 

organic matter and hydrocarbons from the cores of these 

same wells were described by Rice and Claypool (1981) 

and Rice (1984). 

The names, locations, uncorrected bottom-hole 

temperatures, and depths of the cores recovered from 

these wells are listed in table 1. Their approximate locations and a simplified cross-section are shown in figure 

1. The cores recovered from these five wells only included 

the hydrocarbon-productive ''Beecher Island zone.'' 

About 150 samples of chalk and chalky shale and 

15 samples of bentonite beds were collected from core 

from the five wells listed in table 1. Analytical work consisted of porosity measurements, examination of chalk 

textures and mineralogical components by scanning electron microscopy (SEM), HCl-insoluble residue determinations, whole-rock oxygen isotope analyses, and 

X-ray powder diffraction (XRD) analyses of whole-rock 

Table 1. Names, locations, status, core interval, and uncorrected bottom-hole log temperatures of wells studied 

Depth Maximum 

of recorded 

Letter Company and Location Borehole cored bottom-hole 

(fig. 1) well name status interval temperature 

(meters) (degrees C) 

A Kansas-Nebraska Natural Gas, Sec. 32, T. 2 S., R. 43 w.' Gas well---- 447-463 36 

Whomble No. 1-32. Yuma County, Colo. 

B Mountain Petroleum, Sec. 29, T. 2 s.' R. 43 w., Gas well--- 445-452 32 

State No. 1-29. Yuma County, Colo. 

c Kansas-Nebraska Natural Gas, Sec. 33, T. 3 N., R. 47 w., Dry and 853-871 47 

Powell No. 1-33. Yuma County, Colo. abandoned. 

D J-W Operating, Sec. 29, T. 4 N., R. 46 w.' Gas well---- 842-863 46 

Brophy No. 1. Yuma County, Colo. 

E Excelsior Oil, Sec. 14, T. 4 N., R. 58 w., Oil well---- 1 '655-1 ,667 74 

Alice G. Nay No. 1. Weld County, Colo. 

222 Studies in Diagenesis 
en 

a: 

w 

1-

w 

::E 

~ 

:i 

t 

w 

c 

tl 

L

l NIOBRARA CORES 

WYOMING .~ 

jcOwRADQ---·--E~··· .-.ff!. . . .· I 0 I 

I OENVER c B~ I BASIN A•t 

I I KANSAS 

I I 

I I 

I I 

-f-----------.-j_ NEWMEXICO I 

1000 

1500 

100 

B A 

{Beecher Island zone) 

Well {see table 1) 

A Kansas-Nebraska Natural Gas, 

Whomble No. 1-32 

B Mountain Petroleum, 

State No. 1-29 

C Kansas-Nebraska Natural Gas, 

Powell No. 1-33 

D J-WJ Operating, 

Brophy No.1 

E Excelsior Oil, 

Alice G. Nay No. 1 

200 KILOMETERS 

I CORED INTERVAL 

Figure 1. Index map and cross section of wells studied, Beecher Island zone, Niobrara Formation, Denver basin, eastern 

Colorado. 

chalk, shale, and bentonite, chalk insoluble residue, and 

the <2-J.tm and <0.5-J.tm oriented fractions of the residues, 

shales, and bentonites. In addition, about 200 samples 

of bentonite were analyzed from outcrop and other core 

of the Niobrara in the Denver basin and adjacent areas. 

Analytical methods 

About 140 porosity measurements were made on 

2.5-cm (l-inch) coreplugs by bulk-density versus volume 

analyses. Samples collected from cores and outcrops for 

mineralogy and geochemistry were washed to remove 

contaminants, dried, and ground to <35 mesh (0.5 mm). 

Several splits of each sample were taken for various 

analysis. 

Calcite was dissolved in 1 N HCl; the residue was 

filtered and washed immediately after effervescence 

stopped to minimize solution of other minerals. The 

insoluble residue was dried overnight at 65 OC, the weight 

percentage of the residue was determined, and a small 

portion of the residue was spot-checked with 6 N HCl 

for undissolved carbonate. 

The mineralogy of the acid-insoluble residues, 

shales, and bentonite beds was first established by XRD 

analyses of randomly oriented whole rock or residue 

material and oriented aggregates of the <2-J.tm (equivalent 

spherical diameter) fraction. Oriented mounts were prepared by a modified filter-membrane-peel technique 

(Pollastro, 1982). Semiquantitative XRD analysis of the 

whole rock or whole insoluble residue was made by comparing the peak intensities for unknown samples with 

Diagenetic Relationships in Hydrocarbon-Productive Chalk, Niobrara Formation 223 
those from prepared mixtures of concentrated Niobrara 

material, insofar as possible, and following the procedures outlined by Schultz (1964) and Hoffman (1976), 

with minor modification. Semiquantitative analyses of 

the clay fraction were made following the procedure 

outlined by Schultz (1964). 

Composition and ordering of I/S clay was determined on oriented, ethylene glycol-saturated specimens 

of the <0.5-J.Lm fraction. The patterns were interpreted 

using the methods of Reynolds and Hower (1970). Ordering types were defined using the "Reichweite" (R) notation, as described by Reynolds (1980), where "R" 

signifies the most distant layer in an interstratified 

sequence that affects the probability of occurrence of the 

final layer. Diffractograms indicate that most naturally 

occurring I/S falls into one of three ordering types of 

interstratification (Reynolds and Hower, 1970), namely, 

(1) random I/S (R = 0), (2) allevardite or short-range 

ordered liS (R= 1), and (3) Kalkberg or long-range 

ordered I/S (R~3). 

Whole-rock oxygen isotope analysis was performed 

using standard techniques (Craig, 1957). Isotope results 

are reported in per mil deviation from the PeeDee belemnite (PDB) using the standard o notation: 

0180 = IBOfi60 sample - 180/160 standard X 1000. 

180/160 standard 

Relative increases of 180/160 (increased values of o180) 

are referred to in this paper as "more positive" or 

"heavier". Analogously, decreases in the 180/160 ratio 

are termed "more negative" or "lighter". Studies in our 

laboratory and those in cooperation with M.A. Arthur 

(University of Rhode Island, Kingston, Rhode Island) 

have found that the most reliable whole-rock oxygen 

isotope values for evaluating chalk diagenesis are those 

determined from chalks containing less than 35 weight 

percent insoluble residue. Therefore, all isotope analyses 

were determined from chalks having this restriction. 

RESULTS AND DISCUSSION 

Porosity, chalk fabric, and oxygen isotopes 

The acid-insoluble residue of 156 core samples from 

the five wells from the Beecher Island zone (table 1) 

averaged 30 percent by weight. The range of the percentage of insoluble residue as a function of frequency is 

shown in figure 2. Variations in the insoluble residue content within the chalk can be explained by fluctuations in 

the amount of both terrigenous and volcanic detritus that 

was added to the sediment (Pollastro, 1981a) and by temporal variations in biological productivity in the 

Cretaceous water column (Arthur and others, 1985). 

224 Studies in Diagenesis 

0 

10 

20 (/) 

w 

..J 

c.. 

~ 30 <( 

(/) 

LL 

0 

a: 

w 40 al 

~ 

:::::> 

z 50 

60 

70 1-10 11-20 21-30 31-40 41-50 51-60 61-70 

INSOLUBLE RESIDUE, IN WEIGHT PERCENT 

Figure 2. Frequency plot of HCI-insoluble residue from 156 

core samples of Beecher Island zone chalk, Niobrara 

Formation. 

Measurements on 142 samples from the five wells 

described in this report (fig. 3) show that the average 

porosity of the Beecher Island zone decreases from 39 

percent in the shallowest wells at about 460 m depth to 

27 percent at about 850 m depth, to 9 percent in the Alice 

G. Nay well at 1,650 m depth. These porosities are in 

good agreement with those calculated from borehole density logs by Lockridge and Scholle (1978, fig. 6). Porosity 

loss and alteration of the original chalk fabric due to progressive cementation during burial diagenesis are 

illustrated in the scanning electron micrographs (fig. 4) 

of relatively clean chalk from the three depths. Because 

the amount of insoluble residue in these chalks affects 

the porosity and chalk fabric (Scholle, 1977a; Pollastro 

and Martinez, 1985), only chalks containing about 15 

weight percent insoluble residue were used in the comparison. The chalk from the Whomble well at 460 m 

depth (fig. 4A) contained coccoliths and rhabdolith spines 

which showed little or no alteration of their primary 

biogenic textures by cementation. Extremely high 

microporosity (40 percent) was noted at shallow depth. 

The chalk from the Powell well at 850 m (fig. 4B) was 

slightly more altered by pressure solution and cementation. Some of the fossil grains in this sample were 

interpenetrated due to chemical compaction. Rhombs of 

sparry calcite cement as overgrowths were also present; 

the porosity was about 30 percent. The typical fabric of 

the chalk from the Alice G. Nay well at 1,650 m depth 

is shown in figure 4C. Euhedral crystals produced by 

overgrowth cementation are ubiquitous and have reduced the matrix porosity to about 9 percent. Only a 
0.---------~----------~----------~----------.-------~o 

'\7 .... , ..... #. ·~- /6 500 

/ 

2000 / 

/ 

/ 

/ en I- a: '\7/ w w w .,.,,,411fll ... • I- LL. • w 

~ /6 ~ 

ui / ~ 

a: 1000 ui 0 / a: u / 0 

u 0 

I- / 0 

I- :I: 

I- 4000 / :I: a.. / I- w a.. Cl / w 

Cl 

/ 

/ 

/ 

/ 

/ 1500 

'V/ '\7 __ 

.... Average values 

6 --6 

6000 

2000 

0 10 20 30 40 

POROSITY, IN PERCENT 

Figure 3. Plot of porosity versus depth in the Beecher Island zone of the Niobrara Formation 

in the five wells studied. Porosity values were determined from 2.5-cm (1-inch) core plugs by bulkdensity versus volume measurement. Dashed line connects average porosity from each core group. 

Analysis performed on total of 142 samples. 

few recognizable traces of the original biogenic grains 

were noted. 

Plots of porosity versus insoluble residue content 

in chalks of the Niobrara show that porosity is higher 

in relatively clean carbonate beds than in beds with a high 

shale content (Pollastro and Martinez, 1985). At greater 

burial depths, pressure solution (that is, carbonate 

dissolution and transfer) was more active in beds with 

high shale content. As in stylolite formation, the dissolved 

carbonate from shaly beds in chalk was probably 

redeposited in the cleaner, adjacent limestone. Upon continued burial and cementation, this process further 

depleted shaly beds of carbonate and further enriched the 

carbonate-rich chalk beds. This process may have continued until the porosity of chalk was reduced to the 

extent at which the original porosity relations were 

reversed; that is, the adjacent shales may retain higher 

clay microporosities than the adjacent "clean" chalk carbonate at great depths. 

Figure 5 shows the relations between present burial 

depth and whole-rock oxygen isotope values from 32 

samples of chalk. The dashed line connects the average 

porosity from each core-depth group of samples. Oxygen 

isotope values are progressively more negative (or lighter) 

with greater burial and porosity loss. These values are 

a function of the isotope composition of the water and 

the temperature at which the carbonate cements were precipitated (see Garlick, 1969 for review of oxygen isotope 

analysis). The observed shift in isotope values was a result 

of the dissolution of coccoliths and Foraminifera and the 

reprecipitation of this carbonate as cement from pore 

fluids at elevated subsurface temperatures. These shifts 

presumably took place in a semi-closed system (that is, 

one dominated by expulsion of in-place pore fluids). 

Diagenetic Relationships in Hydrocarbon-Productive Chalk, Niobrara Formation 225 
Figure 4. Scanning electron micrographs showing progressive 

cementation of Niobrara Formation chalk during burial. All 

samples contain about 15 weight percent HCI-insoluble residue. 

Typical fabric of chalk from depths of: A, 460 m, porosity about 

40 percent; 8, 850 m, porosity about 30 percent; C, 1,650 m, 

porosity about 9 percent. 

Initial oxygen isotope values of modern marine carbonate 

deposits range from + 2.0 to -2.0 per mil (PDB) 

(Milliman, 1974). The unusually negative starting values 

(projected to zero burial) of the Niobrara most likely 

226 Studies in Diagenesis 

reflect the influence of "fresh", isotopically "light" surface waters draining into the restricted epicontinental 

seaway (Arthur and others, 1985). Oxygen isotope 

analyses of chalks, combined with other geochemical 

data, support the concept that cementation was a gradual 

process throughout the burial history of the rock and that 

these cements were derived internally by selective dissolution (Scholle, 1977a). 

Mineralogy 

Detrital-diagenetic relations 

The composition and degree of ordering of I/S in 

insoluble residues of chalk and chalky shale from the 

Niobrara Formation and shale and calcareous shale from 

the Pierre Shale (Whomble No. 1 well) are shown in figure 

6 along with the composition of the I/S from bentonite 

beds. Fabric, porosity, and whole-rock oxygen isotope 

values of the chalk from the Whomble well indicate that 

the core was not buried to depths much greater than present; therefore, the original clay mineralogy and claymineral composition of liS in these strata probably did 

not change significantly. Core from the Whomble No. 

1 well included about 7 m of the lowermost Sharon 

Springs Member of the Pierre Shale, which immediately 

overlies the Beecher Island zone of the Niobrara Formation. The approximate contact, or transition, between the 

two formations is shown by the dashed line in figure 6. 

The left part of figure 6 represents the weight percentage 

of insoluble residue, or the "shaliness" of the sample. 

The right part shows the percentage of expandable layers 

(or percentage of illite layers) in liS and the general 

nature of interstratification of the liS. 

Illite/smectite in bentonites from both the Sharon 

Springs shale and Beecher Island zone chalk contained 

about 10 percent illite layers. The liS in shale from the 

Sharon Springs Member was randomly interstratified 

(R = 0) and had a composition range of 20-55 percent 

illite layers. The more illitic I/S clay of the Sharon Springs 

Member occurred at the transition between the two formations. The liS in the residues of chalk from the 

Niobrara were typically more illitic and ordered (40-70 

percent illite layers) than the I/S of the overlying Pierre 

Shale. Exceptions (for example, sample 32) include zones 

of chalk which contained a significant admixture of 

altered volcanic ash (see Pollastro, 198la). The relationship between the composition of I/S in the more calcareous versus the less calcareous samples from the present 

study is the opposite of that reported by Nadeau and 

Reynolds (1981a). We suggest that these differences in 

liS are not related to rock chemistry, as suggested by 

Nadeau and Reynolds (1981a), but to the origins of the 

I/S clay, as will be discussed later. 
0~----------~,----------~----------~-----------~---------. 

\7 

• ,., ... ..... - 500 

/6 

/ 

2000 f- / 

/ 

/ 

/ 

1-- \7/ en w .. , ... a: w • w u. 1--

16 w ~ ~ 

w· - 1000 ~ a: I 0 u..i u I a: 

0 0 

1-- I u 

:I: I 0 

1-- 4000 r- 1--

a.. I :I: w 1-- 0 I a.. 

w 

I 0 

I \7....,.... Average values 

r- I ~6 I - 1500 

'VI 

• ••• •• • 6 

6000 -

7000L-----------~'----------~'----------~'-----------~----------~2000 -10.0 -9.0 -8.0 -7.0 -6.0 -5.0 

Figure 5. Plot of whole-rock oxygen isotope ratios (relative to Pedee belemnite standard) versus 

burial depth. All samples contained less than 35 percent HCI-insoluble material. Dashed line connects average isotope ratio from each core group. 

Differences in the composition of liS clay between 

the uppermost Niobrara and lowermost Pierre strata were 

probably due to variations in origin and provenance at 

the close of upper Smoky Hill time to the onset of deposition of the Pierre Shale. Although the Niobrara Formation represents an overall major transgressive cycle during the Cretaceous, the upper Smoky Hill strata were 

deposited during a regressive phase (Hancock and Kauffman, 1979; Hattin, 1981 ). The more illitic (less expandable) liS in the uppermost Smoky Hill strata (fig. 6) 

probably resulted from a dominant influx of reworked 

liS clay as terrigenous detritus deposited from the west 

and, perhaps, a minor component from the east. Little 

volcanic activity occurred during late Smoky Hill (Beecher 

Island) time, as evidenced by the presence of only a few 

thin bentonites and small amounts of altered airfall ash 

in chalk zones containing authigenic kaolinite and highly 

expandable liS (Pollastro, 1981a). 

The start of the next major transgression, the Claggett cyclothem, was coincident with a change from the 

chalky shale and chalk of the Niobrara Formation to the 

dark gray Pierre Shale (McGookey, 1972), represented 

here as the Sharon Springs Member. The basal 10 m of 

the Sharon Springs Member of the Pierre Shale contained 

numerous bentonite beds, including the Ardmore Bentonite Beds. This 2-m-thick bed or series of beds is an 

excellent time-stratigraphic marker unit in the Upper Cretaceous section of this region, and, in the subsurface, produces a distinctive low resistivity reading on electric logs 

(Shurr, 1980). The thickness and wide extent of this 

marker bentonite throughout the northern part of the 

Pierre seaway is evidence of large-scale volcanism, whose 

apparent locus was in the Elkhorn Mountains of westcentral Montana. Major volcanic eposodes during the 

Cretaceous have been linked to major transgressive events 

(Nadeau and Reynolds, 1981b). Increasing volcanism 

Diagenetic relationships in hydrocarbon-productive chalk, Niobrara Formation 227 
447 

• 

• 

PIERRE SHALE 

453 ~;..._ ____ ? ? ------~F------1 

NIOBRARA FORMATION 

(/) 

a: 

w 

Iw 

~ 

~ 

:x.' 

Ia.. 

w 

0 

w 

I-

<{ 457 ~ 

x 

0 

a: 

a.. 

a.. 

<{ 

463 

0 20 40 60 

INSOLUBLE RESIDUE, 

IN WEIGHT PERCENT 

80 100 0 

• 

• 

20 40 

PERCENT ILLITE LAYERS 

IN ILLITE/SMECTITE 

6 

• 

@ 

• 

60 

Insoluble residues 

and shales 

Only random illite/smectite 

Only ordered illite/smectite 

Both random and ordered 

illite/smectite 

Bentonites 

Only random illite/smectite 

Figure 6. Plot of amount of insoluble residue and composition of mixed-layer illite/smectite (1/S) clay versus depth from core 

of Whomble No. 1 well (460 m depth). Dashed lines represent samples for which 1/S ratio could not be determined. 

during early Sharon Springs deposition explains the differences in composition of IIS for these units. Large volumes of highly expandable clay, formed from relatively 

fresh volcanic material over a short marine residence time, 

were deposited in the Sharon Springs Member. Voluminous, highly expandable IIS clay overwhelmed the much 

less expandable 1/S clay (reworked from older, more illitic, 

terrigenous sources) that typically dominated in insoluble residues of chalk from the uppermost Smoky Hill 

strata. 

As in most deeply buried clay-rich rocks, the most 

significant reaction in the clay-rich insoluble residues and 

bentonite beds of the Niobrara chalks was the conversion 

228 Studies in Diagenesis 

of smectite to illite in the 1/S. The bulk mineralogy of 

the acid-insoluble residues from Niobrara chalks typically 

consisted of various amounts of clay minerals, quartz, pyrite, gypsum, anhydrite, plagioclase feldspar, and, locally, 

zeolites. Although the bulk mineralogy of the insoluble 

residues and bentonite beds varied, no depth-related diagenetic trends were evident. Clay minerals in these insoluble residues and bentonites, however, showed signs of 

depth diagenesis. The dominant clay mineral in more than 

90 percent of the samples studied was 1/S. The compositions of the 1/S from the insoluble residues and bentonites 

of the Niobrara Formation and Pierre Shale in these wells 

are summarized in figure 7. IIS in chalk residues of the 
0 

500 

(/) 

a: 

w 

1-

w 

~ 

~ 

:i 

1- 1000 a.. 

w 

Cl 

1500 

Insoluble residues 

80 60 40 

---tJ- Only random illite/smectite 

• Only ordered illite/smectite 

-@- Both random and ordered 

illite/smectite 

20 0 

PERCENT ILLITE LAYERS IN ILLITE/SMECTITE 

Figure 7. Plot of composition of mixed-layer illite/smectite versus depth in residues and bentonites of Niobrara Formation and Pierre Shale from five wells studied. 

shallow ( 450 m) cores was randomly (R = 0) and shortrange ordered (R = 1, R = 2) interstratified, with compositions ranging from 20 to 70 percent illite layers. The bentonites, however, contained only randomly interstratified 

(R=O) liS and few (less than 15 percent) or no illite 

layers. Insoluble residues from chalk of the two wells at 

intermediate depths (860 m) also contained both random 

and short-range ordered liS ranging in composition from 

50 to 75 percent illite layers. In many insoluble residue 

samples both ordered and random IIS were distinguished 

in the XRD pattern of glycolated materials (fig. 8). Bentonites from these wells at 860 m depth contained random IIS having about 30-35 percent illite layers. 

In the deepest core from the Alice G. Nay No. 1 

well (a well at 1,650 m with oil shows), the IIS from chalk 

residues was both short- and long-range ordered (R>O) 

and ranged in composition from 70 to 85 percent illite 

layers. Regularly interstratified CIS was also identified 

in one of the residues. Bentonites from this core contained 

short-range ordered liS having about 65 percent illite 

layers. In four of the five bentonites sampled from this 

well, the liS coexisted with a iron-rich CIS. 

Insoluble residues of the chalk contained liS of 

many different origins. The IIS had variable compositions 

and degrees of ordering. With increased temperature due 

to burial, the range in composition of liS became more 

limited as a result of the extent of diagenetic reaction; that 

is, smectite layers in highly expandable IIS of insoluble 

Diagenetic Relationships in Hydrocarbon-Productive Chalk, Niobrara Formation 229 
30 

DEGREES TWO THETA 

>-

ICi5 

z 

UJ 

1-

z 

Figure 8. X-ray powder diffraction pattern of oriented, 

ethylene glycol-treated, <0.5-ILm fraction from Niobrara Formation insoluble residue (460 m depth) showing mixture of 

random 17-A and ordered (about 13-A} mixed-layer illite/smectite (1/5). ld, discrete illite; K, kaolinite. CuKo: radiation. 

residues were unstable under deeper burial conditions and 

were, therefore, converted to illite layers, resulting in a narrower overall range in composition of liS in deeper cores. 

In a detailed analysis of the abundance and morphology 

of clay minerals in these residues, Pollastro (1985) showed 

that the smectite-to-illite reaction involved cannibalization of smectite layers and the conservation of aluminum 

and potassium, to produce a more illitic liS clay. 

In bentonites, the starting composition of the liS 

precursor was completely or nearly completely expandable 

smectite. With increasing temperature due to burial, the 

liS in bentonites became progressively more illitic, maintaining its narrow composition range, and, at about 60 

percent illite layers, the liS transformed from random to 

short-range ordered. Rettke (1981) reported similar results 

for liS in cores of the Dakota Formation of the Denver 

basin. Thus, the liS that formed in thin, clean, discrete 

bentonites within Niobrara strata, where starting compositions were nearly fixed and where little or no detrital clays 

were present, was found to be the best indicator of the 

extent of diagenetic reaction and, therefore, the most 

useful for clay-mineral geothermometry. 

Nature and origin of regularly interstratified chlorite/smectite 

The five bentonites recovered from the Alice G. Nay 

well at a depth of about 1,650 m were highly kaolinitic 

(fig. 9). The only differences among the four bentonites 

230 Studies in Diagenesis 

that contained CIS and the one that contained no CIS 

were the thickness of the beds and the abundance and 

distribution of associated pyrite. The four bentonites that 

contained CIS were about 2-3 mm thick and contained 

as much as 50-60 percent pyrite distributed throughout 

the beds. The bentonite that contained no CIS was about 

1 em thick and had only 15 percent pyrite that was concentrated chiefly along the outer edges of the bed. 

Kaolinite-rich bentonites from the Niobrara 

typically contained liS having abundant hydroxy-A! 

interlayers. The hydroxy-A! interlayers were recognized on 

XRD patterns of oriented specimens (heated to 300°C) 

by abnormally large basal spacings (fig. 10). In these 

samples, AIH, rather than Na + and Ca2+, was the principal interlayer cation. Gill and others (1972) also found 

that several bentonites from the Sharon Springs Member 

of the lower Pierre Shale and upper portion of the Niobrara Formation in western Kansas are kaolinitic, that the 

smectite comprising the remainder of the clay fraction 

tends to be the beidellitic variety containing hydroxy-A! 

interlayers, and that these features seem tp be most 

strongly developed in thinner bentonite beds. 

The CIS in these bentonites probably formed at 

elevated temperatures (>900C, Hoffman and Hower, 1979; 

Kubler, 1973) and pressures during deep burial from 

aluminous, dioctahedral IIS and (or) kaolinite in the bentonites reacting with iron-rich fluids. Bentonites within 

the chalk probably acted as primary seams for migration 

of solutions, similar to stylolites, during deep burial. Solution transfer of iron-rich fluids along impermeable bentonite beds caused extensive late pyritization and the formation of the CIS phase. Some iron and magnesium may 

have been derived from the octahedral layers of smectite 

during its diagenetic conversion to illite. The samples 

could not be purified for an accurate determination of 

the 060 reflection, however, from the reaction described 

previously, the CIS was probably an iron-rich, dioctahedral variety (tosudite?). Similar products have been synthesized in laboratory experiments by Velde (1977) and 

Eberl (1978). 

GEOLOGIC AND ECONOMIC IMPLICATIONS 

The diagenetic relationships identified in this study 

have been used to produce generalized predictive maps 

for hydrocarbon reservoir and source rock studies of the 

Niobrara Formation. Based on porosity versus presentday burial and paleoburial relationships, porosity can be 

predicted for chalk facies of the Niobrara Formation 

across the Western Interior (fig. 11). Clay mineral assemblages are commonly used as geothermometers (Hoffman 

and Hower, 1979; Weaver, 1979; Pollastro and Bader, 1983; 

Pollastro and Barker, 1984) and, therefore, provide a 

temperature record of a potential source rock. The conversion of randomly interstratified liS {R=O) to liS with 
B 

A 

1/S 

+ 

ld 

28 

K 

I I 

I I 

24 20 

CIS 

1/S 

K 

DEGREES 28 

+ 

1/S 

>- I-

(ij 

z 

UJ 

I-

~ 

Figure 9. X-ray powder diffractograms of air-dried (A), ethylene 

glycol-treated (B), and heated to 550°C (C) oriented specimens 

of<2-~m fraction of bentonite from Alice G. Nay No.1 well (1,650 

m). Abundant kaolinite (K) and a regularly interstratified 

chlorite/smectite (C/5) phase coexist with ordered illite/smectite 

(1/5). ld, discrete illite. CuKa radiation. 

Diagenetic relationships in hydrocarbon-productive chalk, Niobrara Formation 231 
A 

B 

c 

K 

II 

II 

I I 1 

s.8 .A 

j 

20 16 12 

DEGREES 28 

8 4 2 

Figure 10. X-ray powder diffractograms of treated, oriented 

mounts of <2-J,tm fraction of aluminous bentonite in Niobrara Formation strata of the Alice G. Nay No. 1 core showing evidence 

upon heating of abundant interlayer aluminum complexes. A, 

ethylene glycol-treated; 8, heated at 300°(; C, heated at 550°C. 

1/5, mixed-layer illite/smectite; K, kaolinite. CuKa radiation. 

short-range ordered interstratification (R=1) occurs at 

about 1000C, as does the formation of CIS in buried Cretaceous sandstones and shales (Velde, 1973; Hoffman and 

Hower, 1979; Weaver, 1979). The appearance of either or 

both of these ordered interstratified clay minerals and the 

disappearance of randomly interstratified 1/S in thin bentonite beds of the Niobrara can be used as geothermometers. 

Coincidently, the lower temperature limit for the 

generation of thermogenic oil is also about 1000C (Tissot 

and Welte, 1978). The 100°C paleotherm map of figure 

12 was drawn using the composition and degree of ordering of liS in bentonites from numerous Niobrara and 

lower Pierre Shale outcrop and core samples and with the 

232 Studies in Diagenesis 

aid of structure contour and isopach maps of the Denver 

basin. Although this map was produced from a limited 

distribution of samples, it suggests that the Niobrara in 

parts of the Denver basin and adjacent areas was buried 

much deeper in the past (as compared with its presentday depths) and (or) was subjected to a higher geothermal gradient than that which exists today. 

Conversely, bentonites from some outcrop localities 

in figure 12 indicate that the Niobrara remained under 

relatively shallow burial conditions during the development of the basin, perhaps due to the presence of an early 

basin hinge zone or to significant differences in thickness 

of the overlying section. 
MONTANA 

WYOMING 

BURIAL DEPTH 

E3 <900m 

900-1500 m 

td)}\J/~)Jj > 1 500 m 

I 

200 MILES 

I 

300 KILOMETERS 

Figure 11. Regional distribution of predicted porosity in chalk facies of the Niobrara Formation. Patterns 

show thickness of total post-Niobrara Cretaceous overburden. Area buried to 1,000 m or less should contain chalks with more than 35 percent average porosity. Area buried to 1,000-1,200 m should have about 

30-35 percent average porosity. 

SUMMARY AND CONCLUSIONS 

Studies of inorganic diagenesis and reservoir and 

source-rock properties of hydrocarbon-productive chalk 

from the Niobrara Formation present a comprehensive 

model of burial diagenesis. Progressive loss in porosity 

and negative shifts in the whole-rock oxygen isotope composition of the chalk can be explained by progressive 

cementation due to pressure solution and reprecipitation 

of carbonate with increased burial. The composition and 

ordering of 1/S clay in bentonites and insoluble residues 

showed a progressive transformation with increased 

burial depth and porosity loss. Although the starting composition of the 1/S in chalk at the time of deposition, or 

shortly after, was highly variable, the expandability range 

became increasingly limited and shifted toward more 

illitic compositions with greater burial. The liS in bentonites, however, retained a narrow range of composition, became increasingly more illitic with depth, and converted from random to ordered interstratification at 

about lOOOC. Regularly ordered CIS may have formed 

in these aluminous bentonites in the presence of iron-rich 

fluids under deep burial conditions. 

Because of these consistent relations to burial depth 

(or paleoburial depth) and the broad, uniform facies of 

the chalk, reliable paleogeographic reconstructions and 

economic predictions can be made. 

REFERENCES CITED 

Arthur, M.A., Dean, W. E., Pollastro, R. M., Claypool, G. 

E., and Scholle, P. A., 1985, Comparative geochemical and 

mineralogical studies of two cyclic transgressive pelagic 

limestone units, Cretaceous Western Interior basin, U.S., 

in Pratt, L. M., Kauffman, E. G., and Zelt, F. B., eds., 

Fine-grained deposits and biofacies of the Cretaceous 

Western Interior seaway-Evidence for cyclic sedimentary 

processes: SEPM Field Trip Guidebook No. 4, p. 16-27. 

Craig, H., 1957, Isotopic standards for carbon and oxygen and 

correction factors for mass spectrometric analysis of carbon dioxide: Geochimica et Cosmochimica Acta, v. 12, 

p. 133-149. 

Diagenetic Relationships in Hydrocarbon-Productive Chalk, Niobrara Formation 233 
oSterling 

0 Limon 

EXPLANATION 

A RANDOM ILLITE/SMECTITE 

e ORDERED ILLITE/SMECTITE 

PALEOTEMPERATURE> 1 00°C 

25 50 MILES 

40 80 KILOMETERS 

I 

Area of shallow 

iobrara gas fields 

t--------

I 

'(/) 

I~ 

Figure 12. Paleotherm map of Denver basin showing parts of the Niobrara Formation where temperatures may have been greater 

than 1 00°C. Paleotemperatures are estimated from model proposed by Hoffman and Hower (1979) and are based on degree 

of ordering in mixed-layer illite/smectite from bentonite beds in the Niobrara Formation and lower Pierre Shale. 

Eberl, D. D., 1978, Reaction series for dioctahedral smectites: 

Clays and Clay Minerals, v. 26, p. 327-340. 

Garlick, G. D., 1969, The stable oxygen isotopes of oxygen, in 

Wedepohl, K. W., ed., Handbook of geochemistry, v. 2: New 

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Gill, J. R., and Cobban, W. A., 1966, Regional unconformity in 

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234 Studies in Diagenesis 

Gill, J. R., Cobban, W. A., and Schultz, L. G., 1972, Stratigraphy and composition of the Sharon Springs Member 

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Survey Professional Paper 728, 54 p. 

Hann, M. L., 1981, Petroleum potential of the Niobrara Formation in the Denver basin: Colorado and Kansas: Fort 

Collins, Colorado, Colorado State University, M.S. thesis, 

260 p. 
Hancock, J. M., and Kauffman, E. G., 1979, The great transgressions of the Late Cretaceous: Journal of the Geological 

Society of London, v. 136, p. 175-186. 

Hattin, D. E., 1981, Petrology of Smoky Hill Member, Niobrara 

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American Association of Petroleum Geologists Bulletin, 

v. 65, p. 831-849. 

Hattin, D. E., and Cobban, W. A., 1977, Upper Cretaceous 

stratigraphy and paleoecology of western Kansas: The 

Mountain Geologist, v. 14, p. 175-217. 

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of clay minerals in Cretaceous sediments of the disturbed 

belt of Montana: Cleveland, Ohio, Case Western Reserve 

University, Ph. D. thesis, 266 p. 

Hoffman, Janet, and Hower, John, 1979, Clay mineral assemblages as low grade metamorphic indicators-Application 

to the thrust faulted disturbed belt of Montana, U.S.A., in 

Scholle, P. A., and Schluger, P. K., eds., Aspects of 

diagenesis: Society of Economic Paleontologists and 

Mineralogists Special Publication 26, p. 55-79. 

Hower, John, 1981, Shale diagenesis, in F. J. Longstaffe, ed., 

Short course handbook 7: Mineralogical Association of 

Canada, p. 60-80. 

Kauffman, E. G., 1977, Geological and biological overviewWestern Interior Cretaceous basin: The Mountain Geologist, 

v. 14, p. 75-99. 

Kubler, Bernard, 1973, La corrensite, indicateur possible de 

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sediment: Centres Recherches Pau Bulletin, v. 7, p. 543-556. 

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Rockies: Rocky Mountain Association of Geologists, Symposium Guidebook, p. 271-279. 

Lockridge, J. P., and Scholle, P. A., 1978, Niobrara gas in eastern 

Colorado and northwestern Kansas, in Pruit, J. D., and Coffin, P. E., eds., Energy resources of the Denver basin: Rocky 

Mountain Association of Geologists, Symposium Guidebook, p. 35-49. 

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Geologic atlas of the Rocky Mountain region: Denver, Colorado, Rocky Mountain Association of Geologists, 

p. 190-228. 

Milliman, J.D., 1974, Marine carbonates: New York, SpringerVerlag, 375 p. 

Nadeau, P. H., and Reynolds, R. C., Jr., 1981a, Burial and contact metamorphism in the Mancos Shale: Clays and Clay 

Minerals, v. 29, p. 249-259. 

___ 1981b, Volcanic components in pelitic sediments: Nature, 

v. 294, p. 72-7 4. 

Pollastro, R. M., 1981a, Authigenic kaolinite and associated 

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eastern Colorado: Journal of Sedimentary Petrology, v. 51, 

p. 53-562. 

____ 1981b, Clay-mineral diagenesis within a fine-grained, 

marine, hydrocarbon-bearing, carbonate sequence-Evidence from the Cretaceous Niobrara Formation [abs.]: Program and abstracts, 18th annual meeting, The Clay Minerals 

Society, 30th Annual Clay Minerals Conference, Urbana, 

Illinois, p. 13. 

____ 1982, A recommended procedure for the preparation of 

oriented clay-mineral specimens for X-ray diffraction 

analysis-Modifications to Drever's filter-membrane peel 

technique: U.S. Geological Survey Open-File Report 82-

71, 24 p. 

___ 1985, Mineralogical and morphological evidence for the 

formation of illite at the expense of illite/smectite: Clays 

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Pollastro, R. M., and Bader, J. W., 1983, Clay-mineral relationships in some low-permeability hydrocarbon reservoirs and 

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Association of Petroleum Geologists Bulletin, v. 67, p. 536. 

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clay minerals, vitrinite reflectance, and fluid inclusionsApplications to the thermal and burial history of rocks 

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basin, Wyoming, in Law, B. E., ed., Geological characteristics of low-permeability Upper Cretaceous and Lower 

Tertiary rocks of the Pinedale anticline area, Sublette 

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84-743, p. 78-94. 

Pollastro, R. M., and Martinez, C. J., 1985, Mineral, chemical, 

and textural relationships in rhythmic-bedded, hydrocarbonproductive chalk of the Niobrara Formation, Denver basin, 

Colorado: The Mountain Geologist, v. 22, p. 55-63. 

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Cretaceous Niobrara Formation, Rocky Mountains region: 

Oil and Gas Journal, v. 82, no. 17, p. 138-145. 

Rettke, R. C., 1981, Probable burial diagenetic and provenance 

effects on Dakota Group clay mineralogy, Denver basin: 

Journal of Sedimentary Petrology, v. 51, p. 541-551. 

Reynolds, R. C., 1980, Interstratified clay minerals, in Brindley, 

G. W. and Brown, G., eds., Crystal structures of clay 

minerals and their X-ray identification: London, Mineralogical Society, p. 249-303. 

Reynolds, R. C., Jr., and Hower, John, 1970, The nature of 

interlayering in mixed-layer illite-montmorillonite: Clays and 

Clay Minerals, v. 18, p. 25-36. 

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Niobrara Formation, eastern Denver basin-Evidence of 

biogenic versus thermogenic origin of hydrocarbons, in 

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

Hydrocarbon source rocks of the greater Rocky Mountain 

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p. 365-368. 

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sediments-On land and under the sea: International 

Association of Sedimentology Special Publication 1, 

p. 117-148. 

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American Association of Petroleum Geologists Bulletin, v. 

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391-C, 31 p. 

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Technical Report 21, 176 p. 
Burial Diagenesis of Sandstones 

by RAYMOND SIEVER, Department of Earth and Planetary Sciences 

Harvard University, Cambridge, Massachusetts 02138 

CONTENTS 

Abstract 238 

Introduction 238 

Acknowledgment 238 

Permeability and pore-water flow 238 

Analysis of burial history 239 

Application to Paleozoic formations of eastern New York 242 

Conclusions 244 

References cited 248 

FIGURES 

l. Chart showing stages of burial diagenesis of sandstones 240 

2. Sedimentation diagrams for sandstones that have been buried, uplifted, partially eroded, and reburied 241 

3. Time-temperature diagram showing generalized yield curves for slow 

diagenetic reaction 243 

4. Time-temperature curve for pelagic siliceous ooze deposited on spreading sea 

floor above and below carbonate compensation zone 244 

5-7. Sedimentation diagrams: 

5. Paleozoic formations in eastern New York 245 

6. Diagenetic phases in Esopus Shale 246 

7. Diagenetic phases in Slide Mountain Sandstone 247 

F. A. Mumpton, eel., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 237 
Abstract 

The burial diagenesis of sandstones is controlled chiefly 

by the generally permeable nature of sands that allows extensive flow of pore waters at various stages of the process. The 

flow is driven by expression of compaction water from muddy 

sediments in the sands, by infiltration of meteoric water under 

hydrostatic head from an elevated recharge region, by convection transfer from deeper, warmer regions to shallower, cooler 

ones, and by movement of water produced by dehydration reactions. The chemical driving force behind much of the alteration of detrital components is the chemical instability of many 

of these materials, especially those of volcaniclastic origin. A 

second major control is the increase in temperature and 

pressure that accompanies deep burial under a range of geothermal gradients that differ according to the tectonic position of 

the rocks. The use of sedimentation diagrams, the estimation 

of paleogeotherms from tectonic history, and the reconstruction of time-temperature curves has shown that the diagenesis 

(as derived from petrography) of three Paleozoic formations in 

eastern New York State is a function of the duration of and the 

temperatures that prevailed during their burial. 

INTRODUCTION 

An explosion of research has taken place on the 

diagenesis of all types of sediments, ranging from the 

diagenesis of carbonates, which has become an integral 

part of understanding their deposition, to that of clastic 

sediments. Studies of the diagenesis of sandstones have 

been sparked by exploration for oil and gas and the 

necessity to understand reservoir properties to enhance 

oil recovery. As a result, we have learned much that is new 

and rediscovered much that we knew before. New to the 

study of sandstone diagenesis is the emphasis on burial 

effects in various basins, especially on the increase in 

temperature and pressure that affects the course of cementation and authigenesis. This paper emphasizes those 

kinds of changes, although it will of necessity treat in passing other aspects of the subject, knowledge of which is 

required before we can sort out the effects of burial from 

the many other controls. 

Acknowledgment 

Part of this work was done under the auspices of 

NSF Grant EAR-82-12261. 

PERMEABILITY AND PORE-WATER FLOW 

The important aspect of sandstones that sets them 

apart from mudstones and many other lithologies is the 

generally high porosity and permeability that characterize 

the original sediment and that allow the traversal of pore 

waters at various times in the sediment's history, barring 

238 Studies in Diagenesis 

complete cementation or alteration and loss of permeability. We now understand that not only are arenites of all 

kinds permeable, but the original turbidites that later may 

become matrix-rich graywackes are also permeable, at 

least throughout much of their early history. One of the 

major aspects of sandstone diagenesis, therefore, is the 

decrease in permeability that may come from diagenesis. 

One route is the deformation and alteration of rock 

fragments, that, if extensive enough, convert a lithic 

arenite into a wacke or graywacke. Another route is the 

precipitation of mineral cements that occlude pore space. 

Purely physical compaction is minor for most sandstones; 

most of the compaction stems from pressure solution or 

diagenetic dissolution that locally produces as a consequence secondary porosity (Schmidt and McDonald, 

1979). 

Inasmuch as permeability makes possible movements of pore waters, to understand the nature and extent 

of diagenetic reactions we need to know flow rates. These 

rates depend on both permeability and the nature of the 

flow that results from a hydrostatic head or convective 

motion. The hydrostatic head arises typically from the 

recharging rim of a sedimentary basin. Convective flow 

is produced by geothermal gradients that are high enough 

(Wood and Hewett, 1982), most dramatically where an 

intrusion has taken place at depth (Kastner and Siever, 

1983). 

Another source for flow is the water of compaction 

that is expressed from the entire sedimentary column, 

primarily from the muddy, fine-grained sediments. Much 

of this water is expressed early in a sandstone's history, 

when it is near the top of the section where it acts as 

a conduit for still compacting muds and carbonates. 

After moderate burial the rate of compaction decreases 

drastically and the effect diminishes. At greater depths 

the water liberated by dehydration reactions (Surdam and 

Boles, 1979), such as the transformation of smectite to 

illite (Hower and others, 1976), may move through 

sandstones. 

The boundary value for flow rates is zero when there 

is no flow drive or the permeability becomes vanishingly 

small. At this point reactions within a sandstone become 

completely isochemical and dependent solely on molecular diffusion for transport. The flow rate becomes zero 

when the sandstone becomes sealed by an impermeable 

cap and base, which may lead to overpressuring, another 

disturbance in a sediment's history. Here, the sandstone 

itself remains permeable, and the isochemical reactions 

within it are governed by diffusion. The perfect small-scale 

examples of isochemical reactions in an overpressured formation that remain to be studied are sandstone ripples 

encased in a shale matrix. The flow rate will go to zero 

if and when the sandstone itself becomes internally sealed 

by cementation or alteration of rock fragments to matrix. 

As flow rates approach zero, the length of time 

required for cementation rises steeply (Pettijohn and 
others, 1972, p. 397-398). Indeed, the question arises how 

can a completely cemented sedimentary quartzite form 

in any reasonable geological period of time except at 

higher temperatures, where pressure drives may overcome 

the declining permeability. 

Even without overpressuring or complete cementation, the flow rate becomes zero if there is no pressure 

head whatsoever, due to a perfectly level ground-water 

potential surface in the basin. If the geothermal gradient 

is too low, convection-driven flow is absent. A section 

impoverished in fine-grained sediments does not compact 

significantly, and little water will be expressed. 

The chemical driving force of diagenetic reactions 

is the chemical instability of detrital or chemically precipitated components. Reactions between phases that are 

not in equilibrium, such as smectite and potassium feldspar or two different plagioclases, are examples. At most 

sedimentary temperatures chemical reactions are heterogeneous and involve solids and the pore solution. Polymorphic inversions, such as aragonite to calcite or opalA to opal-CT, are affected by the pore solution if the 

mechanism is recrystallization by dissolution and reprecipitation. If pore water does not move, such reactions 

slowly proceed by diffusion. The trend to equilibrium produces a stable mineral assemblage in equilibrium with the 

pore-water solution. Such seems to be true for older shales 

that have a chlorite-illite-quartz-albite assemblage. If such 

completely isochemical changes were common, each 

formation with a different set of detrital components 

would form its own stable assemblage in terms of both 

the phases present and different proportions of the phases 

that reflect different bulk chemical compositions. The fact 

that in most sections we see no such wide variation in 

composition suggests that in sandstones, at least, few 

courses of diagenesis are controlled chiefly by isochemical 

reactions. 

More to the point are those diagenetic reactions 

brought about by movements of pore waters from one 

region to another. Pore waters may achieve a particular 

composition as a result of dissolution of a mineral in one 

place and the flow transport elsewhere of the solute. The 

flow of the solute might produce a precipitation reaction 

in another region. The most common example is that of 

the infiltration of meteoric waters into near-surface formations. Such waters aggressively weather carbonates and 

unstable silicates and sulfides. One result is the dissolution of detrital plagioclase, an important process in many 

sandstones (Land and Dutton, 1978). Invasion by meteoric 

waters may take place shortly after deposition, when an 

alluvial sandstone is still exposed periodically as a recharge 

region for an aquifer and later when, as a result of erosion and unconformity higher in the section, the sandstone is once again brought into the region of downward 

movement of ground water. 

Other movements of water are those related to circulation in deeper formations and those driven by 

compaction, dehydration, or convection. Regardless of the 

cause, these waters move from regions of higher temperature and pressure towards those of lower temperature and 

pressure, higher in the section. In the course of this mass 

transfer, components sensitive to such temperature 

changes, including most of the common elements, will 

move. The most common example of such a component 

is quartz, whose solubility, rate of dissolution, and 

precipitation as a function of temperature are well known. 

Waters at a given depth that are saturated with respect 

to quartz cool as they move upwards in the column and 

thus become supersaturated with respect to quartz. As 

they traverse sandstones with abundant quartz detritus, 

they find the seed surfaces they need for precipitation. 

The opposite may be true of calcite, which decreases in 

solubility with the increase in temperature and pressure 

that follows most geothermal and lithostatic pressure gradients. Thus, the replacement of quartz by calcite at depth 

and the reverse nearer the surface should be expected as 

a consequence of the changes in solubility of these two 

minerals as they are buried. Although these reactions may 

take place, exactly how and where they happen in specific 

formations is still difficult to assess. In the examples cited, 

mass transport is helped by diffusion, which operates in 

the same way to transfer solutes from regions of dissolution to regions of precipitation, although much more 

slowly. 

ANALYSIS OF BURIAL HISTORY 

The stages of burial diagenesis of sandstones are 

shown in figure 1 by means of a flow chart showing how 

various waters affect a detrital assemblage to produce 

diverse diagenetic materials. The course of early 

diagenesis, including the reaction of minerals with the 

water of the environment and the effects of bioturbation, 

was discussed by Berner (1980). Deep burial assemblages 

are affected more by deeper water movements than by 

shallower ones along temperature and pressure gradients. 

Finally, the formation may be part of a later uplifted section and again infiltrated by meteoric water which results 

in reversible changes, such as temperature and pressure 

decrease, decementation, and kaolinization. 

Siever (1983a) quantitatively analyzed depth of 

burial and increase in temperature by use of sedimentation diagrams and by estimation of paleogeotherms. 

Figure 2 shows an idealized example of one such diagram 

drawn for a formation that is buried, compacted, uplifted, 

partially eroded, and then buried again. The sequence of 

geological events for this formation is plotted on the 

diagram, which shows the locus of the formation with 

respect to the sediment-water interface as a function of 

postdepositional time. Plotting the burial history at a 

single point gives a one-dimensional view of the sedimentary column as it changes with postdepositional time. 

Burial Diagenesis of Sandstones 239 
N 

.. 

= 

[ 

~· 

:r 

0 

~· 

i 

~· 

PROVENANCE 

ENVIRONMENT 

Weathering Detrital 

Assemblage 

Precipitation 

Dissolution 

SEAWATER 

Burial 

Assemblage 

FRESH 

WATER 

INITIAL CEMENT: 

Carbonate 

Evaporite 

AIR 

Shallow 

burial 

METEORIC 

WATER 

INTERSTITIAL 

WATER 

FORMATION 

WATER 

Infiltration Compaction 

Early Diagenetic Assemblage 

Mixing Bacterial 

activity 

Bioturbating 

Fauna 

Early Diagenetic 

Cement and Pore-fillings 

Carbonate 

Evaporite 

Opal-A, quartz 

Sulfide 

Smectite 

Kaolinite 

Zeolite 

Sedimentation 

Subsidence 

Compaction 

Convection 

Dehydration 

Deep Burial 

Assemblage 

Uplift 

Heat flow 

Pressuring 

Overpressuring 

Deformation 

Secondary porosity 

Pressure solution 

Illite-Albite-Chlorite 

Rock fragments to matrix 

Feldspar dissolution, 

precipitation 

Heavy mineral dissolution, 

precipitation 

Organic matter maturation 

Quartz, Carbonate, Clay, 

Zeolite cement 

METEORIC 

WATER 

Infiltration 

Outcrop 

Relaxation of 

pressure, 

temperature 

Decem entation 

Kaolinization 

Figure 1. Stages of burial diagenesis of sandstones, showing the effect of various waters on detrital assemblages. Boxes show assemblages produced by the processes 

indicated by the arrows. 
+1 

-1 

(/) 

a: 

w -2 1-

w 

~ 

0 

...J 

:;z 

~ 

:i 

1-

a.. 

w 

c 

+1 

-1 

-2 

TIME. IN MILLIONS OF YEARS 

0 5 10 15 20 25 30 35 

dz 

/ dt s.l. = Otmax.) 

METEORIC WATER INFLUX 

dz 

dt s.l. = Otmin.) 

Zs.l. ~ Zs.i. Zs.l. ~ Zs.i . 

I 

Zs.l. ~ Zs.i. 

1-4-----SEDIMENTATION------_,•14•----EROSION------~- SEDIMENTATIONI 

EXPLANATION 

Sea level 

Sedimentation interface 

Stratigraphic horizon 

~ Unconformity 

dz 

dt s.l. = Otmax.) 

Figure 2. Sedimentation diagrams for sandstones that have been buried, uplifted, partially eroded, and reburied. A, 

plotted with reference to sea level at zero (zs.l.). 8, plotted with reference to the sediment-water or sediment-air interface as zero (z5

_;). (Modified from Siever, 1983a). Times are when values of dz/dt, the differential of depth (referred to 

either sediment-water interface or sea level) with respect to time, are at maxima or minima and are shown correlated 

with relative values of z-s.l. and z-s.i. and burial stages. 

Of particular importance for diagenesis is the time of 

meteoric water influx during an erosional interval. 

To get the background for the causes and effects 

of pore-water movements, the paleogeotherms for this 

Burial Diagenesis of Sandstones 241 
diagram must be estimated (Siever, 1983a). The geothermal gradient is a function of heat flow, thermal conductivity of the rock column, temperature of the thermal 

boundary layer at the sediment-water or air-water interface, and nonequilibrium heat lag during sedimentation 

or uplift. Each of these quantities will be time dependent. 

Heat flow changes in response to tectonic changes. Thermal conductivity changes with diagenetic changes in 

porosity. The temperature of the thermal boundary layer 

changes in response to physical oceanographic variables, 

paleolatitude changes caused by continental drift, and (or) 

global climatic changes. Heat lag changes with sedimentation and subsidence rates, and heat transfer changes by 

convection of ground waters. None of these factors can 

be precisely calculated from current imprecise geological 

data, but they do not vary widely. Hence, we may make 

general estimates of these quantities that allow a calculation of the gradient at any given postdepositional time. 

Perhaps the most important aspect in reconstructing paleogeotherms is the insight we get on the rates of 

slow diagenetic reactions that characterize sandstone 

diagenesis. Most diagenetic reactions take place at less 

than 100°C, and almost all take place at less than about 

200°C. In this temperature region, the transformations 

of clay minerals, zeolites, feldspars, and other silicates 

rarely go to completion except after very long times. Thus, 

reaction kinetics are important in deducing reaction yields 

as functions of time and temperature. Frequently, the 

Arrhenius equation is used, which relates the change in 

a reaction rate constant to an increase in temperature and 

defines the activation energy as an exponential: 

d(ln k) Ea 

dT R'P 

If Ea is not temperature dependent, the integrated form 

may be used: 

Ea 

k =A exp[- RT], 

where k is the reaction rate constant, Ea is the activation 

energy, R is the gas constant per mole, Tis temperature 

(in Kelvins), and A an integration constant (commonly 

called the "frequency factor"). Although the Arrhenius 

equation is convenient and applicable to a remarkably 

large number of common, fast laboratory reactions, it is 

clearly not applicable to many diagenetic reactions in 

which the activation energies may not be constant with 

temperature, that is, where the reaction mechanisms are 

different in different temperature regions. Yet, this equation has been shown to be a good approximation for at 

least one clay mineral transformation that has been 

studied experimentally (Eberl and Hower, 1976). 

Whether or not the specific equation holds, the general form does, and it leads to yield curves such as those 

shown in figure 3, a plot of time versus temperature. Also 

shown in figure 3 are the normal ranges of sedimentation 

242 Studies in Diagenesis 

rates and geothermal gradients, which can also be 

represented on a time-temperature graph because 

temperature and burial are both related to the rate at 

which a formation is buried. On this same kind of graph 

a series of diagenetic transformations together with a 

burial curve translated into time versus temperature has 

been plotted in figure 4. In this figure, the transition from 

opal-A to opal-CT to quartz for 80 percent yields is shown 

for a siliceous oceanic sediment (Siever, 1983b ). This 

reconstruction is for a sediment that was deposited on a 

ridge flank, well above the carbonate compensation depth 

(CCD), then subsided along a general plate-contraction 

path (Parsons and Sclater, 1977) to a position below the 

CCD and then to an abyssal plain, a typical path followed 

by many Pacific Ocean sediments. The temperatures 

shown are deduced from oceanic heat-flow measurements 

adjusted for distance from a ridge and for an average 

pelagic sedimentation rate. The phase transitions are 

drawn from kinetics of the transformations as experimentally determined by Kastner and others (1977) and inferred 

from occurrences in drill holes of the Deep Sea Drilling 

Program (Kastner, 1981). Such a matching of reaction 

kinetics and burial history allows the prediction of phases 

to be expected at any given postdepositional time. 

APPLICATION TO PALEOZOIC FORMATIONS OF 

EASTERN NEW YORK 

The combined use of sedimentation diagrams, 

paleogeotherms, and petrographic data leads to the kind 

of diagenetic history reconstructions shown in figures 5-7. 

These curves are taken for a point in eastern New York 

just west of the Catskill Mountains front. Abundant information is available on the stratigraphic succession in this 

area, and the diagenesis of several formations has been 

mapped from petrographic information (Siever and 

Hager, 1981). Sedimentation diagrams for four formations 

are shown in figure 5, one each for the Late Cambrian 

(Potsdam Sandstone), the Early Devonian (Esopus Shale), 

the Late Devonian (Slide Mountain Sandstone, Rickard, 

1975), and a Pennsylvanian sandstone inferred by 

extrapolation to have been deposited and eroded in the 

late Paleozoic. All these formations are plotted with 

respect to the sediment-water or sediment-air interface; 

the position of sea level is estimated from the environment of deposition of each formation. The major tectonic 

events of the Paleozoic in this region-the Thconic, Acadian, and Allegheny orogenies-are shown above the 

curve. The stratigraphic succession records subsidence and 

sedimentation during the Cambrian and Ordovician of 

a shallow shelf, deepening to a pre-Thconic turbidite basin, 

probably a marginal sea or back-arc basin. After the 

Thconic orogeny, the region was uplifted and eroded, 

followed by a resumption of shallow-shelf subsidence in 

the Early Devonian. This shelf condition gave way to rapid 

subsidence and sedimentation of a thick Upper Devonian 
150 

(/) 

::::> <b 

U5 

....1 ~ ~ w 

u 

(/) 

w 

w ~ 

a: 

(!) f....'t(-"~ w 

c 100 ~ 

~ §-~ 

w <::o«; 

a: 

::::> 

1-

< a: $~ w 

c.. 

:::?! 

w 

1-

50 

TIME, IN MILLIONS OF YEARS 

Figure 3. Time-temperature diagram showing generalized yield curves (in percent) for a hypothetical, 

typical, slow diagenetic reaction. Region of normal sedimentation-burial rates and geothermal gradients 

is stippled. (Modified from Siever, 1983a.) 

clastic wedge associated with the Acadian orogeny. The 

post-Acadian clastic influx continued until the uplift 

associated with the Allegheny orogeny at the end of the 

Paleozoic. 

At the top of figure 5 these tectonic conditions are 

translated into heat-flow and paleogeothermal gradients. 

Following the Allegheny orogeny Triassic-Jurassic rifting 

began. The westernmost extent of the rift zone is within 

100 km of the area considered here and thus might have 

exerted some minor influence in heating the crust in the 

eastern Catskill region. Following the heating and crustal 

extension of the Triassic-Jurassic, the crust has had a long 

history of attenuating heat flow and has evolved slowly 

to its present average continental value. 

The diagenetic consequences of these events for the 

Potsdam Sandstone, based in part on the petrographic 

work of Sahakian (1963), are plotted below the sedimentation curve in figure 5. Carbonate cement was precipitated during early diagenesis and was accompanied by 

a small amount of physical compaction of the sand and 

associated silty and muddy facies. The calcite cement 

seems to have been the result of the inception of carbonate 

precipitation conditions on a low-latitude shallow shelf, 

which, shortly after deposition of the Potsdam, led to the 

deposition of the largely carbonate facies of the Upper 

Cambrian and Lower Ordovician. Authigenic feldspar in 

the Potsdam is one of the early cements and is probably 

the result of burial to as much as 0.5 km in pore waters 

dominated by marine waters. These pore waters would 

have been altered by early diagenesis in which the pH 

remained about 8, the cation activities increased slightly, 

and silica activity increased, making possible the precipitation of authigenic feldspars as proposed by Kastner and 

Siever (1979). At about the same time, with the beginning 

of deeper burial, pressure solution of quartz began and 

steadily increased during later burial. Pressure solution 

is clearly a function of both pressure and temperature 

increase. At some time near the end of the Ordovician, 

perhaps during a higher heat flow regime, burial would 

have exceeded 2 km, and the increased temperatures would 

have initiated the conversion of various species of smectite detritus to illite and chlorite. The reaction may have 

involved some small amount of dissolution of detrital 

feldspar and growth of albite as the stable illite-chloritealbite-quartz assemblage came into being. By the end of 

the Paleozoic, about 300 million years after deposition, 

enough time had elapsed at temperatures characteristic 

of burial to 3-4 km for complete conversion to the stable 

assemblage found today-calcite-albite-illite-choritequartz. Since then, no changes have taken place. 

The next formation, the Lower Devonian Esopus 

Formation, illustrates a somewhat different history (fig. 6). 

Burial Diagenesis of Sandstones 243 
20 

30 

U) 40 :::J 

en 

_, 

w 

(.) 

U) 

w 

w 50 a: 

(!) 

w 

Cl 

~ 

ui 

a: 60 :::J 

~ 

<( 

a: 

w 

a.. 

~ w 70 ~ 

80 

90 

1-------Above CCD----:----~~_;....---B 

ocl.o <o 

0 2 3 4 5 

POSTDEPOSITIONAL TIME, IN MILLIONS OF YEARS 

Figure 4. Time-temperature curve for a pelagic siliceous ooze deposited on spreading sea floor above 

and below carbonate compensation depth (CCD) and beneath abyssal turbidite plain. Phase transitions of silica polymorphs are for 80 percent yields. Stippled area covers time-temperature region in 

which opai-CT is an abundant silica phase. (Modified from Siever, 1983b.) 

Its sedimentation curve is simpler, but the starting detrital 

and chemical assemblage-a siliceous, carbonate-rich, 

siliciclastic mud-predisposed it to a different course 

(Rehmer, 1976). The early events were carbonate cementation, possibly by recrystallized faunal elements, and the 

formation of concretions. At the same time, compaction 

was proceeding and the opal-A of sponge spicules and 

perhaps some radiolarians was converting to opal-CT. 

Rapid burial under the later Devonian clastic wedge 

followed, and in a few tens of millions of years the Esopus 

was buried to a depth of more than 2 km. Physical compaction was largely completed by this time, and opal-CT 

was converted to quartz. Further burial to depths greater 

than 3 km resulted in the growth of the illite-chlorite-albite 

assemblage and the embrittlement of the shale that led 

to its dominant fracture pattern found well away from any 

tectonic deformation zone. By the early Mesozoic, the 

diagenetic changes in the Esopus Shale were complete and 

yielded the present assemblage of carbonate (cement and 

concretions), quartz (cement and chert layers), albite, illite, 

and chlorite. 

The Upper Devonian sandstone had yet another 

diagenetic history stemming from its detrital assemblage 

of quartz, feldspars, and rock fragments of a supracrustal, 

nonvolcanic suite (Burtner, 1964). Its sedimentation 

244 Studies in Diagenesis 

diagram shows a simple burial history of subsidence 

beneath the thick clastic wedge followed by Mesozoic 

uplift and erosion (fig. 7). Hematite cement was one of 

the early diagenetic additions to this deltaic plain sand 

as the sand physically compacted slightly in the first few 

million years after its deposition. The abundant rock 

fragments of argillaceous and low-grade metamorphic 

rocks would have begun to deform and alter as rapid 

burial started. Similar to the other formations in this succession, the conversion of detrital smectite, kaolinite, and 

other clay minerals to the illite-chlorite suite would have 

begun. Carbonate cement seems to have been introduced 

at this time, perhaps as a result of downward pore-water 

movements that transported HC03- and Ca2+ into 

higher temperature regions where they were less soluble 

and precipitated. 

CONCLUSIONS 

The examples given here are representative of the 

effects of burial on three somewhat different, moderately 

deeply buried Paleozoic formations in a continental margin setting. The course of diagenesis was controlled first 

by the composition of the detritus and environmental 
TIME, IN MILLIONS OF YEARS BEFORE PRESENT 

500 200 100 0 

0 

CONTINENTAL COLLISION CRUSTAL THINNING 

THICKENED CRUST RAPID RIFTING. IGNEOUS CRUSTAL THICKENING 

MOLASSE SEDIMENTATION ACTIVITY, HIGH PLATE DIVERGENCE. 

LOW HEAT FLOW HEAT FLOW ATIENUATING HEAT FLOW 

30° C per km 

ALLEGHENY 

OROGENY 

POTSDAM SANDSTONE 

Figure 5. Sedimentation diagram for Paleozoic formations in eastern New York just west of Catskill Mountains front. 

Depths below the sediment interface are shown for four formations: the Upper Cambrian Potsdam Sandstone, the Lower 

Devonian Esopus Shale, the Upper Devonian Slide Mountain Sandstone, and a Middle Pennsylvanian Allegheny sandstone. C, Cambrian; 0, Ordovician; S, Silurian; D, Devonian; C, Carboniferous; P, Permian; Tr, Triassic; J, Jurassic; K, 

Cretaceous; T, Tertiary. Below time scale are place tectonic interpretations for appropriate time intervals of the Catskill 

front region of eastern New York State and estimated geothermal gradient values for these time intervals (in degrees 

Celsius per kilometer). Curve for sea level was inferred from sedimentologic and stratigraphic reconstructions of sedimentary 

environments. Time intervals when sea level was above sediment-water interface, corresponding to time of marine sedimentation, are stippled between sea-level curve and zero reference line for sediment-water interface. Diagenetic phases 

shown are for Potsdam Sandstone. 

chemical additions; that is, kinds and amounts of feldspars, carbonate, clay minerals, and rock fragments. The 

course of early diagenesis was controlled by the chemistry of interstitial water, whether marine or subaerial. 

Under subaerial conditions infiltration of meteoric 

waters may have been significant in the early alteration 

of detrital minerals and the addition of cements. Deeper 

burial diagenesis was controlled by the rate at which the 

sediment was transported to higher temperature and 

pressure regimes. Here, faster kinetics of slow reactions 

took over and, given even relatively short times at deep 

burial, resulted in extensive alterations, such as the formation of the stable silicate assemblage, illite-chloritealbite-quartz. 

The application of sedimentation diagrams and 

paleogeotherms in a context of chemical kinetics is the 

guide to a quantitative as well as a qualitative evaluation 

of the postdepositional time at which a given yield of 

diagenetic product can be achieved. Study of recent 

sedimentary deposits allows us to catch some of the faster 

diagenetic reactions. The examples given here are of nonvolcaniclastic sandstones. The course of diagenesis of 

volcaniclastic sands genetically associated with plate convergences involves many different phases and types of 

diagenetic paragenesis, examples of which were given by 

Galloway (1974) and discussed by Surdam and Boles 

(1979). A complete picture of the range of burial diagenesis of sandstones must include sandstones of all the 

Burial Diagenesis of Sandstones 245 
N 

~ 

~ c 

a. 

;· 

(II 

:r 

S2 

Dol 

' ~ = 

!!!. 

(II 

w +3 u 

<( 

u.. 

a: 

w 

1-

~ 

a: 

w +2 

~ 

~ 

~ z 

w 

~ 

c 

w +1 (/) 

w 

:I: 

1-

T 

~ 0 

...J S.l. w 

aJ 

Q 

z 

<( 

+ 

w 

> -1 

0 

aJ 

<( 

(/) 

a: 

w 

1-

w 

~ -2 0 

...J 

~ 

~ 

en 

(/) 

w 

z -3 ~ 

~ 

:I: 

1-

-4 

550 500 

SEDIMENT INTERFACE 

-

400 

//''(SEA LEVEL 

I \ 

I I 

I I 

J \ 

/ \ 

'"" \ 

\ 

\ 

\ ·~ Its 

\ {s 

\_./· 

TIME, IN MILLIONS OF YEARS BEFORE PRESENT 

300 200 100 

ESOPUS SHALE 

5\-\~\.E // 

_,........---

-~ d\. 

~~ 

~· 

~~ -z. 

~\ 

\ .. ~----~---- •I 

I 

I " / . ·'-...._...· 

./· 

~-0<::, 

<ve:,O 

........- ./ 

/ / 

/..~ /-0 Lv 

' I /~<v~ \. // 

I 

{--'/ 

// 

// 

---~------ / -~-

// 

Figure 6. Sedimentation diagram showing diagenetic phases in Esopus Shale. 

0 
CD 

c 

.. 

!: 

0 

t• 

iii" 

a 

~ 

&. 

s 

m 

..., 

.. ...... 

w +3 u 

<( 

LL. 

a: 

w 

1-

z 

a: 

w +2 

~ j-!. 

z 

w 

550 500 

~ ;) 

~ +1 I \CII w \m w I ~ 

400 

TIME, IN MILLIONS OF YEARS BEFORE PRESENT 

300 200 

SLIDE MOUNTAIN SANDSTONE 

(Rickard, 1975) 

~ I LI I \~ CARBONATE 

~ SEDIMENT INTERFACE / \~""" DECEMENTATION 

100 

0 

u:: 0 ·- -~\Ff· --......---, ,• -----· ~ "'\ \ / \ . HEMATITE '··., .\ .-·p..-.;_E·-- z \ \ 1 \ . CEMENT '··.. ./ \ ~<t.; ./ S\'\ 

::_ \ \ / -. I. PENNSYLVANIAN/· \ G-0 - -<.~~ // 

~ '\ \ I \ I SANDSTONE \ ~<:;) /~~0 // 

// -------

> -1 I · \ <?~ / 

g '\ \_ ~-~O \ ILLITE-CHLORITE / <( )':' \ ' lQ I ~\ • \ / 

/ ------~- ----· /./ / - w ., I ~ \ ROCK FRAGMENT \ / 

t;j I DEFORMATION ' I , / 

,/ /-~~ ~ -2 \ ,1 $\ i l£ '..._/ \ \ / /~ , /.. 

~ 

u) 

w 

w 

~ -3 

~ 

:I: 

1-

-4 

~\ I\ 

0· \ 

a\ z, m\ 

'\ 

I 

COMPACTION 

s(;D ~ 

'\~ARBONATE 

~ I '"'-... __ _,/ 

,1 //<?((.;~ 

L /// 

I-/ 

Figure 7. Sedimentation diagram showing diagenetic phases in Slide Mountain Sandstone (Rickard, 1975) . 

0 
types in different provenances genetically related to plate 

tectonic environments (Dickinson and Suczek, 1979) and 

how the diagenetic history of these deposits is subsequently affected by plate tectonic movement as manifested through paleogeothermal gradients (Siever, 1979). 

To pursue this goal requires the modal mineralogy of 

diagenetic components of sandstones so that reaction 

yields can be compared with products. We need to compare different formations in the same burial sequence to 

see the effects of varying depths of burial. This should 

be paired with studies of one formation at various depths. 

Above all, we must reconstruct the burial history as 

accurately as possible, for that will allow the deductions 

of paleohydrology coupled to reaction kinetics that will 

ultimately lead to a predictive capability for sandstone 

diagenesis. 

REFERENCES CITED 

Berner, R. A., 1980, Early Diagenesis: Princeton, Princeton University Press, 214 p. 

Burtner, R., 1964, Paleocurrent and petrographic analysis of the 

Catskill facies of southeastern New York and northeastern 

Pennsylvania: Cambridge, Massachusetts, Harvard University Ph. D. thesis, 221 p. 

Dickinson, W. R., and Suczek, C. A., 1970, Plate tectonics and 

sandstone compositions: American Association of Petroleum Geologists Bulletin, v. 63, p. 2164-2182. 

Eberl, D. D., and Hower, J., 1976, The kinetics of illite formation: Geological Society of America Bulletin, v. 87, 

p. 1326-1330. 

Galloway, W. E., 1974, Deposition and diagenetic alteration of 

sandstone in northeast Pacific arc-related basins; Implications for graywacke genesis: Geological Society of America 

Bulletin, v. 85, p. 379-390. 

Hower, J., Eslinger, E., Hower, M. F., and Perry, E. A., 1976, 

Mechanisms of burial metamorphism of argillaceous sediment, mineralogical and chemical evidence: Geological 

Society of America Bulletin, v. 87, p. 725-737. 

Kastner, M., 1981, Authigenic silicates in deep-sea sediments; 

Formation and diagenesis, in, Emiliani, C., ed., The sea, 

v. 7, The oceanic lithosphere: New York, Wiley, p. 915-980. 

Kastner, M., Keene, J. B., and Gieskes, J. M., 1977, Diagenesis 

of siliceous oozes-I. Chemical controls on the rate of opal-A 

to opal-CT transformation-An experimental study: 

Geochemica et Cosmochimica Acta, v. 41, p. 1041-1059. 

248 Studies in Diagenesis 

Kastner, M., and Siever, R., 1979, Low temperature feldspars 

in sedimentary rocks: American Journal of Science, v. 279, 

p. 435-479. 

Kastner, M., and Siever, R., 1983, Siliceous sediments of the 

Guaymas basin; The effect of high thermal gradients on 

diagenesis: Journal of Geology v. 91, p. 629-641. 

Land, L. S., and Dutton, S. P., 1978, Cementation of a Pennsylvanian deltaic sandstone; Isotopic data: Journal of 

Sedimentary Petrology, v. 48, p. 1167-1176. 

Parsons, B., and Sclater, J. G., 1977, An analysis of the variation of ocean floor heat flow and bathymetry with age: 

Journal of Geophysical Research, v. 82, p. 803-827. 

Pettijohn, F. J., Potter, P. E., and Siever, R., 1972, Sand and 

sandstone: Heidelberg, New York, Springer-Verlag, 618 p. 

Rehmer, J., 1976, Petrology of the Esopus Shale (Lower Devonian), New York and adjacent states: Cambridge, 

Massachusetts, Harvard University Ph. D. thesis, 288 p. 

Rickard, L. V., 1975, Correlation of the Silirian and Devonian 

rocks in New York State: New York State Museaum and 

Science Service, Map and Chart Series 24, p. 1-16. 

Sahakian, A., 1963, Paleocurrent sudy of the Potsdam Sandstone of New York and Quebec: Cambridge, Massachusetts, 

Harvard University Ph. D. thesis, 94 p. 

Schmidt, V., and McDonald, D. A., 1979, The role of secondary porosity in the course of sandstone diagenesis, in Scholle, 

P. A., and Schluger, P. R., eds., Aspects of diagenesis: 

Society of Economic Paleontologists and Mineralogists 

Special Publication 26, p. 175-207. 

Siever, R., 1979, Plate tectonic controls on diagenesis: Journal 

of Geology, v. 87, p. 127-155. 

Siever, R., 1983a, Burial history and diagenetic reaction kinetics: 

American Association of Petroleum Geologist Bulletin, v. 

67, p. 684-689. 

___ 1983b, Evolution of chert at active and passive continental 

ma~gins, in Iijima, A., Hein, J. R., and Siever, R., eds., 

Siliceous deposits in the Pacific region: New York, Elsevier, 

p. 7-24. 

Siever, R., and Hager, J. R., 1981, Paleogeography, tectonics and 

thermal history of some Atlantic margin sediments, in Kerr, 

J. W., ed., North Atlantic borderlands: Canadian Society 

of Petroleum Geologists Memoir, v. 11, p. 95-117. 

Surdam, R. C., and Boles, J. R., 1979, Diagenesis of volcanic 

sandstones, in Scholle, P. A., and Schluger, P. R., eds., 

Aspects of diagenesis: Society of Economic Paleontologists 

and Mineralogists Special Publication 26, p. 227-242. 

Wood, J. R., and Hewett, T. A., 1982, Fluid convection and 

mass transfer in porous sandstones-A theoretical model: 

Geochimica et Cosmochimica Acta, v. 46, p. 1707-1713. 
Diagenesis of the Morrison Formation, 

Smith Lake Uranium District, 

McKinley County, New Mexico 

By NEILS. FISHMAN, U.S. Geological Survey, Box 25046, DFC, MS 939, Denver, 

Colorado 80225, and 

RICHARD L. REYNOLDS, U.S. Geological Survey, Box 25046, DFC, MS 964, 

Denver, Colorado 80225 

CONTENTS 

Abstract 251 

Introduction 251 

Acknowledgments 251 

Geologic and geochemical setting 251 

Geology 251 

Geochemistry 253 

Organic material 253 

Methods 254 

Paragenetic relations 254 

Alterations preceding emplacement of organic material 254 

Organic material and associated alterations 256 

Alterations following emplacement of organic material 258 

Discussion 259 

References cited 262 

FIGURES 

1-3. Maps showing: 

1. Location of Smith Lake district 252 

2. Smith Lake district and structural setting of Mariano Lake and Ruby 

orebodies 253 

3. Sampling and core localities at Mariano Lake and Ruby deposits 255 

4. Diagram of diagenetic alterations in sandstones that host the Mariano 

Lake and Ruby uranium deposits 256 

5-6. Scanning electron micrographs: 

5. Early authigenic pyrite in embayed areas of detrital quartz 257 

6. Authigenic, smectite-rich mixed-layer illite/smectite 257 

7-10. Photomicrographs showing: 

7. Detrital plagioclase dissolved along crystallographic planes 257 

8. Detrital sanidine in which central core has dissolved to leave 

secondary void 257 

9. Authigenic potassium feldspar overgrowth on detrital microcline 257 

10. Organic material partly filling intergranular pore in host rock 257 

11. Scanning electron micrograph showing distribution of uranium, vanadium, 

silicon, and aluminum in organic material 259 

12-13. Photomicrographs showing: 

12. Organic material coating grains and filling intergranular pores 260 

13. Organic material intimately mixed with euhedral pyrite 260 

F. A. Mumpton, eel., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 249 
14. Histogram of sulfur isotope values of FeS2 minerals from Mariano Lake 

and Ruby deposits 260 

15. Scanning electron micrograph of authigenic vanadiferous chlorite 260 

16. Photomicrograph of dissolved iron-titanium oxide, with titanium oxide 

remnants 261 

17. Histogram of carbon isotope values of authigenic calcite from Mariano 

Lake and Ruby deposits 261 

18. Photomicrograph of kaolinite replacement of calcite 261 

19-20. Scanning electron micrographs: 

19. Authigenic native selenium 261 

20. Probable authigenic goethite coating native selenium 261 

250 Studies in Diagenesis 
Abstract 

Petrographic and geochemical studies have revealed a 

complex diagenetic history for sandstone units within the lower 

part of the Brushy Basin Member of the Morrison Formation 

Ourassic) in the southern San juan basin, New Mexico. 

Underground mine and core samples were obtained from the 

Mariano Lake and Ruby uranium mines, which are about 5 km 

apart in the Smith Lake uranium district. The host sandstones 

of the two deposits have undergone similar diagenetic alterations except that authigenic calcite is abundant at the Ruby 

locality and rare at the Mariano Lake locality. Early in the 

diagenetic history of the host sandstones, solutions bearing soluble organic species deposited noncellular organic material that 

coats detrital grains and fills some voids. The organic material 

localized and concentrated uranium (chiefly as a urano-organic 

complex) to form uranium orebodies that also have high concentrations of vanadium, chiefly as vanadiferous chlorite, and 

sulfur, chiefly as void-filling pyrite and marcasite. The o34S 

values for both iron disulfides range from -29.4 to -41.6 per 

mil. Organic-rich solutions were also probably responsible for 

extensive destruction of detrital titanomagnetite and ilmenite 

grains. 

Diagenetic alterations preceding the introduction of the 

organic material include the precipitation of small (less than 

10 JLm), euhedral pyrite crystals and framboidal pyrite which 

differ only slightly in isotopic composition (o34S values range 

from -11.0 to -38.2 per mil) from the later void-filling pyrite 

and marcasite. In addition, precipitation of smectite-rich, mixedlayer illite/smectite clays, dissolution of some detrital feldspars, 

and precipitation of quartz and feldspar overgrowths also 

preceded emplacement of the organic material. Following 

precipitation of the organic material and associated uranium, 

calcite (o13C values range from -7.4 to -12.5 per mil), barite, 

dolomite, and kaolinite formed as interstitial cements. Subsequent oxidation of parts of the orebodies led to the replacement of iron disulfides by iron oxides and hydroxides, the 

precipitation of native selenium, and a minor redistribution of 

iron disulfide, uranium, and iron oxide(?) minerals. The light 

sulfur-isotope values strongly suggest a biogenic origin for the 

iron disulfide minerals, and the light carbon-isotope values of 

calcite suggest that oxidized organic carbon was incorporated 

into calcite. 

INTRODUCTION 

The sequence of diagenetic alterations in sandstones 

that contain uranium orebodies record geochemical and 

hydrologic conditions during and after the formation of 

these orebodies. In this paper we outline the diagenetic 

evolution of two sandstones containing uranium 

orebodies-the Mariano Lake and Ruby orebodies-in the 

Brushy Basin Member of the Jurassic Morrison Formation along the southern margin of the San Juan basin, 

New Mexico. These two deposits comprise much of the 

mineralization in the Smith Lake district of the Grants 

uranium region (fig. 1), the major uranium-producing 

region in the United States (Chenoweth and Holen, 1980). 

Petrographic observations and chemical data are 

integrated with isotope data to place important constraints 

on the nature and timing of uranium mineralization in 

these sandstones and to help identify factors responsible 

for preservation of the orebodies. Additionally, these 

studies provide a broad picture of diagenetic patterns in 

some Brushy Basin sandstones in the central part of the 

Grants uranium region, a relatively unstudied area. 

Acknowledgments 

We are grateful to Jacques F. Robertson and Martin B. Goldhaber of the U.S. Geological Survey for their 

many helpful discussions of the material in this report and 

to Jacques F. Robertson for providing additional samples 

from the Ruby deposit. We thank Robert A. Zielinski for 

his help in preparing fission-track maps. Mark Stanton 

and Jim Nishi provided valuable laboratory assistance. We 

appreciate access to the Mariano Lake mine granted by 

Gulf Mineral Resources Corporation and to the Ruby 

mine granted by Western Nuclear Corporation. Robert A. 

Zielinski and Christine E. Thrner-Peterson provided many 

helpful comments and suggestions in their critical reviews 

of this manuscript. 

GEOLOGIC AND GEOCHEMICAL SETTING 

Geology 

Uranium deposits of the Grants uranium region 

occur primarily in sandstones of the Westwater Canyon 

and Brushy Basin Members of the Morrison Formation. 

In northwest New Mexico the Westwater Canyon Member 

is composed chiefly of several fluvial sandstone "packages" which were deposited on an aggrading alluvial plain 

(Thrner-Peterson and others, 1980; Turner-Peterson, 1986). 

The overlying Brushy Basin Member is composed chiefly 

of lacustrine mudstones and interbedded lenticular sandstones; many of the sandstones are compositionally and 

texturally similar to those of the Westwater Canyon 

Member (Craig and others, 1955; Thrner-Peterson, 1986; 

Bell, 1986). Brushy Basin mudstone units were deposited 

along the margins and within a saline, alkaline lake 

(Bell, 1986; Thrner-Peterson, 1985; Thrner-Peterson and 

Fishman, 1986). 

The sandstones at the Mariano Lake and Ruby 

localities are, for the most part, the lowermost sandstones 

in the Brushy Basin Member. It is not known if these sandstones are part of a continuous sand body or if they are 

separate bodies. Because sandstones in the Brushy Basin 

Member are typically lenticular and discontinuous, the 

sandstones at the two localities are considered here to be 

two distinct sand bodies. 

Diagenesis of the Morrison Formation, Smith Lake Uranium District, McKinley County, New Mexico 251 
SAN JUAN UPLIFT 

_C_QLO~QQ_ _____________________________________ """"'---------

NEWMEXICO 

• Farmington 

SAN 

\ JUAN 

BASIN 

Gallup • 

Grants • ~ Albuquerque • 

10 20 30 40 50 KILOMETERS ~~~~~~~~--~~~~--LI __ ~il 

10 20 30 MILES 

Figure 1. Index map showing location of Smith Lake district (stippled) in Grants uranium region, San juan basin, New Mexico. 

Lines represent outcrops of the Morrison Formation. 

252 Studies in Diagenesis 
The host sandstones are fairly well to poorly sorted, 

medium- to coarse-grained arkoses. In both deposits, 

uranium is intimately associated with a noncellular 

organic material that partly impregnates the host rock 

(Fishman and Reynolds, 1982; this study). This association is typical of the primary (tabular) uranium deposits 

in the Grants uranium region (Granger and others, 1961; 

Adams and others, 1978) and is expressed as a positive 

correlation between whole-rock abundances of organic 

carbon and uranium and as a directly observed spatial 

association of organic matter and uranium (see later 

discussion). 

The Mariano Lake deposit is in the trough, and the 

Ruby deposit is on the south flank of the same syncline 

(fig. 2). The axial trend of the syncline is approximately 

perpendicular to the present hydrologic and paleohydrologic gradient which is from south to north into the San 

Juan basin (Craig and others, 1955). The structural setting appears to have had different influences on groundwater flow and subsequent diagenetic alterations in the 

two deposits, as discussed later. 

Geochemistry 

Locally, both deposits occur on or near the chemically reduced side of a poorly defined, regional reductionoxidation or "redox" boundary. Such a configuration has 

led previous investigators to conclude that the orebodies 

originated in the late Tertiary by the concentration of 

T 

15 

MARIANO LAKE DEPOSIT 

t 

EXPLANATION 

R 13 W R 14 W 

----FAULT-Dashed where inferred, dotted where concealed + MONOCLINAL FOLD-Showing dip direction 

uranium from oxidation-reduction processes along a 

migrating redox interface (Place and others, 1980; 

Ristorcelli, 1980; Saucier, 1980). Although some deposits 

elsewhere in the region have undoubtedly formed by such 

processes, they are essentially devoid of the organic 

material that characterizes primary (tabular) uranium 

deposits (Granger and others, 1961; Squyres, 1980; 

Fishman and Reynolds, 1983) and are considered secondary in origin, uranium having been remobilized from 

preexisting primary (tabular) deposits (Saucier, 1980; Ludwig and others, 1982). Therefore, the coexistence of 

organic material with uranium at the Mariano Lake and 

Ruby deposits indicates a primary (tabular) origin 

relatively unmodified by later oxidation-reduction 

recyling. 

Organic material 

Organic material in the host sandstones is an abundant authigenic constituent and exerts obvious control on 

the localization of uranium mineralization. The organic 

material occurs as a noncellular (Hatcher and others, 

1986) substance which coats detrital grains and fills some 

primary voids and secondary voids formed by early 

diagenetic processes (Fishman and Reynolds, 1982). The 

organic material imparts a dark brown or black color to 

the host sandstone. The organic material was introduced 

post-depositionally, probably as a soluble or colloidal 

organic species (see Squyres, 1980; Leventhal, 1980) and 

--+- ANTICLINE-Showing trace of axial plane anel direction of plunge 

-+- SYNCLINE-Showing trace of axial plane 

.. MAJOR URANIUM DEPOSIT 

• DRILL HOLE WITH URANIUM MINERALIZATION 

e CORE HOLE 

---4-~ GROUND-WATER FLOW DIRECTION 

Figure 2. Plan map of part of the Smith Lake district (stippled area in figure 1 ), showing structural setting of Mariano Lake 

and Ruby orebodies (modified from Green and Jackson, 1975; Robertson and jackson, 1975; j. F. Robertson, U.S. Geological 

Survey, Denver, Colo., written commun., 1981 ). The location of oxidized rock, reduced rock, and the approximate location 

of the reduction-oxidation (redox) boundary are also shown. 

Diagenesis of the Morrison Formation, Smith Lake Uranium District, McKinley County, New Mexico 253 
was localized in the host rocks in undulatory layers that 

are generally subparallel to stratification and elongate 

parallel to sedimentary trends (see Granger and others, 

1961; Fishman and others, 1985). 

The original chemical structure and isotopic composition of the organic material has possibly been 

modified by radiation bombardment, oxidation, and 

aging. As a result, attempts to identify the exact source 

of the organic material are indeterminate (see Leventhal, 

1980); however, recent studies suggest that the organic 

material was originally a humic substance (Hatcher and 

others, 1986) which was probably derived from the decomposition of vegetal matter (Schnitzer and Kahn, 1972; 

Hatcher and others, 1986). That the original organic material is more likely a humic substance than, perhaps, an 

oil-related substance is also suggested from laboratory 

experiments which indicate that modern humic substances 

concentrate uranium (Szalay, 1964}, whereas oil and some 

oil-related substances apparently do not (Moore, 1954; 

Erickson and others, 1954; Hyden, 1961). The localizing 

mechanism for the organic species is poorly understood. 

Granger (1968) proposed that the organic material 

precipitated at the interface between solutions containing the soluble organic species and a cation-bearing 

ground water. In contrast, Fishman and Turner-Peterson 

(1986) proposed that the cations necessary to precipitate 

the organic material were "scavenged" from detrital 

grains, especially detrital magnetite and ilmenite. 

METHODS 

For this study, 129 samples (mineralized and unmineralized samples) were collected from 26 different sites at 

the Mariano Lake mine (fig. 3A) and 53 samples 

(mineralized and unmineralized) from 19 sites at the Ruby 

mine (fig. 3B). In addition, 100 samples of the ore-bearing 

sandstone were obtained from seven cores drilled by Gulf 

Mineral Resources Corporation. The trend of the core 

fence was approximately north-south; cores 3 and 4 penetrated mineralized ground. Cores 1 and 2 were drilled 

south of and cores 5-7 north of the Mariano Lake 

ore body. 

After drying, part of each mine sample was used 

for chemical analyses which included uranium by a 

delayed neutron technique, vanadium by X-ray fluorescence, and carbon and sulfur by induction furnace techniques. The complete geochemical data are summarized 

in Pierson and others (1983). In addition, a heavy-mineral 

separate of each sample was obtained by established procedures using high-density liquids. A representative portion of each heavy-mineral separate was then imbedded 

in epoxy, and the mount was polished for petrographic 

study using vertically reflected light under oil. Petrographic observations using both transmitted and vertically 

reflected light were also made on polished thin sections. 

254 Studies in Diagenesis 

Clay minerals were identified by X-ray diffraction 

measurements of oriented mounts of the <2-p.m size material. Clay minerals and other alterations were observed 

using a scanning electron microscope (SEM) equipped 

with an energy dispersive X-ray spectrometer (EDX) for 

determining elemental compositions. Additionally, X-ray 

"dot maps:' used to determine the distribution of elements 

at selected sites on polished thin sections and grain 

mounts, were made with the SEM. 

Fission-track images of the concentration and distribution of uranium in 10 thin sections were developed 

in muscovite mica overlays (Zielinski and Rosholt, 1978). 

Sulfur-isotope compositions of iron disulfide 

minerals were determined by mass spectrometric measurements of samples prepared by methods outlined in 

Bailey and Smith (1972) on selected samples characterized by a single authigenic iron disulfide textural type or 

generation (see following discussion). The sulfur isotope 

compositions (o34S) of FeS2 minerals are expressed as the 

per-mil difference between the ratio of 32S/34S in the 

sample to that of a standard (Canon Diablo troilite). The 

carbon-isotope compositions (o13C) of authigenic calcite 

were determined by mass spectrometric measurements of 

C02 released upon digestion of the sample in phosphoric acid. The carbon-isotope compositions (o13C} are 

expressed as the per-mil difference between the ratio of 13C/ 12C in the sample to those of a standard (Belemnitella americana from the Cretaceous Pedee Formation, 

South Carolina). 

PARAGENETIC RELATIONS 

The sandstones that host the Mariano Lake and 

Ruby deposits have had a similar diagenetic history (fig. 

4). It is convenient to divide the paragenetic sequence 

determined for the two sandstones into three groups: (1) 

alterations that preceded emplacement of the organic 

material (pre-ore}, (2) alterations associated with emplacement of the organic material and associated uranium 

(ore}, and (3) subsequent alterations (post-ore). 

Alterations preceding emplacement 

of organic material 

Paragenetic relations reveal that precipitation of 

pyrite was the first alteration in the sandstones. This early 

authigenic pyrite (pre-ore FeS2), identified both 

petrographically and with the SEM, commonly occurs as 

euhedral crystals (less than 10 p.m across) or aggregates 

of crystals perched on detrital grain surfaces or within 

small embayed areas on the grain surface (fig. 5). Framboidal pyrite is also present in both deposits, and some 

may have precipitated early in the alteration history 

of the host sandstones. An origin involving bacterial 
4 

A 1 00 200 METERS 1,_ I I 

EXPLANATION 

• SAMPLING LOCATION 

6.+ CORE LOCATION AND NUMBER 

B 1 00 200 METERS I I 

Figure 3. Map showing sampling and core locations at A, the Mariano Lake deposit (deposit outline 

modified from Sachdev, 1980), and 8, sampling locations at Ruby deposit. Outlined area marks 

the extent of mine workings. 

sulfate reduction is possible for framboidal pyrite (Berner, 

1970); such bacterial control probably also resulted in 

isotope fractionation (Goldhaber and others, 1978) 

yielding the isotopically light sulfur which was incorporated into these pre-ore FeS2 minerals (o34S values 

range from -11.0 to -38.4 per mil). 

Authigenic clays, which coat some detrital grains 

and surround early pyrite, display an interlocking 

honeycomb network with minor crenulated surfaces when 

observed with an SEM (fig. 6). The texture and morphology of the early clays are similar to authigenic 

smectite-rich clays described by Wilson and Pittman 

(1977). X-ray diffraction analyses of host-sandstone 

samples revealed the presence of smectite-rich, mixed-layer 

illite/smectite clays containing 20-30 percent illite layers 

(as determined by the technique of Reynolds and Hower, 

1970), which corroborates the clay identification made on 

the basis of morphology. 

Dissolution of some detrital feldspars followed the 

precipitation of mixed-layer clays. The absence of mixedlayer clays in voids left by feldspar dissolution suggests 

that the clays formed prior to dissolution. Those feldspars 

displaying extensive dissolution features include plagioclase (fig. 7) and some potassium feldspars (fig. 8). Identification of the type of potassium feldspar that was 

dissolved was difficult due to extensive alteration of these 

Diagenesis of the Morrison Formation, Smith Lake Uranium District, McKinley County, New Mexico 255 
ES 

ALTERATION TYPE 

PRE-ORE POST-ORE 

FeS2 PPT 

SMECTITE-ILLITE PPT 

FELDSPAR DISSOLUTION 

QUARTZ PPT 

K-FELDSPAR PPT 

Fe-Ti Ox DISSOLUTION 

ORGANIC MATERIAL PPT 

URANIUM LOCALIZATION 

CHLORITE PPT 

CALCITE PPT 

BARITE PPT 

DOLOMITE PPT 

KAOLINITE PPT --.-?-? 

FeS2 OXIDATION ?-4.111 

SELENIUM PPT ?-4•t~ 

Fe OXIDE PPT ? 

Figure 4. Summary of diagenetic alterations in sandstones that host the Mariano Lake and Ruby uranium deposits. 

Open ellipses indicate uncertain paragenetic relationships between barite and dolomite with respect to calcite. Ppt, 

precipitation. 

grains; however, these grains were probably originally 

unstable under diagenetic conditions, and were possibly 

sanidine. Orthoclase and microcline generally show no 

evidence of dissolution. Potassium feldspar dissolution 

appears to have proceeded from the core of the grain outward, whereas plagioclase dissolution proceeded from the 

grain surface inward, in part controlled by cleavage and 

(or) twin planes. 

Quartz and feldspar overgrowths also formed prior 

to emplacement of organic material. Electron microprobe 

analyses reveal the presence of K, AI, and Si in these 

authigenic feldspars. Void-filling authigenic feldspars are 

morphologically rhombic (observed in SEM studies), 

similar to the authigenic feldspars observed and identified 

by Sibley (1978) as adularia. The overgrowths observed 

in our studies occur on detrital orthoclase and microcline 

256 Studies in Diagenesis 

grains. Additionally, authigenic adularia(?) occurs in 

secondary voids in detrital feldspars (Fishman and 

Reynolds, 1982). Both authigenic quartz and adularia(?) 

enclose previously formed pyrite and are in turn coated 

by organic material (fig. 9). 

Organic material and associated alterations 

Noncellular organic material is extremely abundant 

in ore zones, commonly comprising more than 10 percent 

by volume (Fishman and Reynolds, 1982) of the host rock. 

Petrographic observations indicate that the organic material completely coats detrital grain surfaces, except along 

original intergranular contacts (fig. 10), and fills some 

primary and secondary voids. Because it is an abundant 
Figure 5. Scanning electron micrograph of early authigenic 

pyrite (P) in small em bayed areas on detrital quartz grain. Arrows 

mark positions of pyrite grains that were plucked during mount 

preparation. Material in upper left corner comprises later 

authigenic phases. 

Figure 6. Scanning electron micrograph of authigenic, 

smectite-rich, mixed-layer smectite/illite clays. 

Figure 7. Photomicrograph of detrital plagioclase grain dissolved along crystallographic planes. Dissolved areas are filled 

with later introduced organic material (OM) which also coats 

detrital grain surfaces. 

Figure 8. Photomicrograph of detrital potassium feldspar grain 

in which central core was dissolved resulting in secondary void 

(V). 

Figure 9. Photomicrograph of authigenic potassium feldspar 

(adularia(?)) overgrowth (FO) on detrital microcline (M) grain. 

Note organic material (OM) which coats overgrowth as well 

as detrital grain surfaces. 

Figure 10. Photomicrograph of organic material (OM) partly 

filling intergranular pores in host rock. Note presence of organic 

material along sutured grain contacts near center of photo. 

Original intergranular contacts are few (arrows) and noted by 

absence of organic material. 

Diagenesis of the Morrison Formation, Smith Lake Uranium District, McKinley County, New Mexico 257 
authigenic constituent and commonly fills voids, the 

organic material has probably reduced the transmissivity 

of the host rock in ore zones. The presence of organic 

material between neighboring grains which display long, 

concave-convex, and sutured grain contacts-all of which 

are contacts that resulted from compaction of the host 

rocks-suggests that organic material was emplaced prior 

to major compaction. 

When viewed under vertically reflected light in oil, 

the organic material ranges in color from light gray-brown 

to very dark brown and commonly displays various reflectivities. SEM-EDX dot maps indicate that variations in 

AI, Si, and V contents correspond to organic material with 

lower reflectivities (fig. 11). Fission-track maps (fig. 12) 

and SEM-EDX dot maps (fig. 11) reveal that uranium is 

rather homogeneously distributed throughout the organic 

material that coats grains and fills some voids and thereby 

is not responsible for different organic-material reflectivities. Nearly all the uranium in both deposits is in the 

noncellular organic material; some such uranium is in the 

uranium silicate coffinite, which was identified in the claysize fraction (less than 2 ~-tm) by X-ray diffraction. 

Uranium occurrences other than as coffinite or uranoorganic associations are minor. 

Abundant pyrite and marcasite (ore-stage FeS2 

minerals) are intimately mixed with uraniferous organic 

material (fig. 13), suggesting coprecipitation. Framboidal 

pyrite also occurs within organic material. The isotopic 

composition of ore-stage FeS2 is light ( -29.4 to -41.6 per 

mil; fig. 14) and suggests an origin related to bacterial 

sulfate reduction (Goldhaber and others, 1978). SEMEDX and electron microprobe analyses reveal that FeS2 

minerals associated with the organic material commonly 

contain selenium. Textural relationships between the FeS2 

minerals and the organic material suggest that FeS2 

minerals continued to precipitate after emplacement of 

the organic material. 

Semiquantitative X-ray diffraction analyses reveal 

that chlorite is enriched in zones containing noncellular 

organic material. The chlorite, as observed with the SEM 

(fig. 15), occurs as platelets or rosettes, commonly 

intergrown with microcrystalline quartz, coating organic 

material. SEM-EDX analyses also indicate that the 

chlorite of ore zones ubiquitously contains vanadium. In 

addition, SEM-EDX analyses reveal that uranium is also 

commonly associated with the chlorite. Morphologically 

similar chlorite occurs in unmineralized rock of the host 

sandstone, but it is subordinate to other clay minerals 

(P. B. Blackmon, U.S. Geological Survey, Denver, Colo., 

written commun., 1980) and contains no detectable 

vanadium or uranium. 

In ore zones and in the host rock adjacent to ore 

zones, detrital iron-titanium (Fe-Ti) oxide grains (principally magnetite, titanomagnetite, and ilmenite) have 

been extensively altered; iron has been preferentially 

removed, leaving remnant Fe-Ti-oxide or Ti-oxide. The 

258 Studies in Diagenesis 

residual Ti-oxide typically mimics the crystallographically 

controlled positions of ilmenite (fig. 16) in a parent 

titanomagnetite grain (Ramdohr, 1980). Iron-titanium 

oxide grain alteration undoubtedly resulted from postdepositional processes because the observed delicate morphologies of dissolved grains are unlikely to have survived 

sediment transport and deposition. The altered Fe~fi-oxide 

grains closely resemble those described by Adams and 

others (1974) who suggested that destruction of detrital 

Fe-Ti-oxide grains resulted from solutions containing dissolved organic species. Similar solutions are probably 

responsible for the destruction of detrital Fe-Ti-oxide 

grains in the sandstones that host the Mariano Lake and 

Ruby deposits. Organic material in some partly dissolved 

Fe-Ti-oxide grains indicates that at least some, and 

possibly most dissolution preceded the emplacement of 

organic material. 

Alterations following emplacement of organic 

material 

Alterations following emplacement of organic material include the precipitation of calcite, barite, dolomite, 

kaolinite, FeS2 minerals, native selenium, uranium 

minerals, and ferric oxides (goethite(?)). Most of these 

alterations have been noted in volumetrically minor 

amounts in both deposits with one notable exception. 

Calcite is common at the Ruby locality where it occurs 

in and around ore zones, whereas calcite is rare at the 

Mariano Lake locality and is confined to ore zones. 

Calcite commonly fills remaining primary and secondary 

voids, and single calcite crystals commonly envelop several 

detrital grains. The carbon-isotope composition of the 

calcite from both deposits ranges from -7.5 to -12.8 per 

mil (fig. 17). 

Minor amounts of dolomite and barite are present. 

Dolomite was observed only at the Mariano Lake deposit, 

whereas barite was noted at both deposits. Dolomite 

occurs as small rhombs filling some post-ore voids, 

whereas barite occurs as large masses enveloping detrital 

grains. 

Kaolinite is present in ore zones and in the surrounding rock at both localities. The kaolinite fills some 

interstitial voids and replaces calcite (fig. 18) and other 

authigenic minerals in ore zones. 

The remaining alterations which followed the 

emplacement of organic material have been noted chiefly 

along the southern edges of the deposits and in scattered 

localities throughout the deposits. Such alterations were 

likely in response to the encroachment of oxygenated 

ground waters from the south and southeast (see fig. 2). 

FeS2 minerals have been altered to ferric oxides and 

hydroxides. Native selenium, identified petrographically 

and by X-ray diffraction, and observed with the SEM, 

occurs as a separate phase (fig. 19) intergrown with a third 
Figure 11. Scanning electron micrographs of A, multi reflective organic material; and SEM-EDX and dot maps showing the distribution 

of 8, uranium in the same organic material; C, vanadium; 0, silicon; and £, aluminum. 

generation of FeS2 minerals or intergrown with uranium 

minerals which precipitated late in the alteration process. 

Microcrystalline iron oxides (goethite(?)) coat chlorite, 

kaolinite, and native selenium (fig. 20) in ore zones and 

in oxidized rock. 

DISCUSSION 

Organic material was emplaced in the host sandstones of the Mariano Lake and Ruby deposits early in 

the paragenetic sequence of diagenetic alterations but 

Diagenesis of the Morrison Formation, Smith Lake Uranium District, McKinley County, New Mexico 259 
Figure 12. Photomicrograph (right side) of organic material 

coating grains and filling some intergranular pores and corresponding fission-track map (left side). Note that uranium 

(which produced the "tracks") is evenly distributed throughout 

with the organic material. Remaining voids (V) are noted. 

Figure 13. Photomicrograph of organic material (OM) 

intimately mixed with grain of euhedral pyrite (P). 

followed the precipitation of volumetrically minor preore pyrite, mixed-layer clays, and quartz and feldspar 

overgrowths and the selective dissolution of detrital 

feldspars and iron-titanium oxide grains. The introduction of organic material, however, preceded major burial 

compaction of the host sandstone; an early introduction 

of organic material is consistent with the 100 m.y. 

minimum ages determined for the Mariano Lake and 

Ruby ores by uranium-lead isotope methods (Ludwig and 

others, 1984). This conclusion is also consistent with 

evidence from other orebodies in the Grants uranium 

region indicating that lenses of organic material were truncated by pre-Dakota erosion (Nash and Kerr, 1966) and 

that the emplacement of organic material was partly controlled by pre-Dakota folds (Moench and Schlee, 1967). 

The intimate spatial association of uranium with 

organic material indicates that the organic material played 

260 Studies in Diagenesis 

14 

13 

12 D Mineralized 

11 sample: FeS2 

minerals mixed R with organic 

10 material 

Unmineralized R M 

CJ) 9 sample: FeS2 

w minerals as ...J R M 0.. pyrite euhedra 

~ of frambo.ids <! 

CJ) M R M 

u. 7 0 

a: 

w 

CD 6 ~ 

::) 

z 5 

4 

3 

2 

-1 0 -20 -30 -40 -50 

l•s. IN PARTS PER MILLION 

Figure 14. Histogram of sulfur isotope values of Fe52 minerals 

from Mariano Lake (M) and Ruby (R) deposits. 

Figure 15. Scanning electron micrograph of authigenic 

vanadiferous chlorite. 

a crucial role in localizing and concentrating uranium at 

the Mariano Lake and Ruby localities. The occurrence of 

uraniferous organic material completely enclosed by FeS2 

minerals (fig. 13) suggests a simultaneous precipitation. 

The enrichment of uranium in the organic material after 

the formation of the enclosing FeS2 mineral is difficult 

to envision. 
Figure 16. Photomicrograph of dissolved iron-titanium oxide, 

with titanium oxide remnants (T). Organic material (OM) partly 

fills grain and fills interstices. 

(/) 

w 

...J 

a.. 

~ 

<( 

(/) 

u.. 

0 

a: 

w 

co 

~ 

:::::> 

z 

6 r5 -

4 -

3 r2 r-

-

0 

-6 -7 -8 -g -10 -11 -12 -13 

013C, IN PARTS PER MILLION 

Figure 17. Histogram of carbon-isotope values of authigenic 

calcite from Mariano Lake and Ruby deposits. 

Figure 18. Photomicrograph of kaolinite (K) replacement of 

calcite (C). 

Figure 19. Scanning electon micrograph of authigenic native 

selenium (bladed material), superimposed on previously 

precipitated Q, quartz and C, chlorite. 

Figure 20. Scanning electon micrograph of iron mineral (based 

on SEM-EDX analyses), possibly authigenic goethite, in part 

coating native selenium (expanded view of part of fig. 19). 

The enrichment of chlorite in ore zones of the Ruby 

and Mariano Lake deposits is a distinctive feature commonly noted in other primary deposits in the Grants 

uranium region (Granger, 1962; Brookins, 1975). Also, 

chlorite enrichment is spatially and temporally related to 

uranium mineralization at other localities within the Morrison Formation in Utah (Northrup and others, 1981). 

Chlorite in the uranium mineralized zones of the 

Mariano Lake and Ruby localities and in other uranium 

deposits in the Morrison sandstones is commonly 

vanadiferous (Foster, 1959; Brookins, 1975; Northrup and 

others, 1981). The vanadium probably resides in the 

"brucite" interlayer of the chlorite crystal structure (Northrop and others, 1981). In contrast, the large size of 

uranium or uranyl ions relative to that of the cation sites 

within the chlorite crystal structure probably precludes the 

presence of uranium in the chlorite (C. G. Whitney, U.S. 

Geological Survey, Denver, Colo., oral commun., 1982). 

Diagenesis of the Morrison Formation, Smith Lake Uranium District, McKinley County, New Mexico 261 
The association of uranium with chlorite in the ore zones 

of the Mariano Lake and Ruby localities may possibly be 

due to adsorption of uranium or uranium-rich organic 

material on the chlorite or possibly to the presence of submicroscopic uranium minerals intergrown with the 

chlorite. The enrichment of chlorite in ore zones and the 

anomalously high vanadium content of such chlorite suggests that chlorite precipitation may be spatially and temporally related to uranium mineralization. 

The encroachment of oxygenated ground waters 

from the south during late Tertiary time (Saucier, 1980) 

may account for many of the post-mineralization alterations in both deposits. Oxidative destruction of FeS2 

minerals resulted in the replacement of FeS2 minerals by 

ferric oxides and hydroxides and also in the liberation of 

S, Fe, and Se. The redistribution of these elements led to 

the reprecipitation of FeS2 minerals and native selenium. 

The oxidizing ground waters were probably also responsible for the redistribution of minor amounts of uranium 

and iron as uranium minerals and ferric oxides late in the 

paragenetic alteration sequence. 

The formation of calcite at both localities clearly 

postdated emplacement of the organic material and may 

be genetically related to the Tertiary oxidation event. The 

light carbon-isotope composition of the calcite suggests 

that the carbonate carbon was partly derived from the 

oxidation of organic material (see Hoefs, 1973; Leventhal, 

1980). Calcite formation at the Ruby locality was more 

extensive than at the Mariano Lake locality and may be 

related to more severe oxidation of some of the original 

Ruby ore zone. 

Whatever its origin, the calcite may have been 

important in preserving the Ruby deposit by greatly lowering the transmissivity of the ore-bearing sandstone, 

thereby partly blocking passage of oxidizing ground 

waters. The minor amount of calcite at the Mariano Lake 

deposit suggests different conditions for the preservation 

of that orebody. Stagnation of ground waters or deflection of ground-water flow patterns in the synclinal trough 

may have retarded or diverted the ingress of potentially 

destructive oxidizing ground waters. 

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Thrner-Peterson, C. E., and Fishman, N. S., 1986, Geologic synthesis and genetic models for uranium mineralization in the 

Morrison Formation, Grants uranium region, New Mexico, 

in Thrner-Peterson, C. E., Santos, E. S., and Fishman, N. S., 

eds., A basin analysis case study-The Morrison Formation, 

Grants uranium region, New Mexico: American Association 

of Petroleum Geologists Studies in Geology 22, p. 357-388. 

Thrner-Peterson, C. E., Gundersen, L. C., Francis, D. S., and 

Aubrey, W. M., 1980, Fluvio-lacustrine sequences in the 

Upper Jurassic Morrison Formation and the relationship of 

facies to tabular uranium ore deposits in the Poison Canyon 

area, Grants mineral belt, New Mexico, in Thrner-Peterson, 

C. E., ed., Uranium in sedimentary rocks-Application of 

the facies concept to exploration: Rocky Mountain Section 

of Economic Paleontologists and Mineralogists, p. 177-211. 

Wilson, M. D., and Pittman, E. D., 1977, Authigenic clays in sandstones; Recognition and influence on reservoir properties 

and paleoenvironmental analysis: Journal of Sedimentary 

Petrology, v. 47, p. 3-31. 

Zielinski, R. A., and Rosholt, J. N., 1978, Uranium in waters and 

aquifer rocks at the Nevada Test Site, Nye County, Nevada: 

U.S. Geological Survey Journal of Research, v. 6, no. 4, 

p. 489-498. 

Diagenesis of the Morrison Formation, Smith Lake Uranium District, McKinley County, New Mexico 263 

Diagenesis and Uranium Mineralization in the 

Westwater Canyon Member of the 

Morrison Formation, Grants Uranium Region, 

Northwestern New Mexico 

By PAULA L. HANSLEY, U.S. Geological Survey, Box 25046, 

Denver Federal Center, MS 916, Denver, Colorado 80225 

CONTENTS 

Abstract 267 

Introduction 267 

Geologic setting 267 

Materials and methods 267 

Description of major diagenetic phases and alterations 268 

Epigenetic organic matter 269 

Uranium ore 270 

Sulfides 270 

Clay minerals 270 

Chlorite 271 

Feldspars 272 

Silka minerals 273 

Carbonates 274 

Sulfates 274 

Irolll oxides 274 

Etched grains 274 

Paragenesis and geochemistry of diagenetic alterations 275 

Conclusions 278 

References cited 278 

FIGURES 

1. Regional map showing locations of cored drill holes and major uranium 

deposits in the Grants uranium region, New Mexico 268 

2. Cross section of cored holes showing stratigraphic positions of uranium 

mineralization 269 

3. Photomicrograph of a sandstone impregnated with epigenetic carbona-

':eous uranium ore 270 

4. Photomicrograph of altered iron-titanium oxide grain filled with 

{~pigenetic carbonaceous uranium ore 270 

5. Scanning electron micrograph of botryoidal coffinite and coffinite 

':rystals 271 

6. Scanning electron micrograph of authigenic "honeycomb" illite/smectite 

and adularia 271 

7-9. Photomicrographs showing: 

7. Kaolinite replacement of authigenic barite within detrital plagioclase 

grain 271 

8. Kaolinite surrounding remnant primary uranium ore 271 

9. Chlorite rim on partly dissolved detrital plagioclase feldspar 

grain 272 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 265 
10. Scanning electron micrograph showing the relation between chlorite and 

illite/smectite 272 

11-14. Photomicrographs showing: 

11. Detrital potassium feldspar grain with thin albite rim superseded by 

potassium feldspar overgrowth 272 

12. Potassium feldspar overgrowth that supersedes chlorite rim 273 

13. Stages in the dissolution of a detrital plagioclase grain and formation of a "hollow" potassium feldspar grain 273 

14. Detrital twinned plagioclase grain replaced marginally by potassium 

feldspar 274 

15. Scanning electron micrograph of uranium ore that embays a euhedral 

quartz overgrowth 274 

16. Photomicrograph showing two generations of authigenic silica as 

overgrowths on a detrital quartz grain 274 

TABLE 

1. Paragenesis of major authigenic phases and diagenetic alterations in the 

Westwater Canyon Member of the Morrison Formation near Crownpoint, 

New Mexico 275 

266 Studies in Diagenesis 
Abstract 

Petrologic studies of core from the Grants uranium region 

of the southwestern San Juan basin have revealed that the 

uranium-bearing Westwater Canyon Member of the Morrison 

Formation (Upper Jurassic) has had a more complex diagenetic 

history than has previously been documented. Petrographic 

data suggest that most uranium enrichment occurred after 

precipitation of epigenetic organic matter and before significant compaction. 

Authigenic phases characteristic of ore zones include 

epigenetic uranium-bearing, carbonaceous matrix (primary 

ore), euhedral pyrite, and rosette chlorite. Alterations in the 

stratigraphic interval containing uranium ore, but not necessarily 

restricted to mineralized strata, include leached iron-titanium 

oxides, skeletalized plagioclase feldspars, and authigenic 

titanium dioxide, quartz, potassium feldspar, albite, ordered 

mixed-layer illite/smectite (1/S), edge-to-face chlorite, kaolinite, 

ferroan and nonferroan carbonates, and sulfates. Below the 

mineralized interval, postdepositional alterations are less complex: sulfate, carbonate, and smectite are the most common 

interstitial materials. Above the mineralized interval, smectite 

is the dominant authigenic phase. 

Most diagenetic alterations occurred during primary 

mineralization in Late Jurassic to Early Cretaceous time and during widespread leaching of framework grains and authigenic 

cements in the Tertiary Period. Sandstones in the Westwater 

Canyon Member are characterized by multiple generations of 

authigenic minerals, many of which form distinct alteration patterns throughout cores taken from a line of drillholes that 

transect the Grants uranium region. Alteration patterns suggest 

that mineralizing solutions originated in part from compaction 

of fine-grained units in the overlying Brushy Basin Member. In 

the deeper cores, chemically etched heavy minerals and 1/S 

form a distinct alteration zone that may be the result of reactions caused by the incursion of a warm fluid from deeper in 

the basin. Chlorite-dominated alteration zones in barren sandstone units in the Westwater Canyon Member that are stratigraphically equivalent to ore-bearing units may possibly represent sites o:f previous accumulations of primary (carbonaceous) 

uranium ore. 

INTRODUCTION 

Inasmuch as uranium in the Morrison Formation 

(Upper Jurassic) is present in large part as a result of 

diagenetic: processes, interpretation of the postdepositional history of this formation is crucial to understanding the g(:nesis of sedimentary uranium deposits in the 

Grants uranium region, New Mexico. To study diagenesis 

in the Morrison, core was collected from seven drill holes 

that were positioned in a line perpendicular to the southeastwardly trend of uranium deposits in the Grants 

uranium region and parallel to the northeastwardly direction of th1~ dip of the Morrison Formation (fig. 1). Some 

holes were located on trend with known orebodies, and 

others Wt!re positioned where no mineralization was 

thought to be present. In this way, comparisons could be 

made between diagenetic alterations in mineralized strata 

and those in unmineralized strata; previous studies have 

focused almost entirely on uranium-bearing rocks (Adams 

and Saucier, 1981). The present paper summarizes the 

results of new petrologic studies of the cores; the reader 

is referred to Hansley (1986b) for a description of regional 

alteration patterns and a discussion of models for uranium 

mineralization in the Morrison Formation. 

GEOLOGIC SETTING 

In the study area, interbedded massive sandstone 

and thin mudstone units of the Westwater Canyon 

Member of the Morrison Formation are underlain by less 

massive sandstone units of the Recapture Member and 

overlain by the predominantly fine-grained, tuffaceous 

Brushy Basin Member. Sediments of the Recapture and 

Westwater Canyon Members were deposited by meandering and braided streams that flowed in a northeast to 

southeast direction in the San Juan basin in Late Jurassic 

time (Thrner-Peterson, 1986). The overlying Brushy Basin 

Member is composed of fine-grained, tuffaceous lacustrine units (Bell, 1983, 1986) that interfinger locally with 

the upper part of the Westwater Canyon Member. The 

Morrison is underlain by the Middle Jurassic Wanakah 

Formation (Condon and Peterson, 1986) and is separated 

from the overlying Upper Cretaceous Dakota Sandstone 

by a regional unconformity. 

Because the uranium ore (more than 200 ppm 

uranium) found in the cores was entirely in the Westwater 

Canyon Member (fig. 2), this unit became the focus of 

the present study. Sandstones in the Westwater Canyon 

are chiefly subangular to subrounded, medium- to coarsegrained, moderately to well-sorted lithic arkoses and 

subarkoses with poor to excellent porosity (Steele, 1984). 

MATERIALS AND METHODS 

Paragenetic relations were established with a petrographic microscope by examining polished sections and 

thin sections impregnated with blue epoxy and stained 

with sodium cobaltinitrate for potassium feldspar identification. Selected samples were studied using a Cambridge 250 Mark 2 scanning electron microscope (SEM) 

equipped with an energy-dispersive X-ray unit (EDX) and 

an ARL electron microprobe equipped with wavelengthdispersive and energy-dispersive X-ray attachments 1

• 

Aluminum-coated polished thin sections were examined 

with the electron microprobe to verify the presence of 

carbon. Identification of ferroan carbonates was made 

using the staining techniques outlined by Dickson (1965) 

1

Use of trade and company names is for descriptive purposes and 

does not imply endorsement by the U.S. Geological Survey. 

Diagenesis and Uranium Mineralization, Westwater Canyon Member of the Morrison Formation, Northwestern New Mexico 267 
q 

' ' ' 0 0 

~~ SAN JUAN BASIN 

~ X 

0 UJ 

N ~ 

C2 ~ <C UJ 

z ~ 

Gallup. 

EXPLANATION 

J Morrison Formation outcrop 

f Known uranium ore deposits 

e Cored drill holes 

A Measured sections 

CC-12. 

ae 

7,7A. I 

Grants. \ 

I I 

5 10 

10 20 MILES 

I I I 

20 KILOMETERS 

, 

() 

' 

Figure 1. Locations of cored drill holes and major uranium deposits in the Grants uranium region, northwestern New Mexico. 

and Hutchison (1974). Clay minerals were identified by 

X-ray diffraction with a Picker diffractometer using the 

<1-~tm fraction of each sample (Whitney, 1986), and the 

authigenic nature of the clay minerals was verified by SEM 

examination of selected samples. 

DESCRIPTION OF MAJOR DIAGENETIC PHASES 

AND ALTERATIONS 

The complexity and diversity of diagenetic alterations are more pronounced in the Westwater Canyon than 

268 Studies in Diagenesis 

in the overlying Brushy Basin and underlying Recapture 

Members, both of which have had relatively simple 

postdepositional histories. The Brushy Basin contains few 

sandstone units; these units are characterized by a smectite matrix cemented by minor calcite and silica. The 

Recapture sandstone units contain nonferroan calcite, 

anhydrite, and smectite. The major diagenetic alterations 

in upper fine- to medium-grained sandstones of the 

Westwater Canyon Member are described in this section, 

and interpretations relating to the paragenesis of authigenic phases and to the geochemistry of diagenetic fluids 

are presented in the following section. 
PINEDALE 

EAST 

1--v 

z 

Q 

~ ~ 

IE: 

f2 

z 

~ ~ 

IE: 

i 

I--"' 262 

~~ 

3 

-~ 573 

I 

-· \ 

' ' 

4 5 

----- I 741 765 

..., ..... "" ..... ...-

,/" 

,.._~...-

/ / / 

6 7 7A 8 CC-12 

PJf':J ~ BRUSHY _r 

~ BASIN 

--- MEMBER 

907 

I lSI-

~ 1254 

• 1---

'\ • • 

WESTWATER 

\ CANYON \ MEMBER \ -~ L-

~ ~ 

f.- - / lTD 9B5 

~ 613 r-- / 

t=l - -

"? 666 --- 998 

832 

METERS 

~0 15 

20 

25 

ium D

Uran 

Meters ore 

,.,. 

~ 

- TO 661 

TO 328 

EXPLANATION 

- - - Above line magnetic susceptibility 

ia zero (Reynolds and others, 

1986) 

TO= Total depth, in meters 

c::::::J -TO 705 

-

6 ? 

TO 1110 1338 

RECAPTURE 

MEMBER 

~ TO 1348 

..._ r:::c:l - - TO 871 TO 936 TO 1049 

Figure 2. Cross section of cored holes (see fig. 1) showing stratigraphic positions of uranium mineralization. 

Epigenetic organic matter 

Most epigenetic organic matter in the Westwater 

Canyon Member has been mineralized with uranium. 

The organic matter forms grain coatings, fills primary 

pores, and lacks observable cellular structure. The origin 

of the organic matter has been an enigma. Due to radiation damage, the structure of the organic matter cannot 

be determined by conventional pyrolysis techniques 

(Leventhal, 1980). Recent nuclear magnetic resonance 

data, however, suggest that the organic material was derived from degraded plant material or humate (Hatcher 

and others, 1986). The humate may have been intrinsic 

to the Westwater Canyon Member, which contains 

abundant detrital (cellular) plant material from which 

humic acids may have been derived. On the other hand, 

fluids carrying organic acids may have migrated into 

Westwater Canyon sandstones from adjacent organic-rich 

mudstone facies of the Brushy Basin Member during 

compaction (Turner-Peterson and others, 1980; Fishman 

and others, 1984). In either case, the humate was dissolved 

by moderately high-pH fluids derived from the early 

alteration of abundant volcanic ash in the upper part 

of the Morrison Formation. Subsequently, these humic 

acid-bearing solutions migrated through the sandstones and deposited the humate as matrix in the 

sandstones. 

In sandstones impregnated with uraniferous, carbonaceous matrix, floating framework grains and tangential grain contacts are common (fig. 3), and most irontitanium oxide grains have been totally replaced by humate 

(fig. 4). In strata where iron-titanium oxides have been 

leached of iron and replaced by organic matter, magnetic 

susceptibility approaches zero (Reynolds and others, 1986) 

(fig. 2). Authigenic euhedral pyrite and titanium dioxides 

are associated with these altered grains. 

Diagenesis and Uranium Mineralization, Westwater Canyon Member of the Morrison formation, Northwestern New Mexico 269 
Figure 3. Photomicrograph of sandstone impregnated with 

epigenetic carbonaceous uranium ore. 

Uranium ore 

Two basic types of uranium ore have been recognized in the Grants uranium region: primary ore and 

redistributed ore. For the most part, primary ore is composed of a mixture of uranium and epigenetic humate that 

occurs as grain coatings and fills primary pores in the 

fluvial sandstones (Granger and others, 1961; Fishman 

and Reynolds, 1986). Primary orebodies are usually 

tabular, blanketlike deposits that may be suspended in the 

host sandstone without apparent relation to structure or 

bedding. Commonly, no discrete uranium minerals can 

be detected in the primary ore, leading to the conclusion 

that uranium resides in urano-organic complexes. 

Radiogenic ages (U/Pb) indicate that primary uranium 

mineralization occurred about 130 to 110 m.y. ago or during Late Jurassic to Early Cretaceous time (Ludwig and 

others, 1984). Redistributed (secondary) ore formed by 

the destruction of primary orebodies, that is, by the oxidation of organic material and subsequent redistribution of 

the uranium as secondary deposits. In general, the 

distance from primary to redistributed deposits is only 

a few hundred meters. Most redistribution took place during the late Tertiary as a direct result of Late Cretaceousearly Tertiary (Laramide) uplift, when oxidizing fluids 

migrated from exposed outcrops along the southern 

margin of the basin downdip through Morrison sandstones (Saucier, 1980). 

In the present study, uranium was noted as a mixture of primary (carbonaceous) and redistributed or 

secondary (carbon-poor) ore (Hansley, 1986b). Uranium 

in most primary ore appears to be present as formless 

(even under the SEM), noncrystalline material, and 

elemental X-ray fluorescence (dot) maps of the ore revealed that uranium is concentrated on the margins of carbonaceous material (Hansley, 1986b). Scanning electron 

microscopy of some primary ore zones, however, showed 

270 Studies in Diagenesis 

Figure 4. Photomicrograph of altered iron-titanium oxide grain 

filled with epigenetic carbonaceous uranium ore (U). Relict 

ilmenite lamellae (L) are now titanium dioxide. Post-ore calcite 

cement (C) is at lower left of grain. Sample 3-1866; ore zone; 

transmitted and reflected light. 

that uranium is contained in microbotryoidal material, 

which was found to be a mixture of cryptocrystalline coffinite (<1 J.Lm in size) and carbonaceous matrix (fig. 5). 

X-ray diffraction patterns of this material contain weak 

coffinite peaks, and electron microprobe analyses showed 

that carbon was present. X-ray fluorescence (dot) maps 

revealed that uranium is coextensive with silicon rather 

than carbon. This carbon is presumed to be organic carbon because of a lack of calcium or other cations in the 

X-ray fluorescence spectrum. 

Secondary uranium ore contains less organic matter 

and more coffinite than does primary ore. In cores 7 and 7a, 

groups of tetragonal coffinite crystals are intergrown with 

and apparently embay primary ore, and in core 4, micrometer-size, cigar-shaped coffinite crystals protrude from a 

carbonaceous, uranium-bearing substrate (Hansley, 1986b ). 

Sulfides 

Pyrite is the most common sulfide in these sandstones, 

although minor marcasite and sphalerite are present. Anhedral pyrite in the upper part and locally in the lowermost 

sandstones of the Westwater Canyon Member totally cements millimeter-size areas of framework grains, which contain no observable diagenetic alterations. The most abundant morphological variety of pyrite, however, is found 

almost exclusively in ore-bearing intervals. Here, it occurs 

as euhedral grains that are commonly intergrown with primary uranium ore. Euhedral pyrite also occurs on top of 

corroded barite cement and as framboids on rosette chlorite. 

Clay minerals 

Both smectite and ordered, interstratified illite/smectite (1/S) occur as detrital grain coatings that display a 

characteristic honeycomb texture under the SEM (fig. 6). 
Figure 5. Scanning electron micrograph of botryoidal coffinite 

(Cb) and coffinite crystals (Cc). Sample 7-3259; ore zone. 

When viewed with the SEM, the more illitic clays display 

delicate wispy growths on grain boundaries, and X-ray 

fluorescence spectra show the presence of potassium. Clay 

mineral analyses and identifications presented in this 

report were performed by G. Whitney (see Whitney, 1986). 

The liS was identified by the presence of non-integrally 

spaced peaks that shifted position upon glycolation 

(Reynolds and Hower, 1970). It varies in expandability 

(that is, percentage of smectite layers) from 10 to 70 percent; most samples are 40-50 percent expandable and are 

perfectly ordered. Some X-ray diffractograms exhibited 

superlattice reflections at about 27 A. 

Pure (100 percent expandable) smectite is confined 

to unmineralized intervals in the Westwater Canyon 

Member in core 1 and to the Brushy Basin and Recapture Members in the other cores; ordered liS is sandwiched between these smectite-bearing layers in both 

mineralized and barren sandstone units throughout the 

Westwater Canyon Member in cores 3-7a. The mixed-layer 

clay becomes increasingly illitic in a basinward direction 

in the cores of the present study (Whitney, 1986). 

Coarse plates of kaolinite occur as scattered books 

in pores and are most abundant in cores 1-4 near the basin 

margin. Kaolinite embays most other authigenic phases 

(fig. 7) and contains remnants of primary uranium ore 

(fig. 8). Some skeletal plagioclase grains are filled with 

kaolinite, but some fresh plagioclase grains are surrounded 

by pores containing clumps of kaolinite. 

Chlorite 

Both edge-to-face and rosette chlorite occur in the 

Westwater Canyon Member. Edge-to-face chlorite 

aggregates are present chiefly as grain coatings in both 

mineralized and unmineralized strata. Rosette chlorite, on 

Figure 6. Scanning electron micrograph of authigenic 

"honeycomb" illite/smectite and adularia (A). Sample 7-3259; 

ore zone. 

Figure 7. Photomicrograph of kaolinite (K) replacement of 

authigenic barite (B) within detrital plagioclase grain (P). Sample 7-3303; barren zone; plane-polarized light. 

Figure 8. Photomicrograph of kaolinite (K) that surrounds remnant primary uranium ore (U). Grain coatings (G) are 

uraniferous. Sample 3-1866; ore zone; plane-polarized light. 

Diagenesis and Uranium Mineralization, Westwater Canyon Member of the Morrison Formation, Northwestern New Mexico 271 
the other hand, forms clumps of plates in pores and pore 

throats in both barren and mineralized sandstones. Edgeto-face chlorite (1) lines pores that are filled with sulfate 

or carbonate cements, (2) is overgrown by authigenic 

potassium feldspar, and (3) forms rims on skeletal plagioclase grains (fig. 9). Rosette chlorite occurs on quartz 

overgrowths and in pores associated with micrometer-size 

adularia and idiomorphic quartz. Rosette chlorite and I/S 

commonly coexist in the same pore (fig. 10). 

Both types of chlorite are commonly intergrown 

with and form overgrowths on primary ore. Where ore 

(primary) contains the most organic carbon, chlorite is 

rare; where ore (secondary?) is depleted in organic carbon, chlorite is abundant. Irregular millimeter-sized areas 

in the lower part of the Brushy Basin Member are interpreted to have been plant fragments, because electron 

microprobe analyses show that these carbon-rich areas 

have been partially replaced by iron-rich chlorite (data not 

shown here). 

Feldspars 

Several types of feldspar alterations have taken place 

in the Westwater Canyon Member: plagioclase grains are 

partly to totally dissolved; authigenic potassium feldspar 

occurs as overgrowths and idiomorphic crystals in pores; 

potassium feldspar has replaced the margins of plagioclase 

grains; and albite has replaced plagioclase and formed 

overgrowths on all feldspar varieties. In the upper part 

of the Westwater Canyon Member, particularly in cores 

closest to the basin margin, the dissolution of plagioclase 

has resulted in widespread moldic porosity (Hansley, 

1984). Plagioclase has also been dissolved in the other 

(basinward) cores, but remnants of grains are totally 

albitized. 

Albitized feldspar grains occur throughout the 

Westwater Canyon Member but are most common in the 

upper parts of cores closest to the center of the basin 

(Hansley, 1984, 1986b ). Skeletal plagioclase grains are 

preferentially albitized, and completely albitized skeletal 

plagioclase grains were found in core 8. Thin (<5 p.m) 

albite rims occur on both detrital plagioclase and perthitic 

feldspars in optical continuity with the sodic portions of 

these grains. Irt addition, the albite rims underlie potassium feldspar overgrowths (fig. 11), particularly in cores 

5 and 6. The albite rims are interpreted to be authigenic, 

because they are not present at grain contacts. Albite 

overgrowths on plagioclase are most common in core 8. 

Electron microprobe and universal stage analyses 

of authigenic potassium feldspar overgrowths (fig. 12) 

revealed that the feldspar is a type of adularia, specifically 

a potassic sanidine with a disordered structure and a low, 

variable optic angle (Hansley, 1984; 1986b; Smith, 1974). 

Potassium feldspar overgrowths are most abundant in 

the upper part of the Westwater Canyon Member in 

272 Studies in Diagenesis 

Figure 9. Photomicrograph of chlorite rim (Chi) on partly 

dissolved detrital plagioclase feldspar grain (Pf). Note that 

chlorite rim occurs at grain contacts. V, void. Sample 4-1972; 

barren zone; plane-polarized light. 

Figure 10. Scanning electron micrograph showing relation between rosette chlorite (Chi) and illite/smectite (1/S). Sample 3-

1868; ore zone. 

' . . . . . , 

. ·t . 

·.· . . 0.08 mrlf : 

·< •: I • _1 • "'"'. 'l ~ 4~ 1 

Figure 11. Photomicrograph of detrital potassium feldspar grain 

(Kf) with thin albite rim (A) superseded by potassium feldpar overgrowth (Kfo). Sample 6-2604; barren zone; crossed polarizers. 
Figure 12. Photomicrograph of detrital potassium feldspar (Kf) 

with potassium feldspar overgrowth (Kfo) that supersedes 

chlorite rim (Chi). Later calcite cement (C) fills pore. V, void. 

Sample 8-3129; barren zone; plane-polarized light. 

unmineralized cores 5 and 6 and in mineralized core 7; 

they were not found in cores 1 and 8. Cryptocrystalline 

adularia is commonly intergrown with rosette chlorite and 

idiomorphic quartz in barren sandstone intervals. 

All hollow potassium feldspar grains in the Morrison Formation were identified by Austin (1980) as 

detrital sanidine; however, in the present study, most 

hollow sanidine grains were found to have been overgrowths on preexisting detrital plagioclase grains that had 

been dissolved during diagenesis (fig. 13 A-C). None of 

these hollow grains showed cathodoluminescence, 

indicating that they are authigenic (Kastner and Siever, 

1979), and electron microprobe analyses revealed that they 

are nearly pure, stoichiometric KA1Sip8 (Hansley, 1984). 

Detrital sanidine, which occurs throughout these 

volcaniclastic sandstones, shows extensive dissolution 

features and enrichment of potassium at the expense of 

sodium around dissolution voids. Many detrital plagioclase grains are replaced marginally by potassium feldspar 

(fig. 14), which is composed of intergrown adularia 

crystals that have different optical orientations. 

Silica minerals 

Authigenic silica is ubiquitous in the upper part of 

the Westwater Canyon Member. As many as four generations of authigenic silica cement were observed in the same 

Figure 13. Photomicrographs showing stages of dissolution of 

detrital plagioclase grain and formation of "hollow" potassium 

feldspar grain: 

A, Detrital plagioclase (Pf) with euhedral potassium feldspar 

overgrowth (Kfo) (sample 6-2624, barren); 

8, partially dissolved plagioclase grain (Pf) with dissolution voids 

(V) and potassium feldspar overgrowth (Kfo) (sample 7-3129, 

barren); and 

C, potassium feldspar overgrowth (Kfo) surrounding a void (V) 

where a detrital plagioclase grain (Pf) once occurred (sample 

3-1833, barren). All photomicrographs taken in plane-polarized 

light. 

Diagenesis and Uranium Mineralization, Westwater Canyon Member of the Morrison Formation, Northwestern New Mexico 273 
pore. The most common form of authigenic silica is 

quartz overgrowths that may underlie or, commonly, 

overlie grain coatings of primary uranium ore (fig. 15). 

Double quartz overgrowths have been noted on the same 

detrital quartz grain (fig. 16). Other types of authigenic 

silica are (1) microcrystalline idiomorphic quartz that 

occurs with rosette chlorite and adularia in the pores of 

unmineralized sandstone units, and (2) noncrystalline 

silica that forms spheres (opal-CT(?)) associated with coffinite in cores 7 and 7a. 

Carbonates 

Several generations of nonferroan and ferroan carbonates occur in the Westwater Canyon Member. Minor 

amounts of nonferroan calcite cement occur inside concretions; however, most nonferroan calcite fills pores 

rimmed with primary ore and chlorite (see fig. 9). Nonferroan calcite embays primary ore and is, in turn, embayed 

locally by ferroan calcite and ankerite. Ferroan calcite and 

ankerite contain inclusions of carbonaceous uranium ore 

in cores 7 and 7a. 

Sulfates 

Anhydrite and barite occur as interstitial cements 

and as infillings or replacements of detrital plagioclase 

(see fig. 7). Anhydrite is common in the deeper parts of 

cores 5-7, but it is rare in the other cores. Barite occurs 

sporadically throughout all cores. Both sulfates embay 

other authigenic phases with the exception of kaolinite. 

In core 7 a, barite cement is pink due to oxidized pyrite 

inclusions, and the barite is overgrown with euhedral pyrite 

crystals. 

Iron oxides 

Thin hematitic rims occur on iron-titanium oxide 

grains below the magnetic susceptibility line (see fig. 2}, 

and hematite is present inside reduced mudstone nodules. 

Hematite rosettes occur inside dissolution cavities in 

skeletal plagioclase and on chlorite rosettes in core 4. 

Etched grains 

Deep chemical etching has produced cubic and 

dodecahedral faces on detrital almandine-spessartine 

garnets in certain stratigraphic intervals within the 

Westwater Canyon Member (Hansley, 1986a). Deeply 

etched staurolite grains occur in the same zone: Garnets 

are rounded, not etched, in the overlying Brushy Basin 

and the underlying Recapture Members. The sharpness 

274 Studies in Diagenesis 

Figure 14. PhotomiCrograph of detrital, skeletal twinned 

plagioclase grain replaced marginally by authigenic potassium 

feldspar (Kfa). V, void. 

Figure 15. Scanning electron micrograph of uranium ore (U) 

that embays euhedral quartz overgrowth (Qo). Sample 7-3259; 

ore zone. 

Figure 16. Photomicrograph showing two generations of 

authigenic silica (Qo1

, Qo) as overgrowths on detrital quartz 

grain (Q). Sample 5-2802, barren zone; crossed polarizers. 
of the etched surfaces and the limited distribution of 

etched grains preclude a detrital origin for these textures 

(Hansley, 1983; Hansley, 1985; Hansley, 1986a). The zone 

of deeply etched garnets coincides with the zone of 

ordered 1/S in the core fence. 

PARAGENESIS AND GEOCHEMISTRY OF 

DIAGENETIC ALTERATIONS 

A paragenetic sequence for the Westwater Canyon 

Member based on the information presented in this report 

appears in table 1. The relationships of various authigenic 

phases to each other and to uranium ore are discussed 

in the following paragraphs. The chemistry and temperatures of interstitial fluids are inferred based on the 

observed petrology. 

During early diagenesis, the chemistry of pore water 

in the upper part of the Morrison Formation was controlled largely by the dissolution of tremendous volumes 

of volcanic ash that fell into streams and lakes in the area 

of the San Juan basin during Late Jurassic time. Due to 

hydrolysis and dissolution of the ash, the resultant pore 

waters were buffered to a moderately basic pH. Where 

organic matter was present in mudstones and in fluvial 

channels, the early (pre-compaction) geochemical environment was reducing. The occurrence of ferric oxides in 

many parts of the Morrison, however, suggests that conditions were at least mildly oxidizing in some units. Smectite and minor anhedral pyrite, nonferroan calcite, quartz 

overgrowths, and kaolinite precipitated before significant 

compaction and the emplacement of primary ore. Early 

diagenetic kaolinite is interpreted to be that which occurs 

within volcanic rock fragments. 

Many alterations-especially those involving 

feldspars and iron-titanium oxides-that are most intense 

near the top of the Westwater Canyon Member suggest 

that fluids migrated into the Westwater Canyon from 

overlying Brushy Basin units. The optimum time for 

downward fluid migration was during early diagenesis, 

when the fine-grained Brushy Basin sediments were compacted. Movement of fluids from fine-grained Brushy 

Basin units into the coarser and more permeable sandstones in the upper part of the Westwater Canyon Member 

was first suggested by Turner-Peterson and others (1980) 

based on detailed sedimentologic studies. 

The near-zero magnetic susceptibility that is directly 

related to iron-titanium oxide alteration in the upper part 

of the Westwater Canyon Member has been attributed to 

the early postdepositional movement of organic acidbearing solutions through the sandstones (Reynolds and 

others, 1986; Adams and Saucier, 1981). Euhedral pyrite 

and titanium dioxides, chiefly anatase, were products of 

iron-titanium oxide diagenesis. The paucity of uranium 

within carbonized iron-titanium oxides and the enrichment of uranium on the margins of these altered grains 

imply that most uranium was introduced after emplacement of the organic material. This observation contradicts 

earlier studies that concluded uranium and carbon were 

precipitated at the same time, based on a uranium:carbon ratio of 1:1 (by weight) in many ore zones (Leventhal, 1980; Adams and Saucier, 1981). 

Table 1. Paragenesis of major authigenic phases and diagenetic alterations in the Westwater Canyon Member of the Morrison Formation near Crownpoint, New Mexico 

[Dashed lines indicate tentative interpretations] 

Calcite ••••••••••••••••••••••• 

Smectite •••••••••••••••••••••• 

Illite/smectite ••••••••••••••• 

Quartz •••••••••••••••••••••••• 

Potassium feldspar •••••••••••• 

Albite •.........•........••... 

Heavy-mineral etching ••••••••• 

Fe-Ti-oxide alteration •••••••• 

Titanium dioxides ••••••••••••• 

Pyrite •..........•.••.•••.•... 

Organic matter •••••••••••••••• 

Uranium ore ••••••••••••••••••• 

Chlorite •••••••••••••••••••••• 

Anhydrite ••••••••••••••••••••• 

Barite •••••••••••••••••••••••• 

Kaolinite ••••••••••••••••••••• 

Iron oxides ••••••••••••••••••• 

Ferroan carbonates •••••••••••• 

Late Jurassic-Early Cretaceous 

Primary 

Tertiary 

Secondary 

Diagenesis and Uranium Mineralization, Westwater Canyon Member of the Morrison Formation, Northwestern New Mexico 275 
Primary uranium mineralization occurred during 

early diagenesis when the inferred hydrologic communication between the two members would have taken place. 

Fluids migrating through the Westwater Canyon Member, 

possibly enriched by solutions migrating out of the Brushy 

Basin units, carried soluble humic acids and uranium to 

sites of precipitation. In this way, tabular primary ore 

deposits were formed. Additional uranium enrichment 

apparently occurred after the deposition of epigenetic 

humate as uranium-bearing solutions continued to 

migrate into humate-rich zones, where the uranium was 

continually reduced and precipitated in and around the 

margins of humate-impregnated areas. Multiple generations of authigenic phases in ore-bearing strata reflect the 

extended duration of the mineralization event. 

The presence of coffinite and opal-CT(?) in some 

orebodies suggests that uranium and silicon traveled 

together in solution as a hydroxy-uranyl-silicate complex. 

Such a complex is stable at low temperatures in neutral 

to mildly alkaline solutions in which the carbonate activity 

is less than the silica activity (Yermolayev and others, 

1965). The silica activity was apparently high during early 

diagenesis due to the dissolution of volcanic ash and 

possibly stabilized the hydroxy-uranyl-silicate complex. 

According to Hemingway (1982), a high silica activity (at 

least 10-3

·

5

) favors the formation of coffinite over 

uraninite. The cryptocrystalline coffinite that is coextensive with carbonaceous matrix may have formed in part 

during primary mineralization where silica activity was 

high and in part during later diagenesis when extensive 

dissolution of aluminosilicates occurred. A later generation of coffinite is represented by coarse-grained coffinite 

that has replaced primary (carbonaceous) ore in cores 7 

and 7a. This coffinite may have formed as a result of 

recrystallization of cryptocrystalline coffinite in the 

primary ore. 

The widespread occurrence of altered tuffs and bentonites in the upper part of the Morrison Formation suggest that smectite was an alteration product of rhyolitic 

volcanic ash (Brookins, 1975). Sometime during later 

diagenesis, smectite probably transformed to liS as is suggested by the similarity in textures of the smectite and the 

1/S and by the occurrence of both clay minerals on detrital 

grain rims. Because primary uranium ore and most other 

authigenic phases commonly overlie 1/S (Fishman and 

Reynolds, 1986), the mixed-layer clays probably formed 

during early diagenesis. Random 1/S may form at 

temperatures as low as 60 OC; however, an ordered structure requires temperatures of at least 1000C (Perry and 

Hower, 1970). 

Organic acids may have facilitated the illitization 

reaction at temperatures less than 60°C (Brookins, 1975) 

by reducing ferric to ferrous iron in the smectite, a process that creates a charge deficiency in the structure (Perry 

and Hower, 1970) and that has been demonstrated to be 

a viable process by Eslinger and others (1979). This 

276 Studies in Diagenesis 

negative charge can be balanced by the incorporation of 

potassium as an interlayer cation into the structure. High 

potassium activity due to the dissolution of abundant 

rhyolitic volcanic ash may have also promoted early 

transformation of smectite to liS. Recent studies suggest 

that unusual solution chemistry, such as extremely high 

potassium activity in a saline, alkaline lake environment, 

promotes the transformation of smectite to 1/S at 

temperatures less than 60°C (Singer and Stoffers, 1980). 

Ordering of mixed-layer clays, as mentioned 

previously, requires higher temperatures. The highest 

temperature that the Morrison Formation could have 

attained due to burial diagenesis is about 65 OC according 

to the inferred geothermal gradient of 20°C/km (Scott, 

1986), because the Morrison was buried to a depth of only 

about 2000 m during the mid-Tertiary. Even if burial had 

been deep enough to produce a temperature of 100°C, 

the restricted occurrence of the ordered clays still cannot 

be explained. An alternative explanation is that a warm 

fluid migrated updip through the Westwater Canyon 

Member and caused the transformation of smectite to 

ordered 1/S (Whitney, 1986) during maximum burial in 

the mid-Tertiary when oil is known to have been generated 

in the San Juan basin (Rice, 1983). The presence of a 

warm fluid is also indicated by oxygen and deuterium 

isotope data on the liS clay minerals (Northrop and 

Whitney, 1985). This hypothesis also accounts for the 

restriction of ordered clays to the Westwater Canyon 

Member. Warmer temperatures would also have favored 

albitization of plagioclase, because temperatures of 1000C 

are thought by Boles (1982) to be required for albitization, based on his studies of sediments in the Gulf Coast 

region. The close correspondence between the etched 

garnet and liS zones indicates that the etching of garnets 

may have also resulted from reactions involving a higher 

temperature fluid (Hansley, 1985). 

Several chemical factors may have interacted, therefore, to promote illitization of smectite during early 

diagenesis of these sandstones, but ordering of the 1/S 

probably occurred later in diagenesis due to increased 

temperature. 

Most kaolinite formed late in diagenesis, although 

minor kaolinite formed as an alteration of rock fragments 

during early diagenesis. Some skeletal plagioclase grains 

are partly replaced(?) by kaolinite, but fresh plagioclase 

grains within nests of kaolinite suggest that the alteration 

of feldspars was not the only source for aluminum and 

silica in the kaolinite (Austin, 1980). 

The abundance of pyrite in ore zones has been 

attributed to the reducing environment caused by humic 

material. During early diagenesis, sulfate was reduced by 

anaerobic microbial activity, and the reduced sulfur combined with iron to form pyrite (Leventhal, 1980). The 

paucity of pyrite and abundance of sulfate in co.res 7 and 

7a suggest that oxidation of pyrite may have led to sulfate 

precipitation. Overgrowths of euhedral pyrite on barite 
imply a resulfidization or re-reduction of these sandstones. 

This later reduction may have been associated with 

redistribution of uranium into roll-type deposits along an 

oxidation-reduction front (Rhett, 1980). 

Chlorite precipitated prior to significant grain 

dissolution, because partly dissolved grains have chlorite 

rims that outline the original grain boundaries. Rosette 

chlorite and I/S were probably cogenetic phases (Whitney, 

1986}, although paragenetic relations between the two 

phases are ambiguous. Etched garnets within chlorite 

rims and rosette chlorite on etched surfaces suggest that 

edge-to-face chlorite precipitated before rosette chlorite, 

and therefore, represents an earlier stage of chlorite 

formation. 

The superposition of chlorite rosettes on primary 

uranium ore and intergrowths of chlorite and ore suggest 

that the chlorite formed during mineralization or, alternatively, was a later alteration product of primary ore. 

This latter hypothesis is favored by the following 

petrographic observations: in primary ore zones, ore coats 

grain rims; in barren rock surrounding primary ore zones, 

chlorite coats the detrital grains, thus filling the same 

niche as the primary ore in the ore zones. In high-grade 

primary ore, which is generally deficient in chlorite, 

vanadium was detected in areas where organic carbon was 

present; where chlorite rosettes are abundant, the chlorite 

is vanadiferous (Whitney, 1986). Thus, the vanadium in 

the organic matter was probably incorporated into the 

chlorite structure during diagenesis of the organic matter. 

Inclusions of primary uranium ore within ankerite 

cement suggest that ferroan carbonates also formed during the alteration of carbonaceous uranium ore. Supporting this interpretation by indicating that some of the carbonate carbon was derived from organic carbon are light 

o13C values (-8 to -12) for ankerite (Hansley, 1986b) and 

similar light o13C values for calcite from ore zones in the 

Ambrosia Lake District in the Grants uranium region 

(Leventhal, 1980). 

Delicate remnants of plagioclase grain rims show 

no effects of mechanical compaction; thus, some burial 

was apparently attained before dissolution took place. 

Furthermore, plagioclase skeletalization apparently 

postdated the formation of primary ore, because ore rarely 

occurs within skeletal plagioclase grains, and, in areas 

totally impregnated with ore, plagioclase grains are 

generally unaltered. Chlorite rims on the edges of skeletal 

plagioclase grains and around the edges of moldic pores 

containing plagioclase remnants indicate that plagioclase 

dissolved after chlorite precipitated. Plagioclase skeletalization postdated the formation of potassium feldspar 

overgrowths, because skeletal grains commonly occur 

totally within the overgrowths. 

Recent experiments on the dissolution of silicates 

indicate that organic acids are probably the best naturally 

occurring solvents of aluminosilicate minerals (Surdam 

and others, 1984). During the mid-Tertiary, solutions 

containing organic acids and carbon dioxide (as carbonic 

acid) were produced during the maturation of humate. 

These solutions were probably responsible for extensive 

dissolution of framework grains and cement and the 

development of moldic secondary porosity in the Westwater Canyon Member. Because plagioclase remnants are 

commonly albitized, albitization either predated or was 

coincident with plagioclase dissolution. Albitization of 

the plagioclase remnants probably served to stabilize the 

skeletal grains. If warm fluids did indeed migrate updip 

through the Westwater Canyon Member during deepest 

burial, the increased temperature would have hastened the 

maturation of the organic material and, thus, would have 

facilitated production of organic acids. 

Both authigenic albite and analcime occur in 

dissolution voids within altered plagioclase grains in 

core 8. Because zeolites commonly form as alteration 

products of volcanic glass in saline, alkaline lacustrine 

environments, the analcime may have formed indirectly 

as an alteration product of volcanic glass; however, 

because analcime has not been observed coexisting with 

unaltered volcanic ash, it is therefore thought to form 

from a precursor zeolite, such as clinoptilolite (Hay, 1966). 

During subsequent burial diagenesis, analcime probably 

transformed to albite (Liou, 1971). In geochemical studies 

of the Grants uranium region, Della Valle (1981) found 

that sodium was somewhat enriched in the upper part of 

the Westwater Canyon Member. These data suggest that 

sodium may have been derived from altered tuff beds in 

the overlying lacustrine facies of the Brushy Basin 

Member. Because of the close spatial association of 

skeletal plagioclase and authigenic albite, however, some 

sodium was probably derived locally from alteration of 

detrital plagioclase grains. 

Potassium feldspar probably formed after the formation of the primary uranium ore, because overgrowths 

of this mineral supersede (post-ore) chlorite grain rims (see 

fig. 12). In addition, where primary ore coats detrital 

grains, the ore does not occur on potassium feldspar 

overgrowths (Hansley, 1986b ). 

The various stages of authigenic silica formation 

suggest that pulses of silica-saturated pore waters migrated 

through these sandstones. These pore waters may have 

been derived from the compaction of tuffaceous lacustrine 

sediments in the overlying and intertonguing Brushy Basin 

Member, as has been suggested by Turner-Peterson and 

others (1980) and Thrner-Peterson (1986). Silica may have 

precipitated when these alkaline waters mixed with formation waters of lower pH and/or encountered accumulations of organic matter in the Westwater Canyon Member. 

In addition, some silica was probably derived by the 

dissolution of detrital silicates and volcanic ash in the 

Westwater Canyon Member itself. 

Widespread post-ore oxidation of primary uranium 

ore during the late Tertiary (Saucier, 1980) obliterated 

evidence of earlier diagenetic reactions. At this time 

Diagenesis and Uranium Mineralization, Westwater Canyon Member of the Morrison Formation, Northwestern New Mexico 277 
meteoric water flowed from recharge areas along the 

southern margin of the basin through Westwater Canyon 

sandstone units toward the center of the basin. Oxidized 

ore zones occur in cores closest to the basin margin. 

Petrographic evidence for a late oxidizing event includes 

relicts of carbonate and sulfate cements, remnants of 

uraniferous and carbonaceous matrix, hematite rosettes 

on chlorite, and floating detrital grains in enlarged 

pores. During the leaching process, coarse crystalline 

kaolinite precipitated in pores and within leached grains. 

Some coarse, void-filling coffinite probably formed 

at this time, when the organic matter associated with 

primary ore oxidized to carbon dioxide, and uranium 

formed soluble uranyl complexes, which precipitated as 

secondary ore in reduced zones downdip. Uranium-lead 

apparent ages of near zero indicate that redistribution 

may be occurring at the present time (Ludwig and others, 

1982). 

CONCLUSIONS 

Diagenetic alterations in both mineralized and barren sandstone units of the Westwater Canyon Member 

of the Morrison Formation in the Grants uranium region 

are more complex than alterations in underlying and 

overlying members. The complexity can be attributed 

largely to processes associated with both primary uranium 

mineralization and the later alteration of primary (carbonaceous) ore, because most diagenetic alterations occur 

in mineralized strata or in rocks stratigraphically 

equivalent to mineralized strata. Alteration patterns suggest that some fluids migrated into the Westwater Canyon Member from overlying mudstones in the Brushy 

Basin Member, and the presence of certain authigenic 

phases (such as ordered I/S, etched garnets, and albite) 

confined to the Westwater Canyon implies that warm 

fluids migrated updip through the Westwater Canyon 

sandstones perhaps during the mid-Tertiary. Finally, during the late Tertiary, oxidizing fluids that migrated 

downdip from exposed outcrops caused extensive alteration of primary ore deposits and redistributed uranium 

into organic-poor secondary deposits. Both Tertiary 

events were responsible for the formation of widespread 

secondary porosity through the dissolution of grains and 

cements. 

As the organic matter in primary ore altered during diagenesis due to microbial activity, warm(?) temperatures, and oxidizing solutions, chlorite, ferroan 

carbonates, and coffinite apparently precipitated in its 

place. Therefore, as was first suggested by Squyres 

(1980), chlorite-bearing assemblages in barren strata may 

represent former sites of primary ore accumulation. If 

this hypothesis is correct, primary orebodies may have 

once been more widespread in the Grants uranium 

region. 

278 Studies in Diagenesis 

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Turner-Peterson, C. E., Santos, E. S., and Fishman, N.S., 

eds., A basin analysis case study-The Morrison Formation, Grants uranium region, New Mexico: American 

Association of Petroleum Geologists Studies in Geology 22, 

p. 315-330. 

Y ermolayev, N. P., Zhidikova, A. P., and Zarinskiy, V. A., 

1965, Transport of uranium in aqueous solutions in the 

form of complex silicate ions: Geokhimiya, no. 7, 

p. 813-826. 
Diagenesis and the Sources of the Ore-Forming Fluids, 

Tony M. Vanadium-Uranium Deposit, 

Henry Structural Basin, Utah 

By H. ROY NORTHROP, MARTIN B. GOLDHABER, C. GENE WHITNEY, 

FRED PETERSON, RICHARD L. REYNOLDS, and JOHN A. CAMPBELL, 

U.S. Geological Survey, Box 25046, Denver Federal Center, MS 963, 

Denver, Colorado 80225 

CONTENTS 

Abstract 282 

Introduction 282 

Acknowledgments 282 

Geologic setting 282 

Stratigraphy and lithology 282 

Age of mineralization 282 

Temperature of mineralization 282 

Experimental methods 283 

Results 284 

Mineralogy of the deposits 284 

Stable isotope geochemistry 284 

Discussion 286 

Conclusions 289 

References cited 289 

FIGURES 

1. Map of Henry basin showing location of Tony M. mine 283 

2-3. Charts showing: 

2. Distribution of ore-stage dolomite 284 

3. o180 values of calcite and dolomite 285 

4-5. Histograms of: 

4. o180 values of calcite and dolomite 285 

5. Calculated o180 and oD values of waters in equilibrium with clay 

minerals 286 

6-8. Charts showing: 

6. Calculated o180 versus oD composition of waters in equilibrium with 

authigenic clay minerals 286 

7. Variation of calculated o180 values of water in equilibrium with 

authigenic clay minerals with depth 287 

8. Time versus rate of formation of ore horizon 288 

In Memoriam 

H. Roy Northrop 3/9/48-6/24/87 

F. A. Mumpton, ed., 1986, Studies in Diagenesis: U.S. Geological Survey Bulletin 1578 281 
Abstract 

Tabular-type vanadium-uranium deposits occur in the 

Upper Jurassic Morrison Formation in the Henry structural 

basin, Utah. Mineralization took place at and immediately 

below an interface between a basin-derived brine and meteoric 

water. Based on isotopic compositions calculated from 

authigenic clay minerals, one of the two waters appears to have 

been a connate basinal brine, and was probably derived from 

dewatering of the evaporite-mudflat sediments subjacent to the 

mineralized strata. The other solution appears to have been 

meteoric water that migrated downward through the superjacent sediments shortly after their deposition. Vanadium and 

uranium, soluble in their 4 + and 6 + valance states, respectively, were carried in the meteoric water, were reduced at and 

immediately below the brine-meteoric water interface, and were 

precipitated as coffinite, vanadium-rich chlorite, and vanadium 

oxyhyd roxide. 

INTRODUCTION 

Genetic models for the formation of tabular-type 

vanadium-uranium deposits of the Colorado Plateaus 

province have been proposed by many workers (see 

Thamm and others, 1981; Finch, 1967 for a history). Of 

all these models, only Fischer (1947), Shawe (1956), and 

Granger (1976) proposed ore deposition at the interface 

between a brine and fresh water. Key data to support this 

brine-interface model or any of the proposed genetic 

models have been lacking until our work. 

This paper summarizes chemical, mineralogic, and 

isotopic data from one tabular-type vanadium-uranium 

deposit, the Tony M. orebody in the Henry basin in southcentral Utah. It focuses on data that constrain the sources 

of the waters involved in the ore-forming process and 

emphasizes the role of diagenesis in the formation of this 

deposit. More detailed reports on this subject have been 

published elsewhere (see Northrop and others, in press a, 

in press b; Goldhaber and others, in press; Northrop and 

Goldhaber, in press). 

Acknowledgments 

Plateau Resources Ltd. allowed unlimited access to 

both core and mine samples and shared insight and ideas 

that made this study possible. 

GEOLOGIC SETTING 

Stratigraphy and lithology 

The Tony M. vanadium-uranium deposit is hosted 

by fluvial sandstones of the Salt Wash Member of the Upper Jurassic Morrison Formation in the Henry structural 

282 Studies in Diagenesis 

basin, Utah (fig. 1; and Peterson, 1980). In general, the 

orebody is typical of tabular vanadium-uranium deposits 

of the Colorado Plateaus province and consists of two 

(or locally three) closely spaced vanadium-uranium-rich 

layers separated by an interval barren of uranium but 

enriched in vanadium (Carpenter, 1980). The package of 

rocks containing the uranium ore zones and the vanadium-rich, uranium-poor (barren) zones is referred to as 

the mineralized interval. The tabular (peneconcordant) 

mineralized interval crosses stratigraphic boundaries at a 

shallow angle and climbs stratigraphically to the northeast 

down the inferred paleohydrologic gradient (Peterson, 

1978, 1980). The westernmost part of the Tony M. orebody 

intersects a lower bounding mudstone, the Tidwell 

Member of the Morrison Formation; the remainder of the 

orebody is suspended within sandstones of the Salt Wash. 

The mineralized interval contains anomalous concentrations of organic matter (dominantly detrital plant debris), 

magnesium, molybdenum, selenium, aluminum, and isotopically light pyritic sulfur, in addition to the vanadium 

and uranium (Carpenter, 1980; Northrop, 1982). 

Age of mineralization 

Attempts to date the vanadium-uranium mineralization in the Henry basin have been disappointing. 

Uranium-lead ages are all discordant, apparently minimum ages, that constrain the time of mineralization to 

latest Jurassic or earliest Cretaceous, but allow ages approaching that of the host rocks (K. R. Ludwig, U.S. 

Geological Survey, Denver, Colorado, written commun., 

1982). Rb-Sr ages determined on the authigenic chlorites 

indicate the same time period, but the chlorite appears to 

have remained an open system with respect to strontium, 

resulting in large errors and unreliable ages (K. R. Ludwig, 

oral commun., 1984). Geologic constraints (Peterson, 

1978, 1980; Northrop, 1982) suggest that the vanadiumuranium mineralization took place early in the diagenetic 

history of the host rock. Petrographic observations of 

Goldhaber and others (in press) support this suggestion. 

Temperature of mineralization 

In recent years considerable evidence has accumulated to indicate that vanadium-uranium mineralization 

took place at low temperatures (see, for example, Nash 

and others, 1981). With respect to the Tony M. orebody, 

additional evidence supporting low-temperature mineralization includes the presence of ore-stage bacteriogenic 

pyrite (Carpenter, 1980; Northrop, 1982; Goldhaber and 

others, in press); such sulfate-reducing bacteria live at temperatures less than 80 OC. The occurrence of pre-ore smectite (Northrop, 1982) suggests that these rocks probably 

have never been subjected to temperatures greater than 

about 60°C (Reynolds, 1980). Geologic reconstruction 
110°00' 110°45' 

38°30' /--- -......_., I 

/ " I ' \ UTAH 

I \ Area of 

I \ figure 1 ~ 

I \ 

I \ 38°15' 

\ \ 

\ HENRY BASIN \ 

\ \ 

\ I 

\ 

\ I 

38°00' ~\ I ~, 

~ I ~, 

~ I -o\ 

~, I 

I 

\ TONY M. MINE I 

37°45' 

\ ~ I ' ' / -- __., 

Figure 1. Map showing margin (dashed line) of the Henry basin, south-central Utah, which is a doubly plunging 

syncline, that forms a structurally closed basin. Also shown is location of Tony M. mine, in the Salt Wash Member 

of the Upper jurassic Morrison Formation. Circled numbers, State highways. 

of the sediment pile (Peterson, 1980) places the maximum depth of burial for the mineralized interval at less 

than 1 kilometer. Thus, at any reasonable geothermal 

gradient, vanadium and uranium must have been 

deposited at temperatures less than 50 °C. Temperaturedependent, stable-isotope data are presented here for 

25 °C; however, differences in the actual temperature, 

within the described constraints, will not change their 

interpretation. 

EXPERIMENTAL METHODS 

All samples were obtained from cores drilled below 

the static water table. Bulk samples were disaggregated 

by hand in a mortar and pestle. Oxygen isotope analyses 

of carbonate minerals were made on these disaggregated 

samples. Clay mineral separations were made by soaking 

the disaggregated samples in distilled water and sonifying them to separate the clay-size particles from other 

Diagenesis and Sources of Ore-Forming Fluids, Tony M. Vanadium-Uranium Deposit, Utah 283 
phases. The clay fraction (less than 2 micrometer (mp.)) 

was then separated by centrifugation. Oriented slides were 

prepared for X-ray diffraction analysis using the method 

described by Drever (1973). Clay-size fractions were prepared for oxygen and hydrogen isotope analyses using the 

techniques described by Clayton and Mayeda (1963) and 

Savin and Epstein (1970). Oxygen isotope analyses of 

dolomites were made after first separating calcite from 

dolomite by the extraction procedure of Walters and others 

(1972). All isotope data are presented in the familiar delta 

(o) notation, relative to Standard Mean Ocean Water 

(SMOW) (oxygen and hydrogen), PeeDee Formation 

Belemnite (PDB) (carbon), and Canyon Diablo Troilite 

(CDT) (sulfur). 

RESULTS 

Mineralogy of the deposit 

Based on petrographic, scanning electron microscopic (including energy-dispersive X-ray analysis), and 

microprobe analyses, the uranium in the Tony M. orebody 

is predominantly as the mineral coffinite. Vanadium is present in the mineralized interval both as an oxide mineral 

and in a vanadium-bearing chlorite (Whitney and Northrop, in press), which is the dominant clay mineral in the 

mineralized interval. Vanadium-rich, interstratified 

chlorite/smectite is laterally continuous with the mineralized interval, but areally beyond the economic limits of 

the orebody. Below the mineralized interval the predominant authigenic clay mineral is interstratified illite/smectite containing more than 7 5 per cent expandable layers. 

The predominant clay minerals above the mineralized 

interval are kaolinite in sandstones and smectite in bentonitic horizons. Smectite, in both the Tidwell and unmineralized Salt Wash Members, predates mineralization, 

whereas the kaolinite postdates the mineralization. 

Pore-filling carbonate minerals include pre-ore and 

post-ore calcite and ore-stage dolomite that occurs as a 

micritic cement within the mineralized interval and as 

euhedral rhomboids above the mineralized interval 

(Goldhaber and others, in press). The dolomite content 

of the host rock increases systematically upward from the 

mineralized interval and reaches a maximum about 6 

meters above the mineralized interval. On the other hand, 

the dolomite content of the mineralized interval is low and 

represents a local minimum in the dolomite content of 

the host rock (fig. 2) (Northrop and others, in press a). 

Stable isotope geochemistry 

The o180 of the dolomite cements decreases 

upward from the zone of minimum dolomite in the mineralized interval, through the zone of maximum dolomite 

284 Studies in Diagenesis 

120 

130 

fJ) 140 

a: 

w 

tu 

:E 

~ 

::i 150 

h: 

w 

0 

160 

170 

40 60 80 100 

Dolomite 

Total carbonate X 100 

Figure 2. Distribution of authigenic, ore-stage dolomite 

cement in core 3C, which penetrated the mineralized interval of the Tony M. orebody. Dolomite abundance is shown 

as per cent of total carbonate that is dolomite, as determined 

semiquantatively by X-ray diffraction. 

above the mineralized interval (fig. 3). Oxygen isotope 

values of the dolomite in the mineralized interval are between + 19 and + 22 per mil (SMOW) and decrease to 

between + 17 and + 18.5 per mil in the zone of maximum dolomite content directly above the mineralized 

interval. Values between + 14 and + 16 per mil (SMOW) 

were found in the low-dolomite, unmineralized Salt Wash 

at greater distances above the mineralized interval. 

The o180 composition of pre- and post-ore calcite 

from the Salt Wash ranges from + 15 to + 18.5 per mil 

(SMOW), consistent with a diagenetic environment dominated by meteoric water (Keith and Weber, 1964) (fig. 4). 

In contrast, the o180 composition of the bedded limestones in the underlying marine Summerville Formation 

and in the evaporite-dominated basal Tidwell Member of 

the Morrison Formation ranges from + 16.4 to + 30 per 

mil, more typical of marine and marine-evaporite conditions of formation (fig. 4). The o180 values of the 
f/) 

a: 

w 

t 

:E 

~ 

:i 

t:: 

w 

0 

132 

137 

142 

147 

152 

157 

B180dolomite (SMOW), in per mil 

dolomites in the mineralized interval of the Tony M. 

orebody are also in this range. Even though the range of 

values for each population is fairly large, calcite in both 

the Summerville Formation and Tidwell Member of the 

Morrison Formation and dolomite in the mineralized 

interval are all isotopically heavier than dolomite in the 

unmineralized rock and heavier than the pre- and postore calcites. The existence of these two populations of 

oxygen isotopes in carbonate minerals requires two waters 

of different oxygen isotopic composition. These specific 

isotope signatures imply that one water was meteoric and 

that the other was seawater. 

Figure 3. The 5180 values of dolomite versus depth in core 

3C, Tony M. orebody. Dolomite maximum superjacent to the 

mineralized interval and the mineralized interval are shown for 

reference. SMOW, Standard Mean Ocean Water. 

The oxygen and hydrogen isotopic composition of 

waters in equilibrium with various clay minerals was determined using calculated isotopic compositions of individual clay phases (Northrop and Goldhaber, in press). The 

clay-size fraction itself was always a mixture of two or 

more different clay minerals; hence, the isotopic compositions of these fractions were meaningless. End-member 

isotopic compositions, however, could be solved graphically by plotting o180 or oD versus the abundance of 

either clay phase for two-phase mixtures. End-member 

isotopic compositions for three-phase mixtures were 

calculated using the solutions of the two-phase mixtures 

together with a series of isotope mass-balance equations, 

solved simultaneously. The isotopic compositions of 

waters in equilibrium with clay minerals define two 

distinct populations (figs. 5A, B, and 6). One population, 

having o180 values greater than -4 per mil and oD values 

greater than -30 per mil, included all the samples from 

6 

4 

2 

0 

en 12 14 16 18 20 22 

1&.1 

£ 

I 

~ ~ 

:1!1 

2 

i 

4 

2 

0 

12 

818 0 (SMOW), in per mil 

D 

l§ 

mm 

24 

~ 

• 

~ 

Brushy Basin Member celcite 

S.lt W•h Member calcite 

Salt Wuh Member 

unminerelized dolomite 

28 28 

Summerville Formation 

Tidwell Member bedded celcite 

Selt.W•h Member 

minerelized dolomite 

30 

Figure 4. Histograms of calcite and dolomite 5180 values from the Jurassic Summerville Formation and Tidwell, Salt Wash, and Brushy Basin Members of the Morrison Formation in the Henry 

basin, Utah. SMOW, Standard Mean Ocean Water. 

Diagenesis and Sources of Ore-Forming Fluids, Tony M. Vanadium-Uranium Deposit, Utah 285 
6 

4 

(/J 2 w 

..J 

0.. 

~ 0'---...J---L----L.-.&..::.....1....-.I£-..&.....-'a......;;...;;..&.__--l 

~ -12 -10 -8 -6 u. 

0 

a: 

w 

~ 6 ;:::) 

z 

4 

-4 -2 0 2 

:t_Jlt)rl%r::::·r:·:::·u::,~~~~~ -1 2 -1 0 -8 -6 -4 -2 0 2 

A B 

Mineralized Salt 

Wash member 

~ Tidwell Member 

Unmineralized Salt 

Wash Member 

Figure 5. The o180 (A) and oD (8) composition of waters calculated to have been in equilibrium (at 25°C) with authigenic clay 

minerals from the Salt Wash and Tidwell Members of the Morrison Formation, Henry basin, Utah (Northrop and others, in press b) . 

the mineralized interval and the Tidwell Member below 

the mineralized interval. The other population, having 

o180 values less than -6 per mil and oD values less than 

-60 per mil, included all the unmineralized samples from 

the Salt Wash Member. The boundary between these two 

populations corresponds to the top of the mineralized 

interval in the Tony M. orebody (fig. 7). Over a vertical 

interval of about 2 meters the o180 of the calculated 

waters decreased from about -3 to -9 per mil, and the 

oD of the waters decreased from -25 to -65 per mil. 

DISCUSSION 

During Jurassic time, the Henry basin was a tectonically evolving, doubly plunging syncline. Variable rates 

of subsidence within the basin actively created intrabasinal 

folding (Peterson, 1980, 1984). Continued subsidence 

resulted in a continuous addition to the sediment pile 

within the basin, which was accompanied by compaction 

and eventually dewatering of the deeper sediments of the 

Jurassic Morrison Formation. Evolving intrabasinal structures further influenced sedimentation and basin 

hydrology. Within this sedimentological regime, relatively 

small lakes with anoxic bottom waters existed. These lakes 

appear to have been responsible for the local preservation of detrital plant debris (Peterson, 1980). This plant 

debris later served as a food source for sulfate-reducing 

bacteria, the actions of which appear to have provided 

the reductant for the vanadium and uranium (Northrop 

286 Studies in Diagenesis 

.E 

... 

Q) 

c. 

.!: 

ci 

"0 

0 

-10 

-10 -8 

• 

•••• 

• 

-4 

618 0.in per mil 

• SMOW 

• • 

• • 

~ • 

• 

EXPLANATION 

• Salt Wash Member 

• Tidwell Member 

• Ore Zone 

0 

Figure 6. o1SO versus oD composition of waters, calculated from 

authigenic clay minerals from the Salt Wash and Tidwell Members 

of the Morrison Formation, core 3C, Henry basin, Utah (Northrop 

and others, in press b). Global Meteoric Water Line (GMWL) 

(Craig, 1961) shown for reference. Standard Mean Ocean Water 

(SMOW) is standard against which values are reported. 
132 

137 

(/) 

ffi 142 

tu 

~ 

~ 

:i 147 

b: 

w 

0 

-10 -8 -6 -4 

818 0H2o. in per mil 

Figure 7. Variation of o180 of waters calculated from 

authigenic clay minerals versus depth in core 3C, Tony 

M. orebody. Mineralized interval is shown for reference. 

and others, in press a; Goldhaber and others, in 

press). 

The oxygen and hydrogen isotopic composition 

of waters calculated to be in equilibrium with authigenic clay minerals in the unmineralized Salt Wash have 

values typical of meteoric waters (fig. 6); these values are 

slightly enriched in 180 with respect to meteoric water 

when compared with the global meteoric water line 

(GMWL) defined by Craig (1961). This ''oxygen shift'' 

(see, for example, Taylor, 1976) predictably resulted from 

the migration of meteoric water down through the 

recently deposited pile of sediment containing abundant 

volcanic ash. The difference in hydrogen isotope values 

between the calculated waters and modern meteoric water 

is consistent with the well-documented tectonic extension 

and northerly migration of the western United States since 

the Jurassic (Parrish and others, 1982). The o180 values 

of the pre- and post-ore calcite and the dolomite in the 

unmineralized zone are typical of carbonate cements 

formed from meteoric ground waters (Keith and Weber, 

1964). 

The similarity of the isotopic composition of waters 

calculated from the authigenic clay minerals in the mineralized interval and Tidwell Member to that of seawater 

(SMOW; fig. 6) indicates that seawater was a likely source 

for that water; however, brines coexisting with closedbasin evaporites also have oD and o180 values in the 

same range. The basal Tidwell Member in the structurally 

closed Henry basin contains such evaporites (Peterson 

1978, 1980); moreover, the Tidwell Member has always 

been considered to be of terrestrial origin. Because the 

o34S, oBC, and o180 values of sulfate and carbonate 

minerals in the Tidwell evaporites (Northrop, 1982) are 

sufficiently close to those of seawater, distinguishing 

water source is impossible from the existing data. Regardless of whether the water was of marine or terrestrial 

evaporite origin, the isotopic composition of the authigenic clay minerals in the Tidwell Member and in the 

mineralized interval indicates that these minerals formed 

in equilibrium with the same highly saline connate waters. 

This link between the mineralized interval and the underlying Tidwell can also be seen in the o180 values of the 

dolomites, the magnesium content of authigenic clay minerals, and the sodium content of the pore-filling phases 

throughout the Tidwell and Salt Wash Members (Northrop 1982; Northrop and others, in press a, in press b). 

Thorium:uranium ratios and the vanadium content 

of the unmineralized Morrison Formation in the Henry 

basin indicate that the likely source of the vanadium and 

uranium for these deposits was the overlying Salt Wash 

and Brushy Basin Members (Northrop, 1982; Northrop 

and others, in press a). The oxygen isotopic composition 

of dolomite in the mineralized interval and of the 

vanadium-bearing chlorite indicate that the vanadium and 

uranium-bearing phases formed within the brine. This 

mode of origin requires that the vanadium and uranium 

were transported by the meteoric water and that they were 

removed from solution at the boundary between the two 

waters. The mineralizing process, as reflected by the 

vanadium-uranium content of the rock, was most intense 

immediately below the brine-meteoric water interface and 

decreased in intensity downward for 1-3 meters below 

that interface (Carpenter, 1980; Northrop, 1982). This 

vanadium-uranium pattern, as well as the Mg:Fe ratios 

in authigenic clay minerals (Northrop and others, in press 

b) are repeated at each uranium ore-horizon within the 

mineralized interval in the Tony M. ore body, as well as 

in the other orebodies in the Henry basin (Northrop and 

others, in press a; Goldhaber and others, in press). This 

systematic pattern requires that each uranium ore-horizon 

within the mineralized interval formed at different times 

near a static (or nearly static) interface that moved 

periodically to successively higher stratigraphic levels 

rather than simultaneously during a single mineralizing 

episode. 

Dolomite maxima, one of which is laterally continuous for more than 5 kilometers above each mineralized 

interval in rocks wherein calcite otherwise predominates 

requires either dedolomitization of the entire Salt Wash 

(except for these dolomite-rich peneconcordant layers) or 

dolomitization of thin tabular layers. The sabka model 

of dolomitization, whereby dolomite forms preferentially 

in the fluid-mixing zone (see, for example, Hsu and 

Siegenthaler, 1969; Atwood and Bubb, 1970; Folk and 

Land, 1975; Magaritz and others, 1980) between a 

magnesium-rich brine and meteoric water is our preferred 

Diagenesis and Sources of Ore-Forming Fluids, Tony M. Vanadium-Uranium Deposit, Utah 287 
explanation. The systematic decrease in o180 values of the 

dolomite from within to 10 meters above the mineralized 

interval supports fluid mixing in this zone. Evidence of 

such a mixing interval cannot be seen from the clay mineral 

data because ore-stage clay minerals are absent above the 

mineralized interval. 

The Tidwell and Salt Wash Members in the Henry 

basin contain four identified stratigraphic levels at which 

brine-meteoric water interfaces were stable. Three of 

these contain potentially economic vanadium-uranium 

concentrations (Northrop and others, in press a; 

Goldhaber and others, in press). Based on vanadium and 

uranium contents of natural waters (Granger and Warren, 1978; Langmuir, 1978) and the solubility limits of 

vanadium and uranium in the same solution (Hostetler and 

Garrels, 1962), the time necessary to form an average 

1.0 

0.9 

0.8 

a: 

<( 

w 

> 

a: 

w 0.7 ll.. 

(J) 

a: 

w 

t:; 

::E 0.6 

~ 

w 

~ a: 

~ 0.5 _, 

LL. 

0.4 

0.3 

0.2 

Cf 

\ 

\ 

\ 

\ 

\ 

\ 

\ 

\ 

\ 

q 

\ 

\ 

\ 

\ 

\ 

\ 

\ 

\ 

' \ 

• 

ore-horizon can be calculated. The flow rate of the 

meteoric water that carried the vanadium and uranium to 

the site of mineralization, and the efficiency of the 

mineralizing process to concentrate vanadium and uranium 

are the rate-limiting steps. The rate of change in the position of the brine-meteoric water interface can also be 

calculated, based on Peterson's (1980) suggested tectonic 

movements of 0.01 to 0.02 millimeters per year. This 

calculation obviously assumes that the interface position 

is tied to the geologic evolution of the Henry basin. The 

sample calculation shown in figure 8 illustrates that 

transport times for vanadium and uranium, in the range 

of tens-of-thousands of years, are sufficiently long to have 

formed the ore zone during the time the brine-meteoric 

water interface migrated upwards through the thickness 

of this ore zone. 

• Time versus flow rate 

of the meteoric water 

Q---{) Time versus migration 

of the interface 

~ 

' ' ' ' ' 

2.0 

1.5 

1.0 

0.8 

a: 

<( 

w 

> 

a: 

w 

ll.. 

(J) 

a: 

w 

1-

w 

::E 

ll) 

I 

0 

~ 

z 

0 

~ a: 

(.!) 

~ 

w 

u 

<( 

LL. 

a: 

w 

1-

~ 

a: 

w 

t:; 

::E 

Ln 

0 

0.1 L------'------'-----'----...L.----L------'"-----'--.........;;-....&....---...L.------' 0.6 

0 1 0 20 30 40 50 60 70 80 90 1 00 

TIME, IN 103 YEARS 

Figure 8. Plot showing time versus rate plot of formation of a typical ore-horizon, 0.5 m thick containing 9,750 parts per 

million (ppm) vanadium and 1,500 ppm uranium. Starting water for this calculation contained 85 parts per billion (ppb) uranium 

and 552 ppb vanadium. We have assumed that 75 percent of the vanadium and uranium in the starting water was removed 

by the precipitation reactions. Vanadium to uranium molar ratio shown (6.5:1) is typical of the Tony M. orebody (Carpenter, 

1980; Northrop, 1982). 

288 Studies in Diagenesis 
CONCLUSIONS 

The calculated isotopic composition of waters from 

which authigenic minerals in the Tony M. orebody formed 

indicate that vanadium-uranium mineralization took place 

at the boundary between two isotopically distinct waters. 

The oxygen and hydrogen isotope data place constraints 

on the sources of the waters involved in the mineralizing 

process. The lower fluid appears to have been a connatebasin brine, derived at least in part from the Tidwell 

Member of the Morrison Formation within the Henry 

basin. The upper fluid was meteoric water, migrating 

downward through the superjacent strata. In this structurally closed basin, the brine occupied the deepest 

permeable horizons and formed a stable interface with 

the overlying meteoric waters. A stable interface within 

the same stratigraphic interval must have existed for sufficient periods of time for vanadium-uranium mineralization to have occurred at and directly below the interface. 

Immediately above the interface was a thin (about 10 meters) zone of mixing of the two waters. Most of the authigenic, ore-stage dolomite precipitated in the mixing zone. 

REFERENCES CITED 

Atwood, D. K., and Bubb, J. N., 1970, Distribution of dolomite 

in a tidal flat environment, Sugarloaf Key, Florida: Journal of Sedimentary Petrology, v. 78, p. 499-505. 

Carpenter, D. J., 1980, Elemental, isotopic and mineralogic 

distributions within a tabular-type uranium-vanadium 

deposit, Henry Mountains mineral belt, Garfield County, 

Utah: M.S. thesis, Golden, Colorado, Colorado School of 

Mines, 156 p. 

Clayton, R.N., and Mayeda, T. K., 1963, The use of bromine 

pentafluoride in the extraction of oxygen from oxides and 

silicates for isotopic analysis: Geochimica et Cosmochimica 

Acta, v. 27, p. 43-52. 

Craig, Harmon, 1961, Isotopic variations in meteoric waters: 

Science, v. 133, p. 1702-1703. 

Drever, J. 1., 1973, The preparation of oriented clay mineral specimens for X-ray diffraction analysis by a filter-membranes 

peel technique: American Mineralogist, v. 58, p. 553-554. 

Finch, W. 1., 1967, Geology of epigenetic uranium deposits in 

sandstones in the United States: U.S. Geological Survey Professional Paper 538, 121 p. 

Fischer, R. P., 1947, Deposits of vanadium-bearing sandstone, 

in Vanderwilt, J. W., ed., Mineral resources of Colorado: 

Colorado Mineral Resources Board, p. 451-456. 

Folk, R. L., and Land, L. S., 1975, Mg/Ca ratios and salinity; 

two controls over crystallization of dolomite: American 

Association of Petroleum Geologists Bulletin, v. 59, 

p. 60-68. 

Goldhaber, M. B., Reynolds, R. L., Campbell, J. A., Wanty, R. 

D., Grauch, R. 1., and Northrop, H. R., in press, Genesis 

of the tabular-type vanadium-uranium deposits, Henry 

structural basin, Utah. Part II. Mechanism of ore and 

gangue mineral formation at the interface between brine 

and meteoric water: Economic Geology. 

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290 Studies in Diagenesis 

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basin, Utah: Clays and Clay Minerals, v. 34, p. 488-495. 
Clay Mineral Diagenesis in Lacustrine Sediments 

By BLAIR F. JONES, U.S. Geological Survey, National Center, MS 432, 12201 

Sunrise Valley Drive, Reston, Virginia 22092 

CONTENTS 

Abstract 292 

Introduction 292 

Acknowledgments 292 

Observation and analytical evidence 292 

Mechanisms of clay structural transformation 293 

Chemical and mineralogical changes 293 

Phase relations 295 

Conclusions 298 

References cited 298 

FIGURES 

1. Diagram of stability relations of hydrous magnesium silicates with respect to solution activities as calculated from thermodynamic data 297 

2. Diagram of stability relations of hydrous magnesium and sodium-magnesium 

silicates as estimated from thermodynamic data and natural occurrences 298, 299 

TABLES 

1. Calculated structural formulas of lacustrine mixed-layer magnesium 

smectities 294 

2. Stability reactions for sesquioxide-free hydrous magnesium silicates 296 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 291 
Abstract 

Recent investigations of the geochemical evolution of 

waters and associated sediments in saline lake systems indicate 

that reactions involving silica or silicates remove major cations 

from solution. These studies have confirmed earlier observations by George Mil lot on ancient lacustrine deposits; he found 

that silica and cations in smectite clays. increase relative to 

alumina from basin margin to center. Mineralogical and 

chemical data from Lake Abert in south-central Oregon, the 

Amargosa Desert in southern Nevada, and Great Salt Lake, 

Utah, indicate that solute loss of K, Mg, and Si is commonly 

associated with the formation of a clay mineral component that 

increases with the proportion of solution to detrital solid, that 

is, the proportion of water to sediment inflow. Available evidence suggests that diagenesis in saline and perhaps freshwater 

sediments includes any or all of the following mechanisms in 

general sequence: hydrolysis of primary silicates (especially 

pyroclastic material), exchange and fixation of cations in interlayer positions in preexisting 2:1 silicate layers, and subsequent 

reaction with silica to form wholly authigenic interstratifications. 

The detailed structure of the final product depends on the 

hydrologic environment and its influence on pH, solute silica, 

and alkali/ferromagnesium ratios. 

INTRODUCTION 

As with the marine environment, there has long 

been speculation as to what extent depositional conditions 

in lakes are reflected in the alteration or diagenesis of 

clays in lacustrine sediments. Garrels and Mackenzie 

(1967) presented the classic model of the development of 

alkaline, saline (so-called "soda") lakes by evaporative 

concentration of drainage from crystalline rock terranes. 

They proposed the direct precipitation of sepiolite from 

lake waters as the principal control on the concentration 

of silica (and magnesium) in solution. Similarly, Gac and 

others (1977), following the indications of Carmouze 

(1976), used an evaporative concentration model to 

postulate the authigenic formation of smectite in Lake 

Chad in central Africa. From their examination of the 

<2-~-tm-size materials remaining after the experimental 

evaporation of Chari River inflow water, Gac and others 

(1977) suggested that the amount of smectite had increased over what was in the original suspended sediment. 

Jones and Bowser (1978), however, concluded that definitive mineralogical evidence was lacking for authigenesis 

of clays in freshwater lake sediments. The present paper 

examines the evidence for and the mechanisms of clay 

mineral diagenesis in saline, alkaline, and perhaps also 

freshwater lacustrine environments. 

Acknowledgments 

I thank the manuscript reviewers, Gene Whitney 

and L. V. Benson, and those who offered help and 

292 Studies in Diagenesis 

encouragement to the work: M. W. Bodine, R. G. Deike, 

D. D. Eberl, H. P. Eugster, J. Hower, B. Hudner, 

C. Moss, S. L. Rettig, R. J. Spencer, D. C. Thorstenson, A. H. Weir, and especially Betty Jones. 

OBSERVATION AND ANALYTICAL EVIDENCE 

In his 1964 treatise on the geology of clays, Georges 

Millot reviewed the clay mineral parageneses of alkaline 

lakes and marine basins and concluded that the ''montmorillinoid'' clays found in them were largely diagenetically transformed or authigenic in origin, processes 

that he termed "neoformation by addition." He noted 

a geochemical sequence in deposits of closed basins (or 

"confined environment deposits"): the most aluminous 

clays were at the periphery; they were succeeded by more 

magnesian smectites basinward; and sepiolite was found 

at the center. More recent studies of the geochemical 

evolution of waters in closed basins have documented the 

reaction of solutes with silica or silicates, processes that 

exert a significant influence on the solute budget of lake 

systems (Jones and VanDenburgh, 1966; VanDenburgh, 

1975; AI Droubi and others, 1980; Eugster and Jones, 

1979). Subsequent studies of recent lacustrine sediments 

at Lake Abert, Ore. (Jones and Weir, 1983), and Great 

Salt Lake, Utah (Spencer, 1982), have confirmed the 

sequence of more magnesian clay compositions basinward, and have associated it with increasingly saline 

conditions. 

Although Millot's sequence can be simply related 

to silicate mineral solubility, evidence from modern and 

ancient lacustrine deposits indicates that the authigenic 

formation of clay minerals in a lake environment takes 

place far more commonly by diagenetic alteration of 

preexisting material than by direct precipitation of a new 

phase from solution. This phenomenon seems to be a 

function of the relative insolubilities of the hydroxides 

of major elements that compose the common clay mineral 

assemblages. Thus, colloidal aluminous or iron-rich particles may constitute the fundamental reactive base for 

the development of authigenic silicates. At the same time, 

silica can be accumulated in sediments by organisms or 

by inorganic precipitation as a result of evaporative concentration of a dilute water. Although uncommon, 

examples of direct reaction of sesquioxide with silica in 

relatively dilute water do exist. For example, at Lake 

Chad in central Africa, Pedro and others (1978) suggested 

that smectite formed by the reaction of silica with ironrich detrital particles contributed by the principal stream 

inflow. At Lake Malawi (Lake Nyasa), Africa, Muller 

and Forstner (1973) identified nontronite, which they 

interpreted as a product of the reaction of iron-bearing 

pore fluids with diatomite silica. 

Based on the solubility of pure compounds, the 

next most insoluble precipitate-after alkaline earth 
carbonate-expected from the concentration of most 

natural waters would be some form of magnesium silicate 

(Jones and VanDenburgh, 1966; AI Droubi and others, 

1980; Eugster and Jones, 1979). Detailed mineralogic 

examinations of sediments from modern saline lakes or 

playas in the United States (Droste, 1961; Guven and Kerr, 

1966; Jones and others, 1972) have not identified sepiolite. 

Wollast and others (1968), in their study of sepiolite 

precipitation, noted that if "reactive aluminous phases" 

are present in sediment, a magnesium-aluminosilicate will 

form instead of sepiolite. Transformation of preexisting 

clays seems especially favored where the proportion of 

sediment to water is high and where the solute content 

is significant (that is, in typical saline lakes or playas). 

This type of diagenesis yields the same basinward compositional trend of increasing silica and magnesium that 

is expected from direct precipitation, as shown by the 

sequence of smectite compositions from the Lake Chad 

sediments reported by Carmouze and others (1977). Such 

transformations appear to be facilitated by the high reactivity of cryptocrystalline pyroclastics (Deike and Jones, 

1980). 

Mechanisms of clay structural transformation 

Devitrification of volcanic glass typically results in 

the formation of dioctahedral smectite by complete hydration of the glass and subsequent leaching of variable 

amounts of silica and cations. Depending somewhat on 

the chemical heterogeneity of the glassy pyroclastic 

materials, some interstratification of the smectite may 

result from subaerial weathering and degradation during 

transport to the lacustrine depositional basin (Jones and 

Weir, 1983). Poor crystallinity or structural disorder in the 

cryptocrystalline smectite is probably the result of (1) the 

short-range order of the original volcanic glass, or (2) the 

extensive protonation of fundamental structural bonds in 

2:1 silicate layers (Hem and Lind, 1974). 

In addition to changes related to hydrolysis, the conditions resulting from organic degradation in lake-bottom 

sediments can cause the reduction of smectite iron, thereby 

accounting for the uptake of potassium or other ions in 

interlayer sites. Such a process was proposed by Singer 

and Stoffers (1980) for the formation of illite from smectite in sediment cores from Lake Albert (Zaire-Uganda) 

and Lake Manyara (Tanzania), east Africa; this same process can account for potassium enrichment and incipient 

illitization of smectite at Lake Abert (Jones and Weir, 

1983). Singer and Stoffers (1980) concluded that the 

uptake of potassium into phyllosilicates is accompanied 

by the replacement of Mg for octahedral FeH, similar to 

the mechanism proposed by Drever (1971) for clay 

sediments in marine reducing environments. The 

experiments of Heller-Kallai and Rozenson (1978) with 

magnesium sulfide solutions, however, show that the 

reduction of iron in iron-bearing smectite leads initially 

to simple (and reversible) layer-charge increase (or to the 

formation of a degraded clay phase with increased ionexchange capacity), rather than directly to octahedral substitution of magnesium for iron, as suggested by Drever 

(1971). The lack of such direct octahedral substitution is 

more consistent with evidence for the uptake of potassium earlier than magnesium at Lake Abert (Jones and 

Weir, 1983) and Great Salt Lake (Spencer, 1982). 

Chemical and mineralogical changes 

Along with the ubiquitous carbonate deposition 

that accompanies increased salinity in lakes, the depletion of magnesium relative to solutes that are more conservative in solution has been documented in the geochemical evolution of most closed basin waters (Jones and 

VanDenburgh, 1966; Eugster and Jones, 1979). In many 

basins, magnesium depletion can be attributed to its incorporation in carbonate phases, but, as noted at Lake Abert 

(Jones and Weir, 1983) and Great Salt Lake (Spencer, 

1982), magnesium is generally enriched in lake-sediment 

clay fractions as well. Magnesium solute depletion begins 

by the exchange of magnesium onto clays for calcium 

(which precipitates as CaC03

), and it apparently continues by the interlayer uptake of Mg(OH)+ or by the 

intergrading of chloritic layers in smectite (probably 

facilitated by the reduction of octahedral iron as noted 

previously). Magnesium has been shown to be the predominant exchangeable cation in clays from Lake Abert 

(Jones and Weir, 1983) and the Amargosa Desert (Eberl 

and others, 1982), as well as from Lake Manyara (Singer 

and Stoffers, 1980). After magnesium exchange, these 

magnesium interlayers then react with noncrystalline silica 

to form interstratified kerolite, as has been suggested in 

the Lake Abert and Amargosa Desert deposits. 

This reaction is similar to that of clay mineral 

authigenesis by fixation of silica on a metal hydroxide 

layer, as shown experimentally by Caillere and Renin 

(1961). Indeed, much of the authigenic magnesium 

silicates found in lake deposits probably formed by the 

interstratification of (1) 9.6-A, nonexpanding trioctahedral 

kerolite; or (2) very low charge, nearly trioctahedral, 

stevensite domains and preexisting, perhaps largely dioctahedral smectite. The average structural formulas 

calculated for these interstratified minerals show that most 

of the layer charge arises from octahedral vacancies rather 

than tetrahedral substitution. Hence, the expanding component is termed stevensite rather than saponite, according to the criteria of Faust and others (1959). Brindley 

(1955) first recognized that stevensites may contain talclike layers, and Brindley and Kao (1980) noted the development of expanded basal spacings and (or) mixed layering 

with the experimental addition of equivalent Mg and OH 

to beidellite. 

Clay Mineral Diagenesis in Lacustrine Sediments 293 
Average structural formulas based on the 22 negative charges of a normal smectite formula for suspected 

trioctahedral-dioctahedral mixed-layer materials in both 

modern and ancient lacustrine sediments are listed in table 

1. The hydrochemical environment of these clays ranged 

from highly saline (Great Salt Lake) to somewhat 

brackish (Lake Chad), but presumably all associated 

waters were neutral to alkaline (for example, the pH of 

Lake Abert commonly reaches 10). The most magnesiumrich examples from each location are given for the 

modern lake sediments, whereas a range is presented in 

the formulas for the early Tertiary Green River and the 

Pliocene-Pleistocene Amargosa Desert materials. Larger 

proportions of sesquioxides in smectites from predominantly nonvolcanic terrane suggest a saponite-type 

(tetrahedral) layer charge similar to the Great Salt Lake 

material. 

The modern lake sediments that provided the clay 

compositions given in table 1 are all from closed basins 

maintained by surface water inflow that provides a source 

of colloidal detritus. Spencer's (1982) and the author's 

findings suggest that the Great Salt Lake clay chemistry 

can be obtained simply by the addition of magnesium 

silicate to inflow river clay. Likewise, the formula given 

by Carmouze (1976) was obtained by subtracting the 

Chari River smectite composition from the analysis of 

the Lake Chad mineral. The Green River material of Dyni 

(1976) was apparently also formed in a similar physical 

setting, whereas the stevensite of Tettenhorst and Moore 

(1978) probably resulted from a saline lake environment 

Table 1. Average structural formulas1 for lacustrine mixed-layer magnesium smectite, in atoms 

per 22 negative charge 

[n.r., Not reported] 

Tetahedral 

Si Al Mg 

Ancient sediments 

Green River Formation: 

Dyni (1976) 

Octahedral 

Al Li Sum 

Interlayer 

X+ 

Unadjusted ••••••••••••••• 3.85 0.15 1.67 0.49 0.27 0.04 2.47 0.50 

Adjusted for 

celadonite ••••••••••••• 3.82 .18 1.77 .45 .29 .04 2.55 .39 

adjusted for 

illite ••••••••••••••••• 3.96 .04 2.03 .21 .27 .05 2.56 .47 

Tettenhorst and 

Moore (1978) •••••••••••••• 4.00 .0 2.67 .07 .03 n.r. 2.77 .32 

Bradley and Fahey 

(1962) •••••••••••••••••••• 3.98 .02 2.81 .06 .04 .04 2.95 .04 

Amargosa Desert: 

Fairbanks Butte2 •••••••••••• 3.95 .05 1.77 .41 .27 n.r. 2.45 .46 

Eberl and others (1982) 

Kerolite/stevensite •••••• 3.96 .04 2.72 .07 .03 .09 2.91 .21 

Kerolite................. 3.94 .03 3.01 :o .01 n.r. 3.02 .15 

Morocco (ghassoulite): 

Millot (1954) ••••••••••••••• 3.98 .02 2.69 .08 .05 .10 2.92 .21 

Amboseli basin (kerolite): 

Stoessell and Hay 

(1978).................... 3.91 .04 2.99 .04 .0 n.r. 2.99 .22 

Lake Tecopa: 

Starkey and Blackmon 

(1979) •••••••••••••••••••• 3.90 .10 2.20 .28 .12 .10 2.70 .41 

Modern sediments 

Lake Abert: 

Jones and Weir (1983) ••••••• 3.90 0.10 1.85 0.39 0.30 n.r. 2.54 0.38 

Lake Chad: 

Tardy and others 

(1974) •••••••••••••••••••• 4.00 .0 1.90 .51 .16 n.r. 2.55 .19 

Carmouze (1976) ••••••••••••• 4.00 .0 1.99 .24 .12 3.11 2.46 .45 

Walker Lake: 

Spencer (1977) •••••••••••••• 3.43 .57 1.27 .63 .51 n.r. 2.41 .58 

Great Salt Lake: 

Spencer (1982) •••••••••••••• 3.63 .37 1.67 .51 .36 n.r. 2.54 .41 

Icalculated assuming an anion content o10 (0H) 2• 

2New analysis by S. L. Rettig, U.S. Geological Survey. Mineralogy described 

by Teague ( 1981). 

3Titanium. 

294 Studies in Diagenesis 
that was significantly isolated from detrital influx. On 

the other hand, the more magnesium-rich clays of the 

Amargosa Desert, Nev. (Eberl and others, 1982), and the 

Amboseli basin, Kenya (Stoessel and Hay, 1978), were 

formed under conditions of sustained spring inflow, 

which also provided a minimum of transported detritus. 

The final product was an almost pure trioctahedral 2: 1 

layer phase (kerolite), in addition to sepiolite. Although 

good textural evidence exists for the coalescent nucleation of sepiolite (Khoury, 1979), trioctahedral 2: 1 

phyllosilicates appear to require preexisting sheetstructure substrata, presumably because the true phyllosilicate structure does not as readily accommodate water 

of hydration (see Guven and Carney, 1979). 

Thus, the only silicate phases that nucleate directly 

from natural silica-rich surface waters appear to be 

sepiolite and magadiite, the latter at considerably higher 

salinity than the former. A chemical "combination" of 

these two phases, loughlinite (or "soda-sepiolite"), has 

been described from the saline facies of the Wyoming 

Green River Formation (Fahey and others, 1960), but was 

interpreted as a product of postdepositional diagenesis. 

A facility for direct nucleation of chain-structure clay 

minerals from aqueous fluids has been demonstrated by 

the ready precipitation of sepiolite from laboratory solutions (Siffert and Wey, 1962; Wollast and others, 1968), 

and the observation of magadiite precipitating in playa 

ponds (Rooney and others, 1969), as well as in laboratory 

vessels (Bricker, 1969). 

The only areas where both phases have been reported in the same stratigraphic interval are in the 

Pliocene-Pleistocene beds of ancient Lake Tecopa, Calif., 

downstream from the Amargosa Desert in the Amargosa 

River drainage (Starkey and Blackmon, 1979). By studying these occurrences and those of sepiolite reported in 

the literature, Starkey and Blackmon (1979) concluded 

that precipitation took place only where aluminum (and 

presumably iron also) was low or absent. They referred 

to the observation of Wollast and others (1968) that 

sepiolite was usually associated with carbonate sediments 

where aluminum content was very low. They also noted 

that colloidal silica has great affinity for solute magnesium and suggested that diatom debris may be a substrate for the formation of magnesium silicate. This was 

confirmed in the study of authigenic stevensite formation at Pastos Grandes playa, Bolivia, by Badaut and 

Risacher (1983). Although Starkey and Blackmon (1979) 

related Lake Tecopa magadiite to local soda-rich springs, 

the paucity of detrital or solute sesquioxide and the reaction of Na + with colloidal silica were also likely factors 

in its formation. The bulk of the clay in the Lake Tecopa 

sediments was termed lithium saponite by Starkey and 

Blackmon; however, according to the formula calculated 

from their analysis (table 1), this smectite is similar to 

the stevensite-rich interstratified clay from the Amargosa 

Desert, where it probably originated. 

PHASE RE LA liONS 

Pertinent hydrologic conditions for silicate authigenesis in the lacustrine environment are reflected in 

phase relations for low-temperature magnesium silicates, 

as shown on activity-activity diagrams. The controlling 

equations for the principal phases in the systems Mg0-

Si02-H20 and Mg0-Na20-Si02-H20 are given in table 

2. The first two equations describe the simple precipitation or dissolution of sepiolite and talc (or kerolite), 

respectively. The second pair of equations represents the 

conversion of sepiolite to talc (or kerolite) at constant 

activity of Si02(aq) or H2

0 (for a plot of 2+/a~+ versus 

Si02<aq>' as in fig. 1). The third pair of equations shows 

the sodium addition necessary to form aluminum-free 

smectite (stevensite) from either sepiolite or talc (kerolite) 

at constant activity of Si02(aq) (for the plot of aM 2+/a~+ 

versus a +/a +' as in fig. 2). The final equatio~ illus- Na H 

trates the precipitation of serpentine (chrysotile). For the 

construction of the activity-activity diagrams, available 

thermodynamic data were taken from the sources listed 

for the computer program WATEQ (Truesdell and Jones, 

1974). Although the emphasis here is on silicate solubility 

in the sesquioxide (AI, Fe)-free system, dashed lines and 

parenthetical mineral names have been added to figure 

1 to suggest stability fields for appropriate aluminous 

phases at representative sodium concentrations and at 

saturation with noncrystalline Fe(OH)3• 

The solubility boundaries of figure 1 clearly illustrate the closed-system metastability of sepiolite relative 

to talc even if allowance is made for variable crystallinity (such as the disorder in the hydrated kerolite analog 

of talc) by calculation from the range of free energy of 

formation data determined for talc by Bricker and others 

(1973) and Hemley and others (1977) and for sepiolite 

by Wollast and others (1968). If the precipitation 

mechanism in nature, however, is primarily through addition of magnesium to preexisting colloidal silica, the low 

ratio of magnesium to silicon can be expected to favor 

the initial formation of sepiolite over talc (or kerolite). 

A rise in solution pH accompanying the degassing of 

C02 from spring discharge, organic respiration, or 

evaporative concentration of sediment-free surface waters 

should promote the direct precipitation of either phase. 

Stable ground-water inflow, buffering of an alkaline pH 

by carbonate precipitation, and a preexisting mineral 

substrate, however, apparently aid in overcoming kinetic 

dehydration or nucleation barriers to kerolite formation. 

Long-standing ground-water discharge of appropriate 

composition into an evaporative shallow lake or marsh 

environment has led to substantive formation of both 

sepiolite and kerolite/ stevensite in the Amargosa Desert 

(Khoury, 1979; Khoury and others, 1982; Eberl and 

others, 1982) and the Amboseli basin (Stoessell and Hay, 

1978). Traces of a 7-A reflection on X-ray diffractograms 

from both locales suggest that even chrysotile may form 

Clay Mineral Diagenesis in Lacustrine Sediments 295 
Table 2. Stability reactions for sesquioxide-free hydrous magnesium silicates1 

Equation 

No. Reaction 

2Mg2+ + 3H4Si04 + 4(0H)- = Mg2Si30 7.

5

(0H)·3H20. 

solution sepiolite 

2 3Mg2+ + 4H4Si04 + 6(0H)- = Mg3Si40 10(0H)2 • 3H20 + 7H20. 

solution kerolite (talc) 

3 4Mg2Si30 7.

5(0H) · 3H2

0 + Mg2+ + 20H- = 3Mg3Si40 10(0H)2 • 3H20 + H20. 

sepiolite kerolite (talc) 

4 2Mg2Si30 7.

5(0H) · 3H2

0 

sepiolite 

= Mg3Si40 10(0H)2 ·3H20 + Mg2+ + 2H4Si04 + 20H- + H20. 

kerolite (talc) 

15 1.33Mg2Si30 7.

5

(0H)· 3H20 + 0.33Na+ + 0.330Hsepiolite 

= Na0.

33Mg2.87Si40 10(0H)2 + 3H20. 

stevensite 

16 Mg3Si40 10(0H)2 + 0.33Na+ = Na0.

33Mg2.

87Si40 10(0H)2 + 0.17Mg2+. 

talc (kerolite) stevensite 

7 3Mg2+ + 2H4Si04 + 6(0H)- = Mg3Si20 5(0H)4 + 5H20. 

chrysotile 

1

These equations have been written for end-member octahedral deficiency in sesquioxide-free smectite structures at constant activity of Si02

(a ) to illustrate the stevensite definition according to seat of 

layer charge as discussed in the text. As descr~ed, natural stevensites typically have low layer charge (Faust 

and others, 1959), which is consistent with the apparent interstratification of smectite and talc or kerolite 

layers, and can range through kerolite/stevensite interstratifications to nearly pure kerolite. 

if evaporatively induced pH and solute increase are more 

rapid than mineral precipitation. Despite initial 

metastability in the formation of sepiolite, the boundary 

derived from equation 4 shown in the upper right part 

of figure 1 suggests that the sepiolite may actually be 

stablized at sufficiently high levels of all three factors, 

that is, pH, magnesium concentrations, and silica concentrations at constant aH20 • 

The effects of sodium concentration at silica contents appropriate to glass saturation are considered in 

figure 2. The metastability of sepiolite and lack of 

stability data for pure stevensite makes the precise location of these field boundaries conjectural; however, the 

lower limit of the sepiolite field (and that for noncrystalline talc/kerolite, as well) at -log Si02

(aq) = 3 

was derived from the appropriate curve of figure 1, and 

the upper limit was estimated. The lower concentration 

limit of the stevensite field was drawn to correspond with 

the aNa+/aH+ = 1 condition of equation 5, but the 

higher aNa+/aH+ estimate is based on analyses of 

Amargosa Desert waters. The slopes of the other stevensite field boundaries were derived from the stoichiometry 

of equation 6 and the dissolution of the phase. With the 

addition of sesquioxide, a smectite-chlorite boundary can 

296 Studies in Diagenesis 

be estimated that roughly bisects the stevensite field. 

The phase relations at constant silica depicted 

on figure 2 correspond well with the hydrochemical 

environments observed for magnesium silicate precipitation and diagenesis. Thus, sepiolite precipitation can be 

expected primarily from relatively dilute waters-as in 

the Amboseli basin (Stoessell and Hay, 1978) or the 

Amargosa Desert (Khoury, 1979; Khoury and others, 

1982)-whereas at the same locale kerolite may have 

formed from evaporatively induced, somewhat higher pH 

and Mg-content solutions with decreasing Si02(aq)· 

Kerolite has been identified as a precipitate from Stewart 

Springs, Calif., in waters with aM 2+/a~+ = 18.77 

(pH= 12.2) and log total aqueous sili~a equal to - 4.23 

(1. Barnes, U.S. Geological Survey, Menlo Park, Calif., 

written commun., 1981). 

The normal dominance of sodium in continental 

brackish or saline waters (Jones, 1966) should indicate 

the preference for the expandable smectite (stevensite) 

according to figure 2. Indeed, this result corresponds with 

the paucity of pure magnesium silicates in closed-basin 

lake or playa sediments (Droste, 1961). The suggestion 

of kerolite/ smectite interstratification in Lake Abert 

sediments (Weir and Jones, 1978; Jones and Weir, 1983) 
18 

16 

14 \ \ \ \ \ \ 

+ \ \ 

C\I::I: 12 \ \ 

~ \ \ \ 

N \ " (~ C) \ '-.....~ecb'te ~ \ 

\::$ \ \ ' ........ ~6 10 \ \ ........ ......._?) 

CJ) \ ' 

0 ''\.... ....................... 

\ \ rs.- ......... \ \ 'nect;, 

8 \ tte ........ \ 'SJ:::-' \ \ / ........ 

\ ' 

(Kaolinite) \ \ 

\ \ 

6 \ \ 

\ \ 

\ \ 

\ \ 

\ \ 

4 

-6 -5 -4 -3 -2 -1 

log GH4Si04 

Figure 1. Stability relations for hydrous magnesium silicates with respect to solute activities, as calculated from thermodynamic 

data cited in text. Shaded regions correspond in sesquioxide-free system to supersaturation with solid phases, either Mg-silicate 

(light shade) or noncrystalline (amorphous) silica (darker shade). Field boundaries for sepiolite and talc correspond to equations 

1 and 2, respectively, in table 2 and are drawn for stability of both noncrystalline (amorphous) and crystalline forms with appropriate 

free energy of formation values given in kilocalories per mole. Similar estimate is given in shading of chrysotile boundary, which 

corresponds to equation 7 of table 2. Long-dashed line resulting from calculations based on equation 4 of table 2 is shown in upper 

right of diagram. Double-dashed boundaries, one each for a )a + values of 6.0 and 8.0, are given for sesquioxide-bearing phases 

at saturation with noncrystalline Fe(OH)y assuming AI is c'i)nse1ved. Stability regions for these phases (chlorite, smectites, and 

kaolinite) are labeled on the appropriate side of the two short-dashed boundaries. Smectite-Ab and smectite-BF refer to Kittrick's 

Aberdeen and Belle Fourche, S. Oak., montmorillonites, respectively. Stability constants are modified from Truesdell and Jones (1974). 

Clay Mineral Diagenesis in Lacustrine Sediments 297 
and the recognition of kerolite/stevensite interstratification in the Amargosa Desert deposits (Eberl and others, 

1982) indicate that the upper boundary of the stevensite 

field on figure 2 is diffuse. The slope of this boundary 

is consistent with a trend toward increased interstratification of kerolite with smectite as the salinity and pH rise 

in alkaline saline lakes such as Lake Abert. A corresponding 1: 1 increase in solution magnesium and sodium content on evaporative concentration would mark an oblique path across the talc-smectite boundary. Similarly, the 

evidence for formation of chloritic intergrades prior to 

kerolite interstratification (Jones and Weir, 1983) is consistent with the lower position of a smectite-chlorite boundary depicted for a sesquioxide-bearing system. 

CONCLUSIONS 

In summary, the diagenesis of clay minerals in · 

lacustrine sediments involves primarily the uptake of 

potassium, magnesium, and silica from lake waters 

through the reaction of the dissolved elements with preexisting particulate material. The reactions appear to proceed in sequential steps, as follows: 

1. Hydrolysis of pyroclastics or detrital sesquioxide to 

form hydroxide colloids, intergrades, or incipient, 

largely dioctahedral 2:1 layers; 

2. Exchange of cations for the calcium lost by carbonate 

precipitation, and organically induced reduction of 

iron to increase phyllosilicate layer charge and to fix 

interlayer cations, initially potassium, but also 

magnesium, perhaps as hydroxide polymers; and 

3. Reaction of magnesium (hydroxide?) with silica from 

pore fluids, diatomaceous debris, or glassy 

pyroclastics to form trioctahedral kerolite and (or) 

stevensite interstratification. 

These reactions seem to operate in many hydrologic 

systems over a wide range of salinity. Pure kerolite is 

associated with precipitated sepiolite and carbonate, and 

is apparently restricted to areas of relatively dilute 

ground-water discharge with an absence of reactive 

detritus. The higher solute sodium content of brackish 

or saline lakes favors the development of stevensite, 

especially where detrital input is restricted. 

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1334-1336. 

Droste, J. B., 1961, Clay mineral composition of sediments in 

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Science, v. 113, p. 1928. 

Figure 2. (facing page) Stability relations of hydrous magnesium and sodium-magnesium silicates as estimated from thermodynamic 

data cited in text, as well as from natural occurrences. Fields of supersaturation are labeled according to appropriate mineral phase. 

Solid boundaries outline most probable stability fields, whereas long-dashed lines suggest possible limits for occurrence of low-soda 

minerals. Short-dashed boundaries indicate actual stable extension of talc (kerolite) field, which would preclude formation of sepiolite. 

Region of kerolite/stevensite mixed-layer development is suggested for upper stability boundary of stevensite. For sesquioxide-bearing 

system, chlorite-smectite boundary is shown by short dashes; it was calculated from same data as for figure 1 with log Si02

(aq) equal 

to -3.02. Smectite-Ab refers to Kittrick's Aberdeen montmorillonite in Truesdell and jones (1974). 

298 Studies in Diagenesis 
+ 

C\I:I: 

~ 

20 

18 

16 

14 

NO) 12 ~ 

\:$ 

C) 

0 

10 

8 

6 

~co // Talc (Kerolite) 0 

\'3-'40 / 

~~0// / 

/ / 

----:r // 

I tCh\or\te) ---/---

' ---b) / 

----- -----fSnlecti~~_,-" 

Sepiolite """"7""/ /f.:.:/ .. -Stevensite 

/ ...........____...........____ 

------< ............... ---- ............... ---- ............... -- -- -- ---- -- ----------

Solution 

4L-----~----~----~----~----~----~ 

-2 0 2 4 6 8 10 

Clay Mineral Diagenesis in Lacustrine Sediments 299 
Dyni, J. R., 1976, Trioctahedral smectite in the Green River Formation, Duchesne County, Utah: U.S. Geological Survey 

Professional Paper 967, 14 p. 

Eberl,_ D. E., Jones, B. F., and Khoury, H. N., 1982, Mixed-layer 

kerolite/stevensite from the Amargosa Desert, Nevada: 

Clays and Clay Minerals, v. 30, p. 321-326. 

Eugster, H. P., and Jones, B. F., 1979, Behavior of major ~olutes 

during closed-basin brine evolution: America! Journal of 

Science, v. 279, p. 609-631. 

Fahey, J. J., Ross, M., and Axelrod, J. M. 1960, Loughlinite, a 

new hydrous sodium magnesium silicate: American 

Mineralogist, v. 45, p. 270-281. 

Faust, G. T., Hathaway, J. C., and Millot, G., 1959, A restudy 

of stevensite and allied minerals: American Mineralogist 

v. 44, p. 342-370. 

Gac, J. Y., AI Droubi, A., Fritz, B., and Tardy, Y., 1977, 

Geochemical behavior of silica and magnesium during the 

evaporation of waters in Chad: Chemical Geology, v. 19, 

p. 215-228. 

Garrels, R. M., and Mackenzie, F. T., 1967, Origin of the 

chemical composition of some springs and lakes, in W. 

Stumm, ed., Equilibrium concepts in natural water systems: 

American Chemical Society Advances in Chemistry, v. 67, 

p. 222-242. 

Giiven, N., and Carney, L. L., 1979, The transformation of 

sepiolite to stevensite and the effect of added chlorides and 

hydroxides: Clays and Clay Minerals, v. 27, p. 253-260. 

Giiven, N., and Kerr, P. F., 1966, Selected Great Basin playa clays: 

American Mineralogist, v. 51, p. 1056-1067. 

Heller-Kallai, L., and Rozenson, 1., 1978, Removal of magnesium from interstitial waters in reducing environmentsThe problem reconsidered: Geochimica et Cosmochimica 

Acta, v. 42, p. 1907-1909. 

Hem, J., and Lind, C., 1974, Kaolinite synthesis at 25 OC: Science, 

v. 184, p. 1171-1173. 

Hemley, J. J., Montoya, J. W., Christ, C. L., and Hostettler, P. 

B., 1977, Mineral equilibria in the Mg0-Si02-H20 system; 

I. Talc-chrysotile-forsterite-brucite stability relations: 

American Journal of Science, v. 277, p. 322-351. 

Jones, B. F., 1966, Geochemical evolution of closed basin waters 

in the western Great Basin, in Rau, J. 1., ed., Northern Ohio 

Geological Society, 2nd symposium on salt: v. 2, p. 181-200. 

Jones, B. F., and Bowser, C. J., 1978, The mineralogy and related 

chemistry of lake sediments, ch. 7, in Lerman, A., ed., 

Lakes; chemistry, geology, physics: New York, SpringerVerlag, p. 179-235. 

Jones, B. F., and VenDenburgh, A. S., 1966, Geochemical 

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70, p. 435-446. 

Jones B. F., VanDenburgh, A. S., Deike, R. G., Rettig, S. L., 

and Truesdell, A. H., 1972, Mineralogy of some sediments 

from alkaline saline lakes: International Clay Conference 

Program and Abstracts, Madrid, p. 80-81. 

Jones, B. F., and Weir, A. H., 1983, Clay minerals of Lake Abert, 

an alkaline, saline lake: Clays and Clay Minerals, v. 31, p. 

161-172. 

Khoury, H. N., 1979, Mineralogy and chemistry of some unusual 

clay deposits in the Amargosa Desert, southern Nevada: 

Urbana, Ill., University of Illinois, Ph. D. thesis, 171 p. 

300 Studies in Diagenesis 

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___ Geologie des Argiles: Masson et Cie., Paris ("Geology 

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1034-1043. 

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East African lake sediments: Clay Minerals, v. 15, 

p. 291-307. 

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___ 1982, The geochemical evolution of Great Salt Lake: 

Baltimore, Maryland, Johns Hopkins University, Ph. D. 

thesis, 308 p. 

Starkey, H. C., and Blackmon, P. D., 1979, Clay mineralogy of 

Pleistocene Lake Tecopa, Inyo County, California: U.S. 

Geological Survey Professional Paper 1061, 34 p. 

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of sepiolite and kerolite at Amboseli, Kenya: Contributions 

to Mineralogy and Petrology, v. 65, p. 255-267. 

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lac Tchad; Application aux domains de stabilite des 

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and related deposits from the Amargosa Desert, Nevada: 

Berkeley, Calif., University of California, M.S. thesis, 78 p. 

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program for calculating chemical equilibria of natural 

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Experimental precipitation and genesis of sepiolite at earthsurface conditions: American Mineralogist, v. 53, p. 

1645-1662. 
Zeolitic Diagenesis of Tuffs in an Upper Miocene Lacustrine 

Deposit Near Durkee, Baker County, Oregon 

By ARTHUR j. GUDE, 3rd, and RICHARD A. SHEPPARD, 

U.S. Geological Survey, Denver Federal Center, MS 917, Denver, Colorado 80225 

CONTENTS 

Abstract 303 

Introduction 303 

Geography of the Durkee area 303 

Geologic setting 303 

Stratigraphy 306 

Sedimentation 307 

Materials and methods 308 

Mineralogical results 308 

Zeolites and authigenic feldspar 308 

Analcime and potassium feldspar 308 

Silica minerals 309 

Other minerals 309 

Diagenesis at Durkee 309 

Water 309 

Glass 310 

Clay minerals 311 

Diatoms and diatomite 311 

Phillipsite 311 

Chabazite 314 

Erionite 314 

Clinoptilolite 314 

Mordenite 315 

Analcime and potassium feldspar 315 

Silica minerals 316 

Conclusions 316 

References cited 316 

Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, 

Oregon 318 

FIGURES 

1. Map of late Miocene lacustrine basin showing sampled localities and 

geology at Durkee, Oregon 304, 305 

2. Schematic diagram of a stratigraphic cross section of the Durkee 

basin 306 

3. 

4-6. 

Composite columnar section of Durkee lacustrine and fluvial rocks 307 

Photographs: 

4. View looking northwest from locality 30 across an eroded 

amphitheatre 307 

5. Salt casts and raindrop imprints on underside of an altered tuff at 

locality 7, Durkee, Oregon 308 

6. Ripple-marked bed on surface of an altered tuff at locality 65, Durkee, 

Oregon 310 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 301 
7-16. Scanning electron micrographs: 

7. Partly eroded glass fragment 310 

8. Smectite coating surface of glass fragment 311 

9. Spherical rosettes of phillipsite and clinoptilolite 311 

10. Chabazite on clay lining in relict shards 314 

11. Chabazite on a mat of smectite lining a shard 314 

12. Chabazite from locality SD-4-16, Durkee, Oregon 314 

13. Erionite, chabazite, and clay in cavity between shards 315 

14. Erionite on chabazite 315 

15. Clinoptilolite coexisting with erionite 315 

16. Mordenite ribbons and filaments 315 

TABLES 

1. Paragenesis of authigenic minerals at Durkee, Oregon 309 

2. Chemical and physical data for shards from three fresh vitric tuffs from 

Durkee basin, Oregon 310 

3. Diatom species and varieties at Durkee, Baker County, Oregon 312, 313 

302 Studies in Diagenesis 
Abstract 

Authigenic zeolites derived from siliceous tuffs deposited 

in a late Miocene saline, alkaline lake near Durkee, Oregon, 

formed diagenetically at low pressure and temperature. Here, 

several rhyolitic air-fall and reworked tuffs and an ash-flow tuff 

are interbedded with diatomite, mudstone, sandstone, siltstone, 

and conglomerate. The zeolites chabazite, erionite, and clinoptilolite are common in the altered tuffs; minor amounts of 

analcime, phillipsite, and mordenite have been recognized as 

have authigenic potassium feldspar, silica minerals, and clay 

minerals. The diagenetically formed zeolites and authigenic 

potassium feldspar are restricted to an 18-km2 area in the 

southeastern part of the 120-km2 ancient lake. Altered tuffs 

have been recognized only in the middle and lower 135m of 

the exposed 350-m section in the deepest part of the lake basin. 

Fresh vitric tuffs are interbedded with a thick diatomite in the 

upper part of the stratigraphic section. The 102-km2 area of the 

basin bordering the zeolitic alteration zone contains tuffs 

equivalent to the zeolitic beds but which still contain relatively 

fresh glass. Reactions between silicic vitric tuffs and alkaline, 

saline-lake basin waters (pH >9) probably account for the 

observed generalized paragenetic sequence: glass -smectite 

-phillipsite -chabazite -erionite -clinoptilolite -mordenite, 

and then any of these zeolites -analcime -potassium feldspar. 

Evidence for the required high salinity and alkalinity are the 

saline mineral crystal molds observed in many altered tuffs and 

the Magadi-type chert found in a zeolitic tuff and mudstone 

unit. 

INTRODUCTION 

Zeolites, clay minerals, silica minerals, and potassium feldspar of diagenetic origin are common in many 

Cenozoic volcanogenic lacustrine rocks. A model for the 

alteration pattern of siliceous tuffs in such deposits has 

been developed from studies of the Tecopa-Shoshone area, 

California; the Mud Hills near Barstow, California; and 

the Wikieup area, Arizona (Sheppard and Gude, 1968, 

1969, 1973a). Other lake deposits with similar histories 

are known in this country and elsewhere, such as in the 

rift valleys of east-central Africa (Surdam and Eugster, 

1976). Features which characterize these deposits are: (1) 

beds of siliceous vitric tuffs that were deposited in saline, 

alkaline, closed-basin lakes, and (2) near-surface 

temperatures and pressures that prevailed during the 

alteration of the volcanic ash (glass). 

During mid.:fertiary time, a restricted, intermontane 

lacustrine basin near the present-day site of Durkee, Baker 

County, Oregon, received nearly continuous input of 

volcanic sediment (Prostka, 1967). Sporadically, however, 

the valley outlet at the south end was dammed, forming 

a lake, the level of which fluctuated when the dam 

breached and redeveloped. During the frequent times of 

desert climate, the lake often was shallow enough to allow 

playas to form. Rigorous evaporation conditions in the 

playas concentrated dissolved salts in the water and 

developed the high salinity and alkalinity which elsewhere 

are known to alter tuffaceous material to zeolites (Sheppard and Gude, 1968, 1969, 1973a). At other times the 

lake was fresh to slightly brackish; diatomites and other 

facies characteristic of such lacustrine environments can 

be noted today interfingered and interbedded with the 

zeolitic formations. The actual alteration of volcanic glass 

to zeolites at Durkee was limited to a small, 18-km2 area 

in the southeastern part of the 120-km2 basin. Here, the 

lake ponded at the choked outlet, and the buildup of high 

salinity and alkalinity was not diluted by freshwater mixing. At least 20 distinct zeolitic tuffs have been found in 

this area, one of which is as much as 4 m thick and consists locally of nearly monomineralic chabazite. 

The interplay between fluvial and lacustrine 

environments in the Durkee deposit, and the relations between the zeolitic tuffs and freshwater assemblages of 

diatomites and "unaltered" volcanic glass, are more complex than in most of the saline, alkaline lake deposits of 

zeolites heretofore described. These relationships and the 

unusual thickness of the chabazite beds suggest that a 

detailed analysis of the low-temperature, low-pressure 

mineral reactions that took place in this deposit would 

add substantially to our understanding of diagenesis in 

volcanogenic lacustrine rocks. The present paper describes 

the occurrence and mineral chemistry of the zeolites, clays, 

and other authigenic minerals in the lacustrine rocks of 

the Durkee basin and relates these phases to the diagenetic 

history of the area. 

GEOGRAPHY OF THE DURKEE AREA 

Durkee is well known to zeolite investigators as the 

type locality of erionite (Eakle, 1898; Staples, 1957; Staples 

and Gard, 1959; Sheppard and Gude, 1975). It is a small 

hamlet on the Burnt River in eastern Oregon (fig. lA) 

where Interstate Highway 84 and the Union Pacific 

Railroad transect the basin. The highway provides easy 

access from Baker, 42 km to the northwest. Tributary 

streams to the Burnt River cut badland-type gullies down 

through the soft lake sediments. Veneers of gravel, formed 

during two Quaternary cycles, are draped over parts of 

the exhumed lake sediments and onto the stream benches 

and valley floors. Scattered remnants of unaltered 

Mazama ash are present in perched Holocene stream 

channels. 

GEOLOGIC SETTING 

The geology of the area is shown in detail on a ISminute geologic map of the Durkee quadrangle by Prostka (1967) who delineated all but the westernmost part of 

the late Miocene lacustrine basin and the surrounding 

rocks. The older rocks, which are as much as 650 m above 

Zeolitic Diagenesis of Tuffs in an Upper Miocene Lacustrine Deposit near Durkee, Oregon 303 
A 

Boise 

SD-59/ 

•2 

Iron A 

Mountain 

.106 

1lr25' 

EXPLANATION 

TI, sedimentary rocks with } 

unaltered, intercalated tuffs 

Tiz, sedimentary rocks with . 

zeolitic, intercalated tuffs Late M1ocene 

Ash-flow tuff, welded 

and partly welded 

} 

Pre-late 

Sedimentary, metamorphic Miocene 

and igneous rocks and older 

rocks 

TERTIARY 

TERTIARY TO 

PERMIAN 

\ 

..... 

5 ' MILES 

I 

5 KILOMETERS 

----- Margin of lacustrine 

basin at Durkee 

I 

o o • o •••••• o •• Margin of zeolitic zone 

----- Contacts 

e40 Sampled locality 

..__ Measured section so~ 

Figure 1 (above and facing page). Late Miocene lacustrine basin showing sampled localities (f\) and geology (8) at Durkee, Oregon. 

304 Studies in Diagenesis 
11r3o' 

the basin floor, are Permian argillites and greenschists on 

the north and east and Triassic phyllites and marble on 

the south. Some Triassic gabbro and patches of the 

Miocene Columbia River Basalt Group are present in the 

older rocks. Iron Mountain, which rises as a monadnock 

about 25 m above the valley just west of Durkee Creek, 

is a conspicuous mass of the Columbia River Basalt. 

Figure 1B, which was modified from the maps of 

Prostka (1967) and Gilluly (1937), shows the sampled 

localities, the location of measured sections, and the 

approximate boundary between unaltered, fresh, glassbearing lacustrine rocks and the rocks containing 

1lr25' 11no· 

\ 

..... 

5 MILES 

5 KILOMETERS 

diagenetic zeolites. A few sample localities away from the 

zeolite area, where a trace amount of zeolite was found, 

are not included. These isolated localities represent 

detached microenvironments in which the immediate surrounding water attained the required geochemistry to react 

with the tuffaceous material. The result is the miniscule 

amount of the same zeolites found in the nearby zeoliterich deposits. 

Sediments derived from older rocks are present as 

detrital grains in the lacustrine mudstones, siltstones, 

sandstones, and conglomerates and as minor amounts 

of quartz, plagioclase, and amphibole minerals in the 

Zeolitic Diagenesis of Tuffs in an Upper Miocene Lacustrine Deposit near Durkee, Oregon 305 
intercalated tuffs. Silicic vitric material, commonly referred 

to as "volcanic ash:' was contributed from the fallout 

plumes of nearby volcanic centers and is present as finegrained bubble-wall tuffaceous material in the mud, silt, 

and sandstones and as coarser pumice fragments in some 

conglomerate units. Major Plinian eruptive events formed 

the relatively clean, distinctive vitric tuffs by direct airfall, 

and reworked ash rapidly swept into the lake from the surrounding hillsides. 

These tuff units are mostly thin-bedded layers 1-25 em 

thick. Several composite units that are uniform in appearance are as much as 3m thick. A single "welded ash-flow" 

unit, hereafter called "welded tuff,' as much as 15 m thick, 

occurs at or near the base of the lacustrine beds and is the 

only basinwide unit that can be used as a reliable stratigraphic marker bed. Except where the welded tuff is present, 

correlation of beds across the basin is nearly impossible. 

A set of east-west normal faults with displacements 

of as much as 10 m is shown on Prostka's map (1967). 

Abundant shorter and smaller displacement en echelon 

faults were noted in the field, but are too numerous to 

show in figure 1. Erosion has obliterated outcrops of many 

of the thin-bedded sedimentary tuffs, especially where the 

tuffs are in diatomaceous or mudstone sequences. 

The Durkee zeolite deposit differs from other Tertiary saline, alkaline lakes in that the perquisite saline, 

SOUTHWEST 

METERS 

700 

600 

500 

400 

300 

200 

100 

SEA LEVEL 

G·;·;··············;·~ ·:·:·:=::::;:;:::::·:· 

Pre-lacustrine 

rocks 

Fresh 

glass 

alkaline water formed as an encapsulated body in a fresh 

water or slightly brackish water lake with a flourishing 

diatom population. 

STRATIGRAPHY 

Figure 2 is a schematic diagram across the deep 

southeastern part of the basin. It shows our concept of 

the pattern of alteration at Durkee based on field observations and laboratory analyses. The encapsulated zeolite 

zone is shown before erosion removed the immediately 

overlying rocks. The diatomite and fresh tuffs probably 

covered the core area, because unaltered diatomaceous 

rocks are juxtaposed with zeolitic units in faulted masses 

deep within the central part of the basin. 

A composite columnar section, figure 3, shows 

the approximate sequence of the various rock units. No 

single exposure displays the uninterrupted section. 

Therefore, we were not able to make precise correlation between separated sections. The middle part of the 

diagram depicts a section in which continuity of the beds 

above the zeolite-bearing rock and the overlying 

freshwater units is diffuse and only an approximation. 

Three measured sections-SD-4, SD-32, and SD-8-

were used (fig. 1A). 

EXPLANATION 

Zeolite 

5 KILOMETERS 

Analcime 

and (or) 

Potassium 

feldspar 

NORTHEAST 

Figure 2. Schematic diagram of a stratigraphic cross section of the Durkee basin. 

306 Studies in Diagenesis 
xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

xxxxxxxx 

EXPLANATION 

DIATOMITE 

MUDSTONE 

SILTSTONE, SANDSTONE, 

AND CONGLOMERATE 

lxxxxxxxxl 

TUFF 

WELDED ASH-FLOW 

TUFF 

METERS 

50 

40 

30 

20 

10 

Figure 3. Composite columnar section of 

Durkee lacustrine and fluvial rocks. 

Figure 4. View looking northwest from locality D-30 across 

an eroded amphitheatre. Resistant beds and pinnacles are in 

zeolitized part of the basin. Freshwater diatomaceous section 

is in background, on left. Person in lower middle is shown for 

scale. Numbers are those of units in measured section SD-4 

(see appendix 1). 

SEDIMENTATION 

Tuffs that are altered to zeolites generally form resistant beds which are readily recognized in the predominantly mudstone sequences that filled the basin. Of the 

approximately 350 m of section that was measured, less 

than 25 percent of the total thickness, including the welded 

tuff, is recognizable as tuff beds. Calcareous or silicified 

tuffaceous units found outside the altered area commonly 

mimic the resistant zeolitic tuff beds. In the lower part 

of the stratigraphic section where composite tuff units or 

closely spaced intercalated tuffs are present and the alteration is complete, the resistant beds form prominent cliffs 

and free-standing pinnacles. Figure 4 is a view northwest 

across an eroded amphitheater floored by the welded tuff 

and walled by sheer cliffs and spires composed of the thick 

tuffs measured in section SD-4. Figure 5 is a panoramic 

view showing most of the lacustrine beds north of 

Interstate 84, taken from a site 1 km east of Durkee, looking northeast up Durkee Creek. 

The light-colored rocks in the left background of 

figure 4 are the thick diatomites and diatomaceous 

mudstones and tuffs depicted in figure 2 that overlie the 

prominent resistant beds and pinnacles of the zeolite zone 

seen in the foreground. Several nearly fresh vitric tuffs 

measured in the middle of section SD-8 probably are 

lateral equivalents of tuffs measured in sections SD-4 and 

SD-32 in which they are completely altered to zeolites. 

Correlations, as noted above, were difficult to make, 

especially in the poorly preserved outcrops of unaltered 

rocks such as seen in figure 4. 

Zeolitic Diagenesis of Tuffs in an Upper Miocene lacustrine Deposit near Durkee, Oregon 307 
Figure 5. View to northeast from site about 1 km east of 

Durkee, Oregon showing light-colored diatomite exposed at 

the head of eroding gullies. Arrow indicates amphitheater 

(fig. 4). 

MATERIALS AND METHODS 

Appendix 1 is a compilation of the sample descriptions made from field notes and the mineralogy determined 

in the laboratory. All sample-collecting localities are shown 

in figure 1. Most samples were collected at tuff outcrops 

where the zeolite alteration appeared to be most intense; 

however, some samples were collected away from the main 

zeolite zone to examine the mineralogy of comparable tuffs 

in obvious freshwater areas of the basin. Multiple specimens were generally collected at a single site where 

lithologic differences were visible within a sampled unit. 

In addition, subsamples were prepared in the laboratory 

if close inspection revealed apparent mineralogical or 

lithological differences that were not noticed at the outcrop. 

From about 100 field sites, about 380 separate 

specimens were examined. A conventional X-ray powder 

diffraction (XRD) pattern using monochromatic CuKa 1 

radiation and a horizontal scanning diffractometer was 

made for each specimen handpacked in a back-loaded 

aluminum holder. The relative abundance of each mineral 

identified in a single specimen was estimated in parts-often from comparison of the visually integrated pattern of 

each identified phase on the XRD record. Unaltered glass, 

opaline silica, and some poorly crystalline or very fine 

grained clay minerals were estimated with less certainty because of the broad and diffuse nature of the XRD scattering. 

Selected specimens were examined in thin section, 

in immersion oils by optical microscopy using a polarizing petrographic microscope, and by scanning electron 

microscopy (SEM). For SEM, specimens were mounted 

308 Studies in Diagenesis 

with silver paste on 3.2-cm Cambridge-type stubs and 

sputter coated with gold or gold-palladium. The SEM 

images were produced with Cambridge Stereoscan S-180 

and Stereoscan 250 MK2 instruments 1• Chemical analyses of shards from three vitric tuffs from the Durkee 

basin were made by conventional rock analysis techniques 

by Elaine L. Brandt (U.S. Geological Survey, Denver, 

Colo.). 

MINERALOGICAL RESULTS 

The locality, description, and mineralogical composition of all samples examined are listed in appendix 1. 

Clay minerals 

Although the clay mineralogy in these samples was 

not studied rigorously, three categories of materials were 

recognized based on the prominent basal XRD reflections 

near 10 and 14 A. Some of the 14-A smectite and the 10-A 

mixed-layer clays may be authigenic; however, the 7-A clay 

found in a few samples is most likely detrital in origin. 

The 7-A phase was found only in samples collected away 

from the zeolitic area or in clearly identified detrital-rich 

samples at the boundary between the zeolitic and the 

freshwater units. Samples from mudstones, siltstones, 

sandstones, or conglomerates which showed a 10-A peak 

probably contain detrital mica rather than an authigenic 

illite or mixed-layer clay. 

Zeolites and authigenic feldspar 

Other than analcime, which is discussed later with 

potassium feldspar, the five zeolites derived from the siliceous vitric tuffs were generally present in mixed 

assemblages or, less commonly, in nearly monomineralic 

beds. A group of two or three zeolite varieties in a single 

specimen was common. Seven samples were found to contain four zeolites, and one, sample D-30K, contained five 

zeolites including analcime. Clinoptilolite was the zeolite 

most frequently found; erionite was next followed by 

chabazite, phillipsite, and mordenite. 

Analcime and potassium feldspar 

Five of the six zeolites found at Durkee were authigenically formed by the alteration of the vitric tuffs. 

Analcime has not been shown conclusively to have formed 

1Use of trade and company names is for descriptive purposes only 

and does not imply endorsement by the U.S. Geological Survey. 
from glass here at Durkee, nor in any other lacustrine setting. A precursor zeolite, such as any of the other five 

zeolites present at Durkee, is apparently required for the 

development of analcime (Sheppard and Gude, 1973b ). 

Potassium feldspar, like analcime, was also formed 

from a precursor zeolite at Durkee. The zeolite mineral 

may have been any of the five diagenetically derived from 

glass, or it may have been analcime. Potassium feldspar 

has not been found directly associated with siliceous vitric 

material in any of the saline, alkaline lake deposits 

examined to date. The particular variety of feldspar found 

at Durkee has chemical and crystallographic properties 

which set it apart from feldspars found in more classic 

igneous environments (see Sheppard and Gude, 1973b ). 

These authigenic potassium feldspars may contain as 

much as 1 percent B2

0 3 substituting for Al20 3 (Sheppard 

and Gude, 1965), and have m:arly the stoichiometric 

amount of K20 required for a sanidine; thus, they contain almost no Na2

0. The unit-·cell of these potassium 

feldspars is always distorted in the a-axis dimension that 

is predicted for a nearly pure Banidine (Sheppard and 

Gude, 1973). The euhedral monoclinic crystals formed at 

Durkee are 2-10 ~-tm in size. 

Silica minerals 

Two silica phases present at Durkee that are clearly 

part of the diagenetic process are opal-CT and Magaditype chert. The latter material was found only at locality 

D-107. Some of the quartz found at Durkee is authigenic 

and is easily recognized in petrographic thin sections. 

Most of the quartz determined by XRD is detrital or 

pyrogenic and had no part in the: lacustrine diagenesis of 

the tuffs. 

Other minerals 

Other minerals listed in appendix 1 are mostly 

clastic detritus or late-formed calcite or gypsum. Dolomite 

in a thin bed outside the south end of the zeolite zone 

is probably not part of the glass-to-zeolite diagenetic process. Similarly, apatite found in a thin bed in an isolated 

outcrop, also near the south end of the lake beds, is an 

unexplained anomaly. 

DIAGENESIS AT DURKEE 

"Zeolitic diagenesis" is almost a redundant phrase, 

especially when it is used to denote a process such as that 

in a setting like Durkee. Diagenesis at shallow-lake water 

temperatures and surface ambient pressures yield zeolites 

if the required water geochemistry and fresh siliceous vitric 

Table 1. Paragenesis of authigenic minerals at Durkee, Oregon 

[Modified from Gude and Sheppard, 1978; earliest minerals listed on left] 

Glass, then: 

Phillipsite -- Clinoptilolite + Erionite 

Phillipsite -- Erionite 

Smectite -- Clinoptilolite 

Smectite -- Chabazite -- Clinoptilolite 

Smectite -- Phillipsite -- Clinoptilolite 

Smectite -- Erionite -- Clinoptilolite 

Smectite -- Chabazite -- Erionite 

Phillipsite-- Clinoptilolite 

Phillipsite -- Clinoptilolite + Erionite 

Phillipsite -- Clinoptilolite + Mordenite 

Phillipsite -- Erionite + Chabazite 

Chabazite -- Eronite + Clinoptilolite 

Chabazite-- Clinoptilolite -- Mordenite 

Any of the above zeolites (but not glass or clay minerals) -- Analcime~ Potassium feldspar 

tuff are present. The generalized sequence of formation 

(modified from Gude and Sheppard, 1978) of zeolites 

and associated authigenic silicate minerals is glass - smectite -phillipsite -chabazite -erionite -clinoptilolite -

mordenite and then any of the preceding zeolites -analcime - potassium feldspar. Table 1 gives the detailed 

paragenetic sequence for the authigenic minerals found 

at Durkee. It was assembled from petrographic microscope 

studies, XRD mineral identifications (appendix 1), and 

SEM data supplemented with energy dispersive X-ray 

spectroscopy (EDX). The following discussions are phaseby-phase descriptions for the evidence of diagenesis. 

Water 

No residual formational or interstitial water from 

the late Miocene lake at Durkee has been found. Thus, 

assumptions about the water geochemistry have been 

based on studies of modern lakes such as Lake Magadi, 

Kenya (Surdam and Eugster, 1976); Owens and Searles 

Lakes, California (Smith, 1966); and Teels Marsh, Nevada 

(Surdam and Mariner, 1971); and from extensive work by 

Jones (1965, 1966) and Jones and others (1967, 1977) on 

lakes in northern California and southern Oregon. In 

addition to these present-day analogs, certain inferences 

about water chemistry can be made from field and 

laboratory observations. 

At Durkee and other lacustrine basins (Sheppard 

and Gude, 1968, 1969, 1973a), crystal molds and casts are 

present that have the morphology of gaylussite or 

pirssonite, which are soluble sodium carbonatebicarbonate minerals. Similar minerals have been found 

on playa lake bottoms in saturated brines of modern 

"soda lakes" (Smith and Friedman, 1986). During rapid 

sediment filling of such a lake, clay-size detritus would 

cover the crystals of the sodium minerals, and thus produce the crystal molds. 

Zeolitic Diagenesis of Tuffs in an Upper Miocene Lacustrine Deposit near Durkee, Oregon 309 
Figure 6. Ripple-marked bed on surface of an altered tuff at 

locality D-65, Durkee, Oregon (knife is about 7.5 em long). 

Figure 7. Scanning electron micrograph of partly etched glass 

fragment, from sample SD-4-3GX, Durkee, Oregon. 

Crystals casts of an unidentified mineral found 

on the underside of a tuff were collected at several localities at Durkee. The deposition of these crystals was sufficiently widespread within the zeolite area for this tuff 

unit to serve as a local marker bed. Well-preserved, ripplemarked beds (fig. 6) are additional evidence of a shallowwater, playa-lake deposition. 

Glass 

Early Pliocene volcanic events supplied the silicic 

vitric airfall ash, with each tuff having a slightly different 

composition. Table 2 gives conventional silicate rock 

analyses and measured indices of refraction for three tuffs 

which are representative of those in the Durkee basin. 

31 0 Studies in Diagenesis 

Table 2. Chemical and physical data for shards from three fresh 

vitric tuffs from Durkee basin, Oregon 

[Elaine L. Brandt, analyst] 

Con vent ion a 1 rock ana 1 yses (weight percent) 

Sample No. 1 ••••••••••••••• 2 3 

Si02 73.33 71.67 68.12 

Al 2o3 11.71 11.78 13.77 

Fe2o3 0.64 o. 71 o. 77 

FeO 0.29 1.24 1.04 

MgO 0.11 0.19 0.34 

CaD 0.35 0.81 1.18 

Na 2o 2.62 2.21 3.75 

K20 5.76 5.04 4.14 

H20+ 4.33 5.42 5.57 

H20- 0.24 0.17 0.22 

Ti02 0.11 0.38 0.39 

P205 0.01 0.03 0.07 

MnO 0.08 0.04 0.07 

Cl 0.07 0.03 0.15 

0.07 0.08 0.05 

Subtotal ••••••••••• 99.72 99.80 99.63 

Less 0 ••••••••••••• 0.05 0.04 0.05 

Total •••••••••••••• 99.67 99.76 99.58 

Index of refraction ••• 1.497 1.503 1.508 

Shard morphology ...... 40-50% All platy 10-20% 

pumice. bubble-wall. pumice. 

lField numbers are as on figure 1. 

1. Field No. 90, Lab. Serial No. D-103730. 

2. Field No. SD-8-llC, Lab. Serial No. D-103732. 

3. Field No. SD-8-9A, Lab. Serial No. D-103733. 

The morphology of the glass shards also varied from 

place to place. We were unable to collect matched samples 

from a single tuff unit which showed the compositional 

change from fresh glass to a completely altered zeolitic 

tuff. Thus, no data are available from Durkee to document the change in chemistry during diagenesis. Figure 7 is a scanning electron micrograph showing a partly 

dissolved glass fragment from a specimen taken at the 

top of the welded tuff. The glass material removed by 

the highly saline, alkaline water may have formed a gel, 

which then crystallized to a zeolite (Surdam and Mariner, 

1971). Alternatively, in the absence of direct evidence 

such as entrapped water or gel, the dissolved glass components may have precipitated directly to a zeolite from 

solution. According to field and laboratory observation 

by Surdam and Eugster (1976), the composition of the 

water and the glass and the pH of the solutions may determine which alkalic zeolite forms. The exact paragenesis 
Figure 8. Scanning electron micrograph showing smectite 

coating the surface of a glass fragment, from sample SD-4-

3GY, Durkee, Oregon. 

Figure 9. Scanning electron micrograph showing spherical 

rosettes of phillipsite (Ph) and clinoptilolite (Cp), from sample 

SD8-6C, Durkee, Oregon. 

cannot as yet be predicted in modern closed-lake systems, 

but the series reported here for the Durkee deposits fits 

the patterns studied in other lakes. 

Clay minerals 

Smectite is ubiquitous in the altered tuffs in the 

Durkee basin, where it was generally the first formed mineral phase found as a thin coating on partly altered glass 

shards (fig. 8). This clay layer can be seen petrographically 

to outline relict shard textures after the zeolitic alteration 

was complete and no glass remained. Hay, (1963, p. 249) 

demonstrated that '~teration of glass to montmorillonite 

(is)* * * a hydrolysis reaction" during his work on zeolites in the middle Tertiary John Day Formation. More 

recently, Dibble and Tiller (1981) showed that kinetic processes are important factors in the glass-clay-zeolitefeldspar diagenesis and can explain the brief time needed 

to deposit the thin clay layers. 

Detailed study of clay-mineralogy diagenesis, other 

than determining the presence of three broad categories 

based on the low-angle, strong XRD maxima as 14-A, 

10-A, or 7-A clays has not been done at Durkee. XRD 

methods can detect small quantities of well-crystallized 

phases; however, the clay minerals at Durkee are not only 

poorly crystalline, but also are present in minute amounts, 

thus making detection by XRD in multiphase samples 

nearly impossible. Scanning electron microscopy provides 

a useful means of examining such materials; figure 8 and 

other illustrations for the separate zeolite minerals show 

the typical smectite morphology. 

Diatoms and diatomite 

Diatomite beds and interbedded units of sandstone, 

siltstone, mudstone, and airfall tuff dominate the 

lacustrine sequence in most of the Durkee basin. The lightcolored units visible in outcrops at the eroding heads of 

gullies and in road cuts (fig. 4) are chiefly diatomite. Eight 

samples from three localities were submitted to G. W. 

Andrews for examination and determination of the 

diatoms present. Two of the samples collected in the 

zeolitized area (fig. 1B) contained no diatoms, even though 

approximately correlative units a few tens of meters away 

beyond the limit (fig. 1B) of zeolitic alteration are 

diatomites. Presumably, the alkaline water within the 

zeolite area was capable of dissolving the tests of diatoms 

that may have been present, or diatoms did not thrive in 

such saline alkaline solutions. 

Table 3 lists about 60 varieties of diatoms and their 

relative abundance found in the six diatomite-rich 

specimens (G. W. Andrews, U.S. Geological Survey, 

National Museum, Washington, D.C., written commun., 

1971). Recent paleontological diatom investigations in the 

Durkee basin by J. Platt Bradbury (U.S. Geological 

Survey, Denver, Colo., oral commun., 1985) support most 

of the findings of G. W. Andrews and have provided a 

certain age as late Miocene. 

Phillipsite 

Phillipsite was observed in thin section as the first 

zeolite to have formed in the deposits at Durkee. This is 

the pattern found by Sheppard and Gude (1968, 1969, 

1973a) for such zeolite alteration of volcanic glass in 

saline, alkaline lake deposits. Figure 9 is an SEM 

photograph of rosettes and 10-~tm elongate bundles of 

Zeolitic Diagenesis of Tuffs in an Upper Miocene Lacustrine Deposit near Durkee, Oregon 311 
~ t..J 

til 

2' 

c. 

;;· 

"' :;-

0 

~-

:I 

!l 

;;· 

Table 3. Diatom species and varieties at Durkee, Baker County, Oregon 

[Relative frequency indicated by: A, abundant, C, common, R, rare; leaders(---) indicate none found. Analysts: G. W. Andrews and J. P. Bradbury, U.S. Geological Survey] 

Field No. (see fig. 1, and appendix 1) •••••••••••••••••••••• 

USGS diatom locality No ••••••••••••••••••••••••••••••••••••• 

Achnanthes lanceolata Brebisson •••••••••••••••••••••••••••• 

Aulacosira ................................................. . Amphora ovalis (Kutzing) Kutzing ••••••••••••••••••••••••••• 

Amphora sp •• •••••••••••••••••••••••••••••••••••••••••••••••• 

Caloneis sp ••••••••••••••••••••••••••••••••••••••••••••••••• 

Cocconeis placentula var. lineata (Ehrenberg) Cleve ••••••••• 

Cocconeis placentula Ehrenberg •••••••••••••••••••••••••••••• 

Cymatopleura solea (Brebisson) w. Smith ••••••••••••••••••••• 

Cymbella mexicana (Ehrenberg) A. Schmidt •••••••••••••••••••• 

Cymbe 11 a sp ••••••••••••••••••••••••••••••••••••••••••••••••• 

Diploneis ovalis (Hilse) Cleve •••••••••••••••••••••••••••••• 

Diploneis ovallis var. oblongella (Naegli) Cleve •••••••••••• 

Diploneis smithii (Brebisson) Cleve ••••••••••••••••••••••••• 

Diploneis sp •••••••••••••••••••••••••••••••••••••••••••••••• 

Epithemia sorex Kutzing •••••••••••••••••••••••••••••••••••• 

Epithemia sorex var. gracillis Hustedt •••••••••••••••••••••• 

Epithemia sp A •••••••••••••••••••••••••••••••••••••••••••••• 

Epi themi a sp •. .............•...........•••..........•....... 

Eunotia serra Ehrenberg ••••••••••••••••••••••••••••••••••••• 

Eunotia valida Hustedt •••••••••••••••••••••••••••••••••••••• 

Eunot i a sp •••••••••••••••••••••••••••••••••••••••••••••••••• 

Fragilaria brevistriata Grunow •••••••••••••••••••••••••••••• 

Fragilaria leptostauron (Ehrenberg) Hustedt ••••••••••••••••• 

Fragilaria leptostauron var. dubia (Grunow) Hustedt ••••••••• 

Fragilaria leptostauron var. dubia Grunow ••••••••••••••••••• 

Fragilaria leptostauron var. A •••••••••••••••••••••••••••••• 

Fragilaria leptostauron var. B •••••••••••••••••••••••••••••• 

Fragilaria sp ••••••••••••••••••••••••••••••••••••••••••••••• 

Gomphonema intricatum var. vibrio (Ehrenberg) Cleve ••••••••• 

Gomphonema sp. A ..•........•••........••••......••.•••...... 

(ic)tnphonema sp. 8 • ••••••••••••••••••••••••••••••••••••••••••• 

Gomphonema sp ••••••••••••••••••••••••••••••••••••••••••••••• 

Hantzschia amphioxys (Ehrenberg) Ralfs •••••••••••••••••••••• 

Navicula amphibola Cleve •••••••••••••••••••••••••••••••••••• 

Navicula lacustris Gregory •••••••••••••••••••••••••••••••••• 

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Navicula oblonga Kutzing ·~································· Navicula aff. N. oblonga Kutzing ••••••••••••••••••••••••••• 

Navicula scutelloides W. Smith •••••••••••••••••••••••••••••• 

Navicula sp ••••••••••••••••••••••••••••••••••••••••••••••••• 

Nitzschia sp •••••••••••••••••••••••••••••••••••••••••••••••• 

Opephora martyi Heribana ••••••••••••••••••••••••••••••••••• 

Pinnularia borealis Ehrenberg ••••••••••••••••••••••••••••••• 

Pinnularia brevicostata Cleve ••••••••••••••••••••••••••••••• 

Pinnularia dactylus .. var. dariana (A. Schmidt) Cleve ••••••••• 

Pinnularia major (Kutzing) Rabenhorst ••••••••••••••••••••••• 

Pinnularia aff. P. major (Kutzing) Rabenhorst ••••••••••••••• 

Pinnularia mesolepta fo. angusta Cleve •••••••••••••••••••••• 

Pinnularia viridis (Nitzsch) Ehrenberg •••••••••••••••••••••• 

Pinnularia sp ••••••••••••••••••••••••••••••••••••••••••••••• 

Pleurosigma angulatum (Queckett) W. Smith ••••••••••••••••••• 

Rhopalodia gibba (Ehrenberg) 0. Muller ··············~······ Rhopalodia gibberula var. ventricosa (Ehrenberg) 0. Muller 

Stauroneis phoenicenteron Ehrenberg ••••••••••••••••••••••••• 

Stauronei s sp ••••••••••••••••••••••••••••••••••••••••••••••• 

Surirella ovata Kutzing •••••••••••••••••••••••••••••••••••• 

Su ri re 11 a sp •••••••••••••••••••••••••••••••••••••••••••••••• 

Synedra ulna (Nitzsch) Ehrenberg •••••••••••••••••••••••••••• 

Tetracyclus ellipticus (Ehrenberg) Grunow ••••••••••••••••••• 

Tetracyclus lacustris Ralfs ••••••••••••••••••••••••••••••••• 

Tetracycl us sp •••••••••••••••••••••••••••••••••••••••••••••• 

*Thalassiosiraceae (?} Type A ••••••••••••••••••••••••••••••• 

*Thalassiosiraceae (?} Type 8 ••••••••••••••••••••••••••••••• 

*Thalassiosiraceae, new genus ••••••••••••••••••••••••••••••• 

c 

R 

R 

R 

c 

c 

R 

c 

Mineralogy (see appendix 1). 

Glass ••••••••••••••••••••••••••••••• 

Opal C-T •••••••••••••••••••••••••••• 

7 -A c 1 ay •••••••••••••••••••••••••••• 

10-A clay........................... 1 

14-A clay ••••••••••••••••••••••••••• Quartz.............................. 5 Feldspar............................ 4 

R 

R 

R 

R 

R 

c 

R 

c 

R 

R 

R 

R 

R 

R 

c 

R 

R 

R 

R 

R 

R 

R 

R 

c 

R 

R 

R 

R 

R 

R 

R 

A 

R 

R 

R 

R 

R 

c 

c 

A 

Estimated abundance reported as parts in ten 

1 

2 

4 

3 

1 

5 

3 

1 

5 

1 

1 

2 

1 

2 

3 

3 

2 

5 

1 

2 

2 

Trace 

R 

R 

R 

R 

R 

R 

R 

c 

R 

c 

1 

4 

4 

1 

R 

c 

R 

R 

R 

R 

c 

c 

I{ 

A 

c 

2 

3 

1 

2 

2 

*The discoid centric diatoms in the Durkee deposits are undescribed genera and species of the family Thalassiosiraceae, apparently 

intermediate between the genera Thalassiosira and Stephanodiscus. Similar diatoms are also found in the late Miocene Chalk Hills Formation, 

and in other deposits of similar age. 
Figure 10. Scanning electron micrograph of chabazite (Ch) 

on clay (Cy) lining in relict shards from sample SD-4-2, Durkee, 

Oregon. 

Figure 11. Scanning electron micrograph of chabazite on a 

mat of smectite lining a shard, from sample SD-4-16, Durkee, 

Oregon. 

phillipsite which appear to have formed earlier than the 

perched crystals of clinoptilolite. 

The scarcity of phillipsite is probably due to rapidly 

changing diagenetic conditions. After phillipsite formed, 

the pH of the reacting waters probably increased (Surdam and Sheppard, 1978, p. 149) which favored the formation of the other zeolites in the paragenetic sequence 

shown in table 1. 

Chabazite 

Chabazite from Durkee was described by Gude and 

Sheppard (1978). The SEMs in figures 10 and 11 are new 

to the present study and show chabazite on mats of 

314 Studies in Diagenesis 

Figure 12. Scanning electron micrograph of chabazite from 

sample SD-4-16, Durkee, Oregon, showing twinning habit and 

late-formed smectite (bottom left) encroaching on a chabazite 

crystal (from Gude and Sheppard, 1978). 

smectite which formed by the alteration of glass shard 

walls. Figure 12 (from Gude and Sheppard, 1978) shows 

the twinning habit that is common for chabazite from the 

tuffs deposited in the saline, alkaline lakes. Chabazite beds 

as seen at measured section SD-4 are commonly massive 

and nearly monomineralic through as much as a 15-m section. This is a restricted occurrence, however, and 

chabazite has not been formed in such large quantities 

in outcrops away from this area. 

Erionite 

Durkee, at locality D-29 on Swayze Creek, is the 

type locality (Eakle, 1898) for the original woolly variety 

of erionite. Erionite in tuffs in the Durkee lake beds first 

reported by Sheppard and Gude (1975) is now known to 

be one of the common zeolites in the basin. It clearly 

formed later than chabazite as is seen in figures 13 and 

14. The woolly crystal habit found at the type locality has 

not been seen elsewhere in the Durkee area. Elongate, 40-

1-'m rods of erionite, commonly in bundles of parallel 

needles, is the common form in the lacustrine tuffs as 

shown in figure 13. 

Clinoptilolite 

Although clinoptilolite is widespread in the zeolitic 

areas at Durkee, it has not been found in monomineralic 

units more than a few centimeters thick. Erionite coexisting with clinoptilolite can be seen in figure 15. Clinoptilolite crystals, nearly 1-mm wide, were readily observed 

with a field hand lens at one sample locality, D-66c. Here, 
Ch-+ 

Er Cy 

t t 

Figure 13. Scanning electron micrograph showing erionite (Er), 

chabazite (Ch), and clay (Cy) in cavity between shards from 

sample SD-4-14B, Durkee, Oregon. 

each crystal was coated with iron oxide and 14-A clay. This 

assemblage probably represents isolated kinetic conditions 

wherein smectite formed late. 

Mordenite 

Mordenite was detected by XRD in only five 

samples from Durkee. Sample D79 shows thin ribbons or 

fine filaments draped across earlier formed zeolites 

(fig. 16). Part of the apparent scarcity of mordenite may 

simply be ascribed to the difficulty in finding the very 

fine fibers. From the XRD and SEM evidence, mordenite 

clearly formed late in a paragenetic sequence. 

Analcime and potassium feldspar 

Analcime and potassium feldspar commonly coexist 

at Durkee as they do at other lacustrine zeolite localities. 

The XRD patterns of these assemblages have sharp peaks, 

suggesting well-crystallized materials. This observation is 

supported by petrographic studies in which irregular 

patches of relatively clear analcime replace shard pseudomorphs consisting of clay minerals and other zeolites. 

The preexisting vitroclastic texture commonly seen 

in zeolitic tuffs was apparently destroyed by the development of analcime or analcime + potassium feldspar. It is 

not certain why analcime and potassium feldspar at Durkee 

did not form the euhedral crystals found in similar deposits elsewhere, but it is clear from our optical microscopy 

studies that these minerals were formed from a precursor 

zeolite and did not form from, nor do they coexist with 

relict fresh glass. 

Figure 14. Scanning electron micrograph showing erionite on 

chabazite from sample SD-4-3G, Durkee, Oregon. 

Figure 15. Scanning electron micrograph showing clinoptilolite 

(Cp) with erionite (Er) from sample SD-4-11, Durkee, Oregon. 

Figure 16. Scanning electron micrograph of mordenite ribbons 

and filaments from locality D-79, Durkee, Oregon. 

Zeolitic Diagenesis of Tuffs in an Upper Miocene Lacustrine Deposit near Durkee, Oregon 315 
Silica minerals 

Quartz, opal-CT, and Magadi-type chert have been 

identified at Durkee. Authigenic quartz that formed after 

clinoptilolite has been noted by optical microscopy only 

in sample D-62A. At least 30 other samples may contain 

authigenic quartz without detrital minerals such as plagioclase, amphibole, and clay minerals. Opaline silica was 

noted at many sample localities, but it was generally associated with the diatomite. Magadi-type chert found at a 

single locality (D-107) was not directly related to any tuff 

unit. 

CONCLUSIONS 

Diagenesis at Durkes has produced on 18- to 20-

km2 zeolite deposit in a restricted part of late Miocene 

lacustrine beds which are present in a 120-km2 basin. 

The geochemical requirements of high salinity and 

alkalinity were apparently developed in a part of an 

ancient lake. Diatoms flourished in the surrounding lake 

where freshwater streams entered the lake margins. Zeolitic 

alteration of the siliceous vitric tuffs occurred near the 

narrow sediment-dammed outlet. The preservation of the 

altered zeolitic core area in an otherwise freshwater-lake 

sedimentary basin indicates a rapid kinetic process that 

may have taken place in a sheltered cove of the lake. The 

reaction of saline, alkaline, closed-basin waters with the 

vitroclastic material was able to proceed at temperatures 

and pressures of a desert-floor shallow lake, surrounded 

by a growing deposit of diatoms that coexisted with the 

formation of the authigenic zeolites. 

REFERENCES CITED 

Dibble, W. E., Jr., and Tiller, W. A., 1981, Kinetic model of 

zeolite paragenesis in tuffaceous sediments: Clays and Clay 

Minerals, v. 29, p. 323-330. 

Eakle, A. S., 1898, Erionite, a new zeolite: American Journal 

of Science, v. 155, p. 66-68. 

Gilluly, J., 1937, Geology and mineral resources of the Baker 

quadrangle, Oregon: U.S. Geological Survey Bulletin 879, 

119 p. 

Gude, A. J., 3rd, and Sheppard, R. A., 1978, Chabazite in 

siliceous tuffs of a Pliocene lacustrine deposit near Durkee, 

Baker County, Oregon: U.S. Geological Survey Journal of 

Research, v. 6, p. 467-4 72. 

Hay, R. L., 1963, Stratigraphy and zeolite diagenesis of the John 

Day Formation of Oregon: University of California Publications in Geological Science, v. 42, p. 199-265. 

Jones, B. F., 1965, The hydrology and mineralogy of Deep 

Springs Lake, Inyo County, California: U.S. Geological 

Survey Professional Paper 502-A, 56 p. 

316 Studies in Diagenesis 

___ 1966, Geochemical evolution of closed basin waters in 

the western Great Basin, in Rau, J. F., ed., Second symposium on salt: Ohio Geological Society, v. 1, p. 189-200. 

Jones, B. F., Eugster, H. P., and Rettig, S. L., 1977, 

Hydrochemistry of the Lake Magadi basin, Kenya: 

Geochimica et Cosmochimica Acta, v. 41, p. 53-72. 

Jones, B. F., Rettig, S. L., and Eugster, H. P. 1967, Silicate in 

alkaline brines: Science, v. 158, p. 1310-1314. 

Prostka, H. J., 1967, Preliminary geologic map of the Durkee 

quadrangle, Oregon: Oregon Department of Geology and 

Mineral Industries Geological Map GMS-3, text 8 p. 

Smith, G. 1., 1966, Geology of Searles Lake, in Rau, J. F., ed., 

Second symposium on salt: Northern Ohio Geological 

Society, v. 1, p. 167-180. 

Smith, G. 1., and Friedman, Irving, 1986, Seasonal diagenetic 

changes in salts of Owens Lake, 1970-77, in Mumpton, F. 

A., ed., Studies in diagenesis: U.S. Geological Survey 

Bulletin 1578, 21-29. 

Sheppard, R. A., and Gude, A. J., 3rd, 1965, Potash feldspar 

of possible economic value in the Barstow Formation, San 

Bernardino County, California: U.S. Geological Survey Circular 500, 7 p. 

___ 1968, Distribution and genesis of authigenic silicate 

minerals in tuffs of Pleistocene Lake Tecopa, Inyo County, 

California: U.S. Geological Survey Professional Paper 597, 

38 p. 

___ 1969, Diagenesis of tuffs in the Barstow Formation, Mud 

Hills, San Bernardino County, California: U.S. Geological 

Survey Professional Paper 634, 35 p. 

___ 1973a, Zeolites and associated authigenic silicate minerals 

in tuffaceous rocks of the Big Sandy Formation, Mohave 

County, Arizona: U.S. Geological Survey Professional Paper 

830, 36 p. 

___ 1973b, Boron-bearing potassium feldspar of authigenic 

origin in closed-basin deposits: U.S. Geological Survey Journal of Research, v. 1, no. 4, p. 377-382. 

___ 1975, Distribution of zeolites in Pliocene lacustrine 

rocks, Durkee basin, Baker County, Oregon: Geological 

Society of America Abstracts with Programs, v. 7, no. 3, 

p. 374. 

Staples, L. W., 1957, X-ray study of erionite, a fibrous zeolite 

[abs.]: Geological Society of America Bulletin, v. 68, 

p. 1847. 

Staples, L. W., and Gard, J. A., 1959, The fibrous zeolite erionite, 

its occurrence, unit cell, and structure: Mineralogical 

Magazine, v. 32, p. 261-281. 

Surdam, R. C., and Mariner, R. H., 1971, The genesis of phillipsite in Recent tuffs at Teels Marsh, Nevada: Geological Society of America Abstracts with Programs, v. 3, 

p. 725. 

Surdam, R. C., and Eugster, H. P., 1976, Mineral reactions in 

the sedimentary deposits of the Lake Magadi region, Kenya: 

Geological Society of America Bulletin, v. 87, no. 12, p. 

1739-1752. 

Surdam, R. C., and Sheppard, R. A., 1978, Zeolites in saline, 

alkaline-lake deposits, in Sand, L. B., and Mumpton, F. A., 

eds., Natural zeolites-Occurrence, properties, use: 

Elmsford, New York, Pergamon Press, p. 149. 
APPENDIX 1 

Zeolitic Diagenesis of Tuffs in an Upper Miocene Lacustrine Deposit near Durkee, Oregon 317 
~ 

= 

til 

i 

~· 

:r 

c 

~-

~ ;;;· 

Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, Oregon 

Notes and definitions: 

Sample number (first column) combines the locality number (given in the spanner head) and the unit or bed number (or letter). Thus, sample D-lC-2 is from locality Dl and is unit C plus 

subunit 2. All localities are shown on figure 1. Unit is used here to mean a stratigraphic piece (lithology) of a measured section. Bed is used to mean a change of lithology within a 

unit. All field measurements were made in the English system (yards, feet, inches) and then converted to exact metric equivalents. Because such exactitude may be misleading, metric values in 

the appendix are rounded up or down. For example, a bed measured as 3 ft is given as 90 em. Leaders (---) indicate no finding; Tr, traces; m, meters; em, centimeters; mm, millimeters. 0 

Others: Am, amph~bole; Ap, apatite; Do, dolomite; Gy, gypsum; Ja, ja'rosite; Lg, lignite. The following abbreviations are used for minerals: Gl, glass (noncrystalline); CT, opal-CT; lOA, 10-

angstrom clay; 14A, 14-angstrom clay; Ph, phillipsite; Ch, chabazite; Er, erionite; Cp, clinoptilolite; Mo, mordenite; An, analcime; KF, potassium feldspar; Qtz, quartz; PF, plagioclase 

feldspar; Cl, calcite. 

Sample No. Measured 

(locality thickness 

No. plus shown in m, 

bed or em, or mm 

unit No.) 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

Description of samples 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Locality D-1, NW1/4SE1/4 sec. 23, T. 11 s., R. 43 E. Thick diatomite unit with interbedded fine-grained, gray airfall tuff that appears fresh except in lowest 0.1 em 

D-lB 

D-IC-2 

D-lC-1 

D-lA 

25 em 

50 em 

0.1 ern 

Tuff ••••••••• · ••••••••••••••••• •• • ••• •• •• •• • •• • • •• • • •• • • • • • • • 

Brown, altered hard layer at base of D-lB ••••••••••••••••••• 

White underside of D-lC-2 ••••••••••••••••••••••••••••••••••• 

Diatomite, beneath D-lB ••••••••••••••••••••••••••••••••••••• 

8 

1 

10 

2 

Locality D-2, NW1/4SW1/4 sec. 4, T. 11 s., R. 43 E. Thick tuffaceous exposure including welded tuff and two other ledge-forming tuffs above welded tuff 

D-21 

D-2H 

D-2G 

D-2F 

D-2E 

D-2D 

D-2C 

D-2B 

D-2A 

90 em 

90 em below D-2G 

9.1 rn below D-2G 

Lower part of upper resistant tuff •••••••••••••••••••••••••• 

Top of middle resistant tuff •••••••••••••••••••••••••••••••• 

White bed in middle resistant tuff •••••••••••••••••••••••••• 

Pumice conglomerate at base of D-2G ••••••••••••••••••••••••• 

Blue-gray bed below middle resistant tuff ••••••••••••••••••• 

Yellow tuff with partly altered coarse shards ••••••••••••••• 

Gray tuff with plant fragments •••••••••••••••••••••••••••••• 

Welded tuff, partly altered uppermost part •••••••••••••••••• 

Middle of welded tuff ••••••••••••••••••••••••••••••••••••••• 

8 

7 

5 

8 

5 

7 

2 

10 

Tr 

2 

1 

3 

3 

2 

4 

Locality D-3, SW1/4SW1/4 sec. 19, T. 11 S., R. 42 E. Lacustrine sediments, chiefly diatomite with interbedded green mudstone and thin, gray, unaltered tuffs. Bedding attitude N. 70° W., 30° S. 

D-3C 

D-3B 

D-3A 0.15 ern 

1.5 m above D-3A 

0.15 em above D-3A 

Green mudstone •••••••••••••••••••••••••••••••••••••••••••••• 

Diatomite •••••••••••••••••••••••••••••••••••••••••• • • • • • • • • • 

Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 

2 

Tr 

1 

Locality SD-4, SEI/4SE1/4 sec. 23, T. 11 s., R. 43 E. Measured section. Bedding attitude N. 70° w., 16° N. 

SD-4-16 4.0 m +2.5 rn 

SD-4-15 2.0 rn 

SD-4-14 I. 7 rn 

SD-4-14B +1.4 rn 

SD-4-14A +90 em 

SD-4-13 1. 7 m 

SD-4-12 1.4 m 

SD-4-11 1.8 m 

SD-4-10 3.8 rn 

SD-4-9 8 ern 

SD-4-8 4 m 

Yellow, thick-bedded, massive tuff with vitric shard --- --- --- Tr 

texture; top of unit eroded. 

Yellow-gray tuff with 3-cm pumice fragments ••••••••••••••••• 

Unit 14--Yellow to white tuff, thin to thick bedded, having coarse shard texture: 

White tuff •••••••••••••••••••••••••••••••••••••••••••••••• 

White tuff •••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff with coarse shard texture ••••••••••••••••••••••• 

Brown tuffaceous mudstone, conchoidal fracture; abundant 

white cavity fillings. 

Yellow tuff, medium thick to thick bedded; locally with 

white opal segregations. 

Brown tuffaceous mudstone, crudely bedded; many small 

cavities filled with white crystals. 

Yellow tuff, with one bed consisting of pumice fragments 

as wide as 2 ern. 

Brown tuff, punky with crude bedding, coarser near base; 

many pumice fragments. 

Tr 

7 

Tr 

Tr --- --- --- 1 2 --- TrAm 

3 2 

Tr 

10 Tr 

1 ---

10 Tr Tr 

7 1 Tr 

7 --- 3 

1 

5 

Tr 

3 

8 
SD-4-7 

SD-4-6 

SD-4-5 

SD-4-4 

SD-4-3 

SD-4-3GX 

SD-4-3GY 

SD-4-3G 

SD-4-3F 

SD-4-3E 

SD-4-3D 

SD-4-3C 

SD-4-3B 

SD-4-3A 

SD-4-2 

i SD-4-1 

70 em 

5,2 m 

31 em 

17 m 

16.6 ·m 

0 

0 

1.2 m 

1.2 m 

2.4 m 

7.6 m 

1 m 

2. 7 m 

30 em 

91 em 

4,6 m 

+1.2 m 

+15 em 

+4.1 m 

+30 em 

+90 em 

+2. 7 m 

+15 em 

1, 5 m below sample 

SD-4-3. 

;:;· O Locality D-5, SE1/4SE1/4 sec. 13, T. 11 S., R, 43 E. 

iii' 

0,1 

= tD 

., 

iii' 

a 

-1 

c 

a: 

D-5 36 em 

Locality D-6, NW1/4SW1/4 sec, 24, T. 11 S., R. 43 E. 

D-6D 30 em +4.25 m 

D-6C 25 em +3.4 m 

D-6B 30 em +35 em 

D-6A 30 em At base 

Brown tuff, punky........................................... --- --- --- --- --- 2 --- 8 

Yellow tuff, medium thick to thick bedded, consisting of --- --- --- --- --- 9 Tr 1 

1-cm-wide pumice fragments. 

Yellow tuff, laminated at top, very resistant............... --- --- --- --- --- 10 Tr Tr 

Brown mudstone, poorly exposed with popcorn weathering ••• ,.. --- --- --- 9 --- --- --- ---

Unit 3--Gray or green welded tuff; weathers pink or brown. From bottom to top: lower 3 m is platy, with lowest 30 em partly 

altered.· Angular charcoal fragments at top: 

Remnants of unaltered glassy tuff, "floating" in altered 

tuff (see SD-4-3G). 

White altered tuff along fractures immediately adjacent to 

SD-4-3GX, 

White, completely altered tuff; varies locally from 

1 to 3 m thick. 

Gray, glassy, platy tuff,,,,.,,, •••••••••••••••••• ,, •••••• 

Pale-gray, very fine grained tuff with blocky fracture,, •• 

Pale-gray, very fine grained tuff with blocky fracture •••• 

Olive-green, glassy tuff, locally breaks with blocky 

fracture. 

10 

5 

6 

10 Tr 

Tr 

5 

4 

Gray tuff with black charcoal fragments,.,,,,,,, •••••••••• 10 Tr 

10 

Tr 

Yellow tuff, partly altered .............................. , 

White to yellow tuff, altered; thin-bedded airfall part 

of welded tuff. 

Yellow-brown mudstone; base of mudstone covered; weathers 

with popcorn texture. 

10 

White tuff in gray sandstone sequence, Bedding attitude N. 60° W., 20° S. About 9 m below is a thin, discontinuous diatomite 

White tuff ........................... ••••••• ,,.,,,,,.,.,,.,, 6 --- ---

Sequence of greenish-gray mudstone and sandstone (lower part) containing three interbedded tuffs, Bedding attidute N. 80° E., 14° N. 

Pale-yellow tuff, completely altered ••••••••••••••• ,,,,,,,,, 

Orange tuff, completely altered,,,,,,, •••••• ,,,,,, •••••••••• 

Grayish-yellow tuff, partly altered •••• ,,.,,,,,., ••••••••••• 

Gray tuff .................................................. . 

10 

2 

3 

2 

8 

Tr 

:· Locality D-7, NW1/4SE1/4 sec. 27, T, 11 s., R, 43 E. Sequence of greenish-gray mudstone with interbedded tuff, Road cut on north side of Manning Road. Bedding attitude N, 20° w., l1° s. 

C: = D-7B 

i D-7A ~ 

61 em 

13 em 

+1.5 m Gray tuff with crystal casts (saline minerals?) on tuff 

bottom. 

Orange tuff, calcite cemented .. , • , • , , , .. , .. , • , ..... , ...... . 

4 4 

~ Locality SD-8, NW1/4SE1/4 sec. 23, T, 11 S., R. 43 E. Bedding attitude N, 80° W., 19° N, Lowest part of this measured section is greenish-gray mudstone, containing interbedded 

~ siltstone, sandstone, conglomerate, and altered tuff, At the top are diatomite, diatomaceous mudstone, and interbedded unaltered tuffs 

tD 

= tD 

I» 

n 

c 

=-

~ 

:r 

tD 

0 

t 

m = 

~ 

0 

c 

t 

0 

~ = 

::: 

..c 

SD-8-11 

SD-8-11D 

SD-8-11C 

SD-8-11B 

SD-8-11A 

SD-8-10 

SD-8-9 

SD-8-9F 

SD-8-9E 

SD-8-9D 

SD-8-9C 

SD-8-9B 

44.5 m 

+5,5 m 

+34,5 m 

+10 m 

At base 

6. 7 m +3m 

36.6 m 

+34. 75 m 

+22,8 m 

+9,5 m 

+90 em 

+60 em 

Unit SD-8-11--Poorly exposed diatomaceous mudstone at the eroded top of the section. 

Unaltered diatomaceous tuff •• ,.,., ••••••••••••• ,, •• ,,,,,,, --- --- 2 3 

Gray, 15-cm-thick, una! tered tuff, one of six in upper 

part of this diatomaceous unit. 

Diatomaceous mudstone •••••••••.••.•••.••••••••••.••.•••••• 

Thin (1-cm-thick) yellow tuff at base of unit SD-8-11A .... 

Diatomite and diatomaceous mudstone, medium thick to thick 

bedded with conchoidal fracture, 

10 

6 

4 

Tr 

3 

Several gray or yellow unaltered tuffs are in unit: 

1 3 --- 1 7A 

4 

2 

2 

1 Gy 

Unit SD-8-9--White to pale-brown-gray diatomaceous mudstone. Medium thick to thick bedded, with conchoidal to platy fracture, Numerous 

thin, unaltered tuffs are interbedded: 

Gray, unaltered tuff, 8 em thick ........................ .. 

Diatomite or diatomaceous mudstone ••••••••••• ,.,.,,,.,,,., 

Pale-greenish-gray indurated mudstone about 30 em thick.,. 

Thin (0.2 em) gray, unaltered tuff ....................... . 

Diatomaceous mudstone •••••••••••••••.••.••••••••.••••••.•• 

10 

5 

3 

2 

4 

1 

2 

2 

4 

1 

2 

Tr 
~ 

N 

= 

~ c 

Q. 

~· 

:r 

0 

t• 

~· 

Sample No. 

(locality 

No. plus 

bed or 

unit No.) 

Measured 

thickness 

shown in m, 

em, or mm 

Locality SD-8--Continued 

SD-8-9A 

SD-8-8 

SD-8-8F 

SD-8-8E 

SD-8-8D 

SD-8-8C 

SD-8-8B 

SD-8-8A 

SD-8-7 

SD-8-6 

SD-8-61 

SD-8-6H 

SD-8-6G 

SD-8-6F 

SD-8-6E 

SD-8-6D 

SD-8-6C 

SD-8-6B 

SD-8-6A 

SD-8-5 

SD-8-4 

SD-8-4B 

SD-8-4A 

SD-8-3 

SD-8-3C 

SD-8-3B 

SD-8-3A 

SD-8-2 

SD-8-2D 

SD-8-2C 

SD-8-2B 

SD-8-2A 

SD-8-1 

SD-8-lC 

SD-8-lB 

SD-8-lA 

22.1 m 

5 m 

14.3 m 

4 m 

3.7 m 

1.7 m 

8.5 m 

10.3 m 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

At base 

+21.3 m 

+9.1 m 

+8.2 m 

+7.6 m 

+5.5 m 

+2.4 m 

+9.5 m 

+8.2 m 

+5.8 m 

+4.9 m 

-1.5 m 

+1.5 m 

+5.5 m 

+5.2 m 

+10.2 m 

+8.2 m 

+4 m 

Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, Oregon--Continued 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

Description of samples 

Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Thin (2.5 em) gray unaltered tuff......................... 10 --- --- --- --- --- --- --- --- ---

Unit SD-8-8--Pale-brown-gray mudstone with conchoidal to platy fracture; locally contains abundant ostracods and rare clams. 

Mudstone is medium thick to thick bedded: 

Mudstone •••••••••••••••••••••••••••••••••••••••••••••••••• 

Brown sandstone, fine grained ••••••••••••••••••••••••••••• 

White, upper tuffaceous mudstone with 2- to S-cm-thick 

sandstones; ostracods are in upper parts. 

Brown, lower tuffaceous mudstone, part of SD-8-8D ••••••••• 

Mudstone wit;h ostracods and clams ••••••••••••••••••••••••• 

Thin ( 10-cm-thick) white tuff ••••••••••••••••••••••••••••• 

10 

10 

Brown sandstone, with thin mudstone and conglomerate beds... --- --- Tr 

3 

5 

Unit SD-8-6--Gray to greenish-gray mudstone containing thin siltstone and sandstone beds: 

Pale-orange tuff forms top of unit SD-8-6. This sample --- --- --- Tr 

is at the top of the 35-cm-thick tuff. 

Sample from middle of tuff (see SD-8-61) •••••••••••••••••• 

Sample from base of tuff (see SD-8-61) •••••••••••••••••••• 

White tuff, 1 em thick •••••••••••••••••••••••••••••••••••• 

Uniformly pale-orange tuff, 20 em thick ••••••••••••••••••• 

Gray upper part of 30-cm-thick tuff ••••••••••••••••••••••• 

Yellow lower part of tuff, 5 em thick ••••••••••••••••••••• 

White, very thin (2.5-cm-thick) tuff ••••• ••• •••••••••••••• 

Gray to greenish-gray mudstone •••••••••••••••••••••••••••• 

Olive-green to brown sandstone, medium to coarse grained, 

very poorly consolidated. 

Tr Tr 

1 

Tr 

Tr 

5 

4 

3 

2 

3 

10 

6 

6 

3 

3 

3 

3 

Tr 

Tr 

Unit SD-8-4--Yellow-gray mudstone that has conchoidal fracture when 

Pale-gray tuff, 25 em thick, at top of unit SD-8-4 •••••••• 

fresh. Some thin tuffs in upper part: 

Mudstone •••••••••••••••••••••••••••••••••••••••••••••••••• 

Unit SD-8-3--0range tuff, platy and completely altered: 

Top of tuff ••••••••••••••••••••••••••••••••••••••••••••••• 

Middle of tuff •••••••••••••••••••••••••••••••••••••••••••• 

Base of tuff •••••••••••••••••••••••••••••••••••••••••••••• 

Tr Tr --- --- 3 2 

Tr 

Tr Tr 

1 Tr 

Unit SD-8-2--Dlive-green fissile mudstone; contains several thin, altered tuffs: 

Thin tuff or tuffaceous mudstone in the upper 1 m of --- --- 1 1 

this unit. 

Uppermost of two thin (2.5-cm-thick) orange tuffs ••••••••• 

Lowermost of two tuffs (see SD-8-2C) •••••••••••••••••••••• 

Olive-green mudstone •••••••••••••••••••••••••••••••••••••• Tr 

4 

4 Tr 

3 

6 

Unit SD-8-1--Greenish-gray, fissile mudstone that breaks into small chips. Base of unit is not exposed. 

in the upper part of the unit: 

Orange, sugary tuff near top of unit •••••••••••••••••••••• 

Orange tuff, 10.2 em thick •••••••••••••••••••••••••••••••• 

Mudstone •••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 9 

Tr 10 

1 2 

3 

5 3 

5 4 

2 

5 

Tr 

Tr Tr 

3 1 

3 

Tr 1 

- 1 

5 2 

5 4 

4 

1 

Two thin tuffs are 

Tr 

Tr Am 

3Am 

1 Am 

1 Gy 

Gy 

Tr Ja 

2 Gy 

2 7A 

Locality D-9, SW1/4SW1/4 sec. 14, T. 11 S., R. 43 E. Exposure is a 2.1-m-thick gray tuff stratigraphically above SD-8. Bedding attitude N. 75° E., 5° N. The tuff is fine to medium 

bedded, and cross-bedded. About 9 m of greenish-gray mudstone are above D-9. Below D-9 are about 45.7 m of pale-gray mudstone, sandstone, and diatomaceous mudstone 

D-9E 

D-9D 

D-9C 

D-9B 

D-9A 

+3m 

+1m 

+46 em 

At ba.se 

Brown lignite ••••••••••••••••••••••••••••••••••••••••••••••• 

Gray fresh tuff ••••••••••••••••••••••••••••••••••••••••••••• 

Tuff, partly altered •••••••••••••••••••••••••••••••••••••••• 

White altered tuff •• ••••• •••••• ••• •••• ••• ••• ••• ••••••••••••• 

Unconsolidated sandstone on bottom of tuff base ••••••••••••• 

9 

7 

8 

1 

2 

2 

10 

Tr Tr 

Tr 

2Am 
N 

2. 

:=: 

;:;· 

Locality D-10, SW1/4NE1/4 sec. 24, T. 11 S., R. 42 E. D-10 is at same stratigraphic level as lower part of measured section at locality SD-4 

D-10 Pumice conglomerate (fragments as long as 2.5 em), 18.3 m 

above the welded tuff. 

Locality D-11, SE1/4SE1/4 sec. 36, T. 11 s., R. 42 E. Thin tuff is present in a brownish-gray mudstone and siltstone sequence 

D-liB 

D-llA 

5 em 

25 em 

Black tuffaceous material at top of tuff •••••••••••••••••••• 

Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-12, NW1/4SW1/4 sec. 36, T. 11 S., R. 42 E. Ridge-forming tufa deposit 

8 

9 

8 -- -- -- -- -- -- -- -- Tr 

D-12 Pale-brown tufa that weathers white. Bedding is medium 

thick to thick bedded. 

--- --- --- --- --- --- --- --- --- --- --- --- --- 1 0 

Locality D-13, SE1/4SW1/4 sec. 1, T. 11 s., R. 42 E. Road-cut exposure of the western part of the Durkee lacustrine beds. Bedding attitude N. 10° E., 22° E. 

D-13C 

D-13B 

D-13A 

Gray tuff composed of coarse pumice ••••••••••••••••••••••••• 

Light-gray, fine-grained tuff ••••••••••••••••••••••••••••••• 

Brown tuffaceous mudstone ••••••••••••••••••••••••••••••••••• 

2 

3 

Locality D-14, SW1/4NW1/4 sec. 1, T. 11 s., R. 42 E. Very poor exposure of pumice lapilli tuff, similar to that at locality D-13 

D-14 Pumice tuff in a waxy, clayey matrix •••••••••••••••••••••••• 

10 

7 

7 

Tr 

1 

Tr 

4· --- --- --- --- --- --- --- --- Tr ---

0 Locality D-15, NE1/4NE1/4 sec. 1, T. 11 S., R. 42 E. Very poorly exposed lacustrine sequence of diatomite, green-gray sandstone, and yellow-gray siltstone. Bedding attitude N. 70° W., 20° N. 

~-

:II 

D-15B 

D-15A 

Sandstone ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Diatomite ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr Tr 

7 --- --- --- --- --- (I 

~· 

a Locality D-16, SW1/4SE1/4 sec. 17, T. 11 s., R. 43 E. Two thick tuffs, probably similar to tuffs at D-9. Good exposure in freeway road cut. Bedding attitude N. 40° W., 18° N. 

-t 

c 

~ 

:r 

D-16B 

D-16A 

Upper brown-gray tuff rests conformably on lower tuff; 

locally cross-bedded. 

Gray-brown lower tuff, collected near middle of bed ••••••••• 

10 

10 

AI 

:II Locality D-17, NW1/4NE1/4 sec. 20, T. 11 S., R. 43 E. Excellent quarry exposure. Bedding attitude N. 35° w., 16° N. 

c: 

I D-17F 

D-17E 

~ D-17D 

c$" D-17C 

f'l D-17B 

~ D-17A I'D 

30.5 em 

Gray mudstone above D-17E ••••••••••••••••••••••••••••••••••• 

Gray tuff, 9.1 m above D-17B; tuff is composed of fresh 

pumice shards. 

Gray mudstone below tuff D-17E •••••••••••••••••••••••• •••• •• 

Gray mudstone above tuff D-17B •••••••••••••••••••••••••••••• 

Altered bed. ••• ••••• ••••••••••• •••. •• ••••• ••• ••••••••••• •••. 

Gray mudstone below D-17B ••••••••••••••••••••••••••••••••••• 

5 

10 

Section is lower than the tuffs at D-16 

Tr 

Tr 

1 

5 

2 

Tr 

Tr 

3 

4 

1 

5 

1 

3 

8 

4 

~ Locality D-18, NE1/4SW1/4 sec. 10, T. 11 S., R. 43 E. Sequence of pale-gray to yellow-gray diatomaceous beds. A 2.5-cm-thick yellow bed beneath gray siltstones is at top of hill 

n 

Diatomite, 4.6 m above D-18B ••••••••••••••••••••• ••• •••••••• 

Gray siltstone ••••••••••• ••• •••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

2 

2 

2 

4 

3 

7 

i D-ISC 

.. D-18B 

~· D-18A 

0--

Locality D-19, SE1/4NW1/4 sec. 21, T. 11 S., R. 43 E. t Sequence of pale-gray and yellow diatomaceous sediments similar to those at locality D-18 

:II 

D-19B 

D-19A 

Diatomaceous mudstone above sample D-19A •••••••••••••••••••• 

2.5-cm-thick white bed in lower part of slope ••••••••••••••• 

8 

10 

Tr Tr 

2 

5 

Tr 

1 Am 

Tr 7A 

1 Am 

m 

.. Locality D-20, SW1/4NE1/4 sec. 36, T. 11 S., R. 23 E. Gray and green mudstone and greenish-brown sandstone resting on basalt. Bedding attitude N. 70° w., 35° s. Several interbedded tuffs 

0 

C D-20D f D-20C 

0 

~ 

~ 

N 

... 

91 em Pale-yellow platy tuff •••••••••••••••••••••••••••••••••••••• 

Greenish-brown tuffaceous sandstone; makes up most of 

resistant exposure. 

Tr 

2 

4 

TrAm 
w 

t.,J 

t.,J 

Ill 

e-

~ 

;;· 

"' :r 

0 

£j' 

Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, Oregon--Continued 

~ Sample No. 

= (locality 

Measured 

thickness 

shown in m, 

em, or mm 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

~ No. plus 

iii' bed or 

Description of samples 

unit No.) Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Locality D-20--Continued 

D-20B 

D-20A 

2.5 em 

0.6 em 

Pale-yellow tuff •••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-21, NE1/4NW1/4 sec. 6, T. 12 s., R. 44 E. Poor exposure. Greenish-gray mudstone contains two thin orange tuffs 

D-21B 

D-21A 

7.6 em 

3.8 em 

Orange tuff collected 22.9 m above base of hill. •••••••••••• 

Thin orange tuff collected 13.8 m above base of hill 

(9.1 m below sample D-21B). 

Tr 

8 

5 

9 

10 Tr 

Locality D-22, SE1/4SW1/4 sec. 32, T. 11 S., R. 44 E. Thick-bedded, poorly sorted sandstone, locally caliche cemented. D-82 was recollected at locality D-22 

D-22 Yellow-brown sandstone •••••.•.••••..••.....••••••.••••••.•.• 

Locality D-23, NE1/4SE1/4 sec. 7, T. 12 s., R. 44 E. Poorly exposed, green mudstone containing an interbedded tuff 

D-23B 

D-23A 

15 em Upper part of platy orange tuff ••••••••••••••••••••••••••••• 

Lower blocky part of tuff D-23B •••••••••••••••••••••••••••• 

Locality D-24, SW1/4SW1/4 sec. 8, T. 12 s., R. 44 E. Tuff, lower in section than D-23. Bedding attitude N. 85° E., 28° N. 

D-24C 

D-24B 

D-24A 

45 em Upper, altered yellow tuff •••••••••••• •••••• •••••••••••••••• 

Middle, unaltered gray tuff •••••••••••••••••••••••••• ••••••• 

Lower, altered yellow tuff •••••••••••••••••••••••••••••••••• 

Tr Tr 

1 

10 

10 

10 

3 ---

6 

9 

Locality D-25, SW1/4SE1/4 sec. 8, T. 12 s., R. 44 E. Gray, partly altered tuff low in lacustrine section, about 7-8 m above welded tuff, in a brown mudstone 

D-25B 

D-25A 

Gray tuff ••••••••••••••••••••••••••• • • • • • ••••••••• • • • • • • • • • • 

Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-26, NW1/ 4NE 1/4 sec. 7, sec. 7, T. 12 S., R. 44 E. Thin tuff in a gray or slightly diatomaceous mudstone sequence 

D-26B-2 

D-26B-1 

D-26A 

5 em Brown base of yellow tuff ••••••••••••••••••••••••••••••••••• 

Light-yellow-brown tuff ••••••••••••••••••••••••••••••••••••• 

Grayish mudstone •••••••••••••••••••••••••••••••••••••••••••• 

Locality D-27, SW1/4SE1/4 sec. 6, T. 12 S., R. 44 E. Thick, pale-gray diatomaceous mudstone 

8 

7 --- ---

10 

7 

5 

3 --- ---

Tr 

3 ---

D-27 Diatomaceous mudstone ••••••••••••••••••••••••••••••••••••••• 5 --- --- --- --- --- --- ---

2 --- ---

2 

2 Gy 

TrAm 

TrAm 

Locality D-28, SE1/4NW1/4 sec. 6, T. 12 S., R. 44 E. Poorly exposed green mudstone containing thin, yellow to orange interbedded tuffs. This exposure is stratigraphically lower than D-27 

D-28E 

D-28D 

D-28C 

D-28B 

D-28A 

0.5 em 

7 em 

2.5 em 

15 em 

10 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

White tuff •••••••••••••••••••••••••••••••••••••••••••••••••• 

Pale-gray tuff •••••••••••••••••••••••••••••••••••••••••••••• 

Orange tuff ••••••••••••••••••••• , ••••••••••••••••••••••••••• 

Tr 

10 

10 

8 

10 

10 

Tr 

Locality D-29, SW1/4SW1/4 sec. 36, T. 11 s., R. 43 E. Opal-mine quarry and type locality of erionite. Quarry is in welded tuff and is immediately overlain by a thick sequence of 

tuffs interbedded in mudstone, sandstone, siltstone, and tuffaceous mudstone. Quarry entrance is covered by slumped rubble 

D-29E 45 em Yellow tuff, upper part; separated by 15m of slope 

cover above bed D-29C. 

--- -- --- --- -- --- 10 --- --- -- --- --- --- ---
D-29D 

D-29C 

D-29B 

D-29A 

D-29 

Lower part of yellow tuff in bed D-29E •••••••••••••••••••••• 

Green-gray mudstone and siltstone with a thin fissile 

tuff (tuff sampled). 

Brown and green-brown sandstone, conglomerate, and 

tuffaceous sandstone. 

Yellow lapilli tuff just above welded tuff at quarry 

entrance. 

Miscellaneous collection of specimens of type erioni te 

from quarry rubble and waste dump. 

4 

Tr 

6 

8 

4 

10 

Locality D-30, NE1/4NE1/4 sec. 26, T. 11 S., R. 43 E. Fault-detached sequence of brown mudstone and siltstone with interbedded white to pale-gray or brown tuffs. Most tuffs are coated 

with a white, powdery efflorescenc.e. This sequence appears to be higher in section than that at locality SD-4 and may be equivalent to lower part of sequence at locality SD-8 

;:;· 

~ c 

~· 

I'D = 

Ill 

;;;· 

a 

D-30L 

D-30K 

D-30J 

D-301 

D-30H 

D-30G 

D-30F 

D-30E-1 

D-30E-2 

D-30E-3 

D-300 

D-30C-2 

D-30C-1 

D-30B 

D-30A 

1 em 

75 em 

1.5 m 

1. 75 m 

4 em above bed D-30K Greenish-gray tuffaceous siltstone •••••••••••••••••••••••••• 

7.6 m above bed D-30J Gray tuffaceous siltstone ••••••••••••••••••••••••••••••••••• 

Top of white tuff ••••••••••••••••••••••••••••••••••••••••••• 

12.19 em above bed D-30H Base of white tuff. ••••. •••••• ••• •••• •• ••• ••• ••••• •••• •••••. 

4.57 m above bed D-30E 

1.22 m above bed base 

46 em above base 

Top of white tuff with multiple bedding ••• ••• ••• •••••••••••• 

Middle of white tuff •••••••••••••••••••••••••••••••••••••••• 

Base of white'tuff•••••••••••••••••••••••••••••••••••••••••• 

Pink tuffaceous material •••••••••••••••••••••••••••••••••••• 

Green tuff •••••••••••••••••••••••••••••••••••••••••••••••••• 

Waxy, dark-green tuff ••••••••.•.•••••••••••••••••••••••••••• 

White efflorescent material collected about 9 m north of 

samples collected at D-30A and 30B. 

White efflorescence at D-30B •• •••• ••• •• •••• ••••• ••• ••• •••••• 

Green material in D-30C-2 ••••••••••••••••••••••••••••••••••• 

Lowest tuff in sequence at locality D-30 •••••••••••••••••••• 

Base of lowest tuff bed D-30B ••••••••••••••••••••••••••••••• 

Tr 

Tr 

Tr 

1 

Tr Tr 

Tr 

Tr 

1 

5 

Tr 

Locality D-31, SE1/4SE1/4 sec. 23, T. 11 s., R. 43 E. Thick unit of yellow tuffs equivalent to tuffs at SD-4-13, -14, -15,, and -16 

1 Tr 

2 2 

8 2 

6 2 1 

2 3 3 

6 4 

5 --- 8 2 Tr 

7 2 Tr 

1 3 1 

4 3 3 

1 3 5 

2 4 4 

3 3 3 

Tr 5 3 2 

c;t D-31 Yellow tuff................................................. --- -- --- Tr --- ~ --- --- --- --- --- --- ---

-· 

~ Locality SD-32, SE1/4SE1/4 sec. 23, T. 11 s., R. 43 E. Measured section collected upward from the stratigraphic equivalent position of unit SD-4-12 at locality SD-4 

= Dol 

= c: 

1 .. 

~ 

~-

I'D 

I'D = 

!; 

n 

c 

~ .. 

:r 

I'D 

SD-32-8 

SD-32-7 

SD-32-6 

SD-32-5 

SD-32-4 

SD-32-3 

SD-32-2C 

SD-32-2B 

SD-32-2A 

SD-32-1 

5.5 m 

6.6 m +90 em 

6 m 

2.6 m At base 

1.8 m At base 

60 em At base 

10 em At top 

+4.57 m 

At base 

Gray to yellow, micaceous sandstone, thin bedded to thick 

bedded; fine grained, top eroded. 

Pale-gray tuffaceous sandstone, fine grained, thin bedded 

Poorly exposed, yellow mudstone (no sample collected) ••••••• 

Yellowish-brown sandstone, poorly cemented, fine to 

coarse grained. 

Yellowish-brown fissile mudstone •••••••••••••••••••••••••••• 

Brown sandstone, tuffaceous, angular, poorly sorted, 

medium to coarse grained. 

Top of yellow tuff, completely altered; same as units 

SD-4-13 to -16 at locality SD-4. 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Brown tuffaceous mudstone, same as S4-4-12 •••••••••••••••••• 

2 

3 

5 

5 

4 

1 

2 1 --- --- 4 2 --- ---

2 2 

-- 5 

4 

- --- 4 4 

--- 1 1 

--- --- ---

c 

t 

Locality D-33, NW1/4NW1/4 sec. 25, T. 11 S., R. 43 E. Poorly exposed welded tuff about 5 m above a basalt outcrop. A separate tuff, similar to unit SD-4-2 at locality SD-4, is at 

base of the welded tuff 

D-33-1 

D-33-2 

Yellow tuff •••••••••••••••••••••••••••••••••••••••••••••••• 

White patches in yellow tuff •••••••••••••••••••••••••••••••• 

m 

.. Locality D-34, NE1/4SW1/4 sec. 25, T. 11 s., R. 4-3 E. Welded tuff. Poorly exposed lacustrine sediments below 

c 

c 

~ D-34B 

~ D-34A 

l 

~ 

N 

~ 

10 em -6.10 m 

Yellow altered base of welded tuff similar to D-33 •••••••••• 

Pale-yellow tuff in brown mudstone ••••••••••••••••••••••••• 

Tr 

1 

Tr 10 

9 

6 

8 

Tr 

Tr 

3 

---

---

---

---

---

2Am 

Tr Gy 

TrGy 

Tr Gy 

Tr Gy 

Gy 

1 Gy 

TrAm 

TrAm 

---

TrAm 

TrAm 

TrAm 
CN 

N 

ollo 

~ 

a. 

;· 

Ill 

:r 

0 

~· 

! 

;;· 

Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, Oregon--Continued 

Sample No. 

(locality 

No. plus 

bed or 

unit No.) 

Measured 

thickness 

shown in m, 

em, or mm 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

Description of samples 

Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Locality D-35, NW1/4SE1/4 sec. 23, T. 11 S., R. 43 E. Sequence of thin tuffs similar to SD-8-6C and -60 

D-350 Gray upper part of tuff ••••••••••••••••••••••••••••••••• , ••• --- --- 1 --- --- --- 9 Tr 

D-35C-3 0.5 em Upper brown layer of brownish tuffaceous layer between --- --- 1 --- 3 --- 3 3 

gray and yellow layers, 

D-35C-2 2mm Gray crust .••••••••••.•••••••••..•••••.••.•.••. ••• •••••••.•. --- --- 2 1 2 --- 2 3 

D-35C-l Total rock specimen ••••••••• , •••••••••••••••• ,., •••••••••••• --- --- Tr --- 4 --- 3 3 

D-35B-4 1 mm Yellow lower part of tuff sequence D-35; D-35B4 is gray --- --- 1 3 3 --- 1 2 

top layer of lower yellow part. 

D-35B-3 5 mm Dark-brown layer, •• ,., ••••••• , •••••• ,., •••••••••••• ,,, •••••• --- --- --- 6 2 --- --- 2 

D-35B-2 Middle layer ••••• ,, •••••••••••••••• ,,,,,,, ••••••••• ,,,,,,, •• --- --- --- Tr 6 --- --- 4 

D-35B-l 2mm Base of D-35B tuff sequence, •• ,,,., ••••••• ,,,,,,,,, ••••••••• --- --- 1 3 4 --- --- 2 

D-35A Grayish-green .mudstone below tuff •••••• ,,,, ••• ,,,, •••••• ,,,, --- --- 1 2 --- --- --- 5 

Locality D-36, NW1/4SE1/4 sec. 23, T, 11 S., R, 43 E. Thick tuff sequence of gray and orange beds with subconchoidal, platy fracture, Same unit as unit SD-8-3 at locality SD-8, but 

poorly exposed 

D-36E 

D-360 

D-36C-2 

D-36C-l 

D-36B-2 

D-36B-l 

D-36A 

2,4 m 

5 mm 

1 mm 

+1.98 m 

+1,07 m 

+35.6 em 

Top of tuff sequence D-36.,,,, •••••••• ,, •••• ,,,, ••••••• ,,,,. 

Tuff,,, •••••• ,,,,,,,,,, •••••••• ,,,,,,,, ••• ,,, •• ,,,,,,,,,,,., 

Thin yellow tuff .......... ,,, ... , ........... , ...... , ...... .. 

Gray tuff .................................................. . 

Thin brown tuff •• ,,,, •••••••• ,,,,., , , , •••••••••• , , , • , , , , , ••• 

Lower gray layer •.••.••.•••••••••••••••••.••••••••••.••.••.• 

Brown silty mudstone, just below tuff sequence D-36 ......... 

Tr 

Locality D-37, SW1/4SE1/4 sec. 23, T, 11 S,, R, 43 E. Same tuff as unit SD-8-3. A basalt on the southwest side of this gully drops the tuff below the drainage 

D-37C 

D-37B 

D-37A 

1.5 m Top ••••••••••••••••••••••••••••••••••••••• ••••••••••••••••• Middle •••••••••••••••••••••••••••••••••••••••••• • • • • • • • •. • • 

Base ••• ••••••••••••••• • •• • •• ••• •• • •• • • • •• • •• • • • • • • •• •• •• •• • 

Locality D-38, SW1/4SE1/4 sec. 23, T. 11 S., R. 43 E. Same tuff as unit SD-8-3. The sequence below the tuff is greenish-gray mudstone and sandstone 

D-38C-2 

D-38C-l 

D-38B 

D-38A-2 

D-38A-l 

1.5 m Yellow layer at top of tuff •••• , ••••••••• , •••••••••••• ,., ••• 

Gray layer below yellow layer at D-38C-2 ••••• ,., •••••• , ••••• 

Middle layer •••••••••••••• , •• , •• ,., ••••••••• , ••• , ••••••••••• 

Yellow mottled layer ••••••••••••••••••••••••••••••••••••• ,., 

Base of tuff, gray, manganese stained •••••••••••••••• , •• ,, •• 

Locality D-39, SW1/4SE1/4 sec. 23, T. 11 s., R. 43 E, Same tuff unit as SD-8-3 

D-39C 

D-39B 

D-39A 

90 em Top ••••••••••••••••••••••••••••••••••••••• •••••••••••••••••• 

Middle., •• ,., •• , •••••••••••• , •••• , ••••••••• , •••••••••••• ,, •• 

Base •••••••••••• • •• •• • • •• • • • • •• • • ••• •• • • • • • • •• • •• • •• • • •• •• •• 

Tr 

Tr 

1 

Tr 

Tr 

Tr 

5 

5 

5 

6 

4 

4 

4 

9 

9 

9 

5 

5 

3 

2 

1 

5 

5 

4 

4 

5 

4 

5 

5 

5 

7 

8 

8 

9 1 Tr 

9 Tr 1 

9 Tr 1 

Locality D-40, NW1/4SE1/4 sec. 23, T, 11 s., R. 43 E. Gray, glassy, unaltered tuff in a diatomite sequence equivalent to the upper part of section measured at locality SD-8 

D-40B 

D-40A-4 

D-40A-3 

D-40A-2 

D-40A-l 

30 em 

2 mm 

0.5 em 

3 em 

Gray tuff •• , •••••• , ••••••••••• , •••••••••••••• ,., •••••• ,,, ••• 

Gray, glassy tuffaceous material in diatomite sequence •••••• 

Dark brown layer •••••••••••••••••••••••••••••••••••••••••••• 

Manganese-stained, yellow layer •••••••••••• , •••••••••• , ••• ,, 

Pale-yellow-gray layer •••••••••••••••••••••••••••••••••• , ••• 

10 

10 

9 

5 

8 

1 

2 

Tr 

Tr 

T:t 

Tr 

Tr 

1 

1 

Tr Gy 

Tr Gy 

Tr Gy 

Tr Gy 

TrAm 
;:;· 

~ 0 

t• 

II> iii. 

a 

ct 

a: 

:r 

AI 

:I 

c: 

~ ~ 

~· 

~ 

:I 

~ 

~ c 

l4 

:::!. 

:I 

~ 

i 

::::;: 

:I 

.. 

~ 

0 

c 

t 

0 

~ :I 

w 

N 

Ill 

Locality D-41, SE1/4SW1/4 sec. 23, T. 11 S., R. 43 E. Sequence of greenish-gray and red mudstones with two interbedded tuffs equivalent to bed sampled at SD-8-6E and SD-8-6G to -61 

D-41F 30 em 

D-41E 61 c;m Top 

D-41D +36 em 

D-41C +23 em 

D-41B +15 em 

D-41A-2 5 em Base 

Upper tuff, poorly exposed, probably equivalent to bed 

sampled at SD-8-61. 

Lower tuff, multi bedded, orange ••••••••••••••••••••••••••••• 

Tuff ••••••••• •••. • • ••• •• ••• • •• • • •• •• •• •• • • • ••• • •• • • •• •• •• ••• Tuff •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr Tr 

Tr 

10 

10 

-- 8 

--- --- 2 8 

--- --- --- 10 --- --- --- --- Tr 

D-41A-1 6mm Base bottom 

Tuff •••••••••••••••••••••••••••••• •• • • • • • • • • • •• • • •• •• •• • •• • • 

Tuff ••••••• • • •• •• • • • • • ••• • •••• • • • • • • •• • • • •• • • • • •• • • • •• • • • • • Tuff, hornblende rich ••••••••••••••••••••••••••••••••••••••• 9 --- --- --- --- Tr ---

Locality D-42, SW1/4SW1/' sec. 23, T. 11 S., R. 43 E. Two orange tuffs, same as at locality D-41 

D-42B 

D-42A 

10 em 

10 em 

Upper tuff •••••••••••••••••••••••••••••••••••••••••••••••••• 

Lower tuff, silicified •••••••••••••••••••••••••••••••••••••• 

Tr 10 

4 6 

Locality D-43, SW1/4SW1/4 sec. 23, T. 11 s., R. 43 E. Thick yellow to orange tuff, equivalent to bed sampled at SD-8-3. Bedding attitude N. 10° E., 10° w. Tuff overlies a sequence 

of greenish-gray sandstone and siltstone 

D-43C 

D-43B 

D-43A 

1.2 m Top 

Middle 

Base 

Tuff ••••••••••••••• ••• ••. •• • •• • • • • •• •• • • • •• • • •• • • •• •• • • •• • • 

Tuff ••••••••••••••••••••••••••• •••• •• ••• •• • • •• • • •• • • • • • • • • •• 

Tuff •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-44, NE1/4NE1/4 sec. 27, T. 11 S., R. 43 E. White multibedded tuff, stratigraphically below locality D-43 

D-44B 

D-44A 

30 em Top 

Base 

Tuff ••••••••••••••••••• ••• •••••••••••••••••••••••••••••••••• 

Tuff ••••••••••••••• ••• ••••••••• ••• •••••••••••••••••••••••••• 

Tr 

8 

10 

10 

Tr 2 

Tr Tr 

8 

8 

Locality D-45, NE1/4NW1/4 sec. 26, T. 11 s., R. 43 E. Welded tuff, upper part altered and grades downward in typical gray, unaltered welded tuff. Bedding attitude N. 15° W., 33° W. 

D-45B 

D-45A 3m 

Pale-yellow bedded tuff in brown mudstone; an unknown 

distance above bed D-45A. 

Upper yellow altered part of welded tuff •••••••••••••••••••• 

..;. __ 

Locality D-46, NE1/4NE1/4 sec. 27, T. 11 S., R. 43 E. Tuff, equivalent to tuff at locality D-43, in a sequence of green sandstone and siltstone. 

Many small faults at locality D-46 

D-46C 

D-46B 

D-46A-2 

D-46A-1 

1.1 m Top 

Middle 

Base 

Base bottom 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

White tuff •••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 

Tr 

Tr 

8 

10 

9 

Bedding attitude N. 40° w., 18° SW. 

5 

1 

Tr 

1 Tr 

Locality D-47, NE1/4NE1/4 sec. 27, T. 11 S., R. 43 E. Tuff with blocky to subconchoidal fracture. Stratigraphic relation to other localities is unknown 

D-47B 

D-47A 

36 em Top 

Base 

White to pale-yellow tuff ••••••••••••••••••••••••••••••••••• 

Tuff •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-48, NE1/4NE1/4 sec. 27, T. 11 s., R. 43 E. Several tuffs in a green mudstone sequence 

36 em Upper beds in unit Yellow tuff, highest in sequence •••••••••••••••••••••••••••• 

Tuff •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 

Tr 

3 

Tr --- --- 5 --- Tr 

10 

--- --- --- 10 

--- --- --- 10 

--- --- --- 6 

1 Tr --- 9 

--- --- --- 5 

D-48F 

D-48E 

D-48D 

D-48C 

D-48B 

D-48A 

8 em 

7 em 

15 em 

Base 5 m above bed D-48D 

4.5 m above bed D-48C 

1.2 m above bed D-48B 

Upper part 

Orange tuff with silicified segregations •••••••••••••••••••• 

White tuff with local, dark silicified segregations ••••••••• 

Tuff bed lowest in sequence ••••••••••••••••••••••••••••••••• Tr --- --- --- 10 --- --- --- --- --- Tr 

Base Bottom of lowest tuff (bed D-48B) immediately above 

underlying green silicenous mudstone. 

Locality D-49, SE1/4SE1/4 sec. 22, T. 11 S., R.-43 E. Several tuffs in a greenish-gray mudstone 

D-49K 

D-49J 

46 em 

90 em 

1.5 m above bed D-49J 

0.6 m above bed D-491 

Gray tuff highest in top part of sequence •••••••••••••••••• 

Brown tuffaceous sandstone •••••••••••••••••••••••••••••••••• 

Tr --- Tr --- 6 

3 

5 

--- --- --- --- --- 1 

--- --- --- 3 

--- --- --- 2 

1 Am 

Tr Gy 

TrAm 
~ 

t.) 

="' 

(I) 

[ 

;;;· 

"' ;· 

0 

Ei' 

~ ::II 

rJ 

iii' 

Appendix 1. Mineralogy of lacustrine rooks from the Durkee basin, Baker County, Oregon--Continued 

Sample No. 

(locality 

No. plus 

bed or 

unit No.) 

Measured 

thickness 

shown 'in m, 

em, or mm 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

Description of samples 

Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Locality D-49--Continued 

D-49! 36 em 

D-49H 

D-49G-l 25 em 

D-49G-2 

D-49F 36 em 

D-49E 

D-49D 

D-49C 60 em 

D-49B 

D-49A 

Upper part of· bed D-49H 

2. 7 m above bed D-49G 

5.2 m above bed D-49F 

Base 

5.5 m above bed D-49C 

Top 

Middle 

Base 

White tuff •• •••• •••••••••••••••••••••••••••• ••• ••••••••••••• 

Tuff base ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••• • • • • • • • • •• 

White bottom at base of yellow tuff (bed D-49G) ••••••••••••• 

Yellow massive tuff at top of middle unit in sequence at 

locality D-49. 

Platy material below bed D-495 ••••••• , ••••••••••••••••••••• 

Massive tuff at base of middle tuff unit •••••••••••••••••••• 

Lowest tuff u,nit in locality D-49 sequence, equivalent to 

bed sampled at SD-8-3. 

Tuff •••• ••• ••• ••• ••• ••• •• ••• •• •• • •• •• • •• •• ••• •• •• •••• • •••••• Tu!f •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 

Locality D-50, NEl/4tjEl/4 sec. 27, T. 11 S., R. 43 E. Pale-greenish-gray tuff with conchoidal fracture. Stratigraphic position unknown 

D-50 1.52 m Tuff ••••••••••••••• ••• ••• ••• •••••••• ••• ••••• ••• ••••• •• •••• •• 

Locality D-51, NE1/4NE1/4 sec. 17, T. 11 s., R. 43 E. Tuff probably equivalent to bed sampled at SD-8-3 

D-51C 

D-51B 

D-51A 

91 em Top 

+36 em 

Base 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Tuff •••••• ••• ••• •••••••••••••••••••••• ••• •••• •••• ••••••••••• 

Tuff ••••• ••••••. ••• •••• • •• •• ••• ••• •• •• •• ••• ••• •• • • ••••• • •• • • 

Tr ---

Tr 

6 

Tr 10 

10 

2 

10 

10 

9 

Tr 

10 Tr 

-- 6 

10 

Locality D-52, SE1/4SE1/4 sec. 23, T. 11 S., R. 43 E. Tuff stratigraphically equivalent to SD-8-3. Tuff rests on greenish-gray sandstone and siltstone 

D-52C 

D-52B 

D-52A 

1.38 m Near top 

+81 em 

At base 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• Tr 

Tr 

1 

Locality D-53, SW1/4NE1/4 sec. 27, T. 11 S., R. 43 E. Two tuffs in a green mudstone and sandstone sequence. The tuffs are about 9 m apart 

D-53B 15 em 

D-53A-3 15 em 2. 5 em above bed 

D-53A-2. 

D-53A-2 Be tween bed D-53A-l and 

53A-3. 

D-53A-l 0.1 em At base 

0 

Upper orange tuff, 10-A clinoptilolite peak is unusually 

intense. 

Lower orange tuff, large crystal molds on the tuff base as 

seen at bed at locality D-7A. 

Tuff •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 

9 

10 

9 Tr Tr 

10 

Tr 

--- --- 8 

--- Tr Tr 

5 4 

--- 5 3 

5 5 

---

---

4 ---

--- --- ---

Locality D-54, SE1/4NE1/4 sec. 27, T. 11 S., R. 45 E. Several tuffs in a green mudstone and siltstone sequence. Lowest tuff is same as bed at locality D-7A or D-53A 

D-54K 

D-54J 

D-54! 

D-54H 

D-54G 

D-54F 

D-54E 

D-54D 

D-54C 

D-54B 

D-54A 

28 em 

28 em 

15 em 

At top 

Middle 

At base 

At top 

Middle 

At base 

Upper part 

At base 

1.? m above bed D-54B 

Upper part 

Lower part 

White tuff; uppermost tuff in the D-54 sequence ••••••••••••• 

Tuff •••••• •••••• •••••• ••••• •••• ••• ••• • ••••••. •• • ••• ••• •••• •• Tuff, 2.44 m above D-54H ................................... . 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Tuff •••••••••••••••••• ••••• •• •••••• •• ••• ••• ••• ••••• •• •• ••• •• 

Tuff, 5.79 m above D-54 E ................................. .. 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Tuff, 6.10 m above D-54C. This tuff is the same as D-53B ... 

Gray tuff, poorly exposed ••••••••••••••••••••••••••••••••••• 

Lowest tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Lowest tuff, with crystal molds on bottom •••••••••••• , •• , ••• 

Tr 

Tr 

Tr 

Tr 

10 

6 

9 

10 

10 

10 

10 

3 

6 

10 

9 

--- --- --- --- --- ---

Tr 2 2 1 Am 

1 --- --- 1 Am 

10 

4 

--- --- ---
N 

;:;· 

~ 

0 

~· 

; 

Ill 

iii" 

a 

ooot 

c 

i 

:;· 

Ill 

:I 

c: 

i .. 

~ 

~· 

~ 

~ 

~ ~· 

f 

Locality D-55, SE1/4SE1/4 sec. 27, T. 11 s., R. 43 E. Greenish-gray mudstone sequence, containing a yellow tuff midway up the slope. Bedding attitude N. 50° W., 23° SW. The tuff 

is medium bedded and ripple bedded. This may be the same unit as seen at SD-8-3. Prominent tufa unit caps hill 

D-55F 

D-55E 

D-55D 

D-55C 

D-55B 

D-55A 

5 em 

1.6 m 

2.4 m above bed D-55D 

At top 

+1.3 m 

+76 em 

At base 

Tufa, brown resistant bed capping hill; locally contains 

cavities with quartz crystals. 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 

Tr 

Tr 

Tr 

1 

10 Tr 

1 

10 

10 

4 5 

Locality D-56, NE1/4SW1/4 sec. 27, T. 11 S., R. 43 E. Poor exposure in road cut south of dirt road. Bedding attitude N. 30° w., 25° W. Two tuffs are in this sequence 

D-56C 

D-56B 

D-56A 

31 em 

1.5 m above bed D-56B 

Upper part 

Lower part 

White tuff(?); unknown thickness•••••••••••••••••••••••••••• 

Orange tuff; upper part is silicified ••••••••••••••••••••••• 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

8 

4 2 

3 

3 

6 

Locality D-57, NE1/ 4NW1/ 4 sec. 27, T. 11 S., R. 43 E. Poorly exposed diatomaceous mudstone sequence. Iron-stained float from a tuff(?) is on upper part of hill 

D-57 Iron-stained tuff(?) •••••••••••••••••••••••••••••••••••••••• 10 

Locality SD-58, NW1/4NW1/4 sec. 23, T. 11 s., R. 43 E. Measured section collected upward from valley floor at locality D-9. Bedding attitude N. 50° w., 14° NW. 

SD-58-10 

SD-58-9 5.6 m 

SD-58-8 2.9 m 

SD-58-7 2.1 m 

SD-58-6 5.5 m 

SD-58-5 3.4 m 

SD-58-4 9.6 m 

SD-58-3 25.9 m 

SD-58-3C 

SD-58-3B 

SD-58-3A 

SD-58-2 10.8 m 

SD-58-1 12.3 m 

SD-58-1C 7 em 

SD-58-1B 10 em 

SD-59-1A 

12.04 m 

1.08 m 

+1.2 m 

Middle of this unit 

+7.8 m 

+24.1 m 

+7.8 m 

+5.3 m 

+4.6 m 

+5.2 m 

At base of measured 

section. 

Light-brown, silty and sandy mudstone, poorly exposed; top 

is unconformably(?) overlain by gravels. 

Greenish-gray mudstone, poorly exposed; popcorn weathered 

surface; lignite bed at base of unit D-58-9 was 

re-collected at same place as bed D-9E at locality D-9. 

Diatomaceous mudstone; poorly exposed ••••••••••••••••••••••• 

Pale-gray tuff, mostly pumice shards, multiple bedded and 

cross bedded. (Re-collected from same locality as D-9.) 

No sample collected. 

Brownish-gray, tuffaceous mudstone and siltstone sequence, 

white, 1.5-cm-thick layer collected. 

Gray tuff, poorly exposed, abundant detrital material ••••••• 

Brownish-gray tuffaceous mudstone, with diatomaceous 

mudstone. 

Brownish-gray, very poorly exposed mudstone with an unaltered 

Tuffaceous mudstone, resistant layer, with abundant 

gypsum veinlets; no sample collected • 

Tuffaceous mudstone ••••••••••••••••••••••••••••••••••••••• 

Yellowish-brown, fine-grained siltstone ••••••••••••••••••• 

Diatomaceous, tuffaceous mudstone, with platy to 

subconchoidal fracture. 

Tr 

(8) -- --- Tr 

(10) 

--

3 --- Tr 

Tr 

1 

tuff at base of unit: 

--- --- --- ---

Tr 

6 --- --- 1 

Tr 3 

Pale-gray, diatomaceous mudstone unit; poorly exposed, with subconchoidal fracture: 

Gray unaltered tuff, thin (7 em) at top of unit. No 

sample collected. 

Gray tuff, unaltered, thin (10 em) ••• •••• ••• ••• ••• •••••••• 

Diatomaceous mudstone •••••••••••••••••••••••••••.••••••••• 

8 

3 

Tr 

Tr 

Tr 

3 

Locality SD-59, SW1/4NW1/4 sec. 4, T. 11 S., R. 43 E. Section measured upward from welded tuff. Bedding attitude N. 85° E., 10° N. 

Tr 

10 

Tr 

Tr 

Tr 

Tr 

Tr 

9 

6 

5 

3 

1 

2 

1 

:::;: 

:I SD-59-5 5.2 m Tuff, sandy tuff, and lapilli tuff, very resistant ledge 

former; re-collected from locality D2, bed D-2I; top of 

unit unconformable with overlying gravel; no sample 

collected. 

m 

.. 

0 

C SD-59-4 

f 

0 

i 

~ 

N 

'I 

11.7 m +7.2 .m Yellow, altered sandy conglomeratic tuff; very poorly 

exposed. 

6 --- --- 3 --- --- --- --- --- --- --- ---

Tr Am 

2 7A 

Tr 7A 
~ 

~ 

= 

~ c. 

~· 

:r 

c Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, Oregon--continued 

~· 

:II Sample No. 

IJ (locality 

Measured 

thic\<ness 

shown in m, 

em, or mm 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

iii" No. plus Description of samples 

bed or 

unit No.) Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Locality SD-59--Continued 

SD-59-3 

SD-59-2 

SD-59-2C 

SD-59-2B 

SD-59-2A 

SD-59-1 

SD-59-lC 

SD-59-lB 

SD-59-lA 

10.4 m 

27.3 m 

+23.1 m 

+7.6 m 

+2.7 m 

+1.5 m 

3.3 m below top 

Tuff, lapilli tuff, and sandy tuff, resistant ledge 

former; collected from locality D-2, beds D-2F, 

-2G, and -2H. 

Pale-yellow to gray tuff or sandy tuff beds; altered, poorly exposed: 

White resistant layer, 46 em thick........................ 8 -

Sandy tuff................................................ 7 Tr 

Tuff•••••••••••••••••••••••••••••••••••••••••••••••••••••• 6 1 

Gray welded tuff weathers brown, very resistant, base not exposed: 

Welded tuff, partly altered............................... 10 

Welded tuff, top of resistant ledge....................... 9 

Welded tuff, glassy, perlitic part; from middle of 9 

altered part. 

Locality D-60, NW1/4NE1/4 sec. 5, T. 11 s., R. 43 E. Mudstone containing interbedded tuff between basalt and welded tuff 

D-60D 

D-60C 

D-60B 

D-60A 

77 em 

15.8 m 

Base of welded tuff Welded tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Gray pumice tuff, bedded; at base of welded tuff •••••••••••• 

White tuff ••• ••• ••••••••••• ••••• ••• •••. ••• ••••••••••••• ••••• 

Brown mudstone •••••••••••••••••••••••••••••••••••• • • • • • • •••• 

10 

8 

8 

6 

2 

3 

Tr 

Tr 

Tr 

Tr 

2 

1 

3 

Tr Tr 

1 Tr 

Tr 

Locality D-61, NE1/4NE1/4 sec. 34, T. 11 s., R. 43 E. Thin, very resistant tuff, partly silified, except lowest part which is entirely silicified. Bedding attitude N. 70° w., 12° s. 

D-61 15 em Orange tuff; 2.5 em of gray layer sampled •••••••• •••• •• ••• •• Tr TrAm 

Locality D-62, NE1/4NE1/4 sec. 34, T. 11 S., R. 43 E. Green mudstone sequence, similar to that at locality D-61, but collected tuff is stratigraphically higher than tuff at locality D-61. 

Bedding attitude N. 60° W., 16° s. 

D-62C 

D-62B 

D-62A 

2.5 em 

7.62 em 

Upper part 

At base of bed D-62C 

Lower part 

Orange and white tuff, collected 90 m south of locality 

D-61, and o;tratigraphically lower. 

Tuff •••• •••• ••• •• •••• ••••••• •••••• ••• • •• ••• •• ••• •• •• ••• ••••• Gray, lower part of orange tuff; this tuff is 

stratigraphically above that at locality D-61. 

Locality D-63, NE1/4NW1/4 sec. 35, T. 11 s., R. 43 E. Two thin tuffs in a green mudstone sequence, about 12 m apart 

D-63C 

D-63B 

D-63A 

10.2 em 

7.62 em 

12.2 m above bed D-63B 

Upper part 

At base of bed D-63 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Orange top of lower tuff in the mudstone sequence ••••••••••• 

White tuff ••••••••••••••••••••••••••••••••••••••• • • ••••••••• 

Tr 

Tr 

2 

Tr Tr 

Locality D-64, NW1/4NE1/4 sec. 35, T. 11 s., R. 43 E. Tuff equivalent to that at locality D-7, with crystal casts on underside of tuff base 

D-64 7.6 em Orange tuff, calcareous ••••••••••••••••••••••••••••••••••••• Tr Tr --- ---

10 

4 

8 

10 

5 

4 

7 

Tr ---

Locality D-65, NE1/4NE1/4 sec. 35, T.11 S., R. 43 E. Tuff, with ripple-marked, massive base; thin bedded and ripple bedded at top, in greenish-gray mudstone 

D-65B 

D-65A 

20.3 em Upper part 

At base 

Orange tuff •••••••••••••••••••••.••••••••••••••••••••••••••• 

Tuff •••••• ••••• ••• ••••••• ••• ••• •••• •• •••••••• ••• ••• •• ••• •• •• 1 --- --- -~ 4 

6 ---

Tr Gy 

4 --- --- ---

3 TrAm 
N 

~ ?i' 

c 

~· 

~ 

tD 

"' iii' 

s. 

-1 

c 

=: 

"' 

Locality D-66, SW1/4SE1/4 sec. 26, T. 11 s., R. 43 E. Thick, multiple-bedded tuff overlying siltstone and sandstone. Bedding attitude N. 75° E., 15° S. May be equivalent to bed D-55A at 

locality D-55 

D-66F 

D-66E 

D-66D 

D-66C 

D-66B 

D-66A 

15 em 

30 em 

23 em 

91 em above bed D-66E 

91 em above bed D-66D 

15.2 em above bed D-6 6C 

Yellow tuff in poorly exposed sequence with mudstone 

interbedding. 

Tuff •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Poorly exposed tuff ••••••••••••••••••••••••••••••••••••••••• 

Yellow-green tuff contains "hand-lens"-visible 

equidimensional crystals. 

Tuff •••••••••••••••••• • • •• •• • • • •• •• • •• • • • • • •• • •• • • • • • • • • • • • • 

Tuff •••••••••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 

Tr 

2 

9 

10 

9 

10 

8 

Locality D-67, SE1/4SW1/4 sec. 26, T. 11 S., R. 43 E. Sequence of green siltstone, sandstone, and conglomerate. Bedding attitude N. 85° W., 27° S, Conglomerate is made up mostly of 

gray schist and quartzite 

D-67B 

D-67A 

Brown siltstone ••••••••••••••••••••••••••••••••••••••••••••• 

Coarse sands tone • ••••••••••••••••••••••••••••.•••••...••..•• 

2 

2 

4 

4 

1 

4 

TrAm 

TrAm 

Locality D-68, NE1/4SW1/4 sec. 26, T. 11 S., R. 43 E. Tuffs in brown mudstone sequence. Bedding attitude N. 70° w., 17° S. Tuffs are multiple-bedded; thickness is unknown but is at least 

1,8 m 

D-68C 

D-68B 

D-68A 

Upper part 

Middle part 

At base 

White to pale-yellow tuff ••••••••••••••••••••••••••••••••••• 

Fine-grained tuff ••••••••••••••••••••••••••••••••••••••••••• 

Tuff, coarse grained, silicified •••••••••••••••••••••••••••• 4 

6 Tr 

5 2 

Tr 

Locality D-69, NW1/4SW1/4 sec. 26, T. 11 s., R. 43 E. Collection of fragments of vertabrate fossils from gray siltstone above welded tuff 

D-69 No samples collected, except fossils •••••••••••••••••••••••• 

Locality D-70, NE1/4SE1/4 sec. 27, T. 11 s., R. 43 E. Tuff sequence, poorly exposed 

D-70 5 em Yellow tuff,,, •••••••••••••••••••••••••••••••••••••••••••••• Tr --- --- ---

3 

3 Tr 

6 

10 --- --- --- --- --- --- Tr Am 

S' Locality D-71, NW1/4SW1/4 sec. 26, T. 11 s., R. 43 E. Poorly exposed tuff overlain by green siltstone. Tuff is probably the same as that at locality D-55 

Cl 

~ 

c: 

~ 

~ 

~ 

~· 

tD 

~ 

tD 

rei 

I") 

c 

~ :;· 

tD 

c 

tD 

-g 

~ 

~ 

~ 

c 

c 

f 

0 

~ ~ 

IN 

N 

I.C 

D-71B 

D-71A 

1.5 m Upper part 

Lower part 

Yellow massive tuff ••••••••••••••••••••••••••••••••••••••••• 

Platy tuff •••••••••••••••••••••••••••••••••••• , ••••••••••••• Tr 

10 

Locality D-72, SW1/4SE1/4 sec. 27, T. 11 S., R. 43 E. Several thin tuffs in a green mudstone and siltstone sequence. Bedding attitude N. 60° w., 22° S. 

D-72D 

D-72C 

D-72B 

D-72A 

8 em 

20 em 

10 em 

5 em 

Highest tuff •• , ..... , ....................... ,, ............. . 

Tuff ....................................................... . 

Tuff, locally silicified at base •• ,.,.,,., •• ,,., ••••••• ,., •• 

Lowest tuff, silicified ••••• , ••• , •••• ,.,, ••• , •• , •••••••••••• 

Locality D-73, SW1/4SE1/4 sec. 27, T, 11 s., R. 43 E. Tuff in a green mudstone sequence. Lower part of tuff is silicified 

D-73 15 em Upper part Orange tuff ••••••••••••••••••• ,.,.,,., ••••••••••• , ••• , •• , ••• 

Tr 

Tr 

Tr ---

Locality D-74, SW1/4SW1/4 sec. 35, T. 11 s., R, 43 E. Gray and green mudstone sequence containing thin tuffs. Bedding attitude N, 60° W,, 10° S, 

D-74D 

D-74C 

D-74B 

D-74A 

4 em 

8 em 

13 em Top 

Base 

Thin apatite bed (see also locality D-100).................. (l) 

Tuff........................................................ (4) 

White tuff; lowest tuff in sequence ••••••••••• ,............. (9) 

Orange base of bed D-74B......... .... • ... .. .. .. .. .. • ... .. .. • (8) 

10 

10 

7 

1 0 --- --- --- --- --- ---

2 

Tr 

Tr 

Tr 

Locality D-75, Center of SW1/4 sec. 35, T. 11 S.: R. 43 E. Tuff stratigraphically below tuffs at locality D-74 in greenish-gray mudstone. Bedding attitude N. 60° W., 24° S. Other 

tuffs are above and below the thickest, collected tuff 

D-75C 

D-75B 

D-75A 

15 em 

20 em 

20 em 

Top. 

Middle 

Base 

Gray tuff, thin bedded ..................................... . 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

8 

9 

10 

5 Ap 
~ 

~ 

= 

~ c 

a. 

~· 

:r 

c t' 

;;;· 

Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, Oregon--Continued 

Sample No. 

(locality 

No, plus 

bed or 

unit No.) 

Measured 

thickness 

shown in m, 

em, or mm 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

Description of samples 

Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Locality D-76, SEl/4NWl/4 sec. 35, T, 11 s., R. 43 E. Sequence of yellow and greenish-yellow sandstone, conglomerate, tuffaceous sandstone, and sandy tuff. Bedding attitude N. 35° w., 18° SW. Pre-Tertiary rocks crop out just east of D-76. Exposure is in bank of irrigation ditch where the rancher, Earl Sutter, found vertebrate fossil bones. The authors did not find 

fossils, but borrowed the specimens from Mr. Sutter for examination 

D-76 Sandy tuff •••••••••••••••••••••••••••••••••••••••••••• ,..... --- --- --- --- --- --- --- 10 --- --- --- --- --- ---

Locality D-77, SE1/4SW1/4 sec. 35, T. 11 S., R. 43 E. Thick tuff in a gray mudstone sequence. Re-collected the following year in more detail as D-101 

D-77B 

D-77A 

1.5 m Upper part 

Lower part 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Tuff, ripple bedded ••••••••••••••••••••••••••••••••••••••••• 

Tr 

2 

--- 10 

--- --- --- 8 

Locality D-78, SW1/4SE1/4 sec. 35, T. 11 s., R. 43 E. Thin tuff in greenish-yellow mudstone. Bedding attitude N. 80° w., 15° s. and may be unconformable on the lake beds. The tufa was not sampled. Tuff was re-collected as locality D-102 White to pale-gray tufa is about 3 m above the tuff 

D-78B 

D-78A 

46 em Upper part 

Lower part 

Gray tuff, very calcareous ••••••••••••••••••• , •••••••••••••• 

Orange tuff ............................................... ,. 

Locality D-79, NE1/4SW1/4 sec. 36, T. 11 S., R. 43 E. Thin tuff in gray mudstone 

D-79 5 em Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-80, SW1/4NE1/4 sec. 36, T. 11 S., R. 43 E. Tuff in a brown siltstone and sandstone. Same as bed D-20C, locality D-20 

4 m Tuff ••••••••••••••• ••• ••• •••••••••••••••••••• ••• •••••••••••• 

Tuff ....................................................... . 

Tr 

Tr 

Tr 

--- --- --- --- --- ---

Tr --- 4 

--- --- ---

--- --- --- --- --- --- 10 

--- --- --- --- --- --- 10 

1 2 --- --- --- --- 3 

10 --- --- Tr 

--- --- 1 --- 4 5 --- TrAm 

D-80C 

D-80B 

D-80A 

Top 

Middle 

Base Tuff •••••••••••• •••• •••••• •••••••••••••••••••••••••••••••••• 2 --- Tr --- Tr --- 7 --- --- 1 ---Tr Gy Am 

Locality D-81, SWl/4SWl/4 sec. 31, T. 11 s., R. 44 E. Tuff in greenish-gray mudstone, at locality D-107. Next hill west of D-81 is same sequence as that at locality D-20 

D-81 25 em White tuff, but locally orange stained in a 

greenish-gray mudstone. 

Locality D-82, SW1/4SE1/4 sec. 32, T. 11 s., R. 44 E. See also locality D-22 

D-82 Brown sandstone •.•••••.••••••••••••••••••••••.••.•••••..•••• 

--- --- --- 1 0 

Tr 2 --- --- --- --- ---

Locality D-83, SW1/4SW1/4 sec. 36, T. 11 S., R. 44 E. Several thin tuffs in a greenish-gray mudstone. Bedding attitude N. 10° w., 22° SW. 

D-83C 

D-83B 

D-83A 

0,1 em 

2.5 em 

7.6 em 

9.1 m above bed D-83A 

15 em above bed D-83A 

Yellow tuff •••••••••• , ••••••••••••••••••••••••• , •••••••••••• 

Yellow tuff ••••••••••••••••••••••••• : •••••• , •••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-84, SWl/4NWl/4 sec. 6, T. 12 s., R. 44 E. Thin tuff in a greenish-gray mudstone. Bedding attitude N. 25° w., 26° SW, 

D-84B 

D-84A 

25 em Upper part 

Lower part 

Yellow tuff, ripple bedded •••••••••••••••••••••••••••••••••• 

Yellow tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr 

Tr 

Tr 

Tr 

3 

10 

10 

Locality D-85, SEl/4SW1/4 sec. 6, T. 12 s., R. 44 E. Thick pale-gray to white diatomaceous mudstone and diatomite. Bedding attitude N. 30° w., 14° SW. 

D-85B 

D-85A 

5 em· 

2.5 em 

Iron-stained fresh tuff ••••••••••••••••••••••••••••••••••••• 

Orange layer of thin, very resistant tuff(?) found only 

as float. 

8 

5 Tr 

Tr 

1 --- --- 5 ---

3 

1 Tr 

4 1 

2Am 
i 

;:;· 

c *. = ~ 

"' iii. 

s. 

ct 

a: 

;· 

~ 

= c: 

i .. 

~ 

~-

~ 

= ~ 

~ c 

=- .. 

~-

i 

1 

=-

m = 

.. 

c 

c 

J 

0 

f 

~ 

~ 

Locality D-86, SE1/4NE1/4 sec. 6, T. 12 S., R. 44 E. Thin tuff in a green mudstone and sandstone 

D-86 5 em White tuff•••••••••••••••••••••••••••••••••••••••••••• •••••• 1 0 --- --- Tr --- --- --- --- --- --- --- Tr --- ---

Locality D-87, NE1/4NW1/4 sec. 6, T. 12 s., R. 44 E. Sandstone, 12.1 m thick, coarse grained with cross-bedding. Bedding attitude N. 5° E., 12° E. 

D-87C 

D-87B 

D-87A 

Upper part 

Middle part 

Lower part 

Sandstone •••••••••••••••••••••••••• ••• •• ••• ••.••.••••.•••••• 

C::andstone ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Sandstone, locally cross-bedded ••••••••••••••••••••••••••••• 

Tr 

Tr 

4 

1 

1 

5 

5 

Locality D-88, SE1/4NW1/4 sec. 8, T. 12 S., R. 44 E. Thick (30.4 m) pale-gray to yellow sandstone similar to sandstone at locality D-87. Bedding attitude N. 30° W., 8° NE. 

D-88 Sandstone with rare pumice fragments •••••••••••••••••••••••• 

Locality D-89, SE1/4NW1/4 sec. 6, T. 12 s., R. 44 E. Tuff collected as float in a green mudstone sequence, probably similar to bed D-28A at locality D-28 

2 

4 

2 

2 ---

D-89 Float •••••••••••••••••••••••••• •••• ••• ••• •••••••• ••• • • • • • •• • --- --- --- --- 10 Tr --- --- --- --- --- ---

Locality D-90, SE1/4NW1/L. sec. 2, T. 12 S., R. 43 E. Thick tuffaceous unit that breaks with a conchoidal fracture 

D-90 9.1 m Pale-green tuffaceous unit. 

Locality D-91, SE1/4NW1/4 sec. 27, T. 11 S., R. 45 E. Thin tuffs clearly seen in a poor exposure of partially covered mudstone in lower part of sequence and diatomite in upper part 

D-91D 

D-91C 

D-91B 

D-91A 

em 

em 

em 

7.6 m above bed D-91B 

Gray, resistant sandstone overlying bed D-91C ••••• ••• •• ••••. 

Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Gray, varied calcareous unit, immediately overlying bed 

D-91A. 

Gray tuff in lower part of sequence; rat tooth marks on 

weathered surfaces. 

1 

5 

2 

5 

1 

Tr 

Locality D-92, SW1 I 4SW1 /4 sec. 14, T. 11 S., R. 43 E. Contact between pale-gray diatomite to the south and reddish-brown siltstone, sandstone, and conglomerate to the north. 

appears to be a nearly vertical fault (north side down) with an attitude of N. 80° W., 70° N. 

Contact 

D-92 No sample. 

Locality D-93, SE1/4SW1/4 sec. 24, T. 11 s., R. 43 E. Poorly exposed tuffs in a sequence that is mostly siltstone and sandstone 

D-93B 

D-93A 

4.7 m above bed D-93A Gray tuff, silty, unknown thickness ••••••••••••••••••••••••• 

Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 

Tr ---

Locality D-94, NE1/4SW1/4 sec. 24, T. 11 s., R. 43 E. Thin tuff in a siltstone and sandstone sequence. Bedding attitude N. 40° w., 18° N. 

D-94 10 em Cream-colored tuff, locally cross-bedded •••••••••••••••••••• 

Locality D-95, NE1/4SE1/4 sec. 24, T. 11 S., R. 43 E. Thin tuffs in a mudstone and siltstone sequence 

D-95E-1 2.5 em 4.5 m above bed D-95D White tuff with raindrop prints on base; upper surface 

very irregular. 

D-9SE-2 1 mm At base Thin black layer at bottom of tuff •••••••••••••••••••••••••• 

D-95D 10 em 4.5 m above bed D-95C Orange tuff with crystal casts on base; same tuff as at Tr --- bed D-7A, locality D-7. 

D-95C-1 61 em 6.1 m above bed D-95B Orange to yellow, platy upper part of tuff •••••••••••••••••• 

D-9SC-2 Lower massive part of tuff •••••••••••••••••••••••••••••••••• 

D-9SB 8 em 6.1 m above bed D-95A Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 4 

D-95A 10 em Orange tuff, lowest in sequence ••••••••••••••••••••••••••••• 

Locality D-96, SE1/4NE1/4 sec. 24, T. 11 s., R. 43 E. Thin tuffs in mudstone sequence identical to that at locality D-95 

D-96 Orange tuff with crystal casts as in beds D-7A and D-9SD •••• 1 ---

7 

2 

Tr 

4 

10 -- -- --- --- --- --- --- ---

10 

4 4 

4 5 

10 Tr 

4 6 

--- 3 

10 

--

--

3 Tr --- --- --- --- --- ---

Tr Am 

2 Am 

4Am 

1 Am 

TrAm 

TrAm 

TrAm 

lGy 
~ 

~ 

N 

~ 

a. 

;:;-

fll 

:r 

0 

~· 

; 

fll 

iii' 

Appendix 1. Mineralogy of lacustrine rocks from the Durkee basin, Baker County, Oregon--Continued 

Sample No. 

(locality 

No. plus 

bed or 

unit No.) 

Measured 

thickness 

shown ·in m, 

em, or mm 

Position of sample 

collected above ( +) 

or below (-) base 

of designated unit 

Minerals (relative abundance reported as "parts-in-ten" as estimated from 

X-ray powder diffraction analysis) 

Description of samples 

Gl CT lOA 14A Ph Ch Er Cp Mo An KF Qtz PF Cl Other 

Locality D-97, NW1/4NW1/4 sec. 13, T. 11 S., R. 43 E. Bedding attitude N. 70 w., 24° N. Thick, unaltered tuff overlain by gravels; probably the same as that in unit SD-58-7 

at locality SD-42 

D-97 1.5 m White tuff •••••••••••••••••••• ,,,,,,,,., •• , ........ ,, ...... , 1 0 --- --- Tr --- --- --- --- --- --- --- --- --- ---

Locality D-98, NE1/4NW1/4 sec, 24, T. 11 S., R. 43 E. Several tuffs in a mudstone sequence, Bedding attitude N. 65° W., 25° N, 

D-98F 51 em Upper part White tuff .............................. , •• , ... ,,, ......... , 10 --- --- --- --- Tr Tr 

D-98E 3 m above bed D-98D White tuff with crystal casts on base, overlying a --- --- --- 1 --- --- 8 

sandstone. 

D-98D 25 em 8 m above bed D-98C White and orange mottled tuff •••••••••••••••• ,., •••••••••• ,. --- --- --- Tr --- 3 4 3 --- --- --- --- --- --- D-98C 10 em 5 m above bed D-98B Orange tuff .. · ............................................... --- --- --- 2 --- 3 2 3 

D-98B 36 em 7 m above bed D-98A White tuff with rat tooth marks on weathered surfaces 4 --- --- Tr 3 2 1 

(see bed D-91A). 

D-98A 30 em White tuff, lowest in sequence •••••••. •••••••••••••••••••••• 

Locality D-99, SE1/4SE1/4 sec, 34, T. 11 s., R. 43 E. Thin bed of phosphatic rock in brown mudstone; same as that at bed D-74D at locality D-74. Just north of D-99 an east-west fault 

places the mudstone in contact with a pale gray diatomite 

D-99 4 em Brown phosphate •• , ••••• , ••• ,.,,,,,,,,,, ••••••••••••• ,,,,,,,, 1 --- --- --- --- 1 --- ---

Tr Gy 

8 Ap 

Locality D-100, SW1/4SW1/4 sec, 35, T, 11 s., R. 43 E, Re-collection of bed D-74D, at locality D-74 on top of a small knob, Phosphate bed is in mudstone sequence that has popcorn weathering 

D-lOOB 

D-lOOA 

4 em Thin, brownish-orange phosphate,,,,,,.,,,.,, •••••••••••••••• 

Greenish-gray mudstone,,,,,,,,,,,,,, •• , ••••••••••••••••••••• 

3 

8 

Locality D-101, SW1/4SW1/4 sec. 35, T. 11 S., R. 43 E. Exposure in irrigation ditch bank near Plano Road, in a mudstone-carbonate sequence. Bedding attitude N. 70° w., 24° s. 

D-101N 

D-lOlM 

D-lOlL 

D-101K 

D-lOlJ 

D-101! 

D-lOlH 

D-lOlG 

D-lOlF 

D-101E 

D-lOlD 

D-101C 

D-101B 

D-lOlA 

15 em 

30 em 

20 em 

25 em 

10 em 

30 em 

76 em 

1,8 m 

15 em 

em 

30.5 em above bed D-lOlM Gray tuff, weathers orange.,,,,,,,,,, •••••• ••••••••••••••••• 

Gray tuff ••••••••••••••• , •• , ••••• ,,, •••••••••••••••••••••••• 

61 em above bed D-lOlK Orange and white tuff ...................................... . 

At tuff base Orange tuff at base of thick orange tuff •••••••••••••••••••• 

61 em above base 

30 em above bed D-101B 

Pale-gray massive tuff,,, ••••••••••••••••••••••••••••• ,,,,,, 

Pale-yellow tuff, ripple bedded., ••••••••••••••••••••••••••• 

Pale-greenish-gray tuff,,,,,., •••••••• ,,,,,,,,,,,,,, •• , ••••• 

Pale-gray, ripple-bedded tuff that is the same as bed 

D-77A; section includes sequence from bed D-lOlG 

through bed D-lOlN and is 2 m thick, 

Pale-greenish-gray, calcareous tuff •••••• ,,,,,,,,,,,,.,,,,,, 

Greenish-gray mudstone,.,,,,,.,,.,,.,,,,,, •• ,, •••••••••••••• 

Pale-gray carbonate bed., ••• ,, •• ,,., •••••• , •• ,,,,,,,,,,,,,,, 

Cream-colored tuff,,,,,,,,,,,,,,,,,, •• ,,.,.,,,,,,,,,,,,,,,,. 

Greenish-gray mudstone.,,,,,,,,,,,,,,,,, •••••••••••••••••••• 

Thin tuff near base of exposed sequence •••••••••••••••••• ,,. 

Locality D-102, SW1/4SE1/4 sec. 35, T, 11 s., R, 43 E, Same as locality D-78 

D-102B 

D-102A 6.1 m below 

Calcareous tuff •••••••••••.••••••••••••••••••••••••••.•.••.• 

Green fragment of schistose rock,,,,,,,,, ••••••••••••••••••• 

Tr 

Tr 

Tr 

Tr 

Tr 

Tr 

Tr 

1 

3 

1 

Tr 

1 

Tr 

Tr 

Tr 

---

---

---

---

---

---

---

---

---

---

---

8 Tr 

10 Tr 

--- --- --- Tr 

--- Tr 

--- ---

--- Tr 

--- 1 

--- 1 

--- 5 

--- ---

2 

3 

8 

10 

1 

Tr 

Tr 

1 

1 

9 

4 

---

Tr 

--- 5 

8 

10 

Tr 9 

1 2 

8 

10 

5 Ap 

1 Am 
~ :::0: 

;:r 

c 

~-

~ ~· 

s. 

-t 

c 

~ 

:r 

Dol 

::I 

c 

I .. 

~ 

~-

::I 

I'D 

~ 

i 

.. 

~-

1 

:::0: 

~ Dol 

.. 

c 

c 

t 

0 

~ ::I 

CN 

CN 

CN 

Locality D-103, SW1/4NE1/4 sec. 34, T. 11 s., R. 43 E. Sequence of greenish-gray mudstone containing a thin tuff and a thin dolomite bed 

D-103B 

D-103A 

9.14 m above bed D-103A Gray dolomite, weathers orange •••••••••••••••••••••••••••••• 

Orange tuff ••••••••••••••••••••••••••••••••••••••••••••••••• 

Locality D-104, NEll 4SE1/ 4 sec. 34, T. 11 S., R. 43 E. Sequence of p.oorly exposed mudstone and tuff 

D-104 Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 

10 Tr 

10 --- Tr 

Locality D-105, NW1/4SE1/4 sec. 25, T. 11 s., R. 42 E. Sequence of diatomite and an unaltered tuff. Bedding attitude N. 45° w., 30° s. 

D-105 Gray tuff ••••••••••••••••••••••••••••••••••••••••••••••••••• 10 

___ ___ ___ ___ ___ --- Tr 

Locality D-106, NEl/ 4SW1/ 4 sec. 9, T. 11 S., R. 43 E. Thin, unaltered tuff on the flank of Iron Mountain in a pinkish-brown siltstone, sandstone, and conglomerate sequence 

D-106 10 em Pale-gray tuff •••••••••••••••••••••••••••••••••••••••••••••• 10 --- --- --- --- --- --- --- --- ---

Locality D-107, SW1/4SW1/4 sec. 31, T. 11 S., R. 44 E. Chert bed in mudstone sequence about 90 m east and downslope from locality D-81 

D-107Z 

D-107Y 

15 em 

2.4 em 

Chert bed, Magadi -type (?), dark gray •••••••••••••••••••••• 

Magadi-type chert, float from below bed D-81Z; gray 

translucent, with petrified wood. 

Tr Tr ---

10 

10 

10 Do 

Magadi-Type Chert-A Distinctive Diagenetic Variety 

From Lacustrine Deposits 

By RICHARD A. SHEPPARD, U.S. Geological Survey, Box 25046, Denver 

Federal Center, MS 939, Denver, Colorado 80225, and 

ARTHUR J. GUDE 3rd, U.S. Geological Survey, Box 25046, Denver Federal 

Center, MS 917, Denver, Colorado 80225 

CONTENTS 

Abstract 336 Chemical composition 341 

Introduction 336 Genesis of the chert 342 

Acknowledgments 336 

Occurrences of Magadi-type chert in 

the United States 336 

Significance of Magadi-type chert 344 

References cited 345 

Description of the chert 336 

Textural characteristics 339 

Petrographic characteristics 339 

X-ray diffraction characteristics 340 

FIGURES 

1. Map showing occurrences of Magadi-type chert in sedimentary rocks of 

the United States 337 

2-6. Photographs: 

2. Thin bed of Magadi-type chert in zeolitic tuff near Paradise Valley, Nevada 339 

3. Platy Magadi-type chert from lacustrine facies of Gila Conglomerate 

near Buckhorn, New Mexico 339 

4. Lobate nodule of Magadi-type chert from upper Tertiary lacustrine 

deposits near Paradise Valley, Nevada 340 

5. Magadi-type chert from lacustrine facies of Gila Conglomerate near 

Buckhorn, New Mexico 340 

6. Lobate nodule of Magadi-type chert from "Rome beds" near Rome, Oregon 340 

7-10. Photomicrographs of interior of Magadi-type chert showing: 

7. Aggregate of interlocking, irregular, fine chalcedony 341 

8. Irregular patch of spherulitic chalcedony 341 

9. Vague, ellipsoidal bodies of chalcedony 342 

10. Filling of crystal mold by chalcedony, finely crystalline quartz, and 

coarsely crystalline quartz 342 

11. X-ray diffraction traces of well-crystallized quartz and Magadi-type chert 343 

TABLES 

1. Occurrences of Magadi-type chert in the conterminous United States 338 

2. Semiquantitative spectrographic analyses of Magadi-type chert 344 

3. Minerals associated with strata containing Magadi-type chert 345 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 335 
Abstract 

Chert that formed from magadiite (NaSip13 (0H)3 • 3H20) 

or other hydrous sodium silicate minerals was first described 

in 1967 by H. P. Eugster, who observed it in upper Pleistocene 

lacustrine deposits at Lake Magadi, Kenya. Since then, this 

distinctive chert of inorganic origin has been recognized in other 

Pleistocene lacustrine deposits of the Eastern Rift Valley in Kenya 

and Tanzania and in many lacustrine deposits of Jurassic to 

Pleistocene age in the United States. Magadi-type chert is 

especially common in the upper Cenozoic deposits of the Basin 

and Range province. The chert occurs chiefly as thin, discontinuous beds or as plates or nodules of irregular shape. Typically, 

the chert is dense, homogeneous, and translucent and has a 

thin, soft, white, opaque rind. Most of the chert consists of finely 

crystalline chalcedonic quartz, but the rind may consist of 

chalcedony and (or) various types of silicate or aluminosilicate 

minerals. This chert is known to have formed only in deposits 

of saline and highly alkaline lakes of the sodium carbonatebicarbonate type. The chert did not precipitate directly from 

the lake water or brine; rather, it formed during early diagenesis 

from a primary precipitate of magadiite or some other hydrous 

sodium silicate mineral. The conversion to chert from these 

precursor minerals involved a loss of sodium and water and 

a reduction in volume. The distinctive surface reticulation and 

soft-sediment deformation features of the chert were produced 

during the conversion. Magadi-type chert is an excellent 

indicator of past environmental conditions, and its recognition 

can also be a valuable prospecting tool for locating evaporites, 

zeolites, and fluorite in continental sedimentary deposits. 

INTRODUCTION 

Nodular and bedded cherts are common throughout the geologic record, and many occur in marine 

deposits (Pettijohn, 1975). The origin of most of these 

cherts has been ascribed either to biochemical accumulations or to silicification of carbonate rocks. Magadi-type 

chert, that is, chert that formed from magadiite 

(NaSi7013(0H)3 • 3H20) or other hydrous sodium silicate 

minerals, was first described by Eugster (1967, 1969) from 

the upper Pleistocene lacustrine deposits of Lake Magadi, 

Kenya. Although the cherts at Lake Magadi had previously been recognized and described by Baker (1958), 

their origin was not known until Eugster's investigations. 

Hay (1968) described nodules of Magadi-type chert in 

Pleistocene lacustrine deposits at Lake Magadi and at 

Lake Natron and Olduvai Gorge in Tanzania. 

Magadi-type chert was first reported in the United 

States from the Eocene Green River Formation of Wyoming by Eugster and Surdam (1971) and later from 

numerous localities in lacustrine rocks of Pleistocene to 

Jurassic age in Wyoming by Surdam and others (1972). 

Subsequently, chert believed to have formed from a 

magadiite precursor was reported from Pliocene lacustrine 

deposits in southeastern Oregon and southwestern New 

Mexico (Sheppard and Gude, 1974) and in central Nevada 

336 Studies in Diagenesis 

(Sheppard and Gude, 1980). Nodular varieties of Magaditype chert are prized by rock collectors and lapidaries and 

are appropriately called snakeskin agates. 

Recognition of Magadi-type chert in the Cambrian 

Observatory Hill Beds in south-central Australia recently 

helped White and Youngs (1980) interpret the depositional 

environment as a saline, alkaline lake. This interpretation 

of these Cambrian rocks is the first record of alkaline, lacustrine sedimentation in the Cambrian System of Australia 

and necessitates a reevaluation of existing concepts of Cambrian geography and tectonics in central Australia. 

This present report provides a catalog of the occurrences of Magadi-type chert in the United States; several 

of the occurrences have not been previously reported. In 

addition, the morphological, mineralogical, and chemical 

characteristics of the chert are described to aid further discoveries, and the significance of the chert is discussed with 

respect to the interpretation of the depositional environment of the host rocks and to the potential use of the chert 

as a guide for locating other valuable mineral deposits. 

Acknowledgments 

We thank R. L. Hay for providing samples of chert 

from the east African lacustrine deposits and R. C. Surdam for providing samples of chert from the Green River 

Formation of Wyoming. 

OCCURRENCES OF MAGADI-TYPE CHERT IN 

THE UNITED STATES 

Occurrences of Magadi-type chert are shown in 

figure 1 and are listed in table 1. Except for the locality 

in South Carolina, the known occurrences of Magadi-type 

chert are in the Western United States, particularly in the 

Basin and Range province. Occurrences are especially 

numerous in Wyoming, probably because a special effort 

was made to search for deposits in that State. The chert 

has been identified in rocks that range in age from Jurassic 

to Pleistocene, except that no occurrence of Cretaceous 

age has been found. More than half the occurrences are 

in rocks of Miocene or younger age. 

DESCRIPTION OF THE CHERT 

Magadi-type chert occurs as thin beds or as irregular 

plates or nodules, but nodular varieties are most common. 

Thin beds have been reported from the Green River Formation (Eugster and Surdam, 1971) and the Moonstone 

Formation (Mariner, 1971) in Wyoming and from the 

"Rome beds" in Oregon (Sheppard and Gude, 1974). We 

have also recognized thin beds of Magadi-type chert in a 

lacustrine facies of the Gila Conglomerate near Buckhorn, 
{ 

~ i 

n 

:r 

tD 

::l 

IN 

IN 

...... 

Figure 1. Occurrences of Magadi-type chert in sedimentary rocks of conterminous United States. Data for numbered localities are given in table 1. 
Table 1. Occurrences of Magadi-type chert in the conterminous United States 

[Locality numbers are shown in figure 1. Asterisk(*) indicates localities at which identification of chert as Magadi-type is tentative] 

Locality 

1. Near Durkee, sec. 31, T. 11 s., R. 44 E. , Baker County, 

Oregon. 

~. Near Rome, Malheur Coupty, 

Oregon. 

3. Near Paradise Valley, sec. 17, 

T. 41 N., R. 43 E., Humboldt 

County, Nevada. 

4. Near Dixie Valley, sec. 7, T. 24 

N., R. 38 E., Churchill 

County, Nevada. 

5. Near Eastgate, Churchill County, 

Nevada. 

6. Near Hector, sec. 35, T. 8 N., 

R. 5 E., San Bernardino 

County, California. 

7. Near Bismarck Mesa, sec. 11, 

T. 13 N., R. 6 W., Yavapai 

County, Arizona. 

8. Near Buckhorn, Grant County, New 

Mexico. 

9. Near Coyote, Rio Arriba County, 

New Mexico.* 

10. Near Green River, Sweetwater 

County, Wyoming. 

11. Near Green Cove, Fremont County, 

Wyoming. 

12. Near White Ridge, Natrona 

County, Wyoming. 

13. Near the Bug Ranch headquarters, 

Natrona County, Wyoming. 

14. Near Nelson Springs, Carbon 

County, Wyoming. 

15. Chalk Mountain, Carbon County, 

Wyoming. 

16. Near Buckeye Creek, Albany 

County, Wyoming. 

17. Along Horse Creek, Laramie 

County, Wyoming. 

18. Near Sheep Mountain Table, 

sec. 2, T. 4 s., R. 13 E., 

Pennington County, South 

Dakota.* 

19. Near Orangeburg, Orangeburg 

County, South Carolina. 

20. Near Dragon, sec. 13, T. 12 s., R. 25 E., Uintah County, Utah. 

21. Near Greasewood Gulch, sec. 31, 

T. 1 s., R. 96 w., Rio Blanco 

County, Colorado. 

22. Near Carrizozo, sec. 21, 

T. 9 S., R.10 E., Lincoln 

County, New Mexico. 

338 Studies in Diagenesis 

Occurrence 

Unnamed lacustrine rocks of Pliocene 

age. 

"'Rome beds"' of Pliocene age. 

Unnamed lacustrine rocks of Miocene 

or Pliocene age. 

Unnamed lacustrine rocks of Cenozoic 

age. 

Monarch Mill Formation of Pliocene 

age. 

Unnamed lacustrine rocks of Miocene 

or Pliocene age. 

Unnamed lacustrine rocks of Miocene 

or Pliocene age. 

Lacustrine rocks in the Gila 

Conglomerate of Pliocene or 

Pleistocene age. 

Pedernal Chert Member of the 

Abiquiu Tuff of Miocene age. 

Wilkins Peak Member of the Green 

River Formation of Eocene age. 

Wagon Bed Formation of Eocene age--

Moonstone Formation of Pliocene 

age. 

Bug Formation of Pleistocene age--

Unnamed lacustrine rocks of Cenozoic 

age. 

Lacustrine rocks of the White River 

Group of Oligocene age. 

Morrison Formation of Jurassic 

age. 

Unnamed lacustrine rocks of Miocene 

age. 

Brule Formation of Oligocene age--

McBean Formation of Eocene age--

Douglas Creek Member of the Green 

River Formation of Eocene age. 

Garden Gulch Member of the Green 

River Formation of Eocene age. 

Unnamed lacustrine rocks of late 

Tertiary or Pleistocene age. 

References 

This report. 

Sheppard and Gude (19 7 4) • 

This report. 

R. H. Mariner, U.S. Geological 

Survey, 345 Middlefield Road, 

Menlo Park, California 94025 

(oral commun., 1981). 

Sheppard and Gude (1980). 

This report. 

This report. 

Sheppard and Gude ( 197 4). 

Santos and others (1975). 

Sur dam and others ( 1972). 

Surdam and others (1972). 

Mariner (1971). 

Surdam and others (1972). 

Surdam and others (1972). 

Surdam and others (1972). 

Surdam and others (1972). 

Surdam and others (1972). 

W. H. Raymond, u.s. Geological 

Survey, Box 25046, DFC, MS 921, 

Denver, Colorado 80225 (oral 

commun., 1981). 

Houser (1982). 

R. c. Johnson, u.s. Geological 

Survey, Box 25046, DFC, MS 939, 

Lakewood, Colorado 80225 (oral 

commun., 1982). 

R. C. Johnson, u. s. Geological 

Survey, Box 25046, DFC, MS 939, 

Denver, Colorado 80225 (oral 

Commun., 1984). 

R. H. Weber, New Mexico Bureau of 

Mines and Mineral Resources, 

Socorro, New Mexico 87801 (oral 

commun., 1984). 
New Mexico, and in unnamed lacustrine rocks of Miocene 

or Pliocene age near Paradise Valley, Nevada (table 1). The 

beds are 0.1-15 em thick, but are generally less than 2 em 

thick (fig. 2) and are commonly discontinuous and contorted. Platy chert (fig. 3) is generally less than 1 em thick 

and is locally transitional into beds or nodules. Both 

bedded and platy varieties of Magadi-type chert have external and internal textures that are similar to the more common nodular variety described below. Nodules of chert 

vary greatly in shape and size, although most are irregular, 

lobate, and 2-18 em in the longest dimension (fig. 4). The 

shortest dimension of the nodules is generally perpendicular to bedding. At some localities, the nodules are confined to a single layer; but at other places, the chert nodules 

are distributed vertically through as much as 1 m of 

mudstone or zeolitic mudstone. 

Angular chips of Magadi-type chert occur in a channel sandstone of the Moonstone Formation in the NE ~ 

sec. 30, T. 30 N., R. 89 W., at White Ridge, Wyoming (table 

1, locality 12). At this locality, the same stratigraphic interval contains abundant thin beds and plates of chert in 

brownish-gray mudstone. The chert in place in the mudstone and the epiclastic chert in the sandstone appear identical. These relationships indicate that the conversion to 

chert was relatively rapid, and that the thin, siliceous beds 

in the mudstone were already chert before being eroded, 

transported, and incorporated in a penecontemporaneous 

sandstone. Similar occurrences of chert fragments in 

intraformational gravel or sandstone were described by Hay 

(1968) at Olduvai Gorge, Tanzania, and by Eugster (1969) 

at Lake Magadi, Kenya. 

Textural characteristics 

Most of the Magadi-type chert, whether bedded, 

platy, or nodular, is dense, homogeneous, and milky to 

translucent; it generally has a thin, soft, opaque, white 

coating or rind. Although most of the dense chert is light 

gray, some is dark gray, black, greenish brown, or reddish 

brown. Laminations or concentric color bands are rare in 

the chert. The white rinds are 0.1-5 mm thick, but most 

are 1 mm or less thick. 

Perhaps the most distinctive feature of the Magaditype chert is the surface reticulation (figs. 3 and 4). The 

spacing of the reticulate pattern is commonly 0.5-4 mm, 

and the reticulation generally extends through the white 

rind and into the underlying dense chert. This characteristic 

reticulation led rock collectors and lapidaries to describe 

nodular varieties of the chert as "snakeskin agate" (Sheppard and Gude, 1974, p. 625), and the peculiar texture has 

also been termed "crocodile skin" (White and Youngs, 

1980, p. 1281). Superimposed on the reticulation of some 

cherts is a much larger scale pattern of cracks, such as that 

shown in figure 5. These larger cracks extend deeper into 

the chert than the surface reticulation. 

Figure 2. Thin bed of Magadi-type chert (arrows) in zeolitic tuff 

near Paradise Valley, Nevada (table 1, locality 3). Note thin, white 

rind and dark, dense interior. 

Figure 3. Platy Magadi-type chert from lacustrine facies of Gila 

Conglomerate near Buckhorn, New Mexico (table 1, locality 8), 

showing surface reticulation. Scale is in units of 0.5 mm. 

Crystal molds (fig. 6) are common on the surfaces 

of some chert but are less common in the interior of the 

chert. Although some crystal molds are too deformed to 

be identified, molds from nodules in the "Rome beds" 

resemble gaylussite or pirssonite (Sheppard and Gude, 

1974), and other molds in chert from the Moonstone Formation resemble trona (Surdam and others, 1972). 

Generally, those cherts that have crystal molds on the surface lack the characteristic surface reticulation. 

Petrographic characteristics 

Examination of thin sections of the dense chert shows 

that it consists mainly of an aggregate of interlocking, 

irregular, fine grains of chalcedony (fig. 7). Individual grains 

Magadi-Type Chert 339 
range in size from less than 2 I-'m to about 40 J.tm, but most 

are less than 25 J.tm. All the chalcedony exhibits anomalous 

extinction and is length slow. The mean index of refraction of the chalcedony is betweem 1.53 and 1.54 and is 

distinctly lower than that of quartz. Some chert contains 

rare veinlets or relatively large patches of fibrous chalcedony which is coarser grained than that of the matrix. 

The chert from Bismarck Mesa, Arizona (table 1, locality 

7) contains veinlets and large patches of spherulitic chalcedony (fig. 8). The spherulites are as much as 250 I-'m in 

size and are length slow. 

Figure 4. Large, lobate nodule of Magadi-type chert from 

upper Tertiary lacustrine deposit near Paradise Valley, Nevada 

(table 1, locality 3), showing surface reticulation. Scale is in units 

of 0.5 mm. 

Some of the Magadi-type chert from Rome, Oregon 

(table 1, locality 2) and Eastgate, Nevada (table 1, locality 

5) contain abundant, but vague, ellipsoidal bodies of 

chalcedony that are recognizable under crossed nicols 

(fig. 9). These bodies are 0.03-1.4 mm in their longest 

dimension and characteristically consist of a very fine 

grained core (less than 2 I-'m in grain size) and a coarser, 

fibrous rim. Some ellipsoidal bodies also show a vague concentric structure within the cores. The fibrous chalcedony 

rims of the bodies are length fast, in contrast to all the 

other chalcedony examined. Inasmuch as most of the 

magadiite in lacustrine deposits is spherulitic and about 

the same size as the ellipsoidal bodies in the chert (Eugster, 

1967; Hay, 1968), these chalcedonic bodies are probably 

relicts of the original magadiite. 

Crystal molds in some of the Magadi-type cherts are 

filled or partly filled with a variety of materials. For example, crystal molds in some chert nodules from the "Rome 

beds" of southeastern Oregon (Sheppard and Gude, 1974) 

are filled with chalcedony, quartz, and calcite. A thin layer 

of length-slow chalcedony lines the molds and is followed 

inwardly by finely crystalline, granular quartz, and then 

by coarsely crystalline quartz (fig. 10). Some of the coarse 

340 Studies in Diagenesis 

0 10 20 .30 40 50 60 70 8 0 90 100 

Figure 5. Magadi-type chert from lacustrine facies of Gila Conglomerate near Buckhorn, Nevv Mexico (table 1, locality 8), showing larger-scale pattern of cracks superimposed on surface 

reticulation. Scale is in units of 0.5 mm. 

Figure 6. Lobate nodule of Magadi-type chert from "Rome 

beds" near Rome, Oregon (table 1, locality 2), showing deformed 

crystal molds on surface. Note dark, dense chert beneath thin, 

white rind. Scale is in units of 0.5 mm. 

quartz shows crystal faces where it projects into a still 

unfilled cavity. This sequence is the same in all molds, and 

calcite, where present, postdates the coarse quartz. 

X-ray diffraction characteristics 

X-ray powder diffractometer (XRD) patterns of the 

dense chert confirm the presence of quartz but show some 

significant differences (fig. 11). The ratio of intensities of 

the 101 peak to the 100 peak for well-crystallized quartz 

is about 4, but this ratio for most of the Magadi-type cherts 

is 6-8. Also, the XRD patterns of the dense chert generally 
Figure 7. Photomicrograph of interior of chert nodule showing 

aggregate of interlocking, irregular, fine grains of chalcedony. 

Nodule is from "Rome beds" (table 1, locality 2). Crossed 

nicols. 

show weak, relatively broad peaks at the following d spacings: 4.40-4.44, 3.08-3.ll, 2.86-2.88, and 2.53-2.55 A. 

These peaks have not been identified, but we have also 

recognized them in chert from the Pleistocene deposits 

at Lake Magadi, Kenya, and at Lake Natron and Olduvai 

Gorge, Tanzania. Inasmuch as the cherts consist mainly 

of silica (table 2), this unidentified phase may be a "new" 

silica mineral. Two of the above d spacings are identical 

to those of an unidentified silica(?) phase reported by Hay 

(1970) from chert nodules at Olduvai Gorge. XRD patterns of the dense chert from deposits older than 

Oligocene lack peaks of the unidentified phase, and the 

ratio of intensities of the 101 peak to the 100 peak is close 

to that of quartz. Thus, with increasing age, the chalcedony and the unidentified phase seem to convert to quartz. 

The thin, white rinds on the Magadi-type chert consist mainly of an aggregate of irregular grains of chalcedony 

that are more finely crystalline than the chalcedony of the 

dense part of the chert. Most of the chalcedony is less than 

2 J.Lm in grain size and appears milky in reflected light. Our 

XRD studies of the rinds show that some contain, in addition to chalcedony, one or more of the following minerals: 

calcite, clinoptilolite (a common siliceous zeolite), silhydrite 

(3Si02 • H20), and opal-CT. All these constituents are 

probably diagenetic. Several specimens of platy 

Figure 8. Photomicrograph of interior of chert nodule showing 

irregular patch of spherulitic chalcedony. Nodule is from upper 

Tertiary lacustrine deposit at Bismarck Mesa, Arizona (table 1, 

locality 7). Crossed nicols. 

Magadi-type chert from the Pliocene Moonstone Formation (table 1, locality 12) have rinds that consist only of 

silhydrite. Prior to our recognition of the silhydrite in these 

rinds, the only known occurrence of this mineral was near 

Trinity Center, California, where it has formed from an 

unusual spring deposit of magadiite (Gude and Sheppard, 

1972). XRD studies by Hay (1968) of the rinds of nodules 

of Magadi-type chert from the upper Pleistocene lacustrine 

deposits at Lake Natron, Tanzania, have shown the presence of one or more of the following minerals: magadiite, 

kenyaite (NaSi 11 0 20.

5

(0H)4 • 3H20), quartz, clinoptilolite, 

and erionite (a zeolite). Also, Eugster (1969) recognized 

magadiite and kenyaite in the coatings on chert from the 

upper Pleistocene deposits at Lake Magadi. It is significant that magadiite and other hydrous sodium silicate 

minerals are absent from cherts in deposits older than late 

Pleistocene. Apparently, the magadiite and similar minerals 

have reacted during diagenesis to form silica or aluminosilicate phases. 

CHEMICAL COMPOSITION 

Standard chemical analyses were not performed during this study on any of the Magadi-type cherts, but semiquantitative spectrographic analyses of the chert from six 

Magadi-Type Chert 341 
Figure 9. Photomicrograph of interior of chert nodule showing 

vague, ellipsoidal bodies (arrows) of chalcedony that probably 

are relicts of original magadiite. Nodule is from "Rome beds" 

(table 1, locality 2). Crossed nicols. 

localities are given in table 2. These analyses indicate that 

the cherts contain only minor constituents other than 

Si02 • The most abundant impurity is calcium which is 

probably in minor calcite. The cherts also undoubtedly 

contain a small quantity of water, although no determination or loss on ignition was made. The uranium content 

of the chert is about 1-46 ppm, and these values are not 

considered anomalous (R. A. Zielinski, U.S. Geological 

Survey, Box 25046, DFC, MS 424, Denver, Colorado 

80225, oral commun., 1981). 

GENESIS OF THE CHERT 

Magadiite (NaSi7013{0H}3 • 3H20}, the presumed 

precursor for most of the Magadi-type chert, is known to 

precipitate in alkaline lakes, particularly those of the 

sodium carbonate-bicarbonate variety (Eugster, 1967; 

Rooney and others, 1969; Surdam and Eugster, 1976). Jones 

and others (1967) determined that these brines in closed 

basins in volcanic terranes of Oregon and Kenya have a 

pH as high as 11 and Si02 contents as high as 2,700 ppm. 

Thus, brines of this type are capable of storing substantial amounts of silica in solution. Dilution of these silicarich brines by freshwater is an obvious mechanism for the 

342 Studies in Diagenesis 

Figure 10. Photomicrograph of interior of chert nodule showing 

filling of crystal mold. Thin layer of chalcedony lines mold and 

is followed inwardly by finely crystalline, granular quartz, and 

then by coarsely crystalline quartz. Nodule is from "Rome beds" 

(table 1, locality 2). Crossed nicols. 

precipitation of magadiite (Eugster, 1967). Freshwater 

flooding of silica-rich, alkaline lake water may cause a 

stratified condition. Magadiite could then precipitate at the 

interface between the freshwater and the saline, alkaline 

water because of the lowering of the pH. Solubility studies 

by Bricker (1969) have shown that any mechanism that 

decreases the pH could cause the precipitation of magadiite 

from silica-rich, alkaline brines. Thus, some process other 

than freshwater dilution could cause magadiite to precipitate. For example, biogenically produced C02 could 

reduce the pH of the brine (Eugster, 1969). Hay (1968) also 

suggested that simple evaporative concentration of the 

brine could lead to magadiite precipitation. 

Most of the Magadi-type cherts, including beds, 

plates, and nodules, exhibit features indicative of soft-sediment deformation. Because the precursor magadiite is soft 

and plastic (Eugster, 1967}, it is susceptible to deformation 

under slight loads. 'JYpical features of the chert, such as 

contorted bedding, protrusions on nodules and plates, and 

extrusion forms, were probably produced by plastic deformation of the magadiite prior to its conversion to chert. 

The conversion of magadiite to chert during diagenesis involves a loss of sodium and water as well as a 
(1 01) 

A 

(100) 

(1 01) 

B 

(100) 

50 40 20 

2() CuKa, IN DEGREES 

Figure 11. Comparison of X-ray diffractometer traces 

of well-crystallized quartz (A) with Magadi-type chert 

from Gila Conglomerate near Buckhorn, New Mexico (8) 

(table 1, locality 8). Note difference in ratios of intensities 

decrease in the volume of the solid phase. Two possible 

equations for the conversion are: 

NaSi70 13(0H)3 • 3H20 + H+-7Si02 + 5H20+ Na+ 

magadiite chert 

and 

NaSi70 13(0H)3 • 3H20-7Si02 +4H20+ Na+ +OH-. 

magadiite chert 

The chemical parameters of the interstitial fluids that may 

control the conversion of magadiite to chert are suggested 

by these two equations and include (1) a decrease of the 

pH, (2) a decrease in the activity of sodium, (3) a decrease 

in the activity of silica, and (4) a decrease in the activity 

of water-caused perhaps by an increase in the salinity of 

the interstitial fluids. Eugster (1969) favored a decrease of 

the pH and the activity of sodium to account for the conversion at Lake Magadi. These decreases could be 

accomplished simultaneously by leaching of the magadiite 

by percolating meteoric water. Hay (1970), however, 

attributed the conversion of magadiite to chert at Lake 

Magadi and Olduvai Gorge to a decrease in the activity 

of silica, possibly caused by the sorption or reaction of 

silica with clay minerals. Subsequent oxygen isotope studies 

on east African cherts by O'Neil and Hay (1973) suggested 

that the conversion took place in fluids of widely varying 

salinities and chemical activity (a +:a +) ratios. Inasmuch as the conversion of mag~diiteH to chert is a 

dehydration reaction, relatively high temperatures at or near 

the land surface would also favor the conversion. 

The actual diagenetic process for the conversion is 

still poorly understood, and our studies of the Magaditype chert in lacustrine deposits of the United States have 

not shed additional light on which parameter or combination of parameters was the driving force for the conversion of magadiite or other hydrous sodium silicate minerals 

to chert. Eugster (1969) suggested that kenyaite or noncrystalline silica may be an intermediate product in the conversion, but we found neither field nor petrographic 

evidence to confirm this intermediate step. We, therefore, 

assume that the chert formed directly from magadiite by 

a solution-precipitation mechanism. The conversion 

apparently took place in a relatively short time. The 

minimum time necessary for the conversion in the east 

African lacustrine deposits was probably a few hundred 

to a few thousand years (Hay, 1968; Eugster, 1969). As 

described, the presence of chert chips in an intraformational channel sandstone in the Moonstone Formation of 

Wyoming indicates that the conversion from magadiite to 

chert was rapid. Also, the fact that magadiite has not been 

reported from lacustrine deposits older than late 

Pleistocene is evidence for relatively rapid conversion. 

of 101 peak to 100 peak. Peaks of unidentified phase in Magaditype chert are indicated by X's. 

Magadi-Type Chert 343 
Table 2. Semiquantitative spectrographic analyses of Magadi-type chert 

[Analyst: N. M. Conklin. Results are to be identified with geometric brackets whose boundaries are 1.2, 

0.83, 0.56, 0.38, 0.26, 0.18, 0.12, and so forth but are reported arbitrarily as midpoints of these brackets, 

1.0, 0.7, 0.5, 0.3, 0.2, 0.15, 0.1, and so forth. The precision of a reported value is approximately plus or 

minus one bracket at 68-percent confidence, or two brackets at 95-percent confidence. G, greater than 10 

percent; L, detected but below limit of determination; N, not detected] 

Lab. No •• . . . . 0235863 0235867 0235866 0235868 0235864 0235870 

Field No. . .... FAM-78 LH0-10-15 E5-12J BMK-3 BNM-13B M00-2A 

Locality No. (table 1) • 2 3 5 7 8 12 

Wei (Jht percent 

Si G G G G G G 

Al 0.002 0.07 o. 07 o. 07 0.05 0.03 

Fe .07 .03 .07 .07 .07 .07 

Mg L .003 .003 .03 • 001 .007 

Ca .03 .07 .15 .15 .07 .15 

Na N .07 .07 .07 .07 .07 

K N N N N N N 

Ti • 0003 .007 .005 .003 .003 .003 

p N N N N N N 

Parts per million 

Ag 0.5 N N 0.5 N 0.5 

B 15 70 100 150 300 30 

Ba 30 70 30 30 15 30 

Cr N N 1.5 7 N 1.5 

Cu 15 70 30 30 20 30 

Mn 7 10 15 15 10 7 

S£ N N 7 150 N 15 

u 1. 35 5. 76 4.16 14.3 17.9 46.4 

y 10 N N N N L 

Yb L N N N N L 

Zr 10 L 10 L 15 15 

1 Uranium determined by delayed neutron activation analysis by H. T. Millard, Jr., F. Luman, and B. Keaten. 

Magadi-type chert inherited several features from 

the original magadiite. These features include (1) the 

lobate form and protrusions of the nodular and platy 

chert that were produced by plastic deformation of the 

magadiite prior to conversion to chert, (2) the relicts of 

magadiite spherulites, (3) the molds of saline minerals, 

and (4) the rare rinds containing silhydrite that so far has 

been described as forming only from magadiite (Gude and 

Sheppard, 1972). The characteristic surface reticulation 

of most Magadi-type chert is apparently due to the volume 

reduction during the conversion of magadiite to chert. 

Except for the silhydrite, these features are identical with 

those of chert known to have formed from magadiite, and 

they seem to be unique to chert of this derivation (Surdam and others, 1972). 

SIGNIFICANCE OF MAGADI-TYPE CHERT 

The occurrence of Magadi-type chert in sedimentary rocks is indicative of an alkaline, lacustrine depositional environment. In order to store enough silica in solution for the precipitation of magadiite, the pH of the lake 

water probably has to be at least 9.5 (Surdam and others, 

1972). All the occurrences of chert, magadiite, and 

kenyaite in the Pleistocene deposits of east Africa studied 

by Hay (1968) were attributed to strata of saline, alkaline 

lakes that were rich in dissolved sodium carbonate344 Studies in Diagenesis 

bicarbonate. Theoretical considerations led Drever (1974) 

to suggest that magadiite could precipitate from sea water 

in the absence of silica-secreting organisms; however, 

magadiite has not been recognized in marine sediments. 

Inasmuch as the Magadi-type chert appears to be 

indicative of a saline, alkaline depositional environment, 

certain other minerals, which are common in such 

envronments and have economic potential, may be 

expected to occur in the same strata as the chert. These 

minerals include evaporite minerals, such as trona, nahcolite, and borate minerals, zeolites, potassium feldspar, 

fluorite, and, possibly, lithium-bearing clay minerals. 

Some, such as the evaporite minerals, are primary precipitates from the lake brine, whereas others, such as the 

zeolites and potassium feldspar, are diagenetic and 

generally result from the reaction of volcanic glass with 

the interstitial brine. Those zeolites that have been identified in strata containing Magadi-type chert are analcime, 

chabazite, clinoptilolite, erionite, mordenite, and phillipsite. Table 3 shows the presence or absence of these potentially economic minerals at nearly half of the Magaditype chert localities listed in table 1. Because the chert 

is easily recognized by its distinctive appearance and 

because the chert resists weathering, Magadi-type chert 

is a valuable prospecting guide for locating these minerals 

in lacustrine deposits. It is also an important indicator 

of past geologic environments of lacustrine sedimentary 

formations. 
Table 3. Minerals associated with strata containing Magaditype chert 

[X, present; 0, looked for but not found; leaders(---), not looked for] 

Loca 1 ity Saline minerals Zeolites Potassium Fluorite LithiumNo. or molds of feldspar bearing 

(table 1} saline minerals cla.v 

1 X X X 0 

2 X X X X 

3 0 X X 0 

5 X X X X 

6 X X 0 0 

7 0 X X 0 

8 0 X X X 

10 X X X X 

12 X X 0 0 

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Bricker, 0. P., 1969, Stability constants and Gibbs free energies 

of formation of magadiite and kenyaite: American 

Mineralogist, v. 54, p. 1026-1033. 

Drever, J. 1., 1974, Geochemical model for the origin of 

Precambrian banded iron formations: Geological Society 

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Eugster, H. P., 1967, Hydrous sodium silicates from Lake 

Magadi, Kenya: precursors of bedded chert: Science, 

v. 157, p. 1177-1180. 

___ 1969, Inorganic bedded cherts from the Magadi area, 

Kenya: Contributions to Mineralogy and Petrology, v. 22, 

p. 1-31. 

Eugster, H. P., and Surdam, R. C., 1971, Bedded cherts in the 

Green River Formation: Geological Society of America 

Abstracts with Programs, v. 3, no. 7, p. 559-560. 

Gude, A. J., 3d, and Sheppard, R. A., 1972, Silhydrite, 

3Si02 • H20, a new mineral from Trinity County, California: American Mineralogist, v. 57, p. 1053-1065. 

Hay, R. L., 1968, Chert and its sodium-silicate precursors in 

sodium-carbonate lakes of east Africa: Contributions to 

Mineralogy and Petrology, v. 17, p. 255-274. 

___ 1970, Silicate reactions in three lithofacies of a semi-arid 

basin, Olduvai Gorge, Tanzania: Mineralogical Society of 

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Houser, B. B., 1982, Chert derived from magadiite in the middle 

Eocene McBean Formation, South Carolina: Geological 

Society of America Abstracts with Programs, v. 14, nos. 1 

and 2, p. 27. 

Jones, B. F., Rettig, S. L., and Eugster, H. P., 1967, Silica in 

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Mariner, R. H., 1971, Experimental evaluation of authigenic mineral reactions in the Pliocene Moonstone Formation: Laramie, 

Wyoming, University of Wyoming, Ph.D. dissertation, 133 p. 

O'Neil, J. R., and Hay, R. L., 1973, 180/160 ratios in cherts 

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Pettijohn, F. J., 1975, Sedimentary rocks (3d ed.): New York, 

Harper and Row, 628 p. 

Rooney, T. P., Jones, B. F., and Neal, J. T., 1969, Magadiite from 

Alkali Lake, Oregon: American Mineralogist, v. 54, 

p. 1034-1043. 

Santos, E. S., Hall, R. B., and Weisner, R. C., 1975, Mineral 

resources of the San Pedro Parks Wilderness and vicinity, 

Rio Arriba and Sandoval Counties, New Mexico: U.S. 

Geological Survey Bulletin 1385-C, 29 p. 

Sheppard, R. A., and Gude, A. J., 3rd, 1974, Chert derived from 

magadiite in a lacustrine deposit near Rome, Malheur 

County, Oregon: U.S. Geological Survey Journal of Research, 

v. 2, no. 5, p. 625-630. 

___ 1980, Diagenetic fluorite in the Eastgate zeolite deposit, 

Churchill County, Nevada: U.S. Geological Survey Open-File 

Report 80-506, 8 p. 

Surdam, R. C., and Eugster, H. P., 1976, Mineral reactions in 

the sedimentary deposits of the Lake Magadi region, Kenya: 

Geological Society of America Bulletin, v. 87, p. 1739-1752. 

Surdam, R. C., Eugster, H. P., and Mariner, R. H., 1972, Magaditype chert in Jurassic and Eocene to Pleistocene rocks, 

Wyoming: Geological Society of America Bulletin, v. 83, p. 

2261-2266. 

White, A. H., and Youngs, B. C., 1980, Cambrian alkali playalacustrine sequence in the northeastern Officer basin, South 

Australia: Journal of Sedimentary Petrology, v. 50, 

p. 1279-1286. 

Magadi-Type Chert 345 

Salt-Induced Diagenesis of Argillaceous Sediments 

By CHARLES W. HOLMES, U.S. Geological Survey, MS 972, Denver Federal Center, 

Denver, Colorado 80225 

CONTENTS 

Abstract 349 

Introduction 349 

Geologic background 349 

Methods 351 

Core and sample selection 351 

Water chemistry 351 

Solid-phase chemistry 351 

Results 353 

Pore-water chemistry 353 

Mineralogy 355 

Discussion and conclusions 357 

Summary 361 

References cited 361 

FIGURES 

1-3. Maps showing: 

1. Location of cores used in study 349 

2. Rates of sedimentation in northern Gulf of Mexico 350 

3. Percentage of foraminifera in sand fraction of surficial sediments in 

northern Gulf of Mexico 350 

4. Plot of carbon percentage of surficial sediments and oxygen content of 

waters versus depth in northern Gulf of Mexico 351 

5. Contour map of top of salt structure in northwestern Gulf of Mexico 

showing relationship of boreholes to salt mass 352 

6. Seismic profiles over salt structure penetrated by borehole 66 352 

7. Seismic profiles over salt structure on lower continental slope showing 

positions of boreholes 112 and 123 353 

8-23. Plots: 

8. Grain-size distribution of sediments from core 114 versus depth 353 

9. Chlorinity of extracted pore water from core 114 versus depth 354 

10. Chlorinity of water-leached extract from dry sample of core 114 

versus depth 354 

11. Chlorinity of extracted pore water from core 66 versus depth 354 

12. Chlorinity of water-leached extract from dry sample from core 123 

versus depth 355 

13. Na/Cl ratio of water-leached extract from core 114 versus depth 355 

14. K/Cl ratio of water-leached extract from core 114 versus depth 356 

15. Percentage ~f carbonate of sediments from core 114 versus 

depth 356 

16. Ca/Mg ratio of carbonate fraction of sediments from core 114 versus 

depth 356 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 347 
8-23. Plots-Continued. 

17. Total organic carbon of dry sample from core 114 versus depth 357 

18. o13C in carbonate fraction versus carbonate content and o13C of organic 

fraction versus total organic content of sediments from core 114 357 

19. "Authigenic" iron and manganese of sediments from core 114 versus 

depth 358 

20. Total nitric acid-leached zinc of sediments from core 114 versus 

depth 358 

21. Ratio of total acid-leached zinc and "authigenic" zinc of sediments from 

core 114 versus depth 359 

22. X-ray diffraction patterns of <2-#Lm size fractions of sediments from core 

114 versus depth 359 

23. X-ray diffraction patterns of <2-#Lm size fractions of sediments from core 

66 versus depth 359 

24. Diagram showing metal distribution in ''typical'' nodule in cores immediately 

above salt cap rock 360 

348 Studies in Diagenesis 
Abstract 

Geochemical data from ten 300-m boreholes on the outer 

continental shelf and slope in the northwestern Gulf of Mexico document diagenetic changes caused by processes 

activated by salt intrusions. Two of four boreholes on the shelf 

and a borehole on the slope were drilled in sediments of 

Quaternary age that overlie salt structures. The other boreholes 

were drilled on the flank of salt structures, but also penetrated 

sediments of Quaternary age. In those sediments overlying salt, 

the abundance of expandable clay minerals (smectite) compared with that of nonexpandable clay minerals (illite) decreased with depth. Correspondingly, the total organic carbon 

and carbonate content and the 13C content of the organic 

material also decreased with depth. Clay transformations, 

caused by the heat from the salt, released a significant amount 

of intercellular water (200 liters per cubic meter). This water 

dissolved the leading edge of the salt stock, producing a brine 

and creating a residual of anhydrite between the sediments and 

the salt. The constituents of the brine reacted with the sediment above salt, causing further clay-mineral transformations. 

The sulfate in the brine reacted with the organic material, producing sulfides, carbonate species, and additional water. The 

additional water dissolved more salt and accelerated the 

diagenetic processes. Where the anhydrite cap intruded a layer 

of abundant organic material, such as a hydrocarbon accumulation, the diagenetic processes produced the limestone-gypsumanhydrite sequence that is common in the cap rock of many 

salt structures in the Gulf of Mexico. 

INTRODUCTION 

Sixty-eight cores were obtained by the Shell Development Corporation on the upper slope and shelf edge 

of the northern Gulf of Mexico. Ten of these cores were 

selected for geochemical analysis. The selection was made 

to have the broadest representative sample of the upper 

slope and shelf edge. From these 10 cores, 1,500 samples 

were analyzed for 30 elements by the semiquantative spectrographic method of the U.S. Geological Survey (Grimes 

and Marranzino, 1968). The most notable result was the 

discovery of anomalous concentrations of Zn, Pb, and 

Mn in the sediments above salt structures on the upper 

continental slope. The sediments above salt on the lower 

slope and those immediately adjacent to salt structures, 

however, had concentrations of these and other metals 

in the normal range for marine sediments. To define the 

cause of the anomalies, six cores were selected for further chemical and mineralogical analyses: three boreholes 

penetrated salt (two on the upper slope, and one on the 

lower slope), and three were on the flanks of salt intrusions (fig. 1). 

GEOLOGIC BACKGROUND 

The continental slope of the northern Gulf of Mexico is an area of moderate to high sedimentation rates 

(Holmes and Booth, 1980) (fig. 2). This sedimentologic 

pattern has prevailed throughout the latter part of the 

Quaternary Period although glacial episodes have caused 

local increases. The sediments are dominated by contributions from the Mississippi River and reflect the texture 

and mineralogy of its drainage basin. Much organic 

material is also incorporated within these deposits. Part 

of this material was derived from the watershed of the 

river and is dominated by plant material; however, a 

significant fraction was derived from the water column 

and has a high proteinaceous content. An index of the 

marine component of the organic material is the foraminiferal content of the sand-size fraction of the sediments 

MEXICO 

Figure 1. Location (solid circles) of cores used in study. Depth contours are in meters. 

Salt-Induced Diagenesis of Argillaceous Sediments 349 
(fig. 3), which is greater at the shelf edge and on the upper 

slope than on the central shelf. The oxygen minimum 

within the water column coincides with the sediment with 

the highest total organic content on the upper slope 

(fig. 4). The zone of reduced oxygen varies with sea level, 

oceanic circulation, and climatic changes, but the presence of the unoxidized organic matter in the upper-slope 

sediments suggests that the oxygen-poor zone has been 

about in the same position with respect to sea level 

throughout much of the Neogene Period. 

A thick salt layer exists under the thick wedge of 

clastic material in the northern Gulf of Mexico. Due to 

isostatic disequilibrium and differential loading, the salt 

has been squeezed into the overlying sediments. The 

piercement activity has progressed basinward with time; 

the youngest activity is presently along the outer part of 

the slope. Gussow (1968) argued that the initiation of salt 

diapirism was triggered by heat after a critical loading 

factor was reached. Based on the internal mineralogical composition and the structure of the salt bodies, he 

Figure 2. Rates of sedimentation (heavy contours) in the northern Gulf of Mexico. Rates 

were estimated by the 210Pb method, and are given in millimeters per year. Depths 

(light contours) are in meters. Arrows indicate direction of sediment transport. Data 

are from Holmes and Booth (1980). 

Figure 3. Percentage of foraminifera in sand fraction of surficial sediments in northern 

Gulf of Mexico. Depth contours are in meters. 

350 Studies in Diagenesis 
1-

z 

w 

u 

a: 

w 

D.. 

1.5-

~ 1.0 

z 

0 

CD 

a: 

() 

0.5-

a: 5 w 

1-

::J 4 -

~a: .w 

zo.... 3- W(f) 

(!)a: 

>-w XI- 2~ O::J 1 ::J 

....1 

~ 0 ~ 

0 

I I 

-

~-

-

I I 

I I 

/_ -

-

-

-

I I 

500 1000 

DEPTH. IN METERS 

Figure 4. Carbon percentage of surficial sediments and oxygen content of waters versus depth in northern Gulf of Mexico. Carbon data are based on 150 analyses of slope and outershelf cores. Oxygen profiles were made in summer. Oxygen 

minimum may be higher during this sampling period than in 

late winter and spring when floods in Mississippi River watershed increase productivity levels in the northern Gulf of Mexico (Holmes and Booth, 1980). Thus, maximum carbon 

percentage and oxygen minimum may be in better agreement 

during periods of maximum sediment influx. 

suggested that the emplacement took place rapidly at a 

temperature that exceeded 300 OC. This argument is 

strengthened by the presence of high thermal gradients 

associated with salt structures. Such a thermal environment 

must significantly affect the diagenetic processes in the 

sediments surrounding the intrusion. 

METHODS 

Core and sam pie selection 

Of the six cores selected for detailed analyses, the 

three associated with a salt structure in the northwestern 

Gulf of Mexico received the most attention (fig. 5) 

because of (1) their high lead and zinc contents, (2) the 

availability of good samples of cap rock, (3) the continuity of cores, and (4) the availability of pore-water data 

from the core on the crest of the dome. The borehole 

drilled on the crest of the dome, site 114 (water depth 

= 250 m), penetrated a granular anhydrite cap (Lehner, 

1969). Borehole 113 was drilled on the updip flank and 

borehole 115 on the downdip flank of the same structure (fig. 5). All cores had exceptionally high recoveries 

of 77, 88, and 70 percent, respectively. In the central Gulf, 

site 66 (water depth = 198m) was examined because it 

also penetrated salt (fig. 6), and although only a few 

samples were available for analysis, pore-water data were 

available (Manheim and Bischoff, 1969). The two remaining borings were made on a structure lower on the 

slope; site 123 was made on the crest of the structure, 

and site 122 was on the updip flank (fig. 7). These sites 

were selected because the sediments had been intruded 

by salt in the recent geologic past (Gussow, 1968; Lehner, 

1969). 

Water chemistry 

As mentioned previously, pore-water data collected 

immediately after the cores were taken were available for 

cores 114 and 66 (Manheim and Bishoff, 1969). An 

attempt was made to obtain complementary water chemistry on 600 samples from these cores by measuring the 

amount of water-soluble Ca2+, Mg2+, Na +, K +, and 

Cl-. These analyses were made by leaching a 20-g 

(gram) sample of finely powdered sediments with 500 mL 

(milliliters) of deionized water. The cations were determined by atomic absorption methods, whereas chlorine 

was determined by the standard Mohr titration technique. 

Although the water-leach procedure is subject to complete dissolution or partial dissolution of material not in 

the original pore water, comparison of ratios of ions in 

the water-leach extract and the original pore water (Manheim and Bischoff, 1969) suggests that this effect was not 

particularly significant for sodium and potassium. 

Solid-phase chemistry 

To assess the geochemical variations in the solid 

phases, the metal content of the "authigenic" phases (that 

is, ferromanganese coatings, carbonates, and adsorbed 

phases), the metal content of all nonsilicate phases, and 

the mineralogical composition of the clay fraction were 

determined. Atomic absorption methods were used to 

analyze 550 samples for "authigenic" Fe, Mn, Mg, Ca, 

Zn, and Pb. The samples were dried, ground in an agate 

mortar, and shaken for approximately 12 hours in a 50-

mL solution of 1 M hydroxylamine-hydrochloric and 

Salt-Induced Diagenesis of Argillaceous Sediments 351 
0 

I 

1000 2000 

I I 

Figure 5. Contour (dashed lines) map of top of salt structure in northwestern Gulf of 

Mexico showing relationship of boreholes 113, 114, and 115 (solid dots; fig. 1) to salt 

mass. Data from Lehner (1969). Sparker lines are track lines of geophysical survey ship. 

Contour interval, 25 meters . 

..------...---------,.-SEA LEVEL 

(/) 

a: 

w 

1-

w 

:::?i 

~ 

:i 

h: 

w 

c 

0 

1000 

2000 

30 KILOMETERS 

Figure 6. Seismic profile over salt structure penetrated by borehole 66 (see fig. 1 ). Lines represent 

major reflectors. 

25 percent (V /V) acetic acid. This procedure dissolves 

non-silicate ferromanganese and carbonate minerals 

and extracts the adsorbed trace elements, but it does 

352 Studies in Diagenesis 

not affect detrital silicates or sulfides (Chester and 

Hughes, 1967). Replicate analyses of 10 percent of the 

samples established that the error of this method was less 

than 10 percent. 

Atomic absorption was also used to measure these 

same elements released from the solid phases by hot, concentrated nitric acid. Here, a 5-g, finely ground sample 

was shaken in a heating block for 4 hours. This process 

dissolves the same phases as the preceeding process as well 

as sulfides, sulfates, and some detrital silicates (Holmes 

1977a). Clay mineralogy was determined by the method 

outlined by Carroll (1970). Some samples were also 

analyzed by J. C. Hathaway (U.S. Geological Survey, 

Woods Hole, Mass.). Selected samples were also analyzed 

for the amount of organic carbon and carbonate; the sediment texture was determined by procedures outlined by 

Griffins (1967). The carbon isotope analyses were performed by B. B. Bernard, Texas A&M University, College Station, Texas. 
KILOMETERS 10 

(/) 

a:: 750 

w 

~ 

w 

:E 

z 

:I: 

~ Salt 

a.. 

w 

0 

1500_. ____________________ ~----------------------~--~ 

Figure 7. Seismic profile over salt structure on lower continental slope showing positions of 

boreholes 122 and 123 (see fig. 1 ). Lines represent major reflectors. 

RESULTS 

The slope sediments, as determined by textural 

analyses, are silty clays (fig. 8). Throughout the sediment 

column the coarse fraction (greater than 62 mm) contained both foraminifera and detrital material. For the 

most part, foraminifera dominated except for those layers 

deposited as a single event (turbidites?). At the present 

time, a rich pteropod ooze is being deposited on top of 

the structure from which core 114 was taken (Holmes, 

1977a) and accounts for the greater amount of "sand" 

at the surface of the core. 

Pore-water chemistry 

Pore water extracted from the sediments overlying 

the salt structures increased in salinity linearly with 

depth. In core 114 (fig. 9), the salinity of the pore water 

in the sediments immediately above the cap rock 

approached halite saturation. Chloride values from the 

water-leach experiments (fig. 10) confirm the linear trend 

and further support the concept of linear solid-state 

diffusion of the chloride component suggested by 

Manheim and Bishoff (1969). In core 66, a similar trend 

was found but the chloride concentrations were much 

lower (fig. 11). The pore water at this site, however, 

contained more bromine and potassium than at site 114, 

suggesting an origin from a salt mass containing latestage evaporites. In contrast to these two sites, the 

chloride values from site 123 showed no clear trend 

300 

en ffi 325 1-

w 

:iE 

Z-

:i 

1-

fu 350 c 

375 

contact 

CAP ROCK 

400~--~----~--~~--~--~ 

0 20 40 60 80 100 

SEDIMENT TYPE, IN PERCENT 

Figure 8. Grain-size distribution of sediments 

from core 114 (see fig. 1) versus depth. 

(fig. 12). A computer-generated "best-fit" line through 

the data points suggests a linear trend. 

A plot of the Na/Cl molar ratio in the water-leached 

extract from site 114 against depth shows two trends 

Salt-Induced Diagenesis of Argillaceous Sediments 353 
Figure 9. Chlorinity of extracted pore water from core 114 (see fig. 1) 

versus depth. Diffusion curve is theoretical curve if solid-state diffusion 

were active. 

300 

(J) 

a: 325 w 

~ 

w 

~ 

~ 

::i 

~ 

ll.. 

w 350 0 

375 CONTACT 

CAP ROCK 

----------

• • 

\ 

\ 

\ 

400 \ 

0 10 20 30 40 50 

CHLORINITY, IN PER MIL 

Figure 10. Chlorinity ofwaterleached extract 

from dry sample of core 114 (see fig. 1) versus 

depth. Line is best-fit line. 

354 Studies in Diagenesis 

(J) 

a: 

w 

~ 

w 

575 

~ 600 z 

:i 

t 

w 

0 

625 

0 

I 

I 

NaCI saturation----f 

40 80 120 160 

CHLORINITY, IN PER MIL 

Figure 11. Chlorinity of extracted pore 

water from core 66 (see fig. 1) versus 

depth (modified from Manheim and 

Bishoff, 1969). NaCI saturation is given 

for comparison. 

(fig. 13). In the lower part of the sedimentary column, 

immediately above the cap rock, the Na/Cl ratio was 

similar to that of seawater (about 0.7). From this level to 

about the 325-m level, the ratio increased linearly and then 

was constant at about 1.0 to the sediment-water interface. 

The K/Cl ratio (fig. 14) was the reverse, being constant 

at about 0.06 from the top of the cap rock to the 325-m 

level, from where it increased more or less linearly to 
1800 

fJ'J 

a:. 

w t1i 2000 

~ 

~ 

:i 

li: 

w 

0 

2200 

\ .. \ 

.. \ • • .. • .., . 

. 

.. ' \ .. . . -

\ 

... \ . 

. . \ . . \ : . 

. \ , . ,. \ 

. . \ . . ,. . . ., 

, 

. \ 

' ., 

• . \ 

. . , '

•••• 1 . . \ 

. , . ' 

• 

• • 

• 

• 

_.!:!L.. • ' •• LP e• ~ \ 

I ·'· J·• ' •• • • 

4 6 

• 

8 

• • 

CHLORINITY, IN PER MIL 

• 

10 

Figure.12. Chlorinity of water-leached extract 

from dry sample from core 123 (see fig. 1) versus 

depth. Depth of transition is shown between 

upper Pleistocene (UP) and lower Pleistocene 

(LP) as reported by Lehner (1969). Dashed line 

is best-fit regression line. 

the top of the sediment column. Too few samples were 

examined from core 66 to define trends, and in core 123 

and cores from the flanks of the salt intrusions no changes 

in these ratios with depth were noted. 

The carbonate fraction of the sediment at site 114 

ranged from 1 percent in the lower portions of the 

sedimentary column to more than 20 percent at the 

sediment-water interface (fig. 15). The shape of the curve 

closely parallels the K/Cl curve. The Ca/Mg ratio of the 

carbonate fraction also parallels the K/Cl ratio, increasing from 2 in the lower portion of the sediment column 

to 20 at the top (fig. 16). The organic carbon content, 

an index of the total organic material in the sediments, 

ranged from 0.8 to 1.6 percent and had a distribution 

similar to that of the carbonate component (fig. 17). 

The carbon in the carbonate fraction and in noncarbonate material was isotopically lighter in the sediments 

that had a minimum of each component (fig. 18). Thus, 

both fractions were lighter isotopically in the lower part 

275 

fJ'J 

ffi 325 1-

w 

~ 

~ 

I •• ~N~/CI ·= e m I 

I 

seawater I 

I ' ... 

• • • • 

I 

,:.· 

I • 

I·, •• 

Na/CI 

• • 

Figure 13. Molar Na/CI ratio of the waterleached extract from core 114 (see fig. 1) versus 

depth. Vertical dashed line is molar ratio of 

Na/CI of normal seawater. Horizontal dashed 

line is contact between sediment and cap rock. 

of the sediment column than near the sediment-water 

interface. 

The iron content of the sediments increased linearly 

with depth from a low of 0.2 percent at the sedimentwater interface to 1.6 percent in the sediments immediately above the cap rock. Manganese was nearly constant throughout the core, but increased logarithmically 

in the zone immediately above the cap rock (fig. 19). 

Frank T. Manheim (U.S. Geological Survey, Woods 

Hole, Mass., oral commun., 1982) reported that the pore 

water immediately above the cap rock contained 90 ppm 

(parts per million) manganese, coinciding with the high 

manganese concentrations in the sediments. The maximum total zinc content of the sediments was near the 

325-m level and averaged 1 ,000 ppm. Some samples 

retained more than 2,000 ppm zinc (fig. 20). The ratio 

of hydroxylamine-hydrochloride-soluble zinc to the 

amount of zinc soluble in "hot" nitric acid was minimum 

near the 325-m level, suggesting that the zinc in this zone 

was present in a stable form, probably as a sulfide. 

Mineralogy 

The mineralogy of the reactive minerals, the clay 

minerals, changed with depth at two sites above the salt 

Salt-Induced Diagenesis of Argillaceous Sediments 355 
300-

en 

a: 325- w 

1-

w 

~ 

~ 

:I: 

t w 350- c 

•• 

I • 

• • •• • • 

• •• • • 

• • 

275~--~~~--~~~--~,--~.~~-.~~---,1----~,~ 

•• • • 

300 fen 

ffi 325 r--- 1-

w 

~ 

~ 

::i 

t 

~ 3501-

r--- K/CI 

1 -· in seawater 

I 

I 

I 

• • ,. • • 

• • ,. 

• • •• • • ••• 

' 

• 

• I • : • 

• I •• 

•• • 

• • -

-

-

• 

375r- - t--- _! ------!_ _C~P_RQ.f.!S_ _.!. CONTACT 

400~--~~----~~--~'----'~--~~--~~----~ 0.00 0.04 0.08 0.12 ~ K/CI 

Figure 14. Molar KICI ratio of the water-leached extract from 

core 114 (see fig. 1) versus depth. Vertical dashed line is molar 

K/CI ratio of normal seawater. Horizontal dashed line is contact 

between sediment and cap rock. 

275-.------------------------------~ • 

•• • 

•• 

••• 

• ••• 

• 

• •• •• • 

• 

-1000 

f- 1100 

f-. 1200 

en 1- a: w w w 1- 1.1.. w 

~ ~ 

::i ~ 

1- ::i a.. 

w 1-

a.. c w 

c 

300-

325-

350- • • 

• 

• 

.. 

•• 

• 

• 

• 

• • 

• • • 

• • 

• 

• -1000 

• 

• 

f-.1100 

1-1200 

•• 

375- • - .-• _._--- ------- 375- ---· - - ._ - - - - - ---

1-1300 f-..1300 

CAP ROCK 

CAP ROCK 

400~----------~~----------~~~--~ 400~---------~,--------~,~------~ 

0 10 20 0.4 0.8 1.2 1.6 

CARBONATE. IN PERCENT TOTAL ORGANIC CARBON. IN PERCENT 

1-

w 

w 

1.1.. 

~ 

::i 

t 

w 

c 

Figure 15. Carbonate percentage of sediments from 

core 114 (see fig. 1) versus depth. Horizontal dashed line 

is contact between sediment and cap rock. 

Figure 16. Total organic carbon of dry sample, in weight percent, from core 114 (see fig. 1) versus depth. Horizontal dashed 

line is contact between sediment and cap rock. 

356 Studies in Diagenesis 
2 8 14 

Ca/Mg 

20 26 

Figure 17. Molar Ca/Mg ratio of carbonate fraction of 

sediments from core 114 (see fig. 1) versus depth. Horizontal dashed line is contact between sediment and cap 

rock. 

on the upper slope (figs. 22 and 23). At both sites (114 and 

66), the amount of the 10-A phases increased with respect 

to those clay minerals with higher basal spacings. In core 

114, chlorite was detected in a sample immediately above 

the cap rock (J. Hathaway, U.S. Geological Survey, Woods 

Hole, Mass., written commun., 1974). These mineralogical 

changes are common in deeply buried Tertiary sediments 

of the Gulf Coast (Burst, 1969), but have not been previously reported at these shallow depths. 

In the sediments immediately above the salt intrusion on the upper slope, nodular bodies are common 

(fig. 24). Most of these bodies consist of a rust-colored 

rind with a central core of buff to white crystals. From 

X-ray diffraction examination, the central core contains 

bassanite (CaSO 4 ·1/2H2

0), jarosite (KFe3(SO Jz(OH)J, 

rhombic sulfur, and a small amount of pyrite. Semiquantitative elemental analyses of these minerals and the rind 

indicate abnormal amounts of Mo, Co, and Mn with 

respect to the surrounding sediments. 

DISCUSSION AND CONCLUSIONS 

Diagenic processes act throughout a large temporal 

spectrum, but for conceptual convenience diagenesis is considered to occur either "early" or "late". These terms are 

necessarily vague due to the complex nature of conditions 

that must be present to produce a chemical or mineralogical 

change in the sediment. 

A • • • 

• 

0 • 

• u 

rc 

-1 • • • 

-2 

4 8 12 14 

CARBONATE, IN PERCENT 

B 

• 

-23 • • 

• u • • • • rc • 

-22 • • 

• 

-21 

0.6 0.8 1.0 

ORGANIC CARBON, IN PERCENT 

Figure 18. A, o13C of carbonate fraction versus carbonate 

percentage of sediments from core 114. 8, o13C of organic 

fraction versus total organic content of sediments from core 

114 (see fig. 1). The more negative the number, the lighter 

the carbon mass (that is, the less 13C in sample). 

The degradation of organic material in the sediment 

is one of the most important factors determining the 

environmental conditions of early diagenesis. The complex 

nature of the organic substances which make up the total 

organic content affects their rate of decomposition; for 

example, proteinaceous substances oxidize quicker than 

cellulosic substances. Organic oxidation, however, changes 

the oxidation-reduction potential, the alkalinity of the pore 

water, the sulfate concentration, and the oxygen content 

of a sedimentary environment so that the degree of completeness of the organic degradation directly affects the 

diagenetic processes. Inasmuch as oxygen is necessary for 

this reaction, the limiting factor in the degradation process is the availability of this element. In the marine subaqueous environment, oxygen is supplied by the following 

species (in order of importance): 

0 2 > N03 > MnO, FeOOH > S04 (Berner, 1971). 

In most sediments, free oxygen and the oxygen bound as 

nitrate are rapidly consumed by oxidation. The oxides of 

manganese and iron react slowly, which leaves the sulfate 

concentration as the limiting factor. This species plays 

a dual role, supplying oxygen for oxidation of organic 

Salt-Induced Diagenesis of Argillaceous Sediments 357 
300 

en 

a: 325 w 

1-

w 

~ 

~ 

;£ 

1-

0.. 

w 350 c 

375 

CAP ROCK 

o.o 0.2 0.4 0.6 0.8 1.0 1.2 1.4 1.6 

IRON AND MANGANESE CONTENT. IN PERCENT OF SEDIMENT 

Figure 19. "Authigenic" iron (plus sign) and manganese (solid circle), 

in percent, of sediment from core 114 (see fig. 1) versus depth. Horizontal dashed line is contact between sediment and cap rock. 

358 Studies in Diagenesis 

275-r-----------------------------------. 

300-

en ffi 325- 1-

w 

~ 

~ 

:i 

t 

w 350- c 

• • •• • • • • • •• • • • •• •• .. • 

••• 

• • • • 

• • ' 

I 

I • 

• 

• 

• 

375- '· ------------.... -- _!._- • 

• • CAP ROCK 

-1000 

-1100 

-1200 

-1300 

400~------------~,-----------,~------__. 

0 1000 2000 

ZINC CONTENT OF SEDIMENT. 

IN PARTS PER MILLION 

1-

w 

w 

L.L. 

~ 

::£ 

1-

0.. 

w 

c 

Figure 20. Total zinc (leached with hot nitric acid) of sediments from 

core 114 (see fig. 1) versus depth. Horizontal dashed line is contact 

between sediment and cap rock. 
275-r----------------.-------------~ 

.. • • • • • • • -1000 

• • 300-

. ' • • • •• 

••• .. -1100 

. , • ~ 

w 

w 

••• u.. 

• ~ 

• ~ • ~ 

a. 

•• w 

c • -1200 • 

(/) 

a: 

w 325- ~ 

w 

::?! 

~ 

:I: 

~ 

a. 

w 350- c 

• 

-1300 

400~----p-.---~.--~.~---P-.---~.--~ 

0 1 .0 2.0 3.0 4.0 5.0 6.0 

Figure 21. Ratio of "total" acid-leached zinc (Zn ) and "authigenic" zinc (leached with hydroxylamine-hydrochloride-acetic 

acid) (Zn5

) of sediments from core 114 (see fig. 1) versus depth. 

Horizontal dashed line is contact between sediment and cap 

rock. 

0 

14A 

575 (/) 

a: 

w 

~ 

w 

::?! 

~ 

~ 600 ~ 

a. 

w 

c 

625 

compounds and contributing sulfur for formation of the 

sulfides. In the sediments of the Gulf of Mexico slopecharacterized by a high rate of sedimentation, a high organic content, and a low oxygen supply-organic decomposition is slow, and does not go to completion. The extent 

of the degradation is not known, but the coexistence of 

sulfides and unoxidized organic material is evidence that 

all reactive oxygen has been consumed. The sulfide production resulting from the reduction of sulfate increases 

the concentration of sulfides of metals such as Fe, Zn, 

and Mo in the sediments. Local organic-rich zones, such 

as burrows, also concentrate metals and produce "hot 

spots" such as the arsenopyrite burrow-filling found 

300 

~ 325 w 

~ 

w 

::?! 

~ 

~ 

t:: 350 w 

c 

375 

MIXED 

40 30 20 10 

020, CuKa RADIATION 

Figure 22. X-ray diffraction patterns of <2-~-tm size fractions 

of sediments from core 114 versus depth. MIXED represents 

clay minerals composed of random interstratification of 

smectite and illite. 1 0-A peak probably represents illitic clay. 

Increase in intensity of quartz peak with depth is due to 

increase in crystallinity of quartz and increase of relative 

abundance of illite, which also has a strong diffraction peak 

in this region. Shorter pattern is paired with longer pattern 

below it and is of same sample that was glycolated to expand 

smectite to 17 A. 

40 30 20 10 

020, CuKa RADIATION 

0 

Figure 23. X-ray diffraction patterns of <2-p.m size fractions of sediments from 

core versus depth. Material is probably smectite as it expands to 17 A on glycolation. 10-A material is probably illite. Quartz refers to 3.3-A peak, which is due 

to both quartz and illite in sample. Shorter patterns are glycolated versions of 

sample pattern following. 

Salt-Induced Diagenesis of Argillaceous Sediments 359 
Iron oxide 

PERCENT 

ellow clay 

Iron oxide 

Green clay 

Figure 24. Diagram and metal distribution in "typical" nodule in cores immediately above salt 

cap rock. Calcium and iron are reported in percent; remaining elements are reported in parts per 

million. Data were produced by semiquantitative method of Grimes and Marranzino (1968). 

in one of the cores examined in this study on the upper 

slope. 

Whereas early diagenesis involves degradation of 

organic material, precipitation of cements, and formation of authigenic minerals, the major change in late-stage 

diagenesis is alteration of the original mineral constituents. In argillaceous sediments of the Gulf Coast, this 

change is chiefly the conversion of expandable clays 

minerals (smectite) to nonexpandable clay minerals (illite). 

This process has been discussed by many investigators and 

involves a stepwise conversion with the threshold of each 

step controlled predominantly by temperature; pore-water 

chemistry and pressure are secondary factors (Burst, 

1969; Blattner, 1974). 

Whether or not temperature was responsible for the 

emplacement of the salt domes in the area of the present 

study, as suggested by Gussow (1968}, salt structures are 

known to be ''hot'' compared with the surrounding sediment. Oxygen isotope data of included water in such salt 

bodies confirm the fact that the salt was "hot" at some 

time during its history (Kreitler and Dutton, 1983). 

O'Brien and Lerche (1984) modeled the temperature 

distributions around salt structures and demonstrated that 

the hottest part of a salt structure is that part surrounding the leading edge. Thus, halo kinesis in the Gulf Coast 

sediments may have produced a sufficient heat (> 150 °C) 

to initiate clay-mineral transformations. 

Burst (1969) and Perry and Hower (1972) demonstrated that expandable clay minerals are dewatered in 

three major stages: (1) the pore water of the sediments 

and interlayer water in excess of two water interlayers is 

removed by compaction; (2) the next to last interlayer of 

water is expelled, a process that requires heat (80 °-

1000C); and (3) the remaining water is expelled and the 

layers realign, an extremely slow process in normal 

geologic setting that depends greatly on the thermal 

360 Studies in Diagenesis 

history of the area. The first step reduces the original 

interlayer water content by 50 percent, and the second 

step removes another 25 percent. Bruce (1983) calculated 

that these two processes yield about 200 liters of 

diagenetic water per cubic meter of sediment. 

The effect of major pore-water cations on the dehydration and clay-mineral conversion was investigated by 

Blattner (1974}, who demonstrated that at elevated 

temperatures these transformations are suppressed by the 

presence of calcium and magnesium but accelerated by 

the presence of potassium. A brine formed by the dissolution of salts rich in late-stage evaporities would therefore 

be enriched in potassium and associated brine elements, 

and, thus, would accelerate the clay-mineral transformation. Whereas heat is the dominant factor in the claymineral change, the composition of the pore water determines the rate of the transformation. 

The diffusion of a brine created by the dewatering 

of clay minerals and the dissolution of salt toward the 

sediment-water interface would be influenced by the heat 

of the salt mass and aided by the fracturing caused by 

halokinesis. Lateral diffusion would be inhibited by the 

impervious nature of the sediments, which is not affected 

by the rising heat from the salt mass. Manheim and 

Bischoff (1969) suggested that this diffusion can be 

described by Fick's first law (steady-state diffusion), 

which predicts a concentration distribution linearly 

decreasing from the salt mass to the sediment-water interface. The chloride distributions measured in cores 114 and 

66 were linear and coincided with the distribution 

predicted by the model. The K/Cl and Na/Cl ratios, 

however, did not. With respect to chloride, the relative 

level of sodium increased in the sediments below the 

325-m level in core 114, whereas potassium increased 

with respect to chloride. These trends were apparently 

the result of the clay-mineral transformations, wherein 
potassium was fixed and sodium was expelled to form an 

illitic clay. This trend suggests that the transformation 

from stage two to three, as described by Burst (1969), 

occurred in the sediment near the 325-m level. 

The transformation of the clay minerals and subsequent brine formation also had a significant effect on the 

organic and inorganic species disseminated throughout 

the sediments. Organic material that had been adsorbed 

onto the clay minerals was expelled, and as the transformation proceeded, it was "flushed" toward the sedimentwater interface. The organic material that remained was 

oxidized with the sulfate in the ascending brine. 

Simplified, the reaction in the sediments was as follows: 

where e is the activation energy supplied to this system 

by sulfate-reducing bacteria that are common in these 

sediments and that apparently remain inactive for considerable lengths of time (Freely and Kulp, 1957). White 

and others (1963) suggested that most intraformational 

brines such as these are slightly acidic; therefore, as long 

as the organic material and (or) the sulfate in the pore 

water is dispersed, no carbonate should precipitate, and 

the carbonate in the sediment should be attacked. The 

chemical reaction, however, produced more water, thereby 

diluting the pore water and creating a steeper concentration gradient that accelerated the diffusion processes. As 

a result, more sulfate was introduced into the sediment, 

which led to further oxidation of any remaining organic 

material. Under these conditions, only the most stable 

organic and carbonate species remained-namely, plant 

material for the organic species and high-magnesium 

dolomite for the carbonates. The o13C content of the 

organic material was highly negative, a common 

characteristic of plant material. The Ca and Mg analyses 

suggest that the carbonate material was high in Mg. Thus, 

the data suggest that the ascending fluids created an 

oxidizing environment above the salt structure. If, however, 

the organic material that was "flushed" during claymineral transformations and other organic material were 

concentrated due to their inability to migrate further with 

a decrease in the relative organic to sulfate concentration, 

sulfide should have been concentrated in the sediments. 

Holmes (1977b) and Holmes and others (1974) 

demonstrated that metal sulfides in an oxidizing saline 

environment will dissolve, and metals such as zinc and 

lead will form stable chloride complexes. The presence of 

sulfate nodules immediately above the salt structures suggests that the ascending fluids were oxidizing. The form 

and trace-metal constituents of the nodules suggest that 

these originally were infauna burrows that were mineralized in a reducing environment caused by decaying 

organic material. Subsequently, the ascending sulfaterich brines oxidized many of the metal sulfides and 

formed the present mineral assemblage. As the metal-rich 

ascending brine entered the zone in which sulfide production was prevalent, the metals were again precipitated. This 

dissolution and precipitation formed an enriched zone in 

the sediments that marked the zone in which the most 

diagenetic activity caused by halokenesis occurred. At site 

114, the concentration of zinc and the "authigenic''/ 

"sulfide" ratio near the 325-m level suggest that at this 

level such chemical processes were prevalent. The organic 

content, the carbonate content, and the cations in the pore 

water also changed at this level. All these changes suggest that the 325-m level was the diagenetic front at which 

the major chemical transformations took place. 

SUMMARY 

In the Gulf Coast geosyncline, the intrusion of salt 

into the argillaceous sediments had a significant effect on 

the diagenetic processes. The heat generated at the saltsediment interface dehydrated the clay minerals. The 

expressed water flushed dissolved organic materials toward 

the sediment-water interface and produced brine having an 

increased sulfate concentration by dissolving the salt and 

leaving an anhydrite residue; it also changed the clay 

mineralogy in the sediment column over the salt mass. The 

chemical reactions between the pore water and the 

sediments produced a diagenetic sulfide-rich front. Within 

this zone metals such as zinc and lead were concentrated; 

the metals were derived from the oxidized region below and 

transported into this sulfide zone as chloride complexes. 

If the salt mass-with its anhydrite residue on the 

surface-had been intruded into a sedimentary unit rich 

in organic material, such as a hydrocarbon accumulation, 

common in the Tertiary of the Gulf of Mexico, the resulting 

chemical environment would have been such that limestone 

would have formed above the anhydrite of the salt mass. 

According to the chemical reaction presented herein, one 

product of the reaction, in addition to the formation of 

calcium carbonate, is water. One cubic meter of anhydrite 

can result in the production of 400 liters of water when 

reacting with organic material. This reaction was most 

likely at the carbonate-anhydrite interface and probably 

caused the hydration of anhydrite to gypsum. The result 

of such a process was a limestone-gypsum-anhydrite 

sequence-a sequence commonly referred to as cap rockfound on many salt domes throughout the Gulf Coast. 

REFERENCES CITED 

Berner, R. A., 1971, Principles of chemical sedimentology: New 

York, McGraw-Hill, 240 p. 

Blattner, C. A., 1974, The interaction of clay minerals with distilled 

water and saline solutions at elevated temperatures: Binghamton, New York, State University of New York at Binghamton, Ph.D. dissertation, 264 p. 

Salt-Induced Diagenesis of Argillaceous Sediments 361 
Bruce, C. H., 1983, Relation of illite/smectite diagenesis and 

development of structure in the northern Gulf of Mexico: 

American Association of Petroleum Geologists Bulletin, v. 

67, p. 432. 

Burst, J. F., 1969, Diagenesis of Gulf Coast clayey sediments 

and its possible relationship to petroleum migration: 

American Association of Petroleum Geologists Bulletin, v. 

53, p. 73-93. 

Carroll, D., 1970, Clay minerals-A guide to their X-ray identification: Geological Society of America Special Paper 126, 

80 p. 

Chester, R., and Hughes, M. J., 1967, A chemical technique 

for the separation of ferro-manganese minerals, carbonate 

minerals, and adsorbed trace elements from pelagic 

sediments: Chemical Geology, v. 2, p. 249-262. 

Freely, H. W., and Kulp, J. L., 1957, Origin of the Gulf Coast 

salt dome sulfur deposits: American Association of 

Petroleum Geologists Bulletin, v. 41, p. 1802-1853. 

Griffins, J. C., 1967, Scientific methods in the analyses of 

sediments: New York, McGraw-Hill, 508 p. 

Grimes, D. L., and Marranzino, A. P., 1968, Direct-current 

arc and alternating current spark emission spectrographic 

field methods for semiquantitative analyses of geologic 

materials: U.S. Geological Survey Circular 591, 6 p. 

Gussow, W. C., 1968, Salt diapirism; importance of temperature 

and energy source of emplacement, in Braunstein, J., and 

O'Brien, G. D., eds., Diapirism and diapirs: American 

Association Petroleum Geologists Memoir 8, p. 16-52. 

Holmes, C. W., 1977a, Clay mineralogy, suspended and sea 

floor sediments, in Berryhill, H. L., Jr., and Trippet, A. 

R., eds., Environmental studies, south Texas outer continental shelf: Bureau of Land Management report, 

available from U.S. Department of Commerce National 

362 Studies in Diagenesis 

Technical Information Service, Springfield, VA 22161, as 

Report PB289-144 AS. 

___ 1977b, Effects of dredge channels on trace-metal migration in an estuary: U.S. Geological Survey Journal of 

Research, v. 5, p. 243-251. 

Holmes, C. W., and Booth, J. S., 1980, Impact of the Holocene 

transgression on depositional environments of the northern 

Gulf of Mexico continental margin: American Association 

of Petroleum Geologists Bulletin, v. 64, p. 724. 

Holmes, C. W., Slade, E. A., and McLerran, C. J., 1974, Migration and redistribution of zinc and cadmium in marineestuarine system: Environmental Science and Thchnology, 

v. 8, p. 255-259. 

Kreitler, C. W., and Dutton, S. P., 1983, Origin and diagenesis 

of caprock, Gyp Hill and Oakwood salt domes: Texas Bureau of Economic Geology Investigations Report 131, 58 p. 

Lehner, P., 1969, Salt tectonics and Pleistocene stratigraphy on 

continental slope of northern Gulf of Mexico: American 

Association Petroleum Geologists Bulletin, v. 53, p. 2431-

2479. 

Manheim, F. T., and Bischoff, J. L., 1969, Geochemistry of pore 

waters from Shell Oil Company drill holes on the continental 

slope of the northern Gulf of Mexico: Chemical Geology, 

v. 4, p. 63-82. 

O'Brien, J. J., and Lerche, 1., 1984, The influence of salt domes 

on paleotemperature distributions: Geophysics, v. 49, 

p. 2032-2042. 

Perry, E. A., and Hower, J., 1972, Late stage dehydration in 

deeply buried pelitic sediments: American Association of 

Petroleum Geologists Bulletin, v. 56, p. 2013-2021. 

White, D. E., Hem, J. D., and Waring, G. A., 1963, Chemical 

composition of subsurface water: U.S. Geological Survey 

Professional Paper 440-F, 67 p. 
Sodium-Potassium lon Exchange During 

Smectite Diagenesis-A Theoretical Discussion 

By DENNIS D. EBERL, U.S. Geological Survey, Box 25046, Denver Federal 

Center, MS 404, Denver, Colorado 80225 

CONTENTS 

Abstract 364 

Introduction 364 

Preferential potassium selectivity by illite 364 

Effect of increasing temperature and pressure 365 

Effect of increasing pore-water salinity 365 

Effect of increasing smectite layer charge 365 

Effect of decreasing interlayer water content 366 

Concentration of potassium in illite layers 366 

Conclusions 367 

References cited 368 

FIGURES 

1-4. Plots showing: 

1. Effect of increasing temperature and pressure on exchange free energy 

for sodium-potassium exchange 365 

2. Free energy of exchange as a function of pore-water molality 366 

3. Free energy of exchange for sodium-potassium as a function of 

smectite-layer charge 366 

4. Free energy of exchange for sodium-potassium exchange as a function 

of interlayer molality 367 

5. Plot showing free energy of exchange for sodium-potassium exchange 

calculated from Coulomb's law 367 

F. A. Mumpton, ed., 1986, Studies in Diagenesis, U.S. Geological Survey Bulletin 1578 363 
Abstract 

When smectite reacts to form illite during burial 

diagenesis, the interlayer cation changes from dominantly 

sodium to almost entirely potassium. Approximate thermodynamic calculations based on a model that treats the smectite interlayer as a concentrated electrolyte solution suggest that 

this significant change in interlayer composition is related to 

a dehydration of potassium by interlayer electrical forces in 

response to an increase in smectite layer charge with burial 

depth. The smaller dehydrated potassium ion is strongly 

selected compared to the larger hydrated sodium ion in the 

clay interlayer because adsorption energy is inversely proportional to ionic radius, as shown by Coulomb's law. 

INTRODUCTION 

The work of Perry and Hower (1970) and Weaver 

and Beck (1971), among others, has firmly established that 

smectite reacts to form illite during burial diagenesis. 

Assuming a transformation mechanism, this reaction proceeds via the substitution of aluminum for silicon in the 

tetrahedral sheet of smectite, thereby generating a negative 

charge on the 2:1 layers that is balanced by an increase 

in interlayer cations (Hower and others, 1976). When the 

layer charge becomes sufficiently negative, smectite 

interlayers collapse around the interlayer cations and expel 

interlayer water, thereby becoming illite interlayers. 

This reaction proceeds in a sodium-enriched pore 

water in which interlayer exchange sites of smectite are 

dominated by sodium. Yet, when smectite interlayers collapse to form illite, potassium is concentrated in the illite 

interlayers to the exclusion of sodium. Some simple 

calculations support this apparent anomaly. An exchange 

reaction betweem a sodium smectite (NaX) and a 

potassium smectite (KX) is: 

NaX + K+~ KX + Na+, (1) 

where Na + and K + are each one equivalent of sodium 

and potassium ions in solution. The equilibrium constant 

(Kex) for this reaction is: 

K= ex 

(KX] [Na+] 

[NaX] [K+]" (2) 

For the following approximate calculation, concentrations will be used to approximate activites. The standard free energy of exchange (..1G~) can be calculated 

from 

..1G ~ = -RT In Kex, (3) 

where R is the gas constant and Tis absolute temperature. 

Assuming a [KX] ratio for illite of 77 (Interlake, 

[NaX] 

Montana illite, from Hower and Mowatt, 1966) and a 

364 Studies in Diagenesis 

[Na+] 1 s.- pore water -- ratio of 250 (average va ue 10r 

[K+] 

Cretaceous oil field waters, from Collins, 1975), and 

neglecting the effects of ions other than Na + and K + on 

exchange constants, the free energy of exchange can be 

calculated from equations 2 and 3: 

..10° = -24 kJ/eq (kilojoules per equivalent). ex 

The average ..1G 0 measured at room temperature for ex . data summarized in Eberl (1980) for sodium-potassmm 

exchange on smectite is: 

..1G~ = -1.7 kJ/eq, 

which is more than an order of magnitude too small to 

account for the interlayer chemistry of illite. 

The present paper uses approximate thermochemical calculations, based on a model that treats the smectite inter layer as a concentrated electrolyte solution, to suggest a mechanism that leads to the concentration of 

potassium in illite. 

PREFERENTIAL POTASSIUM SELECTIVITY 

BY ILLITE 

Something happens during the formation of illite 

from smectite to increase the selectivity of the interlayer 

for potassium relative to sodium. Changes that occur 

during burial diagenesis that might affect sodiumpotassium selectivity with burial depth are: (1) increasing 

temperature and pressure; (2) increasing salinity of pore 

water; (3) increasing smectite layer charge; and (4) decreasing interlayer water content. 

Approximate exchange free energies for equation 1 

under these varying diagenetic conditions can be calculated from a modified version of the equation proposed 

by Cruickshank and Meares (1957) for ion-exchange 

resins, and later applied by Eisenman (1961, 1962) to glass 

electrodes and by Eberl (1980) to clay minerals. Assuming that interlayer water is at the same pressure as the pore 

water, the equation for exchange-free energy is: 

..10~ = 2RT(<J>- In')' ±m±)NaX- 2RT(<J>- In 'Y ±m±)KX 

(4) - 2RT(<I> -In')' ±m±)NaCI + 2RT(<J> - In')' ±m±)Kcl' 

where <J> is the practical osmotic coefficient, m ± is the 

mean ionic molality of the smectite interlayer water 

(assuming that the negative interlayer surfaces are the 

anion, the interlayer cation is the cation, and the interlayer 

water is the solvent), m+ is the mean ionic molality of 

the pore water, and 'Y + is the mean ionic-activity coefficient for the subscripted species. The subscripted species 

NaX and KX are each one equivalent of sodium and 
potassium smectite, and NaCl and KCl refer to one 

equivalent of sodium and potassium chloride in the pore 

water. 

The smectite interlayer can be modeled as a halide 

anion having an electric-field strength equivalent to that 

of the interlayer (Eberl, 1980). Hence, equation 4 can be 

solved for room-temperature conditions from the values 

of cp and m ± that were presented as a function of 

molality and alkali halide chemistry by Robinson and 

Stokes (1959), and for diagenetic conditions from similar 

data presented by Helgeson and others (1981, figs. 152, 

154, and 157). 

Effect of increasing temperature and pressure 

The effect of increasing temperature and pressure 

on sodium-potassium selectivity is shown in figure 1. 

Calculations for three interlayer molalities (m) are plotted. Other diagenetic variables are constant: smectite 

layer charge is constant at the equivalent anionic radius 

..... 0 

z 

w 

...J 

~ 

~ -0.1 w 

a: 

w 

Q. 

ffi -0.2 ...J 

::J 

0 

...., 

0 

...J 

~ -0.3 

~ 

TEMPERATURE, IN DEGREES CELSIUS 

Figure 1. Effect of increasing temperature and pressure un 

free energy of exchange for sodium-potassium exchange at 

three interlayer water contents. Pressures correspond to 

liquid-vapor equilibrium for temperatures given. 

of the chloride ion, which corresponds to a layer charge 

of -0.57 equivalents per one-half unit cell, and the pore 

water is dilute. The curves were calculated from equation 4 using the osmotic and activity coefficients found 

in Helgeson and others (1981). 

The curves show that the effect of increasing 

temperature and pressure on sodium-potassium selectivity 

will be small and will tend to increase interlayer preference 

for sodium rather than for potassium. In other words, 

the free energy change for equation 1 becomes more 

positive with increasing temperature and pressure. A 

minimal dependence on temperature (7) and pressure (P) 

also is predicted from the small enthalpy (tli/~ and 

volume(~ V") change for the exchange reaction (as was 

discussed by Helfferich, 1962), according to Le 

Chatelier's principle, and to the following equations: 

(5) 

( 

d In Kex ) = - tli/0 

dP T RT . (6) 

Thus, increasing temperature and pressure alone cannot 

be responsible for the remarkable increase in potassium 

selectivity by illite interlayers. 

Effect of increasing pore-water salinity 

Considering only sodium and potassium, increasing pore-water salinity also will tend to increase sodium 

selectivity (fig. 2). In this figure, temperature is 25 OC, 

pressure is atmospheric, and smectite layer charge is constant at the equivalent anionic radius of the chloride ion . 

Curves were calculated from equation 4 using the 

osmotic- and activity-coefficient data in Robinson and 

Stokes (1959). The trend of increasing free energy of 

exchange with increasing pore-water molality indicates 

that an increase in pore-water salinity with burial depth 

cannot be the answer. 

Effect of increasing smectite layer charge 

As was mentioned previously, layer charge increases 

during burial diagenesis by the substitution of aluminum 

for silicon in the tetrahedral sheet. The theoretical effect 

of increasing smectite layer charge is shown in figure 3. 

In this figure temperature is constant at 25 OC, and 

pressure is constant at 1 atmosphere; interlayer molality 

is 3.5 m, and the pore water is dilute. The curve was 

calculated from equation 4 using equivalent anionic radii 

for chloride, bromide, and iodide ions to represent 

various layer charges (see Eberl, 1980). Osmotic and 

Sodium-Potassium Exchange During Smectite Diagenesis 365 
activity coefficients used in this calculation are from 

Robinson and Stokes (1959). The curve indicates that 

increasing smectite layer charge tends to increase sodium 

adsorption, not potassium adsorption, for the range of 

layer charges shown. 

Effect of decreasing interlayer water content 

The effect of decreasing interlayer water content is 

shown in figure 4. This curve was calculated from 

equation 4 using the data in Robinson and Stokes (1959), 

assuming that the temperature is 25 °C, the pressure is 

atmospheric, the pore water is dilute, and the layer charge 

is constant at the equivalent anionic radius of the chloride 

ion. The data show that a decrease in the inter layer water 

content (that is, an increase in interlayer molality) tends 

to increase potassium adsorption. Here is a possible 

explanation for increased potassium selectivity during 

diagenesis. The effect for hydrated interlayers (fig. 4) is 

minor, but A.G:; does vary in the direction of increasing 

potassium selectivity. Furthermore, the effect can be large 

if interlayer potassium ions dehydrate completely. 

1-

z 

w 

....1 

~ s 

0 

w 

a: 

w 

Q, 

(/) 

w 

....1 

:::::1 

0 

.., 

g 

i2 

z 

;:o 

~ 

a: 

w 

z 

w 

w 

w 

a: 

LL. 

w 

~ 

z 

<( 

:I: 

u 

~ 

NaX+K+~ KX+Na+ 

lnterlayer molality = 0.1 m 

_,0 2 

MOLALITY OF PORE WATER 

Figure 2. Free energy of exchange for sodiumpotassium exchange as function of pore-water 

molality for five interlayer water contents. 

366 Studies in Diagenesis 

4 

1-

z 

w 

...J 

~ 5 

0 

w 

a: 

w 

0. 

0 

NaX + K+ ~ KX + Na+ 

ffl-0.5 

...J 

:::> 

0 

..., 

g 

:;;2 

~ 

> (!) 

a: 

w 

z 

w 

~-1.0 

a: 

LL 

w 

(!) 

z 

<( 

:I: 

u 

~ 

-1.5~------~-------L------~~----~ 

-0.3 -0.4 -0.5 -0.6 -0.7 

SMECTITE LAYER CHARGE (EQUIVALENT PER HALF-UNIT CELL) 

Figure 3. Free energy of exchange for sodium-potassium 

exchange as function of smectite layer charge . 

Concentration of potassium in illite interlayers 

A simple explanation for sodium-potassium selectivity is as follows. The ion preferred by the clay is the 

ion that is adsorbed with the greatest energy. Energy of 

adsorption (Ead) can be calculated approximately from 

the integrated form of Coulomb's law: 

qcqak (7) 

where qc and qa are the charges on the cation (1 +) and 

the anion (1-), k is a constant (1,389), rc is the radius of 

the cation in Angstrom units (A), and ra is the radius of 

the anion in A. 

Holding the radius of the anion (that is, the surfacecharge density of the clay interlayer) constant, equation 

7 predicts that for univalent exchange, the cation with the 

smaller radius is preferred. Several different selectivities 

for sodium-potassium exchange are possible because these 

ions change radius as they change hydration state. Hence, 

a hydrated potassium ion is slightly preferred compared 

with a hydrated sodium ion because hydrated potassium 

has a slightly smaller radius. A hydrated potassium ion 
has a smaller hydrated radius than a hydrated sodium ion 

because the larger potassium ion has a smaller surfacecharge density, and, therefore, attracts fewer water molecules. If a potassium ion dehydrates and a sodium ion 

remains hydrated, the potassium ion is greatly preferred 

compared to sodium in the interlayer. If both dehydrate, 

sodium is preferred compared with potassium, a phenomenon that occurs in some aluminum-rich zeolites (Sherry, 

1969). 

Exchange energies for the above examples are 

depicted in figure 5. Radii used in equation 7 to calculate 

these data were taken from Nightengale (1959), Eisenman 

(1962), and Garrels and Thrdy (1982). The hydrated radius 

of sodium is 3.58 A, that of hydrated potassium is 3.31 

A, that of dehydrated sodium is 0.98 A, and that of 

dehydrated potassium is 1.33 A. The equivalent radius of 

an aluminum-oxygen-silicon anionic interlayer site was 

calculated to be 2.1 A by Eisenman (1962). Exchange 

energies in figure 5 are not meant to be accurate because, 

for example, the "hydrated radius" is a function of how 

the radius is measured and may change between solution 

and interlayer. These values are presented, however, to 

1-

z 

w 

...J 

~ :::> 

0 

w 

a: 

w 

0.. 

en 

w 

...J 

::> 

0 

..., 

0 

...J 

:;;:: 

z 

~-{).5 

(.!) 

a: 

w 

z 

w 

w 

w 

a: 

LL. 

w 

(!) 

z 

<( 

::I: 

u 

X 

w 

NaX+ K+ ~ KX+ Na+ 

-1 ;0 L..--__ __.__ __ ___. ___ ...J_ __ __._ __ _ 

0 -1.0 -2.0 -3.0 -4.0 -5.0 

INTERLAYER MOLALITY Un) 

Figure 4. Free energy of exchange for sodium-potassium 

exchange as function of interlayer molality. 

show the possible magnitude of the effect of dehydration 

on exchange energies. 

The layer charge at which ions dehydrate has been 

presented schematically as a break in slope in figure 5. 

Interlayer cations dehydrate when a critical interlayer 

boundary potential has been exceeded due to an increase 

in smectite layer charge during diagenesis. This boundary 

potential is determined by the layer charge at which the 

energy of attraction of an interlayer cation for its hydration shell is exceeded by the energy of attraction of the 

cation for the smectite interlayer surfaces. The interlayer 

cation is thereby dehydrated by the surfaces; it then 

approaches the surfaces (crosses the "boundary") and is 

fixed, for example, to an aluminum-oxygen-silicon 

interlayer site by an energy determined by the site's electrostatic potential. This explanation assumes that (1) a 

hydrated interlayer cation "feels" the interlayer surfaces 

as planes of nearly uniform charge density, whereas a 

dehydrated cation "feels" a negative charge that is 

localized at a particular site in the 2:1 layer, and (2) smectite layers are converted into illite layers via a 

transformation-style reaction rather than via smectite 

dissolution and illite neoformation (see Eberl, 1984). 

CONCLUSIONS 

A cation in a smectite interlayer dehydrates in 

response to increasing layer charge when the energy of 

attraction between the cation and interlayer surfaces 

exceeds the energy of attraction between the cation and 

1-

z 

~ 100 

~ :::> 

0 

w 50 a: 

w 

0.. 

en 

w 5 0 

0 

..., 

0 

...J 

:;;:: 

z -50 

> (.!) 

a: 

~ -100 w 

w 

w 

a: 

LL. 

w -150 

(!) 

z 

<( 

::I: 

NaX+K+= KX+ Na+ Both ions 

1- dehydrated 

reoth ions hydrated -

1- -

t- -

t- -

K+ dehydrated 

u 

~ - 200 ---LAYER CHARGE, INCREASING-~-

Figure 5. Free energy of exchange for sodiumpotassium exchange calculated from Coulomb's law. 

Sodium-Potassium Exchange During Smectite Diagenesis 367 
its hydration shell. As smectite layer charge increases 

during burial diagenesis, a potassium ion dehydrates 

before a sodium ion can dehydrate, because potassium has 

the smaller hydration energy. Hence, potassium is greatly 

preferred compared to sodium when the layer charge 

required for potassium dehydration has been exceeded. 

The dehydration of potassium by interlayer electrical 

forces leads simultaneously to increased selectivity for 

potassium, expulsion of interlayer water, and collapse of 

the smectite interlayer to form illite. 

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*U.S. G.P.O. 1987- 773-047:46059 

_.