Isotopic Composition of 

Diagenetic Carbonates in 

Marine Miocene Formations 

of California and Oregon

GEOLOGICAL SURVEY PROFESSIONAL PAPER 614-B

Isotopic Composition of 

Diagenetic Carbonates in 

Marine Miocene Formations 

of California and Oregon

By K. J. MURATA, IRVING FRIEDMAN, and BETH M. MADSEN 

SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

GEOLOGICAL SURVEY PROFESSIONAL PAPER 614-B

A discussion of the isotopic composition 

of diagenetic carbonates i?i terms of 

chemical processes that operate in 

deeply buried marine sediments

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1969
UNITED STATES DEPARTMENT OF THE INTERIOR 

WALTER J. HICKEL, Secretary

GEOLOGICAL SURVEY 

William T. Pecora, Director

For sale by the Superintendent of Documents, U.S. Government Printing Office 

Washington, D.C. 20402 - Price 45 cents (paper cover)
CONTENTS

Abstract .________________________________________

Introduction. _______________________________________

Acknowledgments. _____________________________

Mode of occurrence of diagenetic carbonate____________

Factors that control carbon and oxygen isotopic composition. ____________________________________________

Time and circumstance of diagenesis__________________

Methods of preparation and analysis of the samples_____

Petrography of samples____________________________

General aspects of the isotopic composition-___________

Normal marine carbonate____________________________

Diagenetic limestone._______________________________

Bl

1

2

2

4

5

6

7

11

11

12

Oxygen isotopic composition of dolomite_____________ B13

Nature of formation water and its influence on oxygen

isotopic composition of dolomite._______________ 15

Dolomite with variable light carbon and independently 

variable oxygen.______-___-__--___-_-_-_--______. 16

Dolomite with variable heavy carbon and relatively constant heavy oxygen_______________________________ 17

Carbonate with variable heavy carbon and covariable

oxygen __________________________________________ 20

Calcium content of dolomite as related to isotopic

composition. __ _______________________________ 20

References-_ _______________________________________ 21

ILLUSTRATIONS

FIGURE 1. Map showing localities of the carbonates that were studied-____________________________________

2. Photograph of typical outcrops of Miocene diagenetic carbonates-___-_---_-_-------_-___-----_--

3. Photograph of dolomite from the Miocene of Berkeley Hills_____________-_-__--__-__-___-_-_-__-

4. Graph showing covariance of CaCO3 content of dolomite and the d-spacing of [222]_______________

5. Photomicrographs of thin sections of carbonate rocks___________________-_-____-__-___-___-_-.

6. Graph showing relationship between isotopic compositions of carbon and oxygen in carbonates______.

7. Graph showing temperatures indicated by the oxygen isotopes of dolomites from Palos Verdes Hills.

8. Graph showing temperature dependence of O18 content of calcite________-____--_--___-___-__-__.

9. Histograms of carbon isotopic composition of coexisting methane and carbon dioxide._-.___________,

10. Graph showing isotopic temperature and depth of origin of natural gas___________________________

11. Graph showing carbon isotopic composition as related to the CaCO3 content of diagenetic dolomite..

B2

3

4

6

8

12

15

16

18

19

21

TABLE

TABLE 1. Carbon and oxygen isotopic composition of diagenetic carbonates._________________-_-__-_______---_--___ B9

HI

SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

ISOTOPIC COMPOSITION OF DIAGENETIC CARBONATES IN MARINE MIOCENE 

FORMATIONS OF CALIFORNIA AND OREGON

By K. J. MTJRATA, TRYING FRIEDMAN, and BETH M. MADSEN

ABSTRACT

Diagenetic carbonates, mostly protoclolomite, occur as thin 

beds or concretionary zones in marine Miocene shales of California and Oregon. They manifest an unusually wide range of 

carbon isotopic composition (SO13 of 25 to +21 per mil relative to Peedee belemnite standard) and are divisible into a 

light-carbon group (SC13 of 2 to 25 per mil) and a heavycarbon group (SO13 of +5 to +21 per mil). The range of their 

oxygen isotopic composition coincides generally with that of 

normal marine carbonate (SO18 of 26 to 34 per mil relative to 

standard mean ocean water).

Calcite of Foraminifera and other fossils (SC10 of 0±4 per 

mil) in consolidated sediments was dissolved and redeposited 

along certain favorable zones to form diagenetic calcite. The 

formation water involved must have contained at least as much 

Ols as sea water for the SO18 of diagenetic calcite to lie in the 

range of normal marine carbonate. Dissolved in the water were 

carbonate methane, naphthenates, and other substances formed 

through diagenetic alteration of organic matter in the sediments. 

The SC13 of the deposited calcite differed substantially from that 

of the original fossils, the carbon being heavier or lighter depending on whether or not isotopic equilibrium was attained 

between the carbonate ion and methane.

Isotopic exchange equilibration between carbonate ion and 

methane concentrates O13 in the carbonate and is the most likely 

mechanism for producing heavy-carbon carbonates. Without this 

equilibration, the carbonate ion will be rich in O12, in keeping 

with its derivation from organic matter, and will yield diagenetic carbonates of the light-carbon group. The exchange could 

take place between finely porous carbonate and methane gas 

without necessarily involving water; this could lead to variation 

in the carbon isotopic composition among some heavy-carbon 

carbonates without a concomitant variation in oxygen composition or in crystalline texture.

At most places, the first-stage diagenetic calcite was soon 

dolomitized by brines that tended to make the oxygen somewhat 

heavier (maximum SO18 of 35.4 per mil) and the carbon, as well, 

if the carbonate-methane equilibration was operating. Protoclolomite with a calcium content as high as 58 mol percent 

formed first, but in time exsolved the excess calcium to become 

ideal dolomite. Among some heavy-carbon dolomites, these several adjustments led to an inverse relation between the content 

of calcium and either the carbon or oxygen isotopic composition.

The temperature implications of the carbon isotopic composition can be explored in a preliminary way if it is assumed that 

the methane involved in the equilibration had the same isotopic

composition as the methane of natural gas from oil fields (most 

prevalent SC13 values of 52 to 38 per mil). Under this assumption, the heaviest and the lightest carbonates (SC13 of +21 

and +5 per mil) within the heavy-carbon group limit equilibrium temperatures to 34° to 140°C. Existing data on genesis 

of petroleum indicate that transformation of organic matter into 

petroleum proceeds at an appreciable rate only at temperatures 

higher than about 115° C. Most of the diagenetic carbonates 

seem to have formed at somewhat lower temperatures, but tl «ir 

isotopic and mineralogical natures clearly suggest them to be 

byproducts of petroleum-generating reactions.

An interesting variant of diagenetic carbonate is the calcite 

cement of the Miocene oil-bearing sands that occur as coastal 

plain deposits! lapping onto the granite of the Sierra Nevada. 

The formation water in these marine sands is virtually fresh. 

The presence of fresh water accounts for the fact that the carbon and oxygen isotopic composition of the calcite cement is 

similar to that of fresh-water limestone and explains why the 

cement escaped the almost universal dolomitization that occurred elsewhere.

INTRODUCTION

Diagenetic carbonate in the form of indurated bed? of 

concretionary zones is a minor but conspicuous rock in 

the Miocene Nye Mudstone of Oregon and the Monterey 

Shale of California (Suavely and others, 1964; Bramlette, 1946). Recently, interest in this carbonate has 

grown because of the realization that it is mostly dc tostone rather than limestone and that, in many places, 

it is extraordinarily rich in carbon isotope C13 (Murata 

and others, 1967).

In the present paper, we discuss the mineralogy and 

the carbon and oxygen isotopic composition of 79 samples of this and related carbonates from widely scattered localities (fig. 1). The main impetus for the study 

springs from the possibility of gaining new insight 

into the diagenesis of organic-rich sedimentary rocks 

such as the Monterey Shale.

For our purpose, we adopt the broad definition of diageiiesis given in the "Dictionary of Geological Terms" 

of the American Geological Institute (1962) : "Process 

involving physical and chemical changes in sediment

Bl
B2 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

after deposition that converts it to consolidated rock; 

includes compaction, cementation, recrystallization, and 

perhaps replacement as in the development of dolomite." 

Existing data on the formation and diagenesis of limestone and dolostone have been well summarized in two 

recent reviews (Pray and Murray, 1965; Chilingar and 

others, 1967).

Diagenesis of the Monterey Shale has attracted the 

attention of many investigators because the shale is 

unusually rich in opaline silica (diatoms) and because 

it is considered to have been the source bed of much 

petroleum in California. Besides the carbonates that 

constitute the subject of our inquiry, diagenesis has 

produced chert from the opaline silica of diatoms 

(Bramlette, 1946), zeolites and montmorillonite from

Drakes Bay (5-6)

Berkeley Hills (7-9) SAN"' " 

FRANCISCO ,

9Santa Cruz Mts (10-11) 

no Nuevo Point (12)

Rehz Canyon (13-26)

Nacimiento Dam (27-32)

 Round Mtn (41-45)

 Chico Martinez Cr (33-40)

 Temblor Range (46)

 Cuyama Valley (50) 

.San Rafael Mts (51-53) 

.Santa YnezR(54-55) 

" " *Pine Mtn (56-57)

Point Dume (58-60)

Palos VerdesHills(61-75)

San Juan 

Capistrano (76)

San Clemente 

Island (77)

Pismo synchne 

(47-49)

100

L_

La Jolla and Coronada , 

Canyons (78-79)*

0

_ i I

FIGURE 1. Localities of the carbonates that were studied. 

Numbers in parentheses are the sample numbers listed in 

table 1.

volcanic tuffs (Bramlette and Posnjak, 1933), and oil 

from sedimentary organic matter (Trask, 1932; Philippi, 1965). The middle Miocene part of the Monterey 

Shale is highly phosphatic (Gower and Madsen, 1964; 

Dickert, 1966). Although 110 diagenetic effects have been 

noted so far in the phosphorite, future studies will probably disclose such effects even in this relatively stable 

material.

ACKNOWLEDGMENTS

We are indebted to Prof. D. L. Inman of Scnpps Institution of Oceanography for sample 78 from La Jolla 

Canyon. Advice regarding stratigraphy, paleontology, 

and isotope geology of west coast Tertiary rocks was 

generously given by Prof. M. N. Bramlette and Dr. 

Arthur Jokela of Scripps Institution of Oceanography, 

by Prof. R. H. Jahns of Stanford University, by Dr. S. 

R. Silverman of Chevron Research Co., and by our 

colleagues, W. O. Addicott, Ivan Barries, R. H. Campbell, T. W. Dibblee, Jr., D. L. Durham, G. W. Moore, 

J. R. O'Neil, R. O. Rye, J. E. Schoellhamer, Patsy J. 

Smith, P. D. Suavely, Jr., H. F. Tourtelot, and R. E. 

Zartman. Laboratory investigations were carried 

through with the able assistance of R. R. Bruegger, 

Chere N. Barnett, Joy Church, and James Gleason.

MODE OF OCCURRENCE OF DIAGENETIC CARBONATE

Figure 2A shows the typical mode of occurrence of 

diagenetic carbonate. It is found in conformable beds or 

concretionary zones, generally less than 3 fe3t thick, 

spaced 5 to 50 feet apart in the shale or mudstone of the 

west coast Miocene. The associated shale or mudstone 

may be devoid of any carbonate or may be moderately 

rich in calcareous Foraminifera. Figure 25 illustrates 

an occurrence of diagenetic carbonate in a sequence of 

alternating layers of shale and sandstone; the thin sandstone layers are hardened by a calcite cement. A closer 

view of the lowest of the three carbonate beds of figure 

2B is shown in figure 2(7. Although very ordinary in 

appearance, this bed has 'an exceptionally heterogeneous 

composition. Its top is dolomitic; its bottom, calcitic; 

and the middle part, a variable mixture of the two carbonates. It apparently represents an incompletely dolomitized bed of diagenetic limestone; all the cfolostone 

beds may originally have been limestone. Thin sections 

of the dolomitic and calcitic parts of the bed ar^ shown 

in figures W 'and 5F.

In demonstrating the secondary (diagenetic) origin 

of carbonate concretions and lenticular beds in the 

Monterey Shale, Bramlette (1946) called attention to 

the relation of carbonate concretions to the bedding in 

the enclosing strata. Beds continue through the concretions and are thicker inside of them than outside, as 

if the concretions had preserved an earlier thickness of
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON B3

c

FIGURE 2. Typical outcrops of Miocene diagenetic carbonates. 

A, Dolomitic beds and lenses in the Nye Mudstone, north side 

of Yaquina Bay, Lincoln County, Oreg. Sample 2 of table 1 

was obtained from the bed near the man. Photograph by P. D. 

Snavely, Jr. B, Carbonate zones (the three thicker beds) in the 

Monterey Shale of Pine Mountain, Ventura County, Calif. At

B

this locality, the shale contains numerous thin layers of 

calcite-cemented sandstone. The rucksack is 30 inches long. (7, 

Detail of the lowest carbonate bed in view B. The upper part 

of this bed is dolomitic ; the lower part, calcitic; the two parts 

are represented by samples 56 and 57, respectively, in table 1. 

The hammer is 11 inches long.

the beds against the compressive load of the overlying 

sediment. He emphasized the general impurity of the 

carbonate: "All the concretions include more or less 

sedimentary material similar to that forming the adjacent beds, and most of them are little more than carbonate-cemented nodular masses of the sedimentary 

material forming these adjacent beds." In some places, 

impregnation by carbonate protected opaline diatoms 

from being dissolved and glassy pumice from altering 

to montmorillonite.

The dolomite content of the dolostone generally depends on the porosity of the host sediment and varies

among our samples from about 40 weight percent in 

dolomite-cemented sandstone to about 85 percent in 

dolomitized diatomite. The massive, apparently laminated, dolostone shown in figure 3 contains 84 percent 

dolomite without any calcite; its insoluble residue consists almost entirely of tests of diatoms with very few 

detrital minerals. Pure diatomite may have a porosity 

as high as 75 percent and upon impregnation with 

dolomite could yield a rock containing about 80 percent 

dolomite. Thus, this sample of dolostone probably 

formed mainly through impregnation rather than replacement of laminated diatomite, a type of sediment
B4 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

common in the Monterey Shale. A few dolomitized tests 

of Foraminifera in the dolostone further indicate that 

any and all calcite originally present in the diatomite 

would have been converted to dolomite. A thin section 

of this sample is shown in figure 5B.

Two aspects of the diagenetic carbonate are especially 

significant with regard to its origin and nature. First, 

the widely spaced thin beds of the carbonate constitute 

a minor constituent, making up perhaps a percent or so 

of the total thickness of a typical section. Second, with 

a few exceptions, the diagenetic carbonate is dolomite 

rather than calcite. The frequent presence of dolomite 

along sharply delimited horizons in moderately calcareous shale sections indicates that dolomitization was 

effected only along certain favorable zones by solutions 

introduced from outside.

Later in the report, we shall postulate certain isotopic 

reactions between methane (generated by burial metamorphism of organic matter) and the carbonate, in 

order to account for abnormally heavy carbon in the 

carbonate of Monterey Shale. The effect of such reactions would be greatest in a formation like the Monterey

O 1

I I I

2 3cm

i i I

FIGURE 3. Dolomite (sample 8 of table 1) from the 

Miocene of the Berkeley Hills, Calif. This hard 

massive rock originated through impregnation of 

laminated diatomite with dolomite (84 weight percent). Insoluble matter consists largely of opaline 

tests of diatoms with very few detrital minerals. 

Vertical cracks contain very late calcite.

Shale in which the ratio of organic shale to carbonate 

is large and in which the reacting fluids moved only 

along certain zones rather than by diffusing throughout 

the section.

FACTORS THAT CONTROL CARBON AND OXYGEN 

ISOTOPIC COMPOSITION

Because comprehensive reviews of the general geochemistry of carbon and oxygen isotopes have been* 

published recently (Craig, 1963; Rankama, 1963; Silverman, 1964; and Degens, 1965), only data and concepts that are pertinent to the origin of diagenetio 

carbonates need to be considered here. Carbon isotopes 

of sedimentary rock can be discussed in terms of two 

major categories of materials with distinctly different 

isotopic composition: (1) Marine limestone (8C13=0±4 

per mil relative to the PDB, or Peedee belemnite, standard of University of Chicago), and (2) organic substances (such as coal, oil, and dispersed organic matter 

of soils or shales), containing much lighter carbon 

(SCi3= 35 to 8 per mil) and carbon dioxide derived 

through oxidation of these organic substances. The relatively constant isotopic composition of marine limestone results from precipitation of the limestone under 

conditions of isotopic equilibrium with the bicarbonate 

of sea water, which is virtually constant isotopically 

(SC13= -2.5 to -1.3 per mil according to Craig, 1954).

The light carbon that characterizes all organic substances originates in the marked preference shown by 

plants for isotope C12 during photosynthesis. Carbon of 

plant tissues can be deficient in C13 by as much as 30 per 

mil relative to atmospheric carbon dioxide (Park and 

Epstein, 1961). The following table gives the isotopic 

composition of atmospheric carbon dioxide, plant tissue, 

derivative sedimentary organic matter, and other 

related substances. The isotopic composition of organic 

matter of animals is similar to that of the plants on 

which they feed.

Material

Atmospheric COa----- 

Marine bicarbonate. _. 

Marine limestone. 

Marine plants. _______

C»3 (per mU, PDB) Reference

-7.4 to -6,7 Keeling (1958).

-2.5 to -1.3 Craig (1954).

Baertschi (1957).

Deuser and Degens 

(1967).

Craig (1953).

Craig (1953) and 

Sackett (1964).

Craig (1953).

Silverman (1964).

Oana and Deevey 

(1960).

Clayton and 

Degens (1959) 

and Munnich 

and Vogel (1959).

In comparison to sea water and marine limestone, 

terrestrial waters and limestone contain carbonate 

whose carbon is generally lighter and more variable

Terrestrial plants__. 

Organic matter of

marine sediments. 

Coal_____________

Petroleum, ________

Bicarbonate in fresh

water. 

Fresh-water limestone

-4

-30

-29

-31

-27

-34

-23

to +4 

to -12

to 23 

to -14

to 22 

to -22 

to 4

__ -18 to -3
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON B5

isotopically, as indicated in the table. Fresh water contains, besides the normal bicarbonate based on carbon 

dioxide of the atmosphere, substantial and variable 

amounts of much lighter bicarbonate (SC13 of 30 

to 20 per mil), based on the carbon dioxide derived 

from the decaying organic matter of soils. The variation 

of isotopic composition of fresh-water bicarbonate and 

limestone from such light values to those of marine 

limestone depends on the extent to which circulating 

fresh water encounters and dissolves marine limestone 

(Oana and Deevey, 1960).

Oxygen isotopic composition of sea water is fairly constant at SO18 of 0±1 per mil relative to standard mean 

ocean water (SMOW), according to Epstein and Mayeda (1953) and Craig and Gordon (1965). In the hydrologic cycle, water is evaporated from the ocean under 

nonequilibrium conditions to yield vapor of light H2O 

with SO18 of 14 to 10 per mil (Craig and Gordon, 

1965). Water-laden air from lower latitudes moves poleward and precipitates rain or snow on land and sea, as 

it encounters progressively lower temperatures. H2O18 

tends to condense preferentially from the migrating 

vapor, giving rise to changes in the isotopic composition of the precipitation depending upon the latitude. 

As a result, polar ice can be as light as SO18 of 50 to 

 40 per mil (Epstein, 1959). The contrast between sea 

water and fresh water in the middle latitudes would be 

SO18 of zero for the former and of 15 to 5 per mil 

for the latter.

This contrast in oxygen isotopic composition of 

marine and terrestrial waters also appears in marine 

and fresh-water limestones. Besides the original isotopic composition of the water, temperature of the 

water also determines the amount of O18 in carbonates. 

Because oxygen of sea water is nearly constant isotopically, the range of SO18 in marine limestones (mostly 

26 to 34 per mil, SMOW) is largely attributable to 

variations in temperature of the water, and indeed such 

variations form the basis for determining the paleotemperatures of ancient seas (Urey and others, 1951; 

Bowen, 1966). Fresh-water carbonates (SO18 mostly in 

the range of 9 to 23 per mil) are more complex because 

they form from solutions that vary widely in both 

isotopic composition and temperature (Weber, 1964).

With both carbon and oxygen tending to be lighter 

in terrestrial than in marine limestone, isotopic composition would seem to be a useful indicator of the 

depositional environment. The criterion has proved to 

be ambiguous for some samples because of moderate 

overlap of isotopic compositions, but is accurate in some 

80 percent of the samples (Clayton and Degens, 1959; 

Keith and Weber, 1964; Keith and others, 1964). The 

isotopic composition of most of our diagenetic car334-795 O 69 2

bonates indicates derivation from neither fresh water 

nor sea water, but rather from a different category of 

natural water, probably that which is commonly called 

formation water or oil-field brine.

TIME AND CIRCUMSTANCE OF DIAGENESIS

Diagenetic carbonate of the Monterey Shale evidently formed before the final compaction of the host 

sediment (Bramlette, 1946), but neither the circumstance of its formation nor the kind of chemical solution involved is known. Lithification of modern calcareous sediments lying exposed on the sea floor has been 

reported by Fischer and Garrison (1967). However, the 

hard carbonate layers in the Monterey Shale could not 

have formed in this way, because their carbon isotopic 

composition indicates equilibration with solutions very 

different from sea water and because they commonly 

involve carbonate-poor materials, such as diatomite and 

volcanic tuff.

Studies of modern marine sediments and their interstitial waters (Emery and Eittenberg, 1952; Siever 

and others, 1965; Berner, 1966; among others) have 

generally disclosed little diagenetic change in these 

sediments to a depth below the sea floor of 10 meters or 

so, representing a timespan of 105 to 106 years. The 

interstitial water of some modern marine sediments, 

however, is enriched in phosphate and silica and depleted in sulfate as the result of diagenetic reactions 

(Emery, 1960; Brooks and others, 1968).

Even older sediments, such as the Miocene ooze recovered during experimental mohole drilling of the 

Pacific sea floor (Eiedel and others, 1961; Murata and 

Erd, 1964), are soft and porous and have no indurated 

zones of carbonate or chert. The interstitial water in 

this ooze differs very little from normal sea water (Kittenberg and others, 1963; Siever and others, 1965); it 

is, however, somewhat deficient in deuterium (Friedman, 1965). The calcite and dolomite of the ooze have 

carbon and oxygen isotopic compositions in the range 

of normal marine carbonate (Degens and Epstein, 

1964).

Compared to such relatively unaltered sediments, the 

Monterey Shale and overlying sediments accumulated 

to much greater thickness in nearshore basins, underwent moderate burial metamorphism, and at many 

places were uplifted into the zone of weathering during 

Pliocene and later orogenies. The Los Angeles basin 

contains over 25,000 feet of Miocene and Pliocene sedimentary rocks (Barbat, 1958; Yerkes and others, 1965), 

and other basins, such as Chico Martinez Creek (Bramlette, 1946) and Eeliz Canyon (Durham, 1963), have 

more than 7,000 feet of middle and upper Miocene sedimentary rocks which, before much of the cover of
B6 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

younger deposits was eroded away, must have lain at 

depths of 10,000 to 12,000 feet. At such depths, sediments would have been subjected to temperatures 

higher than 100° C, which would 'be sufficient to convert sedimentary organic matter into petroleum 

(Philippi, 1965), accelerate the reconstitution of biogenie opaline silica into chert, and induce many other 

chemical changes in a geologically short time.

Unlike the sea-water-like interstitial solutions found 

in undisturbed deep-sea oozes of even Tertiary age, formation waters (oil-field brines) that occur in the Monterey Shale and related formations vary greatly in 

salinity and composition. They range in salinity from 

fresh water to brines several times more saline than sea 

water (Jensen, 1934; White, 1965). Deeply circulating 

meteoric water becomes converted into saline formation 

water by passage through semipermeable shales and by 

undergoing numerous other chemical reactions that 

alter and reconstitute sedimentary materials (Clayton 

and others, 1966; Anderson and others, 1966; Rittenhouse, 1967). As will be shown later, the isotopic composition of the diagenetic carbonates indicates that formation waters from which most of the carbonates separated resemble neither fresh water nor sea water, but 

rather constitute a different category of natural water.

METHODS OF PREPARATION AND ANALYSIS OF THE 

SAMPLES

To prepare the samples, one or two pieces of each carbonate were first sawed perpendicularly to the bedding 

plane and stained with alizarin red sulfonate (Friedman, 1959). Samples shown by the staining test to be 

finely intergrown mixtures of dolomite and calcite were 

discarded. Many of these consisted of calcareous Foramnif era imbedded in a fine-grained dolomitic groundmass 

and probably represented a stage of alteration of calcareous shale transitional to dolostone in which both the 

Foraminifera and the groundmass are dolomitic. Some 

of the rejected samples were fractured dolostones extensively veined by late calcite.

X-ray diffractometer patterns of all samples that 

passed the staining test were made with nickel-filtered 

copper radiation at a scanning speed of 1° of 20 per 

minute, in order to verify the purity of the carbonate 

and to determine the nature of the associated noncarbonate minerals. Special X-ray patterns were prepared, 

by using a scanning speed of % ° of 2© per minute and

elemental silicon as the internal standard, to determine 

accurately the unit-cell dimensions of 14 dolomite samples by means of the computer method of Evans, Appleman, and Handwerker (1963). The relationship between cell dimension and composition of dolomite 

(Goldsmith and Graf, 1958) indicated that the calcium 

content of the 14 samples varied substantially, ranging 

from 50 to 58 atomic percent.

A crude but useful index of the cell dimension and 

composition was also obtained for all other dolomites 

from the patterns taken at 1° per minute, by determining the ^-spacing of the {222} reflection, using the main 

quartz peak at 3.343A (angstroms) as the internal 

standard. The principal {211} peak of dolomite is so 

broad in these patterns that its spacing could not be 

measured as precisely as that of the {222} peak. The 

working curve that relates atomic percent of calcium 

(determined on the 14° per min patterns) and the d222 

spacing (determined on the 1° per min patterns) is given 

in figure 4, and compositions derived from this curve 

are listed in table 1. Even a crude estimate of the composition is of much interest from the viewpoint of the 

thermal history of the samples, because dolomites containing more than 50 atomic percent of calcium 

(protodolomites) are unstable above about 200°C 

(Goldsmith, 1959).

o o

o

2.680 2.690 

c/222 SPACING, IN ANGSTROMS

FIGURE 4. Covariance of CaCOs content of dolomite and the 

^-spacing of {222}, Equation for the line is as follows: 

OaOOs (mol percent) =322.6(fe, 2.668) +49.0.
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON B7

Carbon and oxygen isotopic compositions were determined on carbon dioxide, which was obtained by decomposing the samples with 100 percent phosphoric 

acid (McCrea, 1950) and was analyzed in an isotope 

ratio mass spectrometer (McKinney and others, 1950). 

Prior to decomposition by acid, many of the samples 

were held at 300 °C for an hour in a stream of helium in 

order to remove the more labile organic matter. Carbon 

isotopic composition is expressed relative to University 

of Chicago PDB standard (Craig, 1957), and oxygen 

isotopic composition, relative to SMOW (Epstein 

and Mayeda, 1953; Craig, 1961). The reproducibility 

of the determinations is ± 0.1 per mil.

The analytical results are given in table 1. Where two 

or more samples are listed from a given locality, they are 

in stratigraphic order, the youngest sample being given 

first. The sampled beds represent from a quarter to a 

half of the total available and are usually spaced 30 

to 100 feet apart with one or more unsampled beds in 

between; a few pairs of samples (such as Nos. 10 and 

11, and 59 and 60) are from beds over 500 feet apart. 

Of the 79 samples listed in table 1,60 represent dolomite 

beds or concretions; 1, vein dolomite; 1, dolomitic cement of sandstone; 8, calcareous fossils or Foraminiferarich shales; 4, limestone lenses; 4, calcareous cement 

of sandstones; and 1, vein calcite.

PETROGRAPHY OF SAMPLES

Photomicrographs of thin sections of a dozen samples 

of different textural type and isotopic compositions are 

presented in figure 5 in order to show the microscopic 

character of the carbonate, whose megascopic appearance was illustrated in figures 2 and 3. The individual 

samples are described in detail in the legends to the 

figures. The photomicrographs were made at a magnification of X100, because at this high magnification inherited textures can be largely ignored and attention 

focused on the crystalline character of the diagenetic 

carbonate itself.

Figure 5 shows clearly that diagenetic calcite and 

dolomite, though differing greatly in isotopic composition, are commonly very fine grained with crystal 

diameters of the order of 5 microns (compare figs. 5A 

and 5Z). The lack of correlation between isotopic composition and grain size of dolostone has already been 

emphasized by Weber (1964). The texture of dolomite 

sample 8, the hand specimen of which was illustrated in

figure 3, is shown in figure SB ; its variability is simply 

inherited from the original varved diatomite. Thin 

sections of the dolomitic and calcitic parts of a single 

bed are shown in figures 5C and 5^, respectively. The 

coarser texture of the dolomitic part suggests that a considerable enlargement of crystal size sometimes occurs 

when fine-grained calcite is dolomitized.

Figure 57 illustrates a type of dolostone that is made 

up of euhedral crystals of dolomite each with a nucleus 

of oily material or a gas, probably carbon dioxide. 

The carbonate-carbon of this dolomite is so light 

(SC13= 11.7 per mil) that it suggests a genetic relationship to the light carbon of the nuclear material (Spotts 

and Silverman, 1966) . Simple impregnation of a carbonate rock with tar or other organic matter, however, does 

not automatically introduce light carbon into the carbonate mineral. Sample 21 of teble 1 (+0.03, 28.4) is a 

tar-soaked coquina of Foraminifera which well illustrates a lack of isotopic exchange between organic 

matter and carbonate. Although organic matter clearly 

has a role in determining1 the carbon isotopic composition of many samples, no general correlation between 

concentration of organic matter and isotopic composition could be established; this agrees with the results 

of Jeffrey, Compston, Greenhalgh, and De Laeter 

(1955).

The state of oxidation of many samples can be 

judged roughly from the relative abundance of small 

blebs of authigenic pyrite and blebs of hematitic pseudomorphs of the pyrite. However, no consistent relationship between the state of oxidation and the isotopic 

composition is apparent.

Primary calcite still retaining to a large degree the 

isotopic composition of normal marine carbonate is 

exemplified by the barnacle shell in figure 5E and by 

foraminiferal debris in the shale in figure 56*. A type 

of secondary calcite, with isotopic characteristics indicating equilibration with fresh water, cements the grains 

of the marine sandstone shown in figure 5/. The calcite 

consists of crystals as coarse as the sand grains and 

very irregular in shape (almost amoebiform).

We have been able to attach genetic significance to 

only a few of the different petrographic features illustrated in figure 5. The figure is presented mainly to 

show typical samples and to indicate concisely the apparent lack of correlation between petrographic character and isotopic composition.
B8 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

D F

K L

FIGURE 5. Photographs of thin sections. (Explanation on facing page.)
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON

FIGURE 5. Photomicrographs of thin sections of carbonate 

rocks, all taken with plain light at magnification of X 100. All 

sections were cut perpendicularly to the bedding. In the legend 

for each sample, the first of two numbers enclosed in parentheses is the carbon isotopic composition relative to the PDB 

standard, a^d the second is the oxygen isotopic composition 

relative to SMOW. Sample numbers are the same as those 

used in table 1.

A. Dolomite sample 33 (+20.8, 33.8). Fine-grained impregnation of a diatom-rich sediment; contains the heaviest 

carbon.

B. Dolomite sample 8 (+19.4, 33.1). Vertical variation in 

grain size in part of a varve layer was inherited from the 

original varved diatomite; rich in diatoms (elongate objects in figure) and formless opal that fills the porous 

parts of each varve. A hand specimen of this sample is 

illustrated in figure 3.

C. Dolomite sample 56 (+12.5, 31.4). From same bed as 

calcite sample 57 of F below. Upper part of bed shown in 

figure 1C. Moderately coarse texture. Foraminifera in 

lower left corner is partly dolomitized and is filled with 

sparry crystals of both dolomite and ealeite.

D. Calcite sample 51 ( + 15.2, 28.8). Fine-grained groundmass. 

Interior of diatoms and Foraminifera are filled with 

sparry calcite, chalcedony, or hematite probably after 

pyrite. Dolomite of A above may have had ealeite like 

this as a precursor.

E. Calcitic shell sample 14 ( 1.4, 27.7) of a barnacle is seen 

in cross section along the bottom of the figure. THy 

euhedral dolomite crystals (+7.0, 30.2) and opal make 

up the light-colored matrix; the matrix also contains 

a moderate amount of detrital minerals, such as those 

lying in the concavity of the shell.

F. Calcite sample 57 (+10.2, 27.3) from the same bed as 

dolomite sample 56 of C above. Fine-grained calc'te 

with much brown organic matter. Foraminifera is filhd 

with sparry calcite. This rock probably was precursor to 

the dolomite of C.

G. Calcareous shale sample 29 rich in foraminiferal debris 

( 0.69, 27.6) and disseminated opal.

H. Dolomite sample 28 (+11.1, 30.8) from a bed in the calcareous shale of G. Unusually large dolomite crystals in 

vuggy parts.

J. Dolomite sample 66 ( 11.7, 31.6). Moderately coarse 

grained; each dolomite crystal has a nucleus of an oUy 

material or of carbon dioxide. This type of dolomite 

was described in detail by Spotts and Silverman (196*>).

/. Caleite sample 44 ( 14.4, 20.8), cement of marine sardstone. Calcite crystals (gray) are as coarse as the sand 

grains (white) and very irregular (almost amoebiform).

K. Dolomite sample 62 ( 15.4, 34.3). Uniformly fine grained 

rock moderately rich in remains of diatoms. Other parts 

have many dolomitized tests of Foraminifera.

L. Dolomite sample 78 ( 25.3, 34.6). Fine-grained dolomite 

enclosing many detrital minerals along with blebs of 

pyrite and organic material; contains the lightest carbon.

TABLE 1. Carbon and oxygen isotopic composition of diagenetic carbonates and calcium carbonate content of diagenetic dolomites in

Miocene rocks of Oregon and California and of comparison materials

[Samples 1, 2, and 3 are from Oregon; samples 4 through 79, from California. Carbon isotopic results are relative to University of Chicago PDB standard; oxygen isotcpic

results are relative to SMOW; CaCOs results are in mol percent]

Sample

1 

2........

3 -----

4. ...... .

5 ---

6 -

7-------

8 ---

9 - 

10 

11 -

12- _- 

13 .-

14 -

15. ------

16 

17 

18 

20 

21 

22. 

23--..--.

24 

25 - 

26 

Lab.

No.

. 3277-4

. 3277-5

3277-6

3277-3

3279-1

3279-7

3277-1

3279-30

3277-2

3279-2

3279-3

3277-31

3277-23

3279-28

3277-22

3277-21

3277-20

3277-19

3277-18

3277-17

3277-16

3277-15

3277-14

3277-13

3277-11

3277-12

Locality

County.

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

County.

County.

County.

Costa County.

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

Santa Cruz County.

Mateo County.

Monterey County.

--.do -- .. .

..-.do - .

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

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

-.-do .-- .

-_._.do 

 -.do 

.-..do- ... . 

 .do.- .

___ ..do. 

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

_ .do 

.--do 

Latitude and longitude

44°^S' TJ 194°ni'W

 do.... ._ .

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

OQOKof -\r 19Q05Q' W

3g°Q9' j»^ fl*>9°5g' W

37°54'N.,122°43'W-.--

070 Cf)t VT 1990-Mf \rr

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

 do........ .

- do..._.. . . .

37°07'N.,122°20'W-_--

36°13'N., 121°17'W- .

.-... do ...._-_--_. .....

 do 

 .do .

.--do _._ 

 -do ._._-... . . -

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

 do . .

- do -. .

 -.do... .

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

..... do .. .

 -.do .

 .do .

Formation

-.... do......, -

 .do....... 

Formation.

 ..do...... 

 do..... 

 do..... 

Formation

(Pliocene) .

 .do 

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

 . do . 

 . do -----

ber of Monterey

Shale.

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

... ..do... . 

-... .do. - -

- .do --- 

 do.. ... -

Rock

2

sandstone.

3_. ------

coquina.

in coquina.

o

3 -------

4 . 

5 ... .

g

Vein calcite in bed 9 ....--

5C13

per mil

+5.8

1 7 ft

+11.0

__ Q K

1 11 °,

+10.2

+17.7

4-1Q 4

+11.3

-2.8

+9 6

+6.4

+7.0

-1.4

+16.5

+17.6

+14.9

+16.2

+17.0

1 17 9

+0.03

 1 Q

+13.5

+11.6

+5.7

-4.1

SO'8

per mil

32.7

00 0

34 0

01 Q

OK A

00 Q

00 1

33.1

90 a

30.4

on A

30.2

27.7

on o

31.3

30.0

on Q

°.1 9

00 0

28.4

28.0

31.9

01 0

97 1

19.5

CaCO"in

dolomite

55

53

51

57

50

49

51

52

58

52

52

57

52

52

52

53

52

52

52

55

Reference

1964.

Do.

Do.

W. 0. Addicott, 1967.

1961; Clark Blake,

oral commun.,

June 1967.

Do.

Do.

1962.

Do.

1961.

Durham, 1963; Durham and Addicott,

1965.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

334-795 O 69-
BIO SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

TABLE 1. Carbon and oxygen isotopic composition of diagenetic carbonates and calcium carbonate content of diagenetic dolomites in 

Miocene rocks of Oregon and California and of comparison materials Continued

Sample Lab.

27

28

29

30

31

32

33

34

35

36

37

38

39

40

41

42-..

43 

44

<!>.

46

47

48

49

50

51

52

53

54

55

56

57

58

59

60

61

62

63

64

65

66

67

68

69

70__

71

72_-

73

74

No.

...... 3277-25

...... 3277-26

...... 3277-27

..--. 3277-28

.-._. 3277-29

_.-.. 3277-30

...... 3279-12

.-.-. 3279-13

..... 3279-14

..... 3279-15

_._.. 3279-16

...... 3279-17

.... . 3279-18

..__. 3279-19

...... 3279-20

. . . 3279-21

.... 3279-22

..... 3279-23

.____. 3276-18

.. .. 3279-26

. .. 3279-24

.... 3279-25

3279-27

... 3276-15

..... 3276-16

.. . 3276-17

... 3279-9

. 3279-10

... ... 3279-32

...... 3279-33

.. . 3279-6

...... 3279-5

... ... 3279-4

..... 3276-3

. 3276-5

... 3276-10

... 3276-2

. 3276-11

.... 3276-6

...... 3277-7

.. 3276-12

..___. 3276-7

__.._. 3276-13

..... 3276-

14D

 . 3276-

14L

...... 3276-8

.._... 3276-9

Locality Latitude and longitude Formation Rock 5C13 SO 18 CaCOain Reference

per mil per mil dolomite

Luis Obispo County.

--- do- -._-._. ..

 do 

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

.....do 

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

Chico Martinez Creek,

Kern County.

 ..do. . . .

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

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

 do .

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

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

field, Kern County.

--_-do -------------

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

--..dO- .-.._-_.____._.

. do . ...

Temblor Range, San

Luis Obispo County.

Pismo syncline area,

San Luis Obispo

County.

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

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

Cuyama Valley, Santa

Barbara County.

San Rafael Mountains,

Santa Barbara

County.

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

 do- .

Santa Ynez River area,

Santa Barbara

County.

 ..do.... .... ..._.. _

Pine Mountain, Ventura County.

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

Point Dume, Los

Angeles County.

.... .do ... 

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

Palos Verdes Hills,

Los Angeles County.

.. ___do- ___ ___ _-_._._-

.....do.. .._-....._..-_.-

-_. .-do-.-... ---------

. ....do.-. _ _- _ ....

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

 do... 

. do-- 

-_. -do_-_ ... .-..-.-. -

 do_._ ... _._ _

 do-- 

35°45'N., 120°52'W ...

 do -_ .

. .do - --.

. do ___- _ .

. do -.. .

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

35°26'N., 119°47'W ...

.-..do- . .

 ..do- _ .

 .do -- .

 ..do __ __.

.- do- _ ....

. do - .

 -do . ..

35°28' N., 118°55' W _.

 do -...

35°29'N.,118°50'W ...

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

 do .. . ..

35°05'N., 119°35'W ...

35°13'N.,120°42'W-..-

35°10'N., 120°42'W-.-..

35°09'N., 120°41' W-.--.

34047' N., 119°36' W---_.

34047' N., 119°36' W

34°47'N., 119°41'W.....

34°35'N., 120°23'W.....

34°35'N., 120°09'W.-_-

34°38'N., 119°23'W.---.

 ..do .. .. .

34°01' N., 118C47' W..-..

. do-- -.

 -do-... . . .

33°48'N., 118°24'W ..

do

33°48'N., 118°21'W__ 

33°47' N., 118°20' W_.__

33°45'N., 118°22' W-. 

.. -.do--_. -----

-... .do-. ._-_._-_-.._._..

 ..do.. .

.....do_-.-------. .

do

33°45'N.. 118°23'W -.

ber of Monterey

Shale.

 __do 

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

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

 do ------

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

Chert of

McMasters

(1947).

...... do ..-..__

 do 

...... do.. . 

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

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

. do -

Silt of Diepenbrock(1933).

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

Godde (1928).

._._-do --------

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

. Monterey Shale.. .

Obispo Tuff

Member of

Monterey Formation.

. Monterey Shale,

immediately

underlying

Pismo Formation.

Santa Margarita,

Sandstone.

. Monterey Shale - 

... ..do... 

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

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

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

. Monterey Shale,

muddy phase.

. do _ -------

. Monterey Shale...

... do... .

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

. Repetto Siltstone

(lower Pliocene) .

Member of

Moaterey Shale.

. Valmoate Diatomite Member of

Moaterey Shale.

Member of San

Pedro Formation (lower

Pleistocene) .

mite Member of

Monterey Shale.

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

Altamira Shale

Member of

Monterey Shale.

Altamira Shale

Member of

Monterey Shale.

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

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

.--..-do. .._-_._._..

... ... do-... 

2- .......

4 . ... ...

2 -- _-_.

3 .-..--.

4 ... _____

5 .-.-...

6 -.--...

7 ... ..

8 .....

stone.

Limestone bed. .___ .__

stone 1.

Calcite cemented sandstone 2.

sandstone 2.

Dolomite bed...... -.._ _.

Dolomite-impregnated

tuff.

Dolomite concretion

zone.

Limestone lens ....

Limestone lens 1.. ........

0

Dolomite bed.... _____. 

Dolomite bed (100 ft

thick).

Dolomitized reef rock (40

ft thick).

Dolomitic upper part of

bed.

Calcitic lower part.... ....

Dolomite bed I..... ......

2.. .. .......

3 ........

Foraminiferal sand .......

Dolomite bed..

Dolomitized diatomite

(soft).

2 (8ft

thick).

Dolomite bed 3. ..........

4.... ...

5

6--.. -.__-_.

7 (dark

core) .

Dolomite bed 7 (bleached

exterior) .

Dolomite cement of glassy

+7.1

+11.1

-0.69

+9.5

+11.3

-1.4

+20.8

+11.1

+11.2

+ 11.2

+14.7

+8.2

+12.1

+11.4

-13.0

-14.7

-15.8

-14.4

-4-° Q

+19.3

-17.3

-2.7

-10.2

-10.1

+15.2

+5.3

-12.5

-13.1

-14.6

+12.5

+10.2

+10.1

+10.2

-5.6

-0.52

-15.4

-8.5

-3.3

-8.3

-11.7

-6.5

-5.2

-8.8

-11.3

-10.5

-10.2

-6.8

-14.9

28.9

30.8

27.6 .--

29.3

31.5

27.4 .-_

33.8

32.0

34.0

32.4

34.6

33.1

32.5

32.9

23.7 ...

30.3 --.

21.1 ...

20.8 ...

30.5 --.

31.6

26.5

27.0

27.4

24.0 --.

28.8 ...

24.6 ...

32.6

30.6

28.0

31.4

27.3 ...

34.9

34.2

30.2

30.8 ...

34.3

33.3

27.5 _._

34.1

31.6

33.4

32.6

32.4

30.9

28.1

28.5

28.9

28.4

56

KA

55

53

49

49

50

52

51

50

51

49

50

49

49

53

56

55

56

55

55

53

55

56

54

55

52

52

55

57

56

50

50

50

50

Do.

Do.

Do.

Do.

Do.

1940; McMasters,

1947, fig. 36.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Godde, 1728; Addicott, If 65.

Do.

Do.

Do.

Do.

J. G. Vedder, oral

commun., September 1967.

Hall and Surdam,

1967.

Do.

Do.

J. G. Vender, oral

commun., September 1967.

Do.

Do.

Do.

Dibblee, 1950.

Do.

Dickinson and Lowe,

1966.

Do.

Durrell, 1954; R. H.

Campl «»11, oral

commun., September 196'7 .

Do.

Do.

Woodring and others,

1946.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

Do.

tuff.
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON Bll

TABLE 1. Carbon and oxygen isotopic composition of diagenetic carbonates and calcium carbonate content of diagenetic dolomites in 

Miocene rocks of Oregon and California and of comparison materials Continued

Sample Lab. 

No.

Locality Latitude and longitude Formation Rock SOW aoi8 CaCQs in 

per mil per mil dolomite

Reference

75........ 3276-1 ...--do-.-....---.--..-.- 33°44'N., 118°21'W..... Intrusive basalt... Drusy dolomite in frac- -18.1 25.1

tures. 

76........3276-4 San Juan Capistrano, 33°35' N., 117°40' W..-_- Niguel Formation Dolomite cement of sand- -15.3 33.5

Orange County. (upper Plio- stone.

cene).

77._ ...3279-8 San Clemente Island.... 33°00' N., 118°33' W__... Monterey Shale-.. Dolomite bed....... ..... -1.8 33.0

78.... ... 3279-11 La Jollasubmarine 32°55' N., 117°34' W..... Cored sediment .....do.---.-------........ -25.3 34.6

canyon. EM3-5(8-17)

(cm).

79........ 3276-19 Coronado submarine 32°31' N., 117°19' W..... Rock broken ofl .....do.................. +12.1 33.4

canyon. canyon wall at

200 fathoms.

54 Do.

53 Vedder and others, 

1957.

57 Olmstead, 1958.

51 Inman and Goldbe g, 

1963.

52 J. G. Vedder, oral

commun., September 1968.

GENERAL ASPECTS OF THE ISOTOPIC COMPOSITION

The extreme values reported in table 1 are tabulated 

as follows :

SC^or &O l*permil Sample Material

Heaviest carbon________ +20.8 PDB 33 Dolomite.

Lightest carbon________ 25.3 PDB 78 Dolomite.

Heaviest oxygen________ 35.4 SMOW 6 Dolomite.

Lightest oxygen._______ 19.5 SMOW 26 Vein calcite.

As previously noted (Murata and others, 1967), the 

most remarkable aspect of these values is the heaviness 

of the carbon in many of the samples. Carbon +10 per 

mil and heavier has rarely been found in carbonate 

rocks. Yet table 1 contains 30 samples with such heavy 

carbon from the Monterey Shale and related formations. Twenty-eight of these samples are dolomite but 

sample 51 (SC13 of +15.2 per mil) is limestone and so 

is sample 57 (SC13 of +10.2 per mil). These two samples 

indicate that heavy carbon is not restricted to dolomites 

and can be generated independently of the dolomitization process.

The total range of carbon isotopic composition (47.1 

per mil) is so large as to be inexplicable in terms of any 

single reaction involving geologically reasonable conditions and materials. Of the several theoretically possible ways of producing such large differences in isotopic 

composition (Kemp and Thode, 1968), we prefer to 

invoke two different reactions, namely, nonequilibrium 

oxidation of organic matter over a wide range of temperature to obtain light-carbon carbonate, and equilibration of carbonate with methane at moderate to low 

temperatures to obtain heavy-carbon carbonates.

Oxygen isotopic composition of most samples falls, 

in the range of SO18 from 26 to 34 per mil. This range 

would be considered normal for marine carbonate (Lowenstam and Epstein, 1954) and would represent seawater temperatures of 30° to O°C. However, for diagenetic carbonates these values do not relate in a simple 

way to the temperature of formation waters, as will be 

discussed later. The oxygen and carbon isotopic compositions of many, but not all, of the samples listed in 

table 1 are plotted in figures 6 A and 65, in order to 

illustrate several modes of compositional variations

that characterize different materials or the same material from different localities.

NORMAL MARINE CARBONATE

Calcareous tests and shells of Miocene, Pliocene, and 

Pleistocene Foraminifera (samples 21, 22, 29, 32, 61, 

and 64) of a Pliocene barnacle (sample 14) and of a 

Miocene pecten (sample 45) were analyzed incidental 

to the study of diagenetic carbonates of several localities. Their compositions are plotted as open circles in 

figure 6J.; all samples lie within the rectangle at a. 

which defines the composition of normal marine limestone. The close grouping of compositions in the range 

of normal marine limestone suggests that the shells 

were secreted in virtual isotopic equilibrium with sea 

water and that the original composition is largely 

retained.

That the original composition may not have been 

retained perfectly is suggested by the rough alineme?it 

of points along the dashed diagonal line in figure QA. 

This line marks a trend of secondary alteration of marine limestone by circulating fresh water, as found in 

the Bermudas by Gross (1964). The identical trend is 

also noted for shells deposited in the Gulf of Mexico 

at various distances from land (Keith and Parker, 

1965). Thus terrestrial waters and organic matter may 

have an appreciable influence on the primary secretion 

of marine carbonate even many miles offshore. Such 

complications discourage the use of oxygen isotopic 

composition to derive precise paleotemperatures. They 

do not obscure the fundamental fact, however, that these 

calcareous fossils represent primary marine carbonate 

and serve as a reference against which secondary carbonates can be compared.

In contrast to the shells, only a few (three out of 71) 

diagenetic carbonates have a composition in the range 

of normal marine limestone. This striking compo^itional distribution divides the diagenetic carbonates 

naturally into a light-carbon group and a heavy-carbon 

group and has led us to treat the two groups in terns 

of different isotopic reactions.
B12 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

-20

+ 22 r-r~

+ 20

+ 16

+ 12

+ 8

+ 4

-16 -12 -4

SO 18 PER MIL (PDB) 

+ 4+6 -20 -16 -12 -4 + 4 +6

-4

-12

-16

-20

-24

Fresh-water 

limestone

45

) Marine 

O limestone

V0*

/ ce «\ C?2 / 56^T )

/ -"/i^/^,3 5:°>^V9

/// * / 

/ /25^ , / '

52 0' ^t/ /

/ / /

/

026

Fresh-water 

limestone

44o

43°

 10 " i

4o

o50 . 

o41 . e

 °42 -

,78

12 16 20 24 28 32 36 8

ao 18 PER MIL (SMOW)

12 16 20 24

B

28 32 36

+ 22 

+ 20

+ 16

+ 12

+ 8

+ 4

-4

-12

-16

-20

-24

-1 -28

6. Relationship between isotopic compositions of carbon and oxygen in carbonates. Black circles denote dohmite with 

positive values of SO13 ; black squares, dolomite with negative values of SO13 ; open circles, calcite. Compositional limits of 

marine and fresh-water limestones are from Graf (1960), and Keith and Weber (1964). The grouping of sampler in parts 

A and B, including a, 6, c, di, fa, and 0, is1 discussed in the text The meaning of dashed lines is discussed in text and in 

the caption to figure 11.

DIAGENETIC LIMESTONE

Beds and concretionary horizons of diagenetic limestone are considered by Bramlette (1946) and others to 

have formed from normal marine carbonate that was 

dispersed throughout a formation dissolved, and then 

redeposited in certain favorable zones. The fact that the 

resultant diagenetic limestone does not have the carbon 

isotopic composition of normal marine carbonate throws 

light on the nature of the formation waters that dissolved and transported the carbonate. Diagenetic. limestone is rare among our samples, but its isotopic composition is of great interest because the limestone was 

probably a precursor of the more abundant dolostone.

The nine diagenetic limestones shown as open circles 

in figure QB differ widely in isotopic composition, which

is to be expected for a precursor to dolostones that have 

equally variable composition. Samples 41, 43: and 44 

from Bound Mountain oil field show the clearest relationship between isotopic composition and geological 

environment. They represent calcite-cemented horizons 

in nearshore marine sandstones that were deposited on 

a shelf marginal to the landmass of the Sierra Nevada 

(Brooks, 1952). A thin section of sample 44 is illustrated as figure 5/. The isotopic composition of the calcite cement is like that of fresh-water limestone and indicates equilibration with, if not deposition from, fresh 

water probably of artesian circulation. At soire places, 

the cemented zones contain abundant macro fossils (Addicott, 1965), including pectens (sample 45 of fig. 6J.) 

with the isotopic composition of normal marine car-
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON B13

bonate. Fresh-water solutions deposited light-carbon 

light-oxygen, calcite all around the fossils shells but did 

not alter the isotopic composition of the shells.

The involvement of fresh water is strongly supported 

by the prevalence of waters of unusually low salinity 

(<4,000 parts per million) in Bound Mountain and 

other nearby oil fields (Jensen, 1934). The low salinity is 

attributed to extensive circulation of runoff from the 

Sierra Nevada through the sedimentary formations of 

these fields. Such a circulation could have begun early 

in the sedimentary history of the region, flushed out any 

connate water from the sediments, and prevented the 

development of highly saline brines like those found in 

most other California oil fields. A present-day example 

of this kind of artesian circulation is provided by the 

fresh water that was found during recent drilling on the 

continental shelf, 45 kilometers off the Florida coast 

(JOIDES, 1965),

Besides accounting for the isotopic composition of the 

calcite cement of sandstone samples 41, 43, and 44, the 

presence of fresh water would also explain why this 

calcite escaped the almost universal conversion of diagenetic limestone into dolostone. According to present 

views (Friedman and Sanders, 196T), dolomitization 

takes place most commonly in the presence of concentrated brines of appropriate composition.

Sample 42 (fig. 6Z?) (also from Round Mountain oil 

field) is from a bed of rather pure limestone associated 

with diatomite instead of sandstone. Although it has 

carbon as light as that of the three calcite cements, its 

oxygen is too heavy (SO18 of 30.5 per mil) to have been 

derived from fresh water. The origin of this limestone 

will be included in a later discussion of the many dolostones that have similar composition.

Another limestone with a fresh-water compositional 

aspect (sample 50) is from the late Miocene Santa 

Margarita Sandstone of Cuyama Valley. Little information is available on the paleogeography of this rock. 

But the Santa Margarita Sandstone is certainly a nearshore marine deposit that, in the region of Cuyama 

Valley, grades into the continental Quatal Formation of 

Hill, Carlson, and Dibblee (1958). These stratigraphic 

relations suggest that fresh water could have easily 

entered the sandstone to form diagenetic limestone.

Sample 26 is coarse sparry calcite that fills cracks in 

a bed of diagenetic dolomite. It appears to be a very late 

secondary product deposited from percolating ground 

water an origin that would account for the freshwater aspect of the isotopic composition.

Unlike the limestones described above, the three remaining samples (Nos. 51, 52, and 57) from muddy 

facies of the Monterey Shale contain heavy carbon with 

SC13 of + 5.3 to +15.2 per mil and thus must be related

to fluids very different from simple fresh water. The 

origin of such heavy-carbonate will be discussed in a 

later section dealing with heavy-carbon dolostoiies.

The isotopic composition of the calcitic materials can 

be summarized at this point. Tests and shells of fossils 

show a narrow range of carbon and oxygen isotopic 

composition that corresponds to that previously established for normal marine carbonate. Shell calcite is a 

primary precipitate whose isotopic composition is narrowly controlled by equilibria between solid calcite and 

sea water, relative to carbon and oxygen.

Dissolution of the primary calcite and redeposition 

in favorable zones during diagenesis form beds of secondary limestone, whose carbon isotopic composition is 

highly variable and generally lies outside the range for 

normal marine carbonate. Fresh water was involved in 

the formation of diagenetic calcite with light carbon 

and oxygen, especially that occurring in sandstone of 

iiearshore facies. Other fluids, to be discussed in the following sections, produced diagenetic limestone with 

normal oxygen and either heavy or light carbon. TT ie 

few samples of diagenetic limestone vary almost as much 

in their isotopic composition as the more abimdrnt 

samples of dolostoiies. This similar variability suggests 

that limestones could be a precursor to dolomite and tl at 

dolomite could inherit the isotopic composition of the 

precursor (Degens and Epstein, 1964). Associated primary calcite and dolomite could likewise have similar 

oxygen composition if their isotopic fractionation 

factors relative to water were virtually the same 

(Friedman and Hall, 1963).

OXYGEN ISOTOPIC COMPOSITION OF DOLOMITE

Figure 6 (A and B) well illustrates the striking fact 

that although the carbon isotopic composition of the 

dolomite (and also of diagenetic calcite samples 42, 51, 

and 57) is decidedly not that of normal marine carbonate, the oxygen composition lies mostly in the range 

of normal marine carbonate (SO1S =26 to 34 per mil). 

One possible explanation for these isotopic relationships 

might be that the diagenetic carbonate inherited the 

oxygen composition of primary marine calcite, but assumed a widely different carbon isotopic composition. 

Many examples of coexisting low-temperature calcite 

and dolomite show the same oxygen (and carbon) isotopic composition in both minerals (Degens and 

Epstein, 1964), and calcite has been converted to dolomite in laboratory experiments without developing a 

difference in oxygen composition (Epstein, Graf, and 

Degens, 1964).

On the other hand, results of other studies (Clayton 

and Epstein, 1958; Weber, 1964; Northrop and Clayton, 

1966; O'Neil and Epstein, 1966; Clayton, Jones and
B14 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

Berner, 1968) suggest that for dolomite and calcite in 

equilibrium at 25 °C the oxygen of dolomite is 4 to 7 per 

mil heavier, and the amount of the difference varies 

inversely with temperature. Figure 6 (A and B) indicates that dolomite on the whole tends to be richer in 

O18 than calcite. To consider this point more closely, we 

tabulated the oxygen isotopic composition of some of 

the most closely associated calcite and dolomite, as 

follows:

Locality

Samples 

in table 1 Material «C13 per mil 5O 18 per mil per 

mil

Reliz canyon..... 21 and 22._ Foraminiferal 1.9to+0.03, 28.0to 28.4, 3.4

calcite. avg 0.9 avg 28.2 

23 and 24.. Dolomite........ +11.6 to+13.5, 31.3 to 31.9,

avg +12.6 avg 31.6 

Nacimiento Dam. 29 and 32.. Foraminiferal 0.7 to 1.4, 27.4 to 27.6, 3.7

calcite. avg 1.0 avg 27.5 

27 and 30.. Dolomite. .... +11.1 to +11.3, 30.8 to 31.5,

avg 11.2 avg 31.2 

Pine Mountain.... 57._.------ Limy part of +10.2 27.3 4.1

bed.

56-_--.-... Dolomitic part +12.5 31.4 

of bed.

*Extent to which 5O 18 of dolomite is greater than that of associated calcite, computed from the averages given in table.

For the Reliz Canyon and Nacimiento Dam localities, pairs of dolostones richest in O18 were averaged 

to compute the maximum difference in O18 content between dolomite and associated foraminiferal calcite. 

At these two localities, beds of diagenetic dolostone 

formed through the dolomitization of certain zones of 

a thick section of foraminiferal shale, and it seems reasonable to postulate that calcite of Foraminifera was 

a major precursor of the dolomite. The two samples 

from Pine Mountain are from different parts of a single 

carbonate bed (fig. 2(7) and thus constitute the most 

closely associated calcite-doloinite pair.

Samples from all three localities show that the oxygen of the dolomite can be about 4 per mil heavier than 

that of the associated calcite. This may be an expression of the isotopic equilibrium between the two minerals, or it may be that dolomite has accurately preserved the oxygen composition of primary marine calcite while the associated foraminiferal calcite had its 

composition altered by the action of isotopically light 

meteoric water. The superior resistance of dolomite to 

such alteration effects has been emphasized by Epstein, 

Graf, and Degens (1964).

Although the oxygen composition of both calcite 

and dolomite is within the range possible for normal 

marine carbonate, the 4-per-mil difference indicates very 

different temperatures (Epstein and others, 1953) 22° 

to 26°C for calcite and 7° to 10°C for dolomite. With 

such Foraminifera-bearing sediments, an independent 

estimate of the original temperature can be derived 

from existing knowledge of depth of habitat of Foraminifera (Kleinpell, 1938; Bandy, 1953; Natland, 1957;

Smith, 1964). According to Patsy J. Smith (written 

commun., July 1967 and February 1968), Foraminifera 

from all three localities indicate a bathyal environment 

with a probable temperature of 7°C or less, clearly 

closer to dolomite temperatures than to those of calcite.

The oxygen of dolomite from Palos Vercles Hills 

(samples 62, 63, 65 to 74 of table 1) tends to be heavier 

the higher the position of the bed in the section (fig. 7). 

Sample 75 (drusy dolomite) is excluded from the figure 

because its oxygen is abnormally light. All carbonates 

collected from the Miocene of this locality proved to be 

dolomite, so that a comparison with calcite is not possible. Translated into paleotemperatures, the oxygen 

composition of dolomite indicates a general lowering of 

water temperature during middle through late Miocene 

time (from the middle part of the Altamir^, Shale 

Member to the Malaga Mudstone Member of the Monterey Shale) ; this trend is in agreement witl the regional trend previously recognized by others (Woodring and others, 1946; Natland, 1957) through studies 

of Foraminifera and other fossils in these sedimentary 

rocks.

The temporal trend of water temperature (8° to 3°C) 

deduced from the nature of fossils (Natland, 1957) is 

shown by the dashed line in figure 7; the oxygen isotopic 

results agree reasonably well with the trend, except 

those of samples 72 to 74 from the lowest pait of the 

sampled section. It is possible that the isotopic composition of the lower dolomite layers was affected by 

nearby underlying igneous intrusions (Macdonald, 

1939; Woodring and others, 1946; Schleicher, 1965) 

a possibility that merits further study. Thus, dolomites 

from Palos Verdes Hills, like those previously discussed 

from other localities, seem to yield roughly correct isotopic temperatures in situations where associated foraminiferal calcites give too high a temperature. However, as will be shown in following sections, dolomite 

temperatures are not always correct; they appear to be 

most accurate for low temperatures (<10°).

Some dolomite cannot be interpreted in terms of inheritance of oxygen composition from a primary 

marine carbonate; therefore this process may not be 

generally valid. For example, the upper Plioce-ne sandstone of sample 76 (table 1) was deposited in shallow 

water (probably less than 400 feet deep according to 

J. G. Vedder oral coinmun., February 1968), yet the 

isotopic temperature derived from the dolomite cement 

of the sandstone is 1.5°C obviously too low to be explained in terms of temperature of depositional or postdepositional environment. Presumably, the original calcite cement had a value of SO18 appropriate for shallow 

waters, and the value increased during dolomitization.
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON F15

c 

 s

|||

Valmonte Diatomite Member

Member ; 

Altamira Shal

QQ. 

QT3

5

25 20 15 10 5 0

!" 

! 

 1

1

I 

1

1

. 1

1 

1

1

1 

/

1 i 1 i 1 i 1 i 1

27 29 31 33 35

3018 PER MIL

FIGURE 7. Temperatures indicated by the oxygen isotopes of 

dolomites from the Monterey Shale of Palos Verdesi Hills. 

Trend of temperature (Natland, 1957), inferred from nature 

of fossils in the Monterey Shale of the region, is shown by 

the dashed line.

Our dolomite samples 6, 37, 58, and 78 have SO18 

values of 34.6 to 35.4 per mil, signifying apparent temperatures of 1° to - 3°C. Kussell, Deifeyes, Fowler, 

and Lloycl (1967) described a dolomite with SO18 of 35.5 

per mil from an offshore environment similar to that of

our sample 78. The indicated temperatures are so low 

as to be questionable. Here again is a suggestion tl x at 

dolomite attains a somewhat higher content of O18 tl an 

calcite a result that could be achieved in more than 

one way. We shall next discuss the possibility that the 

dolomitizing solutions are generally richer in O18 ttaii 

normal sea water. The previously mentioned equilibrium relationship (O'Neil and Epstein, 1966) that 

would automatically make the oxygen of dolomite substantially heavier than that of accompanying calcite 

seems to be too inflexible a mechanism for our purpose.

NATURE OF FORMATION WATER AND ITS INFLUENCE 

ON OXYGEN ISOTOPIC COMPOSITION OF DOLOMITF

The control exerted by the oxygen isotopic composition of water 011 the composition of a carbonate precipitated from it is indicated by the well-known isotopic 

difference between marine and fresh-water limestone. 

The diagenetic carbonates are genetically related to 

some kind of formation water a most variable class 

of natural water, ranging in chemical composition from 

fresh water to brine several times more concentrated 

than sea water (Jensen, 1934; Gullikson and others, 

1961; White and others, 1963). Salinity of format'on 

waters generally increases with depth of occurrence; the 

deep warm brines are thought to be derived from fresh 

water by a process of micropore filtration that occurs 

during passage of water through shales. It is the more 

saline waters (oil-field brines) that are of interest h°,re 

for two reasons: (1) Such brines are involved in the 

genesis of dolomite (Anderson and others, 1966; Fittenhouse, 1967), and (2) the water of many such brines 

is greatly enriched in O18 (up to +8.8 per mil compared 

to 0.0 of SMOW), according to Clayton and otters 

(1966). Brines of the Cymric oil field in California srow 

a SO18 range of +2.9 to +5.9 per mil at depths of 3,000 

to 5,000 feet (Harmon Craig in White, 1965).

The enrichment in O18 in formation waters has b?en 

ascribed to very different processes: (1) To partial evaporation of the water while still at the surface (Degens, 

1965, p. 193; Lloyd, 1966), and (2) to isotopic equilibration of the migrating water with the carbonate nid 

silicate minerals of sediments (Clayton and others, 

1966). The variability in the D/O18 ratio among oilfield brines suggests that both processes may be involved 

(Craig, 1963). Because carbonates are rather scarce in 

the Monterey Shale, formation water more likely equilibrates with silicate minerals such as clays or opaline 

silica, its SO18 varying directly with the temperature 

that prevails at different depths.

According to Clayton and others (1966), the rate of 

equilibration between water and calcite is much farter 

than between water and silicate minerals. This differ-
B16 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

ence raises the interesting possibility that a deep-lying 

water which had attained equilibrium with silicate minerals, might move upward into cooler strata without an 

appropriate reduction of its O18 content and form a 

carbonate that was richer in O18 than would be expected 

from its temperature of formation. Presumably, such an 

effect based on differing kinetics of equilibration would 

be largest at the lower temperatures, say, less than 

100° .

The SO18 of calcite at equilibrium with water of different O1S content (SO18 of -6 to +9 per mil) at different 

temperatures has been computed from the equation of 

Epstein and Mayeda (1953) and plotted in figure 8. 

The compositional range (SO18 of 26.5 to 34 per mil) of 

normal marine carbonate, which includes the composition of most of the diagenetic carbonates, is delimited 

by the two dashed lines. Mean ocean water (SO18 of 0.0 

per mil) yields calcite in this range between 0° and 30°C. 

Heavier water, with SO18 of +6 per mil, yields the same 

calcite O18 composition between 23° and 67°C, corresponding, for example, to temperatures prevailing at 

depths of 100 to 1,000 feet in the sedimentary section of 

the Los Angeles basin (Philippi, 1965). Heavier water, 

therefore, shifts the equilibrium temperatures into a 

higher range, which could be more appropriate for the 

formation and preservation of a diagenetic mineral.

If SO18 of water were as high as +9 per mil (the 

maximum known value), calcite with normal marine 

isotopic composition would have equilibrium temperatures between -±0° and 88°C, which corresponds roughly 

to depths of 2,000 to 6,000 feet in the sedimentary section. Such a water could yield calcite with SO18 of 35.5 

per mil (the largest value on record for any carbonate) 

at 32°C, corresponding to a depth of about 1,000 feet. 

A restriction on considering heavier and heavier water 

is imposed by the fact that such waters increase the

120

100

80

60

40

20

44 42 40 38 36 34 32 30 28 26 24 22 

30 18 OF CALCITE PER MIL

FIGUEE 8. Temperature dependence of O18 content of calcite that 

is in isotopic equilibrium with water of different O18 content. 

The curves were computed from the equation of Epstein and 

Mayeda (1953), and MOW refers to their mean ocean water.

likelihood of producing, at lower temperatures, a carbonate with SO18 greater than 36 per mil, which does not 

seem to exist.

Figure 8 also shows that postulating a high temperature of equilibration (say, greater than 100°C) increases the chance of putting light oxygen (SO18 <20 

per mil) into diagenetic carbonates and of making them 

similar isotopically to carbonates of Irydrothermal origin (Engel and others, 1958; Lovering and others, 1963; 

Rye, 1966). However, the virtual absence of such light 

oxygen among our samples and the general mineralogy 

and organic chemistry of Monterey Shale make clear 

that the diagenetic carbonates were subjected to temperatures well below the hydrothermal range.

The observed enrichment of O18 in oil-field brines thus 

is of interest for at least two reasons: (1) It allows carbonate oxygen to become heavier than 34.5 per mil, 

which is not likely with water of normal O18 content, 

and (2) it counteracts the tendency of higher temperature to lower SO18 values in carbonates and yields values 

in the range of normal marine carbonate at temperatures 

higher than 30°C.

DOLOMITE WITH VARIABLE LIGHT CARBON AND 

INDEPENDENTLY VARIABLE OXYGEN

Dolomite of Palos Verdes Hills (samples 62, 63, and 

65 to 75 of table 1) is characterized by variable light carbon (SC13 of -18.1 to 5.2 per mil) and variable oxygen (SO18 of 25.1 to 34.3 per mil). The individual values 

are plotted as black squares in figure 6A at c. The SC13 

of a dolomite from this locality was previously determined to be -16.4 per mil by Spotts and Silverman 

(1966), who attributed the lightness of the carbon to 

the presence of much carbonate derived through the 

oxidation of petroleum and other organic matter. Thin 

sections of samples 62 and 66 are illustrated in figures 

55T and 5/, respectively.

No regular covariation of carbon and oxyger isotopic 

composition, such as the previously discussed freshwater alteration trend of the dashed line in figure 6J., 

is evident. Single dolomite samples containing light carbon from other localities are plotted as blacl" squares 

in figure QB at e. They collectively show the same irregular variations as the samples from Palo~ Verdes 

Hills, except that sample 78 extends the range of SC13 

down to 25.3 per mil.

Carbon lighter than about 5 per mil in diagenetic 

carbonate is derived most, simply by forming light carbon dioxide through the nonequilibrium oxidation of 

some form of organic matter such as humus or petroleum 

(Thode and others, 1954; Vogel, 1959; Spotts and Silverman, 1966). Virtually all the diagenetic carbonates 

of marine formations studied by Hodgson (1966) con-
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON B17

tainecl such light carbon; he also ascribed this light 

carbon to organic matter. The oxygen of his samples was 

mostly lighter than 26 per mil, suggesting extensive 

equilibration with fresh water. Only one out of 12 dolomites from Palos Verdes Hills contains such light oxygen (fig. 6A ate).

We follow previous investigators in believing that 

the light carbon of the dolomites under discussion was 

derived from sedimentary organic matter. Carbon dioxide, probably so derived and with SC13 as light as 38 

per mil, occurs in marine formations as a component of 

natural gas (Wasserburg and others, 1963) and could 

equilibrate with preexisting marine carbonate or become 

incorporated into carbonate precipitated diagenetically. 

Low-temperature equilibration with meteoric water 

could not have played an important part in the formation of these dolomites because that reaction changes 

the oxygen isotopic composition without greatly affecting the carbon composition (Degens and Epstein, 1962; 

Keith and Weber, 1964; Weber and others, 1965) or 

that reaction induces covariation of oxygen and carbon 

(Gross, 196-4; Gross and Tracey, 1966).

In the next section, we shall discuss how isotopic 

equilibrium between carbon dioxide and methane in 

natural gas leads to very heavy carbon in carbonates. 

Thus both light and heavy carbon carbonates are genetically linked to natural gas and the accompanying formation waters. The key point of difference is whether 

or not isotopic equilibrium prevails between carbon dioxide and methane; when it does not, the carbon in 

carbonate will be light.

When equilibrium prevails, the heaviness of the carbonate-carboii will depend on the temperature, and on 

the relative amounts of methane and carbon dioxide. 

Methane is by far the most abundant component of most 

natural gas, and only rarely would variations in the 

ratio of methane to carbon dioxide be expected to affect 

the isotopic composition of the carbonate.

DOLOMITE WITH VARIABLE HEAVY CARBON AND 

RELATIVELY CONSTANT HEAVY OXYGEN

Dolomites from Berkeley Hills (samples 7 to 9) and 

Chico Martiiiez Creek (samples 33 to 40) exemplify a 

class of diagenetic carbonate in which carbon is heavy 

and variable (SC13 = +5.4 to +20.8 per mil) and oxygen 

is heavy and relatively constant (SO18 = 32.0 to 34.6 per 

mil). These samples, plotted in figure QA at 5, are strikingly different from those of Palos Verdes Hills and 

raise the question of the genesis of heavy carbon.

A few examples of carbon dioxide gas with SC13 as 

heavy as +16 per mil had been found previously in 

natural gas by Wasserburg, Mazor, and Zartman (1963).

The heaviest carbon dioxide found so far in gas of California oil fields was from the Huntington Beach fiVd 

of the Los Angeles basin and had SC13 of +6 per mil 

(S. R. Silverman, written commun., November 1967). 

Because methane is the dominant constituent of natural 

gas, Zartman, Wasserburg, and Reynolds (1961) and 

Wasserburg, Mazor, and Zartman (1963) employed irotopic equilibrium between methane and carbon dioxide 

as a mechanism for obtaining such heavy carbon dioxide 

and hence heavy carbonate ion. Theoretical equilibrium 

constants computed by Bottinga (1968) show that C13 

will concentrate strongly in carbonate. At 25°C, carbon 

in calcite would be 78 per mil heavier than that in 

methane; at 100°C, it would be 53 per mil heavier. Unfortunately, the CHt-COs' 2 exchange reaction has rot 

been investigated experimentally and nothing is known 

about the kinetics of exchange. However, some samples 

of coexisting methane and carbon dioxide, to be discussed below, suggest that these compounds do equilibrate isotopically in nature.

Histograms of carbon isotopic compositions of coexisting methane and carbon dioxide from three different geological environments are presented in figure 9. 

Rows A, B, and G represent methane-rich gases generated through bacterial decay of organic matter at low 

temperature. Rows F and £?, on the other hand, represent gases that are relatively rich in carbon dioxide and 

are emitted from hot springs. Between these two groups, 

methane-rich oil-field gases are plotted in row Z>, and 

carbon from the diagenetic carbonates of this study, 

in row E.

Figure 9 shows that average SC13 of methane is abcnt 

 72 per mil if derived from low-temperature sources; 

about 44 per mil, from oil-field gases; and about 27 

per mil, from gases of thermal waters. These values 

suggest that the isotopic composition of methane alcne 

can serve as a crude index of temperature. The plotted 

points indicate that the contrast in isotopic composition 

between coexisting methane and carbon dioxide generally becomes less with rising temperature, a condition 

agreeing with the theoretical requirement for isotopic 

equilibrium (Craig, 1953; Bottinga, 1968). Hot spring 

temperatures, computed on the basis of the equilibrium 

between methane and carbon dioxide, average around 

300 °C, which falls well within the range observed or 

inferred at these springs. Isotopic temperatures of the 

cool sources of rows J., B, and O in figure 9 are of the 

order of 30°C. These, results substantiate the view tl at 

methane and carbon dioxide do equilibrate isotopically 

in nature and probably play a fundamental role in generating the heavy carbon found in so many of cur 

carbonates.
B18 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

G. THERMAL WATERS 

YELLOWSTONE PARK cD

F. THERMAL WATERS 

NEW ZEALAND

E. DIAGENETIC CARBONATES 

THIS STUDY

D. OIL FIELDS U.S.A.

Dnn tieO

C. GLACIAL DRIFT 

ILLINOIS

D do

B, LAKE MUDS 

CONN. AND N.Y. .-,

n Lb n D D i i n

A. GAS FIELDS n

i \\L

i i i i i i i i i i i

-90 -80 -70 -60 -50 -40

 

 I 

 

i 1 i I i 1 i

-30 -20 -10 (

 3 1

3 2

3 2

q 5 

 3 1

i 1 i 1 i

D +10 +20

3C 13 PER MIL (PDB STANDARD)

FIGURE 9. Histograms of carbon isotopic composition of coexisting methane (open figures) and carbon dioxide (black 

figures) from sources (rows A, B, C, D, F, and (?) of different temperature. Isotopic composition of carton in 

carbonates of the present study are given in row E as black circles. References for the data in the several rov^s are 

as follows:

A. Nakai (1960).

B. Oana and Deevey (1960). Similar data on methane

from Russian swamps are given by Ovsyannikov

andLebedev (1967). 

C. Wasserburg, Mazor, and Zartman (1963).

The isotopic composition of any methane that equilibrated with the cliagenetic carbonate is indeterminate, 

but all criteria suggest that it must have been like that 

of oil-field methane. In order to gain some insight into 

the temperature implications of the heavy carbon in the 

carbonate, we tabulate below the isotopic composition 

of four of the five heaviest carbon dioxides from oil 

fields (row D of fig. 9) along with the composition of 

associated methane. One other sample has been added at 

the bottom of the table for consideration with the four 

samples.

D. Zartman, Wasserburg, and Reynolds (1961) ; Wasserburg, Mazor, and Zartman (1963) ; S. R. Silverman 

(uupub. data, November 1967).

F. Hulston and McCabe (1962).

O. Craig (1953 and 1963).

Depth of SCi'of 5C»of Calc 

gas CH4 per COs per temp, 

source, mil mil in ° C 

in feet

Reference and sample

360 -50.9 +16.4

2,605 -41.5 +12.8

4,885 -39.7 +8.1

8,400 -38.3 -0.5

12,768 -38.5 -8.8

25 Wasserburg, Mazor, and Zartman (1963),

sample 22. 

72 Zartman, Wasserburg, and Reynolds

(1961), sample 38. 

101 Zartman, Wasserburg, and Reynolds

(1961), sample 9. 

162 Zartman, Wasserburg, and Reynolds

(1961), sample 41. 

233 Zartman, Wasserburg, and Reynolds

(1961), sample 39.

Zartman, Wasserburg, and Keynolds (1961) tested 

for isotopic equilibrium between methane and carbon
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON B19

dioxide in their samples of natural gas by using the reasonableness of the computed temperature as a criterion. 

They noted that reasonable temperatures were generally 

obtained for samples containing carbon dioxide that 

was rich in C13, such as those tabulated above; and unreasonably high temperatures, for samples containing 

carbon dioxide poor in C13 . These relations are made 

clear when the computed temperatures are plotted 

against the depth of origin of the gas samples, as in 

figure 10.

In figure 10, the reasonable temperatures obtained 

from the five tabulated samples are seen to be the minimum value for each depth. Collectively, the five samples define a geothermal gradient of 54° C per kilometer, 

which is above average and its matched or exceeded by 

gradients of only a few American oil-fields (Birch and 

others, 1942). Higher gradients involving the other 

samples are thus unlikely; the line in the figure probably signifies virtual attainment of the equilibrium between methane and carbon dioxide, and samples that plot 

away from the line are out of equilibrium. Samples out 

of equilibrium far outnumber those in equilibrium.

KILOMETERS FEET

12,000 -

16,000 L-

\ O 

\ 

\

\ \ \

100 150 200 

ISOTOPIC TEMPERATURE, °C

i I I I i i I 

250 300

FIGURE 10. Isotopic temperature and depth of origin of 

natural gas. Temperature was calculated from the carbon 

isotopic composition of methane and carbon dioxide in 

natural gas, assuming that these compounds were in isotopic equilibrium. Black circles represent the five samples 

tabulated on page B18. The dashed line through the five 

circles defines a geothermal gradient of 54°O per kilometer. 

Four samples with calculated temperatures greater than 

350°C are omitted. All data from Zartman, Wasserburg, 

and Reynolds (1961) and Wasserburg, Mazor, and Zartman 

(1963).

These relationships suggest that heavy-carbon carbonates, such as the ones under consideration, probably 

reached equilibrium with methane; therefore their temperature of equilibration could be determined if th^ 

isotopic composition of the methane were known. In 

figure 9, the great majority of oil-field methanes (row 

D) are seen to have 8C13 values between 52 and 3° 

per mil, a range that includes the five critical samples 

tabulated above. Therefore, using the lowest and highest of the values cited above for both methane and carbonate, we computed equilibration temperature for 

methane-carbonate couplets by means of the Bottinga 

equation (Bottinga, 1968) ; the results are tabulated 

below (values in brackets were computed for the specific temperature of 45°C).

SO" of CHi 

per mil

-52

[-48]

[-52]

-38

-52

-38

SCis of CO3-3 

per mil

+ 21

[ + 21]

[ + 17]

+ 21

+ 5

+ 5

Computed temp, in°C

34

[45]

[45]

76

80

140

The range of 34° to 140°C in the tabulation represents the best estimate that can now be made of the 

temperature of equilibration between methane and the 

heavy-carbon carbonates (8C13 > +5 per mil). Usin^ 

the average geothermal gradient of the Los Angeles 

basin (39.1°C per kilometer according to Philippi, 

1965) for illustration, these temperatures would correspond to depths of 1,100 to 10,000 feet. As the 8C13 of 

the carbonate becomes more negative, the indicated temperature becomes higher; carbonate with 8C13 of 4,0 

per mil and methane with 8C13 of 38 per mil would 

be at equilibrium at 192°C, corresponding to a depth of 

about 14,500 feet.

Judging from the temperature data on couplets of 

methane and carbon dioxide from oil field gas (fig. 10), 

very few, if any, of our carbonates would have been 

subjected to temperatures higher than about 190°C. TH 

conclusion seems inescapable that carbonates with 8013 

more negative than about 5 per mil could not have 

equilibrated with much methane. They, therefore, constitute the separate class of light-carbon carbonate previously discussed, whose carbon was largely derived 

through nonequilibrium oxidation of organic matter 

and whose carbon isotopic composition has no relevance 

to the temperature of formation. An upper limit of 

about 190 °C would also be in harmony with the fact 

that a great many of the dolomites hold excess calcium, 

which would exsolve at around 200°C (Goldsmith, 

1959).
B20 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

The computed temperatures of 34° to 140°C in the 

tabulation given above and the correlative depths of 

about 1,100 to 10,000 feet are only crude indications of 

the conditions under which these dolomites could have 

formed. In detail, no systematic variation of SC13 with 

stratigraphic position is observed at the two localities 

concerned. At the Berkeley Hills (samples 7 to 9), the 

sampled interval was only about 100 feet, corresponding to a temperature difference of only 1°C along a 

normal geothermal gradient, whereas the spread of temperature indicated isotopically (assuming methane of 

constant composition) is as high as 26°C. At Chico 

Martinez Creek (samples 33 to 40) the sampled section 

was 2,200 feet thick, which would allow a temperature 

contrast of around 25 0|C; the difference between minimum and maximum carbon-isotopic temperature is 

41°C.

These apparent temperature discrepancies among 

heavy-carbon dolomites are resolved if it is assumed 

that equilibration between methane and a normal marine 

carbonate either proceeded at variable rates or involved 

natural gas with variable content of methane. Such conditions could yield carbonates manifesting all compositions between some maximum value of SC13 for the locality and the value of around zero for the original 

marine carbonate. Thus, the spread of isotopic values 

at each place would be a measure of the kinetic or concentration limitations on isotopic exchange rather than 

an indication that the exchange process had occurred 

at widely different temperatures.

The heavy carbonates under consideration in this section (points in envelop© b in fig. QA) contain heavy 

oxygen, with SO18 of 32.0 to 34.6 per mil and an average of 33.2 per mil values which would limit the 

maximum temperature to about 45°C (according to 

fig. 8) if oxygen isotopic equilibrium were also to be considered. In this case, SC13 of methane must lie between 

 52 and 48 per mil; SO18 of formation water, between +7 and +9 per mil; and the temperature, between 

34° and 45 °C, as indicated in the tabulation given above. 

Only SC13 values between +17 and +21 per mil in carbonates would then represent equilibrium compositions. 

These conditions are so restrictive that we rejected attainment of oxygen isotopic equilibrium as a necessary 

condition in the final equilibria involving these 

carbonates.

CARBONATE WITH VARIABLE HEAVY CARBON AND 

COVARIABLE OXYGEN

Examples of this isotopic category are grouped together in the upper right side of figure QB. Group d^ 

consist of samples of a slightly variable dolomite from 

the Monterey Shale of Eeliz Canyon (samples 15 to 20

of table 1). Its average SO18 value of 31.0 per mil is too 

light for the group to have been included in the heavycarbon heavy-oxygen group of figure GA at &. Group d« 

consists of samples from a lower member of the Monterey 'Shale in Reliz Canyon and in a section near Nacimiento Dam (samples 23 to 25,27,28, 30, and 31 of table 

1); these samples are isotopically more variable than 

those of group di, carbon varying directly with oxygen. 

Paleotemperatures of the four samples with the heaviest oxygens in group d~ were discussed on page Bll. 

Other representatives of this isotopic category are three 

from Oregon (samples 1 to 3) and pairs of samples 

(Nos. 10 and 11, 51 and 52, 56 and 57, and 59 and 60) 

from several localities.

A simple explanation for the covariation of carbon 

and oxygen isotopic composition would be to have these 

carbonates undergo exchange reactions with meteoric 

water. The presence of late fresh-water calcite (sample 

26) in cracks in dolomite sample 25 would corroborate 

the view that this dolomite was more affected by meteoric 

water than any others of group 6?2 , and that th?, isotopic 

trends within the group wT ere due to the variable degree 

of equilibration of an originally heavy-carbon heavyoxygen carbonate with meteoric water. The same action 

of meteoric water was already used on page Bll to 

explain the discordant paleotemperatures within the 

dolomite-calcite pair of samples 56 and 57. However, 

the calcium content of the dolomites under consideration 

in this section cannot be reconciled with the postulated 

reaction with meteoric water, as will be shown next.

CALCIUM CONTENT OF DOLOMITE AS BELATED TO 

ISOTOPIC COMPOSITION

The calcium content of the dolomites is plotted against 

the carbon isotopic composition in figure 11, and several 

significant relationships are apparent. Calcium appreciably in excess of the 50 mol percent of ideal dolomite is 

common among both heavy-carbon and light-carbon 

dolomites. Graf and Goldsmith (1956) and Goldsmith 

(1959), who gave the name protodolomite to dolomites 

of such nonideal composition, have shown that protodolomite is a widespread and persistent metastable phase 

in sedimentary rocks. It can be made to invert to stable 

ideal dolomite by heating to temperatures greater than 

about 200° C.

Among the samples from Palos Verdes Hills (black 

circles in fig. 11) only samples 72 to 74 are ideal dolomites. Previously, in figure 7, these very samples were 

seen to have deviant oxygen compositions deviations 

which were explained in terms of possible heating 

effects of nearby igneous sills; the same effects could 

have accelerated the inversion to normal dolomite. 

Otherwise, no relationship between isotopic arrl calcium
DIAGENETIC CARBONATES, MIOCENE FORMATIONS, CALIFORNIA, OREGON B21

+ 22

+ 20

O

O

 O

o

o

o o

Q.

O

O

49 50 51 52 53 54 55 56 57 58 59 

CaC03 IN DOLOMITE. MOL PERCENT

FIGURE 11. Carbon isotopic composition as related to the 

CaCOs content of a diagenetic dolomite. Groups tfi and d-s 

are the same as those so labeled in figure QB. Numbered 

triangles, squar.es, and open circles that are linked by tie 

lines denote the same dolomite samples as those linked by 

the dashed lines in figure 65. Black circles represent samples 

from Palos Verdes Hills; open circles without tie lines, 

samples from all other localities.

compositions is apparent among light-carbon dolomites 

in figure 11.

Among the heavy-carbon dolomites in the upper 

half of figure 11, group d2 outlines an inverse trend 

between SC13 and the calcium content. Other heavycarbon dolomites (samples 1 to 3,10 and 11, and 59 and 

60) also show the same trend. The increase in calcium 

content with decrease in SC13 might be interpreted to 

mean that ideal dolomite alters into protodolomite in the 

zone of weathering a direct contradiction of earlier 

work (Graf and Goldsmith, 1956), which indicated 

that ideal dolomite, not protodolomite, is the stable 

phase under near-surface conditions.

The contradiction is removed if the trends in question are illustrating the reverse process, namely the 

transformation of protodolomites into ideal dolomites 

at moderate depths, the carbonate equilibrating with 

the carbon of methane and the oxygen of oil-field brines. 

The lack of comparable trends among light-carbon 

dolomites would be due to the general failure of these 

isotopic equilibria, which causes the protodolomitedolomite transformation to proceed without any reference to isotopic compositions.

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B24 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

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U. S. GOVERNMENT PRINTING OFFICE : 1969 O - 334-795