Geochemistry of Diagenetic Dolomites 

in Miocene Marine Formations 

of California and Oregon

GEOLOGICAL SURVEY PROFESSIONAL PAPER 724-C

Geochemistry of Diagenetic Dolomites 

in Miocene Marine Formations 

of California and Oregon

By K. . MURATA, IRVING FRIEDMAN, and MARCELYN CREMER 

SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

GEOLOGICAL SURVEY PROFESSIONAL PAPER 724-C

The alteration history of deeply buried marine 

sediments as revealed by the chemical and 

isotopic compositions of secondary dolomites

UNITED STATES GOVERNMENT PRINTING OFFICE, WASHINGTON : 1972
UNITED STATES DEPARTMENT OF THE INTERIOR 

ROGERS C. B. MORTON, Secretary

GEOLOGICAL SURVEY 

V. E. McKelvey, Director

Library of Congress catalog card No. 78-189668

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

Washington, D.C. 20402 - Price 30 cents (paper cover)

Stock Number 2401-2062
CONTENTS

Abstract. __________________________________________

Introduction _______________________________________

Acknowledgments__ ___________________________

Methane-carbon dioxide isotopic exchange in natural

gasFormation waters involved in dolomite genesis. 

Chemical analysis of diagenetic dolomites_____

Page 

Cl

1

2

The d(10-4) X-ray spacing and composition. ___________

Interrelation of chemical and carbon-isotopic composiLight-carbon dolomites.. ________

Heavy-carbon-low-iron dolomitesHeavy-carbon-high-iron dolomites 

References cited. _______________

Page 

C6

9

9

10

11

ILLUSTRATIONS

FIGURE 1. Graph showing relation of isotopic temperatures and compositions of coexisting carbon dioxide and methane to depth of 

origin of natural gas-___________-______

2. Ternary diagram showing composition of 

diagenetic dolomites_________________

3. Diagram showing the principal d(10-4) X-ray 

spacing of dolomite and related carbonates_-_______________________________

Page

C3

4. Graph showing relation between the d(10-4) 

X-ray spacing and the apparent CaCO3 

content of dolomite_______--______-__-

5. Graph showing relation between carbon-isotopic composition of dolomite and d(10-4) 

spacing, mol percentage of CaCO3 , and 

mol percentage of FeCO3 _____________

6. Graph showing relation of oxygen-isotopic 

composition of dolomite and the content 

of FeCO3 to the carbon-isotopic composition.-.. __--------------_-_____----_

Page

C7

10

TABLE

TABLE 1. Partial chemical analyses of diagenetic dolomites.

Page 

C6

m

SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

GEOCHEMISTRY OF DIAGENETIC DOLOMITES INiMIOCENE 

MARINE FORMATIONS OF CALIFORNIA AND OREGON

By K. J. MTJKATA, IEVING FEEEDMAN, and MAKCELYN CKEMER

ABSTRACT

The unusually high concentration of C1S (5C13>+5 per mil 

relative to the Peedee belemnite standard) in many diagenetic 

dolomites of California and Oregon Miocene formations was 

brought about by isotopic exchange between methane and carbonate. Computations based on the temperature-dependent 

equilibrium constant for this exchange reaction yield 58°-110°C 

as the approximate temperature of formation of the dolomite.

Chemical composition of the dolomite lies within the following 

limits, expressed in mol percent: FeCOs, 0.0-16.6; MnCOs, 

0.0-0.5; MgCOs, 27.9-51.1; and CaCO3 , 47.&-58.0. The d(10-4) 

X-ray spacing is related to the composition according to the 

equation

d(10-4)angstroms=2.8855+0.00301[ (CaCO3+0.19 FeCO3 

+0.37 MnCOs) 50.0].

Virtually all light-carbon dolomites (SC1S< 5 per mil relative 

to the Peedee belemnite standard) and some heavy-carbon 

dolomites are poor in iron, because of the generally low concentration of iron in interstitial water of pyrite-bearing reduced 

sediments. Other heavy-carbon dolomites contain as much as 17 

mol percent FeCOs, and GaCOs in these samples increases with 

FeCOs, whereas SC13 and SO18 decrease. The chemical and isotopic 

interrelations among high-iron dolomites were probably brought 

about by postorogeny invasion of the reduced sediments by 

initially oxygenated meteoric water.

The overall chronological succession of dolomite types, 

throughout the history of a sediment rich in organic matter, 

seems to be (1) light-carbon dolomite, (2) heavy-carbon-lowiron dolomite, and (3) heavy-carbon-high-iron dolomite. Types 

1 and 2 belong to the long anaerobic phase of sediment history, 

and type 3 marks the first stage of the postorogenic aerobic 

phase.

INTRODUCTION

Diagenetic carbonates in Miocene marine formations 

of California and Oregon have an unusually wide range 

of carbon-isotopic composition and are divisible into a 

light-carbon group (SC13 of 25 to 5 per mil relative 

to the Peedee belemnite (PDB) standard) and a heavycarbon group (SC13 of + 5 to +21 per mil) (Murata and 

others, 1969). The starting material for these carbonates 

is believed to be minor amounts of marine carbonate

(SC13 of 0±4 per mil) in the form of shells of Foraminif era and other calcareous organisms. A long series of 

diagenetic reactions involving decomposing organic 

matter and formation waters converted primary carbonate into lenticular and concretionary masses of diagenetic calcite and dolomite.

Dolomite is by far the most abundant mineral 

among diagenetic carbonates in the California Miocene, 

and it commonly occurs along restricted horizons in 

moderately calcareous shales that were deposited in 

deep marine basins. Dolomitization apparently was 

effected by solutions that migrated mostly along certain 

favorable zones in consolidated sediments. The mode of 

origin was very different from that of early diagenetic 

carbonates, which formed in tidal flats and other coastal 

environments through the agency of evaporated brines 

(Alderman and Skinner, 1957; Degens and Epstein, 

1964; Deffeyes and others, 1965; Mueller and Tietz, 

1966; among others). It may be noted that in the 

Cretaceous marine formations of California, dolomite 

is rare and ferroan calcite is the dominant diagenetic 

carbonate.

Both primary marine carbonate and organic matter 

are important reservoirs of carbon in California Miocene formations. Diagenetic carbonate in these formations rarely has the carbon-isotopic composition of 

marine carbonate, but is lighter or heavier, depending 

on whether or not isotopic exchange equilibrium became 

established between carbon dioxide and coexisting 

methane. Light-carbon carbonates are rich in carbon 

derived through oxidation of sedimentary organic matter (SC13 of 31 to 14 per mil). Such carbonates are 

common in marine formations throughout the world 

(Spotts and Silverman, 1966; Hodgson, 1966; Galimov 

and Girin, 1968, Hoefs, 1970; Fontes and others, 1970; 

among others). None of the California specimens have 

the extremely light carbon (SC13 of 60 to 35 per

Cl
C2 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

mil) that is produced through oxidation of methane 

(Hathaway and Degens, 1969).

Heavy-carbon carbonates are rich in heavy carbon 

produced through isotopic exchange between carbon 

dioxide and methane:

C12O2 + C13H4^±C13O2 + C12H4 . (1)

With sufficient methane, this exchange strongly concentrates C13 in the carbon dioxide, the extent of concentration being inversely related to the temperature 

(Bottinga, 1969). The origin of diagenetic carbonates 

is thus an aspect of diagenesis of sedimentary organic 

matter; the main features of this diagenesis have been 

clearly outlined by Silverman (1965), Philippi (1965), 

andDavis (1967, p. 23).

The factors that promote or hinder isotopic exchange 

between carbon dioxide and methane, and hence determine the carbon-isotopic composition, are not known, 

but a certain consistency in isotopic composition is evident among dolomites from several California localities. 

Dolomites from Palos Verdes Hills contain light carbon 

only, whereas those from Reliz Canyon and from Chico 

MJartinez Creek have heavy carbon only. As will be discussed later, the prevalence of light carbon in dolomites 

of Palos Verdes Hills may be due to the unusually shallow burial (< 1,000 ft) of the Miocene section at this 

locality (Woodring and others, 1946). At several localities of heavy-carbon dolomites, the spread of isotopic 

composition is large enough to suggest that the methane-carbon dioxide exchange equilibrium was approached but not attained by many samples.

ACKNOWLEDGMENTS

We are indebted to J. E. Schoellhamer for samples 

80 and 81 from an exceptionally deep oil well in Kern 

County, Calif. Advice regarding the chemical and isotopic properties of formation waters was generously 

given by Ivan Barnes, J. D. Hem, and J. R. O'Neil; 

help and advice for the analysis of the dolomites, by 

Leonard Shapiro, Harry Bastron, C. O. Ingamells, Neil 

Elsheimer, Brent Fabbi, Chris Heropoulos, and Laureano Espos.

METHANE-CARBON DIOXIDE ISOTOPIC 

EXCHANGE IN NATURAL GAS

Because the isotopic composition of the methane that 

was associated with the diagenetic carbonates studied 

by us is unknown except within broad limits, a direct 

discussion of isotope geochemistry of the carbonates is 

not possible. But the required data are available for 

the closely related system of methane and carbon dioxide gas iii natural gas from various oil fields of the

United States, and these data are presented in figure 1. 

The left half of the figure shows the isotopic compositions of coexisting methane and carbon dioxide plotted 

against the depth of origin of the natural gas (Zartman and others, 1961; Wasserburg and others, 1963); 

the right half, apparent equilibrium temperatures, also 

plotted against depth of origin, computed under the assumption that the methane-carbon dioxide isotopic 

values represent equilibrium compositions (Bottinga, 

1969). Methane is the dominant constituent in all samples and is 12-340 times more abundant than carbon 

dioxide.

The filled points in figure 1 denote samples that contain the heaviest carbon dioxide and that yield the most 

reasonable temperature for their respective depths of 

origin (Wasserburg and others, 1963). The temperature-depth curve defined by the five filled circles corresponds to a gradient of 54°C per kilometer (1°C/60 

ft), the lowest gradient allowed by the data if the top 

of the curve is to be fixed to a surface temperature of 

about 20°C. Although rather high, the computed gradient is equalled or exceeded by gradients of a few 

American oil fields (Birch and others, 1942). Higher 

gradients involving the unfilled circles in figure 1 are 

therefore unlikely. These considerations of geothermal 

gradient suggest that (1) the five samples represented 

by filled points most completely satisfy the basic requirement for accurate isotopic temperatures, which is 

attainment of exchange equilibrium between carbon 

dioxide and methane, and (2) all other samples are out 

of equilibrium.

In keeping with the theoretical expectations, equilibrated samples show the greatest difference in SC13 

values of methane and carbon dioxide at the lowest temperature, the difference decreasing progressively with 

rising temperature. In these samples from marine formations, the SC13 of methane is heavier than 60 per 

mil, and the isotopic difference appropriate to methane 

and carbon dioxide in equilibrium is usually attained by 

the SC13 of carbon dioxide rising above the generally 

prevaling value of about 10 per mil to as high as +16 

per mil.

Isotopic exchange equilibrium between methane and 

dolomite cannot be expressed precisely at present, but 

the data for methane-calcite equilibrium (Bottinga, 

1969) suggest that SC13 in a carbonate would be 10-0 

per mil heavier than in carbon dioxide, within the temperature range of about 30° to 200° C. A finely dashed 

curve in the middle of figure 1 shows this higher content of C13 in calcite as compared with carbon dioxide 

when in equilibrium with the same methane. From all 

these considerations, we conclude that SC13 values in 

dolomite which are greater than about +5 mil signify
DIAGENETIC DOLOMITES, MIOCENE FORMATIONS, CALIFORNIA AND OREGON C3

Kilometers Feet

x 2

UJ 

O

4000

8000

12,000

 A ,

-60 -40 

 13

-20 + 20

O

O

100 200 300

dC" PER MIL (PDB) ISOTOPIC TEMPERATURE, IN °C

FIGURE 1. Carbon-isotopic compositions of coexisting methane (triangles) and carbon dioxide (squares) in natural gas of 

widely scattered American oil fields (Zartman and others, 1961; Wasserburg and others, 1963) and temperatures (circles) 

computed by assuming isotopic exchange equilibrium between methane and carbon dioxide (Bottinga, 1969). All quantities 

are plotted against the depth of origin of the natural gas. Filled points represent isotopic compositions and computed temperatures for samples containing the heaviest carbon dioxide at various depths. The thin dashed curve in the middle is 

SO13 of calcite in exchange equilibrium with methane whose SO13 lies along the curve defined by filled triangles (Bottinga, 

1969).

an appreciable influence of the methane-carbonate exchange and that values less than about 5 per mil indicate the lack of this influence.

We shall attempt to explain the varied isotopic and 

chemical composition of the dolomites in terms of chronological changes of physical-chemical conditions within 

the host sediments. Generally, preexisting carbonates 

are pictured as adjusting their composition to that appropriate to each new condition. As direct exchange of 

elements between solid carbonate and fluids is extremely 

slow (Anderson, 1969), the adjustments probably took 

place through solution-recrystallization of the preexisting carbonate. Because of the many variables involved, 

only broad trends of change can be identified and given 

tentative explanations.

FORMATION WATERS INVOLVED IN 

DOLOMITE GENESIS

The interstitial water of Holocene and Tertiary 

marine sedimentary deposits to a depth of 140 meters or

so below the sea floor is virtually the same as normal sea 

water, except for compositional changes induced by bacterial decomposition of organic matter (Siever and 

others, 1965). The changes of greatest interest in connection with genesis of diagenetic carbonates are the increase in bicarbonate and the generation of hydrogen 

sulfide through bacterial reduction of sulfate. The interstitial water of sediments in the San Pedro basin off 

the southern California coast (Emery and Eittenberg, 

1952; Presley and Kaplan, 1968) contains about 20 times 

more bicarbonate than normal sea water, but only 1/200 

of the sulfate, and the SC13 of the bicarbonate is as low 

as 21 per mil. Methane is present (Emery and Hoggan, 1958), but does not enter into isotopic exchange 

with the bicarbonate, judging from the lightness of the 

bicarbonate. The prevailing temperature is approximately a few degrees centigrade.

According to Presley and Kaplan (1968), the interstitial solution of San Pedro basin further contains 40-110 

ppb (parts per billion) iron, 200-460 ppm (parts per
C4 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

million) calcium, and around 1,300 ppm magnesium. 

(Normal seawater of the region contains 0.8 ppb iron, 

400 ppm calcium, and 1,350 ppm magnesium, according to Brooks and others (1968).) Also, about 8 percent 

of primary sedimentary carbonate is present in the sediment. These proportions suggest that any light-carbon 

carbonate (initially calcite) formed through interaction 

of such a solution and primary carbonate would be poor 

(hundredths of a percent by weight) in iron.

Bacterial decomposition of organic matter declines 

with progressive burial of the sediment, although some 

bacteria, like those that reduce sulfate, seem to function 

at depths of several thousand feet and temperatures of 

approximately 50°C (Davis, 1967). The later degradation of organic matter at temperatures higher than 

about 70°C proceeds largely through inorganic reactions (Silverman, 1964; Philippi, 1965), some of which 

may involve isotopic exchange between carbon dioxide 

and methane. Reducing conditions would prevail at all 

depths, and concentration of iron would be held to low 

levels through precipitation of pyrite (Hem, 1960a) or 

possibly through reactions involving organic matter 

(Hem, 1960b).

Phase relations bearing on dolomitization of calcite 

in deeply buried sediments by formation waters have 

been reviewed recently by Rosenberg, Burt, and Holland (1967) and by Usdowski (1968). The process is 

generally favored by a rise in temperature and salinity, 

which tends to reduce the degree of hydration of magnesium ions in solution and thereby promotes the entry 

of magnesium into an anhydrous carbonate structure. 

At temperatures around 60 °C, solutions with a magnesium/calcium atomic ratio of about 0.33 to 2.00 are at 

equilibrium with dolomite; those with lower and higher 

ratios are at equilibrium with calcite and magnesite, 

respectively.

Magnesium/calcium atomic ratios within the dolomite stability range are very common among oil field 

brines of California (Jensen, 1934; Gullikson and 

others, 1961; White and others, 1963; Weddle, 1968), 

and this fact may be the basic reason for dolomite being 

the predominant diagenetic carbonate in these fields. 

Miocene sedimentary rocks of the Round Mountain oil 

field on the east side of San Joaquin valley are exceptional in having diagenetic calcite rather than dolomite 

(Murata and others, 1969). Although the pore water 

in these formations has a magnesium/calcium atomic 

ratio of 0.54 (Jensen, 1934), its low salinity (<4,000 

ppm) and relatively low temperature (about 45°C; 

Moses, 1962) apparently were not favorable for 

dolomitization.

Certain organic acids have been found to promote 

the entry of magnesium into the calcite structure 

(Kitano and Kanamori, 1966) and could conceivably

have a role in the dolomitization process. Organic acids 

(especially naphthenic acid) are byproducts of petroleum-generating reactions and amount to 10-35 percent 

of the total anion equivalents in brines of some California oil fields (Gullikson and others, 1961). We have 

previously noted that dolomitization is favored by the 

reduction in the degree of hydration of magnesium ions 

induced by increases in salinity and temperature of 

solutions (Usdowski, 1968). According to M. N. A. 

Peterson (written comm., May 1971), the favorable 

effect of organic acids may be essentially the same, 

namely decreasing the degree of hydration of magnesium by means of magnesium-organic acid complex 

ions.

That natural dolomites have a variable content (48- 

57 mol percent) of CaCO3 in solid solution was shown 

by Goldsmith, Graf, Witters, and Northrop (1962), and 

the samples studied by us manifest a comparable variation, as did those described recently by Watanabe 

(1970). In solutions producing dolomite under equilibrium conditions, the factors that lead to calcium carbonate in excess of 50 mol percent seem to be (1) 

magnesium/calcium ratios in the lower part of the allowed range, (2) temperature in the lower part of the 

allowed range (Glover and Sippel, 1967), (3) salinity 

in the lower part of the allowed range (Fiichtbauer 

and Goldschmidt, 1965), and (4) concentrations of organic acids in the lower part of the allowed range 

(Kitano and Kanamori, 1966). As discussed later, under 

some circumstances all four factors could become 

lowered simultaneously to create conditions favorable 

for dolomite with excess calcium, as when meteoric 

water becomes mixed with oil field brines.

Bacterial reduction of sulfate to the sulfide and precipitation of pyrite are well known reactions that hold 

the iron content of formation waters to low levels 

(Davis, 1967). This control of iron concentration operates during most of the postdepositional history of 

sediments; consequently, most diagenetic calcites and 

dolomites are poor in iron.

When orogenic forces deform and raise sedimentary 

formations out of the sea, oxygenated meteoric water 

can circulate through the formations, decomposing 

pyrite and greatly increasing the concentration of dissolved iron to 7 ppm Fe2* or more (Langmuir, 1969). 

Only an edge of the formations need be exposed, as 

illustrated by the seaward-dipping Coastal Plain formations of Florida, which contain fresh water under 

artesian pressure, even 27 miles offshore (JOIDES, 

1965).

The increase in concentration of dissolved iron occurs only where the oxygenated water shifts Eh-pH 

conditions out of the stability field of pyrite into the 

field of ferrous carbonate, but stops short of the field
DIAGENETIC DOLOMITES, MIOCENE FORMATIONS, CALIFORNIA AND OREGON C5

of ferric hydroxide (James, 1966, fig. 5; Hem, 1967). 

Diagenetic calcites and dolomites formed in this highiron environment would be expected to contain substantial amounts of ferrous iron. Analyses of oil field 

brines show iron in the range of 1-100 ppm (Kittenhouse and others, 1969), but how much of this iron 

might be due to corrosion of steel pipes of the wells is 

uncertain.

For some reason, siderite is extremely rare among the 

samples studied by us. Scanty data suggest that it is 

formed in deep-lying brines charged with sufficient 

carbon dioxide to maintain a pH of less than 5 (Carlson, 

1949), under which condition both calcite and dolomite 

may be unstable. Siderite is a common diagenetic mineral in some fresh-water deposits, in which it seems to 

form not far below the water-sediment interface (Fritz 

and others, 1971).

CHEMICAL ANALYSIS OF DIAGENETIC 

DOLOMITES

Our previous report on the California diagenetic 

carbonates included a general description of their mode 

of occurrence and isotopic analyses of oxygen and carbon in 79 samples (Murata and others, 1969). We have 

subsequently made chemical analyses of 24 of these 

samples and two new samples in order to characterize 

the carbonates in greater detail and to establish a correlation between chemical and isotopic composition. 

Sample numbers used in the present paper are the same 

as those in the earlier report, which contains details on 

the nature and locality of the samples. The two new 

samples, 80 and 81, are cored dolomites from Ohio Oil 

Company's deep test well KCL-A-72-4, drilled in an 

extremely thick Miocene section of Kern County 

(Emery, 1956). Sample 80 is from a depth of 12,060 

feet, sample 81, from 15,300 feet.

Because diagenetic carbonates occur as fine-grained 

cement of sandstones and shales, separation of pure carbonate samples for chemical analysis is generally impractical. We deduced the approximate composition of 

diagenetic dolomite by analyzing dilute-acid (0.2^ 

HC1) extracts of whole-rock samples, shown by X-ray 

diffraction patterns to contain no other carbonate, gypsum, or laumontite. A source of error inherent in this 

approach is the possibility of the acid dissolving materials other than dolomite. However, the analytical 

results in table 1 show a general agreement in chemical 

equivalence between carbon dioxide and the sum of Ca, 

Mg, Fe, and Mn; thus errors are relatively small. A 

preliminary estimate of the iron content was obtained 

by staining sawed slabs of the specimens with a solution 

of potassium ferricyanide (Friedman, 1959).

450-13S 72 2

In computing the equivalences, an amount of CaO 

equal to 1.45 times the weight percentage of P2O5 was 

first deducted from soluble CaO to allow for the small 

amount of apatite in the samples. Total dissolved iron 

is reported as FeO. The amount of acid-insoluble matter 

in the samples ranges between 3 and 49 percent.

The strontium content of some of the samples was 

determined through X-ray fluorescence analysis of calcium oxide separated during gravimetric chemical 

analysis of the acid extracts. Semiquantitative spectrographic analyses made by Chris Heropoulos of sample 

10 and its insoluble residue showed that the bulk of the 

strontium is soluble in acid, whereas virtually all of the 

barium (700 ppm) is insoluble.

Dolomite compositions will be discussed in terms of 

mol percent of iron, manganese, magnesium, and calcium carbonates, as computed in table 1. The ranges of 

these components among the 26 samples are as follows:

FeOOi _________________ 0. 0-16.6 

MnCO3 ________________ .0- .5 

MgOO, _________________ 27. 9-51.1 

CaCOs _________________ 47. 8-58. 0

When compositions are plotted on the triangular diagram of figure 2, two compositional groups become 

evident:

1. Solid solutions between CaMg(CO3 ) 2 and CaFe 

(CO3 ) 2 but with 1-5 mol percent excess CaCO3 .

2. Iron-poor dolomites that depart substantially in 

content of CaCO3 from the end-member composition 

of CaMg(CO3 ) 2 , ranging between 48 and 58 mol percent CaCO3 .

CoCOj MgCOj

FIGURE 2. Composition of diagemetic dolomites in mol percent 

CaCOs, MgCO3, and FeCO»; the email amounts of MnCOs have 

been lumped with FeCO3 .
C6 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

TABLE 1. Partial chemical analyses of diagenetic dolomites from Miocene marine formations of California and Oregon

[Analysts: Marcelyn Cremer for samples with CO2 determined; K. J. Murata for samples with COz undetermined; B. P. Fabbi and Laureano Espos for strontium determinations. Localities of samples 1-78 described in detail in Murata, Friedman, and Madsen (1969). Samples 80 and 81 were obtained from Ohio Oil well KCL-A-72-4 in 

Paloma Oil Field, Kern County, Calif., at subsurface depths of 12,060 and 15,300 feet, respectively (Emery, 1956)]

10 24 27 31 33 37

Whole sample (weight percent)

FeO_ ..........................

MnO .. .....

MgO. ..........................

CaO... ...... ..................

SrO....... ..........

C02 ...... .................

PjOs... ...... ..................

Insoluble, 105°C. ..............

Insoluble, 1000° C.-..- . . .

.--. --- 2.2

............ .15

. . . 13.3

............ 24.3

............ 34.78

.-...... .08

 --..--- 21.8

1.1

.18

26.8

fl4

11.4 .

1.5

.02

9 2

16 9

QQ

47.5 .

0.56

.02

1Q f.

26.8

42.52

.14

9.3

2.2

.04

28.0

.036

41.12

.17

10.0

8.8

.28

8.3

23.1

.025 .

09 OQ

.22

23.5

.04

16.1

25.9

.11

12.4 .

0.22

.02

18.9

29.6

.063

.13

6.0

8.0

.26

10.7

24.9

.026 .

.14

16.0 .

0.09

.05

15.0

27.8

38.32

.13

17.0

0.02

.00

18.7

29.3

43.95

.10

6.0

0.31

.02

17.5

23.3

.058

37.23

.15

17.5

0.21

.04

19.2

28.8

.054

43.63

.06

5.7

Acid-soluble dolomite fraction (mol percent)

FeCOj.. ........ ............

MnCOs--- --.. -.

MgCOj. ........................

CaCOs-.......... _..____ _.

d (10-4) angstroms.. ._ ._ _

SO" per mil (PDB). ............

 _-. 3.8 1.7

 .- . .3 .3

 .-. 41.5 46.1

FA. A. Kl Q

2.898 2.891

... - +5.8 +11.0

3 Q

53.7

2.898

0.8

.0

48.9

KO 1

2.886

+17.7

3.3

.1

43.6

53.0

2.896

+11.3

.5

27.9

55.0

2.911

__ O Q

3 9

44.6

51.4

2.892

+9 6

0.3

.0

47.0

52.7

2.891

13.6

.4

32.4

53.6

2.902

+5.7

0. 1 0. 0

.1 .0

43. 0 47. 1

56. 8 52. 9

2. 904 2. 895

+7.1 +11.3

0.5

.0

51.1

48.4

2.882

+20.8

0.3

.1

48.0

51.6

2.891

+14.7

46 47 53 55 62 66 69 74 77 78 81

Whole sample (weight percent)

FeO..................

MnO.. ......

MgO--.. ......... ... .

CaO. ...... ...........

SrO-... ...............

C02..... .......................

PzOs.. .........................

Insoluble, 105° C . . . _

Insoluble, 1000° C .......... .

 ___ .- 1.0

. . - .05

. -.. 19.8

26 2

. -.----- .054 .

............ 42.56 .

. . .. .13

..... .-..- 8.1 .

0.15

04

10.4

15.7

49.3

0.11

20.4

30.0

AOK

.06

3.0

O Qft

AO

15.2

97 4.

.24

19.1

0.07

AQ

17.6

O1 O

m ^

.18

7.1

0.28

AA

17.0

OQ Q

.015 .

.11

10.3 .

0.09

179

28.6

39.60

1 9

11.0

0.07

00

17.6

OA Q

42.60

6.5 .

0.18

.01

10.8

16.7

24.0

.06

46.1 .

0.05

.04

16.1

31.0

41.85 .

19

8.6 .

0.67

.01

14.3

21.6

.023

.20

29.3

5.7

.24

11.7

25.8

.021 ..

.10

18.4

1.2

.06

16.4

26.3

.34

16.0

Acid-soluble dolomite fraction (mol percent)

FeC03 - 

MnCOs. 

MgCOs. 

CaCOs.

1.4

.1

50.7

0.4

.1

47.9

0.1

.0

48.7 

51.2

0.5 

.0

43.5 

56.0

0.1 

.0

43.9 

56.0

0.4 

.0

44.6 

55.0

0.1 

.0

47.0 

52.9

0.1 

.0

44.7

55.2

0.4 

.0

47.3 

52.3

0.1 

.1

41.8

58.0

1.2 

.0

47.7

51.1

9.4 

.4

35.0 

55.2

1.9 

.1

46.0

52.0

d(10-4) angstroms.... 

«C13 per mil (PDB).

2.882 2.889 

+19.3 -17.3

2.887 -2.7 2.903 

-12.5

2. 903 

-14.6

2. 903 

-15.4

2.894 

-11.7

2.901 2.889 

-14.9

2.889 

-25.3

2.904 

+9.6

2.892 

+9.8

1, 3 Yaqulna Bay, Lincoln County, Oreg., Nye Mudstone.

4 Point Arena, Mendocino County, Calif., Skooner Gulch Formation.

7, 9 Berkeley Hills, Contra Costa County, Claremont Shale.

10,11 Santa Cruz Mountains, Santa Cruz County, Mindego Basalt (Oligocene and

(or) Miocene).

24, 25 Reliz Canyon, Monterey County, Sandholdt Member of Monterey Shale. 

27, 31 Nacimiento Dam, San Luis Obispo County, Sandholdt Member of Monterey

Shale. 

33, 37 Chico Martinez Creek, Kern County, Chico Martinez Chert.

46 Temblor Range, San Luis Obispo County, Monterey Shale.

47 Pismo syncline area, San Luis Obispo County, Obispo Formation.

48 Pismo syncline area, San Luis Obispo County, Monterey Shale, immediately 

underlying Pismo Formation.

53 San Rafael Mountains, Santa Barbara County, Monterey Shale.

55 Santa Ynez River area, Santa Barbara County, dolomltized reef rock (40 ft thick).

62 Palos Verdes Hills, Los Angeles County, Malaga Mudstone Member of Monterey 

Shale.

66 Palos Verdes Hills, Los Angeles County, Valmonte Diatomite Member of Monterey Shale.

69 Palos Verdes Hills, Los Angeles County, upper part of Altamira Shale Member 

of Monterey Shale.

74 Palos Verdes Hills, Los Angeles County, middle part of Altamira Shale Member 

of Monterey Shale.

77 San Clemente Island, Monterey Shale.

78 La Jolla Submarine canyon, cored sediment EM3-5(8-17 cm).

THE d(10-4) X-RAY SPACING AND 

COMPOSITION

The substitution of iron, manganese, and calcium in 

the place of magnesium expands the unit cell of dolomite ; calcium by virtue of its large ionic radius has the 

greatest expansive power. In terms of the principal 

^(10-4) spacing, which, in end member dolomite has a 

value of 2.8855 A (angstroms), each mol percent of 

CaCO3 in excess of 50 will enlarge the spacing

0.00301 A (Goldsmith and Graf, 1958); thus

d(10-4) dolomite=2.8855 +0.00301 (CaCO3 -50.0) (2)

The ^(10-4) spacing of dolomite can be precisely determined (standard deviation=± 0.0013) by means of 

an X-ray diffractometer and by the use of the principal 

reflection of quartz (3.3433 A) as internal standard. 

A single parameter like the d ( 10 4) spacing cannot give 

a unique solution to the composition of a multicomponent phase. But it can serve as valuable check on the
DIAGENETIC DOLOMITES, MIOCENE FORMATIONS, CALIFORNIA AND OREGON C7

overall accuracy of the chemical analysis, especially 

with respect to percentage of calcium. Advantage is 

taken of equation 2 by converting the determined percentages of FeCO3 and MnCO3 into equivalent CaCO3 

(in terms of expansive power), in the following way.

The d(10-4) spacing of CaMg(CO3 ) 2 , OaFe(CO3 ) 2 

CaMn(CO3 ) 2 , and CaCa(CO3 ) 2 (calcite), as computed 

by Graf (1961), is shown diagrammatically in figure 3.

0.1504

CaMq(C03) 2 

1,1

2.88

CaMn(COj)z 

,1,

CoCoKOjJj

2.90 2.92 2.94 2.96 2.98 3.00 3.02

d(IO-4) SPACING, IN ANGSTROMS

3.04

FIGURE 3. The principal #(10-4) X-ray spacing of dolomite and 

related carbonates. Data from Graf (1961).

The separation in angstroms of the several peaks from 

that of CaMg(CO3 ) 2 is a measure of the expansive 

power of an ion of Fe, Mn, or Ca when substituted for 

an ion of Mg. The expansive power of Fe and Mn relative to Ca is thus Ca: Fe: Mn: : 1.00: 0.19: 0.37, and the 

determined molar percentages of FeCO3 and MnCO3 

are converted to equivalent of CaCO3 by multiplying 

by 0.19 and 0.37, respectively. The sum of the actual 

percentage of CaCO3 plus the converted percentages of 

of FeCO3 and MnCO3 may be called apparent CaCO3,

and

apparent CaCO3=actual CaCO3

+ 0.19FeCO3 + 0.37MnCO8 (3)

Because the amounts of MnCO3 in the studied dolomites are negligible, only FeCO3 contributes significantly to apparent OaCO3 , and about 5 mol percent 

FeCO3 is needed to displace ^(10-4) to the extent produced by 1 mol percent CaCO3 .

Equation 2, relating the CaCO3 content of dolomite 

to its 0?(10-4) spacing, is indicated by the straight line 

in figure 4. The points in the figure represent the content 

of apparent CaCO3 in the analyzed dolomites plotted 

against their ^(10-4) spacing. For samples having less 

than 2 mol percent FeCO3 (represented by single open 

circles), the <#(10-4) spacing may be considered to be 

determined solely by CaCO3 . The apparent CaCO3 of

samples containing more than 2 mol percent FeCO3 

(filled circles of connected samples) is computed according to equation 3.

Except for a single sample (No. 77), all samples conform to the relation of equation 3, within ± 1 mol percent apparent CaCO3 . Thus, a plot of this kind serves as 

a safeguard against gross errors in chemical analysis, 

especially with regard to calcium, and assures that the

 A Apparent CaC0 3 

O A Actual CaC03

56

. 54

52

50

48

2.880 2.890 2.900 2.9(0 2.915 

</(IO-4) SPACING, IN ANGSTROMS

FIGUBE 4. Relation between the #(10-4) X-ray spacing and 

the apparent CaCOs content of dolomite, where apparent 

OaCOs in mol percent=actual CaCO3 -i-0.19FeCO»+0.37MnCO». 

Single open circles represent actual Ca'COs percentage in samples with less than 2 percent FeOOs. Samples denoted by 

Interconnected open and filled circles contain more than 2 percent F^COa, the open circles representing actual CaCO3, the 

filled circles, apparent CaCOs. Sample A (represented by 

triangles) is a ferroan dolomite from England, described by 

Howie and Broadhurst (1958).

determined components were in solid solution in dolomite and not present as separate phases. Figure 4 also 

emphasizes the rarity of ideal dolomite, approximated 

by sample 7, and how common samples with more than 

50 mol percent CaCO3 are among diagenetic dolomites. 

Samples 33 and 46 exemplify the rare type that contains 

less than 50 mol percent CaC03

INTERRELATION OF CHEMICAL AND CARBONISOTOPIC COMPOSITIONS

'In our previous study of these dolomites (Murata and 

others, 1969), a curious inverse relation was noted 

between the contents of C13 and calcium among samples 

of the heavy-carbon group. The relations did not hold 

among members of the light-carbon group. No chemical 

analyses were made in that study, and the apparent 

calcium content was calculated from the magnitude of 

the d (00-6) spacing. In figure 5, the inverse relation
C8 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

CO

0

OL 

U4 

CL

Ij 

T>

-12

-16

-20

-24

-26

Low Fe

 77

62

2.880 2.890 2.900 2.910 2.920

c/00-4) SPACING, IN ANGSTROMS

47 48 50 52 54 56 

CoC03l IN MOL PERCENT

B

58

E

0 2 4 6 8 10 12 14 16 18 

FeC03 , IN MOL PERCENT

C

FIGUBE 5. Relation between carbon-isotopic composition of dolomite (ordinate) and A, the <Z(10-4) X-ray spacing; B, mol 

percent of CaCO3 ; and C, mol percent of FeCO3. C shows the basis for classifying these dolomites into a light-carbon group, 

an intermediate-carbon group, a heavy-carbon-low-iron (FeOO8 < 2.0 percent) group, and a heavy-carfoon-Mgli-iron (FeCOs 

> 2.0 percent) group. Two or more samples from the same locality are linked by tielines. Numbered samples1 are discussed in 

the text.

between C13 content and the d(10-4) spacing is well 

shown, again among heavy-carbon dolomites only. The 

chemical analyses now allow us to examine the relation 

in greater detail.

The mol percents of CaCO3 and FeCO3 in the dolomites are plotted against the carbon-isotopic composition in figures 5B and 5(7 respectively. Light- and 

heavy-carbon dolomites have similar amounts of calcium, but light-carbon dolomites are poor in iron, generally less than 2 mol percent FeCO3 . The d(10 '4) 

spacing of light-carbon dolomites is controlled solely by 

the content of calcium, according to equation 2.

Heavy-carbon dolomites (excepting samples 80-81) 

show a general inverse trend between the contents of 

calcium and heavy carbon C13, which is not evident 

among light-carbon dolomites. With regard to the iron 

content (fig. 567), heavy-carbon dolomites are of two 

types:

1. Those exemplified by samples 7,24,27-31, and 33-37 

that are poor in iron (<2.0 percent FeCO3 ), resembling light-carbon dolomites in this respect but 

manifesting the inverse relation between calcium 

andC13.

2. The remainder, which holds appreciable amounts of 

iron (>2.0 percent FeCO3 ) varying directly with 

the content of calcium and, hence, inversely with 

the content of C18.

Thus, in place of our previous classification into lightand heavy-carbon dolomites, the chemical data indicate 

a threefold division of (1) light-carbon dolomite, (2) 

heavy-carbon-low-iron dolomite, and (3) heavycarbon-high-iron dolomite. We shall attempt to deduce 

the interrelation among the three types in subsequent 

sections of this report. As seen in figure 567, a region of 

overlap exists between heavy- and light-carbon dolomites in the SC13 range of 4 to 4 per mil. Samples 10
DIAGENETIC DOLOMITES, MIOCENE FORMATIONS, CALIFORNIA AND OREGON

and 77 lie within this range; by using the content of iron 

carbonate as a criterion, sample 10 (16.6 mol percent) 

would be assigned to the heavy-carbon group, sample 77 

(0.1 mol percent), to the light-carbon group.

LIGHT-CARBON DOLOMITES

In sediments rich in organic matter, the absence of 

any isotopic exchange between carbon dioxide and organic compounds is the basic condition for genesis of 

light-carbon dolomites. Without the control exerted by 

exchange equilibria, which inherently are dependent on 

temperature and composition, such dolomites show a 

generally random variation of isotopic composition with 

respect to physical-chemical parameters. Their light 

carbon originates from oxidation of sedimentary organic matter (Spotts and Silverman, 1966), and the isotopic composition varies between the limits set by organic matter (SC13 ^ 30 per mil) and by sedimentary 

marine carbonate (SC13 =0±4 per mil).

The calcium content of light-carbon dolomites is independent of the content of C13 , but correlates weakly 

with the content of O18, as shown in figure 6. We have 

not been able to attribute this correlation to any of the 

four calcium-controlling factors given under "Formation Waters Involved in Dolomite Genesis" (magnesium/calcium ratio, temperature, salinity, concentration of organic acids), for there are many more variables than interrelating equations.

Only a crude picture can be drawn of the conditions 

under which light-carbon dolomites form. In view of the 

demonstrated generation of light-carbon carbonate in 

Holocene marine sediments (Presley and Kaplan, 1968) 

and the samples from the shallow section of Palos 

Verdes Hills as examples, light-carbon dolomites (or 

their precursors) are thought to form at a depth of 

1,000 feet or less in moderately compacted sediments 

rich in organic matter.

HEAVY-CARBON-LOW-IRON DOLOMITES

Heavy-carbon-low-iron dolomites are abundant at 

such classic California Miocene localities as Chico Martinez Creek, Nacimiento River, and Reliz Canyon. Compared with light-carbon dolomite, these dolomites probably formed at somewhat greater depth, where there 

was greater probability of carbon dioxide and methane 

undergoing isotopic exchange.

As previously discussed, exchange equilibrium imposes constraints on the isotopic composition of the involved compounds and allows the apparent temperature 

of equilibration to be calculated from the isotopic composition (Bottinga, 1969). Such calculations yield only 

rough estimates of temperature, but these are good

enough to set highly useful limits on the physicalchemical conditions of diagenesis. Two assumptions necessary for the calculations are that (1) the isotopic 

composition of the methane involved in the equilibrium 

lies on the curve defined by the filled triangles in figure 1 

or has some other parallel values permitted by existing 

data and (2) the equilibrium composition for dolomite 

is the same as that for calcite at all temperatures and 

will lie along the finely dashed curve in the middle of 

figure 1 when the methane composition is defined by the 

curve of filled triangles.

Samples 27, 31, and 33 will be used to illustrate the 

mode of calculation, and the pertinent data are tabulated below. The depth of origin of the dolomite is based 

on the calculated temperature and on the average geothennal gradient of the Los Angeles basin (Philippi, 

1965).

Sample No_______________

3C13 of dolomite, per mil_____

5C13 of methane from fig. 1. 

per mil__________________

Calculated temperature, °C__ 

Calculated depth, feet_______

Chico Martinez 

Nacimiento River Creek

27 31 33 

+ 7.1 +11.3 +20.8

-39

125

3,800

-40

103

7,000

-44

58

3,200

The results for sample 27, namely 125 °C and a depth of 

8,800 feet, seem to be somewhat excessive, whereas those 

for sample 31,103°C and 7,000 feet, are more reasonable 

in terms of the stratigraphy of the Nacimiento Eiver 

region southwest of the Jolon fault (Durham, 1968).

This conclusion is in agreement with the general indication of figure 1 that in any given formation or 

locality, the carbon dioxide (or carbonate) with the 

heavist carbon most closely approached exchange equilibrium. The computed temperature and depth (58°C, 

3,200 ft) for sample 33, with the heaviest carbon from 

Chico Martinez Creek, are also acceptable in terms of 

probable depth of burial of the formation sampled 

(Elliott and others, 1968).

The low-iron content of these dolomites suggests that 

they were derived directly from light-carbon-low-iron 

dolomites through the operation of isotopic exchange 

between carbon dioxide and methane. If the exchange 

reaction and recrystallization occurred with the involvement of a minimum of solution (smallest reservoir 

of dissolved calcium and magnesium), the resultant 

heavy-carbon dolomites could essentially inherit the 

widely ranging calcium content of the precursor lightcarbon dolomite. The decrease in CaCO3 that accompanies the rise in SC13 , as in samples 27-31 and 33-37 in 

figure 55, is probably a gradual adjustment toward 

thermodynamic stability through elimination of excess 

calcium, a process paralleling the gradual attainment 

of isotopic exchange equilibrium.
CIO

HEAVY-CARBON-HIGH-IRON DOLOMITES

Examples of heavy-carbon-high-iron dolomites are 

samples 1, 10, 25, and 80, with 4-17 mol percent FeCO3 . 

In figure 5(7, each of these samples is seen to be linked 

to another from the same locality (3,11, 24, or 81) that 

is poorer in iron and richer in C13. All four of* the latter 

samples have SC13 of about +10 per mil, corresponding 

to a formation temperature of about 110°C, which may 

be near the maximum for diagenetic dolomites. The 

higher temperature corresponding to the lighter carbon 

of iron-rich members of each pair is probably fictitious, 

and the lightness of the carbon is more a measure of 

departure from exchange equilibrium. For example, if 

the SC13 of 2.8 per mil of sample 10 from the Santa 

Cruz Mountains is assumed to be an equilibrium value, 

an excessively high temperature of 180°C is obtained, 

which contrasts with the 110°C computed for sample 

11 from the same locality.

The high concentration of iron in dolomites of this 

category requires a fundamentally different chemistry 

from that of the other categories. And the interpretation of seemingly identical relations, such as that between the contents of CaCO3 and C13, is radically 

different.

A high-iron dolomite like sample 10 could be derived 

from a lower iron dolomite like sample 11 by a dissolutionTreprecipitation process that would increase the 

content of iron and calcium of dolomite as it lowered 

the content of C13. The most direct way to effect this 

would be to introduce oxygenated meteoric water into 

the dolomite-bearing formation. This would most likely 

happen after orogenic movements raise a part of the 

sedimentary formation out of the sea, thereby exposing 

it to fresh water.

The great increase in soluble iron that occurs when 

oxygenated fresh water enters a reduced marine sediment has been well described by Langmuir (1969). 

Although the solubilization of iron was observed at 

depths of only a few hundred feet, in principle the 

reaction should be possible to far greater depths. 

Further effects of the intermixing of meteoric and deeplying waters would generally include a reduction of 

salinity and a shift of the magnesium/calcium ratio to 

lower values; both of these effects would favor an 

increase of calcium in dolomite that was simultaneously 

being enriched in iron. Meteoric water would also raise 

the concentration of light-carbon bicarbonate, derived 

through oxidation of organic matter either in surface 

soils or in the dolomite-bearing formation itself. The 

net result of all of these reactions would be a trend of 

alteration of a given dolomite toward a composition 

richer in iron and calcium and poorer in C13 and O18.

Strong support for a mode of forming iron-rich dolomites through the influence of fresh water is provided 

by the oxygen-isotopic composition (Murata and others, 

1969) of the dolomites (fig. 6). The SO18 of deep-lying

33 f

V-____4 

74 62

ZZ 24 26 28 30 32 34 36

<?0 18 PER MIL (SMOW)

0 2 4 6 8 10 12 14 16 18 

FeC03 , IN MOL PERCENT

FIGURE 6. Relation of oxygen-isotopic composition of dolomite 

and the content of FeCO3 to the carbon-isotopic composition 

(ordinate). Note in the upper half of the figure the symmetrically divergent patterns among certain pairs of heavycarbon-high-iron dolomites (samples 1-3, 10-11, 24-25, and 

80-81).

formation waters can be as heavy as + 9 per mil relative 

to standard mean ocean water (SMOW) (Clayton and 

others, 1966), whereas that of fresh water is always 

lighter than sea water, with SO18 of 15 5 per mil 

SMOW in the middle latitudes (Epstein, 1959). As the 

oxygen-isotopic composition of a formation water becomes lighter through increasing admixture of fresh 

water, the oxygen of dolomite will tend to become 

lighter through isotopic exchange with the water.

The correlation of a higher content of iron with 

lighter oxygen in dolomite is well shown in the upper 

half of figure 6, in the symmetrically divergent pattern 

of several pairs of samples (1-3, 10-11, 24-25, and 80- 

81). The SO18 of these samples decreases at the rate of 

 0.3- 0.7 per mil for every mol increase of FeCO3 , a 

rate too great to be ascribed to any specific effect of iron 

on the equilibrium constant of oxygen exchange between 

dolomite and formation water (O'Neil and others, 

1969). The pair 80-81 differs from the other three by not
DIAGENETIC DOLOMITES, MIOCENE FORMATIONS, CALIFORNIA AND OREGON Cll

showing a decrease in SC13 commensurate with the decrease in SO18. Such a unilateral change could result if 

the introduced fresh water were low in bicarbonate, 

possibly through prior precipitation of a carbonate.

The generally interrelated variations of O18, C18, 

CaCO3 , and FeCO3 in dolomites of the present category 

are consequences of meteoric water gaining access to 

once deeply buried marine formations. If this interpretation is essentially correct, iron in diagenetic dolomites 

of California has a very interesting connotation of time, 

pointing to a postorogeny period of formation for the 

iron-rich samples.

REFERENCES CITED

Alderman, A. R., and Skinner, H. C. W., 1957, Dolomite sedimentation in the south-east of South Austrialia : Am. Jour. Sci., 

v. 255, p. 561-567.

Anclerson, T. F., 1969, Self-diffusion of carbon and oxygen in calcite by isotope exchange with carbon dioxide: Jour. 

Geophys. Research, v. 74, no. 15, p. 3918-3932.

Birch, Francis, Schairer, J. F., and Spicer, H. C., eds., 1942, 

Handbook of physical constants: Geol. Soc. America Spec. 

Paper 36, 325 p.

Bottinga, Yan, 1969, Calculated fractionation factors for carbon 

and hydrogen isotope exchange in the system calcite-carbon 

dioxide-graphite-methane-hydrogen-water vapor: Geochim. 

et Cosmochim. Acta, v. 33, no. 1, p. 49-64.

Brooks, R. R., Presley, B. J., and Kaplan, I. R., 1968, Trace elements in the interstitial waters of marine sediments: Geochim. et Cosmochim. Acta, v. 32, no. 4, p. 397-414.

Carlson, H. A., 1949, Corrosion in natural gas-condensate wells: 

Indus, and Bng. Chemistry, v. 41, p. 644-645.

Clayton, R. N., Friedman, I., Graf, D. L., Mayeda, T. K., Meents, 

W. F., Shimp, N. F., 1966, The origin of saline formation 

waters [Pt.] 1, Isotopic composition: Jour. Geophys. Research, v. 71, no. 16, p. 3869-3882.

Davis, J. B., 1967, Petroleum microbiology: Amsterdam, Elsevier 

Publishing Co., 604 p.

Deffeyes, K. S., Lucia, F. J., Weyl, P. K., 1965, Dolomitization 

of Recent and Plio-Pleistocene sediments by marine evaporite waters on Bonaire, Netherlands Antilles, in Dolomitization and limestone diagenesis A symposium: Soc. Econ. 

Paleontologists and Mineralogists Spec. Pub. 13, p. 71-88.

Degens, E. T., and Epstein, Samuel, 1964, Oxygen and carbon 

isotope ratios in coexisting calcites and dolomites from recent and ancient sediments: Geochim. et Cosmochim. Acta, 

v. 28, no. 1, p. 23-44.

Durham, D. L., 1968, Geology of the Tierra Redonda Mountain 

and Bradley quadrangles, Monterey and San Luis Obispo 

Counties, California: U.S. Geol. Survey Bull. 1255, 60 p.

Elliott, W. J., Tripp, Eugene, and Karp, S. E., 1968, West side 

oil fields and Temblor Range outcrop area, in Karp, S. E., 

ed., Guidebook, geology and oil fields west side Southern 

San Joaquin Valley: Am. Assoc. Petroleum Geologists, Soc. 

Econ. Geologists, and Soc. Econ. Paleontologists and Minerologists, 43d Ann. Mtg., Pacific Sees., p. 104-123.

Emery, K. O., and Hoggan, Dean, 1958, Gases in marine sediments : Am. Assoc. Petroleum Geologists Bull., v. 42, no. 9, 

p. 2174-2188. Emery, K. O., and Rittenberg, S. C., 1952, 

Early diagenesis of California basin sediments in relation

to origin of oil: Am. Assoc. Petroleum Geologists Bull. v. 36, 

no. 5, p. 735-806.

Emery, W. B., II, 1956, The Ohio Oil Company's deep test well, 

Palonia field, Kern County, California: Am. Petroleum Inst. 

Drilling and Production Practice, 1955, p. 336-358.

Epstein, Samuel, 1959, The variations of the O^/O18 ratio in 

nature and some geologic implications, in Abelson, P. H., 

ed., Researches in geochemistry: New York, John Wiley & 

Sons, Inc., p. 217-240.

Fontes, J. C., Fritz, Peter, and Letolle, Rene, 1970, Composition 

isotopique, mineralogique et genese des dolomies du Bassin 

de Paris: Geochim. et Cosmochim. Acta, v. 34, no. 3, p. 

279-294.

Friedman, G. M., 1959, Identification of carbonate minerals by 

staining methods: Jour. Sed. Petrology, v. 29, no. 1, p. 87-97.

Fritz, P., Binda, P. L., Folinsbee, R. E., and Krouse, H. R., 1971, 

Isotopic composition of diagenetic siderites from Cretaceous 

sediments of Western Canada: Jour. Sed. Petrology, v. 41, 

no. 1, p. 282-288.

Fiichtbauer, Hans, and Goldschmidt, Hertha, 1965, Beziehungen zwischen Calciumgehalt und Bildungsbedingungen der 

Dolomite: Geol. Rundschau, v. 55, p. 29-40.

Galimov, E. M., and Girin, Y. P., 1968, Variation in isotopic 

composition of carbon in the process of formation of carbonate concretions: Geochemistry Internat., v. 5, no. 1, p. 

178-182.

Glover, E. D., and Sippel, R. F., 1967, Synthesis of magnesium 

calcites: Geochim. et Cosmochim. Acta, v. 31, no. 4, p. 

603-613.

Goldsmith, J. R., and Graf, D. L., 1958, Relation between lattice 

constants and composition of the Ca-Mg carbonates: Am. 

Mineralogist, v. 43, nos. 1-2, p. 84-101.

Goldsmith, J. R., Graf, D. L., Witters, Juanita, and Northup, 

D. A., 1962, Studies in the system CaCOs-MgCOs-FeCOr 

[Pt.] 1, Phase relations; [Pt.] 2, A method of major-element 

spectrochemical analysis; [Pt] 3, Composition of some ferroan dolomites: Jour. Geology, v. 70, no. 6, p. 659-688.

Graf, D. L., 1961, Crystallographic tables for the rhombohedral 

carbonates: Am. Mineralogist, v. 46, nos. 11-12, p. 1283-1316.

Gullikson, D. M., Caraway, W. H., and Gates, G. L., 1961, Chemical analysis and electrical resistivity of selected California 

oil field waters: U.S. Bur. Mines Rept. Inv. 5736, 21 p.

Hathaway, J. C., and Degens, E. T., 1969, Methane-derived 

marine carbonates of Pleistocene age: Science, v. 165, p. 

690-692.

Hem, J. D., 1960a, Restraints on dissolved ferrous iron imposed 

by bicarbonate, redox potential, and pH: U.S. Geol. Survey 

Water-Supply Paper 1459-B, p. 33-55.

 1960b, Complexes of ferrous iron with tannic acid: U.S. 

Geol. Survey Water-Supply Paper 1459-D, p. 75-94.

1967, Equilibrium chemistry of iron in ground waiter, in

Faust, S. D., and Hunter, J. V., eds., Principles and applications of water chemistry Rudlofs Research Conf., 4th, 

Rutgers Univ., Proc.: New York, John Wiley & Sons, p. 

625-643.

Hodgson, W. A., 1966, Carbon and oxygen isotopic ratios in 

diagenetic carbonates from marine sediments: Geochim. et 

Cosmochim. Acta, v. 30, no. 12, p. 1223-1233.

Hoefs, Jochen, 1970, Kohlenstoff und Sauerstoff Isotopenuntersuchungen am Karbonat konkretionen und umgebendem 

Gestein: Contr. Mineralogy and Petrology, v. 27, no. 1, p. 

66-79.
C12 SHORTER CONTRIBUTIONS TO GENERAL GEOLOGY

Howie, R. A., and Broadhurst, F. M., 1958, X-ray data for

dolomite and ankerite: Am. Mineralogist, v. 43, nos. 11-12,

p. 1210-1214. 

James, H. L., 1966, Chemistry of the iron-rich sedimentary

rocks, in Data of geochemistry [6th ed.]: U.S. Geol. Survey

Prof. Paper 440-W, p. W1-W60. 

Jensen, Joseph, 1934, California oilfield waters, in Problems of

petroleum geology: Am. Assoc. Petroleum Geologists (Sidney Powers Memorial Volume), p. 953-985. 

JOIDES (Joint Oceanographic Institute Deep Earth Sampling

Program), 1965, Ocean drilling on the continental margin:

Science, v. 150, no. 3697, p. 709-716. 

Kitano, Yasushi, and Kanamori, Nobuko, 1966, Synthesis of

magnesian calcite at low temperatures and pressures: Geochem. Jour. (Nagoya), v. 1, no. 1, p. 1-10. 

Langmuir, Donald, 1969, Iron in ground waters of the Magothy

and Raritan Formations in Camden and Burlington Counties, New Jersey: New Jersey Dept. Conserv. and Econ.

Development, Water Resources Circ. 19,43 p. 

Moses, P. L., 1962, Geothermal gradients: Am. Petroleum Inst.

Drilling and Production Practice, 1961, p. 57-63. 

Mueller, German, and Tietz, Gerd, 1966, Recent dolomitization

of Quaternary biocalcarenites from Fuerteventura (Canary

Islands) : Contr. Mineralogy and Petrology, v. 13, p. 89-96. 

Murata, K. J., Friedman, Irving, and Madsen, B. M., 1969,

Isotopic composition of diagenetic carbonates in marine

Miocene formations of California and Oregon: U.S. Geol.

Survey Prof. Paper 614-B, 24 p. 

O'Neil, J. R., Clayton, R. N., and Mayeda, T. K., 1969, Oxygen

isotope fractionation in divalent metal carbonates: Jour.

Chem. Physics, v. 51, no. 12, p. 5547-5558. 

Philippi, G. T., 1965, On the depth, time, and mechanism of

petroleum generation: Geochim. et Cosochim. Acta, v. 29,

no. 9, p. 1021-1049. 

Presley, B. J., and Kaplan, I. R., 1968, Changes in dissolved

sulfate, calcium and carbonate from interstitial waters of

near-shore sediments: Geochim. et Cosmochim. Acta, v. 32,

no. 10, p. 1037-1048. 

Rittenhouse, Gordon, Fulton, R. B., Ill, Grabowski, R. J., and

Bernard, J. L., 1969, Minor elements in oil-field waiters:

Chem. Geology, v. 4, p. 189-209. 

Rosenberg, P. E., Burt, D. M., HoUand, H. D., 1967, Calcitedolomite-magnesite stability relations in solutions The effect of ionic strength: Geochim. et Cosmochim. Acta, v. 31, 

no. 3, p. 391-396.

Siever, Raymond, Beck, K. C., and Berner, R. A., 1965, Composition of interstitial waters of modern sediments: Jour. 

Geology, v. 73, no. 1, p. 39-73.

Silverman, S. R., 1964, Investigations of petroleum origin and 

evolution mechanisms by carbon isotope studies, [Chap.] 8 

in Isotopic and cosmic chemistry (dedicated to Harold C. 

Urey) : Amsterdam, Netherlands, North-Holland Publishing Co., p. 92-102.

 1965, Migration and segregation of oil and gas, in Fluids 

in subsurface environments A symposium: Am. Assoc. 

Petroleum Geologists Mem. 4, p. 53-65.

Spotts, J. H., and Silverman, S. R., 1966, Organic dolomite from 

Point Fermin, California: Am. Mineralogist, v. 51, no. 7, 

p. 1144-1155.

Usdowski, Hans-Eberhard, 1968, The formation of dolomite in 

sediments, in Mueller, German, and Friedman, G. M., eds., 

Recent developments in carbonate sedimentology in Central 

Europe: New York, Springer-Verlag, p. 21-32.

Wasserburg, G. J., Mazor, E., and Zartman, R. E., 1963, Isotopic 

and chemical composition of some terrestrial natural gases, 

in- Earth science and meteoritics: Amsterdam, NorthHolland Publishing Co., p. 219-240.

Watanabe, Masaharu, 1970, Carbonate concretions in the Neogene Tertiary system, Northeast Japan: Tohoku Univ. Sci. 

Repts., 3d ser., v. 11, no. 1, p. 69-110.

Weddle, J. R., 1968, Oilfield waters in southwestern San Joaquin 

Valley, Kern County, California, in Karp, S. E., ed., Guidebook, geology and oil fields west side Southern San Joaquin 

Valley: Am. Assoc. Petroleum Geologists, Soc. Econ. Geologists, and Soc. Econ. Paleontologists and Mineralogists, 43d 

Ann. Mtg., Pacific Sees., p. 104-123.

White, D. E., Hem, J. D., and Waring, G. A., 1963, Chemical 

composition of subsurface waters, in Data of geochemistry, 

[6th ed.]: U.S. Geol. Survey Prof. Paper 440-F, 67 p.

Woodring, W. P., Bramlette, M. N., and Kew, W. S. W., 1946, 

Geology and paleontology of Palos Verdes Hills, California: 

U.S. Geol. Surrey Prof. Paper 207, 145 p.

Zartman, R. E., Wasserburg, G. J., and Reynolds, J. H., 1961, 

Helium, argon, and carbon in some natural gases: Jour. 

Geophys. Research, v. 66, no. 1, p. 277-306.
Shorter Contributions 

to General Geology, 

1971

GEOLOGICAL SURVEY PROFESSIONALPAPER 724

This volume was published 

as separate chapters A C
UNITED STATES DEPARTMENT OF THE INTERIOR 

ROGERS C. B. MORTON, Secretary

GEOLOGICAL SURVEY 

V. E. McKelvey, Director
CONTENTS

[Letters designate the separately published chapters]

(A) Huntington Lake quadrangle, central Sierra Nevada, California analytic data, by 

P. C. Bateman and D. R. Wones.

(B) Normal fatty acids in estuarine and tidal-marsh sediments of Choctawhatchee and 

Apalachee Bays, northwest Florida, by R. E. Miller.

(C) Geochemistry of diagenetic dolomites in Miocene marine formations of California 

and Oregon, by K. J. Murata, Irving Friedman, and Marcelyn Cremer.