Carbonate Rocks: Origin, Occurrence and Classification (Developments in Sedimentology 9A)

DEVELOPMENT S IN SEDIMENTOLO GY 9A DEVELOPMENTS IN SEDIMENTOLOGY 9A CARBONATE ROCKS Origin, Occurrence and Classi$ca...

Author: George V. Chilingar | Harold J. Bissell | Rhodes W. Fairbridge

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DEVELOPMENT S IN SEDIMENTOLO GY

9A

DEVELOPMENTS IN SEDIMENTOLOGY 9A

CARBONATE ROCKS Origin, Occurrence and Classi$cation

EDITED BY

GEORGE V. CHILINGAR Professor of Petroleum Engineering University of Southern California, Los Angeles, Gal$ (U.S.A.)

HAROLD J. BISSELL Professor of Geology Brigham Young University, Provo, Utah (U.S.A.) AND

RHODES W. FAIRBRIDGE Professor of Geology Columbia University, New York, N. Y. (U.S.A.)

ELSEVIER PUBLISHING COMPANY Amsterdam London New York 1967

ELSEVIER PUBLISHING COMPANY

335 JAN

VAN GALENSTRAAT, P.O. BOX

21 1, AMSTERDAM

AMERICAN ELSEVIER PUBLISHING COMPANY, INC.

52 VANDERBILT

AVENUE, NEW YORK, N.Y.

10017

ELSEVIER PUBLISHING COMPANY LIMITED RIPPLESIDE COMMERCIAL ESTATE, BARKING, ESSEX

LIBRARY OF CONGRESS CARD NUMBER

WITH

65-20140

131 ILLUSTRATIONS AND 20 TABLES

ALL RIGHTS RESERVED THIS BOOK OR ANY PART THEREOF MAY NOT BE REPRODUCED IN ANY FORM, INCLUDING PHOTOSTATIC OR MICROFILM FORM, WITHOUT THE WRITTEN PERMISSION FROM THE PUBLISHERS PRINTED I N THE NETHERLANDS

CONTENTS

CHAPTER l. INTRODUCTION R. W. FAIRBRIDGE (New York, N .Y., U.S.A.), G. V. CH!UNGAR (Los Angeles, Calif. U.S.A.) and H. J. BISSELL (Provo, Utah, U.S.A.) . . . . . . . . . . . . . . . . . . CHAPTER 2. MODERN CARBONATE SEDIMENTS W. H. TAFT (Tampa, Fla., U.S.A.) . . . . . . . . . .

29

CHAPTER 3. PETROLOGY AND PETROGRAPHY OF CARBONATE ROCKS Y. GUBLER, J.P. BERTRAND (Rueil-Malmaison, France), L. MATTAVELLI, A. RrzzrNI and R. PASSEGA (Milan, Italy) . . . . . . . . . . . . . . . . . . . . . . . . .

51

CHAPTER 4. CLASSIFICATION OF SEDIMENTARY CARBONATE ROCKS H. J. BISSELL (Provo, Utah, U.S.A.) and G. V. CHILINGAR (Los Angeles, Calif. U.S.A.)

87

CHAPTER 5. ORIGIN AND OCCURRENCE OF LIMESTONES J. E. SANDERS (Dobbs Ferry, N.Y., U.S.A.) and G. M. FRIEDMAN (Troy, N.Y., U.S.A.) .

169

CHAPTER 6. ORIGIN AND OCCURRENCE OF DOLOSTONES G. M. FRIEDMAN (Troy, N.Y., U.S.A.) and J. E. SANDERS (Dobbs Ferry, N.Y., U.S.A.) .

267

CHAPTER 7. CARBONATE OIL RESERVOIR ROCKS J. W. HARBAUGH (Stanford, Calif., U.S.A.) . . . . . . .

349

CHAPTER 8. CARBONATE ROCKS AND PALEOCLIMATOLOGY IN THE BIOCHEMICAL HISTORY OF THE PLANET R. W. FAIRBRIDGE (New York, N.Y., U.S.A.)

399

APPENDIX

433

REFERENCES INDEX

435

SUBJECT INDEX . . .

444

Chapter 1

INTRODUCTION RHODES W. FAIRBRIDGE, GEORGE V. CHILINGAR, AND HAROLD J . BISSELL

Columbia University, New York, N. Y. (U.S.A.) University of Southern California, Los Angeles, Calif. (U.S.A.) Brigham Young University, Provo, Utah (U.S.A.)

INTRODUCTION

Limestones are those rocks composed of more than 50% carbonate minerals, of which 50 % or more consist of calcite and/or aragonite. Limestones may be white, gray, dark gray, yellowish, greenish, blue, and, sometimes, black in color. A small admixture of clay particles or organic matter imparts gray color to limestones. Greenish coloration is due to the admixture of clayey material, glauconite, or finely-dispersed ferrous compounds; whereas reddish or brownish coloration is due to the presence of ferric iron. As pointed out by RUKHIN(1961, p.173), coarse-grained limestones usually have lighter colors (associated with well-oxidized conditions) than the fine-grained varieties (often indicative of deeper, possibly reducing conditions in early diagenesis). Impurities in limestone include a spectrum of magnesium carbonates, dolomite, silica, glauconite, gypsum, fluorite, siderite, sulfides, iron and manganese oxides, phosphates, clays, and organic matter. A rock with somewhat under 50% CaC03 is described as a calcareous sandstone, shale, etc. Its carbonate fraction is often subject to differential weathering, so that it appears like a limestone; the insoluble quartz grains, etc. are often unobtrusive and are easily released during weathering processes. Some calcareous eolianites with less than 30 % CaC03 develop karst features and other “characteristics” of limestones. Dolomites (or “dolostones”) are rocks composed mainly of the mineral dolomite (CaMg(C0s)z). Pure dolomite mineral is composed of 45.7 % MgC03 and 54.3 % CaC03, in weight % (or 47.8 % COz, 21.8 % MgO, and 30.4 % CaO) (see Fig. 1). Impurities in dolomite rock (dolostone) include gypsum, anhydrite, iron sulfides, celestite, opal, chalcedony, iron oxides, magnesite, fluorite, and organic matter. Dolomites are quite similar to limestones in appearance and, therefore, it is difficult to distinguish between the two with a naked eye. (The term “dolostone” is used by some authors to specify the rock composed of the mineral dolomite, in large measure at least; the present authors recognize certain utility in this term. The word is not widely used, however, because the meaning is normally clear from the context.)

2

R. W.

FAIRBRIDGE, G.

V. CHILINGAR AND H. J. BISSELL

20 -n

-s v

l5

Magnesium enriched limestones

Magnesium carbonate

( O h )

Fig.1. Graph showing MgC03 concentration in Paleozoic carbonates (data after CHAVE,1954), ar:d two distinct peaks of “magnesium enrichment”. Note that high-magnesium calcites having 15-35 % MgC03 are notably absent due to their preferential dolomitization, leaving only the limestones with MgC03 content of less than 15 %. There is a progressive dolomitization of all ancient limestones having above 15 % MgC03. (After FAIRBRIDGE, 1957, p.128, fig.1.)

Sedimentary carbonate rocks are, for the most part, products of deposition in marine realms and the principal repositories are the paralia- and autogeosynclines, in neritic environments. Calcium carbonate can accumulate also in deep water, or may be found in coral atoll structures surrounded by an abyssal region, but definitely not on the same scale as those of banks, platforms, and shelves. Little or no evidence is available suggesting that areally-extensive carbonates ac cumulated in deep water prior to the Cretaceous. Data compiled more than 40 years ago by VAUGHAN (1924, p.313) from the “Challenger” Report on modern deep-sea sediments, however, indicate that calcium carbonate comprises on the average 32.2 % of these sediments, largely as Globigerina and pteropod oozes, and as coral mud and sand. As pointed out by TWENHOFEL ( I 943, p.285), “This concentration in the deeper water sediments must be balanced somewhere in other deposits; the sediments with a deficiency of lime may be those of the continental and delta environments.” Probably the deep-sea sediments of the Paleozoic and Precambrian eras were low or deficient in CaC03, because the lime-secreting plankton and nekton were rare or absent. As pointed out by STRAKHOV et al. (1954, p.720), the evolution of carbonate rocks in many respects is different, for example, from that of iron ores (Fig.2). According to STRAKHOV et al. (1954, p.721), inasmuch as carbonate rocks occur preferentially on platforms rather than in orthogeosynclines and with time the shallow platform areas were growing at the expense of deeper geosynclinal zones, the processes favorable to carbonate formation were more intensive during the

3

INTRODUCTION

later periods. It is uncertain, however, whether the rate of carbonate withdrawal has changed greatly through time or not. There are proportionately fewer carbonates in exposed Precambrian sequences, but it should be remembered that the latter are largely preserved today in their deeper geosynclinal facies, whereas their shallow neritic equivalents have had the maximum opportunity to be destroyed by erosion. Over geologic time there have been stages of broad epicontinental seas, favorable to neritic carbonate development (e.g., Ordovician, Cretaceous), and other stages A

1

2

3

4

1

0

5 6

pg

Cr

J

Tr

P C D

S

?

?

?

?

cm 500

1,000

1,500

Fig.2. Schematic diagram showing evolution of carbonate rocks. Absolute time in millions of years is plotted on the ordinate. A . Normal-saline seas and oceans: I = chemical calcite formation; 2 = biogenic calcite formation; 3 = primary chemical dolomite formation; 4 = diagenetic dolomite formation. B. Carbonate lagoons: 5 = primary dolomite formation is a necessary stage during the history of lagoons; 6 = dolomite formation as an optional stage during the history of the lagoon, requiring existence of definite physicogeographical conditions. C. Lakes of arid zone: 7 = lakes of magnesium carbonate type; 8 = lakes of sodium type. (After STRAKHOV et al., 1954, p.720.)

4

R. W.

FAIRBRIDGE, G . V.

I

-7

SITE OF RESTRICTED 1ARRONATE SEDIMENTATION

SITE OF MAXIMUM CARBONATE SEDIMENTATION

I

CHILINGAR AND H. J. BISSELL

I

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(Coastal

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---.,/ NERITIC

i

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300

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miles

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i

400 -

2

500

...- ._ . . . _. - .. -

..- .- . ..

-L-I

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~.

sea - level

BATHYAL ZONE

Fig.3. Diagram illustrating maximum distribution of carbonates during the past at times when the sea level was universally high (thalassocratic stages), and poor development of carbonate sediments 1957, p. 157, fig. I I .) when sea level was low (epeirocratic stages). (After FAIRBRIDGE,

of general withdrawal that were unfavorable for the formation of carbonate rocks (e.g., parts of Upper Carboniferous, Triassic) (see Fig.3). Precambrian carbonate rocks, where preserved from obliteration by intense metamorphism, very frequently contain traces of calcareous Algae (such as Collenia, etc.). Even where the carbonates are totally replaced by silica, it is still possible to detect the unmistakable patterns of calcareous Algae; no tectonic process can duplicate their distinctive “cauliflower” growth forms (FAIRBRIDGE, 1950). Precambrian marbles commonly contain graphite in some sections which perhaps suggests an organic association. It is possible that a high Mg/Ca ratio and lower pH in Precambrian sea waters prevented the formation of hard protective and skeletal structures of organisms, or largely hindered their formation (CHILINGAR and BISSELL,1963; FAIRBRIDGE, 1964). Thus, during Precambrian time the carbonate rocks apparently owed their origin not to the secretion of shells by organisms, but to biochemical control of pH in lagoons by Algae and to direct chemical precipitation out of sea water. Algae became important lime-precipitating organisms quite early in Precambrian times (they are found, for example, in the oldest Precambrian rocks of Rhodesia); but only in Cambrian time did organisms of the animal kingdom begin to participate in significant manner in extracting carbonates out of sea water. Consequently, chemical or biochemical precipitation was over-shadowed by biogenic carbonate formation with time. Accumulations of carbonates in modern seas are produced by mechanical, organic, chemical, and inorganic processes; widespread accumulations of pure to relatively pure limestones, and in particular fine-textured varieties, represent moderately shallow to very shallow marine realms. Some of these deposits have

.I

5

INTRODUCTION

resulted from chemical processes only, whereas others owe their origin to physicochemical and biological processes. Coquinas of hard-parts of organisms are forming at present and there are substantial thicknesses and widespread occurrence of bioclastic carbonates. Dolomite, dolomitic limestone, and magnesium-rich carbonates are not accumulating as thick deposits in present-day seas; and the few occurrences that have been reported by ALDERMAN and SKINNER (1957, pp.561-567), SABINS(1962, pp.1183-1196), H. C. W. SKINNER et al. (1963, pp. 335-336), TAFT (1961, pp.561-562), TAFTand HARBAUGH (1964, pp.58-66), VONDER BORCHet al. (1964, pp.lll6-1118), and by others all indicate shallow marine waters and require very special physicochemical conditions, such as high Mg/Ca ratio of medium of deposition. Recent observations, techniques, and analyses have made the picture of dolomite formation much clearer; however, much work still remains to be done. It is still uncertain whether direct chemical precipitation of dolomite (primary dolomite) under “natural”, low-temperature and low-pressure conditions in the Paleozoic seas was possible or not. Many questions concerning secondary dolomitization, such as the relationship between porosity and degree of dolomitization are still unanswered. CHILINGAR et al. (1966) recently discussed some of these problems in detail, and recognized very early diagenetic, early diagenetic, late diagenetic, and epigenetic (subsequent to lithification) dolomites. FAIRBRIDGE (1 966) recognized syn-, ana-, and epidiagenetic stages. (See Fig.4, 5, 6.) STRAKHOV (1956) recognized the following four types of dolomite facies: ( I ) dolomites forming in marginal, supersaline parts of large marine basins (mainly platforms); (2) shallow, supersaline central parts of seas on platforms;

RAINFALL

pH 7

I

Fig.4. Diagenetic stages: syndiagenesis, anadiagenesis, and epidiagenesis. (After FAIRBRIDGE, 1966).

6

R. W. FAIRBRIDGE, C . V. (’IIILINGAR AND H. J. BISSELL

~~~

ic “ H a r d

P w

~

”

Fig.5. Dolomites belonging to different stages of diagencsis, from an idealized sketch of conditions observed in the Eifel district of Germany. (After FAIRBRIDGE, 1957, p.162, fig.13.) ( I ) Devonian Givetian Dolomite-syndiagenetic; (2) Couvinian Limestone (Middle Devonian) - affected by epidiagenesis corresponding to the pre-Ti iilssic land surfaces (a former “hard pan”), but to a limited depth only; (3) Triassic basal conglomerate, containing two types of dolomitic boulders, plus limestone boulders; and ( 4 ) hydrothermal dolomite, possibly enriched by anadiagenetic (deep-seated) connate solutions moving upward along fault plane.

Water

Pressure Depth

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1,000

100 - 3,000

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______

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FREE ACGCSS I N SEA WATER POROUS WALLS

REDUCING CONDITIONS WITH HIGH ALKALINITY BENEATH LAGOON FLOOR SITE OF CONTEMPORARY DOLOMITIZATION AT ABOUT PRESSURE

!!= ‘:a 20 ATMOSPHERES \

ICOMPLETE=STABiL’TY.

TRUNC-,,,

k!

Discontinuous subsidence since early Tertiary

Fig.6. Diagram of an oceanic atoll with a dolomitized core. During thc periods when the lagoon is largely cut off from the ocean, high-density alkaline solutions enriched in Mg content seep downward and dolomitize reef core. Porous and pcrmeable reef walls permit continued refluxing of Mg-rich solutions until dolomitization is complete. (Aftcr FAIRBRIDGE, 1957, p.148, fig.8.)

INTRODUCTION

7

(3) normal marine dolomites, such as the dolomitized reefs of Upper Triassic of Eastern Alps; and (4) dolomites forming in lagoons and in marine bays of the arid zone, but characterized by freshened (by river) waters. The following important types of dolomite deposits were recognized by TEODOROVICH (1958): ( I ) normal-marine calcareous dolomites and replacement dolomitic limestones; (2) primary “pelitomorphic” (fine-textured) dolomites and anhydrite-dolomite rocks; (3) calcareous dolomite deposits of supersaline seas, reef lagoons and periodically drying intertidal calcareous muds; and (4) primarychemical, fine-textured calcareous-dolomitic muds, which form in continental lakes having dry and hot sometimes seasonal) climates. Dolomites of type 2 have formed in supersaline lagoons and may either be interbedded with strata and lenses of anhydrite-gypsum, or they may be primary more of less homogeneous sulfatedolomite rocks. In reviewing the distribution of Paleozoic dolomites, FAIRBRIDCE (1 957) pointed out that in the offshore direction in a non-evaporite facies as one proceeds seaward the dolomitized sediments are generally replaced by nondolomitized limestones of colder and deeper water character. Where the two alternate, one visualizes a eustatic cyclicity. The near-shore dolomitized carbonate is richer in terrigenous insolubles and initially Mg-rich algal nuclei, whereas the non-dolomitized layers lack such nuclei. The correlation between dolomitization and clastics showing nearness to the shore was demonstrated by many authors as exemplified in Fig.7 (see also Fig. 8, 9). This is a very important criterion when preparing paleogeographic maps. Much work, however, still remains to be done on this subject. In certain deep-water limestones, one may also note a rise in the insoluble fraction, but these rocks do not become so readily dolomitized. The insoluble residue is likely to be authigenic quartz or chert that is secondarily derived from fine-grained siliceous dusts or pelagic sources such as Radiolaria or diatoms (CHILINGAR, 1953). Modern seas are noticeably sparse in carbonate environments. There is no single area where sedimentation of various marine limestones, with a spectrum of evaporites and dolomitic materials, can be seen occurring contemporaneously. For this reason certain lacustrine sites have been examined for some of the answers of “evaporite-suite” sedimentation. Great Salt Lake and the Salt Flats of Utah in the western United States have been the subject of recent studies (GRAFet al., 1959, p.1610; 1961, pp. 219-223; GRIMet al., 1960, pp.515-520; BISSELL and CHILINGAR, 1962, pp.200-210). Although this area represents only present-day and Pleistocene carbonate deposition under lacustrine environmental conditions, it nevertheless gives an insight into some of the chemical, physical, physicochemical, and organic processes which may have operated, with certain modifications and variations, in marine realms in the geologic past. For example, thick deposits of salts of various chemical composition can originate in marine basins of different depths. Inflowing

8

R. W. FAIRBRIDGE, G. V. CHILINGAR AND H. J. BISSELL

90

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CONTENT, %

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VERTICAL S C A L E

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50

%

Fig.7. Relationship between the amount of insoluble residue and dolomite content in carbonate portion of clayey, calcareous dolomites and dolomitic limestones. (After ZELENOV, 1956, p.42.) Fig.8. Relationship between the amount of insoluble, terrigenous grains (near-shore indicators) and the degree of dolomitization. Purer limestones with less insoluble material have lower MgCO3 content, suggesting offshore facies. Based on a Lower Ordovician section near Harrisburg, Pennsylvania. (After FAIRBRIDGE, 1957, p.155, fig.9; analyses by LESLEY, 1879.)

oceanic waters across shallow shelves reach a level of saturation for the various dissolved salts at widely distant points, and, under certain Eh and p H conditions, can create overlapping facies. Above a very shallow shelf, which tectonically may be also rather stable, progressive saturation can lead to a high degree of fractional sedimentation. Under a particular set of conditions, calcium carbonate may not extend into the repository where saline to highly saline salts are accumulating; calcium sulfate may also be negligible or present in rather small amounts. Variation in degree of tectonism can, however, bring about an entirely different set of sedimentological conditions, and alternation as well as interdigitation may follow; for example, one may observe a sequence of micritic and algal limestones, dolomicrites,

9

INTRODUCTION TOP

E D HOI

0

sorroul@Ok

EXPLANATION M3C0,

0caco, INSOL

Fig.9. Rhythmic alternation of dolomitized and nondolomitic limestones of a sequence of Lower Ordovician rocks west of Harrisburg, Pennsylvania. Dolomitization always coincides with the beds containing maximum amount of terrigenous (insoluble) material, and, therefore, presumably closer to the original coastline. (Analytic data by LESLEY, 1879; after FAIRBRIDGE, 1957, p. 156, fig.10.)

soft algal dolomicrites, bituminous limestones and dolomites, argillaceous limestones and dolomites, a spectrum of anhydrites and gypsum (gypstone), rock salt, and other sediments which can accumulate in restricted to semirestricted environments. At times there may be an influx of detrital materials from the adjacent land areas, as shown for example by admixtures of red mud, red silt and red-colored NaC1-rich sediments. Large volumes of ocean water are circulating over depositional sites upon shelves over which water depth is more than a few centimeters, possibly some tens of meters, and also as a rule over those shelves of more unstable (subsiding) tectonism. Passage of water across a marginal shelf, rise, reef, ridge, or other shoal area can promote saturation, thus initiating a sequential order of precipitation of dissolved salts. Commonly calcium carbonate will be deposited first; continued fractionation of salts can occur in sequential fashion with progressively but slowly increasing water depths. The latter depends mostly on the ratio of the depth of water lost by evaporation to total depth of water. This is not to infer that the process operates in open deep-seas but rather over thresholds of parts of the neritic zone on which topographic irregularities occur. Deepening of the closed basin may prevent the most concentrated solutions from mixing with those of the lesser

10

R. W.

FAIRBRIDGE, G. V.


depths and thereby trap the salt. Stagnation could dominate in the deep parts of the basin, and deposits of darker-colored (locally bituminous) rock salts could attain thicknesses of 100 ft. to asmuch as 1,000ft. (See also SCRUTON, 1953, pp.2498-2512; and ADAMSand RHODES,1960, pp.1912-1929.) Students of carbonate petrology and petrography long suspected that there was a difference between the chemistry of modern sea water and that of the barred basin seas of the known Paleozoic and Mesozoic salt deposits, which also contain interlayered and interfingered carbonates. Seemingly, in deep-basin deposits the calcium carbonate and calcium sulfate constituents dominate over the sodium chloride accumulations, and locally gypsum-anhydrite deposits are present almost to the total exclusion of carbonates. In shallow-shelf and platform areas, however, sodium chloride may comprise the most extensive deposits; on the other hand, environmental conditions may have prevailed which permitted deposition of thick massive gypsum (gyprock), which can become altered to anhydrite under moderate burial pressure. Bituminous limestone and/or dolomite and “red beds” may also be present. The red-bed material, however, was derived from a terrigenous source and merely represents admixture of detritals into the depositional site. Potash salts are absent or negligible in such platform deposits. The only logical conclusion that can be arrived at when the details of ancient evaporite suites are carefully investigated is that none of these salt deposits represents complete evaporation of a sea water basin (BORCHERT and MUIR,1964). Thus, it would seem that most of the salt solutions that reached a concentration beyond that of saturation for sodium chloride were transferred back to the ocean at depth as return currents (see Fig. 10). This rule also applies to the sodium chloride solutions in the partly closed basins which deposited Ca and Mg salts such as CaC03, MgC03, and CaSO4. Various platforms and shelves, particularly those which are controlled by mild tectonism (subsidence) contain evaporite suites which were evidently precipitated one after another in the inverse order of their solubility, that is, CaC03, CaMg (CO3)2, CaSO4.2H20, NaCl, MgC12 and KClz, etc. It must be realized, however, that the locus of deposition of each of the salts is diachronic, so that the maximum thickness of each salt formation can only be measured in a profile that allows for lateral displacement through time; in short, the salt beds contain lateral and verticolateral facies variants. Over a shallow shelf where evaporation is high in relation to depth of water, saturation is reached first in the inflowing water near the shelf margin, resulting in a calcium carbonate accumulation which normally forms a belt or band parallel to the shelf boundary. It is commonly succeeded landward by massive gypsum deposits while the heavy chloride solutions move out basinward in bottom currents. Progressive landward evaporation of inflowing waters may also be concomitant with density stratification of various salt solutions in deeper portions of the basin (or unstable shelf), and separation of salts from solution should be according to the ionic ratios of the solution. This would result in chemical carbonate sedi-

11

INTRODUCTION

LABRbOOR SEAor NORTH PACIFIC

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Fig.10. Four standard types of barred basins, showing denser water (cross-hatched). Only type AZ appears to be suitable for the formation of chemical (evaporite) carbonates. (Modified after DIETRICH, 1963.)

mentation along the outer border, sulfates in a marginal zone, and the chlorides in the basin center. Such an ideal distribution is seldom attained, however, and neritic facies often display deposition cyclicity of most, if not all, of the saltswhich occur in that environment; even in the center of the basin all phases can be present in one cycle. In the deeper portions of some basins, the carbonates and sulfates exhibit finely laminated characteristics and may be bituminous, especially in the beginning phases of accumulation. These texture and composition are in a marked contrast to those of the reef carbonates, the “primary” and other dolomites, and the mottled and wavy-laminated gypsum of the shallow-water environment. Some of these concepts are summarized in Table I as an attempt at reconstruction of paleogeologic facies. The Permo-Triassic sedimentary sequence in a platform-to-shelf-to-basin repository in northwestern and southern Arizona of the western United States seems to illustrate the concepts advanced above. They also seem to be compatible with the environment reconstructed by MCKEE(1938, pp.129-132) for the Toro-

12

R. W. FAIRBRIDGE, G . V. CHILINGAR AND H. J. BISSELL

TABLE I PALEOGEOLOGIC DISTRIBUTION OF CARBONATES I N EVAPORITE-SUITE FACIES

Light carbonates

I

Dark carbonates

Light-colored limestone and/or dolomite, interbedded with calcarenites and/or red beds. Dolomite, anhydrite, algal- and coarse-textured oolitic limestones or dolomites. Algal dolomite; algal-plate limestone; bioclastic limestone; coarse to fine oolitic limestones and dolomites. Mottled anhydrite or gypsum with thininter-beds of micrite, dolomicrite, and/or algal limestones. Massive anhydrite with clouds and knots of dolomite. Wavy-laminated anhydrite with dolomicrite and micrite lentils and thin inter-beds; may be bituminous. Knotty and “augen” anhydrite, spongy dolomite and limestone. Thin-bedded anhydrite and dolomite or limestone in an interlayered sequence; bituminous. Varved anhydrite, dolomite, and bituminous limestone. Limestone with thin anhydrite layers. Sediment is bituminous. Bituminous carbonates with little or no anhydrite or gypsum.

Shelf I

I I

J.

Basin

weap Formation and Kaibab Limestone of northern Arizona and southern Utah, where there are excellent cyclothems of carbonate, gypsum, and red beds in the sedimentary sequences. McKee’s ideal cyclothem in these Permian strata is as follows: (5) Red beds and gypsum1 ( 4 ) Chemically precipitated limestones (3) Marine limestones with mollusks or brachiopods (2) Chemically precipitated limestones (1) Red beds and gypsum Workers concerned with petrology and petrography of sedimentary carbonate rocks, including modern accumulations of these sediments particularly in the shallow marine realm, have erected a complex nomenclature of descriptive terms many of which give a suggestion as to environment of sedimentation. Although it is not the purposein this introductorypreface to review all these names, itis noteworthy that researchers no longer must hasten to a glossary of terms to determine the meaning of such now well-known names as calcarenite, calcisiltite, tuffaceous carbonate, coralgal limestone, bryalgal limestone, grapestone, hard algal limestone, micrite, dolomicrite, dolarenite, coquinite, and many others. Nomenclatural terms and rock names are the trade-marks of sedimentary petrographers and it behooves the young worker to familiarize himself with this language. 1McKee showed these with 5 at bottom and I at top of list.

INTRODUCTION

13

Inherent in any study of the carbonate rocks is an attempt at classification and “pigeon-holing”. Today numerous schemes of classification of limestones and dolomites are available in published form, and each plan has particular merit. It is quite evident that each of the proposed classifications has resulted from a personal viewpoint which arose from research in a particular province or geographic area. Geologists, like other scientists, often have a tendency toward provincialism and, therefore, each worker is inclined to consider that his proposed scheme lacks the weaknesses of the proposals which preceded it. The student must learn to recognize at the outset the difficulties he will encounter in future research if he clings with undue passion to only one method of classifying his materials. Experienced petrologists and petrographers have long realized that a classification of modern carbonate sediments cannot follow one rigid outline, nor can a single carbonate classification now in vogue take into account all the possible ramifications of origins, environments, diagenesis and other sedimentary processes of ancient sedimentary carbonate rocks. Seven different chapters have been written by foremost experts on carbonate rocks in this book (Carbonate Rocks, part A ) , and are reviewed by the editors in order of their appearance. Physicochemical aspects of carbonate rocks are stressed in Carbonate Rocks, part B.

MODERN CARBONATE SEDIMENTS

Reviewing the sites and characteristics of modern carbonate deposition, W. H. Taft (Chapter 2) commented that it is possible to find them in almost all types of depositional environment (except for the deepest ocean). The relative abundance and preservation of such sediments, however, is highly variable, subject to rigid physical and chemical limitations. Most modern marine carbonates are demonstrably organogenic; some doubtful cases occur where very fine grain sizes make identification difficult. Results of certain isotopic studies, however, have suggested that some of the calcilutites, at least, are also organogenic. On the other hand, most of the intertidal and continental carbonates, beach rocks (and pelagosite), soil crusts (calcrete or caliche), spring deposits (travertine, tufa), and evaporites are of inorganic origin. The key minerals represented in modern carbonate environments are (a) low-magnesium calcite (characteristic of pelagic Foraminifera and most deep-sea deposits; also present, but not dominant, in warm neritic areas); (b) high-magnesium calcite (commonly found in calcareous Algae and in many invertebrate exoskeletons, especially those of warm habitats); (c) aragonite (typical of scleractinian corals and certain molluscan shells, notably of warm habitat; also occurs as inorganic cement in intertidal beach rocks); and ( d ) dolomite (inorganic product in superheated intertidal localities; also as rare scattered crystals at abyssal depths).

14

R. W. FAJRBRIDGE, G. V. CHILINGAR AND H. J. BJSSELL

Diagenesis of modern carbonates is rapid if they are exposed to fresh water. Thus, beach rock cements that are several thousand years old are converted to calcite. Modern neritic and bathyal carbonates are largely uncemented, but slow changes are taking place, which require further investigation.

PETROLOGY A N D PETROGRAPHY OF CARBONATE ROCKS

The petrology and petrography of the carbonate rocks has been discussed in Chapter 3 by an experienced group of French and Italian geologists: Yvonne Gubler, along with J. P. Bertrand, L. Mattavelli, A. Rizzini and R. Passega. This authorship reflects two interesting trends: ( I ) the steadily growing importance of research by the Institut FranGais de Petrole and AGIP (the Italian National Corporation), and (2) the major role played by carbonates in oil and gas basins of western Europe (notably Aquitaine and Sicily), North Africa and the Middle East. Petrology deals with the geometry of the carbonate rock, its sedimentology, bedding, intraformational structures, and field appearance. In contrast, petrography involves description of features observable on a smaller, microscopic scale: the sedimentary micro-textures, cement and granular detail. Petrographic studies provide feed-back of information to complete the petrologic picture. Statistical integration can be achieved through the method worked out by IMBRIE and PURDY (1961). Carbonate rocks are polygenetic; they are composed of chemical, biochemical and clastic components, generally in various mixtures. The dominant cations in major carbonate minerals are Ca2+, Mgz+, and Fez+; minor components include particulate clays, silica, phosphates, oxides and sulfides. Three fundamental features of carbonate rocks are recognized: (a) Role of orgunisms. Being able to modify the local pco2 directly or indirectly, organisms (especially the photosynthetic Algae and bacteria) play a dominant role in carbonate petrology. Thus, the paleoecology of these organisms is critical t o the understanding of petrology of carbonate rocks. (b) Metastable minerals. Metastable minerals, which are common amongst the newly formed carbonates, control much of the diagenetic evolution of these rocks. High-magnesium carbonate leads the way to stable dolomite; aragonite soon inverts to low-magnesium calcite. Distinguishing primary from diagenetic crystals is not always easy, and the time of diagenesis may range from the earliest t o the latest phase of the rock's history (FAIRBRJDGE, 1967). (c) Past distribution. Past distribution of carbonates was not always the same as today. The fundamental dichotomy took place in the Cretaceous, with an appearance of widespread pelagic organisms (Foraminifera and Coccolithophoridae). Prior to this time the dominant facies were neritic and organisms were

INTRODUCTION

15

mainly benthonic. Phylogenetic position of the principal carbonate rock-builders has changed notably from time to time. As established by GRESSLY, already in 1838, the description of the rock in its total aspect may be termed facies, both lithofacies and biofacies. The interpreted facies is called the petromodel by our European colleagues. On the micro-scale, each individual lamina possesses a characteristic aspect. This is what CUVILLIER (1951, 1952) has termed the microfacies (both the mineral grains and the fossil fragments), which in thin-section or polished (or etched) surface may be employed for correlation (see FAIRBRIDGE, 1954), or may serve to work out the genetic history (CAROZZI, 1950, 1958, 1959, 1961). A succession of microfacies, that is a series of laminae, represents a micvosequence. Brief breaks in sequence (diastems) are recognizable. Micro-crossbedding gives directional indication. Rhythmic sequential repetition, reflecting climatic or tectonic events, may be traced by the geochemical, granulometric or paleontologic studies. Diagenetic changes may appear at any stage of the lithogenesis. They may be due to organic activity (bioturbation, ingestion, etc.), to biochemical effects, or to physical processes. The latter are numerous and include compaction, crystallization, etc. The changes may be brought about by any of three types of waters: by connate waters, by hydrothermal waters, or by meteoric waters. New minerals may form (“neogenesis” or “authigenesis”) and particle size and shape may change (“epigenesis”), with the consequent changes in the porosity and permeability of rocks. Time of tectonism in part controls the time when the various phases are instituted. Inasmuch as the alternation of microfacies leads to anisotropies, marked variations in porosity, etc. are observed; accordingly, the examination of randomly oriented thin-sections is very important. The recent history of carbonate petrography rests virtually in the hands of two men. First, there was CAYEAUX (1935), who, on the basis of vast numbers of thin-sections of samples from all parts of France, and elsewhere, compiled a comprehensive monograph, almost an encyclopedia, of all types of carbonates known to him. (He also prepared comparable works on the phosphatic and siliceous sediments of France and her sphere of influence, interpreted in the broad sense). Secondly, there was SANDER (1930) in Austria, who, on adding stereographic projection to the thin-section technique, first studied carbonate petrofabrics in an attempt to understand the original organization of these rocks, and to distinguish primary from diagenetic features. Sander’s studies opened the way for an evolution from the purely descriptive petrography to interpretive petrology integrated with field study of environments, ancient and modern. Modern development of petrographic techniques resulted in a number of most helpful refinements, extensions, or simplifications. CUVILLIER (1951, 1952) studied the role of micropaleontology in understanding microfacies. CAROZZI (1950, 1958, 1961) applied graphic and statistical methods to the study of distri-

16


bution and frequency of particles and organisms. FRIEDMAN (1959) and others brought out the value of staining methods for mineral recognition. SCHMITT (1963) applied microradiography t o identifying dolomite in fine-grained dolomitic lime(1963) injected colored stones. I n studying porosity and permeability, ETIENNE synthetic resins into carbonate rocks in order to examine the geometry of pore spaces. Anadiagenesis and tectonism, still very little known, result in the reorientation of crystals in a rock; some recent laboratory work in this area has been done by D’ALBISSIN (1963). Evidently the use of conventional microscopes is inadequate for a complete understanding of this phenomenon and developments now are being made in the areas of electron microscopy, X-ray diffractometry and emission spectrography. Statistical data are being processed by digital computers. Yvonne Gubler and her colleagues described a number of field studies in detail. First, they presented studies on the “Pisolitic Limestone” of Vigny, France. They rightly avoided the temptation of allocating a special rock name to every unique rock, but urged the use of Folk’s textural geometric (granulometric) method, e.g., “biosparitic calcirudite”. A second example deals with facies variations in the Triassic-Jurassic reef and basin formations of the Ragusa area of southern Sicily. Much of the primary texture is destroyed by dolomitization. An interesting feature of the area is the persistence of upward growth by a reef, with its shallow platform facies, for a total thickness of 2,500 m (8,200 ft.). This exceeds any measured Cenozoic subsidence of modern Pacific atolls or quasicratonic submarine plateaus. The authors stressed that conventional petrographic examination provides an adequate rock description only in exceptional cases. On the scale of basin dimensions, a serious problem in the way of quantitative integration exists because of inadequate data. Systematic core drilling and standardized data sampling are particularly important to establish comparable profiles and to integrate field descriptions and information obtained on studying thin-sections. This integration is the petromodel of our authors.

CLASSIFICATION OF SEDIMENTARY CARBONATE ROCKS

The question of classification of carbonate rocks has been a long and vexed one, especially when compared with the classifications of other sedimentary rocks. The reason is not difficult to appreciate: other sediments for the most part are fairly easily disaggregated into component grains, which may then be sorted for size analysis; also they respond to fairly straight-forward compositional or mineralogical analysis. On the other hand, calcite, dolomite, aragonite and high-magnesium calcite are all minerals that are highly susceptible to diagenetic change, to partial or total recrystallization, and to selective leaching and replacement. Fossils and distinctive primary syn-sedimentary structures of calcareous sediments may be

INTRODUCTION

17

totally destroyed or largely masked during diagenesis, which may include dolomitization or other fundamental modifications. Many rock types may be classified as biogenic and inorganic, but limestones often constitute a complex mixture, for example, an organogenic framework with an inorganic cement. In Chapter 4, H. J. Bissell and G. V. Chilingar considered the many alternative classifications that have been offered, with special attention to the flood of schemes that have appeared over the last decade or so. They stressed the importance of four parameters in particular: ( I ) Composition. Most authorities subscribe to a basic classification close to that of Cayeaux (1935, interpretation of data) for limestone-dolomite mixtures: limestone (( 5 % dolomite mineral), magnesian limestone (5-10 % dolomite), dolomitic limestone (10-50 % dolomite), calcitic dolomite (50-90 % dolomite), dolomite rock (( 10% calcite). Little support is given to the term “dolostone” although it is attractive; geologists are mostly content to see from the context whether “dolomite” refers to the mineral or to the rock. CHILINGAR (1957) rather favored a system of Ca/Mg ratios for compositional definition: normal dolomites (Ca/Mg = 1.5-1.7), calcitic dolomite (Ca/Mg = 2.03 3 , etc. TEODOROVICH (1958) provided further for a clay fraction: clayey dolomite (35-90 % dolomite mineral, 0-45 % calcite, 30-10 % clay), etc. (2) Texture. Bissell and Chilingar recommended a system of classification broadly based on those of LEIGHTON and PENDEXTER (1962) and on PLUMLEY et al. (1962). Three gross texture categories are identified on the basis of their genesis: (a) Accumulated in place (chemical and biochemical)-non-granular, colloform, etc., lack organic traces, e.g., travertine. ( b ) Accumulated in place (organic frame-builders)-e.g., coral, bryozoan, etc. limestone ( )90 % grains), micritic coralline limestone (50-90 % grains), or coralline micritic limestone (( 50 % grains). (c) Clastic and wave-moved-clastic limestones are further subdivided into those with abraded grains (detrital, skeletal, and pellets in part) and those having accretion-aggregation grains (some pellets, lumps and coated grains). Under each of these categories, Bissell and Chilingar recognized: ( i ) )90 % grains (detrital limestones, including calcarenite, etc., from pre-existing rocks only; skeletal limestone, pelletal limestone, lump limestone, oolitic limestone, etc.); (ii)90-75 % grains (micritic, detrital limestones, etc.); (iii)75-50 % grains (micritic-detrital limestones etc.; (iv) 50-25 % grains (detrital-micritic limestone, etc.); (v) 25-10 % grains (detrital, micritic limestone, etc.). Most of these criteria can be established with the aid of a hand lens and a dropper with dilute HC1. (3) Grain-micrite ratio ( G M R ) . GMR shows the relative amounts of coarse and fine textured carbonate material, theoretically related to wave or current action. For rocks having 90% grains, GMR is 9 / l ; for 50% grains, l / l ; and for 10% grains, 119.

18


(4) Energy index (EI). EI is a quantitative index for the appraisal of the rock and its environment. )90 % grains indicate strongly agitated waters; 75-90 %, moderately agitated; 50-75 %, slightly agitated; 25-50 %, intermittently agitated; and 10-25 %, gently agitated. With regard to secondary dolomites, most of the limestone classifications may be applied if the original texture is not obliterated. Limestones altered to cryptogenic stages may contain “ghost” relics and granular traces. Totally crystalline phases (aphanic or phaneric) can be classified only on basis of simple grain dimensions. Possibly, oxygen isotope studies may enable one to distinguish between primary (direct precipitates out of sea water) and diagenetic-epigenetic dolomites. Some criteria may be used to distinguish between the latter two (diagenetic or epigenetic) dolomites also (CHILINGAR et al., 1967).

ORIGIN AND OCCURRENCE OF LIMESTONES

The origin and occurrence of limestones is analyzed rather exhaustively by John E. Sanders and Gerald M. Friedman (Chapter 5). They have reviewed the major literature and from personal travels were able to report first hand on many examples of both ancient and modern environments of limestone formation. As to origin, they recognized: (a) primary “stony precipitates” (i.e., skeletons, reef structures), (b) lithification of carbonate sediments, and (c) replacement deposits, such as calcite replacing gypsum, with bacterial assistance. The occurrence of ancient limestones, and their modern prototypes is enormously widespread over the earth in both marine and continental settings. They have received in the past certainly more than their fair share of attention as compared with shales and sandstones, although they comprise only 10-15%1 of the continental sedimentary crust. The explanation is partly to be found in the fact that limestones frequently carry a far richer record of organic life than do the noncarbonates; there are also economic aspects, such as the widespread oil reservoirs in carbonate rocks and their world-wide use for cement manufacture. Limestones are being formed today, anywhere from the deep ocean basins to the interiors of continents. In nonmarine environments, limestones are found sometimes in lake depressions and basins; they may develop as primary precipitates or as derived terrigenous fragmentals. Eolian sediments near the coast are frequently rich in carbonate clastic material and readily lithify to form dune calcarenites or eolianite; interdune paleosols often contain calcrete or travertinous calcareous duricrust, comrnonly called caliche in the American southwest. In shallow seas of the warmer continental shelf regions and in analogous lSome estimates go as high as 25 %.

INTRODUCTION

19

seas from offshore platforms (e.g., the Bahamas, the Maldive-Laccadives, the South China Sea and the Coral Sea) in regions of quasicratonic intermediate crust, there are perhaps the greatest diversity of carbonate occurrences. The shallow environment renders them susceptible to rapid facies change under changing energy conditions, due to shore-line changes associated with eustasy, tectonics, or sedimentary processes. Biofacies of great variety here contribute to the lithologic heterogeneity. On the continental shelves there is potentially a considerable mixing with noncarbonate terrigenous material. Pure carbonate deposits are more likely to occur on the intermediate platforms; but here also there is an admixture of clastic, pre-existing terrigenous limestone debris. In the shelf or intracratonic environments, there are numerous cases of mixed evaporite-carbonate sequences; this is particularly true of the past, but not today, because the present sea-level regime is only about 6,000 years old. In deep-sea environments of today, and back to Cretaceous times, there are enormously widespread pelagic carbonates, coccolith ooze, Globigerina ooze; pteropod ooze, etc. The rock equivalents were mainly the chalks and chalky marls. Prior to Cretaceous these were rare or absent, a phenomenon of geologic history closely related to the sudden flowering of pelagic carbonate organisms in the Late Mesozoic, which in turn is probably related to gross changes in paleogeography and marine geochemistry. Particular attention is drawn to the authors’ treatment of the Capitan Limestone of the classic Permian section of West Texas, U.S.A., on which they do not agree. Friedman supports the widely accepted barrier-reef hypothesis; whereas Sanders advocates many of the views of his former student, G. W. Moore, who has raised doubts on many important points of the barrier-reef interpretation and has substituted the rather astonishing alternative that the fine-grained, massive, high-calcium, unfossiliferous limestone of the Capitan Formation is not recrystallized reef rock, but rather originated by wholesale replacement of massive Castile anhydrite. Moore’s alternative contains numerous important stratigraphic and paleogeographic corollaries, which are of more than academic interest to the geology of this petroliferous province. Another subject worthy of comment is the treatment accorded gravitydisplaced, calcareous, deep-sea deposits. In addition to their discussion of calcareous “turbidites”, the authors emphasize the paleogeographic importance of limestone-pebble “Wildflysch” accumulations in defining the location of ancient steep continental slopes. European readers, in particular, may be interested in this subject in view of the ideas about “thrust conglomerates” which have persisted in the tectonic literature of the Alpine chains. A description of true limestone thrust conglomerate of nonmarine origin-a fan deposit described by C. R. Longwell from southern Nevada, U.S.A.-is included by way of comparison. Another feature of this and the succeeding chapter will also be welcomed by European readers: the metric system has been employed.

20


ORIGIN AND OCCURRENCE OF DOLOSTONES

As pointed out by G. M. Friedman and J. E. Sanders in Chapter 6, the “dolomite question” has defied solution until the recent advent of X-ray spectroscopy, which made dolomite identification simple, fast, and definitive. As a result of the X-ray and stable isotope (carbon and oxygen) analysis and field studies of Recent and ancient carbonate sediments, researchers seem to be on the verge of a breakthrough in this field. ADAMS and RHODES (1960) proposed that most dolomite (dolostone) deposits in the geologic record owe their origin to hypersaline brines, and must, therefore, be considered as related to evaporite deposits. The “capillary concentration” and “seepage refluxion” seem to our authors also to be the two main processes responsible for dolomite formation. The former is the effective process in supratidal or intertidal zones or along the margins of basins, whereas the latter is operative under deeper water conditions. By capillary concentration, when sea level is low, the interstitial waters in the sediments transpire upward through the porous tidal flat and coastal sediments and evaporate below the sediment-air interface; this process is similar to that under which some caliche is formed. Refluxion evaporation increases the salt concentration and the density of the water in any restricted lagoon or intermontane basin where evaporation periodically exceeds precipitation. This produces a heavy brine that migrates to the lowest possible topographic depressions, and seeps slowly through the underlying sediments resulting in a progressive doloniitization. Dolomite may also form directly from the brine, with aragonite as a possible transitional phase. The Mg/Ca ratio of the brine, which should be higher than that of the normal sea water, may be raised by the removal of calcium as a result of deposition of gypsum in very shallow water pools where the supply of oxygen is ample. Gypsum tends to be broken down to H2S and iron sulfide by bacteria where the oxygen supply is less. On the floors of such pools near the Persian Gulf today, layers of knife-edged crystalline gypsum can be lifted to expose interfingering layers of black mud rich in H2S. Consequently the layered dolomite forming under these conditions is indicative of a somewhat deeper reflux environment, with gypsum being deposited along the shore. Both capillary concentration and refluxion may operate in the same basin. Friedman and Sanders recognized three different types of mineral dolomite: ( I ) syngenetic, (2) diagenetic, and ( 3 ) epigenetic. These terms imply a genetic sequence and have been used by these authors as a basis of classification of dolomites. Syngenetic dolomite was defined by our authors as a dolomite that has formed penecontemporaneously in its environment of deposition as a micrite or as fine crystals. According to them, diagenetic dolomite forms by replacement of preexisting calcium carbonate sediments following consolidation of sediments or

21

INTRODUCTION

coincident with it. Diagenetic dolomite also may have formed penecontemporaneously by replacement of grains and cement of calcium carbonate sediments. Thus, in borderline cases the distinction between syngenetic and diagenetic dolomites becomes difficult or impossible. FAIRBRIDGE (1966) puts both types into subcategories of syndiagenetic dolomite. Epigenetic dolomite of Friedman and Sanders formed by replacement of limestone, and is localized by post-depositional structural elements. It is closely related to fractures and faults in carbonate rocks; and the term “epigenetic” is synonymous with such terms as “structurally-controlled’’ or “tectonic” dolomites. Many, but not all, epigenetic dolomites are genetically associated with metallic deposits, notably of lead and zinc minerals. These dolomites are the anadiagenetic rocks of FAIRBRIDGE (1967). He recognized a distinct fourth type, related to meteoric water circulation, which he has named epidiagenetic. Dolomite formed after burial, which includes all “epigenetic” and most “diagenetic” dolomites, owes its origin to subsurface brines. These subsurface formation waters are depleted in both so42- and Mgz+ ions, which probably indicates that sulfates are involved in the dolomitization process. Dolomites may revert back to limestones by the process of dedolomitizationl during “epidiagenesis”. A genetic relationship appears to exists between the presence v l dispersed anhydrite and dedolomitization, with MgS04 being carried away (in solution) as a by-product, according to the following reaction: Ca.Mg(CO3)z

+ CaS04 + 2CaC03 + MgS04

CARBONATE OIL RESERVOIR ROCKS

In Chapter 7, role of carbonates as oil reservoirs is discussed by John Harbaugh. This is an aspect of physical petrology that depends largely upon porosity and permeability (mass properties of rocks). The ratio of total pore space to the bulk volume of rock is the absolute porosity, which is of little value economically if the pores are not interconnected. Obviously, it is the ratio of the interconnected pores to the totalvolume of rock (the eflectiveporosity)that is of greatest interest. Inasmuch as the fluids have various viscosities, and the connecting passages have different degrees of tortuosity, a calculation of effective permeability is also required. When two or more immiscible fluids are present in a porous solid, the permeability to any of the fluids is termed the effective permeability of the solid to that fluid. It is often convenient t o convert effective permeabilities into permeabilities relative to some standard base (often the single phase liquid permeability), in which case they are called relative permeabilities. Relative permeability curves for a carbonate reservoir rock are also presented in Chapter 7. 1

Some authors prefer the term “calcitization”.

22

R.

W.

FAIRBRIDGE, G . V.


Primary porosity is that ielated to the initial sedimentary features that are only slightly modified by diagenesis, etc. Several varieties of primary porosity include: (a) Framework porosity, which is controlled by the presence of water-filled cavities in the initial sediments (organogenic, e.g., corals, sponges, calcareous Algae, etc.). (b)Mudporosity, which is due to the original presence of vast numbers of microscopic pores filled with connate water, that was not lost entirely during consolidation. The pores are so small that they normally do not contribute to any economical reservoir. (c) Sand porosity, which owes its origin particularly to the presence of oolites, pellets and fragmented skeletal calcarenites, is often very important. Its relative value is related to the sand/mud ratio, which in turn depends on the rate of deposition and the energy of the regime. Cementation is a negative feature as far as reservoir potential is concerned, because the extent to which cements (carbonate, silica, gypsum, etc.) have filled the voids is inversely related to the porosity and permeability. Secondary porosity usually refers to voids (either primary, or due to fracturing) enlarged by solution. Some of this solution may occur early during syndiagenesis, with the removal of aragonitic shell parts, etc., by the acid waters being generated as a result of bacterial activity. Further solution occurs during epidiagenesis when the entire sequence is subject to circulation of acid meteoric waters. Inasmuch as during the deep, anadiagenetic stage the interstitial waters are normally alkaline, they are saturated with respect to CaC03. Solution is restricted to the early and late phases of diagenesis. Calcite is considerably more soluble than dolomite under comparable conditions. In most reservoir rocks one can observe signs of both solution and reprecipitation, sometimes i n a multiple generation sequence. Circulation of solutions is greatly aided by the presence of unconformities and bedding planes between beds having variable porosities, and the growth of joints and fractures. Particularly the latter process, sometimes called “diaclastic revival”, may occur repeatedly during the later phases of the evolution of a basin. This leads to widespread fluid motion, both from uprising connate waters and downworking meteoric waters. Dolomitization has long been recognized as a valuable potential source of porosity, because a molecular replacement of limestone by dolomite would lead to about 12-1 3 % loss of mineral volume. Some investigators, however, believe that the replacement may also occur on a volumetric basis, leading in fact to reduced porosity. Different sedimentary components, e.g., an easily dolomitized, magnesium-rich algal lime mud may be variably graded or interstratified with pure calcitic skeletal fragments that do not dolomitize as easily. Generally, it may be said that dolomitization increases the porosity; this conclusion is based on extensive statistical studies. Consequently, plotting of isopleths of equal Ca/Mg ratios and directions in which Ca/Mg ratios decrease are of valuein exploration for oil and gas (CHiLImAn, 1953, 1956). At low percentages, the dolomite occurs as scattered euhedral rhombs, producing little porosity change

23

INTRODUCTION

up to 50 % of the original porosity. At higher concentrations of dolomite, porosity rises notably. In one instance studied, however, above about 80 % dolomite, the porosity decreases sharply again. Apparently this is due to the filling of the pore spaces by the precipitates out of late-stage dolomitic solutions. In another case, large calcitic fossil fragments have not been dolomitized, but were later removed by solution, increasing porosity correspondingly. Pore geometry is important and is also subject to systematic dimension analysis. The nature of the liquid (especially its viscosity) is critical in evaluating the permeability. In most reservoirs the pores are lined by a continuous thin envelope of water, whereas the oil occurs as globules in the middle. High interfacial tension between oil and water often prevents the passage of the oil through pore interconnections (“throats”). Migration will occur when the buoyancy and hydrodynamic forces exceed this interfacial tension (also modified by the throat dimensions). These characteristics may be tested in the laboratory by mercury injxtion technique. When oil, gas and water are present in rocks, the nature of movements of fluids becomes a complex problem, as shown in Fig.11. Although Fig. 11 applies best to sandstones, it does show the complexity of the problem. Pore geometry is subject to further study in terms of petrogenesis and GAS, 100 O/o

WATER, 100°/o

% WATER

SATURATION

O I L , 100%

Fig.11. Areas of one-: two-, and three-phase flows in a porous medium. (EREMENKO, 1960, p.530.)

24


paleoecology. Original sedimentary facies variations often had profound effects on subsequent development of porosity and permeability. Consequently, the paleogeographic studies have been stressed, particularly during the last decade or so. Analysis of original environment, e.g., basin, shelf, bank, reef, back-reef, barrier beach, lagoon, etc., is of great help in identification of porosity characteristics and the prediction of porosity trends.

CARBONATES AND PALEOCLIMATOLOGY

In the final Chapter (8), Fairbridge reviews the carbonates as climatic indicators, first at the present-day and then at various stages of the geologic record. “Mineral thermometers” are commonly used by petrologists for appraising the temperature level of metamorphic and igneous events in geologic processes. By the same token, calcite, aragonite and related carbonates can be employed in judging thermal levels of former low-temperature environments; certain element ratios (Ca/Mg, Sr/Ca, Ca/Fe/Ti) and certain isotopes may also be used to help bring a degree of precision to these estimates. The degree of confidence in results decreases as one goes back in time, because there are large numbers of variables; and few, if any, unique solutions arc offered. For Cambrian times, for example, the diameter of the earth is not known, nor is the length of the day, the salinity, alkalinity or pH of sea water; and little is known of the ecologic habits of the organisms prevalent in Cambrian seas. Such looseness might easily lead t o an atmosphere of despair among researchers, yet in fact there are many aspects where the experienced geologist feels intuitively on rather safe ground. He knows the limits of metabolic tolerance in organisms and the stability fields of many sedimentary minerals such as the evaporites. The mean temperature of the earth’s surface may be stated reliably as falling in the vicinity of 20 & 10°C over the last 3 billion years. Paleontologically he can judge with a high degree of reliability certain other environment characteristics from the locomotory habit of organisms (sessile or nektonic, benthonic or pelagic), by studying their appropriate swimming, flying or attachment appendages and by the lithofacies status of the enclosing sediments. Each deduction can be made and appraised on its own basis. The results can then be synthesized so that a “general impression” can be achieved, and certain key indicators can be specified. Of great philosophic as well as practical geologic interest is the fact that a series of critical periods has been worked out, representing great biogeochemical revolutions in the history of the planet earth. These critical stages do not represent any sort of deus ex machina. There is no abandonment of the basic methodology of actualism or the philosophy of uniformitarianism; the past continues t o be judged by the methods and observed processes of today, but a new dynamic approach has t o be adopted. This planet is known to be a member of a

INTRODUCTION

25

solar system and although governed by laws of celestial mechanics, is subject to a systematic evolution through time. Furthermore it is periodically modified by relatively small changes in mass relations and in its energy fluxes, many of which are astronomically predictable, and, if not yet quantitatively, at least qualitatively established. The nature of early carbonate sedimentation can be approached in two ways: (a) by taking a working model of the proto-earth and deductively following through its most probable evolutionary steps; and (b) by taking the known parameters of the present and working backwards. The one may be constantly checked and conpared with the other. Series of models along these lines have already been attempted by various writers: Daly, Kuenen, Conway, Rubey, Urey, Holland, Rutten, Fischer, Berkner and Marshall, and others. A high degree of concordance has now been reached about the general character of the earliest earth and the gradual evolution of our hydrosphere, and the atmosphere of Nz, 0 2 and COZ.In this study, Fairbridge uses the carbonates to demonstrate the threshold stages of this evolution, each major step being designated as a “biogeochemical revolution”. They are: Revolution I. “First Life” (about 3.8 0.3 * lOQyears ago), based on the knowledge that the lime-fixing Algae are isotope-dated in the Bulawayan Dolomite Series as not less than 2.7 x lo9 years old; such Cyanophyte Algae are known today to be quite highly evolved, and as ELSASSER noted at National Academy of Sciences Symposium (1965, p. 1214), even the most simple organisms living today secrete upwards of 2,000 separate enzymes. A long time-span must have extended between their appearance and that of the first life. The latter is visualized as the first self-reproducing organic molecules, possibly of the virus type. During this phase the atmosphere is believed to have been close to the Schmidt-Urey model: CH4, NH3 and HzO with traces of other components notably HzS. The proto-sea was a reducing agency and erosion was essentially limited to mechanical processes. More HzO with COZ,SO2 and HC1 gradually evolved as volcanic gases. Revolution IT. “First Photosynthesis” (about 2.9 i- 0.2 . lo9 years ago) was marked by the beginning of photosynthetic use of COz, with 0 2 as a by-product of immensevital potential. To begin with, all the 0 2 liberated was used up immediately for the oxidation of surface minerals. Lacking a thick atmospheric blanket, as pointed out by BERKNER and MARSHALL (1964), solar UV transmission had split many of these OZ molecules close to their source and produced ozone at ground level. Inasmuch as the latter is a powerful oxidation agent, it contributed to the rapid change of the earth’s superficial environment from a reducing to an oxidizing one. World-wide deposits of SiOz-Fez03 (silica-ironstone) formed during this phase, but never again on this vast scale. Revolution 111.“First Carbonate Shells” (about 6 f 0.3 . 108 years ago) has long been recognized as one of the great enigmas of the geologic record. It is here suggested that the Precambrian atmosphere was richer in COZ (from very active

26


volcanicity) while thepo, built up only slowly. Several authors believe the “Pasteur Level” (poZ at 1 % of the P.A.L. or “present atmospheric level”) was reached at the time of this revolution. If the p H passed from the acid side to the alkaline at this time, the first organisms to form shells simply responded to the difficulty of removing alkali waste products from the body fluids into an environment of high ionic strength and high alkalinity, a serious problem in osmosis. Study of the crustacean carapace formation and the enzyme control of body fluid pH during moulting is a rewarding and illuminating experience. Revolution IV. “Great Coal Age” (about 2.5 i0.3 108 years ago) represents the culmination of the Paleozoic history with its continuous record of vast carbonate removal from the ocean, as limestones, followed by the coal burial of Carboniferous and Permian ages (the latter in the Southern Hemisphere). The result was a lowering of pcoz and achievement of a peak in po2 which favored the mammalian and land reptile respiration during Mesozoic time. RUTTEN(1962, 1965) has suggested that this high 0 2 level is necessary for the respiration and metabolism of the great flying insects of the late Paleozoic, when ancestral dragon flies achieved wing-spans of 3 ft. and more. Revolution V. Marking the first appearance of world-wide pelagic Foraminifera and coccoliths (about 1.0 & 0.2 108 years ago), Revolution V is the last of these great thresholds. As pointed out by KUENEN(1950), this event shifted the fundamental site of carbonate sedimentation from shelves to ocean basins (above the “calcium carbonate compensation depth”). From Cretaceous onwards, the carbonates, instead of being limited to shallow epicontinental basins and in part continuously recycled, began to accumulate with more or less continuous burial. It would appear that the Cretaceous Ocean was more extensively saturated or supersaturated with respect to CaC03 than that of today. The precipitation of CaC03 releases C02 (from the bicarbonate ion); and during the Cenozoic the ocean appears to have slowly become less alkaline, and also cooler. Inasmuch as the recharge rate of CaC03 is very high (it can be totally replaced in about one million years), there is little danger of reverting to an acid ocean. Increased atmospheric CO2 content from combustion of industrial fuel will not achieve this possibility, because there is more than adequate reserve of limestone in the ocean floor to buffer any such C02 imbalance. To conclude, there is a very important “moral)’ that may be drawn from this analysis. It is that a thorough appreciation of the nature of biologic metabolism is essential in the training of the thoughtful geologist. Without fairly intimate knowledge of the nature of metabolic limits, photosynthesis, shell calcification, body fluids, osmosis, enzymes, genetics and other aspects of living organisms, the history of the earth‘s atmosphere, ocean and sediments becomes meaningless. Whereas the biological substrate is certainly an inorganic surface, and organic evolution has certainly been in part a response to the physicochemical environment, that environment itself is essentially the product of biological activity.

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INTRODUCTION

27

REFERENCES

In reviewing various chapters, in many instances the editors used the same references used by the author of a particular chapter; these references are not repeated here.

M. L., 1960. Dolomitization by seepage refluxion. Bull. Am. Assoc. ADAMS, J. E. and RHODES, Petrol. Geologisfs,44: 1912-1920. ALDERMAN, A. R. and SKINNER, H. C. W., 1957. Dolomite sedimentation in the southeast of South Australia. Am. J. Sci., 255: 561-567. BISSELL, H. J. and CHILINGAR, G. V., 1962. Evaporite type dolomite in salt fiats of western Utah. Sedimentology, 1 : 200-210. BORCHERT, H. and MUIR,R. O., 1964. Salt Deposits. Van Nostrand, London, 338 pp. CHAVE, K. E., 1954. Aspects of the biogeochemistry of magnesium, 2. Calcareous sediments and rocks. J. Geol., 62: 587-599. CHILINGAR, G. V., 1953. Use of Ca/Mg ratio in limestones as a geologic tool. Compass, 30: 202-209. CHILINGAR, G. V., BISSELL, H. J. and WOLF,K. H., 1966. Diagenesis of carbonate rocks. In: (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, G. LARSEN and G. V. CHILINGAR pp. 179-322. DIETRICH, G., 1963. General Oceanography. Interscience, New York, N.Y., 588 pp. EREMENKO, N. A., 1960. Petroleum Geology (Handbook), Principles of Petroleum Geology. Gostoptekhizdat, I: 592 pp. FAIRBRIDGE, R. W., 1950. Pre-Cambrianalgal limestone in Western Australia. Geol. Mag., 87 (5): 324330. FAIRBRIDGE, R. W., 1954. Stratigraphic correlation by microfacies. Am. J. Sci., 252: 683-694. FAIRBRIDGE, R. W., 1957. The dolomite question. In: R. J. LEBLANC and J. G.BREEDMG (Editors), Regional Aspects of Carbonate Deposition ( A Symposium)-Soc. Econ. Paleontol. Mineral., Spec. Publ., 5 : 125-178. FAIRBRIDGE, R. W., 1964. Theimportance of limestone and its Ca/Mg content to paleoclimatology, In: A. E. M. NAIRN(Editor), Problems in Paleoclimatology. Wiley, New York, N.Y., pp. 431-530. FAIRBRIDGE, R. W., 1966. Phases of diagenesis and authigenesis. In: G. LARSEN and G. V. CHILINGAR (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, pp, 19-89. FRIEDMAN, G. M., 1959. Identification of carbonate minerals by staining methods. J. Sediment. Petrol., 29: 87-97. GRAF,D. L., EARDLEY, A. J. and SHIMP,N. F., 1959. Dolomite formation in Lake Bonneville, Utah. Bull. Geol. SOC.Am., 70: 1660 (abstract). GRAF,D. L., EARDLEY, A. J. and SHIMP,N. F., 1961. A preliminary report on magnesium carbonate formation in glacial Lake Bonneville. J. Geol., 69: 219-223. A. V., 1960. Clay mineralogy of the sediments of the GRIM,R. E., KULBRICKI,G. and CAROZZI, Great Salt Lake, Utah. Bull. Geol. SOC.Am., 71: 515-520. LESLEY, J. P., 1879. Notes on a series of analyses of dolomitic limestone rocks of Cumberland County, Pa. Second Geol. Surv., Penna. (1876-1878), MM: 311-362. MCKEE,E. D., 1938. The environment and history of the Toroweap and Kaibab Formations of northern Arizona and southern Utah. Carnegie Znst. Wash. Publ., 492: 268 pp. RUKHIN,L. B., 1961. Principles of Lithology. Gostoptekhizdat, Leningrad, 779 pp. F. F., Jr., 1962. Grains of detrital, secondary, and primary dolomite from Cretaceous straSABINS, ta of the Western Interior. Bull. Geol. SOC.Am., 73: 1183-1196. SCRUTON, P. C., 1953. Deposition of evaporites. Bull. Am. Assoc. Petrol. Geologists, 37: 2498-2512. SKINNER, H. C. W., SKINNER, B. J. and RUBIN,M., 1963. Age and accumulation rate of dolomitebearing carbonate sediments in South Australia. Science, 139: 335-336. STRAKHOV, N. M., 1956. About types and genesis of dolomite rocks (status of knowledge). In: and G. I. BusmNsmY (Editors), Types of Dolomite Rocks and their Genesis N. M. STRAKHOV -Tr. Geol. Inst., Akad. Nauk S.S.S.R., 4: 5-27.

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STRAKHOV, N. M., BRODSKAYA, N. G., KNYAZEVA, L. M., RAZZHIVINA, A. N., RATEEV, M. A., D. G. and SHISHOVA, E. S., 1954. Formation of Sediments in Recent Basins. SAPOZHNIKOV, Izd. Akad. Nauk S.S.S.R., Moscow, 791 pp. TAFT,W. H., 1961. Authigenic dolomite in modern carbonate sediments along the southern coast of Florida. Science, 134 (3478): 561-562. TAFT,W. H. and HARBAUGH, J. W., 1964. Modern carbonate sediments of southern Florida, Bahamas, and Espiritu Santo Island, Baja California: A comparison of their mineralogy and chemistry. Statlford Univ. Publ., Geol. Sci., X(2): 133 pp. TEODOROVICH, G. I., 1958. Study of Sedimentary Rocks. Gostoptekhizdat, Leningrad, 572 pp. W. H., 1932. Treatise on Sedimentation. Williams and Wilkins, Baltimore, Md., TWENHOFEL, 926 pp. W. H., 1942. The rate of deposition of sediments: a major factor connected with alTWENHOFEL, teration of sediments after deposition. J . Sediment. Petrol., 12: 99-1 10. T. W., 1924. Oceanography in its relations to other earth sciences. J . Wash.Acad. Sci., VAUGHAN, 14: 307-313. VONDER BORCH,C . C., RUBIN,M. and SKINNER, B. J., 1964. Modern dolomite from South Australia. Am. J. Sci., 262: 1116-1 118. ZELENOV, K. K., 1956. Dolomites of Lower Cambrian deposits of northern slope of Aldan Mountain Range and conditions of their formation. In: N. M. STRAKHOV and G. I. BUSHINSKIY (Editors), Types of Dolomite Rocks and their Genesis-Tr. Geol. Inst., Izd. Akad. Nauk S.S.S.R., 4: 28-74.

Chapter 2 MODERN CARBONATE SEDIMENTS WILLIAM H. TAFT

Department of Geology, University of South Florida, Tampa, Fla. (U.S.A.)

SUMMARY

Modern carbonate sediments are accumulating in almost all depositional environments except for the very deep oceans. Chemical and physical conditions, however, govern preservation and abundance of carbonate relative to non-carbonate material. Most modern marine carbonate sediments owe their origin to accumulation of bioclastic debris. Rarely can inorganic precipitation be actually proven. Mineralogical analysis of these sediments has shown that deep cold-water carbonates are predominantly low-magnesium calcites; whereas shallow, warm-water shelf deposits are composed predominantly of metastable carbonates (aragonite and high-magnesium calcite) with minor amounts of low-magnesium calcite. Significant quantities of supratidal dolomite has been reported in modern sediments. Diagenesis is taking place in modern sediments; however, only rarely can mineralogical or chemical changes be demonstrated in sediments that have not been exposed to fresh water. If modern shelf sediments are good examples of ancient sediments, almost all limestones have recrystallized. INTRODUCTION

Study of modern marine carbonate sediments gained momentum after World War I1 as a result of discoveries of huge accumulations of petroleum in limestones and dolostones in the Middle East. As a result of these discoveries, the petroleum industry, recognizing a general lack of basic knowledge in the fields of shallowwater carbonate deposition, diagenesis and lithification, sponsored and encouraged an extensive 10-year investigation of modern carbonate depositional environments. To date, these studies have concentrated on modern marine carbonates of the Bahama Banks, Florida Bay, Gulf of Batabano, the Great Barrier Reef of Australia, the Persian Gulf, and many specific areas of more limited extent. The purpose of this chapter is to broadly summarize the more significant results of these investigations and, hopefully, to encourage additional investigators to enter the study of modern carbonates. For a useful summary of modern carbonates, one may refer to a review by GRAF(1960).

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W. H. TAFT

DISTRIBUTION OF CARBONATE ACCUMULATIONS

Accumulation of carbonate sediments in the oceans is controlled areally and vertically by physical and chemical parameters such as temperature, pH, pressure, and carbon-dioxide content of sea water (REVELLE and FAIRBRIDGE, 1957; RODGERS, 1957). In addition, the abundance of non-carbonate material determines whether or not the final sediment will be a limestone or a clastic sediment with some carbonate material. Modern sediments with more than 30% calcium carbonate tend to be concentrated between 30 “N and 30”slatitude (Fig. 1). Variations from these somewhat arbitrary limits are common, particularly where warm oceanic currents (such as the Gulf Stream) saturated or supersaturated with respect to calcium carbonate are deflected away from the equator. Four general categories of carbonate sediments are recognized: (1)deep-sea oozes; (2) carbonate turbidites; (3) shelf accumulations of lime sands, silts, and muds; and ( 4 ) organic reefs and reef debris.

DEEP-SEA OOZES AND TURBIDITES

Deep-sea oozes are composed predominantly of Globigerina tests with minor quantities of heteropods, pteropods, rhabdoliths and coccoliths. The preservation of these deep-water carbonate sediments (to a depth of approximately 15,000ft.) is controlled more by the solubility of calcium carbonate in deep cold water than by the abundance of pelagic organisms. The so-called “calcium carbonate compensation depth” occurs at a depth of about 4,000-4,500 m. The continual rain of pelagic Foraminifera tests such as Globigerina through the water mass appears to be rather uniform; however, as a result of increased solubility of calcium carbonate in cold waters rich in carbon dioxide these tests are dissolved, leaving only red clays as the dominant sediment. As pointed out by RODGERS (1957, p.3) this gradation from Globigerina ooze (> 50 % calcium carbonate) to red clay (10 i % calcium carbonate) is very abrupt with increasing depth of water. In the East Indian basins, massive slumps have sometimes carried layers of carbonate oozes well below the normal “compensation depth”. Slumping on both large and small scales is rather common in deep-sea areas and often interrupts the regular sequence of layers in core samples. In some places there are gaps; in others, duplication. Unlike typical deep-water oozes that are situated great distances from accumulations of shelf sediments and are composed of fairly distinct assemblages of tests of plants and animals, carbonate turbidites are mixtures of typical deep-sea oozes and shelf sediments. Carbonate turbidites have been described by BUSBY (1962) and by RUSNAKand NESTEROFF (1964). These sediments are mixtures of

w

c

Fig.1, Distribution of modern marine carbonate sediments; solid black : organic reefs; lined: areas of shallow-water carbonate sediments; 1957. Printed by permisstippled: areas of other sediments, containing more than 30 % CaC03 (especially Globigerina ooze). (After RODGERS, sion of McKinley Publishing Co. and Society of Economic Paleontologists and Mineralogists.)

32 W. H. TAFT

v)

e ‘cf

4

L

W W

Fig.3. Pelletal sand from western margin of Andros Island, Bahamas. Material smaller than 0.062 mm in size has been removed.

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W. H. TAFT

planktonic Foraminifera and pteropods in the ungraded sections of cores, and pelagic and reef-derived sediments in the graded sections.

SHELF SEDIMENTS

Carbonate sediments accumulating on continental shelves may be distinguished on the basis o f ( I ) size distribution, (2) amount and composition of organic constituent, and (3) origin. Facies maps based on these criteria have been used successfully to show distribution of modern carbonates in the Gulf of Batabano (DAETWYLER and KIDWELL, 1959), on the Great Bahama Bank (IMBRIE and PURDY, 1962; PURDY,1963b), and on the Heron Island reef (MAXWELL et al., 1964). In addition, carbonate sediments may be distinguished on the basis of shape of particles.

Fig.4. Hand-picked specimens of grapestone from Yellow-Bank, Bahamas. Note progressive increase in degree of cementation and loss of grapestone texture from top to bottom of photograph.


35

ILLING(1954) emphasized the importance of distinguishing between the skeletal and non-skeletal sediments, as did GINSBURG (1956). A common approach has been that of identifying the relative percentages of skeletal debris. ILLING’S (1954) results showed that there is an abundance of non-skeletal debris in the form of lumps, grains, ooids, aggregates and pellets, the most common of which were lumps. Similar deposits (bahamites) were described by BEALES (1 958). Non-skeletal sand particles also include ooliths. Ooliths tend to form local deposits concentrated in areas where water rises from appreciable depths and spreads out over a shallow bank, such as along the west side of the Andros platform in the vicinity of Cat Cay (NEWELL et al., 1959), and between Pleistocene islands along the western edge of Exuma Sound. Most modern carbonates owe their origin to accumulation of bioclastic debris. If particle grain size is much smaller than 50 p, however, it is extremely difficult to see identifiable organic structure or outlines. Moreover, as grain size approaches 10 p, or smaller, the particles are commonly aragonite needles, which, based on shape alone, cannot be distinguished from aragonite needles artificially precipitated from sea water (REVELLE and FLEMING, 1934). The difficulty of determining origin of modern carbonates is not restricted solely to fine particles. Although the size of ooids, grapestones, pellets and composite grains may be as large as 1 mm, they are generally made up of thousands of fine particles ( < 10 p). These particles owe their origin to organic or inorganic precipitation of CaC03 (principally as aragonite). The physical aggregation of these particles produces sand-sized grains. Examples of some typical carbonate sediments include: (I) foraminifera1 sand (Fig.2), (2) pelletal sand (Fig.3), (3) grapestone sand (Fig.4,5), ( 4 ) skeletal calcareous sand (Fig.6), and (5) sand composed of ooliths (Fig.7).

CHEMICAL A N D MINERALOGICAL COMPOSITION OF CARBONATE SEDIMENTS

Chemical and mineralogical studies of modern shallow-water carbonate sediments (1960, 1962), SIEGEL(1961), STEHLI and HOWER made among others by CHILINGAR (1961), CLOUD(1962), and TAFTand HARBAUGH (1964), demonstrate that these sediments are composed predominantly of metastable carbonate minerals (aragonite and high-magnesium calcite) with minor amounts of stable low-magnesium calcite. Most of the magnesium present in these sediments is substituted for calcium in the calcite structure by solid solution, and occurs in the calcareous skeletons of Algae, Foraminifera, Porifera, Coelenterata, sea urchins, starfishes, Ophiuroidea, Crinoidea, Brachiopoda, Mollusca, Ostracoda, and Cirripedia. Very high Ca/Mg ratios of the skeletons of madreporarian corals, pelecypods, gastropods and cephalopods are due to the fact that aragonitic organisms contain very small amounts of magnesium (seldom over 1 % MgC03).

Fig.5. Grapestone and bioclastic particles from Yellow Bank, Bahamas.

W

-4

Fig.6. Sand-sized bioclastic sediment from Florida Bay, Florida. Note angular nature of carbonate fragments and general absence of whole particles.

w

00

Fig.7. Ooliths from Wide Opening, Bahamas (west side of Euxuma Sound).


39

Detailed chemical and mineralogical studies of the calcareous skeletons of marine organisms (CHAVE,1954a,b, 1962a, b; LOWENSTAM, 1954a, b, 1961, 1963, 1964; KRINSLEY, 1959, 1960a, b; KRINSLEY and BIERI,1959; PILKEYand HOWER, 1960; PILKEYand GOODELL, 1963) show a marked dependence of chemistry on the polymorphic form of calcium carbonate precipitated. For example, the two polymorphs of calcium carbonate, aragonite and calcite, have an inherent preference for certain cations. Strontium substitutes readily for calcium in aragonite up to 4.08 % SrO (NOLL,1934), whereas SrO rarely exceeds 1.O % in calcite. Magnesium, on the other hand, is most commonly found in the calcite structure (up to 30 mol. % M ~ C O ~ , C H A V 1954a); E , little magnesium is found in aragonite. LOWENSTAM (1954b) demonstrated the role of temperature in controlling the polymorphic form of calcium carbonate precipitated in carbonate skeletons. Four animal phyla (Coelenterata, Bryozoa, Annelida, and Mollusca) show a relationship between the calcite/aragonite ratio of their skeletal structures and the temperature of the water in which they are secreted; aragonite skeletons suggest warm water, whereas calcite skeletons-cold water. Because of the effect of temperature on the polymorphic form of calcium carbonate secreted by organisms, this effect also controls the strontiumlcalcium ratios in these organisms. This is principally due to the fact that strontium is concentrated in aragonite skeletons and calcium and magnesium in calcitic structures. DODD(1963) has shown a lack of dependence of Sr content on temperature for young individuals of the pelecypod Mytilus culiforniunus, which exhibits differences in physiology at different growth stages. LOWENSTAM (1964) in a study of the strontium/calcium ratio of the marine biota from Palau, South Pacific, where temperature and strontium/calcium ratio of the waters are constant throughout the year, demonstrated that crystal chemistry plays the primary role in controlling the strontium/calcium ratio. He also showed that the physiologic control is a second order effect controlling the strontium/calcium ratio. CHAVEet al. (1962) demonstrated solubility relationships between aragonite, high-magnesium calcite, and low-magnesium calcite. They determined the following order of decreasing solubility: high-magnesium calcite > aragonite > low-magnesium calcite. Chave and co-workers suggested that because of these solubility relationships high-magnesium calcite and aragonite may be dissolving in the water mass or on the sea floor in many areas, leaving only the more stable polymorph (low-magnesium calcite) as the predominant mineral. JANSEN and KITANO (1963), however, reported a decrease in the solubility of magnesium carbonate in MgC12 solutions. PILKEY and GOODELL (1964) have emphasized the importance of using caution, particularly with reference to recrystallization, in attempting to utilize chemical composition of carbonate skeletons as geochemical and paleoecologic indices. They have stressed the importance of ion migration and stability of ions in skeletons of carbonate-secreting organisms. LOWENSTAM (1963) reviewed the chemical composition of sediments and its relationship to the mineralogic com-

40

W. H. TAFT

position. He confirmed that temperature and chemistry of sea water in which the organisms live affect the aragonite/calcite ratio of the skeletons secreted, and in turn the magnesium and strontium contents as well as the l 8 0 / 1 6 0 ratio of these carbonates. (For details see WOLFet al., 1967.) One of the most significant results of modern carbonate studies in the last 10 years has been the discovery that dolomite in significant quantities is forming in modern supratidal carbonate environments. These supratidal environments are closely associated with evaporitic environments, and should help to further our understanding of ancient deposits that may have formed in a similar manner. Reports of dolomite formation in modern sediments include those by ALDERMAN (1959), ZEN (1959), MILLER (1961), TAFT (1961), DEFFEYES and MARTIN(1962), WELLS(1962), CURTISet al. (1963), SKINNER (1963), ILLING(1964), LUCIAet al. (1964), and SHINNand GINSBURG (1964).

RELATIONSHIP BETWEEN MINERAL AND CHEMICAL COMPOSITION, AND SEDIMENT GRAIN SIZE

CHAVE(1954b) reviewed chemical analyses of modern carbonates by VAUGHAN

(19 18), BRAMLETTE (1 926), and GOLDMAN (1926); he also recalculated and deter-

mined the chemical composition of carbonate sediments from Australia, the Bahamas, and Pago-Pago. The techniques used by all writers were quite similar. The sediments were separated into biological groups under microscope, weighed, and chemical composition assigned to them from data by CLARKE and WHEELER (1922). The chemical composition thus calculated was then compared with the actual chemical analysis. The results of analyses by Bramlette and Vaughan show fairly close agreement between calculated and determined chemistry, whereas Goldman’s results suggest a loss of magnesium to sea water, especially in the finersize fractions. To test Goldman’s results, CHAVE(1954b, p.596) compared the chemistry of eight mechanically analyzed carbonate bottom samples collected from Alaska, Maine, Florida, Bermuda, and Palau. Based on these analyses, Chave concluded that there was no evidence for exchange of magnesium ions between the sediment and sea water for the samples tested. TAFTand HARBAUGH (1964) reported (1) chemical data for size-sorted sediment samples from Key West, Florida Bay, and the Ten Thousand Island areas in Florida, (2) chemical and mineralogical data for sized samples from the Bahamas, and (3) mineralogical data for two samples from Baja California, Mexico. These samples were analyzed to test for gain or loss of magnesium relative to calcium and variations in proportions of high-magnesium calcite in different grade sizes. Any systematic change of chemistry with grain size might be useful in explaining variations of calcium/magnesium ratios as a function of sediment sorting, or susceptibilityof metastable carbonates to recrystallize to the stable mineral calcite.


41

Results of these analyses suggest a progressive decrease of the calcium/magnesium ratio with decreasing grain size in Florida sediments. Samples from the Bahamas show that the coarse fraction (> 0.062 mm) tends to have the highest proportion of high-magnesium calcite and the lowest Ca/Mg ratio, whereas the finest fraction ( < 0.001 mm) has the highest percentage of aragonite and the lowest percentage of high-magnesium calcite, and tends also to have the highest Ca/Mg ratio. Baja California samples, on the other hand, show that the proportion of high-magnesium calcite tends generally to increase as the grain size decreases. These analyses suggest that local conditions of biogenic productivity and susceptibility to mechanical breakdown play a more important role in controlling chemistry and mineralogy of fine-grained carbonate particles than do loss or gain of magnesium after burial of the sediments.

ISOTOPIC COMPOSITION

The relative abundance in sea water of the two stable isotopes of carbon (13C and W)and two of the three stable isotopes of oxygen (lSOand160) has been shown to be fairly constant in open-ocean environments. Nearshore environments, on the other hand, ai-e known to be deficient in I3C relative to I2C and in 1 8 0 relative to 1 6 0 , as compared to mean ocean water. These differences are attributed to: ( I ) higher degree of volatility of 12C and 1 6 0 during evaporation, and (2) dilution by freshwater runoff deficient in 1 8 0 and enriched in IZC. UREYet al. (1951) concluded that l 8 0 / 1 6 0 ratios of carbonates are temperature dependent and are independent of species biochemistry or carbonate mineralogy. Marine carbonates precipitated in equilibrium with the surrounding sea water, therefore, should have carbon and oxygen isotope ratios diagnostic of their depositional environments. Unfortunately, like most other geologic tools, a number of problems arise on attempting to use oxygen isotope ratios, which include the following: ( I ) Although it is fairly well established that carbonates are precipitated by organisms in equilibrium with the surrounding sea water, one must establish that neither recrystallization nor diffusion has altered initial isotope ratios. (2) Temperature at which calcium carbonate is precipitated by carbonatesecreting organisms may be appreciably different from the average yearly temperature of the depositional environment. The precise growth-time temperature is significant. EPSTEIN and LOWENSTAM (1953) demonstrated substantial oxygenisotope ratio variations in different carbonate growth increments of some species of Bermuda mollusks. Some species precipitate CaC03 year-round, whereas others precipitate CaC03 selectively during certain periods of the year. (3) Tolerance of organisms to unfavorable environmental changes, which may either force an animal to migrate to a more favorable environment or, if not

42

W. H. TAFT

mobile, may cause its death, must be considered. Obviously, the mere presence of an organism in sediments is no indication that isotopic ratios of its shell provide absolute indication of temperature conditions of the depositional environment. Organisms, permanently attached to the sea bottom and possessing only slight tolerances to changes in chemical and physical environment, provide ideal working material for determination of isotopic composition in order to arrive at the temperature of depositional environment. Although there are a number of problems associated with interpreting stable isotope ratios of carbonate skeletons, nevertheless this technique is in its infancy and already a number of worth-while applications have been found. For example, LOWENSTAM and EPSTEIN (1957) studied the origin of aragonite needles on the Great Bahama Bank and concluded that the isotopic composition of these needles would best be explained as having originated by organic precipitation. This interpretation, however, was challenged by CLOUD(1962, p.97). GROSS(1 964) compared ls0/16O and I 3 C / l 2 C ratios of diagenetically altered limestone in the Bermuda islands with those of modern sediments in the surrounding marine environment, and concluded that alteration occurred in the freshwater zone. TAFTand HARBAUGH (1964) correlated the marl underlying brackishwater sediments in Whitewater Bay, Florida, with that presently accumulating in the Everglade swamps on the basis of their mineralogy and enrichment in 13C. Isotope ratios are being used as indicators of paleotemperatures (EMILIANI and MAYEDA, 1961) and although some controversy surrounds interpretation of analytical results, judicious application of stable isotopes appears to be a powerful tool for future use.

DIAGENESIS OF CARBONATE SEDIMENTS

Physical and chemical changes are taking place in modern carbonate depositional environments. Diagenesis of carbonates is generally more rapid but may continue longer than is the case with most other sediment types. GINSBURG (1957) called attention to some of these processes in Florida Bay sediments. He noted the following: ( I ) Presence of abundant fecal pellets at the water-sediment interface, but loss of their individuality a few feet below the surface. (2) Importance of organisms in reducing particle size by chemical and physical processes. (3) Reworking of sediments by organisms that bring buried sediments to the surface and obliterate depositional features. (4) Activity of Bacteria in producing hydrogen sulfide, oxidation of organic matter, conversion of organic nitrogen to ammonia, and precipitation of iron. (5) Marked contrasts in p H between the sediments and the overlying sea


43

Fig.8. Underwater photograph of a 36-inch long vertical exposure of modern carbonates on Yellow Bank, Bahamas. Overlying water is 17 ft. deep. Note large blocks in cliff face and near base of photograph. Ruler rests on apparent horizontal layer 36 inches below water-sediment interface (used for 14C dating). Probe is 4 ft. long and extends 3 ft. into the underlying sediments. Tube is a device for removing fines during excavation.

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water are common. In general, sediment p H is at least one full p H unit lower than that of sea water. (6) Compaction, as measured by loss of moisture content, is most prevalent in the first 6-12 inches below the water-sediment interface. (7) No conversion of metastable carbonates to stable forms is evident; lithification is also not noticeable, except in the intertidal zones where alternate wetting and drying allows beach sands to be cemented by aragonite. An excellent petrographic study of modern carbonates by PURDY(1963a,b) demonstrates recrystallization of metastable carbonates on the Great Bahama bank. Recrystallization, however, does not give rise to a more stable mineral species, but rather results in a change from a micro- to a cryptocrystalline variety of the same metastable mineral. According to PURDY(1963a,p.348), “there can be no doubt that cryptocrystalline grains are a recrystallization product, for all stages were observed in the transition of various non-skeletal and skeletal constituents to cryptocrystalline carbonate”. Since GINSBURG’S (1957) excellent paper, a number of papers bearing on the problem of diagenesis in modern carbonate sediments have appeared (SIEGEL, 1960, 1961; STEHLIand HOWER,1961; CLOUD,1962; INCERSON 1962; PURDY, 1963a,b; PILKEY, 1964; TAFTand HARBAUGH, 1964). In general, these authors have shown that modern, warm-, and shallow-water carbonate sediments are composed predominately of aragonite and high-magnesium calcite. There is little or no evidence to suggest either recrystallization to more stable polymorphs since particles were deposited, or lithification of sediments that have never been exposed to fresh water, other than infilling of tests by aragonite and formation of grapestone lumps (ILLING,1954) and flakes (PURDY,1963a, p.348). A notable exception to lack of lithification in modern carbonates is presently being investigated on Yellow Bank, Bahamas, by the author. Fig.8 is a 36-inch vertical exposure of modern carbonates on Yellow Bank in 17 ft. of water. Large blocks (up to 1 cubic ft.) of partially lithified, nonstratified carbonate material underlie a thin veneer of grapestone sediment. These blocks are held together by worm tubes, encrusting Foraminifera such as Homotrema rubrum (LAMARK), and by aragonite (Fig.9). Comparison of mineralogic composition of surface blocks (Table I) and unconsolidated surface samples with the horizontal layer 36 inches below the watersediment interface, however, suggests that these samples have not been exposed to fresh water and have not recrystallized to more stable carbonate minerals.

PROBLEMS

At least three problems continue to plague investigators of modern carbonates:


45

M

3

.-

f

C L

TABLE I CARBONATE MINERAL SAMPLES AND RADIOCARBON AGE DETERMINATIONS FROM YELLOW BANK, BAHAMAS

Sample number

64-81 64-81 64-81 64-81 64-81 64-8 1 64-8 1 64-8 1

Depth beneath

Mineral fractions (%)

sediment-water interface (inches)

aragonite

IoW-Mg calcite

high-Mg calcite

surface 1 4 10

61 67 64 65 86 19 94 98

trace 1 trace 1 trace 1 trace

33 32 35 34

20

23 32 36

1

14 20 6 1

Mol. % Mgco3 in high-Mg calcite

Type of sample

15 17 15 16 15 17 14 13

loose sediment rock rock rock loose sediment rock rock rock

Age (years before present)

2,411 & 120

4,810 i 130


47

( 1 ) There is no known continuous shallow-water sediment record from Recent through the Pleistocene to Pliocene sediments. (2) Widespread recrystallization of unexposed Holocene metastable carbonates to more stable forms has not been found. (3) There is a general absence of lithification in modern (Holocene) carbonate sediments. It seems improbable that a continuous record of shallow-water carbonates, uninterrupted by exposure during the Pleistocene, will ever be located. Owing to eustatic oscillations, the neritic zone boundaries have migrated 50-500 km across continental shelves and back again, resulting in a widespread erosion and displacement of facies. Lack of a continuous record of shallow-water carbonates possibly could account for the apparent absence of widespread lithification of modern carbonate sediments. It is puzzling that depositional environments with 8-10 ft. of modern sediments and 6-10 ft. of overlying water show no evidence of cementation. Inasmuch as most modern carbonates are less than 6,000 years old, possibly this is an insufficient time for the process of cementation to take place. Degree of recrystallization of metastable carbonates to stable forms is probably a function of the magnesium concentration in interstitial waters (TAFT, 1967), and time necessary for solid state recrystallization.

ACKNOWLEDGEMENTS

The lithification study on Yellow Bank, Bahamas, was sponsored by Grant No GP 2527 from the National Science Foundation. (Fig.2-6 are reprinted from 1964, with the permission of the publishers, Stanford TAFTand HARBAUGH, University School of Earth Sciences, 0 1964 by the Board of Trustees of the Leland Stanford Junior University.)

REFERENCES

ALDERMAN, A. R., 1959. Aspects of carbonate sedimentation. J. Geol. SOC.Australia, 6: 1-10. F. W., 1958. Ancient sediments of Bahaman type. Bull. Am. Assoc. Petrol. Geologists, BEALES, 42: 1845-1880. BRAMLETTE, M. N., 1926. Some marine bottom samples fromPago Pago Harbor, Samoa. Carnegie Inst. Wash., Papers Dept. Marine Biol.,23: 1-35. BUSBY,R. F., 1962. Submarine geology of the Tongue of the Ocean, Bahamas. U.S. Naval Oceanog. Ofice, Tech. Rept., 108: 1-84. CHAVE, K. E., 1954a. Aspects of biogeochemistry of magnesium. 1. Calcareous marine organisms. J. Geol., 62: 266-283. K. E., 1954b. Aspects of the biochemistry of magnesium. 2. Calcareous sediments and CHAVE, rocks. J. Geol., 62: 587-599.

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CHAVE,K. E., 1962a. Factors influencing the mineralogy of carbonate sediments. Limnol. Oceanog., 7: 218-223. CHAVE, K. E., 1962b. Processes of carbonate sedimentation. In: N. MARSHALL (Editor), Symposium on the Environmental Chemistry of Marine Sediments. Narragansett Marine Lab., Kingston, R. I., pp.77-85. K. S., WEYL,P. K., GARRELS, R. M. and THOMPSON, M. E., 1962. CHAVE,K. E., DEFFEYES, Observation on the solubility of skeletal carbonates in aqueous solutions. Science, 137: 33-34. CHILINGAR, G. V., 1960. Ca/Mg ratio of calcareous sediments as a function of depth and distance from shore. Compass, 37(3): 182-186. CHILINGAR, G. V., 1962. Dependence on temperature of Ca/Mg ratio of skeletal structures of organisms and direct chemical precipitates out of sea water. Bull. Southern Calif: Acad. Sci., 61(1): 45-60. CLARKE,F. W. and WHEELER, W. C., 1922. The inorganic constituents of marine invertebrates. U.S., Geol. Surv., Profess. Papers, 124: 1-62. CLOUD,JR., P. E., 1962. Environment of calcium carbonate deposition west of Andros Island, Bahamas. U S . , Geol. Surv., Profess. Papers, 350: 1-138. CURTIS,R., EVANS,G., KINSMAN, D. J. J. and SHEARMAN, D. J., 1963. Association of dolomite and anhydrite in the Recent sediments of the Persian Gulf. Nature, 197: 679-680. DAETWYLER, C. C. and KIDWELL, A. L., 1959. The Gulf of Batabano, a modern carbonate basin. World Petrol. Congr., Proc., 5th, N. Y., 1959, Sect.& pp.1-21. DEFFEYES, K. S. and MARTIN,E. L., 1962. Absence of carbon-14 activity in dolomite from Florida Bay. Science, 136: 782. DODD,J. R., 1963. Paleoecological implications of shell mineralogy in two pelecypod species. J. Geol., 71: 1-11. EMILIANI, C. and MAYEDA,T., 1961. Carbonate and oxygen isotopic analysis of core 241 A. J. Geol., 69: 729-732. EPSTEIN,S. and LOWENSTAM, H. A., 1953. Temperature-shell growth relations of recent and interglacial Pleistocene shoal water biota from Bermuda. J. Geol., 61 : 424438. GINSBURG, R. N., 1956.Environmental relationships of grain size and constituent particles in some south Florida carbonate sediments. Bull. Am. Assoc. Pelrol. Geologists, 40: 2384-2427. GINSBURG, R. N., 1957. Early diagenesis and lithification of shallow-water carbonate sediments in southern Florida. In: R. J. Le BLANCand J. G. BREEDING (Editors), Regional Aspects of Carbonate Deposition-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 5: 80-100. GOLDMAN, M. I., 1926. Proportions of detrital organic calcareous constituents and their chemical alteration in a reef sand from the Bahamas. Carnegie Znst. Wash. Publ., 344: 39-65. GRAF,D. L., 1960. Geochemistry of carbonate sediments and sedimentary carbonate rocks, Part I, carbonate mineralogy, carbonate sediments. Illinois State Geol. Surv., Circ., 297: 1-39. GROSS,M. G., 1964. Variations in the 180/160 and 13C/1zCratios of diagenetically altered limestones in the Bermuda Islands. J. Geol., 72: 170-194. ILLING,L. V., 1954. Bahaman calcareous sands. Bull. Am. Assoc. Petrol. Geologists, 38: 1-95. ILLING,L. V., 1964. Penecontemporary dolomite in the Persian Gulf. Bull. Am. Assoc. Petrol. Geologists, 48: 532-533. IMBRIE,J. and PURDY, E. G., 1962. Classification of modern Bahaman carbonate sediments. In: Classification of Carbonate Rocks-Am. Assoc. Petrol. Geologists, Mem., I : 253-272. INGERSON, E., 1962. Problems of the geochemistry of sedimentary carbonate rocks. Geochim. Cosmochim. Acta., 26: 815-847. JANSEN, J. F. and KITANO,Y.,1963. The resistance of Recent marine carbonate sedinients to solution. J . Oceanog. Soc. Japan, 18: 42-52. KRINSLEY, D., 1959. Manganese in modern and fossil gastropod shells. Nature, 183: 770-772. KRINSLEY, D., 1960a. Magnesium, strontium and aragonite in the shells of certain littoral gastropods. J. Paleontol., 34: 744-755. KRINSLEY, D., 1960b. Trace elements in the tests of planktonic foraminifera. Micropaleontology, 6: 297-300. KRINSLEY, D. and BIERI,R., 1959. Changes in the chemical composition of pteropod shells after deposition on the sea floor. J. Paleontol., 33: 682-684.


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LOWENSTAM, H. A., 1954a. Environmental relations of modification compositions of certain carbonate secreting marine invertebrates. Proc. Natl. Acad. Sci. US.,40: 3948. LOWENSTAM, H. A., 1954b. Factors affecting the aragonite/calcite ratios in carbonate-secreting organisms. J . Geol., 62: 284-322. LOWENSTAM, H. A., 1961. Mineralogy, 180/160 ratios, and strontium and magnesium contents of recent and fossil brachiopods and their bearing on the history of the oceans. J. Geol., 69: 241-260. LOWENSTAM, H. A., 1963. Biologic problems relating to the composition and diagenesis of sediments. In: T. W. D ~ N N E L(Editor), LY The Earth Sciences: Problems and Progress in Current Research. Univ. Chicago Press, Chicago, Ill., pp.137-195. LOWENSTAM, H. A., 1964. Sr/Ca ratio of skeletal aragonites from the Recent marine biota at Palau and from fossil gastropods. In: H. CRAIG,S. L. MILLERand G. J. WASSERBURG (Editors), Zsotopic and Cosmic Chemistry. North-Holland Publ. Co., Amsterdam, pp.114132. LOWENSTAM, H. A. and EPSTEIN,S., 1957. On the origin of sedimentary aragonite needles of the Great Bahama Bank. J. Geol., 65: 364-375. LUCIA,F. J., WEYL,P. K. and DEFFEYES, K. S., 1964. Dolomitization of Recent and Plio-Pleistocene sediments by marine evaporite waters on Bonaire, Netherlands Antilles. Bull. Am. Assoc. Petrol. Geologists, 48: 535-536. MAXWELL, W. G. H., JELL,J. S. and MCKELLAR, R. G., 1964. Differentiation of carbonate sediments in the Heron Island Reef. J. Sediment. Petrol., 34: 294308. MILLERJR., D. N., 1961. Early diagenetic dolomite associated with salt extraction process, Inagua, Bahamas. J. Sediment. Petrol., 31 :473476. NEWELL, N. D., IMBRIE, J., PURDY, E. G. and THURBER, D. L., 1959. Organism communities and bottom facies. Great Bahama Bank. Bull. Am. Museum Nut. HiJt., 97: 58-69. NOLL,W., 1934. Geochemie des Strontium, mit Bemerkungen zur Geochemie des Barium. Chem. Erde, 8: 507-560. PILKEY,0. H., 1964. Mineralogy of the fine fraction in certain carbonate cores. Bull. Marine Sci. Gulf Caribbean, 14: 126-139. PILKEY,0. H. and GOODELL, H. G., 1963. Trace elements in Recent mollusk shells. Limnol. Oceanog., 8: 137-148. PILKEY,0. H. and GOODELL, H. G., 1964. Comparison of the composition of fossil and Recent mollusk shells. Bull. Geol. Soc. Am., 75: 217-228. PILKEY,0. H. and HOWER,J., 1960. The effect of environment on the concentration of skeletal magnesium and strontium in Dendraster. J. Geol., 68: 203-214. PURDY,E. G., 1963a. Recent calcium carbonate facies of the Great Bahama Bank. 1. Petrography and reaction groups. J. Geol., 71: 334-355. PURDY,E. G., 1963b. Recent calcium carbonate facies of the Great Bahama Bank. 2. Sedimentary facies. J. Geol., 71: 472497. REVELLE, R. and FAIRBRIDGE, R. W., 1957. Carbonates and carbon dioxide. In: J. W. HEDGPETH (Editor), Treatise on Marine Ecology and Paleoecology. I . Ecology-Geol. SOC.Am., Mem., 67: 239-295. REVELLE, R. and FLEMING, R. F., 1934. The solubility product constant of calcium carbonate in sea water. Proc. Pacific Sci. Congr. Pacific Sci. Assoc., 5th, Victoria, Vancouver, 1933, 3: 2089-2092. RODGERS, J., 1957. The distribution of marine carbonate sediments; a review. In: R. J. LE BLANC and J. G. BREEDING (Editors), Regional Aspects of Carbonate Deposition-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 5 : 2-1 3. RUSNAK, G. A. and NESTEROFF, W. D., 1964. Modern turbidites: terrigenous abyssal plain versus bioclastic basin. In: R. L. MILLER(Editor), Papers in Marine Geology-Shepard Commemorative Volume. Macmillan, New York, N.Y., pp.488-507. SHINN,E. A. and GINSBURG, R. N., 1964. Formation of Recent dolomite in Florida and the Bahamas. Bull. Am. Assoc. Petrol. Geologists, 48: 547. SIEGEL,F. R., 1960. The effect of strontium in the aragonitexalcite ratios of Pleistocene corals. J. Sediment. Petrol., 30: 297-304.

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SIEGEL, F. R., 1961. Variations of Sr/Ca ratios and Mg contents in Recent carbonate sediments of the northern Florida Keys. J. Sediment. Petrol., 31: 336-342. SKINNER, H. C. W., 1963. Precipitation of calcian dolomites and niagnesian calcites in the southeast of south Australia. Am. J. Sci., 162.: 449472. STEHLI,F. G. and HOWER, J., 1961. Mineralogy and early diagenesis of carbonate sediments. J. Sediment. Petrol,, 31; 358-371. TAFT,W. H., 1961. Authigenic dolomite in modem carbonate sediments along the southern coast of Florida. Science, 134: 561-562. TAFT,W. H., 1967. Physical chemistry of formation of carbonates. In: G. V. CHILINGAR, H. J. BIssELLand R. W. F A I R B R I D G E ( E ~Carbonate ~ ~ ~ ~ ~ )Rocks. , Elsevier, Amsterdam, B: 151-167. TAFT,W. H. and HARBAUGH, J. W., 1964. Modern carbonate sediments of southern Florida, Bahamas, and Espiritu Santo Island, Baja California: a comparison of their mineralogy and chemistry. Stanford Univ. Publ., Univ. Ser., Geol. Sci., 8(2): 1-133. UREY,H. C., LOWENSTAM, H.A., EPSTEIN,S. and MCKINNEY,C. R., 1951. Measurement of paleotemperatures and temperatures of the Upper Cretaceous of England, Denmark and the southeastern United States. Bull. Geol. Soc. Am., 62: 399416. VAUGHAN, T. W., 1918. Some shoal-water bottom samples from Murray Island, Australia, and comparisons of them with samples from Florida and the Bahamas. Carnegie Znst. Wash. Publ., 213: 235-288. WELLS,A. J., 1962. Recent dolomite in the Persian Gulf. Nature, 194: 274-275. WOLF,K. H., CHILINGAR, G. V. and BEALES, F. W., 1967. Elemental composition of carbonate H. J. BISSELL and R. W. FAIRskeletons, minerals and sediments. In: G. V. CHILINGAR, BRIDGE (Editors), Carbonate Rocks. Elsevier, Amsterdam, B: 23-149. ZFN, E-AN, 1959. Mineralogy and petrography of marine bottom sediment samples o f fthe coast of Peru and Chile. J . Sediment. Petrol., 29: 513-539.

Chapter 3

PETROLOGY AND PETROGRAPHY OF CARBONATE ROCKS YVONNE GUBLERl, J. P. BERTRANDI, L. MATTAVELL12, A. RIZZINI’ AND R. PASSEGA3

French Petroleum Institute, Rueil-Malmaison (France) Geologist, AGIP, Direzione Mineraria, San Donato, Milan (Italy) Consultant, San Donato, Milan (Italy)

SUMMARY

This paper is basically a synthesis and literature review of the major features of petrologic and petrographic knowledge on carbonate rocks. The authors have devoted special attention to the construction of a petrologic model based on the concept of “facies” as an integration of qualitative and quantitative data situated at different scalar levels.

INTRODUCTION

General

Sedimentary petrology and petrography are two divisions of lithostratigraphy that cover this fundamental branch of stratigraphy. Petrology commonly deals with a rock as a geometric body of varying constitution, modified by sedimentological features such as stratification, orientation, alteration (diagenesis) and resedimentation. Petrography deals with detailed particulars, e.g., constituents and texture. Both disciplines proceed essentially from descriptive science, but are situated at different scalar levels of organization of the rock (GLANGEAUD, 1962). Thus, sedimentary petrology is based on numerous field investigations, regional and local studies, that lead to a more scientific and useful synthesis during petro1961), that is to say, in the laboratory, graphic studies (BISSELLand CHILINGAR, largely by means of the petrographic microscope, and other aids which enable detailed grain-by-grain examination. The final goal is to attempt to construct a petrologic description and explanation from the progressive differentiation of the relationships of all the data collected; the validity of such a petrographic model depends on the techniques employed and the quantitative significance of the data. The usefulness of multidimensional or stereographic diagrams has been pointed out (KRUMBEIN and SLOSS,1951) in drawing up multiple-paleogeographic

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et al.

isopach maps at different scales, as an approach to lithologic models. Most of these maps take into account only a limited number of parameters. As new physical methods are developed, other than the optical ones ordinarily usedin routine work, more and more variables are recorded and quantified; this introduces more selective parameters for identifying a rock. The petrologic model must be worked out in grouping all sets of data, for both field and laboratory investigations. In the case of carbonate rocks, IMBRIEand PURDY(1961) have proposed a simple statistical scheme for identifying discrete sample groupings in a set of data that may be gathered partly in the field and partly in the laboratory. This information is treated by representative vectors; equal weight is given to various properties. Carbonate rocks

Carbonate rocks are a polygenetic group, some of which are true clastics, whereas others are chemical or biochemical precipitates. Most of them are a mixture of clastic and chemical carbonates whose interrelationship is an important parameter of their diagenesis. The major mineral components are calcite, aragonite, dolomite and, less commonly, siderite; the minor constituents are non-carbonate and include clays, quartz, and phosphate, oxide and sulfide minerals. The three prominent and specific features of the group comprise: ( I ) organic content, which can be considered as primary genetic; (2) the crystallinity, either crypto- or diagenetic (SHVETSOV, 1958), or other textural forms of the carbonate particles such as pellets or oolites; and (3) geographic distribution in the past. (I) In regard to the primary petrogenesis, either in marine or lacustrine environments, organisms are providers of carbonates, because they are able to modify, directly or indirectly, the partial pressure of carbon dioxide. Algae, by photosynthesis, are the most important community associated with carbonate precipitation (WOLF,1962). Thus the relationships between organic populations and their distribution (paleoecology) have to be considered as an important set of data in the petrologic model. The organisms fix Ca2+ in the shell, and contribute to the elaboration of stromatolitic or other structures, or they may be represented by transported material (NEWELL and RIGBY,1957; TAFT,1962). (2) Because of the metastability of its crystalline structure, high-magnesium calcite is converted into stable dolomite, and aragonite is transformed into calcite. These transformations can take place very shortly after the sediments are buried. The original structures of carbonate rocks, both mineralogical and organic, thus often become obscure. How, when and where this diagenesis takes place is one of the most difficult problems of the petrology of carbonate rocks. Most dolomites seem to have originated this way (FAIRBRIDGE, 1957; DE CHARPAL et al., 1959). The problem of diagenesis involves the crystalline “sparite” which is considered to be depositional by FOLK(1962), and to be the product of recrystalli-

PETROLOGY AND PETROGRAPHY OF CARBONATE ROCKS

53

zation by BATHURST (1959) and POWERS (1962). Time of crystallization is largely dependent upon the movements of fluid through the rock. (3) The most common carbonate rocks in marine and lacustrine deposits are limestones and dolomites which constitute about 14 % of the sedimentary rocks in the geological column (KUENEN,1941). The problem of their distribution in Recent sediments has been outlined by RODGERS(1957) and in the ancient sediments by FAIRBRIDGE (1961, 1964). In Recent sediments, carbonate oozes are located on the bottom of oceans at depths of less than 3,000 fathoms at low latitudes. Clayey carbonates are known to exist in deeper water or close to coasts in places where amounts of terrigenous deposits are high (ARRHENIUS, 1952). Organic reefs between N 30” and S 30” latitude constitute large carbonate-shelf deposits associated with communities of colonial organisms. Through geologic time the distribution of carbonate rocks has been in large measure dependent upon the evolution of lime-fixing organisms. In the Early Paleozoic they are most often associated with benthonic shelly fossils, characteristic of rather shallow water; whereas in the Cretaceous, pelagic oozes characteristic of greater depths appear with the explosive evolution of pelagic Foraminifera. Many hundreds of publications deal with problems of carbonates, either in Recent sediments or ancient deposits, generally treating one or another of the afore-mentioned three problems. Very few deal with the whole subject. Mostly because “the world over they are familiar host rocks for oil, water and ore bodies, providers of economy” (CLOUD,1962), special attention is devoted to reef environments. The latter contain 90 % calcarenite (NEWELLand RIGBY,1957), which originally had many of the same characteristics as quartzose sandstones, i.e., high porosity and varying degrees of sorting. A brief report such as this chapter on the petrography and petrology of carbonate rocks can be expected to outline only the general fields of research with emphasis on recent work and current ideas, as well as to mention some of the important unsolved problems and to suggest promising directions for the solution of some of them. OBJECTIVES

Establishment of a lithological record As already seen, sedimentary petrology and petrography are complementary. This was the conclusion of CAYEUX(1935, p. 111) in his memoir on “Carbonate Rocks” when he wrote: “. . .j’ai fix6 leurs caract&resminkralogiques, organiques et chimiques, en quoi j’ai fait oeuvre de pktrographe, j’ai donnk A l’analyse de plusieurs roches un diveloppement qui ne cadre nullement avec le r6le qu’elles jouent dans la nature. . ”,

.

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Indeed, the integration of the many megascopic variables contributing to the development of a carbonate rock is part of the original account of the concept of facies expressed in these terms as early as 1838, by Gressly (in: WEGMANN, 1963) with regard to the carbonate rocks of the Swiss Jura: . . . “Le faciks . . . concerne I’aspect pktrographique et l’aspect palBontologique d’une formation . . Une description pttrographique qui traitera de la composition mineralogique et des caractkres pktrographiques de nos roches tels que la structure, la cassure, les couleurs, le ciment, la piite, etc. . . Ces caractkres . . . n’en sont pas moins d’une haute importance dans la dCtermination des divers faciks que prksentent une formation entikre, un groupe, un terrain, ou m i k e une simple couche. . . Une description gtognostique qui traitera des phknombnes de nos roches dans leur ensemble, tels que leur stratification, leur puissance, etc. . . Ces phknomknes indiquent d’une manihe B peu prks constante les divers faciks. La palBontologie nous offrira les principaux caractkres, les grandes divisions des faciks”. The first description deals more with granular feature, the second with the gross characteristics of the rocks that ought to be precisely described and grouped for the purpose of constructing the petrologic model. Through Gressly’s definitions the term “facies” is used on different scales, including in particular the lithofacies.

Lithofacies Lithofacies may be defined in terms of the extent of variations of micro- and mega-characteristics of the particular features introduced during the origin, and other properties of the rock considered as a whole. It reflects a succession of homogeneous or heterogeneous primary deposits (Fig.1) modified by diagenesis during the compaction stage and very often also afterwards. Dirnensional fabrics Microfacies Each primary deposit is the product of the hydrochemical, hydrodynamic and thermodynamic factors and organic materials of the particular environment; these factors and components also control the speed of deposition. A “primary microfacies” can be defined as consisting of laminae, thickness of which is no greater than the largest diameter of the coarsest grain (DOEGLAS, 1962). It seems necessary to point out the geometrical characteristics because the term “microfacies” is used and defined in different ways: ( I ) For the purpose of stratigraphic dating by correlation of the qualitative elements observed in thin sections, paying special attention to the organisms (CWILLIER,1951, 1952; FAIRBRIDGE, 1954). (2) From a genetic point of view, CAROZZI (1950, 1958,1959, 1961) proposed 1

“Lithofacies” and “biofacies” (Editors).


55

this term to express the vertical variations of minute elements (minerals or microorganisms) that reflect successions of primary microfacies. For the petrologic model, the term “primary microfacies” must be reserved for the observations made in sections parallel to the plane of deposition, whether it is flat or undulating. Thus, this term is applied to an oriented section in reference to a geometrical plane in which, in the case of clastic limestones, preferred orientation can be measured by means of petrofabrics (SANDER, 1936). An analysis of the primary microfacies situated in the plane of stratification reveals the hydrodynamic and organic processes responsible for the nature, size, shape, sorting and orientation of the particles. All this can be quantified and treated statistically. The microfacies at the time of deposition is defined by the ratios of its particular components, as seen in Plate IE. Microsequences The vertical succession is a natural sequence that records all the variations during deposition of time fractions, including the period of any diastem (non-deposition; LOMBARD, 1956) or alteration. It coincides with the succession of geometrical planes which reflects the successivevariations of microfacies. These are visiblewhen achange appears in the microfacies, as shown in Fig.1, and can be observed in thin sections or peels perpendicular to the planes of deposition. In the case of deposition of clastics as cross-laminations, these thin sections enable directional measurements to be made of bottom currents (POTTER and PETTIJOHN, 1963). The microsequences, as shown by CAROZZI (1950, 1958,1961),reflectthe history of filling-in of the basin. Whether the thickness of the layer is a few centimeters or about 1 m, the term microsequence can be utilized to define a natural succession of microfacies constituting the layer. No matter what the scale may be, the problem is the same. The succession of microsequences in a geological sequence constitutes large units that may be rhythmic. This rhythmicity can be a function of the deformation or any major climatic changes registered in time in geochemical components (LOMBARD, 1956; STRAKHOV, 1957), granular components, or organic components (LECOMPTE, 1958). Only quantitative analysis of successions of both microfacies and microsequences (FONDEUR, 1964) enables the physical properties of original clastic carbonates to be defined, even if they have been largely obliterated by diagenesis. Diagenesis in carbonate rocks is the most complex problem. As is observed in Recent sediments, it commences immediately after deposition (GINSBURG, 1957; CLOUD,1962). Diagenesis can also appear much later during the history of a sediment and disturb the original composition and structure or grain texture (BATHURST,1958; POWERS, 1962); it can occur even after lithification of the rock (DECHARPAL et al., 1959). FAIRBRIDGE (1967) recognizes “syndiagenesis”, “anadiagenesis” and “epidiagenesis”. Diagenesis is produced by: ( I ) various organic processes (GINSBURG, 1957), the textural and structural aspects of which include

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aggregation, bioturbation, and particle-size reduction; (2) biochemical changes; and (3) physical processes such as compaction, orientation and deformation of crystals and grains (D’ALBISSIN, 1963). The physico-chemical agents include connate water, hydrothermal solutions and downward percolating vadose solutions (DECHARPAL et al., 1959). Diagenesis is accompanied by changes in the nature (authigenesis), size and shape of carbonate particles (epigenesis). Thus it modifies the original pattern of the rock and, therefore, the porosity and permeability (BATHURST, 1959; POWERS,1962). GARRELS et al. (1960) emphasized that reconstruction of diagenetic stages is possible only if the geological history of the area (tectonic history, time of faulting, etc.) is well known. Thus, a complete petrographic record of carbonate rocks must include considerations of all the important scalar phenomena. As an example, Fig.1 and Plate I show a heterogeneous succession of laminae (primary microfacies), each of which has its own characteristics (size, shape, particle and void distribution; in Plate IE the pores are intergranular). These characteristics can be evaluated under a microscope by using a point counter and an integrating Zeiss stage. Each succession of microsequences records the variations in these characteristics. For example, the sample presented in Fig.1, and which is made up of a series of microseyuences, shows a series of anisotropies that are due to an alternation of fine and coarse laminae. These, in turn, are linked to the physical properties of the rock (porosity, permeability, elasticity, etc.). It is known that anisotropy of a rock depends on the type of porosity (intergranular, vuggy, fissural, etc.) and, in the special case of carbonate rocks, on the proportions of matrix and allochems (Plate IE, C). The term “allochem” may be defined as the transported particles: pellets, skeletal detritus, oolites, etc. The relative contents of allochems and matrix are capable of causing considerable changes in the physical properties of a rock. It is, therefore, not without importance that all of the petrographic measurements be made at random on non-oriented samples. This is illustrated in Fig. 1 and Plate I where the components (organisms, oolites, pellets-see Plate IA, B, C) or depositional structures (Plate ID) are true reflection of the depositional environment. On the other hand, the sparite and dolomite between the allochems, of which the matrix is composed, reflect the diagenetic factors (Plate IE) which have modified the original features and properties of the deposit.

TRENDS AND METHODS

Scattered descriptions of carbonate rocks in texts dealing with sedimentary rocks date back to the 19th century. It was during the period between the two World Wars, however, that a comprehensive sedimentary petrographic encyclopedia came into being in France as a result of the research done by CAYEUX (1935), and a pet-

57

- .

Fig.1. Bioclastic limestone in the reef complex of Wolayersee, Carnia, Austria (Middle Devonian). General aspect of a composite layer having thickness of 23 cm. It comprises a succession of microsequences composed of calcarenite and calcilutite; the rock is fissured and scattered dolomite appears in sequences I-2 and 5. MI-MI’ corresponds to a section parallel to the plane of stratification, giving a detailed microfacies. Numbers I-5 correspond to a continuous succession of thin sections showing detailed structure of the rock.

58

Y. GUBLER

PLATE I

et al.


59

rofabric system for carbonate rocks was introduced in Austria by SANDER (1930). CAYEUX (1935) made use of thousands of isolated field samples collected by many workers to prepare his treatise on carbonate rocks, drawing attention to the most varied types in nature and very wisely presenting many problems of their genesis without attempting to solve all of them. This monumental work, based on the microscopic and qualitative examination of rocks using thin sections, is a unique reference book. As Cayeux said himself, his work is confined to the science of petrography. On the basis of systematic field surveys in limited areas of the Alps, SANDER (1936) also using his own thin sections, was able to distinguish all the microstructural characteristics of particles helpful for explaining the original organization of the rock. He introduced methods of stereographic projection which enable the spatial relations of particles (petrofabric) to be determined quantitatively. Sander’s method enables the greatest number of diagnostic characteristics of a rock to be recorded for purposes of description and correlation. His aim was to show how a comprehensive petrographic survey could clarify and add to field data. In doing this, he opened up the path for modern sedimentary petrology. Since then, a considerable amount of research has been done, mainly stimulated by practical considerations of oil finding and oil-reservoir studies (CLOUD, 1962), as mentioned above. The writers have referred to a number of publications which suffice to indicate trends and improvements, some of which contribute to the development of the petrographic side and others to studies of the petrologic fabrics. Petrographic fabrics Particle identijication Organisms. Important qualitative and quantitative advances have been made during the last decade or so in the identification of microorganisms in thin sections (CUVILLIER, 1951, 1952), their frequency (CAROZZI, 1950), and in the idenPLATE I Bioclastic limestone in the reef complex of Wolayersee, Carnia, Austria (Middle Devonian). Thin sections A, B, C, D perpendicular to MI-MI’ (Fig.1). A. Thin section in the middle part of section 5 (Fig.1). It shows from bottom to top: coarse sparite probably filling a “stromatactis” structure, altered surface, and a thin layer of dolomite with structural reliefs (Algae). The upper part is altered too and is covered by a calcarenite made up of crinoidal detritus; x 10. B and C . Both thin sections are in the lower part of section 3 (Fig.1). One can see: difference in the amount of matrix between pellet and oolitic limestones, compaction and deformation of pellets by oolitic deposits (B and C ) and orientation of pellets around an original structure of organism (B); x 10. D. Shows details between sections 2 and 3: recrystallization of oriented calcite and secondary microfolding with stress joints filled with clay minerals: X 10. E. Parallel to MI-MI’ (Fig.1); is made up of allochems (skeletal detritus and pellets), sparite (dolomite or calcite), micrite and intergranular voids, which constitute a “primary microfacies”; x 10.

60

Y. GUBLER et al.

tification of macroorganisms (NEWELL and RIGBY, 1957) or cryptoorganisms (BLACK,1956). This enables a better interpretation to be made of the paleogeography of a deposit.

Mineralparticles. Calcite, dolomite, magnesite, and ankerite are difficult to identify under the microscope without the help of more precise analytical methods, because their optical properties are very close. X-ray diffractometry gives only quantitative data but does not locate the mineral in the rock. The use of staining (1959), is very helpful. Radiomethods, such as those developed by FRIEDMAN graphic methods for identification and localization of dolomite have been described by SCHMITT (1962). Schmitt’s technique is most useful in distinguishing between different carbonates, especially in the case of dolomitic limestones. His method involves study of thin sections of dolomites (including those with variable Mg, Ca, and Fe contents) by contact microradiography, as it is possible to separate carbonates on the basis of their partition efficiency. The values of the mass coefficient of absorption for various carbonates were calculated from the tables for Xrays of V, and the values for Ca from Johnson’s universal function. In microradiography with the radiation K of V, a dolomitic zone appears darker than a calcitic one; the opposite is true with the radiation K of Ca (Fig.2). The contrast is even stronger for magnesite and calcite. The sample studied by SCHMITT (1962) was a dolomite traversed by a calcified fracture and with impurities (Fig.2). The microscopical examination either in natural (Fig.2A) or polarized light (Fig.2B) revealed no optical differences between the crystals. On the other hand, the microradiography with V and Ca showed the localization of calcite crystals, which appeared light-colored in an X-ray of V (Fig.2C), and dark-colored in an X-ray of Ca (Fig.2D). Thus one can observe two types of calcite: one filling up the vein and the other scattered through the rock as relics remaining in the dolomite. Structure identijication Structure identification deals with the orientation of particles and pore spaces with respect to reference planes (Sander’s method of petrofabrics). Pore shape and size can be studied in injecting colored synthetic resin (ETIENNE, 1963). This is of great importance in any study of the physical properties of rocks. In this chapter one should also consider the deformation of crystals which, after a certain degree of pressure, show some response to the stresses applied to them. The crystals become reoriented so that one of their optical axes assumes approximately the same direction as the greatest stress. D’ALBISSIN (1963) used physical methods to study this phenomenon in calcite and to explain the mechanism. In the scale either of an aggregate of calcite crystals or of single crystals, the effect of plastic deformation under great pressure produces intracrystalline slippage by distortion of the crystalline system. At the scale of the aggregate, there is also an


61

Fig.2. Application of contact microradiography. Sample IFP A-604, Robion B.A.; x 30. A. Micrography in natural light. B. Micrography in polarized light. C . Microradiography with vanadium. D. Microradiography with calcium. (After SCHMITT, 1963.)

intergranular slippage, which is accompanied by other phenomena acting on the same scale, such as sliding and stylolitic joints. On either scale, solution and recrystallization phenomena can occur before, during or after deformation. This structural study opens up a new line of research on the physical properties of carbonate rocks. In analyzing the components of a rock, it is evident that microscopic examination is no longer sufficient and that new and more accurate physical methods have to be used to investigate the crystalline structure itself.

Interrelationship between rock coinponents After the components have been identified, it then becomes necessary to establish characteristics for each rock on the scale of thin sections. For clastic carbonate

62

Y. GUBLER

et al.

rocks, CAROZZI(1950, 1958, 1961) using 250 sq. mm thin sections has made statistical investigations of the microscopic components, based on the clasticity index, crystallinity and frequency of the detrital components (either organisms or minerals). The values computed were then plotted graphically as variables expressing a sequence of microfacies that he considered to be indicative of bathymetric changes in the environment. A similar approach was made by PASSEGA (1957, 1960) who considered that the hydrodynamic characteristics of the environment are reflected in the distribution of grain size as represented by two parameters: i.e., C, 1 percentile approach value of maximal diameter; and M , the median. The location of C and M on a logarithmic diagram is indicative of the type of current which prevailed during deposition. In this chapter the authors have not discussed either genetic or descriptive classifications which are dealt with elsewhere in this book; however, it is important to mention here the practical classification introduced by FOLK(1959) whose aim was to establish the spectrum of the textural types which reflect the physical energy of the depositional environments. It is based on the relationship between transported particles (allochems) and chemical particles (micrite). FOLK(1959) developed a carbonate terminology consisting of only a handful of labels for major rock types determined by descriptive and quantitative methods. This system provides a framework within which the sedimentary petrographer may introduce as many variables as necessary, based on field work, e.g., color, hardness, bedding, sedimentary structures, slumping, detailed paleontological data, trace elements, stromatolitic structures, birdseye textures, vugs, etc. Special attention must be devoted to the relationship between certain carbonate minerals, such as dolomite, and minor constituents, such as clay and organic matter (DECHARPAL et al., 1959), which appear to play a role in diagenesis. This conclusion is supported by the studies of Recent sediments by CLOUD ( I 962) and TAFT(1963). The main purpose of this research is to relate any carbonate rock to its depositional environment and thus to reconstruct the original paleogeography.

Petrologic fabrics (sequence analysis) On a megascopic scale, i.e., on the scale of a layer or of a succession of layers, LOMBARD (1956, 1963) introduced the idea of stratonomic analysis based on direct field observations with particular attention to the stratification, bedding characteristics, and the boundary planes between strata. Correlation can be established between the stratum thickness curve and the lithology curve. The rhythmic character of any sequence led Lombard to reach general conclusions cancerning the depositional conditions, i.e., to relate the gross shape and size parameters of carbonate beds to the tectonic deformation of the basin.

Fig.3. Correlations based on lithologic sequences of the “Black Marble” of Golzinne (Middle Frasnian, Belgium); the Golzinne Formation is divided here into nine groups of lithologic sequences and the base of the FalnuQ formation into four. s.v.Z.- s6rie virtuelle locale= local virtual series (Lombard): O = calcareous shales; 1 = shaly limestones; 2 = nodular limestones; 3= biomicrites; 4= pure micrites (“black marbles”); 5= dolomitic limestones; A = quarry of Villeret; B= quarry Marchand; C= quarry Arthur-Etienne; D = quarry Marbres-Pierres-Granites; E= quarry JosephEtienne; F= quarry Merbes-Sprimont; g = quarry Dejaiffe; H and I = quarry of 1’Agasse; J= quarry Deffense; K = quarry “Des Polissoirs”; L= well Puits des Isnes; M = quarry van Rompaye; N = quarry Artoisenet; 0 and P= quarry at Huccorgne. (After MAMET,1963.)

cn w

64

Y. GUBLER

et al.

Lombard’s method was applied by MAMET(1963) who complemented it with petrographic studies. It enabled him to define and correlate a well-known Carboniferous formation, the “Marbre Noir” of Golzinne, Belgium (Fig.3).Using a different method, i.e., the zonation of macroscopic organisms (Rugosa and Tabulata, Porifera, Brachiopoda, Bryozoa, Crinoidea, and Stromatoporoidea) as ( 1958) established the sequences of well as mean grain-size distribution, LECOMPTE carbonate reef environments in the Devonian of Belgium. He interpreted them as a reflection of the progressive “subsidence” of the basin. Lecompte also emphasized that these Devonian reef environments differ in several ways from Recent reefs. He believed that there were constructional phases, detrital and also chemical phases, each related to the turbulence of the respective environments. The utilization of such a large number of parameters in the final model, “petromodel”, is possible only in case of direct and continuous recording of all the data in the form of logs and symbols, such as proposed by BOUMAand NOTA (1960) and by GUBLERand BERTRAND (1965).

Relationship bet ween petrographic fabric and physical propertics No petrographic study can ignore any of the factors which determine the structure of the rock; in particular, grain-size distribution, diagenesis, void distribution and, consequently, the inter- and intragranular porosity. A comparison between the results of granulometric analyses on using thin sections and loose samples discloses a problem, because thin sections are not absolutely representative of the original sediment. In general, it can be assumed that the granulometry deduced from a thin section is closer to that of a loose sample when (f’-f)l is lower as shown in the first example given below. This has led to the devclopment of the technique of making systematic and statistical measurements on thin sections with a point counter and a Zeiss particle analyzer (FONDEUR, 1963). The physical properties of the rock will be different depending on whether the rock has a constant granulometry or not (MURRAY,1960). Relationship between petrology and petrography The above brief discussion shows that the major problem is to identify on a lifesize scale, the original paleogeographic setting of a rock, establishing its origin and any modifications it may have undergone during geological time. This has led to fundamental niineralogical research which is described in other chapters. It includes experimental laboratory studies on physico-chemical processes ( R I V I~RE, 1939; GRAFand GOLDSCHMIDT, 1956; BARON.1960; GRAF, 1960) and on biochemical processes (NESTEROFF, 1955; LALOU,1957). Extensive ~lf’

= porosity

in thin section, andf = porosity of loose sample.


65

fieldinvestigations on the Recent carbonate sediments of the platforms in the Bahamas (ILLING,1954; GINSBURG, 1956, 1957; NEWELL and RIGBY,1957; TAFT,1961, 1962; CLOUD,1962), the Gulf of Batabano (DAETWYLER and KIDWELL,1959), and the Persian Gulf (HOUBOLDT, 1957; EVANSet al., 1964) have produced a general life-size model of the organic communities and mechanical accumulation processes found in such areas. These comprise climatic, hydrodynamic, biochemic and geochemic (including isotopic), and diagenetic processes. On the basis of petrographic and petrologic analyses, and taking into account all the basic data provided by these fundamental studies, the authors of this chapter hope to present a clearer picture of how to identify the former geographic setting of carbonate rocks and their physical properties. Using only standard field and microscopic equipment, the authors tried in two examples to apply in the field the results of quantitative analysis obtained from thin sections. In both cases they used continuous field recording (“petrolog” in the first case and electric log in the second case).

EXAMPLES

“Pisolitic limestone” of Vigny, France In the first example, the authors (Bertrand and Gubler) started with a model considered from the point of view of its reservoir characteristics. Field and laboratory analysis brought out the parameters that control these properties. The final objective was to gain a better understanding of a type of clastic carbonate body from the practical point of view, i.e., location and extent of reservoir properties. The following petrologic and petrographic parameters were considered: (1) composition; (2) stratification, sedimentary structures and granulometry that are direct result of hydrodynamic conditions; (3) color, degree of induration, and minor mineralogical constituents that reflect the environment of deposition and diagenesis; and ( 4 ) biological constituents, trails and burrows that reflect environment and increase primary porosity. The vertical succession of the above characteristics makes up part of the sequential analysis. The field data were collected on a “petrolog” using a scale of 1 : 50 (Fig.4). Vigny is situated 25 km northwest of Paris, France. It is famous for its quarries of the so-called “pisolitic” limestone, whose origin and age have long been controversial. The structural model of the reef complex (Lower Eocene: Montian), about 3 kmzin area, has been described by DESMIDT (1960). There is a superposition of two reef complexes composed of biolithite (Algae and corals), bioclastic limestones (calcirudite and calcarenite), as well as lime-mud (chalk). The present study is only concerned with the section no.2 (C.2 in Fig.5).

67

PETROLOGY AND PETROGRAPHY OF CARBONATE ROCKS E

W +lOOrn,

c.4

r *'O0

c3

C.6 c 2

+

80

+ 80

+

70

* 70 * 60

60

R;g,)Reef

0

Bioclastic ~lmestone D C h a l k

BrokenbiDck

Basement

Fig.5. Section of reef complex showing:distribution:of lithofacies,_Vigny,France.

Although detailed investigations were made in different sections (organic content, granulometry, and cementation: Fig.4), only limited data are given here in an effort to: ( I ) compare the results at different scales of observation (formation, hand specimen, and thin section), with special emphasis on granulometry; and (2) present a rapid examination of the relationships between physical and lithological characteristics. Nomenclature. The writers feel that there is no value in giving a special name to a rock that is specifically determined by its own parameters (Fig.6). In the case of sample No.1253, it could fall within the category of biosparitic calcirudite. CARBONATE ROCKS MICTOSC~PIC data Texture biosparitic General sorting(al1ochems+ micrite) Door jeneity of matrix good

Macroscopic data Structure fine cross-bedded Color: white Fracture: irregular hems Cora Is. echinoids,A Alloct Nature

~~

~~

small coral

algal detritus Iasycladace

Observations

coated sparite detritic particles (7)

normal

broken, round

v

1

0

m

echinoids detritus

lithic fragments (chalk)

normal

micritic deposition in intergranular voids(2)

round

Fig.6. Form for recording petrologic and petrographic data of carbonate rocks.

Fig.4. "Petrolog" of the quarry A. Pisolitic limestones of Vigny, France. Symbols used for apparent granulometry: CG= coarse granular; VC= very coarse; C= coarse; M= median; F= fine; VF= very fine.

68

Y. GURLER

et al.

PLATE I1

Grain size

Granulometric data of sample 1253 in section C.4 (see Fig.5). A. Histograms based on hand specimen (unbroken line) and on thin section (broken line). B. Cumulative curves based on hand specimen (unbroken line) and on thin section (broken line). C . Hand specimen. D. Thin section.


69

Discussion of granulometric results at difSerent sample scales Point examination. Point examination was made on thin sections (Plate IID) prepared from a hand specimen (Plate IIC). The slide was cut perpendicular to the plane of stratification in this case, and can be considered as a typical microfacies because the deposit is homogeneous. The granulometric data are presented on Plate IIA, B. The two curves differ slightly. The hand specimen shows ( I ) a larger size of the grain having maximum frequency, and (2) the best distribution, with good symmetry. It should be noted, as indicated on Plate IIC, that the measurement was made of the same sample from which the thin section was prepared. A zonal extension of measurements would have given different results. Section examination. Section C.4 (Fig.5) has been studied on using microscopic samples (thin sections) and some random hand specimens, in conjunction with M

VC

C

sample! 1254 1253

I

1252

1251

&

!

1250 11249

/

CG

I

1248 1247

1246 1245

1242 1241

Size of most frequent grain

-7

---__

Fig.7. Variations of grade-size of most frequently occurring grain in section C.4, M= median; C = coarse; VC= very coarse; CG= coarse granular; I = measured on macroscopic samples; 2 = measured on slides; 3= based on field data.

70

Y. GUBLER

et al.

continuous observations in the field. Only the most representative grains (most frequent) were studied, and the results are illustrated in Fig.7. The granulometric curves differ as given below. ( I ) The curves of the random hand specimens differ from the continuous granulometric log recorded in the field. This may be explained by non-represen-

C4 C6

C2

Fig.8. Porosity variation in “pisolitic” limestone of Vigny, France.

U 0

c c

1 mm: (2) coarse-grained-0.5-1 mm; (3) medium-grained-0.25-0.5 mm; ( 4 ) finc-grained -0.1-0.25 mm; (5) very fine-grained-0.01-0.1 mm; (6) micrograined - 1) indicates that the rock contains more than 50 % grains, and one with a GMR of 1 defines a rock with about 50% grains and 50% micrite material; finally if

+

106

H. J. BISSELL A N D G. V. CHILINGAR

the GMR is < 1 the rock contains less than 50 % grains. This is shown in the classification of limestones by MOSHER and PINNEY (1963) as modified after the one by L EIGHTON and PENDEXTER (1 962). In any tabular form of rock classification, various “pigeonholes” are usually set up arbitrarily for the grain types, and many rocks consist largely of one textural type, such as micrite, calcarenite, etc. Seemingly, most limestones contain two or more grain types. Where two or more grain types are present in approximately equal amounts in the rock, it is suggested that the two names be hyphenated, such as detrital-micritic limestone. Most generally, however, one grain type predominates over the other. In this case, the limestone is given the name of the predominating grain type with the other as an adjectival modifier, or essential prefix. For example, the carbonate may contain 60% skeletal grains and 40% detrital grains; this would be termed detrital-skeletal limestone. If the rock is composed of 15 % detrital grains and 85 % skeletal grains, however, it is termed detrital, skeletal limestone. In other words, the contents of grains (in %) are arranged in tabular form in the order of increasing percentage in the classification scheme. A rock composed of 10 % detrital grains and 70 % skeletal grains, all embedded in micritic material composing 20 % of the whole, could be named detrital, micritic, skeletal limestone. If detrital grains and micrite are present in about equal amounts, such as 15 % each, and skeletal material comprises the remaining 70%, the rock is named detritalmicritic, skeletal limestone. The name should be euphonious, and pronounceability will dictate the order of arrangement where the hyphen is used. WOLF(1960; see also LEIGHTON and PENDEXTER, 1962) applied a plan of pigeon-holing limestones according to textural types; the vertical and horizontal lines in their charts are arbitrarily assigned for guidance, not limitation. For example, a carbonate may consist of approximately equal amounts (say 30% each) of detrital, skeletal, and pelletal material, cemented with micrite; it could, according to the suggestions advanced by the present writers, be termed detrital-skeletalpelletal limestone cemented with micrite. It may be more pleasing to the ear of a listener to term the rock a detrital-pelletal-skeletal limestone cemented with micrite. It would not be termed micrite; perhaps preference would dictate that it should be identified as a micritic, detrital-pelletal-skeletal limestone. This differs in no fundamental fashion, other than constituents involved, from the naming of an igneous, or a metamorphic rock; for example, quartz-muscovite-albite schist, etc. CLASSIFICATION OF DOLOMITES

General statement

Dolomites have been defined as carbonate rocks composed of more than 50 % by weight of the mineral dolomite (LEIGHTON and PENDEXTER, 1962, p. 53). Some

107


TABLE V CLASSIFICATION OF ROCKS INTERMEDIATE IN COMPOSITION BETWEEN PURE LIMESTONES AND DOLOMITES

(After data by CAYEUX, 1935; CAROZZI,1960, p.264) Content (%)

Rock name

Limestone Magnesian limestone Dolomitic limestone Calcitic dolomite Dolomite

calcite

dolomite

>95 90-95 50-90 10-50 50.1 9.1 -50.1 4.0 -9.1 2.2 -4.0 1.5 -2.2 1.4 -1.5 1.25-1.4 0.80-1.25 0.44-0.80 0.18-0.44 0.03-0.18 0.00-0.03

and thus FROLOVA (1959, p.53) proposed a new classification of dolomite-magnesite-calcite series (see Table VII). On the basis of Ca/Mg weight ratios, CHILINGAR (1957a) recognized the following groups of dolomites: ( I ) magnesian dolomites (Ca/Mg = 1.0-1.5), (2) dolomites (Ca/Mg = 1.5-1.7), (3) slightly calcitic dolomites (CaIMg = 1.7-2.0), ( 4 ) calcitic dolomites (Ca/Mg = 2.0-3.5). The upper limit of calcitic dolomites was chosen because the Ca/Mg ratio of 3.44/1 is the lowest ratio known in skeletal structures of organisms (Goniolithon, a calcareous Algae, has a Ca/Mg ratio of 3.44/1). A pure dolomite has a ratio of 1.648/1, and some dolomites contain an excess of magnesium (magnesian dolomites). Before discussing the classification of dolomites, it is important that attention should be directed to some of the prevailing concepts regarding types of dolomites. For example, VISHNYAKOV (195 1,p. I 12) recognized the following major genetic types: (I) epigenetic, (2) diagenetic, and (3) primary. He considered those dolomites as epigenetic which resulted from the alteration of completely lithified limestones either by downward percolating meteoric solutions or by rising hydrothermal solutions, mainly along fractures. These dolomites have obscure stratification, patchy distribution, non-uniform grain size (clear-grained), relic structure and are cavernous. Diagenetic dolomites are considered to be those of great extent and volume; they were formerly limestones. Fossil relics are commonly present in diagenetic dolomites. Primary dolomites result from direct chemical precipitation out of water, and are recognized when associated with other primary sediments. They are well stratified and are pelitomorphic ( ~ 0 . 0 1or ~0.005-0.003mm, with


109

coagulating centers in some), possess characteristic microlayering, and very rarely contain fossils. Primary dolomites lack appreciable porosity (in fact are normally rather dense) and are commonly interlayered with evaporites, clays, marls, micrite with suspended oolites, gypsum (commonly containing pseudomorphs after gypsum), and oopelletal to micritic limestones with bladed and/or spherical to subspherical algal bodies. TEODOROVICH (1958, p.303) does not believe in very widespread occurrence of epigenetic dolomites because of difficulty in explaining the source of huge amounts of magnesium which is necessary to effect dolomitization. MCKINLEY(1951, pp. 169-1 83) categorized the theories relating to marine replacement of limestones to form dolomites as follows: ( I ) penecontemporaneous replacement, (2) syndiagenetic replacement, and (3)post-diagenetic (i.e., epidiagenetic) replacement. He defined penecontemporaneous as that type of replacement that takes place immediately after deposition of the sediment, before consolidation or lithification into rock, and prior to superposition of any great thickness of additional material. According to him, syndiagenetic dolomitization takes place during diagenesis of calcareous material, and should be considered a factor in transformation of the sediment into rock. Post-diagenetic replacement (which McKinley indicated as occurring in the marine realm) involves alteration of the completely lithified limestone to dolomite by the action of sea water, or of static connate waters. He pointed out that many dolomitized reef limestones appear to have been altered after lithification by continued exposure of the calcareous reef rocks to sea water in an environment conducive to replacement. SKEATS(1918), who studied coral reefs in the south Pacific, concluded that marine dolomitization took place in shallow water (and as deep as 150 ft.), under slightly increased pressure, with abundant COZ,and in porous limestone allowing free circulation of sea water. The present chapter is not a treatise on processes of dolomitization, but rather it relates to problems of classification. The foregoing is, therefore, but a brief discussion serving as a prelude to an organization of these carbonates into a rational scheme of classification. According to LEIGHTON and PENDEXTER (1962, pp. 53-57), the different types can best be accomodated by a descriptive system of nomenclature, one based primarily on compositional grouping and with appropriate modifying textural terms. They recognized two groups: (1) calcareous dolomitesthose containing 50-90 % dolomite, and (2) dolomites -those containing 90 % or more dolomite. For each type, description in parentheses follows the major name to designate the precursor; for example, a rock may be termed calcareous dolomite (calcareous skeletal fragments with calcareous matrix and cement). Schemes of carbonate-rock classification, which include both limestones and dolomites, by their very nature assume that the latter resulted from alteration of the former. This works very well for dolomites which originated through diagenesis af the various types of limestones; the classification scheme then requires only spaces in the table to indicate the degree of alteration. BRAMKAMP and POWERS

110

W u)

2 E 0

5

n W

LL

_I

4 IW

I 0

F z 5

W

0

Clear-grained relic

Micro-grohed relic

Cleor -grained dolomites

Micro-grained dolomites

relics

negofive

wilh

wifh structure

with negative sfructure

Wifh relic. chemical or bio-chemical strucfure

With relic-organic structure

micro-groined

pisolitic

plsolltlc

-

and

pisolitic

bean-shoped

e,

13

-

] 5C8x

8 ._ v

cr

a

3

1 'f

I cd

v

F

18

0

15

1 V6

3

cn

m

m"

lc! a

rn

'0

0

1 .^ %

3

U

-

Fa

9

I A 1 %

19

H. J. BISSELL AND G. V. CHILINGAR

Other

Cryptograined-pelitomorphic from irregular groins

Detrital

Nodules, onlltes, etc

or

Detrital and biomorphic Biomorphic

Oolitic Detrital Biornorphic-detritol

or

Biornorphic oollllc Detrifal

breccia

Biomorphic - detrital

Biomorphic

SlltY Sandy

etc.

Conglomerate and

Radiolife,

Oolitic-like

and

Lumpy (or clotted or nodular1 and micro- lumpy

Oolitic

5

+

3

8 Bio-dctrital Biomorphic - d e t r i l a l

.-

a

Biomorphic


111

(1958), FOLK(1959, 1962), MOLLAZAL (1961), POWERS (1958), TEODOROVICH (1962), CHILINGAR and BISSELL(1963a), RICH (1963, 1964), and others have and POWERS(1958), for successfully demonstrated this possibility. BRAMKAMP example, were among the first to allocate “pigeonholes” in their chart for those limestones that are not visibly altered, those that are moderately altered, and those that are strongly altered, and for those whose original texture has been obliterated. RICH(1963, 1964) modified the classification of BRAMKAMP and POWERS (1958) by assigning percentage parameters to the various degrees of alteration, and by adding such identifying names as “porphyroblastic” and “granoblastic” to two of the columns. MOLLAZAL (1961) indicated relative degree of diagenetic alteration of limestones by assigning objective data relating to size and fabric of crystals. TEODOROVICH (1958, p.295; also in: BISSELL and CHILINGAR, 1961, p.614) devised a plan of classification (Fig.4) which indicates whether the relics are positive (former fossils and minerals identifiable) or negative (former fossils and minerals discernible but not readily recognizable). In his discussion of carbonate rocks, CAROZZI (1960, pp.264-284) was concerned with the petrography (and minor petrology) of dolomitic sediments, and categorized them into the following groups: magnesian limestones and magnesian chalks, dolomitic limestones, bioconstructed dolomitic limestones, bioaccumulated dolomitic limestones, fine-grained dolomitic limestones, dolomi tic chalks, dolomitic lithocalcirudites and lithocalcarenites, dolomitic biocalcirudites and biocalcarenites, dolomitic oolitic calcarenites, dolomitic calcilutites, and dolomites. This approach to the study of these carbonates has merit, and obviously could be a guide to the orderly-minded investigator. It is obvious that dolomites are susceptible to classification utilizing two major parameters, viz.: (I) relic textures (and structures), and (2) crystallinity (which is a function of degree of alteration in diagenetic dolomites, but is commonly lacking in primary types). It may be pointed out that some dolomites which have a high degree of crystallinity (including medium- and coarse-crystalline varieties), lack all vestiges of relics, positive or negative. It is not to be inferred, however, that in all cases high degree of crystallinity is indicative of complete obliteration of relics. It is possible that primary dolomites of the evaporitic and otherwise restricted environments may lose their aphanic to fine-textured character and become medium- to coarsely-crystalline during diagenesis just as limestones do. By utilizing the particle size grade chart (see Table IV), this objectivity in assigning degree of crystallinity can be realized. Furthermore, if the original carbonate was a variety of limestone and has been diagenetically dolomitized, the charts devised by BRAMKAMP and POWERS (1958), POWERS (1962), or the modified version of RICH(1963, 1964) can aid in assigning a percentage parameter to this type of dolomitization. Eventually, petrographers may be able to state more objectively the stage of diagenesis that was attained (such as early or penecontemporaneous, medial, and late).

112


LEIGHTON and PENDEXTER (1962, p.57) stated the following concerning the genetic implications in naming of dolomites: “In spite of attempts to avoid genetic implications in naming dolomites and dolomitic rocks, it is impossible to do so completely.” They cited the example of “dolomitic limestone”, and also pointed out that the association of finely micrograined dolomites, laminated dolomites, and dolomitic breccias with anhydrite, chert, and microcrystalline micritic limestone, has led t o the use of the term primary dolomites in evaporitic sequences. Some TABLE VIII CHARACTERISTICS OF DOLOMITE GROUPS~

(After LEIGHTON and PENDEXTER, 1962; modified by the present authors) _ _ _ _ _ ~ ~ Dolomites in evaporitic sequences

Dolomitized rocks

~.

~~~

-

Generally finely micrograined.

May be very coarse grained (sucrosic).

Commonly aphanic t o finely crystalline with uniform texture.

May be meso- to macrocrysfalline. Associated with limestone sequences.

May be laminated or brecciated.

May contain fossil molds.

May be interlayered with dolarenites.

May contain relic limestone textures and belong to gradational sequences showing increased development of dolomite.

May have admixed clay andfine silt. Associated with anhydrite, chert, and microcrystalline micritic limestone.

May be associated with transgressive or regressive deposits that display truncation of other beds.

May be associated with micrific limestones. (some of which may contain “floating” oolites and pellets).

Positive and negative relics of fossils, pellets, detrital particles, lumps and coated grains in dolorhombic mosaic.

Associated with red shales and siltstones.

Reef and bank deposits retain topographic .form but are nodular, lumpy, vary-grained (and may be cavernous, andlor brecciated) .

Associated with gypsum and anhydrite that contain scattered dolorhombs. Contain no relic limestone texture.

Areally extensive, moderate to large volume, subtle facies changes.

May be interbedded with dololutites some of which are bituminous.

May correlate with tectonic features, fault zones, anticlines, etc.; or with former land surfaces.

Collapse limestone- and dolomite-breccias may be present.

Variation in porosity and permeability, from low to very high values. ~

ICharacteristics in italics were added by the present authors.


113

petrographers avoid the word “primary” when referring to the dolomites in such a sequence, but this appears to be no more than a play on words. If the entire sequence is an evaporite sequence or suite, then the dolomite is possibly an evaporite; if it is “primary” then no justification can be found to avoid the term. It should be remembered, however, that some very early diagenetic dolomites are classed as “primary” by many geologists (see BISSELL and CHILINGAR, 1962). After making a survey among various geologists concerning the terminology of carbonate rocks, RODGERS (1954, p.232) stated: “ . . . common American usage calls dolomite primary only if the particles were dolomite when first formed, as by direct precipitation from sea water; otherwise they are secondary.” SANDER(1936) termed dolomite primary if the particles were dolomite when they reached their present position in the rock fabric, secondary only if they are replacements of some material that occupied the same position. Rodgers has made an excellent point, in that if the terms “primary” and “secondary” are to be used, their meaning must be defined. LEIGHTON and PENDEXTER (1962, pp.57-58) pointed out that each of their two groups of dolomites has particular textural characters and rock associations; these are listed in Table VIIi, with additions of the present authors.

PROPOSED LIMESTONE AND DOLOMITE CLASSIFICATlONS

The proposed classification of limestones in this chapter (Table IX) is intended to demonstrate the utility of application of these parameters: ( I ) composition, (2) texture, (3) grain/micrite ratio (GMR), and (4) energy index (Ei). i f properly applied, these will also be of great value in dolomite classification, particularly for the diagenetically altered limestones. Of utmost importance is the identification of types and relative amounts of carbonate grains (detrital, skeletal, pelletal, lumpal, coated, etc.) in the field with no more than 10 x hand lens, dilute (3 %) HCI, and eye-dropper bottle with glycerin and water (the latter is merely a wetting medium). The field worker (petrologist) should also be able to make a semiquantitative estimate of grain/niicrite ratio, and/or grain/matrix ratio. The classification should also have utility in the laboratory where microscopes, chemicals, and other materials and equipment (such as X-ray, DTA, etc.) can establish greater precision and thereby provide adequate definition. it is argued that any classification should not be a mere play on words which would result in a semantic struggle; the present authors believe that, after certain basic definitions, if a word is not self-explanatory it has no real value to the sedimentary petrologist and petrographer. The grain/micrite ratio shows the relative amounts of coarse and fine-textured carbonate material; this theoretically is controlled by wave or current action and thus gives an insight into the degree of agitation. If the worker can arrive at

114


TABLE IX CLASSIFICATION OF

LIME ST ONES^

(Modified after LEIGHTON and PENDEXTER, 1962; PLUMLEY et al., 1962; and others) Energy index (EI)

Graiiz/ Percentage Moved and deposited by waves micrite of grains and currents ratio abraded grains

detrital

skeletal

Detrital limestone2

Skeletal limestone

-

Strongly agitated (growth and deposition in strongly agitated HzO)

Moderately agitated (deposition in moderately agitated HzO)

9/1

75

Slightly agitated (includes to-andfro HzO action)

Intermittently agitated (alternately agitated and quiet HzO) Relatively quiet (deposition in quiet HzO not necessarily stagnant-may be gently agitated)

90

1/1

50

25

119

10

________

Micritic, detrital limestone3

Micritic, skeletal limestone

Micriticdetrital limestone

Micriticskeletal limestone

Detritalmicritic limestone

Skeletalmicritic limestone

Detrital, micritic limestone

Skeletal, micritic limestone

Micritic limestone

Micritic limestone

Horizontal combinations are useful, as detrital-skeletal, oolitic-micritic, etc. If size connotation is deserved, such terms as calcirudite, calcarenite, calcisiltite, calcilutite, etc. may be used instead. If the particles in the rock are of different orders of size magnitude, and are larger than micrite, the term matrix is used for the smaller individual units that fill the interstices between larger grains.

1 2


115

Accumulated in place _.

~~

accretion-aggregation grains ~

organic frame builders

pellets

lumps

coated grains

Pelletal limestone

Lump limestone

Oolitic, pisolitic, oopelletal, algal, etc. limestone

Colonial corals, stromatoporoids, colonial Algae, bryalgal, bryozoan, foraminiferal, etc. limestone

Micritic, pelletal limestone

Micritic, lump limestone

Micritic, oolitic, pisolitic, etc. limestone

Micriticpelletal limestone

Micriticlump limestone

Micriticoolitic pisolitic, etc. limestone

Micritic coralline, algal, bryalgal, etc. limestone (use appropriate comma and/or hyphen for correct combination)

Pelletalmicritic limestone

Lumpmicritic limestone

Ooliticpisolitic etc. micritic limestone

Pelletal, micritic limestone

Lump, micritic limestone

Oolitic, pisolitic, etc., micritic limestone

Micritic limestone

Micritic limestone

Micritic limestone

Corailine, algal, bryozoan, bryalgal, etc. micritic limestone (use appropriate comma and/or hyphen for correct combination)

chemical, biochemical

a

s

116

H . J. BISSELL AND G. V. CHILINGAR

some measure, even semi-quantitatively, of the level of energy, he can ultimately set up a more objective energy index. Consequently, some measure of success will result in identifying turbulence, energy level, and, of course, the most sound interpretation of particular environments within the depocenter. The grain/micrite ratio (GMR) and energy index (EI) can be used as indicators of the amount of physical (mechanical) energy necessary to transport and deposit the carbonate sediment. These are shown in the left-hand column of the proposed classification (Table IX). The remainder of the chart contains the various compositional-textural groups, and their combinations. Various sedimentary carbonate-rock classifications have been published within the past quarter-century, and many of these are worthy of careful analysis by the student of sedimentary petrology and petrography. The present authors have tested some of these classifications in recent field and laboratory investigations; perhaps certain classifications appear to have more value than others, largely because some proposals relate to a certain province and, therefore, represent the geologist’s particular experience in that area. The classification proposed in this chapter is, therefore, a reflection of the present authors’ researches in the limestone provinces in the western United States. The classification of FOLK (1959, 1962) was proposed for Beekmantown (Ordovician) rocks of Pennsylvania; it has been applied successfully in certain other regions. The classification of LEIGHTON and PENDEXTER (1962) was set up originally in Jersey Production Research Company’s Carbonate Rock Manual and was tested in Montana and contiguous areas, and later was given a rigorous test on Paleozoic carbonates in the Great Basin, Rocky Mountains, and Colorado Plateau regions of western United States. Subsequently, it has been shown to have utilitarian value in many other places. MOSHER and PINNEY (1963), for example, slightly modified this classification, and demonstrated its value for most carbonate rocks. PLUMLEY et al. (1962) developed a scheme of limestone classification which incorporates the parameter of energy index (EI), thereby providing for a certain objectivity in environmental interpretation. The present authors have added this parameter in the classification given in this chapter. DUNHAM (1 962) classified carbonate rocks according to depositional texture; these textures are: ( I ) mudstone (less than 10% grains), (2) wackestone (more than 10% grains), (3) packstone, ( 4 ) grainstone, and (5) boundstone. I and 2 are mud-supported, 3 is grain-supported, and all three contain mud; 4 lacks mud and is grain-supported. These four textural groups d o not show evidence of the original components having been bound together during deposition. Boundstones, by contrast, are those limestones in which the original components were bound together during deposition, as shown by intergrown skeletal matter, lamination contrary to gravity, and other features. If the depositional texture is not recognizable, it was recommended that this type of limestone be termed a crystalline carbonate (DUNHAM, 1962, p.121). PRAYand WRAY(1963) indicated that the classification of DUNHAM (1962) proved


117

useful in their studies of the porous algal limestones of Pennsylvanian age along Honaker Trail in the canyon walls of the San Juan River, southeastern Utah. BRAMKAMP and POWERS(1958) and POWERS(1962) demonstrated the utility of their classification of limestones and diagenetically altered limestones for Arabian Upper Jurassic reservoir rocks. G. E. THOMAS (1962) tested his classification of carbonate rocks on selected Paleozoic carbonate cycles and reef complexes in a western Canada basin; in his scheme of classification, carbonate rocks can be described and grouped into textural and porosity units for mapping purposes. NELSON et al. (1962) used the term skeletal limestones for rocks which consist of, or owe their characteristics to, the in-place accumulations of calcareous skeletal material. Their limestone classification is, therefore, one in which the rock is classified according to the organism primarily responsible for its formation. These workers reviewed usage of the terms reef, bioherm, biostrome, and bank, and related terms, as well as making certain recommendations as to correct usage. Attempts have been made to classify the modern Bahamian carbonate sediments; one such classification is that of IMBRIE and PURDY(1962). Their scheme identifies five discrete sample groups, as follows: ( I ) oslitic, (2) grapestone, (3) coralgal. (4) oolite, and (5) lime-mud facies. They tested their scheme of classification together with that of FOLK(1 959), with substantial success. Their shelf lagoon sands, for example, include the oolite, oolitic, and grapestone facies, and are comparable to the oosparites and intrasparites of Folk. Outer platform sands of the Bahamas embrace their coralgal facies, equivalent to Folk's biosparites and biopelsparites. Their muddy sands of the shelf lagoon are the lime-mud facies, thus equivalent to oomicrites, intramicrites, biomicrites, biopelmicrites, and pelmicrites of Folk. RICH(1963, 1964) proposed a limestone classification, based on his studies of rocks of Late Paleozoic age in southern Nevada; his classification is, however, mostly a modification of the scheme of BRAMKAMP and POWERS (19x9, and of others. Many of the classifications just mentioned were published by the American Association of Petroleum Geologists in 1962 as a symposium Classification of Carbonate Rocks (HAM,1962). This work is generally available, and the charts which it contains need not be reprinted here. BAARS(1963, p.101) pointed out that this ". . . symposium is a monument to the advancement of our understanding of the carbonate rocks and is highly recommended as an introduction to the study of carbonates. However, no particular classification was found to be mutually agreeable to all authors. And so it goes. For every carbonate petrographer there is a unique system of classification." Although BAARS(1963) did not propose a scheme of classification of limestones, he did discuss them in terms of these compositional types: lime mud (= micrite or matrix), particles or grains, cement, and pore space. He recognized five kinds of particles (which are the ones suggested by LEIGHTON and PENDEXTER, 1962), as follows: ( I ) skeletal particles, (2) pellets, (3) coated grains, (4) detrital particles, and (5) composite particles.

118


TABLE X CLASSIFICATION OF DOLOMITES

(Modified after SHVETSOV, 1958) Main genetic groups of dolomite h Y

9 9 2

a

Y

.i;

.e Y

a

v

z Y

E .-

Major part organic

_

Reefs

Coralline, algal, bryozoan, etc.

Banks

Biostromes, layered, coquinites

-~

_ -~

Major part detrital

_

_

_

Pellets -

2;

~

~

~

~.

-

~

Lumps

Moderately to strongly altered (but preserving vestiges of texture and composition of original rock)

~

~~

~

--

~~

Positive pelletal texture

-

~

-

Major part chemical or biochemical

~

~

Dolarenite (original dolomite sand) Dolomitized lithocalcarenite

Lumps and composite

0

~

Dolomitized chalk

_

~

_

~

~

Micrograined

Detrital

-~

-~ ___ -Foraminiferal, sparry criquinites, brachiopods, molluscs, algal

-

Skeletal

(u

n

Characteristics, features, examples, varieties

~~-~~-

Y

-

Main textural types

.~

Megalumps to microlumps

~-

Algal lumpal dolomite

concentric rings

grains ~

-

-~

Superficial oolites and pisolites

Sutureddolorhombssurrounding skeletal-detrital particles Brecciated lumpal dolomite

Dolomitized spherical to subspherical grains with few concentric rings Skeletal, detrital, pelletal, etc. elements as positive relics

-~

-~

Matrix may be micrograined calcite ~

_-

__

Vary-grained, with mosaic of Matrix-micromosaic of dolorhombs sutured incompletely rhombic dolomite surrounding positive Sucrosic dolomites, including sandstones being ieplaced; relics skeletal, detrital material ~

--

Cryptogeniccompletely altered to complete obliteration. Few negative relics may be present, suggesting texture and composition of original carbonate

_-

+

~~~

Reef and bank limestone deposits originally; lumpy, cavernous, patchy, vary-crystalline

~-

~~

Fully dolorhombic; vary-grained Uniformly crystalline; relics rare -

Micro- to finely crystalline

~-

Dolomitized reefs, negative relics

-~

Dolomitized chalks and tuffs -~

Vary-grained lumpal dolomites Nodular, lumpy, brecciated, __ may be cavernous; some rocks are meso- to macro-grained Completely replaced micrites sucrosic, having granular, Calcarenites and quartz sands origitranslucent appearance; nally; dolarenites, sucrosic relics absent

119


Main genetic groups of dolomite

Main textural types

Characteristics, features, examples, varieties

Crystalline (relics completely destroyed, or were never present)

Phaneric

___-

Aphanic

Sparry dolomite (dolosparite)

Micri tic dolomite or dolomicrite (some types are primary dolomicrite, and may be laminated)

Finely crystalline Microcrystalline __________ Cryptocrystalline

___ ~

(uniform textures may be primary; some contain fine silt and clay.

- 0.01

mm

- 0.001

mm

TAFTand HARBAUGH (1964) made a detailed study of the mineralogy and physical chemistry of the modern carbonate sediments of southern Florida, the Bahamas, and Espiritu Santo Island, Baja California; they did not classify the sediments according to any prevailing scheme nor did they present a new classification, but their data contribute in a very substantial manner to an understanding of some of the processes of sedimentation and diagenesis of modern carbonate materials. If judiceously utilized, this information can assist any petrographer working with modern or ancient sedimentary carbonates. Soviet geologists have made excellent contributions to the knowledge of sedimentary carbonate rocks, and particularly to their classification. The two classifications recommended by the present writers include the one by TEODOROVICH (1958, p.291), which is presented in Fig.3 of this chapter, and the one by SHVETSOV (1958, p.292). These classifications were published in English by BISSELL and CHIUNGAR(1961, pp.612 and 615, respectively). KRUMBEIN and SLOSS(1963, pp. 179-1 SO) also called attention to the classifications of these Soviet workers. Obviously, no single classification of the sedimentary carbonate rocks and particularly the limestones will immediately suit taste and fit the needs of each

120

H. I. BISSELL AND G . V. CHILINGAR

sedimentary petrographer. The suggestions made in this chapter are advanced at a time of expanding knowledge of sedimentary petrology and petrography, and should serve as guides for future studies. Some classification plans, such as the ones by BRAMKAMP and POWERS (1958), FOLK(1959, 1962), POWERS(1962), and RICH(1963, 1964), include space for altered limestones and dolomitized limestones. These are herein recommended with the suggestion that each worker, or team of workers, may seek to discover the utilitarian values in each; however, they may have to be modified for particular provenances and perhaps also for certain sedimentary suites. Some have real value for surface work, some for well cuttings and cores, and others may assist the investigator in field and laboratory research (see MOLLAZAL, 196I). The proposed classification of TEODOROVICH (1958, p.295; see also BISSELLand CHILINGAR, 1961, p.612) is recommended for critical analysis and testing; the present authors have discovered that it is of great value. The classification of dolomites proposed in this chapter (Table X) is the outgrowth of field and laboratory work by the present authors with strata of Paleozoic age in the western United States. This classification embodies certain elemental parts from published classifications, which were rigorously tested. Principal parameters are crystallinity, high degree of alteration, and low degree of alteration. No effort was made to allocate particular types to one environmental type only, because this is unrealistic. Careful examination of the two classifications (one for limestones and one for dolomites), proposed by the present authors, however, reveals their utility. If tested in both the field and the laboratory, their strengths and weaknesses can be ascertained, and evaluated accordingly. One of the most comprehensive reviews of various classifications of carboK nate rocks up to the year 1959 is found in the excellent article by M I ~ (1959). A glossary or terms considered necessary for petrologic and petrographic work of sedimentary carbonate rocks is also included in this chapter; this is in no way exhaustive, but does embrace numerous terms that are now scattered in geologic literature. Finally, various plates of photomicrographs illustrate the rock types mentioned in the proposed sedimentary carbonate classifications in this chapter (Plates I-XVI).


121

EXPLENATION OF PLATES PLATE I

A. Detrital limestone (calcarenite). Composed of sand-size particles of limestone, pelletal mate-

rial, and quartz sand grains in finer-grained matrix. Derryan age Oquirrh Formation, Blue Spring Hills, Oneida County, Idaho; x 3. B. Detrital limestone, composed of silt- and sand-size limeclasts, lumps, algal material, and some quartz sand grains. Bird Spring Formation (Pennsylvanian), Lee Canyon area, Clark County, Nev.; x 3 . C. Detrital limestone, composed of silt- and sand-size limeclasts, with some skeletal grains, in finer-grained matrix. Ely Limestone (Pennsylvanian), Burbank Hills, Millard County, Utah; x 3. D. Detrital limestone, composed of oolites, pellets, limeclasts, and some quartz grains, with minor interstitial lime mud. Morrowan age part of Morgan Formation, Duchesne River Valley, Duchesne County, Utah; x 4.6. E. Calcarenitic limestone, composed of sand-size limeclasts, crinoid ossicles, some algal material, and subangular to subrounded quartz grains, embedded in matrix of finer-grained material. Derryan age portion of Wood River (?) Formation, Arc0 Hills, Butte County, Idaho; x3. F. Calcarenite showing graded bedding. Rock is composed of limeclasts, algal material, pellets, and micro-lumps, with quartz sand grains. Cedar Fort Member @esmoinesian age) of Oquirrh Formation, Cedar Mountains, Tooele County, Utah; x 3. PLATE I1 A. Micritic limestone. Rock is composed of very fine-textured to micritic limestone with minor amount of skeletal material. Ely Limestone (Pennsylvanian), Pancake Range, Nye County, Nev.; x 30.8. B. Detrital limestone, composed of abraded and broken brachiopod shells, crinoid ossicles, algal material, lumps, and bryozoan fragments in finer textured particulate lime material and micritic material. Hall Canyon Member (Morrowan age) of Oquirrh Formation, Oquirrh Mountains, Utah County, Utah; x 7.7. C. Skeletal-detrital limestone. Fusulinids (Triticitessp.) in silt- and sand-size particulate limestone, algal material, pelletal material; and quartz silt and sand grains. Missourian portion of Oquirrh Formation, Hobble Creek Canyon, Utah County, Utah; x 7.7. D. Skeletal-detrital limestone, composed of disarticulated brachiopod tests, limeclasts, in matrix of calcilutite containing some quartz silt and sand grains. Garden Valley Formation (Permian), Diamond Range, White Pine County, Nev.; X 3. E. Detrital limestone, composed of sand- and silt-size limeclasts, pelletal material, and some quartz silt and sand grains. Ely Limestone (Pennsylvanian), Leppy Range, northeast of Wendover, Tooele County, Utah; x 3.85. F. Detrital limestone, composed of interlayered calcilutite and calcarenite, with fractures filled by sparry calcite. Ely Limestone (Pennsylvanian), Pancake Range, Nye County, Nev.; x 7.7.

PLATE I11

A. Calcarenite, slightly quartzose. Rock is composed of limeclasts, algal and pelletal material in fine-grained quartzose-calcareous material, and with subrounded quartz grains also present. Derryan age portion of Weber Formation at Pullem Creek, Wasatch County, Utah; x3.85. B. Micritic limestone (calcilutite), composed of lime mud in which quartz silt grains and some silt-size limeclasts are embedded. Virgin Limestone Member (Lower Triassic) of the Moenkopi Formation, Blue Diamond Mountain, Clark County, Nev.; x 61.6. C. Calcisiltite with calcilutite layers. Rock consists of silt-size limeclasts and lime mud, with silt-size and very fine-grained sand-size quartz particles. Summit Springs Member (Medial

122

H. J . BISSELL AND G. V . CHILINGAR

Leonardian age) of Pequop Formation, Southern Butte Mountains, White Pine County, Nev.; x23.1.

D. Dolomicrite with very fine-grained quartz and with few negative fossil relics. Wolfcampian age portion of Weber Formation, Morris Ranch area on south side of Uinta Mountains, Uinta County, Utah; x 61.6. E. Dolosiltite with admixed quartz silt grains. Rock consists of cross-stratified particulate siltsize dolomite, and some interlayered dololutite. Leonardian age portion of Spring Mountains Formation, west side of Kyle Canyon in Spring Mountains, Clark County, Nev.; x 30.8. F. Primary dolomicrite with interlayered fine-grained dolosiltite and some silt-size quartz grains. Incipient diagenesis has occurred in some layers. Leonardian age portion of Spring Mountains Formation, Love11 Wash area in Spring Mountains, Clark County, Nev.; x 61.6

PLATE IV A. Biocalcarenite (= criquinite, or encrinal limestone). Rock consists of crinoid ossicles, bryozoan fragments, and some bladed Algae, with interstitial material (= matrix) composed of sparite and calcisiltite. Riepe Spring Limestone (Wolfcampian age), Moorman Ranch area, White Pine County, Nev.; ~ 4 . 6 . B. Skeletal-detrital limestone, consisting of abraded fragments of crinoid ossicles, bryozoan fragments, pelletal material, and algal material; interstitial material is calcisiltite. Virgilian age portion of Wood River Formation near Hailey, Blaine County, Idaho; x 7.7. C. Reefal limestone with admixed skeletal and detrital material. Rock consists of bryalgal material enclosing tests of schwagerinid fusulinids, crinoid ossicles, algal pellets, and minor limeclasts. Lime mud fills open spaces. Pequop Formation (Leonardian age), Diamond Range, White Pine County, Nev.; ~ 7 . 7 . D. Skeletal-detrital limestone with admixed quartz sand grains. Rock consists of pelecypod shell fragments, few crinoid ossicles, limeclast lumps, and algal material. Virgin Limestone Member (Lower Triassic) of Moenkopi Formation, Bird Spring Range (Cottonwood Pass area north of Goodsprings), Clark County, Nev.; x 7.7. E. Skeletal-detrital limestone, composed of the alga Mizzia sp. with intertwined bryozoans and bladed Algae (= bryalgal limestone), and with admixed silt-size limeclasts. Pequop Formation (Leonardian age), Cherry Creek Range, Elk0 County, Nev.; X 23.1. F. Diagenetically-altered biocalcarenite. Rock consists of crinoid ossicles and intertwined bryozoans with interstitial calcilutite. Rim cement, impingement, and secondary overgrowth are present. Toroweap Formation (Permian), Blue Diamond Mountain, Clark County, Nev. ; x 7.7. PLATE V A. Skeletal-detrital limestone. Rock is composed of fossiliferous-fragmentalmaterial, consisting of lithoclastic carbonate detritus admixed with bioclastic detritus of bryozoans, encrinal material, and algal fragments. Garden Valley Formation (Permian), Diamond Range, White Pine County, Nev.; ~ 3 . 8 5 . B. Skeletal-micrite. Tests of Triticites sp. and some crinoid ossicles (some replaced by sparite), embedded in micrite. Ferguson Mountain Formation (Wolfcampian, Permian), Ferguson Mountain, Elko County, Nev.; x 7.7. C. Calcarenite. Rock consists of crinoid ossicles (some of which are slightly altered diagenetically), with interstitial space filled with calcilutite and calcareous cement. Bridal Veil Falls Member (Morrowan age) of Oquirrh Formation, Provo Canyon, Utah County, Utah; x 3.85. D. Detrital limestone. Rock consists of slightly abraded tests of the fusulinid Pseudoschwugerina sp. in a matrix of organic-rich calcisiltite. Mid-portion of the Ferguson Mountain Formation (Wolfcampian, Permian), Ferguson Mountain, Elko County, Nev.; x 4.6. E. Skeletal limestone. Consists mostly of “tailings” of pelecypod valves, with interstitial space filled by calcarenite and sparite cement. Thaynes Formation (Triassic), Aspen Range, Caribou County, Idaho; X 1.1.


123

F. Skeletal-pelletal limestone. Rock consists of brachiopod shells, crinoid ossicles, fragments of bryozoans, and other skeletal detritus, along with pellets; all in sparite. Brazer Formation (Mississippian), Wellsville Mountain, Boxelder County, Utah; x 3.85. PLATE VI A. Skeletal-detrital limestone, consisting of particulate bryozoans, encrinal material, and probable algal fragments, in part surrounded by sparite and in part filled by sparite. Toroweap Formation (Permian), Spring Mountains, Clark County, Nev.; x 3 . B. Skeletal limestone, consisting of bioclastic detritus, algal and lithic carbonate pellets, brachiopod shells, bryozoan fragments, and minor quantity of millerellid Foraminifera. Hall Canyon Member (Morrowan) of Oquirrh Formation, east of Manning Canyon in Oquirrh Mountains, Utah County, Utah; x 7.7. C. Skeletal-detrital limestone, composed of fossiliferous-fragmental (i.e., bioclastic) detritus, mostly crinoid ossicles and bryozoan material. Calcarenite and calcilutite fill interstitial space. Ely Formation (Pennsylvanian), North Burbank Hills, Millard County, Utah; x 7.7. D. Pzlletal bryalgal limestone. Rock consists of algal and other pelletal debris enmeshed in framework of bryozoans and thread-like to blady Algae which form a rigid framework. Pequop Formation (Leonardian, Permian), northern Cherry Creek Range, Elko County, Nev.; x 7.7. E. Skeletal limestone, consisting of bioclastic debris, largely lioclemid bryozoan fragments, some crinoid ossicles, and minor algal detritus. Dead oil is present in part of the interstitial space; micrite comprises the remainder of the rock. Gerster Formation (Guadalupian, Permian), Medicine Range, Elko County, Nev.; x 7.7. F. Micrite with skeletal material. This rock consists of bryozoans, ostracods, algal pellets, algal filaments, small brachiopods, and lithic carbonate debris in oil-rich micrite. Ferguson Mountain Formation (Wolfcampian, Permian), Dolly Varden Mountains, Elko County, Nev.; x 7.7. PLATE V11

A. Skeletal limestone, consisting of pelecypod shell fragments, some algal filaments and pellets, and calcarenitic interstitial material. Sparry calcite fills some interstitial space. Thaynes Limestone (Triassic) on Dry Ridge, Lanes Creek quadrangle, Caribou County, Idaho; x 3. B. Calcarenite with skeletal material. Rock consists of silt- and sand-size limeclasts and some quartz sand grains, surrounding tests of schwagerinid fusulinids and some crinoid ossicles. Carbon Ridge Formation (Permian), Carbon Ridge, Eureka County, Nev.; ~ 4 . 6 . C. Pelletal-skeletal limestone, composed of algal pellets, limeclasts, bryozoan material, and some brachiopod shells (filled with sparite). Gerster Formation (Permian), Medicine Range, Elko County, Nev.; x 3.85. D. Lumpal (= lump) limestone, consisting of limeclasts, few crinoid ossicles, algal pellets, and Foraminifera (including fusulinids); in calcisiltite matrix. Bird Spring Formation (Pennsylvanian), Spring Mountains west of Mountain Pass, Clark County, Nev.; x 7.7. E. Micrite with skeletal material. Rock consists of silty micrite containing brachiopod shell fragments that have sparite fillings. Hall Canyon Member (Morrowan age) of Oquirrh Formation, Fivemile Pass area, Utah County, Utah; ~ 2 3 . 1 . F. Sparite after skeletal limestone. Rock consists of straparollid gastropods and some crinoid ossicles, with replacement and infilling of sparite. Gerster Formation (Permian), Currie Hills, Elko County, Nev.; x 3. PLATE VIII A. Pelletal calcarenite, consisting of algal pellets, limeclasts, and silt- to sand-size quartz grains. Wolfcampian age portion of Oquirrh Formation, Right Fork of Hobble Creek Canyon, Utah County, Utah; x7.7.

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B. Skeletal calcisiltite. Rock is composed of tests of triticitid fusulinids, few limeclasts, and rare algal pellets, all in matrix of petroliferous calcisiltite. Ferguson Mountain Formation (Wolfcampian, Permian), Ferguson Mountain, Elko County, Nev.; x 7.7. C. Encrinal limestone (= criquinite), composed of abraded crinoid ossicles, rare limeclasts, in matrix of calcisiltite that contains few quartz grains. Encrinal material has been diagenetically altered to a slight degree. Morgan Formation (Pennsylvanian), Weber Canyon, Morgan County, Utah; ~ 7 . 7 . D. Skeletal calcisiltite, consisting of brachiopod “tailings” (= fragniental) that are in part replaced by sparite, in organic rich calcisiltite that contains few fine-textured quartz sand grains. Bridal Veil Falls Member (Morrowan age) of Oquirrh Formation, Provo Canyon, Utah County Utah; ~ 7 . 7 . E. Diagenetically altered skeletal limestone. Rock is composed of algal (?)pellets, crinoid ossicles, bryozoan fragments, and abraded brachiopod shells, all in various stages of replacement by sparite. Toroweap Formation (Permian), Star Range, Beaver County, Utah; x 7.7. F. Calcisiltite with Algae. Rock consists of calcisiltite enclosing the alga Solenopora sp., and smaller algal pellets. Upper Member (Late Leonardian age) of Pequop Formation, Southern Butte Mountains north of Moorman Ranch, White Pine County, Nev.; x 7.7. PLATE IX

A. Sparite with skeletal material. Rock consists of medium- to coarsely-crystalline sparite enclossing crinoid ossicles, brachiopod fragments, bryozoan material, and millerellid Foraminifera. Morgan Formation (Pennsylvanian) at type locality in Weber Canyon, Morgan County, Utah;

x23.1. B. Pelletal-skeletal limestone, composed of fossiliferous-fragmental material, algal pellets and limeclasts, Foraminifcra, in sparite matrix. Brazer Formation (Upper Mississippian), Wellsville Mountain near Deweyville, Boxelder County, Utah; X 7.7. C. Pellet limestone, composed of algal pellets and limeclast pellets, encrinal material, bryozoan fragments, and silt- to sand-size quartz grains. Morrowan age portion of Oquirrh Formation, Gilson Mountain area, Juab County, Utah; x 7.7. D. Skeletal limestone showing effects of diagenesis. Rock consists of fragments of bryozoans, rare crinoid ossicles, and few pellets, with interstitial material composed of calcisiltite. Kaibab Limestone (Permian), Gold Hill District, Tooele County, Utah; x 7.7. E. Coated grains. Rock consists of oolites, pellets, algal plates, and limeclasts in micrite. Virgin Limestone Member (Lower Triassic) of Moenkopi Formation, east of Blue Diamond Mountain, west of Las Vegas, Clark County, Nev.; ~ 7 . 7 . F. Coated grains in sparite. Rock is composed of oolites showing both radial and concentric structure, skeletal material with coatings, and few limeclasts in coarsely crystalline sparite. Lodgepole Limestone (Mississippian), Western Judith Mountains, Fergus County, Mont.; x 7.7,

PLATE X A. Pelletal-detrital limestone, consisting of algal pellets, limeclast pellets and limeclast detrital material, lumps, and scattered abraded skeletal material; much of the material has been diagenetically altered. Morgan Formation, west end of Uinta Mountains, Wasatch County, Utah; x 7.7. B. Lump limestone in sparite. Rock consists of organic-rich limeclast lumps (and possibly algal material) surrounded by coarsely-crystalline sparite. Toroweap Formation (Permian), Bird Spring Mountains west of Arden, Clark County, Nev.; X 15.4. C. Coated grains in sparite. Rock is composed of skeletal and limeclast grains with oolite overgrowths, all in medium-crystalline sparite. Brazer Formation (Mississippian), Weber River Canyon, Morgan County, Utah; ~23.1. D. Pellet limestone, composed of spindle-shaped and rod-shaped pellets (for the most part filled with sparite) in interstitial material of Algae and calcarenite. Pequop Formation (Permian), central part of Cherry Creek Range, White Pine County, Nev.; X23.1.


125

E. Coated grains in micrite. Rock consists of oolites showing both concentric and radial structure; some are distorted. Interstitial material is largely micrite, with local patches of sparite. Hannah Formation (Mississippian), Saypo quadrangle in Sawtooth Mountains, Teton County, Mont.; % 30.8.

F. Lump limestone in sparite. Rock consists of limeclast lumps, possible algal material, algal pellets; in medium-crystalline sparite. Leonardian age portion of Spring Mountains Formation, northern Spring Mountains, Clark County, Nev.; x 3.85. PLATE XI A. Biogenic limestone, consisting of organic frame-building alga, Kumia sp. (by some referred to the stromatoporoids). Interstices consist of algal pellets and limeclasts. Derryan age portion of Wells Formation, South Schmid Ridge, Dry Valley quadrangle, Caribou County, Idaho; x 7.7. B. Diagenetically altered skeletal limestone, consisting of schwagerinid fusulinid tests, algal filaments and plates, bryozoans, and encrinal material, altered in varying degrees to sparite. Pequop Formation (Leonardian age), west end of Leppy Range, Elko County, Nev.; x 7.7. C. Algal limestone (biogenic) in micrite. Rock consists of the alga Mizzia sp. (some surrounding skeletal elements), in micritic limestone in which fine quartz silt grains and occasional crinoid ossicles are present. Pequop Formation (Leonardian age), low hills east of Lund, White Pine County, Nev.; x 15.4. D. Organic-rich micritic limestone with Strornatactis (?).Rock consists of petroliferous limestone containing elongate bodies referred with query to Stromatactis. Riepe Spring Limestone (Wolfcampian, Permian), Rib Hill near Ruth, White Pine County, Nev.; ~ 2 3 . 1 . E. Biogenic-skeletal limestone, consisting of alga (or stromatoporoid) Komia sp., fusulinid tests, algal pellets, crinoid ossicles and limeclasts. Interstitial space partially filled with sparite and remainder by organic-rich calcilutite. Derryan age portion of Ely Limestone, South Schell Creek Range south of Patterson Pass, Lincoln County, Nev.; x 15.4. F. Biogenic limestone in sparite, consisting of the alga Osugia sp., surrounding limeclasts, bryozoans, and other material; in sparite. Some sparite has filled open spaces in skeletal material. Rogers Spring Limestone (Redwall? Limestone), Star Range, Beaver County, Utah; x 15.4. PLATE XI1 In place, organic frame-builders (biogenic). A. Reefal limestone, composed of coraIs, Algae, and sponges, with micrite in interstitial space. Permian of Guadalupe Mountains, Texas; x 7.7. B. Bryalgal limestone, consisting of various bryozoans intermeshed with Algae, and infilled with lime mud. Leonardian age portion of Pequop Formation, Maverick Spring Range, White Pine County, Nev.; x 77. C. Reefal limestone, consisting of brachiopods, Algae, and bryozoans, enclosed in micrite. Garden Valley (?) Formation, east side of Diamond Range, White Pine County, Nev.; x7.7. D. Reefal limestone, composed of bryozoans, sponges, spicular material, and few brachiopod fragments enclosed in lime mud that is organic rich. Leonard Formation, Glass Mountains, Texas; x6.2. E. Reefal limestone, composed of lioclemid bryozoans, Algae, Foraminifera, and crinoid ossicles. Largely fore-reef in situ accumulation. Pequop Formation (Leonardian age), Spruce Mountain area, Elk0 County, Nev.; x 3.85. F. Reefal limestone, consisting of bryalgal material, some of which has undergone diagenetic change to sparite and microsparite. Gerster Limestone (Guadalupian age), Medicine Mountains, Elk0 County, Nev.; x 7.7.

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PLATE XI11 Diagenesis. A. Dolomitized calcilutite; rock consists of dolomite secondary after calcilutite. Original sediment contained variable amount of organic matter and was stratified. Summit Springs Member (Medial Leonardian age) of Pequop Formation, west end of Leppy Range, Elk0 County, Nev.; x 30.8. B. Crystalline criquinite; rock consists of coarsely crystalline material, the fabric having resulted from diagenesis of an encrinal limestone, with the original texture almost completely obliterated in the process. Some algal material is relatively unaltered. Lower part of Kaibab Limestone, low hills east of Ruby Marshes, Elko County, Nev.; ~ 7 . 7 . C. Dolomitized biomicrite, consisting of dolomitized algal filaments and plates (now resembling a form of Stromutuctis), in partially dolomitized micritic matrix. Leonardian age part of Spring Mountains Formation, west side of Kyle Canyon, Clark County, Nev.; x 30.8. D. Dolomitized micrite, consisting of porous rock that originally was a micrite and now is a dolomite with substantial porosity and permeability. Pakoon Formation, east side of Frenchman Mountain, Clark County, Nev.; ~ 4 6 . 6 . E. Dolomitized skeletal limestone, consisting of medium- to coarsely-crystalline dolosparite, secondary after bryozoans, encrinal material, algal filaments and pellets, and possibly Foraminifera. Both positive and negative relics are present. Toroweap Formation, south of Garnet R.R. Siding, Clark County, Nev.; x 15.4. F. Dolosparite, consisting of mosaic of dolospars, secondary after limeclast limestone. Thin unit in Guilmette Limestone (Devonian), Toana Range north of Whitehorse Pass, Elko County, Nev.; X 15.4.

PLATE XIV A. Dolosparite, consisting of cloudy sparry dolomite, secondary after calcarenitic limestone with impure matrix. Lower Kaibab Formation (Permian), Maverick Spring Range, northern White Pine County, Nev.; x 30.8. B. Diagenetically-altered skeletal limestone; rock consists of calcite spar developed in a stratified skeletal-detrital limestone. Unnamed Wolfcampian age (Permian) carbonate sequence, Arc0 Hills, Butte County, Idaho; x 23.1. C. Incipient dolomitization of a calcarenite. Rock is composed of subhedra to imperfect euhedra of dolorhombs and less-altered sparry calcite, with relatively unaltered calcarenite matrix. Loray Formation (Permian), Dead Horse Wash, west of Egan Range, White Pine County, Nev.; x 23.1. D. Sparite, some of which is diagenetically altered to subhedra of dolospar. Rock originally was a calcarenite with impure matrix, and now is a mosaic of sparite and dolosparite. Wolfcampian age part of Weber Formation, Duchesne River area, south flank of Uinta Mountains, Duchesne County, Utah; x23.1. E. Diagenetically dolomitized skeletal limestone, consisting of vary-grained (vary-crystalline) sparry calcite and sparry dolomite, anhedral to subhedral, with patches of less altered material. Original texture strongly altered, some obliterated, but with negative skeletal relics (= crinoid ossicles and bryozoans). Plympton Formation (Permian), Gold Hill district, Tooele County, Utah; x15.4. F. Coarsely crystalline sparite, consisting of calcite spar; no dolomite. Chilliwack Group (Permian), Cascade Mountains, Wash.; ~ 2 3 . 1 .

PLATE XV A. Diagenetically-altered calcilutite, showing development of “eyes” of silicified material (possibly replaced algal material) in dololutite. Some quartz sand grains are present. Park City Group (Permian), near “The Thumb”, 10-12 miles north of Knolls, Tooele County, Utah; x15.4. B. Slightly altered skeletal-detrital liniestone, showing positive relic of fusulinid test, negative

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127

skeletal material, and limeclasts set in altered matrix. Upper Member (Late Leonardian age) of Pequop Formation, west end of Leppy Range, Elko County, Nev.; x 15.4. C. Dolomitized skeletal-detrital limestone, showing almost complete obliteration of original fabric; spots or “eyes” of dolospar are the lighter colored areas, and darker colored spots may be relics of algal pellets. Matrix is dololutite, secondary after limestone. Kirkman Limestone (Wolfcampian age), South Tintic Mountains, Juab County, Utah; ~ 2 3 . 1 . D. Dolomitized skeletal-detrital limestone, showing negative and positive relics of skeletal material, in matrix of calcisiltite and dolosiltite. Grandeur Member of Park City Formation (Permian), Southern Wasatch Mountains, Juab County, Utah; x 15.4. E. Highly-altered diagenetic dolomite, showing negative relics of material not determinable as organic or inorganic, in matrix of less altered dolosiltite. Few relatively unaltered quartz sand grains are present. Plympton Formation (Permian), Granite Mountain area of Northern Confusion Range, Juab County, western Utah; x 15.4. F. Moderately advanced diagenesis in skeletal limestone, showing positive relics of bryozoans, and the stromatoporoid Komia sp. Ely Limestone (Pennsylvanian), west side of Pequop Mountains, southwest of Shafter, Elko County, Nev.; x 15.4. PLATE XVI

A. Sucrosic dolomite, diagenetic after limestone. Rock shows considerable porosity and dead oil in vugs. Wolfcampian age portion of Weber Formation at canyon of Duchesne River, Duchesne County, Utah; x 30.8. B. Diagenetically-altered skeletal-calcarenite, showing first-stage crystallization, but not dolomitization, of limeclasts and algal to pelletal material. Wolfcampian age part of Oquirrh Formation at South Mountain west of Stockton, Tooele County, Utah; x23.1. C. Advanced diagenesis including partial dolomitization of a skeletal-detrital limestone. Rock shows sparite and dolosparite in mosaic of subhedra and euhedra. Vuggy porosity has been developed and dead oil is present. Lower Member of Kaibab Formation, south of Cottonwood Wash and east of Keystone Thrust, Clark County, Nev.; ~ 2 3 . 1 . D. Diagenetic dolomite, secondary after limestone. Rock shows mosaic of anhedra and subhedra of dolomite, with remnants of sparite. Toroweap Formation (Permian), Northern Muddy Mountains just southwest of Glendale Junction, Clark County, Nev.; x 7.7. E. Dolosparite. Rock shows advanced-stage diagenetic dolomitization of an earlier-formed limestone. Well-developed dolorhombs (some by impingement) have formed a fabric having low to moderate intercrystalline porosity. Fairly high permeability. Wolfcampian age part of Weber Formation on south flank of Uinta Mountains near Morris Ranch, Uinta County, Utah; x 38.5. F. Diagenetic dolomite, secondary after encrinal limestone. Rock shows mosaic of anhedra and subhedra of dolomite, with remnants of calcarenite. Suturing has occurred at crystal-grain boundaries. Summit Springs Member (Medial Leonardian Permian) of Pequop Formation, Ferguson Flat, Elko County, Nev.; x 7.7.

PLATE 1

Legend see p.121.

PLATE I1

Legend see p.121

129

PLATE 111

Legend see p. 12 1.

PLATE IV

Legend see p.122.

131

PLATE V

Legend see p.122.

PLATE V1

Legend see p.123.

133

PLATE VII

Legend see p. 123.

PLATE VIII

Legend see p.123.

135

PLATE IX

Legend see p.124.

PLATE X

Legend see p.124.

137

PLATE XI

Legend see p . 125.

PLATE XI1

Legend see p.125.

139

PLATE XI11

Legend see p.126.

PLATE XLV

Legend see p.126.

141

PLATE XV

Legend see p.126.

PLATE XVI

Legend see p.127.

143

I44


REFERENCES

ALDERMAN, A. R. and SKINNERH. C. W., 1957. Dolomite sedimentation in the southeast of South Australia. Am. J. Sci., 255: 561-567. ASCHENBRENNER, B. C. and ACHAUER, C. W., 1960. Minimum conditions for migration of oil in water-wet carbonate rocks. Bull. Am. Assoc. Petrol. Geologists, 44(2): 235-243. BAARS,D. L., 1963. Petrology of carbonate rocks. In: R. 0. BASSand S. L. SHARPS (Editors), Shelf Carbonates of the Paradox Basin, A Symposium-Four Corners Geol. Soc., Field Con$, 4th, pp.101-129. BATHURST, R. G. C., 1958. Diagenetic fabrics in some British Dinantian limestones. Liverpool Manchester Geol. J., 2(1): 11-36. BATHURST, R. G. C., 1959. Diagenesis in Mississippian calcilutites and pseudobreccias. J. Sediment. Petrol., 29: 365-376. BEALES, F. W., 1956. Conditions of deposition of Palliser (Devonian) Limestone of southwestern Alberta. Bull. Am. Assoc. Petrol. Geologists, 40: 848-870. BEALES, F. W., 1958. Ancient sediments of Bahaman type. Bull. Am. Assoc. Petrol. Geologists, 42: 1845-1880. BERRY,W. B. N., 1962. Comparison of some Ordovician limestones. Bull. Am. Assoc. Petrol. Geologists, 46: 1701-1720. BISSELL, H. J.; 1964. Lithology and petrography of the Weber Formation in Utah and Colorado. In: E. F. SABATKA (Editor), Guidebook to the Geology and Mmeral Resources of the Uinta Basin-Ann. Field Con5 Zntermountain Assor. Petrol. Geologists, 13th, pp.67-91. BISSELL,H. J. and CHILINGAR, G. V., 1958. Notes on diagenetic dolomitization. J. Sediment. Petrol., 28: 490498. BISSELL, H. J. and CHILTNGAR, G. V., 1961. Comments on Teodorovich’s “Structural classification of limestones and dolomites”, and Shvetsov’s “Genetic classification of limestones.” J . Sediment. Petrol., 31 : 61 1-616. BISSELL, H. J. and CHILINGAR, G. V., 1962. Evaporite type dolomite in Salt Flats of western Utah. Sedimentology, 1 : 200-2 10. BOUMA,A. H., 1962. Sedimentology of some Flysch Deposits. A Graphic Approach to Facies Interpretation. Elsevier, Amsterdam, 168 pp. BRAMKAMP, R. A. and POWERS,R. W., 1958. Classification of Arabian carbonate rocks. Bull. Geol. Soc. Am., 69: 1305-1318. CAROZZI,A. V., 1960. Microscopic Sedimentary Petrography. Wiley, New York, N. Y., 485 pp. CAROZZI, A. V. and SODERMAN, J. G. W., 1962. Petrography of Mississippian (Borden) crinoidal limestones at Stobo, Indiana. J. Sediment. Petrol., 32: 397-414. CAYEUX,L., 1935. Les Roches Sidimentaires de France, “Roches Carbonaties” (Calcoires et Dolomies). Masson, Paris, 463 pp. CHANDA, S. K., 1963. Cementation and diagenesis of the Lameta Beds, Lametaghat, M.P., India. J. Sediment. Petrol., 33: 728-738. K. E., 1954. Aspects of the biochemistry of magnesium. 2. Calcareoussediments and rocks. CHAVE, J. Geol., 62: 587-599. CHILINGAR, G. V., 1953. Use of Ca/Mg ratio in limestones as a geologic tool. The Compass of Sigma Gamma Epsilon, 30: 202-209. CHILINGAR, G. V., 1955. Soviet classification of carbonate rocks and chemically precipitated CaC03-A review. Bull. Am. Assoc. Petrol. Geologists, 39: 1885-1 889. G. V., 1956a. Note on direct precipitation of dolomite out of sea water. The Compass CHILINGAR, of Sigma Gamma Epsilon, 34: 29-34. CHILINGAR, G. V., 1956b. Dedolomitization-A review. Bull. Am. Assoc. Petrol. Geologists, 40: 762-764. CHILINGAR, G. V., 1956c. Relationship between Ca/Mg ratio and geologic age. Bull. Am. Assoc. Petrol. Geologists, 40: 2256-2266. CHILINGAR, G. V., 1956d. Use of Ca/Mg ratio in porosity studies. Bull. Am. Assoc. Petrol. Geologists, 40: 2489-2493.

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GLOSSARY]

Accretionary: added through overgrowth upon a pre-existing grain and/or organic structure or framework; may be concentric, and may originate through rolling. Aggregation: added through coagulation, flocculation, adherance or other method to result in a composite grain or structural mass. May be algal, foraminiferal, algal-dust, or other materials aggregated together or to another mass, fossil, or grain. May be a method of forming lumps. Algal: relating in any manner to algal limestones or dolomites; more commonly the Rhodophyta or red Algae and some genera of the Cyanophyta or blue-green Algae. Algal dust: a term designated to describe micro-textured to finely textured dark-colored (usually brown and brown-gray) micritic and matrix material in carbonates which also contain discernible algal remains. Allochthonous: an accumulation of transported material; for sedimentary carbonate rocks this amounts to transported grains, fossil fragments, rock fragments, and organic matter into the depocenter from adjacent area(s). Allogenic: term meaning generated elsewhere; in a manner similar to allochthonous, in that the constituents came into existence outside of, and previous to, the rock of which they are now a part. For example, quartz sand blown from land into an evaporite sequence; gravel washed into a basin and mixed with lime mud. Anadiagenetic: this term applies to dolomitization (secondary) that occurs under considerable burial, including most tectonic dolomitization (cf. syndiagenetic and epidiagenetic). (After FAIRBRIDGE, 1966.) Anhedral: individual crystals devoid of crystal boundaries or faces; particularly applicable to dolomites and dolomitized limestones of certain varieties. Aphanic: term to describe the texture of most micritic limestones and dolomicrites; individual crystals and grains less than 0.01 mm in size. A useful field term to describe lithographic and sub-lithographic carbonates. To be used in lieu of uphunitic. Articulate: refers to fossils having two or more parts joined together in their natural relationship; for example, valves of brachiopods or molluscs, fronds of bryozoans, columnals of crinoids, etc. Authigenic: generated on the spot; refers to those constituents that came into existence with or after the formation of the host rock. For example, albite that forms in some limestones, calcite or dolomite rhombs over grains, etc. Autochthonous: pertaining to objects that originated in the places where they now occur, and therefore are in situ; examples are frame-building organisms such as corals, algal masses, various bioherms, etc. Autoclastic breccia: refers to a common structure observed in dolomites; a result of diagenetic shrinkage followed by recementation. 1 To

be used as an aid in describing sedimentary carbonate rock terms. It is not intended to cover all definitions of the rock names.


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Axiolitic: a type of elongated or subspherical oolite or pisolite in which the radial structure(usually acicular needles) develops toward the periphery at right angles to a central axis; may have superimposed concentric structure. Bahamite: name of granular limestones that closely resemble the present deposits of the interior of the Bahama Banks. The texture varies from calcisiltites to calcirudites, in which the grains are accretionary and commonly are composite, consisting of smaller granules bound together by precipitated material into aggregate grains.

Bank: an in situ skeletal limestone deposit formed by organisms which do not have the ecological potential to erect a rigid, wave-resistant structure. Beach-rock: a friable to well-cemented beach sediment consisting of calcareous debris (detrital and/or fragmental skeletal) cemented by calcium carbonate. Bioaccumulated: applied to limestone deposits formed by sedentary but noncolonial organisms and their related ecologic communities. Characterized by a predominance of unbroken fossils, diverse organic components, lack of (or poor) sorting, and scarce fine-grained matrix. Biocalcarenite: special group of calcarenites, mechanically deposited carbonate detritus of sand size (1 /16-2 mm in diameter); predominantly composed of organic fragments of any kind, the products of organic activity which have an internal structure such as foraminifera], algal, and faecal pellets, and recognizable fossil detritus worn to sand size. Numerous encrinal limestones fit in this category. Biocalcirudite: rudaceous, sedimentary carbonate rock, the discrete particles of which are composed of fragmental fossiliferous material; individual particles are larger than 2 mm in diameter. Almost any type of reef-building organisms such as stromatoporoids, branched corals, and calcareous Algae can be broken away, worn to a variable degree, and concentrated into biocalcirudites. Bioclastic: a clastic sedimentary carbonate rock which owes its essential character to organisms (GRABAU, 1913). In view of disparity of definition, it is herein suggested that this term be expanded to embrace fossil detritus that is largely intraclastic, whether of rudaceous, arenaceous, or lutaceous texture, and which originates (or originated in the past) largely by being broken and transported by water currents and waves before coming to rest. Bioclastic limestone, therefore, may be poorly sorted, moderately sorted, or well sorted; and it may be clean or have a matrix of finer detritus. Bioconstructed: term applied to limestone deposits resulting from the vital activities of colonial and sediment-binding organic communities. Algal, bryozoan, bryalgal, stromatoporoidal, coralline, and coralgal colonies are predominant. Biogenic: sedimentary carbonate rock, a deposit of organic material or materials formed through the physiological activities of the organisms. Bioherm: an organic reef or mound built by corals, stromatoporoids, gastropods, echinoderms, Foraminifera, pelecypods, brachiopods, Algae, and other organisms. It is a reef, bank, or mound that is reeflike, moundlike, lenslike or an otherwise circumscribed structure of strictly organic origin, embedded in rocks of different lithology.

Biolithite: a limestone characterized by an organic framework of carbonate laminae that bind grains and skeletal elements as a rigid framework; typical of the cores of some organic reefs.

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Biomicrite: a major group of biogenic limestones containing a significant admixture of fine-textured carbonate material filling the spaces between organic tests and fragments. Biosparite: a foraminifera1 limestone composed largely of tests of bottom-dwelling and floating organisms, and lacking a fine-textured matrix; mostly crystalline. Biostrome: a term for stratiform deposits, such as shell beds, crinoid beds, and coral beds, consisting of, and built mainly by, organisms or fragments of organisms (mostly sedentary), and not swelling into moundlike or lenslike forms. Birdseye: spots or tubes of sparry calcite in limestones (and some dolomites). These “calcite eyes” are common to pelsparites, and may have resulted from one of the following (or certain combinations thereof): ( I ) precipitation of sparry calcite in animal burrows, or in worm tubes; ( 2 ) soft-sediment slumping or mud cracking; (3) precipitation of sparry calcite in tubules resulting from escaping gas bubbles; (4) re-working and rapid deposition of soft sediment containing semicoherent clouds of calcareous mud and spar; (5) recrystallization of calcareous (or dolomitic) mud in patches; and ( 6 ) “arrested” dolomitization. Boundstone: applies to most reef rock, stromatolites, and some biohermal and biostromal rocks in which the original components were bound together during deposition, and remain substantially in position of growth. Breccia (sedimentary) : a rock composed of consolidated angular fragments, most of which are larger than 2 mm in diameter, plus matrix and/or cementing material. CAROZZI (1960) mentioned a crystallization breccia that resulted from the differentiation in place of a homogeneous calcilutite. Crystallization began at numerous points scattered throughout thc rock but was incomplete, and as a result the crystallized patches appear as fragments In a groundmass that was spared by the process. Bryalgal: a term for limestones composed largely of materials constructed in situ by organic frame-building bryozoans and Algae; the word is a contraction of bryozoan-algal framebuilding organisms. Resultant deposits range in thickness from thin units to biostromes, bioherms, patch-reefs, and larger reefs. Calcarenite: a mechanically deposited carbonate rock consisting of sand-size carbonate grains (1 /16-2 mm in diameter); the particulate material in this rock may be of lithoclastic and/or bioclastic derivation, and comprises 50 % or more of the rock. Calcilutite: by decrease in grain size, a calcarenite grades through a calcisiltite into a calcilutite, thus forming a rock composed of 50% or more of clay- (plus some silt-) size carbonate particles; includes biocalcilutites and lithocalcilutites. Calcirudite: the term is used as a general designation for mechanically deposited carbonate rocks that are composed of 50% or more of angular to rounded fragments over 2 mm in diameter, and have matrix and/or cementing material. Culcisiltite: a rock type intermediate between calcarenite and calcilutite, in that it consists mostly of silt-size carbonate detritus that comprises 50% or more of the rock; includes biocalcisiltites and lithocalcisiltites. Calclithite: a limestone containing 50 % or more of fragments of older limestone(s) that experienced erosion and redeposition. The individual particles are termed extraclasts. Calcsparite: see sparite.

Caliche: it is a lime-rich deposit found in soils and is formed by capillary action drawing the lime-


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bearing waters to the surface where, by evaporation, the lime is precipitated. In bajadas, intermonts, alluvial fans and colluvium of parts of the Great Basin of the western United States some of the caliche deposits are dolomitic due to presence of extensive dolomite rubble. Caliche, whether calcareous and/or dolomitc, also cements alluvial fans to form Janglomerale. Carbonate: rocks composed of more than SO%, by weight, of carbonate minerals. For practical microscopic work, area percentages, which approximate weight percentages, are used because they are easier to estimate and measure. Cement: clear to opaque, crystalline material occurring in the interstices between grains and matrix material, or between grains. It may be sparry calcite and/or dolomite, and thus is termed sparite; but more commonly, it is smaller than 0.03 mm in crystal size. Cement commonly is chemically precipitated material into voids and in situ onto the surfaces of the hostframework. The calcareous cement in limestones may be of different crystal size-grades: micrite (often mistaken for detrital matrix), microsparite, and sparite. The morphologic and textural types are cryptocrystalline, microcrystalline, granular, fibrous, blady, and drusy. Carbonate cement often resembles products formed by recrystallization and grain growth.

Chalk: a porous, fine-textured material, light colored, friable to subfriable, largely to wholly calcareous. It may be slightly tuffaceous. Commonly finely grained, not crystalline, and may be composed largely of foraminifera1 tests and/or comminuted remains (notably of Coccolithophoridae). Chalk can also be of partly chemical origin, although it normally represents the “flour” formed by break-down of skeletal, nonskeletal, and pelletoid grains and algal “dust”. It is largzly micro-textured (about 0.01 mm or smaller). ClaJt: an individual constituent of detrital sediment or sedimentary rock produced by the physical disintegration of a larger mass either within or outside the depocenter of accumulation. A limeclast, therefore, may be an intraclast of the limestone particle, or a fragment disrupted from partially consolidated lime mud on the sea floor or lake bottom. Particulate material may also be doloclasts (see extraclast and intraclast). Clasts of all dimensions are recognized in the older literature in part as penecontemporaneous intraformational detritus.

Clastic: particles of either fragmental or chemical origin that have been rolled around and abraded before coming to rest in a sediment. The variety intraclast originates in the depocenter of sedimentation. Coated grains: grains possessing concentric or enclosing layers of calcium carbonate (or dolomitized remnant); for example, oolites, pisolites, superficial oolites, and algal-encrusted skeletal grains. Composite grain: aggregation grains (detrital, skeletal, pelletal, algal, coated grains, etc.) formed from clustering of two or more discrete particles. It may also result from aggregation of lumps. Some resemble grapestones. Compound-pellet: a pellet of silt-, sand-, or granule-size or larger originating from pelletal 01’ pelletoid limestone with micritic or sparry cement, and may also have matrix or interstitial material.

Coquina: carbonates consisting wholly or largely of mechanically sorted fossil debris, weakly to moderately cemented but not completely compacted and indurated; interstitial material does not necessarily fill all interstices. Commonly applied to shell debris. For the finer shell detritus of sand size or less, the term mesocoquina may be applied; microcoqarina usually implies a variety of chalk.

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Coquinite: for the most part, indurated equivalent of coquina. Coquinites are carbonates which are nearly all fossil debris, mechanically sorted in large measure, and may have finertextured matrix together with cement. It is a compact and well-indurated, cemented rock. Most discrete particulate fossil material is larger than 2.0 mm in size, and some fossils may still be articulated. This carbonate is the indurated equivalent of loose shell debris. For the finer shell detritus of sand- to silt-size, the term mesocoquinite is used, and microcoquinite applies to still smaller size fossil debris. Coquinoid (limestone) :a distinction should be made between coquina, coquinite, and coquinoid limestone. Coquina is loosely compact, poorly cemented, and weakly indurated (shell debris), whereas coquinite is its indurated, firm rock equivalent. Coquinoid limestones are autochthonous deposits consisting of coarse shelly materials which have accumulated in place and generally have a finer grained matrix, or may be enclosed in micritic limestone. Coquinas and coquinites experienced substantial to considerable abrasion and transit before reaching the depositional site, but coquinoids have formed largely in situ, and under certain conditions can build up to biostromes. Coralgal: intergrowth of Algae (particularly coralline types) and corals, to form a firm carbonate rock. This may result in a reef rock, a bank deposit, or a biolith which has a lesser degree of framework. Normally, this rock is composed in large measure of frame-building organisms arranged in an interwoven to interlaced arrangement. It is an excellent sedimentbinder. Criquina: coquina of crinoidal debris. Criquinite: indurated equivalent of criquina. Commonly this rock is an encrinal limestone composed wholly or largely of disarticulated crinoid stems and/or plate fragments, is firmly cemented, is compact, and is matrix-bounded. Cryptoclastic: micritic limestone (or dolomicrite) having an aphanic clastic or microgranular texture, discrete particles of which are less than 0.001 mm in size, and under high-power magnification display little or no crystallinity. This is “rock flour”, or extremely finely comminuted carbonate “dust”. Cryptograined is essentially synonymous, although this type of texture could result from chemical precipitation and/or biochemical to physicochemical precipitation or flocculation. Cryptocrystalline: micritic limestone (or dolomicrite) having an aphanic crystalline texture, discrete subhedra and euhedra of which are less than 0.001 mm in size. Some varieties of cryptocrystalline dolomites display a translucent “sheen” when broken, and may granulate rather than flake. Cryptograined: a size term for micritic sedimentary carbonate rocks referring to particles of cryptoclastic detritus (or flocculated, or precipitated), discrete grains of which are less than 0.001 mm in size (some workers prefer an upper limit of 0.004 mm). Crystalline: refers to a texture characterized by interlocking crystals in a mosaic, or to discrete crystals whether juxtaposed against each other or against grains, fossils, or matrix. It may also refer to discrete crystals in finer grained matrix, or in niicrite. Crust (algal): a deposit of algal “dust”, filamentous or bladed Algae, or clots of Algae on larger particulate rocks or fossils, arranged due to accretion, aggregation, or flocculation. May form “biscuit-like’’ encrustations on rocks, fossils, grains, pellets, etc. It is common to fresh-water, lacustrine, and marine deposits, and may form large, bulbous masses or “heads”, such as those formed in sediments of the Lake Bonneville Group (Pleistocene).

Dense: compact, having various parts crowded together. Its use is not restricted to aphanic and


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finely textured rocks, and may be used for phaneric rocks as well because it defines the degree of compactness. Micrites and dolomicrites (in particular the primary dolomicrites) are dense. Depocenteu: basin or other repository of sedimentation; it may be lacustrine, paludal, or fluvial, but more commonly is marine to terrestrial-marine for carbonate sedimentary rocks. Derrital: formed from debris of pre-existing rocks; detrital limestone is one formed from the detritus of older carbonate rocks, whether derived from an extra-depocenter or intra-depocenter source (or both). Petroclastic limestone is considered synonymous with detrital limestone. If composed of discrete particulate fossils or of older fossil fragments, the term skeletal-detrital may apply for the rock name.

Diagenesis: all those processes which change a fresh sediment into a stable rock of substantial hardness, under conditions of pressure and temperature not widely removed from those existing on the earth’s surface in various depocenters. Diagenesis refers primarily to the processes and reactions which occur within a sediment between one mineral and another or among several minerals and the interstitial fluids. It includes all those processes leading up to final induration of the rock but just before incipient metamorphism. Syndiugenesis occurs penecontemporaneously, at the interface in lime ooze or mud, and during early stages of compaction, cementation, and water-expulsion, but before deep burial. Deepburial diagenesis, but still in the realm of temperature-pressure conditions normal to the depocenter, is late diagenetic. Diagenetic dolomite: dolomitized limestone, or dolomitized lime ooze or mud while still in an uncompact (possibly watery) state, and prior to complete lithification. The process involves all those changes leading up to final dolomitic limestone or dolomite (= dolostone), with positive and negative relics still discernible. Dolarenite: dolomite sand, and thus largely “primary” in the sense of being reworked and abra. ded pre-existing rock. Some dolarenite results from clotting, coagulation, and aggregation of dolomite mud, with concomitant and later rolling and shaping into sand-size particulate material. The rock has a sucrosic or “sugary” appearance. Dolurenaceous is a term describing the texture of dolarenites, or dolomite “sand” derived from crystalline dolomite. Dolocast: casts in a dolostone, dolomitic limestone, or in gypsum or anhydrite, indicating former presence of dolorhombs. Dolobtite: dolarenites range downward in size-grade through dolosiltites to dololutites, which are silt-size and silty-clay to clay-size dolomitic rocks, respectively. Dolosiltites and dololutites are common in evaporitic sequences and may be interlayered with dense primary dolomites. Dolomicvite: aphanic to finely crystalline and grained micritic dolomite, resulting from induration of magnesium-rich mud, or diagenetic dolomitization of micritic limestone. Dolomicrites are common to evaporitic sequences. Dolomite: the hexagonal rhombohedra1 mineral, CaMg(C03)~;it may be used by some workers to define a carbonate rock composed of more than 50 % by weight of the mineral dolomite (other workers may prefer the term dolostone for the rock). For practical microscopic work, areal percentages are used instead of weight percentages. Dolomitic: where used in a rock name, “dolomitic” refers to those rocks that contain 5-50 % of the mineral dolomite. Dolomitic can also be used as a general term applying to those rocks which are dolomite-bearing.

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Dolomitic mottling: arrested dolomitization, or arrested (or incomplete) dedolomitization. Common to limestones that have large particulate skeletal or non-skeletal material embedded in finer textured matrix that, under the effects of dolomitization, preferentially replace or alter the matrix but not the large particles. Also common to more or less homogeneous textured limestones that have been incompletely dolomitized, leaving patches, blotches, “birdseyes”, laminae, or other structures and textures unaffected. Dolomitized: refers to rocks or portions of rocks in which limestone (and sandstone) textures are discernible, but which have been converted wholly or largely to dolostone or dolomite rock. If a rock has been dolomitized, positive and negative fossil or grain relics commonly are recognizable.

Dolomolds: molds or partially filled molds in dolomite, dolomitic rock, or in gypsum and anhydrite, indicating former presence of dolorhombs. Dolorudite: as dolarenites pass upward in size grade, they are called dolorudites. They can consist of older rock fragments, subangular to round, of mtraclasts, particularly if reworked, and of dolomitized reef “trash” detritus in the fore-reef tract. Dolomite mud (ofearly diagenetic types) can be disrupted from the floor of the depocenter, reworked and indurated to form dolorudite, such as edgewise conglomerate or flat-pebble conglomerate. Dolosparite: see sparite. Druse (drusy): sparry calcite (or sparry dolomite) lining or filling shells, open spaces such as voids, pore spaces, interstices, cavities, etc. The druse is crystalline. Drusy coating: calcarenite grains, regardless of origin, may be surrounded by a thin layer of needlelike calcite (or dolomite) crystals that grow normal to the grain surface. The coat, composed of tightly packed scalenohedral or rhombohedra1 crystals projecting outward into the intergranular pore space, forms a rind generally not more than 100 ,LA thick. Earthy: refers to a variety of argillaceous to slightly argillaceous carbonate (limestone or dolomite) with earthy texture, generally closely associated with chalky deposits and commonly showing similar porosity values. Contrasts in some respects with the term dense. It is microtextured (0.01 mm and less in size). Some primary dolomites have an earthy texture. Encrinal (encrinite) : containing crinoid stem and/or plate fragments in the carbonate. If the content of crinoidal fragments is more than 10 % but less than 50 % of the bulk, the rock is an encrinal limestone (or dolomite); whereas if there is more than 50 ”/o of such material, it is an encrinite or dolomitized encrinite. Endogenic: refers to components derived from within the depocenter. Energy Zndex: inferred degree of water agitation in the depositional environment. Energy level: the kinetic energy that exists in the water at the depositional interface and a few feet above. This energy of motion may be due to either wave or current action, or to surf surge. Eolianite: sedimentary accumulation formed by wind action. Oolites, oopellets, pellets, and some other particulate material that originally formed in a water environment may subsequently be transported by wind action (such as across a tidal flat, beach area, etc.) and ultimately heaped into dunes and other deposits. Discrete particles may be carbonates, gypsum, and some other materials. Wind-drifted oolite sands, pelletal sands, and gyparenites adjacent to Great Salt Lake and in the Salt Flats of Utah are such examples.


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Epidiagenetic: lithologic changes of a rock under ground or artesian water influence. Applies especially to topographic dolomitization (FAIRBRIDGE, 1966). Epigenesis: this includes all processes at low temperature and pressure that affect sedimentary rocks after diagenesis and up to metamorphism in the depocenter. This gradation may occur: syngenesis +-diagenesis + epigenesis + metamorphism. Some geologists reserve the term “epigenesis” to mineral replacements near the earth’s surface. Equant: equidimensional crystalline fabric, whether aphanic or phaneric. Equigranulur: equidimensional grained fabric, whether aphanic or phaneric.

Euhedral: refers to individual crystals with well-developed crystal boundaries or faces; these can be equant, or inequant. Evaporite-solution breccia: solution breccias are created when intervening soluble evaporites (salt, anhydrite, gypsum, etc.) are dissolved away, letting the carbonate beds crush under the weight of overlying sediments. This rock type is an extremely angular collapse breccia, and the matrix commonly is of the same material as the rock fragments. These chaotic breccias normally are associated with evaporites, and may also be adjacent to reef limestones, which, upon removal of the evaporites, collapse and may be “healed” or cemented by calcareous and/or dolomitic material. Exogenic: referring to components derived from outside, i.e., from either above or below, the sedimentary formation and from an extra-depocenter provenance. Extraclast: fragment(s) of calcareous sedimentary material produced by erosion of an older rock outside of the depocenter in which it accumulated. Fabric: arrangement of discrete particles (grains), crystals, and cement relative to each other in a sedimentary carbonate rock. Fibrous: see cement. Fore reef: the seaward side of the reef trend. The fore-reef sediments, composed primarily of reef detritus, interfinger with the reef and basin sediments. The terms fore reef and back reef apply only to linear reef trends, in contrast to reef core and reefflank, which apply to all types of reefs. The back reef is the landward side of the reef trend, and its sediments are largely reef-derived fossil debris, calcarenite, and calcilutite, which may interfinger with both the reef and lagoonal facies. Fragmental: refers to broken or detached debris. Detrital fragments and those derived from the skeletons of organisms, are included under this term. Fragmental limestone is, therefore, a mechanically-formed rock; fossiliferous-fragmental limestone is common in stratigraphic sections, and commonly is termed bioclastic limestone. Framework: rigid, wave- and current-resisting structures bioconstructed by sedentary organisms capable of erecting a limestone upward and outward in a high-energy environment. Framework-building organisms include sponges, stromatoporoids, corals, bryozoans, Algae, and combinations (such as coralgal, bryalgal, etc.).

Fusulinal: a term denoting presence, in minor to major (and even dominating) amourlth of fusulinid tests in a carbonate sedimentary rock. These may be in a micrite, or other iirriestone, or may remain as relics in dolomitized rock particularly if silicified prior to dolomitization.

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Glomeraclustic: a textural term applied to sedimentary carbonate rocks in which lumpal particles are grouped together in clusters. Grain growth: this process acts in monomineralic rocks of low porosity. The intergranular boundaries migrate causing some grains to grow at the expense of their neighbors. The reaction takes place in the solid state, ions being transferred from one lattice to another without solution. Larger grains tend to replace smaller ones, and a fine mosaic is gradually replaced by a coarser one. As grain growth proceeds, many of the enlarged grains are themselves replaced by their more successful neighbors. In limestones grain growth appears to affect only the very fine mosaics with grain diameters from 0.5 to 4.0 mm. These include calcitemudstones, the walls of Foraminifera, algal frameworks, Bahamite particles, and ooliths (BATHURST, 1958,1959). Grain-supported: rocks in which grains are so abundant as to support one another, with little or no interstitial mud-matrix, but with various cement types. Grains: discrete particles larger than 0.01 mm (for most routine work), but technically particles can be cryptograined and micrograined. They may form the rock framework, similar to sand grains in a sandstone, or they may be subordinate to smaller particles in the rock. Grains include detrital (lithoclastic) particles, skeletal grains (bioclastic), pellets, coated grains, oopellets, and glomeroclastic grains (many of which may have formed by clotting, coagulation, flocculation, aggregation, etc.).

Grainstone: mud-free carbonate rocks, which are necessarily grain-supported, are termed grainstone; some are current laid, some are the product of mud being by-passed while locally produced grains accumulate, or of mud washed out. Granular: applied to sedimentary rocks made up of grains, usually larger than 2.0 mm in diameter. Grapestone: composite grains, clusters of pellets, or irregularly-shaped grains having protuberances resulting from overgrowths, aggregation, flocculation, clotting, coagulation, etc. of lime mud. May be of silt-size, but more commonly are of sand- and granule-size. Grumous: textural feature seen in limestones that experienced pervasive crystallization. Such crystallization (and recrystallization) develops patches of coarsely crystalline carbonate which invade, in an irregular way, shell debris, oolites, and matrix alike. The uncrystallized areas remain dark, dense, and finely textured, and are ultimately surrounded by waterclear coarse crystalline calcite (sparite). Such a rock has a clotted or grumous texture; and in some respects resembles spotted dolomite. This texture is common in diagenetic dolomites. High-energy: the environment of lithoclastic and bioclastic carbonate working and accumulation, mostly in a zone of turbulence created by waves, currents, and surf-surge. Has the highest Energy Index.

Hydroclust: lithoclastic and bioclastic carbonate detritus that is transported, worked, and deposited in a water environment. Hydrolith is the resulting rock. Impingement: a mechanism or process in dolomitization in which dolomite crystals replace limestones, commonly skeletal particles such as crinoid ossicles and plates, but not in optical continuity with the calcite of the original particle. Interface: depositional boundary condition separating two different physicochemical regions. When particles come to rest on the bottom of a depocenter, they form a solid matrix having water-saturated pores. The water has the same composition as the medium above, but marked changes occur once it is sealed from free circulation by confinement in the pores.


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As deposition continues, the lamina of sediment passes from the water-sediment interface to successively lower positions and enters a realm of greater pressure, higher temperature, and of changed chemical and biologic conditions. These new conditions promote the consolidation or lithification of the sediment into a sedimentary rock.

Intergranular porosity: void space between grains, whether bioclastic or Iithoclastic. In sedimentary carbonate rocks the term granular commonly refers to the grains, whether skeletal or nonskeletal. Some geologists, however, regard a granule as a size-grade textural term, with discrete particles larger than 2.0 mm in diameter. Internal sedimentation: allochthonous clastic and/or chemical sediment derived from the surface or from within the rock framework and accumulated in cavities within the sedimentary rock formation. It is a collective term including both mechanical and chemical internal sediments. Interstitial: of, pertaining to, existing in, or forming an interstice or interstices (Standard), In sedimentary carbonate rocks, interstitial denotes the space between grains and/or crystals. Intraclast: fragments of penecontemporaneous, generally weakly consolidated carbonate sediment that have been eroded from adjoining parts of the sea bottom and redeposited to form a new sediment. The particles have been reworked within the area of deposition and within the same formation. Inversion: the process by which unstable minerals change to a more stable form of the same chemical composition (except for a possible change in content of trace elements and/or isotopes) but with a different lattice structure. Limestone: pertains to a carbonate rock composed of more than 50 % by weight of the mineral calcite. For practical purposes in microscopic work, area percentages are used instead of weight percentages. Lithocalcarenite: that variety of calcarenite in which detrital and/or intraclastic fragments predominate; these are devoid of organic structures originating from aggregation processes. This is the lithic calcarenite, or lithoclastic variety of calcareous arenite. Lithocalcarenites contain a predominant number of sand-size carbonate grains, angular to well rounded, which are fine-grained to coarse-grained and are devoid of any internal structure. The name bahumite includes deposits varying from calcisiltites to calcirudites, in which the grains are accretionary and commonly are composite, consisting of smaller particles bound together by precipitated material into aggregate grains. Lithocalcilutite: fine-grained equivalent of Iithocalcarenite, and also the fine-grained equivalent of practically all the allochthonous types of calcareous deposits, particularly of the lithocalcarenites of aggregation origin (bahamites). Many of the original lithocalcilutites appear now entirely crystallized; they have a crystallinity coarser than 0.002 mm and consist of a finely crystalline mosaic of interlocking anhedral to subhedral crystals of calcite or of dolomite (dololutite). Lithoculcirudite: composed of uniform to nonuniform, or composite nonskeletal particles of microcrystalline calcite generated by aggregation processes (bahamites), pisolites, composite oolites, and coarse particles disrupted and torn from any pre-existing autochthonous and allochthonous limestones. Lithodolorudite, like lithodolarenite and lithodololutite, is of CaMg(CO& composition. Lithocalcisiltite: companion term for silt-size carbonate lithic fragments; lithodolosiltite is the dolomitic equivalent.


Lithoclastic: autochthonous or allochthonous carbonate detritus; mechanically formed and deposited carbonate clasts, derived from previously formed limestone and/or dolomite, within, adjacent to, or outside the depositional site. Some linieclasts are intraclasts and were derived from particulate material torn from the sea bottom (or lake bottom) and incorporated in the new unit. Lithographic: pertaining to a compact carbonate rock having about the same particle size and textural appearance as the stone used in lithography. Characterized by conchoidal fracture, extreme smoothness of texture (aphanic), and uniformity of grains. Internally, the rock may be entirely micro- or cryptograined, entirely micro- or cryptocrystalline, or variously combined. Lithographic is a term to describe the appearance, fracture habit, utility in lithography, etc. and applies (in part at least) to the micrites and dolomicrites. Sub-lithographic is a term designating a minor degree of the above features. Lump: in recent sediments, “Iumps” are composite grains typically possessing superficial reentrants and believed to have formed by a process of aggregation; or lumps may result from disruption of partially indurated lime mud or dolomite mud in the depocenter. In ancient sediments, the “composite grain character,” may not be easily distinguished, but “lumpal” material of larger dimensions will be readily discernible. The following criteria are useful in recognizing lumps: ( I ) lobate outline, reflecting superficial re-entrants, (2) grains texturally similar to the material in which they occur, and (3) rock associations. Lumps may originate through clotting, flocculation, aggregation, and through disruption of newly-deposited lime mud or dolomite mud. In the latter case the fragments are reworked and redeposited within the unit from which they were disrupted.

Luster-mottling: sandstones of various compositional types may become cemented with calcite and/or dolomite which assumes crystallinity; each pore may be filled with a single crystal or with several crystals, some of which are up to a few millimeters in size. Freshly broken arenites so cemented by sparry calcite or dolomite display luster-mottling when turned in the light. Many dolomitic quartz sandstones illustrate this phenomenon. Luiite: this is mud, and is thus a combination of clay- and silt-size particles with the former predominating; the term can carry prefixes such as calci-, dolo-, lithocalci-, biocalci-, etc. Mad: semifriable mixtures of clay materials and carbonates. The better-indurated rocks of like composition are marlstones or marlite, which can be considered more correctly as earthy or impure limestones rather than shales. Marl contains 30-70 % of carbonates and a complementary content of clay. Most commonly, the term marl has been used to denote certain friable carbonates (usually earthy) which accumulated in Recent or present-day fresh-water lakes. It may result from precipitation, flocculation, or physical settling out of the water; photosynthesis of plants in lakes, shallow seas, etc. can hasten the precipitation of marl. Certain microorganisms likewise can aid in precipitating, flocculating and settling of the sediment. Matrix: the natural material in which any fossil, rock fragment, crystal, grain, etc. is embedded. In a rock in which certain grains are much larger than the others, the grains of the smaller size comprise the matrix. If the particles in a rock are of different orders of magnitude of size, the term matrix applies to the smaller individual units that fill the interstices between the larger grains. Sand-, silt-, and clay-sized material which is resolvable only by size and shape analysis is included in the definition of carbonate matrix. Micritic material may also be called matrix when it encloses grains or fills interstices between them. Maturity (sediment) :the extent to which clastic carbonate material approaches the end product to which it is driven by the formative processes that operate on it.


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Mechanical: pertaining to particles of sediment brought to their place of final deposition by agents such as water currents, wind currents, or gravity. Megalump: most lump (= lumpal) limestones contain silt-size and sand-size clots, coagulated irregularly-shaped grains, and aggregate grains (bahamite); some, however, contain megalumps, which are coarser-textured and lie in the granule- to boulder-size classes. These are “lithocalcibreccia-” and “lithodolobreccia-” like masses, and may originate through tearing up by waves, currents, and surf-surge of very high energy-index values, and possibly by turbidity currents. Partly indurated lime mud or dolomite mud could be ripped up into fragments which are shaped to resemble rudites, and then incorporated within the unit from which they were derived. In the early literature, these are known as penecontemporaneous intraforniational mud-pebble conglomerates or breccias. Metusomatism (dolomitic) :diagenetic differentiation is the redistribution of the materials within a sediment, leading to segregation of the minor constituents into nodules, concretions, and related bodies. Diagenetic metasomatism involves introduction of materials from without, leading to replacement. Micrite: consolidated or unconsolidated ooze or mud of either chemical or mechanical origin. FOLK(1959) originally stated that the term should be reserved for those rocks that, under the petrographic microscope, are seen to consist almost entirely of microcrystalhne calcite. LEICHTON and PENDEXTER (1962) defined the micritic material as that consisting of particles less than approximately 0.03 mm in diameter. In the present chapter, micrite is employed for material, whether crystalline or finely grained, that is 0.05 mm or smaller in diameter or across faces. Micrite is lime mud or its indurated equivalent, and dolomicrite is dolomite mud or its indurated equivalent. Micritic limestone: a limestone which consists of 90 % or more micrite. Microclastic and microcrystalline limestones are two varieties of micritic limestone (LEIOHTON and PENDEXTER, 1962), the former possessing a clastic texture and the latter a texture of microscopic size interlocking crystals. Aphanic (aphanitic of some authors) limestone, matrix limestone, calcilutite, and lithographic limestone are practically synonymous with micritic limestone. Microcrystalline: usage varies (see Table IV), but it refers to crystallinity in limestones and dolomites between 0.001 and 0.01 mm across crystal faces. Some petrographers prefer the limits of 0.004 to 0.062 mm.

Micrograined: clastic carbonate particles between 0.001 and 0.01 mm in diameter, or it may he preferred (see LEIGHTON and PENDEXTER, 1962) to have an upper limit of 0.0625 mm and a lower limit of 0.004 mm. They apply the terms coarsely micrograined (0.03 to 0.06 mm) and finely micrograined (0.004 to 0.03 mm), with cryptograined particles being smaller than 0.004 mm in diameter. Micropelfetoid: particles of pellet nature, or possibly true pellets (faecal or otherwise), of a fine to very fine grade size, possibly smaller than 0.01 mm in diameter. FOLK (1962) places all pellets in the size range of 0.03 to 0.15 mm.

Mosaic: a textural term, more applicable to dolomites than to limestones (except the non-dolomitized, but diagenetically altered varieties). Secondary overgrowth of dolomite on rhombs produces mosaic-like texture. This destruction of original intercrystalline porosity by continuing growth of dolomite is analogous to the “cementation” phenomena of wellsorted skeletal and nonskeletal limestones, such as encrinal limestones and the like. Mud: silt-clay mixture (the latter size commonly predominant) in water without connotation as to composition. Lime mud dolomite mud, etc. identify the variety.

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Mud aggregate: any aggregate of mud grains, commonly having the size of a sand or silt particle, and usually mechanically deposited. Initially the aggregate may have been a faecal pellet, or a rounded, subspherical aggregate of mud grains cemented originally by aragonite with no signs of organic control, or a fragment of algal precipitate, or a spherical or ovoid growth form of a calcareous alga (BATHURST, 1959). Mud-supported: muddy carbonate rock which contains more than 10% grains, but not in sufficient amount that they support one another; such grains are “floating”, and thus they are mud-supported. Mudstone: muddy carbonate rocks containing less than 10% grains (grain/bulk ratio being 0.1); the name is synonymous with calcilutite, except that it does not specify mineralogic composition, and does not specify that the mud is of clastic origin (DUNHAM, 1962). Nodular limestone: a variety characterized by nodules, lumps, clots, and grapestone accumulations within argillaceous and micritic limestones. Nodules are not necessarily concretions, but may represent lumps, flocculated material, and round to subround aggregations, and similarly-shaped very large coated grains. Most commonly nodules are composed of the same type of material that encloses them. Nonclastic: having a texture showing no evidence that the sediment was deposited mechanically. Nondetrital: minerals that are precipitated from solution by chemical, physical, physicochemical, biochemical, or biologic means. Accumulation occurs at, or generally close to, the site of precipitation. This group also includes the authigenic minerals formed in the sediment after deposition. Nonskeletal limestone: lithocalcarenites and lithocalcisiltites containing subrounded to rounded grains devoid of internal structure, organic or otherwise. These grains have been termed pellets, granules, false oolites, oopellets, intermediate pelletoid grains, pseudo-oolites, etc. Olistolith: exotic blocks of older strata apparently transported by gravity sliding: “calcolistoliths” are limestone exotic blocks of this type, whether transported by turbidity currents or by gravity sliding. “Olistostrome” is an entire formation of slumps and exotic blocks. Oolite (or ooid): spherical or subspherical accretionary grain generally less than 2.0 mm in diameter. In section, oolites display concentric structure, and may also exhibit radial structure. Oolite is a coated grain, and may or may not have a nucleus. Superficial oolite is a type of oolite in which the thickness of the accretionary coating is less than the radius of the nucleus. Some workers prefer the name oii[ith for the rock, and oolite (or ooid) for the discrete coated grain; some petrographers use the two terms interchangeably, regardless of whether it is a rock or a grain. Oolitoid: similar shaped and sized bodies to oolites and ooliths, but which lack the internal structure normally found in oolites; they consist of a fine-grained aggregate of Fe-rich dolomite (GREENSMITH, 1960). Oopellet: spherical or subspherical grain displaying characters of both an oolite and a pellet, and should not be confused with superficial oolite. The internal part is pelletoidal, and thus may be ovoid in shape, but it has an accretionary coating, the thickness of all layers being equal to or slightly greater than the diameter of the pellet which they enclose. Open-space structures: they are structures in carbonate rocks which formed by the partial or complete occupation with internal fillings composed of internal sediments and/or cement of one to several generations.


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Organic lattice: reef-building framework, in situ. Organic linzestone: biogenic limestone owing its origin directly to physiological activities of plants and animals. Organic structures limestones: the major framework of these limestones has been secreted by organisms such as Algae, stromatoporoids, sponges, corals, bryozoans, and combinations thereof (bryalgal, coralgal, etc.); and the fossil remains are still in their approximate growth position. Orthochemical: carbonate sediment or rock equivalent which is of straight or direct chemical origin. Flocculated and precipitated lime mud may form micrites; and primary dolomites are orthochemical because they have been precipitated directly out of sea or lake water. Primary dolomite is, therefore, an orthodolomite. Ovoid grains: pellet-shaped grains having a length two or more times as long as the diameter; these are “football-shaped” grains, and are commonly 0.1-2.0 mm long (though some are up to 5.0 mm in length). They commonly lack an internal structure, although weakly-

developed radial structure, and rarely concentric structure, can be seen.

Packstone: a limestone in which the grains are arranged in a self-supporting framework, and yet contains some matrix of lime mud. If no mud is present the rock is called grainstone; and if grains comprise less than 10% of the rock mass, it is a mudstone. Paragenesis: a general term for the order of formation of associated minerals, textures, and structures in time succession, one after another. Purticulate: discrete particles, grains, fossils, fragments, skeletal material, and crystals. faurocrystalline: lowest size-grade group of the phaneric crystalline carbonate rocks; subhedra and euhedra lie in the size range of 0.01 to 0.1 mm. The termpaurograinedis the clastic or grained equivalent. Pelagosite: this is a deposit (generally white, gray, to brownish with a pearly luster) composed of CaC03 with higher MgC03, SrC03, CaS04.HzO and SiOz contents than those found in normal limy sediments. It is restricted to intertidal spray-formed incrustations a few millimetres thick (see REVELLE and FAIRBRIDGE, 1957). felite: size-grade of lutite, and is of clay- to silt-size material (the former predominant). Pelitomorphic is an all-embracive term for carbonate particles of this size; but more commonly it connotes anhedra to subhedra. Pellet: a grain composed normally of micritic material, lacking significant internal structure and generally ovoid in shape; it may also be sub-ovoid. Most pellets in limestones are of siltsize to coarse sand-size (some are slightly larger). In some respects pellets are pseudooolites, for they are spherical to subspherical to oval bodies with distinct boundaries, and resemble oolites; however, they do not possess comparable internal structure, for example, faecal pellets. Carbonate muds are commonly pelleted, pelletal, or pelletoid, displaying rounded or ellipsoidal aggregates of “grains of matrix” material. These muds are thought to be pelleted either by faecal activity, gas bubbling, or by algal “budding” phenomena. Penecontemporaneous: a term used in connection with the formation of sedimentary rocks, and implies “formed at almost the same time”. fhaneric: textural term for carbonates, particularly limestones, which are crystalline (and/or grained), and the discrete particles of which are larger than 0.01 mm. faurograined (0.01

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-0.1 mm), mesograined (0.1-1 .O mm), and megagrained (1 .O-10.0 mm) are the three textural subdivisions. The term phaneritic is a term applied to texture in igneous rocks in which individual crystals are visible to the unaided eye, and should not be used as a textural term for carbonate rocks.

Pisolite: a grain type similar to an oolite, and generally 2.0 mm or more in diameter. The term pisolite is restricted to crenulated, rounded or semi-rounded, commonly composite carbonate grains or bodies thought to have been formed by biochemical algal-encrustation processes. They could, therefore, become enlarged by accretion to form “algal-balls”, “algal-biscuits”, and the like. Porphyroblastic: textural term to describe limestones and in particular dolomites, in which large crystals (porphyroblasts) are scattered through a crystalline matrix of finer textured materials. Crystallized crinoid ossicles in micrite or in fine-textured matrix would display this texture in limestones; and sparry dolorhombs embedded in finer crystalline or dolomicrite material would be one example among dolomites. Preferred fabric: preferential crystallinity, particularly in diagenetically altered limestones and dolomites; and also a primary depositional fabric in mechanically deposited !ithodastic and bioclastic sedimentary carbonate rocks. Preferred fabric is characteristic in primary dolomites, as deduced with the universal stage and thin sections. For example, the C-axes of crystals in primary dolomites lie parallel to the plane of bedding. Pressure solution: a preferential solution takes place on the higher stressed parts of a grain (or crystal) and deposition of matter 011 surfaces with lower potential energies. The pressure is supplied by the overburden and should result in recognizable grain fabric, with the grains flattened at right angles to the pressure. Regarded as perhaps the most important process in closing the original pore spaces of sediment (BATHURST, 1958, 1959). Primary dolomite: resulting from direct precipitation out of sea water or lake water. It may have a preferred fabric, and is aphanic to finely textured (crystalline and grained). Dolomite “sand”, regarded by some workers as a primary deposit of first-cycle dolarenite, is hardly categorized as primary dolomite, but is a derived clastic carbonate. Pseudobreccia: masses of grain growth mosaic which lie in a “matrix” of less altered limestone: most of these are visible to the naked eye. The “fragments” are irregularly shaped lumps of coarse calcite mosaic usually between 1.O mm and 20.0 mni in diameter, and are dark gray in hand specimen. They lie in the finer, pale-gray, “ground mass” of calcite-mudstone. In thin section, the “fragments” appear light and the “ground mass” dark. (See Bathurst, 1959.) Pseudomorphic replacement: a diagenetic process whereby the original character of a limestone is altered during dolomitization; skeletal material for example (and specifically crinoidal material) is replaced in such a manner that single crystals of dolomite are in optical continuity with the calcite of the original crinoid fragment. The process contrasts with that of impingement, in that in the latter case there is lack of optical continuity of dolomite crystals with the original crinoid fragment. Pseudo-oolites: some varieties of pellets and oopellets. Recrystallization: a term signifying a process wherein original crystals of a particular size and morphology become converted into crystal units with different grain size or morphology, but the mineral species remains identical before and after the process occurs. First-stage crystallization is not to be termed recrystallization. This latter term is usually used today loosely for a number of processes that include inversion, recrystallization sensu stricto, and grain growth, all of which may result in textural and crystal-size changes. Recrystalli-


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zation proper occurs when nuclei of new unstrained grains or crystals appear in or near the boundaries of the old, strained ones. These nuclei grow until the old mosaic has been wholly replaced by a new, relatively strain-free mosaic with a nearly uniform grain (or crystal) size. Recrystallization fabric: mosaic or other crystalline textural features which identify the pattern of any sedimentary carbonate that has undergone recrystallization. This is not the fabric of metasomatically replaced limestones (by dolorhombs). Replacement crystallization is not recrystallization if the mineral species are no longer the same, such as dolomite after calcite. Reef: a structure erected by frame-building or sediment-binding organisms. At the time of deposition, the structure was a wave-resistant or potentially wave-resistant topographic feature. Reefs and the heterogeneous reef-derived materials form the reef complex. The following types of reefs can be recognized: (1) fringing reefs, for those veneering types that lie or were ad,jacent to the pre-existing land; (2) barrier reefs, for sublinear structures that are or were separated from nearby older land by a lagoon; (3) atolls, for composite structures with ring-like outer reefs that surround or once surrounded a central lagoon devoid of pre-existing land; ( 4 )patch reefs, for small, sub-equidimensional or irregularly shaped reefs that are parts of reef complexes; ( 5 ) table reefs, for flat-topped, isolated, characteristically small reef-mounds of the open ocean; (6)pinnacle reefs, which have a very small area and grow almost vertically; ( 7 ) bank reefs, that grow over submerged highs of tectonic or other origin-these reefs are large and have an irregular shape, and the marine bottom water surrounding them is too deep to support growth of reef-forming organisms; and (8) shoal reefs, that grow on the shoals of the fore-reef and back-reef areas; these are smaller in area than the bank or platform reef types, and generally grow on the debris of a larger reef. Reefal: this is purely a descriptive and not genetic term having reference to carbonate deposits in and adjacent to any of the numerous varieties of reefs, and to any or all of their integral parts. Reef milk: matrix material of the back-reef facies, consisting of microcrystalline white and opaque calcite ooze, and derived from abrasion of the reef core and reef flank. Reef tufa: fibrous calcite which forms thin to thick deposits, layered or unlayered, in the myriads of voids in reefs and other organic frame-builders; the fibrous calcite is prismatic in structure and is radial in respect to the depositional surfaces. The fibrous calcite or reef “tufa” is deposited directly upon the framework of the reefs and within the various voids and interstices, from supersaturated water. The mechanism may be largely physicochemical, or, aided by profuse algal growth to extract CO? from the water, may also be biological to biochemical deposition. Development of reef tufa follows and/or accompanies growth of organic frame-builders, and precedes infilling of detritus such as lime mud, calcarenite etc. Relic: vestige(s) of skeletal and nonskeletal material in a sedimentary carbonate rock, commonly dolomites. Crystallization, recrystallization, impingement, and other diagenetic alteration has not completely obliterated these features if they are to be termed relics. There can be two types of relics: ( I ) positive-the skeletal or nonskeletal element can be identified, though it is altered; (2) negative (“dissolved out”)-it is known that there was a skeletal or nonskeletal item present, but it cannot be identified as having been a part of an organic or non-organic species. Rim cement: cement which grows into interparticle voids and is optically continuous on single crystal particles such as crinoid fragments, etc. Thus, the host is a single crystal and the cement forms a single rim in lattice continuity with it. The overgrowth is a continuation of this crystal, and the overgrowth can form by filling the pore space.

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Saccharoidal: a descriptive term meaning “sugary” texture. More specifically it is a result of diagenetic process (or result of dolomitization) in which crystallization or recrystallization gives rise to a new texture. It may be first-stage crystallization, but more commonly is recrystallization that occurs early in a newly-deposited carbonate mud; it does not alter gross primary structures of the sediment such as ripple marks, thin bedded to laminated layers, etc., but does tend to destroy minor structures such as shells of organisms. Saccharoidal texture is recognized by the well-developed rhombs of dolomite of approximately uniform size resting one against the other with point contact and, likewise, commonly separated by exceptionally large as well as small pore openings. The fabric displays loose packing, and suggests that dolomitization occurred when the grains were loose and before compaction altered the original texture (i.e., a packing typical of loose beach and shore-line sands). Recrystallization of the original smaller dolomite grains, or replacement crystallization of the original calcite grains destroys the original particle-size distribution, and substitutes a new, highly restricted, crystal-size distribution ranging from medium- to coarse-sand dimensions. Secondary: a general term applied to minerals and rocks formed as a consequence of alteration. This term is too all-inclusive and ambiguous in detailed studies and should be used only as a very general colloquial term when misinterpretation is absolutely impossible. Ske/etal: pertaining to debris derived from organisms that secrete hard parts and hard material around or within organic tissue. NELSON et al. (1962) defined skeletal limestones as those which consist of, or owe their characteristics to, virtually in-place accumulation of calcareous skeletal matter. These rocks, formed through biologic processes, are contrasted with fragmental limestones. LEIGHTON and PENDEXTER (1962) considered the term bioclastic to be synonymous with skeletal. The term skeletal is thus also used to indicate faunal (or floral) fragments or wholecomponents of these organisms that are not in place of origin. Solution transfer: this is a translation of the German Losungumsatz. It refers to the solution of detrital particles around their points of contact where elastic strain and solubility are enhanced (pressure solution), followed by redeposition on less strained particle surfaces (BATHURST, 1959). Sparite: a contraction of, and therefore synonymous with, sparry calcite. Sparite is a loose descriptive term applied to any transparent or translucent crystalline calcite and aragonite. It can occur in numerous morphologic forms, viz. granular, drusy, fibrous, and blady. Three possible origins are recognized: ( 1 ) physicochemical precipitation, (2) recrystallization, and (3) grain growth. The first is distinguished by adding the genetic prefix ortho-, and the latter two by adding prefix pseudo-. Sparite is larger than 0.02 mm in diameter. Petrographers who prefer to use also the term nzicrosparite set its size limits at 0.005-0.02 mm. The prefix dolo- is used to indicate sparry dolomite crystals, i.e., dolosparite and tlolomicrosparite. Some workers prefer the prefix calc- to distinguish calcsparite from the dolo’ mitic variety, but to some the term sparite is automatically understood to mean the calcareous variety. Sparry: refers to clear, transparent, or translucent, readily cleavable, crystalline particles generally having an interlocking mosaic texture. FOLK (1959) referred to sparry calcite cement which forms grains or crystals lop or more in diameter. The name spar alludes to its relative clarity both in thin section and in hand specimen. Speleothem: a carbonate cave deposit of any sort (also “speleal” limestone). Spergenite: a coquinite (to niicrocoquinite), and/or a biocalcarenite composed of sorted fossil debris, including bryozoan fragments, and Foraminifera (possibly endothyroids), together with carbonate detritus of various types, cemented by sparry calcite; a biogenic calcarenite.


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SpheruZite: a textural term, applicable to limestones and dolomitized equivalents, in which rounded and subrounded, spherical to subspherical ooids are present. These may have either a concentric, radial, or axiolitic structure, or a combination of all three. Spherulitic limestone contains 50 % or more of these spherules; they range in size from about 0.5 to 2.0 mm, although some are up to 5.0 mni in diameter. No particular genetic significance is attached because they may be (1) detrital, (2) coated grains, (3)algal spherules, and ( 4 ) authigenic. According to some workers only a radial structure is indicated; and according to some, a sgherulite is a small spherical or spheroidal particle composed of a thin dense calcareous outer layer with a sparry calcite core. According to PETTIJOHN (1957), spherulites are minute bodies of oolitic nature in which only a radial structure is visible. Stinkstein: “stink-stone”, or smelly rock; among carbonate rocks, both limestones and dolomites, stinksteins are common. Normally these are of three types, or some combination thereof: ( I ) “sweet” or hydrocarbon odor; (2) fetid, or foul odor (common to most criquinites), or (3) “sweet-and-sour”, typical of carbonates rich in organic-phosphatic material, particularly detectable if dilute HCl is applied to freshly-powdered rock.

Stromutuctis: these are open-space structures with horizontal flat to nearly flat bottoms, and are filled by internal sediments and/or cement. Their genesis has been variously interpreted as being caused by the burial of soft organisms which upon decomposition left an open space. More recent studies, however, show that they are most likely syngenetic voids in calcareous sediments, which are or are not changed by subsequent corrosion and corrasion. Algae are only indirectly responsible by overgrowing surface pits and channels, and thus form an internal cavity system. It seems that Strornaiuciis are most common to micritic limestones formed by calcareous Algae, that left little or no evidence in most occurrences thus far reported from Great Britain, North America, etc.

Stromutolite: laminated sediment formed by calcareous Algae, which bind fine detritus and/or precipitate calciumcarbonatebiochemically. The deposit may form irregular accumulations or structures that may remain fairly constant in shape, for example, Colleniu. Subhedrul: refers t o individual crystals exhibiting a few crystal boundaries; the term unhedrul defines those with no well-defined crystal boundaries, and the term euhedral indicates excellent crystallinity. Subhedrul is midway between these latter two varieties. Sucrosic: contraction of saccharoidal, thus meaning “sugary” texture. The term is commonly applied to certain types of dolomites, e.g., dolarenites.

Syneresis cracks or vugs: cracks or vugs formed by a spontaneous throwing off of water by a gel during aging. In some carbonates, precipitation evidently occurs as a colloidal gel encrusting leaves of sea plants (photochemical removal of carbon dioxide from sea water by the plants causes precipitation). The end-result may be the production of cryptotextured limestone which contains “syneresis” cracks and associated contraction vugs. Syngenetic: originating at about the same time; in sedimentary carbonate rocks, it refers to concretions, authigenic minerals, nodules, and other bodies which form at identically, or approximately, the same time as the rock which encloses them. Syndiugenetic dolomitizution, for example, refers to diagenetic dolomitization which occurs at the interface while lime mud is accumulating.

Syntaxid rims: a mechanism of replacement overgrowth, which develops during diagenesis as a syntaxial extension of a detrital single crystal (e.g., a crinoid fragment). Not to be confused with drusy mosaic, which is a convenient term for grain mosaics which have been deposited on the walls of cavities. Whereas drusy mosaic and granular cement are entirely chemically deposited, the mosaic formed by rim cementation processes consists of grains each of which has a core composed of the detrital host. The textural relationships between lime mud and

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calcite overgrowth in some limestones suggest that rim cementation is the dominant process, and commonly dolomitization occurs after rim cementation. Travertine: a massive to often finely layered colloform type of chemical limestone, often with porous interlayers. In the type area, at warm carbonate springs at Tivoli near Rome, Italy, afine type of sound-absorbing building stone is worked and exported all over the world. The Tivoli River of Italy deposits travertine very rapidly because of agitation. Travertine commonly forms about hot springs, faults, in caves, in soil crusts, as soil nodules, etc. Tufa: a spongy, porous rock which forms a thin surficial deposit about springs and rivers, or may form thick, bulbous or otherwhise swelling features in lacustrine environments, particularly around the shore. It has a reticulate structure, and is weak and semifriable. Travertine, by contrast, is fairly dense, banded CaC03, having tan, cream, and white colors. Turbidity limestones: some litho- and biocalcarenites display features indicating resedimentation by turbidity currents; exotic limestones, particularly relatively large blocks (“calcolistoliths”), are now known to have formed by agents other than waves and currents, possibly by gravity-sliding and turbidity currents. Winnow: eolian sorting; should not be used to describe sorting in water, which is “washing”.

1 69

Chapter 5

ORIGIN AND OCCURRENCE OF LIMESTONES JOHN E. SANDERS AND GERALD M. FRIEDMAN

33 Sherman Avenue, Dobbs Ferry, N. Y. (U.S.A.) Department of Geology, Rensselaer Polytechnic Institute, Troy, N. Y. (U.S.A.)

SUMMARY

Limestones originate by three main processes: ( I ) precipitation of calcium carbonate in an initially stony condition, as in travertine and organic reefs; (2) lithification of calcium carbonate sediments, which includes various steps beginning with changes of grain mineralogy, and includes addition of concentric coatings to grains, selective dissolution of matrix and/or grains, precipitation of mineral cement in pore spaces, and may end with recrystallization; and (3) replacement of calcium sulfate or quartz by calcium carbonate. Of these, the second process is by far the most important. Limestones occur in most parts of the world and throughout the geologic column, but they are notably more abundant in Cambrian and younger rocks, reflecting the increased abundance of marine shell-secreting invertebrates during and after the Cambrian Period. Grains of carbonate sediments mgy come from older limestones, but chiefly they originate as first-cycleparticles within the waters of the depositional basin, as skeletal remains of lime-secreting organisms, ooids, superficial ooids, pseudooids, aggregate grains, faecal pellets, and so forth. Carbonate grains are deposited by various vertical and/or lateral sedimentation processes to build up stratigraphic units, distinctive features of which are controlled by the interplay of physiographic environment of deposition and crustal subsidence. “Pure” limestones and limestones mixed variously with terrigenous sediments from Recent environments and the geologic record are grouped broadly into three major environmental suites: ( I ) nonmarine environments from large inland structural basins (lakes and alluvial fans); (2) shallow-water marine and/or marginal marine environments (beaches, lagoons, bays, dunes, and so forth); and (3) open-sea environments. Work on this chapter commenced in the Department of Geology, Yale University, but was completed after the senior author became a Senior Research Associate, Columbia University. Contribution No. 66-5 of the Department of Geology, Rensselaer Polytechnic Institute.

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J. E. SANDERS AND G . M. FRIEDMAN

INTRODUCTION

The origin of limestones, though a complicated subject, may be stated simply: limestones originate in part by direct inorganic or organic precipitation of calcium carbonate, in part by replacement of other substances by calcium carbonate, but chiefly by the lithification of calcium carbonate sediments. The complexities are due to the great variety of textural characteristics and mineralogical and chemical compositions of calcium carbonate sediments, and the varying degrees to which these original variable features are preserved or modified during lithification. The possibilities range from imperceptible changes in original characteristics to complete loss of original characteristics by solution, redeposition, deposition of new material, to thorough recrystallization. The discussion of the origin of limestones in this chapter presupposes the reader’s acquaintance with the chemical and mineralogic composition of the particles in calcium carbonate sediments and the petrology and petrography of limestones (GUBLER et al., 1966), the characteristics of modern calcium carbonate sediments (TAFT,1966a), the physical chemistry of the formation of carbonate rocks (TAFT, 1966b), and the classification of carbonate rocks (BISSELLand CHILINGAR,1966). All these contributions are found in the present two-volume work on carbonate rocks. Beyond these, the origin of limestones embraces the subjects of the processes and products of lithification of calcium carbonate sediments and of replacement of other materials by calcium carbonate, subjects which are properly restricted to this chapter. The occurrence of limestones is here discussed in terms of the stratigraphic principles that govern the relationships of the different kinds of calcium carbonate sediments to each other and to non-carbonate sediments, omitting the less common types of limestones. The stratigraphic principles which apply to calcium carbonate sediments include most of those that govern noncarbonate terrigenous sediments; in addition, calcium carbonate sediments involve certain unique principles which apply only to themselves, owing to the fact that calcium carbonate sediment particles may be derived from the water of the depositional basin and, further, that most of these particles are of biogenic origin. In view of the foregoing, the logical point of departure for a discussion of the origin of limestones would appear to be the calcium carbonate sediments from which most limestones have been derived. Because this topic is discussed by TAFT (1966a), it would be repetitious to organize this chapter in that manner. Accordingly, the arrangement used here is somewhat different, and follows three major headings: ( I ) review o f the classic literature on limestones; (2) origin of limestones, including original stony precipitates, lithification of calcium carbonate sediments, and replacement processes; and (3) occurrence of limestones, including stratigraphic relationships of different kinds of calcium carbonate sediments with each other


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and with noncarbonate sediments, citing examples botn from Recent calcium carbonate sediments and ancient limestones.

REVIEW OF THE CLASSIC LITERATURE ON LIMESTONES

Limestones have long attracted the attention of geologists, perhaps initially because of their early use as building stones and also because of the inherent fascination prompted by the fossils in limestones. LYELL’S classic textbook, which appeared in 1830 and passed through numerous editions for nearly 3 decades, contains extensive descriptions of the Italian travertine deposits and comparisons between recently elevated coral-reef limestones, as in Sicily, with modern living reefs. Publication of DARWIN’S “Journal” (1845) of his voyage on “H.M.S. Beagle” and other writings on coral atolls (DARWIN, 1837, 1889) emphasized the problems connected with the origin of Indo-Pacific coral atolls. Many studies resulted from Darwin’s stimulus; their culmination was in the drilling of many test borings by the U.S. Geological Survey in the decade after World War I1 (LADD et al., 1953). Laboratory study of limestones by means of thin-sections and the petrographic microscope, and also field study of polished surfaces prepared with an abrasive stone or file and then moistened before study, was inaugurated by SORBY(1879). Sorby’s Presidential Address to the Geological Society of London is one of the most penetrating studies of limestones ever made; he applied these methods to a study of the mineralogy and textures of limestones and Recent carbonate sediments with great success. He concluded that limestones are mainly derived from cementation of broken up and decayed shells and corals; hence, a knowledge of the mineralogy and structure of shells is fundamental to understanding limestones. Sorby employed careful specific gravity measurements to distinguish aragonite powder (s. g.=2.93) from calcite (s. g.=2.72); in addition he used optical properties (aragonite is biaxial; calcite, uniaxial) and hardness (aragonite is harder than calcite and will scratch Iceland spar crystals in a direction parallel to the line of the shorter diagonal of the rhombic face, whereas calcite will not do so). Using these techniques he determined the mineralogy of various shell groups and set forth the principles of the different behavior of calcite and aragonite, and discussed how these factors influenced preservation of shells and textural characteristics of limestones. Sorby examined travertine and tufa, oolites, and a suite of specimens collected systematically from the principal limestones of Great Britain. He recognized three kinds of ooids: (1)those with concentric structure; (2) those with radial structure; and (3) those that have been recrystallized. He remarked that “unoriented

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granular matter” (lime mud, etc.) occurs with the oriented crystal layers in the first two kinds. Sorby’s descriptions of the British limestones include numerous references to sand-sized particles of lime mud and lime-mud coatings on shells and as layers within ooids. The full significance of these mud particles and coatings was not realized until 1954, when ILLING’S work on the Bahamian calcareous sands was published. Other important contributors to the petrographic study of limestones include CAYEUX(1897, 1935) in France; HADDING(1941, 1950, 1956, 1958) in Sweden; and SANDER (1936, 1951) in Austria. For the most part, these descriptive studies concerned only ancient limestone rocks; only SORBY (1879) made the all-important transition between Recent carbonate sediments and ancient limestones. Walther, one of the stalwarts of geology of the late 1800’s and early part of this century, wrote a three-volume textbook and devoted the third volume, nearly 500 pages, to “Lithology of the Recent” (WALTHER, 1893-1894). In this volume considerable space is given to Recent carbonate sediments. Walther presented an excellent review of the subject through 1893, and it is surprising how much knowledge had already accumulated by that time. He was an outstanding synthesizer and shrewd observer who had travelled extensively on four continents (Europe, Asia, Africa, and North America) to make geological observations. WALTHER’S study (1888) of the Recent and Pleistocene carbonate sediments and rocks of the Red Sea is many decades ahead of his time. It was Walther who took the term “diagenesis”, which he had borrowed from Von Guembel (WALTHER, 1893-1 894, p.693), and gave it the meaning presently used by most geologists. In his book, Walther made many suggestions for the solution of geological problems and was an early advocate of the experimental approach to problems of carbonate diagenesis. GRABAU (1904, 1932) recognized most of the varieties of limestone that are known today. His insights became obscured, however, because few geologists bothered to penetrate beyond his elaborate genetic terminology, or to follow his scattered discussions of limestone, which are found in eight different chapters of his textbook (GRABAU,1932, pp.272-273, 280-285, 294-297, 331-347, 384-476, 573-578, 692-693, and 645, respectively). One major change from Grabau’s views on limestones, occasioned by subsequent research, is the recognition of the importance of wave working of chemical deposits of calcareous material formed in the sea. Although GRABAU (1932, p.645) indicated that such limestones were “not known”, he did remark that “chemically formed marine oolites when worn by waves may come under this head”. GRABAU(1932) divided all rocks, including “limestones”, into two major categories: ( I ) endogenetic, formed by agents within the rock mass itself, such as solidification from molten material or precipitation out of solution; and (2) exogenetic, formed by agents acting outside of already existing rock masses, as clastic rocks. Endogenetic limestones form in water, hence are hydrogenic; they are called


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calcareous hydrolitlzs in Grabau’s terminology. Many endogenetic limestones have formed by organisms, thus are biogenic; the rocks were called calcareous bioliths and further subdivided into zoogenic types if formed by animals, and phytogenic, if by plants. Endogenetic limestones include the calcareous hydroliths such as: chemical limestone (mostly oolites in the sea, tufa and onyx marble in lakes, rivers, and springs), stalactitic deposits, and calcareous tufa; and calcareous bioliths: organic limestones, coral rock, and shell beds (organic oozes and the like). Exogenetic limestones include varieties broken by the action of water (hydrocalcilytes, or hydroclastic limestones), by organisms (biocalcilytes, or bioclastic limestones), or by the wind (anemoclastic limestones). Grabau further classified the exogenetic rocks into textural groups based on grain size: rudaceous (coarser-grained than sand-size), areizaceous (sand-size), and lutaceous (finergrained than sand-size). He offered a new classification of marine deposits (1932, p.645), in which marine elastic limestones, the hydroclastics and bioclastics, were further subdivided into terrigenous types, if the material composing them was derived from the land, or thalassigenous types, if derived from the sea. It is important to note that in Grabau’s usage bioclastic is not a synonym of biogenic, as many geologists have supposed. This term refers specifically to fragments of older rocks that were broken or fragmented by the mechanical action of organisms, such as parrot fish feeding on corals and nullipores in a reef, grinding up the calcareous material, and excreting it back onto the bottom. Although Grabau’s monumental textbook did not include much discussion based on observations with the petrographic microscope, his treatment of the other aspects of limestones is so thorough that this omission seems relatively minor by comparison. Grabau’s bibliographies are cosmopolitan, but somehow he omitted Sorby’s work on limestones (SORBY,1853, 1855, 1861, 1879), possibly because it is petrographic. Much of the so-called “new” research on limestones that occurred in the United States after 1950 would have been accomplished far earlier had the study of limestones commenced with a combination of the material available in the published works of both Sorby and Grabau. In the 1920’s and early 1930’s, KLAHN (1922, 1923a, b, 192% b, c, 1926a, b, 1928, 1932a, b) was an active worker in the field of carbonate sedimentation and petrology, yet his extensive studies have been largely ignored by modern workers. Most of KYahn’s observations were concerned with fresh-water carbonate sediments, and he was one of the first workers to attempt the experimental approach. Another prodigious student of carbonate sedimentation and petrology was Pia, who between 1912 and 1933 made many contributions to this field. In 1933, in a 420-page book on “Recent Carbonate Sediments”, he summed up the entire field as then known to him (PIA, 1933), and ranged across the spectrum of carbonate sedimentology and petrology. It is still one of the most comprehensive studies on this subject, although now obsolete in many areas. PIA(1933) devoted more than half of his book to the study of Recent nonmarine carbonates, an effort not since equalled.

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J. E. SANDERS AND G. M. FRIEDMAN

On the whole, these early pioneers in the study of carbonate rocks made little impact on the present generation of carbonate petrologists. A breakthrough in this field came during the last 15-25 years. The source of inspiration, energy, and progress, which culminated in this advance, was provided by the major oil companies of the United States. The enormous petroleum reservoirs discovered in carbonate rocks of the Middle East, in the Devonian of western Canada, and in the Pennsylvanian of west Texas prompted the oil companies to increase their efforts in carbonate research. In 1959, a Carbonate Rock Subcommittee was established as a division of the American Assoication of Petroleum Geologists, and a memoir (HAM, 1962) was published on the classification of carbonate rocks which is a major milestone and turning point in the history of the study of carbonate rocks. An annotated bibliography of papers on carbonates from the United States and Canada published in 1953-1958 and referenced by subjects was prepared by SANDERS (1960). ORIGIN OF LIMESTONES

Introduction

Limestones originate by three processes: ( I ) crystallization of calcium carbonate in an initially stony condition by means of inorganic, organic, or combined inorganic and organic processes to form caliche, nari, calcrete, hard skeletal reefs, travertine, and speleothems; (2) lithification of calcium carbonate grains in a manner analogous to the lithification of quartz grains to form sandstones; and (3) replacement of other materials, such as calcium sulfate or quartz, by calcium carbonate. Most limestones have been formed by lithification of calcium carbonate sediments; those formed by the other two processes, though of theoretical interest, are quantitatively insignificant by comparison. Crystallization of calcium carbonate in an initially stony condition

Limestones in which the calcium carbonate crystallized as stony material originate by inorganic and organic precipitation, or by a combination of both types of precipitation. In this discussion the term “inorganic” refers to crystallization outside the tissues of living organisms and “organic” to crystallization within the tissues of living organisms. Some organisms, such as Algae or bacteria, may play a significant role in changing the chemical conditions within the water and thus cause precipitation of calcium carbonate ‘(HASSACK, 1888; CHAMBERS, 19 12). Though such precipitation might appropriately be considered as “organic” in view of the large part played by organisms in it, this type will be treated here as “inorganic” because it took place outside the living tissues and was not a part of the metabolic processes of the organisms.


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Caliche, nari, and calcrete are examples of limestones that originate by inorganic crystallization processes; reef rock and stony biogenic incrustations, of limestones that originate by organic crystallization processes; and travertine, tufa, speleothems, and so forth, of limestones that originate by a combination of inorganic and organic crystallization processes. Caliche, nari, and calcrete Caliche, nari, and calcrete are limestones precipitated from evaporating soil moisture. Caliche forms in a semiarid climate where the predominant direction of movement of soil moisture is upward, owing to the excess of evaporation over rainfall. As carbonate-bearing waters are evaporated from the soil, calcite is precipitated between the soil particles. The term nari (from the Arabic word “nar” which means fire, alluding to its use in limekilns) was first introduced by BLANCKENHORN (1905) to designate a special variety of caliche that forms by surface, or near-surface, alteration of permeable calcareous rocks. Nari occurs in the drier parts of the Mediterranean climatic region, through the combined effects of dissolution and redeposition of calcium carbonate, Nari is distinguished by a fine network of veins, which surround unreplaced remnants of the original rock. The nari often contains clastic rock particles such as flint, hard limestone or dolostone and, locally, shells of sub-Recent terrestrial gastropods. The nari crust in Israel, which reaches a maximum thickness of at least 6 m, develops preferentially on Cretaceous and Tertiary chalks, but is also found on the chalky cementing material of conglomerates. It is thus widespread in Galilee, Mount Carmel, the Shefela, and the Judean and Shomron Mountains. i n regions of arid climate a surface soil crust may form, which is called duricrust (WOOLNOUGH, 1928). The term calcrete has been used where this crust consists of calcium carbonate (LAMPLUGH, 1902). Reef rock and stony biogenic incrustations Reef rock and stony incrustations represent deposits whose initial condition was solid and coherent, owing to the cementing effects of the original organic precipitation of calcium carbonate. i n modern reefs, corals, coralline Algae, and Foraminifera are the chief organic precipitaters of solid limestones. Other organisms, such as modern oysters and serpulid worms and fossil stromatoporoids, sponges, rudistids, and certain brachiopods, also cemented themselves to the substratum to form an initally stony mass (NEWELL,1955). Limestones formed by organic precipitation comprise the constructional limestones of J. H. JOHNSON (1957), the biolithites of FOLK(1959) and WOLF (1960), the residual organic class of R. W. POWERS(1962), and the incrustate limestones of SCHLANGER (1964). In addition to reef-building corals, Schlanger

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listed the Algae Lithothamnium and Lithophyllum, and the Foraminifera Carpenteria, Homotrema, and Gypsina as important genera of incrusting organisms. As noted by R. W. Powers, residual organic limestones are the most valuable paleoecological indicators. H. F. NELSONet al. (1962) referred to this type of limestone as “skeletal limestone”, but the term skeletal has been widely used to refer to the hard parts of organisms, whether in place or not, so that their usage. if adopted, would generate confusion.

Travertine, tufa, and speleothems Travertine and tufa refer to calcareous deposits of lakes, rivers, and springs, in which precipitation of calcium carbonate is a combination of inorganic and organic, chiefly algal, activity. Speleothems are limestones in which the precipitation, essentially inorganic, took place from cave waters. In each type the texture and shape of the deposit are governed by the detailed morphology at the site of precipitation; growth of new calcium carbonate takes place upon the surfaces of the former deposit. Texture, though variable, is characterized by indications of crystal growth on free surfaces and by a complete absence of indications of mechanical shifting of particles prior to final deposition. Travertine is the genetic term for all organic-inorganic nonmarine limestone accumulations formed in lakes, rivers, springs, and caves. It normally is colloform or concretionary, extremely compact in individual layers, but highly porous between layers. It may be wholly or partly hydrothermal. Surface deposits are commonly associated with calcareous Algae. Tufa, or calcareous sinter, is a highly porous spring variety. Onyx marble is a colorful banded variety. Travertine has been extended locally to include caliche and calcrete, but this usage is not recommended. The term travertine is derived from the Italian word, travertino, a corruption of tiburtino, “the stone of Tibur”, which is a former name of the locality now called Tivoli. The Italian travertine from near Rome is known by several local names, which include tufa litoide, the main building stone and material from which the catacombs have been excavated, and tufa granolare. LYELL(1 830, pp.207-210) reported three kinds of travertine here: hard and compact, cellular, and botryoidal-mammillary. Deep gorges at Tivoli show that the travertine is at least 150 m thick and that the diameter of some of the large spheroids with concentric structure reaches 2-2.6 m. Some authors distinguish travertine, as a fairly massive deposit that originated chiefly from deposition from hot spring waters, from calcareous tufa, as a spongy material deposited from ordinary spring and stream waters. According to HATCH et al. (1938, p.179), calcareous sinter (calc-sinter) is a synonym of travertine; both are regarded as products of precipitation from spring waters, especially those in volcanic regions. These authors followed I. C . RUSSELL’S (1885) usage of tufa for the deposits of Pleistocene Lake Lahonton, Nevada.


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GROUT (1940, p.170), on the other hand, defined travertine as “the precipitates from waters on plants”. He used “Mexican onyx” to designate the layered type of calcium carbonate precipitate in which the crystallization took place on flat surfaces. PIRSSON and KNOPF(1947, p.265) defined travertine as “the rock formed by precipitation of calcium carbonate at the orifices of springs”. They mentioned that deposition of carbonate was increased by the action of Algae. Classic localities for such calcareous deposits include Karlsbad, Czechoslovakia, where the famous Sprudelstein includes aragonite pisolites (COHN,1862); Tivoli, Italy; Vichy, France; Mammoth Hot Springs, Yellowstone Park, U.S.A. (WEED,1889);Pleistocene Lake Lahonton, Nevada (1. C . Russell, 1885); and Searles Lake, California (I. C. RUSSELL,1887; SCHOLL,1960). These and other deposits were reviewed by PIA(1933), whose work also includes an extensive bibliography. Dripstone and flowstone refer to speleothems made by dripping and flowing water, respectively. The deposits made by flowing and dripping are texturally distinct and can be easily recognized. Dripstone is limited to the central zones of stalactites; it is characterized by crystals oriented with their long axes parallel to the long axis of the stalactite. Flowstone surrounds the dripstone cores of stalactites and forms the entire mass of stalagmites; it consists of concentric layers of variable thickness in which the long axes of crystals are radially arranged with respect to the long axis of the stalactite or stalagmite (G. W. MOORE,1962). Where the dripping occurs into shallow pools, cave pearls are formed, analogous to ooids (BAKERand FROSTICK, 1947). Most speleothems consist of calcite, but aragonite does occur in warmer climates. Lithijkation of calcium carbonate sediments Lithification of calcium carbonate sediments is by far the most important factor in the origin of limestones. Conversion of the calcium carbonate sediments into limestone rock involves a large number of reactions and processes, which may leave original depositional textures essentially unaltered, on the one hand, or which may so completely rearrange the constituents that original depositional textures are completely obscured or obliterated, on the other. The subject of lithification, which is onlypart of the broader topic of diagenesis, has been much studied recently, both from the point of view of the chemical reactions involved (TAFT,1966b) and of changes in porosity and permeability (HARBAUGH, 1966). Both these contributions are found in the present two-volume work on carbonate rocks. The following discussion traces some of these changes in a sequence that proceeds from the stage of imperceptible alteration of original features of the sediment to the stage of complete obliteration of original features. The subject is conveniently, even if somewhat arbitrarily, subdivided into two parts: (I) changes of individual grains, and (2) changes within the grain mass.

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J. E. SANDERS A N D G. M. FRIEDMAN

Changes of individual grains Many of the changes that influence the individual grains in calcium carbonate sediments do not strictly involve lithification of these grains to form limestone rock; but because these grain changes almost invariably take place during this transformation, they are properly included in this general summary. Two kinds of changes occur: ( I ) changes in mineralogy of the grains without external modification of the texture; and (2) addition of external concentric coatings to the grains. Changes in grain mineralogy. Recent nearshore marine calcium carbonate sediments contain the minerals aragonite, high-magnesian calcite, and low-magnesian calcite. I n ancient limestones, however, of these minerals, only low-magnesian calcite is at all common; the other two occur but rarely (CHAVE,1954a, 1962; STEHLIand HOWER,1961; FRIEDMAN, 1964; see also TAFT, 1967b). An order of LOW- M CALClTir

A

ARAGONITE

HIGH-Mg

CALCITE

@

REEF SEDIMENTS

6

SKELETAL SANDS

Fig.1. Mineralogy of Recent carbonate sediments from Bermuda. (After FRIEDMAN, 1964, p.782; by permission of Journal of Sedimentary Petrohgy.)

179

ORIGIN AND OCCURRENCE OF LIMESTONES LOW-Mg C ALCl TE

ARAGONITE

HIGH-Mg CALCITE SUBMERGED IN SEA WATER

0

FROM MARINE SPRAY ZONE

A

BEYOND REACH OF MARINE SPRAY

Fig.2. Mineralogy of Pleistocene skeletal sands from Bermuda. Note that Pleistocene carbonate sands submerged in sea water or exposed to marine spray contain high-magnesian calcite and that those beyond the reach of the sea water are, with one exception, devoid of magnesium. The absence of magnesium is due to removal by fresh-water leaching at a post-depositional stage of subaerial exposure. (After FRIEDMAN, 1964, p.782; by permission of Journal of Sedimentary

Petrology.)

increasing stability among these minerals is: high-magnesian calcite, aragonite, low-magnesian calcite (STEHLIand HOWER,1961; CHAVE,1962; FRIEDMAN, 1964). The loss of magnesium from reef sands in the Bahamas was first reported by GOLDMAN (1926), long before the mineralogical varieties of calcite were recognized. The exact mechanism by which this change takes place is not known. It may consist of removal of magnesium from high-magnesian calcite to yield low-magnesian calcite, or a solution-and-deposition reaction on a micro-scale, in which high-magnesian calcite is removed and low-magnesian calcite deposited without textural changes in the grains involved (FRIEDMAN, 1964). Coralline Algae, for example, which secrete particles of high-magnesjan calcite, commonly persist in

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limestones with original texture preserved (Plate IB, C ) ; but after diagenesis, in limestones, these grains consist of low-magnesian calcite. The changes in grain mineralogy, as shown by FRIEDMAN (1964), are well illustrated by comparing Recent reef sediments and skeletal, and oolitic sands from Bermuda, the Bahamas, Red Sea, and other areas, with nearby Pleistocene limestones that consist of the lithified equivalents of these sediments. Recent reef sediments consist of aragonite in greater abundance than highinagnesian calcite (Fig. 1), indicating a preponderance of coral material. With only one exception, low-niagnesian calcite is absent in the Bermuda samples studied by FRIEDMAN (1964). The Recent skeletal sands from Bermuda contain these same minerals, but also include a variable amount of low-magnesian calcite, ranging from 10 to I5 %, the variation being due to the presence of terrigenous carbonate grains from Pleistocene limestones on the island (Fig. 1). The mineralogy of Pleistocene limestones formed by lithification of skeletal sands varies according to whether their postdepositional history involved complete subaerial exposure or partial subaerial exposure and partial exposure to sea water (Fig.2). Specimens of limestone that have been partially exposed to sea water contain variable amounts of high-magnesian calcite. Recent oolitic sands from the Bahamas contain abundant aragonite and small quantities of low-magnesian calcite, whereas in Pleistocene oolitic limestones, the abundance of aragonite decreases and that of low-magnesian calcite increases PLATE I Photomicrographs of Pleistocene limestones from Bermuda, which originated by modification and cementation of skeletal sands. (After FRIEDMAN, 1964, p.785; by permission of Journal of Sedimentary Petrology). A. Lime-mud envelopes outline former tests of organisms originally composed of aragonite, such as pelecypod shells or Halimeda, which were subsequently dissolved and the resulting void (mold) was later filled with a mosaic of drusy calcite. Original interparticle porosity has been completely occluded by drusy calcite mosaic. Pleistocene eolianite, Ferry Road, St. George’s Island, Bermuda. B. Skeletal fragments, which were originally composed of aragonite, are outlined by lime-mud envelope. The interior of these organisms has been lime-mud infilled by a mosaic of drusy calcite. Organism in extreme upper left of photograph is Halimedu. Note micritic lining of the internal tubes of Halimedu. Fossils which were originally made up of high-magnesian calcite, as the coralline algal fragment in upper left of photograph, have retained their original texture even though magnesium has been removed from this rock. The original interparticle porosity has been occluded by drusy calcite mosaic. Pleistocene eolianite, Ferry Road, St. George’s Island, Bermuda. Notice difference in texture between “rim cement” on exterior of “lime-mud envelope” and coarser mosaic of the cement. C. Molds with infilled drusy calcite mosaic in upper left of photograph are probably those of Halimeda. Note the well-preserved textural details of the coralline algal fragment (in right center of photograph) which was originally composed of high-magnesian calcite. Removal of magnesium from this rock has not changed the textural characteristics of this fragment. The original interparticle porosity has been occluded by drusy calcite mosaic. Pleistocene eolianite, Ferry Road, St. George’s Island, Bermuda. Notice difference in texture between “rim cement” on exterior of “micritic envelope” and coarser cement mosaic.

ORIGIN A N D OCCURRENCE OF LIMESTONES

PLATE I

181

182


(Fig.3, 4). These and other mineralogic changes associated with lithification are discussed by FRIEDMAN (1964). Paramorphic replacement of aragonite by low-magnesian calcite commonly occurs where the aragonite exists in the form of structureless lime mud, cryptocrystalline grains, or ooids (FRIEDMAN, 1964). Presumably this replacement of aragonite by low-magnesian calcite occurs with an intervening solution-deposition stage on a micro-scale. The result is a complete change of mineralogy of the grains but a retention of their original depositional texture. Thus, for example, ooids and aggregate grains of non-skeletal sand-size particles, which are deposited originally as aragonite in Recent calcium carbonate sediments (ILLING,1954; NEWELLet al., 1960; RUSNAK,1960b), almost always consist of low-magnesian calcite in limestone (LINCK,1903; FRIEDMAN, 1964). This paramorphic replacement contrasts markedly with the large-scale dissolution of aragonite skeletal debris and subsequent deposition of low-magnesian calcite in the empty space formerly occupied by the shells. (See section on Changes within the sediment mass.) The mineralogic composition of calcium carbonate in Recent deep-sea sediments differs from that of nearshore carbonate sediments in containing abunLOW-Mg CALCITE

A

Fig.3. Mineralogy of Recent ooids from Browns Cay, South Cat Cay, Joulters Cay, and Bird Cay (Berry Islands), Great Bahama Bank. (After FRIEDMAN, 1964, p.790; by permission of Journal of Sedimentary Petrology.)

ORIGIN AND OCCURRENCE OF LIMESTONES LOW-

183

Mg CALCITE

Fig.4. Mineralogy of Pleistocene oolites from New Providence Island, Hog Island (near New Providence Island), Andros Island, Long Rock Island (near Andros Island), and North Bimini Island, Great Bahama Bank. (After FRIEDMAN, 1964, p.790; by permission of Journal of Sedimentary Petrology.)

dant low-magnesian calcite and little or no high-magnesian calcite and/or aragonite (STEHLIand HOWER, 1961;FRIEDMAN, 1965b). Carbonate pelagic sediments of the Red Sea at depths of 1,000-2,000 m consist of high-magnesian calcite and lowmagnesian calcite with aragonite being sporadic or absent, whereas those from much greater depth (3,676 m) in the Indian Ocean contain only low-magnesian calcite (FRIEDMAN, I965b). Asample from the East China Sea at a depth of 1,370 m also contained only low-magnesian calcite (STEHLIand HOWER, 1961). The mineralogy of deep-sea carbonate sediments is controlled in part by sedimentation and is a reflection of the composition of the planktonic skeletal remains that are the chief contributors to the sediment, and in part by dissolution of high-magnesian calcite and aragonite. In the deep-sea sediments studied Globigerina is one of the most abundant constituents. Inasmuch as Globigerina tests consist of low-magnesian calcite, that mineral species predominates in the material which falls to the bottom. Coccoliths, which constitute other important planktonic sources of calcium carbonate in deep-sea sediments, also consist of low-magnesian calcite. Removal of aragonite and high-magnesian calcite by dissolution, as a function of

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water depth and grain size, may be an important diagenetic effect. Under deep-sea conditions low-magnesian calcite is the most stable of these carbonate minerals and high-magnesian calcite appears to be more stable than aragonite (FRIEDMAN, 1965b).

Additions of concentric coatings to grains. Grains of calcium carbonate and other minerals, such as quartz or feldspar, found in environments of Recent calcium carbonate sedimentation, commonly show external concentric coatings of calcium carbonate. These occur as envelope(s) of unoriented cryptocrystalline aragonite crystals (“lime mud”, “paste”, or “micrite” of different authors), as layers of clear, oriented needle-like aragonite crystals with concentric or radial structure, or as various combinations of these. It should be added that some grains consist entirely of the same material found as coatings. In such cases, grains are composed either entirely of one type or of combinations of two or more types of coating material. The addition of coatings to individual grains is not strictly a lithification process; in fact, coatings are added while the grains are still in the final stages of sedimentation. Because the fate of original textural features during subsequent lithification processes is greatly influenced by the presence of coatings, particularly of “lime mud”, however, the subject of coatings is included here. A . Envelopes of unoriented cryptocrystalline aragonite (“lime mud”). One of the commonest types of grain coatings consists of unoriented cryptocrystalline aragonite (SORBY,1879; ILLING,1954; NEWELLet al., 1960; RUSNAK,1960b; BATHURST, 1964; FRIEDMAN, 1964; and others). The coatings have been called “micrite envelopes” (BATHURST, 1964)or “micritic envelopes” (FRIEDMAN, 1964), but these imply “rock” by the connotation inherent in “micrite”. These structureless coatings (Plate IA-C) are acquired in the depositional environment, but the mechanism of their formation is not fully understood. Many of them originate by the same accretion process that is involved in the aggregation of tiny aragonite crystals to form sand-size grains, as in the Recent non-skeletal carbonate sands of the Bahamas (ILLING,1954). Other coatings may owe their origin to the perforating Algae or the lime-mud infilling of peripheral zones of grains that have been bored by organisms. The grains may have been the substrate on which microorganisms lived, so that the envelopes formed as a result of physiological processes of these organisms (FRIEDMAN, 1964, p.808). Crusts may also be due to Algae(WoLF, 1965a). Whatever their origin, these envelopes are commonly responsible for the preservation of original shapes of aragonite grains, for they are typically more resistant to dissolution than an aragonite skeletal particle. Thus, they preserve the shape of the coated particle, even if the aragonite of this particle has been removed by dissolution. If the entire grain consists of lime mud, it may be a faecal pellet, intraclast (FOLK,1959), or pseudooid (CAROZZI,1957). The latter is a more general term that can be applied when the origin is unknown; it simply refers to a grain without a clear coating of concentric or radial aragonite crystals. This is a slightly different


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definition than the one thought to be the original (as “pseudo-ooliths”) and which referred to clastic or worn calcite grains or shell fragments that resemble “ooliths” (GRABAU,1932, p.284, after J. G. Bornemann, reference not cited, but possibly BORNEMANN, 1885). B. Clear layers of oriented aragonite crystals with concentric or radial structure. Grains that have acquired a coating of clear aragonite crystals with concentric or radial structure in the arrangement of the crystals have been called superficial ooids or normal ooids (or “true ooids”, or simply “ooids”, which is the term preferred to oolith; oolite is the rock composed of cemented ooids), depending on the number of coating layers. SuperJicial ooids contain only one shell (of variable thickness) with concentric or radial structure, and normal oiiids contain two or more shells (of variable thickness) having concentric or radial structure (CAROZZI, 1957). This definition of superficial ooid is preferred to those based on thickness of the shells, such as “only thin external oolitic coating” (ILLING,1954, p.36) or “thickness of the coating envelope being less than the radius of the nucleus” (LEIGHTON and PENDEXTER, 1962, p.38). Though originally composed of aragonite, these coatings nearly always consist of calcite in limestones. Inasmuch as more than one coating may be present and one or more of the different types may alternate on the same grain, great variation is possible. A central nucleus of foreign material may be surrounded by alternating envelopes of structureless lime mud and clear layers of oriented aragonite crystals with either concentric or radial structure. In other grains, a foreign nucleus may be lacking and the entire grain may consist of material that otherwise forms coatings. The different types of texture of the aragonite have been attributed to the rate of precipitation of aragonite and degree of agitation (RUSNAK,1960b). According to this explanation, rapid precipitation produces envelopes of unoriented cryptocrystalline aragonite; slower precipitation combined with weak agitation produces radially arranged crystal aggregates; and very slow precipitation in combination with vigorous agitation produces a preponderance of mechanical attrition that re-orientates the aragonite crystals from an initial radial attachment to a tangential (concentric) growth position. Whatever their origin, the sequence of coating and particle types may be of great paleogeographic significance in defining ancient shorelines. In Great Salt Lake, Utah, U.S.A., pseudooids form a continuous belt just offshore from the zone of breaking waves. In the breaker zone the particles may consist entirely of normal ooids; if so, then a belt of mixed ooids (and superficial ooids, not separately recognized) and pseudooids intervenes between the normal ooids and pseudooids offshore. In other localities where normal ooids occur in the breaker zone, a narrow landward belt of mixed ooids and pseudooids exists. In still other localities, there is no belt of pure ooids, but only a shoreline zone of mixed ooids and pseudooids (CAROZZI,1957).

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Changes within the sediment mass Changes within the sediment mass involve selective dissolution, which increases pore space; precipitation of mineral cement in pore spaces, which decreases pore space; and recrystallization, the effect of which on pore space is not established. Precipitation of mineral cement between grains or mineral infilling of molds generally does not disturb the texture, but selective dissolution and recrystallization tend to destroy original textures. Selective dissolution. Selective dissolution is the opposite of lithification, but it is included here because it so commonly occurs during or before lithification has been completed. Aragonite, the most important constituent of nearshore marine calcium carbonate sediments, may be dissolved away completely or break down into a whitish powder if the sediments come into contact with meteoric ground water. The relationships are especially well illustrated by the material from the deep borings on Eniwetok and Bikini atolls, in the Pacific Ocean. Unlithified calcium carbonate sediments in which aragonite still persists are found at levels in the borings where the calcium carbonate sediments may be inferred to have been continuously exposed to sea water. Evidently time alone, for spans at least as long as the Cenozoic, is not a relevant factor in the stability of aragonite as long as it remains in contact with sea water. At levels where aragonite has been dissolved away, the calcium carbonate sediment has always been converted to calcite-cemented limestone rock; such levels coincide with indications of subaerial exposure and contact with meteoric ground water (SCHLANGER, 1963, p.997). Locally, aragonite has been preserved even in material that has been subaerially exposed, but, typically, aragonite has been dissolved and a calcite cement precipitated where such exposure has occurred. The dissolved aragonite doubtless provides the material which precipitates as calcite cement (SORBY,1879; see also section on Precipitation of mineral cement). Where skeletal remains consisting of aragonite have acquired a lime-mud envelope, the aragonite fragment on the interior may be dissolved away completely. This leaves an empty mold preserved within the lime-mud envelope, which itself may have undergone paramorphic replacement from aragonite to calcite (FRIEDMAN, 1964). The empty space thus created has been called “moldic porosity”, or “secondary void porosity” (R. W. POWERS, 1962, p.140, pl.IV, fig.3). It may later be filled in with mineral matter (Plate IA-C), the textural arrangement of which indicates that crystallization took place into an empty space, the crystals having grown inward from the marginal walls (BATHURST,1964). On the other hand, the original depositional lime-mud matrix between grains rather than the interior of a lime-mud-coated grain may be selectively dissolved. In the subsurface Devonian crinoidd limestones of the Andrews South oil field, Texas, visible porosity exists in the spaces between crinoid ossicles where an original interparticle lime-mud matrix has been selectively dissolved away. The


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presence of the lime mud inhibited precipitation of clear calcite, which cemented nearby originally mud-free sediments into a non-porous rock (LUCIA,1962). Selective dissolution may also affect the lime-mud matrix of a mixture of lime-sand and lime-mud particles after the growth of a framework of dolomite crystals (R. W. POWERS, 1962, p.185). Soft algal tissue may be removed after initial cementation of the framework grains, leaving empty spaces that may later be filled by mineral cement or internally sedimented clasts (WOLF,1965b). Precipitation of mineral cement in pore spaces. Precipitation of mineral cement in the pore spaces within a mass of calcium carbonate sediments is the most important single lithification process. The pore space may have existed originally between the essentially unaltered grains of the sediment, or it may have been created by selective dissolution of the lime-mud matrix between grains or of the soluble centers of grains within lime-mud envelopes, as previously discussed. The cement commonly begins as clear crystals that grow radially outward from the grains (or inward toward the center of pore spaces) (ILLING,1954; GINSBURG, 1957; R. W. POWERS, 1962; and others). Where crystal growth does not occupy all the available space, the cementation is incomplete and the cement resembles a grain coating; it has been called “drusy coating” (R. W. POWERS, 1962) or “rim cement” (GINSBURG, 1957). Recrystallization of a lime-mud coating, or of original lime-mud matrix between grains, may produce an end product the appearance of which is similar to mineral cement. The cementing minerals closely reflect the chemical environment within the interstitial waters, the composition of which may range from fresh to supersaline. Mineral cements generally have been interpreted in terms of exposure of aragonitic marine sediments to meteoric ground water, but marginally evaporating sea water may diffuse upward through porous intertidal and supratidal sediments and precipitate “evaporitic” cements. Through-flowing connate waters of varying salinity and commonly having high pH, which are released during compaction, may also introduce cements of varying composition. If aragonite grains come in contact with meteoric water, they dissolve forming calcium bicarbonate, according to the reaction: CaC03

+ HzO + CO2 ~2Ca(HC03)z

Precipitation of low-magnesian calcite as a drusy mosaic in pore spaces takes place when the solution becomes saturated and then loses carbon dioxide or water, either of which may easily occur under subaerial conditions where evaporation is possible (FRIEDMAN, 1964, p.809). Plate I1 illustrates the progressive development of calcite cement in pore spaces. As noted previously, at levels in the Eniwetok and Bikini borings where dissolution of aragonite has occurred, calcite cement has always precipitated. By contrast, the Cretaceous chalk, which consists of tiny skeletal fragments

PLATE 11

Photomicrographs of Pleistocene skeletal sands (limestones) from Bermuda. (After FRIEDMAN, 1964, p.785; by permission of Journal of Sedimentary Petrology.) A. Calcite cements skeletal fragments only near grain-to-grain boundaries. Interparticle porosity has been completely preserved except near grain-to-grain boundaries. Somerset Eolianite, Wellington Hill, St. George’s Island, Bermuda. B. Drusy calcite mosaic cements skeletal fragments only near grain-to-grain boundaries retaining interparticle porosity where grains are not in close juxtaposition. Pembroke (?) Eolianite, Grape Bay, Bermuda. C. Drusy calcite mosaic cements skeletal fragments and partially occludes interparticle porosity. Southampton or Somerset Eolianite, top of Park Gates, St. George’s Island, Bermuda.


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that are composed almost exclusively of calcite, has remained uncemented where aragonite was initially absent. In the so-called “hard chalks”, which include larger proportions of organisms with originally aragonitic skeletal debris, however, the aragonite has been dissolved so that only molds remain and a calcite cement is present (HATCHet al., 1938, p.168). Aragonite, or the “evaporite” minerals, such as dolomite, anhydrite, gypsum, barite, celestite, or halite, have also been found as the mineral cement between grains of calcium carbonate deposits; in addition, the “cement” may replace the grains. Such cements may be a great deal more common than is generally realized; they generally do not persist on the outcrop and are preserved in bore-hole samples only where saline drilling fluids are used. One explanation for the presence of these minerals has recently been provided by KINSMAN(1965), who studied the minerals found in the intergranular spaces of calcium carbonate sediments exposed on marginal supratidal flats along the Persian Gulf. Owing to extreme aridity, sea water transpires upward through these porous sediments. With evaporation, the concentration of dissolved salts progressively increases in the interstitial waters, and eventually evaporite minerals precipitate. Kinsman has found that a systematic sequence of evaporite cements exists in a landward direction from aragonite, which precipitates at lower salinity, to halite, which precipitates at higher salinity. This may be visualized as a kind of “marine caliche” process, for the mechanism of upward-moving pore water is exactly the same as that of meteoric ground water, which is the pore water where caliche forms. The difference is that the pore water in the sediments on the supratidal flats is sea water. Deposition of evaporite minerals within the sediment pores may occur even in the absence of associated bedded evaporites, indicating that the salinity of the adjacent body of sea water did not reach extremely high values. An absence of halite from most beach rocks and many littoral flat deposits suggests tidal flushing, a process which involves deposition of halite at low tide and its removal (owing to high solubility) during the next high cycle or seasonal high-water stage. There is a difference in cements within beach rock reported from different localities; this may be explained by considerations of the type of water present within the pores of the beach sediment prior to deposition of the cement. Aragonite has been found to be the cement in many beach rocks (GINSBURG, 1953, 1957; REVELLE and FAIRBRIDGE, 1957; STODDART and CANN,1965), whereas calcite has been claimed as the original cement in others (R. J. RUSSELL,1963a, b). Both aragonite and calcite have been found in Bahamian beach rock (ILLING,1954, p.70). As with other types of aragonite precipitates, however, initial aragonite cement in beach rocks changes eventually to calcite when exposed to meteoric ground water, so that calcite may be the final product even if the initial material was aragonite. Gypsum is found as a beach rock cement in Somaliland today (R. W.

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Fairbridge, personal communication, 1965). Anhydrite is the local intergranular cement in the Queens, Grayburg, and San Andres Limestones (Permian) of west Texas (G. M. Friedman, personal observations), and in the Arab-D Limestone (Upper Jurassic) of Saudi Arabia (R. W. POWERS,1962, p.137, ~1,111,fig.4). Gypsum, barite, and celestite constitute local sulfate cements in some carbonate deposits (G. M. Friedman, personal observations); halite is also known as a cement. Dolomite may also be precipitated as a mineral cement within the pores of calcium carbonate sediments (SHINNand GINSBURG, 1964); this topic is considered separately in Chapter 6. Recrystallization. Recrystallization is the name applied to the vaguely understood process in which original textures are obliterated and a mass of large interlocking crystals forms. The process may operate in either fine-grained or coarse-grained sediments, and it may be selective, involving only parts of the material, or complete, involving all materials. Selective replacement commonly occurs where lime mud is present. The result is a transformation of aragonite particles, which are not easily seen as individuals in thin-sections of standard thickness and consequently appear “turbid”, into a mosaic of coarser, clearer, transparent sparry calcite (BATHURST, 1958, 1959; FRIEDMAN, 1964). The possibility that this transformation occurs is important for limestone classification, so the evidence for it will be considered with additional examples beyond those cited by Bathurst. The Middle Ordovician flysch exposed on the north shore of the Gasp6 Peninsula, Canada, includes a great variety of rocks, among them calcareous wackes (ENOS,1964). The wackes are typically graded, and the lower parts of the graded beds consist of coarser grades of terrigenous fragments set in a mosaic of coarse sparry calcite. The grain size of both terrigenous fragments and sparry calcite decreases upward until at the top the bed becomes a typical calcisiltite composed of silt-size grains of terrigenous (recycled) carbonate fragments in a fine-grained lime-mud matrix. The sparry calcite in the lower part of the bed clearly represents the end product of selective recrystallization of an initial lime mud matrix that was deposited mechanically along with the coarser terrigenous grains. Recrystallization has been impeded near the top of the bed, where the permeability was less owing to the finer grain size. Lime mud has also been thought to have recrystallized to sparry calcite in the Tertiary limestones of Guam (SCHLANGER, 1964, p.D-10). In these rocks an initial lime-mud envelope has been inferred to have recrystallized to form a halo of clear calcite, with crystals oriented perpendicular to the rim of the fragment; first recrystallization took place next to skeletal fragments and other larger grains. An alternative, not discussed by Schlanger, is that the calcite represents a rim cement. This illustrates the problem and the difficulties encountered in trying to formulate a clear-cut interpretation. The most satisfactory results are achieved by following


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the criteria of grain-support and mud-support set forth by R. J. DUNHAM (1962). The indications that a primary lime mud can recrystallize to form coarsegrained sparry calcite, complicates petrographic classification of limestones in which major subdivision is based on the assumption that mosaics of clear sparry calcite always indicate cleanly washed original sediments with numerous primary voids into which the sparry calcite was deposited as a mineral cement (FOLK,1959). Although this assumption may be valid in many cases, it should always be used with caution; a better alternative is to follow R. J. DUNHAM’S (1962) suggestions about major subdivision on the basis of grain-support or mud-support, rather than use sparry calcite alone. Recrystallization of skeletal material in certain limestones from Lau, Fiji, was discussed by CRICKMAY (1945); from Saipan, by J. H. JOHNSON (1957); and from Guam, by SCHLANGER (1964). The following order of decreasing susceptibility to recrystallization was observed: corals (least resistant), Halimedu, mollusks, pelagic Foraminifera, beach-type Foraminifera, larger Foraminifera, echinoids, and calcareous red Algae (most resistant). On Saipan, green Algae showed a more variable susceptibility (J. H. JOHNSON,1957). In the Saipan limestones, recrystallization generally began in the matrix and later attacked skeletal remains, both from inside and outside. The end product is a crystalline limestone with only “dust” outlines, color bands, or textural differences to indicate former skeletal debris (J. H. JOHNSON,1957). In the Guam limestones, the originally aragonite coral fragments have completely recrystallized into a calcite mosaic of varying grain size. The original coral shape is suggested by traces of dark mud in the interseptal spaces, by fine dark lines that preserve the coral outline, and by algal coatings on fragments, which preserve the original outline of the fragments (SCHLANGER, 1964). PURDY(1963) has shown that ooids, corals, and Foraminifera, particularly Peneroplids, recrystallize in the marine environment to cryptocrystalline aragonite. All stages of recrystallization have been observed in the transition of various skeletal and non-skeletal constituents into cryptocrystalline aragonite. Although both the original grains and final recrystallization product are aragonite, recrystallization has destroyed original textures. Experimental work Experimental work planned to duplicate the mineralogical changes that take place during lithification of calcium carbonate sediments was initiated by TAFT(1962) and FRIEDMAN (1964). TAFT (1963) demonstrated that replacement of saline by fresh water played an important role in aragonite conversion to calcite. FRIEDMAN (1964) showed that both the p H and salinity of the waters are important factors which control the mineralogical and, by inference, textural changes in calcium carbonate sediments. According to his experiments, p H is a more important controlling factor than salinity.

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Replacement of other materials, clziefly calcium sulfate and quartz, by calcium carbonate

A few limestones have originated by replacement of other materials, chiefly calcium sulfate and quartz grains. Such limestones are probably not common, but they are important in that their origin differs so drastically from that of other limestones, and because those formed by replacement of calcium sulfate may be associated with free sulfur deposits. The two most common reactions are bacterial conversion of calcium sulfate to calcite and carbonate replacement of quartz grains. i Bacterial conversion of calcium sulfate to calcite Sulfate-reducing bacteria, which feed on organic matter, obtain oxygen by breaking down sulfates, such as calcium sulfate (gypsum or anhydrite). Calcite is formed when carbon dioxide, either liberated in the bacterial reaction from the organic matter, or derived externally, combines with calcium oxide released from the calcium sulfate by the bacteria. A somewhat simplified reaction may be presented as follows:

+ 2(CH20) -+ 2 Ca0 + 2 s + 2 COz + 2 H20 + + 2C02 -+ 2CaC03

2CaS04 2Ca0

0 2

This process is thought to have been responsible for the production of limestone cap rocks atop salt plugs, as in the Gulf Coast region of the United States (FEELY and KULP,1957). It may also have caused the replacement of the Castile Anhydrite (Permian) along faults in the Delaware Basin, west Texas, which has given rise to local limestones that form topographically resistant landforms called “castiles” (ADAMS,1944, pp.1606-1607). Replacement of Castile Anhydrite resulted in a laminated limestone; where massive Castile Anhydrite has been replaced, a massive limestone has resulted (G. W. MOORE,1960). In addition, bacterial processes may have been responsible for the calcite-rich layers in the calcite-aragonite couplets in Recent sediments of the Dead Sea (NEEV, 1963, 1964; see also section on Calcium carbonate lake sediments). A similar bacterial replacement origin has been proposed as an explanation for the thick-bedded, sparsely fossiliferous parts of the Capitan Limestone (Permian) of southeastern New Mexico (G. W. MOORE,1960). According to this interpretation, the main body of the Capitan Limestone is the lateral equivalent of the upper massive Castile Anhydrite, and is not entirely an older barrier reef, as is generally supposed (NEWELLet al., 1953). The massive, unfossiliferous parts of the Capitan Limestone have been explained as the product of recrystallization of the reef rock, but it retains no vestiges of organic remains; rather, it is indistinguishable petrographically from the massive limestone formed on the “castiles” in the Delaware Basin, where massive anhydrite has been replaced.


I93

Carbonate replacerncnt of quartz grains Carbonate minerals, such as calcite, dolomite, siderite, or ankerite, commonly replace quartz grains; a favored site of replacement is near the upper contact or updip pinch-out of sand bodies, particularly lenticular ones. In the Cardium Sandstone (middle Upper Cretaceous) at Pembina, Alta., Canada, siderite-bearing calcareous sandstone forms an impervious zone at the upper contact of the sandstone where it underlies a conglomerate. The calcium carbonate content of analyzed samples ranges from 10 to 50 %. The quartz grains have been markedly corroded and replaced by the calcite cement. Replacement of the quartz grains proceeded from the outer edges inward; locally, calcite ramified through quartz. Several stages of replacement were traced from peripherally corroded and embayed quartz grains to residual quartz embedded in a calcite pseudomorph (G. M. Friedman, unpublished observations). Where replacement has been complete, no traces of the former quartz may be found, or, at best, only a few corroded quartz relicts may remain. WALKER (1 960) noted that the replaced grains are preserved as carbonate pseudomorphs, which can be mistaken for grains of clastic carbonate. Thus, sandstones could, by replacement, be converted into limestones. Although regarded as “uncommon” (NICHOLAS, 1956, p.6), this type of replacement may be rather common, and limestones formed in this manner may be more widespread than is generally realized.

OCCURRENCE OF LIMESTONES

Introduction Limestone occurs in all parts of the world and at all levels in the stratigraphic column, though probably it is less common in ancient Precambrian strata. Limestones have been estimated to comprise 19-22% of available measured stratigraphic sections (LEITHand MEAD, 1915; SCHUCHERT, 1931; PETTIJOHN, 1957). Slightly lower values for total limestone abundance in the continental crust have been calculated from geochemical material balance analyses, ranging from 5 to 14% according to MEAD(1907), CLARKE (1924), or HOLMES (1937). Although outranked by “shales” (44-58 % of measured sections, whereas geochemically calculated values range from 70 to 82 %) and “sandstones” (22-37 % of measured sections and 12-1 6 % based on geochemical calculations), limestones perhaps have received proportionately more than their share of geologic study from the earlier generations of paleontologically-oriented stratigraphers, who were particularly concerned with the great abundance of marine fossils in limestones. Figures for limestone in the deep-sea crust are not yet available. In more recent times, the interest in limestones has continued to expand partly for their fossil content, partly because many limestones are petroleum

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reservoirs (HARBAUGH, 1967), and partly because it is possible to make rather exact environmental comparisons between limestones which preserve information about original calcium carbonate sediments and their Recent analogues (TAFT, 1967a); both of these contributions are found in the present volume. The occurrence of limestones may be considered from many points of view. One view, which became particularly common in the United States in the first half of the 20th century, was to assign all calcium carbonate rocks to the single group, “limestones”, and to contrast them with various kinds of noncarbonate deposits. Thus, “limestone” tended to denote a “chemical” rock, in contrast with “terrigenous” (or "elastic") rocks, which were subdivided on the basis of grain size and/ or mineralogic composition. This approach essentially ignored the variations within limestones that reflect initial variations in the calcium carbonate sediments from which most limestones have been derived. Another approach, which also tended to obscure the variations within the original carbonate sediments, was to subdivide limestones according to different gross tectono-stratigraphic “facies”, such as “platform” limestones, “basin” limestones, and “geosynclinal” limestones (SLOSS,1947). “Platform” limestones, for example, typically are extensive, thin blankets of light-colored rock that commonly contain abundant fossils and have been extensively dolomitized (especially those of Paleozoic age). “Basin” limestones may resemble “platform” limestones lithologically, but locally they are much thicker and pass laterally into reefs and/or the deposits of barred basins. “Geosynclinal” limestones are identified by their occurrence in thick sequences, siliceous composition, dark color, and interstratification with deep-sea shales. More recently emphasis has been placed on analysis of limestones in terms of the different kinds of carbonate sediments from which most of them have been derived, and to use primary lithofacies and biofacies as a basis for paleoecologic reconstructions. This paleoecologic approach is followed and extended here. The occurrence and distribution of limestones is considered to be a stratigraphic problem. Such sediments result from the interplay between the instantaneous lateral pattern of geographic environments of deposition and the vertical distribution of these environments through time, which governs the build-up of the stratigraphic record. Such an environmental-stratigraphic approach applies generally to all sediments; not only to pure carbonate sediments, but also to noncarbonate sediments and mixtures of these two types. Before pursuing this approach further, it is necessary to review the kinds of particles found in calcium carbonate sediments and to explore the interplay between the types of sedimentation processes that influence these particles and that are responsible for creating stratigraphic units.

Kinds of calcium carbonate particles in carbonate sediments Calcium carbonate particles in carbonate sediments consist of two major types:


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(I) recycled particles eroded from older carbonate rocks, and (2) first-cycle particles that originate from the water in the depositional basin itself. Recycled, or allochthonous, calcium carbonate particles belong in the category of ordinary terrigenous grains; they have been derived by erosion of older carbonate rocks on land and delivered to the site of deposition from an external source. Considerable confusion has grown up in connection with the nomenclature of such grains and rocks composed of them. The terms “mechanical” (FERAYet al., 1962, p.41), “detrital” (LEIGHTON and PENDEXTER, 1962, p.35), and “clastic” have all been applied in the same sense as the term “terrigenous”, which is used here; but these other terms are not recommended because they have also been used to refer to the “physical” transportation of carbonate grains by currents as opposed to the “chemical” deposition out of the water. All kinds of carbonate particles, no matter what their origin, behave as quartz grains in a water current, and this behavior has been expressed by applying the terms “mechanical”, “detrital”, “clastic”, or “fragmental”. As pointed out by FOLK (1959), use of the term terrigenous for recycled particles avoids this possible ambiguity. Limestones formed of recycled grains have been called “calclithites” (WOLF,1960). First-cycle, or autochthonous, calcium carbonate particles (allochemical coiistituents of FOLK, 1959; 1962) include skeletal remains of organisms, ooids, superficial ooids, pellets, lumps, aggregate grains, pseudooids (CAROZZI,1957) and intraclasts (FOLK,1959, 1962), all of which owe their origin to various processes from the water or bottom of the depositional basin. Only distinctive skeletal remains are restricted to marine environments; the other types of first-cycle particles may originate in lake water as well as sea water. For further details on the kinds of particles in calcium carbonate sediments the reader is referred to TAFT (1966a) in this volume. In the case of first-cycle carbonate sediment particles, further complication arises due to the fact that they may be deposited in approximate environmental equilibrium, that is, near the place where they originated. They also may be displaced by gravity down subaqueous slopes that typically are steep at the margins of banks, where marine carbonate sediments are forming, and ultimately may be deposited far from their place of origin, in the same manner as recycled terrigenous grains. A remarkable and unique feature of marine calcium carbonate sediments is that first-cycle particles can be produced in enormous abundance from sea water within a very short time. Hence, thick accumulations of pure calcium carbonate sediments may originate even in the absence of any external source of terrigenous grains. This permits sedimentation of calcareous material to keep pace with subsidence of the basin bottom, and this balance may produce thick piles of shallowwater carbonate sediments. For example, approximately 1,500 m of carbonate sediments and limestones derived from calcium carbonate sediments were drilled in hole F-1 on Elugelab Island, Eniwetok Atoll (1 l”40’ N 162”12’ E), the age of

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I. E. SANDERS AND G. M. FRIEDMAN

which ranges from Eocene to Recent. In the Funafuti boring (8 “30’ S 179” E), approximately 330 m of carbonate deposits range in age from Pliocene to Recent; and in the Kita-Daito-Jima boring (26” N 132” E), 435 m of carbonates range in age from Miocene to Recent (SCHLANGER, 1963). In the Bahamas, the well drilled by Superior Oil Company and others at 26’52‘37.2’’ N 78”01’54.7” W, at Stafford Creek, Andros Island, Bahamas, penetrated 5,000 m of carbonates the age of which ranges from Late Cretaceous to Recent (Marie Spencer, in: EARDLEY, 1951). Depositional processes involved in accumulation of marine carbonate- and mixed carbonate- and noncarbonate sediment units In spite of their distinctive origins, first-cycle biological carbonate grains (or other types) respond to various common physical processes of sedimentation just as terrigenous (recycled) carbonate and noncarbonate particles do. Accordingly, they follow the same principles that govern accumulation of almost all stratigraphic units. The physical depositional processes involved may be classified according to the predominance of vertical or lateral sedimentation activities and also on the basis of “energy”. Vertical sedimentation processes. Vertical sedimentation processes involve the more or less simultaneous vertical accumulation of particles that have been derived ( I ) from fallout of skeletal material of planktonic organisms distributed in the water, or suspended load carried by the water; (2) from growth of benthonic shell beds on the bottom, scattered or in the form of in situ biostromes; or (3) from vertical upgrowth of reefs or bioherms. Each of these vertical processes is characterized by upward accretion at the top of the growing layer of sedimentary material. The lateral spreading out or diffusion of the materials involved is not directly included in the depositional process; but rather takes place previously within the water as a result of migration or spreading of living organisms or a mixing of fine-grained sediment particles by turbulence within the water. Lateral sedimentation processes. Lateral sedimentation processes involve the sideways shift along the bottom of sediment already present, or the change in position of environments in which particular kinds of sediments may be restricted. Lateral shifting of sediment along the bottom may be due to bottom currents within the water mass; or it may be due to gravity displacement by turbidity currents, fluidized layers of flowing grains (SANDERS, 1965), or slumps, which cause sediment to move along the bottom independent of water currents. Gravitydisplacement processes require subaqueous slopes, and consequently their effects are more commonly found in sediments deposited in “deep” rather than in “shallow” water. Change of position of environments may take place as a normal consequence


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of sediment reworking, or it may reflect major geographic changes. An example of a change of position of an environment as a normal consequence of reworking is the lateral displacement of creeks on tidal flats due to undercutting of the concave bank and accretion on the opposite convex bank (VANSTRAATEN, 1951; 1952). Examples of changes of position of environments that are due to major geographic changes include those brought about by changes of sea level or changes in abundance of sediment supplied. The shoreline and associated sediment belts may move seaward as a result of seaward progradation of nearshore sediments on beaches or reefs, or because of emergence; these sediment belts may move landward as a result of submergence. These lateral shifts of nearshore sediment belts produce offlapping (“regressive”) and onlapping (“transgressive”) stratigraphic patterns, respectively, in which the patterned vertical sequence eventually formed obeys Walther’s Law of Facies (WALTHER,1893-1894). According to this law the vertical succession at any point is a reflection of the lateral distribution of the adjacent sediment belts at any instant in time.

Physical processes based on “energy”. Another approach to physical depositional processes is to subdivide them on the basis of “energy” into two groups: ( I ) “highenergy” and (2) “low-energy”. Depositional areas of “high energy” are inferred to produce clean, well-sorted, coarse-grained carbonate sediments with a maximum content of broken skeletal remains; whereas areas of “low energy” are held to be responsible for “muddy” deposits that are entirely fine-grained or contain abundant fine-grained matrix between coarser particles, with broken skeletal remains being less common. For further discussion of this approach with several intermediate categories, the reader is referred to the work of PLUMLEY et al. (1962). This analysis is certainly valid to a first approximation; however, the choice of the word “energy” is perhaps unfortunate, for what is implied is not the total energy of the system (which includes also chemical and biological components), but the mechanical effects on the bottom of turbulent motion in the water. In general the surf zone and bottoms subjected to constructed tidal flow (zone of maximum mechanical breakage) are sites of lag concentrates of whatever coarse material is present; any fines that arrive or are generated here by abrasion of coarse particles will remain in turbulent suspension and are moved elsewhere. The interpretation of bottoms subjected to less turbulent water motion is more complex. In the first place, fine-grained sediment may be absent. Inasmuch as particle supply is chiefly dependent on organisms for the most part, only coarse-grained sediment may be available. This produces an anomalous result: coarse-grained sediment in the absence of current turbulence. In addition, particles may be broken by several biological processes, including activity of boring and burrowing organisms, so that breakage need not always signify “high energy ’. Finally, grasses or sediment-binding organisms may trap sediments of finer grain sizes, which may be produced locally by several processes, and thus cause mud

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accumulation even where some currents are present. Examples illustrating these various alternatives have been described from the Florida reef tract by SWINCHATT (1965). Thus, although the “energy” scheme is useful, it should be applied with great caution.

Environmental settings of calcium carbonate deposits Recent calcium carbonate sediments occur in stratigraphic successions that are accumulating in three main environmental settings: ( I ) nonmarine environments, commonly within large inland structural basins; (2) shallow-marine and intermediate-marine platform environments; and (3) open-sea basins. Although certain attributes are common to all the carbonate sediments formed in any of these environmental settings, other attributes are more restricted and hence may be diagnostic characteristics of the environment in which they occur. For example, certain particle types, such as ooids, have been found in deposits from all of these environmental settings; other particle types, such as coccolith remains, however, are generally restricted to open-sea sediments, although they also occur in reef lagoons around the Indian Ocean (R. W. Fairbridge, personal communication, 1965). Aside from particle types, other criteria distinctive of different environments of deposition are now being actively sought by current research; some of these may be found in diagnostic sedimentary structures, geometric shapes of deposits, and stratigraphic relationships with adjoining deposits. Diagnosis is generally achieved on the basis of associations and ratios of constituents.

Nonmarine limestones: calcareous soils, calcareous dunes, and limestones of’ nonmarine successions deposited in large inland structural basins (calcareous lake deposits and,fanglomerates) Nonmarine limestones include: ( I ) those formed in soils, (2) limestones resulting from cementation of calcareous dune sands, and (3) limestones associated with nonmarine successions deposited in large inland structural basins, particularly calcium carbonate lake sediments and fan deposits consisting of recycled limestone particles. Limestonesformed in soils. Carbonate-enriched soil sequences are of minor volumetric importance in the stratigraphic record; but they are climatically controlled, and hence are particularly useful in stratigraphic study of Pleistocene deposits. Carbonate-rich soil concentrates are widespread today in semi-arid and “Mediterranean” climate belts. These include the “caliche” of southwestern United States (BROWN,1956) and “nari” of the Middle East, as mentioned in a previous section; calcrete and calcareous duricrust are more general terms for these deposits. The pisolitic limestone (Pliocene) of the Great Plains, U.S.A., which was


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previously thought to be an algal deposit that formed under water, has been shown to be the result of soil-forming processes (SWINEFORD et al., 1958). Also, the pisolitic limestone in the Permian of the northwestern margin of the Delaware Basin, southeastern New Mexico, formerly thought to be a diagnostic “lagoonal” or “back-reef” deposit, has been interpreted as the product of subaerial processes (R. J. DUNHAM, 1965). Calcareous dune deposits. Calcareous dunes may form wherever a source of loose calcium carbonate grains is available. Limestones formed by cementation of these calcareous grains comprise the “anemoclastic limestones” of GRABAU (1904; 1932) and the calcareous eolianites of SAYLES (1929, 1931). Typically, such dunes form along tropical coasts, where beaches consist of calcareous skeletal debris thrown up by the waves. They are discussed more fully in a later section dealing with sea-marginal deposits. Calcrete paleosol horizons are common in Pleistocene calcareous eolianites. Ancient examples include the Miliolitic Formation of the coast of the Arabian Sea (EVANS,1900) and the Junagarh Limestone of western India, which consists of shallow-water marine organic remains. The latter have been coated by calcium carbonate and then cemented by calcite. The Junagarh Limestone occurs 48 km inland from the sea; it overlies the Deccan traps and its thickness exceeds 60 m (CHAPMAN,1900; EVANS,1900). Calcium carbonate lake deposits. Calcium carbonate lake deposits may be of local or regional extent; they may include most of the varieties of first-cycle particle types found in shallow-water marine environments, such as ooids, superficial ooids, pseudooids, aggregate grains, skeletal debris, bioconstructional material, and so forth. Indeed, calcium carbonate deposits of large lakes resemble marine calcium carbonate deposits in many particulars. They differ chiefly in lacking the variety of distinctive marine organic skeletal remains and in their stratigraphic relationships with surrounding or interfingering nonmarine sediments, which characteristically include coarse-grained alluvial-fan deposits that may encircle the lake sediments on all sides, and, in appropriate climatic areas, nonmarine evaporites. Examples of both Recent calcium carbonate lake sediments and ancient limestones which owe their origin to lithification of carbonate lake sediments are discussed below. A . Recent calcium carbonate lake sediments. Recent calcium carbonate lake sediments are forming in the Dead Sea, Israel; Great Salt Lake, Utah, U.S.A.; Salton Sea, southern California, U.S.A.; and in numerous other lakes. ( I ) Recent sediments of the Dead Sea. The Dead Sea is one of the most hypersaline bodies of water in the world; its salinity ranges from 285 to 330%,. Carbonate minerals, including both aragonite and low-magnesian calcite, and sulfates are forming in it at present (BODENBEIMER and NEEV,1963; NEEV,1963, 1964; FRIEDMAN and NEEV,1966).

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The marginal zone of the Dead Sea is covered by a hard gypsum crust (Fig.5). Sediments and objects in this zone, whether gravel, trees, boats, or rope, for example, become encrusted with gypsum after a few months. Gypsum is precipitated continuously from the surface waters of the Dead Sea. This precipitation is greatest at the surface but extends down to depths of 1-2 m and locally to 3-6 m (NEEV,1964). Although gypsum crystallizes from the water, only calcite accumulates on the bottom of this basin. This calcite is of replacement origin; it forms as a result of the breakdown of caIcium sulfate by sulfate-reducing bacteria (for details, see section on Replacement origin of limestone). Aragonite, on the other hand, is a primary precipitate. Crystallization of aragonite takes place uniformly all over the Dead Sea, causing the water periodically to turn white. Whitings occur at irregular intervals of about 5 years, at times when the temperature of the Dead Sea water reaches its annual maximum; they were observed in August, 1943 (BLOCHet al., 1944; SHALEM, 1949), and in August, 1959 (NEEV,1963, p.153). The high temperature has been suggested as the triggering mechanism responsible for aragonite precipitation (NEEV,1963, 1964). Sediments from the bottom of the Dead Sea consist of rhythmically layered white and black laminae. The white layers contain abundant aragonite and are enriched in heavier carbon and oxygen isotopes and strontium, whereas the black layers lack abundant aragonite, are enriched in lighter isotopes of carbon and oxygen, and contain less strontium. The concomitant precipitation of aragonite and enrichment in heavier isotopes of carbon and oxygen is explained by the con-

Fig.5. West shore of Dead Sea showing 1-2 m zone of gypsum crust along shoreline (white). (Photograph by G.M. Friedman.)


20 1

centration that occurs during intense conditions of evaporation, in which the lighter isotopes are preferentially removed to the atmosphere as carbon dioxide. The evidence independently supports NEEV'S(1963, 1964) observations that high temperatures trigger aragonite precipitation (FRIEDMAN and NEEV,1966). Dolomite does not seem to occur in the Recent sediments of the Dead Sea, despite a report to the contrary (XNGERSON, 1962). Neev did not encounter it and Friednian spent considerable time searching for it unsuccessfully in his own samples and in those collected by Neev. The absence of dolomite in the Dead Sea may be explained by the low Mg/Ca ratio in the Dead Sea water. The Mg and Ca contents of pore waters in Dead Sea sediments have not been determined but the Mg/Ca ratio in many samples of bottom waters is only 2-2.5 (NEEV,1964). Much higher ratios have been measured in the pore waters in areas of modern marine dolomite formation: 3 to more than 10 in the Persian Gulf (ILLINGand WELLS,1964) and approximately 30 in Bonair, Netherlands Antilles (DEFFEYES et al., 1964). Magnesium enrichment in both these localities takes place on large supratidal flats, where evaporation and progressive concentration go hand in hand. Comparable marginal tidal flats are absent in the Dead Sea. (2) Great Salt Luke, Utah. The Great Salt Lake occupies a tectonic depression bounded on the east by the Wasatch Mountains and on the north, south, and west, by various smaller mountain ranges. To the west also lies the lowland of the Great Salt Lake Desert. The lake is nearly split into two subequal divisions by a line of north-south-trending topographic ridges that form the long peninsula of Promontory Mountains on the north and the Oquirrh Mountains on the south; this trend continues across the lake to form Freinont and Antelope Islands. The latitude of the center of the lake is 41 "10' N. The surface area of the lake at its present level is 3,225 km2 and the volume of water is 15.85 km3, giving an average depth of 4.91 m. (computed from figures presented by EARDLEY et al., 1957). Within historic times the level has fluctuated by 5.8 m around an altitude of 1,293 m, ranging from a high of 1,296.4 m in 1873 to 1290.6 m in 1934. The present salinity of the water is 23.0 %; the total amount of dissolved salts, mostly NaCl and N a ~ S 0 4 is , 4,760 . 106 tons. Another 559 * 106 tons are estimated to be distributed in the lake-bottom clays. The rivers bring in an estimated 1 .I . 106 tons of salts per year (EARDLEY et al., 1957). Recent sediments of Great Salt Lake consist of various terrigenous shoreline deposits in the eastern division, and of oolitic and faecal-pellet sands and algal biostroines in the shoreward parts of the western division. Calcareous clays occur in deeper water. The oolitic and pellet sands and algal biostromes represent nearly pure calcium carbonate materials, whereas the clays are mixed terrigenous material and carbonates (EARDLEY,1938). Oolitic and pseudoiilitic sands with well-developed ripples form a nearly continuous band around the western division of the lake from the shore out to depths of 3.60 m. These have been studied in detail from 60 samples collected along 5

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traverses perpendicular to the shoreline, which ranged in length from 200 to 1,760 m and ended in water depths ranging from 1 to 1.80 m (CAROZZI, 1957).Three of the profiles were made west of Antelope Island, one south of the Promontory Mountains peninsula, and one north from the north end of Stansbury Island. Neglecting transported ooids, which occurred on the Promontory peninsula profile out to a distance of 85 m from shore, the size of ooids was found to depend on the amount of agitation of the water and size of the nucleus grain, a substantiation of EARDLEY’s (1938) results (CAROZZI, 1957). Three types of oolitic sediments occur here: ( I ) those consisting entirely of ooids and superficial ooids (not subdivided); (2) those consisting of mixed ooids, superficial ooids, and pseudooids; and (3) those consisting entirely of pseudooids (which are not really oolitic at all). In each profile the sediments of type 3 (entirely pseudooids) form a belt farthest offshore; between this belt and the shoreline is a zone of mixed ooids, superficial ooids, and pseudooids. Depending on conditions of agitation, a belt of purely oolitic sands (including ooids and superficial ooids not subdivided) may exist in the zone of breaking waves. The nuclei of the ooids are composed of terrigenous mineral grains, chiefly quartz, and faecal pellets of the brine shrimp, Artemia gracilis, as noted by EARDLEY (1938). In light of ILLING’S (1954) work in the Bahamas, it would be of interest to know if all the pseudooids consist of faecal pellets or whether some aggregate-type sand-size grains are present. Colonies of the blue-green colonial alga, Aphanothece packardii, occupy approximately 256 km2 in the western division of the lake in a discontinuous belt that extends from the shore out to depths of 3.05-3.66 m. Algae are more abundant along the eastern shorelines of the western division of the lake than along the western shoreline of this division. The thickness of the algal layer is 15 cm; it is underlain by oolitic sand 30.5 cm thick, below which is a substratum of firmer oolitic-argillaceous sand. The shapes of the algal colonies mirror the microtopography of the substratum (CAROZZI,1962). In section, the algal deposit consists of alternating layers of porous white crusts, thought to have resulted from direct precipitation by the Algae, and layers of fragmental sediment composed of broken pieces of the crusts, ooids, faecal pellets, silt, and clay (CAROZZI,1962). An analysis of the algal material shows it to consist of aragonite (66.4 %); dolomite (1 1.8 %); insoluble material composed of silt and sand and clay with brine shrimp pellets (20%); and organic matter (2%) (EARDLEY, 1938). The carbonate content of the calcareous clays from the center of the lake averages 30%; of 200 samples, only 33 exceeded this figure, and only 10 showed a carbonate content of less than 15%. In the longest core collected between Stansbury and Antelope Islands NO.^), the following variations in carbonate content were found with depth in the core: 0-5.49 m, 42% calcium carbonate; 5.49-8.53 m, 22%; and 8.53-13.41 m (bottom of core), 11 % (J. F. Schreiber, 1957, in: EARDLEY and GVOSDETSKY, 1960). (3) The Salton Sea, southern California. The Salton Sea, southern California,

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is a lake which formed in 1905 during a flood and distributary diversion on the Colorado River delta. The latitude of the center of the lake is 33 "20' N. The lake is 56 km long and 16-24 km wide; it occupies an area of approximately 1,100 km2. The altitude of the lake surface is -71.14 m and its deepest closed contour is -85 m. The water, which was initially fresh, has grown progressively more saline and its salinity is now 33.68 %, nearly that of sea water (ARNAL, 1961). Temperature of surface water ranges beween the extremes of 12°C and 35.3 "C (the latter measured in August, 1954), and does not change much with depth. On July 10, 1954, the surface water temperature was 32"C, and the bottom water temperature at a depth of 13.2 m was 29°C. The lake sediments include mostly terrigenous types low in carbonate content, particularly around the lake margins; but where lime-secreting barnacles (Balanus amphitrite saltonensis) are abundant, the carbonate content increases to 40-60%. Carbonate content also increases offshore and is high in the area of clay deposition at the north end of the lake. Calcareous Foraminifera are the other main lime-secreting organisms in addition to the barnacles; their tests have been corroded slightly in places where the p H of the water drops below 7.2, and have been etched extensively where the p H reaches 7.0-6.8. The average p H of the surface water is 8.1-8.4 in the central parts of the lake, but reaches a maximum of 8.93 where plants are abundant in the shallow water. The minimum pH of the bottom waters is 7.34, and that of the sediments is generally 0.5-1.0 pH units lower than the p H of the bottom water. The carbonate budget can be calculated from measurements made on lake sediments and lake water, and by comparing these with the amount of carbonates supplied. The carbonates are derived chiefly from the Colorado River, both as dissolved and suspended sediment load, and secondarily from the drainage within the Salton Basin. An estimated 71.43 . lo6 tons of calcium carbonate are present in the lake sediments deposited in a period of 50 years (1905-1955). In 1955, 1.874 . 106 tons of CaC03 were still in solution in the water, making the total amount of calcium carbonate equal to 73.304 106 tons. This total compares with a total of 71.617 106 tons supplied, which is calculated from the following figures: 29.21 . 106 tons carried in solution in Colorado River water; 30.86 . lo6 tons delivered as suspended load (based on 10 % carbonate content in the suspended fraction); and 11.547 106 tons estimated as the supply from Salton Basin sources (ARNAL,1961). B. Ancient calcium carbonate lake sediments. Ancient limestones formed by the lithification of carbonate lake sediments include the Neogene-Pleistocene formations of the Jordan Valley-Dead Sea graben; Wasatch-Green River-Uinta Formations (Early Cenozoic) of the central Rocky Mountains, U.S.A. ; Horse Springs Formation (Early Cenozoic?) of southern Nevada, U.S.A.; Flagstaff Limestone (Paleocene) of south-central Utah, U.S.A.; Upper Jurassic deposits of Jura Mountains, Switzerland; Triassic deposits of the Connecticut Valley and 1

+

-

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J. E. SANDERS AND G.

M. FRIEDMAN

Virginia, U.S.A.; Mabou Group (Carboniferous), eastern Nova Scotia, Canada; and Coal Measures (Pennsylvanian and Permian) in Pennsylvania and West Virginia, U S A . Additional examples are given by SORBY(1853), CAYEUX (1935), and HATCHet al. (1938). (1) The Neogene-Pleistocene deposits of the Jordan Valley-Dead Sea graben. These represent a complex fill in the Middle Eastern rift valley. The rift valley consists of a northern Tiberias and a southern Dead Sea basin, which at times were connected, but at other times were isolated from each other. Calcium carbonate lake sediments constitute part of the graben filling; these lake sediments interfinger laterally with piedmont facies of redbeds and fanglomerates, fluvial and deltaic sediments, basalt flows, and tuffs (PICARD,1943; BENTORand VROMAN,1960; LANGOZKY, 1960; NEEV,1964). The calcium carbonate rocks derived from ancient lake sediments in the Tiberias basin include platy white calcilutite, pure oolite, calcarenite consisting entirely of pelecypod fragments, oolitic limestone containing abundant to dominant fresh-water gastropods, and calcilutite containing the same gastropods (PICARD, 1932, 1934; SCHULMAN, 1959). Near Beit She’an, the Pleistocene deposits consist of alternating fine-grained white and gray aragonitic, calcitic, and somewhat dolomitic layers and local interbeds of gypsum (PICARD,1943). This alternating sequence resembles the typical cyclic deposits of the Pleistocene of the Dead Sea basin as well as the Recent sediments of the Dead Sea; it, too, presumably reflects cyclic seasonal changes in climate. The graben fill in the Dead Sea basin includes interbedded and intertongued terrigenous sediments and limestones derived from calcium carbonate lake sediments. The age of the lake carbonates ranges from Oligocene through Pleistocene; they consist of deposits from three depositional cycles with a thickness in excess of 4,100 m. The basal member of each cycle is a lacustrine phase; the middle part, a brackish phase; and the upper part, a dry arid phase. The Pleistocene part of the section in the Dead Sea Basin is divided into two formations: the lower Samra Formation, which is composed of oolitic limestone of probable fresh-water origin, marls, calcareous sandstone and silt, gypsum, and conglomerates; and the overlying Lisan Formation, which consists of cyclic layers of finely laminated lightcolored aragonitic and dark-colored calcitic marls, terrigenous sediments, and gypsum. These cyclic deposits are thought to reflect seasonal changes, the lightcolored layers being summer deposits and the darker-colored ones, winter deposits (BENTORand VROMAN,1960, p.68). According to LANGOZKY (1960), the Samra sediments were deposited from a shallow, highly saline lake that was probably somewhat less saline than the present Dead Sea, but more saline than the lake in which the Lisan Beds were deposited. The latter lake was only highly brackish to saline. Fresh-water faunas are found locally in the Samra Beds. During Neogene and Pleistocene times, as at present, the Dead Sea Graben was a topographic lowland which collected the fresh-water drainage derived from


205

the surrounding mountains. Fresh-water carbonates formed in the central lake, whereas terrigenous sediments were dumped into the surrounding marginal areas. The aragonitic layers of the cyclic couplets probably originated during the summer during periods of whiting formation that occurred under intense evaporating conditions, as in the present Dead Sea. The darker-colored calcitic marls were deposited under conditions of less intense evaporation, and presumably also under the influence of sulfate-reducing bacteria, as described under the section on replacement origin of limestone. (2) Eocene Green River Formation of Wyoming, Utah, and Colorado. During Paleocene and Eocene times, a series of large lakes occupied one or more structural basins in Wyoming, Utah, and Colorado, U.S.A. A single large lake covered a wide area (varying in size) of the Uinta Basin, the area of which was 23,000 km2*. In this basin, Tertiary lake sediments and related fluvial and alluvial-fan sediments are 3,600 m thick. Fluvial flats and piedmont alluvial slopes, which surrounded the lake, formed a continuous marginal zone between the lake and the circumferential highlands. The size of the lake and the salinity of its water varied from time to time: when the size decreased, salinity increased, and the fluvial facies spread widely over the former lake bottom; and when the size increased, the salinity decreased, and the lake submerged former fluvial tracts, depositing lacustrine sediments over marginal zone facies. Various stratigraphic classifications have been applied to this complex basinfill assemblage. Initially, the terms Wasatch, Green River, and Uinta Formations were thought to represent distinct lithologic blankets of one type of sediment. A conglomeratic redbed-alluvial-fan unit, the Wasatch Formation of Paleocene age, was thought to be the oldest. It was supposedly everywhere succeeded by the medial lake sediments of the Green River Formation; and the upper fluvial beds, the Uinta Formation, were considered to lie entirely above the Green River Formation. Later, complex interfingering relationships were demonstrated between these three types of sediments, and the names tended to become facies designations that were applied to the sediments from the three major environments with less regard for details of age. A more recent trend is to use time-stratigraphic “formations” bounded by volcanic tuff beds and to recognize various “facies” within this framework. In effect, this trend marks a return to the earlier sequence of Wasatch-Green River-Uinta, but acknowledges that each of these units includes marginal alluvial-fan deposits, intermediate fluvial sediments, and basin-center lake sediments, not just the deposits of one of these environments (DANE,1954; HUNTet al., 1954). The oldest lacustrine carbonates occur as unfossiliferous limestones and “oil shales”, which are interbedded with thicker fluvial sediments in the upper part of the Paleocene Wasatch Formation (time-stratigraphic usage) in the Soldier Summit area, Utah (HUNTet al., 1954). The lake water at that time was relatively

* Area of 23,000 km2 refers to area of Uinta Basin; lake areas varied in size up to this maximum.

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fresh, probably shallow, and non-stratified; it supported a prolific benthonic fauna. The margins of the lake fluctuated rapidly over wide areas. The hydrocarbons deposited at this time were principally ozocerite, a type with long paraffin chains (HUNTet al., 1954). The carbonates are calcitic, with dolomite being absent. The Green River Formation, of either facies or time-stratigraphic usage, includes predominately lacustrine deposits; these have become a standard for comparison with sediments of other large lakes owing to the careful studies of BRADLEY (1926, 1929, 1948, 1964), and others. In the center of the Uinta Basin, the Green River Formation consists almost exclusively of lake sediments, which are 2,100 m thick; these pass laterally into coarser-grained basin-marginal sediments. The lower Green River Formation contains shell marls, algal reefs (as individuals or groups of reefs), algal-pebble beds, and ostracode-bearing and oolitic limestones. The Algae flourished in the lake and built reefs that expanded broadly over the smoothed lake floors. These reefs are remarkably similar to those found in the Miocene lake beds in the Rhine Valley, Germany; and the shore carbonate facies assemblage resembles that of the Recent sediments of the western division of Great Salt Lake, Utah, previously described. Fish, mollusks, crustaceans, and aquatic insect larvae were abundant in the lake water; turtles, crocodiles, birds, and small camels, as well as myriads of winged insects, frequented the lake shores (BRADLEY, 1929, p.204). Hydrocarbons in fissures in lake beds of the lower Green River Formation consist of albertite, with condensed ring structures (HUNTet al., 1954). The middle part of the Green River Formation includes “oil shales” (more accurately dolomitic marls; they are siliceous carbonates composed of calcite and dolomite, fine-grained quartz, clays, and organic matter, according to HUNTet al., 1954), large quantities of halite, trona ( N a ~ C 0 3* NaHC03 . 2Hz0), shortite ( N a ~ C 0 3* 2CaC03), and lesser amounts of other saline and unusual authigenic minerals, such as silicates, borosilicates, fluorides, phosphates, and complex carbonates. Sandstones are more abundant; those in contact with the oil shales are saturated with liquid gilsonite, which is the distinctive hydrocarbon of this interval. Evidently the lake contracted to a small size, the outlet ceased to exist, and the water became more saline and chemically stratified. Bottom water became extremely toxic so that no bottom-dwelling fauna existed and any of the nearsurface organisms that wandered into it were instantly killed. The upper part of the Green River Formation includes the Parachute Creek and Evacuation Creek members. The Parachute Creek member is the most distinctive, widespread, and continuous unit of the formation. It contains the richest and most extensive oil-shale deposits, including the “mahogany marker” a wellknown analcitized tuff bed of the eastern part of the Uinta Basin (DANE,1954). The thickness of this member ranges from 130 to 150 m. Other materials present are muddy marlstones, shale, low-grade oil shales, volcanic ash, varved dolomitic marlstone, sandstone, and limestone. Thin algal limestone beds and nodular masses


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occur locally near former lake margins. The sediments and hydrocarbons are uniform through many hundreds of meters of section, indicating persistence of uniform depositional environment within the lake. Pyrite is common; other minerals include calcite, dolomite, and nahcolite. The size of the Green River lake reached a maximum at this time, as the climate evidently had become humid again and the outlet was reactivated. The overlying Evacuation Creek member is 180-260 m thick (DANE,1954), and consists of more terrigenous sediments, such as bedded shale, mudstone, marlstone, siltstone, and sandstone; oil shale is rare. The intensity of volcanic activity increased during the later Green River time. The lacustrine deposits of the Uinta Formation were formerly considered to belong in the Green River Formation; but on the basis of their lateral relationships and position above prominent tuff beds, they have been reassigned to the Uinta Formation (DANE,1954). These distinctive lake beds, called the “saline facies” (DANE,1954), pass laterally into typical fluvial beds of the Uinta Formation east of the Green River. The carbonate content of the saline facies is higher than that in the underlying beds. In addition, there are calcium, magnesium, and sodium carbonates; the magnesium silicate, sepiolite; numerous crystal cavities; and abundant silica. Hydrocarbons change from aromatic types to the more naphthenic type, wurtzilite, which is high in sulfur and nitrogen contents (HuNr et al., 1954). Evidently the lake shrank again and became progressively more stagnant and saline; it eventually disappeared altogether. (3) Horse Spring Formation of southern Nevada. The Horse Spring Formation (Early Cenozoic?) of southern Nevada is a complex of lake deposits 320-830 m thick that includes limestone, dolostone, magnesite, clay, silt, sandstone, and volcanic ash (LONGWELL, 1928). The light-colored limestone is a conspicuous part of this formation. Near Logan Wash it is 155 m thick and consists of beds that are 60 cm to 18 m thick. Thicker porous varieties, which are medium-grained, alternate with thin layers having banded structure; bedding surfaces are gnarly and hummocky. The limestone is almost pure calcium carbonate; and analysis showed only a trace of magnesium. In White Basin, the limestone is 185 m thick. It is underlain by 90 m of gypsum-bearing beds, and overlain by 185 m of rocks that include white sandstone, thin-bedded magnesite, volcanic tuff, and calcareous clay containing colemanite. ( 4 ) Flagstaf Limestone of south-central Utah. The Flagstaff Limestone (Paleocene) forms the cap rock of the Wasatch Plateau, south-central Utah, U.S.A. At its type locality, Flagstaff Peak, the preserved thickness is 60 m (SPIEKER and REESIDE,1925). The formation is 150 m thick at Ferron Mountain and Sage Flat nearby, and reaches a maximum of 450 m in other parts of the Wasatch Plateau. Its maximum thickness in the Gunnison Plateau to the west is 285 m (LAROCQUE, 1960). The Flagstaff lies nearly horizontal over very large areas of the Wasatch Plateau; it has been extensively dissected and is well exposed in steep cliffs at the crest of the plateau. Along the west side of the plateau, however, it

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dips to the west in the Wasatch monocline, and the overlying Green River Formation is exposed along the western margin of Sanpete Valley (SPIEKER,1946). In the Wasatch Plateau, two units are recognized: a lower unit composed of blue-gray shale and some thin limestones, and an upper unit consisting of massive white limestone, minor shale and channel sandstones, abundant gypsum, and local chert (LAROCQUE, 1960). In the Gunnison Plateau, as many as five units have been recognized (GILLILAND,1951). The minimum area of the lake in which the Flagstaff Limestone was deposited amounts to at least 7,158 km2, and the maximum was probably on the order of 10,240 km2. Three phases in the history of the lake have been recognized; the maximum expansion took place during phase 2. In phase 3, the lake was shallow, and it remained shallow until its final extinction (LAROCQUE, 1960). ( 5 ) Upper Jurassic deposits qf Jura Mountains, Switzerland. The Purbeckian strata (uppermost Jurassic) of the Jura Mountains, Switzerland, include lacustrine limestones that typically are fine-grained, compact, and contain abundant remains of ostracodes and/or charophytes (CAROZZI,1948). A characteristic cyclic sequence is produced by alternations of ostracode-bearing limestones and limestones with charophytes. Terrigenous quartz grains are present in most lacustrine limestone beds. The thickness of the entire Purbeckian rarely exceeds 20 m and in many localities is only 5-10 m. The basal part consists of gypsiferous marls on the northwest and dolostones on the southeast, indicating that the marine shoreline lay to the northwest at that time (see FRIEDMAN and SANDERS, 1967). The middle part consists of two lacustrine limestone tongues which are separated by a thin (1-2 m) unit of marine deposits. An upper unit of marine limestones generally overlies the upper lacustrine limestones, except where the upper marine limestones have been removed by pre-Cretaceous erosion. These Jurassic lacustrine limestones are not deposits from lakes of large inland structural basins, as are the examples previously described, but must have been deposited in shallow lakes that occupied a low-lying coastal plain. Their topographic setting, therefore, more closely resembles that of the lacustrine limestones interbedded with the Coal Measures than that of limestones deposited in large lakes of inland structural basins. (6) Triassic of the Connecticut Valley and Virginia. The thick Triassic strata of central and southern Connecticut, which are predominantly sandstones and conglomerates with three interbedded basaltic lava-flow complexes, include thin lacustrine limestones that are only locally exposed in four places. The extent and stratigraphic relationships of these limestones to each other are not definitely known. They occur in the middle of the succession; those at three localities are interbedded within the lower lava-flow complex, whereas that of the fourth locality overlies the middle lava flow. At two localities, the limestone is only 30-60 cm thick; it is 4 m thick at another locality and its greatest thickness is approximately 8 m. The carbonate rocks contain 7-17 % terrigenous insoluble residue consisting of quartz and feldspar grains, and locally include algal bioherms (KRYNINE, 1950).


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The rock at one locality includes a few layers that preserve small-scale wave ripple marks (5. E. Sanders, personal observations); here also are found many small cavities, that are either empty or filled with white or yellowish calcite crystals or reddish iron-oxide masses (KRYNINE,1950, p.107). No oolite has been reported from the lacustrine limestones of the Triassic in Connecticut, but oolite does occur in the Triassic of Virginia (YOUNGand EDMUNDSON, 1954). (7) Mabou Group (Carboniferous), eastern Nova Scotia, Canada. The Canso and Emery Brook Formations, the fine-grained gray facies of the Mabou Group (Middle Carboniferous), of Cape Breton Island, eastern Nova Scotia, Canada, contain limestones formed by lithification of calcium carbonate lake sediments (BELT,1962, 1964). The maximum thickness of the Canso Formation is 430-770 m. Carbonate rocks represent only a minor proportion of the unit. They consist of (a) thin beds of oolitic limestone, found in the Broad Cove, Ainslie Point, and Cape Douphin sections; (b) coquina, in which the skeletal debris consists of the freshwater mussel, Carbonicola, at the Bay St. Lawrence section; and (c) tan-weathering dolostone with local algal structures, exposed at many sections. The Emery Brook Formation includes only thin beds of calcilutite. According to BELT(1962), the algal colonies and crusts indicate that the maximum depth of water in the Canso lake was 10 m, whereas the encrusting algal colonies indicate a depth of about 15 m. These Middle Carboniferous lake deposits pass laterally into fine-grained terrigenous fluvial facies and these, in turn, into coarse-grained basin-margin fanglomerates. The pattern is identical to that of the Cenozoic deposits of the Jordan Valley-Dead Sea Graben; Green River Formation of the Rocky Mountains U.S.A.; and Triassic deposits of eastern United States. (8) Coal Measures (Pennsylvanian and Permian) in Pennsylvania and West Virginia. Lacustrine limestones appear in the Pennsylvanian cyclothems of northcentral United States. Marine limestones occur with them in Illinois and only a few localities in Pennsylvania; marine limestones are absent in most Pennsylvania localities. Lacustrine limestones (“underclay limestone”) occur in member 3 of the original ten-member cyclothem (WELLER,1930), whereas the marine limestone is member 7. The lacustrine limestones occur as scattered nodules in the clay or as separate beds; they are generally unfossiliferous in Illinois, except for Algae (WANLESS et al., 1963). Lacustrine limestones occur in the two uppermost divisions of the Allegheny Series (Pennsylvanian) of western Pennsylvania, the Kittanning (below) and Freeport (above). They are most widespread in the upper division of the Kittanning (K4 of FERM and WILLIAMS, 1960) and in the middle division of the Freeport (F2), where they occupy the position of the marine limestone found in the analogous Kittanning unit (K2). The most common rock type is nodular calcilutite; insoluble residues vary from 5 to 50%, and average 30%. Fossils include fresh-water

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gastropods, ostracodes, rare pelecypods, and spores. Measurement of trace elements shows 30-90 p.p.m. of boron, with most values falling between 75 and 80 p.p.m. in the lacustrine limestones compared with 133-1,200 p.p.m. (with most values being 200-400 p.p.m.) for marine limestones in the same sequence. Gallium content, on the other hand, is high in the lacustrine limestones (25-110 p.p.m.) and lower in the marine limestones (3-40 p.p.m.1 (DEGENS et al., 1958). Lacustrine limestones are also abundant in the northeastern half of the outcrop belt of the Dunkard Group (Permian) of southwestern Pennsylvania and northern West Virginia, where they are associated with coals. The Lower Dunkard cyclothems include eleven members (CROSS,1950; CROSSand SCHEMEL, 1956; BEERBOWER, 1961). The limestones occur both below the coal and interbedded with it (member 2), and as lateral equivalents of claystones higher in the cyclothems (members 9 and 10). They are extremely variable and are of local extent (BEERBOWER,1961). The lakes in which the limestones associated with coal measures were deposited occupied coastal lowlands rather than intermontane structural basins. Fan or stream deposits consisting of recycled limestone particles. Fan or stream deposits consisting of recycled limestone particles occur where limestone bedrock is being actively eroded in the absence of a continuous soil cover. Most of these deposits accumulate on the downthrown side of an intermittently active fault, where the upthrown side contains limestones. Limestone clasts do not survive extensive abrasion; typically they are limited to areas near their source and only rarely survive more than one or two cycles of erosion and deposition. A single example from Recent sediments and numerous examples from the geologic record are included in the ensuing discussion. A. Recent stream sediments consisting of recycled limestone particles. Modern streams draining the Arbuckle Mountains, Oklahoma, where limestone bedrock is abundant, are small and generally lack much sediment load except during floods; but their flood deposits are similar to Pennsylvanian limestone-pebble conglomerates (JACOBSEN, 1959). During floods these small streams become powerful torrents which transport sand, pebbles, and cobbles together. The floods dissipate rapidly, leaving behind a poorly sorted deposit that is texturally similar to the ancient conglomerates nearby. Although high relief in the source area is almost axiomatically assumed to explain accumulations of coarse-grained recycled limestone fragments, the Oklahoma flood deposits are formingin an area where the relief is less than 60 m. B. Ancient calcareous alluvial f a n deposits. Numerous examples of calcareous alluvial-fan deposits exist in the stratigraphic record. Examples discussed here include the Miocene deposits of the Jordan Valley-Dead Sea Graben, Israel; Overton Fanglomerate (Tertiary?) of southern Nevada, U.S.A. ; Upper Cretaceous of Utah, U.S.A.; Triassic of South Devon, England; Triassic of New York, New Jersey, Pennsylvania, and Maryland, U.S.A.; Pennsylvanian of the Arbuckle

ORIGIN AND OCCURRENCE OF LIMFSTONES

21 1

Mountains and Ardmore Basin, Oklahoma, U.S.A. ; and Bonaventure Formation (Mississippian) of eastern GaspC, Canada. (I) The Miocene deposits of the Jordan Valley-Dead Sea Graben, Israel. These include fanglomerates with limestone pebbles and boulders. Along the fault scarps near the margin of the rift valley of northern Israel the limestone fragments were derived from local Cretaceous rocks; boulders of local basalt flows are also common here. (2) The Overton Fanglomerate (Tertiary?) of southern Nevada. This includes large limestone boulders and a matrix of sand grains that have been cemented by calcium carbonate (LONGWELL, 1928; 1949; 1951). The thickness of this formation ranges from 6-920 m. The upper bed is a single thick layer called a “limestone fanglomerate” (LONGWELL, 1928), which is 65 m thick in Overton Wash and 145 m in Wieber Wash; the strata below include interbedded sandstones and limestone conglomerates. A pebble count near Overton indicated 70-90 % limestone clasts, 5-20 % sandstones, and 0-10 % chert; particle size varies. Some layers contain clasts with an average size of less than 7 cm, and a maximum size of 25 cm which is exceptional. In Kaolin Wash, 12-cm clasts are abundant and a maximum size is 75 cm. Tn Logan Wash fragments range from 10 to 60 cm, and sizes up to 91 cm are common. One layer here, which is only one boulder diameter thick, can be traced for 0.8 km parallel to both the strike and the dip; in it the boulders are packed closely together (LONGWELL, 1928). Near Muddy Creek, limestone blocks with a diameter of 6 m are not uncommon. About 3.2 km north of Muddy Creek a mass of Kaibab Limestone (Permian) occurs that is 0.5 km long; 6.4 km to the south of this is a slab of Paleozoic limestone that measures at least 550 m long and is about 90 m thick (LONGWELL, 1949). Imbrication of flat pebbles indicates transport from the west and southwest near Logan Wash. Clasts of younger formations, such as the Jurassic sandstone and Moenkopi Formation (Lower Triassic), are abundant in the lower parts of the Overton Fanglomerate; clasts of older limestones, such as the Kaibab and the Mississippian and Devonian limestones, predominate in the upper part. The clasts were all derived from the advancing Glendale thrust sheet. The thrust overrode the piedmont slope on which the fanglomerate accumulated and finally, after cessation of movement on the thrust, more of the same debris overlapped and buried the frontal part of the thrust sheet. Accordingly, the Overton Fanglomerate is a true example of a “thrust conglomerate”. It typifies the type of deposit which has been proposed as an explanation for many of the marine limestone pebbly mudstones that occur in geosynclinal tracts; these are discussed in a later section. ( 3 ) The Price River Formation (Upper Cretaceous-Mesaverde age) of central Utah. This formation lying adjacent to the southern Wasatch Mountains, includes limestone pebble and boulder conglomerates, the clasts of which were derived from the Paleozoic sandstones and limestones exposed in the mountains to the west. The Price River Formation represents a piedmont facies that passes eastward

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progressively into inland floodplain, lagoonal, littoral marine, and finally, offshore marine deposits (YOUNG,1955). Petrographic study by Friedman of samples of the conglomerate indicates that the fragments consist of limestone, 65 %; quartz grains, 20 %; chert, 10 %; and metaquartzite, 5 %. The content of limestone particles diminishes rapidly to the east and chert content increases proportionately. Carbonate grains are sporadic to absent in the deposits of the inland floodplain environment. (4) The Triassic of South Devon, England. This includes local limestone scree deposits in which the limestone clasts have been derived from the Plymouth Limestone (Middle Devonian). Good exposures are present in the coastal cliffs at Teignmouth and Petit Tor, northeast of Torquay. The fanglomerates consist of lightcolored clasts of carbonate rocks set in a fine-grained matrix of red mud (shown to J. E. Sanders by D. Laming, on a field trip in July, 1954). ( 5 ) The Triassic conglomerates along the northwestern side of the Newark Basin outcrop belt, New York, New Jersey, Pennsylvania, and Maryland. These conglomerates include accumulations of limestone pebbles and boulders (KUMMEL, 1897, 1898; CARLSTON, 1946). Recycled limestone and dolostone clasts derived from nearby Lower Paleozoic (Cambrian and Ordovician) carbonates comprise a variable proportion of the marginal fanglomerates. The latter were deposited on alluvial fans by short, steep streams that drained southeastward from the intermittently uplifted marginal fault block into the subsiding Newark Basin. Recycled carbonate rocks may exceed 95 % of the clasts in some deposits, whereas they may be absent in nearby deposits. According to the descriptions and map of CARLSTON (1946), the clasts of carbonate rocks are less well rounded than the associated clasts of quartzites and sandstones. Coarse-grained deposits, which include carbonate rock clasts, extend shorter distances into the basin from its margin than do those with quartzite clasts. Carbonate rock conglomerates and breccias, all of them of the “Potomac marble” type, are found along the northwest margin of the Newark Basin from the Hudson River, on the northeast, to the Potomac River, on the southwest. The northeasternmost exposure is just north of Stony Point, New York. A breccia containing almost entirely angular cobbles and boulders of carbonate rocks up to 1.25 m in diameter crops out between Ladentown and Wesley Chapel, Rockland County, New York, 6.4-12.8 km southwest of Stony Point (CARLSTON, 1946, p.1015, p1.2, fig.1). By contrast, pebbles of carbonate rocks are almost unknown in the Connecticut Valley Triassic outcrop belt. Most of the pebbles in this belt were derived from the east, from a terrane where carbonate rocks are rare. An exception occurs in a conglomerate bed near the top of the Shuttle Meadow Formation that is exposed in Saltonstall Ridge, East Haven, Connecticut, where Route 1 cuts through the ridge. Almost all of the pebbles in this bed are noncalcareous: quartzite, granite, and large feldspar crystals constitute 90 % of the pebbles. Carbonate rocks of unknown pro-


21 3

venance comprise about 5 % of the total (J. E. Sanders, personal observations). (6) The Collings Ranch Conglomerate (Pennsylvanian) of the A rbuckle Mountains, Oklahoma. This conglomerate is a limestone-boulder accumulation that is at least 800 m and possibly as much as 925 m thick. It is the thickest and coarsest single deposit of terrigenous debris derived from the rising Arbuckle Mountains (HAM,1954). The particles include boulders and pebbles of fine-grained dolomitic limestone, pellet limestone, algal or stromatolitic limestone, and oolitic limestone from the Arbuckle Group (Cambrian and Ordovician). Most contacts between the Collings Ranch Conglomerate and older bedrock formations are faults; but in one locality it could be established that most, if not all, of the Collings Ranch Conglomerate had been deposited before faulting ceased (HAM, 1954). The Pennsylvanian conglomerates of the Ardmore Basin, Oklahoma, consist of pebbles composed mainly of limestone and chert (99%). The limestone pebbles are confined to a basin-marginal zone that is 4.8-6.4 km wide. The proportion of limestone pebbles diminishes rapidly away from this marginal zone, so that toward the center of the basin only chert fragments persist. Wherever the basin margin is known to be faulted, conglomerates are present at one or more levels; probably all of the Pennsylvanian conglomerates of the Ardmore Basin originated along basin-margin fault scarps (JACOBSEN, 1959). (7) The Bonaventure Formation (Mississippian) of eastern GaspC, Canada. This formation consists predominantly of coarse-grained red sandstone (McGERRIGLE, 1950). On the northwest side of Bonaventure Island, opposite PercC, however, numerous layers of limestone-pebble conglomerate appear. The clasts were derived from the fine-grained Upper Ordovician limestone that occurs on the mainland nearby (SCHUCHERT and COOPER, 1930). The limestone clasts generally are well-rounded and range in size from 3 to 15 cm. Milky quartz pebbles comprise about 2 % of the total and weathered igneous rocks occur in trace quantities (J. E. Sanders, personal observations). The modern beach gravel adjacent to these conglomerates consists of limestone pebbles that have fallen away from the bedrock and are being further rounded by wave action. Shallo w-water marine andlor marginal marine associations Shallow-water marine and marginal marine environments of deposition include a vast array of sub-environments that all fit into the general pattern of an open sea with relatively quiet water on one side, a zone of turbulent (“high energy”) water in the middle where waves break on a reef or barrier beach, and a sheltered lagoon of variable width and depth between the reef or barrier beach and the mainland. In some localities the lagoon is absent, either because it has become converted into a mangrove swamp or tidal marsh or because the beach lies in contact with the mainland itself. In the sites of major accumulations of carbonate sediments there may be no mainland, but only a wide shallow bank or deeper lagoon that grades into a turbulent marginal zone and beyond that into deep water on all sides. The

214


Campeche Bank, on the other hand, is connected to the mainland of Mexico. The sediments deposited in these environmental complexes around the edge of the sea may consist exclusively of first-cycle carbonate materials, or they may consist entirely of terrigenous sediments that may or may not include second-cycle carbonate detritus worn from the land, or they may include mixtures of first-cycle carbonates and terrigenous materials. In addition, where climate is hot and arid and sea-water circulation restricted, evaporite minerals may be precipitated from the sea water. From these numerous possibilities it is readily apparent not only that nearshore limestones theinselves may be extremely varied but also that they can be associated with a wide variety of other types of materials. A common characteristic of all nearshore marine deposits, whether of carbonates, terrigenous sediments, or mixed carbonates and terrigenous sediments, is their rhythmic or cyclic sequential arrangement. Beds of different kinds of sediments occur in a fixed vertical order that is controlled by the lateral distribution of the environments of deposition of the different sediment types and the shifting of the boundaries of these environments in response to subsidence and changing position of land and sea (deleveling). Stratigraphic units thus accumulated are here considered to provide the best examples of lateral sedimentation. Where the sediments deposited initially in zones farther offshore overlie (or onlap) those from more landward zones, relative submergence must have occurred and the shoreline advanced landward in a marine transgression. Where the sediments of the landward zones overlie (or offlap) those from a more seaward position, either an emergence or progradation has occurred. In an oWap of nearshore sediments over offshore sediments, the position of the shoreline may or may not have changed; if it shifted seaward, then a regression took place. The shoreline shifted seaward in most offlaps, but may not necessarily have done so if the changes were restricted to sediment belts that lie further offshore.

Combinations of niarginal niarine deposits including only first-cycle carbonate minerals. Combinations of marginal marine deposits that include only first-cycle carbonate minerals include: ( A ) bank, bank-marginal beaches or shoals, and offshore marine deposits; ( B ) bank, bank-marginal reefs, and offshore marine deposits; ( C )lagoon, fringing reef, and offshore marine deposits. A category (D), ancient biohermal suites in general, is added to encompass the ambiguities found in the geologic record. In all of these combinations, organic skeletal remains constitute an important proportion of the total sediment; the remainder may consist of ooids, pellets, and various other accretion-type particles. The environmental zonation of organisms in the water and on the bottom is a major factor in the deposition of different types of sediments in belts that typically extend parallel to the shoreline. Pure carbonate materials may also accumulate on a low-lying continental landmass where the climate is hot and arid and/or surface rocks are older carbon-


215

ates. The presence of wind-blown calcareous sediment would be diagnostic of a continental area. Nearshore first-cycle carbonate deposits resemble those from marine banks; they may be distinguished, if at all, only by regional paleogeographic analysis. A . Combinations of bank, bank-marginal beaches or shoals, and ofshore marine deposits. Isolated shallow-water marine banks, such as the Great Bahama Bank, or the semi-isolated Campeche Bank, constitute the most important shallowwater setting for deposition of pure lime sediments and limestones. The typical bank sediments include ooids, algal bioherms and biostromes, and sand-size grains that have formed by accretion of aragonite needles as described from the Bahamas by ILLING (1954). Coated grains of various kinds are also typical. The main distinction between an isolated bank and a semi-isolated bank lies in the presence or absence of associated terrigenous sediments. A completely isolated bank lacks terrigenous admixture; whereas a semi-isolated bank may or may not show admixed terrigenous sediments, depending upon the drainage pattern. Although algal bioherms and biostromes are not restricted to bank environments, they are extremely typical of them. The importance of Algae as rock-builders has been realized by only a few students of limestones (e.g., PIA, 1926; MAWSON,1929; FENTON and FENTON,1931, 1933; HOWE, 1932; WOOD, 1941; J. H. JOHNSON,1943a, b, 1946, 1952, 1954, 1961, 1964; ANDERSON, 1950; JOHNSON and KONISHI,1956; ODERand BUMGARNER, 1961; WOLF, 1962, 1965a, b; and FREEMAN, 1964). The algal limestones are so widespread in time and space, however, that they deserve wider appreciation. As noted previously, where subsidence occurs after a marine bank has been established, the enormous carbonate productivity of warm shallow sea water is capable of providing enough sediment to keep the water shallow, and hence to maintain the conditions for further development of lime sediments. The result is that a thick mass of pure lime sediments may be deposited in the absence of any source of terrigenous materials. The Recent carbonate sediments of the Bahamas have been examined by many geologists since they were first noticed by SORBY(1879); for details, see TAFT(1967a). Viewed broadly for the purposes of comparison with more ancient limestones, the sediments of the Bahamas may be divided into (1) bank deposits consisting of muds and sandy muds that have been much influenced by burrowing and sediment-binding organisms such as grasses and Algae; (2) bank-marginal deposits that include beaches and dunes composed of broken skeletal debris, or shoals composed of ooids and various accretionary grains; and (3) offshore marine deposits that consist of skeletal remains of benthonic organisms to which have been added tests of pelagic Foraminifera in increasing proportion farther from shore. Limestones from the stratigraphic record that may have been deposited on marine banks include the “White Limestone” (Eocene-Miocene) of Jamaica;

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Great oolite series (Middle Jurassic) of southern England (KLEIN,1965); Arab-D Formation (Upper Jurassic) of Saudi Arabia (R. W. POWERS,1962); “Mountain Limestone” (Lower Carboniferous) of southern England and Wales; Newman Limestone (Upper Mississippian) of northeast Tennessee, U.S.A.; Rundle Group (Mississippian), and Palliser Limestone (Devonian) of southwest Alberta, Canada (BEALES,1956, 1958); Silurian of Gotland, Sweden (HADDING,1959a, b); Black River Limestone (Middle Ordovician), southern Ontario, Canada (BEALES, 1958); and Ellenburger Group (Lower Ordovician) of central Texas (CLOUDand BARNES, 1948, 1957). During mid-Eocene to Early Miocene time, the south-central third of Jamaica formed a single tectonic unit called the Clarendon Block, which was gradually submerged by mid-Eocene time to form a shallow-water marine bank, the Clarendon Bank. Nearly pure carbonates (limestones and dolostones) were deposited both on the shallow bank and in the deeper water that surrounded it on the east, north, and west. The collective name for these rocks, which cover approximately two thirds of Jamaica, is the “WhiteLimestone”.This unit has been subdivided into various local facies (VERSEY, 1962). The bank limestones consist of calcareous sands derived from reefs (Swanswick Limestone); recrystallized limestones (Troy Limestone); and various dolostones and magnesian limestones. The Newport Limestone is a compact, moderately well-bedded, poorly fossiliferous limestone that contains a few oysters and pectens. Its lower part is 180 m thick in the Williamsfield trough, and a typical upper part is 610 m thick in the Santa Cruz Mountains. The recrystallized limestones (Troy Limestone) occupy an intermediate position between the deeper-water foraminifera1 marginal limestone facies and the dolostones and magnesian limestones of the central part of the bank. The Troy Limestone itself has also been extensively, but patchily, dolomitized. As the Clarendon Bank became more deeply submerged during the Oligocene time, the facies which had formed marginal zones in the Eocene encroached more and more upon the center of the block (VERSEY, 1962). Low-lying continental land masses in a hot, arid climate, such as Western Australia, may become sites of pure calcium carbonate eolian accumulations, owing to their low relief and lack of surface fresh-water drainage. The Coastal Limestone of Western Australia represents a belt of calcareous dune rock 2,400 km long and up to 10 km wide that is locally as much as 10 m thick (on the islands of the Abrolhos group); the dune rock overlies shell debris from beaches (TEICHERT, 1946, 1947a, b; FAIRBRIDGE and TEICHERT, 1953). The geometry of such a limestone may closely resemble that of deposits of isolated marine banks, but the steep foreset beds of typical eolian cross-beds and presence of land snails and other terrestrial animals characterize dune deposits. (See later section for further discussion of Pleistocene eolianite limestones.) B. Combinations of bank, bank-marginal reefs, and offshore marine deposits. The combinations discussed in this section resemble in a general way those dis-


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cussed in the preceding section, but differ in that they include marginal reefs instead of beaches or shoals. The general pattern of sediments remains the same, but a greater variety of skeletal debris will occur in and near the reefs. In addition, in situ organic skeletal frameworks and a variety of initially stony deposits formed by encrusting lime-secreting organisms will be present. Reefs act as sediment source for material shed both landward and seaward, and also as sediment traps for terrigenous material coming from any nearby land; accordingly, they tend to produce nearly pure calcium carbonate deposits even if a source of terrigenous sediment lies close by. The absence of associated terrigenous sediment from the outer parts of a reef tract, therefore, does not necessarily mean the absence of a nearby landmass. The presence of terrigenous sediment, on the other hand, is definitive. The reefs on the north coast of Jamaica (GOREAU, 1959; ZANS,1962), south coast of Cuba (DAETWYLER and KIDWELL, 1959), and east coast of British Honduras (PURDYand MATTHEWS, 1965) are examples of reefs that lie near land masses, but whose seaward sides locally contain only pure carbonate sediments. One of the most extensively developed barrier reefs in the Atlantic Ocean lines the eastern margin of the Great Bahama Bank (NEWELL and RIGBY,1957); it separates the deep water of the Tongue of the Ocean, a branch of the Atlantic Ocean, from the shallow Windward Lagoon on the bankward side of the reef (Fig.9). The barrier is 0.4-0.8 km wide and is discontinuous along its length; it contains gaps opposite the larger creeks and tidal channels that flow east from Andros Island, the largest island on the Great Bahama Bank. In these places, minor differences in salinity, turbidity, and nutrients evidently are responsible for preventing coral growth. Submerged shelves and terraces of Pleistocene bedrock form the foundation of the reef. The upper part of the barrier reef contains branching corals that project upward and oceanward, such as the wave-resistant elkhorn coral (Acroporapalmata) or moosehorn coral ( A . cervicornis), and also other large, massive boulder-shaped corals and coralline Algae. The elkhorn coral is the main frame-builder; it grows to within 1 m or less of the water surface. Seaward, the elkhorn coral is replaced by the staghorn coral and finger coral (Porites). The deeper and outer parts of the reef are dominated by massive heads of star coral (Montastrea), starlet coral (Siderastrea), and brain coral (Diploria); the lower part of the reef consists for the most part of massive corals, interstitial coralline Algae, and reef detritus. The Windward Lagoon behind the reef is about 2.4 km wide. Much of the lagoon floor is swept free of sediment, but is inhabited by attached alcyonarians (sea whips and sea fans), sponges, and green Algae; numerous patch reefs exist along the outer margin of the lagoon close to the barrier reef. The current-swept areas of the lagoon floor constitute a living disconformity or diastem. In relatively protected areas of the lagoon, however, skeletal sand accumulates; here grass and green Algae are common and the fauna consists mostly of burrowifig or mobile organisms.

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Reefs are absent along the western margin of the Great Bahama Bank on the side that faces the Florida Straits. This asymmetrical distribution of reefs (reefs present on the eastern margin, but absent along the western margin) is explained by the predominant wind direction from the east, which is responsible for bringing nutrients to the reef. Hermatypic corals are not present on the shallow interior of the Bahama Bank and their presence along the Bank margin depends on the local topography, where such features as the Windward Lagoon protect them from ebbing bank water (NEWELL et al., 1959, p.212). The south Florida reef tract does not represent a typical marine bank margin in the same sense as the Bahama Banks, in that it does not occupy an isolated position with open sea all around; instead, it is bounded on the north by Florida Bay with low-lying Florida peninsula farther to the north. The carbonate sediments in Florida Bay and the reef tract, however, are inferred to be similar to those that characterize a bank. In addition, the detailed studies of these sediments by GINSBURG (1956) provide a convenient basis for comparison with sediments from the Bahamas and in the geologic record. In Florida Bay, the sediments contain more fine-grained (smaller than 0.125 mm) than coarser-grained particles and the coarser grains consist almost exclusively of skeletal remains of mollusks and Foraminifera; in the reef tract, fines are less abundant and the coarser fractions consist of abundant Algae and corals. Three sub-environments and intervening transitional zones were recognized in the reef belt: back-reef, outer reef-arc, and fore-reef (proceeding from Florida Bay seaward). Halimeda occurs in the back-reef and outer reef-arc; it is rare to absent in the fore-reef zone at depths greater than about 60 m. Coralline Algae are most abundant in the outer reef-arc and are scarce in the back-reef zone, except near shore. They may increase in abundance with depth in the fore-reef zone. Coral debris is most abundant in the outer reefarc and in adjacent transition zones; it decreases more rapidly in fore-reef than in back-reef areas. Foraminifera increase seaward of the lower transition zone, and mollusks are more abundant in the back-reef and fore-reef areas than in the outer reef-arc. An important conclusion reached by GINSBURG (1956, p.2419) is that particles larger than 0.125 mm generally accumulate in the sub-environment where they are produced and do not move far away; hence, the reef tract is not a major source of sand-size skeletal particles for adjacent areas. C. Combinations of lagoon, pinging reef, and ofshore marine deposits. Combinations of lagoon, fringing reef, and offshore marine deposits contrast with the preceding group in that the central body of water enclosed by the reefs is not a relatively flat shoal area where evaporation may cause increased salinity, but a deeper lagoon, where normal salinity typically prevails and the organisms may be zoned bathymetrically. The reefs themselves and their flanking deposits on the seaward side, however, may be alike in the two cases. The best-known Recent examples come from the Indo-Pacific region; their general distribution has been summarized by CLOUD(1958). General discussions


219

of reefs and associated sediments include those of MURRAYand IRVINE (1891), MARSHALL (1931), YONGE(1951), CLOUD(1952), MACNEIL(1954), and GEORGE (1956). Detailed descriptions of carbonate sediments include those from Bikini atoll, Marshall Islands (EMERY,1948; EMERY et al., 1954); Kapingamarangi atoll, Caroline Islands (McKEE, 1958); and Guam (FORMAN and SCHLANGER, 1957, SCHLANGER, 1964). Examples of ancient suites of limestones deposited by and near reefs, characteristics of which resemble those of Recent carbonate sediments from the Pacific, have been described from the Middle East (HENSON, 1950), from the subsurface Tertiary of Louisiana (FORMAN and SCHLANGER, 1957), and from many Pacific islands. Other examples of older “reefs” are mentioned on a subsequent page. The Guam reefs provide typical examples; on the basis of grain size and skeletal organic remains seven different sediment types have been recognized in and near these reefs (FORMAN and SCHLANGER, 1957; SCHLANGER, 1964): back-reef shoal, reef-wall, reef-breccia, off-reef, fore-reef detrital, fore-reef transitional, and basinal types (slightly modified from HENSON,1950). The back-reef shoal areas are characterized by mud bottom and abundant miliolid Foraminifera; these areas were not further subdivided by FORMAN and SCHLANGER (1957) but have been by others (e.g., EMERY, 1948; MCKEE,1958; and SCHLANGER, 1964). Deposits of the reef wall are characterized by corals in growth position, encrusting Algae and encrusting Foraminifera, thick-walled Foraminifera, and a few thin-walled Foraminifera. Much of the rest of the reef-wall material consists of unoriented debris of calcareous Algae and Foraminifera in a sandy limemud matrix. The encrusting organisms cross both corals in growth position and adjoining matrix. Stratification is absent. The material of the reef-breccia area consists of the same organisms found in the reef wall; but in the reef-breccia, all components have been transported, i.e., they are detrital. Their detrital origin is indicated petrographically by the restriction of encrusting organisms to the interiors of transported fragments of reefwall rock in contrast with their function as a binder in the reef-wall material. The internal structure of corals may be cut at its contact with the matrix and the deposit may show vague stratification. The off-reef areas are characterized by well-sorted lime sands that consist of large Foraminifera and algal fragments. The fore-reef detrital area is characterized by large benthonic Foraminifera, which may be present as broken and angular fragments, and remains of corals, echinoids, and mollusks. Sediments in the fore-reef transitional area consist of well-bedded fine lime mud that contains abundant Foraminifera, among which the globigerinids are abundant. Deposits of the basinal area in deeper water consist of a preponderance of planktonic Foraminifera and few benthonic Foraminifera; the generally fragile

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tests of the Foraminifera are set in a matrix of fine lime mud, which may fill the interiors of the tests. D. Ancient biohermal suites in general. Geologists dealing with ancient limestones commonly are not able to make the various distinctions suggested here and deal with “reef” and associated rocks in more general terms. The term biohermal suite is a useful general title for them (FAIRBRIDGE, 1955). Great interest has arisen over these rocks since the discovery of large petroleum reserves in some of them (LAYER,1949; LINK,1950). General surveys of the occurrence of “reef” rucks in the geologic record have been made by DUNCAN (1870), GRABAU (1 903), SMITH (1950), LE(1912), UMBGROVE (1947), LEVET(1950), PUGH (1950), TWENHOFEL COMPTE (1958), and STUBBLEFIELD (1960). A recent summary of bioconstructional limestones (discussed under the title of “skeletal” limestones) is presented by NELSONet al. (1962). An extremely voluminous literature describes the various reef and/or biohermal (CUMINGS,1932) limestones from the geologic record. The accompanying list, though large, is far from complete. Examples are known from the Precambrian of Montana, U.S.A. (FENTONand FENTON, 1933); Lower Cambrian from various parts of the world (summarized in CLOUD,1961); Chazyan (Middle Ordovician) of Champlain Valley, New York and Vermont, U.S.A. (RAYMOND, 1924; OXLEY, 1951; and OXLEYand KAY,1959); Silurian of Gotland (HADDING,1941,1950; Jux, 1957); Silurian of the Great Lakes region, U.S.A. (CUMINGS and SHROCK,1928a, b; SHROCK,1939; LOWENSTAM, 1950, 1957; CAROZZI and HUNT,1960; TEXTORIS and CAROZZI,1964), James Bay lowland, Canada (Gussow, 1953);Devonian ofhlberta, Canada (ANDRICHUK, 1958a, b; BELYEA,1958; and CAROZZI,1961a), New York, U.S.A. (OLIVER,1956), Belgium (LECOMPTE,1954, 1959; reviewed in RUTTEN, 1956) and New South Wales, Australia (WOLF,1965a); Mississippian of Indiana, U.S.A. (CAROZZIand SODERMAN, 1962), southwest Missouri and northwest Arkansas (TROELL,1962), Oklahoma, U.S.A. (HARBAUGH, 1957), northern England (TIDDEMAN, 1889; SHIRLEYand HORSFIELD, 1940; BOND,1950; PARKINSON, 1947, 1953, 1957; W. W. BLACK,1954; and WOLFENDEN, 1958), and northern Ireland (OSWALD,1955; GEORGEand OSWALD,1957; SCHWARZACHER, 1961 ; and CALDWELL and CHARLESWORTH, 1962); Pennsylvanian of Midland Basin, Texas, U.S.A. (ELLIOTTand KIM, 1952; MYERSet al., 1956; and RALLand RALL, 1958); San Juan Canyon, Utah, U.S.A. (WENGERD,1951), Sacramento Mountains, New Mexico, U.S.A. (PLUMLEYand GRAVES,1953); Late Pennsylvanian and Early Permian of New Mexico, U.S.A. (OTTEand PARKS,1963), Permian Scurry reef, 1953), Permian reef Scurry County, Texas U.S.A. (BERGENBACK and TERRIERE, complexes of west Texas basins, U.S.A. (NEWELLet al., 1953); Triassic of Tyrolean VON MOJSVAR,1879); Upper Jurassic of the Grand-Salkve Alps (MOJSISOVICS (Haute Savoie, France), and south of Geneva, Switzerland (CAROZZI,1954; 1955a,b).


22 1

Combinations of carbonate rocks and marine evaporites with minor terrigenous sediments. Carbonate deposits commonly are interbedded with evaporitesl and variable quantities of terrigenous sediments, typically red and green shales, mudstones, and/or siltstones. Examples of Recent marine evaporites are relatively scarce but have been described from the coast of Peru, east coast of the Gulf of California in Mexico, and Persian Gulf. Geologic examples include Mississippian of Alberta, Canada; Permian of west Texas and southeastern Oklahoma, U.S.A., and Upper Jurassic Arab Formation of Saudi Arabia. The classic paper of SCRUTON (1953) has established a firm basis for interpreting evaporite sequences. According to the dynamics of the water circulation pattern required to produce evaporites, the deposits from the most concentrated water lie close to the land at the distal end of the evaporite basin (the proximal end being the open sea of normal salinity); and from this point toward the source of sea water with normal salinity a gradation exists according to the well-established sequence of relative solubilities. Hence, if conditions of halite precipitation are attained, halite will be nearest the shore at the distal end, and it will be flanked by belts of the following deposits in the seaward direction: gypsum (or anhydrite), dolostone, and limestone(s), most likely including oolite and possibly reefs or bioherms and associated varieties. If a gradual transgression takes place without changing the circulation pattern, therefore, the initial transgressive deposit is halite. The admixed terrigenous sediments occur in two places in this arrangement ( I ) along the shore itself, and (2) much farther offshore beyond the zone of chemical precipitation. Commonly the color of the shoreline deposits is red, although rarely it may be black (MOOREand HAYES,1958). On the other hand, the color of the deeper-water deposits typically is green, gray, or black. Examples of Recent evaporites that show this distribution have been reported from the east shore of the Gulf of California in Mexico (MOOREand HAYES, 1958) and the west coast of Peru (MORRIS and DICKEY,1957). Several localities have been described from the Persian Gulf (SUGDEN,1963), adjacent to the Qatar Peninsula (WELLS,1962), and on the coast of the Sheikdom of Abu Dhabi, 320 km east of Qatar (CURTISet al., 1963), and on the Trucial coast in an area covering approximately 2,300 km2 (BUTLERet al., 1965). The Persian Gulf evaporites are associated with carbonate sediments and contain only minor amounts of terrigenous sediments; in addition, they illustrate evaporite precipitation interstitially between sediment grains of tidal-flat areas, which was mentioned previously and is further discussed in Chapter 6. The striking rhythmic or cyclic pattern of ancient evaporite sequences has long attracted the attention of geojogists, but only since SCRUTON’S (1953) work the 1 “Evaporites” as used here designate gypsum, anhydrite, halite, and more soluble salts deposited from saline waters. The evidence that such carbonates as oolites and certain kinds of dolostones should be considered as evaporites has greatly increased in the last few years; this topic is exaniand SANDERS, 1967). ined further in Chapter 6 (FRJEDMAN

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correct basis for their correct paleogeographic interpretation has been established (see also SUGDEN,1963). According to the interpretation advocated here, a sequence from base upward that consists of green shale, oolite, dolostone, and gypsum, for example, indicates an approach of the shore zone to an originally offshore zone (or “regression”); whereas the reverse of this sequence (from base upward: gypsum, dolostone, oolite, green shale) represents a retreat of the shoreline (transgression). If the transgression is rapid, the sequence may be incomplete and the offshore green shale may rest directly on the marginal red shale, for example. The Mississippian Banff, Pekisko, and Shunda Formations of southern Alberta, Canada, include mixed carbonates, evaporites, and terrigenous materials. ILLING (1 959) has interpreted these on the basis of eight environments parallelling the ancient shoreline, ranging from an offshore site of deposition of argillaceous pasty sediment, to nearshore and surf-zone environments where skeletal sands, ooids, and pseudooids accumulated; and to lagoonal environments in which finegrained lime mud, silty fine-grained dolostone, and anhydrite were deposited. The basal Banff Formation consists of the offshore argillaceous and cherty limestone; the overlying Pekisko Formation includes the transition to the surf-zone skeletal sands and oolites; and the lower part of the next higher Shunda Formation consists of silty dolostone and, in the subsurface, anhydrite. Distinctive brecciated rocks found on the outcrop have been interpreted as the products of collapse of overlying materials where solution of an underlying evaporite deposit has occurred (STEARN, 1956, and MIDDLETON, 1961). A nearly comparable sequence, lacking only the anhydrite and including more varieties of nearshore terrigenous sediments, has been observed in the Mississippian rocks of the southeastern part of the Cumberland Plateau, Tennessee, U.S.A. (SANDERS, 1953; and unpublished observations). The Permian stratigraphic sequence exposed around the margins of the Delaware Basin, west Texas and southeastern New Mexico, U.S.A., consists of alternating terrigenous redbeds, halite, gypsum, dolostone, and various limestones, including oolitic, pseudoolitic, grapestone, pisolitic, and skeletal varieties, some of which have been dolomitized (P. B. KING, 1942, 1948; NEWELLet al., 1953; BOYD,1958; and MOTTS,1962) The ancient environments produced the following deposits from northwest (side of former land) toward the southeast (former open sea): red mud, halite, gypsum, fine-grained dolostone, pisolites, and dolomitized coquina and calcarenite. The pisolites have recently been shown to be of subaerial origin (R. J. DUNHAM,1965). These deposits accumulated on a broad shallow basin-marginal shoal area (“shelf” or ‘‘lagoon’’ or “backreef area” of different authors), in which the salinity of the water increased in a landward direction. At times near the submerged edge of the marginal shoal the salinity evidently approached that of normal sea water, because lime-secreting organisms flourished and locally formed bioherms and reefs. As in the Mississippian of Alberta, surface outcrops show distinctive collapse breccias at places in the stratigraphic succession

ORIGIN A N D OCCURRENCE OF LIMESTONES

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where significant solution loss has taken place in evaporite beds that appear in the subsurface; the breccias do not occur in the subsurface where the evaporites are intact (G. W. MOORE,1960). The correct diagnosis of these collapse breccias and their lateral persistence and relationships to surrounding units, as proved by G. W. Moore’s detailed mapping, indicate that several important evaporite beds, now no longer present on the outcrop, formerly extended across the seaward edge of the basin-marginal shoal, and across the places where reefs grew at times when the evaporites were restricted to the more landward parts of the marginal shoal area. The Permian evaporites of southeastern Oklahoma (HAM, 1960), which include only minor carbonate rocks, are cyclic, including an offshore green shale, pellet-ooliticlimestone (now dolomitized), fine-grained, presumably evaporitic, dolostone, gypsum, and red shale. Though in general agreement with Ham’s interpretation that the cyclic arrangement reflects changes of sea level, the senior author here suggests slightly different boundaries of the cycles and different conclusions concerning the relationships of the different units to sea-level movements. Ham used the base of the persistent dolostones to limit the lower boundary of the cycles and supposed that these dolostones are the initial deposits of marine transgressions. An additional interpretation proposed by Ham is that the change from transgression to regression occurs within the gypsum beds. The alternative preferred by the (1953) paper is that the green shales mark senior author in light of SCRUTON’S rapid transgressions, and that the green shale-carbonate rocks-gypsum-red shale upward sequence represents a single regression, and not a transgression (dolostone and lower part of the gypsum) and a regression (upper part of the gypsum and overlying red shale), as supposed by HAM(1960). Dolostone is considered to be transgressive only where it overlies gypsum. According to this alternative interpretation, the cycles begin with green shale and end with red shale. Most of the section, therefore, consists of regressive deposits; only minor transgressive sediments occur. The Upper Jurassic Arab Formation includes interbedded carbonate rocks and anhydrites, but lacks terrigenous sediment admixtures (R. W. POWERS, 1962). Combinations of carbonates with terrigenous sediments (without evaporites). All of the previously discussed kinds of carbonates may be found mixed with various proportions of terrigenous sediment. The terrigenous material may be present as a disseminated admixture within an impure limestone bed, or it may be confined to discrete interbeds of shale, siltstone, mudstone, or sandstong intercalated between beds of pure carbonate rocks. The carbonates present in1these mixed sequences may vary in proportion from nearly pure limestone with scattered grains of quartz, to isolated lenses or patches of calcareous shell debris in the midst of essentially noncalcareous terrigenous sediments. Examples of such admixtures include: ( A ) admixed quartz and carbonate nearshore sands; ( B ) in situ buildups of carbonate bioherms surrounded by terri-

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genous sediment; ( C )transported carbonate shell debris surrounded by terrigenous sediment; and (D)alternations of nearshore and/or offshore carbonates and offshore, nearshore, and/or marginal marine or nonmarine terrigenous sediments caused by variations in amount and direction of supply of terrigenous sediment. A . Admixed carbonate and quartz nearshore sands. Carbonate and quartzose nearshore sands may become mixed if alongshore transportation of one brings it to the site of deposition of the other. Examples of quartz transportation into a carbonate province and of carbonate transportation into a quartz province are known from Florida, U.S.A. A classic example of a quartz “invasion” of the site of otherwise pure carbonate deposition is the east coast of Florida, U.S.A., where quartz migrating southward along shore has entered the carbonate depositional province, in some cases being responsible for terminating reef growth (SHALER,1893, pp.188-189). What may represent a comparable example additionally complicated by Pleistocene changes of sea level and climate lies on the continental shelf off the southeastern United States south of Cape Hatteras. A nearshore zone of quartz sand 30 km wide off Onslow Bay and Charleston lies next to an offshore zone of pure carbonate sands (STETSON, 1938, 1939; GORSLINE, 1963). Any shift in the boundary between these two sands would produce discrete interbeds of sandstone and limestone. Offshore carbonate sediment is being transported westward by alongshore currents into a predominantly quartz-sand environment of deposition on the continental shelf in the northeastern Gulf of Mexico, off the Florida panhandle, Alabama, and Mississippi (G. M. Friedman, personal observations). Off the west coast of Florida a nearshore quartz sand belt about 30 km wide passes seaward into mixed quartz and shell sand and finally into pure carbonate shell sand. Off the northern Florida panhandle and Alabama, the nearshore quartz zone is 80 km wide and carbonate sediments lie to the seaward of it. Particle size is controlled principally by the type of carbonate material present (GOULDand STEWART, 1955). other grain-size studies(G.M. Friedman, unpublished) show that an increaseinmean size takes place both toward the edge of the continental shelf, where the carbonate sediments predominate, and from west to east (Alabama to Florida), which reflects the increase in the carbonate fraction toward Florida. The Cambro-Ordovician carbonates of the Appalachians, Ozarks, and elsewhere in the central and southern United States include interbeds of clean quarts sandstone and scattered quartz grains within carbonate beds. The deposition of this quartz may have been the result of alongshore transportation of quartz into a carbonate depositional province, as along the east coast of Florida. The Ely Limestone (Pennsylvanian) of northeastern Nevada (DOTT,1958) and Jeffersonville Limestone (Devonian) of southern Indiana (PERKINS, 1963), may represent additional comparable examples. Another possibility for producing mixed carbonates and terrigenous sediments exists where pure nearshore carbonate sediments grade offshore into terri-


22s

genous fine-grained sediments that have been supplied from a relatively distant source, or which have by-passed the carbonates and have spread out laterally along the shore. Examples of such nearshore carbonates and offshore terrigenous sediments are not known to the authors from the Recent sediments. Examples from the stratigraphic record include the Mississippian of Alberta and southern Saskatchewan, Canada, and Tennessee, U.S.A. ; cyclic Helderbergian limestones of the Hudson Valley, New York, U.S.A.; and Black River-Trenton sequence (Middle Ordovician) of southern Ontario, Canada. The Mississippian deposits of Alberta and southern Saskatchewan, Canada (ILLING,1959) include nearshore pure carbonate deposits, such as calcarenites and oolites, which grade into finer-grained limestones containing admixed quartz silt. Comparable sediments occur in the Mississippian of the southern Cumberland Plateau, Tennessee, U.S.A. ; in addition, fine-grained sandstones occur as discrete interbeds (SANDERS, 1953, and unpublished observations; PETERSON, 1962). The Helderbergian limestones of the Hudson Valley, New York, U.S.A., include cyclically interbedded pure calcarenites, calcarenitic cherty limestones, and shaly-silty limestones with disseminated quartz. They evidently represent nearshore pure carbonate sands that graded seaward into muddy carbonate sands, and further out, into admixed carbonate mud and quartz silt (RICKARD, 1962). The succession from base upward of Coeymans-Kalkberg-New Scotland (or “Catskill shaly” of earlier workers) and Becraft-Alsen-Port Ewen Formations represents two full cycles, each cycle having been deposited during a gradual submergence. Each cycle begins with a calcarenite (Coeymans, Becraft) and ends with silty limestone (New Scotland and Port Ewen). The Alsen Formation is highly prized, because its mixture of calcite and quartz is extremely valuable for the manufacture of cement. The Black River-Trenton sequence (Middle Ordovician) of southern Ontario, Canada, has been interpreted as the product of a single onlap, with the Black River limestones representing nearshore pure carbonates and the Trenton limestones, offshore mixed carbonate and terrigenous sediments (WINDER,1960). The Black River deposits include four formations about SO m thick: basal clastics of varying thickness (Shadow Lake Formation) overlain by lime mudstones and magnesian limestones (Gull River and Moore Hill Formations) representing lagoon sediments, and upper calcarenites and coquinites (Coboconk Formation) deposited in the ancient surf zone. The overlying Trenton consists of the offshore marine deposits: the basal unit (Kirkfield) comprises interbedded pure limestones and shales from just seaward of the surf zone; and the upper two units (Sherman Fall and Cobourg) contain interbedded argillaceous limestones and shale, from farther offshore. The thickness of the Trenton is 140 m. These units were deposited over an irregular topography carved on Precambrian basement rocks. Some of the former highs project through the limestone today and form hills in the modern topography. The facies patterns around such highs offer excellent opportunities

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to determine the exact depth of water at time of deposition of the various facies. B. In situ carbonate bioherms surrounded by,fine-grained terrigenous sediments. In situ buildups of carbonate bioherms in the midst of fine-grained terrigenous sediments occur in the quiet brackish waters of bays or lagoons, on tidal mud flats, and in deeper offshore waters of normal salinity. Examples from the Recent nearshore sediments include the oyster mounds of the central Texas bays (HEDGPETH, 1953; NORRIS,1953; SHEPARD and MOORE,1955,1960; andLADD et a]., 1957), North Carolina coast (GRAVE, 1901), and Chesapeake Bay (J. E. Sanders, personal observations); Serpula bioherms in Laguna Madre, Texas (G. M. Friedman, personal observations); Mytilus banks of the Dutch Wadden Sea tidal flats (VANSTRAATEN, 1951, 1952); and barnacle bioherms of the Gulf of Maine (K. 0. Emery, personal communication, 1964; HATHAWAY et al., 1965). Deep-water examples include the coral banks of the North Atlantic Ocean (TEICHERT, 1958); crinoid colonies in parts of the Mediterranean Sea (J. Cousteau, personal communication, 1958, and sea-floor movie shown in 1958 at International Oceanographic Congress, New York); and the coral bioherms on the Blake Plateau and shelf edge, northeastern Gulf of Mexico. On the Blake Plateau, living deep-water corals cover an extensive area of 3,100-3,900 km2 (STETSON et al., 1962). In addition to known coral bioherms, other topographic highs with the same morphology as the coral mounds probably also represent accumulations of coral skeletal debris. The deep-water corals occur as scattered colonies or groups of colonies; they contain a great abundance of corals and, in addition, the diversity and numbers of individuals in almost all other groups increases, indicating that the area is the locus of high-density marine populations (STETSON et al., 1962, p.3). The Gulf of Mexico coral structures (LUDWICK and WALTON,1957) are smaller, differently shaped, and contain a different fauna from those of the Blake Plateau. The principal contributors to the Gulf of Mexico coral accumulations are calcareous Algae, Bryozoa, corals, carbonate-precipitating worms, mollusks, and Foraminifera. The absence in them of hermatypic reef-building corals is considered to be evidence of unfavorable temperature and salinity conditions. Such structures are considered to represent a form of marine carbonate deposit that is intermediate between present-day tropical reefs formed by hermatypic corals and the deeper-water coral banks formed by ahermatypic corals (STETSON et al., 1962, p.3). These deep-water coral bioherms are of importance to the geologist who deals with the stratigraphic record, for they indicate that not all fossil reefs and bioherms prove tropical shallow-water conditions, as has been generally supposed (TEICHERT, 1958). DUNCAN (1870) has made the only systematic attempt to apply knowledge of different living coral types to the interpretation of the Cenozoic and Mesozoic formations of Europe. Numerous examplesFof bioherms surrounded by fine-grained terrigenous deposits are known from the geologic record. Some of the most carefully studied


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examples are from the Upper Devonian of southern Belgium (LECOMPTE, 1954, 1959; see also RUTTEN,1956). Others include algal bioherms in the Cambrian of Vermont and Texas; crinoid bioherms of the Lower Mississippian of Indiana (CAROZZI and SODERMAN, 1962); bryozoan-crinoidal bioherms of the Mississippian of southwestern United States (PRAY,1958; TROELL,1962); and Exogyra mounds in the Cretaceous of the western interior of the United States (BERGQUIST and COBBAN,1957). Limestone bioherms thought to be of algal origin occur in at least three Cambrian terrigenous formations in the northern Appalachians, northwestern Vermont (SCHUCHERT, 1937; SHAW,1958): ( I ) in the Parker Slate (Lower Cambrian) north of the Missisquoi River; (2) Skeels Corners Slate (Middle Cambrian), about 60 m above the base in an outcrop southwest of St. Albans Hill and near the top at the Rockledge Estate; and (3) in the lower quarter of the Hungerford Slate (Middle Cambrian), about 1.6 km south of Cutler Pond (SHAW,1958). Other examples occur in the Wilberns Formation (Upper Cambrian) in Mason County, Texas. A particularly striking bioherm approximately 30 m long and 15 m thick is exposed in the south bank of the Llano River, opposite the mouth of Haney Creek (CLOUDand BARNES,1948, pl. 18 and 19A; DUNBAR and RODGERS, 1957, fig.94). Many other examples are included in the Seismographic Service Company’s extensive summary of organic reefs, bioherms, and biostromes (PUGH,1950). C. Transported carbonate shell debris surrounded by terrigenous sediments. Combinations of transported carbonate shell debris surrounded by terrigenous sediment somewhat resemble the deposits described in the preceding section, but differ from them in that the carbonate detritus has been transported and is a lag deposit rather than an organic buildup. Recent examples include the shell pavements formed by the lateral migration of creeks and shell beds formed by the action of burrowing worms on the tidal flats of the Wadden Sea, The Netherlands, and the cheniers of southwestern Louisiana and northeast Texas, U.S.A. Lateral migration of the creeks on the lower tidal flats of the Dutch Wadden Sea results in the formation of shell pavements on the floors of these watercourses. As the creeks migrate laterally, they exume additional dead shells from burrows in the material forming the cut bank and cover up the channel-floor shell pavement with fine-grained terrigenous sediments deposited on the migrating slip-off slope (point bar) (VANSTRAATEN, l950,1951,1952,1954a,b; see also SANDERS, 1957; and KLEINand SANDERS, 1964). On the sandier higher tidal flats of the Dutch Wadden Sea, thinner (approximately 2 cm thick) shell beds are found at a constant depth of 20-30 cm below the surface. These have been shown to be the lag concentrates of material uneaten by the burrowing worm, Arenicola marina; the coarser particles contain numerous shells of the gastropod Hydrobia, which has prompted the name Hydrobia bed

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1. E. SANDERS AND G. M. FRIEDMAN

for them. The coarse debris accumulated at feeding depth after falling down the intake funnel (VAN STRAATEN, 1952). The cheniers of southwestern Louisiana and northeastern Texas consist of linear beach ridges, some of which are composed of transported calcareous skeletal debris of nearshore benthonic invertebrates. These shell beaches lie against the mainland or are separated from it as spits or barriers; on their landward side they are adjacent to swales composed of fine-grained terrigenous muds or marsh peat and they grade offshore into nearshore fine-grained terrigenous sediments. The shell beaches originated when the offshore terrigenous sediment derived from the Mississippi delta was scarce or absent. Under such conditions, the waves scoured the nearshore bottom and washed up on the beach the calcareous shells of dead invertebrates, which lived in or on the nearshore bottom, along with any available quartz sand and silt particles. When the supply of terrigenous sediment from the river was large in the nearshore zone, the shore prograded and built out a mud flat; during such times, no shell beach forms. The various stages of growth of the chenier area and directions of discharge of the Mississippi distributaries have been correlated by means of radiocarbon dates (BYRNEet al., 1959; COULD and MCFAP.L.AN, 1959; KANE, 1959). Little Chenier, southwestern Louisiana, is a good example of a chenier composed almost entirely of shells; it consists predominantly of the mollusks Cuassostrea, Rangia, and Mulinia, with lesser amounts of Anadam, Busycon, Polinices, Crassinella, Cerithidae, Corbula, and Turbonilla, and minor amounts of quartz sand and silt. Little Chenier is approximately 19 km long and its width ranges from less than 30 m t o more than 500 m; its average crestal altitude is 1-1.3 m. This chenier formed about 2,800 years ago and marks a former position of the shoreline with the sea standing at its present level. Oak Grove Ridge is another chenier which lies about 8 km south of Little Chenier and consists of both quartz sand and shells. It is capped by a thin soil zone and is overlapped on both the steep seaward slope and the gentle landward slope by marsh peat, organic silt, and clay (BYRNE et al., 1959). The sand and shell fragments are coarsest along the front slope and crest of the ridge where shells are most abundant. The diverse fauna includes the Pelecypoda Mulinia, Anadara, Dinocaudium, Crassostvea, Donax, Nuculana, and Rangia; Gastropoda Natica, Polinices, Busycon, and Anachis; and Foraminifera Streblus, Elphidium, and Quinqueloculina. A geologic example of transported carbonate skeletal detritus surrounded by fine-grained terrigenous sediments occurs in the “Newberrian” Sandstone in the railway cut about 1 km south of Gummersbach station, Germany (1. E. Sanders, personal observations during a field trip conducted by Dr. R. Thienhaus in April, 1954). Several lenses consisting of large disarticulated crinoid stem plates are surrounded by terrigenous siltstone and fine-grained sandstone, in which fossils are rare. D. Alternations of nearshore andlor ofshore carbonates with various terri-


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genous sediments. Alternations of nearshore and/or offshore carbonates and various terrigenous sediments take place where the amount and/or direction of terrigenous sediment supplied vary and carbonates accumulate when the terrigenous material is absent, or where submergence and progradation alternated. The carbonate materials associated with these terrigenous materials rarely show as much variety as the interbedded terrigenous sediments; furthermore, the carbonates commonly form only a relatively minor proportion of the sequence. The terrigenous sediments may include deposits from the following environments: nearshore marine bottoms, beaches, dunes, marginal marine areas (marshes, lagoons, bays), various parts of marine deltas, coal swamps, and fluvial tracts. Although cores or sections of Recent sediments showing such alternations have not been reported, the physiographic setting required to produce them is thought t o exist in at least two areas: ( I ) the Orinoco Shelf, South America. and (2) east Asiatic shelf. Fine-grained terrigenous sediments form nearshore zones of varying width in both areas. The exposed seaward parts of the Orinoco Shelf include relatively pure calcareous sands and bioherms (KOLDEWIJN,1958; NOTA, 1958), whereas the outer parts of the Asiatic shelf include mixed terrigenous and carbonate sediments, with carbonate content typically about 20 %, but ranging up and EMERY,1961). Both areas contain minor amounts of volcanic to 60 % (NIINO materials locally. Laguna Madre, Texas (HEDGPETH, 1947; RUSNAK,1960a), a long narrow lagoon located in a subtropical arid region on the deltaic plain of southeast Texas, is an area where terrigenous and carbonate sediments are mixing. Abundant terrigenous quartz sand is moved onshore from the offshore marine bottom, across the contiguous Padre Island barrier, and is dumped into the lagoon. In the lagoon, however, carbonate sediment productivity is high; particles include grapestones, ooids, coarse-grained skeletal sands and gravel, serpulid reefs, and algal laminated sediments (FRIEDMAN, 1964, p.804). Carbonate sediments would become predominant if the supply of terrigenous quartz was reduced or diverted for any reason. More ancient examples from the stratigraphic record include subsurface Tertiary deposits of the Gulf Coast, U.S.A.; Cretaceous of Jamaica; Upper Mississippian of Eastern Interior Basin, Illinois, U.S.A.; Pennsylvanian cyclothems of central and eastern United States; cyclic Carboniferous deposits of northern England; and Mississippian of Alberta. The subsurface Tertiary deposits of the Gulf Coast region, U.S.A.,consist of 10,000 m of marine and nonmarine terrigenous sediments that grade eastward into shallow water marine carbonates (RAINWATER, 1963). Detailed lithologic descriptions are scarce owing t o the type of samples available, but the transition zone doubtless resembles the exposed Paleozoic rocks of the Midcontinent region which are described later. The Lower and Upper Cretaceous deposits of Jamaica consist of cyclically interbedded thick conglomerates, marine tuffaceous shales, and thin to t;iicli lime-

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stones which contain rudistids, other mollusks, corals, echinoids, and large Foraminifera. At least twelve cycles are present; each cycle typically begins with conglomerate at the base, includes a middle unit of tufTaceous shales (which becomes the basal unit in the absence of the conglomerate), and ends at the top with limestone. One such cycle is well exposed in the Sunderland inlier, which includes the outcrops on the south flank of an east-west-trending anticline that passes tkrough Johns Hall, 9.6 km southeast of Montego Bay. Johns Hall Conglomerate (at the base) is overlain by the Sunderland and Newman Hall Shales, which are 1,370 m thick; at the top is the Barrettia Limestone, only 7.5 m thick (CHUBS,1962). The Upper Mississippian of the Eastern Interior Basin, Illinois, U. S.A., which is about 500 m thick, includes twelve to fifteen major cyclic dternations consisting of shale (50 %), sandstone (25 %) and limestone (25 %; percentages based on thickness measurements). Some limestones thin out to the north and northeast and grade into terrigenous sediments within the limits of preserved rocks; others are truncated at the edge of preserved rock and originally extended beyond this limit. The limestones are coarser, lighter-colored, and more oolitic along the east and west flanks of the basin than they are in the basin center. The Haney Limestone, a typical example, averages 15 m in thickness, but reaches a maximum of 28 m (SWANN,1964). Swann has presented a strongly documented case for his interpretation that migration of delta lobes controlled the influx of terrigenous sediments, and hence is the basic cause of the cyclic patterns. The Pennsylvanian cyclothems of western Illinois, U.S.A., include the fol1930): lowing typical sequence (WELLER, Top

(10) Gray shale with ironstone nodules; sandy toward top.

(9) Marine fossiliferous limestone and/or calcareous marine shale. (8) Black fissile shale; restricted marine or nonmarine fossils. (7) Marine limestone; commonly absent. (6) Gray shale; fossil plants below but marine fossils locally at top. (5) Coal. ( 4 ) Underclay. (3) Underclay limestone or calcareous nodules in clay; nonmirine ( 2 ) Shale and sandy shale. Bottom ( I ) Sandstone; thin sheet or thicker channel. Entrenched surface WANLESS et al. (1963) have applied physiographic environments of deposition to these different members and have mapped the limits of the different environments in the Carbondale Formation (Summum, St. David, and Brereton cyclothems), Illinois, and Marmaton Group, midcontinent area. The limestones in the cyclothems mapped range in thickness from 1 to 5 m. A notable example is the Myrick Station-Brereton-Providence Limestone, which lies only 1 m above the Lexington


23 1

Herrin (No.6) coal; this limestone was deposited from a sea that transgressed eastward as much as 1,120 km over the former coal swamp. The Pennsylvanian cyclothems become more marine and include a larger proportion of limestone and marine shale to the west (R. C. MOORE,1950), but are more nonmarine to the east of Illinois (CROSS and SCHEMEL, 1956; BEERBOWER, 1961). The cyclic Carboniferous strata of northern England resemble cyclothems (MILLER, 1887; HUDSON,1924; ROBERTSON, 1948; K. C. DUNHAM, 1950; RAYNER, 1953; D. MOORE,1958, 1959; G. A. L. JOHNSON, 1960). The suggested standard cyclothem for the Carboniferous of the northern Pennines is as follows (K. C. DUNHAM, 1950): Top

(7) Coal. (6) Ganister or underclay. (5) Sandstone. (4) Sandy shale, shaly sandstone or grey beds; interbedded sandstone, siltstone, shale. (3) Unfossiliferous (nonmarine?) ferruginous shale. (2) Marine shale. Bottom ( I ) Marine limestone.

A much simpler of alternation, involving only nearshore-offshore marine carbonates and nearshore terrigenous sediments, is found in the Banff Formation (Mississippian) of Alberta (SPRENG,1953). Marginal marine eolian carbonate deposits. Wind-blown accumulations of calcareous grains, typically found in coastal dune fields where the calcareous sediment comes from marine organisms thrown up on the beach by the waves, may occur in pure carbonate sequences or locally may be surrounded by noncalcareous terrigenous sediments. Although strictly speaking such dunes are nonmarine sediments, they are included here with the discussion of the complex of sea-marginal environments. The term eolianite (sometimes spelled “aeolianite”), coined by SAYLES (1929) and defined by him as “all consolidated sedimentary rocks which have been deposited by the wind” (SAYLES,1931, p.390), has been applied most commonly to Pleistocene windblown calcareous sands, and less commonly, to quartzose sands. Pleistocene calcareous eolianites typically contain land snails and show welldeveloped large-scale angle-of-repose cross beds (Fig.6,7); their composition is governed by the characteristics of the calcareous debris on nearby beaches. Sayles’ term grew out of his studies in Bermuda, where calcareous dune sands, first diagnosed by NELSON (1840), consist of the skeletal remains of such lime-secreting organisms as large benthonic perforate Foraminifera; miliolids and hyaline Foraminifera; fragments of crustose and articulate coralline Algae,

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Fig.6. Typical eolianite cross-bedding in three dimensions. Pleistocene eolianite. Skeletal sand, S t . George’s Island, Bermuda. (Photography by G. M. Friedman; in: HAM, in preparation; by permission of American Association of Petroleum Geologists, Tulsa, Okla.)

Fig.7. Truncated cross-beds typical of Bermuda eolianites. Pleistocene eolianite. Skeletal sand, south shore of Bermuda, near Coral Beach. (Photograph by G. M. Friedman; in: HAM,in preparation; by permission of American Association of Petroleum Geologists Tulsa, Okla.)


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such as Amphiroa; mollusks; and other macroinvertebrates (FRIEDMAN, 1964). Bahamian eolianites form island chains along the edge of the banks (Fig.8), as on the Great Bahama Bank (NEWELL,1961); they consist largely of ooids (FRIEDMAN, 1964a). Along the Mediterranean shores of Israel, eolianites are known as kurkar (LOEWENGART, 1928); they consist of variable amounts of quartz and calcareous skeletal debris of mollusks and other invertebrates. The consolidation of these old dunes, which are Late Pliocene to latest Pleistocene in age, has been discussed by PICARDand SOLOMONICA (1936), PICARD and AVNIMELECH (1 937), PICARD (1 943), and most recently by FRIEDMAN (1964a, pp.797-801). Eolianites on Mauritius island consist of mixed coral and invertebrate shell debris (MCINTIRE, 1961, p.44). Other eolianites occur along tropical and subtropical shores of the western Atlantic Ocean, Mediterranean Sea and Indian Ocean (African coasts), and along the coast of Australia. The largest known tract of calcareous eolianites is located in Western Australia, where they comprise the Coastal Limestone mentioned earlier.

Fig.8. Island (cay) composed of eolianite, west edge of Great Bahama Bank. Pleistocene eolianite. (Photograph from approximately 300 m by G. M. Friedman; in: HAM,in preparation; by permission of American Association of Petroleum Geologists, Tulsa, Okla.)

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Open-sea carbonate associations Open-sea environments are also the sites of distinctive associations of carbonate materials, both pure and admixed with terrigenous sediments, as in the shallowwater and nearshore types just mentioned. As used here, “open sea” refers to shallow or deep areas distant from the land. Where they are shallow, open seas give rise to sediments that are governed exclusively by conditions of remoteness from land masses; in them vertical sedimentation alone takes place. Where open seas are deep, the additional factor of gravity displacement of sediments down submarine slopes may be introduced. Where such gravity displacement is possible, both vertical and lateral sedimentation processes may occur; in the absence of lateral gravity displacement, however, only vertical sedimentation may occur, even in deep seas. Both pure calcium carbonate and/or mixed carbonates and noncalcareous terrigenous deposits may result, either from the exclusive action of vertical sedimentation processes, or by the combined action of vertical and lateral sedimentation processes, In the following discussion the pure carbonates are considered first and the mixtures of carbonates and noncarbonates, afterwards. Whereas nearshore carbonates tend to be characterized by cyclic sequences that include many distinctive lithologic types in symmetrical or asymmetrical cycles, open-sea carbonates tend to be characterized by simpler repetitive alternations, generally involving only two lithologic types, and, less commonly, more types. Pure carbonate deposits from open-sea environments. Pure carbonate deposits may result from the exclusive vertical sedimentation of planktonic organisms (pelagic sediments), or from the combined operation of such vertical sedimentation with the lateral effects of gravity displacement, where the displaced material introduced laterally consists of pure calcium carbonate sediments. A. Pelagic carbonate deposits. Pelagic carbonate deposits consist of the tests of planktonic invertebrates or plants and the remains of swimming organisms that fell to the bottom. Benthonic organisins may be present in small quantities, but if they are abundant the term “pelagic” is not appropriate. A low rate of vertical accumulation is a characteristic of pelagic deposits; 1.34 g/cm2/1,000 years is the rate of accumulation of carbonate material in the post-glacial Recent deep-sea sediments from the equatorial Atlantic. A larger rate, 2.80 g/cm2/1,000 years, was found for Pleistocene glacial sediments in the same area (BROECKER et al., 1958). Accordingly, only a few meters of pelagic deposits may represent long segments of geologic time. Pelagic carbonate deposits include the Recent deep-sea oozes; Recent shelf sediments where terrigenous sediment supply is absent; the Jurassic Aptychus limestones of the Alps and Carpathians; Cretaceous chalks of northwestern Europe and parts of the United States; Cretaceous limestones in parts of the Swiss Alps; and Montpellier limestones (Eocene-Miocene) of Jamaica.


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Recent deep-sea calcareous oozes consist of tests of Foraminifera, pteropods, and coccolithophorids. They were first defined and studied in monographic fashion by the “Challenger” expedition (MURRAY and RENARD,1891). Most of the photographs and descriptions of “oozes” in the Challenger reports were made on the basis of the washed residues in which only the skeletal remains were present; they, therefore, tended to foster a misleading impression of the amount and kind of other materials in these sediments. A later systematic description of Recent from the Pacific Ocean and thorough discussion of classification was made by REVELLE (1944). According to the revised definitions in Revelle’s classification, a pelagic sediment is called an ooze if it lacks terrigenous admixture and contains more than 30% skeletal remains; calcareous oozes contain more than 30% calcium carbonate. A useful summary of the distribution of Recent calcareous deep-sea sediments has been made by RODGERS (1957). The Upper Jurassic part of the Czorsztyn series of the Pieniny-Klippe zone, Central Carpathians, Poland, consists of alternating pelagic and neritic pure limestones. The Lower Tithonian includes at the base a red CuZpionella limestone, 2-3 m thick, and an overlying white CuZpionellu limestone that is 4-11 m thick. Both also contain microscopic algal remains of Globochaetes. The Middle Tithonian limestone, 6.2-8 m thick, contains a neritic fauna of brachiopods and crinoids. The Upper Tithonian includes limestones with crinoids and Aptychus and is 0.6-2 m thick (BIRKENMAJER, 1953, 1958). The Aptychus limestones of the Alps are also examples of pelagic carbonate rocks. The Cretaceous chalk of northwestern Europe ranges in thickness from 330450 m. This distinctive deposit consists of more than 98 % calcite and only minute quantities of clay, quartz, and some magnesium carbonate. It was first examined in detail by SORBY(1861) and by HUXLEY (1898), whose conclusions have been substantiated by modern painstaking researches (M. BLACKand BARNES,1959). M. Black and Barnes used centrifuge sedimentation techniques to separate the chalk into its constituent grain sizes and found that the organic debris occurs in sizes that range in diameter from less than 1.0 mm to 0.001 mm (1 p). Very little material occurs in the 1.0-0.1 mm fraction; calcite prisms from the outer layer of Inocerarrtus shells, Foraminifera, and various shell fragments form frequency maxima in decreasing order of size in the 0.1-0.01 mm fraction. The bulk of the fine-grained chalk, however, consists of coccoliths (parts of the cell membrane of unicellular calcareous Algae) in the 0.01-0.001 mm fraction. Aggregates of coccolith particles become abundant in the 6-p size fraction; these tend to break down to individual coccolith particles whose diameter is 1 p, the smallest size present. Some geologists have compared the chalk to Recent deep-sea Globigerina ooze based on the Foraminifera present and have inferred that chalk represents an ancient deep-sea sediment (SUJKOWSKI, 1931); but ecologic studies of the gastropods, sponges, echinoids, and benthonic Foraminifera suggest depths ranging

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from 200 to 800 m (JUKES-BROWNE, 1900,1903, 1904; HATCHet al., 1938). The socalled “hard chalks” contain more abundant organic remains of originally aragonitic shells that lived in shallower water. In the London-Paris Basin, hard chalks are more abundant around the basin margins; these chalks pass laterally into shallow-water shell-fragment limestones and sandstones. The more pelagic soft chalks occur in the center of the basin (M. Black, 10 Dec., 1959, lecture at Yale University; see HATCH et al., 1938). Chalk is best regarded as a pure pelagic deposit the distinctive characteristics of which are due to ( I ) abundance of calcite organic remains, particularly coccoliths; (2) lack of aragonite organic remains; (3) absence of terrigenous admixture and (4) lack of contemporary lateral gravity displacements. Chalk usually is not regarded as a deposit formed under the influence of particularly deep water. The Cretaceous chalks of Alabama, Mississippi (STEPHENSON and MONROE, 1940), Texas, and Kansas, U.S.A., are good examples of undoubted shallow-water chalks; they are both overlain and underlain by and pass laterally into nearshore marine terrigenous sediments. The Upper Cretaceous of the Helvetic nappes, Haute Savoie, France, includes 145 m of white sublithographic limestone and slightly marly and chalky rock. The organic remains of this pelagic deposit include abundant tests of the planktonic Foraminifera, Globigeriita and Globotruncana, and lesser amounts of Iitoceramus prisms, remains of planktonic crinoids and ostracodes, and some benthonic organisms (CAROZZI, 1953). The Montpelier facies of the “White Limestone” (Middle Eocene to Lower Miocene) of Jamaica consists of bedded chalks, generally with flint nodules; or finely crystalline globigerinid limestones, which are the offshore equivalents of shallow-water marine bank deposits of the Clarendon Bank, previously discussed. More than 300 m of Montpelier Limestone are present on the north flank of the Palmyra Hills (VERSEY,1962). Post-Jurassic pelagic limestones differ notably from earlier pelagic deposits owing to the great flowering of calcareous Foraminifera, such as the globigerinids, in the Cretaceous Period and their continued abundance in the seas from the Cretaceous to the present time. B. Combinedpelagic and gravity-displaced pure carbonate deposits. Combined pelagic and gravity-displaced pure carbonate deposits occur only where the opensea environment was deep enough for gravity displacement of pure carbonate materials from an isolated, shallow-water marine bank to take place. Gravitydisplacement processes include slumps, slides, turbidity currents, and flowing-grain sand layers (SANDERS, 1965); they represent a kind of near-instantaneous lateral sedimentation that is responsible for introducing material of shallow-water provenance into deep-water depositional sites. The rate of sedimentation of gravity-displaced material may be very high; beds whose thickness may be 1-3 m (commonly less but rarely more) accumulate in a few minutes or hours at most. Their overall rate


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of sedimentation, therefore, is a function of the frequency and volume of each displacement; the average rate may be in excess of 1,000 cm/1,000 years. The volume of “fine-grained” gravity-displaced sediment that travels in suspension in a turbidity current may exceed that of the“coarse-grained”sediment that is usually associated with this sedimentary mechanism. Every effort should be made to distinguish between fine-grained gravity-displaced (lateral) sediment and fine-grained pelagic (vertical) sediment. Faunal criteria are most useful in this case. Examples of combined pelagic and gravity-displaced pure carbonate sediments include the Recent sediments of the Bermuda Apron; Tongue of the Ocean, Bahamas; and southern part of Hatteras abyssal plain. Ancient examples from the stratigraphic record include parts of the Tertiary deposits of Italy, and the calcareous Jurassic flysch of the Polish Carpathians and Swiss Alps. The Bermuda Apron refers to the steep submarine slopes that surround the Bermuda Islands and lead into the surrounding deep water from the shallowwater platform on which the islands are located. Carbonate sediments of Recent and Pleistocene age sampled by Lamont Geological Observatory have been studied by FRIEDMAN (1 964). These sediments from the apron are essentially identical to the nearshore sediments, but show well-developed graded bedding. The lower part of each graded bed consists of sands containing typical nearshore skeletal debris; mineralogically they consist of aragonite and high-magnesian and low-magnesian calcite. The tops of the graded beds consist of silt or “mud” size grains that are composed entirely of low-magnesian calcite. Whether this low-magnesian calcite reflects environmental, grain-size, or diagenetic control is not known. The graded beds have been attributed to the action of turbidity currents (ERICSON et al., 1961). Some of the deep-sea sediments from the Bermuda Apron contain Globigerina or Orbulina, which have been derived, in part, from the flanks of the Bermuda rise as turbidity currents swept the sediments downslope; and, in part, from vertical settling during periods of quiescencebefore the next graded bed was deposited (FRIEDMAN,1964, p.781). The origin of the deep-sea carbonate sediments of the Bermuda Apron, therefore, is readily understood by these two contrasting depositional processes: ( I ) lateral gravity displacement by turbidity currents; and (2) vertical pelagic sedimentation. The Tongue of the Ocean, a narrow channel more than 2,000 m deep, is surrounded on three sides by the Great Bahama Bank (Fig.9). The contiguous Northeast and Northwest Providence Channel and Tongue of the Ocean form a network of submarine canyons (€€Ess,1933). The very steep rocky walls of the channel are composed of Tertiary and Cretaceous sedimentary rocks (HEEZEN et al., 1959, pp.35-36). The Tongue of the Ocean is underlain for the most part by carbonate sediments with well-developed graded beds that contain displaced shallowwater organisms. In this respect it is similar to many other deep-sea areas that are adjacent to steep slopes and contiguous shallow-water zones.

23 8

J. E. SANDERS A N D G . M. FRIEDMAN

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The nearshore carbonate debris was originally deposited along the marginal escarpment of the Bahamas and was later displaced laterally by gravity, probably by turbidity currents. The sediments contain ooids along with the biogenic carbonate material; the particles are less well sorted and the grading is less obvious than in terrigenous turbidites found in abyssal plains of other oceanic areas


239

(RUSNAK and NESTEROFF, 1964). This characteristic has been related to the small size of the Tongue of the Ocean, which limits the distance over which sorting due to lateral transportation can occur; and to the great diversity of biogenic debris, such as Halimeda plates, pteropods, and Foraminifera, which differ widely in shape and hence also in hydraulic behavior. Moreover, especially in the upper level of the sequence, these sediments are coarser-grained than terrigenous turbidites. Sediments from the Tongue of the Ocean, kindly provided to G. M. Friedman by G. A. Rusnak, were fine-grained, poorly-sorted skeletal sands, some of which contain interstitial lime mud. The fine-grained tops of the graded beds commonly contain a deep-water marine fauna, such as pteropods and Globigerina. The age of the carbonate sediments studied ranges from Late Pleistocene to Recent (25,000 to 800 years; RUSNAK et al., 1963). Turbidity currents probably were active throughout this interval. The sediments contain 30-60 % high-magnesian calcite, a larger amount than in comparable shallow-water sediments of the Bahamas (FRIEDMAN, 1964, p.796). The upper 80 cm of core A 167-49 (collection of Lamont Geological Observatory) from the Tongue of the Ocean consists of lime mud with thin sandy interbeds, and is underlain by 30 cm of sand and 260 cm of lime mud (ERICSON et al., 1961). In this, and other cores from the Northeast Providence channel, the lime mud consists of silt- and clay-sized calcareous particles, planktonic Foraminifera, coccoliths and minute calcareous plates secreted by Protista. The lime muds contain marcasite and hydrotroilite and vary in color from light tan to greenish-gray. The sandy sediments are composed largely of tests of planktonic Foraminifera or shallow-water biogenic particles. On the southern part of the Hatteras abyssal plain, calcareous turbidites derived from the Bahama Banks, are interbedded with deep-sea calcareous ooze (SCHNEIDER and HEEZEN,1965). Parts of the Tertiary section of Italy contain graded limestones. The Eocene calcare alberese north of the Apuane Alps (TENHAAF,1959, p.71), Miocene limestone on highway 80 north of the Gran Sasso massif (KUENENand TEN HAAF, 1956), and early Oligocene brecciole nummulitiche of Tuscany (TENHAAF,1959, p. 91) are examples. The Upper Jurassic (Malm) of the Nappe de Morcles and autochthonous chains, Haute Savoie, France, includes dark pyritiferous limestone 100-1 50 m thick with nine interbeds of calcareous microbreccias composed of reef and nearshore calcareous organic remains (KUENEN and CAROZZI,1953; CAROZZI,1955b). The contemporaneous reefs are exposed around the Aiguilles-Rouges massif; although the reef debris was shed southward into deep water, later thrusting has moved the deep-water deposits to a present position north of the reefs. The Cieszyn Limestone (Lowest Cretaceous of Silesian nappe sequence) of southwestern Poland represents a relatively pure limestone succession of mixed turbidite and pelagic deposits. A single graded bed 3 m thick is exposed in a small

240


quarry south of Goleszow, about 10 km east of Cieszyn (J. E. Sanders, observations on field trip with M. Ksiazkiewicz and S. Dzulynski, Sept., 1960).

Mixed carbonates and noncarbonates from open-sea environments. Carbonates may be mixed with noncarbonates (terrigenous) sediments by vertical sedimentation processes in the open sea or by combined vertical and lateral sedimentation processes. As with the nearshore deposits, the mixed material may be disseminated within individual beds, or be present as discrete layers interbedded with relatively pure carbonate layers, or by some combination of these two. Where terrigenous admixture exceeds 30%, the sediment is no longer an “ooze” but becomes an “organic mud” if the content of organic tests exceeds 30 %, and an “inorganic mud”, 1944). The organic conif the content of organic tests is less than 30 % (REVELLE, stituents of “muds” may be identical to those of “oozes”, however. Organic muds and related organic terrigenous sediments are known as hemipelagic sediments. Hemipelagic sediments may occur alone or in association with gravity-displaced and/or pelagic sediments. A . Hemipelagic sediments. Hemipelagic sediments by definition represent admixtures of terrigenous sediments with otherwise pure pelagic sediments; they occupy intermediate positions between the land and the open sea. Such sediments may exist by themselves, or they may be interbedded with pelagic sediments or with neritic deposits which formed nearer shore. Examples of Recent hemipelagic calcareous organic muds have been described from the Pacific Ocean (REVELLE. 1944). Ancient examples of mixed hemipelagic and pelagic sediments may be present in the West Castleton (Cambrian) and Deepkill (Lower Ordovician) Formations of the Taconic sequence, New York, U.S.A. Both formations include alternations of fine-grained (pelagic?) limestone beds 1-5 cm thick and shale (hemipelagic?) beds 5-20 cm thick (Fig.10). The environmental interpretation of such interbedded successions of finegrained limestones and clays or shales may be extremely difficult, for such sediments may also accumulate in nearshore environments. No detailed studies of this question are known to the authors, but it seems that the problem may be resolved on the basis of faunal and possibly mineralogic criteria. B. Combined carbonate and noncarbonate sediments of hemipelagic and gravity-displaced origin. In the deeper parts of marine basins calcareous and noncalcareous sediments may be mixed in a large variety of ways by the combined action of pelagic and/or hemipelagic deposition and gravity displacement. The carbonate sediment may consist of first-cycle particles in situ, or of displaced shallow-water first-cycle grains, or of gravity-displaced calcareous terrigenous deposit that has been recycled. The noncarbonate material may arrive by vertical settling out from a diffuse suspension or by the various lateral processes of gravity displacement acting along or close to the bottom. The former group invariably consists of fine-grained sediment, but the grain size of the gravity-displaced material may


24 1

Fig.10. Interlayered fine-grained limestone (pelagic?) (white) and siltstone (hemipelagic?), West Castleton Formation (Cambrian), south of Hudson, N.Y. Height of compass is 8 cm. (Photograph by J. E. Sanders.)

range widely. A great variety of sediments may arise from the interplay of these processes and materials. Examples of mixed carbonate and non-carbonate Recent deep-water opensea deposits occur on the northern part of the Hatteras abyssal plain, Puerto Rico Trench, and in the Gulf of Eilat (Aqaba), Red Sea. Examples from the geologic record include the brecciolus (Upper Cretaceous-Eocene) of northern Apennines Italy; Plattin Flysh (Upper Cretaceous) of Switzerland; Upper Cretaceous of Westphalia, Germany; Lower Cretaceous of central Carpathians, Poland; Jurassic breccias of Haute Savoie, France; Permian deposits of Delaware Basin, Texas, U.S.A.; Silurian of central Nevada, U.S.A.; Ordovician flysch of the Gasp6 Peninsula, Quebec, Canada; and Cambro-Ordovician brecciolus at Troy, New York, U.S.A. In addition, the striking deposits known as wildflysck may contain abundant limestone boulders. Examples of these include the Paleocene-Eocene shales, northwestern Venezuela; Mississippian of the Ouachitas, Oklahoma and Arkansas, U.S.A.; and Cambrian and Ordovician deposits of the northern Appalachians in Newfoundland and Quebec, Canada, and in Vermont and New York, U.S.A. The northern part of the Hatteras abyssal plain is the site of sedimentation of quartzose terrigenous sediments, that have been displaced laterally by turbidity currents, and pelagic calcareous ooze that has accumulated between arrivals of the quartzose material (SCHNEIDER and HEEZEN, 1965). The sediments from the floor of the Puerto Rico trench include mixed carbonates and terrigenous grains. Algae, sponge spicules, mollusk debris and Bahamian-type grains probably come from the Bahamas and ooids, from the Puerto

242

J. E. SANDERS AND G.

M. FRIEDMAN

Rico shelf. The terrigenous grains are composed of volcanic materials, quartz, feldspar, pyroxene, amphibole, and olivine, all probably originating from the Puerto Rico shelf (CONOLLY and EWING,1965). The Gulf of Eilat (Gulf of Aqaba), one of the northern branches of the Red Sea, is a narrow, steep-sided tectonic valley the average width of which ranges from 10 to 20 km, with a maximum of 26 km. The Gulf forms part of the Great Rift Valley that extends from Turkey to the Red Sea and beyond through East Africa. This trench is so narrow and steep-sided that Leopold von Buch (in: ROBINSON, 1841, p.673, quoted by GREGORY, 1921, p.17) described it as “a crevasse in the earth‘s crust”. The submarine slopes of the Gulf are fault surfaces; they are virtual precipices with a normal gradient of 60-70 %. The shore cliffs are for the most part equally precipitous. The Gulf sediments sampled in core V 14-126 (collection of Lamont Geological Observatory) reveal diagnostic characteristics of the action of turbidity currents. Graded beds occur, but the grading is less apparent than in most deep-sea terrigenous sediments. The content of carbonate materials ranges from 20 to 75 % of the total. Aragonite is sporadic or absent and the unusual combination of high-magnesian and low-magnesian calcite is present reflecting the mineralogy of the faunal assemblage (FRIEDMAN, 1965a). The sediments of the Gulf of Eilat resemble graded sediments from other deep-sea areas. They differ from the carbonate sediments of the Tongue of the Ocean, Bahamas, mentioned earlier, in that they contain in addition to their carbonate materials terrigenous sediments derived from the Precambrian crystalline rocks of the AraboNubian Shield, which forms the shores of the Gulf (FRIEDMAN, 1965a; and in preparation). The term brecciolas refers to graded limestone breccia beds that alternate with dark-colored shales in the upper part of the scugliu, a formation of Late Cretaceous to Eocene age in the northern Apennines, Italy. The brecciolas have been interpreted as the deposits of turbidity currents, whereas the dark-colored shales have been considered to be offshore deep-water marine sediments (KUENEN and MIGLIORINI, 1950, pp.111-112). The Plattin Flysch (Upper Cretaceous, Maastrichtian) of Switzerland consists of interbedded calcilutites and fine-grained terrigenous sandstones that show large current ripples on their upper surfaces, and current cross-laminae and convoluted laminae within almost every sandstone bed. A particularly good exposure is in the road cut at Jaun Pass, Switzerland. The limestones are interpreted as pelagic deposits; and the terrigenous sandstones, as probable turbidity-current deposits. The currents which last influenced the sand traveled from southeast to northwest (shown to J. E. Sanders by Prof. J. Tercier, on a field trip, 14June, 1954). The Upper Cretaceous of Westphalia includes interbedded graded carbonate beds, some of which contain glauconite, and marls. Their environment has been considered to be neritic (VOIGTand HANTZSCHEL, 1964). The lower 10 rn of the Flysch-Aalenian (Lower Cretaceous) of the Pieniny


243

Klippen zone in the central Carpathians, Poland, include alternations of shale and sandy graded crinoidal limestones 2-20 cm thick and sideritic marly limestone. The crinoidal limestone includes clasts resembling the Middle Triassic limestones and dolostones; some single beds consist of a graded crinoidal limestone in the lower part and the marly sideritic limestone in the upper part, suggesting that both owe their origin to lateral sedimentation in spite of the fine grain size of the sideritic marly limestone. The remainder of this flysch formation consists of mostly noncalcareous terrigenous sediments (BIRKENMAJER, 1957). The stratigraphic sequence of the Nappe de la Br&che,Haute Savoie, France, includes a lower (Middle Jurassic) and upper (Upper Jurassic) unit, in which are intercalated calcareous breccias composed of gravity-displaced material and finegrained pelagic limestones. The two breccia units are separated by 250-300 m of fine-grained terrigenous sediments, presumably of heinipelagic origin. At Pic Marcelly, France, the lower breccia unit is 1,300 m thick and contains blocks of Triassic and Lower Jurassic limestones as large as 8 m long and 2 m thick. Toward the east the thickness of this unit decreases by one-third, to 300400 m, and the maximum particle size, to 20 cm. The upper breccia unit is 200 m thick and its maximum particle size does not exceed 1 m. These breccia beds pass laterally into open-sea limestones, calcareous shales, and silexites (KUENEN and CAROZZI,1953). The Permian rocks in the Delaware Basin, west Texas and southeastern New Mexico, U.S.A., include interbedded sandstones, coarse-grained calcarenites, and calcirudites. These rocks are considered here because of the more widely held view that they are ancient deep-water deposits, the position supported by G. M. Friedman. The senior author of this chapter, however, does not agree with this bathymetric interpretation, according to which the water depth in the Ddaware Basin ranged from 250-300 m when the Goat Seep reefs were active at the northwestern basin margin, and from 360-550 m during the time of the Capitan reef(s). Reefs grew along the seaward margin of an extensive shelf that lay northwest of the basin (based on available outcrops; subsurface data indicate that the shelf extended to the north, northeast, and east, as well). Along the Guadalupe Mountain front two reefs crop out: an older, Goat Seep reef, and a younger, Capitan reef (LLOYD, 1929; LANG,1935, 1937; HILLS,1942; KING, 1942; 1948; NEWELLet al., 1953; NEWELL,1957). Debris from these reefs was shed into the deeper water of the Delaware Basin. The reefs grew rapidly to keep pace with subsidence; detritus from them cascaded down steep submarine slopes and spread out into the basin as fans. The so-called fore-reef beds were deposited on the steep submarine slopes; their dip of 30" is inferred to be an initialone. As a result of storms, earthquakes, or other catastrophic events, the oversteepened unstable sediment mass that built up along the basinward side of the reefs was moved basinward by submarine slides and turbidity currents. Well-developed graded beds, some of them with very coarse pebble-sized debris at their base, are interbedded with the finer-grained, thinlybedded and commonly finely laminated sediments that formed on the bottom of

244

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E.

SANDERS AND G . M. FRIEDMAN

the Delaware Basin. Graded beds increase in abundance toward the edge of the basin, but locally, beds of coarse carbonate debris are found even near the center of the basin. In addition to the graded beds, many slump features and indications of intrastratal flow along limestone beds resulted from the unstable detrital load on the steep gradients. Some of these penecontemporaneous slump features include overturned folds and thrust faults. J. E. Sanders, on the other hand, supports many of the arguments advanced by G. W. MOORE(1959, 1960) based on mapping of collapse breccias in the basinmargin areas, characteristics of the unfossiliferous parts of the massive Capitan Limestone, and sedimentary structures in the basinal terrigenous sediments; which, if correct, cast doubt on many important aspects of the foregoing reef interpretation. The subject cannot be resolved in this brief summary, and will require additional critical appraisal of the outcrop, subsurface, and petrographic data by a new generation of observers who are not involved with the older interpretations. The Roberts Mountain Formation (Silurian) of central Nevada, U.S.A., consists of interbedded graded calcarenites ranging in thickness from a few centimeters to 60 cm (with an average of 15 cm), and fine-grained impure limestones, siltstones, shales, and cherts. The graded calcarenites consist of material derived from the reefs, as in the Upper Jurassic breccias of Haute Savoie, France, previously discussed. The graded beds can be followed for about 0.8 km along strike in some localities. Near the sites of the former reefs, the thickness of individual graded beds increases by an average of 50 %. The fine-grained rocks are considered to be deposits of deep water (WINTERER and MURPHY,1960). The Ordovician flysch of the Gasp6 peninsula, Quebec, consists of an interbedded complex that includes five types of graywackes, two types of calcisiltites, calcareous wackes, silty shales, dolomitic silty shales, dolostones, fine-grained black limestones, and traces of volcanic ash. Carbonate rock clasts appear in only one type of the graywackes, in which their content ranges from 30 to 90 %. Carbonate content in type 1 calcisiltites ranges from 45 to 82% and consists of about equal parts of calcite and dolomite; in type 2 calcisiltites, the carbonate content ranges from 27 to 64% and is mostly dolomite. Carbonate material constitutes 84% of the fine-grained black limestones, 14% of the dolomitic silty shales, and 93 % of the dolostones. The noncalcareous silty shale is the dominant rock type; its abundance ranges from 40 to 85% of the total in the various members of the succession. The dolostone and fine-grained black limestone are considered to be pelagic sediments and the other rock types, except for volcanic ash, are interpreted as products of lateral displacement by gravity (ENOS,1964). Deposits within the West Castleton (“Schodack”) Formation (Lower Cambrian) that are identical with the Italian Tertiary brecciolas are well exposed on the campus of Rensselaer Polytechnic Institute, Troy, New York, U.S.A. Individual beds 3-7 m thick and containing abundant calcareous fragments, local coarse round quartz grains and, more rarely, small clay blebs about 5 cm across, fine-


245

grained light gray limestone fragments, and a few sandy fragments, occur in a brownish-purple, fine-grained sandstone or siltstone matrix (LOWMAN,1961). The shape of fragments in the brecciola beds ranges from irregular to round to tabular. Their average size is 5-12.5 cm with some reaching 45 cm; in addition, a few large slabs are present, the largest of which is 60 cm thick and 240 cm long. The general orientation of fragments is subparallel to the brecciola bed boundaries. Lower Cambrian trilobites (Ellipsocephala asaphoides) were collected by C. D. Walcott (cited in LOWMAN, 1961) in these brecciolas. Brecciolas belonging to the Deepkill Formation (Lower and Middle Ordovician) are also found near Troy. A curious kind of deposit of particular interest in ancient geosynclinal tracts is the so-called “wildflysch”. Wildflysch consists of clasts of varying size, some of them enormous, set in a matrix of fine-grained, typically dark-colored marine shale, siltstone, or mudstone; they are typical examples of diamictites (FLINTet al., 1960). Wildflysch clasts include almost any kind of rocks; only those in which carbonate rocks predominate are discussed here. Characteristically, wildflysch limestonebouldery mudstones are interbedded with terrigenous sequences that lack abundant carbonates and which show other indications of deposition in deep water, such as fauna or repeated intercalations of graded beds. Wildflysch diamictites may be regarded as the marine equivalents of the nonmarine fanglomerates with limestone boulders discussed in a previous section. Three examples include the PaleoceneEocene shales of northwestern Venezuela; Johns Valley Shale (Mississippian) of the Ouachitas in Oklahoma and Arkansas, U.S.A.; and the Cambro-Ordovician of the northern Appalachians in Newfoundland and Quebec, Canada, and Champlain-Hudson Valleys, Vermont and New York, U.S.A. The Paleocene-Eocene marine shales of Lara, northwest Venezuela, about 100 km east of Lake Maracaibo, contain Iarge clasts of pre-Cretaceous, Cretaceous, and Paleocene rocks, including both noncarbonates and limestones. Graded calcareous sandstones are interbedded with the boulder-bearing shales. Blocks of Cretaceous limestone are very large; one measured 100 m long and 30 m thick, whereas another is 1 km long and 100 m thick. The Eocene part of the shale includes boulders of crystalline rocks, Cretaceous limestone, and Paleocene Lithothamnium limestone. The Paraguito boulder bed can be followed for 40 km along strike. These exotic blocks are inferred t o be the deposits of submarine slides in the Barquisimeto trough; some of the blocks were derived from the Maracaibo shelf (RENZet al., 1955). The Johns Valley Shale includes limestone “exotic”’ pebbles and boulders in a zone 200 km long and 40-48 km wide that extends from near Atoka, Oklahoma, eastward to Boles, Arkansas (KRAMER, 1933; R. C. MOORE,1934; CLINE,1960). The shale is 130-250 m thick; it contains two zones of exotics, a lower and an upper, each including several beds of boulders. In size, the exotics range from pebbles to boulders, and some angular limestone blocks are up to 12 m long. The limestones were derived from the contemporaneous and older “Arbuckle facies”. The latter

246


includes carbonate rocks of Cambro-Ordovician, Silurian, Devonian, and Mississippian age, which are restricted to localities north and west of the boulderbearing shales; they also occur tectonically below the shales (CLINE,1960) Cambro-Ordovician marine limestone-pebble diamictites and conglomerates occur in western Newfoundland and near Quebec City, Canada, and in the Champlain Valley, Vermont, and upper Hudson Valley, New York, U.S.A. The spectacular Cow Head Limestone Breccia of western Newfoundland includes blocks up to 30 m in diameter and one that measures 120 x 108 x 6 m (SCHUCHERT and DUNBAR, 1934; DUNBAR and RODGERS, 1957; KINDLEand WHITTINGTON,1958). This unit was formerly thought to be a single layer of Middle Ordovician age at the base of a shale sequence (SCHUCHERT and DUNBAR, 1934); but more recent fossil collections (KINDLEand WHITTINGTON, 1958) indicate that similar breccias are scattered through about 300 m of strata, the age of which ranges from Middle Cambrian to Middle Ordovician, and are not limited to a single horizon. A characteristic feature is that the clasts tend to be flat. They lie in a muddy, rarely sandy matrix, and include mostly different varieties of limestone. Minor amount of clasts are composed of shale and chert; igneous or metamorphic rocks are lacking. Only a limited stratigraphic range is represented by the clasts in any given layer; for example, the upper Cow Head conglomerates include limestone clasts from the Middle Ordovician Table Head Formation (WHITTINGTON and KINDLE,1963). If the clasts were derived from a submarine fault scarp, the relief along it was low (KINDLEand WHITTINGTON, 1958). Somewhat similar rocks are found near Quebec City and LCvis, Quebec, in the Sillery and LCvis Formations (Lower Ordovician) and Quebec City beds (Middle Ordovician) (RAYMOND, 1913; CLARK,1924, 1926; BAILEYet al., 1928; OSBORNE, 1956). As in Newfoundland, the pebbles in each layer are of the same age as (or slightly older than) the formation in which the conglomerates are enclosed. The Sillery conglomerates contain mostly clasts of Lower Cambrian limestone. Those in the LBvis Formation include abundant Upper Cambrian and Lower Ordovician limestones, with a few pieces of gneiss, igneous rock, quartzite, and Sillery sandstones; and clasts in the Quebec City beds include only Middle Ordovician (Black River and Lower Trenton) limestone fragments. In the zone of interfingering of western carbonate rocks and eastern terrigenous rocks in northwestern Vermont, limestone-pebble conglomerates occur in at least ten different formations the age of which ranges from Early Cambrian to Middle Ordovician (M. H. Ross, 1949; SHAW,1958; CADY,1960). Almost all of these conglomerates, however, are local deposits with calcareous or dolomitic matrix, and with pebbles of the same age as the enclosing beds. None of them includes debris from formations with very large age spans, so that they are intraformational conglomerates (WALCOTT, 1894). By contrast, the younger Hathaway Formation (Middle Ordovician) of the Champlain Valley, Vermont, includes large clasts of carbonate rocks, quartz


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sandstone, and coarse-grained graywacke, and small grains that include debris from igneous and metamorphic rocks. These have been interpreted as submarine slide breccias formed along a fault scarp the relief of which was at least 600 m, based on the evidence at the north end of Butler Island (HAWLEY,1957). Ordovician marine limestone-pebble diamictites occur in several scattered localities in the upper Hudson Valley, New York. The rock at Bald Mountain, New York contains rounded and subangular pebbles 60-90 cm in diameter of carbonate rocks from Lower Cambrian to Middle Ordovician formations floating in a black mud matrix. It has been assigned to the Rysedorph Hill Conglomerate on faunal grounds (RUEDEMANN, 1914,1933), even though the matrix of the conglomerate at Rysedorph Hill and other nearby localities is calcareous (RUEDEMANN, 1901, 1930). Some of the limestone-pebble diamictites located near thrusts have not been assigned to the Rysedorph Hill unit, even though they resemble it closely, because they have been interpreted as friction breccias (RUEDEMANN, 1914, 1930, 1933). From this short discussion it should be obvious that limestone conglomerates of many types, ages, and origins are present in the northern Appalachians. Each should undergo careful study in light of newer stratigraphic evidence and recent ideas on sedimentology. The origin of limestone-boulder marine diamictites has been much debated. They have been attributed to glacial action or the results of ice-rafting (particularly in the Ouachitas, U.S.A., by TAFF, 1910; WOODWORTH, 1912; ULRICH,1927; S. POWERS,1928: and others; see summaries by KRAMER, 1933; MISER,1934; R. C. MOORE,1934; and CLINE,1960; and by DAWSON,1888, and COLEMAN, 1926, for the conglomerates near Quebec City); and interpreted as thrust conglomerates (RUEDEMANN, 1914, for the Ordovician of the Hudson Valley, New York; SCHUCHERT and DUNBAR, 1934, for the Cow Head breccia of Newfoundland; and VAN WATERSCHOOT VAN DER GRACHT,1931, for the Johns Valley Shale of the Ouachitas). The glacial theories have lost ground in view of the evidence for warm, tropical seas provided by the limestones, but the thrust-conglomerate interpretation gained many adherents, possibly because of the close spatial association of the deposits in question with particular thrust traces. The major and fatal deficiency of this idea, at least as applied to the Rysedorph Hill Conglomerate of New York, however, was pointed out by KAY(1937): the advancing thrust sheet, which was supposed to have shed the exotic blocks ahead of itself, does not include the kinds of carbonate rocks found in the pebbles; in fact, it scarcely includes any carbonate rocks at all. (See previous discussion of Overton Fanglomerate for a description of a true thrust conglomerate.) The presently accepted idea is that the blocks were displaced laterally by gravity down submarine slopes, as suggested by BAILEYet al. (1928) and OSBORNE (1956) for the Quebec conglomerates; by SCHUCHERT and DUNBAR (1943). DUNBAR and RODGERS (1957), and KINDLEand WHITTINGTON (1958) for the Newfoundland

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conglomerates; by DIXON (1931), KRAMER (1933), MISER(1934), R. C . MOORE (1934), and CLINE(1960) for the Johns Valley Shale; and by RENZet al. (1955) for the Paleocene-Eocene of Venezuela. The fault scarps with which these limestone-boulder diamictites presumably were related probably were “normal” faults. The association of the deposits with thrust faults is thought to be fortuitous and due to later movements, during which the diamictons were passively transported back upslope toward the shallow-water area from which the carbonate rocks were derived. In fact, they may have been transported considerably beyond the steep slope where they originated. This tectonic history closely resembles that of the Upper Jurassic reef breccias of the Nappe de Morcles and autochthonous chains, Haute Savoie, France, described previously. If the submarine sliding and fault-scarp ideas are correct, then such limestoneboulder diamictites are extremely significant for paleogeographic reconstructions: they mark the former bathymetric boundary between the shallow-water banks on which the limestones formed and the deep-water troughs in which the terrigenous sediments accumulated. This boundary may in fact represent ancient steep continental slopes. Analogous Recent marine deposits have not yet been reported; they may exist, but would be extremely difficult to sample by means of existing apparatus.

CONCLUSION

This chapter summarizes the classic literature on the origin and occurrence of limestones. The chief emphasis has been placed on limestones as products of the lithification of calcium carbonate sediments, and upon the environmental approach to these sediments as a key to understanding both the distinctive features of limestones themselves and also their manifold associations with non-carbonate sediments.

ACKNOWLEDGEMENTS

The authors express their thanks to L. V. Rickard for assistance in locating several references, and to George V. Chilingar and Rhodes W. Fairbridge (Editors) for helpful suggestions.


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TAFT,W. H., 1967b. Physical chemistry of formation of carbonates. In: G. V. CHILINGAR, H. J. BISSELLand R. W. FAIRBRIDGE (Editors), Carbonate Rocks. Elsevier, Amsterdam, B: 151-167. TAYLOR, R. E., 1938. Origin of the cap rock of Louisiana salt domes. Louisiana, Dept. Conserv. Geol. Bull., 1 1 : 191 pp. TEICHERT, C., 1946. Late Quaternary changes of sea level at Rottnest Island, Western Australia. Puoc. Roy. SOC.Victoria, 59(2): 63-79. TEICHERT, C., 1947a. Contributions to the geology of Houtman’s Abrolhos, Western Australia. Proc. Linnean SOC.N.S. Wales, 71(34): 145-196. TEICHERT, C., 1947b. Stratigraphy of Western Australia. J . Proc. Roy. Soc. N.S. Wales, 80: 81-142. TEICHERT, C., 1958. Cold- and deep-water coral banks. Bull. Am. Assoc. Petrol. Geologists, 42: 1064-1082. TENHAAF,E., 1959. Graded Beds ofthe Northern Appennines. Thesis, State University, Groningen, 102 pp. TEXTORIS, D. A. and CAROZZI,A. V., 1964. Petrography and evolution of Niagaran (Silurian) reefs, Indiana. Bull. Am. Assoc. Petrol. Geologists, 48: 397426. TIDDEMAN, R. H., 1889. On the concurrent faulting and deposit in Carboniferous times in Craven Yorkshire, with a note on Carboniferous reefs. British Assoc. Advan. Sci., NewcastleeonTyne Meeting, Rept., pp. 600-603. TRACEY JR., J. I., LADD,H. S. and HOFFMEISTER, J. E., 1948. Reefs of Bikini, Marshall Islands. Bull. Geol. SOC.Am., 59: 861-878. TROELL, A. R., 1962. Lower Mississippian bioherms of southwestern Missouri and northwestern Arkansas. J. Sediment. Petrol., 32: 629-644. TWENHOFEL, W. H., 1950. Coral and other organic reefs in geologic column. Bull. Am. Assoc. Petrol. Geologists, 34: 182-202. ULRICH,E. O., 1927. Fossiliferous boulders in the “Ouachita Caney” shale and the age of the shale containing them. Oklalzoma Geol. Surv., Bull., 45: 48 pp. J. H. F., 1947. Coral reefs of the East Indies. Bull. Geol. Soc. Am., 58: 729-778. UMBGROVE, VAN STRAATEN, L. M. J. U., 1950. Environment of formation and facies of the Wadden Sea sediments. Tijdschr. Koninkl. Ned. Aardrijkskirndig Genoot., 67: 354-368. VANSTRAATEN, L. M. J. U., 1951. Texture and genesis of Dutch Wadden Sea sediments. Proc. Intern. Congr. Sedimentol., 3rd, Groningen- Wageningen, 1951,pp.225-244. VANSTRAATEN, L. M. J. U., 1952. Biogene textures and the formation of shell beds in the Dutch Wadden Sea. Koninkl. Ned. Akad. Wetenschap., Proc. Ser. B, 55: 500-516. VANSTRAATEN, L. M. J. U., 1954a. Sedimentology of Recent tidal flat sediments and the psammites du Condroz (Devonian). Geol. Mijnbouw, 16: 2 5 4 7 . VANSTRAATEN, L. M. J. U., 1954b. Composition and structure of Recent marine sediments in The Netherlands. Leidse Geol. Mededel., 19: 1-110. VANWATERSCHOOT VAN DER GRACHT, W. A. J. M., 1931. The pre-Carboniferous exotic boulders in the so-called “Caney Shale” in the northwestern front of the Ouachita Mountains of Oklahoma. J, Geol., 39: 697-714. VERSEY, H. R., 1962. Older Tertiary limestones. In: Synopsis ofthe Geology of Jamaica-Jamaica, Geol. Surv., Dept., Bull., 4: 26-43. VOIGT,E. und HANTSCHEL, W., 1964. Gradierte Schichtung in der Oberkreide Westfalens. Fortschr. Geol. Rheinland Westfalen, 7: 495-548. VONGUMBEL, C. W., 1868. Geognotische Beschreibung des ostbayerischen Grenzgebirge. Gotha, 968 pp. WALCOTT, C. D., 1894. Paleozoic intraformational conglomerates. Bull. Geol. SOC.Am., 5: 191. WALCOTT, C. D., 1896. The Cambrian rocks of Pennsylvania. U.S., Geol. Surv., Bull., 134: 43 p p . WALKER, T. R., 1960. Carbonate replacement of detrital crystalline silicate minerals as a source of authigenic silica in sedimentary rocks. Bull. Geol. SOC.Am., 71: 145-152. WALPOLE, R. L. and CAROZZI, A. V., 1961. Microfacies study of the Rundle Group (Mississippian) of the Front Ranges, central Alberta, Canada. Bull. Am. Assoc. Petrol. Geologists, 45: 1810-1846. WALTHER, J., 1885. Die Gesteinbildenden Kalkalgen des Golfes von Neapel und die Entstehung structurloser Kalke. 2. Deut. Geol. Ges., 37: 329-357.


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Chapter 6

ORIGIN A N D OCCURRENCE OF DOLOSTONES GERALD M. FRIEDMAN AND JOHN E. SANDERS

Department ~fGeology, Rensselaer Polytechnic Institute, Troy, N , Y. ( U.S.A.) 33 Sherman Avenue, Dobbs Ferry, N. Y. (U.S.A.) SUMMARY

This chapter begins with a historical review of the dolomite literature and traces progressive advances in the search for a solution of the “dolomite problem”. The “dolomite question” defied solution until the recent advent of the X-ray revolution, when dolomite identification became simple, fast, and efficient. The general framework of the “question” now appears t o have been largely resolved as a result of intensive field studies of Quaternary carbonate sediments and the application of X-ray and stable isotope (carbon and oxygen) analysis to both Quaternary carbonate sediments and ancient limestones and dolostones. Many problems, however, are still outstanding, but they are on a different level. In the opinion of the authors, most dolostone deposits in the geologic record owe their origin to hypersaline brines. They must, therefore, be considered as evaporite deposits. Exceptions are uncommon and include those that are related t o soil-forming and bacterial processes or have resulted from recycling of pre-existinp dolomite. Hypersaline brines are formed by “capillary concentration” or by “refluxion” in the depositional environment in areas where evaporation exceeds precipitation plus run-off and by as yet unexplained processes in the subsurface. In “capillary concentration”, interstitial waters in the sediments transpire upward through porous sea-marginal sediments and evaporate at the sediment-air interface, a process similar to that which is responsible for caliche formation. This process leads to dolomite formation in supratidal and intertidal environments on broad shallow shelves with interfingering gypsum and/or anhydrite (landward) and marine carbonates (seaward). Under conditions of humid climate, however, anhydrite and gypsum may not develop. Supratidal to intertidal dolostones display many or most of the following structures: mud cracks, “birdseye” structures, burrow mottles, boundinage-like structures, scour-and-fill structures, channels, ripple marks, cross-stratification, and algal structures; these dolostones are commonly unfossiliferous and laminated and are texturally structLI-I”;SS or pellet Contribution No.66-6 of the Department of Geology, Rensselaer Polytechnic Institute. Work on this chapter commenced in the Department of Geology, Yale University, but was completed after the junior author became Senior Research Associate, Columbia University.

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muds. In “refluxion”, evaporation increases the concentration and density of the water in a restricted lagoon or intermontane basin. This produces a brine that sinks and migrates to the lowest possible topographic depressions where it may seep slowly through the underlying sediments, which are progressively dolomitized. Alternately in refluxion, dolomite may form directly from the brine with aragonite as a possible transitional phase, and with the deposition of layered dolomite mud at the bottom of the basin. The Mg/Ca ratio of the brine must be increased from that of sea water to a ratio larger than that which should be in equilibrium with both calcite and dolomite for dolomite to form. This ratio can be raised by the removal of calcium from the brine to form aragonite or gypsum or both; however, gypsum is only deposited in very shallow water where a supply of oxygen is ample, because it tends to be degraded to HzS and iron sulfide by bacteria where the oxygen supply is less. Hence layered dolomite formed under these conditions is indicative of a deeper water environment, of unspecified depth, with gypsum deposited in correlative positions along shore. Precaution is advised, therefore, in interpreting all syngenetic dolomite of supratidal origin. Both “capillary concentration” and “refluxion” may operate in the same basin; thus, at Salt Flat Graben, West Texas, an intermontane basin, the former is responsible for dolomite formation along the margin of the basin and the latter, in its center. Dolomite formed after burial, which includes most diagenetic and all epigenetic (fault and fracture-related) dolomite, owes its origin to subsurface brines. The origin of these brines is unknown, but their salinity is very close to that under which dolomite is formed in the depositional environment. Subsurface waters are normally depleted with respect to both magnesium and sulfate ions which were withdrawn from the brines during dolomitization. Dolostones occur in a variety of stratigraphic associations which can be summarized in the following main types: (1) Syngenetic dolomite, a dolomite which is formed penecontemporaneously in its environment of deposition as a micrite or as fine-grained crystals, and which may also: (a) interfinger with both marine and nonmarine evaporites, with or without associated terrigenous sediments; (b) interfinger with limestones, both marine and nonmarine, with or without terrigenous sediments; (c) interfinger with terrigenous sediments; (d) be in the form of dolomite crystals disseminated in terrigenous sediments; (e)be formed by biological agents, such as bacteria; and ( f ) occur in nonmarine environments as a weathering product in soils and caliche. (2) Detrital dolostone and detrital dolomite is formed by recycling of preexisting dolomite sediments. (3) Diagenetic dolostone is formed by replacement of limestone following consolidation of the sediment or coincident with the consolidation; this type of dolostone may also form by penecontemporaneous replacement, Diagenetic dolostone may form within individual beds or along surfaces of stratigraphic discontinuities.


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( 4 ) Epigenetic dolostone is formed by replacement of limestones with the replacement localized by post-depositional structural elements, such as faults and fractures. Most dolostones are formed by repIacement of pre-existing calcium carbonate sediments, but dolostones may revert back to calcium carbonate rocks by the process of dedolomitization.

INTRODUCTION

The origin and occurrence of dolostones present many of the same problems which were considered in the previous chapter on the origin and occurrence of limestones. The subject of dolostones, however, is more complex than that of limestones, because in addition to almost all of the intricate problems of limestones, it also embraces a large array of additional complications that are associated with the process of replacement of calcium carbonate sediments by the mineral dolomite. Dolostones are abundant in the geologic record; they are most commonly associated with limestones, but also occur with evaporites and terrigenous sediments, both with and without limestones. Volumetrically, Paleozoic dolostones far outrank those from other parts of the geologic column, but dolostones also formed during Precambrian, Mesozoic, and Cenozoic times. Dolostones are important economically, both as sources of magnesium and because many of them are vast petroleum reservoir rocks. The development of porosity in these rocks is closely related to the sequence of diagenetic changes, which also include dolomitization of carbonate sediments, so that these subjects have come under close scrutiny by the research laboratories of the petroleum companies. The “dolomite question” or “dolomite problem” long defied solution by geologists, largely because it could not be resolved by either of the two favored approaches: the experimental, and the comparative, based on conditions under which the material occurs in Recent sediments. No attempts at laboratory synthesis of the mineral dolomite were successful under conditions which could be considered reasonably comparable with those which prevail in nature, and no dolomite had been found among Recent sediments. In fact, most discussions of the subject began with remark that dolomite does not exist in Recent sediments. Most geologists were content with the overwhelming evidence that most dolostones had formed as replacements of calcium carbonate materials, and many of them supposed that, having established this conclusion, nothing further needed to be explained. The level of our knowledge of dolomite occurrence reached a plateau, and most discussions of the subject did little more than add a few new descriptive details and rehash older conclusions. All of this was abruptly changed by the development of X-ray diffraction

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techniques. This development brought about revolutionary changes in understanding of many fields of geology, but nowhere was more dramatic progress made than in the field of dolomite mineralogy and origin. Diagnostic crystal structures of the different carbonate minerals were revealed by X-ray studies; these indicated that mere chemical tests for magnesium, a common basis for the classification of limestones and dolostones, were insufficient to prove the existence of the mineral dolomite. High-magnesian calcite was discovered, in which the magnesium carbonate content was as much as 30%. When Recent carbonate sediments were examined by X-ray diffraction, dolomite was soon discovered. It typically occurs as such fine grains that it would not have been detected by staining techniques. Discovery of dolomite in Recent carbonate sediments opens the way to a full understanding of its origin and occurrence in the geologic record. The importance of discoveries of Recent dolomite is so great that this subject will be discussed as a separate heading under the origin section of this chapter. This treatment contrasts with the attempt of the writers to follow the parallel arrangement used in the foregoing limestone chapter, where examples of Recent calcium carbonate sediments are discussed in the occurrence section alongside examples from the geologic record. This chapter is organized as follows: (1) review of the literature; (2) origin and destruction of dolomite and dolostones, including subsections on (u) lithification of doIomite grains and crystals, (b) Recent and Pleistocene dolomite, (c) replacement of calcium carbonate by dolomite, (d) dedolomitization, (e) description of microscopic textures and fabrics of dolostones, and ( f ) experimental synthesis of dolomite; and (3) occurrence of dolostone, including many types based on mode of origin and associated sediments.

REVIEW OF LITERATURE ON DOLOMITE AND DOLOSTONES

The literature on dolomite and dolostones is enormous. In addition to many papers which describe dolostones as stratigraphic units, a great many others have attempted to solve the “dolomite problem” or “dolomite question”, as it has been widely known. Many of the statements made in these earlier attempts to resolve the enigma of the origin of dolomite and dolostones are no longer applicable as a result of recent discoveries. Nevertheless, it is profitable to re-examine this older literature in light of modern developments with the aim of showing where the earlier ideas were deficient and where they approached presently established concepts. The literature may be subdivided into two periods: ( I ) papers written before X-ray diffraction, and (2) papers written after the X-ray diffraction “revol~ition”. Most of the modern breakthroughs in the dolomite question have resulted directly from studies that have used the X-ray diffraction technique.


27 1

No attempt is made here to review the entire literature concerned with dolomite nor even to cite all of the papers that have reviewed or tried to resolve the dolomite problem; instead, the literature survey is limited to a selection of basic papers. The bibliographies that are included in these basic papers, however, will guide the reader to the world literature on this topic. Any review of the literature of dolomite and dolostone rightfully begins with the monograph of VANTUYL(1916). In this classic paper Van Tuyl assembled all the available literature and examined all possible aspects of the subject. An important early petrographic study that emphasized the replacement origin of most dolostones was made by STEIDTMANN (1917). Shorter summaries of the status of the dolomite question appeared in the Treatise on Sedimentation (VANTWL and STEIDTMANN, 1926; STEIDTMANN, 1926). A particularly good account, emphasizing examples from Great Britain, appeared in the book of HATCHet al. (1938). An excellent summary of the dolomite problem was presented by FAIRBRIDGE (1957) who summarized an extensive literature through 1956. Earlier reviews on this topic are found in papers by LINCK(1909, 1937), and for an almost complete index of the early literature, reference can be made to the article by CLEE(1950). Although the mechanisms of dolomite formation were unknown, the geologic evidence for penecontemporaneous replacement of calcium carbonate sediments by dolomite was sufficiently strong to establish this viewpoint as the leading hypothesis. Primary formation in the original depositional environment was cautiously admitted as a possibility, but most geologists did not accept it as being widely applicable. A major departure from this “standard” American position occurred as a result of the extremely detailed petrographic study of the Triassic dolostones of the Austrian Alps carried out by SANDER (1936, 1951). Sander attempted to classify the carbonate minerals according to their position in the fabric history of the rock. He differentiated between external growth on the upper free surface of the sediment, and internaI growth on the bounding surface of cavities within the sediment mass. Three types of mechanisms of deposition were pointed out by Sander: ( I ) mechanical deposits, including terrigenous detritus (recycled carbonates), crystals, ooids, and so forth, that are moved by currents and finally deposited by them; (2) chemical precipitates, on a free surface, such as sparry crystals growing normal to the surface, bladed crystals or colloform growths (radial fibrous aggregates with rounded, lustrous upper surface), and gels (?) or precipitates not on a free surface, but within the fabric of the enclosing material; and (3)biogenic deposits, such as cheeselike, cavernous crusts without organic patterns, or organic incrustations formed by encrusting organisms. Sander restricted the term biogenic to these organic crusts; he did not use it to refer to mechanical deposits consisting of organic skeletal remains. Sander concluded that dolomite can be ( I ) biogenically deposited; ( 2 )

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mechanically deposited; or (3) chemically deposited. Spathization refers to the crystallization of sparry calcite or dolomite in open or closed fabrics, whether on a free surface, internally on a closed surface, or as authigenic crystals (holoblasts of Sander). Sander emphasized the importance of primary dolomite. He treated dolomite as primary if its position in the fabric of the rock was not previously occupied by calcite. Some of his examples, such as tests of Foraminifera, began as calcite, but were afterward replaced by dolomite. Such a replaced test was assigned to a primary category if it became reworked and transported to a different site of deposition. Other primary dolomite includes sparry crystals, which only rarely grow on the upper free surface of the sediment, but much more commonly grow into internal cavities. In the Triassic dolostones examined by Sander, calcitization of dolomite was inferred to be a more important process than dolomitization of calcite. VON MORLOT in 1847 coined the term “dedolomitization” for the calcitization of dolomite. Later TEALL (1903) used this term to describe the metamorphic destruction of dolomite to form calcite and periclase, but today dedolomitization is known as a sedimentary recrystallization process in which calcite has formed at the expense of dolomite. CLOUDand BARNES(1948) ranged widely over the dolomite problem in connection with their extensive study of the Cambro-Ordovician Ellenburger Group in central Texas. They re-emphasized many of the “standard” points of view of the time: ( I ) that magnesium of bedded dolostones came from the sea and hydrothermal or ground-water types are very minor; (2) that primary precipitation of dolomite is unknown in Recent sediments and, therefore, is only an unsupported theory applied to the geologic record; (3) that most bedded dolostones are penecontemporaneous replacements of calcium carbonate sediments, which took place on the sea floor, or are due to primary deposition of dolomite; and because the former far exceed the latter and no means exist for distinguishing between these two, there is no reason to assume a primary origin when the penecontemporaneous replacement explanation is so satisfactory; ( 4 ) that the grain size and fabric of the coarser dolostones control porosity; (5) that dolostones occur closer to shore than contemporaneous limestones and grade into them in an open-sea direction (but no explanation for this arrangement was offered by them); and (6) that direct precipitation of dolomite may occur in evaporite basins, but this clearly is a special type of sedimentation. Cloud and Barnes remarked that no clastic dolostones occur in the Ellenburger rocks from the Llano area. In admitting the possibility that dolomite might be a primary precipitate from sea water (CLOUD and BARNES, 1948, p.94), they suggested that a current might precipitate calcium carbonate at one locality, and, being depleted in calcium, might then precipitate dolomite nearby, while calcium carbonate continued to be precipitated in the first locality. Finally, in connection with dolomite and porosity, they observed that no porosity was created in the Ellenburger rocks by growth of dolomite rhombs, but was largely


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due to subsequent dissolution of calcite inclusions between dolomite grains; they (1930) who believed that later leaching caused porosity also cited A. N. MURRAY in dolostones. Though he admitted the importance of dolostone formed by the replacement of calcium carbonate sediments, COOPER (1956) argued strongly for the viewpoint that primary dolostone is more widespread than most geologists would admit. His evidence was based on chemical, stratigraphic, and microscopic study of the Cambro-Ordovician rocks in the central Appalachians of Virginia, U.S.A. The Knox Dolostone of the northwestern outcrop belts, especially at Cumberland Gap, Virginia, lacks limestone, but includes beds of chert and numerous interbeds of quartz sandstone. Southeastward the dolostone grades into limestone, and the quantity of silica and A1203 impurities decreases. The SiOz content of the dolostones ranges from 2.5-12 %, whereas that of the limestones, from 0.5-2.5 %. The intermediate outcrop belts show the transition very well, particularly the Chepultepec Formation in the vicinity of Blacksburg, Virginia (middle of the CambroOrdovician sequence). Cooper inferred that dolomitization, both primary and replacement types, occurred virtually at the sea water-sediment interface in many cases. He remarked that if a calcareous mud became sufficiently stiff so that it ruptured into discrete fragments, then it became virtually impregnable to dolomitization thereafter. Because all of the limestone beds are fine-grained and low in silica, Cooper argued that the sandy dolostones of the northwestern outcrop belts did not originate by replacement of calcareous sediments. He concluded: “We must cease to think of dolomitization as demanding impossible conditions no longer met with in shallow-water environments.. .I think there are demonstrated reasons for believing that essentially primary deposition of dolomite is a possible, indeed a likely, process. . .” (COOPER,1956, p.7). DUNBAR and RODGERS (1957) briefly reviewed the dolomite question. They subdivided dolostones into three groups: (I) stratigraphic types, called S-dolostones; (2) dolostones related to fractures, called T-dolostones; and ( 3 ) dolostones formed by ground-water leaching, called W-dolostones. They accepted the idea that primary dolostones can form under evaporitic conditions, where dolomite would be precipitated directly out of sea water along with gypsum, anhydrite, and halite; but they also emphasized the importance and characteristics of penecontemporaneous replacement. They concluded that most S-dolostone originated as a replacement of lime mud and lime sand that is distributed by mechanical processes in warm shallow seas; dolomitization of these sediments may take place if the pH exceeds 9 in the overlying water. In addition, the chemistry of diagenetic fluids may also produce conditions for dolomite to replace calcite or aragonite. The significance of intertidal magnesium enrichment was first emphasized by REULING(1934) and applied by him as a possible explanation for the distribution of dolomite found in the Funafuti boring. FAIRBRIDGE (1957) further emphasized the significance of this process and discussed other examples including

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magnesium enrichment1 in the cement of beach rocks in warm latitudes, citing examples from GARDINER (1898), BRANNER(1904), and THORP (1935). Fairbridge also cited examples in which increased magnesium content coincides with firmly cemented layers in calcareous deposits, as in the Great Barrier Reef previously described by him (FAIRBRIDGE, 1950). Few workers other than Fairbridge appreciated the significance of this process, and even he was not fully aware of its true importance, which was only brought about by subsequent discoveries of dolomite in Recent sediments (see later section of this chapter). TEODOROVICH (1959) summarized the extensive Russian literature on sedimentary dolostone and attempted to relate dolostone occurrences to changing conditions in the earth’s atmosphere with geologic time. He showed that dolomite may form under a variety of conditions, but that the dominant process is replacement; he concluded that the main factors involved in the formation of dolomite are: partial pressure of COz, relative amounts of dissolved salts in water (cations and anions), salinity, temperature of water, and pH. According to Teodorovich, predominantly primary dolomite was precipitated in the Precambrian and Early Paleozoic seas. Primary dolomite did not form in the Late Paleozoic seas; instead, dolomite formed there only by diagenetic replacement of calcium carbonate sediments which were deposited in these seas. Late Paleozoic primary dolomite originated only in salt-water lagoons and large bays of low salinity. The volume of dolostone among the sedimentary carbonate rocks decreased sharply in the Mesozoic and Cenozoic eras and the replacement type became dominant. Teodorovich attempted to explain these conclusions on the basis of systematic changes in the partial pressure of COa in the earth’s atmosphere with time. He inferred that the partial pressure of COe was higher in the Precambrian and Early Paleozoic and that this condition was responsible for precipitation of primary dolomite in the sea. As the partial pressure of COZ became reduced with time, so goes his argument, primary dolomite declined and the replacement type became dominant. Teodorovich also engaged in a debate with STRAKHOV (1956, 1958) on the relative merits of the “basic facts” which Strakhov considered important in the dolomite question. Teodorovich‘s paper may be read with interest; it reflects the thinking of a geologist with 30 years of experience in examining the stratigraphic records of the large carbonate terrains in the Soviet Union and in analyzing the dolomite problem. In the opinion of the writers, the discovery of Recent dolomite, however, casts doubts on his conclusions about the systematic changes of the partial pressure of COZ in the earth‘s atmosphere. GARREAU et al. (1959) summarized French contributions to the dolomite problem under three headings: (I) the petroleum reservoir of Parentis; (2) tectonic

CHILINGAR (1956b) showed the importance of high Mg/Ca ratio of depositional medium for dolomite formation,


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dolomitization in the south of France; and ( 3 ) experimental research and the role of physicochemical factors in the synthesis of dolomite. In part 1 they showed that dolomite in the Parentis reservoir is both stratigraphically and tectonically controlled; the tectonic type is related to the major fracture system. In part 2 they examined the relationship between fractures and dolomitization in Jurassic limestones in the south of France. This interesting treatment recalls the work of HEWETT (1921, 1928) and G. M. Friedman (unpublished) in the United States (see section in this chapter on epigenetic dolostone). In part 3 they described experimental studies in which dolomite was synthesized by using solutions of calcium chloride and magnesium chloride and adding C O i - ions. They also explained in detail the role of temperature, partial pressure of COZ,and other factors in the synthesis of dolomite. INGERSON (1962, pp.830-837) concluded that the problems of dolomite formation centered chiefly around the questions of primary vs. secondary dolomite: the conditions under which primary dolomite is precipitated, the possible sources of solutions that caused formation of secondary (replacement) dolomite, the conditions of replacement, and the time at which replacement took place. He reviewed the literature on primary dolomite and noted the association of dolomite-bearing sediments with shallow water, plant growth, and elevated pH, but he suggested that direct precipitation may not necessarily have occurred. He also suggested that even dolomite in its original depositional environment may have been derived by replacement of calcite and/or aragonite. Ingerson discussed dolostone in geologic time and raised many questions why dolostones are much more prevalent in Precambrian and Paleozoic than in later rocks. Despite many literature citations and quotations of opinions voiced by previous authors no solution was in sight, according to Ingerson. His coverage of the topic was concluded with suggestions for further investigations taken from TWENHOFEL (1932), FAIRBRIDGE (1957), and CLOUD(1962), together with a plea for work on bacterial precipitation of dolomite and on laboratory experiments. Rapid progress in environmental interpretation of dolomite was closely linked with the X-ray revolution which made dolomite identification simple, fast, and efficient. CHAVE(1952, 1954) and SPOTTS(1952) observed a correlation between chemical analyses and X-ray diffraction patterns, which indicated beyond a doubt that the calcite structure can accommodate a large amount of magnesium in solid solution. GOLDSMITH et al. (1955) carried this correlation further and presented diagrams correlating X-ray spacing and magnesium substitution in calcite based on spectrochemical analysis and powder diffraction data. GRAFand GOLDSMITH (1956) noted that a perfectly ordered dolomite is formed experimentally within a reasonable time at somewhat elevated temperatures ( ~ 2 0 "C). 0 At lower temperatures the onIy reaction product the structure of which resembles that of dolomite is protodolomite. Protodolomites are dolomite-like materials which contain up to about 10 mole % excess of CaC03 (GOLDSMITH and GRAP,1958). GOLDSMITH and

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GRAF(1958) noted that compositional and structural features of dolomite of sedimentary rocks, the age of which ranges from Ordovician to Eocene, are similar to those of synthetic protodolomites. Such dolomites frequently contain about 5 mole % excess of CaC03 and their order reflections are weakened relative to those with a 1 /l CaCOS/MgCOs molar ratio. These structural and compositional features were not observed in metamorphic or hydrothermal dolomites, nor in most samples of evaporite sequences. The CaCOs-rich dolomites lose CaC03 and approach equilibrium conditions if held at high temperature in the presence of a flux. The determination of calcite/doloinite ratios by X-ray diffraction methods is now routine procedure (TENNANT and BERGER, 1957; WEBERand SMITH,1961, DIEBOLD et al., 1963; and GRAFand GOLDSMITH, 1963), as is the determination of ratios of calcite and high magnesium calcite (FRIEDMAN, 1964, p.778). The literature describing dolomite in Recent and Pleistocene sediments is included in the section on Origin and destruction of dolomite and dolostones.

ORIGIN AND DESTRUCTION OF DOLOMITE AND DOLOSTONES

The topic of the origin of dolomite and dolostones provides many comparisons and contrasts with that of the origin of limestones discussed in the previous chapter. Before continuing with the dolomite discussion, however, it is necessary to define some terms that appear on the ensuing pages. Particle is used as a general term to designate single grains or crystals. The term grain is used for individual particles, which (a) may be single crystals or crystal aggregates, (b) were once part of a larger object, such as a shell, or aggregate, such as a rock, and (c) have become individuals as a result of the breakdown of this larger object or aggregate. Crystal refers to individual particles with regular internal atomic arrangement, which have grown from a smaller nucleus, either in the environment of deposition or later within the host sediment. These definitions are intended only as first approximations; complications immediately arise. Many grains consist of single crystals; for example, plates of echinoderms. If the echinoderm test disaggregates, the individual plates become the grains of the present usage, even though they consist of single crystals. Also, small individual recycled grains from older dolostones may consist of single crystal particles, which are indistinguishable from first-cycle crystals that have grown in the depositional environment or diagenetically (AMSBURY, 1962). When doubt exists about the origin of the grain, or where both grains and crystals are intended, the term particle is used. The particles (grains or crystals) comprise the framework fabric of a sediment; these may be cemented into a rock by addition of a new mineral in the interstices or by overgrowths of the same mineral which forms the framework. Both framework particles and/or cement may be replaced by the growth of crystals.


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Some dolostones originated by the cementation of dolomite particles, but such dolostones are rare. By contrast, most limestones originated by cementation of calcium carbonate sediments. Most dolostones originated by replacement of calcium carbonate sediments or of limestones. Limestones of replacement origin are known, but they are relatively uncommon. This replacement may take place in the depositional environment; such replacement has been called penecontemporaneous or later. Whether primary dolomite can be precipitated as such or whether all dolomites form by replacement of aragonite or calcite in the depositional environment, or later, has not been established to the satisfaction of all workers. Isotope studies seem to indicate that all sedimentary dolomites, even those formed in the original depositional environment, are products of replacement of calcium carbonate sediments (DEGENS and EPSTEIN, 1964). Finally, no primary dolomite is known which compares with the calcium carbonate organic stony precipitates, such as reefs, or with the distinctive aragonite particles of some calcium carbonate sediments, such as ooids, pseudooids, and the various aggregate and/or coated particles. Although no such primary dolomite analogues of these calcium carbonate materials exist, such calcium carbonate materials may become dolomitized by replacement. In fact, recognition of vestiges of the distinctive fabrics of these calcium carbonate materials in dolostones provides the convincing evidence for the replacement origin of much of the dolostone in the geologic record. The following discussion on the origin and destruction of dolomite and dolostones includes: cementation of dolomite particles, including origin of dolomite particles, and origin of cement; Recent and PIeistocene dolomite; replacement of calcium carbonate sediments by dolomite; dedolomitization; microscopic crystal textures and fabrics of dolostones; and experimental synthesis of dolomite.

Cementation of dolomite particles Dolostone may originate by the cementation of dolomite particles. The dolomite particles may originate in several ways, and the cement may or may not consist of dolomite. Accordingly, the topics of particles and cements will be considered separately, beginning with the particles.

Origin of dolomite particles As with calcium carbonate particles, dolomite particles may be considered under the headings of first-cycle and recycled types. Another type of classification is into “primary” and “secondary”, but, as is explained later, this type of subdivision may be ambiguous and can be confusing, and consequently it is avoided here. First-cycle dolomiteparticles, First-cycle dolomite particles show much less vawty

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G . M. FRIEDMAN AND J. E. SANDERS

than do first-cycle calcium carbonate particles; they generally consist exclusively of euhedral rhombs of varying sizes. These rhombs may have been precipitated directly out of the sea water, either from the main water mass or from interstitial waters, or have formed in sea water at the expense of aragonite, or they may have grown authigenically in calcium carbonate sediments. Some dolomite crystals precipitate out of the marine or nonmarine waters of the depositional basin or are formed there by replacement of aragonite. They settle to the bottom, possibly being drifted by currents en route, or, once on bottom, may be reworked and finally deposited entirely under the influence of currents. Such aggregates of dolomite crystals show stratification features which form by currents, such as lamination or cross-lamination. Other dolomite crystallizes on the surface of tidal and/or supratidal flats and forms crusts of euhedral rhombs (see section on Recent and Pleistocene dolomite). The crystals comprising these crusts may become disaggregated and reworked by currents, or they may become converted directly into rock by in situ cementation without ever having been subjected to current action. The internal morphology of the crystals may reveal clues as to their origin. Dolomite crystals formed in sea water typically lack calcite inclusions. Crystals of authigenic origin may preserve palimpsest relicts of the type of sediment replaced; and recycled grains may include recognizable dolostone rock fragments and/or abraded single crystals. First-cycle dolomite crystals have been found in both skeletal sands and muds in the Pleistocene deep-sea carbonate sediments from the Bermuda Apron (FRIEDMAN, 1964, p.784). The dolomite occurs as flat, rhombic, transparent, sparry crystals, which resemble cleavage fragments. The diameter of most of these crystals ranges from 50-70 ,u; and they are very fragile and easily broken. The host carbonate sediments in which they occur were deposited by turbidity currents (see Chapter 5 on Origin and Occurrence of Limestones). Because the crystals in their present form are too fragile to have survived transportation to the environment in which they were found, from the nearshore environments where the host sediment was originally deposited, the most likely explanation is that they are of early diagenetic (authigenic) origin and grew in place. Dolomite crystals which originate by penecontemporaneous replacement of calcium carbonate sediments are termed “secondary” in most classifications, although, if removed to a new site of deposition, they become “primary” in the sense of SANDER (1936; 1951 transl.). Ordinarily, these crystals are not exhumed and separated from their original surroundings; however they may be separated if the carbonate host material itself has not yet become cemented and thus resistant to reworking. Once freed, however, these “secondary” dolomite particles behave just like their “primary” cousins. The designation “first-cycle’’ for all these particles, which originate within the basin itself, is a useful one in that it avoids the implications of “primary” and


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“secondary” where the evidence for such origin may be lacking. Furthermore, the distinction between primary and secondary, once thought to be the fundamental problem of dolomite formation, has assumed much less significance in light of the newer discoveries of Recent dolomite (see section of Recent and Pleistocene dolomite). Recycled dolomite grains. Previously deposited coarse-grained dolostones commonly break down into grains which consist of individual dolomite crystals, whereas fine-grained dolostones form dolostone rock fragments. The dolomite grains thus formed are exactly analogous to recycled limestone particles; their grain-size may vary from boulders to silt and they may come to rest in any environment of deposition. Fine-grained recycled dolomite rhombs may be indistinguishable from first-cycle dolomite rhombs. The Ellenburger (Ordovician) dolostone in the Llano area, Texas, U.S.A., contributes loose dolomite grains, fragments of microcrystalline dolostone, limestone rock fragments, and chert to modern river sands. Likewise, the Upper Glen Rose Formation (Lower Cretaceous) of south-central Texas contributes fine-grained dolostone rock fragments to the silt fraction of modern rivers in the area (AMSBURY, 1962). Residual sands consisting of dolomite grains also occur over Tertiary dolostone in the Paris Basin and over the Carboniferous and Permian dolostones in England (HATCHet al., 1938, p.185). Dolomite sands are also found on the shores of Mingan Island, Gulf of St. Lawrence, Canada, where an older dolostone formation is being eroded in the coastal bluffs (STEIDTMANN, 1926, p.264). The Gulf of Eilat (Aqaba), one of the northern arms of the Red Sea, is one of the most saline bodies of marine water in the world, hence a likely site for primary dolomite deposition. The Recent sediments from the floor of this gulf were examined with this idea in mind, and dolomite was found. The particles were separated, however, and proved not to be primary, but to consist of recycled grains which may have come from Cretaceous rocks (FRIEDMAN, 1965b). Origin of cement The cement which converts a mass of dolomite particles into a dolostone may or may not consist of dolomite. Where the cement consists of dolomite, it may have been precipitated in open spaces with a radial arrangement in which long axes of crystals project inward from the walls toward the center of the cavity; or it may have grown in optical continuity on the outer surfaces of a nucleus dolomite particle. Such dolomite cement is a “primary” precipitate, even though it may serve to bind together “secondary” dolomite particles. Where the cement does not consist of dolomite, it may be composed of aragonite, calcite, anhydrite, halite, or other minerals. The texture of these cements is

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controlled by patterns of crystal growth within the internal spaces of the sediment. More than one generation of cement may be present. The mechanism of upward transpiration and concentration by evaporation of sea water through the pore spaces of sea-marginal sediments provides a previously unrecognized basis for understanding the mineralogic variety of cements of dolostones (and also other sediments). In a reasonably porous mass of dolomite (or other) particles around the edge of a sea, where evaporation is taking place on marginal flats, a lateral change in cements may be found across the marginal flats which repeats the classical evaporite sequence: halite, gypsum (or anhydrite), dolomite, and aragonite. The halite will occur at the most distant point of upward transpiration of the sea water and the aragonite, on the side nearest the sea. In addition, any return flow of brine toward the sea may be extremely enriched in magnesium, so that it becomes an active agent of dolomitization in the sediments through which it passes (see the following section on Recent and Pleistocene dolomite). Needless to add, cements may form under numerous other circumstances, which are controlled by fluid movements through the sediments; the fluids may be ground water, sea water, or connate waters, and they may be driven by various forces. Recent and Pleistocene dolomite The discovery of dolomite in Recent sediments and study of the conditions under which it occurs has led to a breakthrough in the understanding of dolomite, which extensive study of ancient dolostones failed to accomplish. Because Recent dolomite is treated only briefly in the chapter on Recent carbonate sediments, no repetition is involved by including an expanded discussion of this subject as a separate section of this chapter, where it can serve as a basis for interpreting examples of ancient dolostones from the geologic record. Included here also are Pleistocene examples which fall within the range of the radiocarbon dating technique. Recent dolomite has been found in the sediments, an waters of hypersaline lakes, in both sediments and waters of hypersaline lagoons, and on tidal and supratidal flats. Both “primary” and “peneconteniporaneous replacement” types of dolomite have been found in the same environments, an association which tends to de-emphasize the importance formerly placed on this distinction. Whether the dolomite occurs as the primary framework of a deposit, as a cement of primary framework grains, or as a replacement of primary framework grains and/or cement is still an important topic to establish in textural analysis, but this distinction can no longer be considered as an end in itself. Penecontemporaneous replacement dolomite has been formed where capillary upward movement of sea water, initially of normal or near-normal salinity, flows through sediments on marginal supratidal flats and becomes concentrated


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by evapo-transpiration at the sedimerlt-air interface (“marine caliche” action of the preceding chapter, or “capillary concentration” of MUELLER, 1960).A contrasting process, namely, downward flow of dense brine may occur, either in the main water mass or in the interstitial fluids of sea- or lake-marginal sediments. If the molar Mg/Ca ratio of the brine has been raised much higher than the normal value in sea water (Sjl), owing to evaporation and precipitation elsewhere of calcium carbonate and/or calcium sulfate, then the high-Mg brine may cause dolomite to form. Previous discussion of various aspects of this mechanism includes the “reflux” process (R. H. KING, 1947), the “calcium-precipitation elsewhere” mechanism suggested by CLOUDand BARNES(1948), and the “seepage refluxion” process postulated by ADAMSand RHODES(1960). The discoveries of dolomite related to evaporite brines tend to substantiate the expectations of COOPER (1956) and provide the basis for understanding the process of intertidal magnesium enrichment emphasized by REULING(1934) and FAIRBRIDGE (1957). Recent dolomite has been found in sediments of saline lakes in the U.S.S.R., South Australia and U.S.A. Because of the possibilities of total evaporation of isolated lake-water bodies, most of the attention on dolomite found in saline lakes has been centered on the mechanism of direct precipitation from the water, a process which evidently has been observed in action in South Australia. The capillary concentration process, which was originally conceived as an explanation for nonmarine evaporitic nitrate deposits in Chile, and which has been subsequently demonstrated for dolomite and other “evaporitic” minerals on sea-marginal supratidal flats, may also operate around the edges of saline lakes. An example is the dolomite caliche formed around the margins of a saline lake in West Texas (FRIEDMAN, 1966); other examples may be expected from future investigations. Recent dolomite was first reported from shallow-water sediments of Alakul Bay, in the eastern part of Lake Balkhash, U.S.S.R. (TEODOROVICH, 1946). In the areas of dolomite formation, salinity, magnesium content, pH, and temperature of the water were higher than elsewhere. The pH of the waters ranged up to 9.2-9.4 in Alakul Bay, and varied from 8.9-9.1 in the eastern part of Lake Balkhash. Teodorovich concluded that both penecontemporaneous and replacement dolomites are formed because of the similarities of the solubilities of calcium carbonate and magnesium carbonate in salt water and because of an increase in magnesium content of the solutions with greater salinity and higher temperature. Discovery of Recent dolomite in South Australia, in shallow, isolated lakes and a finger of the sea known as the Coorong Lagoon, sent a spark of excitement through the geologic world (ALDERMAN and SKINNER, 1957; ALDERMAN and VON DER BORCH,1963; SKINNER, 1960, 1963; SKINNER et al., 1963; VONDER BORCH, 1965). The dolomite is formed where the water is very saline, the pH is high and plants are abundant. The plants extract C02 from water during photosynthesis which raises the p H and promotes dolomite sedimentation. Locally, magnesite originates along with dolomite. A mineralogical zonation with increasing age has

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M. FRIEDMAN

AND J. E. SANDERS

TABLE I A MINERALOGICAL ZONATION ACCORDING TO INCREASING AGE

MglCa ratio

Sediment

Younger

Older

+ + +

aragonite Mg calcite Mg calcite Mg calcite Ca dolomite “ordered” dolomite dolomite magnesite hydromagiiesite

5

6 8 10 16

20

been observed as shown in Table I (ALDERMAN and VON DER BORCH,1963). Magnesium content is increased, owing to the annual cycle of inundation of the evaporating basins and selective biogenic withdrawal of calcium carbonate. ALDERMAN and SKINNER (1957) noted that quite commonly the waters of the isolated lakes become turbid with fine, white sediment which slowly settles to the bottom. This suspended white sediment was identified as an extremely fine-grained mixture of high-magnesian calcite and calcian dolomite. As the carbonate sediments accumulate and evaporation of the waters continues, halite and sulfate minerals are precipitated. Even celestite, which never makes up more than 3 % of the total sediment, was observed to precipitate with the carbonates. Calcite and dolomite found in the sediments of the Coorong Lagoon are not the ideal stoichiometric compounds, CaC03 and CaMg(CO3)z. Calcite contains some magnesium in solid solution and dolomite as a rule contains excess calcium in solid solution. The composition of precipitated calcite ranges from Ca77Mg23 to Ca98Mg2 and that of dolomite from Ca50Mg50,apparently stoichiometric dolomite, to Ca56Mg44, a protodolomite. The ratio of calcite t o dolomite collected during active precipitation ranged from almost entirely calcite to 1/1 calcite to dolomite. In sediments found on the lake bottom ratios from l/O (100 % calcite) to 1/4 calcite to dolomite were noted. In one small lake, stoichiometric dolomite was found free of admixed calcite. Better crystallized and partially ordered dolomites are associated with lake waters of significantly higher pH and Mg/Ca ratios than poorly crystallized protodolomite-magnesian calcite assemblages in the area. Total salinities observed in water samples over a two-year period ranged from 1.6 to 18.9 %. At 27.4% salinity, halite began to precipitate and plants began to decay. Yet, plant life tolerated salinities up to more than 20 %. Water temperatures varied between 10.5 “C and 28 “Cand pH from 8.4-10.3. Based on one 14Cage determination, the average accumulation rate of dolomite in one ephemeral lake was about 0.2 mmlyear. By other means, the most realistic estimate of the rate of accumulation was a range of 0.2-0.5 mm/year. Recently, chen has been found to precipitate as a gelatinous opal-cristobalite


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in lakes associated with the Coorong Lagoon (PETERSON and VONDER BORCH,1965). High p H (9.5-10.2) causes dissolution of detrital silicates; lowering of pH (7.0-6.5) and drying of the lakes cause precipitation of chert. Recent dolomite has been discovered in a small playa known as Deep Spring Lake, which is located at the southern end of Deep Spring Valley, a relatively small intermontane basin in northern Inyo County, California, U.S.A. (JONES, 1961). The playa covers an area of about 13 km2; a perennial mass of dense brine occupies approximately two-thirds of this area. The lake water comes from the following sources: ( I ) rainfall into the lake area; (2) surface runoff (3) groundwater runoff from springs located along a prominent fault on the east side of the lake; (4) springs on the north side of the lake; and (5) seepage from below. The depth of water is greatest (about 30 cm) during winter and spring; only a small mass of brine remains during the hot dry summer and fall. The pH of the lake water is about 9.5-10.0 (PETERSON et al., 1963). The saline minerals in approximate order of abundance, and sediments in which they occur (taken from JONES,1961), are given in Table 11. Dolomite is the dominant mineral in the muds throughout most of the playa, except on the far western side, where calcite and/or aragonite predominate over dolomite. The mud is extremely fine-grained. The newly formed dolomite is calciumrich; structural and compositional variations exist within a single sample and within the same size fraction of a single sample. The absence of very small crystals at depths of 30-60 cm in the sediment suggests that recrystallization has occurred (PETERSON and BIEN,1963). The minerals are laterally zoned from playa margin to playa center, as follows: aragonite and/or calcite, dolomite, gaylussite, thenardite, TABLE I1 SALINE MINERALS IN ORDER OF ABUNDANCE IN SEDIMENTS IN WHICH THEY OCCUR

Mineral

Sediment

Dolomite, CaMg(CO3)z Calcite, aragonite, CaC03 Thenardite, NazS04 Halite, NaCl Gaylussite, NazCa(CO3)z . 5Hz0 Burkeite, N ~ ~ ( S O ~ ) Z C O ~ Aphthitalite, NazK(S04)~ Trona, NasH(C03)z. 2Hz0 Pirssonite, NazCa(CO3)z .2H20 Nahcolite, NaHC03 Thermonatrite, NazC03. HzO Glauberite, NazCa(S04)~ Sylvite, KC1 Bloedite, NazMg(SO&

Mud Mud Mud, saline crust, efflorescence Mud, saline crust, efflorescence Mud Mud, saline crust, efflorescence Saline crust, efflorescence Saline crust, efflorescence Efflorescence Mud Efflorescence Saline crust, efflorescence Saline crust, efflorescence Saline crust

284


and burkeite. Only aragonite and gaylussite do not extend from the outer limits to the center of the lake (JONES,1961). Radiocarbon age determinations indicate that this dolomite is virtually of Recent age. Assuming different models of crystal growth, average rates of growth of individual dolomite crystals were calculated to be of the order of hundreds of angstroms/1000 years (PETERSON and BIEN,1963). According to JONES(1961), the distribution suggests that dolomite has formed by precipitation directly from solution. The highest content of dolomite in the mud is found within areas of relatively frequent flooding. Recent dolomite has been found on supratidal flats along the southern shore of the Persian Gulf, in the Bahamas, and in the southern part of Florida. Doubtless, it will soon be found in many other low-latitude localities where evaporation is intense and calcium carbonate sediments predominate. Net evaporation takes place during all seasons of the year in the Persian Gulf, so that increased salinity is typical and values as high as 70 % above that of normal sea water have been recorded in one bay (SUGDEN,1963). Although the hydrology of the Persian Gulf teaches important lessons for the understanding of evaporites, no evaporite deposition occurs directly from the Gulf waters; instead, deposition takes place within the sediments of Gulf marginal areas. Recent dolomite has been found in large areas extending from the west side of the Qatar Peninsula 4 5"

50'

55O

60'

30'

.

25'

::

Oman

0

b B t I i 0KM

20

Fig.1. Trucial coast of the Persian Gulf showing location of Qatar Peninsula, Abu Dhabi, and Oman. Recent dolomite has been found in large areas between the west side of Qatar Peninsula and the Oman Peninsula at the entrance to the Gulf.


285

along the Trucial coast to the Oman Peninsula at the entrance to the Gulf, a shoreline distance of about 900 km (Fig.1). High-salinity water is common in lagoons along the shallow margins of the Gulf; still higher salinities have been found in interstitial waters of marginal sediments. Lagoons and embayments on the west side of the Qatar Peninsula are separated from the main mass of Gulf water by a shallow shelf that is many kilometers long and less than 4 m deep.The tidal range here is about 1 m (WELLS,1962; ILLING,1963; ILLING and WELLS,1964). Salinity of lagoon water ranged from 4.5-5.5 %; it measured 10.0 % just below ths level of diurnal high tide (WELLS,1962). Chlorinity was 30-35 g/l (ILLINGand WELLS,1964). Comparable coastal lagoons and tidal flats extend eastward along the Trucial coast for about 400 km towards Oman. Lagoon waters of this area have a pH of 8.0 (CURTISet al., 1963). In a landward direction the margins of the lagoons consist of an intertidal zone, an algal flat, and supratidal flats, known as sebkhas, the width of which may range up to 11 km. Interstitial waters in the sediments of the algal flats and sebkhas are hot, highly saline, and slightly acid. Measured values from different areas are summarized in Table 111. Minerals from the tidal flats, algal flats, and sebkhas include aragonite, calcite, dolomite, gypsum, anhydrite, and halite; their proportions and distribution vary in different areas. Aragonite predominates in the pelletal muds and sands of the lower tidal flats at Qatar (WELLS,1962); it becomes more abundant on the upper tidal flats. Dolomite appears on the uppermost parts of these tidal flats, between the levels of diurnal and spring high tides, and its abundance increases landward. Landward across the sebkhas, more evaporitic minerals appear with the dolomite, and, still farther landward, these supplant dolomite altogether. These other more evaporitic minerals include gypsum at Qatar (WELLS,1962); anhydrite and halite with no gypsum in one short core from the sebkha at Abu Dhabi (CURTISet al., 1963); and gypsum, anhydrite, and halite in other samples from Abu Dhabi (SHEARMAN, 1963). At Qatar, the consistency of dolomite-bearing sediment is that of stiff mud. In the upper 30-60 cm of sebkha sediments microcrystalline dolomite in rhombs 1-5p in diameter has replaced aragonite, and large (up to 12.7 cm in diameter) gypsum crystals occur. From depths of 60 cm to 1 m, aragonite muds and pelletal sediments lack dolomite and gypsum. Radiocarbon dating of two dolomite samples gave ages of 2,670 and 3,310 years B.P. (ILLING,1963). Dolomite, anhydrite, gypsum, and halite occur above the water table (which is located at a depth of 1.3 m below the sebkha surface) at Abu Dhabi. Anhydrite OCCUTS in bands, lenses, and as layers of nodules in sediments above high tide level; it has not been found in any of the lagoons nor in sediments from the intertidal zone. Some anhydrite bands cut across the foreset beds of the host sediment, indicating that these bands are of diagenetic origin (SHEARMAN, 1963). Gypsum of more than one generation also is abundant at Abu Dhabi; some of it occurs on the sebkha, whereas the rest lies beneath algal mats of the higher tidal flats. Many gypsum

TABLE I11 GEOCHEMISTRY OF WATERS AND MIMERALQGY OF SEDIMENTS OF GULF-MARGINAL AREAS, QATAR PENINSULA AND ABU DHABI, PERSIAN GULF

Locality

Water properties salinity ( %)

chlorinity (gll)

pH

temperature ("C)

Trucial coast Abu Dhabi sebkha

Abu Dhabi Lagoon high Trucial coast sebkhas

30-35 50-130 >I50

3 6.7 6.0 6.2

up to 8.0

>40 (summer) 35 average (Oct., 1962)

>40

(Jan., 1964) 49 (Apr., 1964)

Reference

80 % aragonite 95 % aragonite Dolomite and aragonite

WELLS(1962) WELLS(1962) WELLS(1962)

Dolomite and gypsum

WELLS(1962)

MgICa ratio

West side Qatar Peninsula lower tidal flats algal flats upper tidal flats upper tidal flats 27.5 further landward Sebkha, Qatar Lagoon water, Qatar Algal flat, Qatar Sebkha edge

Minerals in sediments

Core: dolomite overlying aragonite

10

Protodolomite (Ca54Mg46) Core: dolomite, anhydrite, aragonite, calcite and halite; no gypsum, in one, but abundant in others Dolomite, anhydrite, and gypsum

ILLING (1963) ILLING and WELLS(1964) ILLING and WELLS(1963) ILLING and WELLS(1961) CURTISet a]. (1963)

?

g

9

8z

P

2:

CURTISet al. (1963) BUTLER et al. (1964)

tt7i

7

m

5E

6


287

crystals poikilotopically (FRIEDMAN, 1965a) enclose aragonite and calcite crystals and quartz grains. Halite occurs as a cement between quartz grains (SHEARMAN, 1963). Dolomite, anhydrite, and gypsum also occur above the level of high tide in the sebkha sediments along the Trucial coast; as at Abu Dhabi, anhydrite is restricted to levels above the water table (BUTLERet al., 1964). Gypsum replaces anhydrite on the landward parts of the sebkhas, where the less saline ground water promotes hydration. Dolomite has also been found in the Recent carbonate sediments on the west side of Andros Island in the Bahamas. One occurrence is in the predominantly pelletal lime-mud sediments which underlie algal mats at the edge of a mangrove swamp in Wide Opening (FRIEDMAN, 1964, p.797, based on samples analyzed in 1960). At the time of sampling, water depth was 5-8 cm and water temperature (not measured) considerably exceeded that of the deeper water nearby. Tidal levels were not determined, so it is not known whether this occurrence of dolomite should be considered as “intertidal” (as it would be if the water cover was due to normal tidal fluctuations) or “supratidal” (if the water cover was due to a storm surge or other unusual event and was lying above normal high tides). Up to 150 cm of Recent dolomite has formed in the last 5,000 years in large areas on the west side of Andros Island. It comprises up to 80% of surface and near-surface supratidal pelletal muds, which lie a few centimeters above high tide, where algal mats are common and mud cracks and worm burrows are found. The area of dolomite occurrence covers hundreds of square kilometers. The percentage of dolomite increases in more lithified sediments; it also partially replaces gastropod shells and pellets. The dolomite occurs in a single “crust-likeyylayer the upper surface of which has the appearance of an old asphalt road coating. Its occurrence with mud cracks, stromatolites, burrow mottles, and boudinage-like structures parallel the assemblage of features found in many ancient dolostones (see section on occurrence of dolostones). As in the sebkhas of the Persian Gulf, capillary activity and evapo-transpiration of concentrated interstitial brines which began as sea water are thought to have been responsible for the origin of dolomite (SHINN and GINSBURG,1964; SHINNet al., 1965a, b). No anhydrite, gypsum, or halite, however, have been found in the Bahamas. Dolomite has also been found in Florida Bay (TAFT,1961), but a radiocarbon age from dolomite crystals isolated from the Recent sediments of Florida Bay proved to be greater than 35,000 years, suggesting that these crystals have been recycled from older dolostones (DEFFEYES and MARTIN,1962). Recent dolomite has been found in many localities in south Florida (SHINNand GINSBURG, 1964; SHINNet al., 1965a). Recent dolomite also occurs in carbonate sediment from Crane Key, Florida; mineralogic composition was found to be: high-magnesian calcite 47 %, dolomite 23 %, and aragonite 19 % (G. M. Friedman, unpublished). These sea-marginal Recent dolomites and evaporite minerals provide a

288


previously overlooked mechanism for understanding how salinity of sea water may be increased in the absence of “typical” evaporite concentrations in the open water mass. The interstitial water in the sediments begins as sea water of normal or near-normal salinity. Owing to the high temperature and excess of evaporation over rainfall, however, the interstitial waters transpire upward through the porous marginal sediments and evaporate at the sediment-air interface, thus increasing the concentration of salts in unevaporated water. The circulation is identical to that which forms caliche in soils. Dolomite may be the only “evaporite” mineral present, as in the Bahamas and Florida, or “more” evaporitic minerals such as anhydrite, gypsum, or halite, may occur with or near the dolomite, as on the sebkhas of the Persian Gulf. Such “caliche-like” dolomite requires only a porous marginal sediment framework, which need not necessarily consist of calcium carbonate particles, predominantly upward transpiration of sea water through the sediment, and continued access of the sea water to the porous marginal framework. The discovery of dolomite caliche in West Texas (FRIEDMAN, 1966) confirms that this process may also form dolomite in a nonmarine setting. The same principle, called “capillary concentration”, was earlier discussed by MUELLER(1 960) and applied by him as an explanation for the nitrate deposits of northern Chile. Mueller’s experiments showed how precipitates from solutions that moved upward as a result of capillary action may be zoned according to solubility. Although he was particularly concerned with nonmarine evaporites and the influence of slopes, the mechanism proposed by Mueller seems to have operated in sea-marginal settings of flat topography. Interesting results might be obtained by repeating Mueller’s experiments using sea water and lower slopes at the sediment-air interface. Different stratigraphic patterns of units will be produced, depending upon the relationships between submergence, local sediment supply, and rate of dolomite formation. During submergence with counterbalanced local sediment supply, and rate of dolomite formation, the marginal deposits simply thicken by additions at the top and their inner edge encroaches landward, forming a basal transgressive supratidal dolostone. If submergence becomes slightly excessive, the basal transgressive supratidal dolostone will be buried and overlain by non-dolomitic, nearshore, shallow-water marine carbonates. Such basal dolostones should be common where carbonate-depositing seas have submerged adjoining land areas; they may not be conglomeratic, as initial deposits of transgressing seas are supposed to be. The seaward side of an algal flat-supratidal flat complex, however, may expand laterally toward the sea, so that sediments from these environments overlie deposits from tidal flats and lagoons, as at Qatar (TLLINGand WELLS,1964). Here, a stratigraphic succession consists of basal lagoonal carbonate sediments, which are overlain by intertidal skeletal lime muds and lime sands. These, in turn, are covered by stromatolitic algal laminae, and the topmost layer consists of dolomite which


289

originated on the supratidal flats. If slow submergence occurs during such lateral seaward expansion of the marginal sediment zones, then the resulting stratigraphic units will overlap each other in a seaward direction, opposite to that of the transgressing landward side of the sebkha. Not only will the width of the sebkha increase under such circumstances, but the thickness of the distinctive deposits formed on it will thicken by additions at the top. Tiny crystals (about 2,u in size) of Recent dolomite (radiocarbon ages of 1480f 140 and 2190 & 150 years B.P.) have formed by replacement of calcium carbonate sediments on sea-marginal flats on the south end of the island of Bonaire, Netherlands Antilles. Dolomite occurs in most sediments here and comprises as much as 95% of some crusts (DEFFEYES et al., 1964). Although the question of origin of these crusts by capillary concentration of supratidal flats was not discussed, the descriptive data suggest that this process may have operated. Evaporation in isolated lakes and pools enclosed within a ridge of coral rubble has raised the salinity and caused precipitation of aragonite and gypsum; the Mg/Ca ratio of these isolated lakes has also been raised to 30/1 by depletion of calcium. Calculations of the water and material budget of the Pekelmeer, the largest of these lakes, suggest that dense brine escapes downward by seepage through the underlying sediment. The effects of this supposed downward seepage of Mg-rich brine were not directly investigated by Deffeyes et a]., so that this mechanism has not yet been shown to have produced Recent dolomite. Instead, their argument on behalf of this process is based on the: (u)association of Recent dolomite on the south side of the island (which possibly is of supratidal origin); (b) measurements and calculations of the brines in Recent pools on the south side of the island and inferences on downward seepage of the brine from these pools; and (c) application of the downward seepage concept to the older Pleistocene dolostone found along a stratigraphic discontinuity on the north side of the island. Such downward percolation of brines was envisaged as the “seepage refluxion” process by ADAMSand RHODES(1960). Although the belief in feasibility of this process has been greatly strengthened by the work on sediments off the island of Bonaire, the facts are that Recent dolomite formed by it has not yet been demonstrated. It is easy to predict that this demonstration will be made shortly. Adams and Rhodes postulated that evaporation in the overlying water mass produces dense brine, which flows outward along the bottom. If the return flow along the bottom (reflux) of this brine is prevented, then its only route of escape is downward interstitial flow through porous underlying carbonate sediments. The brine is inferred to have been an active agent of dolomitization. The distribution and origin of Pleistocene dolomite in Salt Flat Graben, which extends along the Hudspeth and Culberson County lines, West Texas, U.S.A., for a distance of about 80 km from the Apache Mountains in the south to beyond the Texas-New Mexico state line in the north, have been studied by FRIEDMAN (1966). The graben was formed in the Late Tertiary and has been filled

290


with Pleistocene and Recent sediments. Its average width is 17.6 km; the thickness of its fill is incompletely known. Wells in the graben center which are 404 m and 375 m deep do not fully penetrate the Pleistocene sediments; other wells, however, nearer to the margin of the graben, indicate that the thickness of the fill there ranges from 1.5 to 36 m. The surface sediments of this fill are composed of gypsum, halite, calcite, and dolomite, with subordinate amounts of aragonite and native sulfur. Cores of the upper meter of the fill taken along a transect across the graben reveal that dolomite occurs at stratigraphically persistent levels. One prominent traceable dolomite zone occurs near Highway 62 at a depth of 40.8-51 cm below the surface in the center of the basin; its depth below the surface increases toward the margins of the basin. Other dolomite layers are found at depths of about 61-71.2 cm and 86-91.5 cm. The color of the sediments in the upper 37 cm or so is brown, indicating oxidizing conditions. The color below this depth alternates between gray to black, indicating reducing conditions. The sediments are finely interlayered with laminae of tan and gray colors. These couplets are due to different mineral composition. The finely laminated units, with thicknesses ranging from 4-1 6 cm, consist predominantly of gypsum and halite, with locally abundant calcite, aragonite, or dolomite, and minor amounts of terrigenous quartz and feldspar, arranged in laminae with thicknesses varying from a millimeter to a few centimeters. The grayish-black interbeds, 4-5 cm thick, consist of dolomite and laterally discontinuous sulfur lenses with minor calcite. The distinctive black color of these layers with dolomite facilitates their use as stratigraphic markers. The black color, however, disappears within a month as a result of oxidation if left exposed to the atmosphere. A radiocarbon date of 5,840 f 400 years B.P. was determined for a sample of calcite and aragonite from the surface layer of the old lake-bottom sediments; and a date of 20,300 ,J, 825 years B.P., for a sample of pure dolomite from a depth of 89 cm in the core. The dolomite-bearing layers form hardpans; when these were pierced in coring or trenching, HzS escaped with a swishing sound. Although the black color suggested that the carbon content of these layers is higher than that of the other layers, measurements of organic carbon reveal a range of values from 0.1 1 to 0.37 % and that the dark layers do not contain more abundant carbon than the others. Accordingly, the dark color must be due to iron sulfide. This association of dolomite, H B S ,iron sulfide, and native sulfur suggests that decomposition of gypsum was involved in the history of these layers. A comparison is suggested with Dead Sea sediments in which gypsum has been transformed into calcite by bacterial reduction of the sulfate, with evolution of HBS,some of which reacted with iron in solution to form iron sulfide, the black pigment (NEEV, 1963; see also Chapter 5). The native sulfur of salt domes has been attributed to bacterial reduction of gypsum or anhydrite (BASTINet al., 1926; THODE et al., 1951; JONES et al., 1956).

29 1


TABLE IV ISOTOPIC COMPOSITION OF DOLOMITE FROM SALT FLAT GRABEN, WEST TEXAS,

(Analyses by Roger Ames, Pan American Petroleum Corp.) Specimen

Non-dolomitic layers or layers with subordinate dolomite Layers with abundant dolomite

6 13C

6

'80

Corrected to PDB

Corrected to PDB

-6.4 to -13.8 -4.1 to 0.0

-3.6 to -9.1 -2.0 to + 3 . 3

Regression analysis of the mineralogical data from Salt Flat Graben indicates that a positive relationship exists between increased dolomite and decreased gypsum, halite, and calcite contents. Compared with layers in which dolomite is less abundant or absent, the dolomite layers at Salt Flat Graben are enriched in heavier isotopes; the results are shown in Table IV. Other isotopic studies have shown that carbonate minerals in nonmarine environments where evaporation has not occurred tend to be enriched in the lighter isotopes of carbon ( I T ) and oxygen (160). The process of evaporation, however, tends to concentrate the heavier isotopes, I3C and l*O, because the lighter isotopes are preferentially removed in C02. The enrichment of the heavier isotopes found in the dolomite-rich layers, therefore, indicates intense evaporation. Similar isotopic studies on Dead Sea sediments (FRIEDMAN and NEEV,1966) indicated that aragonite contains heavier isotopes, and calcite, lighter isotopes. These results suggest that the aragonite formed during intense evaporation, whereas the calcite originated as a by-product of the bacterial decomposition of gypsum (NEEV,1963, 1964; see also Limestone Chapter). The calcitic layers, both a t Salt Flat Graben and in the Dead Sea, are related to the activity of bacteria which use lighter isotopes from their carbon source (FEELEY and KULP, 1957, p.1844). Hence isotopically heavy dolomite reflects concentration of brines and the lighter calcite, bacterial fractionation. The sediment couplets of the sediments deposited in the Pleistocene lake in Salt Flat Graben, as in the Dead Sea sediments, probably represent alternations of climate. An important difference between the sediments from these two areas, however, should be pointed out. In the Dead Sea sediments, one half of the couplet, the calcite, formed at the expense of original gypsum. In the Salt Flat Graben lake sediments, on the other hand, the laminated member of the couplet contains gypsum which has not been converted to calcite, although calcite of unknown origin (precipitated or bacterial) is also abundant. The black member contains dolomite and the products of bacterial reduction of gypsum: native sulfur, iron sulfide, and HzS, and some calcite which may or may not be the result of bacterial

292

G. M. FRIEDMAN AND J.

E.

SANDERS

processes. Although the dolomite of Salt Flat Graben is thought to be analogous to the aragonite of the Dead Sea sediments, this dolomite occurs most abundantly within the member of the couplet in which gypsum has been decomposed, and not in the member in which this reaction has not taken place abundantly, as in the case of the Dead Sea aragonite. If this analogy between dolomite and aragonite is correct, then the question arises as to why dolomite formed in the lake in Salt Flat Graben and aragonite in the Dead Sea. The Mg/Ca ratio of the brine was probably the determining factor, as indicated previously by the conditions under which Recent dolomite has formed in sea-marginal settings. The Dead Sea is deeper than the Texas lake and lacks marginal flats; water analyses show that the Mg/Ca ratio is the same as that of normal sea water (NEEV,1964). The Pleistocene lake in Salt Flat Graben was shallow and no information on the Mg/Ca ratio in the medium of deposition at time of dolomite formation is available. Data from some of the shallow wells in the area, however, suggest that at least locally the water has been enriched in Mg content with respect to Ca. Another possibility is that the dolomite of Salt Flat Graben was originally crystallized as aragonite, which immediately or shortly afterward was replaced by dolomite. On the basis of isotopic data, some authors have denied that primary dolomite can be precipitated directly out of the water and, instead, attributed a replacement origin to all dolomite, even that which has been found in the whitish sediment clouds of the South Australian lakes (see DEGENS, 1965, p.119-121). Pleistocene unconsolidated dolomite occurs in lacustrine sediments of Glacial Lake Bonneville near Knolls, Utah, in the Great Salt Lake Desert (GRAFet al., 1961; CAROZZI,1962). The Bonneville lake sediments consist chiefly of halite and calcium sulfate. These salts are thought to have been precipitated from a residual body of water that remained isolated from present Great Salt Lake when the level was lowered from its maximum altitude of 1,566 m to the threshold altitude of 1,287.4 m, which is 6.4-6.7 m above the mean level of Great Salt Lake in historic time (EARDLEY et al., 1957), and then was completely evaporated. Cores show layers composed of calcite, aragonite, dolomite, and quartz, and some layers of nearly pure aragonite. A thin layer composed largely of unconsolidated dolomite crystals a few microns in diameter occurs at a depth of 30 cm. This dolomite is poorly ordered, but the dimensions of its unit cell correspond to those of the ideal stoichiometric dolomite; its radiocarbon age is 11,300 & 250 years B.P. (GRAFet al., 1961). Cores from more shoreward locations revealed the presence of aragonite and magnesite at the same depth as the dolomite from the more lakeward station. Alternating dark and light layers were found in cores from the same general area (BISSELL and CHILINGAR, 1962). The dolomite content of the dark layers varied from 30-50 %, whereas that of the light layers was less than 10 %, and generally comprised only 2-5 % of total carbonates present. This distribution of dolomite is similar to that of Salt Flat Graben just described. Measurements of carbon and oxygen isotopes by E. T. Degens and S. Epstein (reported in BISSELL and CHILIN-


293

GAR, 1962) showed no significant differences between the light and dark sediment layers and between the coexisting dolomite and calcite. This is surprising in view of the clear-cut differences found between the dark and light layers at Salt Flat Graben and between the aragonite and calcite layers in the Dead Sea sediments. Pure, well-ordered, stoichiometric dolomite, ranging in age from 17,400 i600 to 37,000 years B.P., gypsum, and celestite, occur in lacustrine sediments from pluvial lakes in West Texas (REEVESand PARRY,1965). These sediments antedate the maximum advance of Wisconsin glaciers. The well-ordered dolomite contrasts with the protodolomites that characterize most of the other Recent and Pleistocene dolomites, except some of those from Australia. Pleistocene dolomite occurs in the sediments of Lake Eyre, which is situated in a tectonic basin known as the Lake Eyre Basin that forms the lowest point on the 1963; WOPFNERand Australian Continent (KING, 1956; JOHNSand LUDBROOK, TWIDALE,1966). The basin is surrounded by various positive structures; it originated by faulting in the Late Cenozoic. The climate of the basin area is the driest in Australia, but river systems with watershed areas in excess of 1,300,000km2 drain into it, so that innumerable playas occur. During the Pleistocene, gypsite, a sediment composed for the most part of crystalline gypsum, was deposited across the entire basin. Gypsite interfingers with extensive fresh-water limestones which contain dolomite. The limestones vary from dark greenish-gray, microcrystalline, and slightly dolomitic rocks south of Lake Eyre, to yellow-brown, coarsely crystalline, vuggy dolomitic limestones containing abundant casts of reeds and grass near a locality called Coward Springs (WOPFNER and TWIDALE, 1966). Pliocene and Miocene dolomite and dolomitic clay have been found in the center of the basin. No detailed mineralogical studies of Lake Eyre Basin sediments have yet been reported, but the pattern of sediment types is similar to that of other dolomite-bearing carbonate lake deposits. Pleistocene dolomite also occurs in sediments of the Northern Tiberias Basin of the Jordan Valley-Dead Sea Graben of Israel (see previous chapter). The drainage into this basin is internal, as in Salt Flat Graben, Lake Eyre, and other intermontane basins. The absence of dolomite in typical Recent marine carbonate sediments is a question which puzzled geologists for many years. Inasmuch as dolomite is less soluble than either calcite or aragonite, it was naturally expected that dolomite would precipitate readily from sea water of normal composition. Magnesium ions, however, tend to become hydrated, that is, to surround themselves with an envelope of water molecules, in which condition they do not easily enter into crystals. The solution must, therefore, become supersaturated before magnesium crystals precipitate. Dolomite precipitates only when hydration is impeded, such as under conditions of high salinity, high temperature, and high magnesium-ion concentration. When these conditions are realized, whether in nature, or in the laboratory, dolomite originates,

294

G. M. PRIEDMAN AND J. E. SANDERS

Evidence appears to be accumulating that clay minerals may behave as catalysts in dolomite formation by providing a source of magnesium ions or by serving as membranes which effect ionic migration (KAHLE,1965). Clay minerals could provide a source for at least some of the magnesium needed; they could also act as centers of nucleation, by entering into chemical reactions which involve dolomite as one of the products. Replacement of calcium carbonute by dolomite

Most dolomites originated by replacement of calcium carbonate, which was either in the form of aragonite or of calcite in calcium carbonate sediments within the depositional environment; or replacement took place during diagenesis or !ater tectonic activity. Replacement commonly is selective, but it may be pervasive. Replacement dolomite varies over a wide range in its petrographic characteristics: in the size of crystals, degree of crystal development, selectivity of replacement, character of crystal contacts, degree of obliteration or preservation of depositional textures and fabrics, extent of replacement of the original calcium carbonate sediment, and reconversion of dolomite back to a calcite (dedolomitized) rock. The controls of selective dolomitization may be the mineralogical composition of the original sediment in the depositional environment and in a part, commonly a significant part, reflect the faunal distribution, the texture and fabric of the original sediment, and such sedimentary structures as bedding planes. These, in turn, are a function of the physical, chemical and biological nature of the depositional environment. Moreover, inasmuch as not all dolomite forms in the depositional environment, a point which should be stressed after the emphasis in the previous section on dolomite formation in Recent and Pleistocene depositional basins, the effective controls in selective dolomitization may be exercised by the sequence of geological events that follows deposition. Some of these events control lithification, others clearly occur during the post-lithification period. Thus selective dolomitization may take place concurrently with lithification, or long after the rock has been lithified. Selective as well as pervasive dolomitization may proceed in stages; it may begin in the depositional environment and become accentuated by a diagenetic process when magnesium-rich waters react with the carbonate rocks. Much more needs to be learned about selective dolomitization, and commonly no satisfactory explanation can be offered why some limestone interbeds were preferentially dolomitized whereas others were not. Some workers have insisted that all dolomite has originated by replacement. They claim that even dolomite rhombs found in the white sediment clouds of the South Australian lakes, which most geologists would consider to be “primary” precipitates, are instantaneous replacements of aragonite. This argument is based on isotope data, which may be too sparse to be conclusive. As will be discussed presently, replacement dolomite crystals generally preserve some relict textural


295

features of the substance which they replace. The case for the supposed instantaneous replacement of aragonite by dolomite, therefore, would be greatly strengthened if the dolomite rhombs were found to contain relicts of the outline of needlelike aragonite crystals. Many problems concerning the details of replacement remain unsolved. It is not understood whether the results take place by simultaneous solution of the host mineral and deposition of the replacing mineral, or by ionic exchange, or the extent to which or how magnesium is introduced. In the case of dolomite replacement of either aragonite or calcite, crystallographic changes are involved, which are more extensive than simple exchange of ions. Replacement may take place either on a volume-for-volume basis, in which case no single chemical equation is applicable; or on a molecule-for-molecule basis, in which case the theoretical chemical equation is applicable. The molecule-formolecule transformation from calcite to dolomite involves a 12.1 % volume shrinkage, which would be accompanied by an increase in porosity and/or brecciation in the replaced rocks. Such changes do not occur if the replacement occurs on a volume-for-volume basis, which is thought by many geologists to be the commonest type of dolomite replacement. Proof of replacement origin of dolomite consists of: (a) Skeletal remains now composed of dolomite, which were originally secreted as aragonite or calcite; (b) Dolomite crystals which preserve relicts of distinctive calcium carbonate sediments, such as ooids, pseudooids, pellets, coated particles, aggregate particles and so forth; (c) Dolomite patches in limestones; ( d )Preservation in chert of distinctive fabrics of calcium carbonate sediments, where the surrounding dolomite shows only crystal fabrics and lacks vestiges of fabrics from calcium carbonate sediments. The term “replacement” is descriptive; it indicates that mineral a is now in the position originally occupied by b without stating the process which brought about this change. By contrast the German word “Verdrangung”, has a more active meaning. “Verdrangung” implies that a has placed itself in the position of b by the application of force, and has taken the place of b immediately following the application of this force. Strictly speaking, therefore, the meaning of replacement dolomite is not precisely synonymous with the German word “Verdrangungdolomit”. The term metasomatism is commonly used if replacement is accompanied by introduction of new material from outside. Skeletal remains which were originally secreted as high-magnesian or lowmagnesian calcite or aragonite that now consist of dolomite, are convincing proof of a replacement origin of the dolomite. Not all skeletal materials respond equally to dolomitization. Sometimes the original mineralogical composition of the skeletal material determines the sequence of dolomitization, and in many cases dolomitization is guided by the diagenetic history. Wigh-magnesian calcite skeletal materials, such as found in calcareous red Algae, may be dolomitized first, as at Eniwetok Atoll (SCHLANGER, 1957) or Plantagenet Bank near Bermuda (GROSS,1965).

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Yet at Mallorca, one of the Balearic Islands of the western Mediterranean, and at Funafuti and Kita-daito-jima in the Pacific these same fossils resisted dolomitization (FRIEDMAN, 1964, pp.798-799; SCHLANGER, 1963, p.1011). At Mallorca, in Pleistocene limestones, dolomite crystals preferentially replaced drusy calcite mosaics, which originated as void fillings where an original coated aragonite grain had been dissolved (see Chapter 5). Coralline Algae have not been replaced by dolomite crystals in these rocks. A possible explanation is that high-magnesian calcite had been replaced by low-magnesian calcite before dolomitization (FRIEDMAN, 1964, p.799). Shells secreted as aragonite, as in gastropods, cephalopods, and corals, generally are replaced by dolomite before the low-magnesian calcite skeletal remains of brachiopods, ostracods, trilobites, and echinoderms (DIXON, 1907; SANDO,1957). The order just enumerated is based on original mineralogic composition; if thz aragonite shells are first converted to low-magnesian calcite, however, the replaced shells generally are more resistant to dolomitization than skeletal material composed originally of low-magnesian calcite (DIXON,1907). The same is true of the minerals in lime mud and ooids; aragonitic lime mud and ooids are readily dolomitized, but they become more resistant to dolomitization once the aragonite has been replaced by calcite. Dolomite replacement, therefore, may be minutely influenced by the original distribution of aragonite in the calcareous sediments (HATCHet al., 1938; CLOUDand BARNES, 1948). Reefs tend to be dolomitized as a result of one or more of the following three processes: ( I ) the calcareous Algae commonly associated with them consist of high-magnesian calcite that is readily converted to dolomite; (2) reefs and associated sediments are initially porous, so that they become paths of fluid migration; and (3)many (but not all) reefs grew in shallow tropical seas, where the climate commonly promotes increased salinity of the sea water owing to evaporation, both in the main water mass and in the interstitial waters. Such highly saline waters are active agents in dolomitization. The last of the three processes listed, namely strong evaporation of the sea water with resultant formation of hypersaline brines, seems to be on the basis of the evidence presently available, the most logical explanation. The approach shown for the determination of the origin of dolomite in Salt Flat Graben, West Texas (FRIEDMAN, 1966) and for aragonite in the Dead Sea (FRIEDMAN and NEEV, 1966) using carbon and oxygen isotopes has been applied to dolomitization of reefs, Oxygen isotope measurements on dolomite from beneath the atolls of Funafuti, Kita-daito-jima, and Eniwetok in the Pacific Ocean show enrichment in the heavier oxygen isotope (BERNER, 1965). This enrichment can be explained by the reaction of hypersaline brines with the reef material. Dolomitization could have occurred in shallow, restricted backreef lagoons and tidal Aats or below the surface by reflux action. On Plantagenet Bank near Bermuda, poorly ordered Ca-rich dolomite has formed at the expense of skeletal carbonate sands which have the same


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constituent composition as found in the modern Bermuda reef. Oxygen isotope measurements on the dolomitized sediments show enrichment of the heavier isotope (GROSS,1965), which is consistent with an explanation that dolomitization resulted from the interaction of the carbonate sediments with brines produced by the evaporation of sea water. Isotope studies, such as those mentioned, have helped geologists to realize that dolomitization, whether in lakes of intermontane basins, in reef environments, or under tectonic conditions (see section on Epigenetic dolomite) has resulted from the reaction of brines with calcitic or aragonitic carbonate sediments or limestones. Numerous examples of dolomitized reefs are known from the Pleistocene to the Cambrian, both from surface exposures and subsurface samples. The reverse of the usual rule that reefs are dolomitized is found in the Formosa reef limestone (Middle Devonian) of southwestern Ontario, Canada. Here, the reefs, which are composed of flat laminar stromatoporoids and tabulate corals that may comprise up to 50% of the rock, contain nearly pure limestone; their average CaC03 content is 99.13%. The flanking beds, on the other hand, consist of dolomitic limestone (FAGERSTROM, 1961). Distinctive calcium carbonate sediments, most of them originally composed of aragonite, may become dolomitized and the distinctive textural features, if preserved in the dolomite, constitute satisfactory evidence for the replacement origin of the dolomite. Coated spherical particles, such as ooids, pseudooids, and various pellets are known only in calcium carbonate sediments. Their textures are distinct and they commonly persist even where dolomitization has been complete (HOBBS,1957; HAM, 1960; POWERS, 1962). Patches of dolomite in limestone may be related to almost any kind of primary sedimentary structure, either organic or inorganic or to any kind of opening, such as fracture systems or stylolites, found within the deposit. Dolomite may be restricted to a particular part of the deposit as a result of selective dolomitization determined by some primary structure; or it may have spread into the surrounding parts of the mass from some permeable route, which was followed by the migrating solutions that caused the dolomitization. Various examples in the older literature are summarized by VAN TUYL (1916); a review of the types of mottled dolostones and classification of types of mottled features has been drawn up by OSMOND (1956). Five major categories are included: organic, depositional, intrastratal deformational, depositional and diagenetic, and diagenetic, with various subdivisions in each. The sediment filling within burrows may have been dolomitized, whereas the matrix outside the burrows may not have been dolomitized. The reason for this is not known; possibly it relates to grain size or to chemical changes brought about by passage of the sediment through the gut of the burrowing organism. Dolomite may be concentrated along stylolites (OHLE,1951; TOWSE,1957;

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HOBBS,1957), or other openings in the rock. Stylolites tend to be barriers to fluids migrating perpendicular to them, but are channelways for fluids moving laterally along the stylolite surface. Complete dolomitization of very fine-grained lime mud may result in a finegrained dolostone that is almost indistinguishable from fine-grained primary dolomite precipitates (NEWELLet al., 1953; ADAMSand RHODES,1960). Where chert replaced the calcium carbonate sediment prior to dolomitization, however, vestiges of the fabric of the lime sediment may remain in the chert even though they have been destroyed in the parts of the rocks that have been dolomitized (ILLING,1959). The original grain size of the lime sediments that have been replaced by dolomite ranges from fine-grained to coarse-grained. In some cases the dolomite is limited to the fine-grained material; in others, to the coarse-grained sediment; and in still others, it occurs in both. Fine-grained original calcium carbonate sediments were the preferred sites of dolomite replacement in the D-member of the Arab Formation (Upper Jurassic) of Saudi Arabia (POWERS,1962). In these rocks an inverse relationship exists between the amount of dolomite and percentages of grains coarser than sand size. Furthermore, in mixed sediments that include both sand- and mud-size grains, the dolomite first replaced the mud and may have left the sand-size particles unaltered, or replaced them only after complete replacement of the mud had occurred. Other examples of dolomite replacement of original lime mud have been found in the Florena Shale (Permian) of Kansas (IMBRIEand KORNICKER, 1956); Permian of Guadalupe Mountains, southeastern New Mexico (NEWELLet al., 1953); Shunda Formation (Mississippian) of western Canada (ILLING,1959); subsurface Devonian rocks of Andrews South oil field, Texas (LUCIA, 1962); and in originally fine-grained pelletoidal lime mud in the Ordovician of central Virginia (HOBBS, 1957), and Permian of southwestern Oklahoma (HAM,1960). Hobbs concluded that dolomite grew preferentially where initial porosity was high, but original permeability, low. Other dolomite, by contrast, tends to be concentrated in originally coarsegrained sediments. Possibly such dolomite is of later origin (“secondary” as contrasted with “ p e n e c ~ n t e m p ~ r a n e ~ uand s ” ) may be related to fluid migration due to compaction (ILLING,1956). In the Devonian reefs of western Canada, the RimbeyLeduc-Meadowbrook reef chains have been thoroughly dolomitized, whereas isolated reefs such as the Golden Spike, Willingdon, and Redwater have been relatively undolomitized (ILLING,1959). Presumably this relationship is due to the ease of access of through-flowing interstitial waters that were responsible for the dolomitization. Illing attributed these fluid movements to downward escape of waters produced during compaction of the overlying Duvernay and Ireton Shales; accordingly, he inferred that the reefs were dolomitized precisely because they were porous and permeable. Original permeability influences upward migration of sea


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water on supratidal flats, so that all dolomite formed under the influence of original permeability need not be “secondary” in the sense of Illing. i n the Ellenburger rocks of central Texas, “penecontemporaneous” dolomite has been closely controlled by what were thought to have been former zones of greater permeability, even though many of these are no longer permeable as a result of subsequent reactions (CLOUDand BARNES, 1948). Where the dolomite of the Lone Mountain Dolostone interfingers with limestone of the Roberts Mountains Formation (Silurian) in the Pete Hanson Creek area, central Nevada, initially coarser-grained materials have been dolomitized farther from the general facies interface than finer-grained materials (WINTERER and MURPHY,1960). Dolomite occurs in both fine-grained and coarser-grained materials in the Madison Limestone (Mississippian) in the Beaver Lodge oil field, Williams County, North Dakota (TOWSE,1957). Dolomite here is more common along the borders of larger grains and in the matrix next to large grain boundaries, and also in what were streaks of originally coarser-grained sediment. On the other hand, it does occur in very fine-grained material that was initially less permeable. In one wellsorted limestone that consists of alternating layers of very fine-grained sediment and medium-grained sediment, the dolomite has replaced only the very finegrained layers and not the coarser ones. Towse suggested that dolomite replaces fine-grained sediments first where this material lies near coarser and more permeable parts of the sediment. As noted previously, the mineralogical composition of the calcium carbonate may be as important or even more important than grain size per se in controlling selective dolomitization. i n fact, the mineral composition of the calcium carbonate sediments may be related to the grain size, so that both factors may produce the same or opposite tendencies, depending on local conditions. Mineralogy, in turn, in many if not most carbonate sediments depends on the distribution of organisms. The latter reflects environment. Hence, dolomitization is commonly controlled directly or indirectly by depositional environment. In many areas, however, this environmental control of dolomitization is obscured by the introduction of dolomitizing solutions from faults or fractures, a topic discussed in the section on Epigenetic dolomite. Many porous dolostones have formed by replacement of lime muds; lime muds with abundant floating sand-size particles are especially susceptible to dolomitization. An idiotopic fabric (FRIEDMAN, 1965a)is formed by growth of randomly oriented,uniformly-sized dolomite crystals, and the unreplaced calcite is commonly removed by dissolution. Hence, porosity is not formed by dolomitization, as has sometimes been claimed, but by dissolution of calcium carbonate (R. C. MURRAY, 1960; POWERS,1962). This porosity may be obliterated later by precipitation of additional dolomite, forming a xenotopic fabric, or of a calcite cement. Great variation is possible owing to the many mineralogic changes that take place in calcium carbonate sediments and to the influences of mineralogic

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composition on dolomitization. A shell-bearing aragonite lime mud, for example, with originally low-magnesian calcite shells, may be selectively dolomitized and the shells remain unaltered. If later dissolution of calcite occurs, the result is a vuggy dolostone in which only impressions of the fossils remain. On the other hand, if the aragonite of the mud changes first to calcite, dolomite may replace only the originally calcitic shells and may not replace the now-calcitic mud matrix (DIXON, 1907, 191 1). Any aragonite shells that were first changed to calcite may also be as resistant to dolomitization as the calcitized mud. Such mineralogic changes in the calcium carbonate may be responsible for the behavior of lime mud fragments in the Chepultepec Formation (Lower Ordovician) of the central Appalachians near Blacksburg, Virginia (COOPER, 1956). Lime mud that has become stiff enough to break into fragments, thereafter may become very resistant to dolomitization, even though lime mud may be one of the prime sites of dolomite formation prior to fragmentation (HOBBS,1957). In the Funafuti boring, corals at certain levels have been converted to dolomite. According to CULLIS(1904), the change was not directly from aragonite to dolomite; instead, the aragonite of the coral skeletons was first changed to calcite, then primary sparry dolomite was precipitated in the interseptal spaces, and finally, dolomite replaced the now-calcitic skeletal framework. Such changes in susceptibility were used to formulate criteria for distinguishing between penecontemporaneous and subsequent or tectonic replacement dolomite by DIXON(1907, 191 1). In penecontemporaneous dolomitization, aragonite shells, such as cephalopods and gastropods, and lime-mud typically were dolomitized first before calcitic skeletal remains of rugose corals, crinoid plates, and brachiopods. I n secondary dolomitization, which occured after cementation and after aragonite had taken the place of calcite, however, the susceptibility to dolomitization was reversed: the formerly resistant originally calcitic skeletal debris was dolomitized before initially aragonitic skeletal material. These many examples serve to illustrate that definite patterns of replacement exist and that within given geographic and geologic boundaries these may be predictable. Yet, these patterns commonly differ from one area to another or from one formation to another within a given area. Hence, though it may be possible under local conditions to predict patterns of dolomitization, generally valid rules for arriving at such patterns for unknown areas cannot be formulated. It is not surprising that such rules are elusive if one considers the number of petrological variables involved. Dedolornitization

Not only does dolomite replace calcite or aragonite, but calcite may replace dolomite; no examples are known to the writers where aragonite has replaced dolomite. Calcitization of dolomite is called dedolomitization, a term introduced by VON

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MORLOT(1847). This mineralogical replacement of dolomite by calcite may be complete, but commonly it is only partial. Strictly speaking, such dedolomitized rocks are limestones that formed by replacement; therefore, they might belong more properly in the limestone chapter along with limestones formed by replacement of CaS04 and quartz. Because of the close association with dolomite, however, they are included here. Evidence that dedolomitization has occurred is contained in the microscopic fabric of the rocks (see next section). Evidence for dedolomitization includes: (I) remnants of incompletely replaced dolomite within calcite crystals (poikilotopic fabric); (2) calcite pseudomorphs after dolomite; and (3) palimpsest remnants in which ghosts of former rhombic dolomite crystals remain in the form of zones of ferric oxides, or crystal boundaries within a new generation of calcite crystals (see summary by SHEARMAN et al., 1961). Dedolomitization is aided by the presence of sulfate ions, which tend to combine with the magnesium from dolomite to form MgS04 and calcite. The sulfate may come from evaporation of interstitial brine or oxidation of pyrite or other sulfide minerals. According to TATARSKIY (1949) a genetic relationship exists between the presence of interstitial anhydrite and dedolomitization, with MgS04 forming as a by-product, according to the reaction: Ca . Mg (CO&

+ CaS04

--f

2CaC03

+ MgS04

Efflorescent MgS04 has been reported on outcrop surfaces of carbonate rocks which have been dedolomitized. Gypsum interlayered with the dolostones apparently does not aid the dedolomitization process. Carbonate rocks of the Tansill Formation (Permian of West Texas, U.S.A.) illustrate a complex textural history which involves dolomitization, precipitation and dissolution of evaporite minerals, and dedolomitization. All of these reactions could have taken place on sea-marginal supratidal flats. Fossiliferous lime mud which was deposited first, was dolomitized later, and then anhydrite was formed. Still later the anhydrite was dissolved and calcite was precipitated, both as void fillings in the anhydrite crystal molds and as a replacement of dolomite (LUCIA, 1961). Oxidation of pyrite in the weathered zones of limestones in the French Jura is thought to have produced sulfate ions, probably as calcium sulfate and ferrous sulfate, which caused dedolomitization (EVAMY,1963). A similar example may be represented by the carbonate rocks of the Fort Johnson member of the Tribes Hill Formation (Lower Ordovician), near Canajoharie, New York, U.S.A., in which dedolomitization has been widespread. Pyrite is abundant here, but no gypsum or anhydrite were noted (M. Braun and G. M. Friedman, unpublished). Other examples of dedolomitization have been described from many parts of the Soviet Union (TATARSKIY,1949; CHILINGAR, 1956; KHVOROVA, 1957;

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MAKHLAEV, 1957); from the northern Alps (SANDER,1936, 1951); French Jura (SHEARMAN et al., 1961); and Jurassic rocks of the Negev, Israel (Amitai Katz, personal communication to G. M. Friedman). Microscopic crystal textures and ,fabrics of dolostones General statement The original textures and fabrics of calcium carbonate deposits may be preserved by selective dolomitization, but commonly these are obscured or obliterated by dolomitization, and new textures and fabrics originate. The latter bear little or no relationship to the original ones. These new dolostone textures and fabrics are analogous to those of igneous and metamorphic rocks; in both groups of rocks they have formed by crystal-growth processes. Distinctive dolostone textures and fabrics typically form by replacement of nondolomitic carbonate deposits, but dolomite can also grow by replacement of earlier dolomite deposits. In addition, dolomite may be replaced by calcite, a process known as dedolomitization. In this section the typical textures and fabrics of dolostone are defined, classified, and, where necessary, illustrated, essentially following FRIEDMAN’S classification (1965a). The textural terms used are the same as those applied to igneous rocks by Cross, Iddings, Pirsson, and Washington (C.I.P.W.). The fabric terms are modifications of the Rohrbach and Rosenbusch terms based on a “-topic” suffix as applied to dolostone fabrics by FRIEDMAN (1965a). Most dolomites and such evaporite minerals as gypsum and anhydrite are formed by crystallization or recrystallization. This process results in crystals. In the publication referred to (FRIEDMAN, 1965a), the terms crystallization textures and fabrics were introduced. The terms were taken from the process rather than the product. The reason for preference of the longer process than the shorter product term was to retain consistency with metamorphic nomenclature which describes analogous textures and fabrics as crystalloblastic. For the sake of brevity and to make the definitions and classification descriptive rather than genetic, a slight modification is introduced here and the terms are renamed crystal textures and fabrics. DeJinitions and classification Crystal texture refers to the shape of mineral crystals. The terms euhedral, subhedral, and anhedral describe the shapes of individual crystals and are used exactly as defined for igneous rocks, so that it is unnecessary to repeat the definitions here. Idiotopic, hypidiotopic, and xenotopic designate dolostones in which the shapes of the majority of crystals present are euhedral, subhedral, and anhedral, respectively. Crystal fabric refers to the size and mutual relationships between crystals. In equigranular fabrics the crystals are more or less of the same size; whereas inequigranular fabrics are those in which the size of crystals varies. Porphyrotopic


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TABLE V DOLOSTONE TEXTURES AND FABRICS

Crystal textures (I) Individual crystals (a)euhedral (b) subhedral (c) anhedral

Crystal fabrics

(I) Equigranular

(2) Majority of crystals in rock as a whole (a) idiotopic (b) hypidiotopic (c) xenotopic

(2)Inequigranular

(a)porphyrotopic (b) poikilotopic

designates inequigranular dolostone fabrics in which larger crystals (pouphyrotopes) are enclosed in a matrix of smaller crystals of the same or a different mineral. The term poikilotopic refers to inequigranular dolostone fabrics in which larger crystals enclose smaller crystals of a different mineral. Dolostone textures and fabrics may be summarized as in Table V. Being descriptive, these terms are readily applied to rock thin-sections. The different crystal shape types in equigranular dolostones are illustrated in Plate IA-C. Porphyrotopic fabrics may include larger dolomite porphyrotopes in a finergrained matrix of dolomite or calcite, or calcite porphyrotopes which have formed by dedolomitization. Evidence for dedolomitic calcite consists of ferric oxide “ghost” outlines of earlier discrete dolomite rhombs, or of “ghosts” of grain boundaries between earlier dolomite crystals within the calcite crystals. Poikilotopic fabric is also a common product of dedolomitization where later calcite replaced earlier dolomite. As a result, large calcite crystals enclose corroded dolomite relicts (Plate ID). Poikilotopic habit, however, has also been noted in calcite vug fillings, in which dolomite rhombs replaced sparry calcite; in this example, poikilotopic fabric is characteristic of an early stage of dolomitization (Plate IE). Experimental synthesis of’ dolomite A wealth of data is available in the literature on the formation of dolomite under laboratory conditions. Most of these experiments were carried out with compositions, temperatures, and pressures which are far removed from those under which dolomite is formed in the depositional environment. No need is seen to present a historical survey of dolomite synthesis, because the subject is summarized in detail

304 PLATE I



Legend see p.306

305

306

G. M. FRIEDMAN AND J . E. SANDERS

PLATE I (continued)

Photomicrographs of dolomite crystal fabrics. (After FRIEDMAN, 1965a;reproduced by permission of the Journal of Sedimentary Petrology.) A.

B. C. D. E.

Idiotopic dolomite, Lower Permian, southern Tunisia. (After GLINTZBOECKEL and RABATE: 1964.) Xenotopic dolomite. Cool Creek Formation (Ordovician), Daube Ranch, 8 miles southwest of Mill Creek, Oklahoma. Hypidiotopic dolomite. Cool Creek Formation (Ordovician). Daube Ranch, 8 miles southwest of Mill Creek, Oklahoma. Relic, corroded dolomite crystals partially replaced by poikilotopic calcite. This texture is characteristic of dedolomitization. Capitan reef talus, Permian, west Texas (crossed nicols). (Photo by W. W. Tyrell, Jr.) Dolomite rhombs enclosed in poikilotopic sparry calcite. Calcite makes up vug filling and is replaced by dolomite rhombs. Note difference between euhedral shape of dolomite that has replaced calcite in E, and anhedral relic dolomite grains that were replaced by calcite (in D). Capitan reef talus, Permian, west Texas (crossed nicols). (Photo by W. W. Tyrell, Jr.)

ORIGIN A N D OCCURRENCE OF DOLOSTONES

307

in the review articles by VANTUYL(1916), CAYEUX (1935), FAIRBRIDGE (1957), and INCERSON (1962). SIEGEL (1961) briefly reviewed the experimental attempts of the 19th and early 20th centuries; in all of these experiments high temperatures or pressures or both had to be applied before dolomite was formed (CLARKE,1924). CHILINGAR ( I 956b), GRAFand GOLDSMITH (1956), MEDLIN(1 959), BARONand FAVRE( 1959), and BARON(1960) were among the more recent workers to synthesize dolomite, but in these experiments, too, high temperatures, COz pressures or concentrations of reactants were necessary for the crystallization of dolomite from various starting materials to take place in the laboratory. BUDZINSKI (1961) reported synthesis of dolomite by adding concentrated NazC03 solution to 2M (Ca, Mg) Clz solution, in which the Ca/Mg ratio was l/5, and allowing the gel-like precipitate to stand for a few hours. SIEGEL( I 961) precipitated protodolomite in the laboratory with calcium nitrate, magnesium sulfate, and sodium carbonate. Activated charcoal was used to reduce the reaction rate in order to observe the effect of the rate on the final carbonate. The most important factor which affected the precipitation of dolomite was the carbonate ion concentration, through its effect on the pH of the participating medium. Increases in temperature and concentration of the reagents and reduction in reaction rate controlled the ordering and crystallinity of the dolomite precipitated. The presence of sulfate appeared to be essential to the precipitation of the dolomite. This association is commonly found in depositional environments in which dolomite is formed. The conditions of synthesis, in terms of temperature, pressure, and concentration employed, were far less extreme than those previously reported in the literature. Dolomite can be formed experimentally from sea water as a consequence of one of the following factors, or combination of factors: high temperatures, high chlorinity, presence of aragonite or high-magnesian calcite, and an increase of the Mg/Ca ratio in the solutions (SASS,1965). Sea water is supersaturated with respect to dolomite at temperatures above 22"C, but in the presence of calcite it is undersaturated. Spontaneous precipitation generally begins only after a certain amount of supersaturation; yet even calcite is not chemically precipitated in the marine environment so that it is not surprising that dolomite does not form there. A high p H increases the supersaturation of both calcite and dolomite, but again calcite would be the first to be precipitated, thus suppressing the formation of dolomite. For many years geologists have thought that a breakthrough in understanding of dolomite formation would come from laboratory experiments; the breakthrough, however, came from studies in the field. Yet, the evidence available from experimental work bears out the conclusions reached in the field studies on the environmental conditions under which dolomite is apt to form. Now that the field evidence so strongly indicates that hypersaline brines with high Mg concentrations are responsible for dolomite formation, more suitable laboratory experiments can

308


be designed. It should not be difficult to carry out experiments in which density circulation can occur, so that the laboratory conditions of evaporation approximate those in nature. Previous laboratory failures evidently have been due to the fact that in all cases evaporation took place in small containers in which no density circulation and concentration separations were possible.

OCCURRENCE OF DOLOSTONE

Classijication Dolostones may be classified into four major groups: ( I ) syngenetic dolostone; (2) detrital dolostone; (3) diagenetic dolostone; and ( 4 ) epigenetic dolostone. ( I ) Syngenetic dolostone: (a) Dolostone which interfingers with evaporites, both marine and non-marine, with or without associated terrigenous sediments. (6) Dolostone which interfingers with limestones, both marine and non-marine, with or without terrigenous sediments. ( c ) Dolostone which is interbedded with terrigenous sediments. (d) Dolomite crystals which are disseminated in terrigenous sediments. (e) Dolomite formed by biological agents. ( f ) Miscellaneous nonmarine dolomite. (2) Detrital dolostone and detrital dolomite in limestone. (3) Diagenetic dolostone. Dolostone formed by stratigraphic types of replacement of calcium carbonate sediments, within individual beds and along surfaces of stratigraphic discontinuities. ( 4 ) Epigenetic (structurally-controlled) dolostone. Syngenetic dolostone General statement Syngenetic dolostone is here defined as dolostone that has formed penecontemporaneously in its environment of deposition as a micrite or as fine-grained crystals. This type of dolostone is contrasted with diagenetic dolostone, which formed by replacement of pre-existing calcium carbonate sediments during or following consolidation of the sediments. Diagenetic dolostone also may form penecontemporaneously by replacement of grains and cement of calci um carbonate sediments, hence also in the depositional environment. In borderline cases the distinction between syngenetic and diagenetic dolostone becomes difficult or impossible. If dolomite had formed penecontemporaneously by replacement of grains or the tests of organisms, it would be considered diagenetic; but, if dolomite had replaced an aragonitic matrix between the grains or tests, criteria for recognition would be nonexistent, and the distinction between the two classes of dolomite breaks down. This limitation in border-line cases should not detract from the merits of this


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distinction, however, because, as a rule, the separation of these two classes of dolomite is more clearcut.

Interjngering qf syngenetic dolostone with evaporites, both marine and nonmarim, with or without associated terrigenous sediments i n the section on “Origin and destruction of dolostone” it had been demonstrated that dolomite is an evaporite mineral which is commonly formed with other evaporite minerals, such as gypsum or anhydrite, as for instance at the present time in the Persian Gulf or during Early Recent and Late Pleistocene times at Salt Flat Graben, West Texas. At Salt Flat Graben it was demonstrated (FRIEDMAN, 1966) that dolomite formed during periods of intense evaporation from hypersaline brines and that high temperatures were the triggering mechanism that served to bring about formation of dolomite. In ancient environments, therefore, syngenetic dolomite is commonly associated with other evaporite minerals, such as gypsum, anhydrite, and halite. Examples and descriptions of ancient syngenetic dolomite occurrences with evaporite minerals follow. Dolostones interbedded with gypsum, anhydrite, or salt. typically are very fine-grained, show uniform texture, and lack obvious textural features that indicate origin by replacement of calcium carbonate sediments. Such dolostones have been admitted as possible primary precipitates by many, but not by all observers (TARR, 1919; CLOUDand BARNES, 1948; COOPER,1956; EDIE,1956; DUNBAR and RODGERS, 1957; KRYNINE, 1957; ILLING,1959; andHAM, 1960,amongothers, areproponentsof primary origin; NEWELL et al., 1953, and ADAMSand RHODES, 1960, are dissenters). Almost every statement about primary precipitation of dolomite admitted its existence as a theoretical possibility, but at least before 1962, contained a disclaimer that proof of this origin was lacking. As noted on previous pages, however, the missing proof that syngenetic dolomite does form in its depositional environment has now been found, but whether it formed by precipitation or replacement of aragonite is still being disputed. The lack of association with ordinary limestone has been emphasized as a characteristic of evaporitic dolostone (STEIDTMANN, 1926, p.260; DUNBARand RODGERS, 1957,p . 2 4 2 ) . A ~ A ~ and s RHODES (1960, p. 1916), however, considered that the calcite laminae in the Castile Anhydrite (Permian of West Texas Basin) resulted from direct chemical precipitation; they cited these laminae as the basis for the conclusion that “. . .some, probably all, chemically precipitated carbonates were originally non-dolomitic.” An alternative explanation for the origin of these calcite laminae, however, is, as indicated in the discussion of the Salt Flat Graben and analogous Dead Sea carbcnates, that they formed as a result of the decomposition of gypsum. Adams and Rhodes, however, did associate the origin of dolostone with restricted water circulation, even though they doubted the mechanism of primary precipitation of dolomite; they argued that the fine-grained dolostones resulted from replacement of originally fine-grained calcium carbonate sediments,

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as did NEWELL et al. (1953). Although CLOUDand BARNES (1 948, p.94) admitted that primary dolostone might exist in association with evaporites, they remarked that this type resulted from a “special type of sedimentation” and that it does not apply generally to most dolostones. Examination of dolostones associated with evaporites indicates that both syngenetic and diagenetic types may occur. Chief emphasis here is on the syngenetic type. Examples associated with marine evaporites include the Triassic and Permian of Germany; Permian of West Texas, New Mexico and Oklahoma; Mississippian and Devonian of western Canada; Silurian of the New York-Michigan Basin; Mississippian of Indiana; Carboniferous of the Russian Platform; and the Permian of Yorkshire (Great Britain). An example associated with nonmarine evaporites is the Horse Spring Formation (Tertiary) of southern Nevada. The Permian succession of the evaporite basin in north Germany consists of limestones, dolostones, anhydrite, halite, alkali salts, and interbedded clays, shales, and siltstones (BRINKMANN, 1960). This succession is known as the Zechstein; its four main subdivisions, from base upward, are the Werra, Stassfurt, Leine and Aller beds. The Werra beds begin with a basal conglomerate 1 m thick, which is overlain by a thin (0.25-0.3 m) black shale containing copper, the Kupferschiefer. Next follows a skeletal limestone containing brachiopods, clams, and snails, 5-7 m thick, which passes laterally into algal and bryozoan reefs on submarine ridges or in former coastal zones. In north Hanover this limestone (“Zechsteinkalk”) is overlain in ascending order, by a lower anhydrite 30 ni thick, 5 m of salt, and an upper anhydrite 20 m thick. To the southeast, in Thuringia, these Werra anhydrite and salt beds pass laterally into 15 m of dolostone (“Hauptdolomit”) and elsewhere in Thuringia, this whole sequence is represented by 40 m of reef limestone. The overlying Stassfurt beds include 10 ni of shale at the base; then follow upward 2 m of anhydrite, 500 m of salt, and 30 m of covering beds composed of terrigenous sediments. The overlying Leine beds include 5 in of basal gray clay with salt, 30 m of anhydrite (“Hauptanhydrit”), which in Thuringia is represented by 10 m of dolostone (“Plattendolomit”). The middle part of the German Triassic (“Muschelkalk”) also includes limestones, dolostones, anhydrite, and marls, which have been subdivided into three groups; the lower carbonate rocks (“Wellenkalk” or “Wellendolomit”), middle evaporite (anhydrite and halite) and marl unit, and an upper unit (“Haupt Muschelkalk”), which in south Germany (Black Forest and Swabian region) consists of a basal limestone (“Trochitenkalk”), 12-25 m thick, middle shales, 70-80 m thick and an upper dolostone 10 m thick (“Trigomdux Dolomit”). Triassic limestones occur nearer the center of the former marine basin, whereas the dolostones are found nearer the ancient coastal areas, where the salinity of the sea water is thought to have been greater (BRINKMANN, 1960, p.73).


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The cyclic Lower and Middle Permian strata of western Oklahoma (HAM, 1960) are examples of limestone associated with evaporites. Attention here is centered on the presumed syngenetic dolostones. From the base upward, the Permian includes Wellington, Cimarron-Hennessey, Blaine-Dog Creek, and Cloud Chief Formations; the thickness of these sediments is 1,000 m, but only the last two are found as outcrops. The Blaine has been studied in most detail; it is 80 ni thick and consists of 60-70 % gypsum, 17-30 % shale, and 7-15 % dolostone. The anhydrite, gypsum, dolostone, and shale predominate in the northwest, whereas terrigenous sediments predominate in the southeast. The four named gypsum subdivisions of the Blaine from base upward are: Haystack, Cedartop, Collingsworth, and Van Vacter. Very fine-grained, presumably evaporitic dolostones occur in the Mangum Dolostone at the base of the Van Vacter Gypsum. The Mangum Dolostone also includes coarser-grained dolostones formed by replacement of oolitic and pelletoidal calcium carbonate sediments. The Jester Dolostone, which underlies the Cedartop Gypsum, is a persistent bed only 0.3 m thick. It is microgranular and laminated toward the west, but coarser-grained toward the south and east, where it is clearly secondary after pelletoidal calcium carbonates sediments. The Devonian and Mississippian complex of reefs, various limestones, dolostones, anhydrites, and shales of Alberta and southern Saskatchewan locally includes interbedded anhydrite and syngenetic dolostone. Anhydrite is limited to the subsurface; its place on the outcrop is taken by distinctive collapse breccias (ILLING, 1959). Interbedded anhydrite and dolostone occur in the upper part of the Methy Formation (GREINER,1956), and also in the Wabamun (Upper Devonian), Shunda (lower part of the Mississippian) and Mount Head (upper part of Mississippian) Formations (ILLING,1959). Dolostone formed by replacement of finegrained calcium carbonate sediments is also present with the primary type. One of the best-exposed and well-documented examples of dolostones associated with evaporites and minor terrigenous sediments is the Permian succession of the northwestern margin of the Delaware Basin, southeastern New Mexico and western Texas, U.S.A. The basin-marginal area has been termed the shelf; the shoreline in this area trended northeast-southwest, and the facies boundaries are parallel to it. The stratigraphic units have been named Carlsbad Limestone and Queen, Seven Rivers, Yates, and Tansill Formations (KING, 1948; HAYES,1964). Numerous facies belts of variable width and development are present; the general order from southeast (side of normal sea water) to northwest (toward land) is as follows: (1) dolomitized coquina and calcarenite; (2) dolomitic pisolites; (3) finegrained syngenetic dolostones; ( 4 ) sulfates and halite evaporites; and (5) terrigenous sandstones and redbeds. Evidently a very shallow sea spread over the marginal area (shelf) of the basin. Where the water was of normal marine salinity or was slightly hypersaline, both reef- and nonreef organisms flourished. The hypersalinity of the water increased northwestward toward the land; evaporites become the predominant materials in

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this direction. Still farther toward the northwest, quartz sandstones and redbeds were deposited along the edge of a low-lying coastal plain, the surface of which was composed of older Permian strata. This Permian example and others considered previously, both in this chapter and in the preceding chapter on limestones, illustrate a systematic sequence of halite-sulfate-carbonate evaporites, which grade from halite (or “most” evaporitic mineral) at the edge of the sea, to sulfates and carbonates in the direction of sea water of normal salinity, as diagnosed by SCRUTON (1953) and others. In a circular gulf, therefore, the stratigraphic pattern becomes a marginal “halo” of halite, with sulfates and carbonates in successive zones arranged concentrically inward toward the center of the basin and also toward the source of inflow of normal sea water. An exactly reverse arrangement, namely halite in the center with successively outward zones of sulfates and carbonates surrounding it, has been inferred from geologic study of some evaporite basins (STRAKHOV, 1958; SLOSS,1953; RONOV, 1956; BRIGGS,1958). The reasons for these supposedly opposite evaporite zonations are not known; a few examples are considered in the following paragraphs. The theoretical basis for the basin-center halite arangement was proposed by BRIGGS(1958). The “inlet wedge” of the Briggs model agrees with the entire water mass of the Scruton model; taken alone, this inlet (or “influx”) wedge produces an evaporite mineral succession from halite at the margins of the wedge to carbonates in the central part of the wedge. The resulting pattern of evaporites is arcuate and concave toward the inlet (BRIGGS,1958, fig.2), not convex toward the inlet as stated by BRIGGS(1958, p.49). The Briggs model, however, includes an additional feature whose importance has not been established, namely, a basinmarginal zone of water of normal salinity, which grades into saline water at the center of the basin. This additional feature of basin hydrology would seem to require a great influx of fresh water, which at first sight seems incompatible with the conditions of evaporation that produce the high salinity of the influx wedge. One possible explanation is that a major river system discharges into the basin, but derives its water from a distant humid area, so that its continued flow is independent of the arid conditions over the sea water. The Salina Formation (Upper Silurian) of the Michigan Basin and OhioNew York Basin seems to contain a basin-center halite zone with basin marginal carbonate zone, of which dolostone is the predominant carbonate except near the supposed former connections to the open sea, where limestones occur (BRIGGS, 1958). Dolostone of the Russian Platform occurs at the center of the basin and grades laterally outward into partially dolomitized rocks, and beyond these, into limestones (RONOV,1956). The Middle Carboniferous of the Russian Platform, U.S.S.R., consists of interbedded terrigenous deposits and carbonate rocks including limestones and


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dolostones. The dolostones are of two types: (I) “primary” syngenetic dolostones, which contain distinctive brachiopods, and (2) “secondary” diagenetic dolostones. Locally the dolostones include anhydrite interbeds and disseminated crystals of gypsum, fluorite, celestite, and other (unspecified) sulfates. The origin of the dolostones is not clear; they are inferred to have been deposited in a basin whose waters were highly saline (STRAKHOV, 1948). The environment of deposition is thought to have included extremely shallow water and extensive partly emergent shoals, with associated closed lagoons. During hot, dry times, brine is thought to have formed in the lagoons and to have deposited dolomite, fluorite, celestite, and gypsum. In spite of this environmental reconstruction the stratigraphic pattern is presumed to indicate that the high-salinity facies occupied the center of the basin, and the deposits of more normal salinity, the basin-marginal areas (compare with RONOV,1956). The St. Louis Limestone (Mississippian) of southwestern Indiana locally contains dolomitic units that are succeeded by gypsum and anhydrite. A correlation exists between increasingly saline units and increasing segregation into recognizable tectonic units, such as “shelves”, “basins”, and “hinge lines” (MCGREGOR, 1954). Dolomite and anhydrite layers occur in the halite-anhydrite zone of the Permian evaporites of northeast Yorkshire, England. Dolomite is usually subordinate, but in places comprises 70% of the deposit. Petrographic evidence suggests that the magnesite and anhydrite present may have formed at the expense of original dolomite (F. H. STEWART, 1951). Other examples and a review of the literature on limestones and dolostones from evaporite sequences in the United States, Canada, Great Britain, and U.S.S.R. are contained in a paper by GRAF(1960). This paper reviews limestone and dolostone sequences both with and without other evaporite minerals. The limestones are products of deposition in shallow marine waters, and contain abundant fossils which disappear or decrease in abundance as the dolostone is approached; this change in abundance suggests a salinity control on the formation of the dolostone (STRAKHOV,1956; IMBRIE,1957; FOLK,1958). The Horse Spring Formation (Tertiary) of southern Nevada (LONGWELL, 1928) includes variws lake deposits such as limestone, compact clay, sandstone, volcanic tuff, gypsum, dolostone, and magnesite. The dolostone occurs as thin interbeds in soft white magnesium carbonate; other interbeds include pink calcareous sandstone. The thickness of the Horse Spring Formation ranges up to 200 m or more. A diagnostic nonmarine evaporite mineral is colemanite, a source of borax. Interjngering of syngenetic dolostone with limestones, both marine and nonmarine, with or without terrigenous sediments Dolostones most commonly are associated with limestones, either as dolostone interbeds in a predominantly limestone succession, or as dolostone formations

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that are lateral equivalents of limestone units. Both syngenetic and replacement diagenetic types of dolostones are associated with limestones; the primary types are identical with those found associated with evaporites. The primary dolostones occupy a distinctive position with respect to evaporite and limestone associations. In evaporite successions that also include limestones, dolostone occurs between the limestones and the evaporites; in evaporite sequences which include no limestones the dolostones represent deposits that formed further from the shoreline. In general, these same remarks apply to most of the associated dolostone that i s of the stratigraphic-replacement type. Syngenetic dolostone is formed by inorganic processes on broad shelves landward from marine carbonates, usually in an environment hostile to most marine organisms. As discussed in an earlier section, this type of dolostone is formed under supratidal or intratidal conditions and, as in the Persian Gulf of the present day, it interfingers landward with evaporites, such as gypsum and/or anhydrite, and seaward with normal marine carbonates. Typically, syngenetic dolostone which is formed by inorganic agencies is deposited under intense evaporation conditions and genetically must be considered as an evaporite mineral. In the geologic column, it commonly interfingers with gypsum, anhydrite, or both, but this is not always so. Even in the Recent environment of the Bahamas, where dolostone is formed, gypsum or anhydrite are absent. The presence or absence of anhydrite or gypsum is related to climaticconditions. In an arid environment like the presentday Persian Gulf, these evaporite minerals are formed in abundance, whereas under the more humid conditions of the Bahamas they do not develop. Syngenetic dolostone in many areas interfingers with marine or nonmarine terrigenous sediments, and evaporites, such as gypsum or anhydrite, may or may not be present. Numerous examples of syngenetic dolostone can be cited, with an age range from Early Paleozoic to Recent. In this section, only a few selected examples are discussed to demonstrate the depositional environment of syngenetic dolostone, which lacks evaporite association. The Manlius Formation (Lower Devonian of RICKARD, 1962; Upper Silurian of others) in New York State, U.S.A., consists of three distinct sedimentary facies based on grain composition, texture, fossils, and primary structures (LAPORTE, 1964a,b). These facies are time-transgressive and become progressively younger toward the west, indicating a westward migration of environments during submergence. Facies I consists of medium- to thick-bedded pelletal lime mudstone with locally abundant stromatoporoids, which form a compact rock of encrusting and hemispherical colonies. These stromatoporoids appear to have been waveresistant structures composed of framework-building and sediment-binding organisms; many large stroniatoporoid heads have been overturned and abraded, indicating that the water was intermittently strongly agitated. Other fossils include rugose and tabulate corals, gastropods, brachiopods, ostracodes, and codiacean Algae, which are presumed to be green calcareous Algae like the modern Halimeda


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that occurs in subtidal areas of the Bahamas and Florida. Facies 1 is inferred to represent subtidal deposits; it compares with Recent subtidal sediments of the Bahamas and Florida. Facies 2 is composed of thin- to medium-bedded, poorly fossiliferous, pelletal lime mudstones, which locally alternate with pelletal skeletal sands. Scour-andfill structures, ripple-marks, cross-stratification, and lime-pebble conglomerates characterize this facies; ooids are present in minor amounts. Fossils include ostracodes, molluscs, brachiopods, bryozoa, and serpulid worms. The lowest part of this facies commonly contains well-developed algal stromatolites and oncolites (LAPORTE,1963). These algal structures, which make up thin, irregular laminae, encrust free-lying grains (“oncolites”) on the substrate, where they form heads of various shapes and sizes (“stromatolites”). Facies 2 compares with sediments et al. (1964) in the Recent environments of western Australia, described by LOGAN where encrusting mats of filamentous blue-green Algae trap and bind mud and fine sand, at and somewhat below the low-tide level, and form similar stromatolite structures. Facies 2 also compares with the Recent intertidal facies in Florida and the Bahamas, which have been described by GINSBURG (1960) and BAARS(1 963). Facies 3 consists of essentially unfossiliferous laminated dolomitic limestone or dolostone. Individual laminae 1/4-1/2 mm thick, are composed of dolomite spar and rhombs, which grade into calcitic peletal mudstone. The top of the laminae is made up of a very thin bituminous layer which may be the remains of an algal mat. Mud cracks, “birdseye” structures, and burrows are common, but skeletal debris is rare. Facies 3 compares with the Recent supratidal dolomite in the Florida Keys and Bahamas described by SHINNand GINSBURC (1964) and with that from the pellet mud environment of Wide Opening, on the west side of Andros Island, Baha(1964, p.797). mas, described by FRIEDMAN The Lower Ordovician (Canadian) of the Mohawk Valley, New York, U.S.A., consists of similar facies. The Lower Ordovician strata were first examined by EATON(1824), pioneer geologist of the Mohawk Valley and founder and first professor of geology at Rensselaer Polytechnic Institute. These beds have been seen by many other geologists since then; a recent detailed description of them by FISHER (1 954) provides an excellent stratigraphic framework for environmental studies being undertaken by M. Braun at Rensselaer Polytechnic Institute. M. Braun and G.M. Friedman (unpublished) have tentatively recognized two or three distinct facies, which are considered to be deposits of environments similar to those inferred for the Manlius Formation by LAPORTE(1964a,b). These include fossiliferous calcarenite which locally contains: ( I ) worn and abraded pebbles of lime or dolomite mud, thought to be a subtidal deposit; (2) flat-pebble conglomerate in which the pebbles are angular and elongated lime or dolomite mud fragments which locally grade into desiccation cracks, inferred to be of intertidal origin; and (3) aphanitic dolostone, with locally well-developed polygonal mud or desiccation cracks, flat-pebble conglomerate, and abundantly burrowed rock, which are inter-

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preted as supratidal deposits. The flat-pebble conglomerate is composed of lime or dolomite mud, which was broken up by desiccation and then washed by occasional waves into the intertidal zone, or carried into the marine environments, where the pebbles were worn and abraded. Dolostone forms the most landward deposit over what must have been a broad Ordovician shelf, At Hoffman’s quarry, approximately 4 miles east of Amsterdam, New York, the dolomite, known from here eastward as the Gaylor Dolomite, shows typical tidal-flat characteristics with dolomicrite as the principal lithological type and contains algal structures, small channels (probably tidal channels) filled with skeletal sand (biosparite), desiccation cracks in lime and dolomite mud, dolomite pebble conglomerates, and pebbles composed of interlaminated lime and dolomite mud. Landward, the dolomite interfingers with quartz sand (FISHER and HANSON,1951). The Black River Group (Middle Ordovician) crops out on a gentle homocline north and east of Lake Ontario, in New York State, U.S.A., and Ontario, Canada. The regional stratigraphy of these rocks has been summarized by KAY (1937), F. B. YOUNG(1943), and WINDER (1960). The following remarks are based on observations by G. M. Friedman (unpublished). The basal part of the Lowville Limestone near Middleville, New York, includes an aphanitic dolostone, which unconformably overlies the subjacent Cambrian limestones along an irregular surface. Primary sedimentary structures include “birdseye” structures, numerous worm burrows, and typical polygonal mud cracks, such as those found both in the Lower Ordovician and Manlius rocks previously described. Locally, linguoid current ripple-marks and cross-beds indicate unidirectional water flow, such as would be found in tidal channels on tidal flats. No tidal-channel deposits have been positively identified, but such channels may have been responsible in part for the irregular basal depositional surface. The aphanitic dolostone is interbedded with and overlain by calcilutite and pelletal limestone, which are remarkably similar to the Recent sediments of the Bahamas and to the Lower Ordovician rocks and Manlius Formation of New York. Analogous shallow-water and intertidal to supratidal conditions were doubtlessly involved. At the type locality of the Lowville Formation, at Lowville, New York, lime-pebble conglomerates are abundant; they are interbedded with intensely burrowed pelletal calcilutites. Limestone “pebbles” predominate, but dolostone “pebbles” are also present. 7hese “pebbles” must have been derived from areas which were intermittently exposed to subaerial desiccation, such as tidal flats, as in the Lower Ordovician of New York. Near Kingston, and at Napanee, Ontario, calcilutite and lime-pebble conglomerate comprise a major part of the Lowville section; no aphanitic dolostone was observed. At Napanee, lime-pebble conglomerate can be traced through stages from lime mud cut by desiccation cracks, to discrete elongate angular chunks, to well-rounded lime-mud “pebbles”, which apparently were worn by marine waves or currents. Locally oolitic interbeds reflect higher energy conditions in this section. At Napanee, abundant narrow slits or gashes occur in the rocks, as described elsewhere in the


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Lowville Formation by F. B. Young (1943). These slits resemble those which are common in the aphanitic dolostones of the Upper Silurian of Ohio, New York, and Michigan (EHLERS, 1945; SUMMERSON, 1966; and others, summarized in SUMMERSON, 1966); and Devonian dolostones of Indiana (PERKINS, 1963). These slits or gashes have been interpreted as molds of gypsum crystals which have been dissolved away after leaving their impression in the soft carbonate sediments. The layers containing the gash-like gypsum molds in the Silurian rocks of Ohio are unfossiliferous, but stromatolites occur in associated layers. These have been compared with sediments of Recent tidal and supratidal flats of the lower Florida Keys and Bahamas (SUMMERSON, 1966). The gypsum is thought by Summerson to have been precipitated from a brine under conditions of intense evaporation and then to have been dissolved later by successive floods of less saline water. The Lowville association of aphanitic dolostone, linguoid current ripples, cross-beds, desiccation cracks, lime-mud and dolostone pebble conglomerates, worm burrows and gypsum molds is a diagnostic tidal-flat assemblage. These Lowville tidal flats may have been criss-crossed by tidal channels. Intense evaporation evidently occurred from time to time. The sediment-water interface in all of these Paleozoic environments must have sloped gently from the supratidal flats to the submerged subtidal zones, so that extensive areas must have been alternately flooded and drained by the tidal oscillations. In analogous Recent flats, such as those in the Bahamas, or along the Texas Gulf Coast (EMERYand STEVENSON, 1957), tidal waters cut channels and creeks across the flats. At a locality called Greens Corner, approximately 4 miles northeast of Amsterdam, New York, and at a quarry 2 miles north of Tribes Hill, New York, channels, probably tidal channels, occur at or near the top of the Wolfs Hollow Member of the Tribes Hill Formation (Lower Ordovician). The channels cut into a white-weathering lime mud (micrite); they themselves have been filled by gray-weathering skeletal sand (biosparite). The micrite contains abundant dolomitefilled burrows. These burrows are horizontal rather than vertical and many gastropods, especially Oplzileta and Ecculiomphalous(FISHER,1954), are found in this zone suggesting that these burrows were made by gastropods rather than by worms. The channels are filled with skeletal debris as well as with lime mud pebbles and big blocks of lime mud, some of which are over one meter in diameter. These channels must have been high-energy streams that dissected the tidal flats. The Jurassic carbonate rocks of the so-called Makhtesh-Katan, an “erosion cirque” or cove in southern Israel, consist of limestones derived by cementation of pelletal lime sands and both syngenetic and replacement dolostones (GILL, 1966). Composition, texture, and stratigraphy indicate that these materials accumulated in shallow bank environments, parts of which were periodically exposed to subaerial conditions. In the upper part of the dolostone beds erosion channels are found in which Jurassic sediments disconformably overlie the dolostone (M. Goldberg, quoted by GILL, 1966). These minor disconformities were formed

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during intervals of subaerial erosion; they are interpreted by Friedman as being caused by tidal channels which drained the supratidal and subtidal zones of the carbonate banks. Another distinctive feature of these carbonate rocks is the relationship between the limestones and dolostones. The carbonate rocks consist of limestones derived by cementation of pelletal lime sands at the base, but of channeled syrigenetic dolostone in the upper part. This sequence is thought to indicate gradual emergence. Geologic evidence for the origin of dolostone on tidal flats comes from the Cretaceous of Israel. The Quiriat Anavim aphanitic dolostone (Lower Cenomanian; recently renamed Soreq Formation, according to M. Braun, personal communication) occurs in the Judaean Mountains on the west side of the Dead Sea, near Jerusalem. This dolostone complex is 200-250 m thick and is underlain by marine strata of Albian age. At a spot in the village of Beit Zait, near Jerusalem, in an area of about 80 m2, dinosaur tracks have been discovered (AVNIMELECH, 1962). More than 20 footprints occur in a continuous row almost 20 m long; these appear to have been made by a single three-toed theropod reptile, which has been tentatively assigned to the coelurosaurian genus Elaphrosauraus, whose middle toe was 2426 cm long, and side toes, about 20 cm long (Plate IIA). The distance between successive prints made by the same foot is about 160 cm. The hind legs of the reptile are inferred to have been about 120 cm high, and the length of the entire body, 2.5 m or more (AVNIMELECH, 1962). Smaller and less distinct prints occur on both sides of this row of tracks. The entire Cenomanian dolostone complex, though devoid of fossils, was thought to be of marine origin until the discovery of dinosaur tracks. G. M. Friedman interpreted this thick, widespread dolostone deposit as being the product of sedimentation on tidal flats analogous to those of the Recent of the Persian Gulf or of the Manlius and Ordovician deposits of New York State, U.S.A., previously described, even though the Cenomanian dolostone of Israel is much thicker than these Paleozoic deposits of New York. The land reptiles must have moved freely across the tidal flats and, as a result of flooding by diurnal, spring, or storm tides, as in the modern Persian Gulf, the sediments were sufficiently soft for the animals to sink in and leave excellent impressions of their tracks (Plate IIB). In all of these examples, the dolostones and related limestone facies represent bank environments which ranged from supratidal with partial subaerial exposure, to subtidal and complete immersion. The dominant lithologies are syngenetic dolostone and pelletal calcilutites or calcarenites. Such dolostones are the basal deposit of a transgressing carbonate sea. Where they mark such submergences, the dolostones are overlain by fossiliferous marine limestones, as in the Manlius PLATE II Dinosaur footprints in very fine-grained dolostone (dolomicrite), Lower Cretaceous (Cenomanian), Beit Zait, Israel. These tracks appear to have been made in dolomite tidal flat deposits similar to those of the Persian Gulf.

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Formation of New York State, U.S.A., or the Black River Group, southern Ontario, Canada. Marginal marine dolostones may also form as a result of gradual emergence or of seaward progradation of the shore. In such circumstances the dolostone overlies nearshore marine limestones, as in the Jurassic rocks from Israel described previously. An additional stratigraphic characteristic of such sea-marginal dolostones is that they grade seaward into normal marine limestones. Numerous examples of this pattern of distribution have been described from the stratigraphic record; some of these have been mentioned previously. Where dolostone formations pass laterally into limestone units, field relationships suggest that the dolostone formed closer to the ancient shoreline than did the laterally equivalent limestone (DIXON, 1907; VAN TUYL,1918; MCKEE, 1938; HATCHet al., 1938; CLOUDand BARNES,1948; RITTENHOUSE, 1949; FAIRBRIDGE, 1957; BRINKMANN, 1960, among others). This interpretation is further corroborated by the relationship between increased amounts of terrigenous inso-

Fig.2. Stratigraphic diagram across southern Appalachians, showing facies relationships between Cambrian and Lower Ordovician strata. Restored diagram, showing inferred relationships prior and RODGERS, 1957.) to deformation. (After DUNBAR Notice lateral passage from dolostone to limestone in a southeast-northwest direction during deposition of Canasauga Group (Cambrian), but opposite direction of transition (northwest to southeast) in Knox Group (Cambrian and Ordovician). If dolostone represents the deposit formed closer to shore, then a major paleogeographic change took place between the time of deposition of the Conasauga and Knox Groups. The shore line must have lain to the southeast during Conasauga time, but to the northwest, during Knox time. If correct, then the shale of the Conasauga Group represents an offshore accumulation and not a nearshore deposit,


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luble residue and high magnesium content found for the Cambro-Ordovician carbonate rocks in Cumberland County, Pennsylvania, U.S.A., by LESLEY (1 879; see also FAIRBRIDGE, 1957), and by the chemical data assembled for the Lower Ordovician carbonate rocks of the central Appalachians in Virginia, U.S.A., by COOPER (1945, 1956). The stratigraphic relationships of the “Knox” dolostone group, which includes in ascending order Copper Ridge, Chepultepec, Longview, and Mascot Dolostones, found in the northwestern outcrop belts, and ConococheagueJonesboro Limestones, of the southeastern outcrop belts, which illustrate this principle very well, are shown in Fig.2 from DUNBAR and RODGERS (1957, fig.113, p.239). The figure also portrays the distribution of marine shale, limestone, and dolostone in the underlying Conasauga Group (Middle and Upper Cambrian), which is exactly the reverse of that found in the Upper Cambrian-Lower Ordovician rocks just described. The limestones (Rutledge, Maryville, Maynardville) lie northwest of the dolostones (Honaker, Elbrook). Dunbar and Rodgers emphasized that the mud in Conasauga time and non-fragmental silica in Knox time indicate a northwestern source during the entire interval represented by these two groups of strata. If the generalization that dolostone forms closer to ancient shorelines than contemporary limestones is valid, however, then the position of the nearest shoreline must have reversed itself in the area of the diagram: during Conasauga time it lay to the southeast, whereas in Knox time it lay to the northwest. If this analysis is correct, then the shale of the Conasauga Group is an offshore terrigenous deposit flanking a nearshore dolostone deposit, not nearshore terrigenous deposit, as inferred by DUNBAR and RODGERS (1957) and others. If the offshore terrigenous origin is correct, then the Conasauga Group may be added to the other examples of nearshore carbonates and offshore terrigenous sediments discussed in the previous chapter. The Ordovician rocks of the northern Appalachians of western Vermont also provide an example of nearshore dolostone that grades into limestone in the former seaward direction. The Bridgport Dolostone of the Champlain-Richelieu foreland grades eastward into the Beldens Formation of the Highgate Spring-St. Dominique thrust slice, which consists of limestone and dolostone, and this thrust slice grades into the Armand and Corey Limestones of the Phillipsburg thrust slice (CADY,1960). Another example is the Simonson Dolostone (Middle Devonian) of eastern Nevada which grades westward into the Nevada Limestone, in the direction away from the ancient shoreline (OSMOND, 1954). Syngenetic dolostone interbedded with terrigenous sediments In previous sections, syngenetic dolostones associated with evaporites and with limestones have been considered. In both of these associations terrigenous sediments may be present in variable amounts or be absent altogether. A further

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dolostone association is with the terrigenous sediments only, without evaporites or limestones. Typically, such dolostones are interbedded with redbeds. Discrete beds of dolostone are interlayered with terrigenous redbeds in the Permian rocks of Kansas and Oklahoma, Chugwater Formation (Triassic) of Montana and Wyoming, U.S.A. (VANTUYLand STEIDTMANN, 1926, p.253); and New Red Sandstone of England (HATCHet al., 1938, p.186). Units many meters thick that include numerous beds of dolostone are interbedded with terrigenous redbeds of various kinds in the Rome Formation (Lower Cambrian) in eastern Tennessee. The Mangum Dolostone at the base of the Van Vacter Gypsum in the upper Blaine Formation of southwestern Oklahoma extends much farther southeast than the gypsum, so that it becomes interbedded with red shale. Parts of the Mangum dolostone originated by replacement of calcium carbonate sediments (HAM,1960). Thin dolostones that are interbedded with red shales and red and brown sandstones in the Whitehorse Group and Cloud Chief Formation (Permian) of western Oklahoma (SUFFEL, 1930; EVANS,1931;FAY,1962) serve as stratigraphic key beds to subdivide the succession. The basal member of the Whitehorse Group, the Marlow Formation, includes several dolostone beds in its upper part, the Relay Creek Dolostone being the most prominent (FAY,1962). The Marlow Formation is separated from the Rush Creek Sandstone above by the Emanuel Dolostone, and the Rush Creek, from the overlying Cloud Chief Formation by the Weatherford Dolostone. The thickness of these dolostones ranges from 5 cm to approximately 1 m; that of the interbedded terrigenous sediments, from 30-100 m. Thick dolostones are associated with predominantly non-red terrigenous quartzites and slates in the Cambrian rocks of the Rosenburg slice that lies above the Champlain thrust in northwestern Vermont (SHAW,1958). In addition, quartz grains are disseminated throughout many of the dolostones; all variations are present from nearly pure dolostone with a few scattered quartz grains to dolomitic sandstones. Some of the dolostones are overlain by slates with a stratigraphic discontinuity at the contact, whereas others are interbedded in slate formations or form units that are conformable with slates. Stratigraphic discontinuities occur at the contact between the Parker Slate (Lower and Middle Cambrian) and the Dunham Dolostone (Lower Cambrian), and between the St. Albans Slate and Rugg Brook Dolostone (both Middle Cambrian). The eastern exposures of the Dunham Dolostone include dolostone layers 10-90 cm thick interbedded with silty-argillaceous layers up to 15 cm thick. The Parker Slate includes a middle dolostone unit; the Hungerford Slate (Middle Cambrian) includes local dolostone; and the Gorge Formation (Upper Cambrian), which consists of thin-bedded limestones and slates in the middle part and slate in the upper part, includes a basal dolostone unit 170 m thick. The Hungerford Slate, 8-125 (and more) m thick, appears to lie conformably above the Saxe Brook Dolostone, which is 0-215 m thick. Various limestone bioherms and limestone-pebble conglomerates are associated with many


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formations in this succession, so that it would appear to offer a fertile field for petrographic and mineralogic research into the subject of dolomitization. Somewhat similar suites of dolostones and terrigenous sediments with minor limestones and limestone-pebble conglomerates occur in the eastern outcrop belts of the Lower Cambrian in the central and southern Appalachians (RODGERS, 1956). These dolostones (Vintage-Tomstown-Shady) and related rocks constitute the earliest carbonate deposits in the Appalachian geosyncline. The environment of origin of these early dolostones in the Appalachians has not been determined. It is clear from the associations of dolostone with evaporites, limestone or terrigenous sediments, or from their association in different combinations of complex sequences of these rock types, that all dolostone associations are but variations of a single theme: restricted circulation of sea water, evaporation and resultant increased salinity, or highly saline interstitial waters. The question of “primary” (syngenetic) or “secondary” (diagenetic or replacement) is important, but fades into the background; either or both types of dolomite may be present, but they are minor variants in the identical general setting. If dolostone were generally admitted into the ranks of “evaporites”, its connotations would become much more widely appreciated. If this were done, the unifying principle that is evidently responsible for all of these various stratigraphic associations would be emphasized and the geologist would be armed with a powerful tool in attempting to predict stratigraphic patterns. Dolostones interbedded with redbeds or dolostones interbedded with anhydrite, gypsum, halite or alkali salts, are extremes of the various expressions of the evaporitic processes. In the first case, dolostone is the only evaporite present and it indicates a rather low increase of salinity. In the second case, dolomite represents the ‘‘least’’ evaporitic mineral in a sequence where the extreme increase of salinity produced halite or alkali salts. Increased salinity is apparently essential for the origin of dolomite replacements; this increase may take place only within the interstitial waters, however, so that no evaporite precipitation occurs from the mass of sea water, the salinity of which may be “normal”. Saline interstitial solutions may move up or down; they may be in contact only with the upper layers of the sediment mass o r may migrate widely through permeable avenues at greater depths in the sediment mass. Syngenetic dolomite crystals disseminated in terrigenous sediments Syngenetic dolomite crystals have been reported in association with both finegrained and coarse-grained terrigenous sediments. Minute dolomite crystals were first found in the Triassic marls of Gloucestershire and Worcestershire, England (CULLIS,1908,p.506). These dolomite crystals are thought to be primary precipitates from the saline waters in which the marls are known to have been deposited. The precipitated dolomite crystals sank to the bottom and were deposited along with the red mud that had been transported in turbulent suspension.

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Syngenetic dolomite crystals associated with marine terrigenous sediments have been reported from the Upper Cretaceous sandstones of the western interior of the United States (SABINS,1960, 1962). These dolomite crystals, previously not distinguished from calcite, have been identified both by stains (using Alizarin red S dye to color the calcite) and by X-ray methods (the primary crystals show the 2.88 A peak of “ideal” dolomite). Primary dolomite crystals comprise up to 25 % of some Cretaceous marine sandstones, but do not occur in associated nonmarine sandstones; their abundance decreases in a shoreward direction. Sabins has established the association of syngenetic dolomite crystals and marine sandstones in localities from Alaska to New Mexico; accordingly, it is now possible to use the presence of the syngenetic dolomite crystals in the Upper Cretaceous rocks of this vast area as an indication of the marine origin of the enclosing sandstones. The origin of these syngenetic dolomite crystals has not been determined. Sabins has cited them as evidence that dolomite can form from the waters of a basin in which evaporite sediments are absent. In light of the mechanism of upward-transpiration and evaporation of sea water through porous sea-marginal sediments, however, “evaporitic” origin cannot be altogether dismissed merely on the grounds that typical evaporite sediments are absent. Quite possibly the dolomite crystals formed as primary crusts on supratidal flats around the edge of the sea or grew within the sand, while at the same time the salinity of the main water mass did not become abnormally high. Owing to evaporation, the salinity of the interstitial water could have become high enough to precipitate dolomite amongst the quartz grains. Subsequent reworking by the waves would readily remove these dolomite crystals from their place of origin and concentrate them according to their hydraulic characteristics. If the suggested sea-marginal origin is correct, then it should be susceptible of field confirmation, for not all of the loci of interstitial grain growth should have been destroyed, and the sands deposited on the tidal flats and tidal deltas might contain more dolomite than those deposited farther offshore. Syngenetic dolomite formed by biological agents Some syngenetic dolomite has been shown to result from biological activity; this activity involves bacterial processes and not direct organic synthesis of dolomite skeletal material, as noted previously. Crystallites of calcite and spherulites of aragonite and dolomite have formed as films at the surface of aquarium cultures, which contained mud and sea water enriched with glucose (LALOU, 1957a,b). Lalou concluded that these carbonate minerals may be obtained from any sediment if organic matter is present in sufficient amount and the temperature is sufficiently high, and if the bacterial processes take place in shallow water under maximum light and sunshine in quiet and seldomly renewed waters. These conditions are found in tropical lagoons which are isolated from the open-marine environment.


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Dolomite crystals with dark borders, presumably of authigenic origin, have been reported from a depth of 160 m in drill cores of biotite gneiss near Koblenz, Switzerland (NEHRER and ROHRER, 1958). Live bacteria were detected in these rims; the bacteria were isolated and formed new dolomite crystals in the laboratory. The purpose of this experiment was to determine if bacterial action is induced by an enzyme which might contain a characteristic trace metal. If this were true and dolomite containing this trace metal would form, then the presence of the trace metal could be used to indicate a bacterial origin for such dolomite. Dolomite crystals were later synthesized from a pure culture of bacteria with cell size of 0.2-0.3 by 0 . 2 - 0 . 5 ~which ~ grew on a substrate of dissolved gneiss and silica gel at optimum temperatures of 55-58 "C (NEHRER and ROHRER, 1959). This demonstrated the relationship between bacterial activity and dolomite crystal growth. The nitrogen content of the protein of the bacterial cells was lower than usual. This fact was explained on the supposition that a preference existed for the specific substrate used, and, therefore, depended on the activity of the bacteria and possibly on the formation of dolomite. The bacteria referred to by Lalou require light and sunshine, so that dolomite formed by these bacteria in a marine environment would be restricted to water depths shallower than 100 m. The cultures of Nehrer and Rohrer, on the other hand, were active in darkness. Accordingly, different cultures can synthesize dolomite under different conditions (RICOUR,1960). In natural shallow-water marine sediments exposed to sunlight, photosynthesis and the respiration of plants at the water-sediment interface causes diurnal pH changes from 7.4-9.2. During the day, COz is taken up during photosynthesis and the p H is increased; at night, metabolism produces ( 2 0 2 , which decreases the pH (OPPENHEIMER and MASTER,1964). A mixture of bacteria and Algae in experimental aquaria containing quartz and carbonate minerals overlain by sea water subjected to dark and light fluctuations reproduced natural p H changes. After 20 days aragonite had undergone no change, but up to 5 % dolomite had formed in the calcite-containing aquaria (OPPENHEIMER and MASTER,1964). These experiments indicate that dolomite can be formed as a result of biological activity; yet the evidence to date suggests that biological activity has not been directly responsible for the origin of major dolomite deposits. Biological activity, however, may contribute significantly to dolomite formation indirectly; plants may raise the pH of the water, which may be critical in the origin of some dolomite, as in Australia (pHw9-10). It should be added, however, that the supratidal dolomite of the Persian Gulf sebkhas forms under much lower p H ( m 6-8). Miscellaneous nonmavine dolomite Several weathering processes in both humid and semiarid climates have been found to produce dolomite. Examples have been found in Pleistocene deposits. These have not been included in the section on Recent and Pleistocene dolomite in the

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M. FRIDEMAN A N D J . E. SANDERS

Origin part of this chapter, because they are thought to be unrelated to the major dolomite deposits, which are associated with marine sediments. Dolomite was first discovered in Recent soils by investigations carried out in the northwest coastal area of Wales, Great Britain (HARDY, 1921). Coarse-grained or silty calcite and magnesite occur in saline soils in the Zeravshan Valley, U.S.S.R. Calcium carbonate content increases, and magnesium carbonate content decreases with depth (KUGOCHKOV, 1960). Dolomite is formed when these soils are eroded and redeposited. Dolomite has been found within the zone of water fluctuations in the soil profiles formed on silt and very fine-grained sand of glacial lake beds in the Red River Valley, Minnesota, U.S.A. The carbonate in the soils was converted to dolomite by reaction with magnesium-rich waters (SHERMAN et al., 1962). When dolomite was found to comprise 35-79 % of the carbonate in soils developed on glacial till, the process of dolomite origin in calcareous soils was suggested (ALWAYand ZETTERBERG, 1935). Dolomite comprises 90 % of the carbonates in the soils of the fields of the University Schools of Agriculture at Crookstown, Minnesota (HIDE,1935). The identification of dolomite in the subsoils was confirmed by X-ray diffraction analysis (SHERMAN and THIEL,1939). More recently these soils have been studied very extensively; 205 chemical analyses for COa, CaO, and MgO have been made and samples have been analyzed by differential thermal analysis and X-ray diffraction. The proportion of carbonates in the form of dolomite and the CaC03/MgC03 molecular ratios were calculated from the chemical data. The average Fe content of the carbonates was found to be 0.1 % and the average Mn content, 0.03 %. The dolomite in these soils is thought to have formed by the alteration of calcitic carbonates of the original lacusstrine deposits (SHERMAN et al., 1962). Dolomite is abundant as caliche in many parts of the Guadalupe Mountains, West Texas, U.S.A., particularly in and near Salt Flat Graben. This caliche is probably of Pleistocene age; particular care was taken in the study of this deposit to make sure that it was not recycled dolomite, which may have been derived from Permian dolostones in the area (FRIEDMAN, 1966). Dolomite has been reported from caliche in South Africa (T. W. Gevers, personal communication to DEGENS, 1965, p.119). Detrital dolostone atzd detrital dolomite in limestone Dolostones may be derived from cementation of dolomite grains that have been laterally displaced by gravity and resedimented in “deep” water in an environment foreign to their site of origin. Such dolostones resemble comparable limestones; in fact, the point might be raised that they are nothing more than replacements of gravity-displaced calcium carbonate sediments. Whereas this may be perfectly possible, the evidence within the rock itself should provide the basis for assigning


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one origin or the other. The presence of terrigenous (recycled) dolostone rock fragments and abraded dolomite crystals is indicative of re-sedimentation; indications of replacement of calcium carbonate sediments by dolomite are inconclusive, because such replacement might well have occurred prior to gravity displacement. Post-displacement replacement would be indicated by dolomite that post-dates both grains and any earlier cement. Replacement in the deep marine environment, where most gravity-displaced materials have been found, seems unlikely in light of indications that replacement occurs in shallow areas and on supratidal flats (FAIRBRIDGE, 1957; confirmed by recent works previously cited). No examples of Recent gravity-displaced dolomite sediments are known, possibly because of a lack of study of Recent deep-sea sediments by X-ray diffraction, but dolomite grains presumably derived from Cretaceous rocks have been found in the deep-sea sediments of the Gulf of Eilat (Aqaba) (FRIEDMAN, 1966). Two examples can be cited from the geologic record: the Lower Cambrian of the Hudson Valley, New York, U.S.A., and the Middle Ordovician of the Gasp6 Peninsula, Canada. Various amounts of detrital dolomite and dolostone fragments occur in various subdivisions of the Trinity Group (Lower Cretaceous) in central Texas, but dolomite is not abundant enough to form a true dolostone composed of recycled dolomite; rather it is found in the nonmarine sandstones and locally in the middle marine limestone. The amount of detrital dolomite decreases seaward; once the grains become silt-sized, they cannot be distinguished petrographically from silt-sized authigenic dolomite (AMSBURY, 1962). Diugenetic dolostone General statement Diagenetic dolostone is here defined as dolostone that has formed by replacement of calcium carbonate during or after consolidation of the sediment; or it may have formed penecontemporaneously by replacement of grains and cement of calcium carbonate sediments. This type of dolostone is contrasted with syngenetic dolostone that has formed penecontemporaneously in its environment of deposition as a dolomicrite or as fine-grained crystals. The reference to the fine grain size of syngenetic dolostone is important in this definition. Under occurrence of syngenetic dolostone and in the section on Recent and Pleistocene dolomite it was shown that this class of dolostone is formed in the depositional environment as fine mud or crystals. By contrast diagenetic dolomite is a replacement product and can normally (1964) have shown that Recent gasbe identified as such. SHINNand GINSBURG tropod shells and pellets have undergone penecontemporaneous dolomitization in their environment of deposition; they noted that the concentration of dolomite increases as the sediments are progressively lithified. The penecontemporaneous replacement dolomite described by SHINNand GINSBURG (1964) is considered as

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M. FRIEDMAN AND J.

E. SANDERS

diagenetic dolomite, its formation being coincident with lithification; it forms by replacement of pre-existing calcium carbonate sediment. As already pointed out in the definition of syngenetic dolomite, there are borderline cases where the distinction between syngenetic and diagenetic dolomite becomes difficult or impossible. This, of course, is true of any classification and should not detract from the merits of this distinction where it can be made unambiguously. Diagenetic dolostone formed by stratigraphic types of replacement of calcium carbonate sediments and along surjiaces of stratigraphic discontinuities Dolostones formed by stratigraphic types of replacement of calcium carbonate sediments are extremely common. They include two types: ( I ) dolostone formed within beds of carbonate sediments (S-dolostones of DUNBAR and RODGERS, 1957), and (2) dolostones formed along surfaces of stratigraphic discontinuity (“unconformity” type, possibly including some c;f the W-dolostones of DUNBAR and RODGERS, 1957). The topic of replacement origin of dolostone has been thoroughly covered by previous workers, who were concerned with the origin of dolostone, so that little new can be added except interpretations that follow from the latest discoveries of dolomite in Recent marine sediments. The topic of the formation of replacement dolomite has already been taken up in the section on Origin of dolomite so that the present discussion can be kept to a minimum. Diagenetic dolostones formed by replacement within individual beds of calcium carbonate sediments. Dolostones that originated by replacement of calcium carbonate sediments may be of regional or local extent. Although their megascopic aspect may not reveal textural evidence of their replacement origin, typically thinsections show evidence of diagnostic calcium carbonate fabrics. In many dolostones, however, these depositional fabrics have been obscured or obliterated. Examples of dolostones formed by replacement include dolomitized reefs and dolomitized calcium carbonate sediments. The sediments may display ghosts or unreplaced relicts of the original grains, such as ooids, skeletal grains or pellets. Many examples are known of dolomitized reefs. A few include the Lone Mountain Dolostone (Silurian) of Central Nevada (WINTERER and MURPHY, 1960); the Wabash reef (Silurian), Wabash, Ind. (CAROZZIand ZADNIK,1959); the Goat Seep reef and part of the Capitan reef of the Permian of West Texas and New Mexico (NEWELLet al., 1953); the Cooking Lake Formation of Alberta, Canada (ANDRICHUK,1958); and the reefs of the Racine Formation (Silurian) near Chicago, Ill. (WILLMAN,1943). Examples of dolomitized calcium carbonate sediments (from VAN T ~ Yand L STEIDTMANN, 1926, unless otherwise indicated) include: Elbrook Limestone (Cambrian) from West Waynesboro, Pennsylvania; Elvins Formation (Cambrian), Elvins, Missouri; the Hoyt Limestone (Upper Cambrian) at Saratoga Springs, New York; Tribes Hill Formation (Lower Ordo-


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vician), Canajoharie, New York; Monroe Dolostone (Silurian), Sylvania, Ohio; basal Oneota Dolostone (Lower Ordovician) in the northern Mississippi River valley; the Lower Ordovician Dolostone near Blacksburg, Virginia (HOBBS,1957); Jester Dolostone at base of Cedartop Gypsum, Creta Dolostone at base of Collingsworth Gypsum, and Mangum Dolostone at base of Van Vacter Gypsum members of the Blaine Formation (Permian), in southwestern Oklahoma (HAM, 1960); and D Member of the Arab Formation (Upper Jurassic) of Saudi Arabia (POWERS,1962). Diagenetic dolostone formed along surfaces of stratigraphic discontinuity. Dolostones that formed along surfaces of stratigraphic discontinuity are usually called “unconformity” types and their origin is almost universally attributed to the effects of subaerial weathering, as would follow from the connotations generally implied by use of the term “unconformity”. Subaerial weathering may or may not have occurred in the history of the discontinuity, and it may or may not have been involved in the formation of the dolomite. Accordingly, the terminology preferred here is of a more general descriptive nature and refers to “discontinuity” rather than “unconformity”, and does not emphasize subaerial weathering. Stratigraphic discontinuities in marine sediments can form without subaerial weathering, and even where such weathering did occur in the history of a particular contact, it may or may not have been involved in any dolomitization related to the contact. For example, dolomitization along a stratigraphic discontinuity may be caused by downward migration of supersaline sea water (“seepage reflux” of ADAMs and RHODES,1960; a process, incidentally, which is not necessarily restricted to surfaces of stratigraphic discontinuity). The Pleistocene dolomite on the north side of Bonaire, Netherlands Antilles, has been attributed to the influence of downward seepage-reflux of magnesium-enriched waters, magnesium enrichment of which resulted from calcium depletion due to evaporitic deposition of aragonite and gypsum (DEFFEYES et al., 1964). Dolostones of this type extend downward from the surface of discontinuity and tend to form an irregular dolostone capping that cuts across many beds. Subaerial exposure doubtless occurred on Bonaire, but is thought to have been incidental to the dolomitization. The second generation of dolostone in the Devonian of the Eifel, Germany, may have been produced by seepage reflux from the highly saline Permian seas that elsewhere deposited evaporites on a large scale (QUIRING,1913; REULING,1931; see also FAIRBRIDGE, 1957). Because of the stratigraphic discontinuity present, this second generation of dolomite has been ascribed to subaerial weathering, but in light of more recent discoveries, the interpretation that the dolomite formation was related to concentrated sea water merits further investigation. Another possibility is that dolostone formed by upward transpiration of sea water through the subaerially-exposed sediments on a supratidal flat. The basal contact of such a dolostone bed may be irregular, but its top should build up to

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a flat surface, in contrast with the irregularity found both at top and bottom of a capping-type dolostone that extends downward from a discontinuity.

Epigenetic dolostone Epigenetic dolostone is here defined as dolostone that has formed by replacement of limestone with the dolomite being localized by post-depositional structural elements. It is closely related to faults and fractures in carbonate rocks. The term epigenetic is synonymous with such terms as structurally-controlled dolostone, tectonic dolostone, or T-dolostone (DUNBAR and RODGERS,1957). It is consistent with the terms syngenetic and diagenetic which have been introduced earlier in this chapter; the term epigenetic has been adopted for characterizing this genetic class of dolostone to make a consistent classification. The terms syngenetic, diagenetic, and epigenetic imply a genetic sequence and have been used as the basis of classification of dolostone. Many, but not all, epigenetic dolostones are genetically associated with metallic ore deposits, notably of lead and zinc minerals. Examples of epigenetic dolostone related to fractures but unrelated to ore deposit are found in: ( I ) Cool Creek Limestone in Arbuckle Mountains, Oklahoma; (2) Helderberg Group in Hudson Valley, New York; (3) Cambrian and Ordovician limestones of southwestern Ohio; and ( 4 ) Mississippian Limestone along Keystone thrust, southern Nevada. The Cool Creek Limestone (Ordovician) has been extensively dolomitized near Pennsylvanian fault zones, 12 km southwest of Mill Creek, in the Arbuckle Mountains of southern Oklahoma. This area has been mapped by W. E. Ham (unpublished), who provided his manuscript map and pertinent aerial photographs for a study by G. M. Friedman (unpublished). The Cool Creek limestone is a pelletal calcilutite which lacks dolomite; but it has been completely dolomitized and shows typical crystallization or crystal fabric where it is contiguous with the fault zone. The change from limestone to dolostone is abrupt. Thin-section study reveals that the limestone, even in close proximity to the dolomitized zones, retains original depositional textures; nothing in the thin section suggests the possibility of nearby dolomite. The dolostone, on the other hand, invariably shows a completely crystalline fabric, without any traces of original depositional textures of calcium carbonate sediments. The extreme sharpness of the lateral contact between the two rock types and absence of a transition zone between them was confirmed by mineralogical and various geochemical studies. Magnesium and manganese contents increase, strontium decreases, and iron shows no consistent trend on approaching the faults (G. M. Friedman, unpublished). The limestone is of Ordovician age and the faults are of Pennsylvanian age; therefore, dolomitization must have occurred during Pennsylvanian time (HAM,1951). Studies of oxygen and carbon isotopes give a clue to the composition of the fluids that caused dolo-


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mitization. Isotopic analysis of limestone and epigenetic dolostone from the Cool Creek Formation has shown that the dolomite has been enriched in the heavier isotopes (180 and 13C) (G. M. Friedman, unpublished; isotope analyses by W. M. Sackett). The range for 60ls, corrected to PDB, was -7.1 to -7.8 for limestone and -2.7 to -4.6 for dolostone from major fault zones. In dolostone from minor breaks in limestone, the values were closer to those of limestone than to those of dolostone (-6.2 to -7.6). For 6I3C, corrected to PDB, the difference in values between the two lithologies was not as pronounced; S13C for limestone showed a range from -1.4 to -3.9 and that of dolostone from -0.9 to -2.8. Thus, an overlap for 613C between the two lithologies exists; however, of ten limestone samples six have values between -2.6 and -3.9, and of thirteen dolostone samples studied eight show a range between -0.9 and - 1.9, indicating that enrichment in the heavier carbon isotopes has taken place. Enrichment of dolostone in heavy oxygen and carbon isotopes suggests that it formed with the aid of waters having high d18O and 6l3C values, particularly high P O . This enrichment of the original water in heavier isotopes that was responsible for dolomitization must have resulted from strong evaporative processes with the lighter isotopes preferentially removed as part of COz. This isotopically-heavy water was a hypersaline brine. It recalls the observations which were made in the case of the Salt Flat Graben dolomite (FRIEDMAN, 1966), mid-Pacific atoll dolomite (BERNER, 1965), and dolomitized skeletal fragments of Bermuda (GROSS,1965), for all of which a reaction with isotopically-heavy water was inferred. In the opinion of the writers, these observations now permit the generalization that all dolomites, whether syngenetic, diagenetic or epigenetic owe their origin to hypersaline brines. The Helderberg Group and base of the overlying Ulster Group (Lower Devonian) in the Ravena area, New York, 24 km south of Troy, have been cut by a thrust fault, along which extensive dolomitization of limestones has occurred (J. R. Dunn, personal communication). A transition zone 3-6 m wide is present; the dolostone itself may extend up to 10 m away from the fault zone. The Becraft Limestone, studied by Dunn, shows an increase in MgO content towards the fault zone from 0.9-6.2 % in a distance of 3 m. Quarry operators in this area have made many chemical analyses of the carbonate rocks, which show a significant increase in magnesium content near the thrust, reflecting the presence of the mineral dolomite. The transitional contact between dolostone and limestone here contrasts with the abrupt contact between these two rocks in the Mill Creek area of Oklahoma, described above. Replacement dolostones in the Cambrian and Ordovician limestones of the Cincinnati Arch province in Ohio, U.S.A., are locally related to fractures, faults, and solution channels (CALVERT,1964). These dolostones typically are fine to coarsely crystalline and are thought to have been formed by circulation of magnesium-rich waters through the limestones. The magnesium-rich waters are presumed to have moved vertically via deep-seated joints and faults, and then to have spread

332


laterally along porous zones and bedding planes. The upward migration of the magnesian waters was arrested by thick overlying shales. The source of the magnesium was probably subsurface water, which was derived from the underlying Cambrian strata (CALVERT, 1964, p.187). Epigenetic dolostones are related to many lead-zinc deposits in carbonate rocks; such associations are known from widely scattered localities from many parts of the world. These occurrences have been summarized by HEWETT (1928). Examples included here are from: ( I ) Bromide, Oklahoma; (2) Goodsprings, Nevada; and (3) east Tennessee, in U.S.A.; ( 4 ) northern England and southern Wales; and (5) the Eifel, western Germany. The Ordovician and Silurian limestones at Bromide, Oklahoma, U.S.A.,have been-extensively dolomitized along persistent faults (HEWETT, 1921 ; 1928). A progressive increase in magnesium, iron, and manganese contents takes place in the wallrocks going toward the faults. Manganiferous carbonates and hausmannite occur with the dolomite, indicating presence of low concentrations of manganese. The Bromide area is near the Mill Creek district studied by Friedman and mentioned here in a previous paragraph, but contrasts with it in showing a gradual transition rather than an abrupt change from dolostone to limestone. HEWETT (1928) remarked on the existence of abrupt changes from dolostone to limestone, but also noted that in places like Bromide, Oklahoma, and Raible, Italy, a transition zone as much as 15-30 m thick may be present. Near fractures in the Goodsprings quadrangle, Nevada, the limestones have been converted into epigenetic dolostone by destructive replacement. Original depositional textures of calcium carbonate sediments and such fossils as Foraminifera, corals, brachiopods, and gastropods exist in the unreplaced limestone, but have been progressively destroyed in the epigenetic dolostone (HEWETT, 193 1). In the east Tennessee zinc district the ore bodies are located in a coarsegrained tectonic type of dolostone, known locally as “recrystalline” rock. This TABLE VI SUMMARY OF PERMEABILITY MEASUREMENTS TN ROCKS OF MASCOT-JEFFERSON CITY DISTRICT, EAST TENNESSEE (MEASUREMENTS FROM

Rock

_____

Kingsport Limestone Mascot S-dolostone T-dolostone formed by replacement of Kingsport Limestone

No. of samples

OHLE,1951)

Average permeability (millidarcys X lo-@)

11 17

21 81

18

4900

Range of permeability (millidarcys X

160580-

0.25 3.1

39,000450


333

type of dolostone has been formed under two controls: ( I ) faults, particularly small tear faults, which have caused extensive brecciation of the nearby rocks; and (2) stratigraphic level in the Knox Group, the preferred horizon of replacement being the Kingsport Limestone. The Kingsport Limestone is a fine-grained pelletal rock which is similar to the Cool Creek Limestone (Ordovician) of Mill Creek, Oklahoma, mentioned previously. The Kingsport Limestone underlies the Mascot Dolostone, an S-type of dolostone (DUNBAR and RODGERS, 1957),that comprises the uppermost member of the Knox Group (Cambro-Ordovician). A representative example is found in the Mascot-Jefferson City district (ODERand MILLER,1945; BRIDGE,1956; OHLE, 1951). Laboratory measurements have indicated that the coarse-grained epigenetic dolostone formed by replacement of Kingsport Limestone, the host rock for the ore bodies, is the most permeable rock in the fault zones (OHLE,1951). Results of these measurements are summarized in Table VL The average increase in permeability as a result of conversion of Kingsport Limestone to coarse-grained epigenetic dolostone is 18,000%; the extreme range is 2,000,000%. The dolostone has formed in permeable zones outward from the fractures, such as along stylolite seams. Photomicrographs of unaltered Kingsport Limestone, partly replaced limestone, and coarse-grained dolostone are found in the book by DUNBAR and RODGERS (1957). Epigenetic dolostone has formed by replacement of the Carboniferous Limestone in northern England and southern Wales, where it has been cut by faults. In the Pennine chain these faults are part of the Craven fault system. The dolostone borders the faults and extends outward in varying distance into the adjacent limestone along joints and bedding planes. Dolomite crystals also occur along irregular cracks which cut the limestone (HATCHet al., 1938, p.193). The epigenetic dolostones of South Wales and red dolostones of Breedon, Leicestershire, England, include euhedral dolomite rhombs zoned with hematite. This hematite-rich dolomite is believed to have formed during Permian and Triassic time, when redbeds were deposited on a large scale (PARSONS, 1918, pp.255-257). This explanation, however, does not seem very compelling in light of the summary of HEWETT(1928; see below), which reveals a common association of increased iron content with epigenetic dolomite. Local epigenetic dolostones occur with lead-zinc ores along Jurassic faults, which cut the Bunter sandstones (Triassic) in the Eifel district, western Germany (PICARD,1954). The dolostone has not formed at the expense of original limestone here, but occurs in sandstone. In his summary of fault-related dolostone, HEWETT (1928, p.849) also noted that in many areas the dolomitized rocks contain a higher iron content, and locally a higher manganese content than the original limestone. This is not surprising in view of the geochemical similarities between magnesium, iron, and manganese; iron and manganese tend to occur with magnesium under a wide variety of circumstances. The depletion of strontium in the dolostones of Mill Creek, Oklahoma

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(G. M. Friedman, unpublished), is analogously understood in terms of the geochemical similarity between calcium and strontium in such crystal structures as calcite. Where sulfide minerals are associated with epigenetic dolomite, dolomitization of the limestones seems to have preceded precipitation of the sulfides. HEWETT (1928, p.861) concluded that little, if any, of the magnesium which was added during the process of dolomitization of limestone, is of direct magmatic origin. The isotopic data cited for the Cool Creek Limestone of Oklahoma (G. M. Friedman, unpublished) support an origin by hypersaline brines. CONCLUSION: A UNIFYING MODEL FOR DOLOMITIZATION

Three classes of dolostone have been recognized, which have been termed syngenetic, diagenetic, and epigenetic. The purpose of this classification has been for the orderly discussion of the subject and to facilitate descriptive study and genetic interpretation. Yet, it must be recognized that nature transcends these man-made barriers and that in regional studies of dolostones all three classes may occur together and may in fact be genetically related. This is especially true for syngenetic and diagenetic dolostones. The different classes of dolostone are mere variations on a theme: dolomite owes its origin to hypersaline brines, Dolomites, which are related to bacterial origin, are uncommon exceptions indicating that dolomite may form by other processes; however, it seems safe to conclude that all dolostone deposits found in the geologic record, other than those that are recycled, formed under evaporitic conditions. Hence all dolostones, whether syngenetic, diagenetic, or epigenetic, are the result of the action or reaction of hypersaline brines. Thus genetically it makes very little difference whether dolostones are formed in the depositional environment, such as those described in one of the preceding sections (syngenetic dolostone interfingering with limestone) for the Ordovician of New York State, U.S.A., or whether Late Pennsylvanian faults are responsible for dolomitization in earlier, Ordovician, limestones as described in the section on epigenetic dolostone. The solutions to which both of these divergent types of deposits owe their origin are hypersaline brines. Hypersalinity may result from ( I ) the concentration by evaporation of sea water, either in sea-marginal porous sediments or in the water mass itselF; (2) concentration of fresh water by evaporation, as in intermontane basins; and (3) subsurface processes not altogether understood, by which waters are concentrated by diffusion, membrane filtering, or other processes. The salinity (or chlorinity) values of the brines from which syngenetic dolomite is formed in the Persian Gulf (see Table 111),fall within the range of salinity (or chlorinity) values of subsurface waters which have not been diluted by fresh meteoric water (CHAVE, 1960). Hence, brines are available in the subsurface for the formation of both diagenetic and epigenetic dolostone.


335

Hypersalinity in sea-marginal environments and in intermontane basins may result from “capillary concentration” or “refluxion”, both of which were described in this chapter in the section on Recent and Pleistocene dolomite. In “capillary concentration” due to the excess of evaporation over rainfall, interstitial waters in the sediments transpire upward through the porous marginal sediments and evaporate at the sediment-air interface, a process similar to that under which caliche is formed. Dolomite is formed, by evaporation, at the surface and the concentration of the unevaporated wate: is increased. The fact that dolomite is found to form as caliche (FRIEDMAN, 1966; see section on “Miscellaneous non-marine dolomite”) indicates the reality of this process. At Salt Flat Graben, West Texas, dolomite formed at the bottom of the lake in stratigraphic layers, and as caliche around the basin margin. In “seepage refluxion”, a concept developed by R. H. KING(1947), SCRUTON (1953), and ADAMSand RHODES (1960), brines form, as in “capillary concentration”, in areas where evaporation exceeds precipitation plus run-off. Water is lost by evaporation, which lowers the water level on the shelf and increases the concentration and density of the water. The resulting heavy brine sinks and flows seaward down the sloping shelf. Surface currents tend to replenish the lost water by bringing low salinity water from the ocean in respect to hydrostatic head, while at depth oppositely directed currents as a result of density distribution maintain the seaward flow. If the return flow (reflux) of this brine to the sea is prevented by natural barriers, such as reefs or sills, it migrates to the lowest possible topographic depressions and seeps slowly through the underlying sediments, which are progressively dolomitized. The work at Bonaire, in the Netherlands Antilles (DEFFEYES et al., 1964), has greatly strengthened this King-Scruton-Adams-Rhodes theory. At Salt Flat Graben, discussed below, evidence is provided for refluxion to have been active in dolomite formation, but dolomite there formed directly from the brine rather than by the dolomitization of pre-existing bottom sediments. This observation requires a slight extension of the “refluxion” theory to apply both to dolomitization of bottom sediments as well as to formation of dolomite directly from the brine. Formation waters in subsurface sediments, which are responsible for the origin of much diagenetic and all epigenetic dolomite, are considered to be remnants of sea water trapped with the sediments at the time of their deposition. Postdepositional diagenetic changes of these waters led to their concentration and to the formation of brines with salinity values identical to those of dolomitizing brines that have been measured in depositional environments, such as those of the Persian Gulf. Chemical evidence suggests that some of these waters have migrated distances of several hundred miles driven by sediment compaction in the direction of least resistance (CHAVE,1960). Faults and fractures serve as escape hatches for the brines and dolomitization is taking place in the limestone within the fault and fracture zones and contiguous with them. Several lines of evidence independently lead to this conclusion: fault-related dolomitization was controlled by isotopically-

336


heavy brines (G. M. Friedman, unpublished; see section on “Epigenetic dolostone”), subsurface waters undiluted by fresh meteoric water have a salinity identical to that under which dolomitization takes place in the depositional environment, but yet magnesium is depleted in these brines. For some time, workers in the field of geochemistry of subsurface waters have noted that magnesium is removed from these brines in the form of dolomite (WHITE, 1957; VON ENGELHARDT, 1961). For dolomitization to take place, the Mg/Ca ratio of the brines must be increased from that of sea water to a ratio larger than that which would be in equilibrium with both calcite and dolomite (HALLA and RITTER, 1935). The increase of the Mg/Ca ratio of the brines was explained by DEFFEYES et al. (I 964) by the removal of calcium from the brine to form gypsum; BERNER (1965) concurred with this interpretation. An increased Mg/Ca ratio would favor the conversion of calcite or 1965, p.1298): aragonite to dolomite by the reaction (BERNER,

Mgaqz+

+ 2 CaC03 + CaMg (CO3)z + Ca&+

Berner suggested that dolomitization of the calcareous skeletal material of the Pacific atolls had formed by this reaction, and that once buried below the zone of brine reflux, the gypsum would be expected to redissolve in the surrounding interstitial water. This theory is in a sense strengthened by the observations at Salt Flat Graben, West Texas (FRIEDMAN, 1966; see section on Recent and Pleistocene dolomite), but requires modification with respect to the disposition of the sulfate. At Salt Flat Graben it was shown that the degradation of gypsum is genetically related to dolomite formation. Dolomite and gypsum are an antipathetic pair. Formation of dolomite requires: ( I ) brine concentration as indicated by enrichment of the heavy isotopes in dolomite; (2) the reduction of gypsum to HzS, iron sulfide, and native sulfur, with possible presence of calcite as a by-product coincident with dolomite formation; and (3) enrichment of the Mg/Ca ratio of the brine as a result of gypsum or aragonite precipitation. Native sulfur would be removed in solution. Where abundant dolomite is formed, gypsum is not an important phase but its degradation products are. In the Dead Sea, gypsum precipitates continuously, and quantitatively in proportion to temperature (NEEV, 1963). When the temperatures are sufficiently high to reduce the COz content in the water leading to isotopically-heavy water and disturbance of the chemical equilibrium, mass precipitation of aragonite takes place (FRIEDMAN and NEEV,1966). This is accompanied by a drastic decrease in HC03--, SO&-, and oxygen contents of the brine indicating that aragonite and gypsum are formed simultaneously. Yet, the S O P content soon returns to normal as gypsum crystals are reduced by bacterial action at depth to HzS, and near the surface the HzS of the brine is oxidized again to the sulfate ion. The HC03only gradually recovers after aragonite mass precipitation (NEEV,1964). Aragonite mass precipitation across the entire Dead Sea occurs during intense evaporation approximately every five years or so, in contrast to gypsum which is precipitated


337

continuously. Analogously, at Salt Flat Graben, gypsum must have been continuously precipitated. Although gypsum is precipitated all over the Dead Sea, it is preserved mainly along shore where an ample supply of oxygen exists. Below the wave-influenced zone, it is for the most part decomposed by bacteria; the H2S released reduces the Fez03 in the sediments and in suspension, thus blackening the sediment and producing dark laminae (NEEV,1964). Other workers (MOOREand HAYES,1958) have shown that gypsum develops along the margin, and black sediment derived from sulfate is found in the deeper part of the basin. Hence, the dolomite at Salt Flat Graben, with its black coloration due to iron sulfide and reduction products of gypsum must have formed at times when the lake was deeper and gypsum was reduced by bacterial degradation. This combination of inferences shows that lake level fluctuated with a periodic heavy influx of water, which is common in arid regions. Gypsum was formed by continuous precipitation, as in the Dead Sea, and when lake level was low it was deposited at the bottom of the lake. When the lake deepened by influx of water, rapid evaporation around the margin of the lake increased the concentration and specific gravity of the brine. This heavier brine sank to the deeper level of the lake where isotopically-heavy dolomite was formed and gypsum was reduced by bacterial processes. This process, involving dolomite formation from heavy brines and the reduction and partial or complete elimination of gypsum, must be common in areas where dolomite is formed in deeper water. Gypsum is involved in dolomite formation but it decomposes; hence, BERNER’S (1965) model for dolomitization in Pacific atolls, which requires that gypsum should be dissolved once it is buried below the zone of brine reflux, must be modified. In the Pacific atolls, the sulfate must have been reduced by baLteria1 processes within the zone of brine reflux. At Salt Flat Graben in West Texas, as already pointed out, both processes, “capillary concentration” and “refluxion”, were responsible for dolomite formation, the former along the margin of the basin and the latter in the center of the basin. It is interesting to speculate why dolomite was formed at Salt Flat Graben and aragonite in the Dead Sea. The data on the Dead Sea suggest that calcium which is removed by continuous gypsum precipitation must be replenished. NEEV(1964) has shown that the brine is progressively enriched in magnesium content. During the 1959 Dead Sea whiting, the calcium level of the sea fell, but three months later it was back to normal. Yet, the Mg/Ca ratio at any time during the two-year interval that measurements were made by Neev did not deviate significantly from that of sea water. At Salt Flat Graben the formation of dolomite suggests that the ratio was drastically increased. Whether aragonite was an intermediate product or not is not known. The evidence presented in the section on Recent and Pleistocene dolomite indicates that in some areas aragonite forms as an intermediate product, but no certainty exists whether this is always so or not. Epigenetic dolomitization is considered to be analogous to the process proposed for the formation of dolostone at Salt Flat Graben. Subsurface formation

338


water is depleted in both the so42-and Mg2+ion (VON ENGELHARDT, 1961), which indicates again that sulfates are involved in the formation of dolomite. During dolomitjzation the sulfate probably is reduced by bacteria and escapes as HZS. The common association of lead and zinc deposits with epigenetic dolomite suggests that the commonly accepted hydrothermal hypothesis to explain their origin requires reexamination. I n summary, the two main processes responsible for dolomite formation are: “capillary concentration” and “refluxion”. The former is the effective process in supratidal or intertidal zones or along the margins of basins, whereas the latter is operative under deeper water conditions. Caution must be exercised, therefore, in interpreting all syngenetic dolomite as supratidal or intertidal in origin. Dolomite is an evaporite mineral, and most dolomites, syngenetic, diagenetic, and epigenetic, are formed with the aid of brines.

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SKEATS, E. W., 1903. The chemical composition of limestones from upraised coral islands, with notes on their microscopical structures. Bull. Harvard Cull., Museum Comp. Zool., 62: 53-126. SKEATS, E. W., 1905. The chemical and mineralogical evidence as to the origin of the dolomites of southern Tyrol. Quart. J . Geol. Sue. London, 61: 97-141. SKEATS, E. W., 1918. The formation of dolomite and its bearing on the coral reef problem. Am. J. Sci., 45: 185-200. SKINNER, H. C. W., 1960. Formation of modern dolomite sediments in South Australian lagoons. Bull. Geol. Soc. Am., 71: 1976 (abstract). H. C. W., 1963. Precipitation of calcian dolomites and magnesian calcites in the southSKINNER, east of South Australia. Am. J. Sci., 261 : 449472. H. C. W., SKINNER, B. J. and RUBIN,M., 1963. Age and accumulation of dolomiteSKINNER, bearing carbonate sediments in South Australia. Science, 139: 335-336. SLOSS,L. L., 1953. The significance of evaporites. J. Sediment. Petrol., 23: 143-161. SPOTTS, J. H., 1952. X-ray Studies and Differential Tliermal Analyses of some Coastal Limestones and Associated Carbonates of Western Australia. Thesis, Univ. of Western Australia. Perth, W. Austr., 23 pp. STAUFFER, K. W., 1962. Quantitative petrographic study of Paleozoic carbonate rocks, Caballo Mountains, New Mexico. J. Sediment. Petrol., 32: 357-396, STEIDTMANN, E., 1911. Evolution of limestone and dolomite. J. Geol., 19: 323-345. STEIDTMANN, E., 1917. Origin of dolomite as disclosed by stains and other methods. Bull. Geol. Soc. Am., 28: 431450. STEIDTMANN, E., 1926. The origin of dolomites. In: W. H. TWENHOFEL (Editor), Treatise on Sedimentation. Williams and Wilkins, Baltimore, Md., pp.256-265. STEWART, F. H., 1951. The petrology of the evaporites of the Eskdale No.2 boring, east Yorkshire, 2. The Middle evaporite bed. Mineral. Mag., 29: 445475. STOUT,W., 1941. Dolomites and limestones of western Ohio. Ohio, Div. Geol. Surv., Bull., 42: 468 pp. STRAHAN, A,, 1901. On the passage of a seam of coal into a seam of dolomite. Quart. J . Geol. Soc. London, 58: 297-306. STRAKHOV, N. M., 1948. Principles of Historical Geology. Gosgeoltekhizdat, Moscow, 650 pp. (in Russian). Translation by J. KOLODNY and E. ROENTHAL, 1962. Israel Program Sci. Transl., Jerusalem, 689 pp. STRAKHOV, N. M., 1956. Concerning the types and genesis of dolomitic rocks. In: N. M. STRAKHOV (Editor), Types of Dolomitic Rocks and their Genesis-Tr. Geol. Inst., Akad. Naulc S.S.S.R., 4: 374 pp. (in Russian). STRAKHOV, N. M., 1958. Facts and hypotheses in the problem of the formation of dolomite rocks. Izv. Akad. Nauk S.S.S.R., Ser. Geol., 6: 3-22 (in Russian). STRUVE, W., 1952. Zum Problem der Eifler Dolomite. Senckenbergiana, 33: 135-145. SUFFEL, G. G., 1930. Dolomites of western Oklahoma. Oklahoma, Geol. Surv., Bull., 49: 155 pp. SUGDEN, W., 1963. The hydrology of the Persian Gulf and its significance in respect to evaporite deposition. Am. J. Sci., 261 : 741-755. SUMMERSON, C. H., 1966. Crystal molds in dolomite: their origin and environmental interpretation. J. Sediment, Petrol., 36: 221-224. SWINEFORD, A., LEONARD,A.B. and FRYE, J. C., 1958. Petrology of the Pliocene pisolitic limestone in the Great Plains. Kansas, Geol. Surv., Bull., 130: 98-1 16. TAFT,W. H., 1961. Authigenic dolomite in modern carbonate sediments along the southern coast of Florida. Science, 134: 561-562. TARR,W. A,, 1919. Contribution to the origin of dolomite. Bull. Geol. Soc. Am., 30: 114 (abstract). TATARSKIY, V. B., 1949. Distribution of rocks in which dolomite is replaced by calcite. Dokl. Akad. Nauk S.S.S.R., 69: 849-851 (in Russian). TEALL,J. J. H., 1903. On dedolomitization. Geol. Mag., 10: 513. TENNANT, C. B. and BERGER,R. W., 1957. X-ray determination of dolomiteecalcite ratio of a carbonate rock. Am. Mineralogist, 42: 23-39. TEODOROVICH, G. I., 1946. On the genesis of the dolomite of sedimentary deposits. Dokl. Akad. Nauk S.S.S.R., 53: 817-820 (in Russian).


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TEODOROVICH, G. I., 1959. A contribution on the origin of limestone and dolomite. Intern. Geol. Rev., l(3): 50-73 (in Russian). Translation from Tr. Petrol. Znst., Akad. Nauk S.S.S.R., 1955(5): 75-107. TEODOROVICH, G. I., 1961. On the origin of sedimentary calcareous-dolomitic rocks. Intern. Geol. Rev., 3: 373-384. THODE,H. G., KLEEREKOPER, H. and MCELCHERAN, D., 1951. Isotope fractionation in the bacterial reduction of sulfate. Research, 4: 581-582. THOMAS, G. E. and GLAISTER, R. P., 1960. Facies and porosity relationships in some Mississippian carbonate cycles of western Canada Basin. Bull. Am. Assoc. Petrol. Geologists, 44: 569-588. THORP,E. M., 1935. Calcareous shallow-water marine deposits of Florida and the Bahamas. Carnegie Inst. Wash., Publ., 452: 37-143. TOWSE,D. F., 1957. Petrology of Beaver Lodge Limestone reservoir, North Dakota. Bull. Am. Assoc. Petrol. Geologists, 41 : 2493-2507. TWENHOFEL, W. H., 1932. Treatise on Sedimentation. Williams and Wilkins, Baltimore, Md., 661 PP. UDDEN,J. A., 1924. Laminated anhydrite in Texas. Bull. Geol. Soc. Am., 35: 347-354. UDLUFT,H., 1929. Die Genese der flachenhaft verbreiteten Dolomite des mitteldevonischen Massenkalkes. Preuss. Geol. Landesanstalt, Jahrb., SO: 396436. UDLUFT,H., 1931. Ein neuer Beitrag zum Dolomitproblem. Deut. Geol. Ces., Z., 83: 1-13. VANTUYL,F. M., 1916. The origin of dolomite. Iowa, Geol. Surv., Ann. Rept., 25: 251422. VANTUYL,F. M., 1918. Depth of dolomitization. Science, 48: 350-352. VANTUYL,F. M. and STEIDTMANN, E., 1926. Dolomites (dolomite limestones). In: W. H. TWENHOFEL (Editor), Treatise on Sedimentation. Williams and Wilkins, Baltimore, Md., pp. 250-256. VONDER BORCH,C., 1965. The distribution and preliminary geochemistry of modern carbonate sediments of the Coorong area, South Australia. Geochim. Cosmochim. Acta, 29: 781-799, VONENGELHARDT, W., 1961. Zum Chemismus der Porenlosung der Sedimente. Bull. Uppsala Univ. Geol. Znst., 40: 189-204. VONMORLOT, A,, 1847. Uber die Dolomit und seine kiinstliche Darstellung aus Kalkstein. Haidinger Naturwiss. Abhandl., l : 305. WALDSCHMIDT, W. A., 1958. Halite as cementing mineral in sandstones. Bull. Am. Assoc. Petrol. Geologists, 42: 871-892. WALLACE, R. C., 1913. Pseudobrecciation in Ordovician limestones in Manitoba. J. Geol., 21: 402421. WALLACE, R. C., 1927. Recent work on dolomitization. Natl. Acad. Sci.-Natl. Res. Council, Rept. Comm. Sediment., pp.64-70. WEBER, J. N. and SMITH,F. G., 1961. Rapid determination of calcite-dolomite ratios in sedimentary rocks. J. Sediment. Petrol., 31: 130-131. WELLER,S., 1911. Are the fossils of the dolomites indicative of shallow, highly saline, and warm water seas? Bull. Geol. Soc. Am., 22: 227-231. WELLS,A. J., 1962. Recent dolomite i n the Persian Gulf. Nature, 194: 274-275. WEYNSCHENK, R., 1951. The problem of dolomite formation considered in the light of research on dolomites in the Sonnwend Mountains (Tirol). J. Sediment. Petrol., 21: 28-31. WHITE,D. E. 1957. Magmatic, connate and metamorphic waters. Bull. Geol. Soc. Am., 68: 1659-1682. WILCKINS,O., 1929. Die Dolomite der Eifel, Nieder-Rhein. Ges. Nut. Heilkunde (Bonn), Sitz. Ber., Nut. Abt., 1928: 1-62. WILLARD, B., 1961. Stratigraphy of the Cambrian sedimentary rocks of eastern Pennsylvania. Bull. Geol. Soc. Am., 72: 1765-1776. WILLMAN, H. B., 1943. High-purity dolomite in Illinois. Illinois, Geol. Surv., Rept. Invest., 90: 89 PP. WINDER, C. G., 1960. Paleoecological interpretation of Middle Ordovician stratigraphy in southern Ontario, Canada. Intern. Geol. Congv., 21st, Copenhagen, 1960, Rept. Session, Norden, 7: 18-27.

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WINTERER, E. L. and MURPHY,M. A., 1960. Silurian reef complex and associated facies, central Nevada. . I Geol., . 68: 117-139. WOPFNER,H. and TWIDALE,C. R., 1966. Geomorphological history of the Lake Eyre Basin. In: J. N. JENNINGSand J. A. MABBUTT(Editors), Australian Geomorphological Essays. Australian Natl. Univ. Press, Canberra, A.C.T., in press. YOUNG JR., F. B., 1943. Black River stratigraphy and faunas. Am. J . Sci., 241: 144-166,209-240. YOUNG, R. B., 1933. Conditions of deposition of the Dolomite Series. Trans. Geol. Soc., S. Afvica, 36: 121-135. YOUNG, R. B., 1935. A comparison of certain stromatolitic rocks in the Dolomite Series of South Africa with modern algal sediments in the Bahamas. Trans. Geol. Soc. S. Africa, 31: 153-162.

Chapter 7

CARBONATE OIL RESERVOIR ROCKS JOHN W. HARBAUGH

Stanford University, Stanford, Calif. ( U.S.A.)

SUMMARY

Half or more of the world’s petroleum is produced from carbonate reservoir rocks. The oil-reservoir characteristics of carbonate rocks are largely functions of porosity and relative permeability, which, in turn, have been affected by initial composition of the rocks and their subsequent history. Porosity of carbonate rocks may be arbitrarily divided into ( I ) primary porosity (formed during deposition), (2) secondary porosity (formed by solution, fracturing or other changes after deposition), and ( 3 ) sucrose dolomite porosity (resulting from replacement of calcite by dolomite). Primary porosity may in turn be subdivided into (a) framework porosity resulting from pores that remained as a result of the “sheltering” effect of rigid or loosely-aggregated frameworks, (b) mud porosity, consisting mostly of minute pores that remained in partly compacted carbonate mud that was subsequently lithified, and (c) sand porosity consisting of voids between sorted sand and gravel-sized carbonate particles. Most primary pores have been modified by solution (and cementation). Consequently, there is no sharp dividing line between primary porosity and secondary porosity resulting from solution. Sucrose dolomite porosity is important in many oil reservoirs. In these rocks, porosity and permeability have been strongly influenced by composition of the original carbonate sediment and the degree to which the rock has been replaced by sucrose dolomite. For example, in certain Devonian rocks in west Texas, which originally consisted of varying proportions of lime mud and of crinoid-stem fragments, the greatest porosity occurs in rocks that have been most highly dolomitized. Here, the percentage of dolomite tends to be greatest in rocks which originally contained about 45 % lime mud and 55 % crinoid-stem fragments. The performance of carbonate reservoirs depends to a substantial degree on shapes and dimensions of pores and their geometric arrangement with respect to each other. Under oil-reservoir conditions, pores in rocks are generally occupied by either water or oil. Ordinarily, the reservoir rock is water wet, that is, each rock grain is surrounded by a thin film of water, and oil is generally the non-wetting phase. Isolated oil globules ordinarily will not migrate through the rock because the interfacial tension between water and oil is so high that the globules will not

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pass through the throats of pore interconnections. Before the oil can move as a separate phase, the displacement pressure between the oil-water interface must exceed the entry pressure of the pore interconnections. The displacement pressure is influenced principally by buoyancy, whereas entry pressure depends on the interfacial tension between water and oil, and on pore geometry. The minimum height of an oil column necessary for buoyant rise through a water-wet carbonate rock thus partly depends on the diameters of throats of pores and diameters of the interiors of pores. The reservoir performance of carbonate rocks may be predicted by injecting mercury into cores from reservoirs. Mercury, a non-wetting fluid, is forced into the core sample under increasing pressure. A graph of the data, showing injection pressure, versus cumulative volume of mercury injected, is an effective guide to the conditions required for oil to move in the rock. Ancient depositional environments exert strong influence on carbonate deposits formed in them, and, in turn, have subsequent effect on oil-reservoir conditions in carbonate rocks. Many examples could be cited. In Mississippian carbonate reservoirs in southeastern Saskatchewan, oolitic and pseudo-oolitic limestones interpreted to have been formed through chemical precipitation in a barrier bank environment, serve Iocally as excellent, highly permeable oil reservoirs. In west Texas, the Pennsylvanian-Permian Horseshoe atoll is a horseshoe-shaped mass of limestone about 90miles across in an eastwest direction and 70 miles from north to south. It is interpreted to be analogous to modern reef atolls of the East Indies. In the Paradox Basin of southeastern Utah, limestone lenses composed largely of leaflike calcareous Algae serve as oil reservoirs. In Alberta, much oil is produced from Devonian rocks in which favorable reservoir conditions are closely associated with stromatoporoids and calcareous Algae. The geographic outlines of certain oil fields in Alberta, such as Redwater field, are essentially parallel to the trends of ancient organism communities. Thus, there is strong incentive to interpret ancient carbonate environments and organism communities, and to understand their effects on oil-reservoir properties.

INTRODUCTION

The importance of carbonate reservoir rocks is underscored by the fact that half or more of the world’s petroleum is produced from carbonate reservoir rocks. The greatest concentration of large oil fields in the world is in the Middle East, where most of the oil is produced from Mesozoic and Cenozoic carbonate reservoir rocks. Prolific oil production is obtained from Cretaceous limestones in western Venezuela and in many of the large oil fields in Mexico. In the United States, Cretaceous and Jurassic carbonates yield oil in parts of the Gulf Coastal Plain. Elsewhere in the United States, Paleozoic carbonate rocks yield oil in west Texas, in the Midcon-

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tinent fields of Kansas, Oklahoma, and north Texas, in parts of the Rocky Mountain region, and in many small fields in Michigan, Indiana, Ohio, Illinois and Kentucky. I n western Canada, dolomitized Devonian reefs contain much oil and gas, and Mississippian limestones in Saskatchewan and adjacent North Dakota also form prolific reservoir rocks. Obviously, oil in carbonate reservoir rocks has great economic significance and there is stimulus to understand the properties of carbonate rocks in their role as oil reservoirs.

POROSITY AND PERMEABILITY DEFINITIONS

Porosity

An essential feature of any petroleum reservoir rock is porosiry. The rock must not only possess pores to contain the oil or gas, but the pores must also be large enough and sufficiently interconnected to permit oil or gas to move through them. In other words, the rock must be sufficiently permeable with respect to petroleum if it is to serve as a reservoir rock. Porosity is defined as the ratio of pore space to total volume, and is commonly expressed in percent: percent porosity

=

pore volume bulk volume

*

100

The ratio of total volume of pores in rock to bulk volume is absolute porosity, whereas porosity as ordinarily used in reservoir studies is the ratio of interconnected pore spaces to total bulk volume, and is termed eaective porosity. Effective porosity may be calculated by measuring the proportion of mercury or a gas entering the rock at a specified pressure, as for example, at 1,000 lb./square inch. Effective porosity depends in part on the conditions under which it is determined. Permeability

Permeability is defined as the ability of a porous medium to transmit fluids. The conventional unit of measurement is the millidarcy, which is one thousandth of a darcy. The darcy has been standardized by the American Petroleum Institute as follows: “A porous medium has a permeability of 1 darcy when a single phase fluid of 1 centipoise viscosity that completely fills the voids of the medium will flow through it under conditions of viscous flow at a rate of 1 cm/sec/cm2 of cross-sectional area under a pressure or equivalent hydraulic gradient of one atmosphere (76.0 cm of mercury)/cm”. Permeability is related to effective porosity, and in addition, is closely related to the size and shape of pores and to the degree of crookedness or tortuosity of the pore patterns. When more than one fluid is present in the pores of a rock, as is commonly the case in an oil reservoir, the ability

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J. W. HARBAUGH Percent

0

25

Percent

oil s a t u r a t i o n

50 water

75 saturotion

I00

Fig.l. Curves showing generalized relative permeability to oil and to water of typical permeable and CHILINGAR, 1961, p.119.) carbonate reservoir rocks. (After SINNOKROT

of the rock to conduct a particular fluid in the presence of other fluids is termed its effective permeability to that fluid. The ratio of the effective permeability to a given fluid at partial saturation to the permeability at 100 % saturation is termed relative permeability. Examples of porosity and permeability values in carbonate rocks are given by MCCOMAS (1963), and generalized curves showing relative permeability to water and to oil in carbonate rocks are shown in Fig.1.

PRIMARY POROSITY

It is difficult to devise a simple, broadly inclusive classification system for porosity in carbonate rocks. The difficulty arises because porosity is affected by many factors, which include mineralogy, crystal fabric, texture, structural features, and events that took place during deposition, diagenesis and after lithification. Many authors have employed a system in which porosity is classed simply as primary or secondary. MURRAY (1960) has expanded this system to include three general categories: . _ _ _ _ . ~

Fig.2. Example of primary pores occurring within rigid framework of Fuvosites coral. Although skeleton of coral has been completely dolomitized, original voids have been preserved. Length of specimen is about 2.5 inches. (After MURRAY, 1964, p.391.) Reproduced with permission from Approaches to Paleoecology.


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( I ) primary porosity, (2) secondary porosity, and (3) sucrose dolomite porosity. Primary porosity consists of those pore spaces formed during deposition and which have been modified only slightly by subsequent changes, such as compaction, solution and precipitation. There is no sharp distinction between primary porosity and secondary porosity; however, carbonate reservoir rocks generally possess both types. Primary pores may occur within and between fossils and fragments of fossils, between pellets, pisolites and ooIites, between crystals, and at places where shrinkage of unconsolidated carbonate mud occurred during or shortly after deposition. Oil reservoirs whose porosity is mainly of primary origin include oolitic (or

Fig.3. Example of limestone containing both priniery porosity and secondary porosity. Many pores (dark) are closely associated with fragments of leaflike Algae which probably served as miniature umbrellas, allowing primary voids to persist adjacent to their sheltered lower surfaces. Selective leaching by solution has enlarged primary pores and created new pores. Reflected light, polished core. From Pennsylvanian Ismay zone, Paradox Formation, Ismay field, southeastern Utah and southwestern Colorado. Width of that part of specimen shown in photograph is 6 mrn. (After CHOQUETTE and TRAUT,1963, p.171.)


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Fig.4. Primary and secondary pores (light) associated with leaflike calcareous Algae, Ismay zone, Ismay field. Transmitted, unpolarized light, thin section. Width of that part of specimen shown in photomicrograph is 5.5 mrn. (After CHOQUETTE and TRAUT,1963, p.171.)

pellet) limestones and limestone coquinas. Examples are provided by the southern oil fields of Mexico, which yield oil contained in pores in shell fragments, corals, and rudistids, and by Magnolia field and certain other fields in southern Arkansas that produce from the oolitic Jurassic Smackover Limestone, which is exceptionally porous and permeable. It is convenient to classify primary porosity according to three subtypes: (a) framework porosity, (b) porosity in lithified carbonate mud, and (c) porosity in carbonate sand. Framework porosity

Framework porosity is formed when open (water-filled) spaces are left in a depositional framework. Frameworks may be rigid, or may be more or less loosely aggregated when formed. Rigid frameworks are provided by corals (Fig.2), hydrocorallines, some sponges, some calcareous Algae, oysters, and perhaps some bryozoans. More or less loose frameworks are provided by aggregations of loose shells and other skeletal material. Excellent oil reservoirs, for example, occur in some regions in limestones consisting largely of fragments of leaflike calcareous Algae (Fig.3, 4) that formed a loose framework initially (MURRAY,1960, p.66; CHOQUETTE and TRAUT,1963; IRWIN,1963, p.145).

Fig.5. Primary pores (dark) in sorted, calcareous sand. Pores occur between the rock particles. Calcite cement has filled much of original pore space. Porosity is 16.5 % and permeability is 136 millidarcys. Length of part of specimen shown in 1960, photomicrograph is about 3.5 mm. Mississippian Charles Formation, Midale field, Saskatchewan. (After MURRAY, p.64.) Reproduced by permission of J. Sediment. Petrol.


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Mud porosity Some primary porosity occurs in carbonate rocks that were deposited as carbonate muds. The pores are generally so small as to be imperceptible at ordinary magnifications. The pores were, of course, initially water filled, and presumably remained so as consolidation of the mud took place. The pores are generally so small and permeability so low that the rock is excluded from serving as a commercial oil reservoir rock. Larger pores and vugs produced by shrinkage of partly consolidated lime mud may be important locally, however. Sand porosity

Carbonate sediments consisting of sorted sand- or gravel-sized particles may contain appreciable primary porosity. Carbonate sands consisting of oolites or pellets, and sorted skeletal fragments provide good examples (Fig.5). Carbonate sands may be formed by current transport of sand-sized particles, and by competitive sedimentation in which sand-sized particles are formed in situ with silt- and mud-sized particles. The relative rates of formation of the different particle sizes affect the porosity of the sediment, because the porosity is more or less inversely related to the proportions of carbonate mud present in the matrix. The proportion of mud is partly a function of how fast sand-sized material was deposited as compared with the rate of mud deposition. Slow sedimentation of mud relative to sand probably has generally resulted in initially high intergranular porosity. Alternatively, lime mud may be removed from coarser material by current action (DUNHAM, 1962).

Cementation Primary pore spaces tend to be filled by cementation, or by cementation in association with replacement. Cementing and replacing materials in ancient carbonate rocks include calcite, quartz, and anhydrite. Originally much of the calcite cement, however, was probably aragonite that subsequently was transformed to calcite. The extent to which primary pores are filled by cement has a large effect on the porosity that remains.

SECONDARY POROSITY

Secondary pores are formed principally by enlargement of pre-existing voids by solution. Pre-existing voids represent either primary porosity or have been produced by fracturing. It is difficult to draw a sharp distinction between primary pores and secondary voids. Good examples of primary pores that have been at least slightly modified by solution were shown by THOMAS (1962, p.195) and by STOUT

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(1964) and include interskeletal vug porosity, interlump porosity, interoolitic porosity and interpellet porosity (Fig.6). Examples of pores that show extensive solution include: intracoralline vug porosity, interalgal vug porosity, pores resulting from selective solution of shells, selective removal of calcite particles in rock that otherwise consists of dolomite, pores within and between leaflike calcareous Algae and other skeletal material (Fig.7), and large-scale enlargements along fractures and solution of void-filling calcite between mud-lump fragments (HARBAUGH,1960, p.217, 220).

Solution Percolating ground waters of meteoric origin contain dissolved carbon dioxide which forms carbonic acid. In addition, weak organic acids are present. These acids tend to dissolve calcareous rocks if the movement of the water is sufficient

Fig.6. Pelletoid limestone in which primary pores have been modified by subsequent solution. Porosity is 15.1 % and permeability is 63 millidarcys. Pores are dark. Mississippian Mission Canyon Formation, North Dakota. Thin section; crossed nicols. Width of that part of specimen shown in photomicrograph is 10 mm. (After STOUT,1964, p.335.) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.


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Fig.7. Secondary porosity produced by solution of fragments of leaflike Algae and other skeletal material. Pores are dark. From Ismay zone, Ismay field. Thin section; crossed nicols. Width of that part of specimen shown in photomicrograph is 15 mni. (After CHOQUETTE and TRAUT,1963, p. 179.)

to keep the water undersaturated. Much of the solution is presumed to take place above the water table, where ground water tends to be slightly acid. Below the water table, on the other hand, solutions are commonly weakly alkaline. Under normal conditions they are saturated with respect to calcium carbonate and are capable of precipitating rather than dissolving. Many carbonate rocks show evidence of both solution and precipitation. It is possible that solution and precipitation may take place simultaneously, the water being slightly undersaturated at one moment, with power to dissolve, and slightly oversaturated at another moment, resulting in precipitation. If the water is undersaturated, the process of solution goes on as long as the water moves through the rock. Certain crystal grains are preferentially dissolved, forming new channels and exposing additional mineral grains to the dissolving solutions. Other conditions being equal, calcite tends to dissolve about 24 times more readily than dolomite. WEYL(1958) pointed out that the rate at which solutions become saturated with respect to calcium carbonate is critical. As fresh water enters the rock, it will dissolve calcium carbonate until the water is saturated, and the amount of alteration by solution will depend on how far the water enters the rock before it is saturated. Once saturation is reached, the rock can still be altered as changes in

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J. W. HARBAUGH

solubility occur as a result of changes in temperature, pressure and chemical environment. The rate at which the water becomes saturated is a function of the velocity of the water and the size of pores. Thus the solution alteration of calcitic limestone depends on the rate of transport of solute away from the crystal-liquid interface, where the solution is always saturated. Much work needs to be done on the degree of departure of solutions from saturation with respect to calcite and other carbonate minerals. An instrument useful for this purpose is termed a carbonate saturometer (WEYL,1961), and is capable of sensing very slight changes in saturation by use of a pH-sensitive electrode and a reference electrode.

Multiple generations Many carbonate rocks have experienced a complex history in which intervals of pore development alternate with those of pore filling. The history of the Mississippian Fredonia oolite, an important oil-reservoir rock in southern Illinois, provides an example of the sequence of events that some carbonate rocks have undergone. Petrographic study by GRAFand LAMAR (1950) indicated that six generations of calcite precipitation and two intervals of calcite solution have taken place. The first interval of solution is interpreted to have come after precipitation of three generations of calcite, but only the third generation of calcite was attacked by solution during this first interval of solution. Two more generations of calcite were formed after the first interval of solution. In the second solution interval, all five previously formed generations of calcite were attacked by solution. The sixth generation of calcite, consisting of veinlets, postdates the second interval of solution and the other calcite generations. Fractures, bedding planes and unconforniities

The effect of joints, fractures and bedding planes on development of porosity in carbonate rocks was treated at length by WALDSCHMIDT et al. (1956), and earlier by HOHLT(1948, pp.17-18). The principal means by which ground water obtains access to carbonate rocks is via joints, which are generally extensive in carbonate rocks because of their brittleness. Initial solution is more or less restricted to joints that are essentially vertical. Bedding surfaces may also conduct water in some cases and, in association with vertical or steeply dipping joints, form a complex network of conduits for ground water. A very close relationship between meteoric waters and secondary porosity is indicated by the proximity of solution porosity to unconformities which represent old erosion surfaces. In some regions containing thick carbonate sections, conditions favorable for oil accumulation occur only immediately below unconformities.


361

SUCROSE DOLOMITE POROSITY

The dolomitization process Sucrose or “sugary” doloniite is an important type of oil-reservoir rock (Fig.8, 9). The origin of dolomite and the reason for its greater porosity constitutes an important field of research in itself. It has long been known that the molecular replacement of limestone by dolomite would result in a volume shrinkage of 12-13 % because dolomite occupies less space (is more dense) than calcite. The replacement is symbolized by the equation:

2 CaC03

+ MgClz +. CaMg(C03)~+ CaClz

Thus, if replacement of calcite by dolomite occurs strictly on a molecule-by-molecule basis in a rock, and assuming no compaction, a substantial increase in porosity must result. There is, however, considerable evidence that replacement is, in some cases, on a volume-by-volume basis (LINDGREN, 1912) rather than molecular basis.

Fig.8. Microcrystalline dolomite displaying very favorable reservoir characteristics. See Fig.20 for estimate of reservoir performance based on mercury capillary pressure curve. Pores are dark. Reservoir rock in Mississippian Mission Canyon Formation, North Dakota. Porosity is 25.6 % and permeability is 53 millidarcys. Thin section; crossed nicols. Width of the part of specimen shown in photomicrograph is 7 mm. (After STOUT,1964, p.334.) Reproduced by permission of Bull. A m . Assoc. Petrol. Geologists.

z

W

P

Fig.9. Sucrose dolomite containing euhedral dolomite crystals that are 50-75jl in dimensions. Pore spaces (dark) occur between the crystals. Porosity is 11.0% and permeability is 16.5 millidarcys. Length of the part of specimen shown in photomicrograph is about 1.6 mm. Ordovician Red River Formation, Gas City field, Montana. (After MURRAY,1960, P . 70.) Reproduced by permission of J . Serlinieiit. Petrol.

C

2

363


Replacement on a strict volume-by-volume basis, of course, would not cause any change in porosity. The dimensions of crystals in sucrose dolomite vary from less than 5p to more than loop, but crystals in the 25-50p range are most common in oil reservoir rocks (MURRAY,1960, p.67). Effect of original sediment conzposition

The dolomitizing process is markedly affected by the composition of the original carbonate sediment. LUCIA(1962) has shown that porosity in crinoid-rich Devonian carbonate rocks in west Texas that have been dolomitized is markedly affected by the relative proportions of lime mud and crinoid material in the original sediment. Sediments remaining as limestones originally contained high proportions of crinoidal material and less than 35% of lime mud. Sediments that were dolomitized, however, contained relatively high proportions of lime mud and less

u

Percent m a t e r i a l o t h e r than lime mud (mostly crinoid fragments) 30

100 80 60 I I

25 c

C

a L

20

.1

T r a n s i t ion

zone

&

3

15

.-c L1

-

I

20 I

Dolomite

F? 0

40

10

t a,

u L

2

5

0

20

Percent

40

60

80

I 0

o r i g i n a l lime mud

Fig.10. Generalized curve showing relationship of present porosity to original proportions of lime mud and material other than lime mud in Devonian crinoidal limestones in west Texas. (After LUCIA,1962, p.864.) Reproduced in modified form by permission of J. Sediment. Petrol.

364

J. W. HARBAUGH

crinoidal material. The dolomitized sediments may be interpreted as being initially mud supported, whereas the limestones were crinoid-grain supported. In the Devonian sediments that are limestones, the best porosity is developed in those rocks that originally contained between 5-20 7; lime mud (left side of Fig. 10). Rocks that originally contained these proportions of lime mud are interpreted to have had a supportingframework of crinoid fragments. The presence of moderate amounts of lime mud is interpreted by LUCIA(1962) to have inhibited the growth of rim cement around the crinoid particles. Where lime mud was originally present in only very small proportions, primary pores were largely filled by rim cement. As the proportion of original lime mud increased past 20 %, the porosity decreased. This decrease probably reflects compaction of the lime mud, inasmuch as the mud was called upon to support more and more of the weight of overlying sediment while the meshwork ofcrinoid particles provided progressively less and less support.

Fig.11. Dolomite interpreted to have originally consisted of 50-60% of crinoid material. Lightappearing patches identify remnants of crinoid fragments. Specimen from Andrcws South Devonian field, Andrews County, Texas. Porosity is 28 % and pcrmeability is 200 millidarcys. Thin section viewed under plane polarized light. Length of the part of spccimcn shown in photomicrograph is 12 mm. (After LUCIA, 1962.) Reproduced by permission of J . S e d h e n t . Petrol.


365

The relationships between porosity in the dolomitized sediments and proportion of crinoid material are more complex. WEYL(1960) suggested that if the carbonate involved in dolomitization has a local source, then the dolomite formed will occupy a smaller volume than the replaced calcite because dolomite is more dense than calcite. This difference in volume will be compensated b y compaction of the sediment until the time that dolomite crystals form an effective supporting framework. Crinoid fragments would also tend to provide a supporting framework. Up to a point (right side of Fig.lO), the crinoid material would act to reduce the porosity developed in the dolomitizing process because the crinoid fragments would occupy some of the space that otherwise would be occupied by lime mud with porosity. Beyond this point, as the proportion of crinoid material increases, the porosity would increase because the crinoid fragments would tend to form a supporting framework and would inhibit compaction and permit pore space formed in the dolomitizing process to be preserved. An example of a porous dolomite in which crinoid fragments formed a supporting framework is shown in Fig.11. Dolomite- porosity-permeability relationships

In general, porosity varies with the intensity of dolomitization. For example, this is shown by MURRAY (1960, pp.67-76) who determined the porosity and dolomite content in cores of the “Midale beds” in the Mississippian Charles Formation in the Midale field, Saskatchewan. Here dolomite forms 3-90% of the solid part of the rock. Where the dolomite percentage is low, dolomite occurs as scattered, perfect rhombs that replace pre-existing material. As the proportion of dolomite increases, the original material is increasingly obscured, until at 50 % dolomite, only small hazy patches of the original material remain. A plot of porosity versus dolomite content of the “Midale beds” (Fig.12) reveals only slight decrease in porosity as dolomite increases up to about 50% dolomite. Above 50 % dolomite, however, porosity increases rapidly as dolomite increases. MURRAY (1960, pp.7 1-72) suggested that these dolomite-porosity relationships may be an effect of either, or both: ( I ) transportation of carbonate and magnesium in solution from a distant source to the place where dolomite rhombs have formed, and (2) derivation of the carbonate very locally within the rock. Murray also suggested that if dolomite crystals have formed from material brought in from a distant source outside the rock, the porosity should decrease because dolomite rhombs produced in this manner occupy space that was once occupied both by calcite grains and associated fine intergranular porosity. If, on the other hand, carbonate used for development of dolomite was derived locally from within the rock, addition of magnesium from outside the rock was necessary and excess calcium ions must have been removed. Calcite must be dissolved within the rock to provide the necessary carbonate, which results in the dolomitized rock, on

366

I. W. HARBAUGH

the whole, becoming more porous than before. Statistical studies of the relationships between calcium-magnesium ratio and porosity are needed to resolve the problem (CHILINGAR, 1956). The observation that dolomite generally occurs as a replacement (MURRAY, 1960, pp.73-74) suggested that a local source for carbonate is likely and that dolomite growth took place simultaneously as calcite was dissolved. Assuming a local, within-rock origin of carbonate, the rather slight loss of porosity with increasing dolomite in limestones with less than 50 % dolomite may reflect the effects of pore closure by compaction. Above 50 % dolomite, however, the dolomite crystals have been interpreted to have formed a space-supporting framework that resisted further comraction and inhibited further loss of porosity. As a result, rocks containing more than 50 % dolomite commonly display an increase in porosity as the proportion of dolomite rises. WEYL(1960) discussed the quantitative chemical aspects of porosity development through dolomitization. POWERS (1962) showed that the proportions of dolomite in highly doloniitized Upper Jurassic carbonate reservoir rocks in Arabia also affect porosity and permeability. Here the relationships between the proportions of dolomite and other components are somewhat at variance with those in other regions. Powers showed that there is an almost perfect inverse relationship between dolomite and sandand gravel-sized particles. As the percentage of sand and coarser-than-sand par! 00

/

2oi 00-

a, c

.-

E

d o I om i t e

0 60-

0

U

c

40-

\fTrend

below

I

-

50 p e r c e n t dolomite

0 0

10

20

Percent

30 porosity

40

Fig.12. Relationship of porosity to percentage of dolomite in “Midale beds”, Mississippian Charles Formation, Midale field, Saskatchewan. Data points on which curves are based show considerable scatter. (After MURRAY,1960, p.72.) Reproduced in modified form by permission of J. Sediment. Petrol.

-


367

10000

75

85 Percent

90

95

I00

dolomite

Fig.13. Relationship of porosity and permeability to percent of dolomite in Arabian Upper Jurassic carbonate oil reservoir rocks containing more than 75 % dolomite. (After POWERS, 1962, p.187.) Reproduced by permission of Am. Assoc. Petrol. Geologists.

ticles increases, dolomite percentage decreases. In rocks containing less than 50 % dolomite, the proportions of dolomite and lithified lime mud roughly parallel each other. Where the proportion of dolomite in the Upper Jurassic carbonate rocks is below 75 %, it seems to make little difference, as far as reservoir properties are concerned, whether the rocks are strongly dolomitized or not. Where dolomite exceeds 75 %, however, it exerts critical control (Fig.13). At about 77 % dolomite, the network of dolomite crystals begins to open up and effective intercrystalline porosity develops, reaching a maximum at about 80 % dolomite. At this point the rocks have an average porosity of about 19 % and an average permeability of about 300 millidarcys. Progressive increase above 80 % dolomite, however, is accompanied by relatively uniform decline in porosity and permeability. At 95 % dolomite, the rock is essentially impermeable. POWERS (1962, pp. 184-1 86) interpreted the dolomite-porosity-permeability relationships in the Arabian Upper Jurassic rocks as follows: “During the early stages of dolomitization, discrete crystals formed at centers scattered nearly uniformly through the fine-grained part of the rock. The dolomite rhombs that formed

368

J. W. HARBAUGH

are relatively uniform in size (0.1-0.2 mm). As dolomitization proceeded, all lime mud was replaced before sand-sized particles were attacked. Replacement of fine particles having some intergranular porosity by larger dolomite rhombs resulted in some decrease of porosity and permeability, although oil reservoir properties are not markedly affected. As dolomitization approached 75 %,most pores were filled, and the result is very low porosity and permeability. When the dolomite reached 80 %, however, virtually all traces of the original sediment were obliterated and the replacing dolomite crystals formed an open network. Following this stage, depending on circumstances, the rock either ( I ) remained unaltered, (2) underwent solution, resulting in an increase in porosity, or (3) received more dolomite, resulting i n a decrease in porosity and permeability.”

Fig.14. Pores consisting of molds formed by selective removal of crinoid stem fragments. Width of specimen is 3 inches. (After MURRAY,1964, p.394.) Reproduced with permission from Approaches to Paleoeco1og.v.


369

Fig.15. Plastic cast of pores in Arbuckle Limestone, Silica field, Kansas. Width of that part of specimen corresponding to short dimension of photograph is 7 mm. (After IMBTand ELLISON, 1946.)

370

J. W. HARBAUGH

Fig.16. Plastic cast of pores in Arbuckle Limestone, Silica field, Kansas. Note elongate, threadlike pores in right central part and upper left part of photograph. Width of that part of speci1946.) men corresponding to short dimension of photograph is 7 mm. (After IMBT and ELLISON,


37 1

Selective removal of calcite Selective removal of material resistant to dolomitization may result in development of pores. MURRAY (1964, pp.393-394) illustrated (Fig.14) the development of porosity in a dolomite in which the pores are molds of crinoid-stem fragments. The calcite crinoid material was not converted to dolomite and was removed later by solution, the dolomite remaining intact.

PORE GEOMETRY

The complex and varied shapes of pores in carbonate rocks are well illustrated by plastic casts of pores (IMBTand ELLISON,1946). Casts shown in Fig. 15 and 16 were made by impregnating the pores in a rock specimen with liquid plastic, allowing the plastic to become firm upon setting, and then dissolving the rock away with acid, the plastic being unaffected by acid (Nuss and WHITING,1947). Fig. 15 illustrates pores occurring in Cambro-Ordovician Arbuckle Limestone in the Silica field, Kansas. The pores form a fine sponge-like complex which reflects the influence of rounded pellets. Fig.16 illustrates pores in another specimen of Arbuckle Limestone from the Silica field. The pores form a fine honeycomblike structure, although some pores consist of elongate, threadlike tubes.

Fig.17. Diagram illustrating difference between pore interconnections. I. Narrow, elongate interconnection defined as a canal. 11. Constriction defined as a throat. (After ASCHENBRENNER and ACHAUER, 1960.) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

372

J. W. HARBAUGH

Measurement The performance of carbonate reservoirs depends to a substantial degree on shapes, dimensions, and arrangement of pores. Pores may be classified according to the shapes and kinds of interconnections (ASCHENBRENNER and ACHAUER,1960). Connections that are relatively elongate (ten or more times longer than wide) may be termed canals (Fig. 17A), whereas connections that represent less elongate constrictions may be termed throats (Fig. 17B). Pore size can be defined as the diameter of the largest sphere that can be fitted inside a pore, and throat size as the diameter of the largest sphere that will pass through the throat of a pore (SCHEIDEGGER, 1957). ASCHENBRENNER and ACHAUER (1960) tabulated 100 pore size and throat size measurements in 40 thin sections of 1.00 ln L

a,

c

a,

E -

0.50

.-

E

C ._

Y

0

a

.-c

2

n .-

0.IC

L

V

ln

C .-

32 0.05 Y .0

+ ln

a,

P O

-

.+

0 L

W

c

W

E D

.-0 0.01

I

2

5

10

20 30405060 70 80

Cumulative

90 95

98 99

frequency, in percent

Fig.18. Graph showing frequency distribution of pore sizes in Paleozoic limestones in Rocky Mountain and Williston Basin regions. Logarithmic vertical scale represents diameter of largest circle inscribed within pore, whereas horizontal probability scale represents cumulative percent. and ACHAUER, 1960.) Reproduced Individual data points are not shown. (After ASCHENBRENNER by permission o f Bull. Am. Assoc. Petrol. Geologists,

373


2 a,

c

a,

E ._ -

.-

E

.-c

/

c“ 0 ._ c

/’

a, u

/ ’

c c

0

Y

/’

a,

c

.-c

E0

a

L

0

f D ._ 3

E,

.E _ c ._

3 noooi

I

2

5

10

20 3040506070 80

Cumulative

frequency,

in

90 95

98 99

percent

Fig.19. Graph showing frequency distribution of pore interconnection widths measured in PaIeozoic limestones in Rocky Mountain and Williston Basin regions. Logarithmic vertical scale represents minimum width of pore interconnections (throat), whereas horizontal probability scale represents cumulative percent. Individual data points are not shown. (After ASCHENBRENNER and ACHAUER, 1960.) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

Paleozoic limestones in the Rocky Mountain and Williston Basin regions, and found that both pore sizes and throat sizes are essentially lognormally distributed. When the logarithms of pore sizes and throat sizes were plotted as cumulative percentages, the distribution of points formed a line that is virtually straight (Fig.18, 19). It should be pointed out, however, that measurements based on two-dimensional thin sections may not always effectively portray pore systems that exist in threedimensional space. For example, definition of pore and throat sizes in terms of a sphere that just fits in or passes through is somewhat meaningless in case the pore has a tabular shape. Nevertheless, the quantitative approach of Aschenbrenner and Achauer has much to commend it. Permeability and yore geometry

ASCHENBRENNER and CHILINGAR (1960) brought the work on carbonate pore geometry of the Russian geologist, Teodorovich, to the attention of English-speaking readers. TEODOROVICH (1943, 1949) proposed that the permeability of carbonates could be calculated by assigning numerical values to empirical coefficients and then

374

J. W. HARBAUGH

multiplying the coefficients together to obtain permeability in millidarcys. The coefficients are estimated by visual examination of thin sections, and pertain to (a) type of pore space, (b) percent porosity, (c) size of pores, and ( d ) pore shape. In spite of the subjectiveness of this method, permeabilities estimated in this manner were shown to be surprisingly close to permeabilities measured by conventional means of fluid injection in core samples, the average discrepancy between the two methods being about 10 %. CHILINGAR (1957) illustrated some of Teodorovich’s pore space types, as did ASCHENBRENNER and ACHAUER (1960).

Movement of oil and water Under oil reservoir conditions, pores in rocks are generally occupied by two fluids, water and oil. Movement and occurrence of these fluids are controlled by welldefined physical laws which permit their behavior to be predicted (ARPS,1964). Ordinarily the reservoir rock is water wet, that is, each rock grain is surrounded by a thin film of water. Since oil is commonly the non-wetting phase, isolated oil globules ordinarily will not migrate through the rock because the high interfacial tension between water and oil prohibits the globules from passing through the throats of pore interconnections. Before oil can move as a separate phase through water-wet reservoir rocks, the displacement pressure between the oil-water interface must exceed the entry pressure of the pore interconnections. The displacement pressure is influenced by buoyancy (and by hydrodynamic forces if they exist), whereas the entry pressure depends on the interfacial tension between water and oil, as well as the sizes and shapes of pores. Upward migration by buoyancy results when enough oil globules coalesce to form an oil column that is continuous over a certain vertical distance (termed critical height) so that a sufficient pressure gradient exists between the uppermost leading edge of the oil phase and the lowermost rear edge. The critical height for buoyant rise (ASCHENBRENNER and ACHAUER, 1960) may be calculated from the equation:

h=

2T (l/r-l/R) g (@a-eo)

where h = critical height, in centimeters; T = interfacial tension between oil and water in dynesicentimeter; r = radius of sphere which can just pass through the throat of a pore, in centimeters; R = radius of the largest sphere that will fit into the pore, in centimeters; g = gravitational constant (assumed to be 980.7 cm/secZ); ew= density of water in g/cm3; and e o = density of oil in g/cm3.

Mercury capillary pressure curves The performance of carbonate rocks under reservoir conditions may be predicted 1948). Mercury, a non-wetting with mercury capillary pressure curves (PURCELL,

375


fluid, is forced into the rock sample under increasing pressure. Because the pressure required to inject mercury into a pore is greatly affected by throat diameters, the graph of cumulative mercury volume injected, versus injection pressure, is related to the distribution of throat sizes in the rock. The system of largest interconnecting pores is entered at the lowest pressures, and as the pressure increases, mercury enters progressively smaller pores. Because oil is ordinarily the non-wetting fluid in an oil-water system, there is a close relationship between the predicted ratio of oil to water in a reservoir rock at a particular height above the free-water level, and the volume of mercury injected into a specimen of the reservoir rock at a pressure equivalent to height above the free-water level (MURRAY, 1960, p.60). Two examples of predicted reservoir performance based on mercury capillary pressure curves of carbonate rocks (STOUT,1964) are given below. Each graph (Fig.20, 21) represents a plot of percent water saturation (or conversely, percent oil saturation) versus height of an oil column. The height of the oil column is proportional to the pressure gradient due to buoyancy, and has been calculated for the reservoir fluids under reservoir conditions. Toward the bottom of each curve, the total porosity is 100% saturated with water. Little or no oil enters the rock until the displacement pressure is attained, or in other words, until the critical height of the oil column is reached. Above the displacement pressure, the percent Percent oil saturation 100

80

60

40

20

0

C

E

3 0 V

.--

0

L

water saturation

L-

0

Displacement pressure

Total p o r o s i t y

Percent

water

-A2

c

a,

al

LL

saturotion

Fig.20. Mercury capillary pressure curve of crystalline dolomite forming a good reservoir rock. Porosity is 25.6 % and permeability is 53 millidarcys. Mission Canyon Formation, North Dakota. See Fig.8 for photomicrograph of rock. (After STOUT,1964, p.334.) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

376

J. W. HARBAUGH

Percent oil s a t u r o t i o n 100

80

60

40

20

0

512 256 128

64

t32

1

0

I

I

,

20

40

60

Percent

water

=S

E

3

I

80

100

saturotion

Fig.21. Mercury capillary pressure curve of fine-grained dolomite that would not serve as a reservoir rock. Porosity is 17.7 ”/, and permeability is 0.1 millidarcys. Mission Canyon Formation, 1964, p.336 )Reproduced North Dakota. See Fig.22 for photomicrograph of rock. (After STOUT, by permission of Bull. Am. Assoc. Petrol. Geologists.

of oil saturation continues to rise until the irreducible water saturation is reached. After this, increasing pressure has no effect on the fluid saturation percentages. Fig.20 is a mercury capillary pressure curve of a micro-crystalline dolomite (a photomicrograph of the specimen is shown in Fig.8) that forms a good reservoir rock (STOUT,1964, p.334). This rock has a porosity of 25.6 %, and of this porosity, about 95 is available for saturation by oil. The height of the oil column required to exceed the displacement pressure is low. The high effective porosity indicates that both the pores and pore interconnections (throats) are relatively large. Fig.21, on the other hand, pertains to rock that could not serve as a reservoir rock because a very high pressure is needed to force oil into the rock, and the irreducible water saturation percentage is large. In other words, the majority of pores in the rock are either, or both, too small or too poorly connected to contain and transmit oil. A photomicrograph of a specimen of the poor reservoir rock is shown in Fig.22. Methods of evaluating the performance of low-permeability reservoir rocks, including carbonates and sandstones, have been discussed at length by RIECKMANN (1963); and the relationships between pore geometry and petrophysical properties measured by electric well logging have been presented by ARCHIE(1952).


377

Pore geometry and oil migration Under hydrostatic conditions the equation relating critical height of oil column to pore and throat dimensions has been given previously. ASCHENBRENNER and ACHAUER (1960, p.239) calculated that the critical height corresponding to the logarithmic means of pore and pore throats measured by them (Fig.18, 19) is about 7.5 ft. In other words, the oil phase must be continuous over a vertical distance of 7.5 ft. before the oil will move by buoyancy through water-wet carbonate rocks of average pore-space configuration. If oil migrates in the topmost few inches of a carrier bed, the minimum length of continuous oil phase ( L ) sufficient to induce migration, is a function of the angle of slope of the bed (0)and the critical height (h) according to the equation (ASCHENBRENNER and A C H A U E R , 1960, p.240):

L = -

h sin 0

-~

If the critical height is fixed, a long continuous oil phase, combined with gentle

Fig.22. Fine-grained dolomite whose mercury capillary curve is shown in Fig.21. Irregular laminae represent algal structure. Mississippian Mission Canyon Formation, North Dakota. Thin section, crossed nicols. Width of the portion of specimen shown in photomicrograph is 6.5 mm. (After STOUT,1964, p.336.) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

378

J. W. HARBAUGH

50

10

Length

of

100

continuous

500

oil

5000

1000

phase,

in

feet

Fig.23. Graph relating minimum length of continuous oil phase to minimum sIope needed for migration of oil by buoyant rise in water-wet carbonate rocks of average pore configuration and ACHAUER, 1960, p.240.) (critical oil column height of about 7.5 ft). (After ASCHENBRENNER Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

slope, is equivalent to a shorter continuous oil phase and steeper slope as far as oil migration is concerned. A graph relating slope and length of continuous oil phase for carbonate rocks of average pore configuration is shown in Fig.23. Thus, the pore geometry of carbonate rocks affects the critical height of oil column, and in turn, has an important influence on oil migration under hydrostatic conditions. Under hydrodynamic conditions the pressure gradient necessary to move a continuous oil phase, assuming a slope that is roughly uniform, is given approximately by the equation (ASCHENBRENNER and ACHAUER, 1960, p.240): gradp

= _ 2T _

L

(l/r-l/R)&

g sin 0 (em-eo)

where gradp = pressure gradient in dynes/cm2/cm, T = interfacial tension between water and oil in dynes/cm, L = length of continuous oil phase measured in centimeters in direction of flow, r = radius of sphere that can just pass through a pore throat in centimeters, and R == radius of largest sphere that will fit into pore space, in centimeters.

379


If the logarithmic mean values for pore sphere and pore throat dimensions that were given by ASCHENBRENNER and ACHAUER (1960, p.238) are used, and a slope of less than 100 ft./mile is assumed, then the critical pressure gradient is equal to: gradp

44,105 dynes/cm2 L

= __

~

~

Thus, other conditions remaining unchanged, the longer the continuous oil phase (L), the smaller the critical pressure gradient (gradp), and vice versa. The pressure gradient is essentially a function of the rate of flow of the water within the reservoir rock, flow rate and pressure gradient being related by the equation: Q

=

gradp

K

__

P @

where Q = flow rate in cm/sec, gradp = pressure gradient in atm/cm, K = permeability in darcys, ,u = viscosity of water in centipoise units, and @ = porosity expressed as a fraction (not i n percent). For a given permeability and porosity, the minimum length of continuous oil phase is a function of the rate of flow of the water, and vice versa. The relation500

I

0

al

100 L

a,

+

50

0)

a,

Y-

.-c 3- 10

0 -

.+ " - 5 0

al + 0

[II

I

Length of continuous oil

phase, i n

feet

Fig.24. Graph relating minimum length of continuous oil phase to minimum rate of water flow for oil migration under hydrodynamic conditions in water-wet carbonate rocks of average pore configuration. Sloping lines represent three different permeability/porosity ratios, where K = permeability in darcys and @ =porosity expressed as a fraction. (Modified after ASCHENBRENNER and ACHAUER, 1960, p.241.) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

380

J. W. HARBAUGH

ships between minimum rate of flow and minimum length of continuous oil phase, for three different permeability/fractional porosity ratios, are shown graphically in Fig.24. These graphs are based on the logarithmic means of pore sizes and pore throat sizes of Paleozoic limestones in the Rocky Mountain and Williston Basin regions measured by ASCHENBRENNER and ACHAUER (1960, p.238). It is obvious that the permeability/porosity ratio greatly influences oil migration under hydrodynamic conditions. For example, if the permeability is 25 millidarcys and the porosity is 25% (K/@=0.025/0.25 = O.lO), then a water flow rate of at least 20 ft./year is needed if the minimum length of continuous oil phase is 10 ft. For different pore configurations, the permeability/porosity ratio curves would be shifted from those shown.

PETROLOGY AND PALEOECOLOGY OF CARBONATE RESERVOIRS

During the past 15 years, geologists have been actively interested in the petrology and paleoecology of carbonate oil reservoirs. This interest stems from the fact that carbonate reservoirs commonly exhibit complex facies variations, which in turn affect porosity, permeability and other aspects of pore geometry. These facies variations reflect differencesin ancient depositional environments and ancient organism communities. Consequently, exploration for oil in carbonate reservoirs is facilitated by understanding ancient environments. Examples of oil occurrence in specific carbonate reservoirs are described next.

N.E.

S.W. INCREASING BASIN S H ALY S I L I C E O U S L I M E STONE

SALINITY

SHELF

AN0

" BAN K"M A

RG IN

"

FOSSlLlFEROUS - FRAGMENTAL L I M E S T O N E and CHALKY LIMESTONE

(MINOR ZAPHRENTID CORALS and OSTRACODES)

I

I

-

SHOREWARD

BAN K

LAGOON

"

PRECIPITATED LIMESTONE

I

ANHYORITE, EARTHY TEXTURED ARGILLACEOUS DOLOMITE and CHALKY LIME STONE

1

CHALKY

I

LIMESTONE)

Fig.25. Simplified general classification of depositional environments, sediment types, and organism communities of Mississippian rocks in southeastern Saskatchewan.(After EDIE,1958, p.102.) Reproduced by permission of Bull. Am. Assoe. Petrol. Geologists.

381


Mississippian carbovlate reservoirs, southeastern Saskatchewan EDIE(1958) provided a convincing example of the application of petrology and paleoecology in the study of Mississippian carbonate reservoirs in southeastern Saskatchewan. Edie classified the environments in which the Mississippian rocks of this region were deposited, into basin, shelf, barrier beach and lagoon (Fig.25). More or less distinctive suites of organisms and lithological assemblages (Table I) characterize the deposits of each environment, although the boundaries between environmental units are gradational. For example, Edie interpreted the deposition of dolomite to be related to near-shore conditions, a relationship emphasized by CHILINGAR (1960). Oolitic and pseudo-oolitic Mississippian limestones of this region are interpreted to have formed through chemical precipitates in the barrier bank environment. The oolitic limestones tend to be well sorted and highly permeable, much of

INDEX MAP

I

I

6 Miles

I

Fig.26. Map showing oil fields (stippled) and contours on top of Mississippian rocks in part of southeastern Saskatchewan. (After EDIE,1958, p.112.) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

382

J. W. HARBAUGH

+-S

N-t

0

.+ 0 0 d

Oolitic, pisolitic,

or dense limestone

B i o c l o s t i c limestone Shole a n d o r g i l l a c e o u s si I i c e o u s I imestone

and

Fig.27. North-south subsurface section through Nottinghani oil field area, southeastern Saskatchewan. Facies variations in Mississippian rocks are shown by patterns. Note irregular ancient upper erosion surface at top of Mississippian. (After EDIE,1958, p. 11 1 .) Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

the porosity being of primary origin. Locally, the barrier-bank oolitic limestones form good oil reservoirs, not only because of favorable porosity-permeability characteristics, but also because these limestones are relatively resistant to weathering and formed topographic ridges when they were exposed before the overlying Jurassic beds were laid down. In places, such as Nottingham field (Fig.26, 27), the ancient topographic ridges form modern structural highs which serve as oil traps. Pennsylvanian-Permian Horseslzoe atoll, west Texas

Study of the Pennsylvanian-Permian Horseshoe atoll in west Texas by MYERSet al. (1956) provided an outstanding example of the value of interpretation of depositional environments in oil-field development. Application of the term “atoll” to this carbonate complex is interpretive in that it likens the limestone mass (Fig.28) to the present reef atolls of the East Indies and the Australian shelves, and not to the deep oceanic atolls of the mid-Pacific region. The following description of Horseshoe atoll is quoted from MYERSet al. (1956, p.7): “The Horseshoe atoll is a subsurface accumulation of fossiliferous limestone

TABLE I DETAILED ENVIRONMENT CLASSIFICATION OF MISSISSIPPIAN LJMBSTONES IN SOUTHEASTERN SASKATCHEWAN]

(After EDIE,1958, p.103)

5 h

-

Basin

Open marine shelf

-_____

-_____-

Central basin

Dark gray argillaceous calcisiltite containing scattered crinoid fragments and sponge spicules, dark gray dense siliceous limestone (associated with dark gray calcareous shale).

Basin edge

Dark gray calcisiltite containing scattered crinoid ossicles and complete brachiopods in place of growth (minor associated dark gray calcareous shale partings).

Intershoal

Cream-white chalky limestone containing scattered crinoid ossicles, bryozoans, brachiopods, and zaphrentid corals.

Shoal

Cream-white crinoidalbryozoan calcarenites.

Barrier Bank

Lagoon

Shoal open marine edge

Central shoals

Cream-white endothyroidcalcarenites containing pseudooolites, oolites, and crinoid debris.

Light buff pisolitic, oolitic, and pseudooolitic limestones containing algal structures and few gastropods; intraformational limestone conglomerate.

Intershoal lagoon edge

Light buff microgranular limestone containing scattered pseudooolites, and light buff lithographic limestone containing scattered oolites and pseudooolites.

E

C A

E c

EP

Cream-white chalky limestone contaming scattered pseudo-oolites and ostracods, dark gray lithographic limestone (associated with anhydrite and earthy-textured argillaceous dolomite).

w

0 n

E

Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

w

00

W

384

J. W. HARBAUGH

Fig.28. Artist’s reconstruction of northern part of Midland basin of West Texas in late Pennsylvanian time, showing shape of Horseshoe atoll. Vertical relief is exaggerated. Length of block is about 125 miles. (After MYERSet al., 1956, p.6.)

which is as much as 3,000 ft. thick and which was deposited during Pennsylvanian and early Permian time in the northern part of the Midland Basin, in western Texas.It is a horseshoe-shaped mass about 90 miles across in an east-west direction and about 70 miles from north to south. The crest of the atoll is a series of irregular hills and depressions, and the flanks slope gently away to merge with a broad limestone platform on which the atoll rests. Large quantities of petroleum have been produced from reservoirs in the eastern and southern parts of the atoll. The limestone in the Horseshoe atoll is composed mainly of nonbedded clastic calcium carbonate debris of organic origin. It is bonded by lithified lime mud and by calcite cement. The only lithologically distinct units that are traceable from well to well in the atoll are thin beds of dark bituminous shale. Porosity in the reef rocks is almost entirely secondary. Effective porosity ranges from almost nothing to 30%. Studies of cores and well logs have shown that the atoll contains zones of relatively high porosity (more than 4.5 %) having little argillaceous material, alternating with zones of relatively low porosity (less than 4.5 %) having considerable argillaceous material. Chemical analyses indicate the limestone generally contains more than 97 % calcium carbonate. Insoluble residues average about 1.4% of the total limestone by weight.” Pennsylvanian carbonate reservoirs, Ismay Jield, Utah and Colorado Recent studies of the Ismay oil field of southeastern Utah and southwestern Colorado form excellent examples of the application of paleoecology and carbonate

385


petrology to an understanding of the origin of a carbonate reservoir (CHOQUETTE and TRAUT,1963; ELIAS,1963). The Ismay field produces oil principally from lenses or “buildups” of bioclastic limestone formed largely from loosely packed fragments of a leaflike calcareous alga generally termed Ivanoviu (Fig.3, 4). The individual buildups occur stacked one above another in three cyclic intervals in the Ismay zone of the Middle Pennsylvanian Paradox Formation (CHOQUETTE and TRAUT,1963). Although the oil productive trend of the field occurs on a low structural nose and is elongate in a northwest-southeast direction, the individual Ivunovia buildups are elongated in a northeast-southwest direction (Fig.29). The Ivunoviu buildups appear to be flat-bottomed, mound-like deposits whose dimensions are up to several thousand ft. across, at least 2 miles long, and up to 40 ft. thick. The buildups are interpreted to have formed as submerged banks that rose above the surrounding sea floor. Choquette and Traut suggested that the banks probably grew upward to within a few tens of feet below the surface of the sea, perhaps into a zone where the water was somewhat turbulent; but rigorous surf conditions are not thought to have prevailed.

Producing Well

0

4

0

0

Dry Hole

16

=Pelletal

0

Mud Focies

lvanovia F a c i e s

0

Sponge F a c i e s

0

0

Black Shale Focies

0

0

14

21

T

0

T

40

w.

35 N.

0

4 28

4-

23

R

IW

Fig.29. Map of Ismay field showing distribution of Zvanovia facies and other contemporary facies at the end of marine transgression of cycle 1 of Ismay zone. Contour lines show thickness of Zvunoviu facies. Dashed line XX’ marks trace of section shown in Fig.30. (After ELIAS,1963, p.198.)

386

J. W. HARBAUGH

Fig.30. Section XX' through eastern part of Ismay field showing distribution of facies in cycle 1 of Ismay zone. See Fig.29 for location of section. (After ELIAS,1963, p.196.)

Various facies within the Ismay zone may be recognized. ELIAS(1963) distinguished seven facies within a single subdivision or cycle of the Ismay zone (Fig.29, 30) as follows: ( I ) quartz sand and silt facies, (2) black shale facies, (3) fusulinid facies, (4) mud facies, (5) Ivanovia facies, (6) sponge facies, and (7) anhydrite facies. These facies are, of course, intergradational. The Ismay zone facies are repeated in cyclic fashion. ELIAS(1963) interpreted the cyclic variations to reflect rhythmic changes of sea level with respect to land level. He interpreted the depositional history of a single cycle (cycle 1) in the Ismay zone as follows: (a) initially there was regional transgression accompanied by deposition of silt and fine quartz sand (Fig.30), followed by (b) transition to deposition of clay and silt under foul-bottom conditions that resulted in the formation of a black shale facies and then (c) transition to a restricted marine environment in which accumulation of carbonate mud kept pace with rise of sea level. ( d ) As sea level continued to rise, lvunoviu fragments began to accumulate on shoals over oxygenated, calcareous mud bottoms. The accumulation of hanovia kept pace with sea-level rise, forming an asymmetric lens that thickened toward deeper waters (Fig.30). (e) Accumulation of Ivanovia ceased with maximum transgression of the sea, pelletal carbonate mud being deposited over the Ivanovia shoals, and quartz sand and silt lapped against the edges of the shoal. Evaporites were deposited in offshore basins. CHOQUETTE and TRAUT (1963) interpreted the cycles somewhat differently.

387


They interpreted the normal marine build up cycle to be composed of the following units: Pelletal limestone Foraminifera1 limestone Algal calcarenite and calcilutite Shelly calcilutite Dark carbonate mudstone Black shale

top of cycle

base of cycle

They suggested that the black shale facies represents the deepest water phase, formed at maximum transgression of the sea. Deposition of the shelly calcilutite facies and algal facies was accompanied by progressive, gradual shoaling. Following this, pelletal carbonate mud accumulated over the buildup, whereas silty and sandy deposits formed between buildups. The succeeding cycle

0

5

10

Percent

15

20

25

poros i t y

Fig.31. Comparison of porosity-permeability characteristics of Zvunoviu-rich limestones with shelly and dolomitized shelly limestones, Ismay field. Only generalized trends are shown; actual and TRALJT, 1963, 11.182.) data show considerable scatter. (Modified after CHOQUETTE

388

J. W. HARBAUGH

1

I

I

I I

I I

!f i

I

I I I I

i I

I

I

I I

I L.

.’

I

‘.!

SHELF

! 1 i.

I I I I

(. ‘3

L-

- - - - _ _ _ _ _ _ _ !I

Fig.32. Map of Alberta showing location of area of Swan Hills reservoir study, and distribution of reefs of Woodbend Age, including Redwater reef. (After BELYEA, 1960.)

began with renewed transgressions, accompanied by deposition of another blanket of black shale or dark carbonate mudstone. The reader is referred to the original papers of Elias, and of Choquette and Traut to judge the merits of the differing interpretations. Production from the Ismay field is chiefly from porous and permeable zones within the Zvanoviu buildups. Some of the best porosity in the buildups is primary, stemming from incomplete filling by void-filling calcite of spaces between Zvanoviu fragments and other skeletal material (Fig.3, 4). Much of the porosity, however, is secondary, having been produced by solution leaching (Fig.7). Extensive leaching has been largely confined to the buildups, and especially to soluble fossils and other carbonate particles; primary voids appear to have guided the

389


dissolving solutions. CHOQUETTE and TRAUT (1963) pointed out that dolomitization, and infilling of voids by anhydrite have also affected porosity. Dolomitization has affected most of the Ismay zone limestones to some extent. Where dolomitization was accompanied by solution of calcium carbonate, zones of moderate to good permeability were created. Generalized porosity-permeability relationships in rock rich in Algae are contrasted in Fig.31 with finer grained limestones and dolomitized limestones containing considerable carbonate mud.

Devonian Swan HiIls reservoirs, central Alberta The close relationship between oil occurrences and facies variations in carbonate rocks are emphasized when textural and porosity units are mapped. THOMAS (1962) provided excellent examples of the use of different types of maps in relating oil occurrences to lithologic variations in the Swan Hills Member of the Devonian Beaverhill Lake Formation of central Alberta (Fig.32). Oil reservoir conditions in the Swan Hills Member are closely related to facies variations, which reflect differences in ancient depositional environments and organism communities. R12W. 5 MER MT

Fig.33. Isopach of Swan Hills Member carbonate rocks of Beaverhill Lake Formation, central 1962, Alberta. See Fig.32 for general location of area. Contours are in feet. (After THOMAS, p.201.) Reproduced by permission of Am. Assoc. Petrol. Geologists.

w

\D

0

TABLE I1 TYPICAL FACIES AND ASSOCIATED RESERVOIR PROPERTIES OF SWAN HILLS CARBONATES'

(After THOMAS, 1962)

Marginal shoal

Intermediate area

Lagoon

Partially leached stromatoporoids with porous microgranular matrix; 18 % porosity, 60 md.2

Stromatoporoids with poor intraorganic vug porosity in chalky to micrograined matrix. Poorly effective vug porosity. Ineffective matrix porosity; 8 % porosity, 2.5 md. Poorly effective reservoir.

Stromatoporoids with brown, vaguely pelleted cryptograined matrix.

Effective reservoir. Partially leached amphiporids (10%) in very fine granular matrix (reworked stromatoporoidal material); 14 % porosity, 40 md. Effective reservoir. Algal-encrusted, partially leached amphiporid material (10 %) with excellent shrinkage, solution, and intraorganic vug porosity; 7 % porosity, 200 md. Effective reservoir.

Tight.

Skeletal (ostracods) micrograined limestone Amphiporids (20 %) with poor intraorganic vug porosity in chalky to micrograined matrix. with 25 % calcite-infilled amphiporids. Ineffective matrix porosity; 6 % porosity, 2 md. Poorly effective reservoir. Tight. Cryptograined limestone, with vague algal biscuit bodies and calcite-infilled amphiporids (20 %). Tight.

Skeletal (ostracods with poor intraorganic vug porosity) limestone with chalky matrix; amphiporids (15 %) with traces of poor intraorganic vug porosity: 10 % porosity, 1.5 md. Doubtful to ineffective reservoir.

Skeletal (ostracods) pellet limestone, with micrograined matrix; calcite-infilled amphiporids constitute 25 %.

Uncemented skeletal (gastropods) lump (algal) limestone. Excellent intra-skeletal and interlump vug porosity; 19% porosity, 250 md.

Partially cemented skeletal-lump limestone. Gastropod casts infilled with clear calcite. Fair interlump vug porosity; 8 % porosity, 25 md.

Effective reservoir. Well sorted, pellet limestone (0.17 mm) with 20 % partially leached stromatoporoids; 21 % porosity, 220 md. Effective reservoir.

Tight.

Effective reservoir. Well sorted, fine granular limestone. 15 % stromatoporoids and 10% amphiporids with intraorganic vug porosity: 18 % porosity, 60 md. Effective reservoir.

Reproduced by permission of Am. Assoc. Petrol. Geologists. md = millidarcy.

Brown pelleted, cryptograined limestone, “syneresis” cracks and “birdseye” structures infilled with calcite cement (60 %), calcite infilled amphiporids. Tight. Skeletal-lump limestone with 15 % micrograined matrix. Patches of calcite and selenite cement at contact of grains and matrix. Poor blind interlump vug porosity; 4 % porosity, 1.8 md. Poorly effective reservoir. Fine-grained (fine sand and clay-sized intermixture) limestone, 10 % stromatoporoids, 20 % calcite-infilled amphiporids. Tight.

8 P

8 E

-

392

J. W. HARBAUGH

sw

Swan H i l l s Formation

NE +

C

Slice

0 .+ 0

E,

LL

Slice

Fig.34. Cross section through Beaverhill Lake Formation showing distribution of rocks with high effective porosity (black) in Swan Hills Member. Limestone with low effective porosity in Swan Hills Member is diagonally ruled. Shale and shaly limestone are horizontally ruled. Slices 1 and 2 are shown on slice maps (Fig.35 and 36). Trace of cross section is shown in Fig.33. (After THOMAS, 1962, p.211.) Reproduced by permission of Am. Assoc. Petrol. Geologists.

R12W. 5MER

3

i

Effective porosity 30%+ 20% 10%

+qn-

$3 !

I

k " .

cc f

$

Edge of Swan Hills Member, where it grades toshale and shaly limestone

Fig.35. Slice map 1 utilizing contours to show variations in percent effective porosity. (After

THOMAS, 1962, p.218.) Reproduced by permission of Am. Assoc. Petrol. Geologists.


393

THOMAS' (1963, p.217) classification of reservoir properties according to facies of the Swan Hills Member is reproduced in Table 11. In this Table, porosity in percent, permeability in millidarcys (md), and a general appraisal of reservoir performance are listed for different lithologies. The relationships between oil occurrence, thickness and porosity in the Swan Hills Member are brought out in a series of maps. Fig.33 is a map of the thickness of the Swan Hills Member, and reveals that oil occurrence is associated with intermediate thickness. A cross section and series of slice maps (Fig.34, 35, and 36) of effective porosity in the Swan Hills Member show a close relationship between effective porosity highs and oil occurrence.

Devonian Redwater reef complex, central Alberta The Redwater reef complex of central Alberta (Fig.32) provides an excellent examR12W. 5 M E R

T.68 i

E f f e c t i v e porosity 40 %+

010-40°/0

'

c

,..' t

Pig.36. Slice map 2. (After THOMAS, 1962, p.221.) Reproduced by permission of Am. Assoc. Petrol. Geologists.

394

J. W. HARBAUGH

ple of the relationship of oil occurrence to ancient organism communities. The Redwater reef complex is one of a number of reefs that occur in the Upper Devonian Woodbend Group in Alberta. The oil reservoir in the Redwater complex occurs within the Leduc Formation. At a number of other localities, the Leduc Formation has been extensively dolomitized, but at Redwater field dolomitization is less intense and original depositional features are well preserved. The Leduc Formation consists of a carbonate reef complex deposited on a broad carbonate platform formed by the underlying, and, in part, laterally equivalent Cooking Lake Formation. Porosity, permeability, and, in turn, oil occurrence in the Leduc Formation at Redwater are controlled, directly or indirectly, by ancient depositional environments and ancient organism communities. The outline of Redwater oil field (Fig.37) is closely related to facies boundaries. From outside the reef to its interior, in a northeast-southwest direction, ANDRICHUK (1958, pp.77-83) distinguished an outer algal-stromatoporoid facies, an intermediate stromatoporoid facies, an intermediate non-stromatoporoidal facies, and an innermost chemically precipitated limestone facies. In turn, KLOVAN (1964) subdivided the major facies into lesser facies divisions (Fig.38) in terms of lithology, and plotted the distri-

1

6 MILES

I

Fig.37. Map showing outline of Redwater oil field, Alberta, and boundaries between major facies distinguished by ANDRICHUK1958, p.78). See Fig.32 for index map. Reproduced by permission of Bull. Am. Assoc. Petrol. Geologists.

395


+sw

NE

-

- 950 -

+

-1000

-

2W .-c

-1050

-

C

-

0 .-+

-1100

a,

0

-

Fare-Reef Megalodon Facies Reef Detritus Tabular, Strom.Facie

0

Reef Detritus W/Massive Strom. Organic R e e f

a

Non- Skeletal Calcarenites

[B

5

W

0

-1150

2

-

3 cn

-

1200

FocIes

Fig.38. Composite cross section through upper part of Redwater reef complex showing lesser facies divisions as distinguished by KLOVAN(1964, p.55). Reproduced by permission of Bull. Can. Petrol. Geologists,

bution of the principal sediment-contributing organisms (Fig.39). The distribution of facies and organisms strongly suggest that the reef complex is zoned into a forereef environment, which lay adjacent to the open sea, a main organic reef zone, and a back-reef zone. It seems likely the shape of the Redwater reef was strongly influenced by ocean currents and prevailing winds from the northeast (ANDRICHUK,1958, p.85), perhaps much in the same manner as the shapes of modern reefs are affected by winds and currents. Recognition of the response of reefs to ancient regional environmental controls, such as wind direction, should aid in prospecting for oil in other reef complexes.

CONCLUSIONS

In conclusion it may be stated that studies of carbonate reservoir rocks need to be pursued from both the physical-petrological point of view and the biologicalecological point of view. The porosity and permeability of carbonate rocks are controlled directly by their physical properties, but in turn, these properties have been influenced in large part by the environmental conditions under which the rocks were deposited. Carbonate sediments, much more than non-carbonate sediments,

W

396

J. W. HARBAUGH

BACK REEF

P E L E CY PO0

GASTROPOO

STACHYOOES STROY-ALGAL CONSORTIUM

R U 0 0 S E CORAL TABULATE CORA1 TABULAR STROY

Fig.39. Schematic cross section showing distribution of principal sediment-contributing organisms in upper part of Redwater reef complex. (After KLOVAN,1964, p.53.) Reproduced by permission of Bull. Can. Pefvof.Geologists.

reflect the mutual controlling effect of organism communities on depositional environments, and, in turn, the effect of environments on organism communities. It is probably safe to conclude that future exploration for carbonate oil reservoirs will witness an expansion of studies that interrelate regional stratigraphy, paleoecology, petrography, and analysis of reservoir rocks by mercury injection and other laboratory means.

ACKNOWLEDGEMENTS

J. Stout of the CaliforniaOil Company, Ph. W. Choquette of Marathon OilCompany, and R. C . Murray and J. Lucia of the Shell Development Company furnished the photographic prints reproduced in this chapter. M. Malek-Aslani of Tenneco Oil Company discussed certain aspects of carbonate reservoir rock properties during preparation of the manuscript. P. Mary of Stanford University made the drawings,


397

and the manuscript was typed by Mrs. C. G. Trimble and Mrs. B. Knoerle, also of Stanford University. The manuscript was reviewed by R. C. Murray, J. Stout and Ph. W. Choquette. The author is indebted to these people for their assistance.

REFERENCES

ANDRICHUK, J. M., 1958. Stratigraphy and facies analysis of Upper Devonian reefs in Leduc, Stettler and Redwater areas, Alberta. Bull. Am. Assoc. Petrol. Geologists, 42: 1-93. ARCHIE, G. E., 1952. Classification of carbonate reservoir rocks and petrophysical considerations. Bull. Am. Assoc. Petrol. Geolonists, 36: 278-298. ARB, J. J., 1964. Engineering concepts useful in oil finding. Bull. Am. Assoc. Petrol. Geologists, 48: 157-165. ASCHENBRENNER, B. C. and ACHAUER, C. W., 1960. Minimum conditions for migration of oil in water-wet carbonate rocks. Bull. Am. Assoc. Petrol. Geologists, 44: 235-243. ASCHENBRENNER, B. C. and CHILINGAR, G. V., 1960. Teodorovich’s method for determining permeability from pore-space characteristics of carbonate rocks. Bull. Am. Assoc. Petrol. Geologists, 44: 1421-1424. BELYEA, H. R., 1960. Distribution of some reefs and banks of the Upper Devonian Woodbend and Fairholm Groups in Alberta and British Columbia. Geol. Surv. Can., Paper, Can., Dept. Mines Tech. Surv., 59-15: 7 pp. CHILINGAR, G. V., 1956. Use of Ca/Mg ratio in porosity studies. Bull. Am. Assoc. Petrol. Geologists, 40: 2489-2493. CHILINGAR, G. V., 1957. A short note on types of porosity in carbonate rocks. Compass, 35: 69-74. CHILINGAR, G. V., 1960. Ca/Mg ratios of calcareous sediments as a function of depth and distance from shore. Compass, 37: 182-186. CHOQUETTE, PH. W. and TRAUT,J. D., 1963. Pennsylvanian carbonate reservoirs, Ismay field, Utah and Colorado. In: R. 0. BASS(Editor), Shelf Carbonates of the Paradox BasinSymp. Field Con$, 4th, Four Corners Geol. Soc., pp.157-184. DUNHAM, R. J., 1962. Classification of carbonate rocks according to depositional texture. In: W. E. HAM(Editor), Classification of Carbonate Rocks-Am. Assoc. Petrol. Geologists, Mem., 1: 279 pp, EDIE,R. W., 1958. Mississippian sedimentation and oil fields in southeastern Saskatchewan. Bull. Am. Assoc. Petrol. Geologists, 42: 94-126. ELIAS,G. K., 1963. Habitat of Pennsylvanian algal bioherms, Four Corners area. In: R. 0. BASS (Editor), Shelf Carbonates of the Paradox Basin- Symp. Field Conf., 4th, Four Corners Geol. Soc., pp.185-203. GRAF,D. L. and LAMAR, J. E., 1950. Petrology of Fredonia oolite in southern Illinois. Bull. Am. Assoc. Petrol. Geologists, 34: 2318-2336. HARBAUGH, J. W., 1960. Petrology of marine bank limestones of Lansing Group (Pennsylvanian), southeast Kansas. Univ. Kansas Publ., State Geol. Surv. Kansas, Bull., 142 (5): 189-234. HOHLT,R. B., 1948. The nature and origin of limestone porosity. Quart. Colo. School Mines, 43: 1-51. IMBT, W. C. and ELLISON, S. P., 1946. Porosity in limestone and dolomite petroleum reservoirs. In: Drilling and Production Practice. Am. Petrol. Inst., New York, N.Y., pp.364-372. IRWIN JR., C. D., 1963. Producing carbonate reservoirs in the Four Corners area. In: R. 0. BASS (Editor), Shelf Carbonates of the Paradox Basin-Symp. Field Conf, 4th, Four Corners Geol. Soc., pp.144-148. KLOVAN, J. E., 1964. Facies analysis of the Redwater reef complex, Alberta, Canada. Bull. Can. Petrol. Geologists, 12: 1-100. LINDGREN, W., 1912. The nature of replacement. Econ. Geol., 7: 521-535. LUCIA,F. J., 1962. Diagenesis of a crinoidal sediment. J. Sediment. Petrol., 32: 848-865. MCCOMAS, M. R., 1963. Productive core analysis characteristics of carbonate rocks in the Four

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J. W. HARBAUGH

Corners area. In: R. 0. BASS(Editor), Shelf Carbonates of the Paradox Basin-Symp. Field ConJ, 4th. Four Corners Geol. Soc., pp.149-156. MURRAY,R. C., 1960. Origin of porosity in carbonate rocks. J. Sediment. Petrol., 30: 59-84. MURRAY,R. C., 1964. Preservation of primary structures and fabrics in dolomite. In: J. IMBRIE and N. NEWELL (Editors), Approaches to Paleoecology. Wiley, New York, N.Y., pp.388403. MYERS,D. A,, STAFFORD, P. T. and BURNSIDE, R. J., 1956. Geology of the Late Paleozoic Horseshoe atoll in west Texas. Texas, Univ., Bur. Econ. Geol., Publ., 5607: 113 pp. Nuss, W. F. and WHITING,R. I., 1947. Technique for reproducing rock pore space. Bull. Am. Assoc. Petrol. Geologists, 31 : 2044-2049. POWERS, R. W., 1962. Arabian Upper Jurassic carbonate reservoir rocks. In: W. E. HAM(Editor), Classification of Carbonate Rocks-Am. Assoc. Petrol. Geologists, Mem., 1: 279 pp, PURCELL, W. R., 1949. Capillary pressures: their measurement using mercury and the calculation of permeability therefrom. J. Petrol. Technol., 1 : 3948. RIECKMANN, M., 1963. How to evaluate low-permeability reservoir rocks. Oil Gas J., 61 (35): 150-158; 61 (36): 115-118. SCHEIDEGGER, A. E., 1957. The Physics of Flow through Porous Media. Macmillan, New York, N.Y., 313 pp. SINNOKROT, A. A. and CHILINGAR, G. V., 1961. Effect of polarity and presence of carbonate particles on relative permeability of rccks. A review. Compass, 38: 115-120. STOUT,J. L., 1964. Pore geometry as related to carbonate stratigraphic traps. Bull. Am. Assoc. Petrol. Geologists, 48: 329-337. TEODOROVICH, G. I., 1943. Structure of the pore space of carbonate oil reservoir rocks and their permeability as illustrated by Paleozoic reservoirs of Bashkiriya. Dokl. Akad. Nauk S.S.S.R., 39: 231-234. TEODOROWCH, G. I., 1949. Carbonate Facies, Lower Permian-Upper Carboniferous of UraE Volga Region. Izd. Moskov. Obshch. Izpyt. Prirody, Moscow, 304 pp. THOMAS, G. E., 1962. Grouping of carbonate rocks into textural and porosity units for mapping purposes. In: W. E. HAM(Editor), Classification of Carbonate Rocks-Am. Assoc. Petrol. Geologists, Mem., 1: 193-223. WALDSCHMIDT, W. A., FITZGERALD, P. E. and LUNSFORD, C. L., 1956. Classification of porosity and fractures in reservoir rocks. Bull. Am. Assoc. Petrol. Geologists, 40: 953-974. WEYL,P. K., 1958. The solution kinetics of calcite. J. Geol., 66: 163-176. WEYL,P. K., 1960. Porosity through dolomitization; conservation of mass requirements. J. Sediment. Petrol., 30: 85-90. WEYL,P. K., 1961. The carbonate saturometer. J. Geol., 69: 3244.

Chapter 8 CARBONATE ROCKS AND PALEOCLIMATOLOGY I N THE BIOGEOCHEMICAL HISTORY OF THE PLANET RHODES W. FAIRBRIDGE

Columbia University, New York, N.Y. (U.S.A.)

SUMMARY

Physicochemical characteristics of most carbonate minerals are matched by the ecologic distribution of carbonate-secreting organisms to make the carbonate sediments of the present day strongly temperature dependent, and their concentration thus inversely related to degrees of latitude. It is believed that this correlation holds generally for the geological past; refinements added include determination of Ca/Mg, Ca/Sr, Ca/Fe/Ti ratios, the taxonomic habit of the organisms, and so on. Going back in time, however, numbers of difficult problems arise: the size of the earth, the volume of ocean water, and its salinity, alkalinity, and pH. None of these can be uniquely tested at present. The patterns of carbonate sedimentation have changed in very important ways through time, so that a strictly uniformitarian status quo cannot be assumed for the past. Available evidence suggests five great biogeochemical events, each corresponding to a certain threshhold peak: Revolution Z (ca. 3.8 . 109 years), First Life at the atmosphere (CH4, NH3, HzO)-water interface; Revolution ZZ (ca. 2.9 * lo9 years), First Photosynthesis with evolution of 02, followed at an undetermined stage within the next billion years by the first primitive animals; Revolution IZZ (ca. 6 108 years), First Carbonate Shells, appearing on organisms up to highest invertebrate level; Revolution ZV (ca. 3 * 108 years), Great Coal Age, following long Paleozoic history of carbonate removal as limestone and the carbon as coal, marked by secular drop in pco2 and rise in po2; and Revolution V (ca. I * 108 years), First Carbonate Pelagic Foraminifera (and coccoliths), shifting major site of carbonate sedimentation from neritic to abyssal realm and initiating major withdrawal of carbonates from the geologic cycle.

CLIMATIC INDICATORS

It has been a common rule of thumb for geologists to equate carbonates ofthe past with warm climatic conditions; fossil corals, for example, have been particularly

400

R. W. FAIRBRIDGE

tempting “climatic indicators”. It is unfortunate, however, that the organic agents associated with most Precambrian and Paleozoic limestones are in fact Algae, and even the corals of the Paleozoic are taxonomically far removed from the hermatypic corals of today. Thus it is easy to simply reject such indications. But this is a negative and defeatist attitude. The best thing to do is to carefully review the data for every period and taxonomic group, working backwards from the present. Up to the present moment, much of such work has been scattered and unsystematic, and a coordinated program is urgently needed. The best that can be offered here is a synthesis of the available principles. A convergence of evidence in such complex problems very often develops which may be quite helpful. To the physicist or chemist the numerical looseness of this sort of reasoning is very upsetting, but natural science with its infinity of interlocking open systems I S often too complex for simple “elegant” solutions. If geologists in the past had worried unduly about statistical precision, the whole edifice of our science would never have emerged. The human brain is still fortunately able to compare, correlate, and organize fragmentary data that would defy even the most sophisticated modern machine analysis. Thus, although any and every one item of evidence that is marshalled may be pitifully weak, the convergence of such evidence onto a common focus may well leave us with a fairly satisfactory conclusion. This is not a bad state of affairs so long as one realizes its tenuous nature and avoids the comforting urge to permit the working hypothesis to become the dogmatic conviction. If, with Hutton and Lyell, one takes the present as the key to the past, the first approximation for carbonate distribution in the past should be obtained by studying the present ranges of the various carbonate facies. But this approach assumes that the present is geologically “normal”, that is to say typical for long periods of the geologic past. This is most emphatically not true. The present time is the Holocene stage, a waning interglacial stage of the Quaternary Period, an “ice age”, and ice ages recur only rarely in geological time. The present mean world temperature is about 2-3 “C below the mid-Holocene maximum (about 6,000 years ago), and about 5-7°C above the last glacial minimum (FAIRBRIDGE, 1961a, p. 119). The last million years, therefore, stand as exceptional or atypical in comparison with the preceding 200 million years. How can one obtain a numerical yardstick for a geologically “normal” climate? Obviously, this is a shifting concept, but it is possible to use a number of indicators which, for very late geological time, cannot possibly have changed their vaIues. For early (for example pre-Mesozoic) times, soil formations are rarely preserved, but for late Tertiary and Quaternary, there are well-dated and unequivocable examples. For example, lateritic and bauxitic soils are not extensively formed beyond 20”N and S latitude today, whereas Tertiary soils of this sort are found extending to 50”N and S. Fossil plants of “living” genera (and even species) are often found to have

40 1

CARBONATE ROCKS AND PALEOCLIMATOLOGY

similar ranges in the past. Whereas most animals are much more adaptable to climate change than plants, it has long been a subject of surprise that interglacial and late Tertiary deposits at about 50"N and S latitude often contain mammalian fossils of characteristically subtropical types (Hippopotamus, Elephas, Rhinoceros, etc.). No icecap at all is believed to have existed in Antarctica prior to the Pliocene or possibly Miocene (FAIRBRIDGE, 1952), and the same is true for Greenland (WAGER,1933). It has been determined by means of all available indicators that the mid-latitudes (40-50 ") in early Tertiary times were systematically about 10"C warmer than at present (SCHWARZBACH, 1963). In this way the mid-latitude climatic zones of the Tertiary were much more spread out than today, and the major boundaries were displaced about 3000 km polewards of their present limits, respectively. No suggestions of continental drift, global expansion and polar change would satisfy this distribution. It is strictly climatic and world-wide. It is true that along certain coastal belts in the upper middle latitudes, e.g., in the northwestern Pacific British Columbia, and Alaska, or in Japan, the mean climates were only a few degrees warmer during theTertiary (BERRY, 1922; CHANEY, 1940; DORF,1959; URHAM, 1959); in these maritime areas, however, moderating longshore currents and the normally mild nature of such climates would tend to maintain a temperate condition, regardless of world-wide oscillations. This fact is not given adequate recognition by the above-named authors. As far as the tropical ocean is concerned, EMILIANI (1955, 1961) provided an excellent series of180/160isotoperatiosdetermined from pelagic Foraminifera. These suggest that during the Quaternary there has been an oscillation of mean ocean surface temperatures through a range of about 7 "C (Fig. l), where the present temperature is 5 " aboQe the glacial and about 2 O below the interglacial average. On the basis of many different studies (palynological work, etc.), there is general agreement that even during the Holocene (the last 10,000 years) there has been a post-glacial thermal maximum (comparable to an interglacial) that was 2.5-3.0 "C warmer in mid-latitudes than the present (FLINT,1961). For the deep ocean waters, EMILIANI (1956) has carried out oxygen isotope studies on large benthonic Foraminifera. These show a deep bottom-water tem-

P

Warmer

t

Solar radiation

4

Cooler L

200.000

I

100,000

1

Present

Fig. I. Tropical ocean surface temperature (from oxygen isotopes) over the last 200,000 years (EMILIANI, 1955), compared with the effective solar radiation curve of 65 "N (MILANKOVITCH, 1930).

402

R. W. FAIRBRIDGE

TABLE I CLIMATIC ZONES OF THE EARTH, DURING PRESENT, GLACIAL AND GEOLOGICALLY “NORMAL” (NON-

GLACIAL) TIMES

_

_ Present

_

_

Polar Subarctic Cool temperate Warm temperate Subtropical Tropical Equatorial

_

~

T”C

-10

0 5 15 20 25 27

Typical indicator

Ice-sheet Solifluction Podzol Brown soil Desert Laterite Rain forest

_ ~ _ _ _ _ _ _ Approximate latititde --

present

glacial

geologically “normal” times

70-90 60-70 45-60 3045 20-30 10-20 0-10

50-90 40-50 30-40 25-30 10-25 5-10 0-5

75-90 60-75 35-60 30-35 10-30 0-10

N.B. The above table is a gross generalization; it will be modified by such factors as altitude (mountain climates pass upwards through latitude equivalents at about 10” ]at. to 1000 in elevation), continenrality-oceanicif~(distance from and influence of the sea), and position (east or west side of continent). The table should, therefore, be used with caution. Temperature is mean annual. Typical indicator phenomena or soils are given, but likewise only as an “example”. It should be noted that such indicators are “active”; laterite inherited from the Tertiary may be found two zones poleward today and solifluction inherited from the last glacial may be found two zones equatorward today.

perature for the mid-Tertiary periods of at least 8 “C, in comparison with the present figures of about 1-2°C over most of the world (Table I). A plot of the latitudinal shift of temperature zones between glacial and interglacial conditions (Fig.2, FAIRBRIDGE, 1964a) shows that in the equatorial zone the temperature changes are minimal (2-4 “C),whereas inside the North Polar circle they are extremely great (20-40°C). The South Pole was maintained at exceptionally low temperatures because of the high mountainous icecap (FAIRBRIDGE, 1961b, p.561, fig.8). The gentle equator-pole temperature gradient in ngn-glacial times is essentially controlled by thermal balancing mechanisms, which largely become inoperative during periods of extensive ice covers. During warin epochs, due to positive oscillations in effective solar radiation increases, the evaporation in the tropics and cloud cover increase. The latter will then provide a partial radiation screen for the land surface and prevent excessive temperature rise. Decreased evaporation from the oceans during cool epochs will make subtropical lands drier and deserts will vastly expand. Low- to mid-latitude deserts are, therefore, inversely related to total solar radiation influx; deserts are paradoxically smaller in area in the hotter periods (FAIRBRIDGE, 1964b). During the glacial periods high latitudes are covered with snow, or floating ice. Consequently, lacking sources of evaporation, they are arid “white deserts”.


403

Mean sea level air temperature ( O F )

Mean sea level air temperature ('C)

Fig.2. Equator-pole thermal gradients for glacial, present, interglacial and non-glacial periods, based on measurements and calculations for the mean surface water temperatures. (After FAIRBRIDGE, 1964a.)

During non-glacial or interglacial times, however, open water permits normal evaporation (though far less than in mid and low latitudes) and the land will be forested with a cool temperate flora. There is really no support for the ingenious Simpson-Ewing-Donn theory (for objections, see SCHWARZBACH, 1963; FAIRBRIDGE, 1964a). Simpson argued that with any rise of effective solar radiation teaching the 'earth (due to planetary motions and other factors), there would be a rise of evaporation and thus of atmospheric moisture. In high latitudes, this would cause snow to fall, glaciers to grow and an ice age would anomalously result from a rise of insolation. EWINGand DONN(1956, 1958) pointed out that the evaporation from the Arctic Ocean would cease as soon as it froze over, so that the ice epoch could only persist as long as the Arctic was warm enough to remain open. Several fallacies are involved: the freezing of the Arctic is irrelevant, because the moisture feeding the snowfields of Canada, Greenland, Scandinavia, and Central Asia largely comes from the warm latitudes of the North Atlantic; the Arctic Ocean is always cold and would not furnish much moisture under any conditions. Simpson's basic assumption was wrong, as shown by glaciological surveys by AHLMANN (1948) and his students, because glaciers mostly tend to melt when the summer seasons are warm and long, whereas glaciers grow when summers are cool and short and the total precipitation of snow is a relatively minor matter. Glacial and interglacial conditions correlate very nicely with the Milankovitch calculations of celestial mechanics and resultant variations of solar radiation (see Fig. 1). Radiogenic dating by I4C and the uranium series puts precision and a very high degree of probability into what was formerly merely an attractive hypothesis (FAIRBRIDGE, 1961b; BROECKER,1966). The small oscillations of incident solar radiation have a profound effect on an earth which has polar mountains, but very little on an earth

404

R. W. FAIRBRIDGE

without such “catchment areas” for the growth of snowfields and continental ice. To sum up the picture of geological paleoclimatic “normality”, one may generalize as follows: in mid-latitudes the mean ocean and air temperature was about 10“Chigher than today; in equatorial latitudes it was about 3-5 O warmer than today, whereas in polar latitudes it was over 20°C warmer than today; and there were no major continental icecaps. Ocean bottom waters were very generally 8-10 ” warmer than now. RANGE OF MARINE CARBONATES TODAY

On plotting the distribution of marine carbonates today one should distinguish between neritic and abyssal facies, because the former were dominant in the past, whereas the latter dominate the present statistics. Only since Cretaceous times has this been so. Secondly, it is helpful to plot the distribution on an ocean by ocean basis, as well as on a world-wide basis, because certain local features may obscure the significant trends (Fig.3). As shown in Fig.3, the world total of carbonates today is distributed in an essentially symmetrical way about the equator. The neritic plot, however, shows a distinctly bimodal pattern with peaks about 20”N and S. Does this mean that the equatorial zone is really too hot for carbonates? The answer to this ques-

a”

30 0 10 20 -)50% carbonate sediments plotted as ZOO totals Ckm*x106)

Fig.3. Present-day distribution of marine carbonates; heavy lines =pelagic (deep-sea)facies; dotted =neritic (shelf) facies. (After FAIRBRIDGE, 1964a, p.473, fig.12.)


405

tion is “no” because temperatures are ideal for carbonate-secreting organisms. Does it then mean that broad shelves suitable for carbonates are absent from these latitudes? The answer is negative again because the Sunda Shelf, West African Shelf, and Brazil Shelf are each quite appropriate. What it seems to indicate is that equatorial shelves tend to be smothered by terrigenous muds carried down by great rivers, like the Mekong, Solo, Fly, Niger, Congo, and Amazon. It seems to be the high rainfall of the equatorial belts that leads to a high sedimentation rate. The low equatorial carbonate concentration may thus partly be due to the masking of the slowly accumulating carbonates by rapidly accumulating terrigenous muds. Partly, it may also be due to the inhibition of organisms capable of rapid carbonate secretions; for example, the turbidity is not favorable for coral colonization (UMBGROVE, 1947). In the past this bimodality would also be expected, but in an even more clearly marked way. Neritic carbonate sedimentation in the equatorial zone is limited today very largely to reefs and isolated waters of atoll lagoons. But these are mainly oceanic phenomena and are not related to the normal shelf sedimentation patterns of the geological past, which are strictly epicontinental. It should constantly be stressed \ that equivalents to the great coral atolls of the Pacific are never to be found in fossil form on the continents. Although there were shelf atolls in the past, rising in quasicratonic settings, these reefs never rose from volcanic cones resting on a thalassocraton (i.e., a deep-sea floor, with basalt-peridotite foundations) as do the modern mid-oceanic atolls. On considering the world distribution of deep-sea pelagic carbonates today, one may observe, in contrast to the neritic curve, an equatorial peak; however, when broken down to an ocean by ocean analysis, it is seen that this peak is produced by the great mid-Pacific concentration. In the marginal seas of the western Pacific, and in the Indian and Atlantic Oceans, the curves are again bimodal. The reason is the same as for the neritic facies. Inasmuch as the mid-Pacific geotectonic environment has never been found in fossil form in the epeiric and geosynclinal facies of the continents, it is the bimodal pattern that must be accepted as the normal one for the geological past. In the preceding paragraphs the author has discussed marine carbonates in the broadest terms: neritic or pelagic. Certain facies, especially of the neritic group, are very much more precise in their climatic range (REVELLE and FAIRBRIDGE, 1957). Some examples of these include: (a) Odite-concentrically banded, spherical, concretionary grains of aragonite and/or calcite, formed in shallow lagoons or shoal banks. Present range: bimodal, respectively about latitudes 25 ON and S. Water temperature indication: 20-30 “C. Salinity indication: normal or above normal. Nonmarine (lacustrine) oolites also occur in such extreme salinities as those existing in the Great Salt Lake, Utah. (b) Beachrock-an inorganically cemented intertidal calcarenite, that is essen-

406

R. W. FAIRBRIDGE

tially a beach sand rapidly cemented in situ, generally seasonally. The grains are normally composed of fragmented Foraminifera, Mollusca, corals, and calcareous Algae. The cement is usually aragonite, although rapid inversion to calcite occurs; rarely is it gypsum or dolomite. A similar material produced by salt water splash and spray is known as pelagosite. Present range: 30-40 ON and S. Water temperature indication: 15-30°C. Salinity indication: normal, except in the case of gypsum and dolomite cements. (c) Biogenic calcavenites-cemented grains of clastic organogenic carbonate material, formerly called “calcareous sandstone”, or some such term. The grains observed microscopically in thin section or polished surface, provide the basis for a “microfacies” identification, which is often very diagnostic of time, depth, and temperature (FAIRBRIDGE, 1954). Many of the living representatives of those organisms have distinctive climatic ranges. As one goes back in time, however, such analogies become more tenuous. There is room for considerable research in this field, particularly in tracing recognizable types backwards, stage by stage, from the present.

PRESENT RANGE OF CONTINENTAL CARBONATES

Calcareous soils Inasmuch as continental sediments, especially soils, have relatively little chance of being permanently preserved in the stratigraphic column, it might appear to be futile to study their present climatic correlation as a basis for interpreting ancient climates. The rarity of fossil soils, however, does not mean that they are nonexistent, so this is hardly a good reason for ignoring them; and careful studies of soil carbonates may be rather illuminating. Whereas an ancient soil remains or is likely to remain some appreciable time at the earth’s surface or at least in the zone of weathering, its gross features are constantly subject to modification, so that inherited aspects of apaleosol may be increasingly difficult to identify with age. Near the margins of subsiding basins and other diastrophically mobile areas, however, paleosols may become buried, only to be re-exposed long after. Under these conditions a paleosol may be an exceptionally valuable climatic indicator. Carbonates are normally mobilized in modern soils under acid conditions, and the latter are established by the presence of rotting vegetation (COz liberation by bacteria, “humic” acids, etc.); this vegetation in turn requires mild or warm temperatures and moisture. In the absence of hot, dry seasons, and where there is heavy almost constant precipitation as in the cool, wet latitudes as well as in the equatorialzone, the carbonates tend to be completely leached out of the soil and removed by the ground water. If there is a reversal of water movement by means of


407

capillarity, however, such as occurs in the warm temperate, Mediterranean, and subtropical zones, a variety of carbonate soil nodules, pisolites, hardpans, and crusts (“calcrete”, “caliche”, “calcareous duricrust”) are formed. These are the pedocal soils, and are well worth studying as intermediate latitude climatic indicators (not tropical and not cold). Pedocal soils are thus indicative of moderate winter precipitation and strong seasonality (see WOOLNOUGH, 1927; BROWN, 1956; RUTTE,1958; SWINEFORD et al., 1958). Eolian calcarenite

Another example of continental carbonate is the eolian calcarenite (or “eolianite” of SAYLES,1931), sometimes called “coastal limestone”. Its chances of being destroyed by subaerial leaching are high, but it is exceptionally useful as a climatic marker in the Quaternary. It is restricted to coastal belts where carbonate detritals from the beach (Foraminifera, broken Mollusca, corals, calcareous Algae, etc.) are blown up to form littoral dunes. They are found throughout the drier intermediate latitudes (semi-arid belts) of the world, from Bermuda to the Bahamas, North Africa, South Africa, India, and Australia (FAIRBRIDGE and TEICHERT,1953; BRETZ,1960; RUHEet al., 1961). The geomorphic conditions required include a warm, shallow sea (15-30°C) with a vigorous productivity of carbonate shelled organisms, suitable beach conditions, and a relatively long dry season to aid dunebuilding. Large dunes require a rising sea level; during the transgressive process an enormous accumulation of sand is gradually pushed in across the continental shelf to come to rest at such time as a negative fluctuation of sea level occurs, when it is abandoned and becomes stabilized. The loose carbonate grains become cemented by downward percollation of carbonate enriched rain water, and a complex system of crusts, rhizomorphs (fossil roots), rhizoconcretions, and karst pipes develops. With another rise of sea level the now lithified dunes become wave-resistant reefs and islands (e.g., Bermuda). They are often inter-stratified with fossil soils, which may carry additional climatic indicators, in the form of fossil snails, etc., or in the nature of the soil itself, such as terra rossa, chernozem, etc. (FAIRBRIDGE and TEICHERT,1953; see Fig.4).

-prevailingwind

mil with’travertine concretions

I

I

beach-rock c o a r s rained 8 f t 5011 pocket wlth E d t a h dtppmg calcarwus sandstd;le Bdhrlernbryon 32’ with %ells

Fig.4. Pleistocene eolian calcarenites and fossil soils at Hamelin Bay, Western Australia. Length of section ca. 300 m; height ca. 20 m. (After FAIRBRIDGE and TEICHERT, 1953.)

408

R. W. FAIRBRIDGE

Lake mark

Lacustrine carbonates are also good climatic indicators, like calcareous soils, because they are restricted to generally dry climates where the ground waters are not unduly acid. The lakes are often somewhat saline, and thus electrolytes flocculate clays to give marl (a clay-lime mixture). As a result, deposits are partly evaporitic and partly organogenic, the environment being often favorable for freshwater Algae (HUTCHINSON, 1957). Fossils are often in the form of chitinous isopods, Estheria, Chara, and minute gastropods. MODERN GEOCHEMICAL ASPECTS

Certain geochemical attributes of the modern marine carbonate may be particularly helpful for interpreting depths and temperatures of ancient basins of deposition. These relationships are generally concerned with trace elements. It is a feature of the calcite or aragonite crystal lattice that it takes into solid solution magnesium and strontium, respectively; this occurs especially during organic synthesis. Other trace elements, in still minor amounts, are also involved and are currently being subjected to study in several different laboratories. The concentrations are generally expressed as ratios, such as Ca/Mg and Ca/Sr ratios. It is also possible to determine the relationship of CaC03 deposition rate (temperature dependent) to the rate of certain other elements; one can sometimes use, for example, the Ca/Fe(Ti ratio, because the Fe/Ti ratio is not temperature dependent. CalMg ratio

The Ca/Mg ratio should be established for the calcite component of the sediment and should not be merely the gross ratio in the mixture, because the Mg2f is selectively adsorbed by illite clays, and so the gross ratio may only reflect the total illite content or perhaps the degree of diagenesis. Within the calcite fraction, it has been established that Mg2+is selectively favored by certain taxonomic groups of organisms and tends to be positively temperature dependent within any such group (CHAVE,1954; CHILINGAR, 1962). Depending on the water temperature at any particular season of the year, the ratio may reflect the growth period. Different parts of the shell may thus have different ratios (LOWENSTAM, 1954, 1961). Inasmuch as shallow, near-shore waters are systematically warmer as a rule than deep off-shore waters, the gross Ca/Mg ratio of a mixed detrital carbonate sediment reflects temperature-depth-distance from shore relationships (CHILINGAR, 1960; see Fig.5). Owing to the ecologic distribution of certain biotas (which may be of the high or low Mg-secreting taxa), however, there are anomalous concentrations which should not be misinterpreted in terms of temperature, etc. (CHAVE, 1954; CLOUD,1962; see Fig.6).


L

10

409

- 7 - I - - 1 1

0

20 40 60 80 Kilometers from shore

100

Fig.5. Ca/Mg ratio in mixed detrital carbonate sediments plotted as a function of distance From shore in the Bahamas. (Data from CHILINGAR, 1960.) 25 -

20

-

Om

U

' g 0

u 0

+. S

15-

10-

Ol

$

5 -

0

!

/

Barnacles

Environmental temperature

Fig.6. Response of selected members of certain taxonomic groups to thermal range; note increasing Mg content with rise in temperature. (Based on work of CHAVE, 1954.)

CalSr ratio

The Ca/Sr ratio is essentially an attribute of the aragonitic carbonates, for calcite usually carries very little Sr. This ratio is thus a useful relationship only with respect to warm environments, which are favorable to aragonitic organisms and the

410

R. W. FAIRBRIDGE

inorganic precipitation of aragonite. Further restricting its usefulness is the metastable character of aragonite. According t o KRINSLEY (1960), nevertheless, Sr2+ tends to be more stable than Mg2+ with time. The Ca/Sr ratio of sea water seems to be constant, regardless of temperature or salinity. On the average, marine aragonites carry 1.O-0.1 % Sr, marine calcites about 0.01 %, and fresh-water calcites about 0.001 %. In the modern marine realm, as with the Ca/Mg ratio, the Ca/Sr ratio partly reflects temperature-depth-distance from shore (SIEGEL, 1961); again it partly indicates local ecologic concentrations of the aragonite-secreting taxa (e.g., patches of the calcareous alga Halimeda).

CaIFeITi ratio As demonstrated by ARRHENIUS (1952, 1959), the inorganic withdrawal of iron and titanium into ocean bottom sediments today is probably very steady and independent of temperature. Inasmuch as the calcium withdrawal by pelagic Foraminifera is essentially organic and subject to temperature and other metabolic controls, it is highly variable, and the Ca/Fe/Ti ratio is thus a convenient tool for determining it. WISEMANand TODD(1 959) analysed a modern tropical Atlantic core in this way, and the resultant curve matches very closely the curve of Holocene oscillations of mean sea level. The latter are temperature dependent, i.e., glacio-eustatic (FAIRBRIDGE, 1961b. p.556). This equatorial indication of a positive correlation between total calcium productivity and temperature, however, is not reflected in the subtropics and intermediate latitudes. Here taxonomic varieties may change and metabolic activity may be partly related t o nutrients. In the mid-Pacific, ARRHENIUS (1952, 1959) established a “circulation index”, which is inversely related to temperature; that is to say, the lower solar radiation epochs are marked by a steeper pole-equator gradient and the rate of oceanic turnover is thus accelerated. In the intermediate latitudes of the Atlantic and Gulf of Mexico, therefore, the total calcium withdrawal is higher in the glacial phases (BROECKER et al., 1958; WANGERSKY, 1962). Such a relationship probably reflects metabolic stimulation, extra nutrients, etc., for the pelagic Foraminifera, and the distribution of certain species; but the details have yet to be worked out.

GEOCHEMICAL VARIABLES IN THE PAST

Certain recent developments in the areas of cosmogeny and geotectonics have rendered many classical appraisals of ancient carbonate geochemistry, at least partially, out of date (FAIRBRIDGE, 1964b). It is very important to know the salinity and alkalinity of typical sea water at, say, the Precambrian-Cambrian boundary. Simple calculations of present-day productivity rates, and extrapolation back for

41 1


.--• c-, aoo

1'

0

I

I

I

600

I

0

200

x 106yr

Absolute tlme scale

Fig.7. Trace of the mean Ca/Mg ratio in carbonates through time. (Based on datafromCmLmGAR, 1956.)

half a billion years or more, will simply not suffice, when there is little conceptIof the former dimensions of the planet, its volume of ocean water and its atmospheric composition. A number of geochemical indicators can, however, be used to appraise the situation. Ca/Mg ratios for the geological past show a systematic drop prior to the Cretaceous (when the widespread flowering of pelagic Foraminifera upset the environmental patterns: KUENEN,1950). Studies of the Ca/Mg weight ratio et al., 1952; CHILINGAR, 1956) show that through time (DALY,1907; VINOGRADOV the mean ratio was 56 in the Cretaceous, less than 10 in the mid-Paleozoic, and only 4 in the Precambrian (see Fig.7). In contrast the Ca/Sr ratio rises as one goes back into the Precambrian (see Fig.8; and also discussion below). What does this mean? It is possible that it reflects the increasing chance for the older rocks to become dolomitized (Ca/Mg ratio of dolomite = 1.65/1).

0 '

\i K

PC I

0

C I

600

I

S D C P T r J I

I

400

Absolute t i m e scale

I

200

I

Te

I

0

x lo6yr

Fig.8. Trace of the mean Ca/Sr ratio in sediments through time (based on VINOGRADOV et al., 1952); note that the mid and late Paleozoic oscillations are due to including evaporites in calculating the net average.

412

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W. FAIRBRIDGE

It may also reflect an increasing proportion of magnesian calcite-producing organisms going back into the past. For example,in Early Paleozoic and Late Precambrian times, the only carbonate reef or biostromal producers were the calcareous Algae, certain groups of which have a very low initial Ca/Mg ratio; and furthermore, by its metastable condition, the high magnesium calcite is more susceptible to dolomitization than any other mineral. The low Precambrian Ca/Mg ratio might, on the other hand, reflect a more favorable climate for the calcareous Algae involved. It is observed that the richest Mg-secreting organisms today are tropically distributed. And yet again, the localization of the samples used by DALY(1907), VINOGRADOV et al. (1952) and CHILINGAR (1956) (mainly North America and Europe), may simply reflect a former equatorial belt, now displaced by polar migration, global expansion, or continental drift. (1951, 1957, 1958), that the A fifth possibility is that proposed by STRAKHOV atmospheric pco2 in the Precambrian and Paleozoic times was somewhat higher than it is today; it may be deduced that such a condition would lead to a lower ocean p H and would favor primary, syngenetic dolomite production, more or less regardless of temperature. Direct dolomite precipitation is favored by high pcoZ (BARON,1960) and low p H (CHILINGAR and BISSELL, 1963a). High alkalinity could be obtained by means of a higher concentration of borates and SO& ions in solution. An argument in favor of the “acid ocean theory” (of LANE,1906, 1945) is that the Precambrian clay sediments are notably impoverished with respect to CaO (only 17 % of the present mean, according to NANZ,1953). Furtheimore, they were considerably richer in Fe, FeO, A1203, and K 2 0 than modein clays, RANKAMA (1955) pointed out that the FeO/FezOs ratio was lower in the middle Precambrian than it is today. There is, furthermore, no very clear evidence of generally warmer oceans in the Precambrian. (An assumption of this sort has often been made in the past as a consequence of the “molten earth” hypothesis.) Indeed, the feldspars in arkoses 1943), which suggests and graywackes are often very little corroded (PETTIJOHN, low temperature conditions. In addition, in the Late Precambrian time there is evidence of almost world-wide glaciation (SCHWARZBACH, 1963), which could not be made exclusively polar by any amount of ingenious continental drifting or polar shifts, The earlier Precambrian boulder beds which may also be glacial are unfortunately so altered as a rule, that it is not clear if they are really glacial or merely orogenic (the so-called “pebbly mudstone” or “tilloid” facies; see discussion in SCHERMERHORN and STANTON, 1963). (1962) and summarized An interesting experiment performed by CHILINGAR here in Fig.9, shows that if a sea water is made more alkaline by adding bicarbonate, the Ca/Mg ratio in the carbonate precipitate only rises above 10 at very high temperatures (over 60”C), which are hardly realistic. On the other hand, if the solution is enriched with Na+, a high Ca/Mg ratio is obtained at quite low temperatures.

413


Why might one expect the Precambrian sea to be enriched with respect to Naf? Today, this cation is balanced by C1- largely. Na+ would have been liberated by the weathering of sodic feldspars (common in basalt) from the earliest history of earth, but C1- has only been provided by a slow secular emanation from volcanic vents and other sources of mantle volatiles through the whole geologic record. There is no mechanism for the removal of C1- from the ocean except by: (a) granitization, and this can only involve minute fractions of the total ocean water (trapped as connate water); (b) cyclic salt removal as minute water droplets by storm winds, to accumulate in inland soils or return to the sea in rivers; and (c) evaporite deposits, which are also small in relation to the total (BORCHERTand MUIR,1964). It is concluded, therefore, that the NaCl content of the ocean (now 19%J has been steadily increasing with time, but that the Na+ that could be balanced against CO32- in the Precambrian may have been appreciable, thus favoring the lower curve on Fig.9. Returning once more to the Ca/Sr ratio, it was noted that as one goes back to the Precambrian it progressively rises, but only if one is careful to remove certain anomalous figures. Evaporites in the mid-Paleozoic should be excluded from

100

I

,

I

I

I

I-Atmospheric

I

0

Fig.9

20

40 60 Temperature ( " C )

so

160

10-4

10-3

COP

10-2

10-1

Partial pressure of C02 (atrn.)

Fig.10

Fig.9. Ca/Mg ratio in precipitates from sea water saturated (above) with Ca(HCO&, and (below) with a mixture of Ca(HC03)z and Na2C03. Note how the sodium enriched sea water furnishes 1962). a low Ca/Mg ratio at reasonable temperatures. (Data from CHILINGAR, Fig.10. Equilibrium concentrations of bicarbonate and carbon dioxide in water in contact with limestone as a function of the COz partial pressure in the gas phase. (After VOGELand EHHALT, 1963; from DEGENS, 1965.)

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R. W. FAIRBRIDGE

a normal marine curve. VINOGRADOV et al. (1952, 1957) demonstrated a Ca/Sr ratio rise to 6,000 for the Precambrian of the Russian Platform. As in the case of Ca/Mg ratios, this may illustrate: (a) progressive migration of climatic belts; (b) a secular (world) climatic change; (c) perhaps a gradual change in the amounts of ocean salts; (d) a progressive leaching (and removal) of Sr through time; or (e) some combination of the above. If the cause was climatic, it would be in the opposite sense to that suggested by one interpretation of the Ca/Mg curve: the latter has been cited as an indication of a warmer Precambrian, whereas the Ca/Sr curve might imply a cooler Precambrian. It is interesting that BOWEN(1956) was able to draw a somewhat similar Ca/Sr curve, based exclusively on British Paleozoic corals. Evidently, here is another area where there is room for more research. If one visualizes the possibility that in the late Precambrian the ocean may have had a relatively low salinity, but high alkalinity, and a pH suppressed by a high pcoZ,it is appropriate to consider certain carbonate solubility characteristics under such variables. A relatively low salinity would mean low ionic strength and

*L c CO, + H, CO, ( mol /I)

2 0.002

/

HCOj

E

0

zo

c0;-

pH

9.0

8.0

Fig.11. Bicarbonate-carbonate relations in sea water, in relation to pH. (After HARVEY, 1957.) Temp

PH

'OF

I

i

0.0

9.01.o

Calcite Aragonite

solid

I

15

20

25

I 10-3 3 0 molesll

Fig.12. Total COz concentration in sea water as a function of temperature and pH, at varying levels of atmospheric COZ.Note that if the presentpco, (3.5 . atm.) were depressed by algal photosynthesis, precipitation would ensue; alternatively with a higher pco2, a much higher total concentration and probably higher temperature would be needed to cause precipitation. (After REVELLE and FAIRBRIDGE, 1957.)


415

relatively low carbonate solubility. On Fig.10 may be seen the relations between a bicarbonate solution and pcoZ in contact with limestone. If one takes present-day sea water, the p H is controlled mainly by pco2 and temperature (disregarding pressure); and on Fig. 11 one may see the equilibrium of the bicarbonate at low p H (low temperature and high pcoZ). Finally, one may compare the effects of higher or lower temperatures on the calcite or aragonite phases for variations in P C O ~(Fig.12). As WEYL(1966) has remarked, the pH of the ocean, although broadly a measure of the alkali-bicarbonate balance, is a non-conservative measure and it is, therefore, better to consider alkalinity and total CQz. There is a net flux of juvenile CQ2 at 2 . 1.012 gmole/year, compared with 15 times this figure involved in recycling through erosion, river transport and sedimentation, according to modern estimates of the contemporary budget. A similar amount in g equiv. is involved for the alkali flux. If one goes back in time, it is apparent that with changing climates, changing land areas, changing desert-forest ratios, different flux rates must have pertained. Quantitative estimates of these are still awaiting research.

BIOLOGICAL PROBLEMS OF THE PAST AND THE FIRST CARBONATES

There is a broad controlling rule in paleoclimatology that must not be broken. This is the “Law of biological continuity”, which essentially states that inasmuch as life has continued throughout most of geological time, at no stage, not even for an instant, could the metabolic limits of any of the succeeding organisms be transgressed on a world-wide basis. Thus are ruled out many of the wilder flights of imagination, proposed as ad hoc hypotheses. Examples of these involve an earth totally frozen over, or involved in catastrophic planetary collisions or tidal moonbirths; or even eustatic oscillations of such large dimensions that the salinity of the ocean would be seriously affected. “Biologic continuity” does not, however, imply evolutionary stagnation of the geochemical components. Far from it; there is increasing evidence that the natural environments, the earth’s atmosphere and ocean with the present balance of atmospheric gases and ocean salts are in part the result of organic activity. Revolution I There is good reason for believing that the first organisms evolved in and from an atmosphere of methane (CH& ammonia (NHs), and HzO vapor. The appearance of this first life is Revolution l i n earth history (say, about 3.8 & 0.3 * 109years ago). Hydrogen originally present would gradually drift off into the upper atmosphere and away, and this would lead to the instability of methane and ammonia. The latter, in turn, will give rise to free nitrogen, which is the principal gas of our pres-

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R. W. FAIRBRIDGE

ent atmcsphere. Volcanic activity gradually added C1-, SO2 and C02. As there was no soil in this early stage of earth history, rain-water solution would be slight (due to rapid runoff) and the supply of Ca2+and Na+ to the ocean rather meager. Ocean water would be slightly acid and salinity (NaCI) very low. Limestone formation was improbable at first. Any OZ liberated by photo-dissociation of HzO would immediately be withdrawn by mineral oxidation. Bacterial attack on SO$led to production of HzS. Reducing conditions permitted easily oxidized minerals (FeS2, UOZ,etc.) to exist as detrital sediments (RAMDOHR, 1957).

Revolution II At a certain stage, it is evident that primitive chlorophyllic organisms (probably bacterial autotrophs) began to use COz and liberate oxygen as a by-product of sugar synthesis. The beginning of this Revolution 11 (“Eparchean”) probably took place shortly before the time of the oldest radio-isotope-dated stromatolitic limestones, say about 2.9 & 0.2.109 years ago. This removal of COZfrom solution in sea water would raise the pH, especially if it took place in shallow coastal lagoons, and CaC03 would be precipitated. The structures in the early carbonates, however, do not suggest evaporites or primary precipitates; they are probably organogenic rocks. The first carbonate fossil traces are stromatolitic algal structures, and are found all over the world in Precambrian formations, the oldest of which are the Bulawayan dolomites that are over 2.7 * lo9 years old by absolute dating methods. Measurements of the 12C/13C ratios in these rocks may be helpful (WICKMAN, 1956); organic carbonates tend to be enriched in the I3Cisotope. The stromatolites (Collenia, Cryptozoon, etc.) are not skeletal material secreted by multicellular plants, but are wrinkled mats of CaC03 precipitated against the outer surface of unicellular green Algae colonies, just as observed today in Florida (GINSBURG, 1960) and in Shark Bay, W. Australia (LOGAN,1961). The growth of this mat in shallow lagoons may well have sheltered the early cyanophytes from excess UV radiation. This revolution gradually led to the build-up of an oxygen-rich atmosphere, though initially all the 0 2 would have been reabsorbed by mineral oxidation. RUTTEN (1965) suggested that about 1.6 * 109 years ago the 0 2 level reached 0.01 P.A.L. (present atmospheric level). The so-called “Pasteur Level” ( P O , at 1 ”/, of the present atmospheric level) was probably achieved only at the end of the Precambrian (HOLLAND, 1962; BERKNER and MARSHALL, 1964). Nevertheless, it permitted the evolution of the ancestors of all modern animals. Rutten suggested that the first primitive fauna began to evolve about 1.O * lo9 years ago. The end of the Precambrian era is sometimes called the “Lipalian interval” (WALCOTT, 1910) to cover an imaginary epoch, when modern shell-bearing invertebrates werempposed to have evolved. This evolution must have been fantastically rapid and. complex, for the Lower Cambrian discloses already the trilobites, representatives of the highest phylum of the invertebrates together with many other


417

forms. Trilobites are, crudely speaking, ancestral to the modern horse-shoe crab, and such organisms are provided with a highly developed nervous system, brain, eyes, prehensile organs, digestive system, articulated skeletal system, complex musculature and a sophisticated bisexual reproductive system. It has been postulated that such organisms burst out from single-celled primitives in a brief evolutionary explosion at the very close of the Precambrian time to fill a newly created ecologic niche (CLOUDand ABELSON,1961). Modern genetics offer no support for such a revolutionary event. In any case, the very existence of a world-wide stratigraphic gap corresponding to the Lipalian was disproven with the discovery of many fine, unmetamorphosed sedimentary sequences that spanned the whole interval, e.g., the Adelaide System of South Australia (DAVIDand BROWNE,1950; GLAESSNER, 1966). One is forced, therefore, to conclude that the oxygen-breathing invertebrates evolved during an extended time-span of the middle and late Precambrian, although traces of these organisms are extremely sparse and restricted to some impressions, tracks and worm casts. The geochemical environmental conditions appear to have been reasonably acceptable for organisms similar to those living today with but one exception: the sea-water composition was such that they could not secrete carbonate skeletons (SCHINDEWOLF, 1956; CHILINGAR and BISSELL,1963b; FAIRBRIDGE, 1964a). Examples of the fauna are known, and worm tracks are relatively common. The Ediacara fauna of South Australia, with its hydrozoa, etc., appreciably antedates the late Precambrian Sturtian Tillite (GLAESSNER, 1962; GLAESSNER and DAILY,1959). The primitive segmented fauna, presumably arthropods, of the Belt Series of Montana, formerly considered of very late Precambrian or even Lower Cambrian age, is now considered to be a “Middle” Precambrian stage (GILLULY,1963; PFLUG,1965a). Also about one billion years old are the probable Foraminifera reported by PFLUG(1965b). It would appear, therefore, that the middle and late Precambrian sea must have been acceptable for modern invertebrate life in every way, but for this peculiar feature about the carbonate shell secretion. Only the intertidal or lagoonal Algae seemed to have been able to lead to such precipitation. In an acid environment this process would be applicable. In lakes and rivers, modern fresh-water arthropods secrete chitinous shells (with only small clots of calcite) by raising the pH of their own body fluids, regardless of the low p H (5-7) of the environment. It seems possible, therefore, that up till the end of the Precambrian, the oceanic pH did not exceed 7 or a little over. As a matter of interest, all lakes in acid igneous rock areas today have a p H below 7 (HUTCHINSON, 1957); and HOUGH(1958) has demonstrated that the great silica-iron deposits that are known all over the world in the Middle and Late Precambrian, but at IZOother time, could be readily explained by the geochemistry of a lacustrine regime. A predictable low po2 would also favor this deposition (LEPPand GOLDICH,1964) (Fig.] 3). There seems to be good justification for the idea of BERKNER and MARSHALL

418

R. W. FAIRBRIDGE

ZCO,,%

OF PRESENT CONCENTRATION IN CRUST, ATMOSPHERE AND OLEAN

001%

0 1-10

1%

100%rCOL

10%

9

107

'-

~ ~ ~ ~ ~ ~ ~ & i ~ IN - ALL - C SEDIMENTS 0 2 --

\

A L P I N E OROGENY

f

-k

108

\

I

I

\

\

-REVOLUTION -V-(PELAGIC FORAMINIFERA ADDED 7 TO DEEP SEA DEPOSITS) 1

I

-

I

l o "

W

\o

2 c

Earliest Animals

109

IO-~ 0.01%

._______

10.~

10-2

0.1%

1%

10

-'

10%

OXYGEN "1. (or fraction) OF PRESENT ATMOSPHERIC

1

100%

LEVEL

Fig.13. Great biological revolutions of the earth's history. Tentative suggestions of the variation in partial pressures of atmospheric oxygen (largely by photosynthesis; partly by photodissociation of HaO) and COZ(from volcanic emanation).

(1964) that with the extremely thin 0 2 and HzO vapor atmospheric blanket in the early Precambrian, UV-synthesized ozone would be formed directly on the land surface. Inasmuch as ozone, rather like hydrogen peroxide, is a very strong oxidizing agent, the rate of chemical erosion would be very high and continue until such time as the atmospheric blanket is thickened. The production of the enormous silica-iron deposits would be greatly facilitated by such erosion. It must not be forgotten, however, that very important mechanical, unweathered deposits also occur during this period. Inasmuch as the iron deposits are not continuous, a series of oscillatory ozone build-ups may have alternated with less reactive epochs. Their world-wide distribution merely suggests a rather brackish, acid ocean. Alkalinity, however, may have been moderately high owing to the presence of borates and sulphates, as well as limited amounts of the principal sea salt (NaC1) which slowly increased through time. Late Precambrian gypsum and magnesite are known but there is no trace of halite evaporites.

419


Revolution 111

-

Revolution 111 (“Eocambrian”) occurred about 6 & 0.3 lo8 years ago. The first evaporites (halite) appeared in the Cambrian of Jordan, Iran, and India. At the beginning of the Paleozoic, COz was being withdrawn in ever-increasing amounts (raising the pH), adequate concentrations of Ca2+had now accumulated in sea water salts, and unlimited formation of calcite and aragonite shells became practicable for the invertebrates. It is possible that such shells were the inevitable consequences of a rising p H and rising Ca2+ concentration in sea waters, a natural sequence of events that could be predicted as the weathering and solution of silicate rocks on the continents gradually proceeded. Erosion during the Precambrian was dominantly mechanical (cf. graywackes and arkoses), but not exclusively so; during the Paleozoic time chemical erosion (soil development) became more important as the vegetation cover of the land surfaces grew richer and more varied. During Paleozoic times, limestones became very widespread in a belt extending from North America, through western Europe and the U.S.S.R., to Australia. This seems to be a former equatorial belt that persisted (through most of that era), and it is quite unnecessary to propose the existence of climates on a worldwide basis that were warmer than, say, those of the mid-Tertiary. Indeed the mean temperature may have been somewhat cooler; and certainly in the early Paleozoic climates may have been somewhat more extreme, because there was no major vegetation to conserve and transpire moisture. Previously, no doubt, through much of the Precambrian time, there were probably soil bacteria, cyanophytic Algae, fungi, etc. (FISCHER, 1965), but they were probably limited to swampyplaces. Land vegetation (major plants), according to the fossil record available so far, first evolved in Australia in Silurian times, and was widespread already in the Devonian. By Carboniferous times these vast swamp floras were providing for the greatest coal deposits of geologic history. In the Permian the locus of the great coal swamps turned to the Southern Hemisphere, but after the end of the Paleozoic, coal development in general became more restricted. The total coal preserved today is about 6 1012 metric tons, but much has been lost by erosion. Geochemically one may observe that during the Paleozoic, certain animals (e.g., corals) and plants (Algae, and coal swamp flora) were responsible for withdrawing very large quantities of carbon from the atmosphere-hydrosphere, and additional supplies of 0 2 were liberated. Both the C (coal and oil) and CaC03 (limestones) by sedimentological processes became buried on a semi-permanent basis. Thus a new, biologically activated pattern was set up, on a scale that was greater than that which existed during the Precambrian time. There is no reason to imagine that volcanic supply of COZ rose to meet the demand; indeed, it may have been reduced in comparison to the Precambrian rate. On the other hand, the spreading of vegetation onto the land permitted the development of soils and their associated bacteria, which provided for a much more continuous contact of HzCOs

-

420

R. W. FAIRBRIDGE

with rock surfaces than had ever been possible under Precambrian conditions. Thus the soil chemistry was changing to accelerate the supply of the ions of the principal silicate minerals to the sea: Ca2+, Na+, Mgz+, preferentially over K+, Fez+, etc. Revolution ZV The time of Revolution IV (“Permo-Carboniferous”) was 2.5 & 0.3 los years ago. In the ocean it was marked by a rising alkalinity from the supply of alkali metals and, due to the removal of atmospheric and oceanic carbon, there may have been a lowering of the general pcoZ. Thus there was a rise of oceanic pH, perhaps from 7 to about 8. During this transition certain marine invertebrate groups became extinct. The end of the Permian was not a time of wholesale slaughter of Paleozoic life, but rather marked a threshhold or end-point in a long series of geochemical changes that proved too much for certain taxonomic classes. Revolution V It seems that a “modern” atmospheric and oceanic geochemistry had been established early in the Mesozoic. Primitive land mammals, which require oxygen and a low atmospheric pco2 were able to evolve. Nevertheless, in the late Mesozoic a new revolution was developing in the ocean. This concerned principally the evolution of small floating organisms which secrete carbonate tests. This represents, then, the last great biogeochemical event, Revolution V (“Cretaceous”), dated about 1.O rt 0.2 . 108 years ago. For some reason, not well understood, small, pelagic carbonate organisms do not appear to have existed in the Paleozoic. They only gradually emerged during the Mesozoic, and exploded in an immense “flowering” all over the world oceans in the Cretaceous time. These organisms, mainly the Coccolithophoridae and pelagic Foraminifera, are the principal components of the famous Cretaceous chalk formations. They are found from the cliffs of Dover to Australia and from Texas to Antarctica. The high rate of CaC03 withdrawal is reflected by the mean Ca/Mg ratio of the Cretaceous carbonates of about 56 (weight ratio) compared with 16 in the late Paleozoic. Since Cretaceous time it has dropped to about 40. That such an event must have fundamentally changed the earth’s carbonate economy was recognized by KUENEN (1950). The earlier limestones were neritic and epicontinental, and therefore, very liable to be recycled, whereas the pelagic carbonates are oceanic and almost permanently withdrawn from surface circulation (see also CHILINGAR, 1956). It should be remembered, however, that below the “calcium carbonate compensation depth” (now about 4,500 m on the average), all carbonate is redissolved. This level is probably set by a self-adjusting mechanism, related to Ca2+ and HCO3- fluxes.


42 1

Some biological problems

Apart from biogeochemical revolutions, there must have been countless ecological adaptations in the past to changing environments on a paleogeographic basis. With diastrophism new geosynclines have formed; new mountain barriers and archipelagoes were created. And organisms, particularly the sessile ones, such as the corals, became passive “victims” of geotectonic events. Particular difficulties exist for the paleoecologist who would reconstruct a former environment from the evidence of the fossil biota. That the science of paleoecology is a relatively new one may be illustrated by the fact that the Geological Society of America published its joint treatises on Ecology and Paleoecology (edited respectively by HEDGPETH and LADD)as recently as 1957; and the first college textbookonpaleoecologyis the fine work of AGER,issued in 1963. Avery helpful supplement to the latter is the symposium edited by IMBRIEand NEWELL(1964). A Russian view, with useful treatment of methods is given by HECKER (1965). The biogenic carbonates occupy a key place in such studies. The problem is essentially one of understanding the present ecologic patterns and controls, which in itself is not an easy task. Then it is necessary to shift the focus back in time, with an endeavor to understand the multitudinous adaptations and migrations stage by stage in the reverse sense. The wise procedure would not be to jump directly into the Paleozoic, but to take biotas or classes backwards in fairly small steps. Of particular interest to carbonate students are the various reef-forming groups. One has only to study TEICHERT’S (1958) excellent review of cold-water organic reefs to realize the exceptional difficulties that are facing us. Whereas the herniatypic reef-building corals have certain well-marked warm habitats today, it may be noted that the same genera may be present even in rather cool waters. The latter, however, are in the form of small colonies, perhaps 10 cm across instead of 2-3 m. A rapid, upward change in water level and temperature, due for example to a warm climatic swing, would transform these 10 cm “pioneers” to full-fledged reef-builders. Such impoverished forms may be formed both in latitudinally marginal areas (i.e., too far polewards of the subtropical belts), and bathymetriealZy marginal depths i.e., (too deep for vigorous growth and reproduction). The dependence of so many of the major carbonate-secreting organisms on the zooxanthellae, which are unicellular symbiotic green Algae that remove CO2 and waste products from interior mantle folds, means that their hosts are essentially phototropic (YONGE,1940; 1963). These large colonial organisms can nevertheless exist sometimes in an impoverished way, with greatly reduced growth of zooxanthellae, in areas where light is restricted, such as in tropical waters made turbid by mud, in higher latitudes, or in greater than optimum depths. This sort of “metabolic elasticity’’ may thus aid certain groups in surviving the “storms” of geotectonic or paleoclimatic changes. For the paleoecologist,

422

R. W. FAIRBRIDGE

there is the possibility of a sedimentological cross-check on the associated environmental lithology. I n considering past biotic associations, if the lithofacies seems to match what would be reasonable for a modern equivalent, one may have a certain confidence in extrapolating a similar correlation. If the lithofacies seems to be “exotic”, however, caution should be exercised. One has only to consider the curious appearance of layers of broken and jumbled shells of clearly littoral or neritic facies, intercalated in what is lithologically regarded as a turbidite sequence of shallowwater sands and deep-water shales (e.g., in the Hamilton Group of central Pennsylvania). Evidently this is a sedimentologic accident, rather than a biologic adaptation.

PALEOGEOGRAPHIC PROBLEMS

The changing shore-lines and extensive epeiric seas of the geologic past are wellknown features, but in recent years it has become possible to propose certain theoretical guidelines for such paleogeographic events as cyclic transgression and regressions. The importance of the appropriate physical setting to provide most favorable marine environments for the accumulation of carbonates hardly requires further emphasis. Such epicontinental shore-line changes have three main causes: (a) geotectonic revolutions alternating with quiescence; (b) eustatic changes of sea level; and (c) geodetic changes of sea level.

Geotectonic revolutions alternating with quiescence Major diastrophic crescendos in earth history have long been suspected, with names attached to them like the Caledonian, Taconic, Acadian, Hercynian, Appalachian, Sakawa, and Alpine revolutions, together with the large number of “phases”, proposed by STILLE(1924, 1944). Much confusion has been created since the time when CHAMBERLIN (1909) concluded that diastrophism must be the structural basis for historical geology, and thus controlled eustatic changes and provided a world-wide basis for stratigraphic correlation. But it was all too easy to observe an unconformity, and to conclude that here was the evidence for yet another great revolution. ARKELL(1956), GILLULY (1949) and RUTTEN(1949) separately objected. Nevertheless, the increase in the number of isotope datings has shown that there are indeed clear statistical peaks in tectono-magmatic activity (HOLMES, 1960; KULP, 1961), so that the principle of periodic revolution seems to be re-established. If it is generally established that geomagnetic reversals coincide with each diastrophic and stratigraphic boundary event as seems to be suggested by GLANGEAUD and BOBIER (1963), then a cause for organic genetic evolution is offered. The Van Allen radiation screens would temporarily fade out presumably as the mag-


423

netism swung from north to south and back again. Deprived of the usual screen for cosmic radiation, accelerated genetic changes could be expected for those organisms not hidden underground or deeper than about 10 m down in the ocean. From the paleogeographic point of view such revolutionary events also involve the collapse and uplift of old geosynclines and the creation of new. Along rather narrow, but very elongate belts, marked shore-line changes occur. In the orthogeosynclines (the highly mobile belts of eu- and mio-gzosynclinal type), there is a predominance of clastics and volcanic derivatives. In the meantime, far from the sites of violent tectonism, in the cratonic basins or parageosynclines (mainly the exo- andparalia- types of K A Y , 1945, 1947, 1951) quiet conditions are likely to prevail and the carbonates may provide the principal lithofacies. In the latter case, however, they must coincide with favorable paleoclimatic zones. Long periods of quiescence on continents (in equatorial to subtropical belts) will also lead to deep chemical weathering, the leaching of Cat+ from soils and its increased flux to the continental shelves; such conditions correspond to ERHART’S (1956) “biostatic” phases (see also TERMIER and TERMIER, 1963, p.153). Eustatic changes of sea level

Eustatic changes of sea level have been recognized for over a century, but are only just beginning to be incorporated in the working philosophy of stratigraphers. There are many causes for world-wide changes of ocean level (in the same sense), only two of which are of real importance. The “sedimento-eustasy”, or rise of water level due to the sedimentary filling of basins, which stimulated SUESS (1904-1924) to coin the term “eustatic”, is amongst those causes now regarded to be of minor importance (FAIRBRIDGE, 1961a). The two major phenomena are tectono-eustasy and glacio-eustasy. Tectono-eustusy This is a rise or fall of world sea level due to the geotectonic change in shape of basins, large or small. Darwin long ago recognized that if Pacific atolls grew up during a general downwarp of that basin, there would be a resultant drop of world sea level which could explain raised beaches and coastal terraces from Britain to Patagonia. Zeuner believed that the progressive lowering of the raised shorelines through the Quaternary was not an indication that each glacial stage was more severe than the last, but that downwarping of ocean basins was in progress. The general block-faulting and subsidence of the quasicratonic Pacific marginal seas and other “Mediterranean” seas during the Quaternary time was recently brought 1967). out by the present writer (FAIRBRIDGE, Glacio-eustasy Rise and fall of world sea level due to build-up or melting of terrestrial ice sheets,

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this was a direct outcome of the work of Agassiz on glaciation, followed up by MACLAREN (1842) and DALY(1934). Even small changes in the earth’s climate recorded instrumentally during the last century produce measurable changes in mean sea level, as determined by tide gauges (FAIRBRIDGE, 1961a; FAIRBRIDGE and KREBS,1962). Even sunspot cycles are reflected in favorable places (e.g., French Guiana) by rhythmic sedimentation (CHOUBERT and BOY&1959). Discrimination between these two causes in the data provided by world stratigraphy is not easy. Logically it would seem, however, that glacio-eustasy would only be dominant in known “ice ages”, although minor glacio-eustatically controlled cyclic oscillations (of the order of & 5 m) may be expected at any stage in geologic history (HOLLINGSWORTH, 1962). On the other hand, information about world-wide oscillations is still very hard to separate from the vast, but uncoordinated wealth of world stratigraphic data. The so-called “Mediterranean transgressions” of the European Tertiary recognized by SLJESS (1904-1924) may well be tectono-eustatic, and so may be the great stratigraphic guidelines of North America followed out by WHEELER (1963). World-wide demonstration, however, is still needed (see discussion by HALLAM, 1963). The fact that carbonate facies are exceptionally good indicators of minor 1954) may be traced very depth changes, and that their rnicrofucies (FAIRBRIDGE, often for thousands of miles, places them in a good position for helping to elucidate this problem. In reviewing the earth’s stratigraphic record as a whole, one cannot help but .-

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Fig.14. Rise and fall of sea level creates a first-order cyclic sedimentary control over neritic facies in the geologic past. The shallow-seatransgression favors carbonates, the extreme regression does not.


425

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Fig.15. Record of major transgressions and regressions over the last 600 million years. Note coincidence of widespread carbonates with maximal transgressions, and glaciation with maximal re gressions. The general tendency towards a sccular retreat from the continents over this period is interpreted by EGYED(1956, 1957) as evidence for the expansion of ocean basins. (Following (1948) and H. TERMIER and G. demonstrations by EGYED(1956, 1957) using data by STRAKHOV TERMIER (1952); In: FAIRBRIDGE, 1964a, p.469; fig.11.) 0 = Strakhov; 0 = Termier.

be struck by the mega-facies of certain epochs which, on a world-wide basis, appear to have a consistent arithmetic sign. It would be futile to deny that the Triassic, for example, is dominated by terrestrial and terrigenous facies (often red in color) in every continent; its marine members are almost exclusively reserved for narrow orthogeosynclinal belts. By the same token, the Ordovician and the Cretaceous (mainly limestones or chalks) appear to be periods of world-wide transgressive nature, almost regardless of tectonic setting (Fig. 14). There are two distinctive labels for these conditions of world paleogeography: (I) Thalussocratic conditions-when sea level is almost universally high, and epeiric seas are widespread. This is the setting most favorable for carbonates, because the continentality index was lowest, and the shallow seas were widest and warmest. (2) Epeirocratic conditions-when continental shelves are reduced to a minimum. Lowering of base level increases the run-off and terrigenous sediments replace carbonates as the “geo-facies” (see TERMIER and TERMIER,1963, p.333). This is ERHART’S “rhexistatic” phase (1 956), typically shown by red beds (see Fig. 15).

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Geodetic changes of sea level With the recognition of irregular polar shifts, that is to say, displacement of the crust with respect to the geographic polar axis of the Earth‘s rotation, a geodetic consequence must be considered. The ideal spheroid, required by gravity, will be maintained constantly by the hydrosphere, but the lithospheric reaction is slower. Owing to this different reaction time (i.e., viscosity) of the hydrosphere and the solid earth to a sudden change in the relative dipole position, vertical displacements of the hydrosphere will occur, in the positive sense at the new equator, and negative at the new poles. The displacements will be in four quadrants, alternating, with two positive and two negative regions of the globe (FAIRBRIDGE, 1961a). Naturally, this is an ideal condition, omitting continental locations; but it is evident that with the quadrant distribution, it would be possible for an entire continent, such as Australia (5,000 km across) to experience a transgression or a regression at the same moment. If the quadrant coincides entirely with an oceanic area, of course, there may be no trace of the former sea level, except perhaps on isolated seamounts or islands. Along continental margins, however, the effect could be easily confused with a eustatic phenomenon. Only a really representative world-wide survey would suffice to distinguish between eustatic and geodetic phenomena. In their world-wide paleogeographic surveys TERMIER and TERMIER (1 952) have distinguished between “Tethyan” and “Arctic” (“Boreal”) transgressions that have alternatingly affected many regions. In view of the above, the role of geodetic changes would certainly seem worth investigating.

GEOTECTONIC AND PLANETARY PROBLEMS

The geotectonic and planetary problems are of great dimensions and their lack of solution injects into carbonate geochemistry and paleoclimatology the greatest uncertainties. Half a century ago, and even up till quite recently, it was convenient in geochemical calculations t o assume: (a) a stable earth of constant dimensions; (b)an ocean volume and salinity that has not changed much since mid-Precambrian times; and (c) an atmosphere that was essentially the same as the present one. With the philosophic acceptance of a dynamic earth, in a dynamic universe, these concepts of a planetary body of absolutely constant dimensions and characteristics call for remarkable coincidences that today are very difficult to accept. The “contraction theory” was also widely accepted formerly in principle, though admittedly few allowances were made for it. During the last decade, a number of revolutionary discoveries have been made. They include the discoveries in paleomagnetism, which have led to the plotting of former poles, and the conclusions that the continental masses have shifted apart in the course of younger geological time (say the last 0.5 billion years).


427

These migrations have been correlated with revolutionary episodes in the evolution of postulated convection currents in the earth‘s mantle (RUNCORN, 1962), although powerful arguments against convection have also been raised. As a consequence of differentiation within the mantle, there has been proposed an expansion theory (EGYED,1956, 1957; IVANENKO and SAGITOV, 1961; HEEZEN,1962) instead of contraction. The basic dimensions of all geochemical budgets should perhaps be reexamined in this light. Seismic and other recent studies of ocean bottoms have suggested that much of the ocean floor is Mesozoic or younger in age, and not of great antiquity as previously believed. DIETZ(1961) proposed that convection currents are constantly turning over the deep ocean floor material and its cover of sediments, so that a renewal would be possible; but the reflection profiling results offer no support for this idea.

CONCLUSIONS

( I ) Climates of the past are measured, not by meteorologic instruments, but by “indicators”: organisms, processes, soils, etc. Latitudinal shifts of equivalent facies from glacial to non-glacial conditions of several thousand km are recognized. Nonglacial conditions are “normal” for the geological past, with equatorial latitudes 3-5 “C warmer than today, tropical conditions widespread, the poles 20-40 “C warmer than today and the deep ocean floor 8-10°C above the present. (2) The range of marine carbonates today reflects conditions that have only been valid since Cretaceous times, and earlier (neritic) distribution requires a quite distinctive standard for appraisal. Characteristic neritic facies and paleoecologic indicators are available. (3) Continental carbonate indicators are also helpful, but mainly just from the Quaternary period. ( 4 ) Ca/Mg and Ca/Sr ratios have contemporary and historic trends, not all very thoroughly understood. Changes in salinity, alkalinity and pcoZ with time may help to explain them. (5) Five great biologic revolutions have occurred through earth history that have fundamentally shaped the modern geochemical picture. It is significant that organisms have shaped their own environment. Organic COZ and the calcium budget have played critical roles. (6) Rise and fall of sea level (in a relative sense) through time have conditioned the mass of carbonates withdrawn for any one stage, the high sea levels favoring the carbonates and low sea levels being counter-indicators. (7) It must be concluded, finally, that great caution must be exercised about reaching any deductions from long-range calculations, especially those going back to Precambrian, One is not dealing with a closed system in the carbonates. Instead,

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they constitute a vital biogeochemical group that is intimately bound up with the nature and course of early physiologic evolution, of which scientists know little; furthermore this evolution is set in a planetary environment, the basic dimensions of which are not yet agreed upon. REFERENCES

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43 1

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432

R. W. FAIRBRIDGE

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APPENDIX Diagrams for visual estimation of percentages of various components in rock 1954; see also TERRY and CHILINGAR, 1955.) sections. (After SHVETSOV,

Fig.1

Fig.3

Fig.2

434

APPENDIX

REFERENCES

SHVETSOV, M.S., 1954. Concerning some additional aids in studying sedimentary formations. Byul. Mosk. Obshchestva Ispytateki Prirody, Otd. Geol., 29(1): 61-66. TERRY, R.D. and CHILINGAR, G.V., 1955. Summary of “Concerning some additional aids in studying sedime:itary formations” by M.S. Shvetsov. J . Sediment. Petrol., 25(3): 229-234.

REFERENCES INDEX

ADAMS,J.E., 192, 338 ADAMS,J.E. and RHODES,M. L., 10, 20, 27, 249, 281,289, 298, 309, 329, 335, 338 AGER,D. V., 421, 428 H. W., 403, 428 AHLMANN, ALDERMAN, A. R., 40,47 ALDERMAN, A. R. and SKINNER, H. C. W., 5 , 27, 144, 281, 282, 338 ALDERMAN, A. R. and VONDER BORCH,C. C., 281, 282, 338 J. M., 326, 338 ALWAY,F. J. and ZETTERBERG, AMES,R., 291 D. L., 276, 279, 327, 338 AMSBURY, ANDERSON, F. W., 215, 249 ANDREWS, D. A. and SCHALLER, W. T., 338 ANDRICHUK, J. M., 220, 249, 328, 338, 394, 395, 397 ARCHIE,G. E., 316, 391 ARKELL, W. J., 422,428 ARNAL,R. E., 203, 249 ARPS,J. J., 374, 397 ARRHENIUS, G., 83, 410, 428 ASCHENBRENNER, B. C. and ACHAUER, C. W., 92, 144, 311, 372, 373, 314, 377, 378, 319, 380, 397 ASCHENBRENNER, B. C. and CHILINGAR, G. V., 313, 397 AVNIMELECH, M., 319, 338 BAARS,D. L., 294, 95, 117, 144, 315, 338 BAKER, G. and FROSTICK, A. C., 177,249 BAILEY, E. B., COLLET, L. W. and FIELD,R. M., 246, 241, 249 BARON,G., 64, 83, 307, 338, 412, 428 BARON,G. and FAVRE, J., 307, 338 BASTIN,E. S., ANDERSON, B., GREEN,F. E., MERRITT, C. A. and MOULTON, G., 290, 338 BATHURST, R. G. C., 53, 55, 83, 144, 158, 162, 164, 166, 184, 186, 190, 249 BEALES, F. W., 35, 47, 144, 216, 249, 338 BEERBOWER, J. R., 210, 231, 249 BEHREJR., C. H., 338 BELT,E. S., 209, 249 BELYEA, H. R., 220, 249, 388, 391 BENTOR,Y. K. and VROMAN,A., 204, 249

R. E. and TERRIERE, R. T., 220, BERGENBACK, 249 H. R. and COBBAN, W. A., 227, BERGQUIST, 249 BERKNER, L. V. and MARSHALL, L. C., 25, 416, 417 R. A,, 331, 336, 337, 338 BERNER, BERRY,E. W., 401, 428 BERRY,W. B. N., 144 BERTRAND, J. P., 14 K., 235, 243, 249 BIRKENMAJER, BIRSE,D. J., 338 BISSELL, H. J., 144 RISSELL, H. J. and CHILINGAR, G. V., 17, 27, 51, 83, 97, 98, 110, 113, 119, 120, 144, 170, 249, 292, 339 BLACK,M., 60, 83, 236 BLACK,M. and BARNES, B., 235, 299 BLACK,W. W., 220, 250 BLACKWELDER, E., 339 BLANCKENHORN, M., 175, 250 BLOCH,M. R., LITMAN, H. Z. and ELAZARIVOLCANI, B., 200, 250 BRADLEY, W. H., 206, 250 BRAMKAMP, R. A. and POWERS,R. W., 93, 104, 105, 109, 111, 117, 120, 144 BRAMLETTE, M. N., 40, 47 BRANNER, J. C., 250, 274, 296, 339 BRAUN,M., 315, 319 BRETZ,J. H., 407, 428 BRIDGE, J., 339 BRIGGS,L. I., 312, 339 BRINKMANN, R., 310, 320, 339 W. S., 403, 428 BROECKER, BROECKER, W. S . , TUREKIAN, K. K. and HEEZEN, B. C., 234, 250,410, 428 BROWN,C. N., 198, 250,428 BODENHEIMER, W. and NEEV,D., 199, 250 0. B., 339 BOGGILD, BOND,G., 220, 250 BORCHERT, H. and MUIR, R. O., 10, 27, 413, 428 J. G., 185,250 BORNEMANN, BOUMA, A. H., 83, 104, 144 BOUMA, A. H. and NOTA,D. J. G., 64, 83

436 BOWEN,H. J. M., 407,417, 428 BOYD,D. W., 222, 250 BUDZINSKI, H., 307, 339 B~JSBY, R. F., 30, 47 BUTLER,G. P., KENDALL,C. B. ST. C., D. J. and KINSMAN,D. J. J., SHEARMAN, SKIPWITH,P. A. D’E., 221, 250, 287, 339 BYRNE,J. V., LEROY,D. 0. and RILEY,C. M., 228, 250

REFERENCES INDEX

CLOUDJR., P. E. and ABELSON, P. H., 417,428 CLOUDJR., P. E. and BARNES, V. E., 227,251, 272, 281, 296, 299, 310, 320, 339 COHN,F., 177, 251 COLEMAN, A. P., 247, 252 CONOLLY,J. R. and EWING,M., 242, 252 COOPER, B. N., 273, 281, 300, 320, 339 COUSTEAU, J. 226, CRICKMAY, G. W., 145, 191, 252 CROSS, A. T., 210, 252 CADY,W. M., 246, 250, 321, 339 CROSS, A. T. and SCHEMEL, M. P., 210,231,252 CALDWELL, W. G. E. and CHARLESWORTHCULLIS,C. G., 300, 323, 339 H. A. K., 220,250 CUMMINGS, E. R., 99, 145, 220, 252 CALVERT, W. L., 331, 332, 339 CUMMINGS, E. R. and SHROCK, R. R., 220,252 CARLSTON, C. W., 212, 250 CURTIS,R., EVANS, G . and KINSMANN, D. J. J., CAROZZI, A. V., 15, 54, 55, 59, 62, 83, 105, 107, 339 111, 144, 152, 184, 185, 195, 202, 208, 220, CURTIS,R., EVANS,G., KINSMANN, D. J. J. 236, 239, 250,292, 339 and SHEARMAN, D. J., 40, 48, 220, 252, 285 CAROZZI,A. V. and HUNT, J. B., 220, 251 CUVILLIER, J., 15, 54, 59, 83 CAROZZI, A. V. and LUNDWALL, W. R., 83,252 CAROZZI, A. V. and SODERMAN, J. G. W., 144, D’ALBISSIN, M., 16, 60, 84 220, 227, 251 DAETWYLER, C. C. and KIDWELL, A. L., 34,48, CAROZZI,A. V. and ZADNIK.V. E., 328, 339 65, 83, 145, 217, 252 CAYEAUX, L., 15, 53, 56, 59, 83, 107, 144, 172, DALY,R. A., 25, 411, 412, 424, 429 204, 251, 307, 339 DANE,C. H., 205, 206, 207,252 CHAMBERS, C. O., 174,251 DARWIN, C., 171, 252 CHAMBERLIN, T. C., 422, 428 DAVID,T. W. E. and BROWNE, W. R., 417,429 CHANDA, S. K., 144 DAWSON, J. W., 247, 252 CHANEY, R., 401, 428 DE CHARPAL, O., MONTADERT, L., ROUGE,P. CHAPMAN, F., 199, 251 and GUBLER, Y., 52, 55, 56, 62, 84 CHAVE,K. E., 2, 27, 39, 40, 47, 144, 178, 179, D E ~ ~ E YK. E SS., , LUCIA,F. J. and WEYL,P. K., 251, 275, 334, 335, 339, 408, 409,428 201, 252, 289, 329, 335, 336, 340 CHAVE,K. E., DEFFEYES, K. S., WEYL,P. K., DEFFEYES, K. S. and MARTIN,E. L., 40, 48, GARRELS, R. M. and THOMPSON, M. E., 39, 287, 339 48 DEFORD,R. K., 100, 101, 102, 145 CHILINGAR, G. V., 7, 17, 22, 27, 35, 47, 83, 89, DEFORD, R. K. and WALDSCHMIDT, W. A,, 145 92, 108, 144, 145, 248, 301, 307, 339, 366, DEGENS, E. T., 240,292, 326, 413, 429 374, 381, 397, 408, 409, 411, 412, 413, 420, DEGENS,E. T. and EPSTEIN,S., 277, 292, 340 428 DEGENS,E. T., WILLIAMS, E. G. and KEITH, CHILINGAR, G. V. AND BISSELL,H. J., 4, 7, M. L., 210, 252 89, 94, 104, 111, 144, 412, 417, 428 DESMIDT, P., 65, 84 CHILINGAR, G. V., BISSELL, H. J. and WOLF, DIEBOLD, F. E., LEMISH, J. and HILTROP, C. L., K. H., 5, 18, 27 276, 340 CHILINGAR,G. V. and TERRY,R. D., 145 DIETRICH, G., 1 I , 27 CHOQUETTE, PH. W. and TRAUT,J. D., 354, DIETRICH,R. V., HOBBSJR., C. R. B. and 355, 359, 385, 386, 387, 388, 397 LOWRY,W. D., 145 CHOUBERT, B. and BOY&M., 424,428 DIETZ,R. S., 427,429 CHUBB,L. J., 230, 251 DIXON,E. E. L., 248, 252, 296, 300, 320, 340 CLARK,T. H., 246,251 DODD,J. R., 39, 48 CLARKE, F. w., 193, 251, 307, 339 DOEGLAS, D. J., 54, 84 CLARKE,F. W. and WHEELER, W. C., 40, 48 DORF, E., 401,429 CLEE,V., 271, 339 D o n JR., R. H., 224, 252 CLINE,L. M., 245, 246, 247, 248, 251 DUNBAR,C. 0. and RODGERS, J., 227, 246, CLOUD JR., P. E., 35, 42, 44,48, 53, 55, 59, 62, 247, 252, 273, 309, 320, 321, 328, 330, 333, 65, 83, 145,218,219,220,251,275, 339,400, 340 428 DUNCAN, P. M., 220, 226, 252

REFERENCES INDEX

437

FERM,J. C. and WILLIAMS, E. G., 253 FIELD, R. M. and HESS,H. H., 340 A. G., 25, 340, 419, 429 FISCHER, FISHER, D. W., 315, 317, 340 D. W. and HANSON,G. F., 316, 331, FISCHER, 340 FLINT,R. F., 401, 429 A. J., 95, 145, 196, 201, 202, 252 EARDLEY, J. E., and RODGERS, FLINT,R. F., SANDERS, EARDLEY, A. J. and GVOSDETSKY, V., 202,252 J., 245, 253 A. J., GVOSDETSKY, V. and MARSELL, FOLK, EARDLEY, R. L., 52, 62, 73, 84, 93. 99, 100, 102, R. E., 201, 253, 292, 340 105, 111, 116, 117, 120, 146, 161, 166, 175, EATON,A,, 340 184, 191, 195, 253, 313, 340 EDIE,R. W., 340, 380, 381, 382, 383, 397 C., 64, 84 FONDEUR, EGYED,L., 425, 427, 429 S. O., 219, FORMAN,M. J. and SCHLANGER, EHLERS,G. M., 317, 340 253 ELLIAS,G. K., 145, 385, 386, 397 FREEMAN, T., 215, 253 ELLIOTT,R. H. J. and KIM, 0. J., 220, 253 FRIEDMAN, G. M., 16, 19, 27, 60, 84, 178, 179, W. M., 25 ELSASSER, 180, 182, 183, 184, 186, 187, 188, 190, 191, EMILIANI, c.,401, 429 193, 200, 212, 224, 226, 229, 232, 233, 237, T., 42, 48 EMILIANI, C. and MAYEDA, 239, 242, 243, 253, 276, 278, 279, 281, 287, EMERY, K. O., 219, 253 288, 289, 296, 299, 301, 302, 306, 309, 315, R. E., 317, 340 EMERY,K. 0. and STEVENSON, 316, 319, 327, 330, 334, 335, 336, 340, 341 EMERY,K. O., TRACEYJR., J. I. and LADD, FRIEDMAN, G. M. and NEEV,D., 199, 201, H. S., 219, 226, 253 253, 291, 296, 336, 341 ENOS,P. P., 190, 244, 253, 340 FRIEDMAN, G. M. and SANDERS, J. E., 20, 21, H. A., 41, 48 EPSTEIN,S. and LOWENSTAM, 146,208, 253 EREMENKO, N. A., 23, 27 FROLOVA, E. K., 90, 108, 146 H., 423, 425, 429 ERHART, D. B., EWING,M., WILLIN,G. and GARDINER, ERICSON, J. S., 274, 341 HEEZEN,B. C., 237, 239, 253 B., DE CHARPAL, 0. L., MONTAGARREAU, J., 16, 60, 84 ETIENNE, DERT,L., GUBLER,Y. G., ROUGE,P. E., EVAMY, B. D., 301, 340 BARON,G. A. and FAVRE,J. H., 274, 341 EVANS, G., KINSMAN, D. J. J. and SHEARMAN, GARRELS, R. M., THOMPSON, M. E. and SIEVER, D. J., 65, 84 R., 56, 84 EVANS,J. W., 199, 253 GAUB,F., 254 EVANS,N., 322, 340 GEORGE, T. N., 219, 254 EWING,M. and DO", W. L., 403, 429 T. N. and OSWALD, D. H., 220, 254 GEORGE, GEVERS, T. W., 326 J. A,, 297, 340 FAGERSTORM, GILL,D., 317, 341 R. W., 2, 4, 5, 6, 7, 8, 14, 15, 21, GILLILAND, FAIRBRIDGE, W. N., 208, 254 24, 25, 27, 52, 53, 54, 55, 84, 145, 157, 190, GILLOTT,J. E., 341 198, 220, 248, 253, 271, 273, 274, 275, 281, GILLULY,J., 417, 422, 429 307, 320, 321, 327, 329, 340, 400, 401, 402, GLINTZBOECKEL, C. H. and RABATE, J., 341 403, 404, 406, 410, 417, 423, 424, 425, 426, GINSBURG, R. N., 35, 42, 44, 48, 55, 84, 146, 187, 189, 218, 254, 315, 341, 416, 429 429 FAIRBRIDGE, R. W. and KRBBSJR., 0. A., 424, R. N. and LOWENSTAM, H. A,, 146 GINSBURG, M. F., 417,429 GLAESSNER, 429 R. W. and TEICHERT,C., 216, GLAESSNER, FAIRBRIDGE, M. F. and DAILY,B., 417, 430 253, 407, 429 GLANGEAUD, L., 84 FAY,R. O., 322, 340 L. and BOBIER,C., 51, 422, 430 GLANGEAUD, FEELY,H. W. and KULP,J. L., 192, 253, 291, GOLDBERG, M., 317 340 M. I., 40, 48, 179, 254, 341 GOLDMAN, FENTON,C. L. and FENTON,M. A,, 215, 220, GOLDSMITH, J. R. and GRAF,D. L., 275, 276, 341 253 W. G., GOLDSMITH, FERAY,D. E., HEUER,E. and HEWATT, J. R., GRAF,D. L. and JOENSUU, 94, 95, 98, 145, 195, 253 0. I., 275, 341 DUNHAM, K. C., 231, 252 R. J., 94, 116, 145, 162, 191, 199, DUNHAM, 222, 23 1,252, 357, 397 DUNN,J. R., 331 J. W., 401, 429 DURHAM, DZULYNSKI, S., 240

REFERENCES INDEX

GOREAU, T. F., 217, 254 GDRSLINE, D. S., 146, 224, 254 GOULD, H. R. and MCFARLAN JR.,E., 228,254 COULD,H. R. and STEWART, R. H., 224, 254 GRABAU, A. W., 88, 93, 94, 98, 99, 146, 172, 185, 199, 220, 254 GRAF,D. L., 29, 48, 64, 84, 90, 146, 226, 254, 313, 341 GRAF,D. L. and LAMAR, J. E., 360, 397 GRAF, D. L., EARDLEY, A, J. and SHIMP, N. F., 7, 27, 146, 292, 341 J. R., 64,84,275, GRAF,D. L. and GOLDSMITH, 307, 341 GREENSMITH, J. T., 146, 162 GREGORY, J. W., 242, 254 GREINER, H. R., 146, 311, 341 GRESSLY, A,, 15, 54 GRIM,R. E., KULBRICKI, G. and CAROZZI, A. V., 7, 27 GROSS,M. G., 42, 48, 295, 297, 341 GROUT,F. F., 177, 254 GUBLER, Y., 14, 16, 146, 170 GUBLER, Y. and BERTRAND, J. F., 64, 84, 254 GUERRERO, R. G. and KENNER, C. T., 89, 146 Gussow, W. C., 220, 254

HENSON, F. R. S., 219, 255 HESS,H. H., 237, 255 HEWETT, D. F., 275, 332, 333, 334, 342 HILDE,J. C., 326, 342 HILLS,J. M., 243, 255 HOBBS JR., C. R., 146, 297, 298, 300, 329, 342 HOHLT,R. B., 360, 397 HOLLAND, H., 25, 416, 430 HOLLINGSWORTH, S. E., 424, 430 HOLMES, A., 193, 255, 422, 430 HOUBOLDT, J. J. H. C., 65, 84 HOUGH,J. L., 417, 430 HOWARD, W. V. and DAVID,M. W., 342 HOWE,M. A., 215, 255 HOWELL, J. B., 101, 102, 146 HOPKINS, R. H., 342 HUDSON, R. G. S., 231, 255 HUNT,J. M., STEWART, F. and DICKEY, P. A,, 215, 216, 217, 255 HUTCHINSON, G. E., 408, 417, 430 HUXLEY, T. A., 235, 255

ILLING,L. V., 35, 40, 44, 48, 65, 84, 96, 146, 172, 182, 184, 185, 187, 189, 202, 215, 222, 225,255,285, 298, 309, 31 1, 342 ILLING, L. V. and WELLS, A. J., 201, 255, 285, 288, 342 HADDING, A., 172, 21 6, 220,254 IMBRIE, J. and KORNICKER, L., 298, 342 HALL,C. W. and SARDESON, F. W., 341 IMBRIE, J. and NEWELL, N., 421, 430 HALLA,F. und RITTER, F., 336, 341 IMBRIE, J. and PURDY,E. G., 14, 34, 48, 52, HALLA,F., CHILINGAR, G. V. and BISSELL, 84, 96, 105, 117, 146, 313 H. J., 146 IMBT, W. C. and ELLISON, S. P., 369, 370, 371, HALLAM, A., 424, 430 397 HAM,W. E., 117, 146, 174,213,223, 232, 233, INGERSON, E., 44, 48, 201, 255, 275, 307, 342 254, 255, 291, 298, 309, 311, 322, 329, 330, IRWIN JR., C. D., 355, 397 341 IVANENKO, D. D. and SAGITOV, M. U., 427,430 HAM,W. E. and PRAY,L. C., 255 HAMBLETON, A. W., 146 JACOBSEN, L., 210, 213, 255 HARBAUGH, J. W., 21, 146, 177, 194, 220 JANSEN, J. F. and KITANO, Y., 39, 48 HARDMAN, E. T., 341 JENKINS JR., M. A., 146 HARDY, F., 326, 342 JOHNS,R. K. and LUDBROOK, N. H., 293, 342 HARVEY, H. W., 414, 430 JOHNSON, G. A. L., 231, 255 HASSACK, C., 174, 255 JOHNSON, J. H., 147, 175, 191, 215, 255, 256 HATCH,F. H., RASTALL, R. H. and BLACK, JOHNSON, J. H. and KONISHI,K., 215, 256 M., 176, 189, 204, 236, 255, 271, 278, 296, JONES, B. F., 283, 284, 290, 342 320, 333, 342 JONES,G. E., STARKEY, R. L., FEELEY, H. W. HATHAWAY, J. C., SCHLEE, J. S., TRUMBULL, and KULP,J. L., 342 J., 226,255 J. V. A. and HULSEMANN, JUKES-BROWNE, A. J., 236, 256 HAWLEY, D., 247,255 Jux, U., 220, 256 HAYES, P. T., 342 HECKER, R. F., 421, 340 KAFKA,F. T. and KIRKBRIDGE, R. K., 72, 85 HEDGPETH, J. W., 226, 229, 255,358, 397, 421, KAHLE, C. F., 294, 342 430 KALKOWSKY, E., 256 HEEZEN, B. C., 427, 430 KANE,H. E., 228, 256 HEEZEN, B. C., THARPE, M. and EWING,M., KATZ,A., 302 237, 255 KAY,G. M., 247, 256, 316, 342, 423, 430

REFERENCES INDEX

KHVOROVA, I. V., 101, 147, 301, 342 KING,D., 293, 335, 342 KING,P. B., 222, 243, 256, 342 KING, R. H., 281, 342 KINDLE,C. H., 256 KINDLE,C. H. and WHITTINGTON, H. B., 246, 247, 256 KINSMAN, D. J. J., 189, 256, 342 KLEIN,G. D., 216, 257 KLEIN,G. D. and SANDERS, J. E., 227, 257 KLAHN,H., 173, 256 KLOVAN, J. E., 394, 395, 396, 397 KOLDEWIJN, B. W., 229, 257 KORNICKER, L. S., 147 KRAMER, W. B., 245, 247, 248, 257 KRINSLEY, D., 39, 48, 410, 430 KRINSLEY, D. and BIERI,R., 39, 48 KRUMBEIN, W. C. and SLOSS,L. L., 51, 85, 119, 147 KRYNINE, P. D., 85, 102, 147, 208, 209, 257, 309, 342 KSIAZKIEWICZ, M., 240 KUENEN, PH. H., 25,26, 85, 239, 342,411,430 KUENEN, PH. H. and CAROZZI, A. V., 239,243, 257 KUENEN, PH. H. and MIGLIORINI, C. I., 242, 257 KUENEN,PH. H. and TENHAAF,E., 239, 257 KUGACHKOV, D. M., 326, 342 KULP,J. L., 422, 430 KUMMEL, H. R., 212, 257 LADD,H. S., 430 LADD,H. S., HEDPETH,J. W. and POST,R., 171, 226, 257 LADD,H. S., INGERSON, E., TOWSEND, R. C., RUSSELL,M. and STEPHENSON, H. K., 343 LADD,H. S.,TRACEY JR.,J. I. and LILL,C. G., 343 LADD,H. S., TRACEY JR., J. I., WELLS,J. W. and EMERY, K. O., 257 LALOU,C., 64, 85, 147, 324, 343 LAMING, D., 212 LAMPLUGH, G. W., 175,257 LANDES, K. K., 343 LANE,A. C., 412, 430 LANG,W. B., 243, 257, 343 LANGOZKY, Y., 204, 257 LAPORTE, L. F., 147, 314, 315, 343 LAROCQUE, A,, 207, 208, 257 LAYER,D. B., 220, 257 LE BLANC,R. J. and BREEDING, J. G., 27, 48, 49, 84, 85, 145, 146, 147, 148, 259, 341 LECOMPTE, M., 55, 64, 85,220,226, 257 LEIGHTON, M. W. and PENDEXTER, C., 17, 85,

439 91, 92, 94, 95, 96, 99, 100, 105, 106, 109, 112, 113, 114, 116, 117, 147, 161, 166, 185, 195,257 LEITH,C. K. and MEAD,W. J., 193, 258 LEPP,H. and GOLDICH, S. S., 417, 430 LESLEY, J. P., 8, 27, 320, 343 LEVET,M. N., 220, 258 LINCK,G., 182, 258, 271, 343 LINDGREN, W., 361, 397 LINK,T. A., 220, 258 LLOYD,E. R., 243, 258 LOEWENGART, S., 233,258 LOGAN,B., 416, 430 LOGAN, B. R., REZAK, R. and GINSBURG, R. N., 315, 343 LOMBARD, A., 55, 62, 64, 85 LONG,G., NEGLIA, S. and FAVRETTO, L., 80,85 LONGWELL, C. R., 207, 211, 258, 313, 343 LOWENSTAM, H. A,, 39, 49, 99, 147, 220, 258, 408, 430 LOWENSTAM, H. A. and EPSTEIN, S., 42,49,147 LOWMAN, S. W., 245,258 LUCIA,F. J., 147, 187, 298, 301, 343, 363, 364, 397 LUCIA,F. J., WEYL,P. K. and DEFFEYES, K. S., 40,48 LUDWICK, J. C. and WALTON, W. R., 226,258 LYELL,C., 171, 176, 258

MACLAREN, C., 424, 430 MACNEIL,F. S., 219, 258 MAGDEFRAU, K., 343 MAKHLAEV, V. G., 302, 343 MAMET, B., 63, 64, 85 MARCHER, M. V., 147 MARSHALL, P., 25, 219, 258, 343 MATTAVELLI, L., 14, 72 MAWSON, D., 215, 258 MAXWELL, W. G. H., DAY,R. W. and FLEMING, P. J. G., 34, 147 MAXWELL, W. G. H., JELL, J. S. and McKELLAR, R. G., 49 MAYOR,A. G., 343 MCCOMAS, M. R., 352, 397 MCGREGOR, D. J., 313, 343 MCGERRIGGLE, H. W., 258 MCINTIRE, W. G., 233, 258 MCKEE,E. D., 11, 12, 27, 219, 258, 320, 343 MCKINLEY, M. E., 109,147 MEAD,W. J., 193, 258 MEDLIN,W. L., 307, 343 MERRILL, G. G., 258 MIDDLETON, G. V., 222, 258 MILANKOVICH, M., 401, 403, 430 MILLERJR., D. N., 40, 49 MILLER JR., H., 231,258

440

REFERENCES INDEX

MILLER,R. L. and BROSGE, W. P., 343 MISER,H. D., 247, 248, 258 MIS~K,M., 91, 120, 147 MOJSISOVICS VON MOJSVAR, E., 220, 258 MOLLAZAL, Y.,90,91, 100, 101, 102, 111,120, 147 MOORE,D., 231, 258 MOORE,G. W., 19, 177, 192, 223, 244, 259 MOORE, G. W. and HAYES, P. T., 221,259, 337, 343 MOORE,R. C., 147, 231,245,247, 248, 259 MORETTI, F. J., 147 MORRIS, R. C. and DICKEY, P. A., 147,221,259 MOSHER, L. C. and PINNEY,R. I., 90, 94, 95, 106, 116, 147 MOTTS,W. S., 222, 259 MUELLER, G., 281, 288, 343 MURRAY, A. N., 273, 343 MURRAY, J. and TRVINE, R., 219, 259 MURRAY, J. and RENARD, A,, 235,259 MURRAY, R. C., 65,85,147,299,343,352,355, 356, 362, 363, 365, 366, 368, 371, 375, 398 MYERS,D. A,, STAFFORD, P. T. and BURNSIDE, R. J., 220, 259, 382, 384, 398

OHLEJR., E. L., 297, 333, 344 OLIVER JR., W. A., 220, 260 OPPENHEIMER, C. H. and MASTER, I. M., 325 344 OSBORNE, F. F., 247, 260 OSMOND, J. C., 297, 321, 344 OSWALD, D. H., 220, 260 OTTEJ.R, C. and PARKS Jr., J. M., 220, 260 OXLEY, PH., 220, 260 OXLEY,PH. and KAY,G. M., 220, 260

ODER,C. R. L. and BUMGARNER, J. G., 215, 259 ODER,C. R. L. and MILLER,H. W., 333, 344

RAINWATER, E. H., 229, 260 RALL,R. W. and RALL,E. P., 220, 260 RAMDOHR, P., 416, 431

PARKINSON, D., 148, 220, 260 PARSONS, L. M., 333, 344 PASSEGA, R., 14, 62, 72, 85 PERKINS, R. D., 148, 224, 260, 317, 344 PETERSON, M. N. A,, 225, 260 PETERSON, M. N. A. and BIEN,G. S., 283,284, 344 PETERSON, M. N. A,, BIEN,G. S. and BERNER, R. A., 344 PETERSON, M. N. A, and VON DER BORCH, C. C., 283, 344 F. J., 89, 102, 107, 148, 167, 193, PETTIJOHN, 260,412, 430 PFLUG,H. D., 417, 430, 431 NANZJR., R. H., 412, 430 PHILIPPI, E., 344 NEEV,D., 192, 199, 200, 201, 259, 290, 291, PIA,J., 173, 177, 215, 260 292, 336, 337, 344 PICARD,L., 204, 233, 248, 260, 333, 344 NEHRER, J. and ROHRER, E., 325, 344 PICARD,L. and AVNIMELECH, M., 233, 260 NELSON, H. F., BROWN,C. W. and BRINEMAN, PICARD,L. and SOLOMONICA, P., 233, 260 J. H., 95, 99, 117, 147, 166, 175, 220, 259 PILKEY, 0. H., 44, 49 NELSON,R. J., 231, 259 PILKEY,0. H. and GOODELL, H. G., 39, 49 NESTEROFF, W., 64, 85 PILKEY, 0. H. and HOWER,J., 39, 49 NEWELL, N. D., 99, 147, 175, 233, 243, 259 PIRSSON, L. V. and KNOPF,A., 177, 260 NEWELL,N. D. and RIGBY,J. K., 52, 53, 60, PLUMLEY, W. J. and GRAVES JR., R. W., 220, 65, 85, 147, 217, 259 260 NEWELL, N. D., IMBRIE, J., PURDY,E. G. and PLUMLEY, W. J., RISLEY,G. A., GRAVES JR., THURBER, D. L., 35, 49, 148, 218, 259 R. W.and KALEY, M. E., 17, 104, 105, 116, NEWELL, N. D., PURDY,E. G. and IMBRIE, J., 148, 197, 260 182, 184, 259 POTTER,P. E. and PETTIJOHN, F. J., 55, 85 NEWELL,N. D., RIGBY,J. K., FISHER,A. G., POWERS, R. W., 53, 55,93, 111, 117, 120, 148, WHITEMAN, A. J., HICKOX,J. E. and BRAD175, 176, 186, 187, 190, 216, 223, 247, 291, LEY,J. S., 148, 259, 343 329, 344, 366, 367, 398 NEWELL, N. D., RIGBY,J. K., WHITEMAN, A. J. POWERS, S., 260 and BRADLEY, J. S., 148, 192, 220, 222, 298, PRAY,L. C., 148, 227,260 309, 310, 328 PRAY,L. C. and WRAY,J. L., 94, 116, 148 NICHOLAS, R. L., 193, 259 PUGH,W. E., 220, 227, 260 NINO,H. and EMERY,K. O., 229, 259 PURCELL, W. R., 374, 398 NOLL,W., 49 PURDY,E. G., 34, 44, 49, 191, 260 NORRIS,R. M., 226, 259 PURDY, E. G. and MATTHEWS, R. K., 217,260 NOTA,D. J. G., 229,259 Nuss, W. F. and WHITING,R. I., 371, 398 QUIRING, H., 329, 344

REFERENCES INDEX

44 1

RANKAMA, K., 412, 431 SANDO,W. J., 296, 345 RAYMOND, L. R., 344 SARIN,D. D., 148, 345 RAYMOND, P. E., 220, 246, 261 SASS, E., 307, 345 RAYNER, D. H., 231, 261 R. W., 199, 231, 262, 407 SAYLES, REEVES JR., C. C. and PARRY,W. T., 293, 344 SCHEIDEGGER, A. E., 372, 398 RENZ,o.,LAKEMAN, R. and VAN DER MEULEN, SCHERMERHORN, W. I., L. J. G. and STANTON, E., 245, 248, 261 43 1 REULING, H. T., 273, 281, 329, 344 SCHINDEWOLF, 0. H., 417, 431 REVELLE, R., 235, 250, 261 S. O., 148,175,186,190,191,219, SCHLANGER, REVELLE, R. and FAIRBRIDGE, R. W., 30, 49, 262, 295,296, 345 148, 163, 189,261,405,414 SCHMIDT DI FRIEDBERG, P., 75, 85 REVELLE, R. and FLEMING, R. F., 35, 49 SCHMITT, J., 16, 60, 61, 85 RICH, M., 104, 105, 111, 117, 120, 148 SCHNEIDER, E. and HEEZEN,B. c . , 239, 241, RICKARD,L. V., 225, 261, 314, 344 262 RICOUR,J., 325, 344 SCHOLL, D. W., 177, 262 M., 376, 398 RIEKMANN, SCHREIBER, J. F., 202 F., 72, 85 RIGODE RIGHI,M. and BARBIERI, SCHUCHERT, CH., 193, 227,262 RITTENHOUSE, G., 320, 344 G. A,, 262 SCHUCHERT, CH. and COOPER, RIVI~RE, A., 64, 85, 344, 345 SCHUCHERT, CH. and DUNBAR, C. O., 246,247, RIZZINI,A., 14, 72 262 RIZZINI,A. and MATTAVELLI, L., 75, 85 N., 204,262 SCHULMAN, ROBERTSON, T., 231, 261 SCHWADE, I. T., 345 ROBINSON, E., 242, 261 SCHWARZACHER, W., 148, 220, 262 Rocco, T., 75, 85 SCHWARZBACH, M., 401, 403, 412, 431 J., 30, 31, 49, 53, 85, 107, 113, 148, SCRUTON, RODGERS, P. C., 10,27,221,223,262,312,335, 235, 261, 323, 345 345 RONOV,A. B., 312, 313, 345 SHALEM, N., 200, 262 ROSE,G., 261 SHALER, N. S., 224, 262 Ross, C. A., 261 SHAW,A. B., 227, 246, 262, 322, 345 Ross, C. A, and OANA,S., 148 D. J., 285,287, 345 SHEARMAN, Ross, M. H., 246, 261 D. J., KHOURI,J. and TAHA,S., SHEARMAN, ROTHPLETZ, A,, 261 302, 345 RUBEY, W. W., 25 SHEPARD, F. P. and MOORE,D. G., 226, 262 RUEDEMANN, R., 247, 261 G. D. and THIEL,G. A,, 326, 345 SHERMAN, RUHE,R. V., CADY,J. G. and GOMEZ,R. S., SHERMAN, G. D., SCHULTZ,F. and ALWAY, 407, 431 F. J., 301, 326, 345 RUKHIN,L. V., I , 27 SHINN,E. A. and GINSBURC, R. N., 40, 49, RUNCORN, R. K., 427,431 190, 262,287, 3 15, 327, 345 RUSNAK, G. A., 49, 148, 182, 184, 185, 229, SHINN,E. A., GINSBURG, R. N. and LLOYD, R. N., 345 239, 261 E. L., 220, 263 SHIRLEY, J. and HORSFIELD, W. D., 30, RUSNAK,G. A. and NESTEROFF, 239, 261 SHROCK, R. R., 220, 263, 345 A. L. and ~ S T L U N D SHVETSOV, , RUSNAK, G. A., BOWMAN, M. S., 52,85, 118, 119, 148,433,434 H. G., 239, 261 SIEGEL, F. R., 35, 44, 49, 50, 148, 307, 410, 431 RUSSELL, I. C., 176, 177, 189, 261 SINNOKROT, A. A. and CHILINGAR, G. V., 352, RUSSELL, R. J., 261 398 RUTTE,E., 407, 431 E. W., 109, 148, 346 SKEATS, RUTTEN,L. M. R., 422, 431 SKINNER, H. C. W., 27, 50, 281, 346 RUTTEN,M. G., 25,26, 220,226,262, 416,431 SKINNER, H. C. W., SKINNER, B. J. and RUBIN, M., 5, 281, 346 SABINS JR., F. F., 5, 27, 148, 324, 345 SLOSS, L. L., 194, 263, 346 SANDER, B. K., 55, 59, 85, 148, 172, 262,271, SMITH,J. P., 220, 263 278, 302, 345 SOKOLOV, D. S., 149 SANDERS, J. E., 15, 19, 113, 174, 196, 209, 212, SORBY,H. C., 166,171,172,173,184,204, 215, 213, 222, 225, 226, 227, 228, 236, 240, 241, 235, 263 242, 244, 262, 345 SPIEKER, E. M., 208, 263

442

REFERENCES INDEX

H. and TERMIER, G., 423, 425, 426, SPIEKER, E. M. and REESIDE JR., J. B., 207,263 TERMIER, 43 1 SPOTTS,J. H., 275, 346 TEXTORIS, D. A. and CAROZZI, A, V., 229,264 SPRENG,A. C., 231, 263 R., 228 THIENHAUS, STAUFPER, K. W., 149, 346 H. and McELTHODE,H. G., KLEEREKOPER, C. W., 222, 263 STEARN, CHERAN, D., 290, 347 STEHLI,F. G. and HOWER,J., 35, 44, 50, 149, G. E., 105, 149, 357, 389, 390, 391, THOMAS, 178, 179, 183, 263 392, 398 E., 271, 276, 309, 346 STEIDTMANN, G. E. and GLAISTER, R. P., 149, 347 STEPHENSON, L. W. and MONROE, W. H., 236, THOMAS, THOMAS, H. D., 94, 117, 149 263 THORP,E. M., 274, 347 STERNBERG, R. M., 86 TIDDEMAN, R. H., 220, 264 STETSON, H. C., 224, 263 STETSON, T. R., SQUIRES, D. F. and PRATT,R. TOWSE,D., 149, 297, 299, 347 TRACEYJR., J. I., LADD,H. S. and HOFFMEISTER, M., 226, 263 J. E., 264 STEWART, F. H., 313, 346 TROELL,A. R., 220, 227, 264 STILLE,H., 422, 431 TWENHOFEL, W. H., 220, 228, 264, 347 STODDART, D. R. and CANN,J. R., 189, 263 JR., W. W., 306 STOUT,J. L., 357, 358, 361, 375, 376, 377, 398 TYRELL UDDEN,J. A., 347 STOUT. W., 346 UDLUFT,H., 347 STRAHAN, A,, 346 STRAKHOV, N. M., 5, 27, 55, 86, 149,274, 312, ULRICH,E. O., 247, 264 UMBGROVE, J. H. F., 220, 264, 405,431 313, 346, 412, 425, 431 N. M., BRODSKAYA, N. C . , KNYA- UREY,H. C., 25 STRAKHOV, H. A., EPSTEIN, S. ZEVA, L. M., RAZZHIVINA, A. N., RATEEV, UREY,H. C., LOWENSTAM, C. R., 41, 50 M. A., SAPOZHNIKOV, D. G. and SHISHOVA, and MCKINNEY, E. S., 23, 27 VANSTRAATEN, L. M. J. U., 197,226,227,228, STRUVE, W., 346 264 STUBBLEFIELD, C. J., 226, 263 VAN TUYL,F. M., 149,271, 297, 307, 320,341 SUESS,E., 423, 424, 431 VANTUYL,F. M. and STEIDTMANN, E., 271, SUFFEL,G. G., 322, 346 322, 328 SUGDEN, W., 149, 221, 222, 263,284, 346 V A N WATERSCHOOT VAN DER GRACHT, w. A. SUJKOWSKI, ZB. L.. 149, 235, 263 J. M., 247, 264 SUMMERSON, C. H., 317, 346 VATAW, A,, 149 SWANN, D. H., 230, 263 VAUGHAN, T. W., 2, 28, 40, 50, 197 SWINCHATT, J. P., 198, 263 VERSEY, H. R., 216, 236, 264 A., LEONARD, A. B. and FRYE,J. SWINEFORD, VINOGRADOV, A. P., 86, 149 C., 199, 263, 346, 407, 431 VINOGRADOV, A. P., RONOV,A. B. and RATYNSKIY, V. M., 411, 412, 414, 431 S. G., 108, 149 VISHNYAKOV, TAFF,J. A., 263 TAFT,W. H., 5, 13, 28, 40, 47, 50, 62, 65, 86, VOGEL,J. C. and EHHALT, D., 413 149, 170, 177, 178, 191, 194, 195, 215, 247, VOIGT,E. and HANTSCHEL, W., 242,264 263,287, 346 VON DER BORCH,C., 281, 347 TAFT,W. N. and HARBAUGH, J. W., 5, 28, 35, VON DER BORCH, C. C., RUBIN, M. and 40, 42, 44, 47, 50, 52, 119, 149 SKINNER, B. J., 5, 28 TARR,W. A., 309, 346 VON ENGELHARDT, w., 336, 338, 347 VON GUEMBEL, C. W., 172, 264 TATARSKIY, V. B., 301, 346 VONMORLOT,A,, 272, 301, 347 TAYLOR, R. E., 263 TEALL,J. J. H., 272, 346 TEICHERT, WAGER, L. W., 401,431 C., 216, 226, 263, 421, 431 WALCOTT, TENHAAF,E., 239, 264 C. D., 245, 246, 264, 416, 431 TENNANT, W. A., 347 C. B. and BERGER,R. W., 276, 346 WALDSCHMIDT, TEODOROVICH, G. I., 7, 17, 28, 85, 89, 92,97, WALDSCHMIDT, W. A., FITZGERALD, P. E. 98,107,109,110,111,119,120,149,274,281, and LUNSFORD, C. L., 398 346, 341, 373, 398 WALKER, T. R., 149, 193, 264 TERCIER,J., 242 WALLACE, R. C., 347

REFERENCES INDEX

WALPOLE,R. L. and CAROZZI, A. V., 264 WALTHER, J., 172, 197, 264 WALTHER, J. and SCHIRLITZ,R., 265 WANLESS,H. R., TUBEJR., J. B., GEDNETZ, D. E. and WEINER,J. L., 209, 230, 265 WANGERSKY, P. J., 410, 431 WARDLAW, N. C., 149 WARNE, S., 86 WEBER,J. N. and SMITH,G . F., 276, 347 WEED,W. H., 177, 265 WEGMANN, E., 54, 86 WELLER, J. M., 209, 230, 265 WELLS,A. J., 40, 50, 149, 221, 265, 285, 347 WENGERD, S. A., 220, 265 WENTWORTH, C. K., 100, 149 WEYL,P. K., 359, 360, 365, 366, 398, 413, 431 R., 347 WEYNSCHENK, WHEELER, H. E., 424, 432 WHITE,D. E., 336, 347 WHITTINGTON, H. B. and KINDLE,C. H., 246, 265 F. E., 416, 432 WICKMAN, WILCKINS, O., 347 WILLARD, B., 347 WILLMAN, H. B., 328, 347 WINDER, C. G., 225, 316, 347

443 WINTERER, E. L. and MURPHY,M. A,, 244, 265, 299, 328, 348 WISEMAN,J. D. H. and TODD, I., 410, 432 WOLF, K. H., 50, 86, 95, 101, 105, 106, 149, 184, 195, 215, 220, 265 C . V. and BEALES, WOLF,K. H., CHILINGAR, F. W., 40, 52, 175 E. B., 220, 265 WOLFENDEN, WOOD,A., 215, 265 WOODWORTH, J. B., 247, 265 WOOLNOUGH, W. G., 175, 265, 407, 432 C. R., 293, 348 WOPFNERH. and TWIDALE, YONGE,C. M., 219, 265, 421, 432 YOUNGJR., F. B., 348 YOUNG,R. B., 316, 348 R. S., 209, YOUNG,R. G. and EDMUNDSON, 212, 265 ZANS,V. A., CHUBB,L. J., VERSEY,H. R., WILLLAMS, J. B., ROBINSON, E. and COOK, D. L., 217, 265 ZELENOV, K. K., 8, 28 ZELLER, E. J. and WRAY,J. L., 149 ZEN,E-An, 40, 50 ZIRKEL,F., 96, 149

SUBJECT INDEX1

Acadian orogeny, 422 _ _ travertine, 176 Accretionary, definition of, 150 -, lsmay zone, 385, 388, 389 Accumulations of carbonates in modern seas, -, Precambrian, 400, 412 4, 5, 3 0 4 0 -, precipitation of calcium carbonate, 174 Acid ocean theory, 412 -, Redwater reef complex, 394 Acropora cervicornis, 217 - reef rock, 175 - palmata, 21 7 -, susceptible to recrystallization, 191 Adelaide System, South Australia, 41 7 - zooxanthellae, 421 Aegean Sea, 11 Algal, 77, 79-81, 95, 96, 121, 122, 124-126, Africa, eolianites, 407 138, 150, 187, 285,286, 288, 295 Aggregation, definition of, 150 -, definition of, 150 Aiguilles-Rouges massif, France, 239 - dolomite, 77, 79, 80, 81 Alabama, U.S.A., 224, 236 -, dolomitization of, 295 -, Cretaceous chalks, 236 - dust, definition of, 150 Alaska, 40, 401 - flats, 285, 286, 288 Alberta, Canada, 224, 388 - limestone, 95, 96, 138 -, Mississippian deposits, 224 _ - , photomicrographs of, 138 -, map of, 388 - material, photomicrographs of, 121, 122, Albertite, 206 124-126 Algae, 4, 6, 7, 8, 12-14, 22, 24, 25, 35, 52, 65, - tissue removal, 187 77, 78, 93, 94, 96, 98, 108, 109, 174-177, 179, Alkalinity, 24, 26 184, 191, 202, 206, 217-219, 226, 231, 235, Allegheny Series, Pa., U.S.A., 209 241, 296, 314, 315, 354, 355, 359, 385, 388, Aller beds, Germany, 310 389, 394, 400,406-409, 412, 416, 419,421 Allochems, 56 -, barrier reef, 217 Allochthonous, 93, 150 -, coatings by, 184 -, definition of, 150 -, codiacean, 314 Allogenic, definition of, 150 -, coral bioherms, 226 Alluvial fans, 152, 205, 210 -, coralline, 179, 296 --, calcareous deposits, 210 -, cyanophytic, 24, 150, 41 6 _ _, caliche, 152 -, deposition of carbonates, 177 , sediments in Green River Formation, -, environmental temperature, 409 205 -, Florida Bay reef tracts, 218 Alpine orogeny, 422 -, fresh-water, 408 Alsen Formation, N. Y . , U.S.A., 225 -, Great Salt Lake, 202 Amazon River, South America, 405 -, Green River Formation, 206 Amphibole, 242 -, in beachrock, 406 Amphiporids, 390, 391, 396 -, incrusting organism, 175 Anzphiroa, 233 -, in eolianites, 231, 407 Amsterdam, N. Y . , U.S.A., 316, 317 -, - reef-wall material, 219 Anachis, 228 _ - sediments from Puerto Rico Trench, 241 Anadarn, 228

--

The editors are greatly indebted to Herman H. Rieke 111 and Mostafa Karim for their help in preparing the index.

SUBJECT INDEX

445

Anadiagenesis, 16 --/calcite ratio, 40 Anadiagenetic, 5, 150 -, clear coating, 185 -, definition of, 150 -, conversion to dolomite, 300, 336 - stage, 5 -, cryptocrystalline crystals, 184, 185 Andrews South Oil Field, Texas, U.S.A., 186, -, dissolution equation, 187 298, 364 -, formation of, 199, 200 Andros Island, Bahamas, 183, 217, 287, 315 -, Gulf of Eilat, 242 _ - ,Recent dolomite, 287 - in algal matter, 202 - platform, 35 _ - Bonneville Lake sediments, 292 -- chalks, 189 Anhedral, definition of, 150, 302 Anhydrite, 1, 7,9, 10, 12, 19, 80, 112, 189, 192, - _ playa sediments, 283, 284 221-223, 279, 280, 285, 288, 301, 302, 309- _ Recent sediments, 178, 286, 287 311, 313, 323, 357, 389 _ _ sebkhas, 285 - associated with syngenetic dolomite, 309 -, isotopic studies, 291 -, bacterial breakdown of, 192 -, mineralogical zonation, 282 -, Canada, 3 11 -, paramorphic replacement, 182 - cement, 189, 279, 280, 357 -, pisolites, 177 -, formation of, 302 -, precipitation of, 289 - infilling of voids, 389 -, Red Sea, 183 - in sebkhas, 285 -, removal by dissolution, 183, 184 -, pore cement, 357 -, replacement of, 294 -, Zechstein sequence, 310 -, Samra Formation, 204 Ankerite, 60, 88, 90, 193 - shell, formation of, 419 -, replacement of quartz, 193 - shells replaced by dolomite, 296 Annelida, 39 - speleothems, 177 Antarctica, icecap, 401, 402 -, stability of, 179, 186 Antillean sills, 11 -, strontium in, 410 Aphanic, 101, 150 -texture types, 185 -, definition of, 150 Arbuckle Limestone, Kans., U.S.A., 369-371 Aphanitic (see aphanic), 150 Arc0 Hills, Idaho, U.S.A., 121, 126 Aphanothece packardii, 202 Arctic Ocean, 11, 403 Aphthitalite, in sediments, 283 - transgressions, 426 Appalachian(s), 224, 320, 323, 422 Ardmore Basin, Okla., U.S.A., 213 - geosyncline, carbonate deposits in, 323 Arenicola marina, 227 - orogeny, 422 Arkose, corrosion of, 412 -, stratigraphic diagram of, 320 Armand Limestone, Appalachians, U.S.A., 321 Aptychus, 235 Avtemia gracilis, 202 - limestones, Europe, 234 Articulate, definition of, 150 Apuane Alps, Italy, 239 Asiatic Shelf, carbonate sediments on, 229 Arab Formation, Saudi Arabia, 190, 216, 223, Atlantic Ocean, 11, 4042, 226, 241, 404 298, 329 - _ , Bahamas, 4 0 4 2 _ _ , Arab-D Member, 190, 329 _ _ , Bermuda, 40-42 _ _ , dolomitization, 298, 329 _ _, deepwater coral banks, 226 - _ , interbedded carbonate rocks and an_ _, distribution of marine carbonates, 404 hydrite, 223 --, Florida Bay, 40 Arabian carbonate rocks, 104 _ _, Gibraltar Sill, 11 - reservoir rocks, 117 --,Puerto Rico Trench, 241 - Sea, 199 Aspen Range, Idaho, U.S.A., 122 Arabo-Nubian Shield, 242 Atoka, Okla., U.S.A., 245 Aragonite, 1, 13, 14, 16, 35, 39, 40, 42, 44-46, Atoll, 6, 16, 165, 171, 196, 219, 296, 382, 384 52, 88, 90, 171, 177-179, 182-187, 189, 199, -, Bikini, 219 200, 202, 204, 242, 268, 279, 280, 282-287, -, definition of, 165 289, 291,292, 294-296, 300, 324, 336,419 -, dolomitized core of, 6 - cement, 279, 280 -, Horseshoe, 382, 384 -, bacterial origin, 324 -, Indo-Pacific area, 171 -, basic properties, 171 -, Kapingamarangi, 219

446 Atoll (continued) -, Kita-daito-jima, 196, 296 Australia, 29, 33, 40, 216, 274, 281, 283, 293, 401, 416, 417 -, Adelaide System, 41 7 -, Coastal Limestone, 33, 216 -, Coorong Lagoon, 281, 283 -, Coward Springs, 293 -, Ediacara fauna, 417 -, eolianites, 407 -, Great Barrier reef, 29, 274 -, Lake Eyre, 293 -, Shark Bay, 416 Authigenesis, 15 Authigenic, 93, 150 - calcite, 93 -, definition of, 150 Autochthonous, 93, 150 -, definition of, 150 Autoclastic breccia, definition of, I50 Axiolites, definition of, 96 Axiolitic, definition of, 151 Bab el Mandeb, Red Sea, 11 Bacteria, 14, 18, 20, 22, 42, 174, 192, 290, 291, 324, 325, 338, 406, 416, 419 -, decomposition of gypsum, 192 -, formation of syngenetic dolomite, 324, 325 -, hydrogen sulfide production, 416 - in Precambrian soil, 419 -, mobilizing carbonates in soils, 406 -, precipitation of calcium carbonate, 174 -, reduction of sulfates, 290, 291, 338 Bacterial conversion, 192 Baffin Sea, 11 Bahama Islands, Atlantic Ocean, 19, 29, 34, 4042,65, 117, 119, 151, 172, 179, 180, 182184, 196, 202, 218, 233, 242, 284, 314, 315, 401, 409 - -, bank, 29, 34, 42, 151, 218 _ - , calcareous sands, 172 _ - , carbonate sediments, 1 17, 119 _ - , calcium/magnesium ratio in detrital carbonates, 409 _ _ , eolianites, 233, 407 _ _ , loss of magnesium from reef sands, 179 _ _ , mineralogy of ooids, 182, 183 _ _ , Recent dolomite, 284 _ - , - oolitic sands, 180 Bahamite, 35, 151, 159 -, definition of, 151 Baja California, Mexico, 40, 41, 119 _ - , carbonate sediments, 119 Bajadas, 152 Balanus amphitrite saltonensis, 203 Bald Mountain, N. Y . , U.S.A., 247

SUBJECT INDEX

Banff Formation, Canada, 222, 231 , evaporites, 222 Bank, 2, 24, 29, 34, 42, 98, 99, 151, 214, 218 -, Bahama, 29, 34,42, 151, 218 -, definition of, 98, 151 -, first-cycle carbonate minerals, 214 Barite cement, 189 Barnacle bioherm, 226 Barnacles, 203, 409 -, environmental temperature, 409 - in Salton Sea, 203 Barred basin types, I 1 , 194 Barrettia Limestone, Jamaica, 230 Barrier, 19, 24, 29, 192, 217, 274, 383 - bank, environmental classification of limestones, 383 - beach, 24, 361 - reef, 19, 29, 192, 217, 274 _ _ , Australia, 29, 274 _ _ _ , Capitan Limestone, 192 - -, Great Baharna Bank, 217 Basin, 9, 1 I, 12, 16, 18, 20, 22, 24, 30, 64, 81, 116, 152, 194, 199, 204, 220, 241, 312, 381, 383 -, barred, 11, 194 -, Dead Sea, 204 -, Delaware, 199, 241, 3 I 1 -, East Indian, 30 -, environmental classification of limestones, 383 -, Great, 116, 152 -, Michigan, 312 -, Midland, 220 Bathyal, 4, 14 Beach rock, 13, 14, 151, 189, 190, 405 - _ , aragonite cement, 189 - -, definition of, 151, 405 _ _ , gypsum cement, 190 - -, halite cement, 189 “Beagle”, H. M. C . , voyage of, 171 Beaverhill Lake Formation, Alta., Canada, 389, 392 - _ - , cross-section of, 392 Beaver Lodge Oil Field, N. D., U.S.A., 299 Becraft Formation, N.Y., U.S.A., 225, 331 Beekniantown Group, Pa., U.S.A., 11 6 Beit She’an, Israel, 204, 319 Beldens Formation, Vt., U.S.A., 321 Belt Series, Mont., U.S.A., 417 Bermuda, 40-42, 178-180, 231, 237, 278, 331, 407 - Apron, 237, 278 -, changes in grain mineralogy, 180 -, eolianites, 231, 407 -, mineralogy of carbonate rocks, 178, 179 -, Recent skeletal sands, 180

_-

SUBJECT INDEX

Berry Islands, Great Bahama Bank, 182 Bicarbonate-carbonate relations in sea water, 414 Bikini Atoll, 186, 187, 219 - _ , carbonate sediments, 219 _ - , dissolution, 186, 187 Bioaccumulated, 93, 151 -, definition of, 151 Bioarenite, 93 Biocalcarenite, 94, 131, 151 -, definition of, 151 -, photomicrographs of, 131 Biocalcilytes, definition of, 173 Biocalcirudite, 93-95, 151 definition of, 151 Bioclastic, 88, 151, 385 - carbonates, 88 -, definition of, 151 - limestone lenses, 385 Bioconstracted, definition of, 151 Biofacies, 15, 19 Biogenic, 3, 88, 138, 151 - calcite formation, 3 - carbonates, 88 -, definition of, 151 - skeletal limestone, photomicrographs of, I38 Biogeochemical, 24, 25, 399 - events, 399 - revolutions, 24, 25 Bioherm, 99, 151, 215, 220, 223, 226, 227 -, algal, 215, 227 -, barnacle, 226 -, bryozoan,33 I Circulation index, 410 Cirripedia, 35 Clarendon Bank, Jamaica, 216, 236 - Block, Jamaica, 216 Classification, 87-168, 172, 173, 308, 374, 380, 383

-, Edie’s environmental classification of Mis-

sissippian limestones, 383 -, Grabau’s, 172, 173 -, grain size, 100-1 03 - of carbonate rocks, 87-168 - - depositional environments, sediment types and organism communities, 380 _ _ dolomites, 106-1 13, 308 - _ limestones, 101, 113-115, 194 -, Shvetsov’s, 11 8-1 19 -, Teodorovich’s, 97, 107, 110, 1 1 I , 374 Clast, definition of, 153 Clastic, definition of, 153 Clasticity, 72 Clay, 1, 14, 17, 52, 90, 109, 202, 294 - catalysts in dolomite formation, 294 - in algal material, 202 Clayey, 89, 107 - dolomite, 107 - limestone, 89 Climate, 400, 402 Climatic indicators, 399, 408 Cloud Chief Formation, Okla., U.S.A., 31 I , 322

SUBJECT INDEX

Coal Measures, U.S.A., 204, 209, 210 Coastal Limestone, Western Australia, 33, 21 6 Coatedgrains, 17,93,96, 105, 136, 137, 153,215 ___ , definition of, 96, 153 _ - in bank sediments, 215 _ _, photomicrographs of, 136, 137 Coboconk Formation, Ont., Canada, 225 Cobourg Formation, Ont., Canada, 225 Coccolithophoridae (coccoliths), 14, 19, 30, 153, 183, 235, 239, 420 - in oozes, 235 - in deep-sea sediments, 183 Codiaceae, 77, 78 Coelenterata, 35, 39 Coeymans Formation, N.Y., U.S.A., 225 Collenia, 4, 167, 416 Collings Ranch Conglomerate, Okla., U.S.A., 213 Collingworth Gypsum, Okla., U.S.A., 311, 329 Color, 1, 62, 65 Colorado Plateau, U.S.A., 1 16 - River, U.S.A., 203 Compaction of sediments, 44, 298, 364 Compensation depth, calcium carbonate, 30 Composite grains, definition of, 153 Compound-pellet, definition of, 153 Conglomerates (see breccia), 94, 21 0, 21 2, 246, 247, 315, 316 -, carbonate rock, 21 2 -, limestone, 246, 247 -, limestone-pebble, 210 Congo River, Africa, 405 Connate water, 15, 56 Connecticut Valley, U.S.A., 202, 208, 209, 212 _ - , carbonate conglomerates, 212 _ - , Triassic deposits, 203, 206 Conococheague Limestone, Appalachians, U.S.A., 321 Cool Creek Limestone, Okla., U.S.A., 306, 330, 331, 333, 334 Cooking Lake Formation, Alta., Canada, 328, 394 _ _ _ , dolomitized reef, 328 Coorong Lagoon, Australia, 281, 283 Copper Ridge Dolomite, Appalachians, U.S.A., 320, 321 Coquina, 5, 95, 153, 209, 311 -, definition of, 153 -, dolomitized, 3 11 -, Mabou Group, 209 Coquinite, 12, 154 -, definition of, 154 Coquinoid, definition of, 154 Coralgal, 12, 1 17, 154 -, definition of, 154

45 1 Corals, 13, 17, 22, 35, 64, 65, 67, 93, 98, 191, 217, 219, 226, 296, 300, 314, 315, 334, 396, 409 -, barrier reef, 217 -, bioherm, 226 -, environmental temperature, 409 - in beachrock, 406 _ _ _ eolianites, 407 - - reef detrital material, 219 - _ _ replaced limestone, 334 - madreporarian, 35 -, rugose, sediment contributing organism, 396 -, susceptible to recrystallization, 191 - tabulate, sediment contributing organism, 396 Coral Sea, 19 Corbula, 226 Corey Limestone, Appalachians, U.S.A., 321 Couvinian Limestone, Germany, 6 Coward Springs, Australia, 293 Cow Head Limestone Breccia, Nfld., Canada, 246, 247 Crassinella, 228 Crassostrea, 228 Creta Dolomite, Okla., U.S.A., dolomitized sediments, 329 Cretaceous, 3, 19, 26, 53, 82, 175, 189, 196, 211, 227, 234-236, 241, 242 -, Carpathian Mountains, Poland, 234, 235 - chalk, 175, 189, 234-236 - Exogyra mounds, 227 -, Plattin Flysch, 242 -, Westphalia, Germany, 241, 242 Crinoidea (crinoids), 35, 64, 78, 94, 122, 126, 226, 300, 396 -, colonies of, 226 - ossicles, 122 -, photomicrographs of, 126 -, sediment-contributing organism, 396 Criquina, definition of, 154 Cristobalite, 282 Crocodiles, fossils in Green River Formation, 206 Crust, definition of, 154 Crustacean, 26, 206 Cryptoclastic, definition of, 154 Cryptocrystalline, definition of, 154 Cryptograined, definition of, 154 Cryptozoon, 41 6 Crystal, definition of, 276 Crystalline criquinite, plates of, 140 -, definition of, 154 Crystallinity, 1 1 1 Crystallization of calcium carbonate, 174 Crystalloblastic, definition of, 302

452 Cuba, reefs, 217 Cumberland County, Pa., U.S.A., 321 - Gap, Va.-Ky., U.S.A., 273 - Plateau, Tenn., U.S.A., 222, 225 Currie Hills, Nev., U.S.A., 123 Cyanophyte Algae, 25, 150,416 Cycles, normal marine, 387 Cyclothem, 12, 229-23 I Czorsztyn series, pelagic and neritic pure limestones, Poland, 235 Darss/Drogden Sill, North Sea, 11 Davis Sill, Labrador Sea, 11 Dead Horse Wash, Nev., U.S.A., 126 - oil, 123, 127 - Sea, 192, 199-201, 204, 291, 292, 296, 309, 336, 337 _ _ , aragonite precipitation, 336 Dead Sea Basin, 204 --, calcium carbonate lake deposits, 19920 I - _ , dolomite in sediments, 291, 292 _ _ gypsum, precipitation of, 336, 337 - -, isotopic studies, 291 _ _, origin of aragonite, 296 - _ , salinity of, 199 Deccan traps, India, 199 Dedolomitization, 21, 272, 300-303 -, definition of, 272 Deepkill Formation, N.Y., U.S.A., 240 Deep-sea sediments, mineralogy of, 182, 183 Deep Spring Lake, Calif., U.S.A., 283 Deer Mountain Oil Field, Alta., Canada, 389 Delaware Basin, Texas, U.S.A., 192, 199, 222, 243, 244, 31 1 _ - , evaporites, 222 - _ , facies belts, 31 1 - _ , Permian rocks, 243, 244 Dense, definition of, 154 Depocenter, 104, 155 -, definition of, 155 -, energy level, 104 Depositional processes, 196-198 Deposits, 199, 213-223,234-248,271,324-327 -, bank, 215-218 -, biogenic, 271, 324, 325 -, calcareous dunes, 199 -, combinations of marginal marine types, 2 14-223 -, fan or stream, 210-213 -, first-cycle carbonate minerals, 214-223 -, lake, calcium carbonate, 199-210 -, marginal marine, 213-215, 231-233 -, mechanical, 271, 326, 327 -, mixed carbonate and non-carbonate, 240248

SUBJECT INDEX

-, -, -, -,

offshore marine, 215-220 pelagic carbonates, 234-240 reef, 218-220 shallow-water, 213, 214 Depth change, carbonate facies as indicators of, 424 Derryan age, 121 Desmoinesian age, 121 Detrital, 17, 93, 105, 106, 128, 129, 132, 155, 219, 268, 326, 327, 409 -, carbonate sediments, 409 -, definition of, 155 - dolomite, 268, 326, 327 - grains, 93, 105, 106 - limestones, 17, 128, 129, 132 - _ , photomicrographs of, 128, 129, 132 - reef deposits, 218-220 Devonian, 6 , 58, 64, 126, 174, 186, 211, 212, 216, 220, 224, 227, 297, 298, 310, 329, 389 -394 -, Alberta, Canada, 220 -, Andrews South Oil Field, 298 -, Belgium, 220, 227 -, crinoidal limestone, 186 -, Eifel, Germany, 329 -, evaporitic dolomite, 310 -, Formosa reef limestone, 297 -, Palliser Limestone, 216 - reservoirs, 389-394 Diaclastic revival, 27 Diagenesis, 55, 109, 155, 172 -, calcareous material, 109 -, definition of, 155 - in carbonate rocks, 55 -, origin of term, 172 Diagenetic, 3, 5, 18, 20, 21, 93, 108, 111, 143, I55 - crystallization, 93 -dolomites, 3, 20, 21, 108, I l l , 143, 155, 268,294-300, 327-330 _ - , definition of, 155 _ - , photomicrographs of, 143 - -epigenetic, 18 - stages, 5 Diagenetically-altered, I 3 I, 135, 138, 141, 142 - biocalcarenite, photomicrographs of, 131 - calcilutite, photomicrographs of, 142 - dolomitized skeletal limestones, photomicrographs of, 141 - skeletal calcarenite, photomicrographs of, 143 - skeletal limestone, photomicrographs of, 135, 138, 141 Diamictites, 245-248 Diamond Range, Nev., U.S.A., 121, 122, 125

45 3

SUBJECT INDEX

Dimensional fabrics, 54, 55 Dinocardium, 228 Dinosaur, fossil footprints, 318, 319 Diploria, 217 Dismicrite, 94 Dissolution, definition of, 186 Distribution of carbonates, 4, 31, 404 Dogge-Malni, Sicily, Italy, 78 Dolarenite, 12, 112, 155 -, definition of, 155 Dolly Varden Mountains, Nev., U.S.A., 123 Dolocast, definition of, 153, 155 Dololutite, definition of, 155 Dolomicrite, 12, 130, 155 -, definition of, 155 -, photomicrographs of, 130 Dolomite (see dolostone), 1, 3, 5, 7-9, 20, 21, 88, 107, 108, 111, 114, 155, 190, 193, 202, 212, 267-348, 363, 375, 376 (and elsewhere) -, calcian, 282 -, calcitic, 88, 107, 108 -, calcium-rich, 296 -, caliche-like, 288 -, calcitization of (see dedolomitization), 21, 300 - cement, 190 - clasts, 212 - crystals, 278, 363 _ _ , precipitation of, 278 - -, secondary origin of, 278 _ _, size of, 363 -, definition of, 155 -, deposit types, 7 -, detrital, definition of, 327 - distribution of, 7 - facies, 5 - formation, 274, 336, 337 _ - , from heavy brines, 337 - _ ,requirements, 336 -, hypidiotopic, 302-306 -, identification of crystals, 324 -, idiotopic, 299, 302-306 - in algal material, 202 - - Bonneville Lake sediments, 292 -- Florida Bay, 287 _ - hardpans, 209, 290 _ - lacustrine sediments, 281-284, 289-293 - - Lake Eyre, 293 - - playa sediments, 283 - _ Pleistocene graben deposits, 290 - - sebkhas, 285-288 -, insoluble residue, relationship to, 8, 9 -, intertidal, 287 -, isotope studies, 277, 291-293, 330, 331 -, laboratory formation of, 303, 307, 308 -, magnesian, 108

-, mercury capillary-pressure curves of, 375, 376

-, mineral types, 20 -, miscellaneous nonmarine, 325, 326 - particles,

277-279

-, penecontemporaneous, 280, 299, 300 -, Pleistocene, origin of, 289 -, porosity-permeability relationships, 365 -, primary, 108, 109, 114, 275, 280

-, -, origin of, 275 -, -, replacement type, 280

- problem, 269, 270 -, protodolomite, 275, 276, 286, 307 -, radiocarbon age, 284, 285, 287, 289, 290, 292 -, Recent, 281, 283, 284 -, -, in lakes, 280-284, 289-293 _ , _ , - playas, 283 - replacement, 193,289,294-300 _ _ of calcium carbonate, 289, 294-300 --- quartz, 193 - _ , tectonic, 300 -, review of literature, 270-276 -, secondary, 275, 298, 300 -, stoichiometric, 293 -, stylolites, 297, 298, 333 -, supratidal, 287, 315 -, xenotopic, 299, 302-306 Dolomitic, 88, 89, 107, 108, 111, 112, 156 - chalk, I1 1 -, definition of, 155 - limestone, 88, 89, 107, 108, 112 - magnesite, 108 - mottling, definition of, 156 Dolomitization, 2, 5, 6, 8, 9, 16, 17, 21, 22, 77, 78, 93, 109, 294-300, 303, 330, 334-338, 361, 363-365, 389 -, Ismay zone, 389 - model, 334-338 -, original sediment composition, effect of on, 363-365 -, post-diagenetic, 109 - process, 361 -, syndiagenetic, 109 Dolomitized, definition of, 156 -, photomicrographs of, 140, 142 - biomicrite, 140 - calcilutite, 140 - micrite, 140 - skeletal-detrital limestone, 142 - skeletal limestone, 140 Dolomolds, definition of, 156 Dolorudite, definition of, 156 Dolosiltite, photomicrographs of, 130 Dolospar, eyes, 127 Dolosparite, 140, 141, 143, 156

454 Dolosparite, definition of, 156

-, photoniicrographs of, 140, 141, 143

Dolostone (see dolomite), 106-1 13, 267-348, 373 -, association with other rock types, 322323 -, classification of, 106-113, 308 -, definition of, 107 -, detrital, 326, 268 -, diagenetic, 3, 20, 21, 108, 1 1 I , 268, 294300, 313, 327-330 -, -, definition of, 327 -, -, origin of, 328, 329 -, economic importance, 269 -, epigenetic, 268, 269, 330-334 -, -, related to ore deposits, 332 -, evaporitic, 309-3 13 - fabrics, 302-306 -, fault-related, 333 Dolostone interbedded with terrigenous sediments, 321-323 -, mottled, 297 -, origin of, 267-269, 277, 319 - - _ , on tidal flats, 319 -, S-, 273, 328, 332, 333 -, stratigraphic distribution of, 269 -, subdivisions of, 373 -, syngenetic, 20, 268, 308-325 -, -, associated with evaporites, 309-313 -, -, formation of, 314 -, T-, 273, 330 - textures, 302-306 -, unconformity and discontinuity types, 329 -, unifying model of origin, 334 -, W-, 273, 328 Donax, 228 Dripstone, origin of, 177 Druse, definition of, I56 Drusy coating, definition of, 156 Dry Ridge, Idaho, U.S.A., 123 Duchesne River Valley, Utah, U.S.A., 121, 126, 127 Dune rock, calcareous, 216 Dunhsm Dolomite, Vt., U.S.A., 322 Dunkard Group, Pa., U.S.A., 210 Duvernay Shale, Canada, compaction of, 298 Earthy, definition of, 156 East China Sea, carbonate mineralogy of sediment, 183 East Indian basins, 30 Ecculiornphalous, 317 Echinoderms, 191, 219, 276, 409 - in reef detrital area, 219 -, susceptibility to recrystallization, 191 Economic aspects of carbonate rocks, 18

SUBJECT INDEX

Ediacara fauna, Australia, 41 7 Edie’s environmental classification of Mississippian limestones, 380, 383 Egan Range, Nev., U.S.A., 126 Eh, 8 Eifel, Germany, 6, 329 Elaphrosauraus, 3I9 Elbrook Limestone, Pa., U.S.A., 321, 328 - -, dolomitized sediments, 328 Electric logs, 73 Elephas, 401 Ellenburger Group, Texas, U.S.A., 216, 272, 279, 298 - _ , dolomite problem, 272 _ _ , marine bank deposits, 216 - _ , penecontemporaneous dolomite, 298 Elphidium, 228 Elugelab [sland, Eniwetok Atoll, Pacific Ocean, 195 Elvins Formation, Mo., U.S.A., 328 Ely Limestone, Nev., U.S.A., 121, 123-125 Emanuel Dolomite, Okla., U.S.A., 322 Emery Brook Formation, N.S., Canada, 209 Encrinal, 93, 123, 124, 135, 156 -, definition of, 156 - limestone, 93, 135 - -, photomicrographs of, 135 - material, 123, 124 Endogenic, definition of, I56 Endogenic rocks, definition of, 98 Energy index, 18, 101, 113-116, 156 - _ , definition of, 156 - level, definition of, 156 England, 212, 216, 231, 323, 333 -, cyclic carboniferous deposits, 23 1 -, epigenetic dolostones, 333 -, Great oolite series, 216 -, mark, 323 -, Mountain Limestone, 216 -, New Red Sandstone, 322 -, Plymouth Limestone, 212 EniwetokAtoll,PacificOcean, 186,187,195,296 --, dissolution, 186, 187 - -, Elugelab Island, 195 - _ ,oxygen isotope measurements, 296 Environment, supratidal carbonate, 40 Environmental classification of limestones, Edie’s, 380, 383 Eocene, 65, 196, 215, 216, 239, 241 - calcare alberese, 239 -, Jamaica, 216 - shales, Venezuela, 241 -, White Limestone, 215 Eolian carbonate deposits, 18, 23 1-233 Eolianite, 156, 180, 181, 188, 199, 231-233, 333,407

SUBJECT INDEX

-, Africa, 407 -, algal, 231, 407 -, Australia, 407 -, Bahamas, 233, 407 -, Bermuda, 180, 181, 188, 231,407 -, corals in, 407 -, definition of, 156, 231 -, Foraminifera, 231, 407 -, leaching of, 407 -, Mauritius Island, 233 -, Pleistocene, 180, 181, 199, 231, 232, 333 Epeirocratic, definition of, 425 Epidiagenetic, 5, 6, 157 -, definition of, 157 - stage, 5, 6 Epigenesis, 15, 21, 157 -, definition of, 157 Epigenetic dolomite, 269, 330-334 Equant, definition of, 157 Equigranular, definition of, 157, 302 Equilibrium concentrations of bicarbonate and carbon dioxide, 413 Espiritu Santo Island, Baja California, Mexico,

119 Estheria, 408 Euhedral, definition of, 157, 302 Evacuation Creek Member, Green River Formation, 206 Evaporation, formation of hypersaline brines, 296 Evaporite, 7, 10, 12, 40, 113, 157, 189, 214, 221-223, 228, 285, 310-313, 323, 413, 41 9 - breccias, 157, 223, 228 -, Briggs model of, 312 -, Cambrian, 419 - deposits, 40, 221-223, 309-313 - dolomite, 309-3 13 - minerals, (see halite), 189, 214, 285 -, Paleozoic, 413 -, Permian, 222, 223 -, Qatar Peninsula, 221 -, Recent, 221 - solution breccia, definition of, 157 -, - Of, 223 - sequence, 113, 223 - suite, 7, 10, 12 - succession, Zechstein, Germany, 310 - zonation, 3 12 Everglade swamps, Fla., U.S.A., 42 Evolution of carbonate rocks, 2, 3 Exogenic, definition of, 157 Exogyra mounds, 227 Exotic pebbles and boulders, 245 External growth, 271 Extraclast, definition of, 157

455 Exuma Sound, Bahamas, 35, 38 Fabric, 157, 299, 302-306 -, crystal, definition of, 302 -, definition of, 157 -, hypidiotopic, 302-306 -, idiotopic, 299, 302-306 -, preservation by selective dolomitization, 302 -, xenotopic, 299, 302-306 Fan deposits, limestone particles, 210-21 3 Fanglomerates, Israel, 2 11 Favosites, 352, 353 Feldspar, 90, 184, 242, 412 -, calcium carbonate coatings, 184 -, Precambrian sediments, 41 2 Ferguson Flat, Nev., U.S.A., 127 - Mountain Formation, Nev., U.S.A., 122-

124 Ferron Mountain, Utah, U.S.A., 207 Fibrous, definition of, 157 Fish, fossils in Green River Formation, 206 Five Mile Pass area, Utah, U.S.A., 123 Fjords, 11 Flagstaff Limestone, Utah, U.S.A., ancient lake sediments, 203, 207, 208 Florena Shale, Kans., U.S.A., dolomite replacement, 298 Florida, 29, 37, 40, 42, 119, 218, 284, 287 - Bay, 29, 37, 40, 42 -, carbona:e sediments, I19 -, dolomite, 284, 287 - Everglades, 42 - reef tract, 218 Flowstone, origin of, 177 Fluid migration, due to compaction, 298 Fluid movement in carbonate rocks, 374 Fluorite, 1, 313 Fly River, New Guinea, 1, 405 Flysch-Aalenian, Poland, 242 Foraminifera (see Globigerina), 13, 14, 30, 32. 34, 44, 94, 96, 126, 175, 176, 191, 203, 215, 218, 219, 226, 228, 230, 231, 235, 236, 239: 272, 334, 399, 401, 406, 409-411, 417, 420 -, age of, 417 -, benthonic, 219 -, calcium withdrawal, 410 - encrusted skeletal grains, definition of, 96 -, environmental temperature indicators, 409 -, hyaline, 231 - in beachrock, 406 - _ chenier, 228 - _ eolianites, 231, 407 - _ Florida Bay sediments, 2 I8 - _ oozes, 235 - _ replaced limestones, 334 -, pelagic, 41 1

456 Foraminifera (continued) -, photomicrographs of, 126 -, precipitation of limestone, 175 -, replacement of, 272 -, Revolution V, 399 -, Salton Sea sediments, 203 -, susceptibility to recrystallization, 191 Fore reef, definition of, 157 Formosa reef limestone, Ont., Canada, 297 Formation waters, origin of diagenetic and epigenetic dolomite, 335 Fort Johnson Member, Tribes Hill Formation, N.Y., U.S.A., 301 Fossil footprints, dinosaur, 318, 319 Fragmental, definition of, 157 - limestones, 95 Framework, definition of, 157 Freeport division of Allegheny Series, Pa., U.S.A., 209 Fremont Island, Utah, U.S.A., 201 Frenchman Mountain, Nev., U.S.A., 126 Fringing reef, 21 8 Frolova’s classification of doloniite-magnesite-calcite series, 108 Funafuti Atoll, 196, 273, 296, 300 _ - , boring, 196, 273 - -, dolomite, 300 Fusulinal, definition of, 157 Fusulinids, 121, 122 Galilee, Israel, 175 Garden Valley Formation, Nev., U.S.A., 121, 122, 125 Garnet Railroad Siding, Nev., U.S.A., 126 Gas City Field, Mont., U.S.A., 362 Gaspe Peninsula, Canada, 190, 21 3, 324 Gastropods, 35, 210, 228, 300, 314, 317, 391, 396 -, fresh-water, 210 - in a chenier, 228 -, sediment-contributing organism, 396 Gaylor Dolomite, N.Y., U.S.A., 316 Gaylussite, 283, 284 Gela Oil Field, Sicily, Italy, 72, 73, 75-79, 81 Genetic types of dolomites, 108, 109, 268, 269 Geochemical changes, end of Paleozoic, 419 - parameters, variation of, in the past, 410415 Geochemistry of Persian Gulf water and sediments, 286 Geomagnetic reversals, 422 Germany, 6, 206, 228, 241, 242, 310, 329, 333 -, Aller beds, 310 -, Bunter Sandstone, 333 -, Couvinian Limestone, 6 -, Eifel, 329

SUBJECT INDEX

-, Newberrian Sandstone, 228 -, Westphalia, 241, 242 Gerster Formation, Nev., U.S.A., 123, 125 Ghost relics, 18 Giardini Formation, Sicily, Italy, 75, 78 Gibraltar Sill, Atlantic Ocean, 11 Gilsonite, 206 Gilson Mountain, Utah, U.S.A., 124 Givetian Dolomite, Germany, 6 Glacio-eustasy, definition of, 423 Glass Mountains, Texas, U.S.A., 125 Glauberite, in sediments, 283 Glauconite, 1, 90, 242 Glendale Junction, Nev., U.S.A., 127 -thrust sheet, 211 Globigerina (see Foraminifera), 2, 19, 30, 31, 183, 219, 235, 236, 237, 239 - in deep sea sediments, 183 Globochaetes, 235 Globotruncana, 236 Glomeroclastic, definition of, 158 Gloucestershire, England, marls, 323 GMR (grain/micrite ratio), 17, 105, 106, 113, 116 Goat Seep reef, Texas, U.S.A., 243, 328 Gold Hill District, Utah, U.S.A., 124, 126 Golden Spike reef, Canada, 298 Golzinne Formation, Belgium, 63, 64 Goniolithon, 108 Goodsprings quadrangle, Nev., U.S.A., 332 Gorge Formation, Vt., U.S.A., 322 Grabau’s limestone classification, 172, 173 Grain, 93, 100-103, 105, 106, 113, 158, 180, 184, 276, 279 -, concentric coatings of, 184 -, definition of, 158, 276 -, detrital, 93, 105, 106 -/matrix ratio, 113 --/micrite ratio, 105, 106, 113 - mineralogy, changes in, 178-184 -, recycled dolomite, 279 - size classification, 100 -- scales, 102, 103 - terrigenous carbonate, 180 - types, 93 Grainstone, 94, 116, 158 -, definition of, 158 Grain-supported, definition of, 158 Grandeur Member, Park City Formation, Utah, U.S.A., 127 Granite Mountain, Utah, U.S.A., 127 Granoblastic, 111 Gran Sasso massif, Italy, 239 Granular, definition of, 158 Granulometry, 64, 67, 68, 70 Grape Bay, Bermuda, 188

451

SUBJECT INDEX

Grapestone, 19, 34-36, 44, 158 Gravity-displaced sediments, 240 Grayburg Limestone, Texas, U.S.A., anhydrite cement, 190 Graywacke, 244, 412 --, corroded, 412 -, Gasp6 Peninsula, Canada, 244 Great Bahama Bank, Atlantic Ocean, 44, 182, 183, 215, 238 - - _ , index map, 238 _ _ _, shallow water carbonate deposition, 215 - Barrier Reef, Australia, 29, 274 - Basin, U.S.A., 116, 152 _ _ , caliche deposits, 152 - Salt Desert, Utah, U.S.A., 201 _ _ Lake, Utah, U.S.A., 7, 95, 156, 185, 199, 201, 202,216,405 - _ _ ,carbonate sediments, 199, 201, 202 _ _ - , oolites, 201, 202, 405 _ _ _ , pseudooids, 185, 202 _ _ - , salinity of, 201 Greenland, 401,403 Green River Formation, Wyo., U.S.A., 203, 205, 206, 209 Grumous, definition of, 158 Guadalupe Mountains, Texas-N.M., U.S.A., 125, 243, 298, 325 Guadalupian age, 123, 125 Guam, Pacific Ocean, 190, 191, 219 -, carbonate sediments, 219 Guilmette Limestone, Nev., U.S.A., 126 Gulf of Batabano, 29, 34 _ _ California, Recent marine evaporites, 221 _ _ Eilat (Gulf of Aqaba), Red Sea, 241, 242,279, 327 _ _ - , sediments, 242, 327 _ - Maine, barnacle bioherm, 226 _ _ Mexico, 224, 226 ---, coral bioherm, 226 Gull River Formation, Ont., Canada, 225 Gunnison Plateau, U.S.A., 207, 208 Gyparenites, 156 Gypsina, 176 Gypsite, 293 Gypsum, 1, 7, 9, 10, 12, 20, 22, 109, 156, 189, 192, 200, 204, 221-223, 258, 285, 286, 288, 289-293, 301, 309, 311, 313, 323, 336 362, - associated with syngenetic dolomite, 309 -, bacterial breakdown, 192 - cement, 189 -, formation of, 302 - in Blaine Formation, 31 I _ _ Dead Sea, 200 _ _ Lake Eyre sediments, 293

_ _ Pleistocene graben deposits, 292 - _ pluvial lake sediments, 293 - _ Samra Formation, 204 _ _ sebkhas, 285 - _ sediments, 286

290, 291,

-, precipitation of, 289, 336 -, role in dedolomitization, 301

Hailey, Idaho, U.S.A., 122 Hatimedu, 180, 181, 191, 218, 239, 314, 410 -, Florida Bay, back reef area, 218 -, photomicrograph of, 181 -, susceptibility to recrystallization, 191 Halite (see evaporites), 189, 206, 221, 222, 279, 280, 283, 285, 287, 288, 290, 292, 309, 310, 312, 323, 419 - as cement, 189, 279, 280 - associated with syngenetic dolomite, 309 - first evaporite deposit, 419 - “halo”, 312 - in beach rocks, 189 _ _ Bonneville Lake sediments, 292 _ - Green River Formation, 206 - _ Pleistocene graben deposits, 290 _ _ sebkhas, 285 _ _ sediments, 283 -, precipitation of, 221 -, Zechstein sequence, 310 Hall Canyon Member, Oquirrh Formation, Utah, U.S.A., 121, 123 Hamilton Group, Pa., U.S.A., 422 Haney Limestone, Ill., U.S.A., 230 Hannah Formation, Mont., U.S.A., 125 Hard chalks, definition of, 189 Hardpans, 290, 407 -, dolomite-bearing, 290 Harrisburg, Pa., U.S.A., 8, 9 Hathaway Formation, Vt., U.S.A., 246 Hatteras abyssal plain, 237, 239, 241 Hauptanhydrit, Germany, 310 Hauptdolomit, Germany, 310 Haupt Muschelkalk, Germany, 3 10 Hausmannite, 332 Haute Savoie, France, 236, 239, 243, 244 Haystack Gypsum, Okla., U.S. A., 31 1 Helderbergian Group, N.Y., U.S.A., 225, 330, 331 Helvetic nappes, France, 236 Hematite, 333 Hemipelagic sediments, 240 Hercynian orogeny, 422 Hermatypic corals, Great Bahama Bank, 218 Heron Island reef, 34 Heteropods, 30 Highly-altered diagenetic dolomite, photo-

458 micrographs of, 142 High-energy, definition of, 158 Highgate Spring-St. Dominique thrust, U.S.A., 321 Hippopotamus, 401 Hobble Creek Canyon, Utah, U.S.A., 123 Hog Island, Great Bahama Bank, 183 Holoblasts, 272 Holocene, 47, 400, 401, 410 -, oscillation of, 410 -, post-glacial thermal maximum, 401 Homotrema rubrum, 44, I76 Honaker Dolomite, Appalachians, U.S.A., 321 Honaker Trail, Utah, U.S.A., 117 Horseshoe Atoll, Texas, U.S.A., 382, 384 Horse Springs Formation, Nev., U.S.A., 203, 207, 310, 313 ---, ancient lake sediments, 203, 207 Hoyt Limestone, N.Y., U.S.A., 328 Hudson Valley, N.Y., U.S.A., 225, 247, 327, 330 Humic acid, mobilizing carbonates in soils, 406 Hungerford Slate, Vt., U.S.A., 227, 322 Hydrobia, 227 Hydrocalcilytes, definition of, 173 Hydroclastic carbonates, 88 Hydroclast, definition of, I58 Hydrolith, definition of, I58 Hydromagnesite, 90, 282 Hydrothermal, 6, 15, 56 - dolomite, 6 -water, 15, 56 Hydrotroilite, 239 Hydrozoa, 41 7 Hypersaline brines, 267, 296, 331, 334 _ _, dolostone origin, 267 _ - , formation of, 296 _ _ , origin of all dolomites, 331, 334 Hypersalinity, 334, 335, 337 -, capillary concentration, 335, 337 -, causes of, 334 Hypidiotopic, definition of, 302 Ice Age, 400 Idiotopic, definition of, 302 Illite, adsorption of magnesium, 408 Impingement, definition of, I58 Impurities in carbonate rocks, 1 Incipient dolomitization of a calcarenite, photomicrograph of, 141 Incrusting organisms, 176 India, 190, 199, 407, 419 -, Cambrian evaporites, 419 -, Deccan Traps, 199 -, eolianites, 407 -, Junagarh Limestone, 190

SUBJECT INDEX

Indian Ocean, 11, 19, 183, 233, 404 _ _ , carbonate mineralogy, 183 _ _, distribution of marine carbonates, 404

_ _ , Maldive and Laccadive Islands, _ _ , Mauritius Island, 233

19

Indicators of climatic zones, 402 Inequigranular, definition of, 302 Infilling of sand-size grains by aragonite. 45 Inoceranius, 235, 236 In-place organic structures, 93 Insects in Green River Formation, 206 Insoluble residue i n limestones and dolomites, 7-9 Interface, definition of, 158 Internal growth, 271 - sedimentation, definition of, 159 Interrelationship between rock components, 61, 62 Interstitial, definition of, 159 Intertidal magnesium enrichment, 273 Intraclast, 93, I59 -, definition of, 159 Intramicrite, 75 Tntramicrudite, 99 Entrasparite, 75 Intrasparrudite, 75, 93, 99 Inverness Oil Field, Aka., Canada, 389 Inversion, definition of, I59 Invertebrates, oxygen breathing, evolution of, 417 lnyo County, Calif., U.S.A., 283 Ion migration in invertebrate skeletons, 39 Iran, Cambrian evaporites, 419 Ireton Shale, compaction of, 298 Iron, 1, 60, 162, 412 - in dolomite, 162 -, FeOiFeaOs ratio, 412 -sulfate, 1 Trreducible water saturation, 375, 376 Ismay Oil Field, Utah-Colo., U.S.A., 354, 355, 359, 384-389 - _ _ , dolomitization, 389 _ _ _ , facies map of,385 Isotopes, 18-20, 4 0 4 3 , 200, 291, 292, 296, 297, 330, 331, 401, 403, 416 -,carbon, 20, 42, 43, 291, 292, 331, 403, 416 -, composition of dolomite, 291 -, dolomitization of Cool Creek Limestone, 330, 331 -, oxygen, 18, 19, 4042, 200, 291, 292, 296, 297, 331, 401 -, strontium, 200 Israel, 175, 200, 204, 211, 233, 302, 317, 319 -, carbonate rocks, 3 I7 -, Beit She’an, 204, 319 -, Beit Zait, 319

459

SUBJECT INDEX

Israel (con timed) -, Dead Sea, 201,292, 331 -, fanglomerates, 21 1 -, Galilee, 175 -, Judean Mountains, 175, 319 -, kurkar, 233 -, Lisan Formation, 204 -, Makhtesh-Katan, 317 -, Mount Carmel, 175 -, nari crust, 175 -, Negev Desert, 302 -, Samra Formation, 204 -, Shomron Mountains, 175 -, Soreq Formation, 319 -, whitings, 200 Italy, 16, 72-83, 176, 177, 239 -, Gela Field, 75-78 -, Gran Sasso massif, 239 -, Ragusa Plateau, 72-75 Zvanovia, 386-388 Jamaica, 216, 217, 229, 230, 236 -, Bavrettia Limestone, 230 -, Clarendon Bank, 216 -, Cretaceous, 229, 230 -, Johns Hall Conglomerate, 230 -, Montpelier Limestone, 236 -, Newman Hall Shale, 216 -, Newport Limestone, 21 6 -, reefs, 217 -, Troy Limestone, 216 -, White Limestone, 216, 236 -, Williamsfield trough, 216 Japan, 401 Jaun Pass, Switzerland, 242 Jeffersonville, Limestone, Ind., U.S.A., 224 Jester Dolomite, Okla., U.S.A., 311, 329 Johns Hall Conglomerate, Jamaica, 230 - Valley Shale, U.S.A., wildflysch diamictite, 245, 247, 248 Jonesboro Limestone, U.S.A., 321 Jordan, Cambrian evaporites, 419 - Valley, 203, 209, 210, 293 Dead Sea Graben, 203, 204, 209, 210 _ _ , Pleistocene dolomite, 293 Joulters Cay, Great Bahama Bank, 182 Judean Mountains, Israel, 175, 319 Judy Creek Oil Field, Alta., Canada, 389 Junagarh Limestone, India, 199 Jura Mountains, Switzerland. 54, 203, 208 _ _ , ancient lake sediments, 203, 208 Jurassic, 16, 72, 75, 78, 80, 81, 117, 190, 208, 216, 220, 223, 237, 241, 298, 302, 317 -, Arab Formation, 223, 298 -, breccias, Haute Savoie, France, 241 -, carbonate rocks in Israel, 317 -, flysh deposits, 237

-, Great oolite series, England, 216 -, lacustrine limestones, 208 Kaibab Limestone, Utah, U.S.A., 12, 124, 126, 127, 211 Kalkberg Formation, N.Y., U.S.A., 225 Kansas, U.S.A., Cretaceous chalks, 236 Kapingamarangi Atoll, Caroline Islands, Pacific Ocean, 219 Karlsbad, Czechoslovakia, 177 Keystone thrust, Nev., U.S.A., 330 Key West, Fla., U.S.A., 40 Khvorova’s classification of sedimentary carbonate rocks, 101 Kingsport Limestone, Tenn., U.S.A., permeability of, 332, 333 Kingston, Ont., Canada, 316 Kirkman Limestone, Utah, U.S.A., 127 Kita-daito-jima Atoll, Paclfic Ocean, 196, 296 _ _ , drilling on, 196 Kittanning unit of Allegheny Series, Pa., U.S.A., 209 Knolls, Utah, 126, 292 Knox Dolomite, Appalachians, U.S.A., 273, 320, 321, 333 Koblenz, Switzerland, 325 Konia sp., 125, 127 Kupferschiefer beds, Germany, 310 Kuril Sill, Pacific Ocean, I 1 Kurkar, definition of, 233 Kyle Canyon, Nev., U.S.A., 122, 126 Labrador Sea, 11 Lacustrine, 154, 172, 176, 191, 199-210, 29!293,408 -, Algae, 202 -, ancient calcium carbonate sediments, 203210 -, calcium carbonate, Recent deposits, 199203 -, carboniferous deposits, 204, 209 -, climatic indicators, 408 -, Dead Sea, 199 -, dolomite, 292 -, Flagstaff Limestone, 203, 207, 208 -, gypsum, 292 -, halite, 292 -, Horse Springs Formation, 203, 207 -, Jura Mountains, 203, 208 -, Lake Bonneville, Utah, 154, 292 -, Lake Lahonton, 172, 176 -, limestones, 208, 209 -, rnarls, 206, 408 -, Miocene deposits, 206 -, Paleocene deposits, 203, 205, 207 -, Pennsylvanian cyclothem, 209

460 Lacustrine (continued)

-, pluvial lake deposit, 293

- sediments, 204, 209, 291, 293 -, Triassic deposit, 203 -, Uinta Formation, 203, 205 -, Wasatch Formation, 203, 205 Lagoon, 3, 7, 24, 214, 222, 229, 281, 283, 288, 381, 388, 390 -, Coorong, Australia, 281, 283 - deposits, 3, 7, 218, 219, 283, 288 - facies of Swan Hills carbonates, 390 -, first-cycle carbonate minerals, 214 -, environmental classification of, 383 -, Laguna Madre, Texas, U.S.A., 229 Lake Balkhash, U.S.S.R., 281 - Bonneville, Utah, U.S.A., 154, 292 - _ dolomite, 292 - _ Group, 154 _ - sediments, 292 - deposits (see lacustrine) - Eyre Basin, Australia, 293 - Lahonton, Nev., U.S.A., 176, 177 - Ontario, Canada-U.S.A., 316 Laminae, stromatolitic algal, 288 Lau, Fiji, Pacific Ocean, 191 Law of biological continuity, 415 Leaching, 388, 407 - in algal buildups, 388 - of eolianites, 407 Leduc Formation, Canada, 394 Lee Canyon, Nev., U.S.A., 121 Leine beds, Germany, 3 10 Leonard Formation, Texas, U.S.A., 125 Leonardian age, 122, 123, 125-127 Leppy Range, Nev.-Utah, U.S.A., 121, 125127 Levis Formation, Que., Canada, 246 Lexington Herrin (No.6)coal, U.S.A., 230,231 Liassic, 78 Limeclast, definition of, 153 Lime-mud, 184, 187, 364 -, compaction of, 364 -, grain coating, 184 -, recrystallization of, 187 Lime ooze, 96, 105 Limestone, 17-21,57,58,67,87-106,109,113117, 121-143, 159, 171-177, 181, 193, 194, 198, 199, 204, 209, 210, 220, 222, 239, 243, 367, 380, 383, 387, 390, 407 -, anemoclastic, 173, 199 - ,_ ,definition of, 173 - basin, definition of, 194 -, bioclastic, 57, 58, 67, 173 _,- , definition of, 173 -, bioconstructional, 220 -, biogenic, 177

SUBJECT INDEX

-, biohermal, 220

-, classification of, 87-106, 113-117, 194,383

-,

coastal, definition of, 407

-, crinoidal, 243

-,

cryptograined, 390 definition of, 159 detrital, 17, 128, 129, 132 endogenetic, 172, 173 environmentalclassification of, 114,380,383 exogenetic, 172, 173 fan, 210-213 geosynclinal, definition of, 194 hydroclastic, definition of, 173 hydrogenic, 172 incrustate, 175 lacustrine, Pennsylvanian cyclothems, 209 -literature on, 171-174 - lumps, 95 -, nonmarine, 198 -, oolitic, 204, 209, 222 -, origin of, 18-21 -, pebble conglomerates, 210 -, photomicrographs of, 121-143, 181 - platform, definition of, 194 -, porosity of, 367, 387 -, Precambrian, 193 -, pyritiferous, 239 - reef, 220 -, sideritic, 243 -, skeletal, 94, 132-134, 136, 176 -, -, definition of, 176 -, - -detrital, photomicrographs of, 129, 131-133, 142 -, stream deposits, 210 -, tectono-stratigraphic classification, 194 -, terrigenous, definition of, 173 -, thalassigenous, definition of, 173 Lipalian interval, 416 Lisan beds, Israel, 204 Literature on limestones, classic, 171-1 74 Lithification, 44, 47, 177, 178 Lithocalcarenite, definition of, 159 Lithocalcilutite, definition of, 159 Lithocalcirudite, definition of, 159 Lithocalcisiltite, definition of, 159 Lithoclastic, definition of, 160 Lithofacies, 15, 54 -, definition of, 54 Lithographic, definition of, 160 Lithophyllum, 176 Lithostratigraphy, 51 Lithothamnium, 176, 245 Lodgepole Limestone, Mont., U.S.A., 124 London-Paris Basin, 236 Lone Mountain Dolomite, Nev., U.S.A., 299, 328

-, -, -, -, -, -, -, -, -, -, -,

46 1

SUBJECT INDEX

Long Rock Island, Great Bahama Bank, 183 Longview Dolomite, Appalachians, U.S.A., 320, 321 Loray Formation, Nev., U.S.A., 126 Losungumsatz, definition of, 166 Louisiana, U.S.A., 219, 227, 228 -, cheniers, 227, 228 -, Tertiary reefs, 219 Lovell Wash, Nev., U.S.A., 122 Lowville Limestone, N.Y., U.S.A., 316, 317 Lumpal, 93, 105 Lumps, 17, 35, 93, 134, 137, 160 -, definition of, 160 -, limestone, definition of, 95 -, photomicrographs of, 134, 137 Lund, Nev., U.S.A., 125 Luster-mottling, definition of, 160 Lutite, definition of, 160 Mabou Group, N.S., Canada, ancient lake sediments, 204, 209 Madison Limestone, U.S.A., 299 Magnesian carbonates, 107, 108 - limestone, 88, 89 Magnesite, 1, 60, 90, 108, 313 Magnesium, 4, 5, 7, 17, 20, 22, 24, 39, 40, 179, 201, 268, 269, 273, 274, 282, 292, 336, 337 -/calcium ratio (see Ca/Mg ratio), 4,5, 17,20, 22,24, 173, 179,201,269,274,281,282,292, 336, 337 -/-, Bonair, Netherlands Antilles, 201 -/--, Dead Sea, 201, 292, 337 -/-, importance of, in dolomite formation, 274 -/-, intertidal enrichment, 273 -1of brines, 336 -1--, Persian Gulf, 201 -/-, Salt Flat graben, 292 - enrichment, 201 - in calcite (see calcite) - loss from reef sands, 179 Maine, U.S.A., 40 Makhtesh-Katan, Israel, 317 Mammoth Hot Springs, Yellowstone Park, U.S.A., 177 Manganese oxide, 1 Manguni Dolomite, Okla., U.S.A., 311, 322, 329 Manlius Formation, N.Y., U.S.A., 314, 315, 319, 320 Marine bank deposits, 216 Marl, 109, 160, 206, 323, 408 -, definition of, 160 -, England, 323 -, lacustrine, 206, 408

-, Triassic, 323

Marbre Noir, Belgium, 63, 64 Marlow Formation, Okla., U.S.A., 322 Marmaton Group, U.S.A., 230 Marshall Islands, Pacific Ocean, 219 Maryville Limestone, Appalachians, U.S.A., 320, 321 Mascot Dolomite, Appalachians, U.S.A., 320, 321, 332, 333 _ _ , permeability of, 332, 333 Mascot-Jefferson City district, Tenn., U.S.A., 332 Matrix, definition of, 160 Maturity, definition of, 160 Mauritius Island, Indian Ocean, 233 Maverick Spring Range, Nev., U.S.A., 125,126 Maynardville Limestone, Appalachians, U.S.A., 320, 321 Mechanical, definition of, 161 Mechanisms of deposition, types of, 271 Medicine Mountains, Nev., U.S.A., 123, 125 Mediterranean Sea, 11, 226, 296, 424 - -, crinoid colonies, 226 _ _ , Mallorca, 296 - -, transgressions during European Tertiary, 424 Megalump, definition of, 161 Mekong River, Asia, 405 Melanesian Sills, Pacific Ocean, 11 Mercury capillary-pressure curves, 375, 376 Mesocoquina, definition of, 153 Mesozoic, 10, 226, 350 - carbonate reservoir rocks, 350 Metabolic elasticity in organisms, 421, 422 Metasomatism, definition of, 161, 295 Meteoric water, 15, 108 Methy Formation, Canada, 31 1 Michigan Basin, U.S.A., 312 Micrite, 12, 20, 62, 75, 77, 78, 80, 81, 92-94, 105, 106, 109, 129, 133, 134, 161, 184, 317 -, definition of, 161 - envelopes, definition of, 184 -, limestone, micritic, 94, 129, 161 - _ , definition of, 161 - _ , photomicrographs of, 129 -, photomicrographs of, 129, 133, 134 Microcrystalline, definition of, 161 Microfacies, 15, 54, 55 -, definition of, 54, 55 Micrograined, definition of, 161 Micropelletoid, definition of, 161 Microradiography, 60, 61 Microsequence, 15, 55, 56 -, definition of, 55, 56 Midale beds, Sask., Canada, 365, 366

462 Midale Oil Field, Sask., Canada, 356 Middleville, N.Y ., U.S.A., 316 Miliolitic Formation, Saudi Arabia, 199 Mill Creek, Okla., 306 Mineralogical zonation, 282 Mineral thermometers, 24, 408410 Mingan Island, Gulf of St. Lawrence, Canada, dolomite sands, 279 Miocene, 196, 206, 215 -, ancient lake beds, 206 Mission Canyon Formation, N.D., U.S.A., 358, 361, 375-377 Mississippi Delta, U.S.A., 228 Mississippian, 123-125, 21 1, 216, 220-222, 225, 227, 229, 230, 241, 298, 299, 310, 311, 313 -, bryozoan-crinoidal bioherms, 227 -, evaporite deposits, 221, 222 -, evaporitic dolomite, 310, 3 1 1 , 3 13 -, interior basins, U.S.A., 229, 230 -, Madison Limestone, 299 -, Ouachita Mountains, 241 Mizzia, 95, 122, 125 Modern carbonate sediments, 13, 14, 29-50 Moenkopi Formation, Nev., U.S.A., 122, 211 Mohawk Valley, N.Y., U.S.A., 315 Mollazal's classification of limestones, YO, 91 Mollusks, 12, 13, 35, 39, 41, 77, 78, 94, 191, 206, 218, 219,228, 406, 407 -, comprising a chenier, 228 -, Florida Bay sediments, 21 8 -, Green River Formation, 206 - in beach rock, 406 _ _ eolianites, 407 - _ reef detrital material, 219 - susceptibility to recrystallization, 191 Monroe Dolomite, Ohio, U.S.A., 328 Montastrea, 21 I Montpelier Limestone, Jamaica, 236 Moore Hill Formation, Ont., Canada, 225 Moorman Ranch, Nev,, U.S.A., 122 Morgan Formation, Utah, U.S.A., 121, 124 Morris Ranch, Utah, U.S.A., 127 Morrowan age, 121, 122, 124 Mosaic, definition of, 161 Mountain Limestone, England, 216 - Pass, Nev., U.S.A., 123 Mount Carmel, Israel, 175 - Head Formation, Canada, 311 Mud, 94, 116, 161, 162, 240, 299, 405 - aggregate, definition of, 162 -, definition of, 161 -, inorganic, definition of, 240 -, organic, definition of, 240 -, replacement of lime, 299 - supported, definition of, 162

SUBJECT INDEX

-, terrigenous, 405

Mudstone, 94, 116, 162 -, definition of, 162 Mulinia, 228 Multiple generation of pores, 360 Muschelkalk, Germany, 310 Myrick Station Limestone, U.S.A., 230 Mytilus banks, 226 - califarnianus, 39 Nahcolite, 207, 283 Nansen Sill, 11 Napanee, Ont., Canada, 316 Nappa de la Brtche, France, 243 Nappe de Morcles, Haute Savoie, France, 239 Nari, 175, 198 -, definition of, 175 -, formation of, 198 Natica, 228 Negev, Israel, 302 Neogene, 203, 204 Neogenesis, 15 Neritic zone, 4, 14, 47 Nesquehonite, 90 Nevada Limestone, Nev., U.S.A., 321 Newark Basin, U.S.A., 212 Newberrian Sandstone, Germany, 228 Newman Hall Shale, Jamaica, 230 - Limestone, Tenn., U.S.A., 216 Newport Limestone, Jamaica, 2 16 New Providence Island, Bahamas, 183 - Red Sandstone, England, 322 - Scotland Formation, N.Y., U.S.A., 225 Niger River, Africa, 405 Nodular limestone, definition of, 162 Nonclastic, definition of, 162 Nondetrital, definition of, 162 Nonskeletal limestone, definition of, 162 North Bimini Island, Bahamas, 183 - Burbank Hills, Utah, U.S.A., 123 - Carson Oil Field, Alta., Canada, 389 - Confusion Range, Utah, U.S.A., 127 - Sea, 11 Northern Muddy Mountains, Nev., U.S.A., 127 Norwegian Sea, 11 Nottingham Oil Field, Sask., Canada, 381, 382 Nova Scotia, Canada, 204, 209 Nuculana, 228 Ocean floor, age of, 426 Oil, 23, 205, 206, 277, 278, 374, 377-379 - globules, movement of, 374 - migration, 23, 277, 278 - phase, continuous, 377-379 - shales, 205, 206 Oligocene, 204, 216, 239

463

SUBJECT INDEX

-, brecciole nummulitiche, 239 -, Clarendon Bank, 216

Olistolith, definition of, 162 Olivine, 242 Oklahoma, Permian evaporites, 223 Oncolites, definition of, 315 Oneota Dolomite, U.S.A., 329 Onslow Bay, N.C., U.S.A., 224 Onyx, 176, 177 Ooids, 35, 182, 183, 185, 202, 215, 222, 295, 296, 315 - in bank sediments, 215 -, mineralogy of, 182, 183 -, superficial, definition of, 185 -, replacement of, 295 Oolite, 56, 96, 117, 121, 124, 162, 171, 209, 222, 405 -, definition of, 96, 162, 405 -, photomicrograph of, 121 - range in present-day sediments, 405 -, Triassic, 209 Ooliths, 35, 38 Oolitic, 12, 17, 22, 96, 180, 201, 202, 204, 209 -, definition of, 96 Oolitoid, definition of, 162 Oopellet, definition of, 162 Oosparrudite, 99 Ooze, 30-34, 53, 173, 234, 235, 241 -, carbonate, 53 -, deep-sea, 30-34, 234, 235 -, Globigerina, 235 - organic, I73 -, pelagic, calcareous, 241 Opal, 1, 282 Open-space structures, definition of, 162 Ophiuroidea, 35 Oquirrh Formation, Utah, U.S.A., 121-124, 127 Oquirrh Mountains, Utah, U.S.A., 123, 201 Orbulina, 237 Ordovician, 3, 8, 9, 116, 190, 212, 213, 216, 241, 244, 272, 279, 315 -, Arbuckle Group, 213 -, Black River Limestone, 216 -, Ellenburger Limestone, 216, 272, 279 - facies, Mohawk Valley, N. Y . , 3 I5 - flysch, 190, 241, 244 -, Gaspe Peninsula, Canada, 241, 244 -, source of limestone and dolomite clasts, 212 Organic, I , 80, 90, 163, 173, 240 - lattice, definition of, 163 - limestone, definition of, I63 - matter, I, 80, 90 - ooze, 173 Organic structure limestone, definition of, 163 Organisms, 415, 420

-, evolution of, 415

- in

pelagic carbonates, 420

-, sediment-contributing, 396

Organogenic, 13 Origin of dolomites, 108-109, 276-308 - _ limestones, 174-193 Orinoco Shelf, South America, 229 Orthochemical, definition of, 163 Orthogeosynclines, 2, 422 Osagia, 95 Ostracoda, 35, 210, 314, 315, 390, 391, 409 -, fresh-water, 210 -, response to environmental temperature, 409 Ouachita Mountains, 0kla.-Ark., U.S.A., 241 Overgrowth texture, 93 Overton Fanglomerate, Nev., U.S.A., 210, 21 1 Ovoid grains, definition of, 163 Oxygen isotopes, 18, 19, 4 0 4 2 , 200, 291, 292, 296, 297, 331, 401 - _ , enrichment of dolomitized sediments, 297 - -, ISO in Cool Creek Limestone, 331 _ _ , I 8 0 in dolomite, 291, 296 _ _ , 180/160 ratio, 4042, 401 _ _, -, climatic indicator, 401 _ _, - of a diagenetically altered limestone, 42 Oyster(s), 175, 226 - mounds, 226 - stony mass, 175 Ozark Mountains, Ark., U.S.A., 224 Pacific Ocean, 1 1 , 16, 39, 40, 109, 171, 186, 187, 190, 191, 195, 196, 219, 240, 273, 296, 300, 404 _ _ , Bikini Atoll, 219 - _ _ , distribution of marine carbonates, 404 _ _, Elugelab Island, 195 - _ , Eniwetok Atoll, 186, 187, 195, 296 _ _ , Fiji, 191 _ _ , Funafuti Atoll, 196, 273, 296, 300 _ _ , Guam, 190, 191, 219 _ _ , Kapingamarangi Atoll, 219 _ - , Kita-daito-jima Atoll, 196, 296 - _ , Marshall Islands, 219 - _ , organic muds, 240 _ _ , Pago-Pago, 40 - _ , Palau, 39, 40 - _ , Saipan, 191 Packstone, 94, 116, 163 -, definition of, 163 Pago-Pago, Pacific Ocean, 40 Pakoon Formation, Nev., U.S.A., 126 Palau. Pacific Ocean. 39. 40 Paleocene, 203, 205, '207, 245, 248

464 Paleocene (continued) -, ancient lake sediments, 203, 205, 207 - -Eocene shales, Venezuela, wildflysch diamictite, 245, 248 Paleoclimatic normality, 404 Paleoecology, 52 Paleosol, 199, 406 - calcrete, 199 Paleotemperatures, isotope ratios, 42 Paleozoic, 2, 3, 5, 7, 10, 26, 116, 117, 120, 317 - carbonates, 2, 116 - carbonate cycles, 117 - environment, sediment-water interface, 317 Palliser Limestone, Alta., Canada, 216 Pancake Range, Nev., U.S.A., 121 Parachute Creek Member, Green River Formation, Wyo., U.S.A., 206 Paradox Formation, Utah, U.S.A., 354,355,385 _ - , Ismay zone, 354, 355 Paragenesis, definition of, 163 Parageosynclines, 423 Paragiiito boulder bed, Venezuela, 245 Paris, France, 65 Park City Formation, Utah, U.S.A., 126, 127 Parker Slate, Vt., U.S.A., 227, 322 Particle, 10, 59, 60, 102, 103, 276 -, definition of, 276 - identification, 59, 60 - size scales, 10, 102, 103 Particulate, definition of, 163 Pasteur Level, 26, 416 PaurocrystaIline, definition of, 163 Paurograined, definition of, 163 Pebble, 100, 210, 245, 316 Pekisko Formation, Alta., Canada, 222 Pelagic, 19, 183, 234, 235-237, 241, 405 - carbonate sediments, 19, 183, 234, 405 - Foraminifera, 235-237 Pelagosite, 13, 163, 406 Pelecypod(s), 35, 39, 122, 210, 228, 396 -, fresh-water, 210 - in a chenier, 228 -, sediment-contributing organism, 396 - valves in photomicrographs, 122 Pelite, definition of, 163 Pelitomorphic, 7, 163 -, definition of, 163 Pellet, 17, 22, 35, 56, 93, 95, 105, 136, 137, 163, 295 -, definition of, 95, 163 -, photomicrographs of, 136, 137 -, replacement of, 295 Pelletal rocks, photomicrographs of, 132-137 _ _ , bryalgal limestone, 133 _ - , calcarenite, 135 _ - , detrital limestone, 137

SUBJECT INDEX

_ _, skeletal limestone, 132, 134, 136

Pelmicrite, 99 Pembina, Alta., Canada, 193 Pembroke Eolianite, Bermuda, I88 Penecontemporaneous, definition of, 163 Peneroplids, recrystallization of, 191 Pennsylvanian, 117, 121, 123, 124, 127, 174, 210, 213, 220 -, Collings Ranch Conglomerate, 213 -, limestone-pebble conglomerates, 210 -, Midland Basin, 220 Pequop Formation, Nev., U.S.A., 122-125 Permeability, 15, 16, 21, 56, 177, 332, 333, 351, 391, 394, 395 - associated with the Swan Hills carbonates, 391 -, definition of, 351 -, control of, in carbonate rocks, 395 - _ - , - Leduc Formation, 394 -, effective, definition of, 351 - measurements in carbonate rocks, 332 -, relative, 21, 351 - Kingsport Limestone, 332, 333 Permian, 12, 19, 26,96, 105, 121-127, 190, 192, 199, 210, 211, 220, 222, 223, 241, 298, 306, 310, 311 -, Delaware Basin, 199, 241 -, Dunkard Group, 210 - evaporites, 222 - evaporitic dolomite, 310, 31 I -, Florena Shale, 298 -, Kaibab Limestone, 12, 124, 126, 127, 21 1 -, Scurry reef, 220 -, Tunisia, 306 Persian Gulf, 11, 20, 29, 65, 189, 221, 284, 286, 309, 314, 319, 334 _ - , geochemistry of water and sediments, 286 - _ , map of, 284 --, Recent dolomite, 284 --, - marine evaporites, 22 I _ _,salinity of, 334 Peru, Recent marine evaporites, 221 Pete Hanson Creek, Nev., U.S.A., 299 Petrographic model of carbonate rocks, 52-64 - parameters, 65 “Petrolog” of a quarry, 66 Petrologic fabrics, 62-65 pH, 4, 8, 24, 26, 42, 44, 191, 325, 414, 415 -, control of mineralogy, I91 -, differences in sediments and ocean, 42, 44 - of sea water, 4, 8, 24, 26, 44 -- sediment, 44 -, origin of syngenetic dolomite, 325 - Precambrian ocean, 414, 415 Phaneric, definition of, 163

465

SUBJECT INDEX

Phillipsburg thrust, Vt., U.S.A., 321 Phosphate, 1 , 52 Physicochemical agents in diagenesis, 56 Pic Marcelly, France, 243 Pieniny-KEppe zone, Poland, 235, 243 Pirssonite, in sediments, 283 Pisolite(ic), 16, 65, 66, 70, 164, 177, 198, 199 -, definition of, 96, 164 -, limestone, 16, 65, 66, 70, 198, 199 Plantagenet Bank, Atlantic Ocean, 295, 296 Plants, evolution of, 419 Platform, 2, 11, 19, 35, 65, 117 Plattendolomit, Germany, 310 Plattin Flysch, Switzerland, 241, 242 Pleistocene, 7, 35, 47, 154, 172, 176, 177, 179, 180, 183, 203, 204, 224, 231-233, 237, 278, 280, 291, 293, 325, 406 -, ancient lake sediments, 154, 203, 204, 291 -, carbonate sediments, 172, 237 -, deep-sea carbonate sediments, 278 - dolomites, 280, 293 - dunes, 233 - eolianites, 231-233 - gypsite, 293 -, Lake Eyre dolomite, 293 -, - Lahonton, 172, 176 - mineralogy of ooids, 183 - - - skeletal sands, 179 - sea level changes, 224 -, weathering processes, 325 Pliocene, 47, 196, 233 - dunes, 233 Plymouth Limestone, South Devon, England, 212 Plympton Formation, Utah, U.S.A., 126, 127 Poikilotopic, definition of, 303 Polinices, 228 Pore, 23, 92, 360, 371-373, 377 - development, multiple generation, 360 - geometry, 23, 271, 377 - interconnections, frequency distribution of, 373 -, measurements of, 372 - space, classification of, 92 - sizes, frequency distribution of, 373 -, types of interconnections, 371 Porifera, 35, 64 Porites, 217 Porosity, 5, 15, 16, 21, 22, 56, 70, 92, 109, 177, 180, 186, 188, 299, 349, 351, 354, 355, 357, 358, 360, 361, 367, 387, 391, 392, 394, 395 -, absolute, definition of, 351 - associated with Swan Hills carbonates, 391 -, control of, in carbonate rocks. 395 -, --, in Leduc Formation, 394 &

.

-, definition of, 351 -, effective, 351

_ _ in Swan Hills Member, 392

- fracture, 360 - framework, 349, 355 - function of original sediment composition, 363 -, interalgal vug, 358 -, interlump, 358 -, interoolitic, 358 -, interparticle, 188 -, interpellet, 358 -, interskeletal vug, 358 -, moldic, definition of, 186 -, mud, 349, 357 -, pisolitic limestone, 70 -,primary, 109, 349, 352, 354, 355, 358 _ ,- , classification of, 355 _ , _ , definition of, 349 _, _, modification of, 354, 358 -, relationship to permeability, 367, 387 -, - in Zvunoviu-rich limestones, 387 -, sand, 349, 357 -, -, definition of, 349 -, secondary, 186, 349, 357, 358 -, -, definition of, 349 -, -, produced by solution, 358 _ , _ , voids, definition of, 186 -, selective solution of shells, 358 -, sucrose dolomite, 349, 361 -, types of, 349 -, variation with depth, 70 Porphyroblastic, 111, 164 -, definition of, 164 Porphyrotopic, definition of, 302, 303 Port Ewen Formation, N.Y., U.S.A., 225 Potomac marble type, Newark Basin, U.S.A., 212 Potomac River, U.S.A., 212 Pozzillo 1, Sicily, Italy, 72, 75, 77, 78, 80, 81 Precambrian, 2 4 , 25, 193, 220, 225, 242, 400, 410418 - Algae, 400, 412 - Arabo-Nubian shield, 242 - basement rocks, 225 - evolution of invertebrates, 417 -, geochemical variables during, 410-445 - Lipalian interval, 416 -, Montana, 220 -, occurrence of limestones, 193 - sea, 417 -, silica-iron deposition, 417, 418 Precipitates, chemical, 187, 271, Preferred fabric, definition of, 164 Pressure solution, definition of, 164

466 Price River Formation, piedmont facies, Utah, U.S.A., 21 I Primary dolomicrite, photomicrographs of, 130 - dolomite, 3, 5 , 108, 109, 114 - -, definition of, 114 _ _,formation of, 3, 5 Principle of Uniformitarianism, 24, 400 Processes, 196, 197, 236-240, 267, 268 -, bacterial, dolomite origin, 267, 268 -, gravity-displacement, 236-240 -, high energy, 197 -, lateral sedimentation, 196, 197 -, low energy, 197 -, vertical sedimentation, 196 Promontory Mountains, Utah, U.S.A., 201 Protista, secreted plates, 239 Protodolomite, 275, 276, 286, 307 -, definition of, 275 - in sediments, 286 -, precipitated in laboratory, 307 Providence Limestone, U.S.A., 230 Provo Canyon, Utah, U.S.A., 122, 124 Pseudobreccia, definition of, I64 Pseudomorphic replacement, definition of, 164 Pseudo-oolites, 163, 164, 222 -, definition of, 164 Pseudo-ooliths, 185 Pseudoschwagerina, 122 Pteropods, 19, 30, 34, 235, 239 - ooze, 19, 235 Puerto Rico Trench, Atlantic Ocean, 241 Pullem Creek, Utah, U.S.A., 121 Purbeckian strata, Switzerland, 208 Pyrite, 80, 90, 207, 301 - in Green River Formation, 207 -, oxidation of, 301 Pyroxene, 242 Qatar Peninsula, 221, 286 Quarry, petrology of, 66 Quartz, 7, 90, 184, 192, 193, 208,223, 224,242, 287,292, 357 -, calcium carbonate coating of, 184 - in Bonneville Lake sediments, 292 - _ carbonate rocks, 223, 224 -, pore cement, 357 -, replacement by carbonateminerals, 192,193 -, terrigenous grains in lacustrine limestone sediments, 208 Quartzites, associated with dolostones, 322 Quaternary, 400, 401, 423 Quebec, Canada, limestone conglomerates, 246 Queen Formation, Texas, U.S.A., 31 I Queens Limestone, Texas, U.S.A., 190 Quinqueloculina, 228

SUBJECT INDEX

Racine Formation, Ill., U.S.A., 328 Radiocarbon dating, 43, 46, 403 Radiolaria, 78, 81 Ragusa, Sicily, Italy, 16, 72-79 - Plateau, 72-79 Raible, Italy, 332 Rangia, 228 Recent, 29-50, 172, 173, 178-187, 196, 219, 237, 280-294 - carbonate sediments, 29-50, 172, 219, 237 - dolomites, 280-294 - mineralogy of oolds, 182 - sediments, changes in grain mineralogy, 178-187 Recrystallization, 44, 47, 93, 164, 165, 190, 191 -, definition of, 164 -, experimental work, 191 - fabric, definition of, 165 - of metastable carbonates, 44 Red beds, 10, 12, 321 - River Formation, Mont., U.S.A., 362, 363 - _ Valley, Minn., U.S.A., 326 -Sea, 11, 180, 183 Redwall Limestone, Utah, U.S.A., 125 Redwater Oil Field, Alta., Canada, 394 - reef, Alta., Canada, 298, 388, 393-396 _ _ , cross-section of, 395 _ _ , facies map of, 394 - -, sediment-contributing organisms in, 396 Reef, 7, 9, 11, 16, 24, 34, 53, 57, 58, 64, 72, 75, 78, 79, 98, 109, 117, 165, 175, 176, 180, 186, 187, 194, 196, 216-220, 243, 296, 298, 306, 328, 382, 384, 388, 393-396, 421 - bank, definition of, 165 - barrier, definition of, 165 -, Bikini Atoll, 186, 187, 219 -, biohermal, 95 -, biostromal, 95 - breccia, composition of, 219 -, British Honduras, 217 -, Capitan Formation, 243, 306 -, cold-water, 421 -, coral, 109 -, Cuba, 217 -, detrital, 219 -, dolomitization processes, 296 -, dolomitized, examples of, 109, 328 -, fringing, definition of, 165 -, Goat Seep, 243, 328 -, Golden Spike, 298 -, Heron Island, 34 -, Horseshoe Atoll, 382, 384 -, Jamaica, 217 -, Kapingamarangi Atoll, 219 -, Kita-daito-jima, Atoll, 196, 296 -, Louisiana, 219

SUBJECT INDEX

Reef (continued) -, milk, definition of, 165 -, Paleozoic, 117 - patch, definition of, 165 -, pinnacle, definition of, 165 -, Precambrian, 412 -, Recent sediments, 180 -, Redwater reef, 298, 388, 393-396 -, Rimbey-Leduc-Meadowbrook, 298 - rock, origin of, 175, 176 -, Scurry, 220 -, sediment-contributing organisms, 396 -, - deposits, 216-220 - shoal, definition of, 165 - table, definition of, 165 -, tufa, definition of, 165 -, Wabash, 328 -, Willingdon, 298 -, Wolayersee, 57, 58 Reefal, 93, 131, 139, 165 Refluxion, 268, 338 - process in dolomite formation, 338 Relative permeability, 21, 351, 352 Relay Creek Dolomite, Okl., U.S.A., 322 Relic, definition of, 165 - textures, I1 1 Replacement, 109, 182, 294-301, 327-330, 361 - equation, of calcite, 361 of calcium carbonate by dolomite, 294-300, 327-330 _ _ gypsum, 301 -- limestone, 109 - _ quartz, 301 -, paramorphic, 182 -, proof of, 295 Reservoir rocks, 21-24, 349-398 - -, permeability of, 351, 352, 367, 373-380, 387, 389 - _ , Petrology and paleoecology of, 380-396 _ _ , porosity of, 351-380, 387, 389 _ _ , properties, classified according to facies, 390, 391, 393 Revolution, 399, 415419, 420, 422, 423 - I , 399, 415416 - 11, 399, 416418 - 111, 399, 419 - IV, 399, 420 - V, 399, 420 -, geotectonic, effect of, 422, 423 Rhabdoliths, 30 Rhine Valley, Germany, 206 Rhinoceros, 401 Rhizomorphs, definition of, 402 Rhodesia, 4 Rhodophyta, 150 Rib Hill, Nev., U.S.A., 125

-

467 Riepe Spring Limestone, Nev., U.S.A., 122, 125 Rim cement, 165, 180, 187 - _ , definition of, 165, 187 Rimbey-Leduc-Meadowbrook reef chain, 298 Roberts Mountains Formation, Nev., U.S.A., 244, 299 Rocky Mountains, U.S.A., 116, 372, 373, 380 Rogers Spring Limestone, Utah, U.S.A., 125 Rome Formation, U.S.A., 320, 322 Ruby Marshes, Nev., U.S.A., 126 Rudistids, stony mass, 175 Rugg Brook Dolomite, Vt., U.S.A., 322 Rundle Group, Alta, Canada, 216 Rush Creek Sandstone, Okla., U.S.A., 322 Russian Platform, 310, 312, 414 - _ , Ca/Sr ratio during Precambrian, 414 Rutledge Limestone, Appalachians, U.S.A., 32 1 Rysedorph Hill Conglomerate, N.Y., U.S.A., 247 Saccharoidal, definition of, 166 Saint Albans Hill, Vt., U.S.A., 227 - _ Slate, Vt., U.S.A., 322 - David cyclothem, U.S.A., 230 - George’s Island, Bermuda, 180, 188, 232 - Louis Limestone, Ind., U.S.A., 313 Saipan, Pacific Ocean, 191 Sakawa orogeny, 422 Salina Formation, Mich., U.S.A., 312 Saline facies, Green River Formation, Wyo., U.S.A., 207 - minerals, abundance of, in sediments, 283 Salinity, 24, 191, 203, 284-286, 288 -, control of mineralogy, 191 -, mechanism for increase of, 288 -, Persian Gulf, 284-286 -, Salton Sea, 203 Salt (see evaporite and halite), 7-10, 290 - deposition, 10 - domes, 290 - Flat Graben, Texas, U.S.A., 289-293, 296, 309, 331, 335-337 - _ _ , isotopic composition of dolomite, 29 I - Flats, Utah, U.S.A., 7 Salton Sea, Calif., U.S.A., 199, 202, 203 - _ , formation of, 203 - _ sediments of, 202-203 Saltonstall Ridge, Conn., U.S.A., 212 Samra Formation, Israel, 204 San Andres Limestone, Texas, U.S.A., 190 Sand, 33, 35, 44, 179, 201, 202, 224, 225, 232, 278, 279 -, admixed carbonates and quartz, 224, 225 -, beach, 44

468 Sand (continued)

-, foraminiferal, 32, 35

- in algal material, 202 -, oolitic, Great Salt Lake, 201 -, pelletal, 33, 35 Sanpete Valley, U.S.A., 208 Santa Cruz Mountains, Jamaica, 216 Sarah Lake Oil FieId, Alta,. Canada, 389 Saudi Arabia, 117, 190, 199, 216, 223, 242, 248, 298, 329, 366-368 - -, Arab Formation, 190,216,223,298, 329 - -, Arabo-Nubian shield, 242 - _ , Miliolitic Formation, 199 --, reservoir rocks, 117, 366-368 Sawtooth Mountains, Mont., U S A . , 125 Saxe Brook Dolomite, Vt., U.S.A., 322 Scaglia, 242 Scandinavia, snowfields, 403 Schmidt-Urey model, 25 Schwagerinid fusulinids, 122, 123, 131, 134 Scruton model, evaporite zonation, 312 Scurry reef, Texas, U.S.A., 220 S-dolostones, 273, 328, 332, 333 -, definition of, 273 -, Mascot Dolomite, Tenn., U.S.A., 332, 333 Sea level, 410, 423425 _ _,control of neritic facies, 424 - _ , eustatic change, 410, 423 - -, geodetic changes, 425 _ - , thalassocratic and epeirocratic conditions, 425 Sea of Okhotsk, 11 Searles Lake, Calif., U.S.A., 177 Sea water, carbon dioxide concentration, 414 Sebkhas, 285-287,289 -, definition of, 285 -, growth of, 289 Secondary, definition of, 165 Sediment(s), 19, 30, 3442, 44, 119, 172, 183, 202, 203, 216-235, 237, 240, 241, 283, 284, 291, 317, 328, 335, 396, 405, 490, 423 -, abundance of saline minerals in, 283, 284 -, alterations of offshore and nearshore carbonates with terrigenous, 228-23 1 -, bioclastic, 37 -,carbonate, 30, 35-42, 119, 172, 219, 237 -, compaction of, 335 -, contributing organisms, 396 -, couplets, 291 -, deep-sea pelagic carbonate, 405 -, detrital carbonate, 409 -, dolomitized, examples of, 328 -, hemipelagic, 240 -, Orinoco Shelf, 229 -, parageosyncline, 423 -, pelagic, 19, 183, 234, 235, 241, 405

SUBJECT INDEX

-, Persian Gulf, 286 -, playa, 283, 284 -, protodolomite in, 286

-,

Puerto Rico Trench, 241

-, reef, 216-220

-, Salton Sea, 202, 203 -, shelf, 34, 35 -, slumping, 30 -, source of, reefs, 216 -, terrigenous, 221-233 -, volcanic materials in, 229 --/water interface, 317 Sedimentation, neritic, 405 Seepage refluxion, 19, 289, 329, 335 _ _ , concept of, 335 - _ , diagenetic dolostone origin, 329 Sepiolite, 207 Serpula bioherm, 226 Seven Rivers Formation, Texas, U.S.A., 31 1 Shadow Lake Formation, Ont., Canada, 225 Shady Dolomite, Appalachians, U.S.A., 323 Shafter, Nev., U.S.A., 129 Shales, associated with dolostones, 322 Shark Bay, W. Austr., 416 Shelf, 2, 8-12, 19, 24, 34, 35, 117, 222, 381, 383 -, environmental classification of limestones, 383

-, lagoon sands, 117

- sediments, 34, 35 Shell, 227, 296, 300, 417, 419 -, aragonite, 296 -, chitinous, 417 -, dolomitization of, 300 -, formation of, in invertebrates, 419 -, Wadden Sea, beds, 227 Sherman Fall Formation, Ont., Canada, 225 Shoal, 9, 390 Shortite, 206 Shrimp, 202 Shunda Formation, Canada, 222, 298, 31 1 Shuttle Meadow Formation, Conn., U.S.A., 212 Siderastrea, 217 Siderite, I , 52, 90, 193 -, replacement of quartz, 193 Silica (see quartz), 1,4, 14,22,90, 321,417,418 -/iron deposits, Precambrian, 417, 41 8 -, Knox Group, Appalachians, 321 - Oil Field, Kans., U.S.A., 369-371 Sillery Formation, Que., Canada, 246 Silt, in algal material, 202 Silurian, 216, 220, 241, 299, 310, 312 -, evaporitic dolomite, 310, 312 -, Gotland, Sweden, 216 Simonson Dolomite, Nev., U.S.A., 321 Skeels Corners Slate, Vt., U.S.A., 227

SUBJECT INDEX

Skeletal, 39, 93-95, 105, 106, 129, 131-136, 142, 166, 188, 191 -, calcisiltite, photomicrographs of, 135 -, definition of, 166 --/detrital limestone, photomicrographs of, 129, 131-133, 142 - grains, 93, 94, 105, 106 - limestone, 94 --, photomicrographs of, 132-134, 136 - material, susceptibility to recrystallization, 191 - sands, photomicrographs of, 188 Slates, associated with dolostones, 322 Slice map showing variation in effective porosity, 392 Slumping, 62, 244 - Delaware Basin, 244 Snails, 407 Soils, 325, 326, 400, 406, 407 -, calcareous, 406, 407 -, dolomite formation in, 326 -, fossil, 407 -, preserved, 400 Solar radiation, 401, 403 Soldier Summit Area, Utah, U.S.A., 205 Solenoporae, 77, 78, 124 Solo River, Java, Indonesia, 405 Solution, 166, 358, 359 -, acidic ground water, 359 - of carbonate rocks, 358 -, rate of, 359 - transfer, definition of, 166 Somaliland, 190 Somerset Eolianite, Bermuda, 188 Soreq Formation, Israel, 319 South Australian lakes, 292 - Cat Cay, Great Bahama Bank, 182 - China Sea, 19 - Devon, England, 212 - Schell Creek Range, Nev.,U.S.A., 125 - Schmid Ridge, Idaho, U.S.A., 125 - Tintic Mountains, Utah, U.S.A., 127 Southern Butte Mountains, Nev., U.S.A., 122 - Wasatch Mountains, Utah, U.S.A., 127 Sparite, 52, 56, 122, 132, 134, 136, 141, 166 - cement, photomicrograph of, 122, 132 -, definition of, 166 -, photomicrographs of, 134, 136, 141 Sparry, 92, 93, 166 - calcite, 92, 93 -, definition of, 166 Spathization, definition of, 272 Speleal limestone, definition of, 166 Speleothems, 166, 176, 177 -, definition of, 166 Spergenite, definition of, 166

469 Spherulite, definition of, 167 Sponges, 22, 93, 175 Spring Mountains, Nev., U.S.A., 123, 125 _ _ Formation, Nev., U.S.A., 122, 125, 126 Spruce Mountain area, Nev., U.S.A., 125 Sprudelstein, 177 Stachyodes, 396 Stafford Creek, Andros Island, Bahamas, 196 Stalactites, 177 Stalagmites, 177 Stansbury Island, Utah, U.S.A., 202 Star Range, Utah, U.S.A., 124, 125 Stinkstein, definition of, 167 Stockton, Utah, U.S.A., 127 Stony biogenic incrustations, 175, 176 - Point, N.Y., U.S.A., 212 Strait of Hormuz, 11 Strassfurt beds, Germany, 310 Stratigraphic discontinuity, diagenetic dolostone, 328, 329 - patterns, 197, 288 Stream deposits, limestone particles, 210 Streblus, 228 Streppenosa Formation, Sicily, Italy, 75-81 Strom-algal consortium, sediment contributing organism, 396 Stromatactis, 125, 126, 167 -, definition of, 167 Stromatolites, 77, 78, 167, 315, 317, 416 -, definition of, 167, 315 -, origin of, 416 Stromatoporoids, 64, 98, 125, 175, 297, 314, 390, 391, 396 -, Manlius Formation, N.Y., U.S.A., 314 -, massive, 396 -, Redwater reef complex, Alta., Canada, 394 -, stony mass, 175 -, tabular, 396 Strontium, 24, 39, 333, 411 -/calcium ratio, 24, 39, 41 1 - depletion in dolostones, 333 Sturtian Tillite, Australia, 417 Stylolites, dolomite concentration along, 297, 298, 333 Subhedral, definition of, 167, 302 Subsidence, 10, 64, 80 Sucrose dolomite, 361-363 Sucrosic, 143, 167 -, definition of, 167 Sulfates, 11, 199, 200 Sulfides, 1, 14, 42 Sulfur, 290, 291 - in Pleistocene graben deposits, 290 Summit Springs Member, Pequop Formation, Nev., U.S.A., 121, 126 Summum cyclothem, U.S.A., 230

470 Sunda Shelf, Indonesia, 405 Sunderland Shale, Jamaica, 230 Sunspot cycles, 424 Swan Hills Member, Alta., Canada, 388-390, 392, 393 ---, cross-section of the reservoir, 392 - - -, facies of, 390 - - -, isopach map, 389 - - -, slice map, 392 - - Oil Field, Alta., Canada, 388-389 Swanswick Limestone, Jamaica, 21 6 Sylvite, 288 Syndiagenetic calcite, 93 Syndiagenetic stage, 5 , 6, 167 Syneresis cracks, definition of, 167 Syngenetic, 20, 167, 268, 308-325 -, definition of, 167 - dolomite, 20, 268, 308-325 Syntaxial rims, definition of, 167 Tablehead Formation, Nfld., Canada, 246 Taconic orogeny, 422 T-dolostone, 273, 330 Tan Dolomite, Sicily, Italy, 77, 80, 81 Tansill Formation, Texas, U.S.A., 301, 311 Taormina Formation, Sicily, Italy, 75-77, 80 Tasman Sea, 11 Tectono-eustasy, 423 Tennessee zinc district, 334 Ten Thousand Islands, Fla., U.S.A., 40 Temperature distribution in geologic time, 400403 Teodorovich's classifications, 97, 98, 107, 110, 111, 374 - _ of limestones, 97, 98 _ _ - dolomites, 107, 110, 111 _ _ _ pores, 374 Terra rossa, 407 Terrigenous sediments, 221 -231, 309-324 Tertiary, 175, 190, 191, 210, 211,219, 229, 237, 239,244, 310, 313, 400, 424 - chalk, 175 - evaporitic dolomite, 310 -, Italy, 237, 239, 244 -, Guam, 190, 191 -, Gulf Coast, U.S.A., 229 - fanglomerate, 210-21 1 -, Mediterranean transgressions, 424 - reefs in Louisiana, 219 - soils, 400 Tethyan transgressions, 426 Texture, 91-106, 110, 297, 302-306 -, classification of dolomites, 110 - of carbonates, 91-106 -, preserved after dolomitization, 297, 302 Thalassocratic conditions, 425

SUBJECT INDEX

Thaynes Formation, Idaho, U.S.A., 122, 123 Thenardite, 283 Thermal gradients, equator-pole, 403 Thermonatrite, 283 Three-phase flow of fluids in rocks, 23 Thrust conglomerates, 19 Tiberias Basin sediments, 204, 293 Tiburtino, 176 Tidal, 285, 286, 316, 317, 319 -channel deposits, 316, 317 - flat deposits, 319 - flats, 285, 286, 317 - origin of dolostones, 319 Tintinnidae, 78 Tivoli, Italy, 176, 177 Toana Range, Nev., U.S.A., 126 Tomstown Dolomite, Appalachians, U.S.A., 323 Tongue of the Ocean, Great Bahama Bank, 237, 239, 242 Toroweap Formation, Ariz., U.S.A., 11, 12, 122, 123, 124, 126, 127 Tortuosity, 21 Trace elements, 210 Transgressions, during geologic past, 424, 425 Transported carbonate skeletal detritus, 227, 228 Travertine, 12, 17, 18, 168, 171, 176, 177 -, definition of, 168 -, origin of, 176, 177 Travertino, 176 Triassic, 4, 6, 7, 11, 72, 75, 77, 81, 122-124, 203, 204, 208, 209, 21 1, 21 2, 220, 3 10, 323 -, ancient lake sediments, 203, 204, 208, 209 -, conglomerates, 212 -, evaporitic dolostones, 3 10 -, marls, 323 -, Moenkopi Formation, 21 1 -, Tyrolean Alps, 220 Tribes Hill Formation, N.Y., U.S.A., 301, 310, 317, 328 Trilobites, 244, 296, 416, 417 -, Cambrian, 244 -, replacement of calcite in, 296 Trinity Group, Texas, U.S.A., 327 Trificites, 121, 122 Trochitenkalk, Germany, 310 Trona, 206, 283 Troy Limestone, Jamaica, 216 Tufa, 12, 168, 171, 176, 177 -, definition of, 168 -, granolare, 176 -, litoide, 176 -, origin of, 176-177 Turbidites, 19, 30-34, 239, 422 calcareous, 239

-.

47 1

SUBJECT INDEX

Turbidites (continued)

-, Hamilton Group, 422 -, terrigenous, 239

Turbidity, 168, 236, 237, 239, 241, 242

-, definition of, 168 - currents,

236-237, 239, 241, 242 Turbonilla, 228 Turtles, fossils in Green River Formation, 206

Uinta Formation, Utah, U.S.A., 203, 205 - Mountains, Utah, U.S.A., 122, 124, 126, 127 Underclay limestone, 209 Uniformitarianism, 24, 400 Upper Glen Rose Formation, Texas, U.S.A., 279 U.S. Geological Survey, drilling in Pacific atolls, 171 UV radiation, 416, 418 Vadose solutions, 56 Van Allen belt, 422 -Vactor Gypsum, Okla., U.S.A., 31 I , 322,329 Verdrangung, definition of, 295 Verdrangungdolomit, definition of, 295 Vichy, France, 177 Vigny, France, 16, 65, 67, 70 Villagonia Formation, Sicily, Italy, 75, 78, 79, 81 Vintage Dolomite, Appalachians, U.S.A., 323 Virgilian age, 122 Virgin Hills Oil Field, Alta., Canada, 389 - Limestone Member, Moenkopi Formation, Nev., U.S.A., 121, 122, 124 Virginia, U.S.A., Triassic deposits, 204, 208 Volcanic, 229, 242, 244, 313, 416 - activity, 416 - ash, 244 - materials in carbonate sediments, 229, 242 -tuff, 313 Volumetric replacement, 363 Wabamun Formation, Canada, 31 1 Wabash reef, Ind., U.S.A., dolomitized reef, 328 Wackestone, 94, 116 Wadden Sea, 227 Wales, Great Britain, 325, 333 -, epigenetic dolostones, 333 Walther’s Law of Facies, 197 Wasatch Formation, Utah, U.S.A., 203, 205 - Mountains, Utah, U.S.A., 201 - Plateau, Utah, U.S.A., 207, 208 Water covered areas, during geological past, 425

Wave working, importance of, 172 W-dolostones, 273, 328 -, definition of, 273 Weathering processes, formation of dolomite, 325, 326 Weber Canyon, Utah, U.S.A., 124 - Formation, Utah, U.S.A., 121,122,126,127 WedelI Sea, 1 1 Wellenkalk, Germany, 310 Wellington Formation, Okla., U.S.A., 31 1 Wells Formation, Idaho, U.S.A., 12.5 Wellsville Mountain, Utah, U.S.A., 123, 124 Werra beds, Germany, 310 West African Shelf, 405 - Castleton Formation, N.Y., U.S.A., 240,244 - Virginia, U.S.A., 210 Western Judith Mountains, Mont., U.S.A., 124 White Basin, Nev., U.S.A., 207 - Limestone, Jamaica, 215, 216, 236 Whitehorse Group, Okla., U,S.A., 322 - Pass, Nev., U.S.A., 126 Whitings, Dead Sea, Israel, 200 Wilberns Formation, Texas, U.S.A., 227 Wildflysch, 19, 241, 245, 247, 248 -, definition of, 245 -, diarnictites, 245, 247, 248 Williamsfield Trough, Jamaica, 21 6 Willingdon reef, Canada, 298 Williston Basin, U.S.A., 372, 373, 380 Winnow, definition of, 168 Wolayersee reef complex, Carnia, Austria, 57, 58 Wolfcampian age, 122, 123, 126, 127 Wolfs Hollow Member, Tribes Hill Formation, N.Y., U.S.A., 317 Woodbend Group, Alta., Canada, 394 WoodRiverFormation,Idaho, U.S.A., 121,122 Worcestershire, England, mark, 323 Worms, serpulid, stony mass, 175 Wurtzilite, 207 Xenotopic, 302-306 X-ray, 60, 82, 269, 275 - diffractometry, 60, 82 - revolution, 269, 275 Yates Formation, Texas, U.S.A., 31 I Yellow Bank, Great Bahama Bank, 34, 36, 43, 44, 45, 46 Yellowstone Park, Wyo., U.S.A., 171 Zechstein, evaporite succession, 310 Zechsteinkalk, Germany, 310 Zeravshan Valley, U.S.S.R., 326 Zooxanthellae, 421

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