Diagenesis in Sediments

DEVELOP MENTS IN SEDIMEN TOLOGY 8

DEVELOPMENTS IN SEDIMENTOLOGY 8

DIAGENESIS IN SEDIMENTS EDITED BY

GUNNAR LARSEN Department of Geology, University of Aarhus, Aarhus (Denmark) AND

GEORGE V. CHILINGAR University of Southern California, Los Angeles, Calif. (U.S.A.)

II

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CONTENTS

CHAPTER 1. INTRODUCfiON G. LARSEN (Aarhus, Denmark) and G. V. CH!UNGAR (Los Angeles, Calif., U.S.A.) CHAPTER 2. PHASES OF DIAGENESIS AND AUTHIGENESIS R. W. FA!RBRIDGE (New York, N.Y., U.S.A.)

19

...

CHAPTER 3. DIAGENESIS OF SANDSTONES E. C. DAPPLES (Evanston, Ill., U.S.A.)

. . . . . .

91

CHAPTER 4. DIAGENESIS IN ARGILLACEOUS SEDIMENTS . . . . . . . . .

127

G. V. CHILINGAR (Los Angeles, Calif., U.S.A.), H. J. BISSELL (Provo, Utah, U .S.A.) and K. H. WoLF (Canberra, A.C.T., Australia) . . . . . . . . . . . . . . . . .

179

G. MOLLER (Heidelberg, Germany) CHAPTER 5. DIAGENESIS OF CARBONATE ROCKS

CHAPTER 6. SILICA AS AN AGENT IN DIAGENESIS E. C. DAPPLES (Evanston, Ill., U.S.A.)

. . . . . . . . .

. . . . . . . . . . . .

323

. . . . . . . . . . .

343

. . . . .

391

CHAPTER 7. DIAGENESIS OF ORGANIC MATTER E. T . DEGENS (Pasadena, Calif., U.S.A.) CHAPTER 8. DIAGENESIS OF COAL (COALIFICATION) M. TE!CHMiiLLER and R. TE!CHMiiLLER (Krefeld, Germany)

. .

CHAPTER 9. DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS G. C. AMsTUTZ (Heidelberg, Germany) and L. BUBENICEK (Maizieres-les-Metz, Moselle, France) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 417 CHAPTER 10. DIAGENESIS OF SUBSURFACE WATERS E. T. DEGENS (Pasadena, Calif., U.S.A.) and G . V. CHILINGAR (Los Angeles, Calif., U.S.A.) 477 CHAPTER 11. INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS W. VON ENGELHARDT (Ttibingen, Germany) . . . . . . . . . .

.

503

CHAPTER 12. CONCLUDING REMARKS ON DIAGENESIS G. LARSEN (Aarhus, Denmark) and G. V. CHrLINGAR (Los Angeles, Calif., U.S.A.)

523

REFERENCES INDEX

525

SUBJECT INDEX . . .

537

INTRODUCTION GUNNAR LARSEN AND GEORGE V. CHILINGAR

Department of Geology, University of Aarhus, Aarhus (Denmark) University of Southern California, Los Angeles, Gal$ (U.S.A.)

A CENTENARY

The term “diagenesis” is almost a hundred years old. It was introduced by VON GUEMBEL in 1868 as a designation for processes which act on the sediments after deposition (for details see AMSTUTZ and BUBBNICEK, 1967). Thus the recognition of the processes of diagenesis is not new; however, the knowledge about diagenesis has been quite insufficient until the last few years. For example, less than 15 years ago TRASK(1951) in the survey “Dynamics of sedimentation” had to state that the processes of diagenesis were not well understood. In the paper “Diagenesis”, SUJKOWSKI(1958, paper prepared in 1954) opened the introduction with the following sentence: “Few important problems in geology have been so little studied as diagenesis.” Since then, however, the insight into this particular field of lithogenesis has progressed rapidly; and the amount of literature on diagenesis has rapidly multiplied during the last decade. This rapid progress and the appearance of an extensive literature on the subject have created a natural qeed for a survey of the present-day knowledge on diagenesis. Inasmuch as the present book Diagenesis in Sediments is published almost a hundred years after the introduction of the concept of diagenesis, it marks the centenary of this event. The present book contains contributions (by invitation) from a number of authors. The chapter (by R.W. Fairbridge) immediately following the introduction contains a general account of the diagenetic evolution, namely, the successive changes of the type of diagenetic processes taking place in a sedimentary sequence during the sinking and deep burial and the following rise to the erosional level. Furthermore, Fairbridge’s contribution deals especially with the authigenic minerals, which formed during the various stages of the diagenetic evolution. It is to be noted that almost all types of sedimentary materials are discussed in more or less detailed manner in this chapter. After this general account, a number of chapters describe diagenetic phenomena related to special materials. Some of the chapters deal with the main types of sediments: Sands (E.C. Dapples), Argillaceous deposits (G. Miiller), Carbonates (G. V. Chilingar, H. J. Bissell, and K. Wolf), Coal (M. Teichmiiller and R. Teichmiiller), and Sedimentary mineral deposits (G. C. Am-

2

G. LARSEN AND G. V. CHILINGAR

stutz and J. Bubenicek). The diagenesis of materials occurring as more or less subordinate components within the above mentioned main types of sediments are discussed in other chapters: Silica (E. C. Dapples), Organic matter (E. T. Degens), and Subsurfacewaters (E. T. Degens and G. V. Chilingar). There are no individual chapters on evaporites and phosphatic deposits; in these and similar cases, however, reference can be made to the chapter by R. W. Fairbridge where, as previously mentioned, diagenesis of most sedimentary materials is discussed. In addition, the chapter on Sedimentary mineral deposits by Amstutz can be referred to in this connection, because one section in this chapter has a brief introduction to literature dealing with a number of different sedimentary materials. Chapter 1 1 was prepared by W. von Engelhardt, who summarized the present-day knowledge on the subject of interstitial solutions and diagenesis in sediments. In the present introduction only some of the problems related to diagenesis are mentioned.

NARROWING DOWN THE SUBJECT

The diagenesis represents only some of the facets of the history of a sedimentary material. The following brief discussion is presented here in an attempt to elucidate the position of diagenesis within the field of lithogenesis. According to the Glossary of Geology and Related Sciences (1962) sedimentation can in short be defined as ". . . that portion of the metamorphic cycle from the separation of the particles from the parent rock, no matter what its origin or constitution, to and including their consolidation into another rock.. This portion of the metamorphic cycle comprises a large number of richly varied types of events, controlled by factors such as: (a) physical and chemical processes, and (b) tectonic and morphological conditions in both the field of accumulation and in that of denudation. These events can be grouped into a number of more or less distinct stages, as for instance the following three groups: (I) The disintegration and decomposition of the parent material, i.e., the effect of the mechanical and chemical weathering. (2) The separation of the weathering products, i.e., the effect of the processes of erosion, transportation and deposition, leading to the formation of the many different types of sediments. (3) The consolidation and cementation of the deposited material, i.e., the diageneticprocesses. It must be emphasized, however, that the above clear classification is at the same time artificial, because distinct systems and sharp boundaries are scarce in nature. For example, the development of typical weathering profiles (GOLDICH, 1938; WAHLSTROM, 1948) evidently is the result of not only weathering processes but also of removal of weathered material, i.e., leaching of soluble salts started

.".

INTRODUCTION

3

during the weathering itself. Thus, the events which give rise to typical weathering profilesare transitional between the above mentioned groups I and 2. The boundary between groups 2 and 3, i.e., between deposition and diagenesis, also can hardly be drawn clearly. It can be stated that diagenesis does not start until the moment the material is deposited, for instance on the sea floor. At the same time, however, it must be added that the diagenetic processes at work between the newly formed sediment and the overlying sea water do not have to be fundamentally different from the processes already acting between the sea water and the particles suspended in it. The processes which comprise halmyrolysis can possibly be considered to represent the even transition between the processes of deposition and those of diagenesis. STRAKHOV (1953, p.12; 1960) divided the history of sedimentary rock into the following three different and distinct stages: (I) Sedimentogenesis, namely formation of sediment. (2) Diagenesis, or transformation of sediment into sedimentary rock. (3) Catagenesis, which is a long stage of secondary changes in already formed sedimentary rock. In Strakhov’s classification, which is based on recognition of different stages of transformation of sediments into hard rocks, diagenesis is considered as only one of the stages in this development. Thus, one can say that the concept of diagenesis is used by Strakhov in a “restricted sense”. Like many other authors, WILLIAMS et al. (1955), on the other hand, are using the concept of diagenesisin a more “broad sense”. They include the catagenesis stage of Strakhov in diagenesis under the term of “late diagenesis”, which represents a transition to metamorphism. The same point of view is held by SUJKOWSKI (1958), who stated: “Diagenesis, after all, is but the introduction to metamorphism”. Some authors go still further: for example, while speaking of the definition of the metamorphic zeolite facies, COOMBS (1960) stated that this facies includes the products of not only conventionalmetamorphism but also those of hydrothermal activity and diagenesis. An attempt at drawing a borderline between diagenesis and metamoiphism is found in the article by FYFE et al. (1958): “Diagenesis of sandstones results in minor changes in the clay matrix and crystallization of cement minerals in the hitherto open pores. When the coarse clastic grains are also extensively involved in reaction so that the rock becomes substantially recrystallized, the process is classed as metamorphic”. Thus, it is impossible in a pressuretemperature diagram to fix a boundary of universal validity between diagenesis and metamorphism, because some types of sediments are less stable than others and, consequently, during deep burial will cross the boundary faster than the more stable sediments. The editors of this book also consider the concept of epigenesis of utmost importance, especially in the studies of carbonate rocks. Epigenesis includes all processes at low temperatures and pressures that affect sedimentary rocks after diagenesis (used in a “restricted sense”) and up to metamorphism. On the other

4

G. LARSEN AND G. V. CHILINGAR

hand, diagenesis includes all physical, biochemical and physicochemical processes which modify sediments between deposition and lithification or cementation at low pressures and temperatures. It should be remembered, however, that under unusual conditions, diagenesis may grade directly into metamorphism. It is important to mention here that the processes of diagenesis in subaerial environment and in shallow stable seas differ from those in subsiding basins, and were termed “exodiagenesis” by S m s o v (1960). He also pointed out that in addition to dehydration, the coagulation of colloids, rapid growth of crystals (recrystallization), formation of concretions, and preservation of textural properties of sediment (such as fissures and borings) are characteristic features of “exodiagenesis”. In the above brief discussion, the editors attempted to throw some light upon the uncertainties that exist in defining the term diagenesis. It was also emphasized that some authors use the term “diagenesis” in a more restricted sense than others. Consequently, the editors of the present book have found it necessary to leave the defhition of diagenesis to individual contributors.

SOME FBATURES OF DIAGENESIS

The processes of diagenesis, even in the above mentioned “restricted sense” of STRAKHOV (1953, 1960), go on in the marine sediments for hundreds of thousands of years after their deposition (ZAYTSBVA, 1954;cf. CHILINGAR, 1958). The processes of diagenesis include: (I) Formation of new minerals. (2) Redistribution and recrystallization of substances in sediments. (3) Lithiflcation. The diagenetic minerals include sulphides and carbonates of Fez+ and Mn2+, namely siderite, ankerite, rhodochrosite, oligonite, pyrite, marcasite, alabandine, etc. STRAKHOV et al. (1954, p.577) pointed out that their diagenetic origin is evidencedby the fact that Mn and Fe in these minerals are in a bivalent form, which could have formed in a reducing environment that existed in the sediments. Typical diagenetic minerals also include silicates of iron (and manganese), in which iron is present in a bivalent form; all leptochloritesbelong in this category. Dolomite and ankerite also form during diagenesis. Many clay minerals are of diagenetic origin, such as montmorillonite (from volcanic ash), beidellite, mountain leather, and series of zeolites (mordenite, chabazite, phillipsite, and others). Ut e, apparently, can also form during diagenesis. The zonation of new mineral formations in recent seas and in Cenozoic and Mesozoic basins is presented in Fig. 1. The upper 1-2 m of recent sediments usually reveal two early stages of diagenesis. The earliest stage occurs when the upper layer of sediment is situated in the oxidizing or neutral environment. In basins with normal oxygen regime the thick-

5

INTRODUCTION I

I

I

4

'

C

8

9 10

17 12

13 14

-

75

16

17

Fig.1 Zonation of new mineral formations in recent and ancient seas. (After STRAKHOV, 1954, p.585.) A=turbid zone of fine-grained material, and its carrying out from nearshore zone into more central parts of the basin; B=areas of currents usually of circulatory type; C-D =surface zone of agitation and wind currents in central parts of basins; E=deep, quiet (with very slight movements of water) horizons of pelagic part of basins. I =sands; ,?=siltstones; 3=pelites; 4=CaCO3, oolites; 5=biogenous and chemically precipitated CaC03; 6=diagenetic CaC03 (bacterial); 7=various forms of diagenetic dolomite; 8=Fe203, oxides of Mn, &O3; 9=leptochlorites; IO=glauconite; II=carbonates of Fe and Mn (in muds without CaC03 or having very small CaC03 content); 12=sulfides of Fe, Mn (Cu, etc.) in muds with high CaC03 content; 13=biogenically formed SiOz; 14=primary and diagenetic phosphorites; 15=minerals forming through direct precipitation from water; 16=diagenetic minerals; 17=partly primary, partly diagenetic minerals.

ness of this layer is around 10-15 cm and can reach 40 cm and higher (STRAKHOV et al., 1954, p.578). In basins deficient in oxygen this layer is only a few centimeters (and sometimes a few millimeters) thick, or is completely absent in the central portions of such basins. The duration of this stage also varies from thousands of years to several days. Some iron-manganese concretions and crusts, glauconite grains, phosphorites, and some zeolites form during this stage. The second stage of early diagenesis is observed in lower parts of cores (up to 10m)obtained in recent sediments, and is characterized by the reduction of sulphates, oxides of iron, manganese, etc., and formation of new minerals containing lower (-ous) form of these elements. The variation in the ratio of FeZ+/Fe3+with different rHz and Eh values has been studied in details by ROMM (1950).

6

G. LARSEN A N D G. V. CHILINGAR 0 10

8

20

5 30

'c

C

x

.- 40 ii 5 0 n s 60 c 70 5 7 9 1 1 7.1 7.3 z5 7.7

0 2 4 6 0 4 10 16 22

0 0.20.40.6

4 6 8

Fig.2. Variation in composition and properties of the interstitial solutions of the southern Caspian Sea with depth below the depositional interface. (After BRUEWCHand VINOGRADOVA, 1947.) I=station No.16, 100 m depth; 2=station No.26,960 m depth; 3= station No.28,460 m depth.

Only the first two stages of diagenesis can be studied in detail from cores obtained in Recent sediments; and diagrams such as those prepared by BRUEVICH and VINOGRADOVA (1947) are of great help in studying the diagenesis (Fig.2). The third stage is apparently characterized by almost complete termination of bacterial activity (and their ferments), due to accumulation of poisonous compounds and, in some cases, disappearance of organic matter. During the third stage of I

z!

m

FERMENTS REOlSTRlBUTlON OF MATERIAL IN SEDIMENTS WITH FORMATION OF CEMENT AND CONCRETIONS

OEIIYDRATION OF HYDROUS MINERALS AN0 RECRYSTALLIZATION

Fig.3. Diagenetic stages in sediments. (After STRAKHOV et al., 1954, p.596.)

INTRODUCTION

7

diagenesis there is redistribution of newly formed minerals, formation of concretions and local cementation, and recrystallization of previously formed minerals. The fourth stage of diagenesis involves transformation of plastic sediment into rigid compact rock (lithification). The squeezing out of interstitial water occurs to a depth of about 300 m (STRAKHOV et al., 1954, p.595). Dehydration of minerals occurs during compaction. For example, gypsum changes into anhydrite, hydrogoethite into hydrohematite and hematite, opal into quartz, etc. Finegrained clays are recrystallized into coarser grains and aggregates, etc. during this stage. (Different stages of diagenesis are presented in Fig.3.) One of the earliest diagenetic processes involves consumption of free oxygen by organisms, after which the reduction of hydroxides of Fe3+, Mn4+, V, Cr, etc. and. sulphates (S042-) begins. The environment changes from oxidizing to reducing and Eh becomes lower, whereas pH after some initial lowering usually increases (STRAKHOV, 1960, p.79). The solid phases present in the sediment (such as SiOz, CaC03, MgC03, SrC03, etc.) gradually dissolve in interstitial waters. Base exchange occurs between the cations adsorbed on clays and those in the interstitial water. At the same time the organic matter decomposes forming gases (C02, HzS, Hz, Nz, NH3, etc.) and water-soluble compounds, and some complex compounds which remain as solids in the sediment. As a result of these processes the interstitial water becomes devoid of sulphates, and enriched in Fez+, Mnz+, SiOz, organic matter, phosphorous, and minor elements. The 0 2 disappears and H2S, CH4, C02, NH3, H2, etc. accumulate instead. The alkalinity becomes high, Eh sharply decreases (-150 to -330) and pH varies from 6.8-8.5 (STRAKHOV, 1960,p.80). A pronounced exchange of substances occurs between the bottom waters and interstitial waters at this stage. That is why such components as S and Mg are found in higher concentrations in interstitial waters than in originally buried sea water. The eventual saturation of interstitial waters with some components leads to precipitation of diagenetic minerals such as: leptochlorites; siderite; rhodochrosite; sulfides of iron, lead, zinc, etc. The variation in Eh, pH and concentration of various ions in different areas of sediments results in subsequent redistribution (lenses, concretions, etc.) of authigenic minerals. According to STRAKHOV (1960, p.81), the depth at which diagenesis ceases varies from 10-50 m to 200-300 m. It obviously depends on degree of compaction and closing off of supercapillary, capillary and subcapillary pores. Consequently the diagenetic processes practically cease on reaching a certain degree of lithification. It is important to note here that lithification of carbonate rocks is accomplished more rapidly than that of other sediments. Compaction, which is discussed by various authors in this book, is very important in the case of clayey sediments. N.B. Vassoevich (1960, in: KLUBOVA, 1965, p.64) recognized four distinct stages of compaction occurring ( I ) with ease, (2) with difficulty, (3) with considerable difficulty, and (4) with great difficulty. For clays, the porosity varies from 60 to 85% during the first stage; and from 35

G. LARSEN AND G. V. CHILINGAR

70 60 50 40

30 20 10

0

200

400

600

800

1000

30 20

0 10 0 1000 2000 3000 z 4000 5000 4

PRESSURE,

KGICM'

Fig.4. Relationship between moisture content (%) of clays and overburden pressure (kg/cm2). (After V.D. Lomtadze, 1953, in: KLUBOVA, 1965, p.56.)

to 45% during the latter stages. The porosity depends not only on the overburden pressure, but also on the mineralogical composition of clays, chemistry of interstitial solutions, etc. Relationship between the remaining moisture content and overburden pressure is presented in Fig.4. According to RUKHIN(1960, p.307), the average porosity of clayey sediments toward the end of diagenesis (and beginning of epigenesis) is around 40%. This figure, however, can be considerably modified depending on the relative proportions of different clays present (CHILINGAR and KNIGHT,1960; CHILINGAR et al.. 1963). Possibly, at a depth of 400 m the sediments can be considered as fully converted to rocks; epigenetic stage sets in at greater depths. Future research work on compaction of sediments will shed more light on this subject. It is indeed unfortunate that studies on compaction of various sediments have been very few. The sedimentary rocks most susceptible to epigenesis and metamorphism are coals. Brown coals are usually encountered at a depth of 1,000 to 1,200 m. They are associated with plastic clays having specific gravity of 1.4 to 1.9. The coking coals of the Donetz Basin were formed at a depth of 5,000 m, temperature of 160"C and pressure of 1,100 atm (N.F. Balukhovskiy, 1952, in: RUKHIN,1961, p.305). Anthracites, on the other hand, form at a depth of 8,000 m and a temperature of 240" C. Here again, there is no agreement among the geologists as to the boundary between epigenesis (or very late diagenesis) and metamorphism.

9

INTRODUCTION SOME FINDINGS ON DIAGENESIS IN SEDIMENTS

Inasmuch as “the present is the key to the past” statement applies very well to the study of diagenesis, the editors outlined the findings of some scientists on diagenesis in recent sediments. For example, some of the important findings of EMERY and RITTENBERG (1952) on studying the diagenesis of Recent sediments off the Southern California coast can be summarized as follows: ( I ) The water content markedly decreases with depth of burial and is greatest in the finest sediments (Fig.5). (2) Upward discharge of water as a result of compaction transfers some properties to successively younger sediment layers. The compaction of sediments is not limited by the permeability, which is sufficiently great to carry away all the water displaced by grain deformation or repacking. Instead, compaction is limited by the resistance of grains to deformation and repacking. (3) The pH of the sediment surface is slightly higher than that of the bottom water, and generally increases with depth. The zone of lowest pH occurs at a depth of maximum bacterial activity and where sulphates, nitrates and COZare formed by oxidation. Bacterial reduction of sulphates and possibly base exchange give rise to higher pH at depth. ( 4 ) The Eh (oxidation-reduction potential) is generally positive at the sediment surface (oxidizing conditions) and is negative at depth (reducing conditions). The lowering of Eh is due to the withdrawal of dissolved oxygen from the interstitial solutions in upper layers, and to the action of sulphate-reducing bacteria 0 10

20

30 40

50

60 70 90 100

20

40 60 Moisture

80

100 20 I0

40

60

Moisture

80 %

110 20

40 60 Moisture %

Fig.5. Change in moisture content of Caspian Sea sediments with depth below the depositional interface (in cm). (After BRUEVICH and VINOGRADOVA, 1947.) I =station Gx 92, Tyub-Karaganskiy Bay; 2=Station Gx 83, Tyub-Karaganskiy Bay; 3 =stationNo.57, northern Caspian Sea; I=station No.1, central Caspian Sea; 5=station No.26, southern Caspian Sea; 6=station No.28, southern Caspian Sea.

I VARIATION IN COMPOSITION (ALKALINITY AND NITROGEN CONTENT) OF THE INTERSTITIALSOLUTIONS IN SEDIMENTS OF BERING SEA WITH DEPTH

ZAYTSEVA, 1954, p.2W) below depositional (m)

0

I

2

Water depth

Alkalinity (mg-equiv./l)

(4

Total N in upper horizon ( %)

0-150 0-150 0-150 0-150

0.038 0.037 0.14 0.29

2.75 10.9 2.5 41.7

12.8 26.9 8.6 66.2

1,000

0.17

2.9

2.000

-

3,000 3,000 3,000 3,000 3,000 3,000 3,000

0.092 0.084

-

0.086 0.088

-

4

8

16

0

I

2

-

17.5 72.7

3.0

(2.6)' 3.0 (3.5) (3.8) (6.6) 2.6 2.6 (1.73)

(73.7)

3.6 11.4 1.2 45.0

25.4 7.6 111.0

6.3

(21.9)

0.92

4.2

6.0

8.6

5.2 3.8 3.8 7.5 5.6 3.3 10.52

6.5 5.7 5.9 8.8 8.2 4.3

(11.0) 13.5 16.8 9.4 (15.9) 6.5

-

8

-

-

Ammonium nitrogen (mgll)

-

-

4

-

-

-

-

29.2 24.4 12.2

38.5 (32.0)

-

-

-

-

-

-

17.0 159.0

(145.0)

-

1.3

7.5

(23.4)

-

(1.6)

2.5

4.8

7.1

-

0.68 1.7 (2.8) (4.3) 0.24 0.61 (1.6)

2.9

3.8 3.6 5.9 4.9 6.0 3.0

(11.61 10.5 12.4 5.4 (15.0) 5.0

-

1.8

3.2 4.5 5.3 1.9 4.1

-

-

21.5 16.5 (6.4) -

-

(18.6) -

-

values in parentheses indicate that samples were not obtained from the exact localities described. Station No. 553 is located in a bay with productivity.

I1 VARIATIONIN COMPOSITION

(PHOSPHORUS AND SILICA CONTENTS) OF THE INTERSTITIAL SOLUTIONS IN SEDIMENTSOF BERING SEA WITH DEPTH

ZAYTSEVA, 1954, p.290)

belaw surface bottom (m)

0

Total N

1

2

4

8

I6

0

1

2

4

8

-

Silica (mg/l)'

Phosphate P (mgll)]

Water depth (m)

horizon

0-1 50 0-1 50 0-1 50 0-150

0.038 0.037 0.14 0.29

1.6 6.3 0.35 2.5

4.0 3.2 1.8 4.5

-

1.7 4.7

-

0.17

0.25

0.26

0.20 0.25 (0.17)Z (0.43) (0.27) 0.17 0.22 (0.65)

in upper

( %I

0.092 0.084 0.086 0.088

-

-

-

-

(7.5)

-

0.67

-

(0.21)

0.40

0.38 0.17 0.39 0.50 0.17 0.26 1.3

0.43 0.25 0.46 0.79 0.68 0.31

-

-

-

-

14.5 11.0 13.7 26.8

-

-

12.7

16.2

17.2

-

0.77

-

-

(17.7)

16.5

12.6

13.0

-

(1.2) 0.70 4.4 0.69 (1.6) 0.95

1.9 4.4 0.35 -

-

15.5

17.0 14.7 17.7 14.0 26.8 18.0 14.2

16.2 14.7 17.7 14.1 30.9 12.4

(15.2) 17.5 13.8 14.6 29.8 12.7

20.7 14.4 16.8

-

-

3.0 5.4 -

-

-

(17.7) (14.0) 15.5 13.5 (19.0)

contents of silica and phosphate phosphorus were determined by A.V. Fotiev. values in parentheses indicate that samples were not obtained from the exact localities described.

-

-

-

-

12

G. LARSEN AND G. V. CHILINGAR

on dissolved sulphates at greater depth. Sulphates are reduced to sulphides and in some cores totally disappear below about 7 ft. (5) The organic nitrogen and organic carbon both decrease with depth. There is a faster decrease for nitrogen than for carbon, with the sharpest decline in the top layers of sediment. The organic content in the fine sediments is lowest in the basins nearest shore, and highest in the basins at intermediate distance from shore. The organic content is also low in the continental slope sediments and in the basin situated farthest offshore, because the slow deposition of sediment allows oxidation before burial. (6) There is very little difference in the C/N ratio between a depth of 6,000 ft. in the shales of Los Angeles Basin, California, and a depth of 5 ft. in the sediments of offshore basins. Lack of bacterial activity at depth could account for the constancy of C/N ratio at depth. Other contributing factors include absence of oxygen at depth and the more resistant nature of the remaining organic matter. (7) The chloride concentration in the interstitial solutions remains constant with depth of burial, whereas the sulphate content decreases with depth, Ammonia content is greatest at depth in the sediment; there is nitrification and regeneration of nutrients when ammonia is flushed to the surface of the sediment. (8) The content of dissolved silica in interstitial solutions of recent sediments off the Southern California coast increases with depth of burial. ZAYTSEVA (1954), however, found that silica content does not show any systematic increase with depth of burial, and varies within relatively narrow limits for each station. (9) The diagenetic processes are influenced by the character of the overlying basin water.The latter is related to the depths of the basin floor and of the sill of each basin. The vertical distribution of biogenous elements in the interstitial solutions of the Bering Sea sediments is presented in Tables I and 11 (ZAYTSEVA, 1960). The findings of Zaytseva can be presented as follows: (1) The values of alkalinity (which is possibly indicative of the mineralization of organic carbon) and ammonium nitrogen content in the interstitial solutions increase with depth below the depositional interface; and show definite similarity in trend. (2) The quantitative increase in the content of biogenous elements and alkalinity varies markedly from station to station. In some places they increase rapidly with depth, whereas at others, slowly. (3) The interstitial solutions in sediments of shallow-water stations, as compared to deep-water stations (1,000-4,000 m), are characterized by (a) considerably higher values for alkalinity and content of biogenous elements (N, P) and (b)a higher absolute increment with depth of burial as compared to the upper horizons. ( 4 ) In the majority of cases, the phosphorus exhibits systematic increase with depth below the interface. This increase, however, is less pronounced than that for ammonium nitrogen.

13

INTRODUCTION

(5) With increasing organic matter content in the solid phase of sediments, there is a corresponding increase in the content of biogenous elements in the interstitial solutions (see Tables I and 11). ( 6 ) The process of decomposition of organic matter does not stop in the upper layers of sediments (16 m), and continues with decreasing intensity to greater depths.

TABLE 111 GEOCHEMICAL ENVIRONMENTS OF SUBAQUEOUS SEDIMENTS

(MAINLY MARINE)

1946, 1947, 1954) (After TEODOROVICH, .~

~~

~

~~

~

A-I* A-2**

A-3a** A-3b**

A-4** A-5* A-6

-

~

A . Strongly reducing zone

.~

~

Soda facies-Sodium carbonate; primary magnesite, dolomite, and calcite; Camontmorillonite; (sulphides of iron). Calcareous facies-CaCO3 and dolomite afterward (mainly replacement), montmorillonite-Microlaminated calcareous or marly deposits and calcareous clay without benthos, more or less rich in organic matter. Halogenous facies-Gypsum, anhydrite, rock salt, and other easily dissolved salts; sandstone and siltstone with syngenetic gypsum or anhydrite cement, etc. Mgmontmorillonite, FeSz. Replacement dolomite facies with FeSz-Microlaminated dolomites, dolomitic limestones, mark and clays without benthos, more or less with abundant organic matter. Mg-montmorillonite. (FeSz, sometimes chalcopyrite, chalcocite, rarely alabandine.) Pyritic-bauxitic-siliceous facies-Mg-montmorillonite, beidellite, leverrierite, halloysite, ferrihalloysite. Silicidesl without benthos, pyritic bauxites and clays; (FeSz, chalcocite, chalcopyrite). Pyritic-bauxitic-siliceous facies-Halloysite, allophane, kaolinite. Silicides without benthos, pyritic bauxites and clays; (FeS2, chalcocite, chalcopyrite). Kaolinite facies-Dark pyritic kaolinite clays; (chalcocite). ~

B. Reducing zone

B-I* B-2**

B-3a** 8-36**

B-4**

1

_

_ .~

_

~

~

Same as A-1 (carbonates and sulphides of Fe). Calcareous facies-CaCOs and later dolomite (chiefly replacement), montmorilonite-dolomitic limestones and limestones, mark, and clays with ankerite-siderite and FeSz; more or less considerable amount of organic matter (disseminated ankerite-siderite and FeSz are characteristic). Same as A-3a. Sulphide-siderite facies (siderite facies with considerable FeSa)-siderite- or ankeritebearing dolomites, mark, and clays with FeSz, chamoisite-siderite and ankeritesiderite ores with FeSz (Mg-montmorillonite; disseminated siderite, ankerite, and FeSz are characteristic). Sulphidezsiderite facies (siderite facies with considerable FeS2)-siderite, and

Siliceous sediments (chert, flint, novaculite, etc.).

_

14

TABLE XI1

G . LARSEN AND G . V. CHILINGAR (continued)

B. Reducing zone (continued)

B-5* B-6

chamoisite-siderite ores with FeSz; chamoisite or sideritebearing clays and silicides without CaCOa butwith FeSz; siderite-bauxite ores with FeSz (Mg-montmorillonite; disseminated siderite, ankerite, and FeSa are characteristic). Same as A-5.

Same as A-6 (sulphides).

C. Weukly reducing zone C-I* c-2**

C-30** C-36**

C-4**

C-5. C-6

Same as B-I (carbonates of Fe and only partly sulphides of Fe). Calcareous facies--CaCOa and subsequently dolomites (mainly replacement), montmorillonite-limestones, dolomitic limestones, marls and clays with ankerite, siderite, or rhodochrosite, interlayered with benthos fauna; sometimes manganous calcite, mangano-calcite or oligonite (Fe carbonates and only partly sulphides of Fe). Same as 8-30 (carbonates of Fe and only partly suiphides of Fe). 'Siderite facies (Mg-montmoriIlonite)--Siderite or ankerite-bearing dolomites, dolomitic mark and clays, interlayered with benthos; chamoisite; ankerite and ankerite-siderite ores;some phosphatic dolomite rocks,rhodochrosite and dolomite rhodochrosite ores; (siderite, ankerite, rhodochrosite, and only in part FeSz). Siderite facies (Mg-montmorillonite)Siderite and chamoisite-siderite ores; siltstones with syngenetic siderite or considerable siderite and chamoisite cement; sideritebearing clays and silicides without CaCOa; siderite-bawrite rocks; opalrhodochrosite ores; phosphatic-siliceous rocks; (siderite, ankerite, rhcdochrosite, and only in part FeSz). Halloysitic-bauxitesiliceous facies-Allophanc-halloysite, kaolinite-halloysite. and halloysite clays and bauxites; some silicides; some clays and peat with vivianite (small amounts of FeSa or chalcocite). Kaolinite facies-Dark and gray kaolinite clays with small sulphide content; peat (small sulphide content).

D. Neutrd zone D-I D-~o' D-36*

D-4+

0-5*

0-6

Same as C-l. Same as C-30. Leptochlorite facies-Leptochlorite-ilolomite rocks, some dolomites with leptochlorite, leptochlorite rocks with dolomite or with small content of siderite, or phosphatedolomite rocks; some glauconite-bearing phosphorites; (without siderite or with small amount of siderite or chamoisite; iron chlorites are characteristic). Leptochlorite facies-Leptochlorite rocks with aluminum hydroxides and bauxiteleptochlorite ores; leptochlorite formations with small content of siderite or without carbonates; sandstones and siltstones with leptochlorite and siliceous cement; chlorite-bearing silicides; some glauconite-bearing phosphorites and silicides, siliceousphosphorites; (without siderite or with small amount of siderite or chamoisite, iron chlorites are characteristic). Halloysitic-bauxitic-siliceous facies-Bauxites without leptochlorite but with halloysite, halloysitebauxite rocks; halloysitic, allophanehalloysite, and kaolinitehalloysite gray or varicolored clays; halloysite-bearing silicides; sandstones and siltstones with siliceous cement and without ferric oxides; gray and varicolored sandstones and siltstones with halloysite; clays and peat with kertschenite. Kaolinite facies-Gray and varicolored kaolinite clays; clays and peat with kertschenite; gray and varicolored sandstones and siltstones with kaolinite; (kertschenites).

INTRODUCTION

15

E. Weakly oxidizing zone E-I E-2

E-3a E-3h

E-4

E-5

E-6

Same as D-I. Calcareous faciesxertain marine limestones with noticeable content of pigmentary glauconite (or glauconite replacing the skeletal remains of organisms); numerous normal marine limestones with benthos and without glauconite, but with very small amount of FeS2, formed in deep horizons of sediment; limestones with redeposited glauconite. Same as D-3a. Glauconite facies (phosphorite-glauconite subfacies)-Glauconitic-clayey and glauconitic phosphorites with remains of sponges and other benthos interlayered; glauconite-bearing dolomites, marls, and clays; often shales and other rocks with protoglauconite; (glauconite and protoglauconite). Glauconite facies (siliceous-glauconite subfacies)-Glauconitic and clayey-glauconitic rocks; glauconitic silicides with remains of sponges (benthos); phosphateglauconite rocks; glauconitic quartz sands and sandstones or siltstones with siliceous cement and commonly spicules of sponges; “pure” or siliceous bauxites, or predominantly with leptochlorite (higher “-ic”); shales and other rocks with protoglauconite; sometimes manganite ores, interlayered with silicides; (glauconite and protoglauconite). Halloysite-bauxitic-siliceous facies with redeposited glauconite-Sandstones and siltstones with redeposited glauconite; sandstones and siltstones with redeposited glauconite and siliceous cement; halloysite, allophane-halloysite or kaolinitehalloysite clays with redeposited glauconite; silicides with redeposited glauconite and spicules of sponges or subradiolarians; some halloysite-bearing or “clean” bauxites; clays and peat with oxykertschenite; (glauconite redeposited in basin). Same as 0 - 6 (oxykertschenites).

F. Oxidizing zone F-I F-2

Same as E-I (ferric oxides). Calcareous facies-Limestones with normal marine or fresh-water benthos; numerous sandstones and siltstones with syngenetic calcite cement (and fauna); (often with admixture of ferric oxides and hydroxides). F-3a Same as E-3a (but with ferric oxides). F-3b Oxidizing facies of dolomites and silicides of replacement and oxides of manganese and iron-Numerous replacement dolomites with relic organic and oolitic texture; siliceous organic limestones; replacement silicides (usually calcareous) with relic organic texture; some dolomitic limestones and fine-grained calciferous dolomites with remains of calcareous benthos; oxide and hydroxide ores of manganese (“-ic”)-pyrolusite, psilomelane, and vernadyte; some ores of iron hydroxide. F-4 Oxidizing manganese-iron-silica facies (manganese-iron subfacies)-Ores of manganese oxides; silicides with sponge spicules. F-5 Oxidizing manganese-iron-silica facies (iron subfacies)-Manganese and copper oxide ores are absent, and silicides are without sponge spicules; lake hydrogoethite ores and “coating” of iron hydroxides on quartz grains are characteristic. F-4 and F-5 White silicides;red jasperoid, subradiolarian, etc., silicides; iron hydroxide ores; sandstones and siltstones with iron hydroxide cement; sandstones and siltstones with white or red siliceous cement; bauxite-hydrohematite and similar iron-alum i n p oxide (“-ic”) ores; quartz sandstones with iron hydroxide “coating”. F-6 Kaolinite facies-White (“clean”), quartz sandstones; reddish, orange, yellow, etc., sandstones with thin iron hydroxide “coating”; white, yellow, and reddish kaolinite clays; (small admixture of iron hydroxides).

16

G. LARSEN AND G. V. CHILINGAR

The knowledge of physical-chemical environment which exists in the sediments upon deposition is of utmost importance in studying diagenesis. TEODOROVICH (1954) subdivided the geochemical environments of subaqueous sediments (mainly marine) on the basis of oxidation-reduction potential and pH (Table 111). The types of oxidation-reduction potential include the following: ( A ) Strongly reducing or HzS zone (sulphide zone). ( B ) Reducing (carbonates and sulphides of iron zone): the oxidation-reduction potential dividing line on the average is slightly above the surface of sediment. (C) Weakly reducing (siderite and vivianite zone): the oxidation-reduction dividing line commonly coincides with the surface of sediment. (D)Neutral (leptochlorite and kertschenite zone; iron chlorites are both of lower “-ous” and higher “4c” forms): the oxidation-reduction demarcation line is slightly lower than the surface of the sediment. ( E ) Weakly oxidizing (glauconite and oxykertschenite zone): the oxidationreduction dividing line is markedly below the surface of sediment. ( F ) Oxidizing (ferric oxides and hydroxides zone). The following ranges of pH were also selected by Teodorovich: ( I ) Strongly alkaline (soda and similar lakes): pH > 9.0 ( 2 ) Alkaline: pH = 9.0-8.0 (7.8) (3) Weakly alkaline: pH = 8.0 (7.8)-7.2. (a) Saline lagoons; (b) Seas, lakes, etc. ( 4 ) Neutral: pH = 7.2-6.6 (5) Slightly acid: pH = 6.6-5.5 (5.0) (6) Acid (swamps, and some lakes and rivers associated with swamps): pH = 5.5 (5.0)-2.1. In Table I11 the geochemical facies which are most favorable for the formation of bitumens are marked by two asterisks, whereas those peripheral to the petroleum source rocks are marked by one asterisk. The minerals enclosed in parentheses are characteristic of the particular oxidation-reduction conditions. Letters A , B, C, D, E, and F designate Eh, whereas numbers I , 2, 3 , 4, 5, and 6 indicate pH as shown in preceding tabulation, REFERENCES

AMSTUTZ, G. C. and BUBENICEK, J., 1967. Diagenesis of sedimentary mineral deposits. In: G. LARSENand G. V. CHILINGAR (Editors), Diagenesis in Sediments. Elsevier, Amsterdam,

pp.417475. BRUEVICH, S. V. and VINOGRADOVA, E. G., 1947. Chemical composition of interstitial solutions of Caspian Sea. Gidrokhim. Muieviuly, 13 (1, 2). G. V., 1955. Review of Soviet literature on petroleum source rocks. Bull. Am. Assoc. CHLINGAR, Petrol. Geologists, 39(5) : 764-768. G. V., 1958. Some data on diagenesis obtained from Soviet literature: a summary. CHILINGAR, Geochim. Cosmochim. Actu, 13: 213-217.

INTRODUCTION

17

CHILINGAR, G. V. and KNIGHT,L., 1960. Relationship between pressure and moisture content of kaolinite, illite, and montmorillonite clays. Bull. Am. Assoc. Pefrol. Geologists, 44(1): 101-106. JR., J. O., 1963. Relationship between high CHILINGAR, G. V., RIEKE111, H. H. and ROBERTSON overburden pressures and moisture content of halloysite and dickite clays. Bull. Geol. SOC.Am., 74(8): 1041-1048. COOMBS, D. S., 1960.Lower grade mineral facies in New Zealand. Intern. Geol. Congr., 2ist, Copenhagen, 1960, Rept. Session, Norden, 13: 339-351. S. C., 1952. Early diagenesis of California Basin sediments in EMERY,K. 0. and RITTENBERG, relation to origin of oil. Bull. Am. Assoc. Petrol. Geologists, 36: 735-806. F. J. and VERHOOGEN, J., 1958. Metamorphic reactions and metamorphic FYFE,W. S., TURNER, facies. Geol. SOC.Am., Mem., 73: 259 pp. Glossary of Geology and Related Sciences, 1962. American Geological Institute, Washington, D.C., 397 pp. S. S., 1938. A study in rock weathering. J. Geol., 46: 17-58. GOLDICH, T. T., 1965. Role of Clayey Minerals in Transformation oforganic Matter and Formation KLUBOVA, of Pore Spaces of Reservoirs. Izd. Akad. Nauk S.S.S.R., Moscow, 107 pp. PRAY,L. C. and MURRAY,R. C. (Editors), 1965. Dolomifization and Limestone Diagenesis ( A Symposium)-Soc. Econ. Paleontologists, Mineralogists, Spec. Publ., 13: 180 pp. G. V. and ROBERTSON JR., J. O., 1964. High-pressure (up to 500,000 RIEKE111, H. H., CHILINGAR, p.s.i.) compaction studies on various clays. Proc. Intern. Geol. Congr., New Delhi, 1964, in press. ROMM,I. I., 1950. Geochemical characteristics of recent deposits of Taman peninsula. In: Recent Analogues of Petroliferous Facies (Symposium). Gostoptekhizdat, Moscow. RUKHIN,L. B., 1961. Principles of Lithology. Gostoptekhizdat, Leningrad, 779 pp. SHVETSOV, M. S., 1960. Toward question of diagenesis. Intern. Sedimentological Congr., 1960, Rept, Sov. Geologists, pp.153-161. N . M., 1953. Diagenesis of sediments and its significance for sedimentary ore formaSTRAKHOV, tion. lzv. Akad. Nauk S.S.S.R., Seu. Geol., 1953(5): 12-49. STRAKHOV, N. M., 1960. Principles of Theory of Lithogenesis. I . Types of lithogenesis and their distribution on the earth’s surface. Izd. Akad. Nauk S.S.S.R., Moscow, 212 pp. N. M., BRODSKAYA, N. G., KNYAZEVA, L. M., RAZZHIVINA, A. N., RATEEV, M. A., STRAKHOV, SAPOZHNIKOV, D. G. and SHISHOVA, E. S., 1954. Formation of Sediments in Recent Basins. Izd. Akad. Nauk S.S.S.R, Moscow, 791 pp. SUJKOWSKI, ZB. L., 1958. Diagenesis. Bull. Am. Assoc. Petrol. Geologists, 42: 2692-271 7. TEODOROVICH, G. I., 1946. Minerals of sedimentary formations as indicators of physicakhemical environment. In: Questions of Mineralogy, Petrography, and Geochemistry. Dedicated to Memory of A. E. Fersman. Izd. Akad. Nauk S.S.S.R., Moscow. G. I., 1947. Sedimentary geochemical facies: Byuf. Mosk. Ubshchest va lspytatelei TEODOROVICH, Prirody, Otd. Geol., 22(1). TEODOROVICH, G. I., 1954. Toward question of studying oil-producing formations (source rocks): Byul. Mosk. Obshchestva Ispytatelei Prirody, Otd. Geol., 29(3): 59-66. TRASK, P. D., 1951. Dynamics of sedimentation. In: P. D. Trask (Editor), Applied Sedimentafion. Wiley, New York, N.Y., pp.340. VON GUEMBEL, C. W., 1868. Geognostische Beschreibung des ostbayerischen Grenzgebirges, 1-111: 700 pp. WALHSTROM, E. E., 1948. Pre-Fountain and Recent weathering of Flagstaff Mountain near Boulder, Colorado. Bull. Geol. Soc. Am., 59: 1173-1189. H., TURNER, F. J. and GILBERT, C. M., 1955. Petrography. Freeman, San Francisco, WILLIAMS, Calif., 406 pp. E. D., 1954. Vertical distribution of biogenous elements in interstitial solutions of ZAYTSEVA, Bering Sea. Dokl. Akad. Nauk S.S.S.R., 99(2): 289-291.

Chapter 2 PHASES OF DIAGENESIS AND AUTHIGENESIS RHODES W. FAIRBRIDGE

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

SUMMARY

A review of marine sedimentological data leads to a classification of three phases of diagenesis; these are: (a) syndiagenesis (marked by syngenetic authigenesis in two stages, initial or oxidizing and early burial or reducing), that lasts from 1,000 up to about 100,000 years and may extend to depths from about 1 to 100 m; (b) anadiagenesis (marked by hypogene authigenesis, i.e., non-magmatic ascending waters and “natural chromatography”), extending from 103 to 108 years, and 1to 10,000 m depth; (c) epidiagenesis (marked by deep meteoric waters and epigene authigenesis) that may by-pass anadiagenesis, due to tectonism, and may extend from 103 to 109 years, and in depth about 1 to 3,000 m. Many authigenic minerals formed during different stages of diagenesis may be experimentally duplicated, but much remains to be done.

INTRODUCTION

Diagenesis is still a rather poorly understood field of knowledge and, although the term “diagenesis” has been in the technical language for nearly 80 years, it is not even listed in the general index of Encyclopaedia Britannica. The term was introduced by VON GUEMBEL (1868; see also AMSTUTZ and BUBENICEK, 1967) and received “text-book recognition” by the great WALTHER (1894) in his Lithogenesis der Gegenwart. He defined it (p.693) as “. . .all those physical and chemical changes which a rock (i.e., a sediment) undergoes after its deposition, without the introduction of rock pressure or igneous heat”. It is clearly understood that “rock’y was used in the traditional geological sense, which in this case implies an initially soft, unconsolidated sediment. Walther recognized that these soft sediments did not become hard “lithified” rock merely by the action of time. One rather tends to assume that if a rock is old, it is necessarily hard. Therefore, on encountering a friable sandstone or shale in some unmetamorphosed “Proterozoic-type” Precambrian, a geologist may exclaim: “Fantastic: completely unaltered sediments a billion years old!” Admittedly this is unusual, because the older a formation is, the greater are its chances for alter-

20

R. W. FAIRBRIDGE

ation. Such alteration involves the specific chemical and physical processes of diagenesis. Metamorphism is something extra, involving substantial heat and pressure. During the last century and amongst certain scientists even during the present century, metamorphism was taken to include any alteration of any sort to the original material. Thus VANHISE(1904) distinguished between katamorphism, involving near-surface alteration, weathering and the change from complex minerals to simple ones, and anamorphism, the building of complex minerals taking place at depth and under great temperatures and pressures. GRABAU (1913, p.750) called diagenesis “static metamorphism”. Nowadays it is customary to accept only thermal and dynamic metamorphism, and to exclude those reactions taking place in the upper part of the earth’s crust under the influence of the atmosphere and of connate and meteoric waters, that is to say, diagenesis. One should note, however, that diagenesis is not synonymous with katamorphism, which includes weathering. FERSMAN (1922) used the term syngenesis for sediments formed in situ, and katagenesis for any changes to them after burial by even a thin (but distinct) covering layer. The primary process of diagenesis is lith$cation, i.e., “that complex of processes that converts a newly deposited sediment into an indurated rock” (PETTIJOHN, 1957, p.648). (1913, p.75 1) includes the following Lithification, according to GRABAU (with comments by the writer): ( I ) Congelation (e.g. the physical dehydration and hardening of silica or organic gels to form an amorphous or cryptocrystalline solid). (2) Crystailization (the primary reorganization of compounds, as in unstable pyroclastic sediments, but without participation of the interstitial water or of other ions and minerals). (3) Recrystallization (a secondary crystallographic reorganization of the minerals, under increasing stress or other influence, e.g., the inversion of aragonite to calcite). ( 4 ) Compaction, welding and pressure cohesion (as a result of progressive loading, interstitial fluids being squeezed out and grains brought into contact, sometimes leading to local contact solution and redeposition in the voids). (5) Cementation (filling or partial filling of voids by cements, mainly CaC03, Si02, Fez03, derived from circulating waters. Additional factors in diagenesis have been noted by ANDRI~E (191 1). They include: (a) Formation of concretions (both in the pre- or post-lithification phases); and (b) Desalinification (a postlithification phenomenon involving the leaching and sluicing out of connate waters by vadose waters, i.e., circulation of meteoric origin, ground water and artesian circulation. Hydrothermal or pneumatolytic circulation is certainly excluded). Special geochemical factors in diagenetic mineralization, recognized partly by GRABAU (1913, p.750) and others, include:

PHASES OF DIAGENESIS AND AUTHIGENFSIS

21

( I ) Low-temperature metasomatism which embraces mineral replacement (e.g., limestone by silica and vice versa; also dolomitization). (2) Hydration and dehydration (e.g., the transformation of gypsum CaS04 * 2Hz0 to anhydrite CaS04, which is thought to occur under a load of around 100 m of sediments). (3) Ion exchange (typical of the clay and mica families of mineral). ( 4 ) Polymerization and depolymerization (e.g., natural catalytic “cracking” and other organic chemical reactions characteristic of the hydrocarbons). According to ANDRBE(191l), GRABAU (1 913), PETTIJOHN (1957), and others, diagenesis belongs essentially in a geochemical classification, whereas lithificarion includes mechanical factors. Diagenesis is restricted to sediments and sedimentary rocks, but the concept of lithification may refer also to igneous rocks. Diagenesis embraces both post-depositional mechanical modification and geochemical reorganization. It has widespread application; thus, glaciologists speak of the diagenesis of snow, in the sense of compaction, secondary cementation and recrystallization (ANDERSON and BENSON,1963), while organic geochemists speak of the diagenesis of the products of organic metabolism leading to the formation of petroleum hydrocarbons (BREGER, 1960). One may note that the physical state of the latter may be solid, liquid or gaseous. One may turn now to the concept of authigenesis or “neoformation” (also “neogenesis”) as it is known in France (see MILLOT,1953). The term was originally established by KALKOWSKY (1880) to describe the origin of any newly formed or secondary mineral. It is used now only with respect to new minerals in sediments and sedimentary rocks formed in situ. TESTERand ATWATER (1934) emphasized that such minerals must be regarded as discrete crystallographic units, rather than rockforming components. There has thus been some tendency to restrict “authigenesis” to refer to the generation of “exotic” minerals, other than those forming the bulk of the rock (e.g., the clays), but this narrow interpretation is in no way implied by its original definition. A comprehensive listing of such minerals and discussion has been provided by TEODOROVICH (1961). In diagenesis the formation of a new mineral derives from the product of a reaction between the ions of the interstitial water and the primary particles; for example, the simple addition of some new and different ions (adsorption), by the exchange of ions, or the replacement of certain ions by some new ions (metasomatism), but not from the secondary overgrowth of some new ions onto an existing mineral of the same composition. Some of these new minerals are so characteristic of primary, magmatic or high-temperature metamorphic phases that their “exotic” appearance rarely fails to cause surprise. PUSTOWALOFF (1955) has drawn attention especially to zoisite, clinozoisite, epidote and sphene. WETZEL(1955) has noted cinnabar (HgS) and barium minerals, but the question of the metal sulfides and their organic relationships is still open (see discussion on Oxidation and reduction, p.43).

22

R. W. FAIRBRIDGE

PETTIJOHN (1957, p.650) distinguished diagenetic metasomatism from authigenesis; diagenetic metasomatism seems to be part and parcel of his “diagenetic differentiation” (which involves the redistribution of materials within a sediment, such as the formation of nodules and concretions). In such event an existing mineral type, e.g., calcite, may assume a new form or position (as distinct from a new type of mineral), but the solutions can well be supplied from without, i.e., from the motion of connate waters or even vadose waters. KRUMBEIN (1947, p. 170) has noted that at least thirty different diagenetic processes are known, but that of these only about a half-a-dozen are important. One may simplify the problem, by reduction to three categories: ( I ) Mainly physical reactions, such as those leading to compaction and recrystallization. (2) Solution and precipitation phenomena, with simple cementation, decementation and intrastratal solution, including, for example, cone-in-cone, and overgrowths (addition of like ions). (3) Authigenic reactions, interpreted as all those reactions leading to new mineral formation, and incorporating: (a) metasomatism (ionic replacement); (b) ion exchange and adsorption (addition of new, and exchange of different ions, especially base exchange); (c) replacement (complete molecular substitution); ( d ) hydration and dehydration (addition to or release of HzO from the molecule or from solid solution); (e) oxidation and reduction (addition or release of 0 2 or hydroxyl ions); Cf)polymerization and depolymerization (construction and breakup of hydrocarbon chains). Inasmuch as any fortuitous mixture of minerals and ionic solutions such as exists in the fresh sediment is hardly likely to be in chemical equilibrium, a train of events is set in motion to establish such an equilibrium, at rates and in directions that are controlled by the environment. The phenomena of authigenic reactions are discussed after considering the physicochemical boundary limits of diagenesis, especially relative to the marine realm. Nonmarine diagenesis has played only a very minor role in the sum of geological history.

BOUNDARY LIMITATIONS

There are limiting factors for diagenesis, but there are passage zones into metamorphism that vary according to the primary composition of the sediment. Thus the transition from limestone to marble takes place generally at lower temperatures and pressures than that from sandstone to quartzite. The limits of such passage zones may be defined basically in terms of chemistry and physics.

PHASES OF DIAGENESIS AND AUTHIGENESIS

23

Geochemical parameters

In fresh, particulate sediment one deals with two components, namely, the solid, sediment particles and the enclosing liquid (always present initially in marine deposits, but sometimes absent at first from terrestrial sediments such as dune sands). As pointed out by GOLDSCHMIDT (1954), the chief controlling chemical factors in sedimentary petrogenesis are: (a) Hydrogen ion potential (pH). ( 6 ) Oxidation-reduction potential (Eh). (c) Ionic adsorption phenomena. A world-wide study by BAAS BECKINGet al. (1960) has shown that there seems to be virtually no environment found anywhere at or near the earth's surface where the pH/Eh conditions are unacceptable for some form of organic life. As a corollary, one must conclude that there is no environment near the earth's surface (other than volcanic) that is not in some way modified by organic metabolic processes. Inasmuch as COz is the principal by-product of organic oxidation and is also the principal raw material of plant (and much bacterial) photosynthesis, it is to be expected that it plays an all-pervading role. Thus COz reactivity on the earth's crust will be related to the rate of organic metabolism, and inasmuch as the latter is thermophylic within the ecologic limits of the various phyla, provided that adequate water is present, the most reactive regions of the earth's surface will be tropical. Hydrogen ion potential C02 dissolves freely in HzO, creating a bicarbonate ion, and a free hydrogen ion. The hydrogen ion concentration in pure water at 20 "C is lo-' equiv./l (pH 7), but saturated with COZit rises to 10-5 (pH 5). COZ is thus involved with carbonic acid and the bicarbonate ion in the following equilibrium condition, thus: H20

+ C02 + H2C03 + HC03- + Hf + 2Hf + CO&

(in sea water), pH 5, pH 6.3, pH 10.3. In a closed system, this reaction moves to the right as temperature increases. Thus, were it not for organic interference, the tropics would always tend to be alkaline and the polar regions acid. However, under organic control very considerable modifications may be introduced, and the pH range may extend from about 2 to 12 (BAASBECKINGet al., 1960). It is also important to consider the ionicpotentiul of the various components. The ionic potential is defined as the ratio between ionic charge Z and ionic radius r. According to WICKMAN(1944), these potentials fall into three categories: (a) 0-3, soluble cations (i.e., those that stay in true ionic solution, even up

24

R. W. FAIRBRIDGE

to a very high pH; e.g., Na+, K+, Mg2+, Fe2+, Mn2+, Ca2+, Sr2+, Ba2+); their hydroxides have ionic bonds and are therefore soluble. (b) 3-12, elements of hydrolysates (i.e., those precipitated by hydrolysis; e.g., AP+, Fe3+, Si4+, Mn4+, etc.); these have hydroxyl bonds which makes them susceptible to hydrolyzation. (c) Over 12, soluble, complex ions (i.e., those forming ‘‘complexes’’, complex anions containing oxygen; and as a rule give true ionic solutions; e.g., B3+, C4+, N5+, P5+, S6+, Mn7+). These have hydrogen bonds, which also, like group a, lead to soluble compounds. (These relationships have recently been clearly explained for geologists by BARTH,1962.) To quote from BARTH1962, p.29: “Most natural waters go through an evolution of increasing pH, until they eventually empty into the sea, which is slightly alkaline. Silica becomes more soluble with increasing pH, and is therefore often delivered into the sea. But aluminium hydroxide is precipitated in mildly acid solutions near the point of neutrality. . . The difference in behavior of ferric and ferrous iron is of interest. Ferric iron is soluble only in rather strongly acid solutions; it is therefore precipitated before aluminium, but the separation is usually not clean. Ferrous iron remains longer in solution in equilibrium with carbon dioxide in oxygen-free waters. Similarly tervalent and quadrivalent manganese ions are precipitated before bivalent manganese.” In the interplay between high and low pH in natural waters, the two principal players which rank (in total quantitative terms) far higher than all the other elements are silicon and calcium. The reasons for this will appear, on considering the solubilities of the principal elements in the earth’s crust. After oxygen, which is always in combination, there are only seven quantitatively important elements: Si4+, APf, Fez+ or 3+, Ca2+, Na+, K+, and Mg2+, in that order (see Table I). Inasmuch as natural waters are everywhere subjected to organic interference (largely reflected by the concentration of C02) one may observe in cool, humid climates with acid soils (high COBand low pH) that calcium (with Al, Fe, etc.) is mobilized, but that silica, such as comprises quartz sand, remains stable, and becomes progressively cleaner and cleaner (e.g., podzolization). In contrast, in a highly alkaline soil, characteristic of warm rather dry “Mediterranean” climates (low in their supply of HzO and COZ),the pH is high (8-9),so that calcium is precipitated, and results in the well-known lime “caliche” or “calcrete” crusts, whereas silica is mobilized and generally carried into the river system during the brief wet season-partly in colloidal form. In drier places it may simply rise by capillarity to the surface, there to replace calcrete or to be reprecipitated on desiccation as a “silcrete” crust. Thus a specific geological formation, subjected through time to different paleoclimates may be affected by an alternation of ground and artesian waters (both at the surface and at depth) from high to low pH, leading to complex intergrowths and respective replacements of quartz by calcite and vice versa.

25

PHASES OF DIAGENESIS AND AUTHIGENESIS

TABLE I COMMON ELEMENTS IN THE EARTH’S CRUST AND THEIR SOLUBILITY WITH RESPECT TO PURE

(PH 7), OR

MODIFIED BY SOLUTION OF

Coz (PH 5) OR COS ION (PH 9)

Element

Crustal abundance At p H 5 (rnolesll) (at 25°C) (partsllO00, or glkg)

Si Al FelI1 Ca Na K

277 81 50 36

Mg

;:} 21

2 . lO-3* 1.4. 10-7 6 . 10-9 very soluble

H2O

At p H 7 (molesll)

At p H 9 (molesil)

p H at which hydroxide begins to precipitate

4.5. 10-3 1.4.10-13 6 * 10-15 at pH 12

6 * 10-3 1.4.10-19** 6 . 10-21 3.2. 10-1

2 4 2.5 12

very soluble (the hydrogen will not precipitate) very soluble 1.1 * 10-1 10.5

* Approximation, based on curve by CORRENS (1949, p.210). Somewhat different according to SIEVER (1959). SiOz is in the form of the oxide, not hydroxide. ** This value may be too low. Due to the amphoteric nature of AI(OH)3, it begins to dissolve in OHAI(OH)4-. At pH 10, the strongly alkaline solutions, forming the complex AI(0H)a solubility of Al(OH)3 increases sharply. Equally rapid (in the other direction) is the increase in solubility at ca. pH 4.

+

+

Extreme swings, from one absolute pH boundary to the other are to be seen under exceptional conditions in some deserts. Normally the dilution of all solutions by rain water (pure HzO) tends to bring the pH within one unit of neutrality (pH 7), and thus minerals that may develop during diagenesis in the extreme pH ranges are exceedingly rare. Oxidation-reduction (“redox”) potentials In order to understand the boundary limits of pH it is necessary to consider also the oxidation-reduction or “redox” potential (Eh), which is to some extent 1952). reciprocal to the pH, but is influenced by certain other factors (LATIMER, The plot of the pH/Eh relationship for natural environments, as established by BAAS BECKINGet al. (1960), presents a boundary like a distorted shield, with small shoulders (or “ears”) in the low pHjhigh Eh corner and in the high pH/ moderate Eh corner (see Fig. 1). The whole fits between two parallel lines, the slope of which corresponds to -0.059 V/pH unit. These two parallels are absolute barriers (“fences”), representing the equilibrium limits of water at or near the earth’s surface. The upper diagonal bounds the upper limit (HzO/Oz) where the partial pressure of oxygen is equal to 1 atm., and the lower diagonal marks the lower limit (HzO/Hz), where = 1. In the lower diagonal the left-hand end corresponds to pH of 0.0 and Eh of 0.0. (See also Fig.2.) The value of these relationships to an understanding of diagenetic reactions should hardly need emphasizing. Yet as brought out recently by GARRELS (1960,

26

R. W. FAIRBRIDGE

1,000

Fe Sp Oxidation

800

600

400

200

Eh mV 0

-200

-400

I

-600

-\

\

I

%'\\

I

ACID I ALKALINE -800

0

pH

2

4

6

8

\,A/*

10

12

Fig.1. Catenary diagram ihstrating limits of natural environments in terms of pH and Eh especially the sites of syn-, ana-, and epidiagenesis. (Based on works by L. G. M. Baas Becking and R. M. Garrels.)

p.104), it has taken geochemists nearly half a century to recognize this fully. It is fortunate for the sedimentologist that a reliable electric (battery or lineoperated) pH-Eh meter can be purchased for a quite modest outlay, and may then be freely used not only in the laboratory but in the field, so permitting readings to be obtained on the spot in natural media. This is important, for it is not easy to obtain a sample of some gas-saturated mud, for example, and transport this to a laboratory without grossly upsetting the original Eh/pH relationships. TO the sedimentologist, the pH meter is what the field pick or hammer is to the hard rock men. In short, it is absolutely basic to sedimentology. It was through soil studies and bacteriology that pH/Eh relationships were introduced into sedimentology. ZOBELL'Sstudy (1946) brought to this writer his

27

PHASES OF DIAGENESIS AND AUTHIGENESIS

PH

-

7.0 I

8.0 I

I I

Fig.2. “Fence diagram” illustrating principal environments of sedimentation and diagenesis, and GARRELS, 1952.) according to Eh and pH. (After KRUMBEIN

first inkling of their application. Earlier work has been done in France and Belgium, and an English translation of a book by POURBAIX (1949) presented the thermodynamics of dilute aqueous solutions in terms of pH and Eh. These principles have been excellently applied to geology by GARRELS (1960). In nature, oxygen-consuming organisms are the principal agents in lowering the redox potential, but in the atmosphere or at the sea floor there is usually such constant water circulation that the lower half of BAAS BECKING’S(1959) “shield” is not involved. This state of affairs changes, however, as the sediment is buried; in clays, for example, only a few millimeters is sufficient for diagenesis to begin in earnest. The aerobic bacterial attack on buried organic debris quickly removes all free oxygen from interstitial water, and at a pH of about 7, the Eh

28

R. W. FAIRBRIDGE

is about -0.4. The anaerobic bacteria then take over, as their aerobic brethren have literally eaten themselves to death, and they attack the sulfate anion, the most readily divisible ion containing oxygen. SO42- after chloride is the most important anion in the ocean (7.68 % of the total ions). With reduction to sulfite and then to sulfide, the redox potential steadily drops and the pH shifts to higher values, so that in young sediments at a depth of 1-3 m the pH is often up to 9 or more. Any free COZhas long since passed into CaC03, so that only the gas phase is HzS. The setting is now appropriate for the pyrite reaction which is perhaps the most significant in all of diagenesis (see the section on Oxidation and reduction, p.43). The stability fields of Fez03, FesO4, and FeSz have been nicely illustrated by GARRELS (1960, p.145). Under favourable conditions pyrrhotite, galena and other metallic sulfides will start to form. BAAS BECKINGet al. (1960) have beautifully demonstrated them in laboratory-controlled bacterial studies (Fig.3). Ionic adsorption phenomena According to BARTH (1962, p.30), ionic adsorption phenomena “take place at low temperature in colloidal phases or phase complexes that are capable of capturing and binding certain ions through adsorption. One example is the binding of potassium ions by the clayey products of the hydrolysis. In a geochemical adsorption process the binding of the ion to the colloidal surface takes place in competition with the over-all hydration of the ion in the solvent. It can be shown that the degree of adsorption in an ion is a function of radius, charge, polarizability, and normal potential, as well as the nature of the chemical compounds formed at the phase boundary.” “Through the processes of adsorption the natural waters are deprived of many of the rarer elements. Most of the ions of the heavy metals, such as ions of lead, zinc, and copper, as well as complex ions of arsenic and molybdenum, are captured by, and coprecipitated with, the colloidal particles, usually hydrolyzates, and thus are supplied to the sediments.” “The amounts of poisonous metals and metalloids which potentially have been delivered into the ocean from the primary rocks throughout geologic times are so considerable that a serious poisoning of the ocean would have been caused if this process of elimination of poisonous substances had not been in action. Or the evolution of life would have taken a different course, developing organisms not susceptible to our poisonous metals. This statement applies, for instance, to copper, lead, arsenic, selenium, mercury, antimony, and bismuth. In many cases these metals have been removed from aqueous solutions by a means also known in practical medicine, that is, adsorption on freshly precipitated hydroxides of iron. There is considerable concentration of selenium, arsenic, and lead in the sedimentary iron ores. The arsenic content of these ores in most cases is so high that it brings a very notable quantity of arsenic even into iron or steel, from which

29

PHASES OF DIAGENESIS AND AUTHIGENESIS

The equilibrium pH for each dissociation is:

Q so,-

Fig.3. Stability fields of some important naturally occurring non-metallic compounds in terms of oxidation-reduction potentials and pH framed within the limits suggested by this work. Distribution of these compounds in un-ionized states is governed by their dissociation constants. (After BAASBECKING et al., 1960.)

this element is difficult to eliminate by the usual technical processes of refining. Molybdenum is concentrated in manganiferous sedimentary ore deposits." Geophysical parameters

The physical boundaries to diagenesis are defined mainly in terms of temperature, pressure and time. Temperature The mean temperature at the earth's surface through most of geological time for 1967), which there are identifiable indicators, say the last 3 lo9 years (FAIRBRIDGE, has remained at ca. 20 f 10°C. Soil temperatures today at depths of 0.1-1.0 m fall generally within this range. Seasonal, latitudinal and altitudinal variables increase this range from ca. - 100 to 6 0 T , excluding volcanic phenomena, hot springs, and so forth. Certain Algae are adapted to life in hot springs near I O O T , but this is quite exceptional. The mean temperature at the water-sediment interface over most of the deep ocean floor is ca.. 2 "C. At intermediate to shallow ocean depths, the temperature approaches the world mean, noted above. These very moderate to low temperatures greatly influence the geochemistry of diagenesis, for in such ranges crystallization is normally slow and only simple

-

30

R. W. FAIRBRIDGE

compounds form. The complex, mixed crystals commonly involved at the temperatures of formation of metamorphic and igneous rocks (over 100°C) are rarely encountered. The complex mixed lattice of the magnesian calcite-dolomite series is a notable exception to this generalization, but even this reaction is favored in nature by elevated temperatures (30-40“C). In evaporite basins a complex series of halide salts is also favored by somewhat higher temperatures (BRAITSCH,1962; BORCHERT and MUIR, 1964). In geosynclines, that is to say, sedimentary basins or troughs marked by a considerable accumulation of sediments and often extending over periods of the order of 108 years, temperatures below the surface are found to rise, by ca. 0.5 “C/ 100 m, due to the poor conduction of the earth’s internal heat. In certain regions, however, the gradient may be much steeper. There are generally two potential causes of this: (a) Abnormal concentrations of igneous or radioactive heat, as near volcanic vents and major fault lines, notably the celebrated “Mid-Ocean Rift” along which the heat flow may rise to 8 10-6 cal/cm2 sec, in contrast with an average of ca. 1 10-6 elsewhere. (b) Sedimentary accumulations of minerals, which on oxidizing are exothermic, that is to say, they generate heat. The oxygen is generally brought in by artesian water and the heat produced is dissipated by its continued circulation. This water, which may be partly connate (that is primary), tends to migrate upwards and outwards as a basin compacts. One of the principal minerals involved is either marcasite or pyrite (FeS2) and these are commonly present in finely divided particles (but vast quantities) in any shale that was formed under slightly reducing conditions, or in coal seams. The heat generated may spark off a spontaneous coal fire in mines or landslide areas. In artesian wells at Perth, Western Australia, the water from only 300 m comes from a pyritic shale-silt-sand sequence that brings the water temperature at the well head to 90 O F (32 “C).In the deep oil wells on the Texas Gulf coast at depths down to 25,000 ft. (8,000 m), the temperature through normal heat flow should be high, but may exceed 150°C mainly through this same exothermic mineral oxidation.

-

-

Pressure In a gradually accumulating sedimentary basin there is a progressive increase of load pressure (LANE,1922). This has sometimes been called load metamorphism or static metamorphism, as opposed to dynamic metamorphism, which involves tangential stresses as well as simple vertical compressive stress. For this reason, and because temperatures are relatively low, the effect of simple overburden pressure may be considered as “diagenetic”. There is a type of load metamorphism, however, that transcends normal diagenetic changes, because of its complete remobilization of ions and formation

31

PHASES OF DIAGENESIS AND AUTHIGENESIS

of minerals beyond even the limits of metasomatism. The results of this are seen in some ancient evaporites, in particular the Strassfurt deposits of the Permian in north Germany (JANECRE, 1915; RINNE,1920; BRAITSCH,1962; BORCHERT and MUIR, 1964). The question of the dominance of the role of load or of dynamic metamorphism, or of geothermal heat is not yet resolved, but certainly the temperature of the alteration was probably not over 80 “C. Time Time is of course the geologist’s trump card in any argument with physicists and chemists. Some solubilities are so extraordinarily low that they take millions of years to bring about any noticeable effect. It is, however, rather too easy to delude oneself by this line of reasoning, and it is worth bearing in mind that certain seasonal effects are highly episodic, and lead to short peaks of hydrolysis, pH-Eh oscillations, etc., which may pass unnoticed at other times. There are also brief (in geological terms) episodes of diastrophic activity, such as periods of uplift accompanied by massive fracturing, jointing, and faulting, that would favor extensive recirculation of waters which may previously have lain stagnant (or isolated by low permeability) for extended periods of time. Oil geologists are well aware of this characteristic in the history of a basin’s fluid components to be periodically subject to induced flow and interruption. To generalize, one may say that the average geosynclinal basin experiences progressive downwarping and compaction for 107-108 years, and that this is followed by one or more episodes of uplift (with fracturing and faulting), generally in brief spasms, marked by earthquakes of a few hours. Depending on geotectonic factors, and the nature of the underlying crust, the geosyncline may or may not become involved in superficial folding or in deep-seated buckling and compression. In the latter event, the sediments are placed in regions of high heat flow, and then metamorphism, granitization and igneous activity are introduced. Where only superficial folding, however, is involved, in a superficial “skin” that may not exceed 3,000 m, the principal orogenic stress is provided by gravity, and rock alteration is limited to diagenesis except in specific stress-strain zones such as faults. After uplift, extended periods are likely to pass with only episodic and very gentle (epeirogenic) revival of topographic relief. Long-continued exposure to meteoric circulation will be the rule, and some non-metamorphic Precambrian rocks have been so exposed for periods of over 1 109 years.

-

DIAGENETIC EVOLUTION

Diagenesis begins at the moment a sedimentary particle comes to rest, for example, on the sea floor; and it continues to a point in history when either deep burial

32

R. W. FAIRBRIDGE

and orogenic buckling cause the initiation of metamorphism, or when emergence leads to exposure and the initiation of weathering and erosion. It is an almost Davisian evolutionary cycle of youth, maturity, old age, except that it is complicated by rhythmic repetitions and “accidental” alternative courses, introduced by interaction with geotectonic, paleoclimatic and other cycles. One may employ with “diagenesis” the classical prefixes “syn-” (together with, ie., with the sedimentation), “ana-” (again, i.e., lithified), and “epi-” (upper, i.e., modified by surface phenomena). Thus, the three stages may be named as follows (Fig.4): (a) Syndiagenesis (the sedimentation phase). (b) Anadiagenesis (the compaction-maturation phase). (c) Epidiagenesis (the emergent-pre-erosion phase). SIEVER(1959) has called these phases early, middle and late diagenesis, but when DAPPLES (1959) speaks of initial or depositional, then early burial and finally late burial or pre-metamorphic, the meaning is not the same (see below under Syndiagenesis). DAPPLES (1 962) designated three geochemical stages, as follows: ( I ) Redoxomorphic (reactions mainly oxidation or reduction; metabolic control, most effective in the syndiagenetic phase-which includes both initial and early burial sub-phases). ( 2 ) Locomorphic (principally metasomatic, one mineral being replaced by another; important in the lithification of the anadiagenetic phase). (3) Phyllomorphic (characterized by ion exchange associated with clays and micas, to be observed at all phases).

I

RAINFALL

pH 7

I

Fig.4. Idealized profile through a continental margin, showing the sites of contemporary marine sedimentation and the three phases of diagenesis. Note the (1) diffusion potential during syndiagenesis; (2) upward liquid motion in anadiagenesis; and (3) downward motion in epidiagenesis.

PHASES OF DIAGENESIS AND AUTHIGENESIS

33

Inasmuch as it is the anadiagenetic stage that is most likely to be impinged upon by the geotectonic cycles, this episode may be so reduced in the evolution to the epidiagenetic phase that one might almost speak of a short-circuit connection when orogenesis has caused uplift of fresh unlithified sediments and led to their rapid erosion. Alternatively, in the same stage but in a different geographic position in the sedimentary basin similar sediments might be trapped in a downbuckling of the crust and initiated into a metamorphic cycle that might indefinitely postpone or eliminate the epidiagenetic phase.

Syndiagenesis Defined here as the sedimentational, pre-diastrophic phase, syndiagenesis begins at the moment the sedimentary grain touches the bottom, and is marked by the presence of large amounts of trapped interstitial, or connate water, which is expelled only very slowly. The term “syndiagenesis” was first used by BISSELL (1959). In the study of mineral deposits the analogous process is “syngenesis”. However, this term, as originally defined by FERSMAN (1922), was intended only for primary chemical sediment such as oolite, and he used “diagenesis” only as the writer uses “syndiagenesis”. Two stages are recognized in the syndiagenetic phase. These have been called (DAPPLES,1959, 1962): (a) initiaZ stage, controlled by the chemistry of the superjacent water; and (b) early burial stage, controlled by the entrapped, connate water, chemically modified by the bacteria and other subsurface organisms.

Initial stage Buried with the sediment is generally a moderate to large amount of organic matter which provides nutrients for burrowing organisms that greatly disturb the surface layers of the fresh sediment and keep them relatively well oxidized, as earthworms do in soil. Indeed WAKSMAN (1933) described this organic matter as marine humus. BADER(1954) described pelecypod population densities as essentially controlled by what he called the “decomposition coefficient” of the sediment. In basins lacking a free circulation above the sediment-water interface, stagnant, euxinic (“Black Sea”) condition will lead to poisoning of bottom waters, and metazoic benthos will be excluded, resulting in the nice preservation of original finely stratified bedding planes (CHILINGAR, 1956a; CASPERS, 1957). This is typical of the Black Sea, where the trapping of organic debris leads to diagenesis of petroleum hydrocarbons, but this is more a function of rapid accumulation than it is of euxinic conditions (SMIRNOW, 1958). The bacterial population near the sediment surface will, in the well-ventilated basins, belong to the aerobic families and some may even be photosynthetic autotrophs; others will employ the buried organic matter and the oxygen from

34

R. W. FAIRBRIDGE

the connate water. The result will be a sharp rise in thepcoz marked by a drop in pH, which from the surface may pass from 8 to 7 or 6.5 (ZOBELL,1942; DEBYSER, 1952). This zone extends for a few millimeters down to about 30-50 cm depending on factors such as depth of water, amount of organic nutrients, rate of sedimentation, etc. (TWENHOFEL, 1942). The effect on the inorganic sediments is sometimes slight: “cleaning” of quartz sands, for example; or it may be profound: rapid solution of calcite and particularly aragonite grains, destruction of carbonate shelled foraminifera and calcareous spicules, etching of more massive shells and “weathering” of some feldspars and clays. Pelecypods (and other organisms) that inhabit such acid bottoms protect themselves with a chitinous covering (periostracum), e.g., Mytilus; but immediately after death solution begins. Experiments by HECHT(1933) demonstrated that even on the surface, empty shells lost 10-20 % of their weight per year (in ordinary North Sea water). In the richly organic muddy sediments of the Wattenmeer, however, only casts and moulds of shells are normally found. Experiments of burying shells along with the rotting molluscan remains showed losses of up to 25% of the total shell weight in only two weeks. Gypsum crystals sometimes formed on the meat, illustrating the local reduction of marine S042- to HzS and its immediate reaction with the Ca2+ of sea water to form CaS04.2HzO. On tropical coasts, even on many coral reefs, the almost universal presence of mangrove swamps (populated especially by the genus Rhizophora) provides a rich source of organic debris, leaves, branches, etc., so that the pH in the muds (even at the surface) normally drops to 6.5 or less (ORRand MOORHOUSE, 1933). In addition to COZ,it is probable that humic and tannic acids are also liberated. Coral reefs of CaC03 are pocked by giant mud-filled pot-holes up to 5 m in diameter, wherever mangrove trees have been situated. Early burial stage Below the oxidizing zone is a reducing zone (DAPPLES,1959, “early burial stage”). Here anaerobic bacteria become dominant and the pH rises steadily, often to above 9 (ZOBELL,1942). The Eh drops to -0.4 or -0.6. Sulfate reducing bacteria, notably Desulfovibrio desulfuricans, liberate HzS. CaCO3 precipitates freely at a pH of 8.5 and in this Eh range FeS is the stable iron compound (afterwards becoming FeS2; see discussion, under Oxidation-reduction, p.54). Somewhat less commonly, siderite, FeC03 is formed (see stability diagram in GARRELS, 1960, p.130). During diagenesis the chemical reactions are generally governed by the first part of the Van ’t Hoff Law, which states that low temperature reactions usually generate heat (i.e., they are exothermal), and are accompanied by the association of ions. VAN HISE (1898) observed that the operation of this law was characteristic of his “upper physicochemical zone”. In the bacteria-rich reducing

PHASES OF DIAGENESIS AND AUTHIGENESIS

35

conditions of the early burial stage, however, many of the larger organic and inorganic molecules are broken down. Vegetable matter disintegrates and only the most stable parts remain; lignin, the principal residue, is extremely stable in the marine realm and may be a useful indicator for rates of sedimentation (BADER, 1956). An important aspect of the break-down of organic matter is the rapidity of the reaction in the aerobic stage; material that survives this attack and passes into the “early burial” (anaerobic) condition has a much greater chance of preservation in rocks (ABELSON, 1959, p.83), though often further modified to petroleum hydrocarbons and other organic products. After hydrogen, the most important active element in the early burial stage is sulfur. This element is present in sea water as the anion S042-, which represents 7.68 % of the total dissolved constituents of the ocean, and is the most important after Na+ and C1-. “The sediment acts as a chemically open system to the sulfate of the overlying water” (BERNER,1964). Sulfur is also an important member of many organic compounds. It shows a valence change of -2 to $6 during oxidation and reduction. It also has two stable isotopes 32S and 34S, with a 6 % mass differential which is easily measured with modern instruments. During the valence changes, for example, from S042- to S2-, the isotopes are fractionated, so that the sulfide ion is enriched in 32S, the more energetic isotope. The sulfur isotope ratio is therefore a valuable indicator of passage through the early burial stage. Thus HzS and related authigenic minerals formed then show an 32S/34S ratio of 22.1/22.7 (with an average of 22.49 for all sedimentary sulfides), in contrast to a constant 21.76h0.02 for sea water sulfates (AULT,1959); evaporate sulfate figures are similar to sea water, but have a wider spread (f0.2). The mean sulfur isotope ratio for magmatic hydrothermal and meteoritic sulfides is about 22.2, which is readily distinguishable from the mean for sedimentary rocks; but unfortunately the spread of values for the sedimentary rocks makes it difficult to use this device to solve the controversy about the metallic sulfide ore deposits. BAASBECIUNG (1960) and his associates have shown experimentally that, under certain conditions, marine bacteria can synthesize not only pyrite, but also the common ore sulfides. DEANS(1950) reported that Westoll had found fossil fish skeletons diagenetically replaced by galena, sphalerite, chalcopyrite and bornite. There is a considerable controversy, therefore, between those who would attribute all ore sulfides to magmatic sources and those who consider them syndiagenetic. Both sources are possible as demonstrated experimentally. The main problem today is to discover the relative importance of the various sources in the different deposits. While synsedimentary origins are now widely accepted, the localization of very high metal concentrations in sea water, from time to time and in rather limited areas, was probably due to the local thermal springs (“exhalative” magmatic hydrothermal sources) so that ultimately magmatic sources were responsible (DUNHAM,1952; WILLIAMS,1960).

36

R. W. FAIRBRIDGE

The element nitrogen is sometimes forgotten in geological literature, but its role (mainly through ammonia compounds) in the syndiagenetic phase is not unimportant. It can also be a helpful indicator. Whereas the total carbon content drops away sharply with depth of burial, the level of fixed ammonium remains rather constant. Thus the C/N ratio can be determined and used for environmental reconstruction (ARRHENIUS, 1950; STEVENSON, 1960). The reducing zone generally leaves a characteristic mark on the syndiagenetic phase, because it is inevitably the last environment of a sedimentational stage and thus leaves its imprint on sediments for all time. Although its products may subsequently be modified, the evidence is never totally effaced. Indeed the oxidation stage (‘‘initial stage” of DAPPLES, 1959) may be bypassed in the euxinic environments and the reducing stage occurs at and above the sediment surface. If the oxidizing zone has been present, however, the acidizing experience of the sediments is the one that will have had the more striking effects as seen in the light of day, perhaps a hundred million years later; carbonate fossils are absent and the only obvious traces of former life are the chitinous forms such as conodonts. If the sediment is a coarse-grained one, such as a quartzose sand or silt, there is generally a far greater opportunity for oxidation than in clays. Thus entrapped organic matter is totally consumed, under low pH the soluble carbonates and other minerals are destroyed or modified, and the sand is thoroughly cleaned. In this case the reduction zone is left with little nutrient for the bacteria and the populations are thus greatly limited in size and variety. In the case of carbonate sediments, that is where the great bulk of the material, regardless of grain size, is CaC03, no amount of bacterial C02 production will cause the total solution of their substrate. The response of lime sands (“proto-calcarenites”) will be rather similar to that of quartz sands, but lime muds (“proto-calcilutites”) will respond rather like clays, and the resultant limestone may be speckled with marcasite or pyrite concretions. These are quite rare in calcarenites. In the case of fine-grained siliceous material, it is suspected that much of it enters the ocean through rivers in the form of colloidal silica, where it becomes electrolytically flocculated (or adsorbed on to suspended matter, aiding further transport), and accumulates as small globules of gel (EITEL,1954; BIENet al., 1958). These deposits may be augmented by opaline silica from radiolaria and sponge spicules, and those of holothuria and alcyonaria (RIEDEL, 1959). Penecontemporaneous resolution may occur, but is prevented in rapidly accumulating Globigevina oozes. Migration of the silica often seems to occur while the sediments are still quite soft. Indeed much movement may be expected while there is free permeability and electromagnetic response is facilitated. It may be borne in mind that the low pH of the initial stage which may lead to carbonate solution will favor Si02 stabilization but vice versa under the high pH of deeper levels.

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37

Horizons of flint nodules and chert layers in chalks and limestones are often so regularly displayed that one might take them to reflect a sedimentary rhythm. On the other hand, SUJKOWSKI (1958, p.275) speaks of a diagenetic rhythm while admitting that a mild sedimentary rhythm might lead to a very inconspicuous banding of textural character. On these terms migration would be favored by slightly coarser more permeable layers; these would be predisposed to diagenetic rhythmic bedding. The effect of diagenetic rhythm probably goes further than the formation of concretions. According to SUJKOWSKI (1958): “By separating the compounds of an unstable mixture inside a sedimentary series, diagenesis exaggerated the rhythmic differences pre-existing in a deposit. It is also not excluded that in some texturally homogeneous deposits, diagenetic rhythm is quite a secondary phenomenon resulting only from the unmixing of the different chemical components to the limits of diffusion”. Thus it is evident that primary deposits may be more uniform in composition than the rocks derived from them. The thickness and duration of the syndiagenetic phase are defined by a number of variables such as lithology, organic components, rate of sedimentation, aeration and depth of water. Generalizing one may say that the base of the syndiagenetic phase is defined by the lower limit of vigorous bacterial activity, which may be from ca. 1 to 100 m. In terms of organic metabolism, syndiagenesis may be taken to last as long as the food hangs out. In terms of absolute duration this may be for 1,OOOto 100,000 years, but considerablymore research is needed on this aspect. It has been claimed that viable bacteria can be traced back to Carboniferous coal seams (nearly 3.106 years old) but possibilities of contamination are so great that it is very difficult to prove. Anadiagenesis

Anadiagenesis is proposed and defined here as the compaction and maturation phase of diagenesis, during which the clastic particulate sediment grains (or chemical ions) become once again (Greek: ana-) lithified. Diastrophism may or may not be involved: this depends upon the particular geotectonic situation of the sedimentary trough or basin. Characteristically this phase is one of slow compaction and concomitant expulsion of connate water. Rising mineralizing waters are often known as hypogene (with hydrothermal admixture), but it should be emphasized that most anadiagenetic waters are non-magmatic (WHITE,1957). DAPPLES (1959) called this the “late stage’’ of diagenesis, but apparently included with it also the epidiagenetic phase (see section on Epidiagenesis, p.40). During anadiagenesis some of the connate water becomes trapped permanently in the sediments as a result of compaction and cementation to the point of impermeability. It thus becomes “fossil sea water”, though greatly modified from its original form. The name “connate” was proposed by LANE(1909), and

38

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indeed only since the introduction of the term has the importance of this phase of diagenesis been appreciated (see LANE,1927; WHITE, 1957; and CHAW, 1960). The effect and weight of sediment-loading was also studied by LANE(1922). Earlier, VAN HISE(19M), for example, regarded all interstitial water as meteoric. Some of the economic geologists, in contrast, seemed to have regarded it as almost all magmatic (SCHMITT, 1950). Sediments that have passed through anadiagenesis are therefore characterized by cementation, the most common cements being siliceous or calcitic, and more rarely ferruginous. An important “diagenetic fabric” may be studied on polished surfaces or in thin sections. Use of the universal stage microscope permits the identification of the rarer minerals, particularly the sequence of growth (GLOVER, 1963). Etching and overgrowth phenomena are most common. In siliceous sediments this phase may not become well developed until considerable depths are reached, but with carbonates the reactions may be extremely rapid, beginning even at the surface. In orthogeosynclines, it is possible that the progressive downwarp and filling of the trough leads to burial in excess of 10,OOO m, which is the approximate depth at which the geoisothermal level exceeds the normal operative limits of diagenesis (about l0OOC). The system is affected firstly by “load metamorphism”, and secondly, by diastrophism which inevitably tends to take place in any segment of the earth‘s surface which is depressed by 10-20 km. Gravitational sliding and crumpling probably occur on a geotectonic scale, directed at first inwards, while the basin continues to subside. This is followed by vertical readjustment when excessive heating leads to granitization of the roots zone. Uplift results, and further gravitational slides occur, this time externuZZy directed. Through this orogenic evolution, it is evident that, taken as a whole, the sediments involved in the lower part of the trough, and those that slide into it, will become incorporated in the metamorphosed or granitized roots. The superficial sediments, however, are only involved in the “Juratype” displacement and “Alpinotype” nappe slides towards the exterior, escape metamorphism, and should be associated only with lower temperature fields and quite modest dynamic stresses. Ensuing alteration is thus little more than anadiagenetic. Theoretically it might seem possible to distinguish between the strictly compactional and the dynamic phases in such orthogeosynclines, but in practice it is difficult to draw a sharp line, especially inasmuch as much of the sliding and folding is synchronous with the sedimentary accumulation. Ontheotherhand,inpurugeosynclines(cratonic basins, i.e., one of KAY’S, 1951, auto-, paralia-, exo-, zeugo-, taphro- or epieu-geosynclines) one is dealing with a basin that has a rather stable underlying crust, and therefore deep burial is impossible. These basins, on the basis of a world survey (FAIRBRIDGE, 1959) cover 32% of the continents and shelves, with a maximum depth averaging 5,100 m and an average area size of 180,000 km2.

PHASES OF DIAGENESIS AND AUTHIGENESIS

39

It is evident that with such limitations anadiagenesis in parageosynclines is unlikely to pass down into metamorphism, except perhaps in localized zones of intense faulting. It is one of the characteristics of the parageosynclines, however, that they accumulate episodically, that is to say, there is a phase of subsidence, faulting, and downwarping, followed by a period of stability or brief uplift, which in turn is succeeded by renewed subsidence. Thus a number of well-known basins have a two-, three- or four-storied structure, each showing progressively advanced anadiagenesis with depth. Emergence is marked by epidiagenesis and even weathering, only to be succeeded by an unconformable sequence with its own new diagenetic cycle. With each renewed subsidence there will be revived fracturing and jointing. The lower stories thus display multiple generations of joints; this may be called diuclustic revival (from “diaclase”, the classical term for joint or fracture). Geochemically, anadiagenesis is the de-watering stage when connate waters are progressively expelled from the lower levels, moving upwards and outwards, and following the dip of the basin. Gradually pores close, permeability is reduced, and the basin becomes more or less sealed. So it remains until re-activated by diaclastic revival, which may be due to diastrophic motion, ranging from further subsidenceto general epeirogenicuplifts or orogeny. As waters become progressively displaced they must pass through overlying or lateral strata, in general following the predictions of hydrodynamic theory (SCHEIDEGGER, 1957). They are thus subjected to mixing with other generations (and thus potentially differing classes) of connate waters as well as the varied mineral components, some of which may still be metastable. NAGY (1960) has spoken of a “natural chromatography” and others of “clay filtering”, inasmuch as it is essentially the clays that offer ionic adsorption potentials (see further discussion on p.43). The deep waters become progressively but irregularly more saline (CHAVE,1960). While syndiagenesis is mainly characterized by initially acid waters followed by reduction, anadiagenesis is marked by increasing alkalinity but more neutral redox potential. Because in many sedimentary sections pervious sands alternate with less pervious shales, a progressive upwards mixing of expressed waters may lead to curious anomalies in the pressures and salinities in deep basins. Laboratory studies (BERRY,1960) suggest that the clays while being mainly responsible for the geochemical filtering or natural chromatography act also as semi-permeable membranes subject to the law of osmosis. The depth-time limits of anadiagenesis may be broadly defined as extending from the lower limit of syndiagenesis (ca. 1-100 m) to ca. 10,OOO m, and from say 1,O00-100,0o0 years to somewhere between 107 and 108 years. In orogenic belts, however, it may be short-circuited and grossly curtailed by rapid elevation. Numbers of indicators may be used to judge the time range of diaclasis and other phenomena. For example, the dehydration of silica gels seems to be an extraordinarily slow affair, and what may start as an opaline silica, often ends as

40

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chalcedony (the cryptocrystalline dehydrated form) or quartz. Evidence is given in the section on Hydration-dehydration to suggest that the anadiagenesis of primary silica may occupy 104 to 106 years. Epidiagenesis

Epidiagenesis is proposed and defined here as the emergent or post-diastrophic phase of diagenesis. The analogous stage in mineral genesis is associated with the terms epigene or supergene. EDWARDS and BAKER(1951) described pyrite and marcasite nodules which formed in the syndiagenetic phase as “supergene”, but this is not appropriate, because they originated in connate water, whereas supergene refers to descending (meteoric) water. By the time of the onset of the epidiagenetic phase the sediment has been successively exposed to penecontemporaneous environments, to compactional processes and now finally to subaerial controls. In the preceding section it was noted how a brief negative eustatic oscillation could short-circuit the anadiagenetic phases and how the sediment could be exposed immediately after syndiagenesis to epidiagenesis, but this would only be a minor episode or series of episodes in the whole evolution of a subsiding basin. In the epidiagenetic phase, emergence (diastrophic or eustatic) permits deep penetration by ground water, and, in appropriate basins, the establishment of artesian systems, that may in certain regions reach far below the present m.s.1. Since meteoric waters are normally saturated in oxygen and C02, a completely new geochemical cycle is usually initiated. Oxidation becomes very general and the pH will tend to drop, except where the waters are heavily contaminated by connate reserves or where they encounter precipitated soluble salts. The rapid oxidation of such mineral compounds as pyrite (FeS2) will, as discussed early, tend to raise the temperature of vadose waters far beyond the limits normally expected from the geothermal gradient, and further re-solution of certain salts will be facilitated. This general geochemicalrevival permits renewed cementation, and mineralization of fault and joint zones, thus effectively contributing to the “lithification” of the rocks. An important aspect of the epidiagenetic phase is the state of permeability achieved in the preceding diagenetic episodes. Well-compacted and unjointed shales may, for example, be so well sealed that no epidiagenesis is possible. Fresh pyrite crystals may be broken out of them with no trace of oxidation. A porous sandstone on the other hand may be thoroughly sluiced through. The question arises as to where weathering begins and epidiagenesis ends. VANHISE(1 904) subdivided his “katamorphic zone” into two - a“ belt of weathering”, and a “belt of cementation” (Le., the anadiagenetic phase). The “weathering belt” was taken to include all of the zone affected by circulating (vadose) water, which can be classed as epidiagenetic. Certainly there is a link as BLACKWELDER

PHASES OF DIAGENESIS AND AUTHIGENBSIS

41

(1947, p.500) brought out; KRUMBEIN (1947, p.171) had written on “weathering as a diagenetic process”, but evidently referred mainly to soil-forming processes, in other words, to strictly superficial phenomena. He said: “Weathering is essentially a process of delithification, but it is much more than a simple reversal of the reactions and processes of lithification. Weathering is in large part a phenomenon of oxidizing environments, whereas diagenesis proper occurs mainly under reducing conditions.” Interpreted in this light, “weathering” might be considered to embrace the whole phase of epidiagenesis, which, if justified, would do severe damage to generally accepted definitions of weathering: e.g., the group of processes, such as the chemical action of air and rainwater and of plants and bacteria, and the mechanical action of changes of temperature, whereby rocks on exposure to the weather change, decay and finally crumble into soil. It might seem wise, therefore, to keep the term weathering for these surface processes, and to recognize that the oxidation zone, often reaching to depths of 5,000 m or more, is a special (non-reducing) stage of diagenesis. The nature of the depth-time relations of the epidiagenetic phase are controlled by the accessibility of oxygenated waters, provided that the anadiagenetic cementation has not blocked the permeability of the porous sediments. There will be a tendency for meteoric waters to penetrate as soon as a given basin becomes even partially emerged, and thus to set the last phase in motion. Following hydrodynamic theory, pools or pockets of oil and natural gas tend to become isolated, and even to be pushed into tilted reservoirs by the fluid motion. Some oil basins of this sort are now reached well below sea level at depths of over 8,000 m. In mountain ranges, meteoric waters are encountered in fault zones and in pervious strata, and such waters should, theoretically at least, extend to still greater depths. The duration of such exposure is almost unlimited, and the erosional phase of a mountain system may range from 108 to 109 years.

SOME DIAGENETIC PROCESSES INVOLVING AUTHIGENESIS

There have been several attempts in the past to restrict the term authigenesis to a particular phase of diagenesis, to the syngenetic mineralization that occurs in the early or syndiagenetic phase, as distinct from epigenetic mineralization of the late phase (summarized by TEODOROVICH, 1961). PETTIJOHN (1957, p.662) drew attention to a Russian classification by BATURIN (1937) which distinguished these two divisions as: (a) Early diagenesis (“halmyrolysis” of HUMMEL, 1922) marked by authigenic mineralization. (6) Late diagenesis ((‘metaharmosis” of KESSLER, 1922) marked by epigenic mineralization. The suggested synonymy has not been widely accepted, however. “Halmy-

42

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rolysis” (see p.44) is a very useful term but is restricted to reactions directly with sea water, and is thus not so broad as “syngenesis” which affects the whole sedimentational phase (see section on Diagenetic evolution). “Metaharmosis” does not appear to differ essentially from “epigenesis” which has priority. This is not the place for a complete re-study of all diageneticprocesses, which would in any event be handicapped by the lack of experimental data, but it may be useful to review briefly those processes that involve authigenesis more or less in the order of their appearance through the three evolutionary stages of diagenesis. The chemical reactions involved are relatively simple in principle, but are complicated by the extremely “open” type systems, the multiple mixing of constituents, and the uncertain effects of time. As SUJKOWSKI (1958) emphasized, the duration of any reaction may be quite brief, but the diagenetic processes advance in jumps as the chemical media and physical conditions change. Certain metastable mixtures only become reorganized on attaining critical geochemical thresholds. Diagenesis generally tends to lead towards the simplification of the number of rock components. Generally speaking, most of these reactions are those studied under the topic “weathering”, which usually involves the unmixing of complex molecules, though sometimes the reverse is true, i.e., the construction of complex compounds which is the essence of authigenesis. In summary these reactions are: ( I ) Hydration-dehydration. Hydration-dehydration reaction involves the take-up of water on crystallization and recrystallization with loss of water, e.g., the gypsum-anhydrite relationship: CaS04

+ 2H20

$

CaS04 * 2H20

Hydration also occurs during hydrolysis, oxidation, and carbonation. (2) Hydrolysis-dehydrolysis. There is a tendency for water to react with dissolved salts (in the chemical sense). This is hydrolysis. Water plays the role of a base, and yields hydroxyl ions in solution; this is in contrast to an acid which yields hydrogen ions, which on reacting with H2O give HsOf, the hydronium ion. Most silicate minerals are susceptible to hydrolysis. Thus olivine hydrolyzes to serpentine: 2MgzSi04

+ 3H@ + 3Mg0

*

+

2SiO2 ~ 2 H 2 0 Mg(OH)2

Relative susceptibility is useful in identification of unknown minerals and can be demonstrated experimentally by means of the “abrasion pH” (STEVENSand CARRON, 1948; KELLER et al., 1953). The mineral is ground up with distilled water and the pH is measured. Feldspars give a pH of about 10, and wollastonite a pH of 11. It is found that under these conditions metal cations are liberated, e.g., Ca2+, Mgz+, Na+ and K+. Thus orthoclase may break down to kaolinite, liberating K+ and silica. TAMM(1925) demonstrated that rock flour pulverized by glacial grinding is hydrolyzed even at the stage of its incorporation in glacial

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43

melt water; indeed, hydrolysis is the principal reaction of weathering. This continued trend through geological time has largely determined the fact that the ocean is alkaline. Most detrital silicates (notably feldspars, micas and clays) carried into the ocean are in an incomplete state of hydrolyzation. Depending upon the pH of the parent rivers (which vary from extreme acidity to alkalinity), the reaction may be driven to the right, involving further hydrolyzation, or to the left, that is dehydrolyzed. Upon burial the pH in a marine sediment is likely to drop immediately from 8 to 6.5 or lower, driving the reaction to the left, but in the early burial and anadiagenetic phases, the pH rises and the reaction swings to the right again. In addition to this hydrolyzation most of the silicate weathering products also have the capacity of base-exchange and adsorption. (3) Ion adrorption. Ion adsorption is a peculiarity, discussed earlier, of many weathering products such as aluminium hydroxide, ferric hydroxide, and the whole family of clays. After organic dyes, H+ and OH- are most readily adsorbed, generally followed, in the marine environment, by the cations: Cu2+, As+,Zn2+, Mgz+, Ca2+, K+, Na+; and by the anions: Sz-, C1-, S O P . The order of replacement of cations is as follows: H, Ba, Sr, Ca, Mg, Rb, K, Na, Li. For example, hydrogen ion will replace the calcium ion, unless the latter is present at a higher concentration. After the adsorption of H+ or OH-, the adsorbent has a free electrical charge, which is a characteristic of colloidal particles. Inasmuch as clays favor the hydroxyl ion, they are often negatively charged, with the result that they tend to adsorb a whole range of rare metallic cations from ocean water (KRAUSKOPF,1956). It is therefore a vital process in authigenesis. Selective adsorption by aging gels leads to completely new minerals (EITBL,1954, p.458). (4) Cation or base exchange. A long-recognized feature of weathered silicates, such as soil clays, is their capacity to exchange cations with any passing alkaline solution (TAMM, 1925; KELLEY, 1939). An acid solution (rich in COa or “humic acids”) will tend to remove the exchangeable bases, leaving an “acid clay” (GRAHAM, 1941). Means of measuring the “free exchange bonding energy” of the cations in soils have been devised during the last decade (MARSHALL and UPCHURCH, 1933), but there does not appear to have been any attempt yet to carry out quantitative measurements in fresh marine sediments. It is merely deduced that, for example, clay mineral illite, a muscovite degraded with respect to K+, becomes enriched with that ion from sea water, whereas other clays obtain it by base exchange. (5) Oxidation-reduction. As mentioned earlier (in the section on Boundary limitations), the oxidation-reduction balance is closely related to the question of the absolute boundary limitations for diagenesis. Examples are given in a later section (on Hydration-dehydration) and only a few definitions are presented here. Oxidation implies the loss of electrons. The substances that gain electrons

44

R. W. FAIRBRIDGE

are called oxidizing agents or oxidants; and in gaining electrons they are reduced. In weathering, this oxidation is often effected by means of atmospheric or organic oxygen (liberated in photosynthesis); free oxygen is available in normal sea water and surface sediments. In an aqueous medium, however, it is often impossibIe to separate the concept of oxidation from hydrolysis. Furthermore oxidation is also involved when metallic iron is “oxidized” to FeS and then to FeSz. Immediately below the oxygenated layer of sediments is the reducing environment of “early burial”, where the above reactions go to the left. Traces of this reduction may be recognized in sediments after a complete diagenetic cycle.It will be recalled, however, that both the initial and end stages are most generally oxidizing, so that the reducing phase may be deduced only from the survival of certain minerals or by direct sampling (bottom-coring) in contemporary sediments. (6) Carbonutizution. The carbonate ion C03z- (or the bicarbonate HCO3-) is often known to replace silicates during weathering, particularly with CaZ+ and Mgz+. Both being always present in the ocean, they are likely to be organically reprecipitated as CaCOs, “magnesian calcite”, or recrystallized later as dolomite. Decarbonatization normally occurs in the presence of solutions of COz and then of HCOs-. Carbonic acid HzC03 is stabIe only at low pH and thus it is an impossible combination to occur in the ocean. The bicarbonate ion can exist in the initial (near-surface) conditions of open oceanic diagenesis,and again in the epidiagenetic phase when it is reintroduced by meteoric waters. It is rarely to be expected in the early burial (reducing) stage of syndiagenesis,and is equally unusual throughout anadiagenesis. Its role may perhaps to be traced in the early syndiagenetic alteration of brucite Mg(OH)2 to MgCOs in dolomitization, which is discussed under “Metasomatism”. Inasmuch as the processes in such complex mixtures as ocean waters and random sediments are considerably intertwined, it is convenient to consider them in more detail under headings that, at least in part, characterize the site as well as nature of the reactions. Halmyrolysis

The first reactions are those occurring directly between the sediment and normal sea water. This process, sometimes called “submarine Weathering”, was termed hulmyrolysis by HUMMEL (1922), from the Greek roots hali- (sea) and myros(1957 pp. 649, 662) in one place, (unguent), spelled “halmyrolosis” by PETTIJOHN and “halmyrolisis” in another. Chemical reaction with sea water begins, of course, immediately after the sediment reaches the ocean, and continues while it is moved over the bottom or i s swept along by currents as it settles from suspension. On reaching a point of even temporary stability it will be subject to contact with sea water, possibly for extended periods, and may then be retransported to a deeper position. Some parts of the continental shelf and slope, as well as “sills” between

PHASES OF DIAGENESIS AND AUTHIGENESIS

45

basins, submerged plateaux and guyot (flat seamount) tops, are so constantly swept by currents that fine-grained materials remain only in pockets, or where trapped inside the shells of marine organisms. Since sea water takes up elements (e.g., CaySi, P) at the expense of the sediment, the term “submarine weathering” is sometimes quite appropriate (CORRENS, 1950), but the whole picture is more complicated. TEODOROVICH (1954; see review by CHILINGAR, 1955; and discussion by PACKMAN and CROOK,1960) recognized no less than thirteen diagenetic facies that are subject to halmyrolysis or other processes of early diagenesis. They have been defined in terms of pH-Eh and mineral components. Six fundamental geochemical environments are noted: ( I ) Oxidizing (ferric oxides and hydroxides). (2) Weakly oxidizing, Eh = 0 well below sediment surface. (3) Neutral (iron chlorites), Eh = 0 slightly below surface. ( 4 ) Weakly reducing (siderite), Eh = 0 precisely at surface. (5) Reducing (carbonates, scattered FeS, FeSz), Eh = 0 slightly above interface. (6) Strongly reducing (sulfide zone proper), Eh = 0 well above sediment interface, and approaching sea surface. Glauconite Probably the most characteristic of all halmyrolytic phenomena is that of glauconitization. The mineral gluuconite (a hydrolysate mineral related to illite, a hydrous mica rich in iron and potassium)has long been regarded as an exclusively marine product, authigenic and characteristic of certain shelf and slope environments. It seems possible, though rarely, that it may even be anadiagenetic (WERMUND, 1961). CLOUD (1955) has made a useful summary of the conditions of formation, except that he believed that glauconitization was more generally favored by cooler waters; this is not correct, for its most widespread occurrence today is over the continental shelves north of Australia, namely, the Arafura, Sahul and Rowley Shelves, as originally described by MURRAY and RENARD (1891) and later extended by FAIRBRXDCE (1953, p.11). This is one of the most uniformly warm shelf areas of the world, where the sediments are very often tinged green by glauconite, except where they are masked by large supplies of detrital carbonates (near coral reefs), or situated inshore near river mouths. Although MURRAY and RENARD found glauconites down to depths of 4,000 m in the ocean, it seems likely that they were transported there with other sand-size particles, and that the optimum zone of glauconitization is from wavebase (say 15 m) to somewhat beyond the shelf margin (say 500 m). The mineralogy and distribution of glauconite suggest that it is derived from clays, micas, and feldspars by a slow hydration and ion exchange process that is favored by slightly reducing conditions together with a free access of sea water

46

R. W. FAIRBRIDGE

at a pH of around 8. Reducing conditions can be established on the open shelf in a micro-environment such as the rotting interior of a molluscan shell, of certain

Foraminifera, or associated with fecal pellets. Organic participation, even in microenvironments seems to be the rule (OJAKANGAS and KELLBR, 1964), providing the correct Fe3+/Fe2+ratio. The process is evidently very slow, because glauconite is not formed in areas of rapid burial, and is often found at intermediate stages of formation; however, the shelf areas where it is most common have only been inundated during the last 10,OOO years, so that the complete authigenesis of a particle 2 mm indiameter may require something of the order of 100to 1,OOO years. Although one lacks experimental evidence, the glauconite reaction probably lies in the activity field of the heterotrophic anaerobic bacteria (pH 7-8 and Eh 0 to -100 mv), and within a temperature range of 25-5 “C. There is apparently a reduction in the amounts of glauconite in the preMesozoic oceans (CONWAY, 1945), which may be related to the postulated higher pco2 of the Precambrian and Paleozoic times. TUGARINOV and VINOGRADOV (1961) report Precambrian glauconites back to 1.5 * lo9 years. The older the glauconite 1954). the lower is its Fe3+/Fe2+ratio (SMULIKOWSKI, One of the complicating factors about glauconite is that under alkaline conditions it remains very stable so that it may be constantly reworked, over and over again, in marine sands in some areas (e.g., through the Cretaceous and Tertiary of eastern New Jersey), thus rendering it sometimes misleading for K/Ar age determination. As soon as it is exposed in an acid soil, however, it rapidly breaks down, liberating potash (useful as a fertilizer) and iron oxides. This “longevity” of glauconite thus brings it into strange partnerships. CHILINGAR (1956b) has reviewed the Russian literature describing its occurrence in association with chlorite in non-metamorphosed sediments. Such a relationship would normally be incompatible, because authigenic chlorite implies an acid environment with low Eh. KROTOV (1953, quoted by CHILINGAR, 1956b) described a situation where the glauconite sands were formed in shallow marine waters, which he believed became cut off and converted into a lake; oxidation occurred, the glauconites were partly converted to goethite, and then under the low pH of the swampy lake, a chloritic clay cement developed. Experimental measurement of this type of environment showed a pH of 6.9 to 7.15. The low pH led to bleaching of the partly browned glauconite; then secondary pyrite, and siderite formed. A glauconite-chamosite association of rather similar nature was reported from U.S.S.R. by ZAPOROZHTSEVA (1954, also quoted by CHILINGAR, 1956b), and an analogous case seems to exist in the Springer Formation in Oklahoma. There seems no inherent reason why the low pH condition should not be bacterially produced in a submarine environment during the early burial stage. On the other hand this would seem to be impossible at the normal sediment/water interface, for chamosite is not found in contemporary marine sediments.

47

PHASES OF DIAGENESIS AND AUTHIGENESIS

Phillipsite Another probable product of halmyrolysis is the marine zeolite, phillipsite (sometimes given as ( K ~ C ~ ) M & O I ~ . ~ H Z which O ) was first discovered to exist widely in deep-sea environments by the Challenger Expedition (MURRAY and RENARD, 1891). These elegant little crystals, often twinned, make up an appreciablepart of the “red clay” sediments in parts of the Pacific, where they are not masked or inhibited by rapid deposition of other materials. In places they exceed 50 % of the non-carbonate fraction. On land this mineral is found only in association with basaltic lavas, and in the ocean it seems to occur downwind of island volcanoes, the ash showers from which have been widely distributed. It is not found in volcanic ash on land, and so appears to be the result of halmyrolysis, but the details of its formation have never been completely elucidated. Some recent studies suggest that palagonite tuff (basaltic glass) is altered to montmorillonite, and thence to the zeolite. Photographs show phillipsite forming inside palagonite nodules (BONATTI, 1963). Zeolites Contemporary zeolitization is hardly distinguishable from the first stage of metamorphism, the zeolite facies (FYFEet al., 1958; COOMBS, 1959; TURNER and VERHOOGEN, 1960, p.532). This phenomenon may occur at depths down to 20,000 ft. (say 2,000 bars pressure and 100-20O0C), provided that there is suitably metastable parent material, in this case, fresh, rapidly accumulated volcanic ash. Tuffs, deposited from nu6es ardentes, rapidly develop zeolites close to the surface, as around Vesuvius (NORIN,1955). Reactions between connate water and volcanic glass may give analcite (BRADLEY, 1929; EITEL,1954, p.997). There is also the devitrification of glass to heulandite (Ca, Naz) AlzSi6016 5Hzo), and laumontite (CaMzSi401~ 4H20), and the albitization of plagioclase. Indeed, these phenomena are no more than a bulk interaction of compounds normally found scattered through basins that have passed through the phase of anadiagenesis. One might even question the classification of this zeolite facies under the heading of metamorphism. Thus in the West Coast ranges (Oregon and northern California) zeolite facies are regarded by HAY(1962) as simply diagenetic (see also PACKHAM and CROOK,1960).

-

-

Clay minerals A much-studied aspect of halmyrolysis is the authigenesis of clay minerals. According to GOLDBERG and ARRHENIUS (1958), marine pelagic clays account for about 13% of the total sediment laid down over geologic time. Clay and colloidal particles are transported (mechanically) by rivers, in suspension or adsorbed onto organic gels, or by wind, as dust (or loess), to the oceans. When transported in particulate form, the mineral will have its eventual character dictated primarily by

48

R. W. FAIRBRIDGE

the source region (“Heritage” of MILLOT,1953, 1957). Alternatively, when transported as finer colloids and ions, susceptible to electrolysis, flocculation occurs within a few hours on contact with .sea water, and authigenesis can be expected to follow. Upon burial in a largely clayey facies, there is a natural tendency toward the creation of a relatively impervious shield, so that syndiagenetic alteration is restricted by the very limited supply of ions available in the connate waters immediately in contact with each of the clay particles. During early anadiagenesis an almost closed system may be created and clay minerals buried in this environment show little change in a column of over 1,OOO m (MILNEand EARLEY, 1958); with progressive loading and tectonic evolution, fracturing permits recirculation and natural chromatography (see section on Natural chromatography) is facilitated. Burial in a pervious sandy facies, however, permits very extensive authigenesis. Systematic regional studies of deep oceanic clays are still sadly lacking, but the broad pattern that is emerging seems to suggest that there are definite geographical, latitudinal provinces, dependent upon these two factors: (I) inheritance (source materials and conditions) and (2) authigenesis (local materials and conditions). GRIM(1958) has remarked that certain varieties of clay minerals are more “at home” in certain environments than others, and this is the basis of authigenic differentiation. A “concept of dualism is essential to the understanding of clay genesis”, according to WEAVER (1958); he refers here to the inherited characteristics on the one hand and the depositional environment on the other. The various ancient geosynclines and basins of North America are now rather well differentiated on this basis (WEAVER, 1960). Clay minerals most clearly reflect the environment of deposition in young accumulations where the sedimentation rate is low, and thus where time is available for halmyrolytic reactions. In eugeosynclines where the sedimentation rate is high, rapid burial seals in the components, and source region is most clearly reflected by the minerals eventually stabilized. Whereas a very large number of clay minerals are known, only three principal types will be discussed here: (a) Illite, the hydrous mica group of clay minerals, was first named by GRIM, et al. (1937); these are the so-called “low alumina clays” (10-20 % 141203). Crystallographically, they appear to have essentially the same lattice structure as muscovite, but to be degraded with respect to Kf, thus: (OH14 K, (A.b.Fe4.Mg4.Mg6) (S~B-,Al,)

020

where y is around 1-1.5. Illites seem to be formed during weathering in the moderately high pH soils of cool to temperate climatic belts, for example, the partly weathered till and loess

PHASES OF DIAGENESIS AND AUTHIGENESIS

49

plains of Illinois. In the ocean they are either reworked from older illites or authigenically derived from less stable clays; illite appears to be the most stable of the clays in the marine environment. During syndiagenesis it may emerge from the halrnyrolysis of montmorillonite or kaolinite (see (b)and (c) below). Particularly K+ is adsorbed in these reactions, which are favored by the presence of CaZ+, MgZ+, and ferrous iron. Halmyrolysis may well be initiated the moment that the clay colloids reach the saline waters of the ocean. GRIFFIN and INGRAM (1955) showed how, in a single estuary (the Neuse River in North Carolina) that drains a hinterland of relatively uniform metamorphic rocks, kaolinite is by far the most abundant mineral introduced. As the water becomes progressively saltier, it is replaced by “chlorite”, to be replaced in turn by illite near the mouth. Some of the kaolinite gets through this transportation phase, however, because it is present in most oceanic sediments (see p.52). Illite is the commonest clay type in the Paleozoic fold belts of the world, and thus it is not surprising that it is the dominant type in the temperate North Atlantic as well as in the Pacific (GRIMet al., 1949). They suggested, incidentally, that a weathered illite which had lost most of its K+ before reaching the sea could be reconstituted by ion adsorption during early diagenesis. In the tropical Atlantic, the illite-montmorillonitc+kaolinite ratio is 1 :1 :1 (HEEZENet al., 1960), but 8 : 1 :1 in the temperate latitudes. K/Ar isotopic agedating of these North Atlantic illites give mean ages in the 300-400 million year brackets (HURLEYet al., 1963), which strongly suggests simple reworking and transport from Paleozoic sources. It would seem thus that authigenesis plays a rather minor role in the production of illite, although it is certainly possible to interpret this ‘cPaleozoic”age as merely the statistical result of sedimentologically “smoothed” mixing of still older clays with a fraction of modern authigenic products. The supply of K+ available in the connate water would in any case restrict the degree of authigenesis in the halmyrolytic stage. Isotopic dating of Pacific clays (HURLEY et al., 1963) gives a Late Mesozoic age, which also suggests a minor role for authigenic clay components during the syndiagenetic phase; this does not preclude further K+ being taken up during anadiagenesis. It would be interesting to learn if the K/Ar “dates” of deep-sea illites drop as one goes down the core. It is noteworthy that the Kf values in clays rise on going back into the (1945) suggested earlier that there was a Precambrian (NANZ,1953), and CONWAY peak of K+ extraction in the oceans in Late Precambrian time; while he suggested subsequent increased organic activity as the cause, it may perhaps be that progressive rise of oceanic pH is less favorable for illite halmyrolysis today. Also unfavorable to the extensive formation of illites are the widespread desert conditions of today.

50

R. W. FAIRBRIDGE

(b) Montmorillonite, generally given as AlzSi4010 * (OH)2nH20. The structural formula of montmorillonite group of minerals is best expressed as:

-

~ Fe3+)n]04~nHaO (OH18 (R3+, R-~ + )+E- [ S i ~ a -(4, 3

where R3+-Al,

3

Fe3+, Cr; R2+-Fe2+, Mn2+, Mg, Zn, Cu.

In this “expanded” or three-layer lattice clay group, some of the aluminium may be replaced by Mg (suponite), or by Fe (nontronite). A general term for the montmorilloniticclays derived from volcanic ash is bentonite. Because of its expanded lattice, montmorillonite is one of the most reactive types of clay and exhibits a considerable propensity for base exchange and ion adsorption (KELLEY, 1939,p.434). The order of replacement is generally Na+
PHASES OF DIAGENESIS AND AUTHIGENESIS

51

the exchangeable ions available is puzzling, but inasmuch as this region is more or less smothered by carbonate oozes, the surface and immediately subsurface pH is kept near 8 which may stabilize the montmorillonite. It would be interesting to know what becomes of the latter at depths of a few cm below the high bacterial layer (where the pH is less, in spite of the calcium-rich environment). It is noteworthy that in more or less land-locked lagoons, where evaporation becomes dominant and gypsum is precipitated, the pH is generally above 8.5 and montmorillonite is replaced by illite (MILLOT, 1953). In a study of some Permian clays from Kansas, M. F. Norton (personal communication, 1958) found that there was a positive correlation between the abundances of dolomite and the high Mg montmorillonites, and between those of limestone and illite. MILLOT(1953) has remarked that ancient limestones are generally associated with illite. Thus the pH and Eh of the depositional environments were normally high (pH 7.5-8.5), and sediments were poor in organic matter. Montmorillonite appears to be favored by the lower Eh (a feature of many dolomites) as shown by the scattered pyrites in the latter (FAIRBRIDGE, 1957). (c) Kaolinite is the simple two-layer lattice clay, characteristic of the socalled “high alumina clays’’ (2040% A ~ z O S )It. does not appreciably expand under varying water content, and does not exchange iron or magnesium. The typical group formula is: A1203

or:

- 22302

*

2H20

A14Si4010(OH)~ Kaolinite is well known as the stable end-product of laterization or latosols, produced by deep tropical weathering with pH of 5-7; but being well oxygenated, the Eh of environment in which it forms is neutral to high. Its formation is especially favored by aluminous igneous rocks such as granite (Ross and KERR,1931). In this process Ca2+, Mg2+, and Fez+, are leached out; in subtropical regions silica and iron rise to the surface by capillarity to form a ferruginous duricrust, whereas the Ca2f and Mg2+ tend to pass into the ground water and drainage system and then to the sea. Kaolinite itself tends to be broken up only as the result of tectonic and eustatic oscillationsof base level. Under such conditions it is transported to the ocean (as during the Quaternary). In the ocean today, kaolinite is, therefore, mainly a transported, nonauthigenic sediment, and is most widespread in the tropical Atlantic (fed by the Amazon, Orinoco, Congo and Niger), accordingto CORRENS and VONENGELX-IARDT (1938), MILLOT(1953), and HEEZENet al. (1960). It is not widespread, however, in the tropical Pacific (no rivers), according to GRIMet al. (1949). Authigenic kaolinite is found here and there in former marine sediments, but it is evidently epidiagenetic for it is restricted to areas where the interstitial 1964). waters are fresh (SHELTON,

52

R. W. FAIRBRIDGE

GRIM(1951) pointed out that kaolinite and montmorillonite are much less common than illite in ancient sediments, and thus one may suspect that diagenetic replacement by illite has been operative. The possibility of long-term geochemical changes in sea water, however, should not be forgotten (see notice of Conway’s work, p.49). MILLOT(1953) posed a good question: how is kaolinite, which is formed in the acid environment of the continent, preserved from diagenesis in the alkaline environment of the sea floor, with which it seems to be out of equilibrium? ZEN(1960) believed, on the other hand, that the intimate association of calcite-and kaolinite implies an equilibrium condition with sea water. In Paleozoic rocks he noted the five component system calcite-dolomite-chlorite-kaolinitequartz, which seems to be a stable mineral assemblage. Millot made the helpful suggestion that where kaolinite is rapidly accumulating on the ocean floor, it is quickly sealed in by a relatively impermeable layer, and with plentiful bacterial action the pH quickly drops to 6-7 and the H2S lowers the Eh potential. One may conclude, therefore, that in regions of low sedimentation rate and low organic productivity (e.g., the Red Clay areas of the Pacific) there may be regular diagenesis of kaolinitesto illite. Inasmuch as the sedimentation rate here is of the order of 0.1 mm a century, there is a first approximation of the diagenetic rate. To close this section, one may conclude that the limits of clay diagenesis are still rather poorly defined, especially the lower boundary of halmyrolysis. Oxidation and reduction

Inasmuch as normal marine basins are in constant circulation, oxygen saturation is very generally maintained, and the normal open-sea floor is oxygenated at the sediment-water interface. Euxinic environments, such as the Black Sea, constitute special cases, where reducing conditions, even in the liquid medium, exist everywhere below a certain “still depth”. Here, at the sediment-water interface in well ventilated basins, the oxidation of sedimentary particles is another of the earliest possible diagenetic processes. The slower the sedimentation rate, the more complete is the oxidation. Thus the deep water areas far from land, e.g., central Pacific, are the classic sites for what MURRAYand RENARD(1891) identified as the red clay deposition, due to the oxidation of iron. This is also the site of manganese nodule formation which occurs around any sort of nucleus from a glacial boulder to a pelecypod shell. This coating of MnO2 is a function of the time exposed at the sediment-water contact. Divalent Mn2+ present in sea water is apparently oxidized to the tetravalent species Mn4+ (GOLDBERG and ARRHENIUS, 1958). Large boulders do not obtain a coating of MnO2 on their deeply buried side. Such boulders were carried out into the Arctic Basin by icebergs during a time of open water conditions (a controversial question,

PHASES OF DIAGENESIS AND AUTHIGENESIS

53

but the writer suggests about 5,000-10,000 years ago during the “climatic optimum”, FAIRBRIDGE, 1961). After cool conditions returned, the basin was covered by pack ice and further sedimentation was greatly retarded. Inasmuch as the thickness of MnOz accumulated is only 2 or 3 mm, the accumulation rate is around 0.0003 mm per year. The production of red beds in the geological past is a complex problem, depending in part on source materials and in part on diagenesis (PETTIJOHN, 1957; VAN HOUTEN,1961). Red beds are often continental, but include the marine “red muds” (TWENHOFEL, 1950, p.331) of the Orinoco, Amazon and Yangtze; they may extend far out to sea in depths down to 2,000 m, but the color is steadily lost by reduction. According to the theory of biorhexistasy (ERHART, 1956) long-term stable conditions under subtropical weathering conditions (the norm for most of geological time) leads to a general deep weathering resulting in lateritic soils, accompanied by removal of calcium and SiOz. Such times (biostatic) are marked in the marine realm by limestones and cherts (ERHART,1963). A cyclic lowering of the water table which takes place eustatically during periods of cold climate and tectonically lowered marine basins (“bathygenesis” of H. TERMIER and G. TERMIER, 1963) leads to a mechanical break-up of old soils (rhexistasy), to stream dissection and to transport of red iron-rich lateritic debris into continental basins and to the ocean. Such sediments are so rich in iron oxides that, even in deltas (cf. the Devonian Catskill-type deltas of the Appalachians), which normally compromise reducing environments, the red colors often persist, although green layers and mottled and bleached patches bear witness to reduction by HzS (MOULTON, 1926). Buried along with sediment there is always a certain amount of organic matter, which serves as bacterial nutrient, the principal source of geological reducing agents (IRVING,1892). In some areas (as off deltas) and likewise in euxinic basins, where bottom scavengers are inhibited by the H2S poisoning, this amount is very large. Since open-sea floors are normally inhabited by both epiand in-fauna there is a considerable and rapid consumption of much of this organic nutrient. The proportion of organic material that is ultimately incorporated in the sediment (below burrowing depth) depends on two factors: (a) the size and vigor of the benthonic population and (b) the rate of deposition. As TWENHOFEL (1942) pointed out, a slow sedimentary rate permits thorough scavenging. He stated (p.105) that: “It seems a paradox that the more congenial the conditions on the sea bottom for bottom dwelling forms and the more numerous the colonization by organisms, the less likelihood there is of many fossils in the sediments which finally attain entombment. In other words, an abundant bottom population under conditions of slow deposition produces deposits with few complete shells and more

54

R. W. FAIRBRIDGE

or less complete elimination of all nutrient matter. Accumulation of organic materials is not possible under such conditions.” By the same token, the bacterial population in these low sediment rate areas is also low in the sediments just below the surface, and thus the opportunities for authigenesis by reduction under these conditions are strictly limited. Two distinct environments are so characterized: (a) continental shelf and slope regions where the sedimentation and subsidence rates are low; and (b) abyssal plains and rises far removed from continental sediment sources. In contrast, shelf regions near deltas or other sediment source, and particulariy abyssal plains richly fed by turbidity currents, have their benthonic populations constantly smothered by seasonal, or longer cyclic,invasions of sediment that seal off the organic debris and provide a large reservoir of bacterial nutrients. This is the ideal site for authigenesis by reduction. In estuaries, BAASBECKINGand MOORE (1959), have found that the average organic content is 12 %, and may exceed 25 %. In the same samples the iron, initially in the form of FeO(OH), is found to become completely reduced and anywhere that the organic content of the sediment is above 2 % there is an excess of H2S. Typical formulae would be (BAASBECKING, 1959, p.57): Fe(OH)a+CO2+7H(AF298= $33.28 kcal) (1) FeO(OH)+CH4+2H20 -+

FeO(0H)f iC6H1206+H20

--f

Fe(OH)ztCO2+3H(AF298= +0.55 kcal)

(2) The principal minerals so formed are the common ferrous sulfides, marcasite and pyrite, FeS2. These are produced by reduction of iron oxides and the various hydrates (first to FeS and later to FeS2) by the action of H2S liberated by sulfatereducing bacteria and by bacterial breakdown of organic sulfur compounds. It seems that marcasite (the orthorhombic form) is produced under neutral to acid conditions (pH generally less than 7.0), whereas pyrite (the isometric form) is favored by slightly alkaline conditions (ALLENet al., 1912; NEWHOUSE, 1927; TARR,1927). It is observed that pyrite is the common form associated with impervious clays and shales (LOVE,1963), whereas marcasite is most often found in sands, silts, chalks, and limestones, as well as acidic fresh-water swamp deposits such as coal formations. EDWARDS and BAKER(1951) seemed to be under the impression that marcasite was exclusively fresh-water, whereas pyrite was of marine origin. This is not so, however, because bacterial action often lowers the pH of fresh marine sediments into the stability range of marcasite. It may be suggested that this is a function of the permeability of the sediment; the clays under reducing conditions rapidly become alkaline and can preserve a high pH for extended periods of time, whereas the pervious sands, calcarenites, etc., favor the aerobic bacteria, which keep the pH low until most of the organic matter is consumed; the Eh, however, will steadily drop as the organic material breaks down. Even

PHASES OF DIAGENESIS AND AUTHIGENESIS

55

after saturation with H2S (a weak acid) the pH is still low and calcite fossils are generally destroyed (MOSEBACH, 1952). Considering some of the evidence of the geological past, it must be noted that pyritized fossils, such as typically pyritized ammonites in the Mesozoic, are often associated with black shales which lack any benthonic fauna. General poisoning of bottom conditions might occur from time to time, even without an actual barring of the basin, for example by seasonal or cyclic invasions of protista or Algae (“waterbloom”) (RUTTEN,1953), or other sources of mass mortality (BRONGERSMASANDERS, 1957). Reducing environments are strictly syndiagenetic, evidently produced during or soon after deposition (in the time range of about 1 to 1,000 years), with Eh ranging from 0 to -400 mV and pH either slightly above or below 7, as appropriate either for marcasite or for pyrite. Exceptionally, the iron sulfides are also formed at considerable depths (3,000-5,000 m) as observed in deep oil fields, namely under anadiagenetic conditions. In such cases it would appear that sulfur bacteria were living on the petroleum hydrocarbons (BASTIN,1926; GINTER,1938). The fate of the sulfides after the early burial stage is partly illustrated by the 32S/34S isotope ratio (see introductory notes, in the section on Syndiagenesis). The initial ratios of 22.1 to 22.7 in recent sulfides, and figures up to 22.65 for a Precambrian shale from Finland more than 1.8 109 years old (AULTand KULP, 1959), suggest that there may be little change in the nature of the bacterial reduction during that time-span. It has been pointed out by THODEet al. (1960), however, that in general there is a gradual depletion of 3% in rocks of increasing age which may be due to isotopic exchange between pyrite and sulfates in solution. The concentration of these solutions rise with age and the depths of the basin during the anadiagenetic phase. The sulfates below the level of the syndiagenetic phase may be derived either from evaporites, or from the re-oxidation of the sulfides. On emergence into the oxidizing conditions of the epidiagenetic phase, nodules of pyrite and marcasite tend to acquire a coat of limonite, and this in turn is normally externally dehydrated to hematite. The marcasite, in general, seems to be the less stable of the two forms, and the centers of nodules are often found to be ultimately broken down to a grayish powdery form melnikovite. An interesting and unusual example of reduction in the epidiagenetic phase (otherwise very generally oxidizing) occurs in the anhydrite caprock of certain salt domes. The anhydrite was probably laid down as gypsum in evaporite deposits that rose up diapirically during the anadiagenetic phase. The anhydrite becomes concentrated in the caprock by differential solution of the more soluble halite and other salts. The heavily fractured domal structures lead to the migration and accumulation of petroleum hydrocarbons which provide nutrients for sulfatereducing bacteria (Desulfovibrio desulfuricans). HzS is then produced, and on oxidizing (assisted by thiobacteria) gives rise to enormous native sulfur deposits.

56

R. W.

FAIRBRIDGE

The biologic nature of this reduction process is proven by the low 32S/34S ratio of the sulfur with respect to the sulfate (FEELYand KULP, 1957). Furthermore, the 12C/13C isotope ratios of 92-94 in the calcite of the caprock are typical of petroleum and not of marine limestones (89), suggesting that the COi- was derived likewise from the hydrocarbons. Under certain conditions of oxygenation, generally epidiagenetic, and under conditions of heavy tropical rainfall (high COZ 0 2 intake), there may even be a liberation of HzS04 from pyrite or marcasite. Some mine waters actually have a negative pH, which can be explained by the following reaction:

+

2FeS2

+ 2H20 + 7 0 2

--f

2FeS04

+ 2HzS04

It may, however, be better represented (BAASBECKING, 1959, p.61) by the equation: FeSz + 8H2O -,FeS04

+ HS04- + 15Hf + 14e-

Inasmuch as FeS04 is stable only in an anaerobic environment, it is likely to hydrate to melanterite, FeS04 . 7Hz0, and eventually perhaps to coquimbite Fe2(S04)3 * 9H20, in the epidiagenetic stage. This especially occurs along joint planes, where oxidation is followed by desiccation, resulting in the deposition of hematite or limonite in Liesegang rings. These are precipitation patterns that are either planar (in joints) or omnidirectional and roughly concentric (in porous media). They were first observed by R. E. LIESEGANGin 1896 (STERN,1954), who recognized them as diffusion phenomena due to a periodic alternation between solution mobility (diffusion) and supersaturation (nucleation and precipitation). In a poorly stratified sandstone they may sometimes be confused with penecontemporaneous sedimentation phenomena, such as slumping, load casts, and concretions. The Liesegang phenomenon is not, however, limited to the epidiagenetic phases of sediments, being reported in the syndiagenetic phases of sediments, both marine (STETSON,1933) and fresh water (SUGAWARA, 1934). They are also noticed to be diffused in syndiagenetic cherts (BISSELL, 1959). Hydration-dehydration

Inasmuch as one deals with a watery medium in most sedimentary regions, initial hydration or hydrolysis is the rule. The gradual compaction of the basin during anadiagenesis, however, not only raises the confining pressure, but also tends to drive off the interstitial and bound water. This phase is marked by dehydration and dehydrolyzation; in the epidiagenetic phase rehydrolyzation or revived hydration may be expected. Gypsum-anhydrite The most prominent reaction is that between gypsum CaS04 .2H20 and anhydrite CaS04. The hydrated sulfate is produced as a primary evaporite deposit as soon as

PHASES OF DIAGENESIS AND AUTHIGENESIS

57

ordinary sea water is concentrated to somewhat less than 50% of its original volume, as in lagoons along the Texas coast or on the Persian Gulf. It has been stated that when the water temperature exceeds 4 2 ° C anhydrite becomes the stable phase (CONLEYand BUNDY,1958); however, this has not been so observed in nature. Gypsum is also formed in the early burial stage of syndiagenesis by bacterial action, e.g., in the lagoon muds of New Caledonia (AVIAS,1953, 1956). It is also found in certain beachrocks (MACFADYEN, 1950). From field evidence the writer is of the opinion that dehydration from gypsum to anhydrite occurs upan burial to depths of the order of 100 m or so (confining pressures of ca. 50 kg/cm2). Thus, for example, anhydrite has never been-observed forming today in the Persian Gulf, but it is extremely widespread all over the Mesopotamian Basin in Tertiary formations that, near the basin margins, have been buried under little more than 100 m. The gross effect of dehydration is to reduce the volume of the formation; from 100 m of gypsum only 62 m of anhydrite would remain. MACDONALD (1953) believed that it would require an 800 m load of sediment to set this reaction in motion, but BRAITSCH (1962) and BORCHERT and MUIR(1964) believe that even a very slight dynamic metamorphism (with sheer stresses, produced by local faulting and uneven settling of the basin) would greatly reduce this figure. It also seems probable that slightly different conditions of oceanic salinity and somewhat higher mean temperatures in the past may have permitted primary precipitation of anhydrite, so that not all sedimentary anhydrite should be regarded as the result of diagenetic dehydration (FAIRBRIDGE, 1967). Annual layers of gypsum in the Tertiary of Sicily are observed to pass upwards, with increasing grain size, into anhydrite, which is apparently a high summer peak (OGNIBEN, 1955). Likewise, in Mesozoic anhydrites from Texas, primary structures, including graded bedding, have been recorded by RILEYand BYRNE(1961). The question of primary deposition or possible dehydration of gypsum to anhydrite must, therefore, rest on the presence or absence of primary structures. The presence of collapse structures due to the reduction in volume might also be instructive. The reverse process, that is to say, hydration in the epidiagenetic phase, is very well known. The anhydrite expands about 40%, often producing ribbons of intestine-like folds (what GRABAU, 1913, called enterolithic structures). These have a curious symmetry and lack of orientation that distinguishes them from slumping or drag folds. Sometimes the expansion is limited to simple domes or rolls. The “tepee structures” of the Guadaloupe Mountains, New Mexico, and elsewhere in the semi-arid Southwest may be related to this phenomenon (NEWELLet al., 1953). Silica gels Even more important, and certainly more complicated is the dehydration of

58

R. W. FAIRBRIDGE

silica gels, to form authigenic opal, chalcedony and quartz. The opal is not completely dehydrated and is thus somewhat unstable, and will not normally pass through the anadiagenetic phase without dehydrating to chalcedony or quartz. Opal is, however, a very common form to reappear in the epidiagenetic phase, wherever a very high pH permits remobilization of silica (especially in the deep ground waters of semi-arid regions). The introduction of silica into the oceanic sediments occurs in four principal ways : (I) Colloidal form and solution, through river transportation. (2) Dust, by eolian transport from deserts. (3) Volcanic glass and ash, from eruptions. (4) Organic fossils (Radiolaria, diatoms, and sponge, Alcyonaria and holothurian spicules). The annual increment of SiOz to the oceans is 3 108 tons, or 0.23 mg/m3 of sea water (CHILINGAR, 1956c), but the ocean is undersaturated with respect to silica (0.1-4 p.p.m. near the surface and 5-10 p.p.m. near the bottom: perhaps due in part to biogenic removal). Amorphous silica is soluble at 100-140 p.p.m., the solubility rising mainly with temperature, in both normal sea water and distilled water. Quartz, however, is much less soluble (only 7-14 p.p.m., according to SEVER,1957a). The shells of Radiolaria and diatoms appear to have some organic protection from solution (LEWIN,1961), but rapid burial also protects them (RIEDEL,1959). At the present time, silica gels on the sea floor are diStinCtlyrare(TWENHOFEL, 1932,1950; RUSSELL and RUSSELL, 1936); but the present period is not necessarily a favorable one, and it would seem that among the four main sources listed in the previous paragraph there may be considerable quantitative variations from time to time. Ancient cherts have sometimes been correlated closely in time and space with periods and loci of orogenesis and vulcanism (Hoss, 1957). But this may not be the only factor. The supply of silica gels from soil-forming processes may be greater than that from vulcanism, and soil weathering rates are conditioned by temperature and crustal stability. Today it is necessary to consider the exceptionally active erosional history of the Quaternary with its rapid eustatic ups and downs. ERHART (1956, 1963) has brought out that the long stable (“biostatic”) periods would favor liberation of soil silica and calcium (see the thalassocratic stage in the diagram by R. W. Fairbridge in: TERMIERand TERMIER,1963, fig.197 p.333); during the low or oscillatory sea-level phases, there is a break-up of the soil profiles (“rhexistasy”) and mechanical detritus would dominate. As a result silica gels would tend to .be eclipsed by floods of detritals. Indications of such silica gels at times other than the present are numerous. Some have recently been listed by DANGEARD and RIOULT (1961). In the lowermost (1948) found chalcedony forming a cast of Ordovician of Poland, SAMSONOWICZ

-

59

PHASES OF DIAGENESIS AND AUTHIGENESIS

ripple marks. Several generations of ripples were similarly preserved in the Jurassic of Normandy (LEMA~TRE, 1960). A chert containing a worm burrow was noted by DEBELMAS (1959). A fossil Emgyra with its muscle tissue preserved in silica seems to call for genuinely “contemporaneous” replacement (GIDON,1959). WETZEL (1933) found perfectly preserved pollen grains in the Cretaceous flints of North Europe and DEFLANDRE (1936) reported in them delicate FZugeZZatu and hystrichosphaeridae with pseudopodia perfectly preserved. From the perfect preservation of fossils, CAYEUX (1897) long ago concluded that the dehydration of the silica began on the outside, and worked slowly inwards. There are numerous indications of the small-scale coagulation of silica gels apound nuclei or centers of lowered pH during the initial stage of syndiagenesis. The roots of mangrove-type trees, for example, are permanently fixed in marine muds, where the acidity is exceptionally low (pH 5.5-6.5). Cylindrical shaped chert concentrations result and may be termed rhizomorphs (NORTHROP, 1890) or rhizoconcretions (TERS,1961, p.172). “Fossil roots” of this sort are equally well preserved also by travertine. Discussing their form and occurrence in the Jurassic of Normandy (“Pierre de Caen”), DANGEARD and RIOULT(1961) have made a helpful review of this entire field. Such concretions were found forming today in the mangrove swamps of New Caledonia (AVIAS,1956). The question of migration of gel particles towards nuclei presents further difficulty. Mutual attraction of like molecules in solutions are normally attributed to the so-called Van der Waals Forces, assumed forces understood to exist, particularly in liquids, in the absence of which the random motions of electrons would appear to disperse the components. The supposed motion of gels (possibly as solutions) through well packed sediment for distances of a meter or more towards certain nuclei, usually organic, still poses interesting problems. The dehydration of silica gels has not been thoroughly studied, but in general seems to pass through a porous opal stage, eventually to chert and leads to complex brecciation phenomena (TALIAFERRO,1934, 1935). Referring to brecciated chert, CAYEUX (1929, p.371) remarked that here one deals with “a little known subject. , . consolidation of sediments as they are deposited.” GIGNOUX and AVNIMELECH (1937) noted that”. ..on looking closely at the fragments, one often has the impression that they might have been fitted in and cemented together like pieces of a jig-saw puzzle. . as if the debris had been formed by breaking up, scattered and the space between filled with cement.” Such brecciation, filling and rehardening has happened repeatedly with many flints and cherts, evidently due to the slow desiccation of the exterior, crushing and refilling from the reservoir of the still plastic gel. In the hills of eastern Judea (between Jerusalem and the Dead Sea, for example) there are Upper Cretaceous chert beds 30 cm-10 m thick, that have been involved in violent slump folding, 3-10 m high. They are interstratified in soft white chalky limestones extending over several hundred km2 (see Fig.5). The adjacent limestones, however,

.

R.

60

I-----

W. FAIRBRIDGE

EAST

JUDEAN H I L L S (Upper C r e t a c e o u s c h a i k s )

(undeformed )

F i g 5 Sketch of slump structures in Upper Cretaceous chert beds in the Judean Desert, east of Jerusalem. Sliding must have taken place during syndiagenesis while the silica gel was still completely plastic. (Sketch by the author.)

are involved only in very mild undulations (LEES,1928). The silica gel must have been still essentially plastic (externally brecciated only, like lava flows) and thus easily susceptible to sliding before the deposition of the nekt layers of chalky limestones. The movement was thus penecontemporaneous and corresponds in time perhaps to some of the early movements along the Jordan-Dead Sea Rift system. The gels seem to have been buried by a few meters of soft chalk which provided support for the folds; one might place the motion in the earliest stage of syndiagenesis, say within some 100,000 years of burial. Penecontemporaneous slump structures, large and small, are features of almost all chert formations. One can turn now to a second example of a timing indicator. In eastern Denmark, south of Copenhagen, at Steviis Klint and M ~ n Klint s the Upper Maastrichtian chalk is marked by parallel planes of scattered flints, i.e., parallel to the stratification, but there are also vertical joints, and here and there the silica of the flints has flowed up along the joint planes (so-called “flint curtains”, RUTTEN,1957, p.433); there is no sign of brecciation and the flints of horizontal extension appear to be completely continuous with their vertical off-shoots. One must conclude that the dehydration of the silica gel was immediately subsequent to the diaclastic phase. Such jointing must result in loss of gas pressure, release of C02, CH4, etc., and would accelerate dehydration (SUJKOWSKI, 1958). The joints rise to the top of the chalk which is cut off by unconformable layers (“fish-clay” and Cerithium limestone) of Lower Danian age (ROSENKRANTZ and RASMUSSEN, 1960). Evidently the jointing and syndiagenesis occurred after the general compaction and dewatering of the chalk but before the unconformity occurred. The timing of the flint dehydration may have been of the order of 1 . 105 years. Complete dehydration

PHASES OF DIAGENESIS AND AUTHIGENESIS

61

of silica gels may be extremely slow, and SUJKOWSKI (1958) mentioned some flints encountered in deep bores that were still rather soft (see Fig.6). Nevertheless, the emplacement of the silica is usually an early phenomenon. Studies of the 1 * 0 / 1 6 0 isotope ratios in coexisting cherts and limestones suggest early diagenesis under similar marine environments (DEGENSand EPSTEIN,1962). In contrast to the above evidence of very late silica gel mobilization, there is the fact that penecontemporaneous flints are found eroded, reworked, and reincorporated in some conglomeratic beds in the chalk, as for example on an old buried anticlinal ridge joint west of Paris. Evidently these flints were already hard, subaerially exposed and indeed, some show conchoidal fracture from the erosive period. Further there are examples of silicification preceding the penecontemporaneous dolomitization of many midwestern Silurian coral reefs (DAPPLES,1959). What process could cause this apparently accelerated dehydration of silica gels to form penecontemporaneous hard flint? (RUTTEN(1957, p.436) suggested that on the shallow platforms temporary emergence could lead to the desiccation of the gels (which often enclose minutely perfect fossils of sponges, etc.). CAYEUX (1941) in his “Causes anciennes . . .” regretfully came to the conclusion that there were occasionally in geological time conditions which simply cannot be matched today, and chert formation, though formerly penecontemporaneous (as recognized by him in 1897), must be regarded today as non-actualistic. Rutten, though disagreeing in some other respects, concurred that this could be the case with these desiccating silica gels, i.e., that there is simply no comparable shallow carbonate platform today where this phenomenon might reasonably be expected to be operative. But there are some contemporary carbonate platforms-in the Bahamas, for example. West of Andros there is a broad bench of carbonate mud that is slowly drying out, ca. 60 cm above m.s.1. Radiocarbon dates show that the mud is about 2,500 years old (the time of the “Abrolhos Submergence” that appears to be a eustatic and thus world-wide high sea level of up to 1.5 m above the present; FAIRBRIDGE, 1961, p . 169). Yet, unfortunately, this mud carries no desiccating flints. It is undoubtedly true that brief emergences of shallow platform environments are to be expected from the evidence of the eustatic theory and, indeed, the “corrosion zones” or “discontinuity surfaces” that are a characteristic feature of many neritic limestones, are now generally accepted as evidence of brief emergence (WEISS,1958; JAANUSSEN 1961). In such cases the epidiagenetic phase is temporarily applied, the mud becomes rapidly dehydrated, and sometimes even develops a karc,t crust with a trace of red soil (terra rossa). When reburied, no further reactions are likely and the whole syndiagenetic phase is locally short-circuited. Even briefer is the formation of tropical beachrock from calcareous sands or coral or molluscan debris. The beachrock may be loose one year and cemented the next; within a few years it may become sufficiently recemented so as to ring

62

R. W. FAIRBRIDGE

Fig.6~.Illustrating the slow dehydration of silica gel, flints are seen near the top of the Senonian chalk at Stevns Klint, Denmark. Note how the flints occur as concretions both parallel to bedding planes and also merge (continuously) into vertical joint planes; evidently the gel was still mobile at the time when the chalk was sufficiently dehydrated to develop joints.

to a hammer blow and yet may be sharply corroded by its intertidal exposure (REVELLE and EMERY,1957). RUTTEN (1957)felt that in due course, somewhere, contemporary flint formation will be discovered, though he admitted that the present (postglacial, post-orogenic) time is abnormal in the light of historical geology. On the other hand, the evidence presented above suggests rather that one can find flint nodules only by coring to some tens of meters or more. It does not follow from the evidence of the reworked flint layers in the Upper Cretaceous that the flints had only just been formed “penecontemporaneously”, as CAYEUX(1929) and RUTTEN(1957) argued; the chalk accumulated slowly, and the revival of an old anticline could cause the loss

PHASES OF DIAGENESIS AND AUTHIGENESIS

Fig.6b. Detail of F i g . 6 ~Pencil . scale

=

63

15 cm.

of 10-20 m of youthful sediment before the flint layers were reached. The chalk accumulation rate was between 0.1 and 1 mm per year. The level of flint diagenesis may represent a stage 10,000 to 100,000 years older than the contemporary sedimentation, and possibly as much as 106 years (see discussion under the timing of anadiagenesis, p.37). Authigenic feldspar The formation of authigenic feldspar is a phenomenon that has long been known, but little understood; for it is easy to assume that the feldspars are high temperature silicates. However, it now appears that it may be little more than a leaching and

64

R. W. FAIRBRIDGE

dehydration event, probably in the anadiagenetic phase. If one takes, for example, a hydromica (muscovite, or its degenerated form illite) it is not too difficult to forecast the loss of alumina and dehydration to form orthoclase; in simple empirical form:

KzO 3A1203 6SiO2 * 2H20

+ Kz0

A1203 6SiO-2

+ 2 H 2 0 + 2A1203

WEISS(1954) reported three horizons in the Ordovician shales of Minnesota which are now predominantly authigenic orthoclase but still show remnants of hydromica and montmorillonite, apparently due to feldspathization of a bentonitic volcanic ash. Even more common is authigenic albite which seems also to be usually derived from montmorillonite. In other words, the weathering hydrolysis of silicates is reversed. Some of this initial hydrolysis may even take place in the syndiagenetic phase, for the feldspars may not have been subjected to low pH conditions during mechanical weathering and transport in sea water having pH of 8 or more. In this way the marine realm tends to unify sediments and destroy traces of climatic extremes; arid zones (extremely cold or hot deserts) furnish unweathered feldspars, but if feldspars come to rest in a rich organic ooze on the ocean floor, the acid weathering and hydrolysis that was denied them on land can equally well be provided. Attention was first drawn to authigenic feldspars in the Cretaceous chalk of 1897, 1916), and they have been very generally reported the Paris Basin (CAYEUX, elsewhere in England (REYNOLDS,1929), India (SPENCER,1925) and North America (REED,1928; GOLDICH,1934; GRUNER and THIEL,1937;WILLMAN, 1942). An exceptional case was observed by DALY (1917) in Late Precambrian oolite of Glacier National Park, Montana, where up to 40% of the rock was replaced by authigenic orthoclase. Such massive feldspathization, however, is almost unique (BERG,1952). Generally, the crystals occur isolated (as in chalk and limestone), as overgrowths on existing crystals (usually in sandstones) or finely disseminated (as in shales). They may even replace fossils (STRINGHAM, 1940; VANSTRAATEN, 1948). An interesting aspect of this dehydration and recrystallization phenomenon is that adsorbed cations, including various trace elements, are shed during the reorganization. Thus rubidium, a key element in certain geochemical studies, is lost in the diagenesis of illite to orthoclase (HORSTMAN, 1957). Other familiar minerals, that are normally regarded as high temperature forms, are known to be also authigenic. Rutile, brookite and anatase may be derived from biotite (also from ilmenite, etc.; SUNand ALLEN,1957). Further, tourmaline and zircon are known in this authigenic form (BOSWELL, 1933; PETTIJOHN, 1957), but little seems to have been done about investigating their origin. Other examples, sphene, etc., were mentioned earlier.

PHASES OF DIAGENESIS AND AUTHIGENESIS

65

Natural chromatography

In a compacting sedimentary basin, there is a steady hydrostatic head which causes interstitial waters to rise upwards. These waters are for the most part connate, and hence, in a marine basin, essentially modified sea water. There may be some admixture of lacustrine or other fresh waters in a mixed marine-continental sequence; and in volcanic regions there is the possibility of some juvenile water being present, but such mixtures are regarded as quantitatively unimportant in most basins. All sedimentary formations are to some extent porous and additional permeability may be induced by diaclastic action. As the connate waters are progressively squeezed from the more deeply buried horizons they will pass vertically, and to some extent laterally (up-dip), into higher formations in which the chemical equilibrium may be completely different from that of the underlying sequences. Here one must consider the possibility of the operation of selective chemical filtering, that is a “selective adsorption process”. Laboratory use of this as a technique formerly involved the separation of colored substances in a fractionating column; hence the term “chromatography” developed and it is now an important standard method of chemical analysis. The concept of “natural chromatography” has been raised in connection with the evolution of petroleum, and has been applied to sedimentary basins in general by NAGY(1960). The passage of a liquid, heavily charged with organic and inorganic solutions and colloids, through a porous membrane is likely to cause a filtering of the larger molecules and ions of opposite charge to that of the membrane. Nagy devised a simple and adequate little experiment to demonstrate natural chromatography through a quartz sand. Inasmuch as this quartz carries no charge, the separation is mechanical, i.e., differential capillarity. In the case of clays, and most sediments carry at least some admixture of clays, the filtering can become chemical. This is true especially in the case of montmorillonite which has strong negative charges. BREDEHOEFT et al. (1963) suggested that at first the passage of the negatively charged anions will be mechanically restricted and then the corresponding cations (Ca2+, Mg2+, Na+ and K+) will be trapped. Deep basin brines may thus achieve a salinity up to six times that of ordinary sea water. CHAVE(1960) in studying these ancient brines, came to the conclusion that the variability in the nature of the sedimentary membranes was so great that a tremendous variety of brines would result, and that no deductions could be drawn about the salinity of the ancient oceans. The same wide variety is experienced with the petroleum hydrocarbons, exposed as they are to every phase of diagenesis (BREGER,1960; KREJCI-GRAF, 1963a,b). In some basins the solutions are nevertheless very weak, but the extended time factor may effectively permit reactions not otherwise easy to contemplate

66

R. W. FAIRBRIDGE

(IRVING,1892). The average sedimentary basin may accumulate over a period of 106 to 108 years, and by periodic revival (as noted in the section on Anadiagenesis) may obtain a multi-tiered structure. The rising solutions, apart from mechanical filtering, will lead to various reactions which, today, can be deduced in part from fabric studies. The most obvious are expressed by cementation and decementation phenomena. The cements are mainly CaCOs or SiOz, with Fez03 playing a role in some special environments. As outlined earlier, the interplay between high and neutral pH (say 9-7) in the connate waters of normal (alkaline) marine sediments results in calcite or quartz cements, The acid waters required by Fez03 are restricted to certain near-surface bacterial environments in the initial stage of syndiagenesis, or the acid oxygenated waters are introduced in the epidiagenetic stage. Cementation, involving complete silicification, dolomitization and dedolomitization introduce the special problem of sedimentary metasomatism. Diagenetic metasomatism The term metasomatism, coined by NAUMANN (1SSO), was applied essentially to the formation of pseudomorphs, either of the individual minerals or of whole rocks, involving a chemical replacement, atom by atom, but without change in form or volume, and obeying the “volume law” of W. Lindgren (see discussion by HOMER, 1947). This process can be defined as that of a low temperature enrichment of the sediment by new components “from without”, the original ions or molecules being removed in whole or in part. Essentially it is due to change in the chemistry of the enclosed or passing solutions. Metasomatism thus included several processes discussed already, such as oxidation-reduction, hydration-dehydration, etc., but it is convenient to consider under this heading the major ion exchange and replacement reactions. These normally obey the law of mass action and follow the entry of stronger solutions into the sediments. The usual site of such progressively increasing solution strengths is in compacting basins, thus normally during the anadiagenetic phase, though sometimes the epidiagenetic phase (e.g., phosphatization of limestone). Hypogene metasomatism may occur at high temperatures, as in some ore genesis, so that non-hydrothermal metasomatism should be designated as “diagenetic metasomatism”. Some examples of diagenetic metasomatism may be considered briefly as follows. Dolomitization Even after countless research investigations the understanding of this phenomenon is still beset with problems. Individual cases are often difficult to interpret because it occurs in all three phases of diagenesis:

(I) Syndiagenesis. Firstly there is so-called primary dolomite, which is not meta-

PHASES OF DIAGENESIS AND AUTHIGENESIS

67

somatic, but may provide material for later metasomatism. On a large scale it is only known in ancient rocks, by direct precipitation or under early diagenesis (see isotope data and discussion by DEGENS and EPSTEIN,1962). On a small scale syndiagenetic dolomites form today in isolated lakes and lagoons (SKINNER, 1963), in which the role of photosynthetical Algae play a major role in removing COz and thus raising the pH to the point of possible coprecipitation of calcite and magnesite, MgC03, probably mixed with brucite, Mg(0H)z (CHILINGAR, 1962; TEODOROVICH, 1955). Teodorovich remarked that magnesite is not found free; but hydromagnesite does appear, for example, in the South Australian lagoons. With progressive concentration the solubility curves of CaC03 and MgC03 intersect. There is an exceptionally slow ordering from a disordered lattice of “protodolomite” (GRAFand GOLDSMITH, 1956; PATERSON et al., 1963). In the early burial phase the high pH (over 9) is also aided by anaerobic bacterial action. The writer has personally observed a contemporary example of such an environment on the shores of the Persian Gulf, where in the shallow lagoons one may see vigorous contemporary growths of gypsum; dolomite has recently been identified on the floor and in the beach rock here (WELLS,1962). The synthetic coprecipitation of dolomite and gypsum was achieved more than a century ago by HAIDINGER (1848) and also by VON MORLOT(1847; see review by FAIRBRIDGE, 1957). This phenomenon was later observed by the Russian workers in the southern Caspian Sea. The intersection of the CaC03 and MgC03 solubility curves comes just about the point where CaS04 (but not MgS04) is in a saturated state (TEODOROVICH, 1955). A somewhat similar coprecipitation process can be synthetically reproduced, as demonstrated nearly a century ago by SORBY (1879) using a solution of MgClz (a greatly simplified “sea water”) aided by considerable heat and pressure. Indeed, a different MgClz/NaCl ratio would have far-reaching consequences in an ancient ocean (GARRELS et al., 1961). In the presence of COz and only 5 atm. pressure, Sorby’s experiment was repeated by LINCK (1909) thirty years after Sorby’s work, and again, recently, under rigorous modern conditions (BARON,1960), but this is still not quite the natural condition; the reaction was probably as follows: 2CaC03

+ MgClz + CaMg(C0s)z + CaC12

The elevated pressure of COz represents a problem in dolomite synthesis today, (1956, 1957, 1958) that in but there is some support for the theory of STRAKHOV Precambrian and Early Paleozoic times, the principal era of primary dolomites, the atmospheric pcoZ was somewhat higher than today (see also CHILINGAR, 1956~). The large-scale primary dolomites are not known to have formed since the Mesozoic. Also forming today during syndiagenesis is a metasomatic, secondary dolomite. This is initiated as nuclei of Mg in a disordered lattice of calcite, known as magnesium or magnesian calcite (CHAVE,1954; GOLDSMITH and GRAF,1958). It is normally due to organic secretion by various invertebrates and especially by

68

R. W. FAIRBRIDGE

certain varieties of calcareous Algae, the MgC03 reaching a maximum of 25-30 %; however, this is a metastable condition and within a period of time (101-104 years) is likely to invert to calcite, but in the presence of a high Mg2+ concentration dolomite will develop. Probably this is a true metasomatic replacement of Ca2f by MgZ+ (see, for example, photographs of calcareous Algae from an Eniwetok Atoll bore, with euhedral dolomite rhombs beginning to form in the middle; SCHLANGER, 1957). Numbers of syndiagenetic limestones and dolomites are so tightly crystallized that they still retain the fetid odor of decaying organic material, which itself suggests rather rapid diagenesis (LUCAS,1952). The presence of specks of iron oxides, pyrite, marcasite, or ankerite has often been noticed in dolomites; in fact, the Fe/Mg ratio is virtually constant in dolomite (COOPER,1954). This iron concentration seems to call for a reducing subsurface condition and reasonably vigorous benthos (MORETTI, 1957) not likely to be found under evaporite conditions. It seems likely that this metasomatism is not achieved without some elevation of pressure. Whereas it is observed on a massive scale at a depth in Funafuti Atoll corresponding to 20 atm. (SCHMALZ, 1956; FAIRBRIDGE, 1957, p.148), and in isolated crystals in the deep oceanic sediments (ZEN, 1960), such metasomatism is absent from modern stable shelf coral limestones, such as those of the Great Barrier Reef where the base does not exceed 7 atm. pressure today (FAIRBRIDGE, 1950a). Pressure is not everything, however, for the Bikini-Eniwetok bores are only partially affected. It is suspected that it is the primary distribution of the high magnesian calcite nuclei that plays the vital role. Inasmuch as many of the Paleozoic dolomites are alternately banded with limestones, and contain internal evidence (more terrigenous particles, etc.) of deposition closer to the shore than the limestones, it may be that the banding is also an indicator of greater participation by the magnesian-calcite-rich Algae. It would follow then that at least some of the ancient dolomites were metasomatic and quite distinct from those of the great primary precipitate sequences (FAIRBRIDGE, 1957). On the other hand, the banding is attributed by SARIN(1962) to a periodic killing of Algae in the lime facies by invasions of supersaline dolomitic facies. ( 2 ) Anadiagenetic dolomite. Anadiagenetic dolomite is by definition a secondary dolomite, but belonging to a later stage than those listed above. In any metasomatic replacement of this sort a high concentration of Mg2+is a prerequisite. It has been suggested by several workers recently that a refluxing of connate waters from the saline horizons would lead to an ideal setting (ILLING,1959; ADAMSand RHODES, 1960). The hydrodynamic motion can be lateral as well as vertical. In the former, the flux could be from a saline lagoon bounded by a barrier reef; in the Paleozoic sections the barrier reef limestones are almost always dolomitized. MENNIG and VATAN(1959) found a similar barrier reef correlation in the Devonian of the

PHASES OF DIAGENESIS AND AUTHIGENESIS

69

Ardennes; they spoke of “epigenesis”, but the writer suspects that this is merely a different use of the term. Progressivedolomitizationresults in ca. 12 %loss of volume, so that fracturing is favored. Where the results of tectonism have been added to those of simple compaction, there are extensive fault zones and joints which permit more thorough access by these rising waters. On either side of a fault zone there is often a “Christmas-tree” effect produced by the rise and lateral spread (along more porous zones) of such waters (ANONYMOUS, 1959.). Across the Paleozoic of the Midwest it is often noticed that dolomitization is more prominent as one approaches the tectonic belts. (3) Epidiagenetic dolomite, Epidiagenetic dolomite is the superficial phase of secondary dolomitic metasomatism. Such strictly epigene dolomites are not nearly so common as the anadiagenetic ones, for the simple reason that descending waters are generally weak solutions and require some sort of enrichment. This may be achieved near the lower boundary of the soil profile (and here epidiagenesis impinges on weathering), resulting in the formation of an unconformable, dolomitized hard-pan horizon. Such a horizon has developed, for example, in the Eifel district, Germany, where folded Devonian dolomites and limestones are now capped by this epidiagenetic dolomite, without any respect for what was the original lithology (FAIRBRIDGE, 1957, fig.13). It is thus sometimes known as “subsequent dolomitization” (HATCHet al., 1938, p. 193). Dedolomitization Dedolomitization is generally a metasomatic replacement of dolomite by calcite; but it may also result simply from leaching, with a molecule by molecule solution followed by a selective removal of the MgCOs. The result is a loose sandy textured calcite. It seems likely that this phenomenon goes hand in hand with the production of the epidiagenetic dolomite noted above. Where dolomite and gypsum have been interbedded, the dolomite is often fractured and etched (mechanical collapse and solution). VON MORLOT(1847), who first applied the term dedolomitization, suggested the following reaction:

CaMg(CO3))a

+ CaS04 - 2H20 f2CaC03 + MgS04 + 2Hz0

SANDER (1951) has stated that in the Northern Limestone Alps the calcitization of dolomites is really more important today than the original dolomitization. SHEARMAN et al. (1961) have recently described the dedolomitization textures in the French Jura. Hydrothermal solutions may also play a role in dedolomitization (FAUST,1949). A second reaction, possible where MgS04 dominates the brines, may be given as: CaMg(CO3)z

+ MgS04 + CaS04 + 2MgC03

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R. W. FAIRBRIDGE

In this reaction anhydrite is the remaining solid phase; crystals of anhydrite have been observed, in the Permo-Triassic evaporites of England, “eating” into dolomite (STEWART, 1949). Silicification Inasmuch as silica is soluble at normal temperatures only in waters having pH of over 9 (CORRENS, 1949; KRAUSKOPF, 1959), it is rarely mobilized after dehydration of the primary gels in the syn- and early anadiagenetic phases (see Fig. 7). Even low concentrations of dissolved Si02 will be precipitated at the low pH (3-6) existing in some lakes, leading to silicification of wood, as C02 is liberated by 1949), but such conditions are rarely achieved in the bacterial action (CORRENS, ocean. With progressive compaction and concentration of the connate waters during advanced anadiagenesis, however, a very high pH may be obtained. According to calculations by SIEVER(1957b), equilibrium with the CaC03 is reached at pH of 9.8 at 25”C, particularly in a thin film of water containing originally atmospheric COZbetween grains of quartz and calcite. At this point silica dissolves to the limit of the very restricted water film. Pitting (etching) of quartz grains was also observed (DAPPLES, 1959). In quartz sands the usual products of anadiagenesis are eventually quartz overgrowths and quartz cements, which result in quartzose sandstones (or “orthoquartzites”). According to KRAUSKOPF (1959) the first dissociation constant of H4Si04 occurs at pH of 9.8 and, therefore, the often quoted solubility curve of C . W. Correns (in: BARTHet al., 1939, p.129) may be misleading. SIEVER (1959) pointed out, however, that at pH of 9.8 partial dissociaiz 10

4

2

0

0

2

4

6

8

10

PH

Fig.7. Solubility of silica gel, aluminium hydroxide and ferric iron hydroxide in relationship to pH. (After CORRENS, 1939.) Broken line indicates a modified SiOz solubility, following the work of KRAUSKOPP (1 959).

PHASES OF DIAGENESIS AND AUTHIGENESIS

71

tion of H4Si04 produces silicate ions and essentially doubles the quartz solubility at lower pH values. An additional factor is the increased polymerization rate produced by dissolved electrolytes, as found in many deep connate waters. Precipitation, therefore, varies with the ionic strength of the ground water, and solution is particularly favored by rising temperatures (SIEVER,1959), as well as by associations which raise the pH. Hydrothermal waters may assist such temperature rise, for silicification is a marked feature of deep-seated faults. When the silica is removed, molecule by molecule by solutions of high pH, it will tend to be replaced by the least soluble components in such solutions at the given pH. These are generally the carbonates, this progressive metasomatism of chert is described by WALKER (1962), who has further demonstrated multiple replacement reversals. This silicification-desilicification sequence seems to be explicable only by proposing an alternating passage of waters of higher pH (over 9) and lower pH (under 9). Such conditions can be visualized during the gradual dewatering of a compacting basin during anadiagenesis, when successively new connate water sources are liberated from their cement traps by jointing and faulting as the compactional and diastrophic evolution progresses. The same phenomenon may continue during epidiagenesis, but with additional sources of water of lower pH. It is under these conditions that pyrite and marcasite are most liable to become oxidized, sharply dropping the pH and raising the temperature. Silica solubility rises with higher temperature, whereas the reverse is true of calcite (OKAMOTU et al., 1957), and thus quartz-calcite intergrowths occur (DAPPLES, 1959; WALKER,1962). As epidiagenesis proceeds, there is a gradual tendency towards stabilization in the intermediate pH range, where calcite is still slightly soluble and silica completely stable. Thus on old land masses, long-stable tectonically, there is a widespread silicification of limestones. Where, as in Africa, Australia, Peninsular India, and the Brazilian Shield, the climatic record of the last 108 years or so has been tropical or subtropical, there have been immense and almost continuous supplies of Si02 passing into the vadose water system by leaching down from laterized soils. Under such conditions the epidiagenetic stage is one of massive silicification. Solid limestones often become completely silicified down to some hundreds of meters below the weathering zone (FAIRBRIDGE,1950b). It is curious that the Collenia type calcareous Algae of the Precambrian in those regions long went unnoticed because they were found only in what appeared to be primary quartzites. Far-reaching silicification is known in certain coal formations. Whereas normally the low pH associated with the syndiagenesis of coal would inhibit such silicification, a late anadiagenesis with rising waters of high pH would favor it. In heavily folded parts of the Ruhr Carboniferous, nodules, lenses and massive replacements by silica are observed (HOEHNE,1957).

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An interesting phenomenon that has only recently received notice is what in the Paris Basin is called “meulerization” from the French word for millstone (TERMIER and TERMER,1963, p.345). It is believed to be analogous to the formation of siliceous hard pan or “silcrete” and is observed in progress in South Africa and the northern Sahara today (ALIMEN, 1958). Carbonate sands or limestones are locally cemented and partly replaced by opaline or chalcedonic silica; exceptionally gypsum may also be replaced. An opposite reaction is sometimesseenin the epidiageneticphase in temperate latitudes. This has been called “Fontainebleau sandstone crystallization”, to describe the local formation of a calcite cement in a desilicified quartz sandstone (or uncemented quartz sand); the cement is continuous and fracture surfaces of the rock show (by their reflection) that the calcite is in crystallographiccontinuity. Like meulerization the phenomenon was first described from the Paris Basin. One may suggest that this has been a late diagenetic phenomenon dating from the late periglacial (cold-wet) phase when acid ground waters were generated by podzolization. Striking examples of desilicification may be observed in quartz sand-clay mixtures, such as often found in graywackes (see Fig.8). Thin sections may disclose the growth of newly formed illite, eating into the quartz grains.The same phenomenon is associated with stylolite formation. This is unlikely to be simply pressure solution as suggested by many authors (e.g., FAIRBAIRN, 1950; HEALD,1955), but is greatly facilitated by a minor clay fraction (THOMSON, 1959). Stylolites in sandstones or quartzites have long been an enigma, but Thomson showed that a clay layer between sand grains, if subjected to COz-rich waters would liberate KzC03. The latter is a strong alkali, which would mobilize SiO2 at the clay contact, only, to be reprecipitated nearby in the generally acid solution. Phosphatization Another example of diagenetic metasomatism that has provided geologists with considerable problems has been the phenomenon of phosphorization. Generally, phosphatic acids replace the carbonates in limestone. Phosphorus is usually introduced (in reactable form) into the sedimentary cycle by organic agencies. It is present in nucleic acids formed in all living matter. It is contained in many proteins (phosphoproteins); likewise in many lipids and carbohydrates. Its abundance is rather low in sea water (about 0.07 mg/l), so that the limiting factor in phosphatization is the local concentration by organic metabolism. Organisms employ inorganic phosphate to synthesize ADP (the diphosphate) or ATP (triphosphate), which in turn provide fundamental organic energy sources. From almost all periods of geological time, the sedimentary phosphates are of marine origin; and although the secondary terrestrial concentration often occurs, the initial segregation is marine. Certain periods seem to be more favorable

PHASES OF DIAGENESIS AND AUTHIGENESIS

73

Fig.8. Desilicification of quartz sand grains, aided by the base exchange of intercalated clays (from T H O ~ N 1959), , in four stages: A. Nature of quartz and clay shortly after deposition. Clay is between some grains and not others. B. Base exchange of K+ by Caa+and Mgz+begins along edges of sheets. C. Exchange proceeds rapidly along cleavage surfaces. KaCOa is formed and a higher pH develops in clay-rich regions. D. As pressures increase, Si4+ dissolved in regions of high pH migrates to regions of low pH and precipitates.

than others, e.g., Cambrian, Permian, Upper Jurassic, Lower Cretaceous, Upper 1956); these are essentially transgressive, Cretaceous, and Tertiary (GIMMELFARFI, thalassocratic stages. Such stages are associated with the expansion of broad epeiric shelves, which are the preferred environments for photosynthetic Algae; among these the Chlorophyceae (green Algae) are the principal organisms to accumulate calcium and phosphorus (DEMOLON and BOISCHOT, 1948). The open shelf phosphatites are generally nodular or concretionary, and are often associated with coprolites and glauconite, the geochemical evolution of which gives a clue to the phosphatizing environment (VISSE, 1953; RIVIERB and VISSE, 1954). In subsiding basins marked by a higher accumulation rate and probably by somewhat greater depths and poor circulation, the phosphatites tend to be bedded and dark, even black in color. Here they are commonly accompanied by pyrite, reflecting the reducing environment of the syndiagenetic phase; the pyrite in turn is often oxidized to gypsum (anadiagenetic). There is not much data available concerning phosphates in contemporary

14

R. W.FAIRBRIDGE

marine sediments, but in the Caspian Sea, a truncated (formerly) marine gulf of somewhat restricted nature, core studies show that the phosphate concentration drops markedly with depth of sediment. In the Bering Sea, an open basin, however, it rises gradually going down a 16 m core (see CHILINGAR, 1958); the concentration is five times greater in the muds than in sands or silts. Thus the stage for large-scale phosphatization is preset in geological time by certain geotectonic-paleogeographic events. The present time is not one such stage in history. However, the extreme and large-scaled oscillations of sea level during the Quaternary favored a special epidiageneticenvironment that is found in isolated atolls. These offered sites for sea bird sanctuaries, free from predators until recently introduced by man. Here concentrations of phosphate-rich guano accumulated. During high sea level stages, the islands were small and leaching of the soluble organic phosphates by rain water brought them quickly back into the sea. However, during the low sea level episodes, the atolls became emerged and resembled mediaeval castles with limestone walls and floors. This as ideal setting for leaching of phosphoric acids into freshly formed limestone (often composed of porous metastable aragonitic corals), to form, as a rule, collophane (collophanite), Ca3PzOs * HzO, or dahllite, Cas(PO4)4 * CaC03 - +HzO. On small volcanic islands with a barrier or fringing reef, thus incomplete atolls (e.g., Navassa Island, West Indies), or along semi-arid mainland coasts (e.g., southwestern Africa, northern Chile), the same thing may happen, but in such cases, the igneous weathering products may interact during the diagenesis, resulting in viviunite, FesP208 * 8H20, or wuvellite, 4NP04 * 2M(OH)3 * 9Hz0. The geochemistry of phosphatization of atolls suggests a low pH in a freshwater saturated environment, without any unusual temperature or pressure requirements. TEODOROVICH (1954) believed that weakly reducing or neutral environments favor phosphatization. The widespread stratified phosphatites of the past, being so largely marine, might suggest a different setting, but it seems likely that it is the marine condition that is essential only for the organic segregation, which, if followed by a small negative eustatic swing over a shallow shelf region, would lay it bare for the phosphorous mobilization under rain-water leaching. Such shelves are equally favorable for the accumulation of marine carbonate sediments, which provide the necessary “host” rocks. On warm shelves the bulk of the host material is likely to be in the form of aragonite which is particularly soluble in fresh water of only slightly reduced pH (6.5-7.5). Hence such phosphatization may be a good paleotemperature indicator. Such processes on exposed atolls or shelves would both be examples of epidiagenesis, where the syn- and anadiagenetic stages have been rapidly bypassed by a sudden drop of sea level. The duration of such negative eustatic stages during geological history may have been of the order of 5,000 to 50,000 years. The latter thus gives the upper limit of the time required for phosphatization of quite thick formations.

PHASES OF DIAGENESIS AND AUTHIGENBSIS

75

In some atolls, dolomitization has preceded phosphatization, probably during an earlier cycle involving subsidence and re-emergence. Whereas some islands have become phosphatized under eustatic control alone, the majority show also tectonic movement as well. This seems essential for deep phosphatization. A paragenetic relationship between phosphatization and the accumulation of concretionary silica in the Permian Phosphorite Formation of western U.S.A. has been suggested by MCKELVEY et al. (1953). Detailed studies in Morocco, however, show that while bedded cherts there commonly succeed phosphatic sands in the sequence, this does not imply an interrelationship (SALVAN,1955). Sideritization Sideritizadon is a fifth type of sedimentary metasomatic diagenesis. Primary precipitation as siderite, FeCOs, probably does not occur in open marine environments today, but in swamps and other restricted basins the ferric oxide hydrosols in river waters would be reduced to the ferrous state, and removal of COa by photosynthesis of plants would dissociate the bicarbonate ions to cause the direct precipitation of FeC03. Sideritic iron ores are thus associated with some coal deposits. TEODOROVICH (1949,1961) recognized a distinct “siderite facies” of marine environments of the past, which he noted were marked by mildly reducing conditions and by strong fluctuations of pH and Eh, because he believed that it belonged to a sublittoral zone. Indeed, in this position it is a favorable indicator for petroleum. Much more usual, however, seems to be the condition where marine limestones are metasomatically replaced by siderite during syndiagenesis (CAYBUX, 1916; HATCH et al., 1938, p.135). The CaC03 is often in a highly porous state, such as oolites or organogenic calcarenites (often crinoidal), and may also be aragonitic. Thus a pervious and metastable “host” is provided, just as is necessary for dolomitization and phosphatization. In the same way that calcitic fossils (brachiopods, Bryozoa and certain molluscs) in a matrix of carbonate material, believed to have been originally aragonitic or, equally well, metastable organogenic high-magnesian calcite, are often “spared” by the dolomitization, so too they are often found preserved in a matrix totally replaced by siderite. This fact alone is strong evidence for a syndiagenetic origin of such siderites, although they are commonly stated to be epigenetic (e.g., TWENHOFEL, 1950, p.431). On entering the ocean, ferric oxide hydrosols are electrolyticallyflocculated and thus deposition is likely to occur on continental shelves rather than in the 1929, p.507). On semi-tropical sheIves one may deep sea (MOOREand MAYNARD, thus expect the ideal environment. Iron is leached readily out of the laterite sails on the land and brought down to the oceans by rivers. In the warm ocean there is steady concentration of organogenic carbonates, largely in metastable form. It is not surprising, therefore, that the great sedimentary iron ores are, like the

R. W. FAIRBRIDGE TABLE II MORE COMMON AUTHIGENIC MINERALS (EXCLUDING

Mineral

Formula

HALIDES)~ Frequency2 Usual developmenr synMaepidiagenetic diagenetic diagenetic

Ti08 (Tetr.)

R C R Ca(Mg, Fe) (C0s)z Ankerite CaCOs (Orth.) C Aragonite C~s(cOs)z(OH)a R Azurite Bas04 C Barite CusFeS4 R Bomite R TiOa (Orth.) Brookite C CaCO3 (Rho.) Calcite SrSO4 C celestite PbCOs R Cerussite SiOa C Chalcedony R CuFeSs Chalcopyrite Chamosite [(Fe, Mg)Olia[~~Osls[SiOzl~i [HzOlia R C chlorite group C CasPaOe * Hz0 Collophane C Calo(POq)s(COs) HzO Dahllite CaMg(C0s)a C Dolomite R PbS Galena Glauconite Kz(Mg, Fe)&e(Si4010>a(OH)iz C FeaShOlo(0H)e R Greenalite CaS04 2HzO C Gypsum NaCl C Halite FeaOa C Hematite Hydromagnesite 4MgCOs Mg(OH)3 4HzO R Illite (K, Na) (Al, Mg, Fe, Li)e (AlSiaOio) (0H)z C KaOliIlite AlzSiaOs(OH)4 C Leucoxene Ti08 complex C Limonite FeO(0H) n(Hz0) C FeaOs n(Hz0) R Magnesite MgCos Malachite CuaCOs(0H)z R Marcasite FeSz (Orth.) C Montmorillonite (Al,Mg)s(Si4010)s(OH)io12HzOC Muscovite ICAlsSiaOio(0H)a C Natron NaZCOs * lOH20 R Nesquehonite MgCOs * 3Hzo R Opal SiOa * nHz0 C OrthOClaSe KAlSiaOtJ C Phillipsite (Ca, Ba, K., Na)&(Al, Si)z Si10040 15-2OHaO R Anatase

Anhydrite

-

- -

-

1 2

X

c&o4

X

X X X X

X X

X X X

X X X

X

X X

X X X X

X X X X

X X X

X X X X

X X

X X X

X X

X X

X

X X

X

+

-

X

X X X X X

-

X X X X

-

X X

X X

X X

In part, after TWENHOPEL (1950, p.288), and TEOWROVICH (1961). R = rare; C = common.

X X

77

PHASES OF DIAGENESIS AND AUTHIGENESIS

TABLE I1 (continued) Mineral

Formula

Frequency2 Usual development synanaepidiagenetic diagenetic diagenetic

Plagioclase Psilomelane (Wad) Pyrite Pyrolusite Quartz Rhodochrosite Rutile Siderite Sphalerite Strontianite Sulphur Tourmaline

(Ca, Na) (Al,Si)AlSizOs

C

(Ba, HzO) Mn6010 FeSz (Iso) MnOz SiOz MnCO3 TiOa (Tetr.) FeCOa ZnS SrC03 S Na(Mg, Fe)aAle(BOs)aSieO~s

C C C C

(om4

Witherite BaCOs Zeolites (Phillipsite, Heulandite, Laumontite, Chabazite, Natrolite, Analcime) Zircon ZrSiO4

X X X

X X X X X

X

R C R R R

X X

X X

R R C R

X

X X

X X

X X

phosphates, normally concentrated in the more transgressive marine stages (Ordovician, Silurian, Jurassic, Lower Cretaceous). The liberation of large quantities of iron in the periods immediately preceding orogenic episodes was stressed by Cayeux, but the role of soil cycles in concentrating the iron should not be forgotten (ERHART, 1956). Siderite is also found as a scattered authigenic mineral in some formations, but this is rather unusual. The sideritic limestones are often associated with hematite Fez03 and with chamosite, 3Fe0 Ah03 2SiOa HzO. These seem to be mostly primary, though possibly non-marine. In the Wabana (Newfoundland) deposits, the hematite-chamosite oolites appear to be cut through by algal borings, whereas the siderite clearly resulted from diagenetic replacement (HAYES,1915). Equally striking is the syndiagenetic pyritization of the Cleveland chamositeoolite (Lower Jurassic of England), and the incorporation of penecontemporaneous pebbles of pyritized ironstone in the same formation (HATCHet al., 1938, p.135). Moo= and MAYNARD (1929) have explained, and demonstrated experimentally, how ferric oxide and silica can be alternately precipitated from sea water. It is puzzling, however, that the great banded iron ores (FzOs/SiOa) are only Precambrian. They are known from all over the world, but are present only in the older rocks. HOUGH(1958) has suggested that this alternation is a seasonal

-

-

78

R. W. FAIRBRIDGE

phenomenon in giant lakes, and it seems not improbable that the Precambrian sea had a salinity somewhat less than the present (and a lower pH) so that the ocean of that time might be compared geochemically with a brackish lake (FAIRBRIDGE, 1967). Siderites are also known in these deposits. It seems highly unlikely that such a lacustrine condition would favor sideritization of aragonitic sediments. Instead, the primary precipitation of cherty siderite probably occurred under local photosynthetic removal of C02 and elevation of pH. The deposits were modified in some cases by later hydrothermal action. Under what are probably epidiagenetic conditions, some dedolomitization occurs and dolomite is replaced by calcite. Similarly siderite is found to be replaced by calcite in the Jurassic iron ores, forming very similar textures (note by J. H. TAYLOR in: SHEARMAN et al., 1961, p.12). Thus one might speak of desideritization (see Table 11).

CONCLUSIONS

As discussed in the present chapter a distinct c‘stratificatio”’ of diagenetic events takes place in the evolution of a sedimentary basin. These events or stages may be interrupted or bypassed, but in the slow development of a “typical” basin three major categories are recognized. Each phase is indicated by certain types of authigenic minerals (many of which need further study) and each phase may be further subdivided. The most important criterion for such subdivision is the pH-Eh balance of the enclosed liquids. Evidently a considerable amount of research is still needed in the area of mineral stability in natural waters containing numerous ions and complexes in solution. Economic geologists have long recognized the concepts of syngenetic and epigenetic ore implacement; and metamorphic geologists recognize an anamorphic stage (deep burial with rock flowage). The three phases of diagenesis identified here have utilized these valid concepts (but without implications of synonymy) in the erection of a systematic terminology, thus: ( I ) Syndiagenesis is the early sedimentary condition that is instituted at the very moment a sedimentary particle comes to rest on the basin floor. Alteration begins even earlier through halmyrolysis, which involves reactions between the mineral particles and sea water, during their transportation and descent, and continues after burial. Two most common subphases of syndiagenesis were (1960): (a) initial phase, marked by oxidizing conditions; identified by GARREIS and (b) early burialphase, marked by reducing conditions. Under special conditions of isolation the Eh = 0 boundary may be at or above the sediment-water interface, as in euxinic facies, but these are less common; from the authigenic mineral characteristics, TEODOROVICH (1954) recognized no less then six subphases. In this phase ionic motion is both upwards and downwards, mainly by diffusion.

PHASES OF DIAGENESIS AND AUTHIGENESIS

79

(2) Anudiagenesis is the deep burial stage, during which sedimentary basins are progressively deepened and loaded. Rock flowage (anamorphism) is ultimately approached. However, before this stage water is expelled and rises upwards and outwards, the geochemical reactions leading to a “natural chromatography”, which is particularly important economically in the case of petroleum, as was demonstrated by NAGY(1960). With subsidence and subsequent tectonism, various diastrophic structures, joints, fractures and folds, distort and penetrate the sediments. This then permits recycling of solutions and some of the early anamorphic stages of metamorphism. In general in this phase the motion of the liquids is upward or lateral (toward basin margins). (3) Epidiagenesis is the final emergent stage when connate solutions become modified by deep-working ground waters (meteoric or vadose waters) and oxidizing conditions are re-introduced. There is an upper limit to this phase, however, recognized as the level where weathering processes take over and active rock disintegration begins. In this phase the fluid motion is downward or toward basin centers. At each stage distinctive suites of authigenic minerals form, but they may be partly effaced during succeeding events. Some attempts have been made to set rough time limits on these events, some being quite brief and episodic in nature, and others enduring for a very long time. Broad fields of research in this area remain open and inviting.

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61 :45-60. CLOUD JR., P . E., 1955. Physical limits of glauconite formation. Bull. Am. Assoc. Petrol. Geologists, 39: 484-492. CONLEY,R. F. and BUNDY,W. M., 1958. Mechanism of gypsification. Geochim. Cosmochim. Acta, 1557-72. CONWAY,E. J., 1945. Mean losses of Na, Ca, etc. in one weathering cycle potassium removal from the ocean. Am. J. Sci., 243: 583-605. COOMBS, D. S., ELLIS,A. J., F m , W. S. and TAnoR, A. M., 1959. The zeolite facies, with comments on the interpretation of hydrothermal syntheses. Geochim. Cosmochim. Actu, 17: 53-107. COOPER, B. N., 1954. Fundamental problems of genesis of Appalachian dolomites. Virginia J. Sci., 5: 301-302 (abstract). CORRENS,C. W., 1949. Einfiihrung in die Mineralogie. Springer, Berlin, 414 pp. CORRENS,C. W., 1950, Zur Geochimie der Diagenese. Geochim. Cosmochim. Acta, 1: 49-54. CORRENS, C. W., 1963. Experiments on the decomposition of silicates and discussion of chemical weathering. Proc. Natl. Conf , Clays Clay Minerals, IOth-Natl. Acad. Sci. Natl. Res. Council, Publ., pp.443-459. C~RRENS, C. W. und VONENGELHARDT, W., 1938. Neue Untersuchungen uber der Verwitterung des Kalifeldspates. Chem. Erde, 12: 1-22. DALY,R. A., 1917. Low-temperature formation of alkaline feldspars in limestones. Proc. Natl. Acad. Sci. US.,3: 659-665. DANOEARD, L. et RIOIJLT.M., 1961. Observations nouvelles sur les accidents silicieux situb au sommet de la “Pierre de Caen”. Bull. SOC.Gkool. France, 3: 329-337. DAPPLES,E. C., 1959. The behavior of silica in diagenesis. In: Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 36-54. DAPPLES, E. C., 1962. Stages of diagenesis in the development of sandstones. Bull. Geol. SOC. Am., 73: 913-934. DEANS,T., 1950. The Kupferschiefer and the associated lead-zincmineralization in the Permian of Silesia, Germany and England. Intern. Geol. Congr., 18th, London, 1948, Rept., 7: 340-352. DEBELMAS, J., 1959. Une curieuse contribution il l’ktude de la penbe des silex. Trav. Lab. Gkol. Fac. Sci. Univ. Grenoble, 35: 135-136. DEBYSER, J., 1952. Variation du pH dans I’kpaisseur vase fluvio-marine. Compt. Rend., 234: 741-743. DEFLANDRE, G., 1936. Les flagelks fossiles. Acta Sci. Ind., 335: 98 pp. DEGENS, E. J. and EPSTEIN, S., 1962. Relationship between 180/160 ratios in coexistingcarbonates, cherts, and diatomites. Bull. Am. Assoc. Petrol. Geologists, 46: 534-542. DEMOLON, A. et BOISCHOT, P., 1948. Observations sur le cycle du phosphore dans la biosphere. Compt. Rend., 227: 655-656. DUNHAM, K. C., 1952. Age relations of the epigenetic mineral deposits of Britain. Trans. Geol. SOC.Glasgow, 21: 395. EDWARDS, A. B. and BAKER,G., 1951. Some occurrences of supergene iron sulphides in relation to their environments of deposition. J. Sediment. Petrol., 21: 34-46,

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Gallup, New Mexico. J. Sediment. Petrol., 27: 265-270. TALIAFERRO, N. L., 1934. Contraction phenomena in cherts. Bull. Geol. SOC.Am., 45: 189-232. TAL~RRO N., L., 1935. Some properties of opal. Am. J. Sci., 30: 450-474. TAMM,O., 1925. Experimental studies on chemical processes in the formation of glacial clay. Sveriges Geol. Undersokn., Arsbok, 18(5): 1-20. TARR,W. A., 1927. Alternative deposition of pyrite, marcasite, and possibly melnikovite. Am. Mineralogist, 12: 417422. TEODOROVICH, G. I., 1949. Siderite geochemical facies of seas and saline waters in general as oilproducing. Dokl. Akad. Nauk. S.S.S.R., 69: 227-230. TEODOROVICH, G. I., 1954. Towards the question of studying oil-producing formations (source rocks). Bull. Moscow Nut. Hist. Soc., 29: 59-66. TEODOROVICH, G. I., 1955. A contribution on the origin of limestones and dolomites. Trans. Petrol. Inst., Acad. Sci. U.S.S.R., 5. English transl. in Intern. Geol. Rev., l(3): 50-74. TEODOROVICH, G. I., 1961. Authigenic Minerals in Sedimentary Rocks. Consultants Bureau, New York, N.Y., 120 pp. TERMER, H. and TERMER, G., 1963. Erosion and Sedimentation. Van Nostrand, London, 433 pp. TEN, M., 1961. La Vendbe Littorale. ThBse, Centre Natl. Rech. Sci., Paris, 578 pp. TESTER,A. C. and ATWATER, G. I., 1934. The occurrence of authigenic feldspars in sediments. J. Sediment, Petrol., 4: 23-31. THODE, H. G., HARRISON, A. G., and MONSTER, J., 1960. Sulphur isotope fractionation in early diagenesis of recent sediments of northeast Venezuela. Bull. Am. Assoc. Petrol. Geologists, 44: 1809-1817. THOMSON, A,, 1959. Pressure solution and porosity. In: Silica in Sediments-Soe. Econ.Paleontologists Mineralogists, Spec. Publ., 7: 92-1 10. TUGARINOV, A. I. and VINOGRADOV,A. P., 1961. Geochronology of the Precambrian. Geochemistry (U.S.S.R.) (English Transl.), 1961 (9): 789-800. TURNER,F. J. and VERHOOGEN, J., 1960. Igneous and Metamorphic Petrology, 2nd ed. MCGrawHill, New York, N.Y., 694 pp. "HOPEL, W. H., 1932. Treatise on Sedimentation. William and Wilkins, Baltimore, Md., 926 pp. TWENHOFEL, W, H., 1942. The rate of deposition of sediments: a major factor connected with alteration of sediments after deposition. J. Sediment. Petrol., 12: 99-1 10. TWENHOFEL, W. H., 1950. Principles of Sedimentation, 2nd ed. McGraw-Hill, New York, N.Y., 673 pp. VANANDEL,T. and POSTMA, H., 1954. Recent sediments of the Gulf of Paria. Verhandel.Koninkl. Ned. Akad. Wetenschap., Afd. Natuurk., Sect. I, ?0(5): 288 pp. VAN HISE,C. R., 1898. Metamorphism of rocks and rock flowage. Bull. Geol. Soc. Am., 9: 269-328. VANHISE,C. R., 1904. A treatise on metamorphism. U.S., Geol. Surv., Monograph, 47, 1286 pp. VAN HOUTEN,F. B., 1961. Climatic significance of red-beds. In: A. E. M. NAIRN (Editor), Descriptive Palaeoclimatology. Interscience, New York, N.Y., pp. 89-139. VAN STRAATEN, L. M. J. U., 1948. Note on the occurrence of authigenic feldspars in nonmetamorphic sediments. Am. J. Sci., 246: 569-572. VISE, L. D., 1953. Les facies phosphates. Rev. Inst. Franc. Pktrole Combust. Liquides, 8: 87-99. VON GUEMBEL, C. W. 1868. GeognostischeBeschreibung des ostbayrischen Grenzgebirges, Gotha, 968 pp. VON MORLOT, A., 1847. ober die Dolomit und seine kiinstliche Darstellung aus Kalkstein. Haidinger Nat. Abt., 1: 305. S. A., 1933. On the distribution of organic matter on the sea bottom and the chemical WAKSMAN, nature and origin of marine humus. Soil Sci., 36: 125-147. WALKER,T. E., 1962. Reversible nature of chert-carbonate replacement in sedimentary rocks. Bull. Geol. SOC.Am., 73: 237-242. J., 1894. Einleitung in die Geologie a h historische Wissenschaft.Fischer, Jena, 1055 pp. WALTHER, C. E., 1958. Geologic interpretation of argillaceous sediments. Bull. Am. Assoc. Petrol. WEAVER, Geologists, 42: 272-309.

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WEAVER, C. E., 1960. Possible uses of clay minerals in search for oil. Bull. Am. Assoc. Petrol. Geologists, 44: 1505-1 51 8. WEISS,M. P., 1954. Feldspathized shales from Minnesota. J. Sediment. Petrol., 24: 270-274. WEISS,M. P., 1958. Corrosion zones: a modified hypothesis of their origin. J. Sediment. Petrol., 28: 486-489. WELLS,A. J., 1962. Recent dolomite in the Persian Gulf. Nature, 194: 274-275. WEJR., E. G., 1961. Glauconite in early Tertiary sediments of Gulf coastal province. 3uB. Am. Assoc. Petrol. Geologists, 45: 1667-1696. O., 1933. Die in organischer Substanz erhaltenen Mikrofossilien des baltischen KreideWETZEL, feuersteins. Paleontographica, 47: 146-188. WETZEL,W., 1955. Seltene Metallverbindungen in Sedimenten. Geol. Rundschau, 43(2): 464469. WHITE,D. E., 1957. Magmatic, connate and metamorphic waters. Bull. Geol. SOC.Am., 68: 1659-1682. WICKMAN, F. E., 1944. Some notes on the geochemistry of the elements in sedimentary rocks. Arkiv. Kemi, Mineral. Geol., 19(B): 1. WILLIAMS, D., 1960. Genesis of sulphide ores. Proc. Geologists’ Assoc. (Engl.), 71: 245-284. WILLMAN,H.B., 1942. Feldspar in Illinois sands. Illinois State Geol. Surv., Rept. Invest., 79: 87 pp. ZEN,E-AN, 1960. Carbonate equilibria in the open ocean and their bearing on the interpretation of ancient carbonate rocks. Geochirn. Cosmochim. Acta, 18: 57-71. ZOBELL,C. E., 1942. Changes produced by microorganisms in sediments after deposition. J . Sediment. Petrol., 12: 127-136. ZOBELL,C. E., 1946. Studies on redox potential of marine sediments. Bull. Am. Assoc. Petrol. Geologists, 30: 477-513.

GLOSSARY Anudiagenesis (adj.-etic): Lithification or other modification of sediments during deep burial, marked by expulsion and upward migration of connate waters and other fluids (petroleum, etc.), often marked by high pH and low Eh. (FAIRBRIDGE, this chapter.) Anamorphism (adj.-ic): Metamorphism at depth, forming more complex minerals. (VANHISE, 1898; modified to exclude low temperature alteration, i.e., by diagenesis.) Authigenesis (adj.-ic, -ow): Formation of new sedimentary minerals in situ, within the enclosing 1957). sediment, during or after deposition (e.g., PETTIJOHN, Diagenesis (adj.-etic): Physical and chemical changes which a sediment undergoes after deposition and during lithification, without introduction of heat (over ca. 300°C) or great pressure (ca. 1,OOO bars). (VONGUEMBEL, 1868; modified slightly by WALTHER, 1894; FAIRBRIDGE, thischapter) Epidiagenesis (adj.-etic): Lithification or other modification of sediments during and after uplift or emergence, characterized by infiltration of meteoric water and downward migration, usually marked by low pH and high Eh. Near the surface merges with the zone of weathering. (FAIRBRIDGE, this chapter.) Epigene (adj.): As a general term-all processes or phenomena produced at or near the earth’s surface (GEIKIE,1879); specifically for mineral deposits formed later than the enclosing rocks or by secondary alteration. Epigenesis (adj.-etic): Changes in the mineral character of a rock due to external influences Also applied to mineral deposits, as epigene. One may have both epigenetic (A.G.I. GLOSSARY). supergene (with descending waters), and epigenetic hypogene (ascending waters). Halmyrolysis (adj.4ytic): Geochemical modification of sediments during deposition, due to

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reactions with sea water (ionic transfer), originally called “submarine weathering” by HUMMEL (1922), but applies also to ionic rearrangement and replacement Cpernrom, 1957). Hypogene (adj.): Minerals, or changes in rocks, related to ascending waters of magmatic origin,

specifically applied to mineral deposits, but formerly to any deep-seated endogenic processes, 1833; GEIKIE, 1879). involving magmatism and metamorphism (LYELL, Kuturnorphism (adj.-ic): Alteration of rocks, particularly solution and breakdown at or near

the earth‘s surface, due to either supergene or hypogene waters, forming simple minerals from complex (VAN HISE, 1898): it thus included both weathering (upper zone) and near-surface cementation (lower zone), but LEITHand MEAD(1915) excluded the latter.

Kutugenesis: Specifically restricted to the breakdown of rocks (FERSMAN, 1922). Lithification: The complex of processes that converts an unconsolidated sediment into a hard 1957). rock, including compaction, dehydration, cementation and induration (e.g., PETTIJOHN, Lithogenesis (adj.-etic): Synonymous with petrogenesis, relating to the origin of a rock (A.G.I. GLOSSARY). Supergene (adj.): Applied to mineral deposits or enrichment related to descending waters (A.G.I. GLOSSARY). One may have both epigenetic supergene deposits and syngenetic supergene deposits (as in manganese nodules). Syndiugenesis (adj.-etic): Modification of sediments during and immediately following deposition, often by biochemical influences, marked by extreme variations in pH and Eh. (BISSELL,1959; FAIRBRIDGE,this chapter.) Syngenesis (adj.-etic): Formation of mineral deposits more or less contemporaneously with the deposition of the enclosing rocks, i.e., the opposite of epigenesis (A.G.I. GLOSSARY). Specifically refers to the time of geochemical changes, i.e., penecontemporaneous (FERSMAN, 1922).

Chapter 3 DIAGENESIS OF SANDSTONES E. C. DAPPLES

Department of GeoIogy, Northwestern University, Evanston, IN. (U.S.A.)

SUMMARY

This chapter concerns the diagenetic modification of sands following deposition, such modifications occurring as progressive stages. Redoxomorphic (oxidationreduction) reactions involving iron in particular characterize early burial. Locomorphic changes (cementation and mineral replacement) involving primarily silica and carbonates are typical of lithification. The phyllomorphic stage (authigenesis of micas and feldspars) is a late burial feature. Chemical reactions which occur during each of the three stages of diagenesis result in equilibrium mineral assemblages which are considered to identify the pH and Eh of the interstitial fluids. Shifts in the direction of equilibria are indicated by corresponding changes in the mineral assemblage. Although a secondary mineralogy characterizes most of the diagenetic progression there may be also removal of detrital mineral matter to produce a simple residual mineralogy, and special granular intersutured textures.

INTRODUCTION

Sedimentarypetrologists are increasingly in agreement that post-depositional modifications in the mineralogy of sandstones can be very significant, particularly with regard to replacement of one mineral by another. It is equally well established that authigenic minerals may develop as a result of oxidation and reduction, this being manifested principally in rocks as red and green-gray colors and as individual mineral grains, particularly pyrite. Although such modifications are acknowledged, the extent to which such diagenesis has modified the oridnal detritus is a subject of considerable diversity in opinion (PETTIJOHN, 1957, pp.648-650). Are the modifications so profound that the present day mineralogy of the sedimentary rock is only remotely akin to that of the detritus which was deposited? Is it possible that such monominerallic rocks as pure quartz sandstones were never sediments of such composition but are the products of substrata1 solution? If such should prove to be the case, then obviously the existing systems of sandstone classification are not satisfactory devices from which to interpret the genesis of the sediment. It is the purpose of this chapter to clarify some of the uncertainties regarding the importance

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of diagenetic modification of sand detritus, to demonstrate that stages of diagenesis represent the progress of the sediment into a rock, and that once lithification has been achieved additional changes in the mineralogy are to be regarded as counterparts to stages of metamorphic grade. Investigators of strata containing graywackes are aware that their common association with other sedimentary rocks showing low grade of metamorphism indicate the possibility that some of the characteristics may be products of partial metamorphism. For example, the high content of soda is considered related to “fresh” grains of albite-oligodase of good crystal outline recognized in some examples to be authigenic. Common occurrence of authigenic high soda zeolites also has been reported in such rocks and indirectly has helped to establish the zeolite metamorphic facies (COOMBS et al., 1959; ZEN, 1961). Later investigations by PACKHAM and CROOK (1960) suggest the existence of additional facies representative of still lower temperature realms which they assigned to epigenesis or a late stage of diagenesis. In cratonic strata, which can be demonstrated never to have been buried more than a few thousand feet and hence never exposed to temperatures much above those at the earth‘s surface, precipitation of cements and replacement of grains by later cements show relations indicative of equilibrium reactions. With shift in the equilibrium, a mineral assemblage is no longer stable and is rearranged into a different mineral suite. This has led to a proposal by the author to establish certain stages in the progress of diagenesis based upon the prominence of certain types of reactions. Among these some clearly are related to the nature of the particle-size distribution increasing in complexity with an increase in fine-grained matrix content. Other reactions result in simple precipitation of cement. An understanding of the reactions involved concerns the size distribution of the detritus and the various means by which individual grains are bonded into a rock aggregate.

THE NATURE OF SANDSTONE SEDIMENT

Assuming that present-day sand detritus is the compositional and size counterpart of detritus as accumulated in the past, several guiding principles regarding particle-size distribution are well established. In certain sediments, best illustrated by river deposits, the particle-size distribution ranging in size from sand to clay is not log normal but is skewed toward the fine sizes and may includemuch clay sediment. Some of the large particles, however, are aggregates, or pellets, composed of very small individual clay particles. Such pellets moved with sand-sized fractions in the same current velocities and were deposited with sand grains. In the river environment local flood deposits such as sands of point bars, natural levees and the like are accumulated rapidly and under higher velocities and in these a much more restricted particle-size distribution is observed, the clay-size fraction having re-

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mained in suspension. Winnowing action and consequent removal of the clay fraction either by the wind or shore currents characterizes deposition of dune, beach, and nearshore bar sands. These sands approach a log-normal particle-size distribution. In the localities of playas, deltas, and nearshore quiet marine waters the range in transporting current velocities is large although dominated by low velocities. In such environments debris of sand size grades into clay size within short lateral distance and the so-called sand “pinchout” is developed. Primarily the sand bodies are those where sand sizes have been concentrated during rapid deposition and hence are characterized by large scale cross-bedding and upward gradation of grain size in units of as much as 1/3-1 m thick. In directions laterally away from such deposits the sand is diluted by the clay fraction in increasingly greater amounts and the entire texture is changed. The importance of the tectonic control on the composition and texture of the sand debris is also well known and the separation of sediment into shelf, basin, and geosynclinal types is demonstrable. The overall rate of submergence is the dominating influence on the relative fraction of sand, silt, and clay which constitute the ultimate sandstone. Thus, the shelf deposits are the most texturally mature and the eugeosynclinal sediments the most immature. The distinction is not so readily made with regard to the mineral composition. It is the purpose herein, however, to restrict the consideration to the petrographic modifications which are imposed during diagenesis and to omit any discussion concerning the origin and the composition of the detritus as it arrived at the depositional site.

GENERAL PROCESSES OF LITHIFICATION

Lithification is the general process of bonding individual particles together so that an aggregate mass results. The binding of individual grains one to the other can be demonstrated to result from several distinct processes and can be classified into individual types ranging from simple addition of cement to complex interlocking and welding of grains in which cement as such is not significant. Simple cementation Mineral precipitation Common among shelf sandstones are textures in which sand grains touch one another with tangential or point contact and are held in position by simple cement precipitated in the interstitial pore space (WALDSCHMIDT, 1941; TAYLOR,1950). Calcite deposited in the interstitial space of a sand, constituted primarily of quartz and representative of a well-winnowed deposit, is a common illustration of this condition. Another example is hematite precipitated in the pore space between sand particles to produce a red-colored rock. Still another illustration is opal or

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Fig.1. Simple cement bond (crossed nicols); grains of quartz, rock fragments, and other minerals cemented by calcite. Sandstone at base of Platteville Formation (Ordovician) near Minneapolis, Minnesota. Dark areas are grains and calcite in partial light extinction. Note quartz grains are isolated by cement but boundaries tend to be sharply defined. Fig.2. Opal cement (crossed nicols); Ogallala Formation (Paleocene), Smith Co., Kansas. Interstitial dark areas are opal (0).Note the transformation of opal into fibrous chalcedony (f),and small areas of chert (c). Opal is the unstable silica phase in the latest environments.

chalcedony deposited in thin laminae in the intergranular space causing quartz and other grains to adhere (Fig. 1,2). Selected cement and replacement textures are given in Fig.1-6 and Fig.8-10. Precipitation of such mineral matter does not involve reaction with individual detrital grains which are regarded as having behaved as inert bodies. Crystallization of such a cement is a consequence of supersaturation of the permeating interstitial fluid only, the results of a condition which has caused the solubility products of certain ions to be exceeded. The end product of this process is an aggregate of individual particles as a more or less lithified mass, the cement behaving as a glue to cause the particles to adhere. A related but somewhat more complex bond involves addition by overgrowth to selected grains in which such grains have acted as seeds for later precipitation of a like compound from solution. The most common occurrence is addition of quartz to detrital grains to produce geometric outlines on formerly irregularly-shaped particles. The crystal outline results in a texture by which boundaries between grains are lines and interstitial space largely is obliterated (HEALD, 1956,

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Rg.3. Pure-quartz sandstone (crossed nicols) illustrating grain boundaries; a=line; b=concavoconvex; c=simple penetration. Tuscarora Sandstone (Silurian), Virginia, note the slight intergranular interlock with minor quartz weld between grains. The darkened grain ( d ) (right center) is largely secondary quartz and continues as filling between grains. All grains in the photograph are quartz. Fig.4. Sandstone of simple clay bond (crossed nicols). Pleasant-view Sandstone (Pennsylvanian), Rushville, Illinois. Sand size fragments are aggregated by the clay (c) which constitutes a matrix. Some specimens will disaggregate in water; in others, some of the clay is replaced by chert or muscovite and grain bonding is stronger. Dark area (lower left) is very fine clay (illite is dominant in specimen), lighter area contains some authigenic muscovite (m).Note tendency of quartz grains to interpenetrate whereas quartz and clay show very little inter-reaction.

p.16; SIEVER,1962, p.64). Ideally the resulting texture shows slight interlocking of grains and the aggregate ranges in degree of friability in inverse proportion to the amount of precipitated overgrowth (Fig.3). In those examples where much quartz has been added, cementation is pronounced and a quartzite texture is attained primarily by welding of one grain to another as the line boundaries are developed.

Simple clay bond Sandstone textures in which clay-size particles are important either as interstitial fillings or as matrix (i.e., wackes in general) characterize basin deposits. Such sands may be lithified by .a simple bond by which large grains are aggregated through surface cohesion of the clay-sized particles (Fig.4). The nature of the bond is primarily a surface film phenomenon and attains its greatest tensile strength when the clay is dry. Ideally, cementation of this kind involves no precipitation of mineral

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matter and sandstones bonded in such a fashion tend to disaggregate when placed in water, particularly if some of the clay expands. No intermineral reaction is involved and the particles behave as inert constituents; the gluing action is due entirely to the cohesive forces of the clay-size grains and water bonds. Complex bonding Grain solution Fastening individual grains in the aggregate increases in complexity as chemical reactions occur between the mineral grains and invading solutions. Although boundaries between individual grains suggest that solid-solid reaction has occurred, it is considered that this is not the case and that reaction takes place only when a fluid is involved. A most simple example is illustrated by substrata1 solution of quartz grains. The result is an interlock between individual grains brought about by differential solution. Some conditions result in rather broad inter-penetrations resulting in concavo-convex boundaries (TAYLOR,1950), whereas in others the interlock becomes more highly developed and irregular, as illustrated by microstylolitic boundaries (SLOSSand FERAY, 1948; BURMA and RILEY,1955; HEALD,1955). Boundaries between quartz grains, and particle size distribution of flexible sandstones, or itacolumite, most certainly are not representative of the original grains as deposited (Fig.5). Irregular and locally rectangular boundaries, marked intergranular interlock, and the extreme surface corrugation testify to the post-depositional modifications in shape. In these exceptional sandstones individual quartz grains are loose and not bound by cement. Moreover, the random distribution of large and small particles does not suggest deposition of a texturally mature sand. The principal difference between this and other interpenetration textures is absence of an intergranular weld. In the example of the flexible sandstone the grain aggregating process is entirely one of physical interlock which holds grains together as loose pieces of a jig-saw puzzle, and absence of cement is pronounced. The mechanism which produces the texture must be dominated by decementation, that is, removal of mineral matter from some of the pore space and also from grain margins (CAROZZI,1960, p.36). In the particular post-depositional environment quartz must be unstable and soluble in the permeating solutions. This is to be contrasted with other textures in which solution of quartz and reprecipitation on other grains, or in interstitial pores, proceeded simultaneously to produce a granular interlock. Such a condition is attributed by some authors to pressure solution inasmuch as in some examples there appears to be a relation to loading and also to intensity of folding (THOMSON, 1959). Certain clay-quartz boundaries result from decomposition of feldspar in the sediment. Some clay grains preserve the outline of the original detrital grain of feldspar, whereas in others some of the clay has been squeezed partially into interstices during compaction. The processes involved are hydration and hydroly-

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sis of the feldspar without reaction between minerals with which such a grain is in contact. Not infrequently such “defeldspathization” is associated with intersuturing and interlocking of quartz grains which indicates that quartz also has been unstable in the same environment. Intergranular reaction Bonding of grains by reaction between minerals involves examples of one mineral replacing another, development of a third mineral along their common boundaries, or reaction between the clay-size matrix and the larger grains to produce a new mineral. Reactions of these sorts are considered by some authors to be typically diagenetic. As yet, there is no unanimity regarding the stages in the lithification history during which such processes occur, and some investigators regard them as occurring very late (PETTIJOHN, 1957, p.662). More recently, studies of the paragenesis of such minerals has suggested to several authors that there exists a series of diagenetic mineral facies which imply that equilibrium reactions have occurred between minerals (STRINGHAM, 1952;ZEN,1959;PACKHAMandCROOK, 1960). Moreover, it isconsidered that they should have much the same connotation as their metamorphic rock counterparts (PACKHAM and CROOK,1960; DAPPLES,1962). Some of these reactions are those of halmyrolysis, that is, they dominate duringearly burial (prior to lithification). Others should be regarded as part of the lithification process and lead to the bonding of mineral particles one to another. Still others are best developed under conditions of deep burial and relatively strong folding and clearly occur after lithification of the sediment (DAPPLES,1959, 1962). In these the somewhat elevated pressures and temperatures associated with deep burial and folding appear to cause the reactions to proceed at a faster rate or more toward completion.

STAGES OF DIAGENESIS

Processes in the environment of early burial, that is, during initial compaction, ejection of fluids, and prior to lithification, are dominated by oxidation and reduction reactions. Because such reactions are so important this stage has been called redoxomorphic, to signify the existence of reversible reactions dependent upon oxidation potentials primarily manifested in the kind of iron-bearing mineral which crystallizes (DAPPLES,1962). The redoxomorphic stage tends to establish the bulk final color of the rock and is most strikingly noted among sediments in which the total iron content exceeds 7%. Once the reactions have been driven to oxidize or reduce iron, there is much less tendency to reverse the equilibrium during the later history of the rock. Nevertheless “bleached areas” and other later color changes are abundant but are presumed to occur on a small scale. A second and more advanced stage involves significant precipitation of mineral matter in the pore spaces and particularly as replacement of detrital mineral

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Fig.5. Flexible sandstone (crossed nicols) Itacolumite, Brazil. Individual grains in this specimen are loose and can be moved independently of other surrounding grains. Interpenetration of grains is primarily a function of solution although secondary quartz precipitation has occurred also (note boundary a). The large grain (b) is at a position of partial extinction. The abnormal shape of the grain is attributed to repeated solution and reprecipitation of quartz, and is considered not to represent an original detrital grain. Fig.6. Detrital biotite (b) in red sandstone split along cleavage laminae by precipitated calcite (crossed nicols). Minturn Formation (Pennsylvanian), Colorado. Much of the photo shows calcite cement (c) surrounding grains; decomposed oxidized biotite (ob) center; k =border of clay mineral surrounding oxidized biotite.

grains. This is the time of primary cementation and the development of induration. Replacement of interstitial clay by chalcedony, calcite by siderite, and quartz by calcite are common examples. In the usual case the replacement is not pseudomorphic but rather a partial in situ substitution of one mineral for another. This period of important mineral replacement follows in time the principal redoxomorphic reactions, and is identified as the locomorphic stage signifying a change of shape in situ. During the locomorphic stage the sand sediment gradually achieves partial or complete lithification and becomes the initial rock. Any additional diagenetic modification which may occur takes place on a lithified aggregate and for this reason is considered by this author to be associated with its late burial history. A common characteristic of late burial is alteration of clay minerals into mica, and development of well-crystallized phyllosilicate minerals in general.

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For this reason this late episode of diagenesis is termed thephyllomorphic stage.l Certain other minerals, particularly feldspar, appear as authigenic growths during crystallization of the phyllosilicates and at present are included as part of the phyllomorphic stage. In some localities where the phyllomorphic stage is well advanced, there is believed to be gradation into the zeolite and chlorite grade of metamorphism without any currently recognizedlimits (PACKHAM and CROOK, 1960). Redoxomorphic changes Oxidation and reduction reactions can be demonstrated to dominate modification of the sediment during and immediately after burial. During this time, compaction is in progress and fluids are being ejected with concentration gradients toward the depositional interface (WELLER,1959; VON ENGELHARDT and GAIDA, 1963). Principal reactants involved are iron, oxygen, sulfur and carbon. Deposits consisting of a significant fraction of organic matter tend to contain sulfur as well as carbon. Of these, carbon compounds appear to be most rapidly oxidized and may be regarded as contributing electrons to drive the iron into the ferrous state and thus permitting fixation of the sulfur as pyrite. As long as the organic fraction remains important, the gray color will prevail and pyrite will be scattered throughout the rock often in considerable amounts. Among red-colored sandstones essentially two somewhat distinct conditions may dominate the redoxomorphic stage: (1) A situation may exist during which the burial environment is oxygenated by contact with the atmosphere and iron oxides arrive as part of the detritus, or clay minerals with attached iron ions are part of the inorganic suspension 1958). In the burial environment oxygen gathers electrons prinload (CARROLL, cipally donated by iron to form hematite and related ferric oxides or hydrates, which, along with those having arrived as part of the detritus, remain stable. A texture can be observed showing rock fragments, grains of quartz, and other minerals more or less isolated by mixtures of clay minerals and iron oxides as films, matrix, or pore filling. Such a mixture is modified texturally only as a result of differential compaction between the sand and clay sizes. No mineralogic reaction is noted between the iron oxides and the sand grains; the relation is that of inert substances, and grain-matrix boundaries stand out in clearly-defined demarcation (SWINEFORD, 1955, p.155). The detritus composed in part of clay minerals with attached Fe3+ ions would tend to be brown or red after deposition (KLEIN, 1963). In other circumstances the iron could arrive in the reduced form and oxidation would take place during deposition, hence a gray detritus would become red. Inasmuch as the precipitation of iron hydroxide is indirectly controlled by the pH (GARRELS, 1960, pp.116-145), a slightly acid condition of many streams would 1

Considered as an epigenetic stage by some investigators (Editors).

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tend to keep the iron primarily in the reduced form during transportation and red color would develop only after deposition in an environment of higher pH, such as a sea or playa lake. (2) A more common situation is believed to prevail during which the equilibrium between Fez+ and Fe3+ is shifted toward oxidation sometime after burial. In its most obvious form this can be seen in strata in which the red color transects bedding, and coloration follows fractures or permeable positions related to introduction of water carrying dissolved oxygen. Conversely, bleaching of the red color can be demonstrated for masses of irregular outline or for concentric zones about some center. Paragenetic relations between minerals from such zones show rather persistent association between iron oxides, biotite, chlorite, siderite and calcite (see also GREENSMITH, 1957, p.410). The occurrence of biotite, chlorite, siderite and calcite in bleached zones (i.e., where the iron is in the reduced state), and decomposed condition or absence of such minerals where the iron is oxidized, is interpreted to illustrate the following reactions: oxidation biotite

s

iron oxides

+ clay minerals (illite + kaolinite)

reduction biotite (in presence of calcite and in slightly reducing environment) + chlorite In the red-stained parts of the Minturn Formation (Pennsylvanian) of Colorado immediately surrounding bleached zones, calcite which is present in the bleached portion extends a short distance into the red rock. Here it tends to split detrital grains of strongly oxidized and decomposed biotite along cleavage laminae (Fig.6). In the red-stained rock, biotite is obviously unstable and tends to decompose to clay mineral and ferric oxide. Within the bleached zone undecomposed biotite is stable in the presence of calcite. Biotite grains are not invaded by calcite and the two minerals have sharp and regular boundaries. Similarly, chlorite tends to disappear in the red-stained rock, whereas it is significantly present in the bleached zones. The interpretation currently favored is that in an oxidizing environment detrital biotite is unstable and decomposes first to oxidized biotite and later to clay mineral and ferric oxides. Conversely, in a calcite precipitating environment (pH =8) biotite is stable and will form as authigenic crystals if a reducing environment prevails. However, the occurrence of chlorite in greater amounts in the bleached zone in association with calcite, and its absence in the oxidized zone, is regarded as suggestive that chlorite is more stable than biotite in a calcite precipitating environment (Fig.7). Occurrence of secondary biotite is common in the reduced zones of many redbed sandstones, but its distribution is sporadic ranging from local abundance to absence. Crystallization of authigenic biotite within masses of interstitial clay in bleached zones, however, can be established. MILLER(1957) reported that in the center of “reduction zones” in Pierce Canyon (Permian) red beds, biotite is absent

101

DIAGENESIS OF SANDSTONES

onfmorillonite

Eh

I I

I

I

6

I

7

0

PH Fig.7. Generalized stability realms of clay minerals and micas based upon paragenetic sequence of associated minerals and miscellaneous environmental indicators. The intent is to show the Eh and pH conditions under which a clay mineral or mica assemblage tends to form preferentially in sandstones.

whereas secondary calcite is present, but outward from the center biotite has formed in the presence of precipitated gypsum and calcite. In the same reduced zones magnetite, which is abundant in the red rock, is absent. This observation had been previously reported by MILLERand FOLK (1955) in red beds of other ages and localities, and is attributed by them to demonstrate instability of the iron oxide under the conditions which produced the bleached spots. In the Pierce Canyon strata biotite is stable in an outer shell of the “reduced zone” in association with calcite, and the equilibria indicated above have been driven in favor of crystallization of biotite. The implication to be drawn is that the reactions are more sensitive to relative degrees of oxidation and reduction than to pH, but that chlorite and biotite will crystallize from a solution havingapH-8. KELLER (1953) has shown that greencolored reduced zones in certain red beds contain illite rather than chlorite. Perhaps the presence of illite may be important in determining the variety of mica which crystallizes as a stable phase under certain H-ion concentrations. Another equilibrium reaction concerning micas is to be noted between muscovite and biotite. Biotite rather than muscovite appears to preferentially form from illite and kaolinite under reducing conditions and an environment of pH 5 7; i.e., in non-calcite, chert-cemented, carbonaceous, fresh-water deposits. However, it appears that biotite is stable across a wide range of pH from acidic to basic. A possible mechanism which favors precipitation of biotite is suggested by the

102

E. C. DAPPLES

association of carbonaceous fragments with the occurrence of biotite and the equilibrium: oxidation biotite

z

muscovite

reduction is shifted by oxidation of woody fragments and reduction of iron. Siderite is a common, although not abundant, early precipitated cement in many sub-graywacke sandstones. It is especially observed as concretions and individual crystals in sub-graywackes associated with coal beds, and where there is evidence that brackish water prevailed in the depositional site. A necessary item appears to be that enough carbonaceous matter was incorporated to maintain a higher concentration of ferrous than ferric ion. An example is reported by RUSNAK (1957, p.47) in which the order of pore filling in the Pleasantview Sandstone(Pennsylvanian) shows calcite partially replaced by siderite. Both calcite and siderite were later somewhat replaced by quartz precipitated as overgrowths on detrital grains; hence, crystallization of the siderite is intermediate in time of development. The siderite remains as a stable phase as long as reducing conditions prevail; but, wherever ground waters carrying dissolved oxygen have penetrated the rock, the siderite has at least partially oxidized to limonite.

Locomorphic changes Opal and chalcedony In many sediments modifications identified as locomorphic are simple and involve changes in the pore cement only. Such is the case with opal filling in certain sandstones. An example is the Ogallala Formation (Tertiary) where at certain localities opal has been deposited as a primary simple cement (Fig.2; see also SWINEFORD and FRANKS,1959). In some specimens opal has partially filled interstitial pores, whereas in others the amount of precipitation has been large, and detrital grains obviously have been separated. Clearly an increase in total volume from that occupied by the unconsolidated sediment has been brought about by opal precipitation. Opal as a cement is sporadically distributed in sandstones of Tertiary or younger age. In older deposits its occurrence is extremely rare indicating that it is not stable in the late burial environment. An example of the replacement relations was reported by FRIEDMAN (1954, p.239) who interpreted a paragenetic sequence as follows: shell calcite is replaced by opal which in turn is transitional into fibrous chalcedony. Quartz is found only adjacent to the chalcedony which always separates the quartz from the opal. The replacement was evidently volume for volume. Transition from opal to quartz appears to be a product of some aging process, but possibly the ordering in the lattice is inhibited by the presence of some large cations such as Kf or Ca2+ present in the loosely organized opal molecule. During con-

DIAGENESIS OF SANDSTONES

103

ditions of late burial such ions could be ejected along with water as the quartz lattice structure develops (FOLKand WEAVER,1952). Whatever structural modification is involved, it is unidirectional from opal to quartz and appears irreversible within the physical conditions of stability of sandstone. Moreover, the progressive nature of the change as observed is: opal + chalcedony + quartz Precipitation of quartz directly as overgrowths on detrital grains and welding of quartz grains by such precipitation is much more common than precipitation of opal. Based upon the rare occurrence of opal, two physical conditions tend to favor its precipitation. One such condition appears to be supersaturation of water with silica such as develops in hot spring localities or from surface waters leaching volcanic ash beds (KELLER and REESMAN, 1963, p.432). Apparently large, loosely organized molecules of hydrated silica precipitate as a colloid. The latter becomes organized into a structural lattice during late burial, but retains the outward appearance of chert. The other common precipitant of opal is cellulose, or a related carbohydrate, in which by some mechanism substitution of silica and reorganization of the organic molecule occurs as the cellulose is slowly oxidized after burial. Reorganization into chert tends to destroy the details of organic cell structure preserved by the opal. Substitution of chalcedony or chert for clay matrix and interstitial clay is a locomorphic change of common occurrence. The replacement is particularly well observed in sub-graywacke sandstones, which show a simple clay bond as an aggregating agent for sand-sized grains (Fig.4, 8). Locally chert can be recognized to have replaced part or all of such interstitial clay. The nature of this substitution is not entirely clear. Limited investigation has not established whether the silica has been precipitated in the clay aggregates or whether it has in fact replaced part of the clay-mineral lattice by substitution in the octahedral layers. The latter mechanism would require migration of aluminum for which there is no evidence at present. Simple precipitation of silica within the openings between individual clay-mineral crystals currently is favored. Inasmuch as the clay-size interstitial material is known to contain small particles of quartz, these could act as nuclei for precipitation of additional silica. Precipitation of chert in the matrix normally is an early locomorphic process and is considered to follow closely, if not to be contemporaneous with, processes of the redoxomorphic stage. Chert precipitation, however, continues as the locomorphic stage advances and welding of detrital quartz grains gives rise locally to highly silicified sandstones, as, for example, near fault zones. Under such circumstances the process,can be demonstrated to have occurred during the post-lithification stage and obviously late in the sandstone’s history. Chert replacing interstitial clay in some sub-graywackes appears to grade through zones of coarse microcrystalline quartz to be welded to detrital quartz

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Fig.8. Chert (ct) replacement of interstitial clay (c) (crossed nicols); Muddy Sandstone (Cretaceous), Casper, Wyoming. Locally this sub-graywacke has a clay bond partially replaced by chert. Note tendency of chert to weld to quartz grain (4).Dark spot (h) is a hole in thin section left by disaggregation of unaltered very finely divided clay representing simple clay bond. Fig.9. Calcite (cn) replacement of clay matrix (crossed nicols); sandstone in Fort Union Formation (Paleocene) near Wamsutter, Wyoming. Dark and light grains are quartz in various positions of light extinction. Note that outline of quartz grains suggests little replacement by calcite, whereas clay matrix which is present in nearby specimens is completely replaced. Replacement of matrix appears to have occurred without significant shift in position of quartz grains.

crystals or may be transitional into quartz overgrowths. Some examples illustrate recrystallization into progressively coarser individual crystals of quartz. This does not appear to be commonplace in those sub-graywackes in which calcite is an important cement. Indeed it would appear that concentration of calcite tends to inhibit recrystallization of chert to quartz. Calcite replacement of clay A frequent observation among sub-graywackes is replacement of clay matrix by carbonates (RUSNAK,1957, p.41). Primarily such replacement is by calcite but it may also be by dolomite and siderite. The replacement may be so complete as to give the impression that the original particle-size distribution contained virtually no fraction in the clay size (Fig.9). In some sandstones the calcite has actually replaced the clay fraction in the sense that the clay no longer is present as an insoluble residue in the carbonate. In other examples, at least some of the clay remains

DIAGENESIS OF SANDSTONES

105

as a residue within the calcite and can be recovered on solution of the carbonate. The mechanism of such replacement is not understood but it appears that certain clay minerals, known to be primarily illite and kaolinite in the case of some subgraywacke, are flocculated by Ca-ion and occupy less interstitial space allowing the remainder to be filled by the precipitated carbonate. A reaction illustrated below, which was proposed by EADESand GRIM(1960), could explain the development of calcium silicate hydrate or calcium aluminum hydrate and to account for lattice structure changes upon treating such clay minerals with lime. Kaolinite or illite

+ Ca2+ +calcium silicate hydrate or calcium aluminum hydrate

If the new crystals tend to remain very small in size, they could physically move out of the interstitial space with permeating solutions, allowing the precipitating calcite to occupy the former position of the clay. On the basis of what is known of solute precipitation, replacement of clay minerals by calcite is favored by a pH =8 and a high concentration of Ca-ion under which conditions certain clay minerals become unstable (Fig.7). Calcite-aragonite replacement Replacement of aragonite fossil shell material by calcite is an extremely commonplace unidirectional substitution. Precipitation of aragonite appears to be temperature sensitive in that aragonitic shells are more abundant among warm water invertebrates than cold water forms (LOWENSTAM, 1954, p.285). Transition from aragonite to calcite, however, is more rapid under certain conditions than others. Some shells of Mesozoic age still contain aragonite, whereas others of much younger age are completely altered to calcite. Experimental data indicate that transformation of aragonite to calcite is favored by increase in temperature and reduction in pressure (DEERet al., 1962, p.308). In sandstones the transformation appears to be independent of the depth of burial, and fossil fragments are composed of calcite even in very young rocks which scarcely have been buried. At present, knowledge of the causes and conditions of the transformation do not appear to be sufficiently understood to serve as a means of distinguishing between various stages of diagenesis among sandstones. Calcite-dolomite replacement Replacement of calcite by dolomite is to be noted in an important extent in two very distinct groups of sandstones. One is the quartzose group which passes by lateral facies change into limestone (the quartzite series of KRYNINE,1948), which at some later time is subjected to wholesale dolomitization. In such rocks quartz grains may be completely engulfed in carbonate, dolomitization may be limited to replacement of calcite, and quartz is unaffected. Quartz grains are regarded as “floating” within the carbonate and display a well-defined boundary. The interpretation favored at present regarding such a boundary is that no reaction has occurred

106

E. C. DAPPLES

between the carbonate and the quartz, the latter having behaved as inert during the progress of replacement (GLOVER, 1963, p.40). Among sandstones of this type, the calcite characteristically is completely replaced without any seeming increase in individual crystal size. Replacement is also unidirectional and any later addition of calcite appears to be limited to joint or cavity fillings and not to reversal of the reaction calcite-dolomite. A second type is recognized among sub-graywacke sandstones in which the primary calcite cement contains isolated rhombs of dolomite and siderite. The association is one which often is accompanied by carbonaceous matter and in which siderite is more common than dolomite. The presence of Fez+ ions seems to favor precipitation of siderite rather than an iron-rich dolomite, although later investigation may show greater abundance of dolomite than currently recognized. In rocks of these types the amount of available magnesium tends to be insufficient to permit the crystallization of large amounts of dolomite; and the common occurrence is as individual rhombs of good crystal outline within masses of calcite. This occurrence suggests that the dolomite may possibly represent some ex-solution phenomenon rather than the result of introduction of magnesium from some outside source (GOLDSMITH and GRAF,1955; GOLDSMITH et al., 1962).

Feldspar-calcite replacement In certain arkoses which are known to have attained the phyllomorphic stage it is not uncommon to note significant replacement of potash feldspar by calcite precipitated as a cement. In these, both quartz and potash feldspar are partially replaced by the secondary calcite (see also GREENSMITH, 1957, p.410). The typical alteration represented is interpreted as resulting from a process by which solutions rich in Ca2+ and c0s2-ions are capable of destroying the potash feldspar lattice, possibly by causing the silica tetrahedral units to go into solution under the high pH which must characterize the calcite-precipitating solution. HAY(1957) reported that calcite replaces plagioclase in beds of Eocene age in the Absaroka Range, Wyoming; hence, the replacement reaction is not restricted to potash feldspar. Replacement does, however, appear to be related to important precipitation of calcite suggesting that a mass action effect is required. Phyllomorphic changes Crystallization of micas Reactions categorized as phyllomorphic are favored by increase in pressure and seemingly also by increase in temperature. Sandstones which have been subjected to strong pressures either in folded belts or along fault planes develop micas in interstitial openings and along quartz grain boundaries (Fig.10). There is also an abundance effect in the direct proportion between the quantities of secondary mica which are crystallized and the amount of detrital clay which occurs in the sandstone;

DIAGENESIS OF SANDSTONES

107

Fig.10. Vitreous quartzite showing quartz grains (4)with intergranular zone of secondary muscovite (m) (crossed nicols); Ajibik Quartzite (Precambrian), Marquette, Michigan. This bed is exposed on the limb of a strongly compressed syncline in which slaty cleavage has been developed. All grains in photo are quartz in various degrees of partial light extinction. Note absence of interpenetration texture between grains and the significant amounts of muscovite occupying interstices and bordering each grain. Fig.11. Matrix of glauconite in quartzose sandstone showing gradation between clay mineral (c), glauconite (g), and chlorite (ch) (crossed nicols); Lamotte Sandstone (Cambrian), Iron Mountain, Missouri.

the amount of mica observed being independent of the intensity of folding. The greatest development of secondary mica, however, is to be observed in the examples in which the argillaceous sandstones are well folded or have been deeply buried. In general, crystallization of muscovite appears to be preferred over other micas either because its lattice is more readily developed from most of the clay minerals, or perhaps because it is stable under the most common conditions of burial. Where quartzose sandstones have been strongly folded, border zones of muscovite crystals bounding the quartz grains and partially penetrating them is a commonly observed feature. In others, such as shown in Fig.10, little penetration is noted and each grain appears isolated. A third group of folded quartzites show strong intersuture and marked granular elongation, and in these quartz is regarded as having been mobile. The latter textures point toward the existence of an equilibrium between solution and precipitation of quartz as reported by THOMSON (1959). A growing muscovite crystal could readily become surrounded by quartz during

108

E. C. DAPPLES

the process of solution and reprecipitation, and thus appear to have penetrated the grain. In rocks having higher clay content than most quartzose sandstones several micas generally are developed. Mention has been made that biotite appears in sub-graywackes often associated with carbonaceous fragments and in many occurrences having begun crystallization during the redoxomorphic stage. In part this is explained by the frequent presence of iron which may be available either attached to the clay mineral or as Fe2+ ion held by organic compounds. Biotite has not been observed to be as abundant as muscovite except in “reduced zones” in certain red-bed sandstones. If the environment is oxidizing, then the iron forms hematite or limonite and the amount of clay mineral altered to biotite is very substantially reduced, as biotite is unstable and the direction of equilibrium is reversed. Certain sandstones, in which glauconite is distributed as a matrix enveloping quartz grains, indicate a phyllomorphic change involving glauconite and other clay minerals, chlorite, muscovite, biotite and quartz. An example is the upper portion of the Lamotte Sandstone (Cambrian) near Iron Mountain, Missouri (see also O J ~ A N G A1963). S , The rock is a quartzose sandstone which in the extreme upper part‘ contains matrix glauconite. Within the areas of glauconite some poorlypreserved remains of a light gray, non-glauconite clay mineral may be seen in thin section (Fig. 11). Such areas also show silt-size grains of quartz highly embayed and replaced by the glauconite. Large quartz grains are not so affected. The relations involved are interpreted as an equilibrium reaction: clay mineral

+ quartz + K+

glauconite

In the example cited the equilibrium has been displaced toward the right and glauconite is the abundant phase. Presumably this reaction was favored by a slightly reducing environment in which the small amounts of iron were present principally as Fez+ ions. Also the presence of thin layers of carbonate in the sandstone and minor amounts of secondary calcite are interpreted to indicate that a pH of approximately 8 prevailed (Fig.7; see also CARROLL, 1958). Within the mass of glauconite, another equilibrium involving glauconite and chlorite is interpreted as having been established. Irregular masses of chlorite, later in age than the glauconite, indicate that a condition prevailed which resulted in some conversion of the glauconite to chlorite. Presumably, the glauconite-clay mineral equilibrium has had superposed upon it a condition which caused a second and more advanced phyllomorphic reaction to obtain. Hence, the chloriteglauconite equilibrium is regarded as marking the most advanced stage attained by these rocks of essentially horizontal attitude and shallow burial. The equilibrium assemblage suggests also that K-ion is mobile in the system and may be held in the glauconite, or moves out as a free ion where chlorite is crystallized. In other cratonic quartz-glauconite sandstones, muscovite and bio-

DIAGENESIS OF SANDSTONES

109

tite rather than chlorite exist in an equilibrium assemblage with glauconite. This may be the result of K-ion remaining in the lattice of the micas, or it may be the response to a different condition than the chlorite-glauconite equilibrium assemblage. Local patches within the glauconite can be observed to contain secondary crystals of muscovite or biotite. Presumably the presence of small amounts of Fez+ is the factor deciding which one of the micas would crystallize. In this rock muscovite and biotite are very minor constituents whereas glauconite is abundant. Hence, the equilibrium: glauconite

+ muscovite + biotite

is considered to be displaced toward the left and the phyllomorphic stage scarcely has been attained. Glauconite does not appear to remain stable under conditions of well developed folding although it is present in sandstones which have been rather deeply buried. An arkose layer in folded Nonesuch Shale (Keewenawan) near Hancock, Michigan, contains pelletoidal grains now a mixture of secondary chlorite and biotite but which are considered on the basis of the residual shape formerly to have been glauconite (Fig.12). The rock is regarded as an example of the above equilibrium having been displaced in the direction of well-crystallized micas. Chlorite as an authigenic product of glauconite is rather common and is favored over biotite in abundance. Such a preferential development is interpreted as reflecting some general progressive step in development of a generally lower temperature mineral phase than the one which is stable under more elevated temperatures. Muscovite, also, is equally as common as chlorite and most certainly is developed at equally low temperature. A significant association can be demonstrated between the existence of former oxidizing conditions in certain rocks and the crystallization of muscovite from clays. For example, in bright colored red beds, muscovite is quite stable and can be developed authigenically, whereas biotite clearly is unstable decomposing to iron oxide and clay minerals (Fig.13). In those sandstones where chlorite is the stable phase, the predominance of reducing environment is indicated by the common occurrences of authigenic pyrite, preservation of carbonaceous fragments, and dark (organic) color of the associated shales. Biotite appears to be stable under generally similar conditions, but if both chlorite and biotite occur in the same rock the biotite appears within the chlorite mass as an equilibrium mixture. The presence of K-ion in the bulk composition of the rock as it attains the phyllomorphic grade is not considered to be important in favoring crystalIization of biotite over chlorite, inasmuch as the association of glauconite, chlorite, muscovite and authigenic orthoclase is known (Tomah Member, Franconia Sandstone, Cambrian, Hudson, Wisconsin, see Fig. 14). For reasons which are not understood, K-ion appears to be rejected during authigenesis of chlorite even though it appears to have been readily available in the glauconite.

110

E. C . DAPPLES

Fig.12. Altered pellet-shape grain ( p ) and interstitial filling now chlorite (ch) (dark) and biotite (b) (light) (crossed nicols); sandstone bed in Nonesuch Shale (Keewenanan), Hancock, Michigan. Other grains are quartz. Note growth of biotite across boundary separating pellet and interstitial fill.

Fig.13. Crystal of authigenic muscovite (m)grown through grain of oxidized biotite (b) (crossed nicols); red sandstone aspect of Minturn Formation (Pennsylvanian), Colorado. Note biotite, (ob) altered to iron oxides and clay mineral. Muscovite is considered stable in the presence of the latter two minerals.

Among certain arkoses of green-gray color biotite has formed from the interstitial clay. An example is the Stockton Arkose (Triassic), Pennsylvania, in which secondary biotite has grown in abundance replacing much detrital clay mineral (Fig. 15). This is considered to be an example of authigenesis of a preferred mica in the reducing environment in which directed pressure is a minor factor. In the same formation (Stockton Arkose), but lower stratigraphically, and within a portion in which red-colored (hematite) layers are present, the clay matrix has recrystallized into chlorite and to a lesser amount of muscovite (Fig.16). The interpretation of the association of chlorite-muscovite is that the environment of the phyllomorphic stage was slightly more oxidizing, and hence favored crystallization of the muscovite rather than biotite as in the preceding example mentioned. Crystallization of the chlorite, however, is interpreted as indicating that the environment probably ranged from very mildly oxidizing or neutral to mildly reducing. Hence, chlorite was preferentially formed over biotite which is considered to crystallize under

111

DIAGENESIS OF SANDSTONES

Fig. 14. Authigenic growth on orthoclase ( 0 ) (crossed nicols); Tomah Sandstone (Cambrian), Hudson, Wisconsin. Interstitial filling (f) is glauconite partially altered to chlorite. Other grains are quartz with overgrowths earlier in age than the complex phyllomorphic stage equilibrium. Fig.15. Authigenic biotite (b) (light) which has replaced clay mineral (dark) in matrix (crossed nicols); Stockton Arkose (Triassic), Pennsylvania. Note irregular boundaries of quartz (4)and plagioclase (p) indicative of instability of the two minerals.

more strongly reducing conditions. The reaction: chlorite

+ muscovite

reduction

z

biotite

oxidation

is interpreted as illustrating an equilibrium which represents a very mild oxidation in contrast to: reduction

muscovite

+ hematite z

biotite

oxidation

which represents reaction products of stronger oxidation. Crystallization of feldspars Authigenic feldspar has been long reported (PETTIJOHN, 1959, p.664) and is regarded as rather common in sandstones and limestones. Its distribution among sand-

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E. C . DAPPLES

stones is not as yet well understood because it is locally abundant in certain quartzglauconite mixtures of the shelf type as well as in strongly folded graywackes. Excellent examples of such authigenic feldspar-bearing cratonic rocks are the Tomah Sandstone (Cambrian) in Minnesota, and Erwin Quartzite (Cambrian) near Buchanan, Virginia. Each of these sandstones grades laterally into pure quartz sandstones containing very little, or no, feldspar. Certain investigators regard the occurrence of the secondary feldspar as primarily limited to small overgrowths around detrital feldspar grains and do not acknowledge the existence of large individual crystals as products of authigenesis. Existence of crystallization of feldspar in situ is demonstrated by the readily observed boundary between an outer rim, often with crystal terminations, and an inner core (Fig.14). In certain examples the boundary is irregular or concave which has led to the interpretation that the irregularity is aproduct of abrasion of a grain which was detrital. Some of the boundaries, however, are more suggestive of solution rather than abrasion and may represent several episodes of alternate precipitation and solution of grains which originally 1934, p.9 1). Other grains having good crystal were authigenic crystals (GOLDICH, outline are regarded as entirely authigenic not only because of their shape and unaltered appearance, but also because they may show gradational boundaries with clay at one end of the crystal (BASKIN, 1956). Graywackes are associated so often with strata of low-grade metamorphism that certain authors consider partial metamorphism of a sandstone as essential to their development (CUMMINS, 1962, p.65). In this connection the common presence of albite-oligoclase with sharp crystal outline, as opposed to the extremely embayed outline of other grains, suggests that such plagioclase is not detrital. CROOK ( 1960, p.543) described graywackes of the Parry Group (Devonian-Carboniferous) of New South Wales having detrital feldspar cut by veins of zeolites and albite, and, in some examples, completely altered to albite. Other examples show similar development of secondary albite-oligoclase. The sequence of stages

In a preceeding section the three subdivisions of diagenesis have been described as though they constituted distinct episodes, each reaching a culmination in development and subsiding in importance before the advent of the next following stage. Obviously this is not the case as the reactions typical of early stages can occur during the later stages. This is particularly true of reactions of locomorphism of which there may be several generations indicated in the rock. Some replacement reactions occur later than onset of the phyllomorphic stage as indicated by the mineral paragenesis. Certain conditions as yet unidentified result in telescoping of the diagenetic stages and early onset of phyllomorphism as part of the process of lithification. Among shelf sandstones in the most stable sections of the craton this stage is infrequently attained, and the sandstone’s history has been terminated at

DIAGENESIS OF SANDSTONES

113

Fig. 16. Authigenic muscovite (m) and chlorite (c), the equilibrium mixture altered from interstitial clay (crossed nicols); red sandstone aspect, Stockton Arkose(Triassic), Pennsylvania.Surrounding grains are quartz in various positions of light extinction. Fig.17. Tuffaceous sandstone (crossed nicols); Blairmore Formation (Cretaceous), Crowsnest, Alberta. Matrix enclosing large detrital grains has been replaced by coarse chert (c) presumably due to decomposition of original glass. Most of the photo is chert in various sizes of crystallinity and virtually none of the original matrix remains.

the locomorphic stage. In part this is due to the small amount of clay which is present in the interstitial space, but locally where the sandstones intertongue with shales and tend to have matrix clay around the quartz grains, there has been little or no development of well-crystallized mica. Where similar strata are strongly folded, however, the micas are clearly present as a thin border surrounding the quartz grains (Fig.10). Such a relation appears to emphasize the importance of pressure in the development of the phyllomorphic stage. There exist, however, some dramatic exceptions to these occurrences; as for example, the repeated abundance of micas in the more argillaceous sub-graywackes, and the occurrence of important amounts of authigenic feldspar in some of the Cambrian sandstones of the classic stable shelf area of Wisconsin. In these examples appeal to moderately elevated temperatures and pressures cannot be made. Association of very significant quantities of mica, particularly muscovite, in the sub-graywackes is connected to thin films of clay. The occurrence of the micas concentrated on bedding laminae and in the broad troughs of ripple marks has led investigators to consider such grains as universally detrital, and that they have been

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deposited in such positions as they settled out of the current transported load. The proof of some secondary mica, however, exists in its development from the interstitial clays. For this reason, much of the unaltered mica occurring in the thin clay films of the bedding laminae is believed to be of similar origin, but as yet no reliable proof of distinction has been available to the author. In the winnowed sand portions, such as stream channel fillings, muscovite is much less abundant. This is often explained as the result of removal of the mica during deposition, but it may also be the result of the inability to produce secondary mica due to paucity of clay. Oxidation and reduction reactions primarily involving iron and sulphur may occur late in the rock history. Examples of pyrite as isolated crystals or veinlets can be shown by replacement textures to have been among the latest of the authigenic minerals. They are, however, most often related to some earlier existing substance such as a fossil shell, a carbonaceous fragment, or an organic-rich shale streak. Such an association suggests that the sulphur had been previously held near the site of the pyrite in some sulphur complex attached to clay, or perhaps as a sulphate, and that late reducing conditions resulted in crystallization of pyrite. From the nature of many ground waters, however, occurrence of dissolved oxygen and oxidizing conditions are restricted largely to near-surface positions; hence, the late crystallized pyrite remains unchanged at depths not reached by aerated waters. Once lithification has been achieved, a slight tendency for a reducing environment seems to prevail but there is no evidence to point to any extensive shift of the equilibrium in either direction. Rather, the impress of oxidation or reduction during the stage of early burial appears to remain. Hence, red sandstones are known at significant depths, and old red-soil zones developed along unconformities have been preserved to be found now several thousand feet below the earth’s surface. Drilling and quarrying have demonstrated that the variety of brown and buff colors of sandstones at the outcrop do not persist at depth and that gray colors ranging from very light to dark are normal. Such colors indicate that iron tends to remain reduced except near the surface, but neither is there evidence that reducing conditions increase in intensity with increasing depth. There do exist local conditions of strong oxidation or reduction at significant depths as evidenced by such features as the “bleached zones” in red beds. Such “bleached zones” and their opposite counterparts, hematite-stained veinlets, can be demonstrated to be responses to a very localized environment or an association of minerals some of which can be oxidized and others reduced. Here, when the zone is well saturated with waters carrying electrolytes, an oxidation-reduction cell can develop and the reaction will be driven preferentially in one direction. The reader will recognize, however, that the ‘most favorable situation for such conditions exists during the early burial stage. At such a time organic matter is being oxidized and organisms are extracting oxygen from interstitial water and compounds. Conversely, abun-

DIAGENESIS OF SANDSTONES

115

dant oxygen may be present in the interstitial water and oxidation may dominate to influence the oxidation state of iron in particular. As burial progresses and the sediment is compacted, the content of dissolved oxygen diminishes and so also does the organic activity. The oxidation-reduction tendencies are reduced and reactions of this type are less likely to occur. For these reasons, reactions characteristic of the redoxomorphic stage play an increasingly smaller part as diagenesis progresses.

Reaction tendencies of the locomorphic stage Nature of the reactions Replacement of one mineral by another can be classified as one of three types, namely: (I) transition from a metastable mineral phase into a more stable form primarily as a function of time; ( 2 ) solution of one mineral and precipitation of another as a result of change in solubilities with a change in pH or temperature of permeating waters; and (3) shift in an equilibrium assemblage with change in pH, temperature, and oxidation potential. The unidirectional reaction opal-chalcedony-quartz is a good illustration of the first type. The disordered opal lattice becomes increasingly ordered in time and is replaced by either fibrous chalcedony or chert. Once chert is formed, the silica has achieved a relatively stable condition which is not perceptibly altered by moderate increases in pressure or temperature due to folding or deep burial in intracratonic basins. Chert concretions and replacements tend to remain as such. Replacement of interstitial clay by chert can be an important early event of the locomorphic stage and is particularly to be noted among tuffaceous sandstones (Fig.17). In the devitrification process silica is presumed to be released and reprecipitated in an adjacent bed, or within the same specimen, generally in the form of chert. Such chert is accommodated in the interstitial space by crowding or engulfing the detrital clay minerals present. Elsewhere, clay mineral aggregates appear to have been completely replaced which, if true, indicates destruction of the lattice. Similar chertification occurs to a limited extent where sub-graywacke sandstones underlie bentonite beds. In such examples the occurrence of secondary chert is limited to a zone several centimeters thick and is attributed to decomposition of glass and release of silica during alteration to bentonite. Chert occurs replacing clay matrix in graywackes, and much of the durability and tensile strength of the rock is to be attributed to this replacement. In each case where replacement chert is abundant, there appears to be some association with either unstable volcanic material which has decomposed to release silica, or reprecipitation of silica dissolved from detrital grains of quartz. As the reader is well aware, chert is not the common silica mineral phase to be precipitated among quartzose sandstones. Among these, crystalline quartz is

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precipitated on the grains as overgrowths and fillings of interstitial space. The circumstances in which such development can be currently associated are the socalled zones of quartz enlargement or “case hardening” of outcrop surfaces. Crystal quartz is precipitated with optical orientation on grains at the surface to produce an interlocked texture. Some localities where this process is in progress are also those where the sandstones are aquifers and the silica content in the water is not in excess of several parts per million. In the subsurface the textures may be as shown in Fig.3, illustrating that certain grains have received overgrowths whereas others have been dissolved partially. The texture is regarded as representing the existence of an equilibrium between silica solution and precipitation in the same specimen. Chert, however, is not the stable phase which is precipitated. Analyses of interstratal waters from beds in which such an equilibrium exists, indicate a pH of f 7 and general prevalence of oxidizing conditions (MEENTSet al., 1952, p.36). Interlocked textures among quartzose sandstones In a preceding section mention was made of the extreme granular interlock attained in flexible sandstone as a feature developed by solution of quartz along intergranular surfaces (Fig.5). A texture of nearly identical irregularity of granular interlock is developed in certain quartzites in which individual grains are strongly welded to one another by quartz precipitated along intergranular boundaries. Such a texture is an excellent illustration of the former existence of an equilibrium shifting between solution and precipitation of quartz. By this process grains of the original detritus gradually are dissolved, whereas new quartz is precipitated in interstices and on the surfaces of other grains which in the extreme case are authigenic. In certain examples such as the ganister illustrated in Fig. 18, pressure solution as a mechanism promoting such an equilibrium does not appear to be a suitable process. Only part of the sandstone bed has been so modified and overlying and underlying sandstones show little intergranular penetration. Rather, the process appears to be associated with movement of water of very low salinity, but saturated with silica at the low temperatures involved (KRAUSKOPF,1959; SIEVER,1962). Among the prominent characteristics of the pure quartz sandstones is the paucity of minerals other than crystalline quartz. Chert grains, for example, are exceedingly rare and heavy minerals are limited to less than 1% in abundance. Absence of minerals other than quartz is most frequently explained as due to recycling of the sandstone through more than one episode of deposition and lithification, weathering and selective transportation by currents being responsible for the virtual monominerallic composition. Evidence in favor of recycling is based upon the presence of grains showing overgrowths, the crystal edges of which have been rounded, and upon these rounded edges a second younger overgrowth has been deposited (OJAKANGAS, 1963, p.866). KUENEN and PERDOK (1962, p.656) have shown, however, that in the laboratory some of the same types of sur-

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117

Fig. 18. Quartzite (ganister) developed in sandstone below Middle Kittaning Coal (Pennsylvanian) near Pittsburgh, Pennsylvania (crossed nicols). All grains are quartz in various positions of light extinction. Note granular interlock and shape of composite grain (c) all parts of which are in optical continuity. This grain is considered to be a product of solution and reprecipitation of quartz exclusively. Fig.19. Quartz grains replaced by calcite (crossed nicols); Tensleep Sandstone (Pennsylvanian), Bighorn, Wyoming. Thin section has been cut from core of deep drill hole. Quartz grains ( 4 ) are in various positions of light extinction. Grains similar to one marked (c) are calcite cement. Note outline of remains of quartz grain ( a ) remaining as relic inside calcite. Note also extreme irregularity of partially dissolved quartz grain ( b) (dark). Individual quartz grains are isolated by the calcite cement which has replaced the position of quartz. The rock which originally was sandstone is now a sandy limestone locally.

face textures can be produced by solution and reprecipitation of quartz without requirement of recycling. Many of the rocks in question are important aquifers constituting part of artesian systems which extend from the stable craton into adjacent intracratonic basins (examples occur in the central interior of the United States). In such rocks through which an artesian flow is known, solution and reprecipitation textures of quartz are commonplace. The author prefers to regard the mineralogy of such rocks as being in large measure the result of solution or substratal decomposition of such minerals as feldspars, chert, and all but the most stable heavy minerals. In this environment chert is regarded as slightly less stable than a grain of crystalline quartz, and hence dissolves preferentially, but the silica which is reprecipitated is quartz as a secondary deposit on some other grain. According to this interpretation a pure quartz sandstone may not be entirely a product of erosion and redep-

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osition of a pre-existing sandstone; but rather, is conceived as having attained its texture and composition as a result of a special condition of mineral replacement during the locomorphic stage. Calcite-quartz relations Calcite is an extraordinarily abundant cement among sandstones, and only the exceptional thin section does not contain some secondary calcite. In some rocks it can be demonstrated to occur as a primary cement, but in other rocks several generations of calcite precipitation are recognized. Certain examples of precipitation of calcite in large crystals completely engulfing much of the original clastic material is commonplace. Special examples are the so-called “sand crystals” in which siliceous sand grains are cemented into scalenohedrons of calcite habit (PETTIJOHN, 1959, p1.8). More frequently observed are masses of calcite cement precipitated in optical orientation in siliceous sandstones. Such large crystals of calcite with their included grains can be observed to flash with reflected light from outcrop surfaces in desert regions. Precipitation of calcite is directly associated with instability of chert and quartz. In the calcite environment both show prominent solution embayments in the grains and well exhibited calcite replacement textures (WALKER,1962). In some rocks solution of silica is illustrated by extreme irregularity of the grain boundary and “ghost” remains of chert or quartz within the mass of calcite (Fig.19). Some years ago the antipathetic reaction between quartz and calcite was attributed by CORRENS (1950) to the inverse relation in solubility with change in pH. Calcite decreases in solubility and silica becomes more soluble as the p H is elevated. Although the general relations have not been challenged, the silica solubility curve has been shown by additional laboratory data to remain relatively uniform in concentrations of pH below approximately 9; and the important increase in solubility is to be observed only above such values. Inasmuch as p H values of 9 or higher are reputed to be extremely rare in nature, appeal has been made to changes in temperature as having greater control on precipitation of calcite and solution of quartz (SIEVER,1962, p.144). The occurrence of the replacement of quartz by calcite in certain sandstones at the outcrop, particularly in present-day semi-arid climates, suggests to this author, however, that shift in pH is very important in this locomorphic process despite the very small differences in solubility. Local changes in the partial pressure of COZresulting from temperature differences near the outcrop are considered to be the indirect cause of change in p H values. Similar mechanisms could produce comparable replacement in the subsurface. Removal of quartz by calcite replacement is considered to be extremely important in modifying the original percent composition of the detritus as deposited. In exceptional local examples virtually all the quartz has been dissolved and replaced by calcite leaving as a product a limestone in the place of a former sandstone (Fig.19).

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119

Although the most frequently observed relation is replacement of silica by later deposited calcite, there are examples where chert is precipitated and calcite is unstable. Generally the reaction involves replacement of detrital carbonate grains by silica, but some occurrences of a late silica cement replacing earlier calcite cement can be found. In such examples calcite is unstable and quartz is the stable mineral phase. There is reason to regard, therefore, solution and replacement of quartz and calcite as reversible reactions controlled by changes in pH, partial pressure of COZ,and temperature (WALKER,1962). Reaction tendencies of the phyllomorphic stage Nature of the reactions As defined in the sequence of diagenetic modifications the phyllomorphic stage is regarded as most advanced and transitional into the lowest grades of metamorphism (BROWNand THAYER,1963). The basis for establishment of this rank is primarily two-fold: (I) certain large crystals of authigenic mica transect some of the cement precipitated during the stage of locomorphism; and (2) selected authigenic mineral associations, developed during this stage, are considered to be equilibrium minerals representing limited ranges in bulk compositions. In earlier diagenetic stages certain reactions can be considered to represent equilibria, in the sense that they are reversible, as illustrated by the oxidized or reduced states of iron. They represent an equilibrium which is displaced in favor of the oxidized or reduced ion depending upon the presence of other oxidizable or reducible compounds in the detrital mixture. Perhaps the most common example of displacement of such an equilibrium is to be observed where abundance of readily oxidizable organic matter reduces iron to the ferrous state. Reactions characteristic of locomorphism tend to go to completion in one direction; hence, they are not to be considered as representative of typical equilibria. Instead, they can be regarded as reactions displaced strongly in one direction as a result of the primary influence of the ionic concentration of invading solution as a function of temperature, pH, and solubility products. As the conditions which lead to the precipitation of a mineral are altered, the reaction direction may be reversed, a fraction of the precipitated matter may return to solution, and be replaced by whatever else will precipitate under these new conditions. Reactions of this variety can occur during the entire post-lithification history of the rock, although they are most importantly recognized as following the stage of prominent oxidation and reduction and preceding the significant phyllomorphic modifications. Growth of authigenic mica assigned to the phyllomorphic stage begins early in the sediment history probably during the redoxomorphic stage and increases in importance with progress of time. In part this may be attributed to alteration of clay minerals such as illite into micas of more fixed crystal structure and composition. An example of glauconite crystallizing to chlorite and biotite has been pre-

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sented (Fig.11, 12). Also, crystallization of certain micas from clay minerals appears to be influenced by the state of oxidation which prevails and the pH of the permeating waters. Growth of authigenic muscovite, for example, is favored by an oxidizing environment and generally neutral solutions; whereas biotite structure is favored by a reducing condition and a wide range in acidity. With a gradual increase in application of temperature and pressure, clay minerals tend to alter more readily to micas; and the latter increase in crystal size to occupy more of the space formerly filled by clay minerals. Through this process there is an increase in particle size of the matrix material; and in some graywackes distinction between what is to be called matrix and the enclosed larger particles becomes a purely arbitrary decision. Equilibrium assemblages One of the more significant aspects of the phyllomorphic stages is the appearance of what are interpreted to be equilibrium mineral phases. Among the most readily observed are themuscovite-biotite and chlorite-biotite assemblages (Table I). These can be shown as intergrown from former masses of clay in various abundance ratios apparently dictated by conditions not currently recognized. Examples of other assemblages such as albite, quartz-chlorite-prehnite-calcite are reported by BROWNand THAYER (1963, p.414). Growth of secondary feldspar, particularly albite, in graywackes proceeds with advance in phyllomorphism. Such development is considered part of an equilibrium assemblage which appears to favor the crystallization of one mineral over others. TABLE I SUGGESTED NATURE OF SOLUTIONS DURING PRECIPITATION OF EQUILIBRIUM MINERAL ASSEMBLAGES OF THE PHYLLOMORPHIC STAGE

Acid

Neutral

Oxidizing

quartz-muscovite

quartz-muscoviteorthoclase

Neutral

muscovite-biotite

quartz-chloriteorthoclase quartz-chlorite

Reducing

biotitexhlorite

quartz-chloritezeolite quartz-chloritealbite

Basic

chlorite-albite

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Inasmuch as the authigenesis of albite has not as yet been demonstrated to characterize all Precambrian graywackes, there is some reason to suspect that the bulk composition of the original sediment and not low grade metamorphism is important in establishing the ultimate minerals of the phyllomorphic stage. For this reason the author prefers to regard certain mineral assemblages as representing attainment of the phyllomorphic stage despite their mineralogical dissimilarity. Certain of the equilibrium assemblages appear to be responses also to the degree of acidity and the state of oxidation or reduction which exists during the phyllomorphic stage. Such assemblages have been arranged in Table I in what is currently considered to represent the environment favorable for their occurrence. According to the interpretation proposed, one group of minerals is favored over a corresponding one of similar bulk composition if the oxidation state and pH of solution provide a more satisfactory environment. So little is known at present regarding such associations in sandstones and their reason for existence, that much additional information must be gathered before the validity of a table such as herein presented can be authenticated. For the present it should be regarded as the basis for a working hypothesis only.

CONCLUSIONS

The presence of authigenic silicate minerals has been so frequently reported in the mineral assemblage of sandstones that such occurrences can no longer be regarded as accidents of mineral formation. Moreover, the percent occurrence of such minerals in local specimens, or zones, may obviously represent a significant fraction of the total mineral assemblage of the specimen. Common appearance of secondary feldspars and micas normally recognized as representative of high temperature minerals is suggestive that certain progression in the development of such minerals has occurred. Indeed it can be demonstrated that as conditions suitable to their stability were attained they were able to form even at low temperatures at the expense of other minerals no longer stable. Among sandstones those minerals which appear to be most sensitive to change in conditions are silica and clay minerals. Silica alters readily into other forms, dissolves, and reprecipitates repeatedly throughout the history of certain sandstones. Clay minerals appear to be exceedingly stable in certain sandstones, whereas in others they can be recognized to have altered to mica or perhaps feldspar. The thesis which has been presented herein has been to regard such authigenesis as not accidental but rather as part of an orderly sequence of events which in the ultimate ends in a rock which has crossed the boundary of metamorphism. In this regard the stages of diagenesis have been anamorphic and new minerals take the place of some of those in the original detritus. Ideal representatives of such a process are the true and VANHISE,1892, p.306; graywackes (those of the original description, IRVING

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VANHOUTEN,1958; HUCKENHOLZ, 1963). One should not require that such rocks have very narrowly defined limits of composition, but rather that they have attained the upper limits of the phyllomorphic stage. One should look for mineral assemblages such as quartz-chlorite-zeolite or biotite-albite as among those principally diagnostic of this stage. Moreover, one should regard such equilibrium minerals as important in designating some of the primary attributes of the petrography of a graywacke. Diagenesis also moves in an opposite sense, namely, in removal of minerals unstable under the new environment. Such a direction is associated with some aspects of the stage of locomorphism. One mineral is dissolved and another takes its place. In extreme conditions solution is the only dominating trend and a sandstone of virtually monominerallic composition is the ultimate product of such an environment. Textures such as those indicated in Fig.5 and Fig.18 serve to identify s,uch a katamorphic trend of diagenesis and provide the basis for re-interpretation of the origin of certain pure quartz sandstones. Progressive stages of diagenesis provide the basis for regarding certain rocks as approaching an ultimate mineralogy in the particular tectonic environment which they have experienced. Sandstones of the typical stable shelf depositional environment are not to be expected under normal conditions to advance far into the phyllomorphic stage, if at all. Sub-graywacke sandstones typical of basins are likely to show significant attributes of the phyllomorphic stage. Those sediments of the eugeosyncline, which are constituted of important amounts of minerals unstable in the post-deposition environments, are the ones which can be expected to show the most clearly defined influence of diagenesis. One is left to regard all such processes as of significant importance and to consider that the sandstone as observed now in the outcrop, or subsurface, is perforce only partially representative of the mixture as originally deposited.

REFERENCES

BASKIN, Y . , 1956. A study of authigenic feldspars. J. Geol., 64: 132-155. S. E. and THAYER, T. P., 1963. Low-grade mineral facies in Upper Triassic and Lower BROWN, Jurassic rocks of the Aldrich Mountains, Oregon. J . Sediment. Petrol., 33: 41 1426. B. H. and RILEY,CH. M., 1955. Two unusual occurrences of microstylolites. J. Sediment. BURMA, Petrol., 25: 3 8 4 0 . CAROZZI, A. V., 1960. Microscopic Sedimentary Petrology. Wiley, New York, N.Y., 485 pp. CARROLL, D., 1958. Role of clay minerals in the transportation of iron. Geochim. Cosmochim. Ada, 14: 1-27. COOMBS, D. S., ELLIS,A. J., FRYE, W. S. and TAYLOR, A. M., 1959. The zeolite facies with comments on the interpretation of hydrothermal synthesis. Geochim. Cosmochim. Acta, 17: 53-107. C. W., 1950. Zur Geochemie der Diagenese. Geochim. Cosmochim. Acta, 1: 49-54. CORRENS, CROOK,K. A. W., 1960. Petrology of Parry Group, Upper Devonian-Lower Carboniferous, Tamworth-Nundle district, New South Wales. J. Sediment. Petrol. 30: 538-552.

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CUMMINS,W. A., 1962. The greywacke problem. Liverpool Manchester Geol. J., 3: 51-72. DAPPLES, E. C., 1959. The behavior of silica in diagenesis. In: Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Pub& 7: 36-55. DAPPLES, E. C., 1962. Stages of diagenesis in the development of sandstones. Bull. Geol. SOC. Am., 73: 913-934. DEER,W. A., HOWLE, R. A. and ZUSSMAN, J., 1962. Rock Forming Minerals. 5. Non-Silicates. Wiley, New York, N.Y., 308 pp. EADES,J. L. and GRIM,R. E., 1960. The reaction of hydrated lime with pure clay minerals in soil stabilization. In: R. E. GRIM(Editor), 1962, Applied Clay Mineralogy. McGraw-Hill, New York, N.Y., p.268. FOLK,R. L. and WEAVER, C. E., 1952. A study of the texture and composition of chert. Am. J. Sci., 250: 498-510. FRIEDMAN, G. M., 1954. Miocene orthoquartzite from New Jersey. J. Sediment. Petrol., 24: 235-242. FYFE,W. S., TURNER, F. J. and VERHOOGEN, J., 1958. Metamorphic reaction and metamorphic facies. Geol. SOC.Am., Mem., 73: 259 pp. GARRELS, R. M., 1960. Mineral Equilibria. Harper and Row, New York, N.Y., 254 pp. GLOVER,J. E., 1963. Studies in the diagenesis of some western Australian sedimentary rocks. J. Roy. SOC.W . Australia, 46: 33-56. GOLDICH, S. S., 1934. Authigenic feldspar in sandstone of southeastern Minnesota. J. Sediment. Petrol., 4: 89-95. GOLDSMITH, J. R. and GRAF,D. L., 1955. The occurrence of magnesian calcites in nature. Geochim. Cosmochim. Acta, 7: 212-230. GOLDSMITH,J. R., GRAF,D. L. and WITTERS,J., 1962. Studies in the system CaC03-MgC03FeCO3. 1 . Phase relations. J. Geol., 70: 659-687. GREENSMITH, J. T., 1957. Lithology with particular reference to cementation of Upper Carboniferous sandstones in northern Derbyshire, England. J. Sediment. Petrol., 27: 405417. HAY,R. L., 1957. Mineral alteration in rocks of Middle Eocene age, Absaroka Range, Wyo. J. Sediment. Petrol., 27: 3241. HEALD,M. T., 1955. Stylolites in sandstones. J. Geol., 63: 101-1 14. HEALD,M. T., 1956. Cementation of Triassic arkoses in Connecticut and Massachusetts. Bull. Geol. SOC.Am., 67: 1133-1154. HEALD,M. T., 1959. Cementation of Simpson and St. Peter sandstones in parts of Oklahoma, Arkansas and Missouri. J. Geol., 64: 16-30. HUCKENHOLZ, H. G., 1963. Mineral composition and texture in graywackes from the Harz Mountains (Germany) and in arkoses from Auvergne (France). J. Sediment. Petrol., 33: 914-919. IRVING, R. D. and VANHISE,C. R., 1892. The Penokee Iron Bearing Series of Michigan and Wisconsin. US.,Geol. Surv., Monograph, 19: 534 pp. KELLER, W. D., 1953. Illite and montmorillonite in green sedimentary rocks. J. Sediment. Petrol., 23: 3-10. KELLER,W. D. and REESMAN, A. L., 1963. Dissolved products of artificial silicate minerals and rocks. J. Sediment. Petrol., 33: 426-438. KLEIN,G. DEV., 1963. Bay of Fundy intertidal zone sediments. J. Sediment. Petrol., 33: 844-855. KRAUSKOPF, K. B., 1959. The geochemistry of silica in sedimentary environments. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ,, 7: 4-20. K R m , P. D., 1948. The megascopic study and field classification of sedimentary rocks. J. Geol., 56: 130-165. KIJENEN, PH.H. and PERDOK,W. G., 1962. Experimental abrasion 5. Frosting and defrosting of quartz grains. J. Geol., 70: 648-659. LOWENSTAM, H. A., 1954. Factors affecting the aragonite-calcite ratios in carbonate-secreting marine organisms. J. Geol., 62: 284-323. MEENTS,W. F., BELL,A. H., REES,0. W. and TILBURY, W. G., 1952. Illinois oil field brines. Illinois State Geol. Surv., Illinois Petrol., 66: 5-38. MILLERJR., D. N., 1957. Authigenic biotite in spheroidal reduction spots, Pierce Canyon red beds, Texas and New Mexico. J. Sediment. Petrol., 27: 177-181.

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MILLERJR., D. N. and FOLK,R. L., 1955. Occurrence of detrital magnetite and ilmenite in red sediments: new approach to significance of red beds. Bull. Am. Assoc. Petrol. Geologists, 39: 338-345. NANZJR., R. H., 1954. Genesis of Oligocene sandstone reservoir, Seeligson Field, Jim Wells and Kleberg Counties, Texas. Bull. Am. Assoc. Petrol. Geologists, 38: 96-1 17. OJAKANGAS, R. W., 1963. Petrology and sedimentation of the Upper Cambrian Lamotte sandstone in Missouri. J. Sediment. Petrol., 33: 840-814. PACKHAM, G. H. and CROOK, K. A. W., 1960. The principle of diagenetic facies and some of its implications. J. Geol., 68: 392-407. PErTIJoHN, F. J., 1959. Sedimentary Rocks, 2ed. Harper and Row, New York, N.Y., 718 pp. RUSNAK, G. A., 1957. A fabric and petrologic study of the Pleasantview sandstone. J . Sediment. Petrol., 27: 41-55. SIEVER, R., 1959. Petrology and geochemistry of silica cementation on some Pennsylvanian sandstones. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 55-79. SIEVER, R., 1962. Silica solubility, 0-200" C, and the diagenesis of siliceous sediments. J . Geol., 70: 127-151. SLOSS, L. L. and FERAY, D. E., 1948. Microstylolites in sandstone. J. Sediment. Petrol., 18: 3-13. STRINGHAM, B., 1952. Fields of formation of some common hydrothermal-alteration minerals. Econ. Geol., 47: 661-664. A., 1955. Petrology of Upper Permian rocks in south-central Kansas. Kansas Geol. SWINEFORD, Surv., Bull., 111: 179 pp. SWINEFORD, A. and FRANKS, P. C., 1959. Opal in the Ogallala Formation in Kansas. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 11 1-121 . TAYLOR, J., 1950. Pore space reduction in sandstones. Bull. Am. Assoc. Petrol. Geologists, 34: 701-71 6. THOMSON, A., 1959. Pressure solution and porosity. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 92-1 10. VANHOUTEN, F. B., 1958. Contribution to the petrology of the Tanner graywacke. Bull. Geol. SOC.Am., 69: 301-314. VON ENGELHARDT, W. and GAIDA,K. H., 1963. Concentration changes of pore solutions during the compaction of clay sediments. J. Sediment. Petrol., 33: 919-930. W. A., 1941. Cementing materials in sandstones and their probable influence on the WALDSCHMIDT, migration and-accumulation o f oil and gas. Bull. Am. Assoc. Petiol. Geologists, 25: 18391819. WALKER, TH. R., 1962. Reversible nature of chert-carbonate replacement in sedimentary rocks. Bull. Geol. Soc. Am., 73: 231-242. J. M., 1959. Compaction of sediments. Bull. Am. Assoc. Petrol. Geologists, 43: 273-310. WELLER, ZEN,E-AN, 1959. Clay mineral -carbonate relations in sedimentary rocks. Am. J. Sci., 257: 29-43. ZLN,E-AN, 1961. The zeolite facies: an interpretation. Am. J. Sci., 259: 401409. GLOSSARY

Cement, simple clay bond: a cement which forms a hardened aggregate between sand and claysized particles by a process of cohesion between the clay minerals and other clay-sized particles of the matrix of a clastic sediment. Cement, simple mineral: a cement which is precipitated in pores between grains of a clastic sediment to cause an aggregated mass. Intergranu/ar reaction bond: an interlock of grains in a clastic sediment brought about by reaction between adjacent minerals. The process is one of replacement of grains by a newly precipitated mineral. Intergranular solution bond: an interlock of grains in a clastic sediment resulting from differential solution and irregular interpenetration of grains. The interlock may be loose, as in itacolumite, or tightly cemented as in some quartzites.

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Locomorphic stage: the stage of diagenesis characterized by prominent mineral replacement, and typical of lithification of a clastic sediment. This stage is more advanced than the redoxomorphic stage. Paragenetic relations: the order in which minerals have been precipitated when they occur together in rocks as an association. Phyltomurphic stage: the most advanced stage of diagenesis in the progression toward rnetarnorphism, characterized principally by authigenesis of micas and also of feldspars. This stage follows the locomorphic stage. Redoxomorphic stage: the early stage of diagenesis characterized by mineral changes primarily due to oxidation and reduction reactions. The stage is typical of the unlithified sediment and precedes the locomorphic stage.

Chapter 4

DlAGENESIS I N ARGILLACEOUS SEDIMENTS GERMAN MULLER

Laboratorium fiir Sedimentforschung, Mineralogisch-Petrographisches Institut der Universitaf, Heidelberg (Germany)

SUMMARY

Argillaceous muds with an initial porosity of about 80 % undergo a marked physical change during the compaction. The porosity decreases continuously with increasing depth of burial; very rapidly down to about 500 m and more slowly below that depth. An increase in the orientation of the clay-mineral particles vertically to the stress direction accompanies the decrease in porosity. The behavior of porosity, permeability and structure is discussed in relation to grain size (clay content), concentration of electrolytes in superjacent and interstitial waters, mineral composition of clays, and the exchangeable cations present during the pre-burial (initial porosity and structure) and shallow-burial (depth of overburden ranges from 0 to about 500 m) stages of diagenesis. Only grain size, clay-mineral composition and temperature play an important role in the porosity reduction below a depth of about 500 m (deep-burial stage). At this depth the clay mud becomes a mudstone (or shale if fissile) with a porosity of about 30 %; the total volume of the sediment having decreased by about 50%. With a further decrease of porosity in the deep-burial stage, the mudstone (or shale) becomes an argillite with a porosity of only about 5 4 % . The lowest level of the deep-burial stage is probably at a maximum burial depth of 10,000 m. Slate is the product of metamorphism. The chemical and mineralogical changes are quite varied: during the transportation, weathering and the pre-burial stage of diagenesis in the fresh-water environment fine-grained mica minerals are mainly altered into clay micas (illite and ledikite) as a result of loss of K. In the marine environment, the cations in the exchange positions of clay minerals are changed (Mg substituted for Ca). In some marine environments chlorite- and illite-like minerals are formed by the fixation of Mg and K in montmorillonite or degraded illite brought by rivers. Volcanic glass can be altered into montmorillonite, zeolites and SiOz-minerals. In the geologic past palygorskite and sepiolite as well as huge accumulations of analcime may have formed from ionic solutions or from gels under extreme conditions and, probably, also during the pre-burial stage of diagenesis.

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Today glauconite and manganese nodules form on the surface of the sediments in many marine environments having a slow sedimentation rate. The formation of carbonate concretions and sulfides, as well as kaolinite, takes place during the shallow-burial stage. Many processes which start in the preburial stage continue slowly in the upper layers of the shallow-burial zone. During the deep-burial stage illite is formed at the expense of montmorillonite (via mixed-layer intermediate stages); and chlorite (+ illite), at the expense of kaolinite. During these processes large quantities of SiOz are liberated. The deepest stage of diagenesis is characterized by a uniform clay-mineral association “illite--chlorite”; in the transition to metamorphism it changes to a paragenesis “sericite-chlorite”.

INTRODUCTION AND DEFINITIONS

It is generally customary to apply the term “diagenesis” to all changes which take place in a freshly deposited sediment until it reaches the stage of metamorphism. According to READand WATSON (1962): “Diagenesis comprises all those changes that take place in a sediment near the earth’s surface at low temperature and pressure and without crustal movement being directly involved. It continues the history of the sediment immediately after its deposition and with increasing temperature and pressure it passes into metamorphism.” The term “weathering” covers the destruction of rocks and minerals near to or on the land surface, which is not constantly under water, by exogenic forces (such as insolation, frost, water, atmosphere, and organisms) leading to the formation of soils. TAYLOR (1964) compared diagenesis with weathering in the following way: “For the most part diagenetic changes involve increasing lithification, weathering the reverse. In a sense weathering may be regarded as retrograde diagenesis.” The largest part of the products of weathering, together with a certain amount of unweathered rock material, is carried by rivers into the sedimentary basins (mainly into the sea). After some time they are deposited there and become part of the sediment. During this stage of transportation, the rock and mineral particles undergo further mechanical and chemical weathering. It i s to be expected that the chemical weathering of silicates, especially of those which have not undergone the process of soil formation, takes place to a far greater extent here than on land because large amounts of water are available in which the particles are constantly moved (stirring effect). This applies in particular to the most finely grained material because of its relatively large surface area. These particles are particularly strongly reactive. In addition, these particles can remain floating in the marine sedimentary basins for centuries because of the low settling velocity.

129

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

In the transition from the fluviatile to the marine environment, the physicochemical conditions change radically. The still floating minerals go through the same process of adjustment as the particles which have already settled on the surface of the sediment. The kind of weathering which shall hereafter be called “transportationweathering” (Table I) is essentially different from the weathering on the continent which shall hereafter be referred to as “in-situ weathering”. In fresh-water environments it differs mainly through its intensity, whereas in marine and saline environments both through its intensity and through being different. Inasmuch as many physicochemical and chemical processes are similar or even identical during the subaqueous transportation-weathering and the earliest stage of diagenesis, which takes place in the upper layer of the sediment and is largely controlled by the chemistry of the subjacent water, it seems advisable to deal with these processes together and to give them the same name. A possible choice for an overall term for the processes occurring in a marine (or saline) (1 922) (halmyros environment is the term “halmyrolysis” introduced by HUMMEL = salty; lysein = to dissolve). Thus, halmyrolysis can be defined as: all chemical and physicochemical processes which occur during the marine transportation, weathering and the marine pre-burial stage of diagenesis. PACKHAM and CROOK (1960) have already used the term halmyrolysis in this sense. Because an overall term for the physicochemical processes occurring in the TABLE I RELATIONSHIPS BETWEEN PARENT ROCK, WEATHERING, DIAGENESIS AND METAMORPHISM

P A R E N T I

-K

0

1

w

I

I

1

R O C K

in-situ weathering

L

\1

fresh-water t r a n s p o r t a t i o n - w e a t h e r i n g

3,.. ....’ *’ Holmyrolysis”

‘~quotolysis“

’.__

v)”... c (

marine pre-burial stage

fresh-water pre-burial stage

marine shollo\v- burial stage

fresh-water shallow-burial stage

v)

W

: n

\1

I/

deep-burial stage

L.

M E T A M O R P H I S M

130

G. MULLER

fresh-water environment is lacking, the author proposes the term aquatolysis (aqua = water; lysein = to dissolve) for all processes which take place in fresh water during transportation-weathering and in the earliest stage of diagenesis. Thus, aquatolysis may be defined as: all chemical and physicochemical processes which occur during transportation-weathering and pre-burial stage of diagenesis in fresh-water environment. The various stages of diagenesis can be divided into: ( I ) pre-burial stage, (2) shallow-burial stage, and (3) deep-burial stage. In argillaceous sediments, the depth at which the sediments are buried under younger deposits is more essential for the diagenetic evolution than the length of time which passes after the deposition of the sediments. This is especially true of the physical changes which take place in the sediment in the course of diagenesis. This becomes particularly clear by a comparison of a sediment core from Lake Zurich with those from the Santa Barbara Basin and the Black Sea (Table V). Although the sediments in the Santa Barbara Basin and the Black Sea, covered by about 5 m of younger deposits, are at least ten to a hundred times oldel than those in Lake Zurich, the decrease in porosity with increasing depth is almost identical. Obviously, the time during which a sediment is situated in a certain physicochemical environment is important as far as the chemical and mineralogical processes are concerned. The pre-burial stage is easiest to define. It comprises the physicochemical processes (halmyrolysis and aquatolysis) which take place in the upper (youngest) layers of the sediment in the presence of oxygen. To differentiate between the shallow-burial and the deep-burial stages is much more difficult. In the case of argillaceous sediments, the line can be drawn where a soft clay mud becomes an indurated, firm and coherent mudstone (mudstone fissility = shale). This consolidation can be brought about by compaction alone; however, there is often a combination of compaction and cementation. The shallow-burial and deep-burial stages of diagenesis thus correspond to the two principal stages in the lithification of argillaceous sediments proposed by LOMTADZE (1955) on the basis of experimental work: ( I ) the conversion of argillaceous mud to mudstone (or shale), and (2) the conversion of mudstone to argillite (employing the term argillite to mean a non-metamorphosed rock). Slate is a metamorphic rock. In terms of porosity, the point at which a clay mud is converted into mudstone lies at about 30 %. At this porosity the initial volume of the clay mud has decreased by about 50 %. HAMILTON (1959), who studied the relationship between overburden pressure and porosity, observed that no matter how porous the original material, by the time it is under about 100 kg/cm2 of pressure the porosity has usually decreased to about 29 % for various clay-rich sediments. In the Tertiary clay sediments of the Po Basin and of Venezuela the 30% porosity limit is reached when the sediment cover is about 500 m thick.

+

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

131

For subaqueous argillaceous sediments STRAKHOV (1 956) assumed that lithification is almost complete at a depth of about 250-300 m below the depositional interface. As compared to the near-surface sediment layer the increase in pressure can amount up to 80 atm and the increase in temperature up to an average of 9 “C. At the boundary between shallow- and deep-burial zones, the depth and pressure are around 500 m and 50-100 kg/cm2, respectively. Laboratory experiments on clay-water systems showed that at pressures over 50 kg/cm2 the influence of the electrolyte content and the exchangeable ions is insignificant. There are essential differences between marine and fresh-water sediments, because the chemical-mineralogical and partly also the physical changes in a sediment during diagenesis depend mainly on the chemistry of the water (subjacent and interstitial) in contact with the mineral particles. This applies at least to the first two stages of diagenesis; in the third stage it can be assumed that the pore solutions in most cases have already become similar in composition. In this chapter “argillaceous sediments” are defined as fine-grained sediments (with an average grain size in the range of about 1 - lop) which mainly consist of silicate clay minerals (chiefly layer-silicates). These types of sediments are usually to be found in the deeper offshore regions of marine and lacustrine basins.

INITIAL COMPOSITION OF RECENT ARGILLACEOUS SEDIMENTS

In order to understand the diagenetic changes in inorganic and organic matter, it is essential to be familiar with the initial composition of a clay sediment. Among the most important allogenic components which come from outside the sedimentary basin are: (1) clay minerals including gibbsite, (2) quartz, (3) feldspars, ( 4 )carbonates, (5) amorphous silica and alumina, (6)pyroclastic material, and (7) organic matter. In addition, the biogenic carbonates, biogenic amorphous silica and organic matter are formed in the basin itself. The clay minerals and quartz are most stable against changes in the physicochemical environment, whereas the pyroclastic materials are the least stable. The important allogenic and biogenic materials of the modern clay sediments are described below. Clay minerals (including gibbsite) The clay minerals of the Atlantic and parts of the Indian Ocean were thoroughly studied by BISCAYNE(1964); and those of the Pacific, by GRIFFIN and GOLDBERG (1963). In the Atlantic Ocean, illite, kaolinite, chlorite, montmorillonite, a random mixed-layer mineral, and gibbsiteform the essential part of the clay fraction ( t2p).

132

G. MULLER

Of these “kaolinite, gibbsite, mixed-layer minerals and chlorite contribute the most unequivocal information because they have relatively restricted loci of continental origin” (BISCAYNE, 1964). As a rule, illite is the dominant clay mineral. In the northern Atlantic the illite content amounts to more than 90 % of the clay mineral fraction. Kaolinite is the second most important clay mineral. A large increase in kaolinite content (up to more than 50%) and of gibbsite (more than 10%) can be observed in sediments adjacent to tropical rivers in Africa and South America. Extremely high gibbsite contents (more than 30%) are to be found in the Indian Ocean sediments around Madagascar. As a rule, the chlorite content in the clay-mineral fraction is less than 20 %, although larger amounts can be found in Antarctic regions as well as off Newfoundland. In the North Atlantic area montmorillonite accounts for up to 20 %, whereas in the South Atlantic montmorillonite constitutes up to 40% of the total clay minerals. Extremely high values (over 80%) are to be found in the southwestern Indian Ocean. BISCAYNE (1964) pointed out that in this region an in-situ formation, i.e., a diagenetic formation, from volcanic material is possible. There are no mixed-layer minerals east of the Mid-Atlantic Ridge. West of the Mid-Atlantic Ridge, their frequency decreases from north to south; the d-value of the first basal spacing increases. The first basal spacing lies between 10.7 and 13.3 A in the non-glycolated samples. Biscayne gave no further details on the composition qf the mixed-layer minerals, which probably consist of illite-montmorillonite components. Because of the strong geographic control of the mixed-layer minerals distribution, Biscayne assumed continental origin rather than in-situ diagenetic origin. In the newly formed sediments of the Atlantic Ocean zeolites were observed in small quantities. In the Pacific Ocean, the illite, montmorillonite, chlorite and kaolinite, as well as halloysite (to a much lesser extent), are the main clay minerals according to GRIFFINand GOLDBERG (1963). In the North Pacific area, abundant illite was found in all samples, whereas montmorillonite (mainly nontronite), chlorite and kaolinite are generally present and their abundance is a function of geographic location. Montmorillonite is generally more abundant in nearshore sediments, whereas chlorite increases with increasing latitude in nearshore sediments. Kaolinite is confined to nearshore areas. In the South Pacific area, montmorillonite (mainly nontronite) is the most abundant clay mineral; illite, kaolinite and chlorite are less abundant. In strong association with montmorillonite is phillipsite. In this chapter the enumeration of the clay minerals occurring in recently formed marine argillaceous sediments must be restricted to those which occur most frequently. Naturally, it is possible that local accumulations of relatively

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

I33

rare layer-silicates are to be found (talc, pyrophyllite, serpentine etc.), but in such a case their occurrence always depends greatly on the presence of these minerals on the adjacent continent. For example, there are considerable amounts of pyrophyllite in the basins adjacent to Central America and the northern part of South America, pointing to source areas on the northern coasts of Columbia, Venezuela, French and British Guiana, Surinam and Brazil, and the southern parts of Cuba, Haiti and the Dominican Republic (BISCAYNE, 1964). The clay-mineral composition of lake sediments is determined to an even greater extent by the hinterland. The most important clay minerals here are also illite, kaolinite, chlorite and montmorillonite, as well as random mixed-layer illite-montmorillonite.

Quartz and feldspars Quartz and feldspars (particularly plagioclase) are the most important allogenic minor constituents of argillaceous sediments. REXand GOLDBERG (1958) demonstrated a regional regularity in the quartz distribution in the Pacific sediments, which they interpreted as an indication of fall-out of dust from the high altitude jet streams. In the Atlantic the broad zone, with relatively high content of quartz, west of North Africa is attributable to eolian quartz from the Sahara (RADCZEWSKI, 1939). The increase in quartz content near Antarctica and in the northern Atlantic probably stems from glacial outwash (BISCAYNE, 1964); a similar tendency for feldspars in these areas has also been generally noted and explained in the study of Biscayne. For the Atlantic, RADCZEWSKI (1937) and LEINZ(1937) showed that acidic rocks must have been the main source of feldspar in the Cape Verde and the Guinea basins. The majority of the plagioclases showed a composition of h 3 0 , whereas the dominant plagioclases in most pacific sediments range in composition between labradorite and oligoclase. The feldspar distribution in South Pacific pelagic sediments was investigated by PETERSON and GOLDBERG (1962). Most of the feldspars are of volcanic origin, and several source areas could be established (especially for basic and acidic groups). The composition of the plagioclases ranges between oligoclase and anorthite. Alkali feldspars (sanidine, orthoclase) also show a wide distribution.

Carbonates Allogenic carbonates (such as calcite and dolomite) from outside of the sedimentary basin can be assumed to play only a very minor role in argillaceous sediments as compared to the biogenic carbonates (mainly calcite and aragonite). Between the argillaceous muds with no or little carbonate content (red clay) and pure biogenic oozes (Globigerina ooze covers enormous areas in the Atlantic, Pacific and Indian Oceans) all transitions are possible.

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G . MULLER

Amorphous silica and alumina

I n argillaceous marine deposits amorphous silica (and alumina in some cases) may become very abundant. In the sediments of the Atlantic Ocean, amorphous silica content ranges from 1 to 56 % (BISCAYNE, 1964). Most of this silica is biogenic and was produced by planktonic diatoms, radiolarians and silicoflagellates, and some benthonic sponges. An alumina content of 0-3 % found in the same sediments by Biscayne may be accounted for by the varying gibbsite contents of the samples; gibbsite is strongly attacked as a result of the leaching procedure for the determination of the amorphous silica. The presence of amorphous alumina, therefore, cannot be definitely confirmed. Sediments with high biogenic opaline silica are found in the Subarctic Convergence, the Equatorial Divergence and the divergences along the west coasts of the continents (ARRHENIUS, 1963). Accumulations of non-biogenic amorphous silica and alumina (up to 60 %) were reported by MOBERLEY (1963) from sediments adjacent to Hawaii, which are derived from tropically weathered basalts. In fresh-water lakes small amounts of biogenic opaline silica are mainly produced from diatoms; yellow-brown Algae and siliceous fresh-water sponges are of minor importance. During the Ice Ages diatoms played a much more important role in sedimentary processes in lakes. Pyroclastic material The main location of recent volcanism is the area around the Pacific Ocean (circum-Pacific Circle). More than 3 of all active volcanoes are situated here. According to Sapper (cf. BRINKMANN, 1961), between 1500 and 1914 about 18 km3 of lava and 330 km3 of pyroclastic material were erupted from the 339 active volcanoes in the circum-Pacific Belt. During the eruption (in 1883) of the Krakatau in the Sunda Straits, about 18 km3 of pyroclastic material were hurled up to 50 km into the atmosphere. Extremely fine-grained volcanic ash circled the earth several times. In addition to this terrestrial volcanism mainly limited to the rims of the continents, there is also a submarine volcanism the dimensions of which can hardly be estimated. It may be assumed that submarine volcanism has played an important role, especially in the Pacific area. Huge amounts of palagonite, which spread over enormous basin areas associated with the seamounts and guyots, are a sure indication for submarine volcanism. The tuffs of terrestrial volcanoes, mainly consisting of volcanic glass, either land directly in the sea during eruption or are carried into the sea from already existing tuffs on the continent. Acid porous glass (pumice) because of its low specific weight due to the high porosity can remain drifting in the sea for a relatively long period of time.

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

135

In the case of submarine eruptions, the conditions of genesis of the partially devitrified glass (palagonite) are completely different, because it is not the high gas content of the magma but the interaction of hot basic lavas with cold sea water which is decisive. According to BONATTI(1963): “Palagonite has a high water content (10-30%) and its only crystalline components appear to be nontronite and occasionally goethite. During a submarine volcanic eruption the rapid chilling of the hot lava in the water appears to cause granulation or pulverization of the hot lava itself. The part of the magma directly reacting with the water forms palagonite, which often encloses and protects fragments of lava from direct contact with water, thus allowing their rapid cooling to form sideromelane glass with eventual crystalline inclusions.” In the Pacific, the palagonite distribution generally coincides with the area of phillipsite occurrence. In the recently formed sediments of the Atlantic area, volcanic glass occurs much less frequently; accumulations of glass ought to occur particularly in the immediate neighbourhood of volcanic areas (Mediterranean, Cape Verde Islands, Canaries, Azores, Central American volcanic provinces, etc.). For example, the Recent sediments in the Gulf of Naples are almost exclusively formed from volcanic glass and its alteration products (MULLER, 1961). Organic matter

The organic matter content in Recent marine and fresh-water argillaceous sediments ranges from about 1 to 2 %, most of the samples containing 2-5 %. According to DEGENS (1967), the organic matter in clay muds consists mainly (60-80 %) of chemically undefined “organic residue” plus less than 10 % of amino acids. INITIAL (PRE-BURIAL)

POROSITY AND STRUCTURE OF ARGILLACEOUS SEDIMENTS

Initial porosity

The initial porosity and thus the water content1 of argillaceous muds is very much higher than that of sands. Clay muds from Recent sea and lake bottoms as a rule have a porosity of 70-90% corresponding to a water content of about 50-80%. In sands porosity is only 30-50 % which corresponds to 20-30 % water content. Fig. I shows the dependence of the water content of Recent clay muds on the amount of the clay fraction (<2,u) in different sedimentary basins: with increasing clay content the water content and porosity increase. Sands essentially consist of more or less rounded particles of quartz and feldspar, the geometrical arrangement of which can be compared with packings of spheres. With decreasing grain size, the amount of layer-silicates increases. 1 In

this chapter the water content is expressed as a percentage of the wet weight.

G. MULLER

20

40 '1. F r a c t i o n

60 2p

80

Fig.1. Correlation between water content and percentage of clay fraction (4,~) in different environments; I = Lake Constance (after G. Miiller, unpublished data); I1 = Zuider Zee (after 1956); 111 = Mississippi delta (after SHEPARD and MOORE,1955); Wiggers and Smits, cf. SEIBOLD, IY = Rockport area, Texas (after SHEPARD and MOORE,1955); V = California Basin (after EMERY and RITTENBERG, 1952). Environments: I = fresh water; 11-IV = brackish; v = marine.

From the purely geometrical point of view clay minerals could aggregate to a compact sediment with only a very low porosity. The fact that this is not the case with argillaceous sediments shows that their porosity cannot be understood only on the basis of simple geometrical models as is qualitatively possible with sands. The special behaviour of layer-silicates is due to their extremely large surface areas. The sediments of the different sedimentary basins with the same amount of clay-size fraction (or the same average grain size), have different water contents. According to MEADE(1964), this may be due to the following important factors: (I) influence of clay-mineral composition, (2) influence of interstitial electrolyte solutions, (3) influence of exchangeable cations, and ( 4 ) rate of sedimentation. In addition, alkalinity or acidity and associated organic matter can also play a role. These influences, however, have not yet been studied in detail. Injluence of clay-mineral composition If pure clay minerals are allowed to settle in water, it becomes evident that in the resulting sediments montmorillonite retains more water than illite, and illite in turn retains more water than kaolinite. These differences are mainly due to the different grain-size distributions and to the differences in the specific surface areas.

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

137

According to MEADE (1964), the specific surface area of montmorillonite is 800-600, of illite 100-65, and of kaolinite 30-5 m2/g. The effect of specific surface area on water sorption is that a montmorillonite particle, extremely fine-grained, adsorbs a water envelope of the same thickness as that on a relatively large kaolinite particle; therefore, in comparison to its mass more water is adsorbed by montmorillonite than by kaolinite.

Influence of electrolyte solutions In natural and artificially produced sediments there is a relation between the water content of a freshly deposited sediment and the electrolyte content of the depositional medium. With increasing electrolyte content, water content decreases. Clay-rich fresh water sediments, therefore, as a rule contain more water than comparable marine sedimentsl. This behaviour can be explained by the different forces (Van der Waals attractive forces between the clay plates and repulsive forces due to the presence of diffuse electrical double layers on the charged clay plates-"osmotic swelling"), which have a mutual effect on the neighbouring clay particles. For details on this subject the reader is referred to the work of OVERBEEK (1952), BOLT (1956), LAMBE (1958a,b, 1961), MARTIN(1962) and MEADE(1964). High electrolyte contents bring about a coagulation of the particles. This is important for the behaviour of clay minerals in the transition from fresh water to salt water. The coagulation effect of the ions increases rapidly with increasing valency (Schulze-Hardy valency rule). The concentrations of uni-, di- and trivalent cations, which are necessary for flocculation are in the ratio of 500/10/1. The sequence of the flocculation intensity is: AI>Ca>Mg> K > N a

Influence of exchangeable cations Experiments carried out by SAMUELS (1 950) on relatively coarse-grained kaolinite showed that the effect of exchangeable cations on the initial water content of the sediment is only slight. Kaolinites saturated with Al have larger water contents than Ca- and Na-saturated kaolinite, but the differences between the latter two clays are extremely small. With relatively fine-grained montmorillonite the differences are much greater and the situation is reversed: Na-saturated clays contain more water than Camontmorillonites, and these in turn, contain more water than Al-montmorillonites. For further discussions on this subject see MEADE(1964). Rate of sedimentation Up to now no comprehensive studies were made on the dependence of initial According to BOLT(1956), the effect of electrolyte concentration seems to change with particle size. In very fine-grained clays (<0.2p fraction of Fithian illite), the larger porosities are associated 1960, and MEADE,~ 964.) with the lower concentration of electrolyte. (For discussion see MITCHELL,

138

G. MULLER

porosity on the annual rate of sedimentation. Observations made by FUCHTand REINECK (1963) in the southern part of the North Sea seem to indicate that with a high rate of sedimentation the porosity is greater than with a slow rate. Clay muds from a Recent bay (rate of sedimentation up to 50 cm/year) had a porosity of 83 % on the sediment’s surface, whereas clays from the Wadden Sea and the foreshore had a porosity of only about 70 %. Also the relatively very high porosities in the uppermost few centimeters of fresh-water sediments in lakes (Table IV, Fig.7), with an estimated sedimentation rate of about 1-5 mm/year, seem to indicate that with this high rate of sedimentation an abnormal porosity exists only for a short period of time. As the porosities of the lower layers show, they are reduced to the normal values after a very short period of time. BAUER

The combined efSects of different factors in natural sediments In natural argillaceous sediments, several clay minerals occur together, as a rule, and each clay mineral can occur in different grain-size classes; furthermore, varying amounts of organic substances and non-clay minerals can be present with the clay minerals. Consequently, the combined effects of several single factors described above and their interrelationships are very complicated and cannot be investigated without exact knowledge of the sediments in question, which in most cases is not available. As shown in Fig.1, in comparison to the marine sediments the fresh-water sediments of Lake Constance, with a clay content of less than 40 %, show the expected high porosity. With a clay content of more than 40 %, however, this situation is reversed. This behaviour could be explained if both sediments had a high content of particularly fine-grained clay minerals, as explained earlier (effect of electrolytes). One cannot ascertain, however, whether this is the case, because only the fraction < 2p was determined, without making a further subdivision. Initial structure

The dependence of primary porosity on particle size, type of clay mineral, electrolytecontent of the depositional medium, and exchangeable cations, as shown in the previous section of this chapter, determines the structure of the newly formed clay sediments. It is generally assumed that the structure of freshly deposited argillaceous mud is an open-work arrangement of particles dominated by contacts between corners and planes (edge-to-surface contacts) as shown schematically by TAN (1958). ROSENQUIST (1962) showed this fabric in electron micrographs of several marine clays (Fig.2). This structure known as the “house of cards” structure is due to the fact that the surfaces of clay particles bear a net negative charge but, due to broken bonds, the edges tend to be positive. Inasmuch as the electrostatic bond be-

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

139

Fig.2. Electron micrograph showing mutual arrangement of minerals in the blue Oslo clay. (After ROSENQUIST, 1962.)

tween edge and surface is rather weak and would result in a rather unstable association of particles in a mud, MUNGANand JESSEN(1963) proposed a new structural scheme: the “hinge” structure. This structure comprises clay particles oriented randomly to form a network in which each particle is “hinged” to other particles through shearing a bound particle-water. In fresh-water clays ROSENQUIST (1962) observed a greater degree of parallel orientation between clay particles. He related the random orientation of marine clays to the flocculating effect of the electrolytes in sea water. From observations on the sedimentation volume of illite and kaolinite with variable electrolyte concentrations made by ROSENQUIST (1955) and HSIand CLIFTON (1962), MEADE (1964) concluded that the tendency toward random orientation of illite and kaolinite particles is a direct function of the electrolyte concentration. In case of montmorillonite clay, HOFMANN and HAUSDORF (1945) found the opposite behaviour. The equilibrium volume of the settled-out montmorillonite decreased with increasing electrolyte concentration (NaC1 and KC1 solutions). For further discussion of this subject see MEADE(1964).

140

G. MULLER

Fig.3. Two-dimensional model of the aggregate structure. (After VON ENGELHARDT and GAIDA, 1963.)

From the behaviour of permeability during the compression of clays with different electrolyte contents in their pore solutions, VONENGELHARDT and GAIDA (1963) concluded that clays which settled from solutions rich in electrolytes have a heterogeneous or aggregate structure (Fig.3). Porosity is due to the internal pore spaces between aggregates. Clays deposited from solutions poor in electrolytes will consist of small aggregates or of free primary particles. Orientation, therefore, will be good and permeability low because the channels for fluid flow are very small. With higher electrolyte concentration, aggregates will be large, orientation poor, and permeability high.

CHANGES IN CHEMISTRY AND MINERALOGY DURING DIAGENESIS

Changes during transportation-weathering and pre-burial stage (aquatolysis and halrnyrolysis) A quatolysis The clay minerals carried by the rivers and deposited in lakes and oceans are derived from soils, or directly from the outcropping sedimentary, igneous and metamorphic rocks in the hinterland (Table I). The minerals of soils and, to a large extent, also those of the sedimentary rocks are largely adjusted to the conditions on the earth surface. This, however, does not always apply to those minerals of igneous and metamorphic rocks which have not undergone the soil-forming weathering processes. Analyses for the K content of clay-grade mica minerals, suspended in the Alpenrhein (Rhine river before it enters Lake Constance) and deposited in

141

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

Lake Constance, and which may have been derived from micas of outcropping igneous and metamorphic rocks in the Alps, show a remarkable loss of potassium reaching up to 20 % of the initial K content. There is a corresponding increase in the water content (hydroxonium) (G. Miiller, unpublished data). In this case there is a very rapid process of the formation of clay minerals from micas during the subaqueous transportation. The mechanical breakdown and aquatolysis result in the formation of illite (dioctahedral) from muscovite; and ledikite (trioctahedral), from biotite. It is probable that after deposition these processes also continue during the pre-burial stage. According to GRIM (1958), it is possible that mixedlayer structures could develop from micas by aquatolysis, but apparently no such process has yet been described. A pronounced aquatolysis also occurs in the case of K-rich feldspars. In the clay minerals already derived from soils or sediments, or other layersilicates of the metamorphic rocks, no considerable change is to be expected to occur during the subaqueous transportation. Halmyrolysis Cation exchange and dissolution of clay minerals. In the transition from fresh water to marine and saline environment such marked changes (qualitative and quantitative) take place in the electrolyte content of the water, that one must carefully study the cations absorbed on the clay-mineral surfaces. Pons (1959, cf. KELLER, 1963) immersed clay (mainly montmorillonite) suspensions of the Missouri River in sea water for 36 and 86 h. The data, showing a marked increase in amounts of exchangeable Mg, Na and K ions and a decrease in total exchangeable ions, are presented in Table 11. In a more compreTABLE? I1 EXCHANQEABLE CATIONS OF MISSOURIRIVER CLAY TREAIZD IN SEA

WATER^

(After P m , 1959) Cations

Untreated chy

After 36 h in sea water

After 86 h in sea water

Largest increase or decreases

ca

60.7 20.1 1.7 I .4

38.3 29.7 1.8 1.4

17.3 39.3 3.4 2.0

decrease to 34x5 increase to 195% increase to 200 % increase to 140%

Mg

Na

K

1 All figures are

given in mg eq./100 g dry material. original amount of exchangeable cations.

a With respect to

8 sample calculation:

17.3 60.7

100 = 34%.

142

G. &LER

hensive study, CARROLL and STARgBY (1960) immersed samples of montmorillonite, a mixed-layer mineral (mica and montmorillonite), illite, and kaolinite in sea water for 10 days. Additional samples of the first three clay minerals were immersed for 150 days. The results are given in Table ID,showing: (I) a marked increase in the amount of exchangeable Mg ions; (2) a decrease in total exchangeable ions and exchange capacity in the three-layer clay minerals, montmorillonite and illite; and (3) an increase in exchangeable ions and exchange capacity in the two-layer clay minerals, kaolinite and halloysite. The decrease in total exchangeablecations and the cation exchange capacity in river-borne montmorillonite (F‘om, 1959) and mixed-layer montmorilloniteillite (CARROLL and STARKEY, 1960) was interpreted by KELLER(1963) as being due “to conversion of easily exchangeablecation sites to fixed cation sites or to some on which tightly bonded ions, such as K, are being held too tightly to be exchanged by ordinary cation exchange capacity reaction”. The possibility of fixation of K in expandable three-layer clay minerals, that at a later stage can lead to the formation of illite, is, therefore, strongly indicated. (1960) on the solubility of clay Studies carried out by CARROLL and STARKEY minerals in sea water showed that small quantities of Si02 and Al203, totalling less than 1% were removed from the following minerals in a decreasing order: (I) montmorillonite, (2) mixed-layer minerals, and (3) kaolinite and halloysite. Greater quantities of Si02 and were removed in 150 days than in 10 days. Transformationof clay minerals into illite, sudoite and chlorite. Transformation and neoformation of clay minerals, which already take place when detritic clay minerals are transported from fresh water into the marine environment, have been reported by many authors. Frequently, these observations are based only OD a comparison between the clay minerals found in a sedimentary basin and those transported there from the hinterland. The interpretations deduced from this must be treated with extreme caution because several other factors in addition to authigenesis could be responsible for differences in the composition of the clay (1959) emphasized this fact strongly and among these factors minerals. WEAVER included preferential flocculation, current sorting, floods, periodic variations in composition and concentration of river detritus, “and probably numerous factors of which we are not aware”. A good example of transformation of clay minerals as a result of the change of environment is the transformation of montmorillonite into chlorite described by GRIMand JOHNS (1954). In the Rockport area in the Gulf of Mexico a small river (Guadalupe River) forms a delta in a sheltered bay (San Antonio Bay). The predominant clay mineral in the Guadalupe River is Ca-montmorillonite. On entering the initially brackish and then increasingly marine milieu, the montmorillonite is transformed stepwise into chlorite with increasing salinity. First the Ca-montmorillonite is saturated with Mg through ionic exchange, then

143

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

TABLE J I l SOME EXPEUMF3l"T CLAY MMERALS UNDER VARIOUS CONDITIONS IN SEA WATER

(After CARROLL and STARKEY, 1960,p.87) Mineral

Montmorillonite ( A ) Natural Sea water, 10 days2 H-form(a)a Sea water, 10 days4 Sea water, 150 days5 Mixed-layer mineral (B) Natural Sea water. 10 days* H-f~rm(a)~ Sea water, 10 days4 Sea water, 150 days6

Exchangeable cations (meq.llW g) Ca2+

Mg2+

Na+

total

Cation Percentage of exchange capacity exchange1 (determined)

11 7 8 7 8

15 34 14 27 32

54 16

80 57 22 55 61

89 93 66 91 76

90 61 33 60 80

26 3 3 9

7 12 1 8 7

33 17 4 19 17

33 31 34 21 28

100 55 11 90 61

20 20 12 19 16

20 20 24 21 28

100 100 50 90

0.9 2.6 0.9 3.1

5 8 13 16

18 32 7 19

4.1 10.3 2.6 10.1

11 47 42 31

37 22 6 32

5

UIllite"(C)6 Natural Sea water, 10 days2 H-form(a)a Sea water, 10 days4 Sea water, 150 days6

17 12 10 9 8

Kaolinite (0) Natural Sea water, 10 days2 H-form(a)S Sea water, 10 days4

0.5 0.9 0.9 1.4

-

Halloysite (E) Natural Sea water, 10 days2 H-f~rm(a)~ Sea water, 10 days4

1.8 2.7 1.8 4.1

1.7 7.1 0.8 5.4

1 Positions

-

21 21

2

2 5

0.4 1.7 1.7 0.6 0.5

-

0.6

51

filled by cations other than H+. Natural mineral soaked in 50 ml of sea water. 8 Clay mineral (montmorillonite, kaolinite, etc.) treated with Hf ions. All exchangeable sites are occupied by H. Prepared by pouring a slurry of clay mineral and water through a column of Amberlite IR-120resin in the H-form. 4 H-form(a) soaked in 50 ml of sea water. 6 H-form@) soaked in 50 ml of sea water. H-form@): as the treatment with resin did not give the minimum pH expected for some of the minerals, treatment in 1 :3 HCI (a. 3 N) at room temperature (25"C)was used subsequently. 6 "IUite" (c) contained about 2% calcite as impurity; the figurcs given above have been corrected for Ca due to calcite. a

144

G. MULLER

brucite is precipitated as isolated islands between some of the silicate layers, and montmorillonite-chlorite mixed-layer minerals form. As the formation of brucite between the layers continues, isolated islands of brucite grow laterally until finally a chloritic mineral comes into being. JOHNS (1963) proved the authigenesis of chlorite (which was still doubted 1959) by the determination of a regular increase in the content of by WEAVER, chlorine, incorporated in the brucite layers of the chlorite, with increasing salinity and increasing conversion of montmorillonite into chlorite. Relationship between the mineral content, salinity, and the chlorine distribution in the fraction < 1,u is presented in Fig.4. VONENGELHARDT (1961) pointed out that the chlorites which formed in sea water from montmorillonite (dioctahedral) should deviate in composition from the normal chlorites (combination of a trioctahedral talc-like sheet with a trioctahedral brucite sheet), and should have the combination of dioctahedral

‘L 300

6 200 100 0

0

0

0

0

c

/

/

0

/

/

/

0

0

0

/

0

0

/

1

Fig.4. Chlorine content and mineral distribution in fraction
DIAGENESIS IN ARGILLACEOUS SEDIMENTS

145

(montmorillonite- or pyrophyllite-like) with trioctahedral brucite-like sheets. GRIMand JOHNS (1964) observed a dioctahedral bo-parameter for the freshly formed chlorites. Inasmuch as the a-b parameters are the most important criteria for differentiation of di- from trioctahedral minerals, this chlorite must be classified in the sudoite series which together with the chlorite series form the sudoitechlorite group (MULLER,1963). Indications of recent chlorite formation were also observed by POWERS (1954) on the east coast of the United States (Chesapeake Bay) as well as by PINSAK and MURRAY(1960) in the Gulf of Mexico. GRIMet al. (1949) suggested that degraded illite was slowly rebuilt in sediments from the Pacific Ocean off the California coast and from the Gulf of California by the aquisition of K and Mg ions from the sea water. GRIMand LOUGHNAN (1962) described the formation of illite in the marine environment of Sydney Harbor, Australia. On passage from the fresh-water environment of the Lane Cove and Parramatta Rivers into the harbor, the interlayer ferric iron of vermiculite (14 A mineral) is reduced and the mineral becomes degraded. In the marine environment illite develops from this degraded material. The transformation of montmorillonite into chlorite or sudoite and illite was con6rmed by laboratory experiments. WHITEHOUSE and MCCARTER (1958) treated montmorillonite suspensions with artificial sea water, and after 3 years observed a growth of illitic and chloritic materials. The transformation of illite or kaolinite into chlorite was not observed. Formation of clay minerals in supersaline environment (chlorite, talc). During the evaporation of sea water, the already high Mg content of the sea water is increased. In addition, the deposition of carbonates and Ca-sulfates leads to a shift in the Ca/Mg ratio. When a mineral passes from fresh water into the supersaline environment, the differences are even greater in comparison to the normal marine milieu. The very high chlorite content in some salt clays seems to indicate that in this case, too, chlorite was formed from other clay minerals, probably above all from montmorillonite-vermiculite (FUCHTBAUER and GOLDSCHMIDT, 1959; BRAITSCH,1962). Without doubt the genesis of talc in anhydritic rocks of the Zechstein (FUCHTBAUER and GOLDSCHMIDT, 1959) must have occurred in an early stage of diagenesis @re-burial or shallow-burial stage). SiOz-gels deposited on the sea bottom reacted with the Mg-enriched sea water. According to Fiichtbauer and Goldschmidt: “Because of the deposition of talc plates parallel to the bedding plane one can assume that their neoformation was of synsedimentary origin and was concluded in an early stage of diagenesis, when the anhydrite still had a continuous pore space.” In a laboratory experiment, HOHLING(1958) obtained talc in MgClz solutions with excess of SiOz-gels and admittedly at elevated temperatures. According to FUCHTBAUER and GOLDSCHMIDT (1959), serpentine,pyrophyl-

146

G. M h L E R

lite, montmorillonite and corrensite, sporadically occurring in salt rocks, might also be of synsedimentary to early diagenetic origin. In clay muds of recent continental salt lakes (Great Salt Lake, Utah: GRIM et al., 1960; several salt lakes in California: DROSTE, 1961) no neoformation of clay minerals could be observed. Inasmuch as waters in these lakes are enriched in Na and poor in Mg, transformation of detritic minerals into Mg-rich minerals does not occur. Authigenesis of palygorskite and sepiolite. Palygorskite (attapulgite of the older literature) and sepiolite are silicate clay minerals (but not layer-silicates), which without doubt cannot be formed by normal weathering. Thus,their appearance in subaqueously deposited, predominantly marly and carbonate sediments indicates that they were formed in the sedimentarybasin itself and did not come from outside. MILLOT’Scompilation (1964) of the most important occurrences shows that palygorskite and sepiolite were formed mainly in marine environment but could also form both in fresh and in supersaline waters. Millot pointed out the following concerning the lake deposits: “prdsence de carbonates, frdquence des silex, sursalure facultative. ll s’agit, en un mot, d’une sddimentation chimique basique.” Millot also listed similar characteristics for the marine formation of these minerals: association with chemically precipitated carbonates, chert layers and nodules, and, often, with phosphates. Frequently palygorskite-bearing argillites and marls are interstratified with carbonates, cherts and phosphates. Clastic components are almost completely lacking. These facts strongly point to the neoformation in the sedimentary basin under extreme chemical conditions. The prerequisite for this was supply of large quantities of silica (for the formation of cherts as well as of the palygorskite and sepiolite), which were carried into the oceans from the continent in a dissolved or colloidal form. This can only be explained by extreme weathering conditions on the continent (deep-reaching weathering; lateritic soils). The first occurrence of Recent sepiolite in a deep-sea environment on the Mid-Atlantic Ridge has recently been reported by HATHAWAY and SACHS(1965). The sepiolite is intimately associated with laminated clay and coccolith ooze, both containing clinoptilolite; one sample contained chert. The sepiolite seems to have formed by the reaction of Mg in solution with Si02 liberated during alteration of silicic volcanic ash. The formation of glauconite. “La glauconie est caractkristique du milieu marin: ceci appartient A l’alphabet de tout gdologue” (MILLOT,1964). The most familiar and characteristic product of halmyrolysis is glauconite. It forms during the pre-burial stage of diagenesis and can be found today in many oceans at a depth of about 20-700 m in areas with a decelerated rate of sedimentation, such as for instance on the outer edge of the shelf in the Gulf of Mexico and

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

147

off Trinidad. According to SEIBOLD (1964), the green muds off Guinea at a depth of about 100 m can contain up to 50% glauconite. Since the Paleozoic period, glauconite is found in almost all formations. The average composition of 32 glauconites ( H ~ D R ~ Cand K S ROSS, 1941) as compared with the average composition of iUite (GRIMet al., 1937) is as follows:

mte: K0.~8(Ali.~~Fe3+0.~7Fe2+0.04Mgo.s4) (~~)2(~~8.41~0.59~~10 Glauconite:

(K,Cao.a,Na)o.sr(Alo.47F~+0.s7Fe2+0.leMgo.40) (OH)2(Si3.eaAlo.ss)Oio The main differences are the higher ratio of di- to trivalent iron in glauconite as well as the higher total iron content at the expense of the Al in the octahedral layer. In addition, there is a possibility of the substitution of Ca and Na for K. The Al content in the tetrahedral layer is considerably less than in illite and the K content is higher. There is no need for a detailed description of the formation of glauconite, because MILLOT(1964) recently gave a comprehensive account of our present knowledge on the subject, and FAIRBRIDGE (1967) also described the formation of glauconite. The works of BURST(I956,1958a,b) should also be consulted. As shown by the extensive literature on glauconite, several varieties of glauconite can be distinguished: (I) well-ordered glauconite (1 M-polytype), (2) disordered glauconite (1 Md-polytype), and (3) random mixed layer glauconitemontmor illonite. Millot gave 10 different possibilities for the mode of occurrence of glauconite which include: (I) the jilling of recent foraminifera1shells; (2) granular glauconite, showing no relationship to organisms and already existing minerals; (3) the replacement of spicules of siliceous sponges and opal or carbonate fossil debris; (4) the authigenesisin fissures of feldspars, as quartz incrustations, and as envelopes on phosphorite nodules; (5)replacement of coproliths, etc.; and (6)the transformation of biotite, illite and, possibly, other layer-silicates. From this it can be concluded that glauconite can be formed both through transformation as well as through a complete neoformation, and that its genesis extends beyond the preburial stage into the shallow-burial stage. BURST(1958a,b) considered redox potential as critical, suggesting that glauconites are commonly formed in locally reducing milieu associated with decaying organic matter, in generally oxidizing environment. This, however, does not apply to all glauconites. VANANDEL and POSTMA (1954) have recorded glauconite formations in the Gulf of Paria in moderately to highly oxidizing environments. Formation of zeolites. BONAITI’S studies (1963) in the Pacific show that phillipsite “is one of the most abundant mineral species of the upper layers of the earth’s

148

G. m L L E R

crust”. In extensive areas of the Pacific Ocean, phillipsite concentrations are greater than 50%. According to Bonatti, it does not seem likely that these zeolites crystallize directly from the ions commonly dissolved in deep-sea water, but are transformation products formed at the depositional interface. Microscopic observation shows that the crystals are invariably accompanied by more or less altered palagonite, a special kind of devitrified volcanic glass. All alteration steps can be observed from the palagonite grains without apparent zeolite growth, to the formation of small zeolitic nuclei on the grains, and finally to the development of well-formed crystals at the expense of the palagonite. The sedimentation rate in areas with a high phillipsite content is extremely low. In the whole Quaternary period less than 1 m of sediment has been deposited here and the sediment surface thus remained over a long period of time in the pte-burial stage. In the Recent sediments of the Gulf of Naples, Italy, analcime was found by MULLER(1961) together with newly-formed opal, quartz, and clay minerals which owe their origin to the halmyrolitic transformation of the sediment mainly composed of volcanic glass. Analcime had also been observed by NORIN (1953) in sediment cores of the central Tyrrhenian Sea, which contained ash layers. In comparison to the predominantly basaltic composition of the Pacific pyroclasts, these glasses have a trachyticleucitic chemistry. This could possibly be the reason why analcime and not phillipsite was formed. The three samples examined by BISCAYNE (1964) in the Atlantic and the Indian Ocean, which contained zeolites (phillipsite, clinoptilolite, and heulandite), were not of Recent age. Attention should be drawn to the occurrence of unusually huge deposits of analcime in the Central Congo Basin of Central Africa. VANDERSTAPPEN and VERBEEK (1964) have observed that Upper Jurassic and Lower Cretaceous deposits in this area, composed of alternating series of subaqueous, mainly non-marine beds of sandstone and pelitic rocks with no traces of volcanic material, are characterized by a high concentration of analcime. Locally this mineral is the essential constituent of beds attaining a thickness of tens of meters (“analcimolite”). The analcime forms well-defined, authigenic crystals 2040 p in size, spherulitic aggregates and interstitial fillings. Petrographic examinations and geologic conditions point to a syngenetic origin of the analcime for the formation of which a set of rather special physicochemical conditions are required. In the author’s opinion, this includes well-defined relative concentrations of soda, alumina and silica and an alkaline medium. Total retention of the soda by alkaline lakes without outlet to the ocean was a determining factor. The analcime must have crystallized from ionic solutions or from alumina-silica gels. Thus, there are certain similarities between the formation of palygorskite and sepiolite in Mg-rich environments. Already in 1959, JOULIA et al. described large deposits of analcime in Lower Cretaceous sandstones in the Central Sahara (“Analcimolites de Karafou”), which might have formed under similar conditions.

DIAGENESIS IN ARGJLLACEOUS SEDIMENTS

149

In Late Cenozoic lake sediments in the western United States, the zeolites (phillipsite, erionite, clinoptilolite and analcime) are products of the low-temperature reaction between the sediments (usually containing glass shards) and the lake waters of restricted basins (DEFFEYES, 1959). Theformation of montmorillonite. In samples collected southeast and east of Madagascar, near the Mascareignes, BISCAYNE(1964) observed greater amounts of montmorillonite in the uppermost layers of the sediments. Microscopic examination of the <20p fraction revealed the presence of considerable amounts of partially devitrified glassy fragments with translucent peripheries and opaque centers. “Correlation between montmorillonite abundance and glass abundance was not perfect, although the highest incidence of both occurred in the same sample. ‘This may represent alteration of local volcanic material represented by the dark glass fragments, perhaps detectable because of low sedimentation rates. Phillipsite, however, was not detected in these samples.” Montmorillonoid mineral of nontronite type occurs in significant quantities around centers of volcanic activity, like the Cape Verde (Atlantic) and the Hawaiian Islands (ARRHENIUS, 1963). In the Pacific, accumulations of nontronite coincide with the occurrence of phillipsite and palagonite. Very probably this nontronite was not formed through the relatively slow transformation of already cooled glass, which had undergone a long process of transportation, but through the direct contact of hot lava with sea water. In this case, therefore, the nontronite is more the product of a hydrothermal reaction in the contact zone between basalt and the sea water than of a halmyrolyticprocess. The neoformation of SiOa-minerals in pelitic sediments derived .from pyroclastic material. During the transformation of volcanic glass and the resulting formation of zeolites and clay minerals, large quantities of silica must be liberated because the alteration products are always poorer in silica than the initial material. If the transformation of pyroclasts in sea water takes place very quickly, all the silica liberated cannot be dissolved in the sea water; and, therefore, there is at least a partial h a t i o n of SiOz in the sediment. The sediments in the Gulf of Naples (MULLER,1961) can be mentioned as an example of recent halmyrolitic authigenesis of SiOz-minerals from line-grained glassy pyroclastic particles. The latter are mainly derived from the volcanic tuffs surrounding the coasts of the Gulf. Mineralogic and X-ray examinations showed that in the volcanic glass particles, with a decreasing grain size the content of newlyformed quartz and chalcedony increases. In the 0.02-0.002 mm fraction of one sample it amounted to 16.2 %. The highest content in the whole sample was 10 %. The content of opal present could not be quantitatively determined; however, it is probably considerable. The grain size of the sediments decreases more or less

150

G. I&LLER

5 km

,<\.

P

Capri

Fig.5. Quartz content of Recent sediments in the Gulf of Naples. (After M ~ ~ L E1961.) R,

continuously from the coast to the middle of the Gulf of Naples (M~~LLER, 1961, 1964); on the other hand, the quartz and chalcedony contents increase (Fig.5). The transformation observed here leads to an extreme chemical change in the sediment. Inasmuch as large quantities of SiO2 are also released during the transformation of feldspars into clay minerals, SiOz-minerals might also be formed in a similar way. It is probable, however, that this process only plays an important role during the later stages of diagenesisand is important for cementation. Formation of manganese nodules. Among the most unusual neoformations on the sea floor, are the concentrically structured manganese nodules, which cover enormous deep-sea areas in the Atlantic, Indian and Pacific Oceans. An extremely low sedimentation rate exists in these areas. AU the facts indicate that these are not concretions, which have been uncovered by the water action from older strata, but nodules which have grown on the surface of the sediment. The mineralogical composition (at least three different manganese minerals) and the chemistry are very complicated and differ locally. For the ultimate source and mode of transportation of manganese and associated elements, according to ARRHENIus (1963) several processes have been suggested to explain the subsequent accretion of the manganese oxide minerals: ( I ) various inorganic reactions, (2)organic (bacterial) mechanisms, and (3) removal of manganese from the bottom water by catalytic oxidation of manganous ion

DUGENPSIS M ARGILLACEOUS SEDIMBNTS

151

by colloidal ferric hydroxide at the sediment-water interface. Attention is drawn to the comprehensivetreatment of the subject by ARRHBNnrs (1963) and MERO(1965). Changes during the shailow-burial stage

As soon as a sediment layer at the depositional interface is covered by a younger layer, the physicochemical conditions in the interstitial water change. In particular, hydrogen-ion concentration and redox potential are iduenced. As these problems are treated in detail by FAIRSRIDGE (1967) in this book, and in addition a comprehensive account on oxidation-reduction in sediments was published by BUBENICEK (1964), this subject is not covered in this chapter. Changes in the chemistry of interstitial water EMERYand RI”ENBERG (1952) in their study of the sediments from the basins off the southern California coast found that the topmost sediment layer had an average pH of 7.59 compared to an average of 7.52 for the bottom water. This slight increase commonly continued downwards, pH values ranging up to 8.5 at a depth of about 8 m, and with most of the values falling between 7.5 and 8.0. This downward increase of pH was not observed by S ~ et al. R (1965), who determined the pH in several hundred samples of modern oceanic sediments in 22 cores from six general areas in the Atlantic and Pacific Oceans. The pH values of core sampleswere uniformly lower than that of surface sea water. The total range is from 7.00 to 7.85, most values clustering in the range of 7.2-7.7. “The most likely explanation for the pH being lower than sea water is an increase in C02 pressure in the sediment over that in equilibrium with the atmosphere. Corroborating this hypothesis are results of pH measurements on squeezed waters brought into equilibrium with the atmosphere by aeration; uniformly these measurements gave values 8.1-8.2, normal for sea water.” The increase in CO2 pressure in the sediments is most likely the result of bacterial oxidation of organic matter. The indications are that this effect is still operative at 10 m below the ocean bottom. Examination (SEVERet al., 1965) of the interstitial water of the same material showed that C1, Na, Ca and Mg contents are similar to those in sea water, but Mg is slightly depleted as a result of uptake by chlorites and possibly incipient dolomitization. A slight increase of the chlorinity with increasing depth is probably due to the semipermeablebehaviour of compacting clays (see DEGENS and CHILINGAR, 1967). In most samples, interstitial water is enriched in K content as a consequence of continuous hydrolysis of K-feldspars. There is consistently higher concentration of Si02 in pore water, which is traceable to dissolution of diatoms. These results obtained from sediment cores having a length of several meters also apply to the pore solutions of the over 150-m long experimental Mohole core, in which the deepest sediment layers are about 15,000,000 years old. If the composition of the pore solutions of these cores is compared with sea

152

G. MULLER

water and the pore waters of deeply buried ancient sediments, “the conclusion would be inescapable that sea water and recent sediment waters are closely related but that both are greatly different from ancient sediment pore waters. Yet the direction of the changes is such as to make modern sediment waters look if they may be in early intermediate stages in the often hypothesized transformation from sea water to sediment brine.” (SEVERet al., 1965, p.71.) The depth at which a more pronounced change in chemical composition of the pore solutions takes place is not known. The author does not know of any complete profile in which a complete transition has been observed. SIEVER et al. (1965) did not determine the sulfate content in their cores. In the sediments of the California basins EMERYand RITTENBERG (1952, p.789) found a decrease in the sulfate content with increasing depth. In core “5” the sulfate/chloride ratio decreased from 0.131 at 0-6 inches to 0.030 at about 70 inches, and at the bottom of the core, 83-87 inches below the surface, no sulfate could be detected. In two other cores changes were not as pronounced. In most cases, cores from the western Gulf of Mexico (CHAVE,1960) also showed a downward decrease in the sulfate/chloride ratio. The decrease of the sulfate content is brought about by sulfate-reducing bacteria in anaerobic environment. If the sediment contains more organic matter than can be decomposed by aerobic processes in the water containing oxygen (positive Eh values), a depletion of oxygen and, finally, a complete lack of oxygen can be observed. Thus, a reducing environment with anaerobic conditions and negative Eh value would result. This applies both to marine and non-marine environments. The transition from aerobic to anaerobic milieu can take place at some depth in the sediment; however, frequently it is already observed in the uppermost few centimeters of the sediments. In the Black Sea, anaerobic conditions predominate even in the water layers over the sediment, and thus also in the topmost layer of the sediment. Within the sequence of sediment layers, anaerobic zones can be intercalated among aerobic ones depending on the content of organic matter. Only little information is available on the changes of the chemistry of pore solutions of fresh-water sediments in relation to the burial depth. Inasmuch as from the very beginning only very low electrolyte contents were present in these interstitial solutions, the relative changes are probably much greater here. Formation of su&des During the activity of sulfate-reducing bacteria, H2S is produced which in the presence of dissolved iron or Fe-hydroxides is transformed into black hydrotroilite (FeS.nHz0). This explains the typical black coloring of these sediments. After a very short time, stable iron sulfides (predominantly pyrite) are formed from the unstable hydrotroilite in clayey sediments (LOVE,1964). The formation of hydrotroilite is not limited to marine sediments; many

153

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

sediments in fresh-water lakes (for instance, Lake Constance) locally contain several per cent of this sulfide. In addition to pyrite (and marcasite, the formation of which is favored by lower pH), in clayey sediments of older formations very small quantities of other heavy-metal sulfides occur (mainly sphalerite, galena, and chalcopyrite). In pyriterich, bituminous, dark clay-marls of different Mesozoic strata, MULLER (1955) and H A U S S ~ ~and L L MULLER (1963) observed numerous idiomorphic wurtzite crystals which were formed during the early diagenetic stage. The conditions for the formation of such heavy metal sulfides have not yet been completely clarified; however, in all probability the heavy metals are derived from the surrounding sediment itself and were brought to the formation site in ionic solutions. Changes in clay-mineral composition It is to be expected that most of the processes starting in the pre-burial stage also continue in the shallow-burial stage, even if greatly decelerated, because a single mineral grain is now in contact only with a limited amount of water. For example, the cation substitutions and the neoformation of chlorite-sudoite (and probably illite) from montmorillonite continue in the marine environment. The early diagenetic formation of illite in clay-rich sediments, containing volcanic ash, was shown in the Mediterranean Sea by NORIN(1953) and MULLER (1961). The nature of the partly worm-like crystals completely excludes the possibility of transportation; and they cannot be alteration products of already existing minerals. Their development from ionic solutions seems to be very probable. According to FUCHTBAUER and GOLDSCHMLDT (1963), the formation of kaolinite is likely to occur during the shallow-burial stage of argillaceous rocks. The formation of this kaolinite stops at a depth of several hundred meters and is followed by the kaolinite-chlorite transformation during the deep-burial stage. This view is supported by the clay-mineral studies of shales that have been compared with shales of the same sediment but in which diagenesis has been hampered by oil accumulation or carbonate cementation. The clay-mineral content of some carbonate concretions consists of about 85 % illite plus about 15 % kaolinite and chlorite, which may be considered as the initial (detritic) material. Around the concretions the amount of kaolinite chlorite of the clay fraction had doubled during diagenesis; the increase is greater in the shales poor in carbonates than in the carbonate-rich ones. The kaolinite and chlorite content in the outer rim of the concretion is just as great as in the sediments outside the concretion. On the basis of this, it can be assumed that the outer rim formed considerably later than the center. In addition, the kaolinite/chlorite ratio is always highest in the outer rim. This indicates that during the early stage of diagenesis the chlorite authigenesis is less than at a later stage. MURRAY and HARRISON (1957) were able to show that the uppermost sediment layers in young deposits in the Gulf of Mexico contain slightly crystallized

+

154

a. Mh im

illite and chlorite. In deeper layers the same minerals were found in a much greater state of crystallization. In a fresh-water environment it is probable that a further decomposition of illite and ledikite takes place (removal of K). The alteration of clastic biotite (or ledikite) into chlorite already beginning in a very early and shallow stage of diagenesis will be treated in the section on the deep-bmial stage of diagenesis. Carbonate concretions Mudstones and shales poor in carbonates commonly contain concretionary bodies (mainly consisting of calcite, and to a lesser extent also of siderite). It seems probable that most of the concretions started to form in the shallow-burial (and early) stage of diagenesis,because the enveloped relics of organisms are commonly not deformed. The concretions frequently contain laminae which may continue into the surrounding shale. In the concretion, however, the laminae are several times thicker; they represent an earlier stage of compaction. As relics of organisms are frequently to be found in the center of the concretions, it seems likely that organic matter played a role in the formation of concretions. LPPMANN(1955) explained the genesis of concretions as follows. Ammonia resulting from the decomposition of organisms or amines gives rise to a strongly alkaline reaction in the vicinity of the animal (or plant) embedded in the sediment, and the pH is increased. As the solubility of the carbonates decreases with increasing pH, they are precipitated on the fossil from the interstitial solutions, which have been saturated with carbonates by dissolving the disseminated calcareous material (also present in predominantly argillaceous sediments). Thus here the carbonate concentration of the pore solution decreases in comparison to the surrounding environment, and because of the difference in the concentration, more carbonate is constantly diffused toward the fossil. This process, accompanied by a constant growth of the concretion, continues until the production of ammonia stops, or until there are no more dissolved carbonates available in the vicinity. As reducing conditions can predominate in the immediate vicinity of a decomposing organism, siderite is also a possible concretion-forming material. The precipitation of the carbonates takes place in the water-filled pore spaces without the latter changing in size to any marked degree (e.g., by the clay minerals being forced apart). Thus, the water content (and consequently also the porosity) at the time of the formation of the concretion can be calculated from the volume ratio of carbonate to non-carbonate minerals in the concretions. For instance, LIPPMANN (1955) and SEIBOLD (1962) made such calculations on Lower Cretaceous and Liassic concretions of argillaceous sediments from Hoheneggelsen (northwestern Germany) and Wutach-Schlucht (southwestern Germany). According to Lippmann, the calculated water content is about 55%, corresponding to

DIAGBNESIS IN ARGILLACEOUS SEDIMENTS

155

a porosity of about 75 %. Seibold’s values are slightly lower: about 70 % porosity. From this, both authors concluded that the concretions must have formed in the uppermost few meters of the sediment. Changes during the deep-burial stage of diagenesis According to TAYLOR(1964), with increasing depth of burial hydrogen-ion concentration and oxidation-reduction potential decreasein importance in controlling the diagenesis. Temperature and overburden pressure, both on mineral grains and on pore fluids, together with the partial pressures of the main components of these fluids, are likely to be the significant factors. Changes in the chemistry of interstitial water The chemical composition of interstitial solutions of ancient sediments differs greatly from that of sea water and Recent sediments. In oil-field brines electrolyte contents range from less than that of sea water up to ten times as much. In many formations an increase of the electrolyte content with increasing depth can be observed (cf. VONENGBLHARDT, 1960). Compared with sea water, the ancient pore waters show a decrease in the Na/Cl, K/Cl, Mg/Cl, HCOa/Cl and S04/C1 ratios and an increase of the Ca/Cl ratio. The main reasons for this are discussed in detail by VONENGELHARDT (1960, 1961), and in this book by DEGENS and CHILINGAR (1967). The increase in concentration of salts can be attributed mainly to the argillaceous sediments, i.e., ionfiltration by charged-net clay membranes (cf. DEGENS and CHILINGAR,1967). Changes in clay-mineral composition The changes in the clay-mineral composition during the pre-burial stage and the upper shallow-burial stage can become apparent to a certain extent by means of a comparison of material delivered from a known hinterland with material from the sedimentation basin. In older sediments at greater depths this comparison is not possible, and thus the unknown factors correspondingly increase. Alteration of montmorillonite into illite. Among the best-known alterations, is the transformation of montmorillonite into illite in the two first stages of diagenesis. It is possible that this takes place through a simple fixation of K from the interstitial water or as a result of K release during the decay of feldspars. According to P o w (1959), this fixation of K in the crystal lattice is preceded by a preferential adsorption of K instead of Mg. The Mg/K ratio of 5/1 in sea water explains the observed exchange preference of clays for Mg in sea water and in interstitial waters of shallow-burial stage. The preferential adsorption of Mg over K with depth cannot continue beyond a certain Mg/K ratio. That level below which K is adsorbed to a greater extent

156

G. MULLER

than Mg may be called the “equivalence level” for Mg and K. The depth of this level will depend upon the difference in the ionic potentials of the two ions and their relative concentrations, and should not exceed several hundred feet. Frequently, a gradual transformation of montmorillonite takes place; i.e., first mixed-layer minerals with a small illite content are formed, and this transformation increases until pure illite emerges. From the study of numerous strata in the United States, WEAVER (1959) came to the following conclusion: “There is good likelihood that below a depth of approximately 10,000 ft. montmorillonite gives way to a mixed-layer illite-montmorillonite with the portion of illite increasing with depth. Around 15,OOO-17,000 ft. the mixed-layer illite/montmorillonite ratio is approximately 7/3. Samples that are believed to have been at one time buried to approximately 25,000 ft. have mixed-layer ratios of 9/14/1. It may be that deeper burial would cause the complete disappearance of montmorillonite layers. This process may, in part, account for the lack of montmorillonite in the older Paleozoic sediments.” In the Wilcox Formation (Eocene) of the Gulf Coast area, BURST (1959) observed the transformation of montmorillonite to illite at an even higher level. Montmorillonite, a common mineral of Wilcox outcrop material, becomes less evident below 3,000 ft. and is not normally found in an unmixed state below 9,00010,OOO ft. of overburden. At depths between 3,000 and 14,000 ft., montmorillonite lattices are commonly interspersed with illite, the abundance of which increases with depth to a virtual elimination of montmorillonite swelling characteristics below 14,000 ft. In the Upper Cretaceous of Cameroun, DUNOYER DE SEGONZAC (1964) observed the following diagenetic changes with increasing depth: ( I ) <1,500 m montmorillonite, (2) 1,300-3,800 m mixed-layer montmorillonite-illite, and (3) >3,500 m illite. The transformation of montmorillonite via a mixed-layer mineral into illite was also described from strata in the Soviet Union by KOSSOWSKAJA (1960). It takes place at depths ranging from 3,000 to 4,000 m. A completely different kind of illite formation from montmorillon‘ite was described by KELLER (1963). The Brushy Basin member of the Morrison Formation (Jurassic) on the Colorado Plateau, a bentonitic mudstone composed predominantly of montmorillonite derived from alteration of volcanic ash, at one locality (on Lone Tree or Blue Mesa) is composed of illite. These illitic clays are anomalously blue in color instead of pink to red. They are stratigraphically continuous and interfinger with the gray and yellow-red bentonite surrounding them. The source of the K is postulated to be from solutionsfrom evaporites, which have risen through salt-cored anticlines in the vicinity of Blue Mesa.

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

157

Transformation b-axis disordered kaolinite

X-ray investigations of “kaolinite-coal-claystones” of the Carboniferous of the Ruhr district, carried out by ECKHARDT (1965), show that there is a relation between the b-axis disorder of the kaolinite and the degree of coalificafion in the surrounding coal seams. By applying coal-petrographic data one can fur the ranges of temperature forming the different types of coal. It is shown that the b-axis disorder is continually decreasing under the influence of increasing temperature of the diagenesis, when the primary kaolinite has a relatively higher b-axis disorder. Regarding a temperature of more than 200” C, no b-axis disorder of the kaolinite has been found. The transformation of biotite

The diagenetic behaviour of clastic biotite (or ledikite) in platform and geosynclinal successions of the Soviet Union (marginal part of the Siberian platform; Verkhoyano-Kolimskaya syncline) has been studied by KOSSOVSKAYA et al. (1965). This behaviour is completely different from the transformation of biotite into vermiculite during the process of weathering. Under diagenetic conditions modifications of the biotite up to the montmorillonite and kaolinite can be traced going through a series of intermediate stages of regular and irregular inter-stratified phases depending on environmental conditions and the rate of immersion of the sediments. In platform succession clays (total thickness 500 m) with increasing depth the following phases derived from biotite were detected: (I) trioctahedral micachlorite; (2) chlorite-“mobile” chlorite and (3) “mobile” chlorite-montmorillonite. The succession of the changes may be pictured as follows: (I) potassium replaces Mg, which forms brucite layers (appearance of chlorite); (2) gradual hydration and decomposition of the brucite layer (appearance of “mobile” chlorite); (3) presence in the interlamellar spaces of only water molecules and cations of the exchanged complex (formation of montmorillonite). Besides the listed minerals, kaolinite and illite are formed. The formation of kaolinite represents the final stage of the decomposition of biotite. Kaolinite clays are usually encountered at the bottom of a coal bed. Their formation occurred under the conditions of very slow sedimentation and lengthy diagenesis in acid medium. In geosynclinal succession clays (total thickness 3,500 m) the following associations can be pointed out: (I) association with chlorite and mixed-layer formations montmorillonite-hydromica; (2) chlorite-dioctahedral hydromicaceous association with relics of the transitional phase, and (3) association of chlorite and dioctahedral illite.

158

G. MmLm

The formation of tetraphormic clay minerals (chlorite, sudoite). Relatively uniform illite-chlorite parageneses are almost always formed from the different primary clay-mineral associations as they approach the boundary between diagenesis and metamorphism (ECKHARDT,1958; GRIM,1958; F ~ ~ C H T Band A~R GOLDSCHMIDT, 1963; MILLOT, 1964). The increase in the chlorite content takes place more slowly than that of illite (ECKHARDT, 1958). Chlorite forms in larger amounts only under low-temperature metamorphism; at the same time illite is transformed into sericite (ECKHARDT,1958). Both montmorillonite and kaolinite, and occasionally illite, are source material for the formation of chlorite during the deep-burial stage. Great difficultiesare encountered in the interpretation of the chemical changes because the normal chlorite is trioctahedral, whereas the supposed initial minerals are dioctahedral. Bivalent cations (above all Mg and Fe) must replace the trivalent ones (mainly Al,and to a lesser extent Fe). In the case of kaolinite devoid of Mg, the chemical changes should be the largest. In addition to this, there is another problem of explaining the structural changes which again are most extensive in the case of kaolinite: kaolinite is diphormic, whereas chlorite is tetraphormic. Chemical and structural changes of this magnitude are so improbable, that a direct transformation of kaolinite into chlorite is hardly possible. Either there are intermediates ta g e e a s yet unknown to us-or the formation of chlorite takes place via the chemical decay of kaolinite by hydrolysis in the presence of Mg in the interstitial water; i.e., it essentially evolves from ionic solutions. The sprouting of chlorite crystals observed in argillaceous sediments by ECKHARDT( 1958)supports the latter mode of genesis. In thecase of montmorillonitel however, a transformation into a tetraphormic layer-silicate can take place fairly easily through fixation of a brucite layer in the interlayers (cf., the formation of chlorite-sudoite in the pre-burial stage); but the mineral formed in this manner would be dioctahedral (sudoite). In fact, such dioctahedral tetramorphic layersilicates (sudoites) have been described as diagenetic neoformations in argillaceous sediments. WEAVER(1959) found that sudoite (“dioctahedral chlorite”) is a common mineral in the Middle Ordovician K-bentonite beds from Pennsylvania, Tennessee, Virginia, Alabama and Kentucky. Volcanic ash presumably altered to a dioctahedral montmorillonite-like mineral; the presence of Mg in interstitial water enabled precipitation of brucite between the layers. VONENGELHARDT et al. (1962) found that sudoites are important clay minerals in weakly consolidated clays and mark of the Middle Keuper of southwest Germany. The sudoite content increases with decreasing amount of kaolinite present. This led KROMER (1963) to believe that sudoite forms from kaolinite, be1 The transformation of (trioctahedral) biotite or ledikite into chlorite (trioctahedral) has already been mentioned in the foregoing section.

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

159

cause primary differences in the kaolinite distribution are not very probable, even though they cannot be excluded. One must presume that there is a considerable addition of Mg for the formation of trioctahedral chlorites. Ross and HENDRICKS(1945) observed the following average composition for montmorillonite:

1.7Al20~.0.6Mg0.8Si02.2H~O

In addition to these components, there is also a maximum amount of 0.3 Mg which can be absorbed. By adding the amounts of MgO and H2O necessary for chlorite formation the following equation is obtained (ECKHARDT, 1958): 1. ~ A ~ z O ~ . O . ~ M ~ O . ~ S ~ O Z . ~ H =~10.1Mg0.1.7Al~0~.6.4 O+~.~M~O+~H~O Si02.8Hz0+1.6SiOa

According to ECKHARDT(1958) the first theoretical intermediate product which should be obtained is saponite, then corrensite, and finally, chlorite. It is doubtful that these intermediate products occur in nature. Random mixed-layer montmorillonite (or vermiculite)-chlorite minerals, however, are frequently found in nature. Eckhardt's calculations show that the Mg content in the interstitial water is not suilicient for the formation of pure Mg-chlorites. Yet, if the divalent iron which is present in the sediment in smcient quantities is included, then the formation of chlorites containing iron in the later stages of diagenesis can easily be explained. In fact, most sedimentary chlorites have a high Fe content. On the other hand, the chlorites formed by halmyrolysis in the fist stages of diagenesis should be Mg-rich chlorites or sudoites.

Fig.6. Influence of the maximum depth of burial on porosity and the kaolinite/chlorite ratio in shales from northwestern Germany. (After Ftk-um and GOLDSCHMIDT, 1963.)

160

G. &LLER

The increase of the chlorite content at the expense of kaolinite with increasing and depth of burial and decreasing porosity was clearly shown by FUCHTBAUER GOLDSCHMIDT (1963) in sediments from northwestern Germany (Fig.6). In shales of Albian, Liassic and Dogger ages the kaolinite/chlorite ratio decreases rapidIy with increasing depth, which means that the absolute chlorite content increases at the expense of the kaolinite. The increase in the chlorite content is always accompanied by an increase in the amount of quartz. This is to be expected, however, because chlorite is poorer in Si02 than kaolinite. Transformation sepiolite-talc. WCHTBAUER and GOLDSCHMIDT (1963) pointed out that sepiolite and palygorskite frequently occur in younger evaporites, whereas talc is more frequent in older ones. Under certain conditions there could have been a transformation of sepiolite to talc, and it would have to be investigated to what degree such a structural change is possible. Formation of authigenic feldrpars According to TAYLOR (1 964), authigenic orthoclase may be extensively developed in shales containing pyroclastic material. In the Tertiary John Day Formation of Oregon, it occurs over an area of about 600 sq. miles. The results of potassiumargon age determinations from both authigenic and pyrogenic constituents suggest that orthoclase was formed at burial depths ranging from 400 to 2,200 ft. and at a temperature of 20-40”C. Also a great proportion of the so-called “Tonsteine” of the Upper Permian of the Saar-Nahe district, western Germany, consists of shales composed of orthoclase-quartz mixtures, which can be derived from the devitrification of volcanic glass (HHM, 1960). Some types of “Tonstein” also contain idiomorphic albite. WEBS (1954) described Ordovician shales in southeastern Minnesota containing up to 70 % orthoclase, largely authigenic. Cementation of argillaceous sediments during diagenesis

Cementation is not a prerequisite for the consolidation of a clay-rich sediment. In fact, numerous clay sediments without any pore cement are found at considerable depths, the clay minerals and other clay-sized material being kept together by cohesion. As most argillaceous sediments initially contain allogenic or biogenic carbonates, feldspars, SiO2-minerals, etc., it is very probable that partial or a complete cementation can occur at an early or a late stage of diagenesis. Calcite is by far the most impoitant pore cement, which enables cementation to proceed even at the beginning of the shallow-burial stage, as indicated by the carbonate concretions. The cementation of argillaceous sediments, however, generally takes place during a later period in the deep-burial stage.

161

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

The second most important pore cement is quartz. In addition to the quartz and opal present initially in the sediment, additional silica is released during most of the diagenetic mineral alterations that can serve as a pore cement in situ. In this book, DAPPLES(1967) draws attention to the possibility of an early replacement of clay minerals by silica in argillaceous sediments rich in SiOz, and the formation of siliceous shales.

CHANGES IN POROSITY AND STRUCTURE DURING DIAGENBSIS

The loosely packed structure of freshly deposited clays becomes more dense under the weight of new layers of sediment, with a simultaneous reduction of the water content. This process is called compaction. Compaction is d e h e d as a “decrease in volume of sediments, as a result of compressive stress, usually resulting from continuous deposition above them” (AMERICAN GEOLOGICAL INSTITUTE, 1957). The compaction of argillaceous sediments can be directly or indirectly observed in many areas with a younger sediment cover. Some examples of this have been presented by VON ENGELHARDT (1960). According to UMBGROVE (1951), the sinking of the Dutch coastal area is mainly due to the gradual compaction of argillaceous sediments. In the region of the Po river delta (Italy), the land surface is constantly sinking to a considerable extent. Here in the Quaternary alone 2,400 m of mainly clay-rich sediments were deposited. In all old towns in this area it can be observed that over the centuries buildings have sunk into the ground by compressing it by their weight. For example, in Ravenna the foundations of the Theoderich burial monument and the churches of the 6th century sank into the ground.

TABLE IV WATER CONTENT OF YOUNG LAKE SEDIMENTS

(After ZULLIG,1956) Zuger See depth of burial

(cml

0 1.2 2.4 3.6

Luke Constance water content ( %I

83.6 74.0 74.2 70.6

calculated porosity ( %I

92 87 87 85

depth of burial

water content

(cm)

( %)

calculated porosity

0 -0.5 0.5-1 .O 1.2-1.5 2.3-2.8 4.0-4.6

73.8 72.2 68.0 61.7 60.3

87 86 84 80 79

( %)

162

G. m L E R

Changes in porosity Most studies of the relationship between porosity of argillaceous sediments and 1960) show that with an increasing overburden burial depth (see VON ENGBLHARDT, of the younger sediments the porosity decreases continuously. Apparently, even a very thin layer of younger sediments causes compaction. For example, according 0

rn

45 1

2

3

4

5

6

7

a

Water content C/.wet.w

100 300 500 700 900 Compressive strength (dyn/cm'.981)

Fig.7. Water content and compressive strength of sediments from Lake Z ~ C H(After . ZULLIQ, 1956.)

163

DIAGENESIS IN ARGlLLACEOUS SEDIMENTS

TABLE V DECREASE OF POROSITYwmi DEPTH IN SUBAQUEOUS A R ~ ~ A C E O USEDIMENT+ S

Depth (m)

Porosity (%) Black Sea2

0.00 0.20 0.50 1.00 2.00 3.00 4.00 5.00 6.00 7.00 8.OO 1 Porosities in

Santa Barbara Basin, Calif.8

79 73 72 71 70 65

82 81 79 77 75 74 73

Lake Zurich, Switzerland4 88 78 77 75 73 71 68 66 64 62 60

percent calculated from water content.

a After Sawelejew, cf. RUCHM(1958). 3 4

After EMERYand RITTENBERC~ (1952). After ZijuIo (1956).

to ZULLIG(1956) in the Zuger See even after the formation of a sediment layer 3.6 cm in thickness, the water content decreases from 83.6 to 70.6%; in Lake Constance the decrease in water content is of the same order of magnitude between 0 and 4.6 cm below depositional interface (Table IV). The water content and the calculated porosity of clay-rich sediments of a 8.20 m long core from Lake Zurich (Switzerland) studied by Z~~LLIG (1956), the lowest layers of which have an estimated age of about 5,000-10,000 years, are presented in Fig.7. The individual measurements show marked deviations probably due to the variation in the grain size and composition of the core in the different layers; however, on the whole the decrease in porosity with increasing depth can be clearly seen. Between 0 and 8 m the porosity of the sediment decreases by about 30% with the most marked decrease in the uppermost few centimeters. With decreasing porosity, the compressive stress increases considerably, which clearly demonstrates the process of consolidation. The porosities of argillaceoussedimentsdetermined from sediment cores from three completely different environments (Black Sea, Santa Barbara Basin and Lake Zurich) are compared in Table V. In spite of these differences in environment, mineral composition and the completely different ages of the comparable layers (the lowest layers of Lake Zuiich core are about ten to a hundted times younger than the lowest layers of the other two cores), there is a strong resemblance in the trend of porosity decrease with depth.

164

G. MULLER

Highly colloidal clays

Silts

1

10

100

E f f e c t i v e overburden load (kg/cm2)

B o

10

1

100

Pressure (kg/cm2)

'1 C

1

10

100

Pressure (kg/crn2)

D +o

1

10

100

Pressure ( kg/cm2 1

M NaCl

E

0

1

10

1

100

Pressure (kg/cm2) Fig.8. Relationship of void ratio to other factors, observed in natural sediments and in laboratory experiments. Void ratio is ordinate in all graphs; note different void ratio scales. (After MEADE, 1963.) A. Generalized relation to effective overburden load and particle size in sediments. (Modified 1953, p.55.) after SKEMPTON, B. Experimentally-determined relation to pressure and clay-mineral species. (Modified after CHILINGAR and KNIGHT,1960, p.104; to show their results to 100 kg/cma.)

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

165

The unusually high water contents in the deeper core layers of the experimental Mohole (&”BERG et al., 1963) seem to contradict the assumption of a decrease in porosity with increasing depth. The red clay samples taken at a depth of about 1-3 and 85 m had a water content of 52-50 and 45 %, respectively. Whereas here a slight dependence on the depth seems to be indicated, the water content at a depth between about 103 and 141 m remains constant at about 50%. In the latter case, however, the sediments are calcareous-siliceous and siliceous-calcareous oozes and, for this reason alone, a direct comparison is not possible. Inasmuch as cores of experimental Mohole exhibit an increase in strength with depth, which is not due to an increase in compaction, Rittenberg and coworkers suggested that “the greater stiffness of sediments at depth results from a weak cementation or the establishment of some kind of bonding between the clay minerals, such as the synaeiesis observed by WHITE (1961) in other argillaceous sediments and rocks”. The influence of the different factors during compaction in the shallowburial stage of diagenesis, which are also largely responsible for the formation of the primary porosities, is shown in Fig.8. With increasing pressure (= overbuiden load = burial depth), the influences of the various factors become similar. From this, MEADE(1964) concluded that at overburden pressures greater than about 50 kg/cm2 the important influences on the water content of clayey sediments seem to be only particle size, type of clay minerals and temperature. Compression experiments (30-3,200 kglcmz), carried out by VON ENGELHARDT and GAIDA(1963) on pure montmorillonite and kaolinite clay muds treated with NaCl and CaClz solutions and distilled water, showed that the equilibrium porosity with a distinct void ratio which is reached at a certain pressure does not depend on electrolyte concentration. MEADE(1964) concluded from this and other observations that “apparently the physicochemical influences of the different cation types and electrolyte concentrations do not affect the amount of water held by a clay unless the amount exceeds a certain minimum necessary to develop diffuse double layers around the particles. This minimum amount seems to be about 50% by weight in very finegrained ( t O . 1 ~ ) Na-montmorillonite; it should be somewhat less in coarser grained and more silty clays. When overburden loads have reduced the amount of water in a clayey sediment below this minimum, the forces that must be oveIcome in order to compact the sediment further are more conveniently thought of as forces

C. Experimentally-determined relation to pressure and adsorbed cations in <0.2p fraction of montmorillonite.(Modified after BOLT,1956, p.91.) D. Experimentally-determinedrelation to pressure and electrolyte concentration in unfractiona1960, fig.M3.) ted Fithian illite (about 60%by weight coarser than 2.4. (Modified after MITCHELL, E. Experimentally-determined relation to pressure and electrolyteconcentration in <0.2,u fraction of Fithian illite. (Modified after BOLT,1956, p.92.)

166

G. MthLER

of hydration-the attractions between the clay surfaces and water or between cations and water-rather than as forces of repulsion or attraction between particles.” VAN OLPHEN(1963) differentiated between two stages of compaction: one in which the particles are relatively far apart, and a second one in which they are separated by only a few monomolecular layers of water. In the first stage, compaction is primarily due to double-layer repulsion (“osmotic swelling”). The pressures range from a fraction of 1 atm to tens of atmospheres. In the second stage, compaction is largely controlled by the forces of adsorption of the water layers on the clay surfaces. The removal of adsorbed water from between the surfaces of clay particles as well as from quartz particles, will generally require extremely high compaction pressures-higher than those usually encountered in nature. The release of adsorbed water, however, will be facilitated by an increase of temperature. Pressure-temperature curves for dehydration reactions of montmorillonite and ROY,1959) show that minerals saturated with different cations (CROWLEY water can be removed from clay surfaces at fairly low temperatures. MEADE (1964) concluded that “temperature might be as important a factor as pressure in removing the last increments of water from clay”. There are detailed studies on the compaction (or density) of argillaceous sediments which are not influenced by tectonics, at depths up to 3,000 m for the 1936), the Tertiary sediments of the Tertiary deposits of Venezuela (HEDBERG, Po Basin (STORER,1959), and the Liassic of Germany (VON ENGELHARDT, 1960). As shown in Fig.9 a relation exists between the void ratio and the logarithm of the depth, which can be expressed by the equation: E = El E=

-b

Ig

t

e

100-e

TABLE VI CONSTANT+ C A L ~ T E DFROM CURVES IN

ma9

(After VONENOELHARDT, 1960)

Tertiary, Venezuela Tertiary, Po Basin Liassic, northwestern Germany

El = void ratio at a depth t = lm;b depth at which)heoretically e = 0.

1

1.844 1.700 1.160

0.527 0.481 0.317

= compressibility of

65 63 54

3.160 3.500 4.570

clay; e = porosity; to = calculated

167

DIAGBNESIS IN ARGILLACEOUS SEDIMENTS Void

ratio

0.2

10

20

30

40

50

Porosity

(*M

61

Fig.9. Porosity and void ratio of argillaceous sediments related to depth of burial. (After VON

ENOELHARDT, 1961.)

where E = void ratio at a depth t (in m), EI = void ratio at a depth t=lm, b = index of compressibility of a certain clay, and e = porosity (in %). If the above equation is valid until the amount of pore space reaches 0, the depth to can be calculated at which porosity no longer exists (e = 0):

Ei

Ig to = b TABLE VII DECREASE OF VOLUME OF ARGILLACEOUS SEDIMENTS WITH INCRBASINct DEPTH

(After VONENGELHARDT, 1960) ~~

Depth (m)

Decrease in volume (in % of original volume) Tertiary, Venezuela

Tertiary, Po Basin

Liassic, northwestern Germany

50.0 55.5

48.0

39.6 44.0 48.5 51 .O

61.1 64.4

53.4 58.7

61.7

168

G. m L L E R

The constants given in Table VI apply to the compression curves presented in Fig.9. Von Engelhardt pointed out, however, that this simple relationship does not apply to all depths. Deviations are to be expected both for small and for great depths. Based on previous discussion, it is probable that at greater depth the pole space increases more sIowly than the empirical equation indicates. The calculated depth to, therefore, has no real meaning. The decrease of the initial volume of argillaceous sediments with increasing depth of burial is considerable (Table VII). The rate of decrease is very high down to a depth of 500 m, but with increasing depth of buiial it becomes less. It is important to note here that once a certain overburden load has been reached through compaction, the process is irreversible.This means that, even after a later uplift of the sediment with erosion of the upper layers (and consequent release of pressure), the porosity reached at the maximum burial depth remains Void ratio

Depth

5

10

15

20 25 Porosity

30

35

Fig.10. Porosity and void ratio of Liassic shales as related to present depth. The arrows indicate 1960.) the maximum depth of burial. (After H. Fiichtbauer, cf. VON ENGELHARDT,

169

DIAGENESIS IN ARGILLACEOUS SEDIMENTS

constant. Thus, from the porosity of unweathered argillaceous rocks the maximum depth of burial experienced by the sediment in question can be estimated. The porosity of argillaceous sediments of Liassic a from deep bore holes in northwestern Germany in relation to the present depth are presented in Fig.10 (Fiichtbauer in: VON ENGELHARDT, 1960). Whereas the porosities of argillaceous rocks in regions I-Vare more or less what is to be expected at the present depth of overburden, the sediments in regions A, C, Ho and He have too high a poiosity in relation to their present depth. In these latter areas, during the uplift of salt plugs the Liassic sediments were raised from a level shown by the arrows in the diagram. Structural changes Laboratory experiments on clay-mineral-water systems prove that with increasing pressure, there is an increase in the orientation of the clay particles: basal plains lie in the plain vertical to the compressive stress (MITCHELL, 1956; VONENGELHARDT and GAIDA,1963). VONENGELHARDT and GAIDA(1963) also observed that with coarser kaolinite particles the resistance against orientation is lower, and at the same time better orientation is obtained than with montmorillonite clay. The electrolyte concentration influences the degree of preferred orientation considerably. With increasing electrolyte content, degree of orientation decreases. In the previous section it was pointed out that the equilibrium porosity at a distinct pressure does not depend on the electrolyte concentration of pore solutions. On the other hand, the experiments did show that with the same equilibrium porosity the permeability in the clays treated with electrolyte solutions was higher than in those containing little or no electrolytes. Thus, VON ENGELHARDT and TABLE VIII BULK DENSITY AND ORIENTATION RATIO OF ILLITE IN CRETACEOUS SHALES COMPARED WITH THOSE IN RECENT CLAY AND SLATE

(After MEADE,1964)

Recent clay Cretaceous shale Cretaceous shale Cretaceous shale Cretaceous shale Slate 1

.

After MEADE (1961).

Present depth below surface (ft.)

Bulk density

0 4.340 5,136 6,776 9,120

1.9 2.4 2.5 2.6 2.7 2.9

-

(glcms)

Orientation ratio'

1.5 2.8 2.2 3.2 5.0 32.0

170

G.

M~~LLER

GAIDA(1963) concluded that these differencescan be explained on assuming aggregate structures with large channels instead of homogeneous structures (Fig.3). According to Von Engelhardt and Gaida, during compaction these aggregates will be destroyed, and at very high pressures completely or almost completely homogeneous structures should result, consisting of only primary particles. Based on conclusive results of laboratory experiments, one should expect to find similarly clear situations in nature. This does not seem to be the case, however. (1959) in flat-lying Cretaceous shales of The structure studies by -BERG North America, buried under 2,700 ft. or more of overburden showed that the preferred orientation of illite increases with depth of burial. From X-ray analyses given in Kaarsberg’s report, MEADE (1964) computed the “orientation ratio” (Table VIII). As shown by MEADE(1961), the larger the ratio, the greater the preferred orientation of clay particles parallel to the bedding plane. In a study of a 2,OOO-ft. thick section of unconsolidated,nonmarine sediments from San Joaquin Valley, California, MEADE(1963) did not find a progressive increase in the orientation of montmorillonite clay parallel to the bedding with increasing depth. Unfortunately there are still too few combined studies on mineralogy-granulometry-texture of natural sediments in order to show clearly influences other than compaction on the orientation of the clay particles. Natural sediments are not monomineralic systems and, as a rule, also contain considerable quantities of quartz, feldspar, and other detritic components alongside the clay minerals. Because of their more isometric shape, these components have an adverse effect on the preferred orientation of the clay minerals. GRIMet al. (1957) showed that in fine-grained shales (most of the particles were smaller than 2p in size) the orientation was better developed than in comparable coarser grained shales that contained larger amounts of non-clay grains. CORRELATION OF MECHANICAL AND CHEMICAL-MMBRALOGICAL CHANGES WITH DBPTH OF BURIAL, PRESSURE, TBMPERATURR AND DURATION OF BURIAL

In Fig.11 an attempt is made to correlate the main mechanical and chemicalmineralogical changes occurring during diagenesis with the depth of burial, increase of temperature and pressure, and the duration of burial. The figures given in the graph should only be considered as average values; for any particular argillaceous sediment, the true values may vary within greater limits. THE TRANSITIONAL ZONE BETWEEN DIAGENESIS AND METAMORPHISM

At greater depth, argillaceous sediments are subjected to elevated temperatures. Some new minerals formed under these conditions are unstable at shallow depths.

---

- --- -- - -

concrellons

Ieldspar from volcanic g l a s r

kaolinite

sultldes

carbonate

talc In solt c l a y s

--- - ---

r e o l l I * s ( t r a m volcanlc g l a s s )

m o n l m o r i l l o n i l * (from v o l c a n i c g l a s s ) .

palygorsklte and seplolite

g l a u c o n l t e a n d m a n g a n e s e nodules

(treoh w a t e r ) - i i l i t r + l e d l k l l e f r o m m i c a s

amorphous + organic material--

volconk glass

------

I

I I

-

172

G. I&LLER

The critical transition temperature between diagenesis and metamorphism is 300 “C and is almost independent of pressure (WINKLER, 1965). In burial metamorphism at normal geothermic gradients (30 O C/km) this temperature is reached at depths of ca. 10,000 m. In regional metamorphism connected with large-scale orogenesis, additional thermic energy is produced so that critical temperature occurs at much shallower depths. These are the p-t conditions of the greenschist facies whose low-temperature subfacies are characterized by the assemblage quartz-albitemuscovite-chlorite. If the original argillaceous material was poor in K and rich in Al, pyrophyllite becomes the characteristic mineral of the greenschist facies above 400 C (WINKLER, 1965): 1 kaolinite

+ 2 quartz = 1 pyrophyllite + 1 H2O

In recent years the so-called “zeolite facies” has been described from New Zealand, Australia, the Soviet Union and North America. It is the typical mineral facies of the burial metamorphism of sediments containing zeolites (for literature see WINKLER, 1965). The critical zeolite of this facies is not just any zeolite but laumontite (CaAlzSi4012. 4H20), which never occurs in unmetamorphosed sediments. Winkler, there fore, more accurately calls the zeolite facies the laumontite-prehnitequartz facies. The sedimentary assemblages of both analcime plus quartz and heulandite (or clinoptilolite) disappear under the same pressure and temperature conditions and are replaced by albite and laumontite;

+

+

analcime quartz = albite HzO heulandite = laumontite 3 quartz

+

+ 2H20

At only slightly higher temperatures are these reactions followed by a reaction between laumontite and calcite in which prehnite is formed, which is also characteristic for the zeolite facies: laumontite

+ calcite = prehnite + quartz + 3H20 + COZ

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173

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175

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VONENGELHARDT, W., MULLER,G. und KROMER, H., 1962. Dioktaedrischer Chlorit (“Sudoit”) in Sedimenten des Mittleren Keupers von Plochingen (Wiirtt.). Naturwissenrchaften, 49 (1962): 205-206. WEAVER, CH. E., 1959. The clay petrology of sediments. Clays Clay Minerals, Proc. Natl. Con$ Clays Clay Minerals, 6: 154-187. WEISS,M. P.,1954. Feldspathized shales from Minnesota. J. Sediment. Petrol., 24: 270-274. WHITE, W. A., 1961. Colloid phenomena in sedimentation of argillaceous rocks. J . Sediment. Petrol., 31: 560-570. U. G. and MCCARTER, R. S., 1958. Diagenetic modification of clay-mineral types WHITEHOUSE, in artificial sea water. Clays Clay Minerals, Proc. Natl. Con$ Clays Clay Minerals, 5: 81-119. WINKLER, H. G. F., 1965. Die Genese der metamorphen Gesteine. Springer, Berlin-HeidelbergNew York, 218 S. ZULLIG,H., 1956. Sedimente als Ausdruck des Zustandes eines Gewassers. Schweiz. Z. Hydrol., 18: 5-143.

Chapter 5

DIAGENESIS OF CARBONATE ROCKS GEORGE v. CHILINGAR, HAROLD I. BISSELL AND KARL H.

WOLF^

University of Southern California, Los Angeles, Calif. (U.S.A.) Brigham Young University, Provo, Utah (U.S.A.) The Australian National University, Canberra City, A.C.T. (Australia)

SUMMARY

Diagenesis of carbonate rocks comprises more than thirty different processes which are controlled by both local and regional factors, and which can alter the composition and texture of the sediments. Lithification is either physico-chemical or biochemical, and the controversial beach-rocks, for example, can be cemented by either process. Destructive processes include corrasion, corrosion, solution, decementation, and disintegration. Textural changes of limestones are often the product of inversion and several types of mechanisms, collectively termed “recrystallization”. Under favorable conditions internal sedimentation and chemical infillings can form complex open-space structures (e.g., stromatactis). The sparry calcite is divisible into granular, drusy and fibrous types of which there are various genetic types and some reflect the conditions of formation. Of the processes that form micrite, in particular “grain-diminution” is a relatively new concept. The replacement of carbonates may be slight to extensive and is either by other carbonates (e.g., dolomitization) or by non-carbonate minerals (e.g., silicification), and may or may not follow a predictable sequence. In general, the paragenesis of carbonate sediments can be grouped into predepositional, syngenetic, diagenetic and epigenetic processes, products and stages with numerous useful subdivisions related to Recent and ancient deposits. The sum total of all the syngenetic to epigenetic features can be used as an indicator of environmental conditions, in particular where each stage left some recognizable evidence. Hence, diagenesis is of practical applicability in exploration. The numerous possible syngenetic and diagenetic aspects discussed indicate that limestone and dolomite classification schemes should not be used indiscriminately in environmental reconstructions.

1

Present address: Oregon State University, Corvallis, Ore. (U.S.A.).

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INTRODUCTION AND DEFINITION

The widely accepted division of sedimentary processes into syngenesis, diagenesis, and epigenesis is inadequate for detailed research. The stages at which most of the individual processes and products occur cannot be distinctly demarcated. Two or more processes may be active simultaneously. They may overlap or the termination of one may mark the commencement of another process; and still other alterations may occur independently in both space and time. Many of the interpretations depend on the scale-thin-section, handspecimen or outcrop-at which observations are made. Hence, “pigeon-holing” of processes without contradictions is difficult, sometimes even impossible. Difficulties in genetic interpretations occur in particular in monomineralic rocks such as limestones. For these and other reasons it is not surprising that no general agreement has been reached on the definition and extent of diagenesis (see review by TEODOROVICH, 1961). Diagenesis has been restricted to those processes that cause lithification. Such a limited application, however, is arbitrary, artificial and impractical (NEWELL et al., 1953; GINSBURG, 1957). Not only are there several distinctly different lithification processes which are frequently difficult to recognize and separate, but they are so gradational as to defy precise definition and can occur at any stage during the early history of sediments. It is virtually impossible, therefore, to exclude other early alterations. It is preferable to apply diagenesis in a wider sense to processes that affect a sediment after deposition and up to, but not beyond lithification and/or filling of voids. Although these two processes can, and usually do, take place at different times within a sedimentary formation, especially if composed of different facies, the final stage of lithification and/or filling of voids appears to be the most convenient time at which diagenesis can be terminatedl. Hence, the following rather all-inclusive definition, in general agreement with the concepts of GINSBURG (1957) and KRUMBEIN (1942), has been adopted here: diagenesisincludes all physicochemical, biochemical and physical processes modifying sediments between deposition and lithification at low temperatures and pressures characteristic of surface and near-surface environments. Post-lithification processes grade into epigenesis, and epigenesis passes into metamorphismz. Epigenesis near the

1 This approach has been found to be. of particular use in coarse-grained or open textured limestones, but may be more difficult to apply in micritic rocks. Nevertheless, the paragenetic model used here is convenient, because after lithification (=infilling of voids by cement) of a rock the intrastratal fluids related to surface conditions cannot penetrate readily the rock framework. In cases where limestones maintain their porosity and permeability for a long period of time, even after being far removed from the original depocenter, the intrastratal fluids occupying the cavities can also be looked upon as either of syngenetic, diagenetic or epigenetic origin. Strictly speaking, epigenesis as defined here, passes into metamorphism only if an increase of pressure and/or temperature o m s .

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depositional environment is called juxta-epigenesis (“juxta-” meaning near), and epigenesis remote from the surface is named apo-epigenesisl (“apo-” meaning far, remote). Most limestones have some small voids which have been partly or wholly filled by one or more generations of cement. Hence, it is possible to divide diagenesis into pre-, syn-, and post-cementation stages as was the case in the study of a Devonian algal reef complex (WOLF, 1963a, 1965a, b). In other paragenetic investigations, however, the diagenetic-epigenetic boundary may have to be based on some other criterion to be determined by the individual investigator concerned. No definite rule is possible. As long as the boundary is precisely defined by certain fabric or structural relations, little confusion should occur. A diageneticepigenetic boundary established on the basis of a few thin-sections of a local outcrop, however, may have to be revised and shifted up or down the paragenetic scale as soon as the petrologic and petrographic information of the whole formation is available, or it may be found that the termination of diagenesis in one area may be completely unrelated to that of other localities. Not all diagenetic stages are present in limestones. For example, pre-cementation dolomitization may completely alter a limy deposit resulting in an elimination of syn- and post-cementation stages. The raw material of diagenesis, as KRUMBEIN (1942) called it, consists of organic and inorganic sediment of allochthonous and/or autochthonous origin, interstitial fluids and other components subsequently formed or introduced into the system. In general, it is possible to subdivide the components that interact during the diagenetic processes into the following (WOLF, 1963b): ( I ) Diagenetic-endogenic (2) Diagenetic-exogenic (a) supergenic-exogenic (b) hypogenic-exogenic This is merely an expansion of AMSTUTZ’S (1959) division: syngenetic-supergenic, syngenetic-hypogenic, epigenetic-supergenic, and epigenetic-hypogenic. In most cases, diagenesis derives its raw material from both endogenic (within the sediments) and exogenic-supergenic (outside source-from above) sources. One or the other may prevail. Under unusual conditions, however, a volcanic, i.e., exogenic-hypogenic, source may supply components for diagenesis without a marked increase in temperature. This would be particularly true for siliceous material introduced into a geosyncline (or other depocenter); during diagenesis of the sediments, quartzose arenites can be converted to orthoquartzites, carbonates can become siliceous, and fossils can be replaced prior to dolomitization. Diagenesis may express itself in a number of different ways. KRUMBEIN (1942) mentioned that a total of about 30 separate diagenetic processes have been de1 Weathering not related to the original depositional environment of the sediments is not included in diagenesis apd epigenesis as defined here.

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scribed in the literature. They may result in mineralogical changes, addition and removal of material, and textural and structural modifications and alterations ranging from slight to extensive or complete. Several generations of diagenesis may each leave evidence, or each successive one may obliterate or destroy the products of earlier processes. In many cases, however, diagenesis may appear to be absent if only visually obtained information is consideredl. The division of sedimentary processes into several subdivisions (see below) is suitable to our present purpose and state of knowledge. Even the pre-, syn-,and postcementation subgroups, however, will become arbitrary and artificial with an increase in our understanding of diagenesis. To ascribe to certain products merely a genetic term such as “precementation-diagenetic”, without relating it to the sediment’s history as a whole, invites criticism. It is more accurate to relate all processes and products to a paragenetic sequence. In other words, a paragenetic scheme furnishes less ambiguous information in cases where it seems impossible to define exact syngenetiediagenetiopigenetic boundaries. The absolute time of formation may be impossible to determine, but the textural and structural relationships permit the interpretation of relative time of formation. The widely used terms “primary” and “secondary” have very little meaning in diagenetic investigations unless precisely defined, although they may be quite useful in a very general colloquial sense. Paragenetic interpretations are relatively easy and non-controversial if the investigation is made on the scale of one thin-section or handspecimen. Syn-, dia-, and epigenetic processes, however, are not only gradational, and overlap in time and space on a microscopic scale, but especially do so on a regional scale. Regional diagenetic studies may be rather tedious and resemble structural analyses, for example, in that the micro-, meso-, and macro-scopically examined features are assembled step by step. The so-called pre-depositional (or pre-syngenetic) processes and products can be deduced from limestone rock fragments (=calclithite fragments of FOLK, 1959; extraclasts of WOLF, 1965b; see below) which were derived from older limestones that had undergone diagenesis, e.g., lithification, recrystallization, dolomitization and silicification, before erosion and transportation, as has been possible in the Nubrigyn-Tolga algal reef complex of New South Wales (WOLF, 1963a).

Note that diagenesis is not part of the petrographic (=descriptive) stage but belongs to the subsequent stage of petrology and petrogenesis (=interpretive). Reliable diagenetic reconstructions cannot be made, therefore, on a few local thin-sections, but must be based on as much geochemical, petrographic and stratigraphic information as circumstancespermit to be obtained.

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DIAGENESIS OF LIMESTONES

Factors controlling diagenesis

Certain factors will initiate diagenesis, and the same or other factors will perpetuate the old and/or cause commencement of new diagenetic processes. The sediments have a tendency to adjust to new physical and chemical conditions and would, theoretically, reach equilibrium. The micro- and macro-environmental conditions above and within the sediments, however, change continuously. Sometimes, equilibrium may be established, as, for example, in cases where limestones are completely replaced by iron oxide, silica or dolomite. In many cases, however, the physical and chemical conditions shift so rapidly that only a small fraction of the reactions involving the limestone framework reach equilibrium. In particular during the early diagenetic stages numerous successive and overlapping processes will be acting at a relatively fast rate on both micro- and macro-scales, when movements of interstitial fluids are at a maximum, biological activity is producing chemically active substances, maximum pore space is available, temperature change is more or less sudden due to diurnal exposure, and so forth. The following list of factors influence diagenesis of carbonate sediments: (I) geographic factors (e.g., climate, humidity, rainfall, type of terrestrial weathering, surface water chemistry); (2) geotectonism (e.g., rate of erosion and accumulation, coastal morphology, emergence and subsidence, whether eugeosynclinal or miogeosynclinal); (3) geomorphologic position (e.g., basinal versus lagoonal sediments, current velocity, particle size, sorting, flushing of sediments); ( 4 ) geochemical factors in a regional sense (e.g., supersaline versus marine water, volcanic fluids and gases); (5) rate of sediment accumulation (e.g., halmyrolysis, ion transfer, preservation of organic matter, biochemical zonation); (6) initial composition of the sediments (e.g., aragonite versus high-Mg and low-Mg calcite, isotope and trace element content); (7)grain size (e.g., content of organic matter, number of bacteria, rates of diffusion); (8) purity of the sediments (e.g., percentage of clay and organic matter, base exchange of clays altering interstitial fluids); (9) accessibility of limestone framework to surface (e.g., cavity systems permit replacements); (10)interstitial fluids and gases (e.g., composition, rate of flow, exchange of ions); (11) physicochemical conditions (e.g., pH, Eh, partial pressures of gases, C02 content); (12)previous diagenetic history of the sediment (e.g., previous expulsion of trace elements will determine subsequent diagenesis). The numerous large-scale environmental parameters listed above influence in one way or another the more local environments and these in turn influence the micro-environments. There is a complete gradation and overlap of these macroand micro-factors as one example below illustrates (WOLF, 1963b): Climate

particle

J. [size Geomorphology

amount and rate of pH type of +type of --+ --f and+ replacement diagenesis Eh bacteria

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The actual processes that lead to diagenetic alterations and modifications of limestones are divisible as follows (among other): ( I ) Physicochemical processes: solution, corrosion, leaching, bleaching, oxidation, reduction, reprecipitation, inversion, recrystallization, cementation, decementation, authigenic mineral genesis, overgrowth, crystal enlargement, replacements, chemical internal sedimentation, aggregation and accretion. (2) Biochemical and organicprocesses: accretion and aggregation,particle-size reduction, corrosion, corrasion, mixing of sediments, boring, burrowing, gas-bubbling, breaking down and synthesizing of organic and inorganic compounds. (3)Physicalprocesses: compaction, desiccation, shrinkage,penecontemporaneous internal deformation and corrasion, and mechanical internal sedimentation. Many of the above processes are commonly considered syngenetic. As they can occur within the sediments and directly alter and influence diagenesis, however, they must be considered as part of diagenesis. It is the total or collective influence of all factors that must be examined in a final analysis. As KRUMBEIN (1942) pointed out, variations in the diagenetic end-products may occur either with different sediments in the same environment, or with the same kind of sediment in different environments. Compaction

Compaction of sediments is the process of volume reduction expressed as a percentage of the original voids present. The process affects mainly loose, unlithified limestones and, of course, other sediments not considered here. Autochthonous limestones such as reefs do not undergo much compaction. The intergranular spaces of allochthonous deposits are eliminated by closer packing, crushing, deformation, expulsion of interstitial fluids, and possibly corrosion of the grains. KRUMBEIN (1942) gave the following values of porosities or amounts of fluid content of freshly deposited material: sand = 45%, silt = 50-65%, mud = 80-90%, and colloids (less than l p ) = approximately 98% water. Lime-mud apparently behaves similarly to clay minerals. The degree of compaction, in general, depends largely on the ratio of fine to coarse material and on the character of the sediment framework. Fine-grained sediments undergo the highest degree of compaction in the first foot (GINSBURG, 1957). As he suggested, the negligible weight of overlying sediments cannot cause compaction during this early stage. Ginsburg believed that mixing by organisms, the gel-like character of the sediment and the escape of bacterial gases contribute to rapid packing. The burial pressure of sediment accumulation becomes effective somewhat later to produce some sort of physical cohesion between the particles. The expulsion of interstitial fluids and gases during compaction may be predominantly vertical or horizontal. Even freshly deposited sediments have different degrees of permeability and, although unlithified, some may act as

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“cap” rocks to cause very early horizontal fluid movements. Differential compaction may determine the direction and rate of fluid and gas migration. NEWLLet al. (1953) proposed, for example, compaction-induced lateral movement of CaCOsrich basinal fluids toward the reef-talus and reef, with a subsequent precipitation of carbonate cement, to explain the well-lithified state of these sediments. Interbedded carbonate and clay mineral accumulations may both have at the beginning interstitial water that is identical in composition. Differential adjustment of the fluids to their new micro-environmentswill soon take place. In particular the clay minerals will motivate chemical changes. For example, the Ca-Na relationships seem to depend partly on the base exchange properties of clay minerals (VONENGELHARDT, 1961). During compaction the interstitial fluids of a number of different physicochemical horizons intermingle and cause reactions that may have significant results. Clay mineral layers between limestone beds may cause filtration of certain cations when fluids pass through them. WERNER (1961) believed that during compaction the fluid-movements from a clay into iron-oolite deposits caused filtration along the clay-oolite boundary. Certain cations were held back and remained below the boundary, whereas others passed freely until the concentration was sufficiently high within the oolite sediment to result in precipitation. The poorly lithified oolite beds exhibit compression features in contrast to the well-cemented oolites which lack any signs of deformation. This seems to indicate to Werner that the precipitation of the CaC03 cement must have been early diagenetic and occurred during compaction when the interstitial fluids assured sufficient quantities of chemicals. Werner’s explanation may well apply to limestone deposits that contain clay-rich beds. Shelf-to-basin facies of Pennsylvanian and Permian carbonates in the Cordilleran miogeosyncline are examples. Compaction processes can alter textures and structures of carbonate rocks. Poorly cemented faecal pellets, for example, have been reported to form a textureless lime-mud a few inches or feet below the surface due to merging of the individual grains. Movements of connate waters during compaction may form tubes, channelgand bubbles (CLOUD et al., 1962) which, when filled by carbonate cement, resemble the so-called birdseyes (dismicrite of FOLK,1959; dispellet of WOLF,1960). An excellent example in the geologic record is found in lower limestones of the Ely Group in the Hamilton district of White Pine County, Nevada. On the other hand, lime-mud, more so than its clayey or muddy terrigenous counterpart, may lack compaction especially if early cementation took place. The result may be a loose, sponge-like dismicrite with minute sparite-filled voids. However, relatively large open-space structures in micrite and pellet limestones, for instance, cannot be explained by lack of compaction. It has been suggested that soft-bodied organisms became buried and upon decomposition left voids. Although this is possible in some cases, most cavities in micrite limestones are probably of inorganic origin and Algae were responsible only in an indirect way (WOLF,1965a). Rate of cementation, degree of compaction and pressure-solutionare close-

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ly related. If the former varies on a regional scale, the latter may follow the same pattern. For example, loosely packed, birdseye-rich shallow-water Nubrigyn algal calcarenites of New South Wales must have undergone early cementation in contrast to the basinal algal, graded-bedded deposits that exhibit tight packing and extensive pressure-solution. Structures believed to have been the product of compaction were described by TERZAGHI (1940); and early diagenetic formation of cone-in-cone structures 1963; may be related to compaction of clayey micrite limestones (USDOWSKI, see also the section on recrystallization). Lithijcution Lithification is the process that changes unconsolidated sediments into weakly to strongly consolidated rocks. Lithification may occur through cementation, recrystallization, replacements (i.e., dolomitization), crystallographic welding of limemud, and by other processes such as desiccation. Only cementation will be considered in this section. Cementation, as understood here, is the process of openspace filling by physicochemical and biochemical authigenic precipitates, and excludes allogenic internal sediments. In cases of allochthonous limestones, cementation causes lithification; but autochthonous carbonate rocks may merely undergo a decrease in porosity and permeability without marked consolidation.’ Cementation of limestones is very often associated with solution, corrosion, leaching and replacement phenomena, and can form a number of generations until the available open space is completely eliminated. Some limestone bodies are cemented stratum by stratum, whereas others are cemented “en bloc” (KAYE, 1959). Precipitation of carbonate cement can take place in littoral environments and subaerially; within sediments but above the water-table; at or near the watertable; below the water-table; in zones where fresh water mixes with marine water, and normal marine with supersaturated waters; and under a thick overburden. As JAANUSSON(1961) indicated, however, no undisputable evidence of recent submarine calcium carbonate cementation at or close to the sediment-water interface has been reported in contrast to the widely occurring cementation processes above low-tide level. Calcium carbonate is the principal cement in limestones and occurs as aragonite and various types of calcite in Recent and near-Recent sediments, and as calcite in older rocks (see inversion). The CaCOs source for cement may be either endogenic or exogenic, i.e., the carbonate may be derived from within the formation or brought to the site of precipitation from an outside source. Calcite is also one of the most significant cement types in terrigenous sediments. It is interesting to note, therefore, that whatever the original conditions of the depositional environSee discussion by CRICKMAY (1945).

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ment may have been, subsequent convergence of the physicochemical conditions appears to permit the formation of CaCOs cement. The explanation of calcite cement in a wide variety of rocks may lie in the independence of both Ca2+and COa2- of the Eh parameter (KRUMBEIN and GARRELS, 1942). Precipitation of calcium carbonate can be brought about by numerous factors discussed below (NIGGLIand NIGGLI,1952). When considering the problems of CaCOs solution and precipitation one has to distinguish between: ( I ) changes in solubility when the COZ content remains constant or CO2 is absent, and (2) changes in solubility when there is a possibility of C02 decrease or increase. (I) The solubility decreases if the free C02 content remains constant or if free C02 is absent (and when the pressure of carbonic acid remains constant), i.e., C a C a is precipitated from a saturated solution: (u) With increasing temperature, if C02 is not present. (For example, in COz-free sea water the solubility product constant of CaCOs at 0" C is 8.3 10-7 and at 30°C it is 4.4 10-7.) (b) With decrease in the hydrostatic pressure associated with a decrease in dissociation of carbonic acid at constant COZcontent. (c) With a decrease in soluble NaCl or Na2SO4, etc. (so-called salt content) at constant gas pressure of the carbonic acid in the gaseous phase. This is due to the changes of the dissociation constant of the carbonic acid and the constant K (solubility product), where (Ca2+)(C0s2-) = K. (At 20°C and a salinity of 35%,, K = 6.2 * 10-7; at 20°C and no salt content K = 0.5 * 10-8.) (a) With addition of Ca ions (for example, bonded to Sod, Cl), or if these are present and not bonded to carbonate and will form new (NH&COs as a result of organic decomposition. (e) When only calcite can exist and not the unstable aragonite or vaterite. df)When the water evaporates. (2) Water, to which C02 has been added, increases the solubility of CaCOs because of formation and dissociation of H2COs into Hf and HCOs-. The added H+ will combine with the COs2- ions already present to form the more stable HCOs-. In order to reach equilibrium at a constant solubility product, therefore, more Ca has to go into solution. On the other hand, because of this phenomenon a decrease in CaCOs solubility occurs, i.e., CaCOs is deposited, when the carbonic acid content decreases. That is the case: (a) When the partial pressure of CO2 in sea water, that is in equilibrium with the COz of the atmosphere, decreases. (The C02 content in the atmosphere is increased by volcanic activity, respiration of animals, decomposition of organic substances, etc. On the other hand, CO2 is removed from the atmosphere by photosynthesis.) (b) When the pressure decreases (at constant temperature, CO2 escapes into the atmosphere). (c) When the temperature increases (at constant pressure and constant par-

-

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tial pressure of C02, with increasing temperature COz is freed into the gaseous phase). (a) When organic substances are formed by plants in the marine waters (in contrast to the animals which exhale C02). (e) When the formation of COz through organic decomposition is reduced or made impossible. (f)When the salinity increases, because less COz can be dissolved in marine than in fresh water. In general, the above conditions conducive to CaC03 cement precipitation are controlled by three main processes discussed further below, namely, physicochemical, bacterial and decompositional, and algal processes. Little work has been done on rock cementation and much of the following presentation is confined to theories that attempt to explain the formation of Recent and Pleistocene beachrocks. Least of all is known about the factors that control genesis of fibrous, drusy and granular carbonate cement (see carbonate types). Very little precise information is available on the parameters that control precipitation of aragonite in preference to high-Mg and low-Mg calcite. Many of the experimental results appear to be contradictory. GOTO’S(1961) experiments suggest that slow reaction, higher pH value of solution, diminished solvation effect of water, and balanced proportions of Ca2+ in relation to c032-are responsible to some extent for the formation of aragonite and are less favorable to calcite genesis. The presence of Mgz+, Sr2+, and Ba2+ seems to be unfavorable for aragonite formation. On the other hand, experiments by ZELLERand WRAY(1956) suggest that aragonite formation is favored by low Mn2+ but also by high Sr2+, Ba2+ and Pb2+ contents, which differs from Goto’s conclusions. Other factors evidently are responsible to cause such seeminglycontradictory results. In the present-daycalcareous sediments aragonite is formed especially in shallow-waterenvironmentsin tropical and semitropical regions suggesting that temperature is significantlyinfluential in the genesis of aragonite. Physicochemical precipitation GINSBURG (1957) mentioned that extensive cementation of beach-rocks occurs in those young carbonates that are subaerially exposed or located in zones of meteoric waters. Those sediments still in a marine environment or above the groundwater table are very friable. This agrees in general with the observation made by KAYE (1959) who stated that cementation of beach-rocks of Puerto Rico is not coincidental with high tide but extends up to 3 ft. above it. The upper limit is possibly controlled by capillary action or by splash. The lower limit lies slightly below low spring tide and is probably controlled by the lowest level of wave trough at low spring tide. These beach-rocks are mainly cemented by calcite rather than aragonite. As Kaye examined very recently formed sediments, it seems that inversion or recrystallization from aragonite to calcite is unlikely. In many other localities, however, beach-rock is cemented by aragonite. ILLING (1954) suggested

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that calcite is precipitated from fresh, and aragonite from salt water. Some of the Puerto Rican beach-rocks were locally cemented by iron oxide which was derived from a rubbish dump containing iron objects. KAYE(1959) pointed out that this dump is not more than 80-100 years old, which is indicative of relatively rapid alterations. EMERY et al. (1954) reported that lithification of beach-rocks may be due to the consolidation of interstitial pasty lime-mud matrix. They believed that consolidation results from precipitation of microcrystalline carbonate within the paste, or from crystallization of the paste probably during daytime periods at low tide when the tidal waters are warm, and when algal processes utilize much COz. Inasmuch as most beach-rocks appear to lack a matrix, however, this explanation is of restricted application. Cementation of littoral, beach, and dune limestones has also been explained by the action of fresh water that dissolves part of the carbonate framework and later precipitates it upon evaporation, aeration and possibly with the help of organisms such as bacteria. RUSSELL (1962) supported this theory by showing different degrees of corrosion of carbonate grains in relation to the fresh ground-water table (see corrosion). Cementation of these Puerto Rican beach-rocks investigated by Russell seems to take place in the vicinity of the fresh-water table. The beachrock is thickest where seasonal contrasts in sea level are most pronounced. It seems that with the changes of sea water level, the fresh-salt water table is displaced correspondingly, with the lighter fresh (or brackish) water floating above more saline water. Thus, the zone of cementation is shifted causing thickening of the beachrock. The cement is almost wholly calcite with subordinate amounts of aragonite. Although an endogenic origin of cement is likely where there are signs of internal corrosion, this theory is not applicable in cases where no solution of the carbonate sediments has taken place. In some instances, therefore, the calcium carbonate must have come from an exogenic source. This is supported by the carbonate cementation of beach-rocks composed wholly of terrigenous material, i.e., quartz, which could not have supplied any endogenic CaC03. CRICKMAY (1945) discussed some interesting examples in limestones of Lau, Fiji. Related to the above is the theory that weak acids, in particular humic acids, percolating down into the fresh ground-water lower its pH. Thus, the fresh water dissolves carbonate from the surrounding medium, and a t its contact with the underlying salt water precipitation of the calcium carbonate cement may occur. Seasonal and yearly fluctuations of the fresh-salt water interface may cause a thick zone of cementation. However, MAXWELL (1962), for example, found this theory inapplicable to the beach-rocks of Heron Island, Great Barrier Reef. Capillary action is a third likely process that supplies CaC03 in some cases. In particular, limestones subaerially exposed may be heated to the extent that fluids are brought up from a lower level by capillary forces, and deposit carbonate cement on evaporation and aeration. Certain tufa, travertine and caliche deposits

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are the result of this process: for example, the areally-extensive tufa and travertine deposits of Pleistocene and Recent age within alluvial fans, slope-wash, and even lacustrine sediments of parts of the Great Basin area of western United States (in particular southern Nevada). Another hypothesis is based on the premise that fresh ground-water brings dissolved calcium carbonate from the hinterland, which is followed by precipitation as this water seeps out through the coastal sediments. Various objections have been raised against this process because it is thought that it does not explain the sporadic occurrence of beach-rock, for example. Also, cementation of sediments takes place where the hinterland is devoid of carbonate source material. It has been pointed out that present-day accumulations of tufa along the shores of Utah Lake west of Provo, Utah, consist of beach-rock, shell heaps and other materials more or less tightly cemented by calcium carbonate (BISSELL, 1963). It was also noted that beach-rock and massive tufa deposits formed along the shores of Pleistocene Lake Bonneville through combined action of wave splash and Algae in releasing C02 and thus precipitating CaC03 (BISSELL, 1963). GINSBURG (1 953a) believed that, because the beach-rocks he investigated are exposed to saturated marine waters and provide abundant nuclei for CaC03 precipitation, cementation occurs due to heating and evaporation of interstitial fluids. At low tide, water remains as intergranular films and permits a more complete exchange of Cog between atmosphere and solution, inducing a more rapid equilibrium and precipitation from supersaturated solution. According to REVELLE and FAIRBRIDGE (1957), the water temperature of splash pools just above tide limit in Western Australia varies from 13°C at night to 24°C in daytime, with a change of pH from 8.2 to 9.4. During the night the pH remains equal to 9 and is ample to account for large precipitation of carbonates. The above authors also mentioned the formation of superficial pelagosite crusts and pore-space fillings formed by the action of spray and evaporation. WOLF(1963a) has similarly explained the numerous open-space calcite patches in internal channels and cavities of Devonian littoral algal bioherms. The internal voids must have undergone a sharp temperature increase during low tide when the reef-structures were directly exposed to sunlight. The films and small patches of intrastratal fluids that remained behind at low tide then reached supersaturation and precipitated CaC03 on the cavity walls. The restriction of penecontemporaneous carbonate cementation to particular localities, e.g., intertidal zones, whereas it is absent below low tide, is explained by GINSBURG (1953a) by the sluggishness of the equilibrium between solid and dissolved calcium carbonate, and the inhibition of this equilibrium by organic matter. According to experiments, it may take 6-8 h under laboratory conditions for a system to reach equilibrium (HINDMAN,1943; MILLER,1952; both in GINSBURG,1953a). DALY(1924, in GINSBURG, 1953a; and KAYE,1959) thought that beach-

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rock cementation occurs in two stages. An initial precipitation of CaC03 from sea water was a result of ammonifying action of decaying organic matter originally incorporated in the sediments as detritus. The chemical reaction consists of ammonia combining with C02 to form ammonium carbonate, which then reacts with calcium salts in solution to form CaC03. Daly envisioned a second stage of cementation during which the precipitation of CaC03 from marine water is caused by aeration and surf agitation, and their effect on the COZ partial pressure in the water. The first stage provides nuclei essential to the deposition of CaC03 in the second stage. The varying distribution of detrital matter can, therefore, explain the localization of cementation and formation of beach-rock. As Kaye pointed out, however, numerous localities rich in organic matter lack beach-rock genesis. Kaye discussed at length the physicochemical factors and had to reject them as an explanation for beach-rock cementation. He believed that the Puerto Rican sediments were formed most likely by microbiological processes. One has to conclude that the problems of physicochemical cementation have not been solved. Whatever the factors are, they cannot be of equal importance in all environments. It seems that temperature is one of the most important parameters and restricts beach-rock genesis to tropical and sub-tropical localities. Other factors, however, must be of equal significance as beach-rock formation does not take place at many localities where temperature, evaporation and other conditions seem to be favorable for cementation. Bacterial processes and decomposition Bacterial processes and decomposition of organic matter are closely linked and are inseparable in the study of organic influences on CaC03 precipitation, and other diagenetic processes. Recent calcareous sediments may contain bacteria in concentration from about 10 to 10,000 billion/g in contrast to the water above the sediments that contains only 10 to 1,000 organisms/mm3. ZOBELL(1942) has reported even higher concentrations of bacteria in sediments. The quantity of bacteria is a function of sediment grain size and presence of organic matter. The heterotrophic bacteria depend on organic matter as a source of carbon or energy. Not much organic matter, however, is required to assure the presence of some bacteria. Small quantities of organic particles adsorbed on sand grains and cavity walls suffice. The type of bacteria and availability of oxygen control the depth to which oxidation of organic matter to COZcan take place. For example, GINSBURG (1957) mentioned that in the reef and back-reef deposits investigated by him most of the organic matter is relatively rapidly removed and the sediments are, therefore, light colored and have only a slight odor of HzS at depth. On the contrary, the shallow-water calcareous muds may contain three to six times as much organic compounds. Here, only the upper fraction of an inch is light colored and the sediments as far down as 8 ft. smell strongly of H2S. The limy muds of the Red Ssa

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reef complex, although light colored, give off a slight odor of HzS and, when placed in a closed bottle, become black and form minute pyrite crystals 1 mm long. Lagoonal white carbonate muds of Pacific atoll reefs become black and saturated and TERMIER, 1963). All these occurwith HzS only an inch or so down (TERMIER rences are attributable to bacterial processes. Based on these observations, GINSBURG (1957), among others, concluded that shallow bays can form a type of barred or stagnant basin in which the original organic matter need not be diagenetically removed. Bacterial decomposition of proteins leads to the formation of carbon dioxide, ammonia, hydrogen sulfide and a variety of intermediate products, and many carbohydrates are converted into carbon dioxide, carbon monoxide, methane and organic acids (ZOBELL,1942). Some species oxidize, for example, organic calcium salts, thereby increasing the Cazf concentration. On the other hand, the autotrophs obtain energy from oxidation of inorganic substances such as ferrous iron, manganous manganese, hydrogen sulfide, hydrogen, carbon monoxide, methane or ammonia. These latter bacteria are aerobes, i.e., require free oxygen, and thus predominate mainly in the surface sediments. The bacteria may also produce significant amounts of biocatalysts or enzymes which can activate numerous chemical reactions. Some of these catalysts continue to be functional after death of the bacteria (ZOBELL,1942). The biologically mobilized material in turn stimulates and controls diagenesis by changing the pH, Eh, partial pressure, composition of interstitial fluids, temperature, and so on. For example, when normal sea water (chlorinity 19x0 at 25°C) is in equilibrium with solid CaC03, the pH remains stable at approximately 7.5 due to bicarbonate-carbon dioxide reactions. This carbonate buffer system can be changed by bacterial decomposition of organic matter and by addition of acidic or basic components. COz addition acidifies water, whereas sulfate reduction may produce either an acid or an alkaline effect depending on the organic matter and products of decomposition. Addition of ammonia increases the pH to alkaline conditions, and oxidation of ammonia to nitrate and nitrite causes a decrease in pH. CLOUDet al. (1959) reported that an increase in hydrogen ions, with an accompanying decrease in pH, is caused by the bacteria that produce COz. Together with HC03- addition from the reactions: COZ HzO+HzC03+Hf HCOs-, the combination of new H+ with CO32- to produce still more HC03causes a reduction in the CO32- component of alkalinity, and allows both Ca2+and alkalinity to reach high values without precipitation. The interstitial fluids are, therefore, in a condition favoring precipitation whenever they may be exposed to an environment of higher pH and higher C032- content. This condition may be fulfilled when, for example, bacterial production of ammonia causes an increase in pH, waves or organisms stir up bottom sediments, burrowers transfer sediments, or when lateral movements of interstitial fluids bring them to a higher pH environment; CaC03 precipitation and cementation may then result. In the latter case of

+

+

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lateral movements of fluids, the transfer to a higher pH locality, results in loss of H+ from HC03- and in increase in CO32- content which causes precipitation of CaC03. This may explain, as CLOUDet al. (1959) pointed out, the seaward increase (toward the bank margin) of aragonite lithification of pellets and algal grains, and the cementation of pellets to form lumps (grapestones). Measurements by these authors indicate that this lithologic change from lime-muds to pellets and lumps is accompanied by rising pH and Eh as pore space and circulation of oxygenbearing water increase. It seems, therefore, that lateral movements of interstitial waters, in addition to other factors, may cause regional variations in limestone lithology, textures and structures. PURDY (1963) mentioned the possibility of cementation of bahamite sediments by decomposition of organic detritus by ammonifying and nitrate-reducing Bacteria which produce ammonia. The latter reacts with the calcium bicarbonate in the immediate surrounding water and causes CaC03 precipitation. Such a process can occur on a very small scale within pellets, for example, and cause cementation and preservation of very friable material. REVELLE and FAIRBRIDGE (1 957) concluded from the available published evidence that bacterial precipitation of calcium carbonate in Recent marine environments seems to be strictly limited in scope. This may be correct if one speaks of the total volume of carbonate sediments but, as the considerations above suggest, it does not exclude the possibility that bacterial activities lead to cementation and stimulate other diagenetic processes. Algal cementation Algal cementation is one of the most important lithification processes in shallowwater limestone genesis. The lime precipitates of blue-green, green and red Algae can occur as crusts and are then considered more or less of syngenetic origin. As it seems possible, however, that microscopic cells and filaments can exist for short periods of time to some depth within sediments, and as Algae in general cause chemical modifications in surface and interstitial waters, the Algae must find a place in a discussion on cementation and diagenesis in general. Calcareous Algae play a dual role: on one hand they dissolve lime possibly by use of oxalic acid or by indirectly acidifying the water; and on the other, the same organisms cause CaC03 precipitation. GINSBURG (1957) listed the corrosive action of Algae as one of the major diagenetic processes changing extensively Recent calcareous sediments on the Bahama Bank and in Florida. Precipitation of CaC03 by Algae occurs through respiration processes. At night the Algae give off C02 into the microenvironment which may result in corrosion and solution, whereas in daylight COz is utilized during photosynthesis leading to alkaline conditions sufficiently strong to cause precipitation of CaC03 and pelagosite. It seems then that corrosion and precipitation may alternate, and the predominance of one over the other depends on local circumstances. EMERY et al. (1954) mentioned the

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possibility that boring Algae which dissolve chemically surface layers might cause precipitation of calcium carbonate a distance down within the sediments. Another indirect process of algal cementation may be due to the detritusbinding Algae. Schizophytu filament and cell colonies, for example, when covered by a sudden influx of detrital sediment are not necessarily killed but move upward through the layer and re-establish themselves above the sediment. Recent stromatolites are formed by similar processes. As Algae give off oxygen and use COZ during their life processes, it may be possible that they stimulate diagenetic changes within the uppermost sediment accumulations while they move upward. The micro-environment beneath algal mats may also be conducive to corrosion and precipitation. The waters beneath the mats are relatively isolated from the overlying sea water as indicated by emission of HzS odor from the mats upon disturbance (REVELLE and FAIRBRIDGE, 1957). The metabolic activity of Algae and the decay of organic matter cause marked changes in C02 content and other properties of the water. pH variations from 6.5 to 8.7 in fluids collected from beneath algal mats in Tahiti are sufficient for solution and precipitation of CaC03. In addition to the indirect algal influence, some genera are capable of precipitating carbonate on their surfaces and/or internally, which is partly due to absorption of C02 from the water medium. This biogenic carbonate is commonly dense cryptocrystalline in thin-section (WOLF, 1965a). It is quite conceivable that the so-called umbrophile, i.e., shade-adapted, Algae can exist within the upper layers of sedimentary frameworks and cause precipitation of thin crusts around detrital grains and on walls of voids. For example, WOLF(1963a) described beach-rocks of Portuguese Timor composed of skeletal and algal debris and volcanic rock fragments which are circumcrusted by layers of dark brown, nearly opaque, cryptocrystalline to very dense calcium carbonate (Plate I). This carbonate is identical to the algal debris forming part of the detrital framework, and where this debris is encrusted by the brown layers, the two merge completely and can be distinguished only with difficulty in thin-section. Wolf has also shown that algal calcareous deposits of algal pisolites (Plate 11) and algal crusts (Plate 111, IV) of very recent origin can change relatively quickly from a thin upper layer containing clear algal cellular structure to a dense textureless cryptocrystalline calcium carbonate a few millimeters below. Intermediate cases with faintly preserved cellular features are also present. Remarkably similar changes have been reported in a Devonian algal reef complex (Plate V). The exact mechanism of this alteration is not clear, but it seems, that the mere absence of algal features does not preclude a floral origin for the cryptocrystalline circumcrusts such as those that lead to the cementation of Portuguese Timor beach-rocks. Further research is required to confirm these observations, in particular because it is quite likely that other processes can give rise to cryptocrystalline carbonate cement.

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PLATE I

A section of a beach-rock of Portuguese Timor composed of a dense algal fragment ( I ) and

skeletal fragments (2). All have been surrounded by a crust of micrite (3). The remaining white spaces are pores. Petrographic studies indicate that the dense algal grain has most likely formed by “degrading recrystallization” (see text). It is, therefore, a fragment of pseudomicrite. The origin of the micrite crust or “cement” is not certain-it may have been formed by direct biochemical precipitation or by degrading recrystallization of either physicochemical or algal calcite (see text). Note the merging of the dense fragment and the crusts.

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PLATE I1

Section of a Recent algal pisolite exhibiting distinct cellular features ( I ) grading into dense cryptocrystalline calcium carbonate (2) toward the nucleus. The latter may be a product of “degrading recrystallization”or disintegration (see text). (Thin-section supplied by Mr. J. Standard.)

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PLATE 111

Recent algal biolithite composed of cells ( I ) and dense cryptocrystalline carbonate (2), which may be micrite formed directly by algal precipitation, by “degrading recrystallization”, or disintegration of algal cellular material (see text). Note the resemblance between the material in Plate I1 and 111 of recent origin and that of the Devonian in Plate V.

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PLATE IV

Recent algal micrite crust. It is a product of direct algal precipitation (= orthomicrite), or formed by “degrading recrystallization” or disintegration of an algal cellular colony (= pseudomicrite). Important to note is that whatever its origin, it is automicrite, i.e., formed in situ. Therefore, it is either an ortho-automicrite or a pseudo-automicrite (see text and Table 111). Many of the Devonian algal bioherms, Nubrigyn Formation, N. S. W., consist of identical dense algal automicrite-biolithites.

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PLATE V

Micrite and algal filament-biolithite of a Devonian algal bioherm, Nubrigyn Formation, N.S.W. One stromatactis ( I ) and a recrystallized stromatoporoid crust (2) are present. The filaments are Rothpletzellu (3). Most of the dense micrite is automicrite of which most of the 300 knoll-shaped bioherms are composed. Note the resemblance with the material in Plates I11 and IV. Compare it with the micrite circumcrusts in Plates I, XV, XVI and XIX; and with algal pellets and grains in Plates I, X, XIIL, XV, XX and XXIV. (The filaments have been identified by Professor J. H. Johnson.)

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Corrusion, corrosion, solution, decementution, disintegration

Several processes can alter and obliterate limestones during syngenetic, diagenetic and epigenetic stages. Some of the products may have (resemble) present-day, near-surface weathering features. Corrasion and corrosion may occur very extensively in littoral and subaerially-exposed limestones and can form diagenetic “micro-karst’’ structures. Little is known as to how significant these processes are in sub-littoral carbonate rock environments. They deserve considerably more attention because they may be excellent paleogeographic indicators, may illustrate formation and destruction of porosity and permeability, and may influence diagenesis to a large degree, Internal cavity systems originate in a number of ways. Many of them were originally surface depressions of various kinds which became part of the limestone framework. For example, the pitting of algal-encrusted limestone surfaces reported by KAYE (1959) resulted in depressions and irregularities from a few millimeters up to a few feet in size with shapes changing systematically corresponding to the environment. A relative rise of sea level causing spreading, or transgression, of the encrusting Algae over the smaller pits would result in an internal cavity system. A similar process is mentioned by J. W. WELLS(1957) who described recent surge channels with upper “eaves” of calcareous algal deposits. These spread until the channels are completely roofed over and constitute part of a tubular labyrinth within the limestone body. From this stage onward, diurnally surging waters and solutions can penetrate the limestone to flow internally below the surface and cause internal corrosion and abrasion, internal sedimentation, replacements, and chemical precipitation, all described later. As many of the voids have been inherited from the surface, they are mostly horizontally oriented and are often concentrated along specific beds. Hence, the internal fluids will continue to flow predominantly horizontally and any further corrosion occurs mainly sideways and to a lesser degree downward, commonly resulting in flat-bottomed voids. Internal open-space structures in Devonian algal bioherms (Plate VI-XIV) have been reported by WOLF(1963a, 1965a,c), among others. Analogous to the recent occurrences, many have flat bottoms with irregular upper parts and are predominantly horizontally oriented. In serial thin-section and polished section studies it can be demonstrated that two or three horizons of cavities meet laterally to form part of one system (Plate VII). Many of the cavities, all of which have been completely filled by secondarily introduced material, exhibit solution and/or abrasion features (Plate VI, VIII, IX, XII). Where extensive alterations occurred, it is impossible to determine the original factors that localized the fluids, for evidence of the last process only is present. In less altered void structures, however, it can be shown that many were originally inter-biolithite spaces (i.e., spaces between colonies) and channels with algal overgrowths. Some of the algal filaments and cells clearly line and follow more or less concentrically the outlines of the channels.

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PLATE VI

Complex section of a Nubrigyn algal bioherm showing algal colony (I), and detrital skeletal fragment enveloped by algal micrite layer (2). A large cavity, with some features of differential solution has been filled by allomicrite internal sediment (3) and some clear granular orthosparite (4). A stylolite cuts across the slide (5). Several patches of dense algal automicrite (6) are recognizable.

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PLATE VII

Part of a Devonian Nubrigyn algal bioherm consisting of lower filament colony (I) overlain by spongy algal growth and one coral fragment. The intra-biolithite cavity is filled by dense automicrite bottom sediment (2). Two horizons of open-space structures merge to form part of one cavity (3-4).Detrital internal sediment composed of algal pellets and micrite is distinctly recognizable (3). Most of the calcite cement is of the clear granular spante type (4).

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PLATE VIII

Open-space structure of a Nubrigyn bioherm composed of a rim of brown fibrous orthosparite (1) lining both the micrite framework and the skeletons extending into the cavity. The remainder of the space is occupied mainly by internal allomicrite (2) and clear granular orthosparite (3). The host-rock is composed of algal automicrite (4) rich in detrital skeletal fragments bound by filaments and cells, and some encrusting Foraminifera of the genus Wetheredella (identified by Professor J. H. Johnson).

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PLATE IX

Micrite limestone of an algal bioherm with detrital fragments. The stromatactis open spaces were clearly formed by differential solution of the micrite. The crinoid ossicles ( I ) and the shell (2) remained unaffected. Note the hematite impregnation of the bottom, and the thin film of iron oxide on the exterior of one valve extending into the cavity.

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PLATE X

Detrital skeleton and algal grain limestone of the Nubrigyn Formation, N.S.W., cemented by brown fibrous orthosparite. A cavity left after fibrous sparite precipitation was partly filled by red iron oxide internal sediment and clear granular orthosparite. Note two well-preserved algal stem segments ( I ) . The large skeletal fragment is thinly encrusted by dark algal micrite. The paragenetic sequence is: detritus accumulation-solution(?) cavity-fibrous sparite hematite internal sediment-clear granular sparite. In this case it is not quite clear how the cavity has been formed. If it is of primary origin, it is difficult to see how the grains could have supported the fragments. It may be possible that a slightly cemented calcarenite underwent solution resulting in cavities similar to those shown in Plate XVIII.

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PLATE XI

Nubrigyn algal micrite limestone with some skeleton and algal fragments, and flat-bottomed stromatactis. Note that the bottoms of two large stromatactis have been impregnated by red hematite. Some others have thin films of iron oxide. The cavities have been filled by orthosparite. The vertical fracture is filled by calcite, which is synchronous with that of the stromatactis cavity into which the fracture passes.

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PLATE XI1

Section of a Nubrigyn bioherm with algal filaments ( I ) , recrystallized stromatoporoid crusts (2), and automicrite (3). A flat-bottomed cavity, i.e., a stromatactis (4), is filled by detrital red-brown hematite and granular orthosparite. The cavity shows corrosion features (5). Note that the sparite of the fractures is contemporaneous with that of the cavity above the hematite. The longest dimension of that portion of the sample represented by the thin-section is equal to 8.25 mm.

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PLATE XI11

Open-space structure in an allomicrite limestone of the Nubrigyn Formation. Note the numerous brown fibrous orthosparite generations each separated by a layer of hematitic pellets, or micrite and calcareous pellets, or films of hematite. The remaining space is filled by clear granular orthosparite. The allomicrite host-rock contains numerous algal grains and pellets and skeletal fragments. It seems that the cavity was formed by solution. The paragenesis is: Syngenetic: micrite accumulation. Pre-cementation-diagenetic: solution cavity. Syn-cementation-diagenetic: fibrous calcite and internal sediments. Post-cementation-diagenetic:granular sparite.

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PLATE XIV

Algal micrite limestone of the Nubrigyn Formation, N.S.W., with skeletal and algal fragments and stromatactis. The large open-space structure is lined with numerous generations of brown fibrous orthosparite ( I ) . The dark brown crystalline patch is dolomite (2) formed by internal chemical precipitation. Subsequently, fracturing of the host-rock permitted solutions t o precipitate clear granular orthosparite. The flat-bottomed stromatactis are filled with fibrous, granular, or both types of orthosparite. The paragenesis is: Syngenetic: framework. Pre-cementation-diagenetic: solution cavity. Syn-cementation-diagenetic:fibrous calcite. Post-cementation-diagenetic:dolomite, fracture and granular calcite. There are numerous dark micrite circumcrusts around many crinoid ossicles and other skeletal debris. The Iayers are most likely formed by Algae.

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The flat-bottomed character of these cavities is identical to the so-called stromatactis structures described, for example, by BATHURST (1959a) and OTTEand PARKS (1963). The above discussion suggests that no “soft-body burial” hypothesis is required to explain the genesis of stromatactis, and that these controversial structures are more likely caused by a combination of both syngenetic and diagenetic inorganic processes, and that Algae play only an indirect role. (A “soft-body burial” origin is not completely impossible in the genesis of some other open-space structures, however.) It has been reported that the characteristics of the stromatactis change with location in a reef complex. This is in agreement with KAYE’S (1959) observations of a systematic environmental change of the shape of surface pits before they become part of an internal cavity system. Not all surface pits are incorporated into the sedimentary rocks as open spaces. Under conditions other than those described above, the depressions may be completely filled by detritus, especially if the “eaves” of encrusting Algae are not present. JAANUSSON (1961) mentioned pits in limestone, some of which are flat-bottomed. All are completely occupied by fine-grained detritus forming part of the overlying bed. Interesting to note is the bleaching of one bituminous limestone parallel to the pitted surface. Probably diagenetic oxidation of the bituminous substance resulted in bleaching of the upper part of the sediment. Corrosion, solution and leaching of argillaceous limestone subsequent to its accumulation can form patches, lenses, laminae and beds of marls or clay (LINDSTROM,1963). Removal of calcareous skeletons by solution may leave internal casts if the internal parts of organisms were filled by less soluble or insoluble material. Both direct and indirect organic processes may lead to significant corrosion, solution and disintegration of calcareous sediments. The bacterial processes leading to corrosion and solution of CaC03 have been reviewed by REVELLE and FAIRBRIDGE (1957). As soon as organisms die and become buried, bacteria concentrate to decompose the soft parts. It has been illustrated, for example, that during the decompositional process shells lost 10-24% of their CaC03 in 1-2 months (in one case, 25 % in 2 weeks); and merely traces of insoluble chitinous material remained after complete removal of the carbonate. Etching, corrosion and solution of calcareous material occurs in mangrove environments with high organic content and rapid decay processes. The processes are not well known but it has been suggested that carbonic acid produced by decomposition of organic matter and other acids are the principal agents. Similar changes of calcareous components may take place in freshly accumulated sediments where bacterial oxidation of organic matter produces C02 and lowers the carbonate concentration and pH. Revelle and Fairbridge mentioned estuarine and lagoonal muds in France and Africa with a pH as low as 6.5 and 5, respectively. Particularly in the latter case, carbonate shells were found to dissolve with great rapidity. Both concentration of organic matter and porosity-permeability of the sediment

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influence greatly the amount and rate of CaC03 solution. The smaller the organic matter content and permeability of sediments, the greater is the possibility that shells remain unaffected. Sulfate-reducing bacteria may cause an increase of pH up to 8.5 due to replacement of the strongly acid sulfate radical by sulfide, a weak Bronsted base. CaC03 solution will not take place then, and the skeletons may be preserved. If free iron is not available to react with the sulfide, however, dissolved HzS will diffuse upwards to the surface where it becomes reoxidized to the sulfuric acid and thus reduces the pH causing solution of CaC03 (REVELLE and FAIRBRIDGE, 1957). Pyrite associated with the corrosion horizons may indicate that HzS04 produced by the oxidation of H2S may have been responsible for solution. Corrosion and solution of calcareous particles can occur while sediments pass through the digestive system of organisms. DAPPLES(1942) mentioned that this is indicated by the pH of the fluids of the alimentary tract which may range from 4.75 to 7 before feeding and increases to 7 when the gut is filled with calcareous material. Limestones may be reduced up to 1 inch and more in thickness annually. Dapples gave examples where holothurians dissolved up to 414 g/year. On the Aua reef flat at Samoa, 290,000 individual holothurians destroy 104 tons of sand and lower the entire reef flat 0.2 mm/year. GINSBURG (1957) reported that material larger than sand-size is broken down by boring and burrowing organisms. Worms, molluscs, sponges, and Algae chemically bore into limestones from hightide level to a depth of a few hundred feet below sea level. Differential attack by Algae also has been reported from ancient sediments. In Devonian algal reefs of the Nubrigyn Formation, New South Wales, in particular crinoid ossicles exhibit algal corrosive surfaces (Plate XV, XVI) beneath thin cryptocrystalline calcite circumcrusts, and occasional shells are riddled with bore-holes formed most likely by Algae (Plate XVII) (WOLF,1963a, 1965a). The selected examples of organic corrasion and corrosion indicate that the cumulative diagenetic effects of organisms control to a considerable degree the growth, porosity and permeability of reefs. The constant organic corrosion and abrasion weaken the reef limestones and make them more susceptible to mechanical erosion by waves and currents. Further research on the textures and structures caused by diagenetic organic processes may reveal some useful environmental criteria. Subsurface physicochemical solution and corrosion may be closely related to the water table as indicated by carbonate grain morphology. RUSSELL (1962) has shown one example where the grains of beach sand above the ground-water table are polished and devoid of corrosion features, whereas the grains within the groundwater zone show pitting and various degrees of dissolution caused by the undersaturated fresh water. Precipitation of CaC03 apparently occurs in the vicinity of the water table resulting in coating of the detrital fragments. Selective solution and leaching is widespread in some carbonate rocks.

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PLATE XV

Algal circumcrusted crinoid-sparite-calcarenite of a Devonian reef complex. A number of the crinoid ossicles have been. nearly completely destroyed by algal corrosion (see Plate XVI) as shown by minute specks of crinoid fragments (I) left in some of the dense micrite grains. Once a nucleus has been completely destroyed, the result of corrosion with simultaneous formation of a micrite crust is an algal pellet or grain composed of dense micrite; and resembles algal debris directly derived from abrasion of algal micrite bioherms.

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PLATE XVI

Section of a Nubrigyn calcarenite composed of a greatly enlarged crinoid ossicle (I) with a thick algal micrite circumcrust exhibiting irregular algal corrosion features (2).

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PLATE XVII

Bore-holes (I), most likely of algal origin, in a shell surrounded by allomicrite. The latter contains a small patch of algal filaments (2).

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LUCIA(1962) suggested the possibility that the presence of lime-mud in the detrital rocks studied by him inhibited the genesis of initial calcite cement. The limemud was later selectively leached to form pores. A similar selective matrix removal process has been postulated by GRAFand LAMAR(1950). LUCIA(1962) presented evidence that the best porosity is found in those crinoidal limestones that originally contained 5-10% of micrite matrix. These rocks all had a supporting framework and the lime-mud merely occupied the available intergranular spaces; thus, solutions found access easy. As the lime-mud increases above 20 %, the porosity decreases. Lucia concluded that once the matrix exceeded this value it supported more and more the overburden and this resulted in compaction of the micrite. The sediment became less permeable to the fluids and prevented the removal of the matrix. Selective removal by internal corrosion has also been shown to be widespread in the algal bioherms of the Nubrigyn reef complex (WOLF,1963a). Relatively large proportions of the micrite have been dissolved to leave cavities and tubes. In many cases, Bryozoa and brachiopod fragments extend deeply into the solution voids (Plate VIII, IX). On the other hand, in some rare occurrences the micrite matrix and Bryozoa remained unaffected, whereas gastropod shells and crinoid ossicles were selectively removed leaving external molds. Little has been published on decementation of carbonate rocks, a process suggested for terrigenous rocks by PETTIJOHN (1957). MURRAY (1960) described anhydrite cement which partially replaces both fossils and matrix of limestones. Leaching of the anhydrite left characteristically shaped vugs and increased the porosity of the sediments. If under similar conditions intergranular anhydrite cement is leached, it seems conceivable that a large section of a limestone may undergo decementation. Later, possibly during epigenesis, recementation by calcite would form a limestone without evidence of its previous cementation and decementation history. A case in point might be the reefal limestones of the Devonian Guilmette Formation in parts of western Utah and eastern Nevada of the Great Basin region, which illustrate this process in a superb fashion. Pressure-solution of allochthonous carbonate grain accumulations is possible, especially in cases where the deposit is not exposed to warm, saturated water, and consequently remains uncemented for a relatively long period. It can be demonstrated, for example, that in the Nubrigyn-Tolga reef complex, New South Wales, the shallow-water calcarenites exhibit pressure-solution features in contrast to the graded-bedded basinal deposits (WOLF,1963a, 1965a). This paleoregional change is most likely a function of different degree of saturation and pH at the time of sedimentation: the near-shore waters were more saturated and aerated, caused early cementation and prevented pressure-solution; whereas the basinal waters were undersaturated and reducing, delayed cementation and permitted pressure-solution. The basinal fluids may have migrated upwards and reefwards during compaction and pressure-solution, thus removing the dissolved CaC03 and allowing a continuation of the pressure-solution process.

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G . V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

EMERY et al. (1954) observed in the subsurface limestones at Bikini that coral and mollusk fragments have disintegrated to a chalky powder, and noticed a general disintegration of crystalline calcareous faunal and floral skeletons into microcrystalline and cryptocrystalline material without obvious change in the mineralogy and chemical composition. The original fibrous aragonite crystals of organisms can be seen in thin-sections to have altered to microgranular material. Especially Halimeda segments underwent alterations from a microcrystalline to a denser cryptocrystalline calcium carbonate. Unaffected Halimeda in thin-section exhibit a mat of minute aragonite needles approximately l p in size, which changes into a brown isotropic matrix with scattered needles of high birefringence. A similar alteration occurs in the fibrous calcite tests of small Foraminifera, for example, which change to brown isotropic material. Interstitial microgranular lime-mud shows a similar change to brown marly isotropic micrite. These alterations increase with depth in an irregular to regular fashion. This change of organic carbonate structures to cryptocrystalline micritic material is identical to those observed in Australian algal pisolites and crusts mentioned earlier (Plate I, IV; WOLF, 1963a). HADDING(1958) suggested that bacteria may obliterate or destroy algal components. It is not known, however, to what extent bacteria are responsible for the disintegration mentioned above. It appears, therefore, that crystalline substances can change relatively rapidly into dense calcium carbonate by inorganic and/or organic processes as yet poorly understood. An interesting question arises as to whether inorganically formed aragonite needles can also convert into a brown cryptocrystalline mass or not. If this is so, it may explain the controversy between those who have observed cryptocrystalline crusts surrounding carbonate fragments in beach-rocks and believe them to be of algal origin, versus those who recorded acicular or fibrous aragonite cement and insist that it is of physicochemical origin. If both algal and physicochemical aragonite needles are identical, as shown by LOWENSTAM (1955), and the former can change into a cryptocrystalline mass, then there seems to be a distinct possibility that the latter does the same as suggested previously (WOLF,1963b). Hence, a modification in our concepts of diagenesis, and more refined techniques than thin-section studies only, will be necessary to solve these problems. Our present information, however, suggests that both algal and physicochemical, and possibly other, processes can cement carbonate sediments. Inversion, recrystallization and grain growth Inversion is 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 isotopes) but with a different lattice structure. MAYER (1932) mentioned that some organisms form a gel-like CaC03 which quickly changes into vaterite. The latter is very unstable (TERMIER and TERMIER, 1963), but may remain

DIAGENESIS OF CARBONATE ROCKS

217

unaltered up to almost 1 year before inverting into aragonite. In the sequence gel-vaterite-aragonite-calcite the latter two are the most stable. STEHLIand HOWER(1961) reported that of recent high-Mg calcite, aragonite and low-Mg calcite, the first is very unstable, whereas the other two may persist for a long time under natural conditions. They suggested that the increase of volume during change from aragonite to low-Mg calcite and from high-Mg calcite to low-Mg calcite (the Mg content in the latter two is approximately 8 % and 1 %, respectively) may affect the porosity, cementation and dolomitization of the sediments. Inversion of aragonite may be relatively rapid or slow, i.e., it may occur within 12 months or require tens of thousands of years. LOWENSTAM (1954), for example, mentioned that more than 50 % of the aragonite laid down by some of the marine invertebrates inverted to calcite within 1 year. On the other hand, aragonite has been identified from rocks as old as Late Paleozoic. DEGENS (1 959) believed that aragonite may be found in traces in rocks as old as the Cambrian. TAFT(1963) reported that aragonite and high-Mg calcite of Florida Bay sediments, determined by 1% dating to be 3,600 years old, exhibit no evidence of recrystallization. Taft stated that recrystallization rates appear to be controlled by concentration of a particular cation in the surrounding liquid. As experiments suggest, magnesium chloride solution and Mg in sea water seem to prevent recrystallization; whereas solutions of calcium and strontium chloride, and distilled water, cause recrystallization of aragonite and high-Mg calcite at different rates. Taft suggested, therefore, that marine carbonates tend to remain unstable for long periods until they are exposed to Mg-deficient water. It is quite obvious, then, that the inversion process is both early and late diagenetic as well as epigenetic, and may even be due to burial metamorphism. In general, the exact causes that initiate and perpetuate inversion, recrystallization, and grain growth of limestones are not well known. In particular, calciteto-calcite conversion is largely an unsolved enigma. Numerous parameters have been thought to be conducive to secondary alterations: trace elements, associated organic and inorganic impurities, unstable mineralogic composition, physical and chemical conditions of interstitial ffuids, degree of compaction, degree of solubility, permeability, differential pressure and distortion, availability of nuclei or seeds, temperature variations, and others. Near-recent fossils have been shown to be susceptible to recrystallization in a certain order (CRICKMAY, 1945; EMERY et al., 1954), namely corals, mollusks, Hulimedu, thin-walled pelagic Foraminifera, thick-walled Foraminifera, larger Foraminifera, echinoids, and Lithothamnion. Some evidence suggests that corals may break down into microgranular aragonite before changing into a mosaic of calcite. The segments of Hulimedu are rarely completely recrystallized and are altered first in the central areas (pores) and boundaries. The fibrous calcium carbonate of some organisms changes into coarsely crystalline calcite and may be more or less radially oriented. This indicates that fibrous aragonite may invert to granu-

218

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CHILINGAR, H.

J. BISSELL AND K. H. WOLF

lar calcite and need not form pseudomorphs as suggested by USDOWSKI’S (1962) experiments. It is interesting to note here that the recrystallization or inversion to coarser calcite is contrary to the “disintegration” into cryptocrystalline material mentioned earlier. The term recrystallization has been used loosely for a number of processes that commonly cause a change in crystal or grain size, predominantly an enlargement and occasionally a reduction in size, without causing a chemical alteration except for changes in isotope and trace element concentrations. Some prefer to include inversion and grain growth, whereas others prefer a restricted use of the term recrystallization. BATHURST (1958) stated that “grain growth s. str. is nowadays distinguished from primary recrystallization which may precede it and from secondary recrystallization which may follow it. Grain growth acts in monomineralic fabrics 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 and a fine mosaic is gradually replaced by a coarser. As grain growth proceeds, many of the enlarged grains are themselves replaced by their more successful neighbors.’’ Recrystallization, as defined by Bathurst occurs when “nuclei of new unstrained grains appear in or near the boundaries of the old, strained grains. These nuclei grow until the old mosaic has been wholly replaced by a new, relatively strain-free mosaic with a nearly uniform grain size. Its coarseness depends on the density of the initial nucleation. Where the nuclei are widely spaced there is an intermediate porphyroblastic stage.” Grain growth and recrystallization should be accepted as two distinct processes wherever possible. BATHURST (1958) considered the following processes that may cause grain enlargement: solution of supersoluble small grains with redeposition on larger grains (aragonite is 3-9 % more soluble than calcite, CHILINGAR, 1956c), solution transfer, primary recrystallization, inversion of aragonite to calcite, and grain growth s. str. It is possible, however, that at least one of these can cause a relative decrease in crystal size. If recrystallization of a coarse crinoid ossicles accumulation occurs, starting with nucleation in the interior of the ossicles, minute calcite crystals will replace each larger crinoid crystal and may spread to consume the whole limestone. WARDLAW(1962) suggested, therefore, that under favorable conditions a calcarenite may be converted into a limestone composed of silt-sized calcite crystals, i.e., microsparite. BATHURST (1958) mentioned syntaxial replacement rims which are similar in appearance to those formed by what he termed syntaxial rim cementation. For example, crinoid ossicles in contact with lime-mud may undergo grain growth at the expense of the fine matrix. The result is a calcite rim in optical continuity with the ossicles. Calcite deposited from interstitial solutions onto free surfaces of crinoids may have the same result. The two processes, however, are very different. A similar syntaxial replacement phenomenon has been described by FOLK

DIAGENESIS OF CARBONATE ROCKS

219

(1962a). He mentioned an oolite with an echinoderm fragment as nucleus, and rays of sparry calcite in optical continuity with the echinoderm. To make a clear distinction between the so-called open-space calcium carbonate precipitated in voids on one hand and that formed by inversion, recrystallization, and grain growth on the other, the latter have been named pseudosparite by FOLK(1959), and the former orthosparite by WOLF(1963b). The term “sparite” is purely descriptive, therefore. The processes that cause crystal enlargement have been collectively called “aggrading recrystallization” by FOLK(1956; 1959); and those that cause a decrease in crystal size were named “degrading recrystallization” by FOLK (1956), “degenerative recrystallization” by Dunham (in FOLK, 1956), and “grain diminution” by ORMEand BROWN(1963). The disintegration change of crystalline coral and algal material to brown cryptocrystalline calcium carbonate described in the foregoing sections may be considered as “grain diminution” although the actual causal factors are not known. Inversion, recrystallization and grain growth vary not only in sign and extent but also in position and resultant grain or crystal morphology, as indicated by FOLK(1956). The synthesis below is based on Folk’s work and is presented here with slight alterations, with his permission: Sign: (I) marked increase in crystal size, (2) marked decrease in crystal size, and (3) no or very little change in crystal size (e.g., formation of pseudomorphs). The extent and position of the processes discussed here can be divided into phases a, 8, and so forth. This is useful as it may save time and eliminate repetitious descriptions in preparing logs and reports, for example. It is understood, of course, that all phases are completely gradational. a phase. Limestone is unaffected by inversion, recrystallization, or grain growth. /3 phase. Limestone is slightly affected. A few of the allochemical grains and possibly small portions of the micrite matrix or sparite cement is “recrystallized”. Inversions of originally aragonite fossils, e.g., most pelecypods, many gastropods, some Algae, to calcite have occurred. y phase. The limestone has undergone major alteration but the original nature of the matrix is still discernible, and the rock can still be described and classified according to Folk’s scheme and the modified version given in this chapter. The allochems are still recognizable and may range from unaltered to completely recrystallized. This phase passes into the next phase when the original matrix is completely recrystallized. 6 phase. Limestone is extensively altered. Inversion, recrystallization and/or grain growth of an original cryptocrystalline to microcrystalline calcite or aragonite matrix and cement resulted in microsparite. This is probably the most common process according to Folk. It agrees with BATHURST’S (1958, 1959b) idea of the existence of a “universal threshold state at which fabric evolution stops and beyond which it can, but need not, continue”. The microspar consists of calcite crystals

220

G. V. CHILINGAR, H. I. BISSELL AND K . H. WOLF

5-20,u in diameter in contrast to the original micrite size of 1-5p. The fragments, i.e., allochemical grains, or colonial growths in autochthonous limestones, remain largely unaffected in this phase. In handspecimens it is impossible to distinguish micrite from microsparite, but in thin-section they are easily discriminated. If recrystallization of the matrix is incomplete, patches of the matrix may “float” in the microsparite. The contact between micrite and microsparite may be very gradual or sharp. Sometimes the contact is oblique or vertical to bedding planes. Hence, the microspar is not merely a coarser crystalline detrital material as pointed out by Folk. If the alteration of the matrix is complete, the detrital grains may “float” in the microsparite. Open-space sparite may be more resistant to changes in contrast to micrite matrix and clear calcite cement may be found in a microsparite due to preferential alterations (WOLF,1963a). E phase. Alterations affected matrix, cement and grains or framework of the limestone. The patches formed by recrystallization and/or grain growth may be very irregularly shaped, may occur as “fronts”, veins, or transgress the whole rock. Faint relics (“ghosts”) of grains are still recognizable. The criteria for partial recrystallization are the same as those employed to determine other replacement phenomena. phase. Limestone alteration is complete. None of the original textures and structures of allochemical grains or colonial growths, matrix and cement are recognizable. The rock is composed of microsparite or sparite only. The genetic nomenclature of different types of micrites and sparites is presented below (WOLF, 1963b). FOLK (1956) defended the viewpoint that the products listed in phases E and 5 may be common locally, but their overall volumetric significance is small. Phases a-y are the more common ones. It must be noted that the above phases are based on the assumption that there is an increase in crystal size during alteration. Although this seems to be the case in the majority of recrystallization and grain growth occurrences one should not loose sight of the “degradation recrystallization” and disintegration possibilities mentioned earlier. It is not known how significant they have been in the geologic past, but it seems possible that many of the problematic micrite knoll-reefs may have been formed by grain diminution of algal and faunal colonies. The morphology of the crystals or grains after inversion, recrystallization and/or grain growth may be granular, drusy, fibrous and/or bladed. USDOWSKI’S (1962) experiments indicate that inversion of aragonite may result in the formation of pseudomorphs and does not necessarily destroy the original fibrous nature. The small volume changes are apparently insignificant in obliterating the original texture. Hence, it seems reasonable to conclude that any primary textures and structures of aragonite, whether granular, drusy, fibrous, oolitic or spherulitic, may be preserved as suggested in Table I. On the other hand, it has been observed that during inversion a change of crystal morphology can take place, and the likely

221

DIAGENESIS OF CARBONATE ROCKS

TABLE I POSSIBLE GRAIN OR CRYSTAL MORPHOLOGY CHANGES DURING INVERSION, RECRYSTALLIZATION AND GRAIN GROWTH

(After WOLF,1963b) Inversion ~~

~-

-

originally -

finally

_ _

~~

Granular aragonite Drusy aragonite Fibrous aragonite

-~ -

~

~~

~

Granular aragonite Drusy aragonite Fibrous aragonite

~~

-+

Granular calcite pseudomorphsl --f Drusy calcite pseudomorphsl + Fibrous calcite pseudomorphsl Granular calcite Fibrous calcite (?)

Recrystallization and grain growth ~~

~~

originally

finally

Granular aragonite or calcite Drusy aragonite or calcite Fibrous aragonite or calcite

Granular calcite Fibrous calcite Bladed calcite

~

~~~

_

~

_

_

_-

1 These types of pseudomorphs are usually referred to as paramorphs because no change in composition occurs during inversion (except for possible changes in trace element and isotope contents). *2 Referring to possible processes that cause change of drusy aragonite to either granular or fibrous calcite; fibrous to either granular or blady calcite; and so on.

possibilities have been given in Table I. Changes in textures attributed to volume increase during inversion have been reported by BATHURST(1959b). More research is required on these aspects. Recrystallization and grain growth have been shown to form granular, fibrous (FOLK,1962a; ORME and BROWN,1963) and blady (HARBAUGH, 1961) calcite crystals or grains. Hence, the latter two forms are not always indicative of openspace formation. The powerful displacing ability and forces of both open-space and replacement fibrous calcite have been demonstrated by FOLK(1962a). He described fibrous calcite overgrowths in optical continuity with an ostracod shell. The spar began as open-space calcite growth which completely filled the cavity, but then appears to have continued to enlarge and force the shell apart. In other cases, articulate ostracods were originally filled by clay. Fibrous calcite overgrowths then grew inward from the upper and lower valves into the former body cavity forcing the clay to the center by the pressure of crystallization. The fan-like overgrowth of

222

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

fibrous sparite caused in some cases an expansion of the sediment to two to three times its original volume. The spar must have crystallized when the sediment was at sea-floor level or upon very slight burial to permit such expansion. That some of the fibrous calcite is definitely of the replacement rather than the open-space type is further indicated by cases where fibers grew preferentially downward from both the upper and lower valves of fossils. The lower ones definitely replaced host-rock matrix material. FOLK(1962a) mentioned unaltered micrite intraclasts within a microsparite “matrix”. He suggested that the differential recrystallization was due to differences in compaction and permeability. The micrite limestone fragments, which were derived from some locality within the depositional environment, were less permeable as they had undergone some compaction before dislocation and transportation. Jn the new area of deposition these clasts were mixed with a less compacted micrite matrix. The latter was more permeable to solutions, and possibly had other features conducive to recrystallization. On the other hand, faecal pellets, possibly because of presence of organic matter, have recrystallized at the same rate and/or extent as the matrix and the end-product is a fairly uniform microsparite with “ghosts” of pellets outlined by organic specks and other impurities. The degree of inversion and recrystallization may change on a regional scale. NEWELLet al. (1953) mentioned, for example, a regional trend in degree of recrystallization: the basinal sediments are least affected, whereas the reef and lagoonal limestones are most extensively recrystallized. The fore-reef talus deposits take an intermediate position and exhibit slight recrystallization near the basin and grade into more altered limestones close to the reef. Hence, the fossils are generally better preserved near the basin. Somewhat similar conditions typify some of the Pennsylvanian and Permian limestones (Callville and Pakoon) of the shelf facies near the shelf-to-basin transition along the Las Vegas Line of southern Nevada (BISSELL, 1959). These micrites and algal to pelletal limestones that have a fine-textured matrix have been recrystallized (and dolomitized) on a much greater regional scale than have their basinal equivalents (Bird Spring and Spring Mountains Formations). Furthermore, the fossils (and in particular fusulinids) are better preserved in the basinal sediments. Some recent sediments show evidence of extensive recrystallization, whereas others lack it. The causes are poorly known. Inversion seems to be more rapid in subaerial environments and in zones of meteoric waters. According to FOLK’S (1962a) experience, it seems that brackish-water micrites recrystallize more readily than either normal marine or lacustrine fresh-water micrites. STEHLIand HoWER (1961) indicated that minerals of shallow and deep-water environments are different and their diagenesis susceptibility must differ accordingly. In some recent shallow-water areas at least 70% of the carbonate consists of metastable CaC03: aragonite and high-Mg calcite. The deep-water sediments appear to be composed of low-Mg calcite and are, therefore, more stable. Thus, diagenesis should in

DIAGENESIS OF CARBONATE ROCKS

223

general be initiated earlier and be more pronounced in shallow-water type carbonates. Changes in contents of trace elements and isotopes are commonly associated with inversion, recrystallization and possibly grain growth. These processes are accompanied by expulsion of trace elements (e.g., Mg2+, Sr2+,Mn2+, Ba2+)to the interstitial fluids, matrix or cement, where they are available for other diagenetic processes either in the immediate vicinity or at remote localities. As high-Mg calcite is the least stable among the aragonite and calcite sediments, it seems that Mg is one of the earliest elements available. Quantitatively, it is also more significant than the other elements and this Mg may form, therefore, the raw material for early diagenetic dolomitization. In the sediments examined by STEHLIand HOWER(1961) the elements present in the calcium carbonate lattice are in the following order: Mg2+> Sr2+>Mn2+> Ba2+. In every case diagenetic alterations appear to have resulted in a marked decrease in trace element concentration. SIEGEL(1960) stated that whenever the Sr/Ca ratios of Recent corals are compared, it becomes clear that when Sr is present in amounts that indicate that none or only very little of it has been lost from the original aragonite, the inversion to calcite has hardly begun. On the contrary, however, where the amount of Sr has been appreciably reduced, the alteration from aragonite to calcite has reached a point that appears to be directly related to the degree of Sr removal. Siegel suggested, therefore, that the presence of Sr, not as SrC03 but rather in substitution for Ca in the aragonite lattice, inhibits and, therefore, prevents or slows the rate of inversion under natural conditions. Siegel further proposed that the inversion occurs only when much of the Sr has been removed. Many scientists, however, maintain that inversion merely expels Sr, i.e., loss of Sr is an effect and not a cause of inversion. LOWENSTAM (1954) reported that a distinct decrease of the Sr/Ca ratio occurs during recrystallization. For example, unaltered corals with about 1.4 % strontium carbonate recrystallize to microgranular calcite having about 0.7 %, and to more coarsely crystalline calcite containing 0.2 % strontium carbonate. USDOWSKI (1962) advanced an interesting theory to support the idea of inversion of aragonite oolites to calcite. This author assumed that the composition of the water medium remained the same from the time of oolite formation until cement precipitation. Therefore, both must have had the same trace element composition. Analyses indicate, however, that the cement has a much higher content of Mg, Fe, Mn, and Sr; Usdowski suggested that during inversion the oolites expelled the trace elements, which were either incorporated into the calcite cement or were removed by interstitial solutions. In a subsequent publication on early diagenetic cone-in-cone, structures USDOWSKI (1963) supported his theory of trace element expulsion during recrystallization. The limestone beds which underwent recrystallization, resulting in cone-in-cone features, have an Sr content of 247 p.p. m. The unrecrystallized beds are richer by a factor of 0.4. A similar relationship

224

G . V. CHILINGAR, H. J. BISSELL AND K . H. WOLF

exists for the Mg content which is 0.7 % and 4.7 % for recrystallized and unaltered limestones, respectively. DEGENS’ (1959) studies show that recent fresh-water limestones havz a lower content of strontium than marine carbonate sediments, caused presumably by the lower amounts of Sr in fresh water. With an increase in age of limestones, however, the difference in Ca/Sr ratio between the fresh- and marine-water sediments appears to diminish; and Paleozoic carbonates, independent of facies, do not deviate much from the average value of 500 p.p.m. of Sr. Hence, fresh-water limestones must have gained and marine limestones lost Sr during the diagenetic-epigenetic geologic history. Ross and OANA(1961) concluded that both the environment of deposition and the diagenetic history of limestones determine the carbon-isotope distribution in carbonates. Their work indicates that biosparites (terminology of FOLK,1959) with a large amount of sparry calcite cement have 613C values between +1.0 and -1.0. On the other hand, biomicrites or biomicrosparites have either distinctly positive or negative d13C values. As the amount of sparry calcite decreases, the S13C value becomes either more positive or negative. This suggested to the two authors that limestones which underwent little recrystallization have a wider spread of 613C values than do those which exhibit evidence of considerable recrystallization or introduction of calcite cement. The effect of recrystallization is to shift the 613C values toward the range of 1 to - 1. More research on different types of grains, micrites of diverse origins, and sparry calcites formed by different processes is necessary, however, in order to check the validity of these interpretations.

+

Internaljlling and internal sedimentation

Internal filling and sedimentation processes of physicochemical, biochemical, and physical nature cause partial to complete filling of voids within sedimentary frameworks and form the so-called open-space structures (Plate VI-XIV; Fig. 1 ; WOLF, 196%). Many of the cavities that are formed diagenetically are interconnected and give the sediments a very high degree of primary permeability. Most of the larger systems are open at some points to the upper surface and tidal waters can penetrate the system t o deposit detritus (Fig.1, top). The same fluids may also chemically precipitate a number of substances or cause wall-rock alterations as, for example, in the Nubrigyn algal reef complex (Table 11; WOLF,1963a). These components form a complex paragenesis due to cyclic deposition. The detrital internal sediments form minute lenses, patches and thin layers in the voids, and smooth out irregularities of the floor of the cavity (Plate VII). In thin-section, the internal sediments are either dense and structureless, or are laminated and graded on a microscopic scale. In most cases, the internal sediments are different both in texture and/or composition from the host-rock. In some occurrences, however,

225

DIAGENESIS OF CARBONATE ROCKS

. sediment . .

'.

,

.

.

*

. , .

.

'

e of such open-space structures varies from m i c r o - t o mesoscopic i n scale(incks to feet).

'

Three generations of internal sediments. TWO generat ions of brown fibrous orthosparite. Colourless granular orthosparite in cavity and fracture. X Internal sediment slipped 'into fracture.

. '

_

Colourless granular. Coarsely crystalline dolomite cavity filling.

. : . . .' . . .. . . '

.

..

,

.

.-==-==T,, .:. . ,

. . . .

.

.

Brown fibrous orthosparite (numerous generations ).

.

. .

. .- .- .- . . . .

. . . ., _ .

iron oxide replacing framework ide "fronts" mit of oxide "halo"

Fig.1 . Open-space structures in Devonian algal bioherms, Nubrigyn Formation, N.S.W. (After WOLF, 1963a.)

they blend. Occasionally, cavities are completely filled by internal sediments; but in the majority of internal sediments they are confined to the lower part of the voids and it is the subsequent deposition of clear colorless sparite that filled the upper spaces. Numerous cavities show wall-rock alterations prior to internal sedimentation and calcite cement precipitation (Plate IX, XI). Either iron oxide replacement, or leaching and bleaching, occurred in a semiconcentric fashion around the voids or was limited to the area near the floor of the openings. Some of the lower parts of the cavities were differentially leached, corroded and oxidized, resulting in red iron oxide rich pellets that are easily mistaken for detrital internal sediments.

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

TABLE 11 DIAGENETIC MODIFICATIONS I N DEVOMAN ALGAL BIOHERMS, NEW SOUTH WALES

Detrita! internal sediments

Chemical internal fillings

Wall-rock alterations

Lime-mud

Fibrous calcite

Leaching

Pellets

Drusy calcite

Bleaching

Fine algal and skeletal debris

Granular calcite

Solution

Iron oxide

Iron oxide

Iron oxide replacement (irregular and as “fronts”)

Clay

Dolomite

They were formed in situ, however, and constitute a residual product on a microscopic scale. On the other hand, in most cases the internal open-space iron oxide was directly precipitated from solution and/or mechanically deposited. Although it may have originated at the same time as the iron oxide replacing the wall-rock, it is of a different origin. In numerous occurrences, fibrous calcite precipitation encrusted the walls of the cavities before the internal sediments, iron oxide, dolomite, and/or granular sparite were deposited (Plate VIII, XIII, XIV; Fig.1). In addition to the above-mentioned cavities, minor open-space structures beneath large faunal fragments are common. They usually lack a complex paragenetic history, however, and are only filled by internal sediments, fibrous and/or granular sparite. It seems that they were isolated and out of reach of oxide- and dolomite-precipitating solutions.

Morphologic and genetic calcium carbonate types

As illustrated in Table I, the three basic morphologic types of calcium carbonate can be formed by a number of primary and/or secondary processes. Recent research in carbonate petrology has resulted in valuable information that permits the discrimination of the numerous aragonite and calcite types formed by open-space precipitation, recrystallization and grain growth, Hence, it is possible to present in Table I11 a scheme that attempts to cover all likely occurrences ranging from a simple descriptive to a more complex genetic nomenclature of cryptocrystalline to coarsely crystalline carbonate. Fig.:! gives a diagrammatic illustration of the numerous possible fabrics or textures. All have been listed also in Table I11 except for the two types of syntaxial rims. They are either of open-space or grain growth origin and need no special pigeon-hole. It should be emphasized that the writers feel

227

DIAGENESIS OF CARBONATE ROCKS

as strongly as others who oppose the “game” of semantics. Analogous to nuclear physics, however, the more we learn about the minute, often subtle, and yet important differences, the more terms will be required for precise, unambiguous communications, and in order to eliminate long repetitive descriptions. Also, both descriptive and genetic types of terms are necessary if confusion is to be avoided. The former can be easily changed to the latter by merely adding prefixes. The following criteria and discussions represent a modified summary of the works of BATHURST(1958, 1959b), HARBAUCH(1961), FOLK(1962a), ORME and BROWN(1963), and WOLF(1963a).

Granular, drusy, and fibrous open-space calcium carbonate The following features may be characteristic (4 and 5 may also occur in graingrowth products): ( I ) The crystals or grains are in contact with surfaces that were once free, i.e., surfaces of voids. The contact may be horizontal, vertical or oblique, and GENESIS

recrystallization)

I

FABRIC

Micri te (Pseudo-)

-organic

structure

Fig.2. Diagenetic fabrics (WOLF,1963b). Modified after ORMEand BROWN(1963).

I11 DESCRIPTIVE AND GENETIC NOMENCLATURE FOR MICRITE-SPARITE RANGE OF ARAGONITE AND CALCITE

WOLF,1963b) Descriptive indicating ,crystal size

Microsparite

(often called calcilutite, ooze, lime-mud) Cryptocrystalline7

Footnotes see p.229.

Genetic descriptive

Origin

Open-space precipitation, i.e., void fillings

approx. size

indicating crystal morphology and size

appron. proportions

> 0.02 mm

Granular sparite

Equidimensional

Orthosparite

Drusy sparite (size and morphology change distally)

Elongate

Pseudosparitel

Fibrous sparite6

Elongate

Granular microsparite

Equidimensional

Orthomicrosparite

Drusy microsparite4

Elongate

Pseudomicrosparitel

Fibrous microsparite6

Elongate

0.005-0.02 mm

< 0.005 mm

Too small to observe visually morphologic differences except by use of electronmicroscope

e

.-* .-w

Recrystallization grain growth (BATHURST,1958)

c

+-d

2

Open-space precipitation

29

k s Recrystallization grain growth

0

8.9

8

2

& w

Degradation recrystallization (= grain crystal diminution) ~

Orthomicrite3.5

“Genuine primary” micrite5

Allomicrite3

-d

Automicrites

a 9 .E

.*

(I)

s

.5

Allochthonous micrite Autochthonous

DIAGENESIS OF CARBONATE ROCKS

229

crystal growth may occur preferentially upward, downward, or in any other direction (Plate VIII, X, XIII, XIV, XVIII, XIX; Fig.1). (2) If the cavity is not completely filled, the remaining space may be filled by succeeding generations of the same calcite type, or any of the other two types, or by detrital or chemical internal sediment, or any combination of the above. Euhedral terminations extending into the voids are frequent. (3) Some of the cavities underwent pre-cement modifications: the host-rock may have been slightly replaced or was leached, bleached, corroded; or internal detrital bottom sediments smoothed out irregularities before precipitation of cement (Plate VI, VII, IX, XI; Fig. I). ( 4 ) There is usually an abrupt contact between calcite mosaic and host-rock. (5) The mosaic-filled region has the obvious form of a cavity, but may be very complex in shape, or too laqge to be recognized in one thin-section. (6) The intergranular boundaries of the mosaic are usually planar (Plate XVIII, XIX; Fig. 1). (7) In many cases there is an increase in grain or crystal size of the mosaic away from the wall: this is the so-called drusy carbonate (Plate XVIII). In other cases, fibrous calcite of uniform length forms a relatively wide crust on the surfaces of open spaces, and no changes of crystal size need occur (Plate XIV, XIX). Similarly, granular sparite may fill open spaces without systematic grain size change. (8) Drusy and fibrous calcite show a preferred orientation of the longest grain axis normal to the surface of the host-particle (Plate XVIII, XIX). (9) Drusy and fibrous calcite grains and crystals are preferentially oriented with the optical axes normal to the surface. Occasionally, it may be a type of overgrowth, e.g., fibrous calcite on a shell in optical continuity with the shell’s surface. (f0)Most commonly the drusy and granular calcite is clear and colorless. The fibrous calcite, however, has been frequently reported as light brown in color (Plate XVIII, XIX). (I I ) In many cases, early diagenetically cemented limestones are reworked by intraformational processes. In the Nubrigyn Formation, N.S.W., fibrous sparite is present as intraclasts, indicating that it is of very early diagenetic origin. It is uncertain at the present time whether recrystallization or grain growth can form a drusy fabric. Recent algal colonies appear to change to cryptocrystalline material by an unknown diagenetic process (WOLF,1965a, b). Commonly called “matrix” in contrast to sparite cement. Grades rapidly into sparite size-grade. Orthomicrite is a collective term for unaltered micrite and includes both allo- and automicrite. Fibrous sparite has been called “drusy” by mistake. Fibrous carbonate consists of acicular needles of roughly uniform length. In contrast, drusy carbonate changes from small blady or acicular grains or crystals to larger ones toward the center of the cavity. This change in size is accompanied by a gradual change of morphology to equidimensional (granular) sparite (BATHURST, 1958). Micrite when not resolvable by a petrographic microscope is cryptocrystalline in appearance.

230

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

PLATE XVIII

Skeleton and algal grain-orthosparitexalcarenite from a cross-bedded eolianite, Lord Howe Island, Australia (specimen collected by Mr. J. Standard). Note the well-developed drusy orthosparite, filling solution channels. The drusy carbonate is in contrast to the acicular (= fibrous) sparite of beach-rocks. The large cavity was formed by subaerial solution of a slightly cemented calcarenite. The paragenesis is: Syngenetic: calcarenite accumulation. Syn-cementation-diagenetic:thin film of cement. origin Post-cementation-diagenetic:solution cavity and drusy cement.

DIAGENESIS OF CARBONATE ROCKS

23 1

PLATE XIX

Two crinoid ossicles ( I ) , circumcrusted by dense algal micrite (2), are surrounded by brown fibrous orthosparite(3). The central void is occupied by clear granular sparite (4). The fibrous morphology has been slightly obliterated, most likely as a result of diagenesis.

232

G . V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

PLATE XX

Nubrigyn calcarenite, N.S.W., composed of skeletal fragments and algal grains ( I ) surrounded by pseudosparite. A crinoid ossicle circumcrusted with micrite exhibits syntaxial rim cement (2).

DIAGENESIS OF CARBONATE ROCKS

233

Syntaxid rim cementation Syntaxial rim cementation is characterized by some or all of the following: ( I ) A detrital core is present. It is usually a single crystal, most commonly a crinoid ossicle, but other fossils have served as nuclei. (2) The core may be recognized by its inclusions or outer rim of impurities, in contrast to the clear outer overgrowth. In some cases, relatively thick algal circumcrusts around the ossicles did not prevent syntaxial growth (Plate XX; WOLF, 1963a). (3) Host and rim are syntaxial, i.e., in optical continuity. LUCIA(1962) stated that rim cement grows on single-crystal fragments, such as crinoid ossicles, and multicrystalline (dog-tooth) cement grows on multicrystalline hosts. There is little doubt that this is possible. But it must be remembered that not all calcite cement deposited on uni- or multicrystalline components forms overgrowths or is in optical continuity. In the Nubrigyn Formation, N. S. W., much of the crinoid debris is cemented by open-space fibrous calcite that is not in optical continuity with it. It seems then, that more than the mere presence of suitable nuclei controls the genesis of syntaxial rims. ( 4 ) The outer boundaries of the rims are mostly in contact either with other rims or granular cement or with detrital particles, but seldom with a micrite matrix. According to BATHURST (1958), contacts with a micrite matrix are typical of syntaxial rims formed by grain growth or replacement. The presence of lime-mud, however, need not exclude the possibility of open-space syntaxial deposition. For example, a crinoid ossicle calcarenite overlain by lime-mud has been shown to develop syntaxial rim cement preferentially downward, because the lime-mud prevented overgrowth along the upper ossicle-lime-mud contact. Also, as FOLK (1962b) has shown, a matrix saturated with fluids may cause formation of fibrous calcite which, due to the crystallization force, may merely force the matrix aside without actually replacing it. (5) Boundaries between the rim and adjacent cement are planar. (6) Host grains are in contact with each other in three dimensions. (7) The mosaic resulting from overgrowth has plane intergranularboundaries. (8) The mosaic may be rather equidimensional resulting from the syntaxial growth on well-sorted detrital particles, which is in contrast to the rather more heterogeneous grain-size pattern of granular cement and grain growth on recrystallization mosaics. (9) The mosaic may be arranged in layers which differ in composition and coarseness. (10) The mosaic may contain patches of skeletal fragments, pellets, oolites, and so on, with components similar to those of the mosaic grains. ( I ] ) The longer axes of the detrital particles plus their syntaxial rim may be arranged sub-parallel to the original substratum.

234

G. V. CHILINGAR, H. J . BISSELL AND K. H. WOLF

Grain growth The calcium carbonate formed by grain growth has the following features: ( I ) The grain or crystal diameter ranges upwards from about 5p. Diameters between 50 and loop are common and larger grains occur. The coarse mosaic has often been confused with granular open-space cement (Plate XXI). (2) The contact between fine and coarse mosaic may be abrupt; it can also be very gradual and the intermingling of fine and coarse grains makes it difficult to draw a definite boundary (Plate XXI). (3) Grain growth may be very selective or preferential (Plate XXI). ( 4 ) The grain size in the coarse mosaic varies irregularly and may change from place to place even over distances of 0.5 mm. This irregular pattern of grain size is distinct from the vectorial variation in drusy and fibrous mosaics and from the well-sorted mosaics of rim-cemented detritus. Porphyroblasts are possible. Grain growth can also form rather equidimensional grain mosaics. (5) Boundaries between grain growth and unaffected material may cut depositional features, e.g., laminations. (6) Grain boundaries in the coarse mosaic vary generally from curved to consertal. Implicate boundaries appear among the larger grains; and the plane boundaries so typical of open-space sparite occur less frequently. (7) Some large marginal grains in the coarse mosaic embay the adjacent fine mosaic causing it to have a “nibbed” appearance. Many embayments are plane sided. Once convex curved boundaries, e.g., of pellets, are now locally concave. Fine-grained mosaic may occur as wisps or threads in the coarse mosaic. Fossils may be extensively interrupted to leave only disconnected relics of the original skeletons. In addition, pseudo-breccias may form (BATHURST, 1959b). (8) Some detrital components, e.g., patches of lime-mud, oolites, pellets, sparite cement, are entirely surrounded in three dimensions by the grain growth mosaic (Plate XXI). (9) Although well and extensively developed drusy and fibrous carbonate is usually indicative of open-space precipitation, both fibrous and blady calcite may form by grain growth (HARBAUGH, 1961; ORMEand BROWN,1963). (10) Grain growth may occur under favorable conditions unidirectional. (11) Grain growth may cut textures formed during preceding generations of diagenesis. (12) Presence of impurities between crystals is common. (13) Spherulites may be formed by grain growth (Plate XXII). (14) Patches of pre-grain growth open-space sparite cement remain often unaffected as they are more stable (Plate XXI).

DIAGENESIS OF CARBONATE ROCKS

235

PLATE XXI

Incipient recrystallization product of a Nubrigyn algal bioherm. The micrite matrix has been preferentially recrystallized to pseudomicrosparite (I), which surrounds unaffected open-space granular sparite (2), and a recrystallized coral fragment (3). The fact that orthosparite (2) was only slightly affected by recrystallization indicates that it is more stable than micrite. The paragenesis is: Syngenetic: algal automicrite framework with detrital coral fragment. Syn-cementational-diagenetic: orthosparite as “birdseye” patches. Epigenetic(?): preferential recrystallization.

236

G . V. CHILINGAR, H. I. BISSELL AND K. H. WOLF

PLATE XXII

Recrystallized Tertiary(?) limestone of Portuguese Timor composed of pseudomicrosparite and pseudo-spherulites formed by grain growth. Note that the latter are quite distinct from the algal genus Calcisphaera.' The pseudo-spherulites are composed of sparite formed by grain growth or recrystallization, whereas the latter are open-space structures filled by micro-drusy calcite in the Nubrigyn Formation. The longest dimension of that portion of the sample represented by the thin-section is equal to 1.3 mm.

DIAGENESIS OF CARBONATE ROCKS

237

Syntaxial grain-growth rims Syntaxial grain-growth rims usually have some or all of the following distinct features: ( I ) The syntaxial grain-growth rim resembles superficially a cement rim but otherwise has quite different fabric relations. The host is most commonly a crinoid ossicle. ‘ (2) The rim, or the hosts where no rimming occurred, is in contact with a matrix of lime-mud (unlike a cement rim) or with other rims, or other detrital particles. (3) The adjacent lime-mud matrix may include detrital particles. It is this micrite which is interrupted by the rims. ( 4 ) A rim may interrupt the fabric of a skeleton or embay the surface of pellets. (5) Unlike the open-space syntaxial rim, the grain-growth rim has a highly irregular outer boundary, part of which may be produced into small spires often plane sided. These may be relatively wide, e.g., lS30,u in examples reported by BATHURST (1958), and taper generally distally to a point. Other extensions of the rim taper proximally, having swollen distal ends. (6) This kind of syntaxial rim is commonly associated with the coarse graingrowth mosaic, and genuine open-space sparite cement may be absent from the rock. (7) The nuclei that underwent syntaxial enlargement may “float” in the spar. Grain diminution Calcium carbonate can “recrystallize” at low temperatures and low pressures resulting in a relative decrease in crystal or grain size to form features such as the ones listed below: ( I ) Small grains or crystals replacing coarse material of crinoids, Bryozoa, and algal colonies (Plate 11-IV). This process may explain the controversial knollreefs in various parts of the world, which are composed of micrite and finegrained particles. If preferential grain diminution of an algal bioherm framework occurs and internal sediments and calcite cement remain unaffected, the resultant limestone is composed of dense material (lacking any evidence of previous organisms) and some patches of open-space fillings (WOLF,1963a, b). (2) Both granular and fibrous grains or crystals can form. (3) Patches of fine grains may be irregularly distributed. ( 4 ) Selective replacement may be common. (5) The mosaic formed by grain diminution may increase by grain growth to form a coarser mosaic. If this occurs, many of the features associated with grain growth are applicable here also. Non-calcareous replacement

Non-calcareous replacement or substitution of one mineral by another of differ-

23 8

G. V. CHILINGAR, H. J. BISSELL A N D K. H. WOLF

ent composition in limestones may be ( I ) absent, (2) very local, (3) regional, ( 4 ) partial, (5) preferential, and (6) complete. Dolomite replacement, which is the most common of all, is treated in the second part of this chapter. As other replacement phenomena are considered in detail by other authors in this book, a few

HUNDREPS

TENS

ONES

F

in

K

0 I0

2 u)

3

0 a

3 > m

n W

z

TEN THOUSPNDS

THOUSPNDS

I K

W

HUNDREDS

IW 0

TENS

u)

a

ONES

I>

2 m

_I 3

in 0

TMDUSbNDS1

~

1

1 /

.D

HUNDREDS u)

W z 0

a I 0

Fig.3. Solubility of the most important chemical components of sediments (A); their type of solution @); their general variability depending on physicochemical conditions of the depositional medium (C); and the specific factors pH@), Eh (E), and COZcontent (F). (After RUKHIN,1961, p.275, 211.)

239

DIAGENESIS OF CARBONATE ROCKS

remarks on regional differentiation and local conditions that lead to silica, iron oxide and pyrite formation in limestones will suffice here. The parameters responsible for non-carbonate replacements are similar to those listed for other diagenetic processes but some have a more intense significance. For example, in the monomineralic limestone all or most of the raw material for diagenesis may be of endogenic origin. The components necessary for non-carbonate replacement, however, must have had an exogenic source in many cases. The iron and silica for extensive limestone replacements must have come from an outside source; this could have been from the continent or from intra-basin highs such as volcanic archipelagos (BISSELL,1959). Suitable climatic and geomorphic conditions are a prerequisite to assure a supply from an outside source. As indicated in Fig.3, the components in solution react differently to physicochemical conditions, and inasmuch as these solutions pass through various natural environments (for example, from the continent through near-shore to deep-water environments) chemical differentiation is possible (Fig.4-6). The precipitation of components from solution depends on their solubilities (Fig.3A). In the sequence Al-Fe-MnOxides

q Iron

oxide

,-Silicates --

p

FeO

ManganeseSiOZ

oxide

-

Carbonates

I

2g

._ UI

salts

CaC03

Sulphates and haloid salts

CaSOL

+ ._

-,E

NaCl

KCI

! MgC$

(MgSOL)

Fig.4. Generalized deposition sequence during transportation leading to chemical differentiation. (After L. V. Pustovalov in: RUKHIN, 1958, p.323.)

c m

F i g 5 Generalized sequence of precipitation of oxides of iron, manganese and silica with distance from source and related to coast-line. (After N. M. Strakhov in: RUKHIN,1961, p.382.)

240

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

SiO~-Pz05-CaC03-CaSO4-NaCl-MgClz, the solubility gradually increases from A1 to MgClz: from a fraction of a milligram to lo6 mg/l at atmospheric pressure. In normal, near-shore, marine environments concentrations may be high enough for dolomitization. Special conditions, however, are required for a high degree of supersaturation and deposition of sulfates and halides to permit diagenetic reaction between them and limestones. Under coastal conditions such as depicted in Fig.5 and 6 , concentrations are usually not high enough, and closed basinal environments such as those in the Red Sea are necessary. The differences in solubility of various components is related to a change from a colloidal state to a true solution (Fig. 3B), i.e., components which are only slightly soluble have a tendency to form colloidal solutions and coagulate readily, whereas highly soluble ones form true solutions. Sea water has a coagulating effect on certain constituents in solution, which explains the near-shore concentration of a number of minerals (Fig.5 and 6). The COZ content greatly affects the solubility of colloids, but it causes less drastic relative effects on components in true solution (Fig.3F). In addition, the pH, Eh and chemical composition of the solutions and environments are the most important factors influencing precipitation, and thus replacement diagenesis. Iron under certain conditions is predominantly transported as a bicarbonate and is primarily deposited in shallow water due to oxidation and early coagulation.

C0,-Zone

p H : 6-7.5

/

.-’ /,,.d”‘2 occurrence

Most active agent:HC03

CL-, SO-,; H2S-Zone

PO?‘

,NOT

pH:7.2-9 E h : -0.2 to -0.5

organisms ,ated

NOTE :Limestone facies extends through ail environments from shore to basin

No benthonic orqanisms Rich i n bacteria

Fig.6. Diagrammatic presentation of physico-chemical zones related to diagenesis of carbonate sediments. @iagram by Borchert in BRAUN,1961, p.472. Reprinted with permission of Drs. H. Braun and H. Borchert and the Z . Erzbergbau Metallhuettenw.)

DIAGENESIS OF CARBONATE ROCKS

24 1

Only small quantities may reach the deeper parts of the basin for pyrite genesis. After its precipitation in shallow water it will be subject to hydraulic factors, similar to clay; in other words, iron oxide may be just another detrital component undergoing mechanical reworking from the time of precipitation. This may explain the occurrence of red-brown hematite, predominantly as clay-sized detritus, in cavities within the Nubrigyn algal bioherms (Plate X, XII; see also WOLF, 1965~). On the other hand, the small proportion of iron in true solution may have been responsible for the occasional wall-rock impregnations and replacement in the vicinity of the cavities (Plate IX, XI) and of the matrix in detrital limestones (Plate XXIII). Silica is transported in true solution and as a colloid in natural waters (Fig.3B), and has, therefore, a better chance to reach deeper waters in contrast to iron oxide (Fig.5 and 6). On the other hand, silica may have an endogenic origin, e.g., siliceous skeletal debris such as spicules, that upon solution may be precipitated as chert or chalcedony. NEWELLet al. (1953) suggested that horizons having different physicochemical attributes, in particular pH (some changes are possibly caused by bacteria) are conducive to the migration of SiOz and CaC03. The former moves to layers of lower pH and the latter to higher pH environments during diagenesis. In agreement with Fig.4-6, NEWELL et al. (1953) and WOLF(1963a) report silicification in deeper water limestones. The former described fossils preferentially replaced by silica, and silica that occurs as geode-like cavity fillings, post-dating the sparry calcite cement, and as nodules and crusts. These silica formations are typically absent from the reef and lagoonal limestones. The selective silicification of fossils is in particular obvious in the lower slopes of the reef talus, but may extend into small reefs buried within the basin sediments of the Capitan reef complex. Current studies by one of the writers of this chapter (H. J. Bissell) on the Permian Kaibab Limestone in southern Nevada in the shelf-tobasin transition zone indicate that greatest silicification of limestones occurred near the hinge-line, particularly slightly basinward. Selective silica replacements of calcareous fossils in preference to the limy matrix of the rock is not unusual. The sequence, from most readily to least readily silicified groups of fossils listed by NEWELL et al. (1953) seems to agree in general with observations made elsewhere, and is as follows: ( I ) bryozoans, tetracorals, tabulate corals, punctate brachiopods; (2) impunctate brachiopods; (3) mollusks (replacement is usually spongy and imperfect); ( 4 ) echinoderms (replacement is usually limited to the surface); (5) Foraminifera; and (6) calcareous sponges and dasycladacean Algae. Silicification beyond material (6) is not selective and will affect the matrix also. Selective replacement of fauna and flora is most probably analogous to other replacement processes. The relative solubility or rate of solution of the particular skeletons concerned may control the differential replacement. Space vacated by the dissolved carbonate is immediately, or sometimes later, occupied by

G . V. CHILINGAR, H. .I. BISSELL AND K. H. WOLF

PLATE XXIII

Nubrigyn skeletal-pellet-limestone. The different color shades of the pellets are due to differential hematite impregnation. The pellet patch (I) shows a gradual downward decrease in degree of oxidation. The pellets in direct contact with the oxidizing fluids passing through the upper void became more intensely affected before cementation. The limestone consists of gastropod, coral, bryozoan and algal fragments and some brachiopods. Under high magnification, the pellets are composed of loosely compacted light-brown micrite unless impregnated with red iron oxide.

DIAGENESIS OF CARBONATE ROCKS

243

PLATE XXIV

A brachiopod shell fragment within the basinal Tolga calcarenite, which is the deep-water facies of the Nubrigyn-Tolga algal reef complex, N.S.W. The fragment is outlined by “sooty” pyrite ( I ) , and partly replaced by minute patches of chalcedony (2). One algal pellet (3) is present.

The cement is composed of thin layers of light-brown fibrous orthosparite (2) and clear granular orthosparite (3). The paragenesis is: Syngenetic-syndepositional: skeletal and algal debris. Syngenetic-prediagenetic: secondarily introduced pellets. Pre-cementation-diagenetic: hematite impregnation of pellets. Syn-cementation-diagenetic:sparite cement.

244

G . V. CHILINGAR, H. J. BISSELL A N D K. H. WOLF

silica (NEWELL et al., 1953). Although aragonite is more soluble than calcite under ordinary conditions, aragonite skeletons are not necessarily silicified prior to calcitic ones. Newel1 et al. suggested that the protective nitrogenous conchiolin which pervades the mollusk shells, for example, releases sufficient ammonia to raise the pH within the immediate micro-environment to prevent solution. Under certain conditions, therefore, aragonite may be more stable than calcite and can resist replacement. The basinal Tolga calcarenite of New South Wales (WOLF,1965a) has two very distinct chalcedony types: one detrital and derived from an older limestone source, and a second type formed authigenically by replacement. Only the latter is of interest here. Silicification is quite often selective to the extent that only a very small central portion of brachiopods is replaced by silica, whereas the outer portions are encrusted by “sooty” pyrite. The silica and pyrite are separated by recrystallized shell calcite (Plate XXIV). On the other hand, along many laminae of the Tolga calcarenite beds there are irregular small patches of chalcedony up to 2-4 mm in thickness which replace grains and matrix. These chalcedony-rich layers are parallelled by upper and lower portions that have a distinct leached and recrystallized appearance in thin-section. The silicifying fluids apparently have also affected the host-rock to some extent to form a “halo” parallel to the siliceous laminae. The inverse physicochemical relationship between the silica and carbonate precipitation, i.e., the former precipitates whereas the latter dissolves, may explain the widespread pressure-solution in the basinal Tolga calcarenite in contrast to its absence in the shallow-water Nubrigyn detrital accumulations. The chemical conditions of the basinal interstitial fluids favored solution of CaC03 and precipitation of SiOz, whereas possibly no such reactions could occur in the shelf environment. Minute (up to 8p) authigenic quartz crystals with well-developed hexagonal cross-sections are present in both shelf and basin limestones of the NubrigynTolga reef complex, N.S.W. Significant to note is their restriction to algal cryptocrystalline calcite and their absence in faunal products and detrital matrix. It is not possible to determine the cause and stage of formation of this authigenic quartz. Most likely it is of late diagenetic-epigenetic origin. On the other hand, TERMIER and TERMIER(1963) reported fairly early-formed euhedral quartz in recently emerged reefs buried in mud. Similar to the SiOz, pyrite is confined to the deep-water limestone facies (Fig.6) of the Nubrigyn-Tolga complex. Three morphologic pyrite types have been noticed here: ( I ) the most widespread are the fine films of “sooty” pyrite on faunal skeletons penetrating even the most minute surface pores on brachiopods, for example; (2) minute pyrite spheres occurring in clusters; and (3) cubes. The formation of most of the pyrite must have been pre-cementation-diagenetic, because it occurred distinctly before CaC03 cement precipitation in most cases. Only the cube-shaped pyrite is of post-cementation origin as it replaces both detrital grains and calcite cement. The occurrence of silica with pyrite suggests that both formed in reducing,

DIAGENESIS OF CARBONATE ROCKS

245

low-pH conditions characteristic of the basinal euxinic black limestone and marl facies. Although it is impossible to state with absolute certainty the origin of the pyrite, a bacterial origin is possible under euxinic conditions. Anaerobic bacteria attack organic matter, extract oxygen and release hydrogen. The latter combines with sulfur derived from sulfates to form HzS, which is a toxic gas readily soluble in sea water. H2S attacks soluble iron compounds to form FeSz, which is highly insoluble and is precipitated in the form of pyrite or marcasite. Other anaerobic bacteria attack sulfates to obtain oxygen needed in metabolism and free sulfur for the genesis of H2S required. The precipitation of iron sulfides by anaerobic bacteria may take place as finely divided dark pigment or it may replace shells or form nodules. Although pyrite genesis may be restricted in some localities to the deep-water facies, it is important to remember that euxinic shallow-water environments may also lead to pyrite and marcasite formation. Textures, structures and diagenesis Diagenesis may lead to formation or destruction of textures and structures. Some of these have been mentioned already in this chapter and others are so well known that a few remarks will suffice. Laminations of certain types, discontinuitysurfaces, stromatactis, birdseyes, club-shaped stromatolites, cone-in-cone, certain spherulites and oolites, faecal pellets, pseudo-breccias, and early fractures may all be of syngenetic-diagenetic or purely diagenetic origin. BOTVINKINA (1960) pointed out that lamination and stratification can be the result of ion transfer and differential precipitation of iron oxide, silica, carbonates, and others, at horizons with corresponding favorable pH and Eh values. Such laminations may be very similar to those formed by ordinary detrital accumulations. Other diagenetic products are concretions, lenses and beds. Stratification or bedding may also be the result of solution and corrosion, sometimes forming discontinuity-surfaces (JAANUSSON, 1961). Residual clay, iron oxide, phosphate, glauconite, corrosion and bore pits, burrows, and bleaching of the underlying sediments characterize hiatuses formed by near-surface diagenetic alterations. Early cementation may control the shape and preservation of stromatolites. LOGAN(1961) suggested that the domed and club-shaped stromatolites are a function of acicular aragonite cement precipitation, which has to occur very early to prevent collapse of these high relief structures in the turbulent littoral environment. USDOWSKI (1 963) presented evidence that cone-in-cone structures, composed of fibrous calcite, are the result of early diagenetic recrystallization of lime-mud beds shortly after the sediment accumulated and was still in an unconsolidated state and saturated with interstitial fluids. If this interpretation is factual, further research may indicate that cone-in-cone structures are valuable paleoenvironmental criteria.

246

G. V.

CHILINGAR, H.

J. BISSELL AND

K.

H. WOLF

Spherulites, other than those shown in Plate XXII and formed by recrystallization, have been reported as products of diagenetic bacterial processes (MONAGHAN and LYTLE,1956; LALOU, 1957).uYE(1959)mentioned that precipitation of calcium carbonate may occur as a gel, and coagulation may mechanically entrap particles of non-colloidal size and form alternate bands of colloidal and entrapped material. Laboratory experiments indicate that spherulites composed of vaterite or one of the several hydrates of calcium carbonate form during initial crystallization. These have not been found in nature, however, and it is most likely that their unstable nature caused a change to aragonite and calcite spherulites soon after formation. CLOUD et al. (1962) stated that the tendency of bacteria to adhere to surfaces may be conducive to the genesis of some types of oolites. Accretion of successivelayers by aggregation of sedimentary particles around successive slimy or gelatinous bacterial sheaths surrounding the initial nucleus may be a likely process. It may conceivably occur up to a few feet within the sediments. The process, however, requires further study. EARDLEY (1938) believed that radial, in contrast to the concentrically laminated, structures of oolites in Great Salt Lake of Utah are a diagenetic feature formed during inversion of aragonite to calcite. Early diagenetic fracturing of the algal micrite bioherms of N.S.W. resulted in calcite veins that post-date internal sedimentation and fibrous calcite, but predate or are contemporaneous with granular sparite. From the structural relations described elsewhere (WOLF,1963a, 1965c),it appears that the solutions that precipitated the granular sparite reached the voids only after fracturing took place. Paragenesis

In a general way, though not always, the sequence of diagenesis takes place in the following order: ( I ) biological and biochemical, (2) physicochemical, and, (3) physical (TERMIER and TERMIER,1963). These processes, of course, overlap to a large extent. With time, there is a decrease in rate of these processes. Little information is available that would permit a paragenetic reconstruction of diagenesis on a regional scale, although it may be a valuable tool for paleogeographic reconstructions. Most of the data are of local value, of meso- and microscopic dimensions. The previously mentioned difficulty of making clear distinctions between syngenetic, diagenetic, and epigenetic processes and products even on a micro-scale is illustrated in the following paragenetic example. Diagenetically formed cavities first have been lined by diagenetically precipitated fibrous calcite. The central cavity was then filled by dolomite which was precipitated from saturated surface waters penetrating the limestone framework (Fig.1; Plate XIV; WOLF, 1963a). (1962a) reported similar dolomite infillings. Both SCHWARZACHER (1961) and FOLK The former calls it syngenetic or primary dolomite. This, although correct, is confusing as it suggests that syngenetic products may be preceded by diagenetic cement, for example. An identical situation occurs when cavities are diagenetically

247

DIAGENESIS OF CARBONATE ROCKS

encrusted by fibrous calcite and the central cavity is filled by detrital internal sediments, brought into the system from the surface that is exposed to syngenetic processes (Plate VI-VIII). From these two examples it seems clear that under certain circumstances one has to expect cyclic formation of syngenetic and diagenetic products. WOLF(1963a) has distinguished, therefore, between syngenetic, diagenetic TABLE IV PARAGENESIS OF DIAGENETW FEATURES (EXEMPLIFYING POSSIBLE REGIONAL TREND)

Littoral algal biohermsl

Fore-reef talus2

Basinal “turbidite”1

Paragenesis I (a) Detrital internal sediment (b) Fibrous sparite (c) Granular sparite

Paragenesis I (a) Calcite cement

Paragenesis I (a) Pressure-solution (b) Granular sparite

Paragenesis 2 Paragenesis 2 (a) Hematite replacement of (a) Calcite cement framework concentrically around voids (b) Detrital internal sediment (b) Chalcedony open-space filling (c) Fibrous sparite (d) Granular sparite Paragenesis 3 (a) Fibrous sparite (b) Chemically precipitated, coarse dolomite, openspace filling (c) Granular sparite

Paragenesis 3 (a) Calcite cement (b) Chalcedony replacement

Paragenesis 4 (a) Hematite open-space filling (6) Granular sparite

Paragenesis 5 (a) Fibrous sparite (6) Chemical and/or detrital internal sediment (c) Granular sparite; a and b alternate to form up to six generations prior to c

Simplified after WOLF(1963a). et al. (1953). After NEWELL

Paragenesis 2 (a) Minute fringe of microdrusy sparite (b) Granular sparite

Paragenesis 3 (a) Granular sparite

Paragenesis 4 (a) Any of the above with (b) Pyritization and/or (c) Silicification prior to or succeeding cementation

248

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

and epigenetic processes and products on one hand and stages on the other. The same complexities occur on a regional scale. Within a limestone formation or a reef complex, one geomorphologic environment may be still in the syngenetic stage, whereas others are undergoing rigorous diagenetic alterations. Or, if all sections of a formation are exposed to similar diagenetic processes, then various parts may be characterized by distinct paragenetic histories as a synthesis in Table IV indicates. If both are present, i.e., characteristic diagenetic products as well as complex paragenetic histories, then the two combined will be a valuable tool for environmental studies. Significant paragenetic relationships may exist between different types of sparry calcite cement. It has been noticed, for example, that light brown fibrous TABLE V PARAGENETIC MODEL OF LIMESTONES

(After WOLF,1963a)

(1) Pre-depositional stage, processes and products (Based on limestone rock fragments, i.e., calclithite detritus, derived from an older carbonate source that underwent diagenesis before erosion.) (2) Syngenetic stage, processes and products (a) syndepositional (e.g., framework accumulation) (b) pre-diagenetic (e.g., purely physical reworking; mechanically deposited internal sediment) ( 3 ) Diagenetic stage, processes and products1 (a) pre-cementationz (e.g., chemical internal

sedimentation, replacement and corrosion of the framework) (6) syn-cementation (e.g., deposition of calcite in open cavities; chemical and mechanical internal sedimentation alternating with generations of cement) (c) post-cementation (e.g., early fracturing permitting deposition of granular calcite)

( i ) above high tide, i.e., subaerials

I

-+ (ii) intertidal3

(iii) below low tide3

( 4 ) Epigenetic stage, processes and products (a) juxta-epigenetic4 (b) apo-epigenetics

In detailed studies of Recent and Pleistocene carbonates it may be possible to subdivide diagenesis further into i , ii, and iii. The suffix “cementation” can be replaced by “lithification” (STRAKHOV, 1958) depending on the product permitting subdivision. In the present case, the first generation of brown fibrous calcite was used. Other subdivisions can be used depending on what type of sediments (i.e., marine or nonmarine, etc.) are under investigation (WOLF, 196%). 4 “juxta-” meaning “near” or “close-by”. “Apo-” meaning “far”, “remote”.

249

DIAGENESIS OF CARBONATE ROCKS

sparite always precedes colorless granular calcite, where both are present in the Nubrigyn-Tolga reef complex of New South Wales (WOLF,1963a, 1965a,c). The fibrous calcite is restricted to the shallow-water shelf bioherms and associated algal calcarenites. From one to six generations of fibrous sparite, sometimes separated by internal sediments such as minute pellets and iron oxide, fill open spaces in algal reefs and form the cement of the calcarenites (Plate VIII, X, XIII, XIV, XIX). If any voids remained, they were subsequently occupied by clear granular sparite (Plate V, VIII-XI, XIII, XIV, XIX). In numerous instances it can be demonstrated that fractures terminating in the voids permitted solutions to deposit the granular sparite (Plate XI, XII, XIV). These distinct paragenetic relationships, which remain constant throughout the reefs, permit the subdivision shown in Table v. The brown fibrous sparite forming the cement of the shallow-water limestones marks the syn-cementation stage and separates, therefore, the pre-cementation from the post-cementation stage. The clear granular calcite belongs to the postcementation period as it has not contributed to cementation of the shallow-water sediments, and was formed by different processes only after fracturing of the rock occurred. Obviously, there is a hiatus between the fibrous and granular sparite formation. The deep-water Tolga calcarenite, on the other hand, was cemented by clear granular sparite at a much later stage than the equivalent Nubrigyn shelf deposits as indicated by considerable pre-cementation pressure-solution. In other words, while the shelf deposits were in the syn-cementation stage, the basinal limestones were undergoing pre-cementation diagenesis. For regional paragenetic reconstructions of diagenetic and epigenetic alterations it may be important to find features that overlap in critical areas to permit a “time-correlation”. For example, if hematite and pyrite geneses are confined to TABLE VI SIMPLIFIED EXAMPLE ILLUSTRATING CORRELATION OF DIAGENETIC PRODUCTS

(After WOLF,1963a) -~

Shelf sediment

Intermediate sediment

Basinal sediment

(I) Limestone accumulation

(I) Limestone accumulation

(I) Limestone accumulation

(2) Hematite debris

(2) Hematite debris and “sooty” pyrite on fossils in places

(2) “Sooty” pyrite on fossils and minute pyrite spheres

(3) Fibrous calcite cement

(3) Occasional fibrous sparite, (3) Pressure-solution but mainly granular sparite (4) Granular sparite cement

250

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

shallow and deep water facies, respectively, then one needs some criteria to prove that both did actually occur during the same paragenetic stage. Table VI shows a simplified example where it is possible to demonstrate that hematite and pyrite geneses took place more or less penecontemporaneously, because both occurred before cementation and pressure-solution. If, for example, the pyrite in the basinal limestones had distinctly replaced both fossils and cement, the FeSz would have been of a later origin compared to the hematite. (Both pre-cementation and post-cementation pyrites are present in the Tolga calcarenite-the former is “sooty” and the latter occurs as cubes.)

Paleogeographic environment-indicating diagenetic products Early diagenesis is controlled by surface and near-surface factors and its products are indicative of the environment. Some may reflect only very local conditions; others, however, may be useful criteria to interpret paleogeomorphologic and paleogeographic conditions. The petrographic discussion below is based on an Australian Devonian algal reef complex (WOLF,1963a, 1965a)and serves as an example. It should be emphasized that although only the diagenetic products are used here to illustrate their usefulness in environmental interpretation, all other paleontologic, structural and stratigraphic criteria support the reconstructions made. The following early diagenetic features were found to be indicative of a littoral environment for the algal bioherms: (I) internal open-space structures, i.e., incorporated former surface pits and surge channels, and the so-called stromatactis; (2) extensive internal detrital sedimentation; (3) certain internal chemical sediments, e.g., red iron oxide and dolomite; (4) fibrous calcite cement; (5) travertine; and (6)complex paragenesis. The open-space structures, detrital sediments and chemical internal precipitates have been described earlier, and the latter are listed in Table I1 and IV. The chemical composition of the interstitial fluids must have changed relatively quickly as indicated by the successive and alternating generations of iron oxide, calcium carbonate and dolomite precipitation; and bleached, leached or corroded host-rock walls. The paragenetic picture is very constant from bioherm to bioherm within the same unit. It seems very unlikely that such a complex paragenesis could occur either in a sublittoral environment, or under supralittoral conditions. First, extensive internal channel systems are not likely to form under sublittoral conditions. It is true that they can occur in subaerially-formed limestones such as eolianites, but these have mainly vertically oriented channels in contrast to the horizontal ones of the littoral carbonate sediments. Second, the diagenetically formed internal sediments, such as iron oxide and dolomite, are typical of littoral origin. If one admits the possibility that both iron oxide and dolomite internal cavity fillings could occur in limestones below low tide, then there is still a third factor, the complex cycles of fibrous calcite and internal sediments to explain. For sediments to

DIAGENESIS OF CARBONATE ROCKS

25 1

penetrate into a limestone framework, turbulent conditions and surging powerful currents, unlikely to occur in sublittoral environments, seem to be necessary. Admittedly, density currents and/or turbidity currents can transport sediment into and across the sublittoral zone. Under conditions below low tide, sediments would merely settle and, at the most, drift to and fro; but it seems unlikely that they could penetrate into a complexly channeled sediment framework. Superficial observations made on Recent and Pleistocene limestones tentatively suggest that well-developed and extensive fibrous, and possibly drusy, calcite and aragonite development is confined to shallow-water and supralittoral environments. It is interesting to note that many beach-rocks have acicular, i.e., fibrous, carbonate cement, whereas subaerially cemented eolianites of Lord Howe Island, for example, show mainly drusy sparite (Plate XVIII). These observations suggest that the Nubrigyn bioherms and associated calcarenites, which are characterized by fibrous calcite cement, were formed in a littoral environment. In a number of thin-sections it has been observed that filamentous, unicellular algal mats and algal micrite colonies abut against overlying dense laminated travertine crusts which are composed of fibrous sparry calcite. Similar colonies in turn encrust the travertine layers. It appears that the travertine could have been formed only by exposure of the algal bioherms above sea water at low tide, whereas solution, evaporation and precipitation formed the sparry calcite crusts after dissolution of part of the algal colonies. One can conclude from these discussions that, based on the diagenetic products alone, the Nubrigyn algal bioherms were formed most probably in a littoral environment (see WOLF,1965c, for more details).

Diagenesis and limestone classification The foregoing information on diagenetic alterations imposed on syngenetic limestone textures (Fig. 7-14) makes it clear that it is very difficult to follow one simple nomenclature and classification scheme for carbonate rocks, particularly for limestones. A scheme is necessary that suits both the practical and research geologists and is applicable in superficial and super-detailed studies, if a common meetingground of ideas can be realized. Perhaps no such ideal classification is available. As shown in Table VII (WOLF,1963b), the descriptive and genetic stages can each be subdivided into two sub-stages based on method and accuracy of the investigation carried out. During superficial studies only the size-nomenclature may be required. With the use of a binocular microscope it is possible to determine the grains or framework and the micrite matrix/cement ratio and textures, but it may not be possible to make a genetic interpretation of these components. This has to await detailed thin-section investigations. In the final analysis, when all depositional, stratigraphic and paleontologic information has been assembled, a paleogeomorphologic reconstruction is possible, and the sediments can be named

252

G . V. CHILINGAR, H. J. BISSELL A N D K. H. WOLF

Fig.7. Diagenetically altered organic-rich micrite, showing few centers of growth of slightly larger microcrystalline calcite. Darker areas are possibly “dead oil”. Loray Formation (Permian) from outcrop in Dead Horse Wash, White Pine County, Nevada; x 40.

Fig.8. Diagenetically altered skeletal limestone, showing formation of sparry calcite within brachiopod and gastropod shells, and in the matrix material as well. Loray Formation (Permian) from outcrop in Dead Horse Wash, White Pine County, Nevada; x 10.

DIAGENESIS OF CARBONATE ROCKS

253

Fig.9. Diagenetically altered bryozoan-encrinal limestone, illustrating selective diagenesis. Crinoid ossicles show authigenic overgrowths, but the lioclemid bryozoans and interstitial matrix are relatively unaffected. Gerster Formation (Permian) at type locality near Gerster Gulch, Tooele County, Utah; x 5.

Fig.10. Advanced stage of diagenesis of a criquinite, showing mostly relics of crinoid-stem fragments and development of calcite. Hall Canyon Member (Morrowan) of Oquirrh Formation in Oquirrh Mountains, Utah County, Utah; x 5.

254

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

Fig.11. Early to medial stage diagenesis of calcarenite (criquinite variety), illustrating alteration of encrinal material and to a lesser degree the finer-grained, calcarenitic, interstitial matrix material. Hogan Formation (Desmoinesian) west of Wendover, water reservoir, Tooele County, Utah; x 5.

Fig.12. Early diagenesis of a calcarenite, showing calcareous overgrowths on lime-pellet grains. Meadow Canyon Member (Derryan) of Oquirrh Formation, Cedar Mountains, Tooele County, Utah; x 20.

DIAGENESIS OF CARBONATE ROCKS

255

Fig. 13. Coarsely-crystalline sparite, illustrating advanced stage of diagenesis of a calcarenitic criquinite. Ely Limestone (Derryan), Moorman Ranch area, White Pine County, Nevada; X 30.

Fig.14. Incipient dolomitization of a skeletal-detrital limestone, showing diagenesis of encrinal material and to a lesser degree other skeletal elements. Fusulinid test was silicified first, but later was partly replaced by dolomite. Ferguson Mountain Formation (Wolfcampian) in outcrop near top of the “Bear’s Claw” north of Wendover, Tooele County, Utah; x 20.

256

G . V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

TABLE VII PETROGRAPHIC AND PETROLOGIC STAGES OF CARBONATE INVESTIGATIONS

(After WOLF,1963b)

I

Descriptive

Genetic

-

~

Hand-lens investigation

Binocular microscope investigation

Petrographic microscope investigation

Total petrographic summary including depositional structures, stratigraphy, and paleontology

Size nomenclature only, i.e., calcirudite, calcarenite, calcisiltitel, calcilutitel (= micrite)

Morphologic and size nomenclature

Genetic, morphologic and size nomenclature

Genetic, morphologic and size nomenclature with geomorphologic terminology

e.g., calcarenite

e.g., pelletsparite-calcarenite

e.g., algal pellet-orthospariteecalcarenite

e.g., algal pelletorthosparite-beach calcarenite

e.g., micrite

e.g., micrite

e.g., algal auto-micrite biolithite

e.g., algal automicrite knoll reef

__

-

The term calcilutite is usually used for both clay- and silt-sized particles in the descriptive stage as they may not be distinguishable. In thin-section work, however, discrimination is possible.

1

accordingly, e.g., skeleton-orthosparite-beach calcarenite. For such a step-by-step build-up, the descriptive and genetic nomenclature must be kept separate, as repeatedly emphasized. Hence, terms that satisfy requirements in both stages are given in Tables 111 and VIII. The classification scheme (Table IX) is descriptive; the only difficulty lies in the recognition of a micrite biolithite in handspecimen, and it would be called in most cases “micrite limestone” until thin-section work furnishes more detail. Two examples on how the descriptive names can be easily changed into genetic terms by adding prefixes are presented in Table VII. Due to the diagenetic alterations of the matrix and cement in limestones, the micrite/sparite ratios may not be a true reflection of turbulence and washing (“winnowing” of some geologists) of the depositional environment, and FOLK’S (1959) concepts should be considered with care in this regard. His classification scheme, and the modified version given here, based on the syngenetic “grain or framework-matrix/cement ratio” is still useful and need not be discarded because

257

DIAGENESIS OF CARBONATE ROCKS

TABLE VIII COMPONENTS OF ALLOCHTHONOUS LIMESTONES (SOME ARE IN SITU PRODUCTS)

(After WOLF,1963b) __ Descriptive-morphologic

~~

~-

~~

Genet ic-morphologic

pellets

faecal pellets bahamite pellets algal pellets

limeclasts

intraclasts extraclastsl (if rock is composed of more than 50% of extraclasts = calclithite)

oolites pisolites (= concentric fabric^)^

physicochemical algal weathering2

lumps

physicochemical oolites and pisolites algal oolites and pisolites weathering oolites and pisolites2

skeletons (floral and faunal)

e.g., coral, Bryozoa, Brachiopoda, and Algae

micrite

allomicrite orthomicrite automicrite pseudomicrite3

sparite and microsparite

orthosparite and orthomicrosparite pseudosparite and pseudomicrosparite3

*

}

Of pre-depositional origin, i.e., from an older limestone source (WOLF,1965b). Mostly a Recent or Pleistocene residual weathering product. Diagenetic to epigenetic product; in ancient rocks it is a penecontemporaneous product. Includes superficial oolites and circumcrusted particles.

of diagenetic alterations. In the scheme outlined above, i.e., the four sub-stages leading from descriptive to genetic stages, the syngenetic, diagenetic and epigenetic characters of limestones have been included (Table IX). Hence, the lower endmember in Table IX is either an unaltered micritic limestone or a crystalline limestone composed of pseudosparite or pseudomicrosparite. (Those interested in details of classification may wish to consult HAMand PRAY,1962, for example.) M icvites Limestones herein classed as micrites are those rocks originating from diagenesis of calcareous mud or ooze. Lime ooze may have a lower plasticity than clay, yet readily (and probably early in the diagenetic process) forms sets and systems of

258

G . V. CHlLlNGAR, H . .I. BlSSELL AND K . H . WOLF

TABLE IX LIMESTONE AND DOLOMITE CLASSIFICATION SCHEME

(Modified after FOLK,1957, 1959; and WOLF,1960) Limestone

____.__

micrite and/or sparite ( %) __

~~

~

limeclasts

~-

~

50

90

- -

-

-

. ~ _ ~

~~

~~~

~~

~~

skeletons

-.

skeletonlimestonel

10 -

~__

~

_ _ _ skeletonmicritelimestonel

oolites pisolites

~~

~

limeclastlimestone _

pellets

_ - ~ ~ pelletlimeclastmicritemicritelimestone limestone

organic in situ growths

.~

~~

pelletlimestone

lumps

~~

oolite(pisolite-) limestone

lumplimestone

coral(algal-, etc.) bioli thite

oolitemicritelimestone

lumpmicritelimestone

coral-micritebiolithite, algal-micritebiolithite, etc.

or

or

or

or

or

or

skeletonsparitelimestonel

limeclastsparitelimestone

pelletsparitelimestone

oolitesparitelimestone

lumpsparitelimestone

coral-sparitebiolithite etc.

mi~rite-~ skeletonlimestone

micritelimeclastlimestone

micritepelletlimestone

micrite oolitelimestone

micritelumplimestone

micrite3-coralbiolithite, micrite-algalb iolithite, etc.

or

or

or

or

or

or

spariteskeletonlimestone

sparitelimeclastlimestone

sparitepelletlimestone

spariteoolitelimestone

sparitelumplimertone

sparite-coralbiolithite etc.

___

__

~-

~~

~

-

micrite limestone3 ~

. -

~

__

~~

-~

- .-

~

~

-

-

micri tebiolithite3 -

or sparite3s4(= crystalline) limestone Use size-nomenclature, i.e., calcarenite, dolarenite, etc., instead of “limestone” wherever possible. State composition of impurities, i.e., quartzose, etc. 3 Sparitelmicrite ratios do not necessarily indicate degree of washing because the sparite may be pseudosparite. Also automicrite can form wave-resistant growths, e.g., algal micrite bioherms. Tufa, travertine, and caliche are often sparite limestones formed in situ. 5 Note preferential dolomitization of matrix, etc. 6 These columns are examples only. In fact any grain, colonial growth, matrix, and sparite can be replaced. 7 Similar to limestones, dolomites range from dolomicrite to dolosparite. 8 Dolomicrosparite and/or dolosparite may be used instead.

1 2

. .

259

DIAGENESIS OF CARBONATE ROCKS

~Dolomitized limestone and dolomite ~~

~~~

~~

~~

-~

.~

grains absent,

allochemical grains present ~~

partially replaced by extensively replaced586 dolomite5 96

impurities present ( I 0-50 %)

~~

completely replaced5 (> 90%)

(10-50%)

(SO-90 %)

dolomiticskeletonlimestone, pelletlimestone, etc.

calcareous skeleton- skeleton-dolomite, dolomite, pelletpellet-dolomite, dolomite, etc. etc.

completely replaced (> 90%)

~~~~

_~_____

~

-

~_______

e.g., sandy2 skeletonmicrite-limestone, silty2 dolomitic oolite-sparite, etc.

dolomitic pelletmicrite-limestonel

calcareous pelletmicrite-dolomite’

pellet-micritedolomite1

or

or

or

dolomitic pelletsparite-limestonel

calcareous pelletsparite-dolomitel

pellet-sparitedolomite1

dolomitic micritepellet-limestone

calcareous micritepellet-dolomite

micrite-pelletdolomite

or

or

or

dolomitic sparitepellet-limestone

calcareous sparitepellet-dolomite

sparite-pelletdolomite

dolomitic micrite7

calcareous dolomicrite’

dolomicrite

dolomicrite

dolosparite7

“primary”?

e.g., sandy2 dolomitic pelletlimestone, silty-sandy2 skeleton-dolomite etc.

~~

dolomitic sparite7

..

pebbly, gritty, sandy, silty, clayey, skeleton-dolomite, pellet-dolomite, etc.

~

e.g., clayey-micrite, dololutite, etc.

joints (diaclases), the development of which provides avenues for gas and liquid transfer. Open channels are created, and what was hydrostatic pressure becomes directed pressure, and changes can occur with greater rapidity at the interface. Before a discussion of diagenetic effects upon micrites can be expanded, it is necessary that certain concepts and nomenclature should be clarified. FOLK (1959) termed micrite the lime-mud component (very fine-grained ooze or paste). Mud is very

260

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

fine-grained (or crystalline) dense material which geologists have described as “lithographic”, “cryptocrystalline”, “cryptograined”, “microcrystalline“, “micrograined”, etc. LEIGHTON and PENDEXTER (1962) arbitrarily set the upper limit of the mud component at 0.031 mm, but some prefer this limit to be 0.005 mm (Table 111). An exact size limitation is not too critical (BAARS, 1963). Numerous of the calcilutites and some calcisiltites fit into the category of micrites, or micritic limestones depending on the limit set for “micrite”. A prevalent tendency among some petrographers is to apply the textural term of “aphanitic ”to the micrites, at least to those which have a micro- or cryptotexture. The term aphanic was proposed by DEFORD (1946) as a textural term for carbonates, particularly limestones, which are crystalline (and/or grained), and the discrete particles of which are smaller than 0.1 mm. Microcrystalline (also micrograined) and cryptocrystalline (also cryptograined) are the two textural subdivisions. The term aphanitic is more loosely defined, and it is herein rejected as a textural term for carbonates, with the proposal that aphanic should be adopted because of its precise definition. Aphanic as a textural term has been applied successfully to petrographic studies of limestones (MOLLAZAL, 1961) and of dolomites (OSMOND, 1956). Many geologists, particularly sedimentary petrographers, use an upper limit within the medial silt-range to define micritic texture; as pointed out by BAARS(1963), for most cases this is the practical limit for particle recognition. The origin of lime-muds cannot be determined with accuracy in all cases, and it is probable that several mechanisms contribute and operate for its formation (Table 111;WOLF,1962, 1965b). Discrete particles may be chemically or biochemically precipitated fine crystals, or finely comminuted “clastic” particulate material derived from originally larger particles and other sources. LOWENSTAM (1955) and LOWENSTAM and EPSTEIN(1957) recognized the possibility that mud-sized needles of aragonite on parts of the Great Bahama Bank may be derived from calcareous Algae. BAARS(1963) pointed out that several codiacean (green) Algae secrete clay-sized aragonite needles within their tissues; the genera Penicillus, Rhipocephalus, and Udotea were cited as examples. These aragonite needles disintegrate to produce lime-mud upon death of the Algae. It is because these Algae (and others) are important in modern lime-depositing seas and because of their short life cycle, that they may be extremely important sources of lime ooze. Furthermore, the small particle size and resultant small size of interparticle pores renders the lime-muds particularly susceptible to diagenesis, especially pressuresolution and simple interparticle cementation. As will be pointed out, these characteristics also make the lime ooze amenable to diagenetic dolomitization early in the sedimentary history. LOWENSTAM (1955) has stated that some calcilutites attributed to physicochemical precipitation have formed by breakdown of calcareous Algae, particularly the poorly calcified forms. In referring to the “white reef” in certain areas of Alberta, BELYEA(1955) stated that much of the reef mass is fine-grained commi-

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nuted organic debris, and much of it is white dense limestone probably formed largely by lime-trapping Algae. Some lime ooze may precipitate on or near algal plants and form “algal slime”, because the plants extract carbon dioxide from immediately adjacent sea water (PRAY,1958). THOMAS and GLAISTER (1960) mentioned that in some Mississippian carbonate sequences, which they studied in the Western Canada Basin, microgranular carbonates graded vertically and laterally into chalky micrograined carbonates. They regarded part of the carbonates to be of chemical origin, but much of it represents “flour” formed by disintegration and abrasion of fossil debris and algal growths which developed in a shelf environment. HAMBLETON (1 962) indicated that in Missourian-age carbonate rocks in Socorro County, New Mexico, the dominant matrix material of back-reef facies is microcrystalline calcite ooze and “reef milk” (the latter being very fine-grained, white and microcrystalline calcite), derived from abrasion of the reef core and reef flank. EDIE(1958) recognized chalky (micritic) limestones in carbonates of Mississippian age in southeastern Saskatchewan, suggesting that their origin may be attributed to “flour” formed by disintegration and abrasion of fossil debris and algal growths under intense wave action in shoal areas. He stated (EDIE,1958, p.105) that this flour “. . .might be expected to settle in the quiet-water environments of lagoons, intershoal areas of the shelf, and in the basinal areas.” Some micrites possibly originated from “algal dust” at least in part. WOOD(1941) first called attention to certain finer-grained varieties of Carboniferous limestones which he ascribed to an “algal dust” origin, thus coining the term. In applying this descriptive and genetic term, CAROZZI (1960) noted that the grains themselves are angular with a diameter reaching 2-3p; he favored usage of the term “algal dust” when the finegrained limestone contains associated algal tubes, or when it is clearly derived from algal material. Later CAROZZI and SODERMAN (1962) pointed out that petrographic studies of Mississippian limestones in Indiana suggested that certain calcilutites developed from “algal dust” produced by phytoplankton. Algae are capable of precipitating micro- and cryptocrystalline calcite which, attendant upon attrition, abrasion and disintegration, yields aphanic-textured detrital lime particles which are in a sense “algal dust” (or algal allomicrite, Table 111). The extent to which bacteria can precipitate directly lime ooze, which ultimately will result in micrite, is not fully understood. In his studies of bacterial precipitation of carbonates in sea water, LALOU(1957) emphasized that perhaps the role of bacteria is largely one of changing the physicochemical conditions of the medium, increasing its concentration of C02 up to saturation, enriching it in calcium and giving rise to an escape of H2S by reducing sulfates. The effect of such reactions is to change the alkaline reserve of the medium, the pH, etc. It was his interpretation that the formation of carbonates by bacteria may be obtained if: (I) there is presence of assimilable organic matter in sufficient quantity, (2) the temperature is sufficiently high, (3) there is maximum light and sunshine, and ( 4 ) the waters are quiet and are seldom renewed. These conditions, he believed, are

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to be found in the lagoons and portions of the tropical sea water most isolated from the open seas. . During compaction of lime-mud differential strain may result, and this can vary from one depositional site to another, i.e., whether a lagoon, bank, miogeosyncline, etc. Nuclei of recrystallization will be set up giving rise to ultimate crystalline mosaic. It was pointed out by WARDLAW (1962) in his studies of diagenesis of Irish Carboniferous limestones, that during recrystallization nuclei of strain-free grains originate at several points, the number of points increasing with time, and the strain-free grains may grow until they completely consume the matrix. Obviously, to produce a finely crystalline micrite or microsparite under such circumstances requires a large number of sites where new nuclei can develop. Lime ooze has a high fluid content in the interparticle pore spaces. THOMAS and GLAISTER (1960) studied porosity and facies relationships of some Mississippian carbonates in the Western Canada Basin and called attention to the fact that lime-mud, which formed in quiet-water environments of lagoons and shoal areas and which is chalky to clay-like in grain-size, has a low oil-wetting ability and a high connate water saturation. Diagenetic dolomitization proceeds relatively fast in such ooze, and it would appear that dolomitization processes are strongly controlled by the presence of fluids in intergranular and intercrystalline pore spaces, particularly in those which have a high fluid content. It is noteworthy that calcium carbonate mud which precipitated as a colloidal gel, encrusting leaves of Algae, normally has a high fluid content; during diagenesis a crypto- or micrograined limestone will form first and commonly “syneresis” cracks, joints, and primary contraction vugs will develop. Magnesium ions present in the original algal material may now be disseminated in the “alga1 dust” and will serve as nuclei for diagenetic dolomitization. With sufficient concentration of Mg2+ ions in the interparticle pore fluid, the transfer of Ca2+ions out, and Mg2+ ions in, through the intergranular film is hastened and wholesale diagenetic dolomitization of the limemud can occur, particularly if additional magnesium ions are added at the interface or from the lime ooze beneath. Crypto- to microtextured chalky lime-mud that is rich in comminuted shell material, and/or cryptocrystalline or microcrystalline tests of calcareous composition, is also normally high in mapesium (from trace up to 12% and, exceptionally, more; cf. CORRENS,1939). Percolating waters dissolve the calcium much faster than the magnesium (in accordance with the law of mass action) from a deposit of lime-mud composed of such detritus, and the relative amount of magnesium increases with progressive diagenesis. As noted by SUJKOWSKI (1958),the proportion of Mg/Ca approachesslowlyto 1/1,and accompanying replacement will give dolomite. He believed that such a diagenetic dolomite results in a much greater reduction of volume than takes place in the diagenesis of calcareous mud leading to limestone. Numerous fine-textured limestones which petrologists may term calcilutites and calcisiltites in the field may, upon petrographic examination, be defined as mi-

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crites (Table 111). Originally, the sediment may have been crypto- or microcrystalline; such finely divided material (whether crystalline or grained, or both) can recrystallize by pressure-solution into a mosaic of larger crystals by the solution of the smallest, supersoluble grains and redeposition on the larger grains, or by grain growth (BATHURST,1958). Pressure-solution is the transfer by solution of ions from a point of intergranular contact (where the crystal lattice is strained) by diffusion down the ion concentration gradient to a point of deposition on a crystal where there is no strain (STAUFFER, 1962). Grain growth in limestones is defined by BATHURST (1958) as the ion transfer from one crystal lattice to another without any intervening solution. The process of ion migration in the solid state leads to the enlargement of the larger grains at the expense of the smaller. During diagenesis of a lime-mud to form micrite, particularly one containing particulate skeletal material (such as echinoderm ossicles), there will be transfer of ions with concomitant enlargement of the skeletal material. Inasmuch as crinoid and other echinoderm fragments consist of single large calcite crystals, they are commonly enlarged by the deposition of calcite in crystallographic continuity with the fragments (STAUFFER, 1962). It should be pointed out here that if the lime-mud consists of finely comminuted material (by some geologists termed “matrix”) in which there are embedded larger fragments, including skeletal material, diagenetic dolomitization (if such occurs) will affect the matrix material first; the particulate larger skeletal material is most resistant. SIEGEL (1963) has noted that a factor that may influence diagenetic dolomitization of micrite is the polymorphic form of the calcium carbonate that is precipitated. Aragonite, because of its metastable state, should react more readily than calcite to magnesium-bearing waters to form dolomite. Vaterite is a more metastable form of calcium carbonate than aragonite and would, therefore, be even more likely to form dolomite. ZELLERand WRAY(1956) have demonstrated with laboratory studies that certain elements such as strontium and barium cause calcium carbonate to precipitate in the form of aragonite under conditions where the carbonate phase would normally be calcite. As pointed out in a preceding section, SIEGEL (1960) found that the alteration of aragonite to calcite in natural samples was inhibited by the presence of strontium, and that strontium might have to be removed before an alteration could take place. As he (SIEGEL, 1963) pointed out, a strontium-bearing aragonite might not react with sea waters to form dolomite, and the lithologic association observed in the geologic rock column would be limestone and gypsum, which is a relatively common pairing. “The role of the impurity ion must, then, be considered when speaking of the susceptibility of calcium carbonate to either early diagenetic or metasomatic alteration” (SIEGEL, 1963). The mechanism of cementation of lime-mud during diagenesis to form micrite (as well as certain other limestones) presents numerous problems. In studying certain limestones in Indiana, NITECKI (1960) suggested these two

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possibilities: ( I ) dissolution at points of high compressive stress and reprecipitation at points of low stress; and (2) dissolution of the organically formed calcite because it is unstable for reasons other than stress, i.e., because there is an unstable amount of MgO present as impurities in the organic calcite, reprecipitation of stable calcite in pores will occur. In the first case, Nitecki noted that as long as the pore space is filled with water the grain-to-grain contact is limited; the existing pressures are, however, hydrostatic except at the points of grain-to-grain contacts. The pores begin to fill gradually with precipitated calcite, giving rise to cement. As the process proceeds, the pressure becomes geostatic. Newly precipitated cement is nearer to the thermodynamic state of equilibrium than the organically precipitated, metastable calcite of the fossils. The result is a further growth of cement-like calcite in preference to the pre-existing organically precipitated crystals. NITECKI (1960) believed that, because the hydrostatic pressure is dependent upon the depth of the overburden, the solubility of limestone is higher at greater depths than at lesser depths (lower pressure). The CaC03 in solution will migrate to areas of lower pressure (lesser depths) where it will precipitate, fill the pores, and cement the sediments. The process of cementation will thus “proceed upward and will be generally accelerated because the pressure will be more geostatic in character” (NITECKI, 1960). These conclusions harmonize, in general, with statements presented herein, and add further credence to the suggestion that fluids highly charged with dissolved carbonate minerals can migrate toward the shelf area from the basin (greater depths and greater overburden) and effect diagenetic changes in the transition, hinge-line, or shelf lime-muds. Dolomitization does not occur because of lack of a copious supply of magnesium ions (and other factors as well), and the resultant diagenetic effect is cementation leading to lithification. As has been noted herein, dolomitization (if it does occur) does not necessarily occur in the same sediment, but the magnesium ions can migrate considerable distances through the interparticle fluids to cause diagenetic dolomitization in another realm. Perhaps this explains the presence of more areally extensive dolomites in the carbonates of Pennsylvanian and Permian age along and in the immediately adjacent shelfward portion of the Las Vegas Hinge Line in the three-corners area of Nevada-Arizona-Utah. The occurrence of syngenetic pyrite and/or marcasite in micrites has been noted by many authors. KRUMBEIN and CARRELS (1952) pointed out that pyrite and calcite can form and be stable in an environment in which the pH is approximately 8.0 and in which the Eh is approximately -0.3. It is to be noted that the negative Eh value does not imply a stagnant environment. As emphasized by MORETTI (1957), marine open-circulation conditions may exist down to the depositional interface, whereas below this level there may be a tendency toward a reducing environment due to depletion of oxygen. Thus, lime-muds beneath neritic normal marine opencirculation environments may have the property of the euxinic environment (KRUMBEIN and GARRELS, 1952). MORETTI (1957) stated that if one assumes the depositional interface and the zero Eh level to be coincident, i.e., above the depo-

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sitional interface oxidizing conditions exist whereas below the depositional interface reducing conditions prevail, then the decomposition of entombed organic matter would be anaerobic. Such decomposition would yield various products, including HzS, and a reducing capacity would be rendered the environment and the S ion would be provided. Pyrite could form if sufficient amount of iron was introduced to the sea at the time of accumulation of the lime-muds. Syngenetic pyrite and/or marcasite would form under these conditions, and diagenetic iron sulfides could form at a later date. Iron monosulfides, such as hydrotroilite, could accumulate syngenetically, but under the effects of diagenesis would change to pyrite. It should be remembered that various strains of bacteria can cause iron to be taken into solution (such as at the provenance site) and be transferred to the depositional site where it is subsequently precipitated to react and form syngenetic products and possibly diagenetic minerals. Diagenesis of pure lime ooze usually leads to fairly homogeneous micrite or micritic limestone as a result of compaction, with accompanying expulsion of water, and filling of pore spaces by micrite and by sparry cement. Presence of clay minerals retards the process of crystallization, and this is reflected in the texture of the indurated material. Perhaps influx into a sedimentary basin of pure lime-mud of micrograined texture for a prolonged period of time is an unusual circumstance. By the same token uninterrupted accumulation of microcrystalline lime ooze from supersaturated waters is an anomalous sedimentary feature of depocenters. Yet, limestones and “primary” dolomites (or dolomites of the restricted or evaporitic suite) of this category form thick and areally extensive members and formations in rocks of Precambrian to Pleistocene age in the Eastern Great Basin area. These finely textured dolomites, dolosiltites, dolomicrites, micritic limestones, and micritic limestones with oolites are of particular significance in certain Permian units of the hinge-line area of southern Nevada. Thin-sections of some micrites, however, reveal presence of finely-divided organic matter (in some instances “dead oil”) and micro-textured silica (not necessarily cement), indicating that other sediment was also introduced; the process of diagenesis did not obliterate the evidence. Skeletal limestone Rocks herein classified as skeletal limestones include those fragmental clastic and detrital rocks that have been given a number of names: bioclastic, fossiliferousfragmental (e.g., criquinites), skeletal-detrital, and others. No single set pattern of diagenesis has been established for these sediments; grain growth, introduction of rim cement, and numerous processes collectively lumped under the catch-all term “recrystallization” occur with apparent rapidity in some skeletal limestones, and with variable speed and direction in others. In his studies of the petrography and facies of some Upper Vistan (Mississippian) limestones in North Wales, BANERJEE (1959) differentiated five limestone types, as follows: ( I ) shelly calcite-mudstone, (2) shelly calcite-siltstone, (3) co-

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quina-lutite, (4) bioclastic calcarenite, and (5) crinoidal calcarenite. Petrographers will readily recognize that types I and 2 are transitional from the micrites on the one end to the skeletal limestones (3, 4 and 5 ) on the other end. His pure calcitemudstone (grain size 0.5-4.0,~) is the micrite of some geologists. A limestone consisting dominantly of calcite-mudstone, but with some skeletal debris, is modified by the term “shelly”. His “coquina-lutite” is a limestone, the dominant component of which is skeletal debris of sand and silt grade, which imparts to the rock a coarser texture than that of the shelly calcite-mudstone, though calcite-mudstone is present as matrix. In some respects this usage corresponds to certain nomenclature of DUNHAM (1962) who applied terminology of grain-support versus mudsupport. Thus, if the skeletal limestone is grain-supported with minor amounts of lime-mud interstitial material, it will react to diagenesis differently than if it was mud-supported and skeletal particles were in the minority. BANERJEE (1959) for example, divided his coquina-lutites into three subtypes on the basis of particle orientation and grain size, as follows: Type I : without preferred planar shape orientation of skeletal particles; coarse-grained. Type 2: skeletal particles with preferred planar shape orientation parallel to the bedding plane, and having roughly the same grain size as Type 1. Type 3: the finest-grained of the three with more calcite-mudstone, and with a preferred planar shape orientation of skeletal particles parallel to the bedding plane. As will be pointed out herein, some of these parameters of grain- and skeletal-orientation exert a significant influence on the processes of diagenesis of skeletal limestones. Skeletal detritus is, of course, subject to abrasion and disintegration in high-energy environments which are typified by wave and current agitation and surf surge. Some of the skeletal particles may be reduced to sand-, silt-, and even clay-size grades in lower-energy environments, and the process may be to some degree syngenetic and to a degree diagenetic. For example, DAPPLES (1938) suggested that the continued size reduction of skeletal debris by scavengers might have produced the structureless calcilutites which are common in the Paleozoic. GINSBURG (1957) pointed out that boring blue-green Algae, although small, are extremely abundant in carbonates, and tiny filaments penetrate shell fragments. He stated that in the modern seas these organic destructive agents attack skeletal debris differentially; coral skeletons are most susceptible, and the dense skeletons of red Algae are most resistant. Detritus feeders such as holothurians, worms, crustaceans, echinoids, and others are instrumental in churning up sediment and in reducing coarse- and medium-textured skeletal detritus to fine-textured lime-mud. GREENSMITH (1960) pointed out that in some Scottish limestones, a common feature of the fossiliferous carbonaceous varieties is the presence of early diagenetic microspheroidal and nonspheroidal pyrite which replaces the calcite shells and the carbonate of the matrix. He contended that their formation and the replacement reaction probably took place soon after burial because lenticular aggregates in the matrix sometimes show subsequent warping caused by compaction.

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Reef-flank skeletal limestones are amenable to diagenetic changes; noteworthy among these are introduction of sparry calcite cement, skeletal grain growth, pressure-solution, and emplacement of rim cement. For example, HAMBLETON (1962) noted that in some Missourian age rocks of New Mexico the reefflank deposits contain a profusion of gastropods, brachiopods, pelecypods, and cephalopods. A sparry calcite matrix cementing the fossil allochems suggested to him that strong local currents removed much of the microcrystalline calcite ooze. The dominant matrix material of the back-reef facies is microcrystalline calcite ooze and “reef milk” (very fine-grained, white and opaque microcrystalline calcite) derived from abrasion of the reef core and reef flank. It should be emphasized that in other occurrences of this “reef milk”, the microcrystalline material can be preserved in the matrix rather than being washed out, and is, therefore, subject to diagenetic changes, including dolomitization in the proper environments. Particulate material comprising newly deposited skeletal limestones does not react uniformly to diagenetic changes. As pointed out by LAPORTE (1962), the skeletons of many marine invertebrates consist of small masses of crystalline carbonate (calcite or aragonite) intimately intermixed with organic tissue. Details vary from one taxa to another. When the organic matter of skeletal material begins to decompose through oxidation or bacterial activity, the imbedded crystalline fraction is freed. Aragonite secreting corals, for example, produce upon total decomposition a type of sediment somewhat different than that produced by mollusks which may yield larger, hexagonal prisms. Diagenesis will affect one to a differenL degree than the other. Not to be overlooked in this assessment of diagenesis of skeletal limestones is the effect of Algae in secreting aragonite needles (LOWENSTAM, 1955). If a skeletal limestone consists in large measure of bioclastic algal detritus such as algal grains or lime clasts (not necessarily “algal dust”) described by WOLF(1962, 1965a, b) and the debris contains an abundance of aragonite needles, it becomes obvious that the path of diagenetic alteration will be different than in a brachiopod skeletal limestone, for example. Furthermore, presence of strontium in the aragonite may inhibit diagenesis in the algal bioclastic limestones. Another implication is “. . .that fossil calcilutites attributed to physicochemical precipitation or mechanically reduced skeletal carbonates may have been partially or largely derived from algally-secreted aragonite needles from ancestral Algae” (LOWENSTAM, 1955). Many of the criquinites of the geologic rock record have resulted from the diagenesis of coquinas of echinoderm debris (= “criquinas”). The most obvious diagenetic process is the formation of optically continuous calcite overgrowth on crinoid or other echinoderm fragments, particularly the ossicles. Individual plates of modern echinoderm skeletons are made of optically oriented calcite crystals containing large interstices which become solid single crystals after death. The overgrowth is a continuation of this crystal as described in earlier sections. According to BATHURST (1958), this overgrowth can form by filling pore space or by

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replacing the lime-mud surrounding the crinoid fragments. He called the pore-filling overgrowth “rim cement” and the replacement overgrowth “syntaxial rims” (Fig.2). LUCIA(1962) made a study of diagenetic effects in a crinoidal sediment in Devonian rocks of Texas, and in adhering to the usage of Bathurst, stated that the textural relationships between lime-mud and calcite overgrowth suggest that rim cementation is the dominant process in diagenesis. Of particular significance in Lucia’s studies is a consideration of the effect of dolomitization of the crinoidal sediment. He noted that the original character of the sediment which was dolomitized can be reconstructed by noting how the crinoidal material was replaced; these two mechanisms were suggested: ( I ) a single crystal of dolomite in optical continuity with the single calcite crystal of the original crinoid fragment, a process referred to as pseudomorphic replacement, and (2) dolomite crystals not in optical continuity with the calcite of the original crinoid fragment, a process referred to as impingement. The most commonly observed mechanism in Lucia’s studies is pseudomorphic replacement, and he noted all stages from partial pseudomorphic replacement to complete pseudomorphic replacement with none of the original calcite left. Furthermore, he pointed out that the tendency for dolomite to replace singlecrystal crinoid fragments with single dolomite crystals of the same crystallographic orientation suggests that it is difficult for dolomite to nucleate within a solid calcite crystal. The research of Lucia, which appears to be borne out by studies of many other petrographers, suggests that the sequence of dolomitization of crinoidal sediment (as observed in thin-sections) proceeds from dolomitization of the intercrinoid areas to dolomitization of the crinoid fragments. Lucia indicated that none of his thin-sections showed any dolomitization of the crinoid fragments unless the intercrinoid areas were entirely dolomite, with the exception of the small amount of impingement on their edges by the external doIomite crystals. Pseudomorphic replacement of crinoid fragments appears to take place mostly after the formation of internally impinging dolomite crystals. Lucia found no case in which the dolomite was composed solely of crinoid fragment pseudomorphs. In the rocks which he studied, the evidence proved that dolomitization occurred after rim cementation. If any of the dolomites had been composed solely of crinoid fragments and rim cement at the time of dolomitization, they would appear as dolomites composed essentially of crinoid pseudomorphs. LUCIA(1962) stated that: “The presence of randomly oriented 0.1-mm dolomite rhombs between the crinoid fragments therefore discredits the argument that the intercrinoid areas had been filled with rim cement, and it implies that the intercrinoid areas were filled with a finer matrix.” These arguments can be applied equally to other limestones of this category (= bioaccumulated skeletal detritus) in which bioclastic material consists of larger clasts in a matrix of smaller comminuted material. Non-dolomitization diagenetic effects will include crystallization (as well as recrystallization) of the matrix material first, followed by crystalline overgrowth (with or without optical continuity on the larger clasts). If lime-mud is also present, however, and the sed-

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iment is modally tripartite (i.e., consists of larger skeletal clasts such as crinoid ossicles, with a matrix of smaller skeletal debris, and impalpable lime-mud), the diagenesis may in some circumstances affect first the finest grade size material and then the larger particles. Certain clay minerals, and some clay-size particles, in the lime-mud may inhibit dolomitization there, but permit diagenesis to proceed directly to the matrix material and finally to the ossicles or other skeletal elements. Furthermore, leaching of the lime-mud may occur after rim cementation, thereby indicating that the interparticle lime-mud remained permeable to water. LUCIA (1962) stated this thusly: “Where interparticlelime-mud was present, it inhibited the development of the calcite overgrowth and was available for selective leaching to form the visible porosity. The leaching process was not as effective where the limemud was supporting the load as where the crinoid fragments were supporting it.” Consequently, skeletal grain-supported limes would differ in diagenetic effects from those that are mud-supported. The amount of porosity that develops during (or through) dolomitization may be related to the ratio of mud to crinoids or other echinodermal material, that is, the sediments containing the most echinoderm bioclastic material have the highest resultant porosity (see LUCIA,1962). In his studies of the Mississippian carbonate deposits of the Ozarks, MOORE (1957) pointed out that diagenetic effects were not limited merely to development of interlocking grains and to infiltration of fine calcareous mud, but were largely accomplished by precipitation of crystalline calcite out of solution. He demonstrated that none of the edges of crinoidal and other grains indicate the effect of solution, and thus concluded that most, if not all, of the cementing calcite was derived from the interstitial waters and not from the grains themselves. Secondary calcite was observed to occur as an approximately equigranular mosaic which lacks crystallographic continuity with adjacent crystalline echinoderm fragments. MOORB(1957) added: “Lithification has been effected by compaction and calcite welding, not by recrystallization, although some secondary crystalline calcite is identifiable in various rock samples.”

Lithodastic (=detrital) limestone Rocks not classified with skeletal types or with micrograined micrites can be termed lithoclastic (= “detrital” of LEIGHTON and PENDBXTER, 1962; “intraclasts” and “calclithite” fragments of FOLK,1959; “limeclasts” of WOLF, 1963b, 1965b;) if the components are of calcareous composition and have been worn or reduced by attrition to yield a clastic texture. Some sediments that obviously formed by aggregation have been added to this category by FOLK(1959), for example. Many calcarenites (particularly nonskeletal types) are to be classified in this group, and certainly many of the calcirudites and dolorudites are included here. During the various processes of diagenesis, newly deposited sediment of this large group is converted to pre-lithified and juxta-lithified equivalents by introduction, cementation and compaction of interstitial material, formation of crystalline material during authi-

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genesis, development of coated grains and overgrowths of allogenic sediment, and possibly near-complete to complete recrystallization leaving only vestiges (= relics or “ghosts”) of skeletal and/or litho-detrital material. Lithoclastic, detrital, or limeclast limestones have been termed mechanical limestones by some workers. ANDRICHUK (1958) studied Late Devonian sedimentary carbonate rocks of central Alberta, Canada, and pointed out that the calcarenites, calcisiltites, and calcilutites were formed primarily by two main processes, as follows: (I)mechanical disintegration of organic skeletal material and redeposition as bioclastic limestone at, near, or a considerable distance from the original site of organic growth; and (2) chemical or biochemical precipitation of calcium carbonate in quiet or agitated waters in association with, or separate from, sites of active organic growth. He contrasted the two types, and compared the latter variety with the baharnites of BEALES(1958), which are present-day deposits of the interior areas of the Bahama Banks (or ancient counterparts) and which are considered to consist predominantly of precipitated material that has aggregated into granules and composite grains (see also ILLING,1954). ANDRICHUK (1960) termed some of the calcarenites “pseudo-oolites”, and indicated that: “pelletoid or pseudo-oolitic calcarenites and calcilutites may have formed by precipitation in a slightly supersaline environment in the interior of a bank as compared with the more normal salinity of waters in which bioclastic limestones were deposited” (cf. ILLING, 1954; BEALES,1956). Among the diageneticdolomites which ANDRICHUK (1960) recognized are those that have microsucrosic to coarse textures; he believed that diagenetic dolomitization of calcisiltites and calcarenites (whether of bioclastic or lithoclasticorigin) accounts for these varieties. He stressed their significance by stating (ANDRICHUK, 1960): “The coarser dolomites with crystal sizes greater than 1/16 mm are considered to be of secondary origin where dolomitization occurred penecontemporaneously with deposition or at any time thereafter. These dolomites comprise the potential petroleum reservoirs.” In discussing diagenetic effects of carbonate rocks of Mississippian age in the Lisbon area of the Paradox basin of the Four Comers area, BAARS(1962) stated: “There, diagenesis has greatly increased the reservoir potential because of the solution of crinoid columnals. Diagenesis is of primary importanceto petroleum geologistsbecause of this close relationship with porosity.’’ BATHURST (1959b) studied diagenetic effects in Mississippian calcilutitesand pseudobreccias in limestones of England and Wales, and indicated that in any limestone the matrix is an accumulation of one or more types of grain mosaic. He noted the presence of three dominant mosaics, as follows: (I) granular cement and drusy mosaic, (2) rim-cemented single crystals, and (3) grain growth mosaic. He stated: “The calcilutites in their simplest form are rim-cemented carbonate muds or silts. Commonly, however, the original sediment was composed of aggregates of mud or silt, either faecal pellets or “grains” similar to ILLING’s (1954) Bahaman sands.” He also noted that in some limestones grain growth mosaic is common

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and forms the pseudobreccias where the “fragments” are masses of grain growth mosaic which lie in a “matrix” of less altered limestone. BATHURST(1959b) also defined a mud aggregate “. . . as any aggregate of mud grains, usually having the size of a sand or silt particle, which has been mechanically deposited. Initially the aggregatemay have been a faecal pellet (EARDLEY, 1938;ILLING,1954), or a rounded, sub-spherical aggregate of mud grains cemented originally by aragonite with no signs of organic control (as ILLING’S, Bahaman sands, 1954, et seq., whichlithify to yield the Bahamites of BEALES,1958),or afragment of algal precipitate (WOOD, 1941; and EPSTEIN,1957; WOLF, GEORGE,1954, 1956; LOWENSTAM, 1955; LOWENSTAM 1965a, b), or a spherical or ovoid growth form of a calcareous alga (ANDERSON, 1950).” A review of the diagenetic processes of chemical deposition, solution transfer, and grain growth can be found in the excellent articles by BATHURST (1958, 1959b). Detrital, lithoclastic or limeclastic limestones, which contain larger clasts embedded in a matrix of finer detritus, commonly display variation in diagenetic effects, particularly those of dolomitization. Crystallinity of the matrix of calcarenites and calcirudites normally is coarser than that of the grains. Dolomitization appears to select the matrix in preference to the grains which may remain unaltered. BEALES(1953) observed this effect in studying dolomitic mottling of Devonian limestones of Alberta, Canada, and considered that dolomitization took place at a time when the grains were still embedded in relatively porous mud. He stated: “The Palliser formation, laid down as limestone that was possibly magnesian, was subsequently altered to dolomitic limestone at an early stage in diagenesis. Secondary alteration and recrystallization produced the dolomitic mottling now so conspicuous in the rock.” It was his contention that dolomitization began in the more susceptible centers, triggered further diffusion, and permitted dolomitizing solutions to spread. Dolomitization, he discovered, in the lower beds was localized along certain bedding laminae and spread irregularly from them; higher in the succession “worm burrows” and “Algae” were most affected. Many calcilutites have apparently been diagenetically altered to a microcrystalline mosaic of interlocking anhedral crystals, from about 5 to 20p in average crystal size (= microsparite, Table 111). Calcarenitescan alter to similarly-appearing rock, the crystalline mosaic of which is coarser textured (= sparite, Table 111)than that of altered calcilutites. In other words, detrital (lithoclastic) rocks can under the effects of diagenesis become finer textured but will have an interlocking anhedral mosaic as suggested earlier. These effects, however, are common (but not limited) to more or less equigranular calcilutites and calcarenites. Diagenetic effects on these clastic carbonates, in which relatively larger grains are embedded in a “matrix” of finer texture, differ in that the matrix normally crystallizes to subhedral and euhedralforms that are not necessarily interlocked. Impingement and suturing may occur, however. Grain growth, authigenic overgrowth (including optical continuity

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with originalclasts), rim cement, pressure-solution, and syntaxial rims are common diagenetic effects. CHANDA(1963) studied the effects of cementation and diagenesis of the Lameta Beds (Turonian) of Lametaghat, M. P., India, and noted that silicification starts as advancing fronts from the peripheries of the detrital grains and continues to grow at the expense of interstitial calcite, in the case of calcareous sandstones. In sandy limestones, however, silicification was not as extensive or as systematic. Chanda pointed out that the Lameta limestones are sandy microsparites, where the microsparites did not result from primary precipitation but have formed by aggrading recrystallization of micritic calcite. Diagenesis of these limestones, he noted, involved both selective and what he termed “perversive recrystallization”. The microspars which developed on the floating clastic grains are always water-clear, whereas areas free of clastics are in places occupied by coarsely crystalline anhedral cloudy microspars. Perhaps many lithoclastic limestones have experienced various degrees of diagenesis, not necessarily in uniform process, or as a continuum. FOLK (1959)and CHANDA(1963) offered certain criteria as evidences of recrystallization of microcrystallinecalcite; because they are applicable to many lithoclastic limestones (as well as some micrites and skeletal limestones), they are repeated here: (1) the looseness of packing of clastic grains requires aggrading recrystallization of microcrystalline calcite; (2) uniformity of size of the microspars of calcite; (3) patches of microspars grading by continual decrease of grain size into areas of normal microcrystalline ooze; ( 4 ) microspars have a radial fibrous form oriented perpendicular to the surface of clastic particles as an outwardly advancing aureole of recrystallization; and (5) relic patches of microcrystalline calcite and partially warping quartz grains, embedded in a mass of mosaic of microspars.

Pelletal and coated grain limestone Limestones herein classed as pelletal and coated grain types include the faecal pellet and other pelletal limestones, and various oolitic and pisolitic types. Although many of these are intimately associated with reefal limestones on the one hand, and with detrital (lithoclastic) limestones on the other, they are treated separately here because diagenesis does not necessarily affect them as it may the other two. They may form in extremely shallow to moderately shallow waters, and normally develop best in agitated waters although some subtypes may form in quiet water environments. Many oolitic limestones develop in or adjacent to the reef complex where they are subject to movement by trans-reef currents as well as to surf-surge and wave activity. Diagenetic effects range from simple boring by Algae to complete obliteration of primary features during dolomitization, as well as complete silicificationwith faithful reproduction of internal details. Some of the particulate material in these limestones displays excellent concentric and/or radial features, whereas others are pseudo-oolites, sub-round to sub-spher-

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ical pellets, and superficial coated grains. Each reacts quite differently to diagenesis. Pelletal (= pelletoid) and pseudo-oolitic limestones may undergo a certain spectrum of diagenetic effects yielding a final product not too dissimilar to certain calcarenites; in fact, petrographic distinction may be difficult in some instances. Some may, in verity, resemble the “mud-aggregate” limestones of BATHURST (1959b). One of the first effects of diagenesis on oolitic limestones is the development of water-clear to semi-transparent sparry calcite. Subsequently this orthosparite can be altered to an interlocking anhedral to subhedral mosaic of pseudosparite. This mosaic can, by impingement, invade oolite envelopes (i.e., peripheral rings) and ultimately may take over all the rock. A “negative” relic of the oolite may remain, however, and display a dusty ring around the unaltered to slightly altered core or nucleus of the ovoid. Petrographers are particularly concerned with dolomitization of pelletal and coated grain (= oolitic) limestones, if for no other reason than to evaluate reservoir potentialities. When carbonate sands composed of oolites and/or pisolites retain their original interparticle porosity, they are excellent reservoir rocks. If lime-mud, sparite, or other material binds and cements the particles, the resultant limestone may be devoid of effective porosity. Some oolite units have been subjected to leaching, and although the “rind” remains relatively unaffected, the nuclei are removed by solutions and a rock having considerable porosity results. BEALES (1958) made a careful study of various ancient carbonate rocks of Canada, and compared them to Bahaman type limestones. Aragonite is more susceptible to alteration than calcite, and so he pointed out that present-day Bahaman deposits which are aragonitic, are subject to recrystallization. It was his contention that oolites show varying susceptibility to dolomitization; the matrix is most readily altered, followed by bahamite cores, oolitic envelopes, and coarsely crystalline skeletal cores, in that order. Regarding these bahamites and oolites, BEALES(1958) had this to say: “Direct precipitation of calcium carbonate from sea water resulting in the formation of bahamites, or under more active water conditions of oolites, has probably formed very considerablethicknesses of limestone that occur throughout the geologic column from Late Precambrian to Recent time.” Beales argued that a theory of aragonite needle agglutination for oolite growth is more satisfactory than one of direct precipitation. If true, dolomitization of such oolites may proceed with rapidity in some instances. It is to be remembered, however, that the conversion of aragonite to calcite may be a slow process under certain conditions. If protected by a covering of stable calcite, aragonite may be stable for a long period before inverting to calcite. In areas of incipient dolomitization, oolites sometimes have dolomite rhombs concentrated in their nuclei (BROWN,1959). If oolites are encased in a calcarenitic matrix, then perhaps during recrystallization of this matrix material the periphery of the aragonitic oolites was converted to calcite. According to BROWN(1959),

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when this protective layer was formed, the aragonite in the centers of the oolites remained unaltered until the oolites were fractured during compaction. Metastable aragonite material in the nuclei of the oolites then constituted natural foci for dolomitization. Brown’s studies were concerned with diagenesis of a Late Cambrian oolitic limestone in Montana and Wyoming, but the principles are nonetheless worthy of consideration in petrographic investigations of other diagenetically altered oolitic limestones, EDIE(1958) made rather intensive studies of sedimentation of the Mississippian Mission Canyon and Charles Formations in southeastern Saskatchewan,Canada, concluding that these four environmental types are represented: ( I ) basin, (2) open marine shelf, (3) barrier bank, and (4) lagoon. Pisolitic, oolitic, and pseudooolitic (= pellet) limestones characterized the barrier banks. He observed that the pseudo-oolites are calcareous pellets composed of cryptocrystalline material and are similar in size to oolites but lack concentric layers. He believed that some of these pellets are chemical precipitates formed on the sea floor under moderately agitated water conditions similar to the calcareous sands of the Bahama Banks described by ILLING (1954); but that some, if not most, of the pseudo-oolitic limestones may be largely of algal origin, and possibly represent both accretionary algal grains and “bioclastic” material formed by the fragmentation of algal colonies in areas of intense wave action. WOLF(1963a, 1965a, b) arrived at similar conclusions. If dolomitization affects sediments of the type described by EDIE (1958), a rock having an earthy to sucrosic texture would likely result. It would have intercrystalline and interparticle porosity and commonly would contain dolomitized positive relics of fossils or fossil debris as well as relic oolites containing dolorhombs in the nuclei. The petrographic studies of the Oil-Shale Group limestones of West Lothian and southern Fifeshire, Scotland, by GREENSMITH (1960) are quite informative. Oolitic texture is very common in all limestones of the group, and the carbonate of the ooliths appears to be an iron-rich dolomite (o= 1.679-1.683), commonly set in a matrix of similar nature. Partial breakdown of this mineral to limonite during weathering gives many of the beds a distinctive light brown surface color. Coarse calcite euhedra are not common in the matrix, but internal pressure-solution effects have produced micro-stylolites. Evidence for dolomitization is almost negligible and is expressed in the form of irregularly shaped, small transgressive vugs up to 1.2 * 0.10 mm in size. These sporadic cavities have a lining of coarse subhedral dolomite and often have a subsequent infill of a kaolinite-like mineral. In the words of GREENSMITH (1960): “In thin-section the coarse clear dolomite is seen to grade into the Fe-rich dolomite grains of the matrix which suggests that it represents a localized solution and reprecipitation effect hardly akin to the wholesale metasomatic changes associated with true dolomitization.” It was noted by GREENSMITH (1960) that intimately intermingled with the ooliths in many of the limestones are similarly shaped and sized bodies to which the term “oolitoid” was applied. They

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lack the internal structure normally found in the oolites and consist of a finegrained aggregate of iron-rich dolomite. Greensmith ruled out an origin due to recrystallization of oolites, as well as one associated with faecal pellets, but rather considered them to represent cross-sections of tubes that presumably resulted from activities of organisms such as worms. Seemingly, the presence of pelletal, oolitic, pisolitic, algal circumcrusted (WOLF,1965b), and other noncoated, coated, and superficially coated grains in limestones, even to the point of comprising most of the rock, has been a point of dissension among some petrographers concerning diagenetic changes. Some regard one type as the alteration product of another. Many geologists regard each type as a distinctive sedimentary product, penecontemporaneous with sedimentation of the host rock. Terms like “spherulite”, “axiolite”, “ooloid”, “oolith”, etc. have been coined to define some of these coated grains. It is important, however, to make a distinction between a “superficial oolite” in which most of the particulate material consists of “superficial ooliths” with only thin external oolitic layers, and a true oolite which is dominantly composed of “ooliths” with well developed concentric structure. Most of the modern Bahaman oolitic sands are composed of superficial ooliths (ILLING,1954). In the course of geological investigations of sedimentation in the Bimini, British West Indies region, KORNICKER and PURDY(1957) discovered an area in the Bimini lagoon in which at least 90 % of the sediment is composed of a single type of faecal pellets. The delicate faecal pellets were preserved because of extremely low agitation and current activity, scarcity of scavengers, and bacteriological precipitation of aragonite within the pellets. Of real significance, however, is the fact that during emergence at low spring tides desiccation results in permanent hardening of the pellets. The studies of Kornicker and Purdy, though suggestive, point up the importance of bacteriological precipitation of aragonite, and hardening through desiccation, in early diagenesis of various carbonate sediments. One can readily appreciate the importance of these early diagenetic changes leading to various types of limestones. NEWELLand RIGBY(1957) pointed out that faecal pellets, ooliths, ovoids, and a variety of grains termed “lumps” by ILLING (1954) make up most of the bottom sands over great areas of the Bahama Banks. Some of the friable aggregates are bound together by algal mucus, others are held together by calcium carbonate cement. ILLING (1954) has identified these particles in all stages of cementation. As indicated by NEWELL and RIGBY(1957): “They become h e r by precipitation of aragonite cement within the aggregate and as they are rolled about they lose their irregular shape, and the final grain, composed chiefly of cryptocrystalline aragonite, shows but little evidence of the original composite nature. Recrystallization of the fine detrital constituents takes place concurrently with cementation, quickly destroying the original texture.” NEWELLand RIGBY(1957) stated that THORP(1936) probably was the first to record the large quantity of faecal pellets in Bahamian sands and muds around Andros Island. When fresh, the pellets are

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friable aggregates of fine detritus held together by mucus. They very soon become firmly bound together by aragonite cement, deposited perhaps through bacterial activity (ILLING,1954). They become finely crystalline, however, as a result of crystallization of the finest material. RUSNAK(1960) studied Recent oolites forming in the hypersaline environment of the Laguna Madre along the southern Texas coast and considered that the rate of carbonate precipitation and mechanical reorientation may very well be the controlling factors of primary crystalline orientation within oolites. He indicated that with rapid precipitation, individual needles may not assume a preferential orientation on the nucleus and thus will result in unoriented carbonate deposition. But, with slower rate of precipitation they may become oriented radially, as in artificially precipitated oolitesor spherulites(cf. MONAGHAN and LYTLE,1956; LALOU, 1957). It has been contended that where precipitation rate is very slow, crystallites become attached tangentially to the nucleus by rolling or agitation, or even become bent by mechanical rubbing. RUSNAK (1960) stated: “. ..tangentially oriented oolite layers must be subjected to relatively high crystalline strain during the bending process. These strained crystallites may thus be more susceptible to recrystallization by diagenetic processes in response to a release of acquired strain.” Published reports on petrography of some limestones contain references to what is known as “granular” limestones. Many of these are not in the real sense of the word “grained” as pertains to lithoclastic or detrital limestones, but actually represent a stage of diagenesis of what were originally pelletal, oolitic, pisolitic or superficially coated granular limestones. Some such limestones are well sorted, and do, it is true, contain a significant (but not dominant) proportion of crinoidal and algal material formed by attrition. Oolitic and related coated-grain material that formed in current-agitated waters is a dominant component. Perhaps floating, calcareous, planktonic Algae (Coccolithophoridae) contributed to some of the finely-divided, even microgranular, matrix material in which the oolites and ovoid bodies are embedded. If certain skeletalelementscomprise the spacebetween packed granules, porosity and permeability values may be high (THOMAS and GLAISTER, 1960). Limestone of this category is particularly susceptible to diagenetic changes, and when dolomitized gives rise to a rock having a crystalline-granular texture, sometimes with a reduced porosity and permeability. The generalized term “recrystallization” has often been applied to diagenetically altered oolitic, pisolitic, and pelletal limestones, without specific reference to details of the alteration. BATHURST(1958) recognized two types of cement, for example, depending on whether the cementing material grew into void space or replaced the carbonate mud. Cement which develops into interparticle voids may be optically continuous on single crystal particles (e.g., crinoid ossicle), and thus be termed rim cement. This may be difficult to determine petrographically on some oolites and pisolites, however. The cement may consist of small crystals commonly oriented perpendicular to the void walls, and give rise to fibrous and/or drusy

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cement (Fig.2). This should be looked for in coated-grain limestones, particularly in those where the matrix has been dolomitized and created additional space in which the fibrous and/or drusy cement forms in the next step or phase of diagenesis. Furthermore, if two generations of dolomitization of the matrix material are represented, continued development of cement may preferentially form a coarse mosaic. Limestones of the type herein discussed may have lime-mud also present, and if the cement occupies space previously taken up by the carbonate mud (but which has been washed out, or leached away), and continues to enlarge pellets, granules, oolites, pisolites, etc., then grain growth may be instituted in some cases. Again, such phenomena are to be looked for in thin-sections. Reefal limestone Diagenesis of reef limestones, bioherms, biostromes, and comparable rocks built by wave-resisting organisms in the marine and lacustrine environments normally includes introduction of interstitial material, as well as many of those changes indicated for limestones discussed above. Because reefal limestones have already constructed a hard and relatively compact framework, diagenesis may be somewhat different in contrast to other limestone types. It is also true that some finely comminuted material (i.e., calcilutite,calcisiltite, calcarenite) may still remain in or near the reef framework after abrasion and disintegration of some of the reef rock, and this material is particularly susceptible to diagenetic changes. LOWENSTAM (1955) has stated that some of the calcilutites attributed to physicochemical precipitation may have formed by breakdown of calcareous Algae, particularly the poorly calcified forms. In writing about the “white reef” in certain Devonian reefs of Canada, BELYEA(1955) stated that much of the reef mass consists of fine-grained comminuted organic debris, and that much of it is white dense limestone probably formed in large measure by lime-trapping Algae. Various references have been made to the “aphanitic” reef limestones (HADDING 1941,1950;HENSON, 1950; WENGERD, 1951; NEWELL et al., 1953; WOLF,1962, 1965a, b, c). Possibly some of this fine-textured aphanic limestone within or near reef cores represents chemical or biochemical precipitates, and products of recrystallization that brought about loss of the original texture as pointed out earlier. Various workers have investigated dolomitized Devonian reefs in Alberta, Canada; ANDRICHUK (1958) stated: “. . .the threshold between a dolomitizing and non-dolomitizing environment appears to be very subtle and sensitive, and only a slight change in one of the factors affecting dolomitization may be sufficient to promote complete dolomitization in a limestone province. ..less intense agitation and aeration and a less oxidizing environment would be more suitable for penecontemporaneous dolomitization. . .” A sufficient body of factual information is available concerning direct precipitation of calcite from marine and lacustrine waters, that it need not be reviewed here. Thus, the source of the calcium carbonate precipitated early in primary pores

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and interstices of reef limestone can be easily accounted for. NEWELL (1955) stated: “Surface waters, which are supersaturated with calcium carbonate, are warmed in the daytime over shallow reef flats several degrees above the waters of the open sea, and the solubility of the carbonate is further reduced by photosynthetic activity of reef plants. During ebb tides these reef-flat waters form a hydrostatic head a few inches above the surrounding sea . . . Part of this water escapes seaward by sinking through the myriads of pores which riddle the reef flat, and calcium carbonate probably is deposited in transit.” In the reefs which Newel1 studied there is an abundance (locally as much as one half of the rock mass) of fibrous calcite that is deposited over the surfaces of the frame builders. The prismatic structure of the calcite is radial with respect to the depositional surfaces. NBWELL (1955) noted that, in practically every example, deposition of the fibrous calcite clearly occurred in primary voids of the reef frame at a time when prevailing conditions prevented simultaneous accumulation of detrital sediment. Calcite was deposited directly from solution, and any remaining voids were filled by detritus. Identical features have been studied in detail by WOLF(196%). It may be argued by some workers that the above-mentioned processes are not to be classed as diagenetic, but are syngenetic. Still others may favor a term such as “syndiagenetic”, but this is to a certain degree only a play on words. NEWELL et al. (1953) stated: “Diagenesis, as illustrated by the Capitan reef complex, is chiefly the result of interactions between sediments and the fluids contained within them. Other factors, largely responsible for these reactions but also partly contributing independently to diagenesis, are biotic activity within the sediments, compaction, and the migration of ions and fluids.” These workers indicated that the changes which take place below the temperature and pressure levels of metamorphism s. str. are considered to constitute the processes of diagenesis. Furthermore, because organic frame builders are in a sense lithified prior to most of the diagenetic processes, NEWELLet al. (1953) indicated that lithification is only one result of the processes which bring about post-depositional changes; it is too gradual a process, they pointed out, to restrict diagenesis (insofar as reefs are concerned, at least) to those changes which affect a sediment after deposition and up to, but not beyond, lithification (see discussion in the Introduction). It is beyond the scope of this chapter to review all facets of diagenesis of the Capitan reef complex and associated rocks, and so the interested reader is referred to NEWELL et al. (1953, chapter 6). Petrologists studying the fabric of some reefs have called attention to the presence of “reef tufa”, a particular variety of which is called “stromatactis” (WOLF, 196%). PARKINSON (1957) studied Lower Carboniferous reefs in northern England, and stated: “The most characteristicfeature of the reefs, apart from the anomalous dips and the non-bedded nature of the calcite mudstone which comprises much of the rock, is the abundance of fibrous calcite, the “reef tufa”. He made reference to occurrence of “reef tufa” in Permian reefs of western Texas, as reported by NEWELL (1955), but had this to say of the English reefs: “When it is recognized that algal

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remains are readily obliterated by recrystallization or disintegration, it seems possible that such organisms may have been of some importance in the English reefs. In this connection it is noteworthy that calcite muds, according to WOOD(1941), might have originated from disintegrated algal deposits.” BLACK(1954) referred to reef-like structures in various parts of the world where evidence of frame-building organismsis slight, but in which calcite mudstones are prominent. Furthermore, he suggested that recrystallization of algal skeletons could give rise to calcite mudstones of the knoll-reefs of England. From the foregoing, it is apparent that petrologists and petrographers must exercise extreme caution in assessing diagenetic changes of reefal limestones. For example, one worker may term interstitial sparry calcite, that fills voids in frame building organisms, recrystallized calcite, and it actually may be open-space sparite. Evidence for the latter conclusion is based largely on the work of BATHURST (1958) and was discussed in detail in earlier sections. Lime-mud, commonly detrital rather than directly precipitated crystalline cement, usually accompanies sparry calcite or granular cement development in voids of reefal limestones. Some of this lime-mud undoubtedly is the detrital infilling of remaining pore space as discussed and some may actually represent aragonite needles in detail by WOLF(1965~)~ secreted by Algae in an organic framework (LOWENSTAM, 1955); or it is an algal slimeformed on or near algal plants: CaCOs precipitated as the plants extracted carbon dioxide from immediately adjacent sea water (PRAY, 1958). The latter conclusion is worthy of further investigation, particularly for the bearing it may have on dolomitization of material within the reef between the originally formed organic framework (see WOLF,196%). SCHLANGER (1957) pointed out that during deep drilling operations on Eniwetok Atoll a dolomitized core was recovered from a depth of 4,078-4,lOO ft., and subsequently was determined to be of Eocene age. Dolomite in much of the core is restricted to rod-shaped segments of articulate coralline Algae, identified as Corallina. SCHLANGER (1957) stated: “The dolomite crystals are definitely restricted to the Algae and their growth seems to be controlled by the shape of the rod. The area near the axis of the rod is occupied by a finegrained mosaic of anhedral dolomite that grades outward into coarser, more euhedral crystals.” He further noted presence of fine detrital filling (finely comminuted algal particles) in interseptal areas of unaffected corals in the core, and termed this “paste” fill which probably served as centers for dolomitization. Schlanger argued that there may be regions within the algal fragment in which the Mg-ion concentration is in excess of the “average” for the entire fragment. He further stressed the findings of CHAW(1954) that considerable variation exists within a single algal colony, possibly due to seasonal temperature variation, or influence of metabolism of the Algae in inducing short-term high uptake of magnesium. Thus the Alga, as it grows, may contain disseminated dolomite nuclei whose size limits are too small to be detected even by X-rays. With time the dolomite nuclei, which evidently are more stable than the surrounding Ca-Mg solid solution

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under existing conditions, enlarge by diffusion of the originally adsorbed Mg ions in the structure and by addition of Mg ions from sea water (SCHLANGER, 1957). SCHWARZACHER (1961) studied the petrology and structure of some Lower Carboniferous reefs in northwestern Ireland, and mapped “knoll-reefs” in detail. He noted that the reef limestone consists of clotted fine-grained calcareous mud (which he termed bahamite) that contains larger mud pebbles. Bryozoans were the only frame builders of significance. It was pointed out that during early diagenesis a cavity system was formed, probably due to sliding movements on the reef talus. This was soon filled with calcite and dolomite crystals, and at a stage when the reef was still a mound on the sea floor. Lithification set in at a later stage, and led to a preferred orientation of calcite grains. The relatively large volume of the cavities was filled first by calcite, but as pointed out by Schwarzacher: “Almost all cavities show some dolomite; if the cavities are lined with fibrous calcite then the dolomite crystallizes later; if fibrous calcite is missing then the dolomite may form the first lining on the calcite mudstone wall. Most dolomite occurs in well defined rhomb-shaped crystals whereby a definite growth relation of the crystals to the wall exists. . . . Most commonly, the “c” axis is parallel with, and the longest diagonal of the rhomb is at right angles to, the wall.” WOLF(1965a, c) has presented details on the Devonian algal reef-knolls of the Nubrigyn complex. Not to be overlooked in any discussion of diagenesis of reefal limestones are the results of collapse attendant to removal of soluble materials such as evaporites in or adjacent to the reef. Some of these “evaporite-solution breccias” can be confused with “reef-edge breccias” (GREINER, 1956) and only through detailed petrologic and petrographic studies can the correct assessment be placed on diagenesis. Both types, it is true, are subject to dolomitizing solutions that migrate “updip” from the basin adjacent to the reef tract. The chaotic jumble of the blocks comprising the “solution breccia”, however, do not resemble the rubble of frame-building organisms, which grade perceptibly into bioaccumulated calcarenitic material of the “reef-flank” or “reef-edge” breccias. As has been pointed out herein, the voids of these breccias can be filled partially or wholly by fibrous, drusy and/or granular sparry calcite. Should any space remain, “reef milk,” “algal dust”, or detrital, chemical, biochemical, or physicochemical carbonates may fill it. Complete to nearcomplete dolomitization of the entire rubble is not uncommon in some of the reeftalus material and collapse-breccias. Sparry limestone Petrographers have differences of opinion regarding origin of certain crystalline limestones and dolomites?particularly if lithologic association (e.g., reefs) does not contain suggestive data. To some workers, the formation of coarse sparry limestone is a function of metamorphism? and thus outside the realm of diagenesis. When a limestone having such a texture is found interbedded in a sequence of limestones and other sedimentary rocks which are not metamorphosed, however,

DIAGENESIS OF CARBONATE ROCKS

28 1

and some in verity are only slightly diagenetically altered, the evidence clearly suggests that diagenesis can create such a texture in limestones. Sparry calcite cement, particularly in cores of reefs or adjacent thereto, and as infilling of mollusk and brachiopod shells, is quite common and was treated in large measure in preceding pages. Units adjacent to micrite reef cores commonly are composed of skeletal detritus (including reef-talus breccia) that is cemented by sparite, In some reefs the coarsely crystalline material is composed of dolorhombs (dolosparite) and if of the open-space variety it may line cavities or void walls. Farther from the cores, however, an intermixing of sparry calcite-cemented skeletal detritus and sparite-in-calcarenitemay occur, and still farther out micrite may be found (CRONOBLE and MANKIN,1963). The energy factor is an important one in development of sparry calcite: if one disregards the reef cores, the progression away from the cores through sparite (i.e., sparry cemented talus) to interdigitated talus, calcarenitic material, and finally micrite indicates in most instances a decrease in the energy in the depositional environments (CRONOBLE and MANKIN, 1963). Grain growth and recrystallization, pressure-solution, syntaxial rim cementation, and drusy and fibrous sparite cementation may work not only independently, but also one in harmony with at least one of the other, to form sparry limestone from rock which was not originally a sparite. One example may be the encrinal limestone (= criquinite) in which sparry calcite ultimately develops, particularly under slight differential load-stress, engulfing the crinoid fragments. The end result may be authigenic overgrowth only; or it may continue to near-completerecrystallization (i.e., combination of grain growth, recrystallization and cementation, all aided by pressure-solution), with obliteration of all but “negative” relics of the echinoderm fragments. A similar process possibly operates on some faecal pellet limestones, and upon bahamites. Cements may completely engulf areas of fine carbonate mud, in particular the so-called “drewites” or algal-precipitated aragonite needles, with fine interparticle porosity. Coarse sparry calcite can result from this diagenetic process. If intercrinoidal voids are filled by sparry calcite which is enlarged by impingement or forms continuous overgrowths on monocrystalline carbonate fragments, such as crinoid ossicles, sparite will result. Should dolomitization be the process, complete obliteration of the original texture can ensue, or it can be “arrested“ in some stage, giving a mottled or even coarse sucrosic appearance to the rock; fossils and/or other particulate material may be “positive” (recognizable as to organic remains, or inorganic fragment), or they may be “negative” (= strongly suggestive of a vestige of a former fossil or other fragment, but proof lacking). A most important consideration in the process of dolomitization to form dolosparite is that in most instances the rhombs grow by replacement as opposed to growing as a cement. Exceptions are to be noted, however. If, for example, a lime-mud contains particulate skeletal (i.e., crinoidal) or detrital (i.e., fragmental

282

0.V. CHILINGAR, H. J. BISSELL A N D K. H. WOLF

limestone clasts) material “floating” or embedded in the ooze, the mud is preferentially replaced first, and the particles may follow and be dolomitized in the order of their susceptibility to dolomitization. Cements occupyingspaces previously occupied by carbonate mud may grow by replacement in optical continuity with a large host or by grain enlargement of the smaller particles to form a coarse mosaic of anhedral to euhedral crystals (MURRAY, 1960). Reduction in porosity by pressure-solution may be an important mechanism, yielding a limestone (or dolomitic rock or dolomite) that has an interlocking texture. If it is an interlocking mosaic of anhedra, in all likelihood it is a calcspar; whereas if it is composed of subhedral to euhedral grains, it may be a dolospar. It is important to note that diagenesis does not always operate to form larger crystals, and thus create sparry limestones. Disregarding for a moment the process of recrystallization, one should be cognizant of the strong possibility of some calcarenites to be transformed during diagenesis into a rock with silt-sized calcite crystals (= microsparite) and realize that the calcisiltite is a product of alteration and not a primary detrital (that is, lithoclastic) limestone (WARDLAW, 1962). The petrographer should also be aware that recrystallization of large calcite crystals, whether infilling of voids or as interstitial sparry cement in reef framework, does not always result in fine-grained textures. Grain growth can readily produce sparite; some thin-sections show strained, twinned calcite crystals that have untwinned, unstrained rims of calcite in optical continuity with one of the sets of twin lamellae. Certain bioclastic as well as lithoclastic limestones, particularly biocalcarenites and lithocalcarenites, are susceptible to pressure-solution that results in cementation by a sparry calcite, or ultimately, under proper conditions, by dolosparite. During the diagenetic processes, the larger fragments become etched and corroded around their boundaries and crystallized into subhedral or euhedral mosaic, or into sparite the discrete rhombs of which are in optical continuity with the host particles. Some grains invade other grains on contact points (TOWSE,1957). THOMAS and GLAISTER (1960) in discussingfacies and porosity relationshipsof certain Mississippian carbonate rocks of western Canada, stated: “With regard to the relation of dolomite development to textural features of original limestones, it has been observed that it preferentially occurs in open pores or in matrix (chalky, granular, and carbonate mud) material that surrounds the larger skeletal or nonskeletal grains. These larger grains are generally the last to show conversion to dolomite, Many skeletal fragments remain as calcite even when the remainder of the rock may be dolomite. The final type in this sequenceis a dolomite with fossil casts.” Among Pennsylvanian and Permian fusulinid-bearing limestones that have been dolomitized, one can commonly discover bank deposits and hinge-line accumulations of fusulinidcoquinitesin which interlocking subhedra of dolosparite are riddled with spindle-shaped cavities (molds of the Foraminifera). Interestingly enough,

DIAGENESIS OF CARBONATE ROCKS

283

when silicification precedes dolomitization (as it often does), the fusulinids are faithfully replaced even to details of cell wall, whereas the matrix material has been converted to dolosparite. PERKINS (1963) made a detailed study of the petrology of a Middle Devonian limestone in southeastern Indiana, and mapped different carbonate facies. He indicated that the interstitial sparry calcite of the pelsparite facies is considered to be granular cement and not recrystallizedmicrite; the evidence supporting this conclusion was taken from the work of BATHUR~T (1958). Perkins consideredthe cement to be “granular” where the host particle, usually multigranular, lies in a mosaic of cement. Where the host is a single crystal, such as a crinoid fragment, and the cement forms a single rim in lattice continuity with it, the cement is called “rim cement”, following the usage of BATHURST (1958). Perkins also observed that sparry limestones commonly are not dolomitized, whereas micritic rocks were more susceptible to dolomitization due to their greater porosity and the greater surface area of the minute micrite grains. This again, bears out a well-established principle observed by many petrographers that calcspar resists dolomitization, whereas micrite and matrix (finely comminuted biocalcisiltite and biocalcarenite, also lithocalcisiltite and lithocalcarenite) may readily respond to dolomitizing solutions high in magnesium content. Possibly, some “catalysts” initiate nuclei of dolomitizing centers. Sparry calcite possibly can also form in various cavities, for example in animal burrows, gas-bubble pockets, worm-burrows, etc., yielding “eyes”. Thus an ultimate lithification of the sediment can result in a “birdseye” limestone. The origin of “birdseyes” is discussed in detail by WOLF (196%). It is possible for dolorhombs to form in association with evaporite suites of sediments; and with continued growth, a dolosparite can result. MILLER(1961) noted clear, pale pink dolomite rhombohedra up to 0.14 mm long, disseminated with subhedral aragonite and calcite, in sludge dumps of salt extraction processing plants at Inagua, Bahamas. The rhombohedra are intricately associated with the other evaporite minerals; Miller suggested that in all probability the dolorhombs formed in minute voids where magnesium-rich brine accumulated or filtered through the waste sludge. Such a process conceivably could have operated in the geologic past within certain evaporite suites, and by so-called ‘Wter-pressing” could have removed disseminated dolomite from one stratum only to concentrate it in another as dolosparite. The process, so it seems, would not of necessity be limited to dolomite but could have worked equally well with calcite. Calcspariteand dolosparitepossibly are morecommon in limestonesanddolomitized limestones of the geologic record than published literature would suggest. Extensive cementation has been noted in young carbonate deposits which are subaerially exposed or are in the zone of meteoric waters along the Florida coast (GINSBURG, 1957). The Late Pleistocene Miami Oolite is thoroughly cemented by calcite at the exposed surface and below the ground-water level. But where it is

284

G. V. CHILINGAR, H. J. BISSELL AND K. 13. WOLF

still in the marine environmentor abovetheground-water table, it is friable and poorly cemented. GINSBURG (1957) noted that the oolite has a clear mosaic and partially recrystallized ooliths. Sparry calcite as a term has been employed by STAUFFER (1962) ". .for that calcite which has been deposited from solution on a free surface" (orthosparite in Table 111). If the term is enlarged to embrace sparry dolomite as well, most of the criteria for recognition of sparite listed by Stauffer are applicable. For those dolosparites which have replaced other carbonate materials, criteria normally are readily available for recognition of the host rock. The following criteria, which are more or less directly indicative of sparry calcite, are taken from STAUFFER (1960): ( I ) crystals in contact with a once free surface, such as oolites or inside of shell chambers; (2) crystals in the upper part of a former cavity which was partly filled with more or less flat-topped detrital sediment; (3) an increase in crystal size away from the wall of an allochem; ( 4 ) a decrease in the number of crystals away from the wall; (5) preferred orientation of the optic axes of crystals normal to the wall; (6) preferred orientation of the longest diameters normal to the wall; and (7) plane boundaries between crystals. In addition, Stauffer listed eight more criteria that he considered suggestive of open-space sparry calcite. He also presented criteria that are indicative and criteria that are suggestive of recrystallization in calcite. The interested reader is referred to the comprehensive and detailed article by Stauffer for additional information bearing on the subject of sparry calcite and recrystallized calcite.

.

FORMATION OF CARBONATE CONCRETIONS DURING DIAGENESIS

The escape of C02 during diagenesis appears to be one of the main driving forces for the formation of carbonate concretions. This can be seen on examining the following system of equilibriums presented by STRAKHOV (1954; see also BISSFJLL and CHILINGAR, 1958):

CO&H2CO3~(Ca,Mg,Fe,Mn)[HC03]2~(Ca,Mg,Fe,Mn) coa (1)

(2)

(3) liquid phase

(1)

(4) solid phase

During the first and second stages of diagenesis (as explained by LARSEN and CHILINGAR, 1965), as a result of energetic bacterial activity, the amount of C02 in the interstitial waters is increasing. Thus, the above reaction goes to the right (from 2-2-3) and solid carbonates dissolve, resulting in higher alkalinity of interstitial waters. During the third stage of diagenesis, with decreasing amounts of C02, the reaction goes to the right (from 3-4) and especially close to the avenues (such as sandy layers) along which C02 can escape. As aragonite or calcite changes to dolomite, the avenues of increased porosity also enable C02 to escape and thus carbonates (CaC03, FeC03, MgCOs, M n C a , or their mixtures) can precipitate.

DIAGENESIS OF CARBONATE ROCKS

285

This also occurs in sandy layers inside clayey deposits. As a result of precipitation of carbonates and decrease in alkalinity of interstitial waters, the bicarbonates from adjoining clayey layers will move in to compensate for this created deficiency. As the new portions of COz escape, additional carbonates are precipitated. This precipitation commonly occurs along certain horizons and around certain centers, giving rise to series of concretions. Uniform precipitation gives rise to sandstones cemented with various carbonates. The carbonate concretions are also commonly found inside clayey deposits close to the ventilation avenues along which escape of COz can occur. Calcite concretions are found in highly calcareous sediments, whereas siderite concretions occur in sediments poor in calcareous material. The findings of VITAL’(1959) indicate that many calcite and siderite concretions are found in sediments having low COz content (<1 %, and rarely 2-3 %). Possibly this is due to secondary leaching-out of carbonates from rocks studied by Vital’. Wide variations in physicochemical conditions within the sediment also account for considerable redistribution of substances during the third stage of diagenesis. For example, if a portion of sediment is characterized by higher pH (>8.0-8.5) whereas the other part by lower pH ( -7), then CaC03 will move toward the area of high pH and the dissolved SiOz (from diatoms, sponge spicules, etc.) will move toward the zone of low pH where it will precipitate (NEWELL et al., 1953, p.165; STRAKHOV et al., 1954, p.593). Certain workers (Brodskaya and Timofeeva, both in STRAKHOV, 1959) studied carbonate concretions and their origin. During late diagenesis as a result of first stage crystallization, or of crystallization of primary colloidal material with concomitant contraction, fractures form inside the concretions. These fractures do not reach the surface of concretions and end in V-pointed terminations. This observation led VITAL’(1 959, p.236) to believe that crystallization and lithification start at the periphery, because, as the loss in moisture and attendant volumetric contraction was reaching the central portions of concretions, the outer crust was already solid. The writers have observed that fractures are arranged parallel to the surface of the crust (concentric) or are diametrical, thus cutting the inner mass of concretions into sections. It is noteworthy that VITAL’(1959, p.224-227) on studying the amounts of minor elements (Ni, Co, V, Cr) inside concretions and in surrounding sediments found much higher concentrations of these elements within the latter environment. He also observed that the amount of minor elements (with the exception of Cu) increases with increasing concentration of insoluble residue. This probably indicates that minor elements did not migrate during diagenesis and are included inside concretions by mechanical means (together with captured particles of sediment). SUJKOWSKI (1958, p.2704) pointed out that flint nodules in some beds of the English White Chalk contain a small amount of chalk, commonly in a state of loose aggregation, in their centers. The enclosed chalk is composed chiefly of shells of microorganisms. Some layers contain flints that are empty. This observation led CAYEAUX (1897) to postulate that flints grew inward, and not outward as was gen-

286

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

TABLE X RELATIVE EFFECTS OF DIAGENESIS ON LIMESTONES

(WOLF,1963b; modified after KRUMBEIN, 1942, table 11) -

Compaction

+++ +

Surface texture

+

-

+

-

Mineral composition

-

-

Recrystallization

Solution

+++

++

++

++

++

+++

+ bytallization

-

++

+ expulsion + expulsion + of trace of trace ++

CW-

elernents

+ ++

+++ Permeability

Inversion

++ +++

Shape and roundness

Porosity

Cementation

+ ++

Particle size

Particle orientation

Pressuresolution

+ ++

+++

sand silt mud sand silt mud

++ ++

Color

+

-

Paleoenvironmental indicator

only of indirect value

poor indicator

elements

++

?

+?

++

++S

?

+?

++

T

-

too little information available

excellent to poor indicator

++S

morphology of CaCO3 cement; good indicator

?

+

Explanation of figures: + = small to moderate effect, ++ = moderate to large effect, + + + = most strongly affected, - = negligible effect, ? = uncertain.

1

erally believed. Though the concretions are siliceous, the analogy to calcareous concretions should be pointed out. Probably most chert nodules, Liesegang concretions, and calcareous concretions in carbonate rocks originated during diagenesis.

287

DIAGENESIS OF CARBONATE ROCKS

Cavity Internal formation sedimentation

Reworking

Dolomitization

Non-carbonate replacement

Authigenesis

++

-

-

-

+

.-

+

+

+

+

++

-

-

+

+++

+++

+++

-

++

+

tt-t

+++

-k

t t S

+++ + ++

may be excellent excellent indicator

Particle size Shape and roundness Surface texture Particle orientation Mineral composition

+?

+ ++

Porosity

+?

+ ++

Permeability

+

+ ++ +++ + ++ +++ +

+ ++

Color

some are good indicators

fair to good; internal fillings excellent

good to excellent; good to Paleoenvisilica, pyrite, excellent, e.g., ronmental hematite, etc. glauconite indicator

+

DIAGENESIS OF DOLOMITES

Diagenet ic dolomitization

As pointed out by BISSELLand CHILINGAR (1958, p.493), a study of the processes of diagenetic dolomitization should involve consideration of porosity changes.

288

G . V.

CHILINGAR, H.

J. BISSELL AND K. H. WOLF

The replacement of calcite by dolomite involves a contraction (increase in porosity) of about 12-13 % (CHILINGAR and TERRY, 1954) if the reaction proceeds as follows: Mgz++CaMg(CO& Ca2+ 2CaC03 Obviously, this contraction will occur only if there is no additional precipitation of carbonates in the pores and there is no subsequent compaction. The majority of carbonate petrologists agree that dolomitization gives rise to porosity providing a solid framework is available which will minimize the effects of subsequent compaction. For example, the findings of LUCIA(1962) indicate that porosity in the dolomite facies in Devonian crinoidal rocks in the Andrews South Devonian Field (Andrews County, Texas) was formed during diagenetic dolomitization. He found that porosity is generally related to the ratio of the lime-mud to crinoid fragments. CHILINGAR and TERRY (1954) also showed that a definite relationship exists between porosity and degree of dolomitization. T. F. Gaskell of British Petroleum Co., Ltd. (persona1 communication, 1963) determined the porosity and density of carbonate reservoir rocks in southwestern Iran. The average density values for the different oil fields, grouped in ranges of porosity of 0-4.0, 4.1-8.0, 8.1-12.0, and >12.1 %, are presented in Table XI and Fig.15. The mean values were weighted according to the number of observations for each oil field. A certain amount of the density scatter may be due to impurities in the limestones, variation in the amount of initial, primary porosity, variation in the degree of secondary cementation subsequent to dolomitization, etc. The gradual trend of density from 2.70 g/cm3 at the low porosity to 2.80 g/cm3 for the high porosity group indicates that dolomitization gives rise to porosity. Inasmuch as at 20°C the density of calcite is 2.71 and that of dolomite is equal to 2.87, the average values given in Table XI1 correspond to the percents of dolomitization given

+

+

TABLE XI RELATIONSHIP BETWEEN POROSITY AND DENSITY OF IRANIAN CARBONATE ROCKS

Name of oil

field

Porosity range' ( %) 04.1

4.1-8.0

8.1-12.0

3 12.1

Haft Gel Naft Khaneh Gach Saran Agha Jari Naft Sefid Lali Mi-S

2.68 2.62 2.71 2.74 2.67 2.74 2.67

f 0.04 (14) f - (1) f 0.09 (7) 0.09 (7) f 0.04 (3) f 0.08 (6) f 0.12 (11)

2.73 2.77 2.77 2.74 2.73 2.74 2.71

f 0.08 (3) i 0.08 (10) f 0.09 (10) f 0.04 (3) f 0.03 (6) 0.11 (8)

2.75 2.81 2.78 2.73 2.76 2.79 2.73

f 0.14 (8) f 0.05 (9) f 0.08 (11) & 0.06 (5) & 0.06 (9) f 0.04 (2) 0.12 (9)

2.78 2.83 2.79 2.81 2.76 2.79 2.80

Mean

2.70

(49)

2.74

(49)

2.76

(53)

2.80

& 0.09 (9)

& 0.12 (20)

f 0.05 (24) f 0.08 (8)

i 0.09 (9)

f 0.08 (12) 0.03 (4) i 0.06 (12) (89)

figures are mean square errors of the average values, and the figures in parentheses are The the numbers of observations.

1

289

DIAGENESIS OF CARBONATE ROCKS

in Table XI. These results (Table XI) are in close accord with those obtained by CHILINCAR and TERRY (1954). Diagenetic dolomitization may include two major stages, viz, ( I ) early diagenetic and ( 2 ) late diagenetic. In the case of early diagenetic dolomitization, due to subsequent compaction the porosity resulting from dolomitization is appreTABLE XI1 RELATIONSHIP BETWEEN POROSITY AND DENSITY -~ -

Porosity ( %)

0-4.1 4.1-8.0 8.1-12.0 2 12.1

._

Density (glcm31

2.70 2.74 2.76 2.80 2.84

.-

Dolomitization ( %)

P

26 32 58 82

ciably lowered (or disappears completely). TEODOROVICH (1958, p.306) observed that in marine replacement dolomites, the extremely fine-grained (0.01-0.05 mm) dolomite forming the relics of calcareous skeletons, with original pelitomorphic (< 0.005 mm) or micro-grained ( < 0.01 mm) texture, apparently formed at the same time as the central cores of dolomite grains in the main rock mass. The major part of dolomite grains (that is, crystals) formed during the second phase of dolomitization.

t

POROSITY, PERCENT

Fig.15. Relationship between density and porosity of Iranian carbonate rocks. (After T. F. Gaskell, British Petroleum Co. Ltd., personal communication, 1963.)

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G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

Fig. 16. Dolomitic limestone with complex texture. The dolomite rhombohedrons are composed of core and thin rim. The cores are composed of very fine-grained calcite with dolomite background, whereas the rims are pure dolomite. Areas between rhombohedrons are filled with medium-grained calcite. Insoluble residue: 1.70 %, CaC03: 75.37 %, CaMg(C03)~:22.88 7:. Upper Carboniferous, Samarskaya Luka. (After KHVOROVA, 1958, fig.175, p. 127.) Diagenetic dolomite, later altered during epigenesis; x 45.

Fig.17. Dolomitic limestone having complex texture. Dolomite rhombohedrons have cores with very he-grained calcite and light dolomite rims. In places where dolomite rhombohedrons are in close contact, the rim disappears and dolomite crust is created which borders several adjoining rhombohedrons (a). .The areas between aggregates of rhombohedrons, which apparently previously constituted pore spaces, are filled with medium-grained calcite (b). Thin-section was colored with K2Cr04. Upper Carboniferous, Samarskaya Luka. (After KHVOROVA, 1958, fig. 177, p. 127.) Diagenetic dolomite, later altered during epigenesis; x 45.

DIAGENESIS OF CARBONATE ROCKS

29 1

Fig.18. Strongly dolomitized algal (Dvinellu) limestone.In the mass of very fine-graineddolomite, one can observe numerous poorly preserved algal tubes (Family Eereselleue). Podol’skiy horizon, 1958, fig.181, p.128.) x 45. Onega. (After KHVOROVA,

TEODOROVICH (1958, p.306) also observed dolomite crystals showing threestage formation. The peripheral, clearer rims apparently belong to the third stage of dolomitization. In some very fine-grained (0.01-0.1 mm), fine-grained (0.10.25 mm), and vary-grained dolomites (0.25-0.5 mm = medium-grained; 0.5-1 .O mm = coarse-grained; > 1.0 mm = very coarse-grained) one can observe not only transparent, colorless(in thin section), rhombohedra1 grains with sharp contours but also cloudy crystals with irregular contours, often having zonal structure. Possibly, the latter crystals formed earlier in more recent sediments, whereas the former crystals developed later in muds which started to lithify. Dolomite selectively replaces fine-grained CaC03, but coarse-textured patches of CaC03 remain relatively unaffected. The order of dolomitization in organisms appears to be as follows: ( I ) Foraminifera, (2) brachiopods, (3) corals, and ( 4 ) crinoids (CHILINGAR, 1956~). In diagenetic dolomites the c-axes of dolorhombs do not appear to lie within the plane of bedding, and exhibit random orientation (HOHLT,1948; CHILINGAR and TERRY,1954). Examples of diagenetic dolomites are given in Fig.16-21. Diagenetic dolomites formed in the past, and STRAKHOV (1953, p.24) believed that most of these accumulated in Late Paleozoic time. The probability of high C02 content in the atmosphere and high Mg/Ca ratio of sea water during the Precambrian and Early Paleozoic times convinced some writers (STRAKHOV, 1953; CHILINGAR, 1956b) that dolomite precipitated directly out of sea water, in certain realms at least, during various stages of sedimentation. SUJKOWSKI (1958, p.27 15-27 16), however, made this statement: “Since Precambrian time, sediments have been more or less of the same type as recent ones, and, as temperature, air composition, and other geographical factors were not far removed from those known

292

G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

Fig. 19. Fine-grained, very porous dolomite. The pores are intercommunicated. Upper Carboniferous, R. Pinega. (After KHVOROVA, 1958, fig.244, p.139.) x 20.

Fig.20. Fine-grained dolomite with the green Algae Siphoneae (a). The walls of the latter are replaced by finer crystals than those of the inner portions and the main rock mass. Non-dolomitized remains of crinoid (b). Upper Carboniferous, Samarskaya Luka. (After KHVOROVA, 1958, fig.255, p.141.) x 45.

today, it is only logical to assume that diagenesis followed the same lines as today. This is shown by identical rock types throughout the geological column.” Chemistry of diagenetic ‘dolomitization

Numerous mottled dolomites are reported in the literature as examples of “ar-

293

DIAGENESIS OF CARBONATE ROCKS

Fig.21. Fine-grained dolomite with remains of micro-grained calcite and non-dolomitized organic material. Upper Carboniferous, Samarskaya Luka. (After KHVOROVA, 1958, fig. 186, p.129.) x 20.

rested diagenesis”. According to STRAKHOV (1956, p. 18), some dolomite could also precipitate during sedimentation and later upon dissolving could be redistributed during the diagenetic stage, giving in places a spotty appearance to dolomitic limestones. TEODOROVICH (1958, p.306), however, disagrees stating that at high pCOz conditions dolomite is less soluble than CaC03. Some dolomites possibly can also form through the interaction of CaC03 with some MgC03 salts [xMgC03 * yMg(0H)z * zHzO] during the process of diagenesis (CHILINGAR, 1956a; BISSELLand CHILINGAR, 1958, p.495; HALLAet al., 1962). Probably, diagenetic dolomitization occurs in strongly reducing to weakly reducing environments having high alkalinity: ([A] = [HC03-] 2 [C032-]) and a pH> 8 (up to 9 and higher). TEODOROVICH (1958, p.93-103) suggested that replacement dolomites (with some glauconite and oxides and hydroxides of iron) also form in weakly oxidizing to oxidizing environments having pH of 8 (7.8)-9. He believed that many replacement dolomites with relic organic and oolitic textures formed in oxidizing environments that have a pH of 7.2-7.8. Furthermore, CHILINGAR and BISSELL (1963a) assigned a pH value of Q 8 to the environment necessary, in their evaluation, for the formation of “primary” sedimentary dolomites. TEODOROVICH (1 958, p.305) believed that Haidinger’s reaction could account for some dolomitization in waters saturated with CaS04, according to this reaction:

+

+ MgS04+MgC03 + CaS04 (in solution) CaC03 + MgC03 = CaC03.MgC03 4

CaC03

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G . V.

CHILINGAR, H.

J. BISSELL AND K. H. WOLF

The second part of this reaction (2), however, does not occur at high concentrations of MgS04. Thermodynamics of this reaction were discussed by HALLAet al. (1962). On the basis of experimental data, VALYASHKO (1962, p.47) presented the main reaction which occurs between calcium bicarbonate and sulfate solutions: Ca(HC03)z

+ MgS04eCaS04 + Mg(HC03)z I

hydration

I

hydrolysis

4 4 CaS04.2H20

I

(3)

xMg(0H)z. yMgCO3. ZHZO4 With small concentration of MgS04 in solution, the above reaction practically does not occur, but instead there is decomposition of calcium bicarbonate with formation of calcite, as follows: Ca(HC03)z+CaC03

+ HzO+COZ

(4)

As the MgS04 increases, first gypsum starts to form, followed by basic carbonates of magnesium. At this time, CaC03 practically disappears from the bottom phase (see also CHILINGAR and BISSELL,1963b). In addition to the reaction (3), reaction (5) also occurs very slowly: 2Ca(HCO&

+ MgS04+CaMg(C03)~+ CaS04 + 2H20 + 2C02

(5)

VALYASHKO (1962, p.55) obtained individual, rhombohedra1 crystals of dolomite (identified by crystallo-optical analysis) in the laboratory at atmospheric conditions (low COz pressure). Possibly this reaction could account, in a measure at least, for certain extensive dolomite formation at higher COZpressures. Methods of introduction of Mg2+ ions into carbonate ooze

The present-day concentrations of Mg2+ and Ca2+ in the ocean are 1,297 and 408 p.p.m., respectively. Geologists, however, cannot be sure that the Mg2+ concentration of ancient seas was not higher at times than it is at present (CHILINGAR, 1956b, p.2263). Influx of Mg2+ ions into carbonate muds may occur as a result of reduction of the sulfate ion if Mg2+ is combined with S O P ( ? ) . According to STRAKHOV (1956a, p.19), if the sulfate present in the mud, having a moisture content of 75 %, undergoes complete reduction, it will give rise to 0.26 % sulfur which can form pyrite in the sediment. Thus, the amount of pyrite in dry sediment should be present in an amount of about 0.5 %. Actually, the normal content of pyrite in clayey marine sediments ranges from a small fraction of 1 % to 1 % (rarely higher). Thus, the sulfate ions used in the formation of pyrite probably come mostly from the initial

295

DIAGENESIS OF CARBONATE ROCKS

(connate) water trapped in the sediment, and there seems to be no appreciable diffusion of sulfate ions from the bottom waters into the sediment. But together with SO42- anions, Mg2+ and Ca2+ cations are also trapped in the mud. The amount of magnesium thus can be calculated as follows: at the complete reduction of sulfates in interstitial water, for 0.26 % of sulfur there will be 0.13 % magnesium (considering that only two-thirds of sulfur is combined with magnesium). This will constitute about one-third of the total magnesium contained in the interstitial water at the time of burial, which is equal to 0.4 % of dry sediment. If it is assumed that all of the magnesium produced as a result of the reduction of sulfates goes to form dolomite, there should be a decrease of magnesium content in the interstitial water in the amount of one-third as compared to its content in the bottom water. STRAKHOV (1956a, p.19) suggested that this differential alone could not create a strong influx of magnesium from the bottom water into the mud. Thus, the deposition of 0.13 % magnesium would give rise to only 1 % of dolomite (i.e., 1 %on dry-weight basis). Strakhov also pointed out that the amount of organic matter in carbonate muds he studied is so low, that it is doubtful that any appreciable reduction of sulfates occurred there. Another reason for the influx of magnesium from bottom waters into sediment would be the decomposition of organic matter by bacteria with generation of C02. As a result of this process, first the alkalinity ([HC03-] 2[C0s2-]) would rapidly increase and then would decrease with loss of COz from interstitial water of the sediment. With increasing alkalinity dolomite could reach a saturation value and precipitate. Removal of Mg2+ in this manner could result in its additional influx from bottom waters. This alone, believed STRAKHOV (1956a, p.20), could not create any appreciable influx of Mg2+from bottom waters because the amount of organic matter (generating COz) is negligible in the carbonate oozes. The process very likely would give rise only to: ( I ) individual crystals of dolomite, ( 2 ) rare and small spots (mottled dolomite), and (3) concretions. Yet, as pointed out by STRAKHOV (1956a, p.20), average dolomitization of spotty metasomatic dolomites reaches 30-70 % and much more in some instances. It would appear, therefore, that to create such diagenetic dolomitization as is envisioned by geologists, the above process cannot of itself fulfil the requirements. Possibly, as diagenetic replacement by dolomite progresses, the impoverishment of interstitial waters in magnesium content would enable additional influx of Mg2+ ions from bottom waters. Some porosity which results from secondary dolomitization could furnish additional space. This porosity developed during the change of calcite or aragonite to dolomite is possibly sufficient to provide necessary diffusion centers from which Mg2+ ions can migrate through the intergranular film (BISSELL and CHILINGAR, 1958). It is interesting to note here that interstitial water in the dolomite beds of Tertiary carbonates from Experimental Mohole (Guadalupe Island, Mexico) still contains the same Mg2+ levels as that of the modern sea (RITTENBERG et al., 1963).

+

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CHILINGAR, H. J.

BISSELL AND K. H. WOLF

DEGENS and EPSTEIN (1964) suggested that possibly this may serve as an indication that sea water entrapped in the sediment is not a significant source for magnesium in the dolomite. According to FAIRBRIDGE (1957), calcareous Algae enriched in MgC03 are geochemically reorganized into low-Mg calcite and dolomite during diagenesis as exemplified by Funafuti dolomite. Possibly the magnesium carbonate present in solid solution in the skeletal and protective structures of organisms (CHILINGAR, 1962a) comprises one of the sources of Mg necessary for dolomitization. The fact that this MgC03 is thermodynamically unstable is partly supported by the findings O f CHAVE (1954), LOWENSTAM (1961), CHAVE et al. (1962), CHILINGAR (1962b) and et al. (1964), however, found that the reorganization into SEIBOLD (1962). EPSTEIN calcite and dolomite does not seem to have any significant effect on isotope ratios of Funafuti dolomite; the values of the dolomites are close to those originally fixed in the calcium carbonate. Epigenetic dolomites

The epigenetic dolomites result from alteration of completely lithified limestones by downward percolating meteoric solutions or rising hydrothermal solutions (VISHNYAKOV, 1951, p. 112). Epigenetic dolomites are cavermous, have obscure stratification, patchy distribution, non-uniform grain size, and relict structure. Idiomorphic rhombohedrons of dolomite with nucleus and zonation are common. The texture is not uniform and the original fauna remains in the form of molds. The Ca/Mg ratio of epigenetic dolomites varies widely over short distances, both vertically and horizontally. It should be remembered, however, that many geologists call epigenetic dolomites “late diagenetic”. According to STRAKHOV (1956b, p.199), the apparent increase in degree of dolomitization due to secondary removal of CaC03 (an epigenetic process) does not seem to be important. For example, in 1 dm3 of limestone (5% dolomitized; s. g. = 2.7) there is 13.5 g of dolomite. Creation of 10% porosity would necessitate the removal of 27 g of CaC03 from 270 g of rock; and consequently the percentage dolomitization of the remaining mass would be:

13.5 270-27 ~

100 = 5.5%. In other

words, it will increase by a negligible amount of only 0.5%. An increase of porosity to 20% would give an apparent increment in degree of dolomitization of 1.2%; to 30%-2.1%; to 40%-3.3%; and to 50%-5%. In the latter case, the rock would have a sponge-like appearance, and yet the apparent degree of dolomitization would increase by 5% only. The above calculations should be kept in mind, because secondary removal of CaC03 is frequently used to explain high degree of dolomitization.

DIAGENESIS OF CARBONATE ROCKS 180/160

and

13C/12Cratios

297

of carbonate rocks as a tool in studying diagenesis

It has been noted by Degens and Epstein (in BISSELLand CHILINGAR, 1962), that in the case of co-genetic precipitation, the calcites and dolomites should be different by 8 in their 180/'60 ratio. Thus, if the 6180 of dolomite is, for example, only 1 heavier than that of calcite which occurs together with it (evaporite type dolomite in the salt flats of western Utah; BISSELLand CHILINGAR, 1962, p.207), then the dolomite is probably of diagenetic origin. In a comprehensive study of coexisting dolomites and calcites, DEGENS and EPSTEIN (1964) showed that dolomites from Recent and unconsolidated marine sediments have 6180 values of -0.8 to +4.9 ($2.1 average) and d13Cvalues of - 1.2 to 1.4 (-0.2 average). The associated calcites have P O of -0.4 to +4.6 (+1.8 average) and 613C of -0.8 to +1.2 (-0.4 average).l Thus, the coexisting calcites and dolomites are quite similar isotopically. Consequently, DEGENS and EPSTEIN (1964) concluded that recent dolomites are not precipitated directly out of aqueous solutions, but are formedasaresult of metasomatism from pre-existing crystalline calcite carbonate. They also stated that geologic and 1% data indicate that dolomitization can start shortly after deposition; and that dolomitization has to proceed under solid-state conditions by amole for mole exchange of Ca2+forMg2+,without chemically altering the CO$- unit, because no isotope fractionation (either for C or 0)occurs during the alteration of the CaC03 precursor material to dolomite. They ruled out the participation of bicarbonate phase during dolomite formation. EPSTEIN et al. (1964) demonstrated the reluctance of dolomite to adjust isotopically to changes in temperature and 1 8 0 / 1 6 0 ratio of formation waters. This presents a possibility of using dolomites for the evaluation of paleotemperatures of ancient seas. It was also demonstrated that 613C apparently is not significantly affected by the dolomitization process (DEGENS and EPSTEIN, 1964).

x0

x0

+

PRACTICAL APPLICABILITY OF DIAGENESIS

It is beyond dispute that a clear concept of the paleoenvironments is most important in the search for both non-metallic and metallic economic deposits in calcareous sediments. As many secondarily introduced accumulations are controlled by primary and secondary porosity and permeability, the investigation of diagenesis will, therefore, help in elucidating: ( I ) the locality and type of porosity and permeability; (2) permeability pinch-outs; (.?) the factors that cause some non-reef limestones to be good reservoir rocks in one locality, whereas under seemingly similar conditions identical deposits are cap-rocks; (4) the conditions that make The data are reported as permil deviation relative to the PDBI Chicago Belemnite Standard (CRAIG,195 7).

1

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G. V. CHILINGAR, H. J. BISSELL AND K. H. WOLF

shallow-water limestones either good or poor source rocks; (5) the likelihood of reef-flank and reef-core deposits occurring in one direction over others, and so forth. A useful approach for the reconstruction of paleoenvironments in oil exploration is the plotting of equal sparite, microsparite and micrite contents or ratios, (1962), for example; or plotting lines among other parameters, as done by STAUFFER of equal Ca/Mg ratios as proposed by CHILINGAR (1953, 1956~).In one case, diagenetic and syngenetic features were used by WOLF(196%) to prove a littoral environment of a reef complex. All this is, of course, based on a thorough understanding of diagenetic processes and products, for it is important to distinguish between cementation and recrystallization products, for instance. Detailed work on carbonate diagenesis may also facilitate our understanding of certain metallic deposits in limestones as pointed out by CLOUDet al. (1962), who mentioned that the location of particular lead, zinc and manganese deposits in carbonate rocks may well reflect some intrinsic chemical, biological, earlier diagenetic or textural property of the rock. DANCHEV and OL'KHA(1959), for example, studied uranium-bearing limestones to determine the parameters that controlled mineralization and the location thereof. They concluded that organic content controlled localization of the minerals, and that mineral dissemination predominates where the rock is least affected by recrystallization and leaching. The ore is of diagenetic origin and has been epigenetically redistributed.

ACKNOWLEDGEMENTS

Plate I-XXIV have been prepared with the kind assistance of the Medical Department, Illustration Section, University of Sydney, and form part of the work of WOLF(1963a).

REFERENCES

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DIAGENESIS OF CARBONATE ROCKS

299

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GLOSSARY

Algal dust: angular to subangular medium- to dark-colored grains or crystals of carbonate, commonly 1-5p in diameter, derived from breakdown of algal felts, algally-precipitated aragonite needles, algal slime, and comminution of phytoplankton; associated with algal tubes, algal nodules, and other Algae or definite evidence thereof. Term proposed by WOOD(1941), with certain details added by CAROZZI (1960). ‘‘Algal dust” has also been called algal micrite (WOLF,1965b) which occurs as allo- and auto-micrite types (Table III). Algal paste: dark gray to,black finely-divided flecks, micrograined, microcrystalline, or cryptocrystallinein texture, forming a rather dense micritic limestone or dolomite, and associated with organic frame-builders such as corals, sponges, bryozoans, etc. Common, but not restricted, to the reef core. May actually represent compact, dense, diagenetically altered 1957.) dust. (Term used in a loose sense by SCHLANQER,

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Allo-: as used here, a prefix derived from allochthonous and indicating that the material has been transported before accumulation (Table III). Allochthonous: a term used here to designate sedimentary constituents which did not originate in situ; they were derived from outside of or within the area of depositional site and underwent transportation before final accumulation. Allogenic: a term meaning “generated elsewhere,” and applied to those constituents that came into existence outside of, and previously to, the rock of which they now constitute a part, e.g., extraclasts. Aphanic: a term proposed by DEFORD (1946) for the texture of carbonates, particularly limestones, in reference to crystalline (and/or grained) textures, the discrete particles of which are smaller than 0.005 mm (Table 110. Microcrystalline (also micrograined) and cryptocrystalline (also cryptograined) are the two textural subdivisions. Aphunic is used here to replace the term “aphanitic,” which is loosely defined and is not utilitarian for carbonate rocks. Apo-epigenesis: as used here, epigenesis affecting the sediments after diagenesis while they are far remote from the original environment of deposition under a relatively thick overburden. With an increase in temperature and pressure it grades into metamorphism (WOLF,1963a). Aragonite: a mineral, orthorhombic CaCO3, dimorphous with calcite. Authigenic: generated on the spot. Applied to those constituents that came into existence with or after the formation of the rock of which they constitute a part. Auto-: as used here, a prefix derived from autochthonous and indicating that the material was formed in situ (Table 110. Autochthonous: a term applied here to sedimentary rock components which originated and formed in situ, without undergoing prior transportation. Buhamite: name proposed by BEALES (1958) for the granular limestones that closely resemble the present deposits of the interior of the Bahama Banks, described by ILLING(1954). The texture varies from calcisiltites to calcirudites, in which the grains are accretionary and commonly composite, consisting of smaller granules bound together by precipitated material into aggregate grains. Many misinterpretations of this rock type have been made (WOLF,1965a, b). Bunk: a skeletal limestone deposit formed by organisms which do not have the ecologic potential to erect a rigid, wave-resistant structure. Contrasts with reef, which is a skeletal limestone deposit formed by organisms possessing the ecologic potential to erect a rigid, waveet al., 1962). resistant, topographic structure (NELSON Beach-rock: a friable to wellcemented beach sediment consisting of calcareous debris cemented 1953). by calcium carbonate (GINSBURG, Biolithite: a term applied to faunal and/or floral organisms that grew and remained in situ (FOLK,1959). Birdseye: spots or tubes of sparry calcite in limestones (HALL,1847). PERKINS (1963) pointed out that 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 spany calcite in tubules resulting from escaping gas bubbles; (4)reworking and rapid redeposition of soft sediment to produce a rock with very vaguely defined proto-intraclasts, semicoherent clouds of calcareous mud, and irregular patches of spar;

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and (5) recrystallization of calcareous mud in patches. For alternative interpretation see WOLF (1965~). Bfady calcite: see cement.

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 1962). substantially in position of growth (DUNHAM, Breccia: a rock made up of angular rock fragments, most of which are larger than 2.0 mm in diameter. CAROZW(1960) mentioned a recrystallization breccia that results from the differentiation in place of a homogeneous calcilutite. Recrystallization began at numerous points scattered throughout the rock but was incomplete, and as a result the recrystallized patches appear as fragments in a groundmass that was spared by the process. It is here recommended that the definition of Carozzi be expanded to include other carbonate rocks such as calcisiltites, calcarenites, etc. Bryulgaf: term proposed by BISSELL (1 964) for organic frame-building combination of bryozoans and Algae whichcreate arigid, wave-resistant limestone mass that forms banks, and is reefal, or at least is intimately associated with reefs. The deposit is in situ; in some occurrences, one organism encrusts the other. Algae, for example, may encrust bryozoans; they may also encrust corals, stromatoporoids, sponges, and other framebuilders. Culcureous: as used here, referring to calcitic and aragonitic material. Calcilutite to calcirudite: a range of terms suggested by GRABAU (1904; 1913) for limestones indicating the size of the calcareous components as given below: Calcilutite = clay-sized calcareous particles, Calcisiltite = silt-sized calcareous particles, Calcarenite = sand-sized calcareous particles, Calcirudite = gravel-sized calcareous particles. (Compare with dololutite to dolorudite range.) Calcite: a mineral, calcium carbonate, CaC03, hexagonal-rhombohedral, dimorphous with aragonite. Cafclirhite: a limestone containing 50% or more of fragments of older limestone eroded and redeposited (FOLK,1959). The individual fragments are called extraclasts (WOLF,1963b, 1965b). (Compare with intraclasts.) Cufiche: it is a lime-rich deposit found in soils and formed by capillary action drawing the lime-bearing waters to the surface where, by evaporation, the lime is precipitated (PETITJOHN, 1957). In bajadas, intermonts, alluvial fans and colluvium of parts of the Great Basin of Western United States some of the caliche deposits are dolomitic due to presence of extensive dolomite rubble. Caliche, whether calcareous and/or dolomitic, also cements alluvial fans to form fungfomerute. Culcsparite: see sparite.

Cavbonute rock: a sedimentary rock composed of more than 50% calcite, aragonite, and/or dolomite. Cement: chemically prFipitated 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 morphological and textural types are granular, fibrous, blady, and drusy. Carbonate cement often resembles products formed by recrystallization and grain growth. The cryptocrystdline

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carbonate is too small to be resolved by an ordinary petrographic microscope and appears as a dense mass. The granular sparry cement consists of more or less equi-dimensional crystals. The fibrous sparite occurs as very thin elongate fibres, whereas the blady type has somewhat wider elongate crystals. The term drusy sparite does not refer to a single crystal but to the textural relation of aggregates composed of crystals that increase in size and elongation with increasing distance from the host. Drusy calcium carbonate usually changes into a granular type. (See Table 111 and Fig.2.) Cementation: the process of chemical precipitation of material into voids and in situ onto the surfaces of the host-framework. (Compare with internal chemical sedimentation.) Clust: an individual constituent of detrital sediment or sedimentary rock produced by the physical disintegration of a larger mass either within or outside the basin of accumulation. (See extraclast and intraclast.) Coated grains: grains possessing concentric or enclosing layers of calcium carbonate; for example, oolites, pisolites, superficial oolites, algal-encrusted skeletal grains (LEIGHTONand PENDEXTER, 1962), and circumcrusts (WOLF, 1962, 1965b). Cone-in-cone:a concretionary structure occurring in marls, etc., characterized by the development 1928). of a succession of cones one within another (HOLMES, Contemporaneous: existing together or at the same time in contrast to penecontemporaneous. Coquinu: carbonates consisting wholly, or nearly so, of mechanically sorted fossil debris. Most commonly applied to the more or less cemented coarse shell debris. For the finer shell 1957). detritus of sand size or less, the term microcoqiiina is more appropriate (PETTIJOHN, Coquiizite: indurated equivalent of coquina. Criquinu: coquina of crinoidal debris. Criquinite: indurated equivalent of criquina. Cryptocrystalline: crystalline material that is so fine that it cannot be resolved by a petrographic et al., 1955). Electron microscope studies, however, show distinct microscope (WILLIAMS crystalline features. Cryptocrystalline carbonate forms the finest part of the micrite (Table 111). Dense: compact; having its parts ciowded together. Not necessarily restricted to fine-textured carbonate rocks, although commonly applied in this sense by many petrographers. Depocenter: contraction of depositional center; refers to an environment of sedimentary deposition, without particular restriction as to whether it is a basin, bank, shelf, trough, etc. (MURRAY, 1952). Geosynclines, particularly of the miogeosynclinal type, consist of basins, troughs, swells, banks, welts, incipient-to-prominent highs, accessways, thresholds, reefs, barriers, lagoons, hinge-lines, and various other repositories of sediment accumulation. Centers of carbonate deposition, i.e., depocenters, make up these repositories of geosynclines, intra-cratonic basins, platforms, etc. (BISSELL,1962). Detrital limestone: limestone composed of fragments that have been transported before accumulating. (Detrital is synonymous with “allochtonous”.) Detritus: transported material not formed in situ. (Detritus is synonymous with allochtonous material, allochtonous fragments, and debris.) Diagenesis: it includes all physicochemical, biochemical and physical processes modifying sediments between deposition and lithification, or cementation, at low temperatures and

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G. V. CHILINGAR, H. I. BISSELL AND K. H. WOLF pressures characteristic of surface and near-surface environments. In general, diagenesis is divisible into pre-, syn-, and post-cementation or lithilication processes. Diagenesis, as defined in this chapter, takes an intermediate position between syngenesis and epigenesis, the former grading into diagenesis by syndiagenesis, and the latter grading into metamorphism. Under unusual conditions, however, diagenesis as defined here may grade directly into metamorphism(see epigenesis). Because reef limestones, and other limestones which are constructed in situ by organic frame-builders, are largely in and of themselves lithified to a degree, the definition must be expanded for this particular group of limestones to include the interactions between sediments and the fluids contained within them below the temperature and pressure levels of metamorphism sensu stricto, and in a similar sense between fluids and framework, W e d detritus framework, and combinations thereof.

Dololutite to hlorudite: a range of terms applied to sedimentary dolomites composed of constituents ranging in size from clay to gravel, similar to those in limestones, as follows: Dololutite = clay-sized dolomite particles, Dolosiltite = silt-sized dolomite particles, Dolarenite = sand-sized dolomite particles, Dolorudite = gravel-sized dolomite particles. Dolomite: ( I ) a mineral, CaMg(COa)z, hexagonal rhombohedral. (2) A carbonate rock composed predominantly of the mineral dolomite; in normal routine petrographic work, dolomite (or dolostone of some geologists) is a carbonate rock composed of more than 50% by weight of the mineral dolomite. More practically, 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, as cement and/or grains or crystals. Dolomitic can be applied to the large spectrumof sedimentary rocks that are dolomitebearing, and also to limestones which have been dolomitized to a degree but not completely. Dolomitic mottling: incipient or arrested dolomitization, or arrested (or incomplete) dedolomitization. Common to limestones that have large particulate skeletal or nonskeletal material embedded in finer-textured matrix. Under the effects of dolomitization there is a preferential replacement or alteration of the matrix but not of the large particles. Also common to more or less homogeneous textured limestones that have been incompletely dolomitized, leaving patches, blotches, laminae, or other structures unaffected. Dolomitized: refers to rocks or portions of rocks in which limestone textures are discernible, but which have been changed to dolomite. Dolosparite: see sparite. Drusy: see cement. Earthy: refers to a variety of slightly argillaceous carbonate with earthy texture generally closely associated with chalky deposits and commonly showing similar porosity values. 1962). Microtextured (0.01 mm and slightly less) (THOMAS,

Endogenic: as used here, referring to components derived from within the sedimentary formation. Eolianite: sedimentary accumulation formed by wind action. Epigenesis: as used here, it includes all processes at low temperature and pressure that affect sedimentary rocks after diagenesis and up to metamorphism. Epigenesis has been subdivided into juxta- and apo-epigenesis (WOLF, 1963b, 1965~).It is possible that under unusual conditions syngenesis and diagenesis grade directly into metamorphism. For example, unconsolidated sediments may be exposed to volcanic high temperatures and metasomatic material and undergo metamorphism before diagenesis. Also, sediments

DIAGENESIS OF CARBONATE ROCKS

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partly undergoing cementation may be metamorphosed by shallow intrusions causing an increase of temperature and possibly pressure before epigenesis could occur: Syngenesis --+ Metamorphism Diagenesis Epigenesis

J.

Metamorphism Evaporite-solution breccia: solution breccias are created when intervening soluble evaporitessalt, anhydrite, gypsum, etc.-are dissolved away, letting the carbonate beds crush under the weight of overlying sediments. An extremely angular collapse breccia results, in which the matrix is of essentially the same material as the rock fragments ( S w and LAIRD, 1947; GREINBR, 1956). These chaotic breccias normally are associated with evaporites, and may also be adjacent to reef limestones which, attendant to removal of the evaporites, collapse and may be "healed" or cemented by calcareous and/or dolomitic material. Exogenic: as used here, referring to components derived from outside, i.e., from either above or below, the sedimentary formation. Extraclast: fragment of calcareous sedimentary material produced by erosion of an older rock 1963b, 1965a,b). (Compare with outside the depocenterin which it accumulated (WOLF, intraclast and calclithite.) Fibrous: see cement, Flour: chalky-appearing, finely comminuted material in limestones or dolomites, generally formed by disintegration and abrasion of fossiI debris and algal growths under intense wave action, surf-surge, and current action in shoal areas. It may represent clay-sized particulate carbonate mud formed through attrition, or may result from chemical flocculation, biochemical activity, or through other means. These micrograined, chalky carbonates may be due to disintegration and abrasion of fossil detritus on banks and shelves that are subject to the high energies of waves and currents. 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 6ne mosaic is gradually replaced by a coarser. As grain growth proceeds, many of the enlarged grains are them1958). In limestones grain selves replaced by their more successful neighbors (BA-T, growth appears to affect only the very fine mosaics with grain diameters ranging from 0.5 to 4.0 p. These include calcitemudstones, the walls of Foraminifera, algal frame1959b). works, bahamite particles, and ooliths (BATHURST, Grainstone: mud-free carbonate rocks, which are necessarily grain-supported, are termed grainstone; some are current laid, whereas others form as a result of mud being by-passed while locally produced grains accumulate, or of mud being washed (= winnowed) out (DUNHAM, 1962). Grain-supported: carbonate sedimentary rock in which grains are so abundant as to support one another, just as they do in mud-free rocks (DUNHAM, 1962). Granular: see cement. Grumous: a term signifying clotted, aggregated, flocculated. As applied to sedimentary carbonate rocks it refers to micro- and macroscopic aggregation of lime-mud particles and other flocculated or otherwise clotted and aggregated, irregularly-shaped material. In a sense comparable to bahamite, but commonly of smaller dimension.

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Halmyrolysis: the chemical rearrangements and replacements that occur while the sediment is still on the sea floor (PBTTUOIIN, 1957). It is sometimes called submarine weathering. Hypogenic: a term applied to material that is derived from within the earth interior in contrast to supergenic components. (See supergenic.) Impingement: a mechanism or process in dolomitization in which dolomite crystals replace limestone, commonly skeletal particles such as crinoid ossicles and plates, but not in optical continuity with the calcite of the original particle (LUCIA,1962). Internal filling: a collective term including both internal sediments and cement that fill cavities within a sedimentary.formation (WOLF,1963a, 1965~). Internal chemical sediment: allochthonous chemically precipitated sediment both formed and deposited intraformationally in cavities (WOLF,1963% 1965~).(Compare with internal mechanical sediment and internal filling.) Intergranular porosity: void space between grains, whether bioclastic or lithoclastic. In sedimentary carbonate rocks the term granular commonly refers to the grains, whether skeletal or nonskeletal. Internal mechanical sediment: allochthonous clastic sediment brought in from the surface, or derived by intraformational abrasion, and deposited in cavities within the sedimentary formation (WOLF,1963a, 1965~). Internal sedimentation: allochthonous 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 (WOLF,1963a, 196%). Interstitial: of, pertaining to, existing in, or forming an interstice or interstices. Intraclast: fragment of more or less consolidated calcareous sedimentary material produced by erosion within a basin of deposition and redeposited there (FOLK,1959). (Compare with extraclast.) Intruformational: formed by, existing in, or characterizing the interior of a geological formation. Zntragranular porosity: pore space or voids within individual particles, particularly skeletal material. Of significancein leached ostracodal, Foraminiferal, algal, and oolitic limestones, but, like intergranular porosity is sometimes adversely affected by diagenetic processes. Inversion: the process by which unstable minerals change to a more stable form of the same chemical composition (except for a possible change in contents of trace elements and/or isotopes) but with a different lattice structure. Juxta-epigenesis: epigenesis affecting the sediment after diagenesis while it is near the original environment of deposition either under a relatively thin overburden or, if regression occurred, while exposed above sea level (WOLF, 1963a). (Compare with epigenesis and apo-epigenesis.) Limestone: a sedimentary rock composed of at least 50% calcium carbonate material. For practical microscopic work, it is a carbonate consisting of 50% or more, by areal percentage, of &lcite or calcareous material. Lithifcation: that complex of processes that converts a newly deposited sediment into an indurated rock. It may be contemporaneous with, shortly after, or long after deposition.

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Lithoclastic: autochthonous and 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. (See also detrital, limeclast, intraand extraclast, and calclithite.) Lithographic: pertaining to a compact carbonate rock having about the same particle size and 1946). Characterized by textural appearance as the stone used in lithography (DBFORD, conchoidal fracture. Numerous micro- and crypto-textured micritic limestones and dolomites are lithographic. Littoral: belonging to, inhabiting or taking place on or near the shore between low-tide and high-tide level. (See sub- and supra-littoral.) Lump: a descriptive term applied to an aggregate grain composed of two or more pellets, oolites, skeletons, etc., or fragments thereof. The aggregation or accretion can form by physicochemical, algal, and weathering processes. Calcareous grains such as pellets lying in contact with each other on the sea bed tend to become cemented or welded together and 1954; WOLF, 1965b). form lumps (IL.LIN~, Matrix: if the particles in the calcareous rock are of different orders of size grades, the term matrix is used for the material that fills the interstices between the larger grains. Matrix is thus the material in which any sedimentary particle is embedded. The matrix may be either microtextured or granular. With an increase in matrix percentage, a limestone grades into a deposit composed solely of micrite, of calcisitite, or of calcarenite. Granular matrices tend to become more poorly sorted as particle size increases. Some prefer to restrict “matrix” to clay-sized or micritic components surrounding coarser material. Metamorphism: this term refers to the mineralogic, textural and structural adjustment of solid rocks to physical or chemical conditions at higher temperatures and pressures than those under which the rock in question originated. Micrite: a descriptive term for calcareous crystalline and/or grained material less than 0.005 mm in diameter (Table TII) as used here. (FOLK,1959, used 0.004 mm, whereas LEZGHTON and P E ~ E X T E1962, R , drew the limit at 0.031 mm). Micrite that is so finely crystalline that it cannot be resolved by a petrographic microscope is called “cryptocrystalline”. It is consolidated or unconsolidated ooze or lime-mud of either chemical or mechanical origin, and possibly of biologic, biochemical, and physicochemical origin. It is used by some geologists as synonymous with caldutite (clay-sized particles). The exact range of both micrite and calcilutite, however, has been differently placed by other workers. (See orthoand pseudo-micrite, and Table III.) Microgruined: a grain-size term pertaining to carbonate particles smaller than 0.02 mm and larger than 0.005 mm in diameter; microckustic is more or less synonymous. Microsparite: see sparite.

Mud aggregate: any aggregate of mud grains, usually having the size of a sand or silt particle, which has been mechanically deposited. Initially, the aggregate may have been a faecal pellet, or a rounded, sub-spherical aggregate of mud grains cemented originally by aragonite with no signs of organic control, or a fragment of algal precipitate, or a spherical 1959b). or ovoid growth form of a calcareous alga (BATHURST, Mudsfone: muddy carbonate rocks containing less than 10% grains (10 % grain-bulk); the name is synonymous with calcilutite, except that it does not specify mineralogic composition, 1962). and does not specify that the mud is of clastic origin (DUNHAM, Mud-supported: muddy carbonaterock which contains more than 10 % grains, but not in suilicient

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amount to be able to support one another; such grains are “floating,” and thus they are mud-supported. Oiilite or odith: a spherical to ellipsoidal body up to 2 mm in diameter which may or may not have a nucleus, and has concentric or radial structure or both. It is accretionary. The term is descriptive. At least three genetic possibilities exist identical to those mentioned for pisolites. If particles lack concentric or radial features, one should refrain from calling them false or pseudo-oolites, but name them pellets. (See pisolite and pellet.) 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 (WOLF, 1963a, 1965~). Organic lattice: reef-building framework, and some bank deposits constructed by organic frame-builders, in situ. Orthomicrite: it is a genetic term applied to micrite that has not undergone secondary changes such as recrystallization and grain growth. Two types are recognizable: allochthonous and autochthonous dcrite named allomicriteand automicrite, respectively (WOLF,1963b). (See micrite and pseudomicrite, and Table m.) Orthosparite: see sparite. Paragenesis: a general term for the order of formation of associated minerals, textures, and structures in time succession, one after another. Pelagosite: it is a deposit generally white, gray to brownish with a pearly luster, composed of CaCOa with higher MgCOs, SrCOa, CaS04.Hg0 and SiOz contents than in normal limy 1957). sediments (RHVELLE and FAIRBRIDGE, Pellet: a spherical, sub-spherical, ovoid, to irregular-shaped small particle composed of claysized to fine silt-sized material and devoid of any internal structure. Micrite pellets have been called pseudo-oolites, false oolites, etc. Threegenetictypes appear to be of significance: faecal, bahamite, and algal pellets. Penecontemporaneous: a term used in connection with the formation of sedimentary rocks, and implies “formed at almost the same time”. (Compare with contemporaneous.) Pressure solution: a preferential solution takes place on the higher stressed parts of a grain and deposition of matter on surfaces with lower potential energies. The pressure is supplied by the overburden and should result in a 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 space of sediment (BATHURST, 1958, 1959b). Microcrystalline calcite can recrystallize by pressure solution into a mosaic of larger crystals by the solution 1962). of the smallest, supersolublegrains and redeposition on the larger grains (STAUFFER, Primary: characteristic of or existing in a rock at the time of its formation. This definition is too all-inclusive and vague in detailed studies and its use should be discouraged. It can be used unambiguously as a very general colloquial term in COMectiOn with genetic discussions only if the context leaves absolutely no doubt. (See Secondary.) 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 irregularlyshaped patches of coarse calcite mosaic usually between 1 and 20 mm in diameter, and are dark gray in handspecimen. They lie in the finer, pale-gray “groundmass” of calcite-mudstonr. In and thin-section the “fragments” appear light and the “groundmass” dark (DIXON VAUOHAN, 1911; BATHURST, 1959b).

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Pseudomicrite: it is a genetic term applied to micrite formed by secondary changes such as “grain diminution” or “degenerative recrystallization” of faunal and floral material (WOLF, 1963b). The causes of this process are still poorly understood. (See micrite and orthomicrite, and Table ILL) Pseudomorphic replacement: a diagenetic process whereby the original character of a limestone is altered during dolomitization; skeletal material, and specifically crinoidal material for example, is replaced in such a manner that single crystals of dolomite are in optical continuity with the calcite of the original crinoid fragment. The term contrasts with the process termed impingement, which does not give rise to the optical continuity of dolomite crystals with the original crinoid fragment (LUCIA,1962). Pseudo-odlites: calcareous pellets which have cryptocrystalline and/or microcrystalline internal texture, and are of similar size and shape to oolites but lack concentric structure. These particles can form as faecal, bahamite, and algal pellets, whereas others are formed by the abrasion of micritic limestones. In general, “pseudo-oolite” is a synonym of “pellet”. Pseudosparite: see sparite. RecrystaZZization: this term is usually used loosely for a number of processes that include inversion, recrystallizationsensu stricto, and grain growth, all of which may result in textural and crystal-size changes. Recrystallization 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 size. Its coarseness depends on the density of the initial nucleation. Where the nuclei are widely spaced there is an intermediate porphyroblastic 1958). As used by FOLK (1959), recrystallization is a process wherein stage (BATHURST, original crystal units of a particular size and morphology become converted into crystal units with different grain size or morphology, but the mineral species remains identical (1958, 1959b) presented criteria for before and after the process occurs. BATHIJRST recognition of various diagenetic fabrics and made a plea for the elimination of the term “recrystallization”in favor of specific recognition of the individual process. Aggradation recrystallization results in the enlargement of the crystals, whereas degradation recrystallization gives rise to a relative decrease in size of crystals or grains. The latter process has also been termed “grain diminution” and “degenerative recrystallization” (see text). 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. A reef is thus a skeletal deposit. By contrast, a bank is a skeletal limestone deposit formed by organisms which do not have the ecologic potential to erect a rigid, waveresistant structure. Reef and bank deposits, therefore, denote origin, whereas the terms btoherm and biostrome denote shape (LOWENSTAM, 1950; Cmm, 1952; NELSON et al.,

1962). Reef complex: the aggregate of reef, forereef, back-reef, and inter-reef depositswhich are bounded on the seaward side by the basin sediments and on the landward side by the lagoonal et al., 1962, for an exhaustive treatment of skeletal limestones, sediments. (See NELSON including reef terminology.) 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 (€€AMBLETON,

1962).

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 isprismatic in structure and is radial in respect to the depositional surfaces. The fibrous calcite, or reef

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“tufa” is deposited directly upon the framework of the reef and within the various voids and interstices, from supersaturated waters. The mechanism may be largely physicochemical, or, aided by profuse algal growth to extract COa from the water, may also be biological to biochemical. Development of reef tufa follows and/or accompanies growth of organicframe-builders,and precedes infilling of detritus such as limemud, calcarenite, 1955; PARKINSON, 1957; WOLF, 196%). Many types of “reef tufa” have etc. (NEWELL, been called “stromatactis”. Reefal: as used herein, a purely descriptive and non-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. Rim cement: cement which grows into interparticle voids and is optically continuous on single crystal particles such as crinoid fragments. Thus, the host is a single crystal and the cement forms a single rim in lattice continuity with it. The overgrowth is a continuation 1958). of this crystal, and the overgrowth can form by filling the pore space (BATHURST,

Sacchuroidal: a descriptive term which, in general, means “sugary” texture. More specifically it is a product or result of dolomitizationin which crystallizationor recrystallization effects a new texture. It may be first-stage crystallization,but more commonly is recrystallization that occurs early in the newly-deposited lime-mud. It does not alter gross primary structures of the sediment such as ripple mark, thin bedding, 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 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 shoreline sands). Recrystallization of the original calcite grains destroys the original particle-size distribution and substitutes a new, highly restricted crystal-size distribution ranging from medium- to coarsesand dimensions @APPLES, 1962). Secondary: a general term applied to rocks and minerals 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. (See primary.) Skeletal: pertaining to debris derived from organisms that secrete hard material around or within organic tissue. The term biochtic is considered to be synonymous with skeletal. (NELSON et al., 1962, use “skeletal” in a somewhat different sense. See also LEIGHTON and PENDEXTER, 1962, for discussion of term skeletal and skeletal limestone.) Solution tranrfer: this is a translation of the German Losungsumsatz. It refers to the solution

of detrital particles around their points of contact where elastic strain and solubility are enhanced (pressuresolution), followed by redeposition on less strained particle surfaces (BATHURST, 1959b).

Sparite: it is an abbreviation of, and is therefore synonymous with, sparry calcite. Sparite, as used here, is a descriptive term applied to any transparent or translucent crystalline calcite and aragonite. It can occur in numerous morphologic forms, namely, granular, drusy, fibrous, and blady. Three possible origins are recognized: (I) precipitation into open voids, (2) recrystallization, and (3) grain growth. The frst is distinguished by adding the genetic prefix ortho-, and the latter two by pseudo-. Microsparite ranges in diameter from 0.005 mm to 0.02 mm, whereas sparite is larger than 0.02 mm (Table III). The prefix dolo- is used to indicate spany dolomite crystals, i.e., dolomicrosparite and dolosparite. Some workers prefer the prefix calc- to distinguish calcsparitefrom the dolomitic variety, but to some sparite is automatically understood to mean the calcareous variety.

DIAGENESIS OF CARBONATE ROCKS

32 1

Sparry: see sparite. Sphrulite: as used here, a small spherical or spheroidal particle composed of a thin dense calcareous outer layer with a sparry calcite core. It can originate by recrystallization or grain growth and the central sparite is then a typical pseudosparite. On the other hand, spherulites can be formed as minute bodies by biological processes and the open space is then filled by orthosparite, e.g., Calcisphaera of algal origin. As defined by PETITJOHN (1957), however, spherulites are minute bodies of oolitic nature in which only a radial structure is visible. The surfaces of such bodies, unlike those of oolites, are somewhat irregular. Stromatactis: these are open-space structures with horizontal flat to nearly flat bottoms, and are filled by internal sediments and/or cement. They have been termed “reef tufa” by some. Their genesis has been variously interpreted as being caused by the burial of soft organism 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 (WOLF, 1963a, 1965~).It seems that Sfromafactisare most common in micritic limestones formed by calcareous Algae, that left little or no evidence in most occurrences in Great Britain, North America, etc., but are well preserved in one Australian locality (plate I-XXIV). SCHWARZACHER (1961) described the fabric of some Lower Carboniferous reefs of northwestern Ireland, and noted that in some places calcite grows into what at one time must have been hollow spaces; this was interpreted to represent either recrystallization phenomena or remains of frame-building organisms, i.e., Stromutuctis. Schwamcher referred to BATHUR~T (1950) who recognized the cavity nature of structures he described under the name of Stromatactis, and tentatively interpreted them as hollow molds of organisms which presumably disappeared at an early stage in diagenesis. LOWENSTAM (1950) regarded Stromatactis as a rigid frame-building organism. (See “reef tufa”.) Stromatolite: laminated sediment formed by calcareous Algae, which bind 6ne detritus and/or calcium carbonate precipitated biochemically. The deposit may form irregular accumulations or structures that may remain fairly constant in shape, e.g., Collenia. Subaerial: formed, existing, or taking place on the land, in contrast to subaqueous. Sublittorul: belonging to, inhabiting or taking place in the bottom environment extending from low-tide level to approximately 100-1 50 ft. below low-tide level. Sucrosic: contraction of saccharoidal, thus meaning “sugary” texture. Supergenic: a term applied to those processes and products caused by material derived from descending fluids and gases. (See hypogenic.) Supralittoral: belonging to, inhabiting or taking place in the near-shore region above high-tide level. Syndeposition: see syngenetic. Syneresis cracks or vugs: cracks or vugs formed by a spontaneous throwing off of water by a gel during aging. THOMAS and GLAISTER (1960) pointed out that in some Mississippian carbonates of the Western Canada Basin calcium carbonate evidently was precipitated as a colloidal gel encrusting leaves of sea plants (photochemical removal of carbon dioxide from sea water by the plants, causing precipitation). The end-result was the production of cryptograined limestone which contains “syneresis” cracks and associated primary contraction vugs. When these vugs are filled by sparite, they resemble “birdseyes”.

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K. H. WOLF

Syngenesis: as used here, the processes by which sedimentary rock components are formed simultaneously and penecontemporaneously. Syngenesis has been subdivided into syndeposition and pre-diagenesis. The former comprises processes responsible for the formation of the sedimentary framework, whereas the latter is responsible for those parts that were introduced subsequently but before the principal processes of diagenesis began, i.e., internal mechanical sedimentation. The latter does not constitute part of diagenesis because its products are formed by ordinary sedimentary deposition and do not alter the sedimentary framework as such. Syntuxiul rim: a mechanism of replacement overgrowth, which develops during diagenesis as a syntaxial extension of a detrital single crystal (e.g., a crinoid fragment). During recrystallization or grain growth some of the newly formed crystals become opticaUy oriented with a detrital grain, commonIy a crinoid ossicle, and form the so-called syntaxial rim. It is not to be confused with similar optically oriented overgrowth formed by chemical precipitation of calcium carbonate in voids (BATHIJRST,1958, 1959b). In studying diagenetic effects of a crinoidal sediment, LUCIA(1962) observed that the textural relationships between lime-mud and calcite overgrowth suggest that rim cementation is the dominant process; furthermore, dolomitization occurred after rim cementation. Terrigenous: land-derived; refers particularly to sediments resulting from erosion of the land. Travertine: calcium carbonate, CaCOa, usually of light color and commonly concretionary and compact, deposited from solution in ground and surface waters. It is a more dense and often banded variety, in contrast to tufa. (See tufa and caliche.)

Tufa: a chemical, spongy, porous sedimentary rock composed of calcium carbonate, deposited from solution in the water of a spring or of a lake, or from percolating ground-water. (See reef tufa; travertine; caliche.)

Vuterite: it is a metastable hexagonal form of calcium carbonate, CaCOa. It is doubtful if it occurs in the geologic column, but if so, such occurrences are rare. Welding: term used in reference to crystal welding, in which discrete crystals and/or grains become attached one to another during compaction and in large measure through diagenesis. Pressure-solution, and solution transfer are likely the operative processes. Welding can continue beyond normal diagenesis to epigenesis. Winnow: the Old English word is windwiun, and has reference to exposure to the wind such that lighter particles are blown away, thus winnowing grain, and the word winnow is a contrivance of winnowing grain. In this chapter winnowing can apply only to eolianites, and not to water-moved limestones. The term washed is preferred in the latter case.

Chapter 6

SILICA AS AN AGENT IN DIAGENESIS EDWARD C. DAPPLES

Department of Geology, Northwestern University, Evanrton, Ill. (U.S.A.)

SUMMARY

In this chapter the precipitation of silica is examined from the viewpoint of solubility values obtained in the laboratory, and its occurrence as a deposit in rocks. Approximate ranges of solubility in natural waters are compared with similar data gathered in the laboratory which show solubility in s i m c a n t amounts to be directly proportional to temperature. Precipitation of chert in cratonic carbonates is shown to post-date accumulation of the sediments, and is presumed to be part of the lithification process. Eugeosynclinal deposits such as siliceous shales and porcellanites indicate early replacement by silica of clay minerals and carbonates, and this is considered to be the mechanism of chertification. In some carbonates the time of precipitation and replacement follows dolomitization, whereas in many stable cratonic occurrences deposition of chert and quartz precedes widespread dolomitization.

INTRODUCTION

Despite the enormous quantity of silica to be found as various deposits in sedimentary rocks, understanding of the mechanisms by which it is precipitated in nature is still shrouded in uncertainty. Except for the laboratory data which have been supplied through many carefully controlled analyses, the actual solubility of quartz in natural waters is as yet not precisely established. Such is the case because of the relative insolubility of the substance and the very slow rate of attainment of the solid and solution equilibrium. Moreover, the distinct difference between the solubilities of amorphous silica and quartz confound the problem further by the uncertainty concerning the amount of silica which is colloidally suspended, a factor which tends to increase the values of “soluble” silica above that in true solution. Recent studies have served to strengthen the long held concept that much of the silica held in solution in river waters must be precipitated upon entering the oceans. Despite the increased tempo of bottom sampling in coastal and deepwater oceanic areas, however, no localized concentrations as masses of primary gel

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silica have been discovered. Among geologists there are some proponents of the concept that primary accumulation of silica in amorphous globules will be found. Nevertheless, there is increasing reason to favor the interpretation that precipitation of silica occurs as very small discrete particles along with clay minerals and carbonates in clay-size dimension. Concentration of silica in the form of concretions is increasingly regarded as an example of accumulation following deposition of the enclosing sediments; and the process of concretionary growth as a migration of “ions” to some center of precipitation. The occurrence of exceedingly siliceous sediments associated with former eugeosynclines has been supported by additional observations, but the problems of the source of the silica and the time of its precipitation remain unsolved. Such sediments not only are argillaceous and arenaceous, but also comprise those, the composition of which indicates that they are lacking in land derived aluminum-silicate debris (for example, beds of novaculite). Are these sediments to be regarded as replacements of former carbonate deposits? Carbonates of the stable craton rather commonly are chert-bearing, the chert occurring in the form of nodules, lensing beds, and veinlets. All such occurrences locally are restricted to thin beds separated by intervals of “chert-free” carbonate in which the silica content tends not to exceed 10%. Among such shelf sediments the chert zones are restricted to the carbonates, whereas the shales and sandstones uniformly are free from concretionary silica. Similarly, cherty carbonates commonly constitute some ancient interreef deposits, whereas the associated carbonate reefs and banks tend to be characteristicallylow in chert. A parallel case occurs with evaporites. Bedded anhydrites, salts and associated carbonates are low in silica, whereas some of the carbonates rimming the evaporite basin are cherty. In terms of quantity of silica precipitated, the range clearly is lowest among cratonic sediments, increases in the so-called miogeosynclinal carbonates (but remains low in shales and sandstones), and attains a maximum in the eugeosynclinal strata where sandstones, shales, and carbonates are prominently silicified. Except for the ideal eugeosynclinal sediments, local distribution of silica is sporadic; and highly siliceous strata are interbedded with strata of low silica content despite no apparent change in environment of deposition. All presently available evidence leads to the conclusion that outside the limits of eugeosynclinal sediments silica was not deposited rhythmically as concretionary nodules to be buried with the carbonate sediments; whereas in the eugeosynclinal belts continuous precipitation and concentration of silica appears to have been possible. The aspect of the precipitated silica, however, and the ways and means of its emplacement still remain in the realm of uncertainty. Petrographic studies as a whole tend more and more to favor interpretations of a diagenetic origin rather than primary precipitation; and the important processes are those of complex replacement by silica of pre-existent mineral matter. The reader must be aware that insofar as demonstrated mechanisms are concerned, there has been little advance during the past 50 years in describing the

SILICA AS AN AGENT IN DIAGENESIS

325

origin of flints and geodes, as well as nodular and bedded cherts. One is, however, in a better position to limit concepts of origin on the basis of restrictions imposed by laboratory studies, and by paragenetic relations between silica and its associated minerals. The principal purpose of this chapter is to outline the restrictions on the precipitation of silica in order to clarify for the reader the currently preferred mechanisms of deposition. SOLUBILITY OF SILICA

Laboratory studies

Ranges of solubility of amorphous silica have been prepared independently by a number of investigators with sufficiently close agreement to permit fixing the limits of silica in solution either as Si02 in tetrahedrally-bondedunits or as the monomeric acid H4Si04, and most probably in some equilibrium distribution between the two. In this connection, the saturated solutions of silica have been prepared by two methods essentially: ( I ) allowing supersaturated solutions to come to equilibrium by precipitating Si02, and (2) allowing silica to dissolve from suspensions until equilibrium is reached (KRAUSKOPF, 1956; 1959, fig.2). By either method, ranges of 100-150 p.p.m. silica in solution have been attained as equilibrium values at temperatures of 22-27°C in about 70 days. Although this appears to be a somewhat sluggish equilibrium with respect to laboratory experimentation, the rate virtually is instantaneous from the viewpoint of geologic rates of deposition. As the temperature is elevated, the solubility rises and at 150°C over 600 p.p.m. of silica are in solution. Natural hot springs also may contain silica in amounts commensurate with those obtained in the laboratory. In some there is evidence of the presence of colloidal silica, or supersaturation, as silica is precipitated from the solution on standing. Silica is also precipitated from such highly concentrated solutions when passing over previously deposited siliceous sinter; whereas Algae do not appear to exert any important control on the precipitation (WHITE et al., 1956, p.39). Certain ground waters rich in silica are saturated approximately at the temperatures of such waters, but most ground waters are considerably undersaturated at the existing temperatures (Table I, Fig. 1). There is reason, therefore, to consider that only exceptional natural waters are saturated with respect to silica, and some very special mechanism must be envisaged as responsible for removal of silica from stream and oceanic water. Quartz is not readily soluble and is acknowledged to be very sluggish in reaching the equilibrium solution value. SIEVER (1962, p.135) considers it most likely that the attainment of equilibrium may be in fact in the order of spans of geologic time. The presence of authigenic growths of quartz on present-day sands off the New Jersey coast and Lake Michigan shores suggests that perhaps Sever’s statement may be somewhat liberal. One would agree that some natural waters

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may not be in equilibrium with quartz, whereas they most assuredly are with respect to amorphous silica. Sufficiently careful work has been done to establish with small error the range in solubility of quartz with change in temperature. The slope of the curve has the same general trend as that of amorphous silica, but at the same temperature the solubility is much lower pig.1). According to SIEVER (1962, p.133), the slope of the quartz solubility curve tends to rise somewhat more steeply than that for amorphous silica, and the two are considered to join near the critical temperature of water. TABLE I PARTIAL ANALYSES OF NATURAL WATERS (AFI'ER HEM,

1959)

Variety of water

Issuing from rock at

P.P.M.

Silica

Flowing well Rhyolite

99 103

0.04 0.0

1.4 9.2 1.1 6.7

50 38

Owyhee Co., Idaho Rio San Antonio, New Mexico

60 49 38

0.01 0.01 0.24

7.5 12 7.8 6.6 -

-

48

0.03

20

-

-

Harney Co., Oregon Umatilla Co., Oregon Yakima Co., Washington Greenlee Co., Arizona

71

0.0

8.8 7.9

-

Albuquerque, New Mexico

62 29

0.01 0.33

4.4 8.1 1.7 7.1

15

Dunning, Nebraska Burke Co., North Carolina

43 7.6 2.0 7.4 8.4 6.4

13 22 14

Columbus, Ohio Memphis, Tennessee Blount Co., Tennessee

17

6.7

-

Elizabeth City, North Carolina

1,330

7.4

-

Okmulgee co., Oklahoma

15

Ishperning, Michigan Weighscale, Pennsylvania Ransom Co., North Dakota Fulton, Mississippi

Rhyolite, possibly basalt Basalt Basalt Volcanic rock Conglomerate and sandstone Sand? (Loup River) Mica schist Iron

Sand and gravel Sandstone Dark shale Calcareous sands and clays Sandstone (oil well brine)

sfoz

pH Fe

20 12 26

2.3 2.9 10

41

8.1

9.1

32

8.1

Mg

Temp. Locality

("0

17

Iron mine (siliceous) Stream

21

0.34 15

17 68

7.5 3.0

-

Sandstone

23

4.8

35

-

-

1.5 6.3

17

Sand

7.9

11

327

SILICA AS AN AGENT IN DIAGENESIS

TABLE I (continued) Variety of water

Issuing from rock at

Magnesium Dolomite

P.P.M. SiOz 8.4

Fe

7.4

14

1.4

40

6.7

12

42

8.2

12

614

-

18

71

8.3

-

Carlsbad Cavern, New Mexico

7.0

-

Wisconsin River Grant Co., Wisconsin Calhoun Co., Alabama

Serpentine Dolomitic limestone

80

5.6

11

0.04

4.2

Jefferson City, Tennessee Sidney, Ohio

22

18

Large stream in dolomitic limestone (locally)

Temp. Locality ("C)

0.24

Dolomite Olivine tuff breccia

31

pH

Mg

-

0.13

8.5

Quartzite Limestone (primarily)

12

0.03

12

7.6

18

13

0.01

43

7.4

-

Dolomite

18

0.39

33

8.2

10

Navajo Indian Res., Arizona Colusa, California

Chaves Co., New Mexico Milwaukee Co., Wisconsin

Influence of p H

Since the classic study of C O R ~ N(1941) S on the relative solubility of silica with change in pH, additional data have provided points for construction of a reasonably precise curve showing this relation. The slope of the curve tends to approximate horizontality and to pass through essentially the same values between the pH limits of 2 and 8.5. At higher pH the solubility rises abruptly attaining values approaching 5,000 p.p.m. at pH 11 (KRAUSKOPF, 1959). Such extremely basic waters are most exceptional near the earth's surface; hence, only in very localized environments do pH values approach such magnitude. Where an Algaematte is growing on carbonate sediments, periodically exposed to the atmosphere, interstitial waters attain high pH values; and in these waters concentrations of silica could be exceptionally large (BAASBECKINGet al., 1960). Observations in the natural state

Recent data supplied by KELLER et al. (1963) on concentrations of silica in solution as a result of grinding of natural silicates show rather high values (7.5-8.6 p.p.m. for quartz and as much as 68.6 p.p.m. for pumice). Their analyses of the silica content of various glacial milks, however, show values which are considerably

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500

400

h

E

4 300 Q

Y

U 0 ..v)

200

100

T e m p e r a t u r e (OC)

Fig.1. Solubility of silica with change in temperature of solutions. The range of values for amorphous silica are those obtained in the laboratory by Alexander, Okamoto, Krauskopf, and others (KRAUSKOPP, 1959, p.5). The curve for amorphous silica in hot-spring water is based on data taken from WHITEet al., 1956; whereas the data on solubility of quartz are principally obtained fromsiever, Van Lier, and others (SEVER,1962, p.134). The single points showing values for the et al. (1963); maximum and mean for silicate rocks, and for quartz are those reported by KELLER and are solubilities obtained by crushing the specimens in distilled water. Scattered values for natural waters are selected analyses of ground waters in silicate rocks (Table I; HEM, 1959). Note that such waters are supersaturated with respect to quartz and some tend to approach values of solubility of amorphous quartz.

lower, approximately 0.4-17 p.p.m. (KELLER and REESMAN, 1963). Presumably, abrasion during erosion is not to be regarded as an important mechanism for releasing silica for solution. A representative group of analyses of surface and underground waters in the United States tend to reveal that the silica concentration at the temperatures recorded may be somewhat above that of the sblubility curve of quartz, but clearly below that for amorphous silica (Table I). For example, Mississippi River water near the delta mouths (temperature approximately 20 "C) averages 4-7.5 p.p.m. of silica (BEN et al., 1959); whereas the average concentra-

SILICA AS AN AGENT IN DIAGENESIS

329

tion in the surface waters of the Gulf of Mexico is less than 1 pap.m.(0.1 1 p.p.m. as reported by BIENet al., 1959). Elsewhere, ocean waters tend to contain no more than 4-5 p.p.m. of silica. Because these values clearly are lower than the solubility of amorphous silica at such temperatures, one can assume the surface waters of the oceans to be under-saturated with silica. Certain authors point out that it is likely that at the temperatures which prevail in oceanic bottom waters, the latter are approximately saturated with respect to quartz (SEVER, 1962). Numerous authors have expressed the opinion that the low values for silica in solution in the seas are due to its extraction by organisms (SEVER,1957). Opaline tests of organisms do not return to solution readily as indicated by commonly well-preserved radiolarian and sponge remains. Also, experimental work (D.E. White, as quoted by KRAUSKOPF, 1959, p.9) showed that opal did not completely saturate a solution of sea water even after 2 years. Many opaline remains of organisms, therefore, are likely to be buried before being dissolved. Some siliceousfossil remains commonly found in carbonates are in an excellent state of preservation, whereas others clearly have corroded boundaries and have been partially dissolved. The carefully controlled analyses of BIENet al. (1959) show that Mississippi River water loses silica upon entering the Gulf of Mexico in a quantity greater than that which can be attributed to removal by organisms. According to their interpretation, electrolytes cause silica to be absorbed on the surface of suspended inorganic particles; or facilitates co-precipitation of silica with such inorganic particles to cause a most significant depletion in dissolved silica. Accordingly, oceanic waters would be kept under-saturated with respect to silica, because concentration of the latter would be lowered below the saturation level in the turbid near-coastal waters. Beneath the depositional interface, the tendency for the chemical system to be closed is greater than in the open waters; hence, interstitial waters have a greater opportunity to come to equilibrium with the siliceous material incorporated in the deposit (DAPPLES, 1959, pp.37-38). Although the data still are scanty, the concentration of dissolved silica in interstitial fluids in uncemented marine sediments tends to increase with depth. Such values are considerablyin excess of the solubility of quartz, but less than those for amorphous silica (EMERYand F~ITTENBERG, 1952). One has reason to believe that below the levels currently analyzed, but in sediments as yet unlithified, local concentrations of silica in pore-water may approach the saturation level of amorphous silica.

BEHAWOR OF SILICA IN EARLY DIAGENESIS

Silica in carbonate reefs

Indirect evidence of originally low concentration of silica precipitated, or accumu-

330

E. C. DAPPLES

lated, in carbonate deposits appears from analyses of reef “core” and “flank” bedded calcarenites. Sufficient analyses are now available to demonstrate that these carbonates tend to be unusually low (&2%) in silica. The accumulations represent the organically precipitated carbonates constructed as rocky, wave-resistant structures, and the wave and current eroded debris of such original accumulation. In essence, such carbonates reflect the absence of inorganic precipitation of silica at such sites. For example, NEWELL et al. (1953, p.67) report that in the carbonate association of the Capitan Reef, fossils in the basin calcarenites have been selectively replaced by silica, but those of the inclined reef talus rarely are silicified, and the silica content of the coarse talus is low. Important amounts of silica as fossil replacements, nodules, and crusts, however, occur basinward in the talus, and micritic basin-limestones are cherty. In the interreef deposits containing land-derived clastic sediments the chert content rises; and tongues of such sediments projecting into the reef complex carry nodules and thin beds of chert. In evaporite basins rimmed by reefs the higher concentration of salts in sea water necessary to produce salt beds has not concomitantly resulted in the precipitation of silica (DELLWIG, 1955). Indeed, one is to conclude that the concentration of silica in such waters was so low that virtually no silica could be precipitated even under conditions of high concentrations of electrolytes. Time-stratigraphic equivalents of such strata, primarily carbonates, but containing some interbedded shale or a significant clay insoluble fraction contain chert. The association of chert and clay-size, land-derived, detritus appears to be of considerable significancein the diagenesis of silica. Experimental work by JAVERING and PATTEN (1962) suggests that cold, acid solutions supersaturated with silica should, on contact with limestone or dolomite, generate COZ and precipitate silica. A cold, neutral solution supersaturated with silica but without Na-ions should be capable of moving long distances through carbonate rocks until making contact with either C02 or Na-ions at which point silica should precipitate. Conceivably interstitial waters in sediment containing very finely-divided silica could become saturated with respect to silica. Migration of such waters into positions where Na-ions or higher C02 concentrations occur, could result in precipitation according to the mechanism suggested by Lovering and Patten. Bedded siliceous deposits Bedded chert Certain rather typical strata of eugeosynclinal sites are identified as siliceous shales and bedded cherts. They are distinct from other siliceous accumulations because of the large volumes of silica involved and the strict parallelism of the siliceous strata to bedding. Many of these are exceedingly thinly laminated suggesting slow deposition in quiet water. Investigators of such strata are uniformly impressed by

SILICA AS AN AGENT IN DIAGENESIS

331

the characteristics which point to very early precipitation of silica, either as a primary accumulation with the clastic sediment or as a diagenetic product of very early appearance (TALIAIWRRO, 1934, p.196). Siliceous shales Typical examples of siliceous shales occur as units within the Stanley Shale (Mississippian) of the Ouachita Mountains in Arkansas and Oklahoma (GOLDSTEIN,1959). The siliceous strata are dark, poorly to conspicuously laminated, silty, micaceous, carbonaceous, and pyritiferous. Where conspicuously laminated in thin section, crenulated lenses of clay minerals intermingled with cryptocrystalline or isotropic silica can be recognized. Sponge spicules, Radiolaria, spore exines and similar fossils are characteristic, and range from well to rather poorly preserved. Locally, cherts are fractured or brecciated and recemented by silica as a veined mass. Porcellanites All gradations exist between porcellanites, siliceous shales, and cherts. The term porcellanite is not uniformly applied by all authors, but, as used herein, it refers to a siliceous bed which has the outward, somewhat porous appearance of unglazed porcelain, and which is neither a true chert nor a siliceous shale (TALIAFBRRO, 1934, p. 196). As the amount of aluminous impurities increases, the characteristic wax-like luster of chert gives way to a more granular aspect, which because of its increased porosity, assumes a dull porcelain-like appearance. In the more extreme examples of development, the chert appears to have recrystallized to elongated quartz crystals often with well-developed crystal terminations and arranged in what is considered to be a random orientation. Individual crystals are interlocked into a framework which is either porous or dense, depending upon the amount of granular interstitial chert present. The texture appears to be primarily developed by recrystallization of the chert. Porcellanites are gradational also into fine-grained tuffs in which the glass has altered to chert for the most part. Such porcellanites have much the same outward appearance as others; but in thin section one can note a significant percentage of highly angular fragments of unaltered feldspar, In many such cherty tuffs the percentage of detrital quartz is low; whereas where the volcanic mixture is diluted by land-derived detritus, gradation into a tuffaceous sandstone occurs. With such a transition, the typical appearance of porcellanite tends to give rise to that of a quartzitic sandstone or siltstone. Transition of porcellanite into siliceous shale occurs through dilution of the chert by clay minerals and some micas. This transition is accompaniedby appearance of well-developed lamination and corresponding loss of wax-like luster. In thin section the relation between chert and clay minerals is difficult to see; but the toughness of the rock and the tendency for smallpieces to break with conchoidal

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fracture suggests that granular interlock exists and that some diagenetic recrystallization may have occurred. The relation of porcellanites to carbonate-rich beds is poorly understood. Some porcellanites contain rhombs of siderite or dolomite which have replaced part of the silica; whereas other examples show that chert has replaced what most certainly must have been carbonate lenses or thin beds. Among certain of the shale sequences occur beds which can be recognized as former limestones now intensely silicified. Where they are as yet calcareous, they may contain as much as 20 % silica which may have been deposited penecontemporaneously with the carbonates. Elsewhere, similar beds can be seen to be for the most part replaced by secondarily deposited chert. Interbedded with such former carbonates are cherts and siliceous shales which by many geologists are considered to represent primary deposits, or to have become siliciiied very early in their burial history. A currently preferred mode of origin is diagenetic;namely, the high percentage of silica is the product of alteration of volcanic ash deposited in association with land-derived detritus (GOLDSTEIN and HENDRICKS, 1953, p.441). In the process of alteration of the ash, local bottom waters are considered to have had in solution an abnormally high content of silica. Such waters are considered a highly suitable environment for the growth of silica precipitating organisms such as sponges and Radiolaria; hence, the abundant presence of their well-preserved remains. The excellence of the state of preservation of such tests argues against their being the direct source of the chert in the sediments. TIME OF SILICIFICATION

Dolomitization and precipitation of silica

Certain carbonates, which constitute part of the so-called miogeosynclinal strata, contain significant quantities of chert as nodules, thin beds, and veins engulfed within the carbonate mass. Often such strata are extensively dolomitized and the entire formation may be dolomite over distances of many square kilometers. That such dolomitization is a secondary process is well established on the basis of such evidence as the presence of dolomitized fossil remains, which as living organisms precipitated shells of aragonite and calcite. Uncertainty exists as to the time when the dolomitization occurred, but an increasing amount of evidence points to the process as being early in the sediment’s history, principally prior to general lithification, and promoted by slightly positive tectonism in the depositional site. Paragenetic relations between chert and dolomitization are extremely useful in helping to establish the time duringwhichsilicificationof the carbonate occurred. Individual euhedral crystals of dolomite are found in some cherts. This relation is interpreted to indicate that dolomitization preceded precipitation of the silica as a post-depositional process.

SILICA AS AN AGENT IN DIAGENESIS

333

Some outcrop exposures of dolomite reveal broken edges of thin nodules and lenses of chert to overlap in a sort of imbrication. Rupture of the chert obviously has occurred after its solidification inasmuch as the broken edges of adjacent fragments may be joined into a single piece. Excellent examples of such an occurrence have been described by DIETRICH et al. (1963, p.649) in which the continuous unbroken overlying lenses of chert prove that fracture of the chert is not due to any tectonic forces. They attributed rupture of the chert to syneresis of the gel during a period of exposure to air, drying, and shifting of fragments before deposition of overlying carbonates. In similar examples observed by the writer, it is possible to interpret the fracturing of chert to have occurred after deposition of perhaps several centimeters of carbonate material; thin laminae of carbonate are continuous but sag into the gap in the pieces of chert, suggesting that rupture occurred before lithification of the carbonate. The presence of clear dolomite rhombs within ironoxide stained chert was regarded by DIETRICH et al. (1963, p.650) to suggest that some dolomitization occurred prior to the development of the chert lenses. Moreover, the pelletoidal texture and gradational boundaries of the chert indicate a replacement of calcium carbonate sediment. Similar paragenetic sequences can be substantiated in other strata elsewhere where remnants of dolomite exist within the chert, and skeleton outlines of chert nodules have marginally replaced and surrounded masses of the dolomite. Overgrowths on quartz grains in carbonates

Certain thin beds of quartz sandstone interbedded with carbonate rocks contain quartz grains of exceptional roundness. Some of the grains show additions of quartz as well-terminated overgrowths, but rarely such grains are completely surrounded by chert deposited as a rind around the grain (DIBTRICHetal., 1963, p.660). Generally the time of precipitation of such silica is difficult to date, but Dietrich and his associates have been able to show rhombs of dolomite included within the quartz overgrowths. The order of paragenesis is not always the same as shown by examples in which secondarily enlarged grains of quartz are engulfed and partially replaced by calcite. Such calcite was precipitated as a cement, but has by replacement of quartz grains virtually destroyed the original sandstone texture (WNDA, 1963). Silica in cratonic carbonates Chert and chalcedony Chert is commonplace among shelf and intra-cratonic basin carbonates as nodules, lenses, thin beds, and veins with much the same appearance that is observed in miogeosynclinal strata. Some occurrences are lensoid bodies concentrated in the axial positions of small, recumbent, primary folds, which generally parallel the

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bedding; whereas other chert masses partially replace the carbonate strata (NEWELL et al., 1953,p.162). Evidence for replacement of carbonate by chert is abundant, and boundaries between the carbonate and silica are irregular and transitional between the two rock types. Irregular masses of microcrystalline quartz (chert), commonly only 1-3 mm in greatest leDgth, appear as partial replacements of fossils, or are associated with detrital quartz grains. Replacement relations visible on etched polished surfaces demonstrate that such chert nodules are secondary; and inasmuch as their shape is generally elongated parallel to bedding, they are considered to have been precipitated as amorphous silica before compaction of the sediments. The source of the silica is not evident and does not appear to be derived from solution of associated detrital quartz grains although commonly their surfaces are pitted and irregular (LAMAR, 1950, p.30, 32). Fibrous quartz (chalcedony with quartz crystallites oriented with C-axis perpendicularly to the length of the fibers) is also a common precipitate in cavities in shelf carbonates (PELTO,1956). Such precipitation appears to distinctly follow lithification of the carbonate rock. In some occurrences isolated carbonate crystals within the chalcedony, and along the borders between chalcedony and the host rock, are dolomite primarily. Such crystals have different optical properties and composition than limestone of the host rock. They tend to have well developed crystal outline and cut primary structures, such as oolites. Silicification slightly preceding widespread dolomitization in time, is noted in certain shelf carbonates in which well-preserved fossils occur in the chert; whereas their counterparts in the dolomite show very poor outline. In the same rocks, dolomite is observed to partially replace chert along boundaries of the nodules (DAPPLES, 1959, p.49). For reasons which currently are unexplained, certain cratonic limestones show a paragenetic sequence by which silica is deposited on detrital quartz grains as overgrowths of clear quartz rather than as chert. Such a case has been observed by GLOVER (1963, p.40) in the Septimus Limestone (Carboniferous, Western Australia) the composition of which is 6 % sparry calcite, 40 % quartz, and 53 % dolomite. The paragenetic sequence of importance which he reported is formation of zoned dolomite, precipitation of quartz overgrowths, followed by crystallization of sparry calcite. Precipitation of silica as quartz rather than chert can be substantiated in other carbonate rocks, but only a very crude correlation is noted between precipitation of quartz as overgrowths where detrital quartz is abundant, and precipitation as chert where quartz grains are rare. Geodes

Occurrence of geodes in shelf carbonates is far more common than in the correspondingbasin sedimentsand is not associated with bedded cherts. Geodes characteristically have an exterior rind of chalcedony of subspheroidal outline, which

SILICA AS AN AGENT IN DIAGENESIS

335

presents a rather sharp boundary with the enclosing carbonate. The relations suggest that beginning with some central nucleus, such as a shell containing an initial opening, chalcedony was deposited on the outer margins and in fissures to expand the dimensions of the opening. Such precipitation was in progress when the enclosing carbonate was still in an unconsolidated condition; hence, it could be pushed aside. In other examples, actual replacement of the carbonate by silica was accomplished. HAYES(1964) has presented evidence for replacement of carbonate concretions by chalcedony to form the rind of the geode. After precipitation of the outer chalcedonicring, megacrystallinequartz was depositedto completetheinternal aspect of the geode. In this connection, there appears to be a relation between the precipitation of varieties of silica and the nature of the surface on which precipitation occurs. Megacrystalline quartz is precipitated in cavities lined with chalcedony but not against the carbonate rock per se. Chalcedony and chert are preferentially deposited against the carbonate but rarely upon macrocrystalline quartz. Replacement of fossils Selective replacement of fossils by silica is a common occurrence, but the mechanism of replacement and the localities where it can be expected to occur are known only on an empirical basis. An example is the occurrence in the Capitan Reef (Permian, New Mexico) where in the reef mass itself virtually no silicification is reported (NEWELL et al., 1953, p.173). In the flank deposits, however, replacements are abundant and follow a general order (Table 11). CAROZZI and SODERMAN (1962, p.401) reported that in a Mississippian crinoidal limestone in Indiana calcilutite and spar-calcite matrix between coarse fragments of fossils is replaced locally by silica, whereas the fossils rarely are altered. Table I1 lists a few examples where an order of preferential silicification has been reported. If these examples are representative of what would be observed on the basis of future extensive studies, one might be led to conclude that bryozoans, brachiopods, and corals generally are more sensitive to replacement by silica than gastropods, cephalopods, and echinoderms. On the basis of some of the reversals in the orders shown in Table 11,however, it may be more reasonable to consider that the preferential replacement is a function more of the distance from centers of silica precipitation, and the quantity of silica being precipitated, rather than the crystal habit or composition of the carbonate constituting the fossil shell. Replacement of calcareous oiilites Among good examples of siliceous oolites is the occurrence in Cambrian strata near Bellefonte, Pennsylvania. Here siliceous beds usually only a few centimeters thick are separated by several meters of calcareous oolitic beds, but there are all gradations between siliceous and calcareous oolites. Commonly, where the oolites are embedded in a siliceous matrix, they appear to “float” in the chalcedony, but

Iz SILICWCATION OF FOSSILS

Ordovician central Nevada (HINTZE, 1953) (HINTZE,1953)

Ordovician Permian southeastern Capitan Reef (NEWELL et al., Nevada

1953)

Bryomans, corals,punctate brachiopods Trilobites Impunctate brachiopods Mollusks

Gastropods

Brachiopods, Brachiopods bryowans, and ostracods Gastropods Trilobites

Echinoderms Cephalopods Forams Algaeandcalcareoussponges

Silurian %rnton Retf,

Pennsylvanian northeastern Nevada (Dm,1958)

Ordovician Fish Haven Dol., Idaho (GmBs, 1960)

Ordovician Tanners Creek Formation, Indiana (Fox, 1962)

IndioM

Brachiopods

Bryozoans

Colonial corals

Punctate brachiopods, corals

Solitary corals

Brachiopods

Stromatoporoids

Spirifers, productids, fenestrate bryowans

Bry ozoans, cephalopods Crinoids

solitary corals

(ZNGELS,

1963)

Echinoderms Forams

SILICA AS AN AGENT IN DIAGENESIS

337

elsewhere they are in contact with one another. In other examples, individual oolites are fractured and the fissures are sealed by quartz, or carbonate particles of irregular outline are isolated in the silica matrix. Such textures are interpreted by CHOQWTTE (1955) as being illustrative of the effects of a process whereby the silica is a replacement of an original carbonate matrix and not a primary precipitate. Siliceous crust Late precipitation of chert clearly associated with weathering surfaces has been reported at numerous localities.Primarily,suchan occurrenceconsists of crusts on the weathered surface and of veins filling joints which were enlarged by near-surface weathering. These occurrences establish without question that chert can be precipitated along openings intimately related to the present-day weathering surface. In contrast to other occurrences, however, such late chert has been deposited along the walls of an opening; hence, the problem of making room for the chert in the host rock does not exist. Moreover, the boundaries with the host rock tend to be sharp, following joint planes and weathering surfaces, and the amount of replacement of carbonate by silica is at a minimum. Nevertheless, the boundary between carbonate and chert, when examined in thin section, does show silica to be invading the carbonate. The degree of the replacement, however, is small in contrast to the replacement which can be demonstrated to be early, namely, that which precedes or immediately follows dolomitization. Moreover, it is the early precipitated chert which is overwhelming in its abundance in contrast to the late chert occurring as crusts and veins.

OPAL AS A CEMENT

Opal has been reported as a prominent cement and as a vein material in sandstones which were exposed during important stages of weathering in semi-arid regions. FRANKS and SWINEFORD(1959) reported such an opal in the Ogallala Formation (Paleocene), Kansas. Their analyses showed this opal to contain significant amounts of Al,O3, Fez03, CaO, MgO, K20, and NazO. Moreover, the internal lattice shows spacings which are in accord with those of a disordered form of low cristobalite as reported by FLOW (1955). In thin section, the opal can be observed to be gradational into chalcedony and microcrystalline quartz; indicating that as a cement in sandstones, opal tends to become more crystallo1967, graphically ordered with aging under near-surface conditions (DAPPLES, fig.2). Progression of opal-chalcedony-quartz has long been observed and frequently reported (TALIAFERRO, 1934, p.206-209). Except for siliceous tests of organisms, occurrence of opal in nodular cherts of shelf' carbonates virtually is

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unreported; and its occurrence in significant proportions is associated with the thick, bedded-cherts in eugeosynclinal deposits. In the latter deposits evidences for precipitation of exceptional amounts of silica exist, and in many localities an association with some volcanism can be established. If the precipitation of opal at hot spring deposits can be accepted as an analogous mechanism, there is reason to believe that the high concentrations of silica furnished at the sites of bedded cherts has been due to a process which results in the removal of silica from solution below the saturation level. It is suggested that the equilibrium is controlled by the solubility of quartz (WHITEet al., 1956, p.52); or a combination of electrolytes and finely-divided aluminum silicates cause precipitation of silica below the normal saturation level (BIEN et al., 1959).

SUMMARY OF RELATIONS OF OCCURRENCE

As more specific information has become available with regard to the tectonic control of sedimentary facies, a deeper insight has also been gained into the distribution of diagenetic silica. In this connection, it appears that a minimum amount of such silica is to be found in strata constituting part of the evaporite basins. In these, carbonates tend to be chert-free; sandstones are cemented by calcite, or are aggregated by a simple clay bond; whereas shales tend to remain soft, argillaceous, or calcareous. In the evaporite association there is virtually no indication of early precipitation of silica; and only in strata marginal to the occurrence of evaporites are to be observed minor quantities of chert. Cavities in dolomitized beds left by solution of fossil shells and in ancient organic banks and reefs tend to remain unfilled by late precipitation of chalcedony or chert. CHILINGAR (1956) suggested that the environment favorable for the precipitation of carbonateshaving low Ca/Mg ratio is not favorable for the precipitation of silica. Strata marginal to the evaporite basin, in particular those deposited on the stable craton, tend to contain considerably more chert, always in the form of irregularly-shaped nodules, lenses, veinlets, and thin discontinuous beds. Relations of such chert to the enclosing strata indicate clearly that the silica is diagenetic in origin and that its precipitation probably occurred during the early burial stage as lithification was in progress. In this connection abundant examples exist of silicification of the beds actually preceding dolomitization. Other examples, mentioned previously, show that some dolomite was formed before the silica was deposited. In either type of occurrence, the mineral paragenesis tends to be simple, but the chert clearly replaces carbonate which was deposited as the initial sediment. Another general association which tends to be chert-free is the coal-basin assemblage. In general, carbonates tend to be uncommon in many such lithologic associations, but the shales and sandstones show only minor amounts of secondary silica. Among the sandstones calcite is the typical cement; whereas secondary

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silica appears as generally feeble overgrowths on detrital quartz grains. Nodules, lenses, and other similar occurrences of chert are noticeably absent even in those coal basins where limestones constitute an important part of the Stratigraphic sequence. Indeed, the presence of chert nodules, or silica as an abundant interstitial cement, or fossil replacement, most certainly represent exceptional cases. An abundant occurrence of chert occurs in thin carbonate units which have been deposited on a portion of the craton which has been generally positive. In the United States a classic example is the Ozark Dome (Missouri) which existed as a slightly positive feature during the Paleozoic. Strata exposed on the dome are significantly cherty; the time of deposition of chert was after the deposition of the carbonates, and primarily during the early burial stage. Abundant information can be gathered to demonstrate replacement of the carbonate by silica, and deposition of chalcedony and quartz, in that order, as alings in previously-developed solution cavities. In this region chert is precipitated also as a recent accumulation, inasmuch as crusts and fillings in joints occur associated with the present-day weathering surface. Considerably more chert is believed to occur per unit volume of carbonates in miogeosynclinal strata than their counterparts of the cratonic region (BISSELL, 1959). In the cratonic margins of the miogeosynclinal belts much nodular chert is found to interrupt bedding and clearly is of epigenetic origin. In stratigraphic sections composed of thin units of limestone in a dominantly shale matrix, the carbonate may be replaced by chert. An example of this was observed in the Fayetteville Shale (Mississippian) of the south flank of the Ozark Dome passing into the miogeosynclinal strata of the Ouachita Mountains, Arkansas (OGREN, 1961, p.24). Here, silicification of thin limestone beds can be traced progressively. Along the craton margin only a thin irregular zone of chert has replaced upper and lower surfaces of carbonate units, whereas the interior remains unaltered. Basinward, in the dxection of the miogeosyncline, chert replacement of the limestone increases until all the carbonate is silicified, but the shale remains unaffected. Where last observed, somewhat deeper into the miogeosyncline, portions of the shale are silicified, chert having been precipitated in the pore space. In the examplecited, deposition of the silica post-dates accumulation of the sediments, and has selectively replaced carbonate in advance of shale. Presumably the replacement was accomplished during the early burial stage, but as yet this has not been demonstrated. The condition deeper within the miogeosyncline is not known as the strata do not crop out and drilling has not penetrated such depths. However, strata of equivalent age in the eugeosynclinal belt of the Ouachita Mountains are more intensively silicified; hence, there is reason to consider that silicification becomes more pronounced in the shale on approaching the eugeosyncline. In this connection the aspect of the silica is worthy of note. Whereas chert in cratonic strata of the Ozark Dome occurs as nodules and lenses which transgress bedding, and in other ways manifests its epigenetic origin, silica in the Fayetteville Shale is of re-

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placement origin. Such replacement carefully follows bedding, first selectively replacing the carbonates followed by at least partial silicification of the shale. It is important to note that the general aspect which ultimately is produced is complete replacement of the carbonates by silica. An investigator, who could observe only the final product, would be tempted to conclude that precipitation of the silica was a syngenetic process inasmuch as he would fail to recognize the former presence of the thin, semi-lenticularcarbonate beds. D o n (1958, p.7) has called attention to massive replacement of carbonates early in the burial history of somewhat similar sediments; whereas BISSELL (1959, p.177) favors the possibility of the chert being essentially syngenetic and that it certainly precipitated early in the sediment’s history. The siliceous shales and bedded cherts of the ideal eugeosynclinalassemblage show no obvious transgression of bedding; and the consensus is that the silica was deposited contemporaneously with the clastic sediments, or that it was precipitated immediately after deposition of the detritus. Although the possibility of the silica as a replacement deposit of carbonate and to a lesser degree of shale has not been eliminated by certain investigators, the proof has not been regarded as satisfactory. There exists agreement, however, that the quantity of silica associated with such eugeosynclinal strata is enormous, and far exceeds the quantity associated with similar volumes of cratonic sediments. Moreover, an appeal must be made to a special source for the exceptional amounts of such silica, and this must have been extracted from marine waters which under normal conditions are known to contain a very low proportion of silica in solution. There appears to be in current favor the concept that localized volcanism may have supersaturated the waters sufficiently to result in precipitation of the required enormous amounts of silica. According to such a concept the silica is precipitated as a primary constituent of the sediment, and may under certain circumstances constitute the entire sediment being accumulated. There appears to be no modern counterpart in recent sediments in volcanically active localities, hence the possibility exists that this concept requires modification. The petrographic evidence suggests that the precipitation of most chert occurs immediately following deposition of the sediments as a product of reactions which occur in the interstitial fluids. One can expect important local differences in ionic concentrations to exist below the depositional interface, and under such conditions precipitation of silica may occur. One must await investigation of such interstitial fluids before further advance in our knowledge of the diagenesis of silica is to be achieved.

REFERENCES

BAASBECKING,L. G. M., KAPLAN, I. R. and MOORE, D., 1960. Limits of the natural environment in terms of pH and oxidation-reduction potentials. J. Geol., 68: 243-284. B m , G. A., CONTOIS, D. E. and THOMAS, W. H., 1959. The removal of soluble silica from fresh

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water entering the sea. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., I : 20-35. BIGGS,D. L., 1951.Petrography and origin of Illinois nodular cherts. Illinois State Geol. Surv., Circ., 245: 25pp. BISSELL,H. J., 1959. Silica in sediments of the Upper Paleozoic of the Cordilleran Area. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 150-185. CAROZ~, A. V., 1960.Microscopic Sedimentary Petrography. Wiley, New York, N.Y., pp. 291-343. CAROZZI,A. V. and SODERMAN, J. G. W., 1962. Petrography of Mississippian (Borden) crinoidal limestones at Stobo, Indiana. J. Sediment. Petrol., 32: 391415. CHANDA,S. K., 1963. Cementation and diagenesis of the Lameta Beds, Lametaghat, M. P., India. J. Sediment. Petrol., 33: 728-138. CHILINGAR, G. V., 1956.Distribution and abundances of chert and flint as related to the Ca/Mg ratio of limestones. Bull. Geol. SOC.Am., 61: 1559-1562. CHOQUETTE, P. W., 1955. A petrographic study of the “State College” siliceous oolite. J. Geol., 63: 331-348. CLOUDJR., P. E., 1962. Environment of calcium carbonate deposition west of Andros Island, Bahamas. U.S.,Geol. Surv., Profess. Papers, 350: 138 pp. DAPPLES,E. C., 1959. The behavior of silica in diagenesis. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., I : 36-54. DAPPLES,E. C., 1967.The diagenesis of sandstones.In: G. LARSEN and G. V. CHILINGAR (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, pp.91-125. DELLWIG,L. F., 1955. Origin of salina salt of Michigan. J. Sediment. Petrol., 25: 83-1 11. DIETRICH,R. V., HOB= JR., C. R. B. and LOWRY, W. D., 1963. Dolomitization interrupted by silicification, J. Sediment. Petrol., 33: 646-664. D o n JR., R.H., 1958. Cyclical patterns in mechanically deposited Pennsylvanian limestones of northeastern Nevada. J. Sediment. Petrol., 28: 3-15. EMERY,K.O. and RITTENBERG, S.C., 1952. Early diagenesis of California basin sediments in relation to origin of oil. Bull Am. Assoc. Petrol. Geologists, 36(5): 135-806. FL~RKE, 0.W., 1955. Zur Frage des Hochcristobalit in Opalen, Bentoniten und Glasern. Neues Jahrb. Mineral., Monatsh., 1955( 10): 21 7-233. Fox, W.T.,1962.Stratigraphy and paleoecologyof the Richmond Group in southeasternIndiana. Bull. Geol. SOC.Am., 73: 621-642. FRANKS, P. C. and Swnmom, A., 1959. Character and genesis of massive opal in Kimball Member, Ogallala Formation, Scott County, Kansas. J. Sediment. Petrol., 29: 186-196. GIBBS,R. J., 1960. The Stratigraphy and Paleontology of the Fish Haven Dolomite of South Central Idaho. Thesis, Northwestern Univ., Evanston, Ill., unpubl. GLOVER, J. E., 1963. Studies in the diagenesis of some western Australian sedimentary rocks. J. Roy. SOC.West Australia, 46: 33-56. GOLDSTEIN Jr., A., 1959, Cherts and novaculites of Ouachita facies. Silica in Sediments-Soc. &on. Paleontologists Mineralogists, Spec. Publ., 7: 135-149. GOLDSTEIN, A.and HENDRICKS, T. A., 1953.Siliceous sediments of Ouachita facies in Oklahoma. Bull. Geol. SOC.Am., 64:421442. HARRIS, L. D.,1958.Syngeneticchert in the Middle Ordovician Hardy Creek Limestone of southwest Virginia. J. Sediment. Petrol., 28:205-208. Ihw,J. B., 1964.Geodes and concretions from the Mississippian Warsaw Formation, Keokuk Region, Iowa, Illinois and Missouri. J. Sediment. Petrol., 34: 123-134. HEALD,M. T., THOMSON, A. and Wncox, F. B., 1962.Origin of interstitial porosity in the Oriskany Sandstone of Kanawha County, West Virginia. J. Sediment. Petrol., 32: 291-298. HEM, J. D., 1959. Study and interpretation of the chemical characteristics of natural water. U.S.,Geol. Surv., Water Supply Papers, 1473:269 pp. HINTzE, L. F., 1953. Silicification of Ordovician fossils in Utah and Nevada. BUN. GeoL SOC. Am., 64: p.1508. INGELS, J. J. P., 1963. Geometry, paleontology and petrography of Thornton Reef Complex, Silurian of northeastern Illinois. Bull. Am. Assoc. Petrol. Geologists, 47: 40540. KELLER, W. D. and REESMAN, A. L., 1963.Glacial m i l k s and their laboratory-simulatedcounterparts. Bull. Geol. SOC.Am., 14: 61-16.

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KELLER, W. D., BALGORD,W. D. and REESMAN, A. L., 1963. Dissolved products of artiticially pulverized silicate minerals and rocks. J. Sediment. Petrol., 33: 191-205,426438. KRAUSKOPF, K.B., 1956.Dissolution and precipitation of silica at low temperatures. Geochim. Cosmochim. Acta, 10: 1-26.. KRAUSKOPF, K. B., 1959. The geochemistry of silica in sedimentary environments. Silica in Sediments-Soe. Econ. Paleontologists Mineralogists, Spec. Publ., 7:4-20. LAMAR, J. E., 1950. Acid etching in the study of limestones and dolomites. Illinois State Geol. Surv., Circ., 156: 47pp. L~VERINO, T. G. and PAITEN,L. E., 1962.The effect of COa at low temperature and pressure on solutions supersaturated with silica in the presence of limestone and dolomite. Geochim. Cosmochim. Acta, 74: 787-796. NEWELL, N. D., RIGBY,J. K., FISCHER, A. G., WHITEMAN, A. J., HICKOX, J. R. and BRADLEY, J. S., 1953. The Permian Reef Complex of the Guadalupe Mountain Region, Texas and New Mexico. Freeman, San Francisco, Calif.,pp.160-182. OGREN,D. E., 1961.Stratigraphy of the Upper Mississippian Rocks of Northern Arkansas. Thesis, Northwestern Univ., Evanston, Ill., unpubl. PELTO,C.R., 1956.Chalcedony. Am. J. Sci., 254: 32-50. PITIMANJR., J. S., 1959.Silica in Edwards Limestone, Travis County, Texas. Silica in SedimentsSOC.Econ. Paleontologists Mineralogists, Spec. Publ., 7: 121-134. PROKOPOVICH, N.,1953. Silicification in the Oneota dolomite. J. Sediment. Petrol., 23: 174-179. RIEDEL, W. R.,1959. Siliceous organic remains in pelagic sediments. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 80-91. SIEVER, R., 1957.The silica budget in the sedimentary cycle. Am. Mineralogist, 42: 821-841. SIEVER, R., 1959.Petrology and geochemistry of silica cementation in some Pennsylvanian sandstones. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 55-19. SIE~ER, R.,1962. Silica solubility, 0-200" C.,and the diagenesis of siliceous sediments. J. Geol., 70: 127-150. TALLAFERRO, N. L., 1934. Contraction phenomena in cherts. Bull. Geol. SOC.Am., 45: 189-232. TALIAFERRO, N.L., 1935.Some properties of opal. Am. J. Sci., 30:450-474. THOMSON, A., 1959. Pressure solution and porosity. Silica in Sediments-Soc. Econ. Paleontologists Mineralogists, Spec. Publ., 7: 92-110. VANTWL, F. M., 1918.The origin of chert. Am. J. Sci., 41: 449-456. WHITE, D. E., BRANNOCK, W. W. and MURATA, K. J., 1956.Silica in hot-spring waters. Geochim. Cosmochim. Acta, 10: 27-59.

Chapter 7

DIAGENESIS OF ORGANIC MATTER EGON T. DEGENS~

Department of Geology, California Institute of Technology,Division of Geologic Sciences,Pasadena, Calif. (U.S.A.)

SUMMARY

Geochemical data are presented in this chapter which shed some light on the type of alteration organic matter undergoes in the course of diagenesis and epigenesis. In general, the formation of fossil organic matter is accomplished by a three-step process: ( I ) the microbial and chemical(hydrolysis)destruction of former biochemical macromolecules in the early stage of diagenesis; (2) the condensation of metabolic and hydrolysis products, resulting in the formation of “heteropolycondensates” (humic materials); and (3) a slow inorganic maturation of the heteropolycondensates (e.g., loss in functional groups), with thermal degradation being the number one alteration factor. It is tentatively suggested that, given sufficient time, all fossil organic matter is eventually reduced either to aromatic condensates resembling graphite, or to light paraffinic hydrocarbons; temperature will accelerate the reactions. The occurrence of original biogenic matter such as amino acids, sugars, fatty acids, or the bases of the purines and pyrimidines in ancient rocks may serve as an indication that diagenesis and epigenesis of organic matter has not yet reached its final stage. Their presence may eventually help to outline the thermal history of the host rock. Aside from a discussion of the geologic history of biogenic matter, some ideas are expressed on the origin and diagenetic fate of abiotic organic matter that once has been synthesized on a prebiotic earth or on the meteorite parent bodies. Hydrocarbons and other organic volatiles in meteorites are assumed to have been generated from the finely disseminated organic matter. This diagenetic formation occurred rather late in the history of the meteorites, i.e., after the incident of asteroid collision and distintegration.

INTRODUCTION

An answer to the question of how organic molecules which constitute today’s California Institute of Technology, Division of Geologic Sciences, Contribution No.1231.

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living matter respond-upon death of the organisms-to diagenesis is important for a number of basic research topics, such as ( I ) biochemical evolution, (2) origin of petroleum and coal, and (3) thermal stability of biochemical constituents. A fourth topic concerns the diagenetic fate of terrestrial and extraterrestrial organic matter which once has been synthesized on a prebiotic earth or on the meteorite parent bodies. Its occurrence can be inferred by the presence of non-combustible organic carbon in sediments and meteorites as old as 3 or 4-5 billion years, respectively. The elucidation of the diagenetic history of this material may reveal important details concerning the origin of life on earth and perhaps elsewhere in the universe. Over the past three decades, beginning chiefly with the classical work of TREIBS (1934, 1936), much has been learned of a quantitative nature of fossil organic matter present in various terrestrial and extra-terrestrial rock materials and natural waters. These studies have evolved far beyond the estimation of total organic carbon and nitrogen to the isolation and measurement of specific kinds of organic matter, for example, the amino acids, carbohydrates, heterocyclic compounds, and lipids. In the early organic-geochemical studies, the presence of sugars, amino and fatty acids, or the bases of purines and pyrimidines in hydrolysis liquors of ancient rocks were used as indication that intact polysaccharides, polypeptides, nucleic acids, and other biogenic polymers could survive diagenesis. The present consensus, however, is that during diagenesis new polymers or condensates are synthesized from breakdown products of former biochemical macromolecules. Namely, the bulk of the fossil organic matter has acquired its molecular framework in the course of diagenesis, whereas biogenic macromolecules, such as proteins or nucleic acids, are eliminated rather radically in the early stages of diagenesis. It is the objective of this chapter to briefly outline the principal mechanisms responsible for the formation of the more common organic compounds presently encountered in sediments and meteorites. Special emphasis is given to those features that elucidate the fate of organic matter throughout geologic time.

GEOCHEMICAL BALANCE

Most of the organic debris that has come to rest on earth is found in sediments. HUNT(1962) determined the total organic matter in more than 1,000 rock specimens collectedfrom 200 formations in 60 major sedimentary basins. Shales averaged 2.1 %, carbonates 0.29 % (GEHMAN, 1962), and sandstones 0.05 %. Based on these data and the figures reported by WEEKS (1958) on the total volume of sediments and the ratio of shales to carbonates to sandstones, the organic matter entrapped in sediments is around 3.8 1015 metric tons. The overwhelming part of it, namely 3.6 1016 metric tons, is present in shales. Consequently, most of the organic

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matter occurs in a finely disseminated state and is associated with fine-grained sediments. For comparison, the coal deposits of the world have been estimated to total about 6 * 1012 metric tons. This is 1/500th of the disseminated organic matter. Estimates on the ultimate primary petroleum reserves run at 0.2 1012 metric tons. This represents only 1/16,000th of the total organic matter incorporated in sediments. The amount of dissolved organic substancesin the sea is only a few milligrams per liter (GOLDBERG, 1961). Fresh waters and some interstitial waters may carry considerably higher quantities, but when compared to the total mass of organic matter in the lithosphere and the ocean, their volume is small. The only extraterrestrial materials to which geochemists have access are the meteorites. According to WIIK(1956), the carbon content in fourteen of the twenty known carbonaceous chondrites ranges from 0.19 % (Ornans) to 4.83 % (Ivuna). Ordinary chondrites contain on an average only a few tens to a few hundred parts per million in organic carbon. Carbonaceous chondrites represent about 2 % (in mass) of all meteorites which have fallen on earth, and the mass fraction of organic matter in carbonaceous chondrites is about 0.5 %. This infers that about 0.01 % of all meteorite material is made of organic matter. The total mass of the meteorite parent bodies (asteroid belt) is roughly equivalent to a sphere of terrestrial density about 1,000 km in diameter (SAGAN, 1961).The mass of this sphere is less than 1/1,OOOth the mass of the earth. Namely, the asteroids contain about 1014 metric tons in organic matter, which is only 40 times less than the total amount of terrestrial organic matter. The actual difference might even be smaller, because in transit to the earth, meteorites might preferentially lose organic matter as a result of distillation and fractionation processes. In summary, one is reasonably well informed about the range in total organic matter present in various terrestrial rocks, meteorites, and natural waters. More interesting from a geochemical point of view, however, would be a knowledge of the precise chemical nature and distribution pattern of individual organic species. Also a general understanding of processes and mechanisms that are responsible for the preservation, alteration, or destruction of organic matter throughout geological history will have greater scientific application than a mere study of the total organic carbon content.

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CLASSIFICATION OF ORGANIC MATTER

More than 500 organic constituents have so far been identified in various geologic materials. This number does not include all the individual organic compounds that have been isolated from crude oils (ROSSINI, 1960). No classification can reasonably take care of all these individual organic

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molecules separately. They have to be grouped into distinct classes of compounds and treated individually. It is of interest to observe that most organic compounds of interest to geology bear some structural resemblance to former living matter. Even constituents once synthesized on a prebiotic earth or another planet are not too widely remote both chemically and structurally from the individual monomeric building blocks that constitute today’s living things. A geochemical classilkation of organic substances, therefore, is somewhat similar to a biochemical one. But, whereas the proteins, carbohydrates, and lipids are quantitatively more important in the plant and animal kingdom, the phenolic “heteropolycondensates”, hydrocarbons, and asphalts are of greater significance in geological materials. A classification scheme, which includes the major biogeochemical compounds, is proposed below: (I) Amino acids and related substances (2) Carbohydrates and derivatives (3) Lipids, isoprenoids, and steroids ( 4 ) Heterocyclic compounds (5)Phenols, quinones, and related substances (6) Hydrocarbons (7)Asphalts and allied substances Information on the structural composition and the general geochemistry of organic constituents has been presented in some detail by DEGENS (1965). The classification of fossil organic matter can also be based on its diagenetic behavior. Namely, fossil organic matter can be grouped into compounds that are (I) survivors of diagenesis, and (2) products of diagenesis. The first group includes all those organic molecules which are chemically similar or identical to living matter. Compounds that are, in general, biochemical building blocks of plants and animals, but which are formed from pre-existing organic debris sometime during diagenesis, are representatives of the second group. Metabolic waste matter may be regarded as an intermediate product, inasmuch as certain metabolites such as urea may also arise via thermal degradation in the course of diagenesis.

STABILITY OF ORGANIC MATTER IN DIAGENETIC ENVIRONMENTS

General concepts The prime question in all biogeochemical studies concerns the stability of the organic debris in a diagenetic environment. As data accumulated, it became apparent that microorganisms play a key role in the (I) destruction, (2) alteration, and (3) posthumous formation of organic matter. Although microbes undoubtedly control to a large extent the fate of organic matter in the early stagesof diagenesis, it is surprising how little is known concerning their occurrence, type of population,

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distribution, and biochemical activities in various geological materials. This feature can partly be linked to inadequate analytical preparation techniques. There is no doubt that the number of microbes in soils and sediments sharply decreases with depth of burial. For example, the bacterial population in surface sediments of marine origin may amount to 100 million bacteria per gram and more. This would represent approximately 10-3 g of organic matter or 1 mg/g of sediment. Sediments at a depth of 50-100 cm have considerably lower population densities, i.e., a few thousand bacteria per gram of sediment. Thus, the following question arises: do microorganisms further decrease in abundance to zero at some depth of burial below which sterile conditions would exist, or are they still present in ancient rocks as old as the Precambrian? The answer to this question is undoubtedly of great significance in understanding the problem in hand. There have been periodic claims in the geologic literature on the isolation of viable bacteria from geologic ancient materials such as coal, fossil dung, petroleum, connate waters, sulfur domes, salt deposits, various sediments, igneous rocks, and even meteorites. In many such instances the investigators have stated or implied that the bacteria had survived over the geologic ages involved. One should not review these claims here; it is sufficient to point out that the problem of procuring and sampling such materials aseptically is a very difficult one and, in most instances, a critical appraisal of the published work reveals gross inadequacies in the experiments.The pertinent publications include: LIPMAN (1928),WAKSMAN et al. (1933), RTTENBERG (1940), PORTER (1946), ZOBELL(1946a,b, 1952, 1963), EMERY and RITTENBERG (1952), NEHER and ROHRER (1959), BAASBECKING et al. (1960), LIM~BLOM and LUPTON(1961), OPPENHEIMER (1961, 1963), SISLER(1961), and RITTENBERG et al. (1963). It is highly problematic whether one will ever understand all of the microbial activities in sediments. But the fact that the alteration of organic matter in the early stages of diagenesis is largely controlled by the activities of microorganisms and burrowing animals should stimulate geochemists to learn more about the functions of the living populations within a sediment and soil. After microbial activities have virtually ceased in the strata at some shallow depth of burial, the diagenetic alteration predominately proceeds in the following manner: ( I ) inorganic maturation, i.e., polymerization and condensation with the associated mineral and organic matter, (2) redistribution of pre-existing organic molecules, e.g., as observed during sediment compaction and formation of crudeoil deposits, and (3) thermal degradation, partly catalyzed by clay minerals etc., which may eventually eliminate certain organic compounds, or else yield new (authigenic) organic molecules. In knowing the stability ranges of the individual organic compounds, the presence, absence, or ratios among the various organic species may reveal valuable information regarding the thermal history of the host strata. In view of the fact that theoretically more than five hundred different organic constituents might be

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present in any one rock specimen, there is no further need to stress the importance of this type of approach to aspects of diagenesis and metamorphic petrology. In order to evaluate alterations in biochemical compounds immediately following death of plants and animals, one naturally has to be familiar with the principal organic building blocks that constitute the molecular framework of all living organisms, i.e., (1) proteins, (2) carbohydrates, (3) lipids, and (4) heterocycles (nucleic acids and pigments). A fifth group, the lignins, generally occur only in the higher organized plant kingdom, in which they are common and widespread. For details on chemistry, structure, and physiology of these compounds, one can consult WHITEet al. (1959). The stability of biogenic macromolecules is of rather short duration. Practically all of the original biogenic polymers are eliminated during the early stages of diagenesis. Nevertheless, biogeochemical alterations that proceed during these early periods will be the key for the understanding of the biogeochemistry of fossil organic matter in sediments and natural waters. It has been stated before that shales contain most of the fossil organic matter, limestones some, and sandstones the least. One way to account for this phenomenon is to assume that recent clay muds are also higher in organic compounds by a factor of ten relative to recent carbonates and sands. This assumption is certainly correct for the sandstones, but does not hold true for the carbonates. As can be seen in Fig. 1, total organic matter is about the same for recent I

Total organic matter

63A m i n o

compounds

Organic solvent extract

Ancient shale organic matter)

Shell Carbonate lMyti/us C a l ifornianus)

Limestone

ff/orida Bay)

Clay

(Son ffiego7ioughl

Fig.1. Distribution of major organic fractions in recent marine carbonates and clay muds. Organic matter in recent sediments usually ranges from about 0.5 to 10%. Most of the samples, however, do contain about 1.5-3 % of organic matter.

DIAGENESIS OF ORGANIC MATTER

349

Fig.2. Distribution of amino acids in recent and fossil shell proteins of Mytilus calvornianus. (After HARE, 1962.) Explanation of abbreviations: ALA = alanine; ARG = arginine; ASP = aspartic acid; CYS = cysteic acid; DIAM. PIM. ACID = diaminopimelicacid; GLU = glutamic acid; GLY = glycine; HIS = histidine; HYFXO = hydroxy-proline; LLEU = iso-leucine; IS0 = iso-leucine;LEU = leucine; LYS = lysine; MET = methionine; MET SUL = methionhe sulphoxide; ORN = ornithine;PHE = phenylalanine;PRO = proline; SER = serine; THR = threonine; TYR = tyrosine; VAL = valine.

shell carbonates, limestones, and clay muds. All three samples are representative for a great number of studied specimens obtained from different marine environmentsl. Whereas a mean value of 2 % organic matter is approximately maintained in shales of all geologic ages, ancient limestones have on an average ten times less organic matter than their recent counterparts. This is probably due to the chemical differences in the original organic matter. Amino acids

Amino acids constitute 90% and more of the total organic matter in recent shell materials and carbonates, compared to less than 10 % in the clay mud. In all three specimens (Fig.l), most of the amino acids come from intact proteins or peptides. Carbohydrates and lipid materials are about evenly distributed in the samples under consideration. The missing 60-80 % in the case of clay muds is made up of chemically undefined "organic residue". This portion represents both the humic acid fraction and the so-called kerogen. The ratio of these materials is about 1/1. It appears, therefore, that the organic-residue fraction in a shale environment has greater chances of diagenetic survival than the protein fraction in a limestone environment. Outstanding experimental work on the thermal stability of proteins and their

350

E. T. DEGENS

individual amino acids has been done by VALLENTYNE(1957b), CONWAY and LIBBY (1958), and ABELSON (1959). Their data indicate that proteins which survive microbial attack are most likely hydrolyzed in an aqueous environment within hundreds or thousands of years, but that some of the released amino acids are stable over millions of years under low-temperature conditions. Radiocarbon-dated shells of Mytilus californianus,ranging in age from 400 to 5,500 years, as well as specimens from an upper Pleistocene marine terrace, show a signilicant decrease in total organic matter, when compared with recent Mytilus carbonate shells (HARE,1962). After a few thousand years, the protein concentration drops from about 1 %/g (recent) to about 0.03 %/g (Pleistocene) of shell carbonate. The amino acid spectra obtained are presented in Fig.2. Identical results were observed in sediment samples from the Cariaco Trench (LINDBLOM and LUPTON,1961). The protein content in a near-surface sediment was about 1 %/g of sample. With depth of burial, the proteins decreased gradually to a value of 0.09 % in probably Pleistocene sediments, 23 ft. below the sediment-water interface. Similar trends were established in a number of other 1963). marine deposits (LINDBLOM and LUPTON,1961; LINDBLOM, These data may serve as an indication that proteins are eliminated in rather short periods of time. Intact polypeptides will probably never be found in sediments older than the Pleistocene, except where unusual environmental conditions are established, as, for example, in a dry climate (desert), or in cases where organisms and sediments become impregnated with asphaltic materials shortly after burial (La Brea asphalt pits, Los Angeles). The complexity, in addition to the limited stability of proteins and peptides with geologic time, makes geochemical studies of amino acids more attractive. This is so because a great number of amino acids are still present in ancient rocks, natural waters, and meteorites. Furthermore, the analytical identification and quantitative estimation of amino acids is well advanced. In general, one may distinguish betweenfree and combined amino acids. Free amino acids are those that are dissolved in natural waters or that can be extracted with water or ammonium acetate from solid rock materials. In contrast, combined amino acids are those which require acid or alkaline hydrolysis for their final release. Ancient sediments generally lack free amino acids. Exceptions include some fossil shell carbonates which occasionally contain small amounts of free amino acids. This is due to the fact that proteins in shell carbonates can escape microbial decomposition in the early stages of diagenesis; microbes are too large (> 1 p) to have access to protein films embedded in the carbonate matrix. On the other hand, in the presence of water, proteins will eventually become hydrolyzed, and free amino acids are simultaneously released. Recent sediments and soils yield reasonable quantities in free amino acids (JONES and VALLENTYNE, 1960; STEVENSON,1960; DEGBNS et al., 1963a). In both

0

2

<

GLYCINE 5

VALINE

102c

-:.-

1

2

:

3

-

E 4

Basin (reducing). (Aftex DEOENS et al., 1961,1963.)

LYSINE

L GLUTAMIC ACID 2

2 sediments,

5 1 0 2 0

off California.

I4

352

E. T. DEGENS

Ocean

a

Sediment

I00

( d r y weight)

I50

200

50

Id0

L 4 i r - d F

Fig.4. Depth distribution of amino acids in ocean waters (,ug/l) and in the underlying sediment (pguglg dry weight). (After R~TTENBERQet al., 1963.)

instances, the size and nature of amino acids appear to be closely related to the (1961), soils low in level of microbial activities. As shown by PAULand SCHMIDT microbial activities contain free amino acids in concentrations ranging from about 0.05 to 0.5 pg/g, whereas concentration in well-populated soils may reach peak values of 100-200 pg/g. Most free amino acids in soils are excretory or autolytic products of microorganisms. Some may have also been added by root excretion. The rapid destruction by microorganisms, however, prohibits large accumulations of free amino acids in soils. Another reason is that free amino acids can interact with the surrounding mineral and organic matter, resulting in the formation of complex condensation products. Recent fresh-water and marine sediments which were deposited under oxidizing conditions contain about the same amount of free amino acids as the average soil. As the microbial activities decrease with depth of burial, so do the free amino acids. Below 3-5 m the level of concentration is about 0.5 pg/g. A completely different picture is obtained in recent reducing sediments. Here free amino acids constitute most of the amino acid fraction. Values in the neighborhood of a few hundred up to 1,OOO pg/g are quite common (DEGENS et al., 1961). Furthermore, considerable fluctuations exist between contents of the individual amino acids in systematically taken cores (Fig.3). By calculating the mean value from several samples of the upper few meters of deposit and comparing it with that of the ,plankton in the sea, however, a surprising uniformity of the amino acid spectrum is observed in both sets of samples. The following explanation is offered to account for this phenomenon. In the initial stages of diagenesis, plankton protein supplied from the sea

DIAGENESIS OF ORGANIC MATTER

353

above becomes gradually hydrolyzed. Inasmuch as microbial activity is low and burrowing organisms are absent, the generated free amino acids are not consumed and may survive diagenesis; but they redistribute themselves by means of ionexchange chromatography during compaction and dehydration of the strata. Adsorption and resorption of free amino acids occur along the clay-water interface. The degree of separation between individual amino acids depends on ( I ) type of clay mineral, i.e., chemistry, size, structure, (2) salinity, (3) pH, and (4) temperature and pressure in the clay-water-amino-acid system. The final result is the irregular distribution pattern observed in Fig.3; namely, isolation and purification of amino acids is achieved. Combined compounds represent the bulk of the fossil amino acids extractable from sediments of all ages. In order to illustrate the progressive alterations of these materials in the course of diagenesis, a sediment core obtained during Experimental Mohole drillings (RIEDELet al., 1961; RITTENBERG et al., 1963) was examined. This core penetrated about 170m of hemipelagic ooze and covered 15 million years of earth history. With depth of burial, the amino acid concentration sharply decreases (Fig.4). Amino acids biogenically produced in the mud at or near the water-sediment interface gradually decrease relative to the amino acids formerly contributed by the Sea(DEGENSet al., 1964b). By taking an average of the first 2 m and comparing it with that of the lower 70 m of the section, the relationship, schematically plotted in the Fig.5 was obtained. Ornithine, serine, glutamic and aspartic acids, the leucines, threonine, glycine, and alanine are diagenetically more stable than the rest of the amino acids. It is of great significance that these are the same amino acids that have been found in natural waters from the Paleozoic and younger formations, and in most of the ancient sediments investigated so far. It is conceivable that the highly metastable arginine and tyrosine are altered into ornithine and serine, respectively. The last reaction would yield an additional benzene ring. Amino acids are found in sediments of as early as Precambrian times (ABELSON, 1954, 1957, 1959; BARGHOORN, 1957; SWAINet al., 1958; DEGENSand BAJOR, 1960; SWAIN1961;HARRINGTON, 1962;H A R R I N G T OCILLIERS, N ~ ~ ~ 1963;and others). In Table I the order of magnitude of amino acid concentrations in various geologic materials is presented. In the case of recent shell carbonates and sediments, most of the amino acids obtained are hydrolysis products of proteins and peptides. Amino acids in ancient materials, however, are for the most part incorporated in non-proteinaceous complexes. From the published data it can be inferred that older sediments contain progressively less amino acids than geologically younger ones. This decrease is largely a result of thermal degradation over millions of years, and some amino acids appear to be less stable than others. At room temperature, the reactions proceed slowly. Higher temperatures of a few hundred degrees Celsius are needed to measure experimentally (that is, within reasonable time) the thermal stability of the indi-

I OF AMINO ACIDS IN VARIOUS OEOLOQIC MATERIALS

wncmtrates

carbonates

sediments

waters

Material

Geologic age

Concentration (in P 8 k )

humic acids (91Peat (2) lignite (19) kerogen (3) bituminous coal (9) anthracite (3)

2QoO0 10,000 300

petroleum (4)

Recent Recent Miocene Pennsylvanian Tertiary Jurassic-Cretaceous Paleozoic

shell carbonate (10) marine limestones (4) shell carbonate (1) shell carbonate (1) marine limestones (4) fresh-water concretions (3) marine limestones (2)

Recent Recent Pleistocene Miocene Miocene Miocene Pennsylvanian

10,OOo 10,Ooo 300

soils (6) marine argilIaceous mud (5) marine argillaceous mud (2) marine shales (5) marine shales (5) chert (2) manganese nodule

Recent Recent Pleistocene Miocene-Cmtacmus Pennsylvanian Precambrian Recent

10,Ooo 5,000 1

petroleum brine water (15) ocean water (10) rain water, Pasadena (1)

Paleozoic Recent Recent

carbonaceous chondrites (8)

Meteorites

chondrites (4) Number in parentheses

=

number of studied samples.

200 20 1

0.5

50 20 5 5

400 20 10 5

0.20 0.05

o.oO05

100 (combined) 30 (free) 20 (combined) 10 (free)

355

DIAGENESIS OF ORGANIC MATTER

II

+Or

nit hine

+Serine

u a LL W

I tGI yci n e +Aspartic acid

Threonine

=Glutarnic

_ _ - - - - - - - - - ---

acid

A,anine-

t L y s i n e

,Valine fl% t0-Alanine ,/* - *c Proline

%3%. .

114

Y

=?--;O,t;,ee

Fig.5. Relative stability of amino acids dunng diagenesis; wet oxiaimg conaitions. (Arter ~ N B E X Q et al., 1963.)

vidual amino acids incorporated in the humic acid or kerogen complexes. DEGENS and HUNT(1964) and SELLERS(1965) subjected sediments, humic acids, and kerogens to temperatures up to 500°C for hours or days. Their results agree with nature; namely, the amino acids most commonly found in ancient rocks are the same ones which are least affected by higher temperatures. In order to illustrate the effect of higher temperatures on the amino acid distribution in fossil organic complexes, sedimentsof the CretaceousPierce Formation, cut by a basic 1m dike, were investigated (Fig.6). With distance from the dike, the amino acid spectrum changes systematically. For instance, in approaching the dike the acidic amino acids drop down in abundance to practically zero, others seem to be little affected, e.g., proline; and some even increase in concentration (urea). Similar trends are established in the case of the hydrocarbons, sugars, and phenols (HUNT, 1962; DEGENS and HUNT, 1964). The ultimate goal of this type of research

356

E. T. DEGENS

tc m 0 l . h AMINO ACIDS

AMINC SUGAR:

0.76

0.06

3.75

0.03

3.27

0.04

'4.39

0.05

Fig.6. Amino acid spectra of wall rock sediments (1 m basic dike); Pierre Shale, upper Cretaceous. (After DEOENS and HUNT,1964.). *Distance from 1-m basic dike. Explanation of abbreviations: see legend Fig.2.

is to find a number of organic compounds that can be used as a geoIogicaI thermometer. The fact that certain organic constituents survive temperatures of 500 "C for periods of days is very encouraging. Inasmuch as most sediments and low-rank metamorphic rocks derived from sediments contain organic matter, differences in the temperature sensitivity or organic molecules should allow more detailed insight into geological aspects of regional metamorphosis, diagenesis, ore formation, and others, where differences in temperatures are critical. Carbonaceous chondrites yield about ten times as much amino acids as do ancient sediments having about the same amount of organic carbon. In addition, meteorites contain both free and combined amino acids in a ratio 113. A comparison of amino acid spectra of terrestrial and extraterrestrial materials reveals interesting results. The amino acids most frequent in ancient sediments and kerogens are in some approximation also the more dominant ones in the meteorites (Fig.lo). Namely, serine, glycine, alanine, the leucines, ornithine, and the acidic amino acids are present in greater abundance. Also cysteic acid, urea, and perhaps taurine appear in greater concentration, particularly in the meteorites. Of great significance is, furthermore, the presence of free amino acids in the meteorites. This can mean that amino acids were formed relatively late in the history of the meteorite, inasmuch as the thermally unstable amino acids are the most common ones. The other alternative explanation is that the free compounds were actually released from the organic residue during analytical treatment with boiling water (KAPLANet al., 1963). In view of the greater abundance of water-soluble salts in meteorites, one may reasonably suspect that a partial hydroIysis was already

DIAGENESIS OF ORGANIC MATTER

357

accomplished during a 24-h reaction period. In any event, the fact that a boiling water treatment can release measurable quantities of amino acids in meteorites and not in shales and limestones suggests that the formation of “heteropolycondensates”1is perhaps less advanced in the meteorites than it is in the ancient sediments. The higher amino acid yields in the carbonaceous chondrites, when compared to ancient sediments, would support this assumption. KAPLANet al. (1963) concluded that the amino acid fraction is indigenous to the meteorites and was present at the time of falling to the earth. The lack of optical rotation and a number of other criteria suggested that the amino acids were formed (1953, by abiotic rather than known biogenic processes. The studies by MILLER 1963), CALVIN(1961, 1963), 0 ~ 6 and KIMBALL (1962), PALMand CALVIN (1962), and 0 ~ (1963a, 6 b), and others, offer ample evidence for the formation of complex organic molecules from water, ammonia, and methane (or hydrogen cyanide), and on utilizing different forms of energy. In summary, most of the diagenetic alteration of the proteins occurs within the zone of biological activity. Proteins, for example, supplied by the sea above, serve as convenient food source for organisms living in marine sediments. The disappearance of microorganisms below the first few meters of sediments can be linked to the depletion of biochemically available food sources supplied by plankton and various organisms living in the sediments. Combined amino acids found in ancient sediments are incorporated in non-proteinaceous condensates. It is most likely that the data in the literature on the distribution of amino acids in various rocks and natural waters represent only minimum values, because hydrolysis with 6N HC1 does not, in general, release all amino acids present in a rock specimen (DEGENS and HUNT,1964). Amino acid analyses of about 500 ancient sediments and natural waters show no apparent trend with geologic time, which can be used for stratigraphical correlations. The amino acid pattern of a sediment largely reflects the thermal conditions the rock has been subjected to in the course of its geologic history. Even meteorite spectra are not too much different from those of terrestrial materials. In other words, amino acid compounds synthesized on the “prebiotic” meteorite parent body, and possibly on a prebiotic earth, are of the same nature as the amino acids generated by living things and preserved during biogenic decay. They differ only in one little detail: the first ones are racemic mixtures, whereas the latter ones are optically active. It is conceivable that the number of aliphatic groups in the extraterrestrial organic complex is far in excess of aromatic groups. In contrast, the bulk of the fossil terrestrial heteropolycondensates is organized in aromatic nuclei, at least the ones formed from former biogenic matter. One may speculate that the aliphatic molecules in the meteorites are arranged in chains which are coiled so as to give stable conformations. A particularly favorable arrangement, found in proteinaceous materials, is the a-helix; and perhaps similar structures exist in the case of abiotic primordial complexes.

358

E. T. DEGENS

Carbohydrates Independent of the type of recent marine sediment, polysaccharides are less plentiful than the proteins. Sugar carbon constitutes only a small percentage-generally less than 5 %-of the total organic carbon. As shown by RITTENBERG et al. (1963), in Experimental Mohole sediments, the ratio drops sharply from about 4 % at the surface to less than 0.1 % at a depth of about 40 m. Below this depth, little variation is observable for the next 130 m of the core section. This relationship indicated to them that carbohydrates are even more rapidly eliminated than other components of the organic matter in the first stage of diagenesis, in the oxidizing environment which prevails in these sediments. More detailed information on the stability of individual carbohydrates (sugars) in oxidizing and reducing marine sediment environments was presented by PRASHNOWSKY et al. (1961) and DEGENS et al. (1963b). The relationships among sugars are illustrated in Fig.7. Total sugars drop from about 2,000 pg/g in the most recent muds to about 100 pg/g (oxidizing) and 500 pg/g (reducing) in sediments at a depth of about 3-4 m below the present sediment-water interface. Noteworthy are the significant fluctuations between the individual sugars in the case of the reducing sediments, compared to the smooth pattern established in the oxidizing sediments. Free sugars represent less than 1% of the total sugars in the recent oxidizing environment. In contrast, most of the sugars in the reducing sediments are present in a free form. This may explain the irregular distribution pattern. Namely, HEXOSES

PENTOSES

Fig.7. Distribution of sugars in recent marine sediments, off California. 0 San Diego Trough et al., 1961; and DEOENS (oxidizing); m Santa Barbara Basin (reducing), After PRASHNOWSKY et al., 1963.)

DIAGENESIS OF ORGANIC MATT'ER

359

free sugars are chromatographically separated during natural compaction of the strata, in a fashion similar to their corresponding free amino acids, which have previously been discussed. In view of the low microbial activity in the reducing strata, free sugars may survive the early stages of diagenesis. Under oxidizing conditions, however, they are rapidly consumed by the biologicaI population living in the sediment. Terrestrial soils and sediments may yield considerably higher quantities of carbohydrates. Sugar concentrations in the order of 1-5 % of the sediment or soil are rather common. Most of these carbohydrates, however, are also eliminated biochemically in the early stages of diagenesis. Some, perhaps, may survive bacterial consumption by becoming incorporated in clay structures or by forming nutritionally unattractive polymers (GREENLAND, 1956; LYNCHet al., 1957a; and BADERet al., 1960). The rate of decomposition of organic matter in a soil has been studied by means of 14Clabeled barley straw (SPIRENSEN, 1963).Among 14C active substances added to a soil were straw, hemicellulose, cellulose, compost, lignin, and watersoluble materials extractable from straw. Rate of decay of the individual compound was measured by means of radioactive C02 evolved from the soil. Finally, the activities of organic extracts recovered from the soil after a period of about 100 days were measured. Of the activity initially added, 90 % or more could be accounted for by the recovered C02 and the combined fractions of humic acids, fulvic acids, and humins. In Fig.8 the values of evolved carbon dioxide and total organic matter-adjusted to 100%-are presented. It can be seen that hemicellulose and cellulose become more rapidly eliminated from the soil in the form of C02 than do lignin materials, which actually produce only small quantities of C02. The diagenetic fate of cellulose has also been studied geochemically in some detail, particularly in connection with the origin of coal (GOTHAN, 1922; H a s and KOMAREWSKY, 1928; STAUDINGER and JURISCH,1939; BARGHOORN,1949; BARG HooRNand SPACKMAN, 1950; THEANDER, 1954; MET^ et al., 1960). Cellulose has been isolated from lignites as old as early Tertiary. But the degree of polymerization in the Tertiary deposits is about ten times less than in present-day cellulose. The depolymerization into about ten smaller units may have been accomplished by the action of microbes; another alternative is that the peat-bog environment may have favored the hydrolysis of cellulose into smaller fragments. Summing up all information on sugars, it appears that cellulose is one of the more stable polysaccharides under geological conditions. In contrast, starch, hemicellulose, or alginates are rapidly eliminated in the very early stages of diagenesis under both oxidizing and reducing conditions. Due to a prolonged interaction of free sugars in reducing sediments with the associated organic and inorganic matter, the chance of formation of rather stable complexes is here greatly enhanced. On the other hand, free sugars in an oxidizing environment are rapidly eliminated microbiologically. The small amounts of combined sugars, i.e., 0.1-10

360

E. T. DEGENS

25

50

75

I00 Hemi cel lulose

Cellulose Water-soluble

su bstonces Straw, ext. Compost

PER CENT14C ACTIVITY RECOVERED IN CO, AND SOIL ORGANIC MATTER

Fig.8. Distribution of 14C active organic residues and carbon dioxide, recovered after about 100 days decomposition in a soil; in per cent of original input. (After SBRENSEN, 1963.)

p.p.m., reported from ancient sediments as old as the Precambrian (PALACAS, 1959) are probably part of the kerogen molecule. Free sugars do not occur in ancient materials. It is presently not fully understood whether a prolonged hydrolysis with mineral acids may yield greater amounts of combined sugars from fossil matter. Free and combined sugars, in a ratio of about 1/2, have been extracted from stony meteorites (DEGENS and BAJOR,1962; KAPLANet al., 1963). On an average, total sugars are five times less abundant than the correspondingamino acid fraction. The presence of free sugars and the significantly higher concentration of total sugars in the meteorites when compared to ancient sediments may find the same interpretation as was offered in the case of the amino acids. Namely, “free” sugars were already released from a predominately ahphatic complex as a result of a partial hydrolysis, when the meteorites became extracted with boiling water. Lack of optical rotation of the meteorite sugar fraction would support the inference made earlier that the organic matter presently encountered in meteorites was synthesized on a prebiotic parent body. Among the sugar derivatives,glucosamine and chondrosamine are of particular geochemical interest. Both are found in recent sediments and soils in concentrations ranging from a few hundred to a few thousand pg/g of dry material. Based on limited data, their content in ancient sediments is of the order of a few

DIAGENBSlS OF ORGANIC MATTER

361

tens to a few hundreds pg/g (RITTENEIERG et al., 1963). Comprehensive studies and Smw (1954) and STBVENSON on hexosamines in soils are those by BREMNER (1957, 1960). The data suggest that most of the mucopolysaccharides are of bacterial origin. Insight on the path of microbial destruction in soils and sediments may be obtained by analyzing not only for the total amount of hexosamines, but separately for chondrosamine, glucosamine, and glucuronic acid (GRAVELAND and LYNCH,1961), inasmuch as all three of these monomers are the principal building blocks of one or the other of the mucopolysaccharides. Unfortunately, no detailed investigation has been made on the occurrence of these sugar derivatives in the stratigraphical column. Considering the fact that intact chitin from wing remains of Eocene insects is known (AE~ERHALDEN and H m s , 1933), and, in addition, many reports are published on the tentative identification of chitin-type materials from arthropods, graptolites, and some other invertebrates from as early as Cambrian deposits, there is much to be learned about the diagenetic history of mucopolysaccharides. As far as extraterrestrial organic matter is concerned, there is no indication for the presence of both chondrosamine and glucosamine in stony meteorites (Fig. 10). Lipids, isoprenoids, and steroids

Lipids, isoprenoids, and steroids constitute a group of biochemical compounds that are water-insoluble, but are very soluble in certain organic solvents such as ether or hydrocarbons. A reason for discussing all three substances simultaneously is their close biochemical relationship to one another. In the geological literature, the term lipid is used in a very broad sense. A variety of chemically unrelated substances such as certain heterocyclic compounds, fats, hydrocarbons, inorganic sulfur, and many others are combined and referred to as the “lipid fraction”. All these compounds have only one feature in common, i.e., they can be extracted with organic solvents. Biochemically speaking, however, the term lipid is restricted to esters of long-chain carboxylic acids. As in the case of proteins and polysaccharides, there is virtually no data available on condensed lipid compounds, e.g., glycerides, in geological materials. All informationis based on hydrolysisproducts such as fatty acids or certain alcohols. Fatty acids have been isolated from soils, recent and ancient sediments, peats, lignites, waxes, petroleum, natural waters, and meteorites ((SCHREINER and SHOREY,1908, 1910a,b; TANAKA and KUWATA,1928; CAWLEY and KMG, 1945; WILLIAMS, 1961; COOPER,1961; SLOWEY et al., 1962; ABELSON et al., 1963; and NAGYand BITZ, 1964). Most noteworthy is the observation that acids having odd numbers of carbon atoms are found along with those having even numbers of carbon atoms. This is in contrast to nearly all biological systems, where fatty acids are even numbered.

362

E. T. DEGENS

In a few selected samples, COOPBR (1961) presented a comparative analysis on the distribution of fatty acids in recent and ancient sediments, and petroleum brine waters (Fig.9). Independent Qf age, sediments show a preference of evennumbered fatty acids. Most abundant are palmitic and stearic acids. The relative abundance of odd-numbered acids, however, increases with geologic time. In the case of the petroleum brine water, concentration differences between neighboring odd- and even-numbered acids become small. In addition, a nearly straight-line decrease in relative abundance from c14 to Cso acids is developed. Principally, one might think of two possibilities for generation of oddnumbered fatty acids: (I)biological production, and (2) decarboxylation. Some biogenic products are known to contain small amounts of odd-numbered fatty acids. 20

RECENT

SEDIMENT

(SANTA BARBARA

BASIN)

ANCIENT SEDIMENT (EAGLE FORD SHALE)

PETROLEUM RESERVOIR WATER (PANHANDLE FIELD)

o

~

"

'

~

'

'

'

"

"

'

'

'

Carbon atoms in acid

Fig.9. Comparison of the distributions of fatty acids in a m t sediment, an ancient sediment, and in a water from a petroleum reservoir. (After CooPER, 1961.)

DIAGENESIS OF ORGANIC MATTER

363

Assuming that microbial activity selectively removes even-numbered fatty acids, there will be a relative increase in odd-numbered fatty acids. The work of SILLIKER and RITTENESERG (1952) on the selective use of even-numbered fatty acids by microbes would support this inference. Although microbial activity cannot be fully excluded as a cause for the relative enrichment of odd-numbered fatty acids with time, decarboxylation seems to be a more likely explanation to account for the observed odd-even pattern. The increase in abundance of odd-numbered acids during diagenesis apparently matches the generation of even-numbered p a r a f i s (BRAY and EVANS,1961). This parallelism suggests related processes for the formation of these acids and parafis. COOPER (1961) proposed that intermediate products are formed by decarboxylation of a fatty acid. Intermediates of this type could yield mixtures of odd-numbered acids and paraffins as reaction products. The odd-numbered acids produced would react similarly to form even-numbered acids and paraffins. The conversion of the naturally even-numbered carbon containing fatty acids in a sediment by chemical decarboxylation would cause the introduction first of odd-numbered acids and paraffins, and later even-numbered parafis. Also, fossil waxes in peat and lignite deposits contain a variety of characteristic even- and odd-numbered fatty acids and alcohols. Whereas information is plentiful on the type and concentration of wax compounds in various coals (FRANCIS, 1961;VANKREVELEN, 1961), only limited data are available on waxes in sediments. As a matter of fact, the only comprehensive study is that of MEINSCHEIN and KENNY(1957) on soil waxes. The principal constituents of soil waxes are normal aliphatic acids, normal primary aliphatic alcohols, and sterols. The types of acids and alcohols present show close resemblance to beeswax. As shown in Fig.1, ether-extractable lipids (largely fats and waxes) constitute about 5 % of the total organic matter in recent carbonate materials and marine clay muds. In contrast to the interesting trends of protein and carbohydrate decrease with depth of burial (Fig. 3 and 7), the so-called “lipid fraction” stays constant or, occasionally, even slightly increases with depth of deposition (LINDBLOM and LUPTON,1961; LINDBLOM, 1963). It is presently not f d y understood whether the constancy in yield with age is due to the excellent geological preservation of the original biogenic lipid compounds, or to the diagenetic generationof ether-extractable lipids. The production of organic solvent-extractable matter can proceed (I) biologically, i.e., by organisms living in the sediments, or (2) chemically, i.e., by thermal degradation and maturation of organic debris. Taking all sources of information into consideration, the second alternative seems to be more likely. According to it, the constancy in yield is attributed to the diagenetic generation of new (authigenic) ether-extractable compounds, which apparently matches or sometimes exceeds the decomposition of the original fats and waxes. A series of long-chain fatty acids, c14 to CSO,has been isolated from the

364

E. T. DEGENS

Orgueil carbonaceous chondrite (NAGYand BITZ, 1964). The relative abundance of even- and of what appear to be odd-numbered chain acids resembles that of fatty acids in ancient sediments on earth. Various theories, i.e., organic and inorganic ones, have been offered to account for their presence. Among the isoprenoids studied most intensively so far are the compounds of the carotenoid family (TRASKand Wu, 1930; Fox, 1937; Fox et al., 1944; VALLENTYNE, 1956, 1957a; ERDMAN, 1961; and SCHWENDINGER and ERDMAN,1963). Plant tissues contain on the dry weight basis approximately 0.1 % carotenoids. Inasmuch as soil and sediment microorganisms do not generally metabolize isoprenoids, their diagenetic stability will be largely controlled by inorganic processes. Indeed, biogenically unaltered carotenoids have been extracted from terrestrial and marine sediments as old as 100,000years. Older sediments, however, are free of them. According to SCHWENDINGER and ERDMAN(1963), the ratio of xanthophylls to carotenes in unconsolidated sediments is 1.3/3.6, compared to that of living plant materials of about 3/10. Namely, xanthophylls-the oxygencontaining compounds- are disappearing at a faster rate than the carotenes, the hydrocarbon compounds. Environmental differences are established in the way that the marine sediments show highest carotenoid concentration, up to 800 p.p.m. relative to organic carbon. This feature is possibly a result of the larger percentage of carotenoid pigments in planktonic matter. 1963; and It was suggested (HANSON,1959; ERDMAN,1961; MEINSCHEIN, MULIKand ERDMAN,1963), that isoprenoids are potential sources for lowmolecular weight hydrocarbons. This inference is supported by thermal degrada(1963) and MULIKand ERDMAN (1963). A mild tion studies of DAYand ERDMAN thermal treatment will release benzenoid hydrocarbons from /3-carotene as well as recent marine sediments. But in view of the high quantities of aromatic hydrocarbons (benzene, toluene, xylene) released from a mud slurry during the mild thermal treatment, other sources in a sediment besides carotenoid pigments have to be considered. Likely precursors of benzene are polyunsaturated fatty acids. From similar non-isoprenoid branched polyunsaturated chains o-xylene might arise. Toluene and m-xylene may originate from isoprenoids other than carotenoids. The presence of greater amounts of pentacyclic and hexacyclic compounds, possibly triterpenes, has been demonstrated by MEINSCHEIN and KENNY(1957). Benzenoid hydrocarbons may also be generated from aromatic amino acids (phenylalanine, tyrosine) which are present in considerable quantities in recent sediments. An indication toward this supposition is given by work of J. M. Hunt (personal communication, 1964). A variety of compounds being or resembling the steroids structurally are found in coals, waters, sediments, and soils. Most thoroughly studied are resins such as amber, kauri, or colophony. They are plant exudates and are associated with many coals, or can occur as detrital constituents in sediment deposits

DIAGENESIS OF ORGANIC MATTER

365

(e.g., the Baltic Sea). Among steroids reported from recent and ancient sediments and natural waters are the animal-derived cholesterol and the closely related phyto-sterols (SCHREINER and LATHROP, 1911, 1912; TRASKand Wu, 1930). Especially noteworthy is the presence of vitamin Dz (calciferol), in view of its importance in biological mineralization processes. In summarizing all geochemical data on fossil lipids, isoprenoids, and steroids, it becomes apparent that our information regarding the diagenetic fate of these classes of compounds is rather limited. A total of less than 100 reliable analyses is all that has been published on these constituents. More work is urgently needed in this area, particularly because lipids and related compounds appear to be rather stable, when compared to the metastable proteins and carbohydrates. They further represent a prospective source for hydrocarbons. Heterocyclic compounds A wide spectrum of heterocyclic compounds has been isolated from coals and petroleum. The nitrogen members received particular attention largely because of their detrimental effect on the storage ability of petroleum products, and their poisoning effect on catalysts. In addition, coal tars and crude oils serve as a potential base for the industrial production of many heterocyclic compounds such as pyridine, carbazole, acridine, quinoline, and many others. The use of heterocycles for the elucidation of the petroleum problem, or the formation of coals, has been recognized relatively early. Pyrolle compounds (porphyrins) bearing structural relationship to chlorophyll and hemin were among the first organic constituents extracted from bituminous sediments, and were thoroughly studied from a geochemical point of view (TREIBS, 1934, 1936). Porphyrins, present in bituminous materials, are preferentially complexed with either vanadium or nickel. Fossil porphyrins, possessing other metals, for instance uranium, or those present as metal-free chlorins, are comparatively less frequent. In contrast, recent sediments contain only chlorin pigments, and lack porphyrins. It is generally agreed that most of the porphyrins in sediments are derived from chlorophyll, and to a lesser extent from hemin or its derivatives. For this reason, it is of interest to know the mechanism by which the chlorophyll molecule, a dihydroporphyrin, is diagenetically altered to a metal-complexed porphyrin. Various suggestions have been made as to the most likely alteration mechanism (BLUEWR, 1950, 1962a, b; GROENNINGS, 1953; DUNNING et al., 1954; HODGSON and BAKER, 1957; ORRand GRADY, 1957; VALLENTYNB and CRASTON, 1957; ORR et al., 1958; HODGSON et al., 1960; and BLUMBR and Om", 1961). According to some of the authors, alteration of the original chlorophyll may proceed by a direct replacement of magnesium by another metal. The other alternative involves a twostep process: (I) magnesium is expelled to give rise to a free chlorin pigment, and

366

E. T. DEGENS

(2) the “secondary” chlorin is re-complexed with a metal. Geological data (VALand CRASTON,1957; ORRet al., 1958; RITTENBERG et al., 1963) and experimental data (LAMORT, 1956; HODGSON and HITCHON,1959; and BLUMERand OMENN,1961) indicate the ease by which chlorophyll is converted to pheophytine. Actually, pheophytine is the predominant chlorin pigment in recent sediments, and its concentration ranges from 1-100 p.p.m. The conversion of chlorins into porphyrins may be accomplished by the reduction of the vinyl into an ethyl group, and the elimination of the carbonyl group. This reduction causes dehydrogenation at positions 7 and 8 with the simultaneous formation of another ethyl group. Thus, the chlorin is converted to a porphyrin pigment. There are two conceivable ways by which pheophytin is altered subsequently to a metal-complexed porphyrin. In the first case, reduction and dehydrogenation of the chlorin precedes the metal complexation; in the second case, the metal is introduced into the chlorin structure prior to the porphyrin formation. Neither free porphyrins nor vanadium and nickel intermediates are known from recent sediments. Namely, there is a missing link between the occurrence of pheophytin in recent and metal porphyrins in ancient sediments. Experimental data by HODGSON et al. (1960), showing the ease by which a metal-free chlorin can be altered to a metal-complexed chlorin, may support the second alternative. It is noteworthy that nickel apparently complexes more readily with the chlorin than does the vanadium. The complexity and sometimes reversibility of the reactions involved in the chlorin-porphyrin system has been clearly demonstrated by BLUMERand OMENN (1961) in studies of uncomplexed chlorins of Triassic age. The ratio of hydrocarbons to pyrolle compounds in recent reducing sediments is mostly less than 1011, as compared to lO,OOO/l in crude oils. Diagenetic generation of hydrocarbons (factor l,O00), a preferential loss of pigments, or both factors, may be responsible for the observed distribution pattern in crude oils. A positive redox potential may favor the elimination of pigments from the strata, as can be inferred from the general low pigment content in recent oxidizing sediments when compared to their reducing counterparts. For reasons of thermal stability, vanadium and nickel are two of the heavy metals more likely to become associated with porphyrins. The quantities of both elements necessary to complex all available chlorins in a sediment can easily be obtained from the surrounding organic and mineral matter (KRAUSKOPF, 1955; KEITHand DEGENS, 1959). The ratio of nickel and vanadium porphyrins can fluctuate appreciably from one formation to the other. Perhaps future studies may show the usefulness of porphyrins for stratigraphical correlations. Tryptophan is the main parent compound for indole acids, neutral indoles, indoxyls, and other metabolic breakdown products such as kynurenine. A great LE”E

DIAGENESIS OF ORGANIC MATTER

367

1952; SAUBR et al., 1952). variety of indoles are reported from crude oils (LOCHTE, The total yield in indoles and other nitrogen heterocycles, including carbazoles, pyrolles, quinolines, and pyridines, may range from about 10 to l,OOOp.p.m. of the petroleum. Information on indole compounds in sediments and soils is lacking. In view of the metabolic importance of indoles and their apparent stability, as indicated by their presence in crude oils, a careful geochemical examination might be rewarding. The occurrence of purine and pyrimidine bases in soils, peats, and a few sediments has been established by various investigators (SCHREINER and SHOREY, 1910a, b; SCHREINER and LATHROP, 1912; SHOREY,1913; BOTTOMLEY, 1917; WRENSHALL and MCKIBBIN,1937; WRENSHALL and DYER,1941; ADAMSet al., 1954; ANDERSON, 1958, 1961; PRASHNOWSKY et al., 1961). The type of bases recognized are characteristic of deoxyribonucleic acid and, to a lesser extent, of ribonucleic acid. The total values for the bases range from about 10 to 100 pmoles /1OOg soil.The proportions of the bases indicate that they were probably in the form of polynucleotides derived mainly from bacterial DNA. The bacterial origin is inferred from the fact that cytosine and guanine are in excess of thymine and adenine, respectively (ANDERSON, 1961). In addition to the afore-mentioned bases, uracil and the two oxypurines, hypoxanthine and xanthine, have been found even in Paleozoic sediments. All published results on the distribution of purine and pyrimidine bases in fossil material, however, are too fragmental to allow a critical evaluation concerning their diagenetic stability. Some important vitamins are heterocyclic compounds; and those isolated from natural waters, soils, and sediments include thiamine (vitamin Bl), riboflavin (vitamin Ba), nicotinic acid (vitamin Be), folic acid (vitamin Blo), cyanocobalamin (vitamin BIZ),and biotin (vitamin H) (LILLYand LEONIAN,1939; HUTCHINSON, 1943; HUTCHINSON and SETLOW, 1946; Romm and SCHOPFER, 1950; SCHMIDTand STARKEY, 1951; DROOP, 1955; and BURKHOLDER and BURKHOLDER, 1956). Benzo derivatives of the pyrones (coumarins, flavones), as well as anthocyanins, have not been geochemically investigated. It is conceivable that the first constituents, like p-carotene, may become degraded to benzenoid hydrocarbons upon thermal treatment. They further are very reactive and may polymerize with other materials. In conclusion, it can be said that certain heterocyclic compounds have been isolated from geological materials. With the exception of the porphyrins, however, our knowledge concerning origin, stability, and eventual use of heterocycles for the interpretation of geological phenomena is rather poor. In view of their biological abundance, their structural composition, and stability during diagenesis, heterocyclic compounds are expected to be widespread in recent and ancient sediments.

368

E. T. DEGENS

Phenols, quinones, and related substances Major biological sources for aromatic oxygen compounds and quinones in organic matter that is found in soils and sediments are ligmns, tannic substances, aromatic amino acids, and monomolecular propylphenols. In the living plant material, the propylphenols act as respiratory chromogens, and can be regarded as the likely precursors of lignins. Recent biochemical studies have even put forward the idea that lignins are also loosely linked with the respiration of the cell, rather than being only physical impregnations of the maturing cell. During diagenesis and soil formation, aromatic oxygen compounds, originally supplied by plants and animals, may undergo a number of compIex chemical transformations. These alterations are in most instances linked to microbial activity or are a result of non-biogenic maturation processes. Eventually, they lead to the formation of a variety of geochemical substances known as humic acids, hymatomelanic acid, crenic and apocrenic acids (fulvic acids), humic substances, humin, ulmin, and kerogen. All of the listed organic geochemical compounds refer largely to certain organic fractions obtained by specific analytical extraction techniques. In most instances, a knowledge of the molecular structure is lacking, or rather hypothetical. The idea has even been put forward (KONONOVA, 1956, 1961) that all of the forementioned humic materials up to the ulmins are chemically and structurally not too widely remote from each other. The conception that all humic materials have a similar if not identical molecular framework has also been extended to the kerogens (DEGENS, 1965). Namely, humic substances syngenetically produced in soils and sediments are regarded as the most likely precursor of kerogen materials, which represent 95 % or more of the finely disseminated organic matter in ancient sediments. From the data presently available, it appears that all humic materials are high-molecular weight condensation products having phenols, quinones, and amino compounds (amino acids, amino sugars, urea, and other amines) as their principal building blocks. Other biochemicals, such as sugars, fatty acids, or heterocyclic compounds (e.g., indoles, purines, pyrimidines, or pyrolle derivatives) participate in the formation of kerogens, humic acids and related substances; however, their contributions are considerably smaller. It is conceivable that the presence of carbohydrates accountsfor the reducing capacities of humic materials. The number of structural units, the mode of monomer arrangements, and the type of linkage established in humic heteropolycondensates may vary considerably. Much of the difference between the humic materials can be attributed to the variation in the ratio of aliphatic to aromatic structures. Increase in aromaticity, for instance, results in a decrease of the hydrophilic characteristics of the humic compounds. Although most of the humic materials formed-or, more precisely, preserved-over the last 500 million years are predominately aromatic in character;

DIAGENESIS OF ORGANIC MATTER

369

but non-aromatic humic compounds resembling browning reaction products may also be encountered in nature. They may arise via amino acids and sugars, but little is known concerning their structural composition and stability. Previously, it was tentatively suggested by the writer that amino acids and sugars in meteorites are part of chain molecules which are coiled so as to give stable conformations. Perhaps early forms of humic acids, humins, or kerogens that have once been synthesized on a prebiotic earth or on the meteorite parent bodies were predominately aliphatic in nature. The absence of organisms allowed the preservation of this “bouillon-type” humic material. On the other hand, this material is assumed to have served as a potential substrate for the primordial organisms. One may even speculate that the primitive living matter actually copied the molecular framework of the pre-existing aliphatic humic compounds. Comparative studies on the structure and chemical composition of humic acids and related substances extracted from meteorites and recent and ancient sediments as far back as the early Precambrian may shed light on this interesting subject. Soils and unconsolidated sediments of marine and continental origin contain significant amounts of humic acids. Very frequently, this fraction constitutes 50 % or more of the total organic matter present in the rock material (KONONOVA, 1961; DEGENS, 1965). Most scientists agree that microbial activity is largely responsible for the formation of humic acids. There is, however, still the lignincellulose controversy, particularly in conjunction with the origin of coal. This feature has to do with the ultimate source of the humic acid material and the intermediate reaction products. The most widely accepted viewpoint at present is that, although lignin is probably the major progenitor of humic acids and coals1 (BREGER,1958), other biochemicals such as cellulose and protein degradation products can contribute to this substance. The ratio of aliphatic to aromatic constituents in humic acids may depend on parameters such as the chemical nature of the original precursor material, the type of environment, and the metabolic pathway of the micro-organisms present. If, for example, wood is the only starting material, as is the case for most coals, the resulting humic acids may contain more aromatic nuclei; but if, on the other hand, plant or animal proteins and lipids are the major precursors, the corresponding humic acids may have a more aliphatic character. Upon destructive distillation, prolonged alkaline or acid hydrolysis, and other analytical means, many of the individual building blocks of humic acids can berecovered(Fig.10; DRAGUNOV, 1948; BREMNER, 1949,1951,1955,1958;HAYASHI and NAGAI,1955; LYNCH et al., 1957b; COULSON et al., 1959; KONONOVA, 1961; and SAVAGEand STEVENSON,1961). Especially noteworthy are the high yields in aromatic oxygen compounds, quinone derivatives, and amino acids.

The diagenesis of coals is separately discussed by TJXHM~~LLER and TEICHM~~LER (1967).

370

E. T. DEGENS prnollg

#MINO mINO PCIDS SUGAR:

2.23

-

1.64

-

0.58

-

1.52

32.2

57

003

1.5

37.1

1.88

tr.

8.14

1.56

Fig.10. Amino acid spectra of various terrestrial and extra-terrestrial materials. Explanation of abbreviations: see legend Fig.2.

In analogy to humic acid analysis, kerogens, subjected to alternate acid-base treatments for longer periods of time, yield a wide spectrum of components

(DEGENS and HUNT, 1964). In comparing kerogen and humic acid hydrolysates,

the correspondence in type of the released constituents is rather striking. In the case of the amino compounds, there is no basic difference between a kerogen and a humic acid hydrolysate (Fig. 10). That the first humic acid extract shows a somewhat more complex pattern can be attributed to adsorption phenomena. Namely, most of the initially recovered amino acids are only loosely attached to the humic acid molecule. The structurally incorporated amino acids, as obtained during the

DIAGBNESIS OF ORGANIC MATTER

371

second hydrolysis, require a more rigid hydrolysis and for a longer period of time. Rather interesting is the presence of urea, serine, taurine, and cysteic acid, and the ratio of glycine to alanine. This spectrum suggests that humic acids, as well as kerogens, incorporate a greater number of metabolic products which are most likely of microbial origin. Inasmuch as urea is slowly degraded during hydrolysis, its role in kerogen formation cannot be completely evaluated at present. But, considerin‘gthe fact that urea, and for example, polyfunctional amides, may condense with aldehydes to form stable resins (plastics), this may serve as indication of the role urea can have in the formation of kerogen-type materials. The high abundance of serine in ancient kerogens is unexpected in view of the results of VALLENTYNE (1957b) and AEIELSON (1959). Their data suggested that serine, among a small number of other amino acids, is metastable, and should not be found in fossil materials. Inasmuch as their thermal degradation experiments were performed on “free” serine and not “complexed” serine, the data obtained by different authors do not necessarily contradict each other. Some of the serine in kerogen could possibly have been derived from tyrosine. More than twenty different phenolic constituents are known from both the humic acids and kerogens. Vanillin, p-hydroxybenzaldehyde, syringaldehyde, guajacol, or the corresponding phenolic acids are particularly abundant (MORRISON, 1958; KONONOVA, 1961; DEGENS et al., 1963b).The same phenolic constituents are also recognized in microbial (HENDERSON, 1955) or analytical (BRAUNS, 1952) degradation products of lignins. In the past, this relationship was interpreted to mean that humic acids are merely alteration products of lignins, unaccompanied by major structural changes. This viewpoint is certainly incorrect. Instead, it is more likely that during decomposition,lignins undergo a series of severebiochemical and chemical modifications, resulting in the production of monomolecular aromatic oxygen compounds. Upon oxidation of phenols, the formation of quinones can be initiated. Both phenols and quinones may take part in a condensation reaction with other simple molecules such as amino acids, urea, sugars, and heterocyclic compounds, to form the nucleus of the humic acid molecule. Aside from lignins, aromatic oxygen compounds or quinone derivatives can be supplied by tannic substances, aromatic amino compounds, and microbial propylphenols. Kerogen, therefore, appears to be in some approximation a denatured form of humic acid. The loss in oxygen and nitrogen, and the gain in carbon, when compared to the elemental composition of humic acids (Fig. 11) may be linked to ( I ) dehydration, (2) decarboxylation, (3) loss of carbonyl and hydroxyl groups, and (4) deamination processes. Due to the resulting increase in aromatic structures, the organic residue becomes less soluble in aqueous media. The tight fixation with the associated mineral matter by ion exchange or chemisorption will further enhance the stability of the organic residue, and may partly account for the poor extraction characteristics which this material exhibits when treated with mineral acids or bases (FORSMAN and HUNT,1958a,b). Upon increase in grade of meta-

372

E. T. DEGENS

Oxygen

Hydrogen

Carbon

+

7.5

30

3

5

20

2

60

2.5

10

1

50

0

0'

0

80

Nitrogen

4r

9Or

z W

" 70 (1I

w

a

Soil Hurnin

[LI] Shale Kerogen

Limestone Kerogen

Fig.11. Carbon, hydrogen, oxygen, and nitrogen in soil humin, and kerogens. (After KONONOVA, 1961; and FORSMAN and HUNT, 1958a, b.)

n

Hydrogen

Nitrogen

6

3

4

2

2

1

0

0

100

60

80

40

60

20

40

0

c

c "

0 L

Woad

Peat

Lignite

Law-rank Bituminous Coal

Highwnk Bituminous Coal

Anthracite

Soil

Humic Acid

Shale Kerogen

Fig.12. Comparison of carbon, hydrogen, oxygen, and nitrogen content in various humic materials. (After KONONOVA, 1961; FORSMAN and HUNT, 1958a,b; and FRANCIS, 1961.)

DIAGENESIS OF ORGANIC MATTER

373

morphosis or heat treatment, the total yield in nitrogen, oxygen, and hydrogen gradually declines, and in its final stage, the kerogenous material seems to acquire the characteristics of graphite. The relationships between the elemental composition of different humic acids and kerogens are presented in Fig.12. The chemical nature of kerogen appears, therefore, to depend largely on four parameters, that is (I) the chemical composition and the molecular size of the humic acid precursor, (2) the redox potential, particularly in the early stage of diagenesis (phenol-quinone relationship), (3) the thermal history of the host rock, and (4) the type of sediment in terms of mineral composition and water content. With regard to the first parameter, there are certain chemical fluctuations which can be observed even in present-day environments. Although most humic acids are metabolically related to microorganisms, yet the original biochemicals, as supplied by plants and animals, will exercise some control on the chemistry of the metabolic end-products termed humic acid, humin, kerogen, etc. The paleontological record shows ample evidence for the wide variation in plant and animal species throughout the stratigraphical column. Consequently, the humic acids, for example, in Cambrian time, are expected to show some differences when compared to those formed earlier or later in the history of the earth; kerogens, consequently,will also differ. Very promising in this respect are the phenols and quinones, because lignins and tannic substances appeared rather late in the evolution history of biochemicals. Perhaps the limited number of phenols present in Precambrian kerogens, when compared to Recent and Tertiary kerogens, may be a result of this biochemical phenomenon. The spectrum of amino acids in marine kerogens of the last 1.5 billion years, however, does not show formational differences. This may be attributed to the fact that the source material (e.g., plankton) did not change over the geologic periods involved; also microbial activity must have been uniform. Another alternative explanation is that only a few amino acids are selectively incorporated into the humic acid molecule. An indication toward the last supposition is given by the analysis of a recent marine humic acid. Although a wide spectrum of amino acids is available (Fig.10: Istextruct), only some, and not necessarily the most abundant amino acids, are structurally fixed (Fig.10: 2ndextract). But these combined amino acid spectra are about the same as those found in most fossil marine organic matter. Similar amino acid patterns are obtained even from extra-terrestrial materials (Fig.10), which supposedly are abiotic in origin and a few billion years old. In contrast, meteorites show pronounced differences in distribution of phenolic constituents when compared to terrestrial materials, both qualitatively and quantitatively. In carbonaceous chondrites m-hydroxybenzoic acid exceeds in abundance by far (by a factor of 10 or more) any of the other fifteen phenolic compounds (KAPLAN et al., 1963). On the other hand, p-hydroxybenzoic, vanillic, and syringic acids are more dominant (by a factor of 5-10) in marine sediments

374

E. T. DEGENS

than the remaining phenolic compounds (DEGENSet al., 1963b). Furthermore, ten of the extra-terrestrial phenolic constituents have not been identified in any of the terrestrial sediments studied so far. In conclusion, humic acids formed from biogenic materials can be both aromatic and aliphatic in nature. The aromatic ones appear to be diagenetically more stable. Humic acids are metabolic products of microbial activity and constitute most of the organic fraction in soils and recent sediments. During diagenesis, humic acids, as well as their less soluble humins and ulmins, are altered into kerogen-type materials. Most significant is the loss of oxygen functions during diagenesis. Methoxyl groups are probably the first ones to become eliminated. As evidencedby the presence of quinones and phenols, carbonyl and, in particular, hydroxyl functions appear to be rather stable. Most kerogens still contain small amounts of carboxyl groups which are missing in coals having the same carbon content (rank). As far as nitrogen and sulfur functions in humic acids and kerogens are concerned, it is tentatively suggested that most of the nitrogen is present in the form of amines and heterocycles. Thioether and thiophenic groups may account for the organic sulfur. More work is urgently needed to determine the type, abundance, and distribution of functional groups in kerogens. Analogous to coal (FRANCIS, 1961; VANKREVELEN, 1961), this then may allow the determination of the diagenetic "rank" of the kerogen under consideration. Supplementary work on the identity of the individual amino acids, amines, phenols, quinones, heterocycles, fatty acids etc., may eventually permit the outlining of the diagenetic fate and the thermal history of the studied specimen.

Hydrocarbons Hydrocarbons can be synthesized by plants and animals (SMITH,1954; BLUMER, 1961, 1962b; ERDMAN,1961; MEINSCHEIN, 1963). In general, the concentrations do not exceed a few tenths or a few hundredths p.p.m. of the total living matter. The same level of concentration is also found in most recent and ancient sediments. Only petroleum pools represent natural hydrocarbon concentrates. For this reason, the origin of crude oil usually receives particular attention. The compositionof petroleum varies widely among the various oil-producing areas throughout the world. Saturated hydrocarbons from C1 to CSO constitute a substantial part of the crude oils. Naphthenes (cyclopentane and cyclohexane derivatives) and aromatic hydrocarbons also make up a significant portion of petroleum. Analyses of representative petroleums have been presented by SMITH (1952) and ROSSINI(1960). Sedimentological, petrographical, and tectonic factors that control the distribution of crude oil are now relatively well understood. One is less certain, however, about the ultimate source of the hydrocarbons and associated compounds,

DIAGENESIS OF ORGANIC MATTER

375

as well as the mechanisms that cause and govern the migration and accumulation of the oil phase. Although the number of opinions expressed on the petroleum problem are numerous, most hypotheses on this subject are related to either one of the two concepts: (I)hydrocarbons in crude oils are for the most part biogenic in origin, and (2) the bulk of the hydrocarbons are products of diagenesis, i.e., formed during diagenesis from the finely disseminated organic matter. Indeed, there are striking similarities established betweep hydrocarbons found in biogenic materials, soils, sediments, and crude oils which would support the first theory. Accordingly, chemical conversions of organic matter to petroleum, or significant changes in petroleum composition, do not occur during the postdepositional history of petroleum. But, on the other side, such pronounced differences exist between these compounds that it is rather difficult to believe that petroleum deposits are simply accumulations of biogenic hydrocarbons as proposed by MEINSCHEIN (1963) and others. Some facts in support of the diagenetic theory are as follows: ( I ) It has been demonstrated by VEBERand TURKELTAD (1958), SOKOLOV (1959), ERDMAN (1961), and HUNT(1962) that light hydrocarbons in the C3 to c14 range are absent from present-day organisms and recent sediments. They occur in ancient sediments, however, and constitute about 50 % of the average crude oil. (2) Ancient shales of non-petroleum areas contain an average of about five to ten times as much hydrocarbons as recent sediments of the same lithology (HUNTand JAMESON, 1956; HUNT, 1961). (3) Crude oil tends to get lighter with depth of burial, i.e., oils in a single system (genetically related) become more and more paraffinic, which means enrichment in the molecules that are more stable at reservoir conditions (MCIVER et al., 1963). (4) Recent, and to a much lesser degree ancient, sediments show a preference for odd- over even-numbered paraffins; in contrast to petroleum paraffins that exhibit no such preference (BRAYand EVANS,1961), as shown in Fig.13. Although biogenic hydrocarbons cannot be completely dismissed as a source for the hydrocarbons in petroleum deposits, the bulk of the hydrocarbons appears to be diagenetically produced from pre-existing fossil organic matter present in sediments, interstitial waters, or the petroleum deposits themselves. Precursor materials for the light hydrocarbon fraction (C3 to c14)include, for example, polyunsaturated fatty acids, amino acids, or isoprenes (ERDMAN, 1961; HUNT, 1962; ABELSON and HOERING,1963; DAYand ERDMAN, 1963; and MULIKand ERDMAN, 1963). They may also be generated during thermal degradation of kerogens which have incorporated some of the afore-mentioned organic molecules or some other hydrocarbon functions. Consequently,this reaction would take place within the source sediment; it may be catalytically influenced by the presence of clay minerals. Namely, the keynote of this hypothesis is that only

'""i

376

E. T. DEGENS RECENT SEDIMENT (WEST CORTEZ BASIN)

4 00

u-

2 E

+

Ic

0

0

'

T

'

'

l

l

~

'

'

~

'

'

J

PENNSYLVANIAN SHALE

400-

(SCURRY CO, TEXAS)

300-

L

0)

5 200-

z 2! .c

s

d

1000

-

"

"

"

"

"

"

'

CRUDE OIL

Fig.13. n-Paraffin distributions for a recent sediment, a Pennsylvanian marine shale, and a 1961). Pennsylvanian crude oil; r = odd- to even- carbon-number ratio. (After BRAYand EVANS,

simple chemical reactions such as decarboxylation, deamination, or elimination of alkyl groups are involved in the diagenetic transformation of former biochemicals into more stable aromatic and aliphatic hydrocarbons. Other investigators assume that before the light hydrocarbons are generated, a so-called "proto-petroleum'' concentrate accumulates, consisting only of highmolecular weight compounds (BROOKS,1952; ANDREEV et al., 1958; BOGOMOLOV and PANINA,1960; BOGOMOLOV et al., 1960; DOBRYANSKIY, 1961; MCNABet a]., 1962; and SILVERMAN, 1962, 1964a, b). Thus, most of the light hydrocarbons are formed in the crude oil deposit itself, i.e., from the pre-existing high-molecular weight hydrocarbons. The inference is that petroleum and its precursor concentrate is a dynamic mixture of chemical constituents that are subject to change thoughout the post-depositional history. The question as to which one of the theories is correct cannot be fully answered at the present time. It appears, however, that the light hydrocarbons can be generated from former biochemicals in the source sediment, and from heavy hydrocarbons in the petroleum reservoir as well. Analyses of numerous recent sediments indicate that molecules having odd

DIAGENESIS OF ORGANIC MATTER

377

carbon numbers predominate in the heavy normal paraffins in sediments of all environments (BRAY and EVANS,1961). The predominance of odd-numbered chains may indicate a metabolic process in which a p-keto acid has been decarboxylated and the resulting ketone has been reduced. Ancient sediments exhibit a smaller ratio of odd- to even-carbon number n-paraffins, and petroleums show no such preference (Fig.13). The odd-even pattern may be explained by the data of HUNT(1961) on the distribution ofhydrocarbons in recent and ancient sediments. Inasmuch as ancient sediments contain about five to ten times as much hydrocarbons as their recent counterparts, one has to assume that additional amounts of hydrocarbons were generated in situ from the residual organic matter. Migration from outside sources into the sediment is also possible, but only in oil-producing areas. In any event, both processes would dilute the original biogenic mixture and change the distribution to give the small odd- to even-number carbon ratios. The relationship that exists between odd- and even-carbon number fatty acids and paraffins has already been discussed in this chapter. For structural and chemical reasons, lipid materials are generally regarded as the most likely precursor material for the heavy hydrocarbons (CISand higher) (BREGER,1960). According to Breger, biological degradation and associated DielsAlder reactions may convert fatty acids and related substances into straightchain compounds with both odd and even numbers of carbon, to alicyclic compounds, and, upon dehydrogenation, into aromatic substances. Many more reaction schemes have been presented (HANSON,1959; ERDMAN, 1961), all of which stress the importance of lipid materials as a convenient precursor of some petroleum hydrocarbons. Ancient sediments of all environments yield considerable quantities of hydrocarbons. HUNT(1961) has estimated that hydrocarbons of all non-reservoir sediments of the world would produce more than a hundred times as much petroleum as is estimated to represent the world’s crude oil reserves. Namely, it is necessary to extract only small quantities of sediment hydrocarbons, but over a large area, to form petroleum pools. Keeping this in mind, one may assume in following BAKER’S theory (1960) that hydrocarbons are extracted from a sediment during compaction and dehydration only in quantities which are soluble in the interstitial aqueous phase. As a general rule, increase in molecular weight will decrease the solubility of hydrocarbons in aqueous solutions. On the other hand, increase in salt concentration and geothermal gradient will enhance their solubility. Consequently, the extracted high-molecular n-paraffins will have no odd- over even-number preference, inasmuch as only small portions of the individual hydrocarbons-that is, those soluble under the given environmental set-up-are released from the sediment. This would fit the odd-even pattern of the n-paraffins (c15 c80) observed in crude oils. Due to the ion-filtration on charged-net clay membranes and compaction phenomena, the interstitial solutions carrying the hydrocarbons in a dissolved or

378

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emulsified state will eventually become supersaturated. As a result, they have to release the hydrocarbons and related substances in the form of an organic phase (oil-droplet formation). From here on, the formation of crude oil deposits just depends on the general geological setting. Much has still to be learned about genesis of hydrocarbons and the processes responsible for their final accumulation. The outlined ideas represent the first step toward unraveling the petroleum problem. Both aliphatic and aromatic hydrocarbons have been isolated from various meteorites. The most extensive survey has been made by NAGYet al. (1961) and MEINSCHEIN (1961b). These authors concluded that the organic material was of biogenic origin and an indication for extraterrestrial life. According to ANDERS (1962) and FOWLER et al. (1962), hydrocarbons can be synthesized in outer space at some time during the formation of the solar system. KAPLANet al. (1963) also concluded that the hydrocarbon fraction isolated from a number of different meteorites is abiotic in origin. DEGENS(1964) stated that hydrocarbons and other volatile organic constituents in meteorites are exclusively products of diagenesis. They were formed when the fragments in orbits, obtained during collision and disintegration of asteroids, became exposed to higher temperature (not exceeding 370 OK).Namely, the finely disseminated organic matter in meteorites served as the source for the relatively young hydrocarbons and related substances. Inasmuch as fragments of “secondary” bodies stay on an average only a few million years in orbit, there is sufficient time and an ideal diagenetic environment to generate any of the found hydrocarbons. Thus, similarities between terrestrial biogenic and extraterrestrial abiogenic hydrocarbons do not necessarily imply that the latter ones have also been synthesized by organisms. It is more reasonable to assume that the first organisms evolving on earth utilized and copied the molecules of the pre-existing abiotic matter. Asphalts and allied substances As in the case of the hydrocarbons of the gas and gasoline fraction, asphaltic constituents have never been identified as direct biogenic products; also, they have not been found in recent sediments. Thus, it appears that asphalts are a product of slow inorganic maturation, generated in some fashion during diagenesis. Various alternatives have been offered as to the most likely source of asphaltic constituents. (1961, 1963), the evidence for fused ring centers According to ERDMAN strongly suggests that cellulose may have been a contributor, and the more aliphatic matrix was provided by lignin. It may be that n-heterocycles are linked to biogenic purines and pyrimidines. Sulfur is assumed to be supplied later in the process of genesis. SILVBRMAN (1964b), on the basis of stable carbon isotope studies, tentatively

DIAGENEEIS OF ORGANIC MATTER

379

concluded that the asphaltenes are polymerization products of unsaturated compounds that developed during the generation of light paraffins from high-molecular weight parent molecules (proto-petroleum hypothesis). According to PFEIFFER and SAAL (1940), the asphaltenes consist of highmolecular weight hydrocarbons of predominately aromatic character with a comparatively low hydrogen content; they were formed by condensation and dehydrogenation of aromatic-naphthenic hydrocarbons of low-molecular weight. This model suggests that asphaltenes are centers of micelles. The important single factor that controls the composition and molecular structure of bitumens1 seems to be the salinity, as evidenced by the study of HUNT (1963). With increase in salinity, the dominant molecular structure of the asphaltic materials changes from paraffin chains to aromatic rings and to chain and ring compounds high in sulfur and nitrogen. Namely, in the case of the investigated Uinta bitumens, the sequence ozocerite-albertite-gilsonite-wurtzilite developed with increase in salinity. Weathering is another significant factor controlling the chemical composition of asphalt. For example, whereas unweathered solid grlsonite contains only traces of oxygen, the same product, when weathered, yields a few per cent oxygen. Inasmuch as this enrichment in oxygen is unaccompanied by changes in carbon hydrogen, nitrogen, and sulfur contents, the formation of stable oxygenated compounds, rather than increased polymerization, may be assumed. High redox potentials may produce similar alterations during diagenesis.

DISTRIBUTION OF CARBON ISOTOPES IN ORGANIC MATERIALS

Organic carbon exhibits a range of 11% in W/12C ratio, Factors causing this variation have been considered by WICKMAN (1952), BAERTSCHI(1953), CRAIG (1953,1957), BOWEN (1960), PARKand DUNNING (1961), PARKand EPSTEXN (1961a, b), and others. From these studies, it is apparent that photosynthesis exercises major control either directly or indirectly on the distribution of carbon isotopes in plants and animals. 13C/12C ratio ranges of common organic materials are presented in Fig.14. The data are largely obtained from WICKMAN (1952), CRAIG (1953), LANDERGREN (1954), JEFFREYet al. (1955), KRFUCI-GRAF and WICKMAN (1960), ECKEZMANet al. (1962), and S I L ~ R M A(l962,1964a, N b). Marine organisms are about lo%,heavier than Iand plants. This difference can be largely attributed to the fact that during 1 Bituminous substances can be grouped into the organic materials that are soluble and insoluble in carbon disulfide. Petroleum, asphalts, coals, and kerogen are four of the more common bituminous substances. Bitumens have been classified by ABRAHAM (l960,1961,1962a,b, 1963) and HUNT(1963).

380

E. T. DEGENS

Average Values

Atmospheric C 0 2 Marine Algae

rine Petroleum I

1

-35

-30

-25

-20

-15

-10

-5

s ’%

I

0

1

t5

Fig.14. 613C in various organic materials.

photosynthesis marine organisms utilize carbonate and bicarbonate ions in the sea, whereas land plants thrive on the isotopically lighter C02 in the atmosphere. In contrast, fossil organic matter, i.e., coal, kerogen, and petroleum, are surprisingly uniform in carbon isotope composition. About 90% of all samples investigated so far fall within a narrow range of about 5 X0. For example, in the case of the coals, there is no correlation between isotopic composition, degree of coalification, and geologic age of the coals. Terrestrial coals of all formations show a similar isotopic composition of about-25X relative to the PDBI Standard (CRAIG,1953, 1957). This was interpreted to mean that no isotopic fractionation occurred during diagenesis, and that the former land plants had essentially the same isotopic composition as modern wood specimens. More difficult to explain are the 5-10%, lighter values of marine kerogen and petroleum, when compared to recent marine organisms. In the case of petroleum, some investigators (SILVERMAN and EPSTEIN,1958; PARK and EPSTEIN, 1961a,b; SILVERMAN, 1963) attributed this feature to the fact that only the lipid fraction contributed to the petroleum; lipids are generally only a few permil (up to about lo%,) lighter than the rest of the organic matter. Inasmuch as the total organic matter (kerogen) of a petroleum source rock is isotopically similar, if not identical, to the generated petroleum (KRETCI-GRAF and WICKMAN,1960; ECKELMAN et al., 1962), one may assume that (I) the marine kerogen is also a former lipid concentrate; (2) petroleum and kerogen became isotopically lighter during diagenesis, and at the same rate; or (3) land derived organic matter, deposited under marine conditions, was the major precursor for marine kerogen and petroleum. Most geologists believe that the marine plankton is the source material for

DIAGENESIS OF ORGANIC MATTER

381

crude oils (VASSOEVICH, 1958; KREJCI-GRAF,1960). But it was shown that microorganisms and burrowing animals clean up rather rapidly the nutrients supplied from the sea. Their metabolic products (for simplicity, termed humic acids), however, may survive diagenesis. It is this material and the land-derived humic acids formerly carried into the sea in solution, or via clay minerals, that eventually end up as marine kerogen and perhaps generate most of the hydrocarbons and asphalts in the course of diagenesis. Some of the 12C-enrichment can, therefore, be explained by terrestrial contributions. Another way to decrease the 13C/12Cratio has been shown by A~ELSON and HOERING(1961). Their work indicates that amines, formed by decarboxylation of amino acids, can be enriched in 12Cby as much as 20%,, relative to the original amino acids. More research work such as that performed by ABELSONand HOERING (1961) should be conducted on fatty acids, hydrocarbons, and other organic materials, in order to show what effect the elimination of carboxyl and other functions will have on the isotopic composition of the organic residue. Aside fromW/12C, studies on DH/H2,180/160, and 15N/14Nratios may assist in the elucidation of the diagenetic history of organic matter. Inasmuch as the distribution of carbon isotopes is governed by a rather complex array of factors, it is inferred that carbon isotope data of petroleum and organic matter have to be used rather cautiouslyfor any geochemical interpretation regarding the source, environment, or chemical transformation mechanisms of organic constituents.This also appliesto the other isotopes used in biogeochemical studies.

CONCLUSIONS

Biochemical macromolecules, for instance, proteins, polysaccharides, fats, or nucleic acids, are rapidly eliminated in the early stages of diagenesis. The decomposition is largely accomplished by ( I ) microbial activity, and (2) chemical hydrolysis. Breakdown products, such as amino acids, sugars, the bases of the purines and pyrimidines, fatty acids, and phenols, can interact and give rise to constituents commonly termed humic acids, fulvic acids, humins, and ulmins. Inasmuch as most of these complex humic materials are no longer of nutritional value, they may accumulate as organic residue, even in the zone of microbial activity.Interaction with the surrounding mineral matter(e.g., clay minerals) may stabilize the organic complex or catalyze certain reactions. After termination of bacterial action, which is largely a result of the depletion in available food materials, the diagenetic history of the organic matter is that of a slow inorganic maturation and redistribution (migration). Loss of functional groups may result in the formation of hydrocarbons, phenols, amines, etc.,

382

B. T. DEGENS

and cause a reduction in aliphatic side-chains in former humic acids. The resulting coal or kerogen-type material will become more and more aromatic in nature. The number one factor causing the diagenetic alteration is thermal degradation. Although some hydrocarbons in the CI to Ca and c15 to c40 range can be produced by plants and animals, most of the hydrocarbons and all asphalts encountered in ancient sediments and crude oil reservoirs are products of diagenesis. In addition to fatty acids, a number of other biochemicals, including isoprenoids, steroids, amino acids, and cellulose and lignin degradation products are prospective source materials for hydrocarbons and asphalts. Some of these constituents also form part of the kerogen molecule. It is tentatively suggested that geologically young and, therefore, less maturated kerogens are the organic source from which hydrocarbons, asphalts and some minor constituents encountered in petroleum have been generated in the course of diagenesis. Coals differ from marine kerogens in that they have a uniform lignin-cellulose source, whereas kerogens have, in addition, a predominately protein-fat-alginate source. Perhaps this may be one of the reasons why marine kerogens yield more hydrocarbons than do the coals of the same “rank”. Thus this may mean that practically all former biochemical matter, independent of chemical nature, origin, environment of deposition, etc., is diagenetically reduced to either coal, kerogen, or the various crude oil components. If given sufiicient time, the organic residues will gradually acquire the structural characteristics of graphite, whereas the petroleum will become more and more paraffinic. Increase in temperature will accelerate the reactions. That some original building blocks of living organisms can still be recovered from ancient rocks as old as 2 billion years suggeststhat in these cases the postdepositional alteration of organic matter has not yet reached its final stage. REPBRBNCES

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DIAGENEPIS OF ORGANIC MATTER

387

graphic analyses of the carbohydrate constituents of the acid-insoluble fraction of soil organic matter. Soil Sci., 84: 40541 1. M. 0. and JAMIESON,G. W., 1963. Maturation of oil, MCIVER,R. D., COONS,C. B., DENEKAS, an important natural process. Bull. Geol. SOC.Am., Abstr., 1963. JR., P. V. and B m , R. L., 1952. The evolution of petroleum. Znd. Eng. MCNAB,J. G., SMITE Chem., 44: 2557-2563. MEHTA,N. C., DUBACH, P. and DEUEL, H., 1961. Carbohydrates in soil. Advances in CarbohyaFate Chemistry. 16. Acad. Press, New York, N.Y., pp. 335-355. M E I N S C G. ~ , W.,1961a. Significance of hydrocarbons in sediments and petroleum. Geochim. Cosmochim. Acta, 22: 58-64. M m s c m , G. W.,1961b. Hydrocarbons in the Orgueil meteorite. Proc. Lunar Planetary Exploration Coll., 2: 65-67. G. W., 1964. Origin of petroleum. In: Italian Encyclopedia OfPetroleum and Natural MLZ", Gas. University of Rome, Rome, in press. MHNSCHEIN,G. W.and KENNY,G. S., 1957. Analyses of chromatographic fraction of organic extracts of soils. Anal. Chem., 29: 1153-1161. MILLER, S. L., 1953. Production of amino acids under possible primitive earth conditions. Science, 117: 528-529. MrwLw, S. L., 1963. The possibility of life on Mars. Proc. Lunar Planetary Exploration Coll., 3 : 1-7. MIUER, S. L. and UREY,H. C., 1959. Organic compound synthesis on the primitive earth. Science, 130: 245-251. MORRISON, R. I., 1958. The alkaline nitrobenzene oxidation of soil organic matter. 2. Soil Sci., 9: 130-140. M a x , J. D. and ERDMAN, J. G., 1963. Genesis of hydrocarbons of low-molecular weight in organierich aquatic systems. Science, 141:806-807. NAGY,B. and Brrz, M. C., 1965. Long-chain fatty acids from the Orgueil meteorite, in press. G. W. and HENNESSY, D. J., 1961. Mass spectroscopic analysis of the NAGY,B., MRINSCHEIN, Orgueil meteorite as evidencefor biogenichydrocarbons. Ann. N. Y. Acad. Sci., 93: 25-35. E., 1959. Dolomitbildung unter Mitwirkung von Bakterien. Eclogue NEHER,J. und ROHRER, Geol. Helv., 51: 213-215. OPPBNHEIMBR, C. H., 1961. Bacterial activity in sediments of shallow marine bays. Geochim. Cosmochim. Acta, 19: 244-260. 036, J., 1963a. Experimentalorganic cosmochemistry. The formation of biochemical compounds. Proc. Lunar Planetary Exploration Coll., 3: 9-28. 0 ~ 6 J., , 1963b. Studies in experimental organic cosmochemistry. Ann. N.Y. Acad. Sci., 108: 64. -6, J. and KIMBALL,A. P., 1962. Synthesis of purines under possible primitive earth conditions. 2. Purine intermediates from hydrogen cyanide. Arch. Biochem. Biophys., 96: 293-313. ORR, W. L. and GRADY,J. R., 1957. Determination of chlorophyll derivatives in marine sediments. Deep-sea Res., 4 263-271. Om, W.L., EMERY,K. 0. and GWY, J. R., 1958. Preservation of chlorophyll derivatives in sediments off southern California. Bull. Am. Assoc. Petrol. Geologists, 42: 925-962. PALACAS, J. G., 1959. Geochemistry of Carbohydrates. Thesis, Univ. of Minnesota, Minneapolis, Minn., 103 pp. PALM, C. and CALVIN,M., 1962. Primordial organic chemistry. 1. Compounds resulting from electron irradiation of 14CH4. J. Am. Chem. SOC.,84: 2115-2121. PARK,R. and DUNNING, H. N., 1961. Stable carbon isotope studies of crude oils and their porphyrin aggregates. Geochim. Cosmochim. Acta, 22: 99-105. PARK, R. and EPSTEXN, S., 1961a. Metabolic fractionation of lac and lacin pIants. Plant PhysioI., 36: 133-138. PARK,R. and EPSTEIN,S., 1961b. Carbon isotope fractionation during photosynthesis. Geochim. Cosmochim. Acta, 21: 110-126. PAUL, E. A. and SCHMDT,E. L., 1961. Formation of free amino acids in Rhizosphereand Nonrhizosphere soil. Soil Sci. SOC.Am. Proc., 25: 359-362.

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389

SLOWEY, J. F., JEFFREY, L. M. and HOOD,D. W., 1962. The fatty-acid content of ocean water. Geochim. Cosmochim. Acta, 26: 607-616. SMITH,H. M., 1952. Composition of United States crude oils. 2nd. Eng. Chem., 44: 2577-2585. SMITH, P. V., 1954. Studies on origin of petroleum. Occurrence of hydrocarbons in Recent sediments. Bull. Am. Assoc. Petrol. Geologists, 38: 377-404. SOKOLOV, V. A., 1959. Possibilities of Formation and Migration of Oil in Young Sedimentary Deposits. Gostoptekhizdat, Leningrad, pp.59-63. SBRENSEN, H., 1963. Studies on the decomposition of 14C-1abeledbarley straw in soils. Soil Sci., 95: 45-51. STAUDINGER, H. und JURISCH,I., 1939. Uber makromolekulare Verbindungen; uber den Polymerisationsgrad der Cellulose in Ligniten. Papier-Fabr. (Tech.- Wiss. Teil), 37: 181-184. STEVENSON, F. J., 1957. Investigationsof aminopolysaccharidesin soils. 2. Distribution of hexosamines in some soil profiles. Soil Sci., 84: 99-106. STEVENSON, F. J., 1960. Some aspects of the distribution of biochemicals in geologic environments. Geochim. Cosmochim. Acta, 19: 261-271. SWAIN, F. M., 1961. Limnology and amino acid content of some lake deposits in Minnesota, Montana, Nevada, and Louisiana. Bull. Geol. SOC.Am., 72: 519-546. SWAIN,F. M., B L ~ N T A L A. S , and PROKOPOVICH, N., 1958. Bituminous and other organic substances in Precambrian of Minnesota. Bull. Am. Assoc. Petrol. Geologists, 42: 173-1 89. TANAKA, J. and KUWATA, T., 1928. Higher fatty acids in petroleum. J. Fac. Eng., Univ. Tokyo, 17: 293-303. TEIC&LLER, M. and TEICHMULLER, R., 1967. Diagenesis of coal (coalification). In: G. (Editors), The Diagenesis of Sediments. Elsevier, Amsterdam, LARSENand G. V. CHILINGAR pp.391-415. THUNDER,O., 1954. Studies on sphagnum peat. 3. A quantitative study on the carbohydrate constituents of sphagnum mosses and sphagnum peat. Acta Chem. Scand., 8: 989-1000. TRASK,P. D. and Wu, C. C., 1930. Does petroleum form in sediments at time of deposition?. Bull. Am. Assoc. Petrol. Geologists, 14: 1451-1463. TREIBS, A., 1934. Chlorophyll- und Hhinderivate in bituminosen Gesteinen, Erdolen, Erdwachsen und Asphalten. Ein Beitrag zur Entstehung des Erdols. Ann. Chem. Liebigs, 510: 42-62. TREIBS, A., 1936. Chlorophyll-und Hhinderivate in organischen Mineralstoffen. Angew. Chem., 49: 682-686. VALLE"E, J. R., 1956. Epiphasic carotenoids in post glacial lake sediments. Limnol. Oceanog., 1 : 252-262. VALLENTYNE, J. R., 1957a. The molecular nature of organic matter in lakes and oceans, with lesser reference to sewage and terrestrial soils. J. Fisheries Res. Board Can.,14: 33-82. VALLENTYNE,J. R., 1957b. Thermal degradation of amino acids. Carnegie Inst. Yearbook, 56: 185-1 86. VALLE"E, J. R. and CRASTON, D. F., 1957. Sedimentary chlorophyll degradation products in surface muds from Connecticut lakes. Can. J. Botany, 35: 35-42. D. W., 1961. Coal. Typology-Chemistry-Physics-Constitution. Elsevier, AmsterVANKREVELEN, dam, 514 pp. VASSOEVICH, N. B., 1958. Probleme der Erdolgenese. Z. Angew. Geol., 4: 512-515. VEBER,V. V. and TURKELTAUB, N. N., 1958. Gaseous hydrocarbons in recent sediments. Geol. Nefri, 49: 682-686. WAKSMAN, S. A., CAREY,C. L. and REUSZER, H. W., 1933. Marine bacteria and their role in the cycles of life in the sea. Biol. Bull., 65: 57-79. WEEKS,L. G., 1958. Habitat of oil and factors that control it. In: L. G. WEEKSmitor), Habitat of Oil. Am. Assoc. Petrol. Geologists, Tulsa, Okla., pp. 1-61. WHITE,A., HANDLER,P., SMITH,E. L. and S T E ~ NDEW., , 1959. Principles of Biochemistry. McGraw-Hill, New York, N.Y., 1149 pp. WICKMAN, F. E., 1952. Variations in the relative abundance of the carbon isotopes in plants. Geochim. Cosmochim. Acta, 2: 243-254.

390

E. T. DEGENS

Wm, H. B., 1956. The chemical composition of some stony meteorites. Geochim. Cosmochim. Acts, 9: 279-289. W m u w s , P. M., 1961. Organic acids in Pacific Ocean waters. Nature, 189: 219-220. WRENSHALL, C. L. and DYER, W. J., 1941. Organic phosphorus in soils. 2. The nature of the organic phosphorus compounds. A. Nucleic acid derivatives. B. Phytin. Soil Sci., 51: 235-248. C. L. and MCKIBBIN,R. R., 1937. Pasture studies. 12. Can.J. Chem., B, 15: 475. WRENSHALL, ZOBELL,E., 1946a. Studies on redox potential of marine sediments. Bull. Am. Assoc. Petrol. Geologists, 30: 477-513. BELL, E.,1946b. Murine Microbiology. Chronica Botanica Co., Waltham, Mass.,240 pp. ZOBELL,E.,1952. Bacterial life at the bottom of the Philippine Trench. Science, 115: 507. ZOBELL,E.,1963. Geochemical aspects of the microbial mdication of carbon and its compounds. In: U. cOLoMB0 (Editor), Advunces in Organic Geochemistry. Pergamon Press, London, pp.330-356.

Chapter 8

DIAGENESIS OF COAL (COALIFICATION) MARLIES T E I C m L L E R AND ROLF TEICJ3dkLER

Geologisches Landesamt Nordrheitz- Westfaletz, Krefeld (Germany)

SUMMARY

Plant matter exhibits a sensitive response to temperature increase, which occurs with the burial of peats and coals at greater depths. Criteria indicative of the rank of coal (e.g., moisture, volatile matter, carbon and hydrogen contents) permit a better evaluation of different stages of coal diagenesisl than is possible for the diagenesis of other sediments. The rank of coal increases with depth the more rapidly the rock temperature rises during the subsidence of the coal. Moreover, the duration of the heating process is of great importance, and the geochemical coalification requires temperatures of more than 50” C. Higher temperatures accelerate these processes. A strong heating of short duration, however, can have the same effect as low heating over longer periods of time. Thus, the rank of a coal is indicative of temperature only if the duration of the heating of the coal is known. Pressure may indeed change the physical properties of the coal (porosity and optical anisotropy), though generally it impedes the chemical reactions.

INTRODUCTION

In the stage of hard brown coals there are already such severe physical and chemical changes that it becomes virtually impossibledespite the prevailing low temperatures and pressures-to talk about diagenesis, and one must denote these changes as metamorphic ones. Generally, the termination of coal diagenesis is seen at the boundary between soft brown coal and hard brown coal (Table I). The so-called coalification or carbonification includes the process of diagenesis, and the metamorphic changes of the coal. The diagenesis of coals is often referred to as “biochemical coalification”, whereas the metamorphism of coals is called “geochemical coalification”. The coalification begins with peat and ends with highly metamorphic anthracite. Graphite is formed only at pressures and temperatures prevailing during the metamorphism of rocks. SCHULLER (1961) proposed that the diagenesis of rocks should be considered as terminated with the transforCoalification: Some epigenetic and metamorphic effects are included Wditors).

392

M.

TEICHMULLER

AND R. TEICHM~~LLER

TABLE I THE BROAD STAGES OF COALIFKCATION DISTINGUISHED BY MICROSCOPICAL AND CHEMICAL CHARACTERISTICS~

Coal rank

Carbon

of vitrite

Volatile matter (d.a.6) of vitrite

(010)

(010)

content

(d.a.f.)2

Moisture content .insitu (010)

Calorific value

Reflectance

'

of vitrite (kcal./kg)

of vitrinite

Plo)

Peat

m

$

5

e

Hard

a brown coal

-g 0

ca. 53-

dull

ca. 71 --a. 49&@r- ca. 77 -ca.

Bituminous hard coal

is

-

-60

Soft brown coal

low rank

high rank ~

87

-29

91

-.__

8-

___

35 ___ 4,000 -ca. 0.3 __ 25

-5,500 __ -

42 ___ 8-10-

7,000 ___ ca. 0.5

___

8.650 _ _ 1.1 ___ 8,650 -2.5 ___

Anthracite Scientific classification of coal. The right side shows the applicability of various indicators of rank. (After PATTEISKY and TEICHMULLER, 1960; and according to the INTERNATIONAL COMMIITEE FOR COAL PETROLOGY, 1963.) d.a.f. = dry, ash-free. 3 a.f. = ash-free.

mation of anthracite to graphite. Inasmuch as carbonification is a very complex process one must use different properties to measure the rank of a coal. The values of the different rank parameters change with the degree of carbonification.

THE PARENT MATTER OF COAL

Most coals originate from peats of low moors with plant associations of forests or reeds. Plants usually decompose after their death, i.e., under the influence of

393

DIAGENESIS OF COAL

TABLE I (continued)

Important microscopic characteristics

Applicabi1it.v of the different rank parameters

large pores details of initial plant material often recognizable, free cellulose

.-am .-E:

I

I

Y

no free cellulose,

plant structures still

v

5

*

recognizable, cell cavities frequently empty marked gelification and compaction takes place; vitrinite is formed

low reflecting exinite exinite becomes markedly higher in reflectance (“coalification jump”)

m ._

2

T

I

I

g

1

I

$1

xt? .

‘E

exinite no longer distinguishable from vitrinite (20% volatile matter) pronounced reflectance anisotropy

diffraction

oxygen they are converted into gaseous compounds and water. In swamps with a high water table and lack of aeration, however, the plant residues are in a reducing environment, which is conducive to the process of peat formation. One can differentiate between humic coals and sapropelic coals. Humic coals originate from true peats, which develop through the accumulation of dead plant matter at the site of the peat forming plants. Sapropelic coals, on the other hand, are formed-from organic muds, which are deposited on the floor of poorly aerated quiet-water lakes and ponds. They contain many allochthonous elements. The sapropels do not undergo peatification, but pass through a process of “sapro-

394

M. TEIC-LER

AND R. TEICHMULLER

ficationyy,which is characterized by putrefaction processes under anaerobic conditions. Inasmuch as the sapropelic coals are relatively rare, only the humic coals will be discussed here. Most humic coals originated from forest peats, and thus mainly from wood and bark substances, leaves, and roots of swamp vegetation. I n swamps with herbaceous plant associations (e.g., reed marshes), the roots of the sedges and grasses play an important role in the formation of peat. During the process of peat formation, wood, bark, leaves and roots are usually almost completely transformed into humic substances, which are characterized by a relatively high oxygen content. These humic substances eventually form vitrinite. Exinite is formed from the chemically resistant and relatively hydrogen-rich plant components such as pollen and spores, leaf epidermis (cuticles), resins, and waxes. These substances are of minor importance in comparison with the humic matter. Besides, relatively carbon-rich maceralsl occur in coals such as fusinite and micrinite, which are combined under the term inertinite. These components originate from the strong aerobic decomposition of plant residues at the Peat surface, with the exception of a part of the fusinite which is derived from fossil charcoal. Chemically the parent material of coals consists mainly of cellulose, hemicelluloses and lignins, with minor amounts of proteins, sugars, pentosanes, pectines, tannins and bitumens. The bitumens comprise such substances as fats, oils, waxes, resins, sterins, sporopollenins, cutine and suberine. The inorganic components of the coal originate partially from the plants. Most of them, however, were transported by water or air into the swamps (clay, silt, sand); or they were precipitated syngenetically or epigenetically from solutions in the peat or coal (pyrite, quartz, calcite, siderite, dolomite, etc.).

THE FORMATION OF PEAT

A prerequisite for the formation of peat is stagnant ground-water, in which the plant residues are not decomposed. The peat formation begins at the surface under oxidizing conditions. Fungi and aerobic bacteria play an important role in this process. Gradual sinking and covering by younger peat layers, however, produces reducing conditions, and Fungi and aerobic bacteria are substituted by Actinomycetae and anaerobic bacteria. The microbial activity decreases with increasing depth. Certain bacteria of the fluorescent group which are pressure resistant, however, survive up to the soft brown coal stage (BECKand POSCHENRIEDER, 1957). This phase of coalification is therefore called a biochemical one. Typical products of peat formation are the humic substances which are responsible for the dark color of peat. The decomposition of plant matter is most rapid Maceral = smallest petrographicalunit of coal, comparableto the mineral of rocks (see Table IV).

DIAGENESIS OF COAL

395

at, and immediately below, the peat surface. One can divide peat formation into: (I) a primary phase, which is effective in the “peatigenic layer” (KURBATOV, 1963) at and immediately below the surface, and which is characterized by fast oxidation processes; and (2)a secondary phase with much slower conversions in a reducing 1963). The primary phase controls the degree of decomenvironment (KURBATOV, position or of h d c a t i o n of a recent peat. A great part of the original plant structures disappear, depending on the intensity of the microbial and chemical decomposition and on the type of parent material. The amount of colloidal humic matter increases. The chemically-resistant bitumens are usually also preserved in their original form and structure (pollen, suberine, etc.). Parenchymatic tissues not protected by antiseptic agents (e.g., as in the case of sphagni) are more easily destroyed because of their high cellulose content than are lignin-rich wood tissues. A great part of the cell structure is usually conserved, even if the lignin fraction is subsequently completely converted into humic matter. One can, therefore, find well-preserved plant tissues even in coals. The more easily hydrolyzable substances of the chemical components of the swamp vegetation, such as cellulose, hemicellulose, sugar, pentosanes, pectines and proteins, are decomposed first. Lignin, tannins, and bitumens are thereby concentrated. But even lignin is attacked, mostly under the influence of woodrotting Fungi and aerobic bacteria, and is gradually converted into humic substances. The most important chemical process during the formation of peat is undoubtedly the formation of humic acids, which is enhanced by the access of air (e.g., caused by a lowering of the water table and a drying of the peat surface), by higher temperatures (e.g., in tropical climates), and by alkaline environments (e.g., as a result of the addition of lime). Humic acids, however, are not formed only at the peat surface, in the presence of air oxygen, but according to WELTE (1952) the formation of humic acids from lignin is an auto-oxidative condensation process, which takes place in a weakly acid environment under microbiological influence. In neutral or alkaline environments, however, it can also proceed as a result of purely chemical reactions. Lignin is not the only parent substance of humic matter, as has often been assumed. Tannins and decomposition products of cellulose and other carbohydrates,and of proteins as well as metabolic products of microorganisms, are also involved in the formation of humic acids. According to RAKOWSKI et al. (1963), a large part of the humic acids seems to have been formed from carbohydrates. Thus, reed peats, for example, may contain a large amount of humic acids (>55%), although the original vegetable matter contains very little lignin.Recent peats generally contain, depending on their degree of humification, 1540% of humic acids in their dry substance. The humic acid content increases or decreases with depth, depending on the primary degree of decomposition of the individual peat layers. The average water content of recent peats is about 90 %.The water-retaining capacity of peats with natural bed moisture may even reach 98%. Tropical and

396

M. T E I C W L L E R AND R. T E I C W L L E R

subtropical peats contain less water than do peats from moderate climates (VAN MOLENand SMITS, 1962). Only a relatively small part of the water fUs the large pores and can be squeezed out. By far the greatest part is absorbed and the capillary water can only be removed by drying. Air-dried peat still contains about 12% of water. Relatively little is known about the diagenesis of peats with increasing depth. Undoubtedly the pore volume decreases and the density increases with depth of burial, whereas the water content decreases. According to ZAILER and WILK(191l), the water content of air-dried p a t s increases with an increasing degree of decomposition. The chemical changes in the anaerobic zone are relatively insignificant; whereas the carbon content (dry, ash-free) in the “peatigenic layer” near the surface may increase from 45 up to 60 % as a result of the decomposition of cellulose and hemicellulose and the concentration of lignin and humic matter, there is only a small increase in carbon content at greater depths (PIGULEVSKAJA and RAKOWSKI, 1963). RAKOWSKI et al. (1963) observed up to 40% undestroyed carbohydrates in primarily weakly decomposed interglacial peats. DER

BROWN COAL

The transition from peat to brown coal (“Braunkohle”) is a gradual one. KAYSER (1952) set the boundary at a moisture content of 75%. The limit for the carbon content is often given as 60%. RAKOWSKI et al. (1963) set the boundary between peat and brown coal at the point where the carbohydrates (which can amount up to 70% of the original plant material of the peats) disappear. According to the petrographic properties in Middle Europe, the brown coals are differentiated into soft brown coal (“Weichbraunkohle”), dull brown coal (“Mattbraunkohle”), and bright brown coal (“Glanzbraunkohle”) (Table I). Dull brown coal and bright brown coal are combined in the term hard brown coal (“Hartbraunkohle”). This classification corresponds to that used in the U.S.A.: brown coals (“Weichbraunkohlen”), lignites (“Weichbraunkohlen” and “Mattbraunkohlen”), and subbituminous coals (“Mattbraunkohlen” and “Glanzbraunkohlen”); but the boundary lines are considerably shifted. Soft brown coal

Soft brown coal (“Weichbraunkohle”) is light brown to dark brown, dull and earthy. Petrographically, soft brown coal still resembles peat, but is more solid and is denser. In many soft brown coals individual plant remains, such as wood remnants, leaves and fruits, are macroscopically recognizable. The microscope reveals a relatively good preservation of many plant tissues. As a result of thin sediment cover, these structures are only slightly deformed. The cell lumens are often empty or filled with water. The humic detritus surrounding the recognizable

DIAGENESIS OF COAL

397

plant remains is loosely packed. Microbedding is not usually, or only obscurely, developed (subaquatic deposits such as gyttja and other organic muds constitute an exception). The reflectance of the humic substance in polished sections is very weak (< 0.3%, using oil immersion), and their absorption in thin sections is accordingly low. A tension anisotropy which is oriented parallel to the bedding plane does not yet exist. On the other hand, cellulose can be determined in thin sections by its optical anisotropy. It has been preserved in such tissues where lignin or cutine incrustations protected it from microbial and chemical decomposition.

Hard brown coal Dull brown coal Soft brown coal (“Weichbraunkohle”) is similar to peat, whereas dull brown coal (“Mattbraunkohle”) is more similar to the bituminous coals (“Steinkohlen”). Dull brown coal is more solid, darker in color and already well-bedded. Plant remains are only occasionally recognizable megascopically, Under the microscope one can observe a distinct compaction and homogenization of the humic substances. This phenomenon is especially striking in the case of xylites (stems): in soft brown coal they hardly differ from recent stems, whereas in hard brown coal they have become layers of vitrite. In dull brown coal the empty cell cavities disappear, and the rigid cell walls become homogenized, plasticly deformed, and compressed. One can simulate these processes to a great extent by compressing soft brown coals at room temperature, and squeezing the water out. The homogenization and compaction of the humic substance is called “gelification” (“Vergelung”) (M. TEICHM~~LLER, 1950). It has not yet been determined whether or not peptization of the humic colloids occurs during this process. The experiments of H. STACH(1948) indicate that peptization does occur, because he was able to convert soft brown coal, by peptization with alkali and subsequent drying of the coal-gel, into bright brown coal (“Glanzbraunkohle”). These colloidal-physical processes lead to a significant compaction of the humic constituents, and a consequent reduction of pore volume. The individual constituents are arranged into the bedding plane by the overburden pressure. This pressure also rearranges the smallest elements of the humus colloids, the micelles, and gives rise to a tension birefringence. These petrographical changes at the boundary between soft brown coal and hard brown coal are more striking than those between R, other coal ranks ( T E I C ~ L L E and R T E I C ~ L E1954). Bright brown coal At the stage of the bright brown coal (“Glanzbraunkohle”) the gelification is completed. Bright brown coals can not be distinguished petrographically from bituminous coals and, therefore, in the U.S.A. they are included among the

398

M. TEICHldhLER AND R. TEICHhdhLER

"subbituminous coals" (A.S.T.M. classification). Bright brown coal is solid, black, bright, and under the microscope shows the typical microlithotypes of the bituminous coals (these are obviously different in Tertiary coals and in Carboniferous coals). Wood and bark have been converted into layers of vitrite and the originally loose detritus has been compressed into compact layers of clarite and durite. The reflectance of vitrinite varies from 0.3 to 0.5 % (oil immersion). Chemical changes in the brown coal stage On the basis of the ultimate and proximate analyses, the chemical changes which accompany the significant petrographical changes at the boundary between soft brown coal and hard brown coal are relatively small. The average carbon content (dry, ash-free) of soft brown coals is 60-69%, and increases to 71-77% in the bright brown coals. The content of volatile matter varies considerably in brown coals. It decreases roughly from < 65 % (d.a.f.) in soft brown coals to > 40% at the boundary between bright brown coal and bituminous coal. The decrease in moisture and the corresponding increase in heating value are quite pronounced. When mined, a soft brown coal contains 75-35 % water, a dull brown coal 35-25 % water, and a bright brown coal 25 to about 8 % water (Table I). This decrease in water content is partially caused by the reduction in porosity. In addition, there is a progressive splitting of hydrophilic groups from the humic molecules, which react very sensitively to temperature rises. The decrease in water content slows down with increasing rank: in soft brown coals with about 60% water, the moisture of the coal as mined decreases by 1%/25-30 m of increase in depth; in dull brown coals by l%/lOOm of depth; and in bright brown coals by 1%/100-200 m of depth. Little is known yet about the chemical changes which become explicit in the functional group analysis. According to BLOMet al. (1960) the concentrations of methoxyl (-OC&), carboxyl (-COOH), and carbonyl (: C = 0) groups

TABLE II DISTrmCITON BETWEEN BROWN COAL AND BlTUMlNOuS COAL

Brown coal ("Brawrkohfe")

Bituminous coal ("Stetnkohlc")

Streak (color)

brown, seldom black

black, seldom brown

Behavior on boiling with KOH (detection of h d c acids). Behavior with dilute HNa (oxidizability)

brown solution

no color

red solution

no color

I11 COMMERCIAL CLASSIFICATIONS FOR BITUMINOUS COALS AND ANTHRACITES AS USED IN THE INTERNATIONAL E.C.E.-SYSTEM AND

IN DIFFERENT COUNTRIES

VANKREVELEN, 1961) Classes of nationbl systems the international system parameters calorific value Belgium Germany France (calcrdated to matter standard mois content ture content)

Maigre

3-6s 6.5-10 10-14

3

14-20

2 Gras

20-28 28-33

> 33 (32-40) (3244)

Gras I

8,450-7,750 7,750-7,200

(34-46) (36-48)

< 6,100

Anthrazit

Anthracite

Maigre Magerkohle Demi Esskohle gras Fettkohle

Gras B courte flamme

Gaskohle

Gras proprement dit

GasFlammkohle Flammkohle

Flambant

United Kingdot

Italy

Antraciti speciali Antraciti comuni

Meta-antracyt Anthmcite

ciet

Low volatile bituminous Medium volatile bitum.

Carboni semi-grassi Carboni grassi corta fiamma Carboni grassi media fiamma

I koksowy Gazowo

Carboni da gas vlam-

gr?lS

Carboni secchi Flambant sec

Metaanthracite Anthracite semianthracite

Carboni magri

Carboni grassi da vapore

United States

Vlamkool

Gazowoplomienny Plomienny

High volatile bituminous High volatile

High volatile bituminou;

High volatile bituminous Subbituminous

400

M. TEIC&LFZR

AND R. mcHMiiLLBR

decrease rapidly with increasing carbon content in the brown coals. The hydroxyl (-OH) concentration, on the other hand, varies and shows a significant drop only during the very first sMges of the formation of brown coal, and then only once more in the early stage of the bituminous coal formation. Lignin is gradually completely converted into humic substances. In addition, cellulose, originally present in small concentrations in the soft brown coal stage, eventually disappears. Starting from the hard brown coal stage, the humic acids gradually lose their acid character, and thus become insoluble in alkali. The molecular complexes of the humic substance are enlarged, and consequently the color and streak of the coal become darker. The oxidizability by nitric acid decreases. The boundary between brown coal and bituminous coal The transition from brown coal to bituminous coal is, like the transition from peat to brown coal, a gradual one. The boundary between brown coal and bituminous coal is drawn differently. In central Europe three characteristics are used in distinguishing between brown coals and bituminous coals. Two of these distinguishing features must be present in order to call a coal “brown coal” or “bituminous coal” (Table II). On the basis of these characteristics PATTEISKY and T E I C ~ L L (1960) ER have drawn the boundary between brown coal and bituminous coal at concentrations of about 8-10 % water, 77 % carbon (dry, ash-free), 16 % oxygen(d.a.f.), and with a heatingvalue of 7,000 kcal./kg (ash-free), all values being based on vitrites. COMMITTEE FOR This classification has been accepted by the INTERNATIONAL COALPETROLOGY (1963) (Table I). In the commercial classification of coals (E.C.E.-system) the boundary is drawn at a calorific value of 5,700 kcal./kg (ash-free). Thus, using this classification, the bright brown coals (“Glanzbraunkohlen”) would belong to the bituminous coals.

Bl+”OUS

COAL AND ANTHRACITE

The stages of bituminous coal (“Steinkohle”) and of anthracite are also defined differently. Table I11 shows the well-known commercial classifications. Depending on the range of rank they are based on different “rank measures” (calorific value, volatile matter content). The microscopic picture of a low-rank bituminous coal resembles that of a bright brown coal (“Glanzbraunkohle”): the individual macerals show distinct differences in their reflectance, color and relief in polished sections. The macerals of the exinite group ‘(protobitumen) reflect the incident light only weakly. The humic matter, which has been converted to vitrinite in bituminous coal, has a stronger reflection, appears light grey, and has a less pronounced relief. The

DLAGENESIS OF COAL

40 1

macerals of the inertinite group (e.g., fusinite, micrinite) show the highest reflectance, appear white, and have a strong relief. These differences diminish with increasing rank. The reflectance of vitrinite increases continuously (Fig.1, Table I) and so gradually that it is possible to determine rather exactly the rank of coal by reflectance measurements of vitrinites. The inertinite, on the other hand, hardly shows any changes. At first, the optical properties of exinite change only slightly. Only in the stage of the fat coal (“Fettkohle”), with 29 % volatile matter, the exinite becomes distinctly lighter in incident light. E. STACH(1958) called the beginning of this relatively rapid change “coalification step” (“Inkohlungssprung”). In the stage of the high-rank fat coal, with about 22-20 % volatile matter, the exinite becomes indistinguishable from vitrinite by optical and chemical means (Fig.2). Thus the microscopic picture of a high-rank bituminous coal shows much less variation. With increasing rank the reflectance of vitrinite increases and approaches that of the inertinite; the same holds true for the vitrinized exinite. In highly metamorphic anthracites only the strongly reflecting fusinite is easily contrasted to vitrinite. The increase in reflectance corresponds to an increase in refractive and absorption indices (Fig. 1). The increasing absorption becomes clearly noticeable in thin sections: vitrinite becomes increasingly darker, so that with increasing rank of the coal the sections must be made continuouslythinner& order to obtain transparent thin sections. These optical changes are correlated with chemical changes, i.e., increasing aromatization and condensation of the molecular groups. The growing optical anisotropy, on the other hand, can be attributed more to physical changes: with increasing overburden pressure the more or less laminated molecular complexes are arranged into the bedding planes. The tension anisotropy, which is at first only visible in transmitted light, starts to become noticeable at the fat coal (medium volatile bituminous) stage also in reflected light, and becomes more pronounced. In anthracites the tension anisotropy is very noticeable: in a highly metamorphic anthracite with 3 % volatile matter, the reflectance of vitrinite (in air) in a crosssection in polarized light may fluctuate between 18.1 and 10.5% on rotating the microscope stage (DAHME and MACKOWSKY, 1951). Although the anthracite resembles graphite, as is shown by the reflection anisotropy, the X-ray diagram, and the electric conductivity, it is still far from being a crystalline substance. Even a high-rank anthracite is still colloidal in nature; its macromolecules-consisting of aromatic lamellae-are merely oriented into the bedding plane and thus yield diffraction patterns similar to those of graphite. The formation of a true graphite lattice, however, requires pressures and temperatures which exist only during the metamorphic stage of phyllitization (SCH~~LLER, 1961). The physical and chemical changes of the coal, in the stage of the bituminous coals and anthracites, are better known than the changes in the brown coal stage. New knowledge has been gained on using colloid-chemistry techniques, the

medium volatile 'bituminous coal 1-22% voi.m.1

aromacity

ring condensation

.

max.internal moisture 'heat ofwetting [in methanol] [internal surface, porosity1

dimension of aromatic clusters lcrystallites--I

.

hardness

.

free radicals

solubility in ethylenediamine

.

grindability

Irecipr. strength I

fluidity during carbonisation

refractive index

absorption index

reflectance

403

DIAGENESIS OF COAL

/

/

/-micrinite

exinite

[soores1

lprimarily highly decomposed humus1

Fig.2. The chemical assimilation of the macerals during coalification. (Development lines based on DORMANS et al., 1957.)

structural analysis of petroleum, spectrometry (e.g., infrared, nuclear magnetic resonance, electron magnetic resonance), X-ray analysis, differential thermal analysis, electron microscopy and solvent extraction (Fig.1). The chemical changes of the coal during the formation stage of the bituminous coals and anthracites consist of an increase in carbon content and a decrease in oxygen and, subsequently, also in hydrogen content. Fig.2 shows these changes in terms of the H/C and O/C atomic ratios for the various macerals. Moreover, the content of volatile matter, which is liberated during the crucible-coking of the proximate analysis, decreases with increasing rank. As far as the chemical structure is concerned, the humic complexes become more and more aromatized and condensed (Fig.1); the aromatic clusters of low-rank bituminous coals are still relatively small. Long chains and bridges of oxygen-rich and hydrogen-rich groups (-OH, -COOH, -CHs) are attached to these aromatic clusters. With increasing rank these groups are gradually split off, which at fist yields mainly C02 and

Fig.1. Molecular structure and physical alterations of vitrite during the coalification process. hydrogen bonding, - molecular bonding) is presented The molecular structure of vitrite (-0above, whereas the positions of the elementary particles in a section perpendicularto the bedding plane are shown below.The lenses represent the aromatic clusters, and the lines depict the nonand MAGGS,1944; BLAYDEN et al., 1944; KING aromaticfringes and bridges. (After BANGHAM and WILKINS,1944; DULRUNTY, 1950; DRYDEN,1951; DRYDEN and GRIFFITH, 1953; HUCK and 1953; BROWNand H ~ c H 1955; , AVSTEN and INGUM, 1958; KARWEIL, 1953; VAN KREVELEN, and others.) Vol. m. = volatile matter.

404

M.

TEICHM~~LLW AND

R. TEICIIM~~LLW

HzO and then increasing quantities of CH4. The concentration of aromatic carbon increases at the same time, and the aromatic clusters grow. At concentrations of 29% volatile matter and 87% carbon (dry, ash-free), the rate of oxygen loss suddenly decreases. Instead, hydrogen is given off in increasing amounts as CH4. At concentrations of about 20 % volatile matter and 89 % carbon (d.a.f.), many physical properties of the coal are changing rather rapidly. At this stage the lowest microporosity and internal surface area (as determined by the heat of wetting), the lowest density, and the minimum strength are attained; whereas plasticity and coking qualities become maximum (Fig. 1). Also, the moisture content of the coal reaches a minimum of about 0.5 %. The intermolecular spaces at this stage are particularly small, because bulky bridges and chains have been degraded to a great extent. The relatively heavy oxygen has been given off for the most part, and the light hydrogen has been concentrated. The low strength results from the lack of the original web of chains and bridges.

PHYSICAL-CHEMICAL INDICATORS OF COAL RANK

One must use different physical-chemical indicators of coal rank depending on The range of rank which is to be investigated. It is always necessary to perform comparative rank investigations on the same petrographical constituent of the coal, because the chemical changes proceed differently in the different macerals (Fig.2). Vitrinite is best suited for comparative rank investigations, inasmuch as it is the most abundant component of the coal, and because it can be isolated relatively easily. In addition, it exhibits continuous changes. Table I shows the applicability of different indicators of rank to the individual rank ranges, as has been shown by binary diagrams of the individual parameters (PATTEISKY and TEICHM~LER, 1960) and by rank sections of bore-holes (Fig.3). In the stage of the brown coals and the low rank bituminous coals (up to 30 % volatile matter, d.a.f.) the moisture content or the calorific value of vitrinite, as mined, is the best measure of rank. Inasmuch as the calorific value of low rank vitrites is dependent on the moisture content, the determination of the moisture content is usually suEicient. The analysis should be performed under constant conditions (temperature, air humidity, etc.). Instead of the moisture content of the coal as mined, it is also possible to choose as a measure the maximum internal moisture content (water-retaining capacity). In the range from the bright brown coals (“Glanzbraunkohlen”) to the gas coals (“Gaskohlen”), with 30 % volatile matter (d.a.f.), the carbon and oxygen1 contents are also applicable as rank param1

Inasmuch as oxygen usually is not determined directly in the ultimate analysis of coals, and

in this rank range the oxygen content is inversely proportional to the carbon content, the carbon content should be used preferentially.

405

DIAGENESIS OF COAL

40

35

30

15

10

volatile matter 1% d . a f I

5

-80

5

moisture l%l

calorific value IkcaVkg a.f.1

85

90

carbon[%d.a.f]

5

hydrogenl%d.a.tI

Fig.3. Increase of rank with depth on the basis of different parameters (volatile matter, moisture, carbon and hydrogen contents, and calorific value). The bands are drawn on the basis of vitrite analyses of coal samples taken from deep bore-holes and shafts. (Based on P A ~ I S Kand Y TEICHM~LER, 1960.)

eters. The concentrations of volatile matter and hydrogen, on the other hand, are unsuitable as rank parameters in this range because their values vary considerably and scarcely show any correlation with depth (Fig.3). The rank of more highly metamorphosed bituminous coals (< 30 % volatile matter) is best determined by using the volatile matter content. As shown in Fig.3, the volatile matter of these coals decreases rather rapidly with depth down to the anthracite boundary (8% volatile matter content). In the Carboniferous coals of the Ruhr Basin this rate of decrease is equal to 1.9 % volatile matter per 100 m stratigraphic depth. Volatile matter is, therefore, used as a parameter in the commercial classification of bituminous coals and anthracites (Table 111). The contents of carbon, hydrogen, and moisture show only small changes in coals having 8-30% volatile matter. Starting with the anthracite stage the hydrogen content decreases considerably, whereas the carbon content increases slowly. The concentration of volatile matter decreases also relatively slowly over this range. Because of this, the hydrogen content is the best-suited indicator of rank for the anthracites. The reflectance of vitrinite in polished sections is also applicable as rank measure for bituminous coals, especially because it may be determined quickly

406

M. TEICHlbdhLER AND R. TEICHMULLER

and requires such small sample quantities, which would be insufficient for chemical analyses. The rank of even microscopically small coal inclusions in various sediments can be determined in this manner. The examination of other indicators of rank and their possible utilization is a problem for the future. Of interest would be measures which depend upon the changes in the chemical structure during coalification, e.g., the content of certain functional groups of the molecules (-OH, X! = 0, -COOH, -OC&), the content of aromatic carbon, as well as the degree of condensation and the size of the aromatic clusters. SIGNIFICANCE OF PRESSURE, TEMPERATURE AND TIME IN COALIFICATION

Pressure

In the past it was assumed that pressure plays an important role in the process of coalification. The following observations were used to support this assumption: ( I ) The rank increases with increasing depth of burial (and thus with increasing overburden pressure) - Hilt’s rule. (2) The rank of coal in the folded part of a coal basin is usually higher than in the non-folded region. The obvious conclusion, therefore, was that the coalification process was accelerated by the tangential (folding) pressures. The experimentsof HucK and PATTEISKY (1964; cf. HUCKand KARWEJL, 1962) have definitely shown, however, that the static pressure does not enhance but instead inhibits the chemical processes of coalification. The increase of rank with the depth of burial is also readily explained by the increase of rock temperature with depth. J ~ T G E Nand KARWEIL(1962) succeeded in obtaining a distinct increase of rank by applying dynamic pressure when grinding coal in a rocking ball mill. The authors themselves, however, emphasized that this experiment is not comparable to natural conditions, for, as a rule, the tectonic movements proceed so slowly that the term “dynamic pressure” (i.e., pressure which changes constantly in magnitude and direction) or even acceleration is not applicable. The small increase of rank which occasionally and locally occurs in the immediate vicinity of tectonic thrust planes could as well have been caused by the heat of friction generated during the tectonic movements. It could be shown that in the Carboniferous deposits of the Subvariscan foredeep (Ruhr Basin; Belgium) and in other coal-seam-bearing foredeeps, the carbonification was completed to a great extent before the folding processes started (R. T E I C ~ L E1952). R , Consequently, the carbonijkation cannot possibly be an efect of the folding. Inasmuch as the layers in the landward part of the foredeeps generally were buried to a greater depth, more intensive carbonification of the coal seams in this part of the orogeny is easily explained by the higher rock temperature with a greater depth of burial.

DIAGENESIS OF COAL

407

Whereas pressure obviously does not have a significant effect on the chemical process of carbonification, it can undoubtedly change the physical structure of the coal. Pressure can close the pores and arrange the micelles. Intensively folded coals are sometimes anomalously poor in moisture content because of their relatively low porosity (DUNNINGHAM, 1944; EDWARDS,1948; BERKOWITZ and SCHEIN,1952); and their optical anisotropy may have been affected by tangential pressure (WILLIAMS, 1953; PETRASCHECK, 1954; MONOMAKHOFF, 1961). The decrease in porosity of low-rank coals with depth is associated with a decrease in moisture content (Fig.3). Thus the question arises as to whether this dehydration is an effect of heating during subsidence of the coal-seam bearing strata, or an effect of pressure. Temperature

The effect of high temperature on the chemical reactions during carbonification is undeniable (see, e.g., DULHUNTY, 1954; HUCKand KARWEIL,1955; DRYDEN, 1956; VANKREVELEN, 1961). Each magmatic contact of the coal demonstrates this as well as all the experimental work. The effect is more pronounced the larger the intrusive body and the higher its temperature. Narrow vertical magmatic dikes dissipate relatively little heat; consequently their zone of influence ranges only from a few inches to a few feet. The heating of coals above large intrusive bodies (batholiths), on the other hand, is still distinctly noticeable over distances of several thousand meters, even where the host rock does not seem to be any longer affected (TEICHMULLER and T E I C ~ L E 1950). R, The effect of temperature plainly follows Hilt's rule, i.e., the increase of rank with depth. The maximum depth of burial, and the corresponding maximum temperature to which the coal was exposed for long periods of time, determines its rank. Based on experience, as a rule, this maximum temperature in the great foredeeps is attained prior to the initiation of folding (preorogenetic). This, for example, is the case in the Ruhr Basin ( T E I C ~ L Eand R TEICHMULLER, 1949). The Carboniferous strata commonly bear bituminous coals, whereas Tertiary sediments generally contain only brown coals even though the depth of burial of the seams might have been the same. This variation may be caused-at least in some areas-by a considerably lower geothermal gradient (m/"C) during the Carboniferous time, as compared to that of today (KUYLand PATLJN,1961; R , This, for example, applies to the Carboniferous deposits R. T E I C ~ L L B1962). of the Ruhr Basin, where coal pebbles from high-rank coal seams of the uppermost Namurian and the lowermost Westfalian already appear in the Westfalian C sandstones. The coal seams of the latter have been coalified merely up to the flame-coal stage (>40 % volatile matter). Inasmuch as only a period of approximately 10-20 million years was available for the coalification of the coal seams of the lowermost Westfalian, one has to assume that the temperature gradient

408

M.

TEIcHM~~LER AND

R. TEICHMULLER

(“C/m) during Westfalian time in this region was rather large. The geothermal gradient at that time may have amounted to 10-15 m/”C (today it is around 30 m/”C). This may well be caused by the occurrence of magmatic intrusions in the subvariscan foredeep. Intrusions like the granite porphyry of Elberfeld and, perhaps in connection with this, the hydrothermal mineralization (epi- and mesothermal, and locally even catathermal) may serve as examples. Such intrusions and ore dikes are lacking in the Subalpidic foredeep of Upper Bavaria, in which the coal seams merely attained the stage of the bright brown coal (“Glanzbraunkohle”), although depth and duration of burial were the same as those in the Subvariscan foredeep. Time

Inasmuch as the coals of the Lower Carboniferous in Moscow are still in the stage of hard brown coals, the conclusion has been drawn that time does not play an important role in the process of carbonification. This undoubtedly applies to those coals which have never been buried deep enough during their geological history and, therefore, have never been subjected to intensive heating. In this area the minimum temperature required for the formation of bituminous coals has not been attained. Some low-rank Carboniferous coals, however, were later buried in Mesozoic and Cenozoic times at greater depths (2,000-4,OOO m), as in the Lower Saxony Basin of northwestern Germany. In this case the coals have been subjected to temperatures of 80-140°C for periods of far more than 100 million years. This resulted in an intensive posterior coalification of the Carboniferous coal seams; the flame-coals (“Flammkohlen”) which formed during Westfalian C time were subsequently converted into forge coals (“Esskohlen”) and meagre coals (“Magerkohlen”). The fact that today the coals of the Carboniferous period are found to be predominantly bituminous coals, implies that they have been also affected by the post-Carboniferous coalifkation. The latter is important wherever the Carboniferous strata have been subsequently buried at greater depths. Thus, time contributes to the process of carbonification only if temperatures are sufficiently high. The same coal rank can be produced by either a short intensive heating or by low heating over long periods of time. KARWHL(1956) has presented a diagram which shows a correlation between duration of heating, rock temperature, and the rank of coal based on the volatile matter content.

GASEOUS PRODUCTS OF COALIFICATION

The coalification of brown coal results in release of water and carbon dioxide. In the stage of the low-rank bituminous coal (with about 29% volatile matter)

409

DIAGENESIS OF COAL.

mainly COZ is produced, whereas subsequently CH4 is the predominant coalification gas (p.404). Furthermore, small amounts of heavier hydrocarbons are split off. This is based on the observations of J ~ G E and N KARWEIL (1962) on experimental coalification. On grinding coal samples, GEDENK et al. (1964) discovered heavier hydrocarbons as adsorbed gases. Furthermore, they found that with increasing rank, hexane, pentane, butane, propane and ethane disappear successively from the coal (Fig.4). It is interesting to note that anthracite contains only methane as an adsorbed hydrocarbon gas. Considerable quantities of gas are released during the coalification process starting from a flame-coal (“Flammkohle”), containing 40% of volatile matter, up to a semi-anthracite with 10% volatile matter. According to JUNTGBN and KARWEIL (1962), about 100m3 of carbon dioxide and about twice as much methane are produced per ton (1,OOO kg) of final coal. PATTEESKY (1950) obtained even higher values in his calculations. The carbon dioxide is dissolved in water and rises with it upward along joints and cracks. Part of the methane likewise starts to migrate; it accumulates in pores and joints of the host rock of the coal seams or in porous sandstones of the overlying sediments, and can form natural gas reservoirs of commercial significance. Part of the methane, however, remains depth ft. 6,000

the increase o f rank o f coa [decrease of volatile mattei

.i

“ ‘

he disap

B’b

I

the decrease o f pore volume i n sandstones

i e increase o f seismic interval velocity

heme

7,000

pmfanc

too0

I

9,000 I

mpoo

10

20

301

volatile matter d.a.f

1

5

”

’

p o r e volume

.

,

‘

‘

‘

.

I

LpWmlrec

Wo:eirmie velocity

Fig.4. The changes of coal rank; the decrease of heavier hydrocarbons (ethane and higher homologues related to &-free and COa-free gas) in the rest-gas, and the consolidation of rocks in the flat-lying Upper Carboniferous beds of the bore-hole Miinsterland 1. (After GEDENK, 1963; LICHTENBERO, 1963; M.TEICHM~LER, 1963;and Tij”,1963.)

410

M. TEICHMULLEEAND R. TEICHMULLER

absorbed in the coal and is released only on decompression, which occurs in boreholes or during the mining of coal seams. In the coal basin of Upper Silesia, mining district of Karwin, 1 ton (1,OOO kg) of mined coal yields more than 100 m3 of methane (PATTBISKY, 1950). Such large quantities of gas, however, are rarely found in other places.

COMPARISON OF DIAGENESIS OF COALS WITH DIAGENESIS OF THE HOST ROCKS

Soft brown coals are associated with weakly consolidated clays and sands, whereas hard brown coals occur with more indurated sediments. Bituminous coals are associated with consolidated rocks; therefore, the diagenesis of coal, as a rule, runs parallel to the diagenesis of the associated sediments. An example of this is presented in Fig.4 and 5 for the range of bituminous coals and anthracites found in the Munsterland 1 deep bore-hole (in which the present depth of the Carboniferous deposits approximately corresponds to the original one). Fig.5 shows that a stronger sericitization of the feldspars and an increasing ankeritization and silicification of the sandstones begin in the range of the anthracites. At this point the seismic velocities in these layers show an abrupt increase (Fig.4). This parallelism, however, is not found everywhere. In the broad vicinity dlagenesis of t h e sandy -clayey sediments “C

ft. rn

gmo W-- 69

-- 75

w

r’

z,?w-- 80

. 1-

-.86

i

-.97 c

102

medium volatile bit.

Fettko h Ie low yol. bltum In ous Esskohle semi-

\

2 p - 41

,qmo 3,mo-.

Gaskohle

I

..108

Magcrkohle

anthracite

Anthrazit

anthracite

3,UQ--113

iz,om

--lZo

’*

3,Bw--m -.135

-.refiettancr rapid increase of anirotmpy

I@m 4m-W--I%

“ L

1w

’/.

6

4

. Z

kaolinite- illite-zone beginning of: stronger sericitlzation o f feldspars; ankeritization,and silification of sandstones, chioritization of blolites: pronounced crystallinity of clay minerals

beginning of pronounted neomineralization of r u t h mobilization of quartz,chlorite and apatite; beginning o f blastesis of sericite; no preferred orlentation of texture

0 reflectance

Fig.5. The diagenesis of coal and of host rocks with increasing depth and increasing temperature in the Upper Carboniferous sedimentsof the borehole Miinsterland 1. (After HEDEMANN, 1963; LENSCH, 1963; SCAERP,1963; and STADLER, 1963.)

DIAGENESIS OF COAL

411

of magmatic intrusions, the coal sometimes has been highly metamorphosed, whereas the host rock does not yet show any intensive metamorphic effects. In such cases the coal proves to be a very sensitive geologic thermometer, provided the duration of the heating is also known.

ACKNOWLEDGEMENT

The writers are greatly indebted to Dr. H. Reuter (Julich) for the translation of the German text. REFERENCES

AUSTEN,D. E. G. and INGRAM, 0. J. E., 1958. Electron resonance in coals. Brennstof-Chem. 39: 25-30. BANGHAM, D. H. and MAWS,F. A. P., 1944. The strength and elasticconstants of coals in relation to their ultra-fine structure. Proc. Conf. Ultrafine Struct. Coals Cokes, London, 1944, pp.118-130. BECK,TH. und POSCHENRIEDER, H., 1957. Drucktoleranz, ein Kriterium fiir den autochtonen Charakter der Braunkohlenmikroflora.Zentr. Bakteriol.Parasitenk., Abt. ZZ, 110: 534-539. BECK,TH., POSCHENRIEDER, H. und BUKATSCH, F., 1956. Untersuchungen iiber die Bakterienflora der Oberpfiilzer Braunkohle. Zentr. Bakteriol., Parasitenk., Abt. ZZ, 109: 201-225. BERKOWITZ, N. and SCHEIN, G., 1952. Some aspects of the ultrafine structure of lignites. Fuel, 31: 19-32. BLAYDEN, H. E., GIBSON,J. and b y , H. L., 1944. An X-ray study of the structure of coals, cokes and chars. Proc. Con$ Ultrafine Struct. Coals Cokes, London, 1944, pp.176-231. BJBM, L., EDELHAUSEN, L. and VAN KREVELEN, D. W., 1957. Chemical structure and properties of coal. 18. Oxygen groups in coal and related products. Fuel, 36: 135-153. BROWN, J. K. and HIRSCH,P. B., 1955. Recent infra-red and X-ray studies of coal. Nature, 175: 229-242. D-, A. und MACKOWSKY, M. TH.,1951. Mikroskopische, chemische und rontgenographische Untersuchungen an Anthraziten. Brennstof-Chem., 32: 175-186. DOWNS, H. N. M., HUNTJENS, F. H. and VANKREVELBN, D. W., 1957. Chemical structure and properties of coal. 20. Composition of the individual macerals (vitrinites, fusinites, micrinites and exinites). Fuel, 36: 321-339. DRYDZN, I. G. C., 1951. Einige neue Fortschritte zum VerstZindnis der physhlischen Struktur und chemischen Natur von GIanzkohlen. Brennstof-Chem., 32: 321-324. DRYDEN, I. G. C., 1956. How was coal formed? Coke Gus, 18: 1-11. DRYDEN, I. G. C. and GRIFFITH, M., 1953. Quantitative estimation of the changes in chemical structure of coals during metamorphism. Fuel, 32: 199-210. J. A., 1950. Relations of rank to inherent moisture of vitrain and permanent moisture DULHUNTY, reduction on drying. J. Proc. Roy. SOC.N. S. Wales, 82: 286-293. DULHUNTY,J. A., 1954. Geological factors in the metamorphic development of coal. Fuel, 33: 145-152. DUNNINGHAM, A. G., 1944. The inherent moisture in coal. Proc. Con$ Ultrafine Struct. Coals Cokes, London, 1944, pp.57-70. EDWARDS, A. B., 1948. Some effects of folding on the moisture content of brown coal. Ausfrulusian Inst. Mining Met. Proc., 150/151: 101-112. GEDENK, R., 1963. Die Zusammensetzung des Restgases in Kohlen und Nebengestein der Bohrung Miinsterland 1. In: Die Aufschlussbohrung Munsterland I-Fortschr. Geol. Rheinland Westfalen, 11 : 205-237.

412

M. TEICHI&LLER AND R.

TEICMLLER

GJDENK,R, HEDWA”, H. A. und R m ,W., 1964.Oberkarbongase, ihr Chemisrnus und ihre Beziehungen zur Steinkohle (Untersuchungsergebnisse aus Nord- und Westdeutschland). Congr.Avan.EtudesStratigraph.Gdol. Cwbong&re,Compte Rendu,5, Paris, 1963,2:431-40. HFDEMA”, H., 1963. Die Gebirgstemperaturen in der Bohrung Miinsterland 1 und die geothermische Tiefenstufe. In: A ufschlussbohrung Miinsterland I-Fortschr. Geol. Rheinland Westfalen, 11 : 403-418. €€EDEMA”, H. A. und TEI-UR, R., 1964. Das Oberkarbon der Bohrung Miinsterland 1, seine Diagenese und seine Bedeutung fiir die Palaogeographie. Fortschr. Geol. RheinlandWestfalen.In press. HUCK,G. und KARWEIL, J., 1953. Versuch einer Modellvorstellung vom Feinbau der Kohle. Brennstof-Chem., 34:97-102, 129-135. HUCK,G. und KARWEIL, J., 1955. Physikalisch-chemische Probleme der Inkohlung. BrennstofChem., 36: 1-11. HUCK,G. und KARWEIL, J., 1962.Probleme und Ergebnisse der kiinstlichen Inkohlung im Bereich der Steinkohlen. Fortschr. Geol. Rheinland Westfalen, 3(2): 717-724. HUCK,G. und PATTEISKY, K., 1964.Inkohlungsreaktionenunter Druck. Fortschr. Geol. Rheinland Westfalen, 12:551-558. INTERNATIONAL COMMITTEE FOR COALP E T R O ~ Y1963. , International Glossary of Coal Petrology, 2 ed. Centre National de la Recherche Scientiiique, Paris. J ~ ~ T G EH. N ,und KARWFZL, J., 1962. Kiinstliche Inkohlung von Steinkohlen. Freiberger ForSChWSh., A, 229: 27-36. KAYSER,H., 1952. Die Veredlung der Braunkoble und der geologisch jiingeren Brennstoffe. In: K.WINNACKER und E. WEINQAERTNER. Chemische Technologie.ZZZ. Organische Technologie. Hanser, Miinchen, pp.123-124. KARWEIL,J., 1956. Die Metamorphose der Kohlen vom Standpunkt der physikalischen Chemie. Z. Deut. Geol. Gar., 107: 132-139. KING, J. G. and WILKINS,B. T., 1944. The internal structure of coals. Proc. Conf. Ultrafine Struct. Coals Cokes, London, 1944, pp.46-57. KURBATOV, J. M., 1963.Zur Frage iiber die Genesis des Torfes und der Torfhuminsauren.Intern. Peat Congr., Lmingrad, 1963. Preprint, 8 pp. KUYL,0. S. and PATUN,R. J. H., 1961.Coalification in relation to depth of burial and geothermic gradient. Congr. Avan. &tEtudes Stratigraph. Gdol. Carbonifpre, Compte Rendu, 4, Heerlen, 1958,2: 357-365. LENSCH,G., 1963. Die Metamorphose der Kohle in der Bohrung Mbsterland 1 auf Grund des optischen Reflexionsvermagens der Vitrinite. In: Die Aufschlussbohrung Miinsterland I-Fortschr. Geol. Rheinland Wesrfaen, 11 : 197-203. LICHTENBERG, K., 1963. Die Geschwindigkeitsmessungenin der Bohruug Miinsterland 1. In: Die Aufschlussbohrung Miitasterland I-Fortschr. Geol. Rheinland Westfalen, 11 : 387-388. MONOMAKHOFF, C.,1961.La tectonique tangentielle dans les bassins houillers de la France et sa r6percussion sur la continuitt et le comportement de ces gisements. Congr. Avan. Etudes Stratigraph. Gdol. Carbonif&re,Compte Rendu, 4, Heerlen, 19S%,2: 423435. PATIELSKY, K., 1950.Die Entstehung des Grubengases. Bergbau Arch., 11/12:5-24. PATT~KY K. , und TEIC~&LLER, M., 1960. Inkohlungs-Verlauf, Inkohlungs-Masstabe und Klassiiikation der Kohlen auf Grund von Vitrit-Analysen. Brennstof-Chem., 41: 79-84, 97-104,133-137. F’ETRASCHECK,W.E., 1954. Zur optischen Regelung tektonisch beanspruchter Kohlen. Mineral. Petrog. Mitt., 4(1954): 232-239. PIQULEWSRAJA, L. V. und RAICOW~, V. E.,1963. Die hderung der chemischen Zusammensetzung einiger Torfarten in Abhhgigkeit von ihrem Alter. I. Alter und hderung der Torfkomponenten. Xr. Znst. To& Akad, Nauk,Beloruss. S.S.R., 1963,6:12-31. RAKOWS~, W.,BATURO,W. und R Q ~ W S W JL.A , 1963. , Die Humusbrennstoffe und ihre Bildung. Intern. Peat Congr., Leningrad, 1963. Preprint,35 pp. SCH”DER, S., 1963. Chemische und stratigraphische Untersuchungem an Hochmoorprofilen von Nordwestdeutschhd. Intern. Peat Congr., Leningrad, 1963.Preprint, 22 pp. SCHBRP,A., 1963. Die Petrographie der palitomishen Sandsteine in der Bohrung Miinsterlanu

DIAGENESIS OF COAL

413

1 und ihre Diagenese in Abhhgigkeit von der Teufe. In: Die Aufschlussbohrung Miinsterland I-Fortschr. Geol. Rheinland Westfalen, 11: 251-282. ScHijLLBR, A., 1961. Die Druck-, Temperatur- und Energiefelder der Metamorphose. Neues Jahrb. Mineral., Abhl., 96: 250-290. STACH, E., 1958. Der Inkohlungssprung im Ruhrkarbon. BrennSroff-Chem., 34: 353-355. STACH, H., 1948. Experimentelle Beitrage zur Frage der Brikettierbarkeit von Weich- und Hartbraunkohlen und der Quellung und des Zerfalls von Braunkohlenbriketts. Braunkohle, 1: 3 5 4 .

STADLER,G., 1963. Die Petrographie und Diagenese der oberkarbonischen Tonsteine in der Bohrung Miinsterland 1. In: Die Aufschlussbohrung Miinsterland I-Fortschr. Geol. Rheinland Wesrfaen, 11: 283-292. STOPES, M. C., 1919. On the four visible ingredients in banded bituminous coal. Studies on the compositions of coal. Proc. Roy. SOC.(London), Ser. B, 190: 470-487. TEIcHM~~LLER,M., 1950. Zum petrographischen Aufbau und Werdegang der Weichbraunkohle (mit Beriicksichtigung genetkcher Fragen der Steinkohlenpetrographie). Geol. Jahrb., 64: 429488. TEIC~&LLEX, M., 1962. Die Genw der Kohle. Congr. AVM.l h d e s Stratigraph. G b l . Carbonif&e, Compte Rendu, 4, Heerlen, 1958, 3: 699-722. TEICEIM~~LER, M., 1963. Die Kohlenflaze der Bohrung Miinsterland 1. In: Die Aufschlussbohrung Miinsterland I-Fortschr. Geol. Rheinland Westfalen, 11: 129-1 77. TEI&LLER, M. und Tni,R.,1949. Inkohlungsfragen im Ruhrkarbon. Z. Deut. Geol. Ges., 99: 40-77. TEIC&LLBR,M. und TEICHMULLER,R., 1950. Das Inkohlungsbilddes niederskbsischenWealdenbeckens. Z. Deut. Geol. Ges., 100: 498-517. TEICHMULLER, M. und TEICEIM~LLER, R.,1954. Die stomche und strukturelle Metamorphose der Kohle. Geol. Rundschau, 42: 265-296. TEIC-LER, R., 1952. Zur Metamorphose der Kohle. Congr. Avan. l?tudes Stratigraph. Gkol. Carbon@+e, Compte Rendu, 3, Heerlen, 1951, 2: 615-623. T E I C X MR., ~ ~1962. , Zusammenfassende Bemerkungen uber die Diagenese des Ruhrkarbons und ihre Ursachen. Fortschr. Geol. RheinlPnd Westfalen, 3 (2): 725-734. T h , W.,1963. Auswertung der Bohrlochmessungen der Bohrung Miinsterland 1 unterhalb 1870 m Teufe. Gesteinsphysikalische Daten und Sandsteine. In: Die Aufschlussbohrung Mlinsterland I-Fortschr. Geol. R h e i n l d Wesrfalen,11: 239-242. VAN DER MOLEN,W. H. und S m , H., 1962. Die Sackung in einem Moorgebiet in NordGriechenland. Intern. Torf Kongress, Bremen, 1962, Ber., 1lpp. VANKREVELEN, D. W., 1953. Physikalische Eigenschaften und chemische Struktur der Steinkohle. Brennstoff-Chem.,34: 167-182. VAN KRBVBLEN, D. W., 1961. Coal. Elsevier, Amsterdam, 514 pp. WELTE, E., 1952. ober die Entstehung von Huminskuen und Wege ihrer Reindarstellung. 2. PflanzenerniuU., Diingung, Bdnkunde, 56 (101): 105-139. W n ; u ~ ~E., s , 1953. Anisotropy of the vitrain of South Wales coals. Fuel, 32: 89-99. V. und WEWC, L., 1911. Der EinfluB des Vertorfungsprozesses auf die Zusammensetzung von Carextorf. Z. Moorkultu Torfverwertung,9: 153-168.

APPENDIX I PETROGRAPHICAL COAL CLASSIFICATION

STOP= (1919) classified coal into four lithotypes which may be distinguished megascopically:vitrain, clarain, durain and fusain. Each is composed of microscopic entities termed maceruls such as collinite, telinite, sporinite, resinite, fusinite and micrinite. Macerals which are similar in their optical and chemical behavior

414

M. TmCHM8LLER AND R. TErCHMhLER

TABLE IV CLASSIFICAlTON OF COAL, LITHOTYPES, MICROLmOTYPES AND MACERALS

(After the hTJlR"TERNAnNAL Maceral

Collinite Telinite

COMMI'ITEE FOR COAL, PFXROLOOY,

Maceral-group

Vitrinite

Microlithotype (and its principal maceral-group constituents)

Lithatype

Vitrite (vitrinite)

Vitrain

Vitrinertite (vitrinite

+ inertinite)

+ liptinite) Duroclarite (vitrinite + liptinite + inertinite) Clarite (vitrinite

Cutinite Resinite

1963)

Liptinite (= Exinite)

Clarain

Sporinite Alginite

Clarodurite (inertinite

Micrinite Semifusinite

Durite (inertinite

Fusinite Sclerotinite

+ liptinite + vitrinite)

Durain

+ liptinite)

Inertinite Fusite (inertinite, except micrinite)

Fusain

are divided into the three maceral groups; vitrinite, liptinite (= exinite) and inertinite. Under the microscope it is possible to distinguish layers of characteristic maceral associations which are called microlithotypes (vitrite, clarite, durite, fusite, etc.). Table IV shows the classification of the Stopes-Heerlen Nomenclature according to the INTERNATIONAL COMMITTEE FOR COAL PETROLOGY(1963).

GLOSSARY

Anthraxylon: humic matter of coals, largely derived from leaves, roots, stems and bark (lignin and cellulose). Because of impregnation with colloidal humic matter, the cell structures are more or less flattened by means of the overburden pressure. Megascopically bright, glossy, or jet-like bands or lenses (vitrain). Attritus: component of coal consisting of finely divided organic matter. According to the classification of Thiessen Bureau of Mines divided into translucent attritus (humic matter, protobitumina) and opaque attritus (fusinite, micrinite). Megascopically dull.

DIAGENESIS OF

COAL

415

Bone coal: argillaceouscoal or carbonaceous shale. Cannef coals: the product of lithification of sapropelic (bituminous) organic muds, deposited in still-water lakes and ponds. Coals composed of a very fine attritus (with anthraxylon largely absent) with a pronounced microbedding, may be differentiatedas (1) cannel coals s. str. (composed of spores and similar organic materials, vitrinite and inertinite); (2) humic or pseudocannels (largely composed of vitrinite); and (3) boghead cannels (fatty Algae are the characteristic ingredient). Clarite: microlithotype ( m a d association) of banded hard coals, consisting of vitriniteand exinite. Durite: microlithotype (maceral association) of banded hard coals, consisting of inertinite and exinite. Exinite (= lipfinite): group of coal macerals, comprising the sporinite (spore exines), cutinite (surface layer of leaves), resinite (resins and waxes), and alginite (Algae). Characterized by a relatively high hydrogen content and a low reflectance in polished sections. Fusinite: rnaceral of coal, relatively high in carbon content and in reflectance. It has no coking power and has lower contents of volatile matter, oxygen and hydrogen than all other components of coal. Derived from wood or other plant tissues which are highly carbonized. (Fusain = fossil charcoal, mineral charcoal or mother-of-coal = lithotypc of coal. which is megascopically visible.) Gyttja: black organic mud deposited in poorly aerat lakes and ponds. Organic matter is more or less determinable. Inertinite: group of coal macerals, including fusinite, semifusinite, micrinite and sclerotinite (fungal remains). Characterized by a relatively high carbon content and a high reflectance in polished sections. More or less inert during coking processes. Maceral: smallest micropetrographicunit of coal. Recognized by virtue of physical and chemical similarity. Micrinite: maceral of coal, relatively high in carbon content and in redectance; without cell structure. Derived from highly decomposed plant matter. Phyteral: coal particles of microscopic size based on recognition of plant fossil parts and pieces (for instance, spore, cuticle, bark, etc.). Protobitmen: see exinite. Vitrinite: major maceral of most hard coals, relatively high in oxygen. Coalified humic matter, largely derived from lignin and cellulose of the plants.

Chapter 9 DIAGENESIS I N SEDIMENTARY MINERAL DEPOSITS1 G. C. AMSTUTZ AND L. BUBENICEK

Mineralogisch-Petrographisches Institut der Universitat Heidelberg, Heidelberg(Germany) Institut de Recherches de la Sidkrurgie, Station d’Essais, MaiziPres-16s-Metz, Moselle(France)

SUMMARY

The role of diagenetic processes in the formation of mineral deposits is reviewed and their main traits are summarized. Diagenetic features are important criteria for the determination of deposits of sedimentary or exhalative sedimentary origin discussed in this chapter. In thefirst section the problem is approached from a historical angle and an answer is sought for the absence of considerations of diagenesis as a factor in ore genesis in “economic geology” textbooks until 1963. The second section shows that the concept of diagenesis (and the knowledge of diagenetic processes) has been very well developed before finding its way into the literature on mineral deposits. Again, a historical reason is sought. The keen interest of sedimentologists and stratigraphers in diagenesis since about 1868 is stressed. (The present volume is almost a “centennial of diagenesis”). The third section contains an outline of geometric evidence for diagenesis in mineral deposits. Various typical examples are briefly described and a number of them are documented by figures. This section is considered to be the “pitce de rdsistance” because the geometric evidence is the most direct and basic evidence. It is also the most powerful tool of the exploration geologist. Geochemicalevidence, which is also very important, actually consists of an indirect geometric study, i.e., the distribution patterns of elements in space (through abundance curves, phase diagrams, etc.). Basically, genetic studies are always investigations of isomorphism or symmetry of geological bodies. The fourth section refers to the essential processes of a geochemicalnature, pointing out that they are no different from those existing in common rocks. A brief literature review is given. Thefifth section lists a number of important types of mineral deposits, and for each type, a number of pertinent publications on problems or observations on diagenetic processes. 1 The first part of this chapter is presented by Prof. G. C. Amstutz, the second part on the role of diagenesis in the formation of oolitic iron-ore deposits was written by Dr. L. Bubenicek.

418

G . C. AMSTUTZ AND L. BUBENICEK

The sixth and last section is a case study, made by the second author, of diagenetic processes involved in the formation of oolitic iron ore deposits.

HISTORICAL REVIEW

The history of the recognition of diagenesis as a process in ore mineral formation reflects the trend of ore-genesis theories as a whole. These theories essentially moved back and forth between two extreme patterns of thought relating to time and space of rock and ore formation. The basic patterns involved are shown in Table I and are naturally present in any field of human endeavor: the arts, the sciences, and the religions. Table 11 is a schematic summary of the prevailing theories on ore and rock genesis. It is easy to see that the trends run parallel to the general historical trends of human culture. The prevailing tendencies during syn-endo-periods meant an urge to first look for causesfrom within, before attempting to explain observations by assumption of forces acting from the outside. During epi-exo-periods it was fashionable to speak in terms of injection or invasion and replacement by fluids, often emanating from unknown sources at depth, or at least by fluids moving around after the time of formation of the enclosing rock, instead of first concentrating on a geometric and geochemical study of the host rock involved. Naturally, the prevailing pattern of thought during epi-exo-periods kept the geologist from becoming interested deeply in the time of formation of sediments. TABLE I BASIC PATTERNS OF GENETIC INTERPRETATIONS OF ROCKS AND ORES

Time

Space

syngenetic formation

endogenousformation

contemporaneous with the enclosing rock =

epigenetic formation

= origin same as, or from within the host rock

exogenous formation

formation later than that of the host lock =

possible combinations: syn-endo syn-exo epikndo epikxo

= ore matter originates from without the host rock

419

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

TABLE II HISTORICAL EVOLUTION OF ORE-GENESIS THEORIES

Emphasis on syn - endo or at least on syn-exo

Historical periods

L

Albertus Magnus (1 193/ 1205-1280) Congenerationists (about 1300-1500) Agricola (epigeneticists) (1494-i 555)

Ettphasis

OII

" p i - exo or at

least on epi-endo I

Old Greek, Roman, Arabic, lndian and Chinese theories

1

Balanced theories with observational classi@ations

various theories displaying all basic pitterns of thought

!

I

I

earth

Hg-

metals

I

everything is congenetic -

and

+S

ore minerals I

, II I

\

fire is main agent

(about 1830-1890) Neo-epigeneticists (Poxepnf, Lindgren in part)

I

1960: Detailed space- time differentiations and determinations (symmetry, i.e., fabric studies, esp. in sediments, of decisive importance)l

+

fire water main agents

classifications main basis for genetit interpretation:

Geometric and geochemical distribution patterns (primary fabrics and compositional histograms) inherent to the host rock are first compared before causes from without are assumed.

Consequently, discussions referring to the period of deposition, crystallization, compaction and induration were rare during epi-exo-periods and more common during syn-endo-periods. Of course, detailed descriptions of depositional and compaction textures do not abound in the literature before publications of VON GUMBEL (1868, 1883,1888), WALTHER (1894), and HEIM(1917), but may be found in a number of detailed descriptions of fossil localities and in accounts on litholog1930, ic sequences in early works, even before Werner. Lye11 (cf. HAARMANN, fig.67, p.156) pictured changes taking place through creep under the load of overlying sediments. It should be pointed out that the evolution of theories on the formation processes of mineral deposits has been extremely slow and may be regarded to

420

G. C. AMSTUTZ AND L. BUBENICEK A

Fig.1. Schematic representation of ore genesis theories according to the conventional and the new patterns of thought. A. This figure shows the domination of the myth of epigenetic replacement and of the unknown, depth (“deep-seated sources”). Epigenetic introduction from the outside is an axiomatic condition for the formation of most ore deposits. This pattern corresponds essentially to the creationistic pre-Darwinian beliefs in paleontology. B. This figure pictures the pattern of ore genesis theories according to the new “petrographic” or integrated theory, according to which ore deposits normally formed at the same time and essentially within or very near the observed host rock. Just as man and animals in the evolution theory of paleontology, ore deposits are, in the new theory, considered a normal integral part of rock evolution. I = igneous intrusive rocks (known!); I1 = igneous extrusive or subvolcanic rocks (known); 111 = metamorphic igneous rocks or migmatites; I V = metamorphic sedimentary rocks; V = sedimentary rocks (non-, or partly metamorphic); VI = introduction from the (unknown) outside source is assumed; VII = some migration probable, possible, or (?) questionable. List of major types of ore deposits for which a syn-endo as well as an epi-exo origin has been proposed. In sediments and volcanic rocks: I = Arkansas type barite deposits; 2 = Mississippi Valley type deposits (including the barite and fluorspar deposits in the same type of sediments); 3 = Rammelsberg and similar deposits; 4 = magnesite, rhodochrosite and siderite deposits of the Alps and elsewhere; 5 = Kupferschiefer and/or Red Bed copper deposits as well as various disseminated to massive copper-lead-zinc deposits, for example of the Kuroko type; 6 = Blind River, Witwatersrand and similar deposits; 7 = propylitic deposits of copper, gold, and othzr metals; 7a and 76 = deposits of sulfides, oxides and native elements (Cu, Ag, Au) in or near volcanic rocks (often with spilitic phases); 8 = Mina Ragra type vanadium deposits; 9 = Colorado Plateau or “sandstone type” uranium deposits; 10 = iron deposits of the Lake Superior type; I1 = Ducktown, Broken Hill, Outukumpu, Falun, and similar deposits in metamorphic belts. In and adjacent to igneous rocks: a = porphyry copper deposits in and around intrusions(inc1uding the Climax molybdenum deposit); 6 = Granite Mt., Utah, deposits of magnetite, and similar deposits; c = tin deposits in and around intrusions; d = contact deposits, pipe deposits, perimagmatic vein deposits; e = chromite deposits;f= pegmatites.

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

42 1

€ig.l (continued). Legend see p.420.

lag behind that of zoology, paleontology, crystallography, and botany by a time period of about 100 years. The most interesting reasons for this delay have been analyzed on various occasions (AMSTUTZ,1959a, 1960, 1963a). In the frame of the present condensed discussion only two additional basic historical facets can be reported. First, it should be noted that an excellent start in the direction of an observational foundation of the science of mineral deposits was made in the second half of the last century, when VON GRODDECK (1879) published his book, entitled Die Lehre von den Lagerstatten der Erze. Von Groddeck’s System der Erzlagerstatten considers the “geognostic” or observational parameters first and proceeds from there to the suggestion of a “geo-genetic” theory. Von Groddeck‘s classi$cation is given in Table 111. This observational approach, which is indeed the sine qua non of any scientific procedure, was largely neglected. Even in some “modern” textbooks (e.g., BATEMAN,1950) geognostic descriptions are subordinate to genetic headings. Consequently, the description of a deposit has to be in agreement with a preconceived theory on its genetic history. This explains why in the conventional textbooks the terms diagenesis, compaction, diagenetic recrystallization, etc. cannot be found. It also explains why the interest in the enclosing rock was practically zero in studying many layered deposits, for example, most Mississippi Valley area deposits, the barite deposits of Arkansas, and many other similar deposits. To the

422

G . C . AMSTUTZ AND L. BUBENICEK

TABLE 111 CLASSIFICATION OF ORE DEPOSITS

(After VONGRODDECK, 1879, p.84)

A . Bedrock deposits I. Layered deposits 1 . massive ore strata 2. coprecipitation ore strata (of disseminated ore matter) 3. lenticular ore layers (or strings) 11. Massive (non-layered) deposits 111. Cavity fillings 1 . fissure fillings or dikes

a. dikes in massive rocks

b. dikes in layered rocks 2. fillings of caves

IV. Metamorphic mineral deposits

B. Weathering deposits (detrital deposits)

proponents of a panepigenetic trend which took a strong hold on ore-genesis theories from about 1890to 1940or 1950,it was “beyond any doubt” (these words were used) that the layered disseminated deposits were formed by replacement long after the formation of the sediments. Many of Von Groddeck‘s descriptions of layered deposits contain reference to depositional features (“Lagerung”) formed during diagenesis, although the terms diagenesis and compaction were not used as such (to the knowledge of this author). Concretions, for example, are often described as having formed during the deposition of the sediments, for instance in the “Knottensandsteine” of Commern (Eifel; Bleiberg between Call and Mechernich). The galena in them is considered to represent the original cement and to have crystallized, consequently, during diagenesis, together with the cementing quartz. This assumption is supported in many samples by modern microscopic work. Von Groddeck’s observational approach which was based on actual properties of the rock and not on genetic assumptions, was largely neglected through the influence of works of VONCOTTA(1870) and PoSEPN~~ (1902), although the latter was somewhere in between the two schools of thought of syn-endo- and epi-exopatterns. In the Americas, and in part also in originally nomadic parts of Asia, the theories of PoSepnL prevailed for ethnological reasons to a large extent until today, as pointed out in more detail by AMSTUTZ (1959,1963a). Historically, the departure from the observational foundations laid down by

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

423

VON GRODDECK (1879) is a reversion to older, less conscious and less integrated patterns of thought, probably caused by ethnological factors. The exaggeration of exo-epi-patterns by some recent workers, led to the situation pictured in Fig. 1. The majority of the ore deposits are considered to have been brought in from the outside, mainly from largely unknown igneous sources a t depth or from unknown “source beds” of regional extent. Consequently, petrographic studies of sediments for other purposes than for the determination of epigenetic hydrothermal alterations were, and still are, considered to be futile and were not carried out. This extreme tendency was also supported or accompanied by the segregation of so-called “economic geology” from other fields of geology; and the unfortunate division into “hard-rock” and “soft-rock” geology has contributed to the “schism”. The geology of mineral deposits was considered to be reserved for “hard-rock geologists” and “soft-rock’’ interests were taboo. Such tendencies are in direct contradiction to recommendations of integrated approaches like that of VON GRODDECK (1879, p.2). During the past 10-20 years, a major change has taken place in the trend of thought on ore genesis. The pattern of this new approach is pictured in Fig.lB. Although the objective approach initiated by some during the 19th century was never completely abandoned, it was not accepted by a large proportion of “economic geologists” until recently. It is most interesting to note that the trend changed 1959a). Again, contemporaneously in many countries (as described by AMSTUTZ, ethnological factors appear to have been, and still are, controlling this change, very similar to the change which took place during the acceptance of evolution. The present change to syn-endo-patterns is indeed basically the same as the acceptance of evolution as a major theory, as pointed out by AMSTUTZ (1959a, 1964a, b). The fact that only rather recently diagenetic processes became accepted as being of importance in ore genesis can be understood only within the historical-ethnological context, and this called for a brief historical review. To sum up, diagenesis was not recognized as an important factor in ore genesis until very recently, because the branch of geology often termed “economic geology” has in many ways lagged far behind the other fields of the earth sciences. As a matter of fact, at a time when physics has matured to the point of losing its old narrow image and of being able to integrate both “objectivity and intersubjectivity” (MATSON, 1963), the field of geology of mineral deposits is still struggling toward objectivity. But perhaps it is possible to leap ahead fast and to catch up withphysics. This goal can only be achieved if most “economic geologists” are willing to accept the criticism and to recognize their anachronistic place in the history of science. But, as Barzun(in: SIMPSON, 1964)pointed out: “the scientific culture lacks an equivalent of the criticism so characteristic of humanistic culture”. Wherever this is the case it will be difficult to move ahead because the new approaches will be repressed as undesirable and uncomfortable,

424

G . C. AMSTUTZ AND L. BUBENICEK

MODERN APPROACHES

The “schism” between the field of “economic geology” and most other fields of geology is illustrated, as briefly mentioned in the first section of this chapter, by the following observation: the term diagenesis is not found in books on mineral deposits until 1944 when SCHNEIDERH~HN introduced it in the first edition of his Erzlagerstiitten (p.208). Even in his fourth edition (1962) he only briefly mentioned that “descendent and diagenetic changes occur in deeper zones below the oxidation and cementation zone”. The second and only other reference made to diagenesis is found in the section on coal (SCHNEIDERH~HN, 1962, p.274). To the knowledge of this author ROU~HIERS’ textbook (1963) is the first one to list and to adequatcly cover the role of diagenesis in ore formation. He refers to diagenetic processes on about 20 pages. Of the following standard textbooks in use during the past 50 years, none lists diagenesis in the subject index: LINDGREN (1933), MCKINSTRY(1949), RAGUIN (1949), BATEMAN (1950), M~GNUSSON (1953), BATES(1960), and PETRASCHECK (1961). Quite in contrast to this record displayed by books on economic geology is the very early and extensive coverage found in books on general geology, stratigraphy, lithology, and sedimentary yerrology. As a matter of fact, WALTHER (1894) dedicates a chapter of nineteen pages (WALTHER, 1894, pp.693-711) to the topic of diagenesis alone and offers extensive discussions on approximately 30 other pages of his Einleitung in die Geologie a h historische Wissenschaft. This was only 25 years after VON GUMBEL (1868) had first introduced the term, and 5 years after his large textbook had been published. According to GRABAU (1924, p.748) and FISCHER (1961, p.lll), the term diagenesis was introduced first in 1868, whereas WALTHER (1894, p.693), JOHANNSEN (1939, p. 173), and MURAWSKI (1963) give the date of 1888. It appears that VON GUMBEL (1868, p.383) has used it first in his Geognosrische Beschreibung des Ostbayrischen Grenzgebirges, which appeared in 1868 and not in 1888; the latter is the date of publication of his textbook Grundziige der Geologie. Von Giimbel defined the term in exactly the same way as it is used today, but was still of the opinion that the “crystalline schists” were also directlyprecipitated from waters like the normal sediments (VON GUMBEL, 1868, pp.57, 334), and that their compaction was a result of high-temperature diagenesis (VONG U M BEL, 1868, pp.381, 492, 1056-1058) which led to the present rock. He did not yet accept the modern idea of gradual metamorphism but did discuss it at length on p.1056. Some books, e.g., GRABAU (1924, p.748) and FISCHER (1961, p.lll), report that Von Giimbel applied the term diagenesis to embrace metamorphism as well. This is a misinterpretation as can be seen readily from the original text (VON GUMBEL, 1868, pp.57,334,381,492, 1056-1058). As a matter offact he distinguished clearly between depositional or original transformations, which he called diagenetic on the one hand, and later changes, which he included in metamorphism on the

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

425

other. Indeed, diagenesis seems to be one of the few geologic terms which has not undergone any significant change of meaning since its introduction. A careful description of diagenetic processes is included in all major textbooks (1 894). In 191 1 and 19 1 5, ANDREE devoted on general geology ever since WALTHER two extensive papers to discussions of sedimentation and diagenesis. BEHREND and BERG(1927, pp. 514-520) and CORRENS (1939) gave comprehensive summaries of diagenetic processes. Since about 1940, a major portion of the sedimentological literature deals with diagenetic processes, and a wealth of information on mineral genesis is available, which can be summarized only in part, as follows: ( I ) Fabric studies relevant to mineral deposits are briefly summarized. ( 2 ) Geochemical or compositional processes affecting the nature of sedimentary deposits are examined in cross-section. (3) A reference list of various examples of diagenetic processes in ore deposits is presented, and finally, a case history of diagenetic processes in oolitic iron ores is given by Dr. Bubenicek.

FABRIC CHANGES DURING DIAGENESIS OF SEDIMENTARY MINERAL DEPOSITS

By way of introduction it may be pointed out that, naturally, fabric and compositional changes are intimately related, and that the subdivision offered in this section and the section on geochemical studies is, therefore, largely artificial. A classijcation of the different fabric or geometric changes during diagenesis of mineral deposits may be presented as given below. ( I ) Individual grains. ( a ) Grain and pore size. ( 6 ) Grain and pore shape. (c) Grain and pore orientation. ( 2 ) Grain aggregates (“Korn-Verband”). ( a ) Aggregate-net unit(s) size (grain-group or pore-group size). (b) Aggregate-group shape (or pore-group shape). (c) Aggregate-group orientation (pore-group orientation). Grain and pore size in many deposits changed first of all and mainly during diagenetic crystallization. The grain orientation is influenced by the forces active during compaction. The rate and time of these changes are a function of various factors, e.g., the concentrations of the crystallizing material, the pH-Eh relations, the salinity and other properties of thc pore solution, etc. Accordingly, thecrystallization process leads to different generations of grains, and the time position of a phase in the system determines the geometric pattern to a large degree. The time of diagenetic crystallization of ore minerals is variable not only for different ore minerals, but also for one and the same mineral. Pyrite is perhaps

426

G. C. AMSTUTZ AND L. BUBENICEK

the best illustration of an extreme case because it can be of early and of late origin within the same rock, as illustrated by AMSTUTZ et al. (1964). Most sulfides, however, are diagenetically late, a fact which has caused a good deal of the confusion with “hydrothermal alterations”. The aggregate patterns of rock textures resulting from diagenetic crystallization of ore minerals are basically the same as those of common minerals. Some unusual textures, however, may result, which in many cases led to the assumption of epigenetic influences from outside sources again. Recent work has shown that a diagenetic interpretation of many of these ore structures requires only about half or one-third of the assumptions necessary for a hypogene replacement origin; but, in addition, a number of primary sedimentary textures make any other interpretation extremely far-fetched. Descriptions are given in the captions of the frontispiece of this book and of many figures. The aggregate development of common sediments, as well as ore minerals, during diagenesis may proceed in essentially three main directions. These are illus-

2

I

10

.A

4

5

6

7

9

8

Fig.2. Geometric classification of rock fabrics, free of genetic connotations. Patterns I, 2, 3, and 10 = stromatiticorlinearfabrics; pattern 4,5, and 6 = merismitic or network fabrics; pattern 7,8, and 9 = ophtalmitic or disseminated fabrics; pattern I 1 = massive or homogeneous fabric (to be placed at the fourth corner of a tetrahedron of which this drawing shows one face only). (This purely geometric classification and nomenclature of rock fabrics is a systematic modification 1948,in Rocks andMineraZ Deposits, for chorismatic, polyschematic of patterns pictured by NIGGLI, GEOLOGICAL INSTITUTE, rocks and mineral deposits; and from Data Sheet 21 of the AMERICAN 1960. Additional adjectives may be used in order to designate transitional patterns: pattern 3 may be called phlebitic stromatite; pattern 4a, phlebitic merismite; pattern 8b, miarolithic ophtal1959b.) mite; and pattern 10, nebulitic stromatite, After AMSTUTZ,

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

427

trated by Fig.2. The diagenetic differentiation may produce either a layered or stromatite pattern, a merismitic network or an aggregated or ophtalmite pattern. When different scales are considered, various patterns may combine, one pattern being the smaller scale subpattern of another pattern. Mineral deposits exhibit particularly often patterns 4-9 within pattern 1. Also, quite often pattern I or 2 contain one or more individual layers. These textural relationships have been known to sedimentologists for a long time. Only very recently, however, has this knowledge started to be applied to ore deposits. Before, the ore solutions were, and in some cases today still are, assumed to have produced perfect epigenetic replacement textures. On analyzing these theories one can recognize that a great number of “tacit” assumptions are made in explaining any late introduction of “ore fluids” from unknown depth or from some sort of a source bed. A study of known facts on diagenetic processes in common sediments shows quite clearly that the same criteria also apply to other material considered rare because of shortcomings of human technology. The material value of a mineral subconsciously influences the genetic interpretation attached to it by the human mind. The assumption of epigenetic hydrothermal replacement has been made solely on the basis of a superficial mineralogical similarity with deposits in intrusives, and a large-scale resemblance with the zoning found occasionalIy in large layered deposits. Zoning is, however, entirely normal for deposits of sedimentary origin. The identity of ore mineral patterns with common diagenetic aggregate fabrics can be tested by detailed studies of size, shape and orientation of both the ore minerals and the gangue. The comparison of the results for both affords a sort of symmetry or congruency comparison. If the agreement is good, this points to the same or a similar origin, e.g., the formation of both groups of minerals, valuable and useless, by the same or a similar process. The orientation study may be done in great detail on using SANDER’S (1936, 1948-1950) petrofabric methods. SCHACHNER-KORN (1954), VON GEHLEN (1960), SIEMES (1961), and others have applied X-ray goniometer methods to the study of ore minerals, mainly in metamorphic zones. The same methods ought to be applied more to non-metamorphic material in order to establish the effects of diagenesis. Among the many typical geometric features of diagenetic fabrics, the paragenetic sequence patterns (includiag diagenetic replacement patterns), the density-gravity features and the motion-disruption features are perhaps the most useful in the study of theediagenesis of mineral deposits. These classes of criteria are considered here briefly. Paragenetic sequences

The various generations of diagenetic crystallizatio~~ can be recognized from simple

428

G. C. AMSTUTZ AND L. BUBENICEK

Fig.3. Fredonia Limestone, Southern Illinois Fluorspar district. Diagenetic generations which can be seen here are as pictured in Fig.6. Sphalerite (black) occurs within bituminous rings of oolites and is later accumulated as stylolitic recrystallization residue. Average ovoid is 0.5 mm in diameter.

Fig.5. Stylolite zone with residual sphalerite in the same oolitic rock as pictured in Fig.3 and 4. Stylolite formation is seen from ( I ) the partial elimination by solution of oolites, (2) the formation of coarsely recrystallized calcite differing in size, shape and orientation from all previous calcite generations, (3)the accumulation of black bituminous material, ( 4 ) the accumulation of residual sphalerite(white)with thesame size and shape as that of the oolitic ring (note also the ring structure in the sphalerite fragments), and ( 5 ) the formation of calcite twins within the stress range of the stylolite plane; x 50.

Fig.4. Details from the oolitic limestone pictured on Fig.3. The diagenetic generations of the paragenetic chart in Fig.6 are clearly seen. Sphalerite (white) shows (1)confinement to the bituminous rings of the oolites, (2) solution pitting by the corona-calcite, and (3) complete absence in the cemeniingcalcite, which is the last generation, despite the presence of tar seams. These observations date the sphalerite formation as being contemporaneous or immediately subsequent to the formation of the oolites. Diameter of the large oolite is 8 5 0 , ~ ;x 60.

Legend see p.428.

430

G. C. AMSTUTZ AND L. BUBBNICEK

geometric superposition patterns (“Anlagerungstexturen”), exactly in the same way as in igneous rocks. Recrystallization processes can of course obliterate the primary fabric, but more often parts of it are still recognized, as pointed out by many authors (e.g., SCHNEIDER, 1964). In Fig.3, recrystallization along stylolitic seams and along a vertical gangue veinlet has destroyed the carbonate oolites, but sphalerite and quartz are left over as an “insoluble residue”. In the same sections (Fig.4 and 5) the sphalerite is seen to have been emplaced in the bituminous rings of the oolites, where it forms idiomorphic faces toward the inside but follows the rim of the oolites on the outside. Also, the sphalerite is corroded on many outside borders by second generation carbonate, the clear corona of calcite as seen under higher magnification (Fig.4). The paragenetic sequence deduced from many identical samples occurring widely apart (several kilometers) is pictured in Fig.6. A number of cores of the oolites consist of fossil fragments. The age of quartz varies and its third and last generation is accumulated in stylolite seams as prisms oriented parallel to the seams. The diagenetic formation of quartz in rocks of this type was discussed by FUCHTBAUER (1961) and GRIMM(1962). Pyrite shows an almost homogeneous distribution throughout the rock, but is somewhat more abundant within sphalerite, where some of the dot-like inclusions may also consist of chalcopyrite. A late generation of pyrite also occurs within stylolite seams.

1 s t diagenctic etage (deposit~onal)

2 nd divgrtielic stage (carly burial )

3 r d diagenvtic stage (pve-metamorphic)

I

I

I

C l r a i carbonate core

I

SiO (quartz.quartzite)

Bituminous o6litic rings

-- c

Sphalente (2”s) Clear corona c a l c ~ t e Clear cement calcites Pyrite (FeSz) Galena (PbS) Fluorite (CaF2)

I

I I

-- -

Stylolite formation

Cornpactlo”

- - ?? , - < - -

Fig.6. Diagenetic crystallization sequence (paragenesis)in oolitic limestone of the Southern Illinois Fluorspar district. Fluorite is mostly contemporaneous with the clear cement calcite.

DIAGENESIS IN SEDIMENTARY MlNERAL DEPOSITS

43 1

Fig.7. Typical diagenetic load cast from the Pb-Cu-Ni-Co-deposit of Fredericktown, Mo. The geopetal nature is clearly seen. The country rock-greywacke facies of the transition zone between the Upper Cambrian Lamotte Sandstone and the Bonneterre Formation.The sultides consist of marcasite, pyrite, siegenite, chalcopyrite, and galena; x 4. (After AMSTIJTZ,1963a; and EL BAZ, 1964.)

Fig.8. Diagrammatic cross-section of the Leduc Reef chain, showing the escape paths of connate water squeezed out of the surrounding compacting Ireton and Duvernay shales. (After ILLING, 1959, fig.4.) (Some of the heavy metals may have been extracted “chromatographically” from the hydrocarbons in the connate water, while moving through zones having different pH-Eh values and consequently of different bacterial content.)

432

G . C. AMSTUTZ AND L. BUBENICEK

Galena and fluorite are located mostly within the “cement” spaces as last generations, but may locally replace earlier generations. These neat paragenetic relations are pictured in Fig.6, where the time of stylolite formation is placed within the period of compaction for the many reasons (1964). outlined by PARK(1962) and PARKand AMSTUTZ In the Lead-Belt of Missouri sulfides pictured in Fig.7 and the frontispiece of this book, the paragenetic pattern is very consistent and, combined with evidence of the gravity-density features (discussed below), offers excellent criteria for a diagenetic crystallization of the sulfides. Concentric patterns with consistently the same sequence of crystallization, diagenetic fractures and breccia spaces filled with galena, which is the latest mineral in the nodules and clusters, demonstrate quite clearly the role of diagenetic processes. In a brief outline on the Mississippi Valley type of deposits, AMSTUTZ (1963a) showed contraction cracks along algal finger structures; in these cracks late diagenetic galena has collected, as well as in other later diagenetic loci. This makes an epigenetic replacement process quite impossible. If the diagenetic circulation paths of pore solutions in a reef structure are

F

syngenefic

epigenetic

I

Da Supergene o r e minerals deposited by

la Supergene: o r e minerals deposited contem-

supergene solution and replacement, I e., by epigenetic lateral solution, migration and Secretion caused by groundwater movements, the source of the o r e m a t t e r is the Same.or some adjacent sedimentary bed

poraneously i n and with the sedimepts;the o r e matter.Ba.Pb.Zn.Fe.Cu.Ni,Co,S.etc,Is of erosional and thus of supergene origin.

t

I I b Hypegene o r e rnin$al;

deposited contemporaneously in and w i t h the sediments. t h e ore matter is of exhalative-volcanic and thus of hypogene-hydrothermal origin

1

a b Hypogene o r e minerals deposltea by s o l u t i o n s a n d replacements In t h e course of upwards percolations of telernagmatlc hydrothermal o r regenerative hydrothermal fluids o r emanations. either along faults o r fractures.or through pore spaces.along grain boundaries. both t r o m unknown sources .7 7 a.t depth

I

Fig.9. The four basic theories on the genesis of the Mississippi Valley type deposits.Diagenetic fabrics point to a syngenetic mode of formation.

I

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

433

Fig. 10. Erosion channels which are filled with rhythmically precipitated coarse-grained fluorite (light layers) and a finer grained mixture of fluorite, clay minerals, carbonates with shells, and variable amounts of sphalerite. Diagenetic stylolite seams are seen to separate the individual layers and the two sets of layers. Base line = 120 cm. Southern Illinois fluorspar district. This photograph (1962; fig.l2b, p.523), who attributed all these features to a replacecorresponds to that of BRECKE ment mechanism. (Photo by G. C. Amstutz and W. Park.)

contemplated (Fig.8), then a good explanation for the high content of galena in porous flanks and crests of reefs has been found. The criteria for the presence of a diagenetic crystallization sequence of the sulfides, together with the gravity-density features, provide the strongest evidence for placing the Mississippi Valley deposits in the group l a or Zb (Fig.9). BERNARD (1958, 1964) described diagenetic parageneses in French deposits of the same type. For comparison, the four basic theories on the genesis of the Arkansas barite belt deposits are given in Fig.12. A most valuable and complete textural description of the similar Pb-Zn-Cu deposits in Triassic rocks of the northwestern part of the Balkan region was given recently by RENTZSCH (1963). Sedimentary ore rhythmites which are so-to-say paragenetic sequences drawn out in time are frequently observed in the coon tails of the southern Illinois Fluorspar deposits (Fig.10), and parts of the Rammelsberg ores (Fig. 11). Diagenetic paragenesis sequences are also displayed in barite nodules (Plate I), where the nodules represent early and the matrix late generations. The many new observations on these support theory l a or Ib of Fig.12. Valuable paragenetic evidence is also available from the fossil matter fre-

434

G . C . AMSTUTZ AND L. BUBENICEK

Fig.11. Polished-section photograph of a diagenetic load-cast fossil feature (gastropod) in Rammelsberg lead-zinc ore. The small spheres and groups of spheres are typical framboidal pyrite; x 50. White = pyrite; grey = sphalerite; black = gangue; note the mica flakes and carbonate rhombohedra. Bottom is to the right.

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

435

quently present in many ore deposits in sediments, for example, the many Red-Bed copper and/or Kupferschiefer deposits (Mansfeld, Corocoro, etc.). Similarly, the transitions from this type to the “sandstone type uranium deposits” on the one hand and to the complex Cu-Pb-Zn-Ba ores on the other hand, afford good examples of differential emplacement of sulfides and oxides both in space and in time. Fig.11, 13 and 14 are typical examples for this type of host-rock material. The space sequence shown in Fig.14 may well be a time sequence also. BARTHOLOMB (1962) described parageneses from Katanga (Congo, Africa), which may be understood best as diagenetic generations. Similar features were observed and described in the Rhodesian Copperbelt by NEWLANDS and TYRWHITT (1964): “The effect of differential compaction during diagenesis is demonstrated in a specimen of siltstone, where the pebbles within a gritty layer have been forced down into the underlying siltstone causing distortion of the bedding laminae. The interstitial sulphide within these distorted siltstone laminae is more densely distributed below the pebbles than elsewhere in the same lamina. This suggests that the sulphides were present before diagenetic compaction. Both pyrite and chalcopyrite have recrystallised at the same time as quartz overgrowths on detrital quartz grains. In the absence of any replacement textures between the sulphides, the mineralization is provisionally considered to be syngenetic with slight modifications during diagenesis.” Authigenic overgrowths like that of hematite, quartz, pyrite and galena in the Lamotte Sandstone (Plate 11) also afford good examples of paragenetic sequences; and the lengthy discussions on the genetic meaning of overgrowth, for example, in the Blind River or Witwatersrand areas (RAMDOHR, 1953; LIEBENBERG, 1955,1957), show how essential these features may be. The problem of authigenic generations is an old one. Recent discussions, however, such as the ones by TOPKAYA (1950), STRACHOV (1953,1956,1959), TEODOROVICH (1958,1961), VON ENGELHARDT (1960, 1961), and FEDIUK (1962), offer clear evidence of much value for genetic studies of ore deposits. In Fig.15 the major types of sulphide deposits affected by diagenetic changes are drawn schematically.

Gravity-density features The discovery of gravity-density features has perhaps started the change in the trend of thought outlined at the beginning of this chapter more than any other observations. The frontispiece of this book shows the first geopetal features recognized as such. in the Mississippi Valley deposits of Missouri. The two most complete collections of pictures and descriptions of gravitydensity features are probably contained in the classic book by SHROCK(1948), entitled Sequence in Layered Rocks, and the one by POTTER and PETTIJOHN (1963).

PLATE I

A

B

437

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

S y n g e n e t i c

As

4

y

nt=o

Ep i g e

n e tic

Af=n

la Sedimentation of material originating

10BaS04 endogenous to the

Ib Volcanic-exholativeBa and/or S mxed

l l b Bas04 exogenous to the sedi-

from erosion only; Ba, S, 0 endogenous to the rivers and to the ocean waters

andprecipitated with normal sediments; thus, exogenous source for Ba and/or S.

sediments, but concentrated by circulating groundwate/:

ments,introduced from the Cretac.Magnet Cove ring dike.

Fig.12.Thefour basic theories on the genesis of the Arkansas barite belt. S.S.

=

sandstone.

SANDER (1936, 1948, 1950) used the name “geopetal” for the distinctly polar gravitydensity features. The family of geopetal or top-bottom features is large and varied. Only a few examples are pictured in this chapter: Fig.7 and the frontispiece of this book show common load casts from Fredericktown ore; Fig.11, a microscopic load cast of a gastropod from Rammelsberg ore; and Plate IA, barite load casts from Arkansas. In Fig.17 pieces of dense quartzose dolomite rest as load casts in the sandstone; the latter shows flow streaks and must have been soft at the time of breccia flow and when it came to rest. In a general sense, stylolites are also gravity-density features, because they originate as a result of pressure exerted in most cases by the mere weight of the overlying strata (the pressure caused by gravity). Stylolites are common beddingplane indicators and thus can be considered as “isobars”, although they can also be caused by lateral or inclined pressures. Indirectly, stylolites have genetic significance too, because most of them formed during diagenesis (PARKand AMSTUTZ, 1967). Consequently, if diagenetic stylolites affect ore minerals, these ore PLATE I

Nodules and lenses of diagenetically early barite from the upper part of the ore body of the Chamberlain Creek syncline, Arkansas. Direct photographs of thin sections. South flank of the ore body. (Compare ZIMMERMANN and AMSTUTZ, I961,1964a, b.) A. Current bedding around load cast nodules in a shaly siltstone. The lower one is broken during diagenetic stage and the lower half is moved to the left. x 4. B. Various shapes of zoned nodules or layers broken or bent during diagenetic stage. Black matrix -shale. x 6 ,

438

G. C. AMSTUTZ AND L. BUBENICEK

Fig.13. Differential localization of sulfides in wood, which most likely occurred during diagenesis. Compare with diagram of Fig.14. Mitterberg Mine, Austria; x 100. Py = pyrite; Tetr = tetrahedrite; Cu = chalcopyrite; As = arsenopyrite,

439

DIAGENESIS IN IN SEDIMENTARY SEDIMENTARY MINERAL MINERAL DEPOSITS DEPOSITS DIAGENESIS

minerals are considered to be of diagenetic origin. This is the case in many places, although little attention has been paid to stylolites in ores. This author has observed good and abundant stylolites which most likely formed during diagenesis in the following mineral deposits and their host rocks: barite in Arkansas (cf., ZIMMERMANN and AMSTUTZ, 1961, 1964a, b); barite in Missouri (near Potosi, in many of the barite pits and road cuts); in the mud volcano-flow breccia of Decaturville, between two different flow breccias; in the southern Illinois fluorspar district, in many types of materials (Fig.3,5,6, 10,20); in many other Mississippi Valley type deposits in the Leadbelt proper, in the southern Tri-State area, and in the Tennessee zinc mines; in the barite of Meggen; and in some other mineral deposits.

?x

I

Fig.14. Differential localization of sulfides in fossil wood, which probably occurred during diagenetic crystallization. X signifies possible destruction of wood textures during crystallization. This illustration corresponds to the photomicrograph given in Fig.13. Mitterberg, Austria.

I I I I--

I

I

-----A _..4:,.

-_

_--

-. ...... .... /L .......... 1.

+ K

intra-.

peri-,

volcanic

or

apo-

kryptoor

or

purely

tole

- v o L can1 c

sedimentary

Fig.15. Schematic drawing of types of mineral deposits (especially sulfide deposits) which are affected by diagenetic changes. To the left (A’-H) are types of deposits which are clearly associated with volcanic rock, and on the right ( H ’ L ) are deposits which may or may not be connected with volcanic exhalative activity. K is largely contained in late diagenetic compaction fractures in and near organic reefs. Types A’-F’ merely refer to common types of distribution patterns in lavas, whereas H-L are located within sediments, with which they were formed.

440

P P 0

PLATE

G . C. AMSTUTZ AND L. BUBENICEK

c

I and 2. Lamotte Sandstone (or quartzite: MATHUR,1959) cemented by overgrowth of quartz, despite the iron-oxide ring, which here served as a guide to outcropping Precambrian iron deposits. Southeast Missouri, near Fredericktown. Diameter of sand grains is about 0.5 mm. 3 and 4. Transition zone between the Lamotte Sandstone and the Bonneterre Formation in the Fredericktown mine.

m

2

5x

3. Detrital quartz shows overgrowth generations of pyrite, quartz, and galena, which fill the interstitial spaces between the grains. Solid pyiite coatings on the detrital surfaces have obviously inhibited authigenic quartz overgrowth. (Compare e.g., 1963, fig.2, 3, and 4.) 4. Detrital quartz shows overgrowth generations of siegenite (intermediate relief) and pyrite (marked relief) before the authigenic idiomorphic quartz overgrowth sets in. As in 3, galena (light gray, scratched) is the latest generation. In other portions of the carbonates, chalcopyrite and some sphalerite also join into the parageneticsequence of diagenetic crystallization; x 300.

442

G. C. AMSTUTZ AND L. BUBENICEK

A few examples of gravity-density features are reported in the literature by MAUCHER (I957), BERNARD(1958), SCHULZ(1960), and SCHNEIDER (1964) from lead-zinc deposits of the Alps and from central France. A transition of gravity features to motion features was observed by Chico (as quoted by AMSTUTZ, 1962, fig.2).

A

B Fig.l6A, B. Diagenetic micro-tectonic fabrics from Mt. Isa, Australia; x4. White = galena; et al., 1964.) grey = sphalerite, pyrrhotite; dark grey or black = gangue. (After AMSTVTZ

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

443

Motion and disruption features Motion and disruption features are also quite common in sedimentary ore deposits, but have not been given much attention until recent years when some literature on coal was given by Lye11 (see HAARMANN, 1930, p.156). Motion features include diagenetic folding, slumpage patterns, flow and flowage features (ripple marks, cross-bedding, etc.). Disruption includes fracturing of various degrees to the extreme of breccia formation. A combination of motion and disruption takes place when fracturing is followed by flow in the form of intrusions or injections (Fig.16-20). A beginning of disruption is also evidenced by the broken barite nodule of Plate1A;andthe barite of Plate IB displays both flowage and breakage. Fig.16A and B also show a combination of both; as the competent layers break, their pieces float in the still soft matrix of the non-consolidated material. Small thrusts and folds show how rigid or plastic a particular bed was when the movement took place.

Fig. 17. Early Paleozoic mud-volcano stream. Unconsolidated sandstone is mixed and injected by the typical poorly sorted mud-volcano mass. The black matrix of the mass contains up to 30 % FeS2, PbS and ZnS. The broken white fragments are dolomite and dolomitic sandstone. Flow and geopetal features are conspicuous, Center of Decaturville, polygonal uplift, central Missouri. Length of knife i s 9 cm,

444

G. C. AMSTUTZ AND L. BUBENICEK

Rock layers which consolidate after younger, overlying beds are of much interest. If the roof breaks, which occurs commonly because of the constant supply of seismic energy in most areas, fissures, intrusions and extrusions would form. The most common form of such “sedimentary volcanicity” are the mud volcanoes. They are diagenetic phenomena of double interest in the study of mineral deposits: first, they occur more abundantly in or near oil fields, and secondly, they often contain high amounts of sulfides. An outstanding example of a fossil mud-volcano field is exposed in outcrops and many drill holes of the Decaturville area, Missouri. A wealth of flow patterns, load casts, and two- and three-fold brecciation is found in this area; and the mud has been injected in cross-cutting and “lit-par-lit” joints of the glauconitic Davis shale. Fig.17 and 18 give an idea of the features in this area. A heterogeneous collection of breccia fragments was carried along in the mud. The sulfides which consist of FeS2 (pyrite and marcasite in alternating layers), PbS and ZnS show the same brecciation features as the rock minerals, except fqr about 15 % of the pyrite-marcasite crusts or botryoidal masses which are rarely broken or overgrown as later generation crusts on the rock fragments.

Fig. 18. Mud-volcano dykes through the glauconitic Davis shale, “sulphide pit”, Decaturville (500 m southwest of the pegmatite outcrop). The dikematerial i s thesame as that shOwn in Fig.17.

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

345

Fig. 19. Diagenetic fracture filling is folded and congested by compaction of the surrounding dolomitic limestone. It supported the country rock, the volume of which was reduced by compaction to a lesser degree in the neighborhood of the stiff vein filling. Vertical extent of veinlet is about 50 cm. Southeast Oxford Mine. (This feature was also pictured by BRECKE,1962, p.527.) Many such features are also reported from the Jefferson City Dolomite by AMSTUTZet a]. (1964).

Of the numerous papers on mud intrusions, sand dikes and similar features, (1957), BAKER(1961) and by GANSSER (1959) should the ones by GILLand KUENEN be mentioned. The paper on the gas content of sediments by EMERY and HOGGAN (1958) may explain why some basins may have quietly active volcanoes, whereas others are erupting violently. In many mineral deposits the disruption features stand out extremely well, as for example in the mud-flow breccia shown in Fig.17. The strong color contrast between breccia fragments and matrix, or also between the wall rock and sedimen-

446

G. C. AMSTUTZ AND L. BUBENICEK

tary dykelets with sulfides, has produced the erroneous impression that brecciation occurs only or predominantly in ore deposits. This impression has been created on the one hand by the fact that breccias of common rock matter, although many thousand times as abundant as ore breccias are hard to see because the fragments and the matrix are almost always identical in color. On the other hand, the abundance of outcrops in mining areas (mine walls, drill holes, etc.) is also responsible for this wrong impression which has found its way into most textbooks. It may also be said again that the erroneous impression was of course also supported by the strong traditional inclination toward genetic explanations by epigenetic and “exogeneous” sources; and, therefore, breccias were often used erroneously in support of the myth of replacement from unknown depths. In the literature, sedimentary motion and disruptionfeatures are only scantily described for mineral deposits. At Meggen, excellent features belonging to the diagenetic period were found and described in detail by ZIMMERMANN and AMSTUTZ (1 967). These were previously considered to be of mechanical metamorphic origin. SANDER(1936, 1948-1950), BATHURST (1958), SCHULZ(1960), SCHNEIDER (1964), and others described a number of motion and disruption features in various rocks and ores. SELLEY (1961) observed penecontemporaneous deformations of heavy mineral bands in sandstones. Other examples could undoubtedly be added. In concluding this section on geometric patterns characteristic of the period of diagenetic crystallization and compaction, the following statement made at the beginning should be repeated: any subdivision into classes or types of features is always to a large extent artijicial. Not only are the transitions gradual, but the

bottom

Fig.20. Diagenetic rearrangement of marl-mud in thin coal layers of the Peissenburger Molasse beds; formation of ptygmatic folds during diagenesis; xO.5. Compare with Fig.19. (By courtesy of Prof. R. Fischer, Munich, Germany.)

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

447

three types of features are intimately connected. The following example from large ore deposits of the Lead belt of Missouri may suffice (see figures in AMSTUTZ, 1963a; or EL BAZand AMSTUTZ, 1963): the contraction fissures along algal fingers are disruption features filled with the last crystallizing phase. In the load casts, diagenetic tension and pressure have disrupted the early, large marcasite grains which show pressure twins and also cracks filled by the late galena. Fig.19 pictures a diagenetically formed and filled fissure in a competent carbonate layer, also in the southern Illinois Fluorspar district. In Fig.20, marly mud intrudes relatively competent coal layers. In establishing criteria, one has to make use of the excellent detective value of these mutually supporting observations. They create excellent cross-references and one feature confirms the assumption made on using another. In this manner the crystallization times of ore minerals can be determined, and detailed observations lead to a firm theory on an epigenetic or a syngenetic origin of an ore deposit. With regard to all three classes of features in ore deposits, viz. ( I ) theparagenetic crystallization sequences, (2) the gravity-density features, and (3) the motion and disruption features, one common statement can be made at this time: the literature is full of good descriptions of patterns which very probably belong to these sedimentary diagenetic features, but many were not recognized yet as such. During the corning years, it will be a main task in the study of mineral deposits to look at all conformable, congruent deposits in sediments with the eyes of a sedimentologist, and to determine the degrees of “isomorphism”, i.e., of symmetry between the enclosing “normal” rock and the ore minerals, into which geologists have projected an epigenetic origin, following the tradition of the myth of replacement by fluids from hidden outside sources.

GEOCHEMICAL CONSIDERATIONS

The information on the geochemical relations and conditions of the formation of sedimentary ores is well documented in the literature. In recent years a good number of laboratories have carried out successful experiments on bacterial sulfide formation and on direct inorganic precipitation of sulfides. The conditions reported are so closely related to those of common rock formation, that all essential details are contained in the chapters on the diagenesis of carbonates, silicates and especially clays. Consequently, this section will be kept brief. For the sake of completeness, however, a few comments on available literature are presented here. Valuable information on the geochemical conclusions to be drawn from the paragenetic relations in Cu-Pb-Zn ores was given by STANTON (1958, 19641, and on Cu-Co ores by BARTHOLOM~ (1962). Manganese-nickel nodules and their environment of formation on the present-day ocean floor was reported by MERO (1962a, b, 1965) and ARRHENIUS et a]. (1964). The geochemical environment of

448

G . C. AMSTUTZ AND L. BUBENICEK

sea-floor phosphorite formation, also in Recent sediments, was reported by MERO (1961, 1965). Geochemical reports on pH-Eh conditions, salinities, densities, profiles, etc. in soft sediments are numerous and their number has increased ever since the “Challenger” Expedition (about 1870), and especially in recent years. ZOBELL’S work (1942, 1946) will remain a classic in this field. The pH-Eh profile above and within ocean-floor muds and partly consolidated sediments answers many problems on diagenetic parageneses in sedimentary mineral deposits. Trace-element work on fossils by NODDACK and NODDACK (1939) and recently by LEUTWEIN (1963), indicate the receptivity of living matter in the ocean for base metals and other elements of economic significance. One of the most outstanding case studies is the group of papers on the geochemistry of the Black Sea. The papers of SHISHKINA (1959), STARIKOVA (1959) and ZAITSEVA (1959) contain pertinent information on sulfide formation. Bacterial action was recognized early to be the possible major single environmental factor in ocean-bottom muds. It was investigated in many different areas of direct interest to ore deposits. BAASBECKING (1955, 1957, 1958, 1959) was one of the pioneers. WESTOLL (1955) and SCHWARTZ (1957) offered excellent summaries on the role of Bacteria, whereas STRACHOV (1953, 1956, 1959), MARKEVICH (1 957, 1960), and NICHOLLS (I 958, 1959) provided biogeochemical facies discussions. DEGENS (1964) studied the problem of the origin of fossil organic matter and its genetic significance. A study of biogenic constituents of ore deposits should provide an excellent clue to the possible functions of various biogenic assemblages in the fixation of base metals and other elements. Hydrocarbon derivatives of metals as outlined by PAULSON (1962) may lead to the discovery of significant collection processes. The research on syngenetic mineralizations (WALPOLE, 1961) needs to be pushed in this direction. The result may well show that most of the accumulations considered to be due to the epigenetic movement of groundwater or hydrothermal water by SEIDL(1958), POUSTOVALOV (1959), and GERMANOV (1961) are of diagenetic origin. It is certainly remarkable that the accumulations of PbS, ZnS and FeS2 in the Krakow Leadbelt as well as in the northern Tri-State and in other areas occur immediately above petroliferous or at least bituminous shales (cf. GRUSZCZYK and WAZEWSKA-RIESENKAMPF, 1960; GRUSZCZYK, 1961; and TZSCHORN, 1963). A most significant paper regarding the role of diagenetic processes in the formation of sulfide minerals published in the last few years is that by STANTON and BAASBECKING (1962) on the formation and accumulation of sedimentary sulfides in seaboard volcanic environments. The oxidation-reduction cycles and the role of bacterial action are reviewed in terms of the availability of base metals from submarine volcanic exhalations. Evidence for ore formation from volcanic exhalative sources is abundant. Major references on exhalative origins of disseminated and massive sulfide deposits are: BORCHERT (1957), CISSARZ (1957), OFTEDAHL (1 958), MILLER(1960), SCHNEIDERH~HN (1962), STANTONand BAASBECKING

IV ON THE CHEMICAL CHANGES TAKING PLACE IN THE DEEP GROUND-WATER ZONES

TALJPITZ,1955) processes oxidation (oxygen is carried down by downward moving waters)

reduction (oxygen has been usedup due to oxidation of organicsubstances) ~

Sulfides-sulfates

efects

change

effects

Metals mostly very soluble; carbonates very soluble. Ba is precipitated; soluble Fez+ precipitates as Fe3+.

is reduced by organic substances (possibly through bacterial action); COZ present.

Metal fixation as sulfides; Ba as sulfide soluble. Fe3+ as Fez+ soluble but precipitated as sulfide.

alkaline

COz":

By oxidation of organic matter.

Through NH4:

HzS04:

Through Ca(HC03)z:

HzS:

Mainly through oxidation of sulfides; also s04'- from gypsum, anhydrite and salts. From organic compounds and reduction of sulfates.

C1-:

From salts.

Through NaHC03:

Concentration of ions and gases in solution high

low

Many metals and Ba readily solubie. PbS (La. sulfides) more soluble. Ba precipitated as BaS04; many heavy metals readily soluble. Carbonates more soluble.

Ore minerals less soluble. PbS (La.) less soluble.

Solubility of HzS and COz in water increases with increasing pressure.

Carbonates less soluble.

From organic substances in a environment. From limestones. Reaction with Ca(HC03)z involving base exchange (Ca replacing Na), in clay minerals (see FOSTER,

450

G . C. AMSTUTZ AND L. BUBENICEK

(1962), ROUTHIER (1963), and SCHNEIDER (1964). (See also Bernauer as reported in AMSTUTZ, 1959.) A collection and discussion of some Eh-pH diagrams applicable to lowtemperature, low-pressure processes in rocks was first published by GARRELS (1 960). A more complete collection of equilibrium diagrams without text was made available in 1962 by the Geology Club of Harvard University (SCHMITT,1962). Recently BUBENICEK (1964) also presented a summary to which the reader is referred. The review on the role of diagenesis in the formation of mineral deposits by T A u P I T z (1965) concentrated on geochemical solution and redeposition processes essentially after the rock has been formed. Some workers may perhaps not want to include these sometimes definitely postdiagenetic ground-water effects in the class of diagenetic processes. The valuable summary of TAUPITZ (1 965), however, certainly also applies equally well to strictly diagenetic crystallization. His tabulation of processes is, therefore, reproduced in Table IV. It may be stressed again that not all diagenetic processes are “Umlagerungen”, the closest translation of which is “re-deposition”. Diagenesis of mineral deposits just as all sedimentary rocks is essentially a “trilogical” process consisting of crystallization, recrystallization and metasomatism. All three steps occur with or without introduction and removal of material (dissolution of solid phases or replacement by poresolution material). The relative time at which the rock and ore minerals formed can commonly be determined by their geometric position and shape, as outlined in the previous section. A relatively new theory on diagenetic mineral formations, which affects the ideas on the origin of some ore deposits, especially the native copper deposits of the Lake Superior type, has been advanced by COOMBS et al. (1959) and CROOK (1960). Their basic idea is that some minerals formerly considered to be of hydrothermal-volcanic, i.e., deuteric origin, may have formed during the diagenesis of tuffs and lavas. The fact that mineralogical distributions of the copper and its associated minerals in the Michigan lavas are perfectly congruent to primary features, has lead CORNWALL (1951) and AMSTUTZ (1963) to reject any epigenetic-hydrothermal theory and to propose an essentially contemporaneous formation for the copper and the typically associated minerals albite, chlorite, zeolites, prehnite, carbonates, quartz and epidote. Refined methods will have to be used in establishing whether many of these minerals may in fact form as diagenetic facies markers. The distribution patterns shown by COOMEIS et al. (1959) and by CROOK (1960, e.g., fig.1, p.79) can be interpreted to represent original compositional differences as well. The occurrence of a certain mineral at a certain depth does not, without matching positive evidence supported by different independent parameters, prove that this mineral would not occur higher in the section if it had been deposited in a lava bed or tuff as a deuteric mineral.

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

45 1

As a matter of fact, extensive lava layers of similar composition which were also deeply buried do not reveal the same “facies” (for example the Permian spilitickeratophyric tuffs and lavas of the Glarus Freiberg). This problem of diagenesis as related to ore deposits in and near basaltic rocks is as yet unsolved and more refined methods must be applied before the mineral assemblages involved may be assigned a deuteric or a diagenetic origin. It seems, however, that both possibilities exist, that transitions may be abundant, and that refined observational methods must be applied to differentiate between these two modes of formation.

DIAGENESIS IN DIFFERENT TYPES OF DEPOSITS

The foregoing sections consist of a brief review of the historical aspects, the geometric and geochemical factors and observations. Many individual mineral deposits were mentioned, but no systematic listing was given of papers relating to the diagenesis in various groups or types of deposits. The following incomplete tabulation is meant to be a mere guide to some of the pertinent literature. Not all of the works listed discuss diagenesis. Many, especially older ones, may have interpreted certain features as of metamorphic or otherwise epigenetic origin (e.g., SEIDL, 1958;PousTovALov,1959;GERMANOV, 1961,1963). All the papers, however, describe features or processes which, in the light of modern research and with the application of modern schools of thought, are directly or indirectly produced by, or responsible for, diagenetic changes in mineral deposits. This list is presented in alphabetical order. A short tabulation of references

Authigenic minerals TOPKAYA (1950), RAMDOHR (1953, 1960), STRACHOV (1953, 1956, 1959), TEODOROVICH (1958, 1961), VONENGELHARDT (1960, 1961), and FEDIUK (1962). Bacteria See biochemistry. Biochemistry NODDACK and NODDACK (1939), ZOBELL(1942, 1946), STRACHOV (1953, 1956, 1959), BAASBECKING(1955,1957,1958,1959), WESTOLL (1955), SCHWARTZ (1957), SHISHKINA (1959), STARIKOVA (1959), ZAITSEVA(1959), WALPOLE (1961), KUZNETsov et al. (1962), PAULSON(1962), LOWENSTAM (1963), DEGENS (1964), FRIEDMAN (1964), FUCHTBAUER (1964), JOHNS(1964), and LEGATE and JOHNS (1964).

452

G. C. AMSTUTZ AND L. BUBENICEK

Coal and underclays or claystones THIESSEN (1945), JACOB(1954, 1955), MACKOWSKY (1955), MARSHALL (1955), and T E I C H ~ L L (1958). ER Copper ENTWISTLE and GOUIN(1955), STANTON(1958, 1964), STEINBRECHER (I 959), KOBE(1960), PIEKARSKI (1961), BARTHOLOM~ (1962), GARLICK (1964), NEWLANDS and TYRWHITT (1964), and MACQUAR and TREUIL (1965). Exhalative volcanic origin BORCHERT (1957), CISSARZ(1957), OFTEDAHL (1958), AMSTUTZ(1959), MILLER (1960), SCHNEIDERH~HN (1962), STANTON and BAAS-BECKING (1962), and SCHNEIDER (1964). Facies STRACHOV (1953, 1956, 19-39), LOMBARD (1956), MARKEVICH (1957, 1960), MAUCHER (1957), BERNARD (1958, 1964), NICHOLLS (1958), SUJKOWSKI (1958), WELLER (1959), LOMBARD and NICOLINI(1960), LOWENSTAM (1963), POTTER and PETTIJOHN (1963), ROUTHIER (1963), and MACQUAR and TREUIL(1965). Fluorite KRUGER(19611, PARK (1962), and PARK and AMSTUTZ (1967). Gypsum-anhydrite RICHTER-BERNBURG (1953). GoId-uranium RAMDOHR (1953, 19601, LIEBENBERG (1955, 1957), ROUTHIER (1963), and SCHIDLOWSKI and TRURNIT (1966). Iron deposits TAYLOR(1949, 1955), JAMES(1955), OSTROUMOV and SHILOV(1956), ANDERSON and HAN(1957), HOUGH(1958), BUBENICEK (1961), BRAUN(1963), MOHR(1963), (1964). and PETRANEK Lead, zinc, barite KONSTANTINOV (1952,1954), MAUCHER (1957), SCHUELLER (1958), STANTON (1958), GRUSZCZYK (1960a, b), GRUSZCZYK et al. (1961), WILSON(1961), ZIMMERMANN and AMSTUTZ (1961, 1964a, b), ROUTHIER (1963), TZSCHORN (1963), GARLICK (1 964), PUCHELT and MULLER(1964), STANTON (1964), and HAGNIand GRAWE,(1964). Manganese EPPRECHT(1946), MOHR(1959, 1963), MERO(1962a, b), and ARRHENIUS et al. (1964).

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

453

Mud volcanoes GILLand KUENEN(1957), EMERYand HOGGAN(1958), GANSSER(1959), and BAKER(1961). Nickel-cobalt AMSTUTZ(1958b, 1963a), AUGUSTHITIS (1962), EL BAZand AMSTUTZ (1963), and AMSTUTZet al. (1964). Petroleum SUJKOWSKI (1958), CORBETT (1959), ILLING (1959), andTHoMAsandBLUMER (1964). Phosphorite MERO(1961, 1965). Placers (in general) RAMDOHR (1953, 1960), and SELLEY (1964). Pyrite B ~ ~ ~ ~ ( 1 9 6 1 , p p . 2 9A , 5M 2 )S, T U Tal. Z ~(1964), ~ Lov~(1964),H O N J Oal.~(1965), ~ and Low and AMSTUTZ (1966). Quartz-sandstone AMSTUTZ (1957), MATHUR(1959), and DAPPLES (1962). Salt deposits STEWART (1951, 1955), BRAITSCH (1962), JOHNS(1963), and EUGSTER and SMITH (1 965). Snow BADER(1939), and GRAENICHER and JONA (1960). Uranium RAMDOHR (1953,1960), LIEBENBERG(1%5,1957),AMSTUTZ (1962),ROUTHIER(1963), FINCH (1964), and SCHIDLOWSKI and TRURNIT (1966). Final remarks

Before closing this analysis and outlook for future research, a word of caution should be added, i.e., not everything can be explained by diagenetic processes. As a matter of fact diagenesis is only one of very many processes involved. Also, this author knows of no ore deposits named in this chapter which do not show some obvious or most probable epigenetic features. The genesis of most ore deposits is complex, and virtually all deposits show signs of more than one stage of geological

454

G . C . AMSTUTZ AND L. BUBENICEK

processes. The task of the scientist is not to succumb to that negative trait of “western cultures”, i.e., to think in terms of mutually excluding opposites, and to terminate all further scientific investigation with a “credo” for one of the extremes. The historical evolution of rocks is also in itself a superposition process and an analysis needs to “peel off” the later processes in order to uncover the very first period of formation. The fixation to a genetic dogma is a psychic process which consists of fixations to assumptions. JUNG (1959, p.18) made the following statement on this fixation: “Under the influence of scientific assumptions not only the psyche but the individual man and, indeed, all individual events whatsoever suffer a levelling down and a process of blurring that distorts the picture of reality into a conceptual average. We ought not to underestimate the psychological effect of the statistical world picture: it displaces the individual in favor of anonymous units that pile up into mass formations.” As long as this fixation is in the open and recognized in open dogmatic statements it is not difficult to overcome. However, if it hides behind mechanistic, pseudo-scientific facades in the form of statistics or enormous calculations based on hidden assumptions, not in any way combined with differentiated thought, it certainly causes a lot of confusion. This situation has to be mastered in regard to the problem of diagenesis, too. This is the “inside front of science” which lies inside man and not within the object itself. No problem of science of this large a scope can evolve today without a keen open mind on both fronts of science, the “objective” one, and the “subjective” one. Diagenesis, the “soil science of the oceans”, can be studied only with a keen realization of the various forms of early fixations to premature assumptions.The pitfalls ofpseudo-scientific,mechanistic approaches without thought as the checking guide was perhaps best described by RITTENHOUSE (1959, p.1501): “All too often, experience is substituted for thought-not made its partner. Because certain types of observations or measurements have contributed to the successful solution of problems in the past, they are selected and applied indiscriminately to new problems. Here we have one type of ‘shot gun’ approach, based on a philosophy which, in effect, says: ‘If we can make enough observations on enough rocks, put them in a machine and turn the crank, something useful may come out.’ To some, this is the ‘modern statistical approach’ to geological problems. This maligns statistical methods-which can be valuable in geology-though not as a substitute for thinking. Since this approach does not reach the objective in a minimum of time, if at all, it is wasteful of time, manpower, and money. I prefer the ‘there is a reason’ philosophy.” “In selection of features that may be significant in rocks, we are faced with ever increasing knowledge in geology and related sciences, and ever increasing specialization, In looking at rocks and at problems involving rocks there is an

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

455

increasing tendency to look at specialized aspects of the rocks, rather than at the rock as a whole. The sands ‘belong’ to one specialist; the clays, the carbonates, the isotopes, the trace elements and various fossil groups, to others. Specialization has great potentialities to our science, since it can lead to a better understanding of the physical, chemical, and biological factors that have made rocks what they are today. But it has great hazards as well. Let’s not forget how the elephant appeared to the three blind men-like a wall, like a post, like a snake. Just as we need to consider the entire elephant, we need to consider the entire rock-putting each type of observation and deduction in its perspective.”

THE ROLE OF DIAGENESIS IN THE FORMATION OF OOLITIC IRON ORE DEPOSITS-A

CASE

STUDY

Introduction-de$nitions Iron ores of sedimentary origin have been found with many different physical and chemical properties and in all types of rocks and from all stages of sedimentary evolution. Four fundamental categories, however, may easily be distinguished, in order of importance: (1) The Lake Superior ore type, including low-grade banded ores, B.H.Q., taconites, etc., i.e., the different iron-bearing formations (JAMES, 1965). (2) The oolitic ores, with very varied parageneses. (3) The glauconitic ores. (4) The sphaerosiderite ores. The stages of evolution undergone by the sedimentary iron ores before reaching the present state are numerous and varied. Attempts have always been made to understand the mode of formation by also trying to explain the apparent anomaly of the iron mineralization. The diversity of the proposed classifications reflects the complexity of the phenomena. If one attempts to distinguish the role of the main sedimentary processes and in particular the role of diagenesis, however, it is necessary to study all the stages of transformation. Many authors working mainly on pre-Hercynian ores that have been considerably modified, have often denied all possible effects of diagenesis, even of epigenesis. They explained the mineral associations solely by the processes of syngenetic precipitation during deposition. Taking into consideration the present state of knowledge, however, it is impossible to ignore diagenesis. It is an extremely important process, particularly in the case of iron, an element whose chemical behavior varies greatly according to its ionic state. At present, the ores of the Lake Superior type have still not been sufficiently studied in terms of their origin and subsequent changes to allow an approach to such a precise subject as their diagenesis (JAMES,1955). The importance of sphaero-

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G . C . AMSTUTZ AND L. BUBENICEK

siderites is mainly historical, because these ores were the mainstay of the British iron industry during the 19th century; but no new exploitation seems likely at present. Diagenetic segregation of the siderite from diffused material in clays appears to be the most common present interpretation (WILLIAMS et al., 1954; KAZAKOV, 1957). The glauconite iron ore deposits that could possibly be worked are very rare; however, the frequent occurrence of these ores has led to some research. The problems in this case are similar to those of the oolitic ores as they have many common features. These are mainly: texture, nature of associated minerals and depositional environment. The great peculiarity of these ores is that the glauconite is of primary origin, and several hypotheses have been proposed to explain this. The oldest involves the activity of Foraminifera. Also the frequent association of glauconite with biotite led to the hypothesis that glauconite is a weathering product of biotite. More recently, some Russian authors assigned a diagenetic origin to this mineral in a slightly oxidizing environment (KROTOV,1952; KAZAKOV, 1957; HOWER, 1961). This brief general survey shows that the research is not sufficiently advanced to allow a satisfactory definition of the diagenetic history of the iron ore deposits, except for the oolitic ores. All possible evolutionary stages may be noted in the latter ores, which occur in all formations from the Precambrian up to the Lower Quaternary. It is in these ores that the writer attempted to outline the role of diagenesis. In a publication to appear soon this author will show that the evolution of the Precambrian banded iron ores can be explained in much the same way as that of the oolitic iron ores; modified, however, to the conditions which existed at that time. Historical review of literature on diagenesis in oolitic iron ores

In the earlier publications on oolitic iron deposits, emphasis was laid upon the process of diagenesis in the formation of the deposit, that is to say in the concentration of the iron. Thus, numerous hypotheses have been suggested, involving mineralizing solutions and the replacement of pre-existing rocks. As a result of progressive accumulation of new data, however, it was concluded that the concentration of iron in a deposit is a phenomenon related to differentiations antedating the deposits. Many different genetical schemes have, therefore, appeared to explain the existing textures. The role assigned to different processes varies a great deal from author to author, and according to which observational element is regarded as decisive. The different theories may be classified as follows: ( I ) The iron has been concentrated and deposited during diagenesis. The iron (and silica) has been brought in by mineralizing solutions of various origins (marine, submarine, thermal springs, etc.) with replacement of oolitic and cal-

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

457

careous detritus by iron solutions containing SiOz and Fe (CASTANO and GARRELS, 1950). ( A ) Without reworking phenomena: the different minerals were directly formed during the replacement processes (VILLAIN,1899, 1902; AREND,1933). ( B ) With reworking phenomena: the different mineralogical types were formed by oxidation during the reworking process (CAYEUX, 1909, 1911, 1922; DEVERIN, 1945). (ZZ) The iron ore deposit resulted from a sedimentary differentiation process predating the deposit, i.e., predating sedimentation. ( A ) With direct precipitation of iron along with other components at the bottom of the sedimentary basin. ( I ) Direct formation of different iron ore minerals during precipitation. Formation of oolites in situ as a result of diagenesis. The mineralogical distribu1940; CAILtion depended on the distance from the shore line (POUSTOVALOV, L ~ R Eand KRAUT,1954, 1956; BRAUN,1962; PETRANEK, 1963). (2) Formation of oolites in situin minerals having reduced iron (chlorite). Formation of limonite due to oxidation during a reworking stage, with detrital deposition of oxidized oolites. The siderite is considered by some authors to be of diagenetic origin (BERG,1924,1944; HALLIMOND, 1925; BICHELONNE and ANGOT, 1939; TAYLOR,1949; KROTOV,1952; BUSHINSKY, 1956. Without details: BRACONNIER, 1883). ( B ) Formation of oolites in dynamic conditions before deposition. ( I ) The oolites were directly formed with various mineralogical types of iron minerals, and then deposited. (a) In a primary cement (BORCHERT, 1952). (b) Alone: cement or a few minerals only being of diagenetic origin (POPOV,1955; TOCHILIN,1956; COURTY,1959, 1960; FORMOSOVA, 1959; DUNHAM, 1960; TEODOROVICH, 1961). (2) The oolites are of an oxidized nature: the different reduced iron minerals are the result of diagenetic processes (BROWN,1943; HARDER, 1951, 1957; CORRENS, 1952; KOLBE,1953, 1960; BUBENICEK, 1961, 1963). These various hypotheses reflect the treading of new paths of research, as new data were obtained. Every genetical theory, however, should not only give an explanation of the facts but also allow practical conclusions and inferences to be drawn. A look will now be taken at the present knowledge on this subject. Present knowledge of the diagenesis of oolitic iron ores Conditions of deposition The data available at present leads one to believe that the sedimentary differentiation of the iron is due to two main processes. The first one, of a pedological nature, is associated with the evolution of the continent. The second process takes

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G. C. AMSTUTZ AND L. BUBENICEK

place in the basin itself, where iron is precipitated and oolites are formed. It appears, therefore, that there is no direct relation between the influx of iron and the evolution of the basin. Because of the oolitic form of iron it is deposited together with detrital grains, quartz grains or fragments of various shells. It is at this stage that the last possible concentration of the iron takes place by variation of the relative proportions of detrital particles. Consequently, there are deposits within primary structures of current-bedding and cross-bedding, which reflect current effects, whereas in clay-muds penecontemporaneous deformation features, such as slumping and animal burrowing, are common. The conditions of deposition are such that the iron-bearing oolites contain the iron in its most oxidized form (Fe3*) in association with oxides of aluminium, phosphorus and manganese. The constant composition of the material precipitated in the oolitic envelopes (limonite), and the fact that generally in almost all known oolitic iron deposits subsequent changes did not lead to the subtraction or addition of major components, allows a preliminary interpretation of the chemical composition of these ores. Some conclusions can be drawn on the basis of the almost constant amounts of A1203 and P, along with the dominant iron, especially in the original limonite: ( I ) These ores always have A1203 and phosphorus contents1 proportional to the amount of iron (the more common A1203/Fe ratio is 0.10-0.12 in Recent deposits and 0.05 in Paleozoic deposits) in the case of the non-clay ores. This is valid for all the iron contents, which depend on the original dilution with quartz grains or calcite from shell debris. (2) An excessive amount of A1203indicates clay ores.

Diagenetic evolution The limits of diagenesis. When discussing diagenesis it is always necessary to define the limits of this process. The present author assigns to diagenesis all the processes which act on the sediment after the end of mechanical movements of the particles, whether in deposition or in seeking a mechanical stability through slumping, and in a milieu having lost all direct relation to the medium of precipitation. Thus, on the one hand diagenesis can start in fine clay sediments that are only a few millimeters thick, whereas in very permeable sands circulation of water coming from the surface of deposition can preserve syngenetic conditions over a certain period of time. Generally, however, cessation of the mechanical movement of the particles involves isolation of interstitial solutions. The communication between the depositional environment and these solutions exists only through diffusion or by very

For the whole ore body; taking into consideration the whole thickness of a layer, and not only sections where segregation phenomena could be involved. l

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

459

slow circulation, which depends on the permeability of the sediment and on possible charge differences. Circulation occurs on the beaches where the water of waves returns to the sea through the deposited sand. The end of diagenesis is more difficult to establish. The writer places it at the beginning of epigenesis, when hydrated minerals are transformed into less hydrated forms (limonite to hematite; silicates to micas) with compaction of the rock and loss of its permeability. It is between these limits, in fact, that the oolitic iron ores acquire the essential characteristics of their mineralogical facies. The role of diagenesis in the paragenesis of iron-bearing minerals Nature of the diagenetic environment. The diagenetic environment mainly consists of two phases: first, the solid phase, i.e., the detrital sedimentary phase where the dominant iron-bearing components are present in the most oxidized forms; and second, the liquid phase, namely the pore solutions filling the pore spaces between the grains or the micropores and microfractures of these particles. These solutions are at first identical to those of the sedimentation environment (practically always sea water). They are, therefore, rich in salts, organic matter and organisms of all kinds. As a result of lack of oxygen, the environment becomes progressively more reducing; this is usually associated with a notable change of pH through temporary acidification (BUBENICEK, 1964). The diagenetic environment has some oxidation-reduction potential and capacity (a function of the quantity of the confined organic matter and of its possible renewal through slow movements of solutions). At a given pH, there are organisms which are able to live under reducing conditions, in particular the sulfatereducing Bacteria. Furthermore, the solutions contain an important reserve of a variety of different ions in variable concentrations. This new environment establishes itself more or less rapidly and deeply in the deposit, depending on the type of material and other conditions. In order to define these ideas more accurately, one has to consider the level at which Eh = 0 with respect to the depositional interface. ( I ) If the level at which Eh = 0 is within the sediment, its depth will depend on the conditions given above. In this case the changes will be of a diagenetic nature. (2) If the level at which Eh = 0 is above the top of the sediment, deposition of products by direct precipitation from solution could occur. In this case one may only apply the term syngenesis (as defined in sedimentology). Nature of the transformations. The passage from one environment to another, having different physico-chemical characteristics, produces a reorganization of the chemical elements in order to establish a new and more stable equilibrium. In this process, the minerals previously formed become unstable and the new minerals

I

4

8

6

PH

1

0

1

2

1

L

0

2

L

8

6

I

-1.0

I

1 0 1 2 1 4

PH

Stability diagrams of the main iron minerals under various experimental conditions;.t Absence of silica and COz, CS=10-6. CS =10P. Absence of silica, CCOz= Presence of silica, CCOz= 10-2, CS=10-6.

C I

I

r

0

2

4

I

8

6

1

0

1

PH =

25°C; p = l atm. (After GARRELS, 1961.)

2

1

4

3

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

461

100

s

h

-.

2 .-

50

c

I

0

0

A

100

E

v Y,

Z

. a C

50

0 3

0

Fig.22. Diagrams of the evolution of limonite-quartz-calcite ores during diagenetic reduction. These ores have not undergone metamorphism (Lorraine; central Great Britain; deposits of the 1963.) Jurassic of Germany; etc.). (After BUBENICEK, A. Evolution when quartz is in excess. B. Evolution when limonite is in excess.

are precipitated in a stable form, accompanied by exchange of chemical elements (addition or subtraction) with the interstitial solutions. (1 961), who defined The formation possibilities were discussed by GARRELS the theoretical conditions of stability of the principal iron minerals (Fig.21). Although theoretical, these diagrams show that all pH-Eh variations involve crossing of stability boundaries, and this happens during diagenesis. The following should be noted, however: (1) These diagrams should be completed by the addition of the areas where the mineral species can exist in a metastable state. (2) The stability diagrams only give an imperfect picture of the reactions,

462

G. C. AMSTUTZ AND L. BUBENICEK

Fig.23. Calcite concretions. There is a continuity of the Iaminae between the concretion and the inter-concretionary ore. The inflection of laminae along contact with the two margins indicates a settling of the inter-concretionary ore material and increase in volume of the concretions during the process of CaCO3 displacement; x0.25.

because, if a closed environment is supposed to exist, the concentrations would vary in accordance with the displacement of the reactions. Consequently, threshold to other reactions would appear (Fig.22).

Main transformations. The main reactions which can occur, as already SUSpected on studying many deposits, have been made clear by the study of the Lorraine iron ore deposits. The first fundamental reaction for the non-clay ores is: limonite quartz --f siderite chlorite The determination of the relative proportions of siderite and chlorite is

+

+

TABLE V CHEMICAL COMPOSITION OF THE IRON-BEARMG PHASES OF THE MINEWE OF LORRAINE

limonite chlorite siderite

Fe

SiOz

CaO

MgO

Ah03

P

HzO

52.0 31.9 31.7

4.0 29.1 -

0.5 1.5

1.2 6.5

6.0 9.9

0.7 0.2

11.5 10.9 39.2

5.3

5.6

-

-

463

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

made possible by the invariance of A1203 and iron contents and by the knowledge of the A1203/Fe ratio in the original limonite. This ratio is equal to 7/52 in the case of the Lorraine limonite. It seems that this ratio is applicable to all Lorraine-type oolitic iron ores. Taking the average chemical compositions of the phases given in Table V, the reaction becomes: 6.05 limonite 1 quartz + 4.72 siderite 4.26 chlorite These values only represent the first attempt to roughly estimate the proportions of the constituents. In this reaction, the retained constituents are Ak-03, Fe, Si02, Mn, and P; MgO and S are added, whereas H20 is subtracted. The origin of the C02 is open to discussion; however, in the siliceous sediments where calcite is not abundant, pyrite is formed. This would suggest that at least this COz is derived from the calcite, and that there would be a corresponding loss of CaO with the gradual destruction of the calcium carbonate. Depending on the initial quartz/limonite ratio, a first threshold would appear when one of these components had been completely consumed in the reaction forming chlorite and siderite. New transformations then intervene: ( I ) When the limonite is in excess, there is formation of hematite, siderite and then magnetite; and (2) when the quartz is in excess, there is destruction of the first generation of chlorite and formation of siderite, chlorite, and secondary quartz. Equilibrium is reached when the limonite/quartz ratio is equal to 6.05. The various parageneses and their evolutionary relationships are presented diagrammatically in Fig.22. The limonite and the quartz transformations as a result of fundamental reduction reaction can come to a standstill in the case of exhaustion of the reducing agent or of one or the other of the original components. If in the initial stage these two components are consumed, an excess of quartz or of limonite would be the result. The boundaries, beyond which new parageneses occur may be determined through calculations.

+

+

The role of diagenesis in modifying structures. The primary and fundamental structures within the oolitic iron ores are, as noted earlier, current-bedding, crossbedding (most common), and contorted structures. Diagenesis can modify these in two different ways. ( I ) By the general volume shrinkage due to formation of more compact new minerals. This is followed quite frequently by the development of upright cracks, which can be filled mechanically. (2) By change in the distribution of components (Fig.23); usually by segregation of some primary component (e.g., calcite) or of components produced during diagenesis (e.g., siderite, pyrite). These segregation phenomena correspond to the reorganization of the constituents, in order to reach a greater degree of stability. As RAMBERG (1952) pointed out, due to surface-energy differences, the free energy of concretion material is lower when the material is concentrated than when it is dispersed. As the stability of a mineral increases with a decrease of free

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G . C. AMSTUTZ AND L. BUBENICEK

energy, this phenomenon is frequent for minor components (SEIBOLD,1953; PETTIJOHN, 1956). Differences of structure or texture can control the arrangements of the concretions, which may be scattered in beds rich in concretions or may be perfectly localized and rounded. The presence of centres of preferential attraction and the role of the permeability in the emplacement of concretions and aggregates should also be considered. It must be noted that migration of components may appear to modify the above rules at the level of individual samples. The role of diagenesis in the formation of textures. The final texture of the rock reflects all the above described changes and also those which arise as a result of metamorphism and the effects of meteoric waters, which largely tend to destroy earlier textures. Nonetheless textural studies alone often enable one to determine the history of the physical and chemical reorganization of the components. Three groups of textures due to three fundamental processes may be distinguished: (1) Filling textures (cementation textures) due to displacement of material at the time of segregation. At the actual time of crystallization an increase in volume frequently occurs. On the other hand, in impoverished zones physical reorganization is brought about by compaction, quite often with breaking of

Fig.24. Development of chlorite from the limonite of oolites. The chlorite appears as a pellicular cement around the limonitic oolites. The advanced transformation of the limonite into chlorite appears clearly on some oolites. Lorraine ore; grey bed = chloritic paragenesis; x 150. Natural transmitted light.

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

465

Fig.25. Corrosion of a quartz grain by siderite. The two residual areas of quartz show indented rounded outlines, and have the same optical orientation. The outlines of the quartz-siderite association preserve the shape of the detrital quartz grains; x 150. Natural transmitted light.

oolites or plastic deformation. There is frequently secondary growth in the zones of enrichment and also sometimes of saturation, especially in the case of quartz. (2) Solution textures are frequently associated with corrosion textures, and the disappearing minerals are clearly different from the new ones. The disappearance can take place in a progressive or alternating way, as in the case of the transformation of the oolitic limonite into chlorite. (3) Growth textures are the fundamental textures which distinguish the growth (by precipitation of whatever origin) of a mineral at a given point and include: (a) Incrustation-cementation on all minerals; the pores may remain open or may be filled completely (the most usual case for chlorite, Fig.24). (b) Corrosion: with the new mineral filling up gaps left as a result of the disappearance of destabilized minerals (frequent for siderite corroding quartz grains, Fig.25). (c) Authigenic, with euhedral shapes: the mineral crystallizes into euhedral shapes and forces away the surrounding material. Siderite crystals could develop crystalline faces against clay minerals, but would only adjust their shape to that of an oolite or its envelope. At the interface between the authigenic crystal and the surrounding material the effect is comparable to that of corrosion, and siderite shows this very often against clay minerals and sometimes against calcite. In

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G. C. AMSTUTZ A N D L. BUBENICEK

ores with very little quartz, magnetite (euhedral grains) formed as a result of diagenetic reduction. The effect of diagenesis on the chemical composition of ores. As indicated above, the diagenetic reactions which occur after deposition essentially do not involve exchange of the original Fe, A1203, P and Mn with interstitial solutions. At the most, the role of these solutions is to take part in displacements of material, which are considerable in the case of the formation of concretions and slight in the case of authigenic mineral formation. It is different for other elements and compounds, which may indeed undergo exchange with solutions present in the pores of the rock. These include particularly MgO, H20, C02 and CaO. The movements of these materials can considerably modify the mass relationships between the elements and greatly change the iron content of the ore. Thus, two ores with the same original iron content could be quite different if diagenesis affected them differently. In both cases, however, the relationships between the elements which were not exchanged will remain unaltered. This permits an adequate comparison of different ores. In order to reduce them to identical conditions it is advisable to compare them after an ignition loss test. The Fe2+/Fetotalratio is used to determine the evolutionary stage of the ore. In the case of the Lorraine ore, exact knowledge of the changes and of the minerals that are present, has permitted the calculation of the chemical and mineralogical compositions solely on the basis of the determination of Fetotal, Fez+, Si02, A1203 and CaO contents (BUBENICEK,1963). Conclusions

The present analysis of the role of diagenesis in the formation of iron ores shows the complexity and also the intensity of the phenomena which affect the sediments immediately after their deposition. Many Mesozoic and Tertiary deposits have not been so intensely modified by epigenesis as the older formations and they, therefore, give a clear picture of the fundamental process of diagenesis. The older deposits especially, and all those which have been deeply buried or affected by an orogeny, have been subjected to important modifications which frequently mask the earlier textures. It would appear, nevertheless, that the chemical characteristics established during diagenesis, and particularly the reduction index Fe2+/FetOta1,are preserved for a long time if the deposit has not been subiected to weathering, recent or ancient.

ACKNOWLEDGEMENTS

The assistance of Dr. L. G . Love, Sheffield, in the translation of the text by L. Bubenicek is gratefully acknowledged.

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

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REFERENCES

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STRAATEN (Editor), Deltaic and Shallow Marine Deposits. Elsevier, Amsterdam, pp. 3 19-322. PETRASCHECK, W. E., 1961. Lagerstattenlehre, 2 Aufl. Springer, Berlin, 374 pp. F. J., 1956. Sedimentary Rocks, 2 ed. Harper, New York, N.Y., 718 S. PETTIJOHN, PIEKARSKI, K., 1961. On the origin of the copper ore deposit at Miedziana Gora near Kielce. Acta Geol. Polon., 3: 43-68. POPOV,B. P., 1955. Au sujet des carbonates et des silicates dans les minerais de fer de la presqu'ile de Kertch. Tr. Inst. Geol. Nauk, Akad. Nauk Ukr. R.S.R., Ser. Petrogr., Mineral. Geokhim., 61: 97-100. P o S E P ~F., , 1902. The genesis of ore deposits. In: F. P O S E Pet~ al. (Editors), The Diagenesis of Ore Deposits, 2nd ed. Am. Inst. Mining Engrs, New York, N.Y., pp.1-281. POTTER, P. E. and PETTIJOHN, F. J., 1963. Paleocurrents and Basin Analysis. Springer, Berlin, 296 pp. POUSTOVALOV, L. V., 1940.Pe'trographie des Roches se'dimentaires. 1-2. Gostoptekhizdat, Moscow. L. V., 1959. Uber die Formation der Lagerstatten sedimentarer Bodenschatze. POUSTOVALQV, Eclogae Geol. Helv., 51(3): 712-716. PUCHELT, H. und MULLER,G., 1964. Mineralogisch-geochemische Untersuchungen an Coelestobaryt mit sedimentarem Gefiige. In: G. C. AMSTUTZ(Editor), Sedimentology and Ore Genesis. Elsevier, Amsterdam, pp. 143-1 56. RAGUIN,E., 1949. Ge'ologie des Cites Mineraux. Masson, Paris, 641 pp. RAMBERG, H., 1952. The Origin of Metamorphic and Metasomatic Rocks.1. Univ. Chicago Press, Chicago, Ill., 317 pp. RAMDOHR, P., 1953. Neue Beobachtungen an Erzen des Witwatersrands in Siidafrika und ihre genetische Bedeutung. Abhandl. Deut. Akad. Wiss. Berlin, Math.-Naturw. Kl., 3: 43 pp. RAMDOHR, P., 1960. Die Erzmineralien und ihre Verwachsungen.Akademie Verlag, Berlin, 1089 S. RENTZSCH, J., 1963. Zur Entstehung der Blei-Zink-Kupfer-Lagerstatten in Triassischen Karbonatgesteinen des Nordwestbalkans. Freiberger Forschungsh., C, 166: 100 pp. RICHTER-BERNBURG, G., 1953. Uber salinare Sedimentation. 2. Deut. Geol. Ges., 105: 593-645. RITTENHOUSE, G., 1959. Presidential address, 33rd Annual Meeting of the Society of Economic Paleontologists and Mineralogists, Dallas, Texas, March, 1959. Bull. Am. Assoc. Petrol. Geologists, 43: 1500-1502. ROUTHIER,P., 1963. Les Gisements me'tallij2res. Ge'ologie et Princiges de Recherche. Masson, Paris, 1282 pp. SANDER, B., 1936. Beitrage zur Kenntnis der Anlagerungsgefiige. Mineral. Petrog. Mitt., 48: 27. SANDER, B., 1948-1950. Einfuhrung in die Gefugekunde der Geologischen Korper. Springer, Wien 1: 215 S., 2: 409 S. SCHACHNER-KORN, D., 1954. Ein Wachstums-und ein Rekristallisationsgefiige in Bleiglanz aus einer rheinischen Lagerstatte. Mineral. Petrog. Mitt., 4 (Sander - Festband): 111-1 16. SCHIDLOWSKI,M. und TRURNIT, P., 1966. Drucklosungserscheinungen an Gerollpyriten aus den Witwatersrand-Konglomeraten. Ein Beitrag zur Frage des diagenetischen Verhaltens von Sulfiden. Schweiz. Mineral. Petrog. Mitt., 46: 332-342. SCHMITT,H. (Editor), 1962. Equilibrium Diagrams for Minerals at Low Temperature and Pressure. Geol. Club Harvard, Cambridge, Mass., 199 pp. H.-J., 1964. Facies differentiation and controlling factors for the depositional leadSCHNEIDER, zinc concentration in the Ladinian gyosyncline of the eastern Alps. In: G. C. AMSTUTZ (Editor), Sedimentology and Ore Genesis. Elsevier, Amsterdam, pp.2947. SCHNEIDERH~HN, H., 1962. Erzlagerstiitfen, 4 Aufl. Fischer, Stuttgart, 371 pp. SCHULLER, A., 1958. Die Metallisation im Kupferschiefer und Dolomit des Unteren Zechsteins in den Bohrungen Spremberg 13 El57 und 3/54. Geologie(Berlin), 7(3-6):651-675. SCHULZ,O., 1960. Beispiele fur synsedimentare Vererzungen und paradiagenetische Formungen im alteren Wettersteindolomit von Bleiberg-Kreuth. Berg- Hiittenmiinnische Monatsh., Monfan. Hochschule Leoben, 105(1): 1-1 1 . SCHWARTZ, D., 1957. Die Bakterien des Schwefelkreislaufes und ihre Lebensbedingungen. Freiberger Forschungsh., C,44: 5-1 3. SEIBOLD, E., 1955. Zum Phosphat-, Eisen- und Kalkgehalt einiger Horizonte des siiddeutschen Jura, Geol. Jahrb., 70(1955): $77-610. VAN

414

G. C. AMSTUTZ AND

I,. EUEENICEK

SELDL,K., 1958. Dolomitisierung und Erzbildung in Karbonatgesteinen unter der Einwirkung von Salzsolen. Neues Jahrb. Mineral., Monatsh., 1958: 25-55. SELLEY, R.C., 1964. The penecontemporaneous deformation of heavy mineral bands in the Tor(Editor), ridonian sandstone of northwest Scotland. In: L. M. J. U. VAN STRAATEN Deltaic and Shallow Marine Deposits. Elsevier, Amsterdam, pp.362-367. 0.V., 1959. Metamorphization of the chemical composition of muddy waters in the SHISHKINA, Black Sea. In: N. M. STRAKHOV (Editor), Toward Knowledge of Diagenesis of Sediments (Symposium). Izd. Akad. Nauk S.S.S.R., Moscow, pp.29-50. SHROCK, B. R., 1948. Sequence in Layered Rocks. McGraw-Hill, New York, N.Y., 507 pp. SJEMES, H., 1961. Betrachtungen zur Verformungund zum Rekristallisationsverhaltenvan Bleiglanz. Thesis Technische Hochschule, Aachen, 250 S. SIMPSON, G. G., 1964. The glorious entertainer. Science, 144(3614): 38-39. (Redakteur), ffundbuch der MikroSTACH,E., 1952. Braunkohlenmikroskopie. In: H. FREUNLI skopie in der Technik, 2(1): 483-686. STANTON, R. L., 1958. Abundances of copper, zinc, and lead in some sulfide deposits. J. Geol., 66: 484-502. STANTON, R.L., 1964. Textures of stratiform ores. Nature, 202(4928): 173-174. R. L. and BAAS-BECKING,L. G. M., 1962. The formation and accumulation of sediSTANTON, mentary sulphides in seaboard volcanic environments. Koninkl. Ned. Akad. Wetenschap., Proc., Ser. B, 65(3): 236-243. STARIKOVA, N. D., 1959. Organic matter in the liquid phase in Black Sea deposits. In: N. M. STRAKHOV (Editor), Toward Knowledge of Diagenesis of Sediments (Symposium). Izd. Akad. Nauk S.S.S.R., Moscow, p.72. STEINBRECHER, B., 1959. Die Sedimentation im Saaletrog im Bereich des ostlichen Harzvorlandes wahrend des Zechsteins 1 und 2. 2.Angew. Geol., 9: 381-385. STEWART, F. H., 1951. The petrology of the evaporites of the Eskdale, No.2 coring, east Yorkshire.2. The middle evaporite bed. Mineral. Mag., 29: 445475. STEWART,F. H., 1955. Deposition and metasomatism of salt deposits. Inter-Univ. Geol. Congr., 3rd, 1955, Durham, Proc., 46 pp. N. M., 1953. La diagenhe des skdiments et son importance pour la metallogenese STRACHOV, skdimentaire. Izv. Akad. Nauk S.S.S.R., Ser. Geol., 5 : 12-19. STRACHOV, N. M., 1956. Vergleichendes lithologisches Schema authigener Sedimentbildung in den Meeresbecken. Z. Angew. Geol., 2-3: 119-130. STRACHOV, N. M., 1959. Schkma de la diagenkse des dkp6ts marins. Eclogae Geol. Helv., 51(3): 76 1-767. SUJKOWSKI, ZB.L., 1958. Diagenesis. Bull. Am. Assoc. Petrol. Geologists, 42: 2692-2717. TAWITZ, K. CH., 1955. ober Sedimentation, Diagenese, Metamorphose, Magmatismus und die Entstehung der Erzlagerstatten. Chem. Erde, 17: 104-164. J. H., 1949. Petrology of the Northhampton Sand Ironstone Formation. GeoZ. Surv. Gt. TAYLOR, Brit. Mem. Geol. Surv. Gt. Brit., Engl. Wales, 1949: 111 pp. TAYLOR, J. H., 1955. Concentration in sediments. In: Natural Processes of Mineral Concentration -Inter-Univ. Geol. Congr., 3rd, 1955, Durham, Proc., pp.15-20. TEICHM~LER, M., 1952. Die Anwendung des polierten Diinnschliffes bei der Mikroskopie von (Redakteur), Handbuch der Mikroskopie Kohlen und versteinerten Torfen. In: H. FREUND in der Technik. I I ( 1 ) . Mikroskopie der Steinkohle, des Kokses und der Braunkohle, 2(1): 235-310, TEICHM~~LLER, M. und T E I C ~ L L E R., R , 1958. Inkohlungsuntersuchungenund ihre Nutzanwendung. Geol. Mijnbouw, 20: 41-66. TEODOROVICH, G. I., 1958, 1961. Authigenic Minerals in Sedimentary Rocks. Consultants Bur., New York, N.Y., 120 pp. THIESSEN, G., 1945. Forms of sulfur in coal. In: H. H. LOWRY(Editor), Chemistry of Coal Utilization. 1.Wiley, New York, N. Y., p.430. THOMAS, D. W . and BLUMER, M., 1964. Pyrene and fluoranthene in manganese nodules. Science, 143: 39. Tocmm, M. S., 1956. Geochemistry of authigenic siderites. Vopr. Mineralog. Osad. Obrazov., L'vovsk, Gos. Univ., 3 4 : 203-21 1.

DIAGENESIS IN SEDIMENTARY MINERAL DEPOSITS

475

TOPKAYA, M., 1950. Recherches sur les silicates authigenes dans les roches skdimentaires. Bull.

Lab. Gdol., Mindral. Gdophys. Musee Gdol. Univ. Lausanne, 97: 132 pp. G., 1963. Zur Geologie und Petrographie des Zechsteinkalks im Raum SprembergTZSCHORN, Weisswasser. Z. Angew. Geol., 9(11): 561-568. VOIGT,E., 1962. Friihdiagenetische Deformation der turonen Plaenerkalke bei Halle/Westfalen, als Folge einer Grossgleitung unter besonderer Berucksichtigung des Phacoid-Problems. Mitt. Geol. Staatsinst. Hamburg, 31 : 146-275. VONBURGER, K., 1962. Die Kaolin-Kohlentonsteine der unteren und mittleren Essener Schichten des Westfal B im mittleren Ruhrrevier. Fortschr. Geol. Rheinland Westfalen, 3(2): 563: 580. VONCOTTA,B., 1870. A Treatise on Ore Deposits, 2 ed. Van Nostrand, New York, N.Y., 575 pp. VON ENGELHARDT, W., 1960. Der Porenraum der Sedimente. Mineralogie und Petrographie in Einzeldarstellungen.1. Springer, Berlin, 207 S. VON ENGELHARDT, W., 1961. Zum Chemismus der Porenlosung der Sedimente. Bull. Geol. Znst. Univ. Upsala, 40: 189-204. VON GEHLEN, K., 1960. Die rontgenographische und optische Gefugeanalyse von Erzen, insbesondere mit dem Zahlrohr-Textur-Goniometer.Beitr. Mineral. Pefrog., 7 : 340-388. VONGRODDECK, A., 1879. Die Lehre von den Lagerstatten der Erze. Von Veit, Leipzig, 351 S. VON GUMBEL,C. W., 1868. Geognostische Beschreibung des Ostbayrischen Grenzgebirges. 1-3. Fischer, Kassel, 700 S. VON GUMBEL,C. W., 1883. Beitrage zur Kenntnis der Texturverhaltnisse der Mineralkohlen. Munchen, S.109-2 16. VONG U ~ E LC., W., 1888. Grundziige der Geologie. Fischer, Kassel, 1144 S. WALPOLE, B. P., 1961. Research on syngenetic mineralization. In: Syngenesis in Ore Deposition -Symp. A.N.Z.A.A.S., Brisbane, 1961, pp.11-13, unpublished. WALTHER, J., 1894. Einleitung in die Geologie als Historische Wissenschaft. 3. Lithogenesis der Gegenwart. Fischer, Jena, 1036 S. WELLER, J. M., 1959. Compaction of sediments. Bull. Am. Assoc. Petrol. Geologists, 43: 273-310. T. S., 1955. The biosphere as an agent in concentration of elements. Inter-Univ. Geol. WESTOLL, Congr., 3rd, 1955, Durham, Proc., pp.28-33. H., TURNER, F. J. and GILBERT, C. M., 1954. Petrography. Freeman, San Francisco, WILLIAMS, Calif., 406 pp. WILSON,A. F., 1961. Two problems for both syngeneticist and magmatist at Mount Isa. In: Syngenesis in Ore Deposition-Symp. A.N.Z.A.A.S., Brisbane, 1961, pp.28-32. unpublished E. D., 1959. Alkalinity and biogenic elements in the seabottom deposits in the north ZAITSEVA, (Editor), Toward Knowfedge of Diagewestern part of the Black Sea. In: N. M. STRAKHOV nesis of Sediments (Symposium). Izd. Akad. Nauk S.S.S.R., Moscow, pp.51-71. ZIMMERMANN, R. A. and AMSTUTZ, G. C., 1961. Sedimentary structures in the Arkansas barite belt. Ann. Meeting Geol. SOC.Am., Cincinnati, 1961, Abstracts-Bull. Geol. SOC.Am., 68: 306-307. ZIMMERMANN, R. A. and AMSTUTZ,G. C., 1964a. Small-scale sedimentary features in the Arkansas (Editor), Sedimentology and Ore Genesis. Elsevier, barite district. In: G. C. AMSTUTZ Amsterdam, pp.157-163. ZIMMERMANN, R. A. und AMSTUTZ,G. C., 1964b. Die Arkansas- Schwerspatzone, neue sedimentpetrographische Beobachtungen und genetische Umdeutung. 2. Erzbergbau Metallhiittenwesen, 17(7): 365-371. ZIMMERMANN, R. A. und AMSTUTZ,G. C., 1967. Diagenetische Texturen in den Erzen von Meggen. im Druck. ZOBELL,C. E., 1942. Changes produced by microorganisms in sediments after deposition. J . Sediment. Petrol., 12: 127-130. ZOBELL,C. E., 1946. Studies on redox potential of marine sediments. Bull. Am, Assoc. Petrol. Geologists, 30: 477-51 3,

Chapter I0

DIAGENESIS OF SUBSURFACE W A T E R S EGON T. DEGENS AND GEORGE V. CHILINGAR

Division of Geological Sciences, California Inrtitute of Technology, Pasadena, Calif. (U.S. A.) University of Southern California, Los Angeles, Cali/ (U.S.A.).

SUMMARY

In the present chapter, the writers are concerned with the origin and geochemical make-up of subterranean waters. The chemical properties of surface waters are only considered as long as they have some bearing on the subsurface water cycle. Aside from the common electrolytes present in natural waters, the nature and geological significance of dissolved organic molecules, and stable and radioactive isotopes are briefly discussed here. Special consideration is given to the mechanisms that in one way or another may alter the geochemistry of water during its residence in the lithosphere. Some ideas on the significance of water studies for the elucidation of the petroleum problem are also briefly outlined.

INTRODUCTION

In studies of rocks and minerals one has always to consider the action of water. This is so because water is (I)the main agent during physical and chemical weathering, (2) the transporter and carrier of matter in ionic, colloidal and the particulate state, and (3) the environment of life processes. Furthermore, water makes up two-thirds of all living matter and is essential in balancing the structure of a great number of sedimentary minerals. Namely, life and most sediments would not exist on earth, without water being available. As long as only rocks from surface environments were available for hydrogeochemical studies, no conclusive information was available regarding the nature of fossil waters syngenetically entrapped in ancient sediments. The original water, that had “survived” compaction and diagenesis, has been gradually replaced by atmospheric and biogenic gases or meteoric water at the time the sediments were uplifted or became exposed to the earth surface by tectonic activities or erosional processes, respectively. As a result of petroleum exploration, however, geochemists now have access, to some of these interstitial solutions, and there is no longer a physical limitation on geochemical studies of deeply buried waters. California Institute of Technology, Division of Geological Sciences, Contribution No.1230.

478

E. T. DEGENS AND G. V. CHILINGAK

From the data so far obtained, one can infer that there are porous sediments even at depths of several thousand meters (VONENGELHARDT, 1960, 1961; MEADE, 1963). The pore spaces are occupied by waters rich in electrolytes and occasionally by natural gases and petroleum. The question therefore arises: Where do the brines ultimately come from? In geosynclines, where the rate of deposition is quite rapid, large quantities of waters are continuously extracted from the hydrosphere during sedimentation. Most of the buried water thereby occupies the pore spaces of sediments. Recent muds, for example, may contain up to about 80% water by volume; but upon compaction of the strata, the connate waters’, as they are often termed, will be gradually expelled to the next environs. The speed at which the water is released from the original bed rock not only depends on overburden pressure, but is also a function of mineral composition, texture, and structure of the sediment. It has to be emphasized, however, that all the deeply buried waters so far available for geochemical studies are derived from highly permeable sediments such as sandstones and certain carbonates. Thus, there is a lack of knowledge concerning the geochemistry of waters present in shales. But with the recent advances in pressure and dilution techniques for extraction of waters from clay minerals and shales, hydrogeochemical data which are being accumulated now throw considerable light on the chemical make-up of waters in less permeable rock materials (LOMTADZE, 1954; VONENGELHARDT, 1960, 1961; SIEVEK, 1962; RITTENBERG et al., 1963; VONENGELHARDT and GAIDA,1963; RIEKE et al., 1964).

CLASSIFICATION AND CHEMICAL COMPOSITION OF SUBTERRANEAN WATERS

Waters can be classified in a number of ways. Most commonly they are grouped according to ( I ) origin in terms of meteoric, connate, or juvenile waters, (2) chemistry, e.g., bicarbonate, sulfate, or chloride waters, and (3) total salinity, i.e., fresh water, salinized water, or brine water. Many chemical classifications have been proposed or discussed by TOLSTIKHIN (1932), DESITTER (1947), DUROV (1948), SULIN (1948), VASSOEVICH (1954), CHEBORATEV (1955), KREJCI-GRAF et al. (1957), GORRELL (1958), RAINWATER and WHITE(1958), CHAVE(1960), and EREMENKO (1960); to mention just a few investigators. The subject of classification of waters has been reviewed by CHILINGAR (1 957, 1958) and CHILINGAR and DEGENS (1964). The term connate water is frequently employed with different notations. Some investigators use this term for all waters syngenetically incorporated in sediments no matter whether they are of fresh-water or marine origin. In the present context, however, only sea water will have the prefix “connate”.This is done because all fresh waters in sediments are ultimatelymeteoric in origin and cannot be differentiated into those syngenetically entrapped and those infiltrated later on into the rock strata.

479

DIAGENESIS OF SUBSURFACE WATERS ,4100 c o

YlOO

I I

100

I

0_ - -

Na+K

*)

0

d

+

0

V

+ V

v)

*)

0 0

r 0

Ca t M g

100

Fig. 1. Graphical representation of chemical composition of various oil-field waters from the 1954, p.212.) U.S.S.R. (Modified after TOLSTIKHIN,

100 Na

50 Ca

Fig.2. Cations in petroleum brine waters of the United States. ( I ) Woodbine Sand, Texas (Cretaceous); (2) California (Tertiary); (3) Kansas and Oklahoma (Paleozoic); (4) Appalachian (Missis1947, sippian); (5)Appalachian (Upper Devonian); and (6) Arkansas (Jurassic). (After DESITTER, 1960.) and VON ENGELHARDT,

480

E. T. DEGENS AND G . V. CHILINGAR

As an example, the classification scheme of N.I. Tolstikhin (in: VASSOEVICH, 1954, p. 112) is presented in Fig. 1. It is based principally on the distribution of the most abundant cations (Na+, Mg2+, Ca2+) and anions (HCOs-, C1-, Sod2-). In following this scheme, one may chemically classify a water, for instance, as sodiumbicarbonate type, or calcium-chloride type, depending on the position the prospective water occupies in the diagram. The positions of representative oil-field waters of the U.S.S.R. are indicated by the stippled areas. Additional information on the chemistry of petroleum brine waters and their genetic coefficients is presented in Tables I, I1 and Fig.2.

CHEMICAL ALTERATIONS OF SUBSURFACE WATERS

Meteoric and connate waters represent most of the water present in the lithosphere. Juvenile waters are exceedingly rare even in areas of recent volcanic activities. Meteoric water contains initially an average of only a few hundred p.p.m. of salts as compared to 35,000 p.p.m. in the ocean. Furthermore, sea water is uniform in chemical composition, whereas composition of meteoric water can vary considerably. But as soon as water comes in contact with minerals and organic matter, reactions and interactions take place that may cause ( I ) desalting of water, (2) concentration of salts, or (3) preferential increase or decrease of some dissolved mineral and organic species. This diagenesis or metamorphosis of water is often so pronounced that the present chemistry no longer reflects the original chemical make-up of the water at the time of deposition or infiltration. In principle, both genetic types of waters will behave in a similar fashion in the same diagenetic set-up. But, inasmuch as ocean waters are chemically identical, in contrast to meteoric waters, the degree and type of diagenetic alteration of subsurface waters can best be evaluated in cases of former sea waters. The widespread occurrence of marine sediments in most of the sedimentary basins, which are of interest to the petroleum industry, suggests that the pore waters presefit in the petroleum-bearing strata are ultimately derived from the sea. Similarities in the solutes also point in this direction, if one believes that the presentday ocean is a representative geochemical standard for all geologic ages. On geological and biogeochemical grounds, many scientists reasonably assume that steady-state conditions were already reached in the ancient sea as far back as late Precambrian time. Yet there are systematic chemical differences developed both qualitatively and quantitatively between the ancient and modern connate waters. Magnesium, highly abundant in the ocean, is present only in minor amounts in oil-field waters, whereas the opposite is true for calcium. Calcium-chloride waters are actually not formed in any surface environments, whereas they are widespread among the petroleum brines. This feature can possibly be linked to dolomitization processes. Magnesium may also proxy in chlorites or certain mixed layer clay

TABLE I

pp.481-486

CHEMICAL COMPOSITION OF SOME OIL-FIELD WA‘ERS

(After EREMENKO, 1960, pp.152-157) Petroliferous region and deposit

Stratigraphy

CI-

Western Ukraine, Borislav Sakhalin Island, Okha, bore-hole 92 Central Asia (Fergany) Palvantash Western Turkmeniya, Cheleken Emba, southeastern part of Munayli Azerbayjan, Neftechala Tersko-Dagestan, Izberbash Kuban-Black Sea, Khadyzhy Ural-Volga, Twmw Ural-Volga, Ishimbay Ural-Volga, Krasnokamsk Ural-Volga, Krasnokamsk

Chemical composition (% equiv.)

Chemical composition (mg-equiv.10

Menilit horizon

HCOa-+ c0a2-

S042-

Mga+

Na++K+

Cr

A1

Aa

NaICI

(Na-CoI

so4

(CI-Na)/ Mg

24.68

-

0.06

0.75

-

5.18

-

75.76

4.48 0.04

4.93 0.83

-

190.40

-

a3

0.03

9.99

2.38

37.63

75.26

0.2 60.72

34.4 5.0 0.557

1.4 196.7

0.8 81.91

46.9 1,063.27

98.2 2,683.76

9.68 647.72

0.20 2.26

35.03 5.09 0.02

1.43 7.33

0.81 3.05

47.76 39.62

19.76 79.24

1.94

0.59

0.16

49.25

96.12

-

2.38

1.5

1.02

1.42

0.87

47.11

92.68

-

2.74

4.58

1.13

4.13

34.71

69.42

30.54

-

0.04

0.70

-

8.8 18

3.9

1.1

328.2

666.4

48.06

8.0

3.1

1.9

104.2

218.4

42.31

4.03

3.66

49.7

0.28

0.02

1

708

0.5

3.4

30.4

526.7

155.7

2.1

3,598

79.7

1,224

102.9

3,994

7.6

-

rC1-rNa rMg

> 1);

(2) Sodium bicarbonate (3)Sodium sulphate type

rC1-rNa rMg

< 1).

-

12.9

339.3

(After SULIN,1948; see CHILINGAR, 1958, p.168.)

Sa

0.01

Upper Maykop Lower Permian Lower Permian Lower Permian

< 1 and

Na++K+ SI

49.96

51.21

(4) Magnesium chloride type

Mga+

9,819.34

5.51


Caa+

3,695.44

155.75

~

coe

233.49

“B” formation

( ;E

HCOa-+

980.74

3,152

(I) Calcium chloride type waters

so4a-

3.0

Stratum I

Upper Devonian

CI-

Genetic coefficients1

1.28

4,905.39

3 strata, 9.5 VIII formation Cretaceous 1,280.6 (chalk) Red Bed forma320.3 tion, depth of 1,248 m Senonian 92.4

Cazf

Characteristic coefficients (Palmer’s salinity and alkalinity values)

262

2,201

6,342

11.6

20.12

-

1.34

-

(SOr.lOO)/

CI

0.02

2.66

2.01 4.74

-

0

-

9.52

3.63

0.56

-

3.55

2.57

0.14

208.73

424.94

36.65

1.3

12.05

0.40

0.48

49.12

75.9

-

22.34

1.76

1.34

16.9

295.9

686.4

49.43

0.07

0.50

4.43

2.46

43.11

86.22

12.78

-

1.00

0.87

41.1

46.2

591

1,369

38.48

11.37

0.15

3.00

3.38

43.62

87.24

12.46

-

0.30

1.13

7.0

96

46

3,543

7,370

48.82

1.08

0.1

1.31

0.62

48.07

96.14

3.66

-

0.20

0.98

-

1.22

2.21

6.3

73.4

51

1,208.8

2,666

45.9

3.86

0.24

2.75

1.91

45.34

90.68

8.84

-

0.48

0.99

-

0.29

8.4

832.3

254

2,865

7,904

49.9

0.10

10.53

3.22

36.25

12.26

27.74

-

-

0.72

4.38

0.20

1.71

2.03

-

9.59

0.45

d

-

29.5

II COMPARISON OF MARINE AND OIL-FIELD WATERS

M x m m v , 1956, p.303) Palmer's salinity and alkalinity values1

Percentequivalents (r) Na+fK+

1 ' 5

s2

A

a

0.69

49.23

89.76

-

10.52

1.52

0.81

1.01

48.18

60.06

-

36.38

3.56

0.42

4.44

5.35

40.21

80.40

18.80

Cl-

S042-

HCOa-

Caz+

Mgz+

KarachukhwOIKTI)

-

0.39

5.82

0.08

BalakhanvSabunchi-

31.77

-

19.23

49.58

-

R o m y (IIK)

(W stratum)

Neftechala

Sea

49.74

0.23

0.03

11.02

4.70

3428

66.96

32.98

45.22

4.62

0.16

1.77

8.92

39.31

78.6

21.1

34.58

14.51

0.91

3.81

13.81

32.28

64.56

33.62

XNa+-f-K+>XC1-SS04a; a-+rS04z-]; 2[r(Na+ K+)-r(Cl-+ S042-)]; = 2(rCa2++rMgz+).

+

ZNa++K+<X~-+S04z; arNa++rK+l; 2[r(CI-+S048-jr(Na++K+)I; 2(rHCO~+rcOa&).

-

-

(3)If CCI-+S042->XNa++K++Caz++Mgz+, SI= 2r(Na++K+); SZ = 2r(Caa+-tMg2+);

S3 = 2[r(C1-+SOa2-)--r(Na++K++Caa++Mgz+)l.

0.80 0.06

0.3 1.82

488

E. T. DEGENS AND G. V. CHILINGAR

minerals for alkalies, iron and calcium. The variations established in the cases of the Na/K and Ca/Na ratios between fossil and present sea waters can also be accounted for by adsorption and exchange phenomena. Another example is the decrease in sulfate content, which can be linked to microbial activities. The other alternative is to ascribe the decrease in sulfate content to inorganic precipitation. Inasmuch as many brines have greater than 10% salinity,the solubility product of calcium sulfate is commonly exceeded, and consequently results in the formation of anhydrite or gypsum. In summary, the elemental variations between ocean waters and petroleum brine waters have reasonable geochemical explanations. More difficult, however, is to account for the increase in salinity up to tenfold of connate waters remaining in the compacting rock formation. There are two principal ways by which an increase in salinity can be achieved: ( I ) by evaporation in a surface environment, and (2) by compaction of sediments during their burial, with the interstitial water being squeezed out. Experiments have shewn that clays upon moderate compaction in the 100 p.s.i. range release water somewhat enriched in electrolytes relative to the original interstitial water, whereas a gradual increase in pressure from 100 up to 200,000 p.s.i. yields water that exponentially decreases in electrolyte content (VON ENGELHARDT, 1961; VONENGELHARDT and GAIDA,1963; RIEKEet al., 1964). In order to illustrate the geochemical alterations interstitial waters undergo during compaction of the rock strata, the data by &EKE et al. (1964) are presented here (Tables 111,IV; Fig.3, 4). It can be seen that the mineralization of solutions squeezed out during the different stages of compaction is a function of overburden pressure. Namely, the resistivity of squeezed-out solutions increases with increase TABLE III THE PERCENTAGE INCREASE IN THE RJX4JSl'MTY OF SOLUTIONS SQUEEZED OUT OF MARINE MUD WITH INCREASING OVERBURDEN PRESSURE

(After RJEKEet al., 1964) Overburdenpressure (p.s.i.1

Increase in resistivity (%), as compared to resistivity of solution squeezed out at 500 p.s.i.

2.3- 6.5 3.5-15.2 10.5-19.6 16.3-32.0 18.6-37.0 23.245.6 25.6-48.0

489

DIAGENESIS OF SUBSURFACE WATERS

TABLE IV AND CONTENT OF VARIOUS IONS IN SOLUTIONS SQUEEZED OUT AT DIFPBRBNT OVERBURDEN PRBSSURBS FROM MONTMORUONITB CLAY (NO. 25, UPTON, WYO.) SATURATED WITH SEA WATER

MplERALIzAnON

(After RreKE et al., 1964) Overburden pressure (p.s.i.) #

Remaining mois- Percentage of the concentration in solution squeezed out at ture content (% 100 p.s.i. of& weight) Total c1Nat Caa+ Mgs+ so$- mineralization

81 60 50 40 32 21 14

100 400

1,OoO

3,000

10,Ooo 40,000 9OYOOO

100 91-95 7w3 40-82 36-61 36l

-

100 93-95 841 25(7)-87

371 -

100 75-84 67 50-62

251 -

100

100

-

-

84-95

80 60l

-

67-81

-

I

38l

-

-

100

-

201

Only one trial.

1

I

I

i

CI'

so:0

OIOO

tpoo

Remaining

moisture content

r0,ooo

l00,000

Overburden pressure in p s i .

Fig.3. Content of various anions in solutions squeezed out at different overburden pressures from a sea-water saturated montmorillonite;dashed line represents remaining moisture (%dry weight) content. (Modified after RJEKB et al., 1964.)

490

E. T. DEGENS AND G. V. CHILINGAR

in overburden pressure (Table 111). The degree of reduction in mineralization with enforced compaction and the content of various ions in the free aqueous phase obtained from sea-water-saturated montmorillonites is presented in Fig.3,4, and Table IV. All these results support the findings of KRYUKOV et al. (1962) that the mineralization of interstitial solutions in shales is less than that of waters in the associated sandstones. It appears that, in general, the concentrations of the principal cations and anions decrease at about the same rate under pressure. This further suggests that (I) the ions being removed represent interstitial electrolyte solution and do not include the desorbed cations, and (2) the analysis for a single ion in the effluent (for example, C1-) might reveal as much as the analysis for all of the ions. Studies on interstitial waters incorporated in recent marine sediments of the Black Sea (SHISHKINA, 1959) and in the Atlantic Ocean (SIEVER et al., 1961) indicate that the buried waters are geochemically different from the sea water above. Shishkina's data, presented in Table V, show a general decrease in electrolyte concentration with depth, except for calcium. Most pronounced are the changes with depth below depositional interface in the case of sulfate. The findings of SEVERet al. (1961) indicate a higher salinity for the interstitial waters when compared to sea water salinity (Fig.5). In the light of the foregoing and ensuing discussions, it is questionable,

Na+ EI Ca2+ Mg2+

h

20_1 0 100

1p00

10,000

100,000

Overburden pressure in psi.

Fig.4. Content of various cations in solutions squeezed out at different overburden pressures from a sea-water saturated montmorillonite, (Modified after RIEKEet al.. 1964.)

49 1

DIAGENESIS OF SUBSURFACE WATERS

TABLE V CONTENT OF vmous IONS IN SOLUTIONS SQUEEZED OUT FROM BLACK SEA S B D I M E ~ AT, P.S.I. (mg-equiv./l). CORE ~ 0 . 1 3 ,1,301 CM LONG, SEA DEPTH OF 2,122 M.

5,700

(After SHISHKINA, 1959, p.37) Depth in cm

Cl-

soh2-

Na+

Ca2+

Mg2+

Kf

Black Seal 0-18 139-180 400-430 576610 763-780

271 349 337 278 231 193

27.2 33.4 26.2 2.1 3.8 3.0

232 302 292 216 175 131

12.3 9.2 14.7 25.7 31.8 39.7

53 66 61 40 n.d. 27

5.3 7.5 4.8 1.5 n.d. trace

(1953, p.269). Composition of sea water changes with depth; for Average values, after ALEKIN 1953, example, C1 content is 10.27%,at the surface and 12.64%,at a depth of 2,000 m (ALEKIN, p.270).

however, whether the observed differences between the ocean and connate waters are real in the sense that they reflect the true chemistry of the interstitial solutions as a whole. It is possible that the distribution pattern is caused-at least partlylby the compaction procedures employed during the extraction of the water phase. Namely, the waters of the Black Sea sediments were obtained by high-pressure techniques ( 5,700 p.s.i.) and, therefore, ion-filtration by charged-net membranes must have been effective, causing the lower concentration of electrolytes in the waters squeezed out. That the deepest samples analyzed are the least mineralized ones is no contradictio in adjecto,inasmuch as this feature coincideswith the natural decrease in moisture content with increasing burial depth. The Atlantic samples, on the other hand, were subjected to pressures in the 100-200 p.s.i. range (SIEVER et al., 1961; SIEVER,1962).The somewhat higher salinities in the squeezedout solutions agree with the statement made before that pressures below 100-200 p.s.i. produce waters somewhat enriched in electrolytes relative to the original solution. In conclusion, it is suggested that data on the chemical composition of interstitial waters in marine sediments obtained by high- or low-pressure techniques have to be used rather cautiouslyfor any interpretation regarding natural diagenesis of connate waters or paleosalinities. N

Some of the chemical alteration may have been caused by microbial activities, or the generally reducing environment in the case of the Black Sea. Reducing conditions, for example, may have been responsible for the low sulfate content in deeper buried interstitial solutions (Table V). One should also not lose sight of the possibility that chemistry of Black Sea waters was changing with time.

492

\.

B. T. DEGENS AND

(1.

V. CHILINGAR

bottom sso water

0OJ

160 240

sediment-wofsr interfoca

Stondord seo woter chlorinlty' 19.374%0

-

400 480 560 320

640- , . , ,

, ,

Fm.5.

at aL,

No such restriction exists in the case where the connate water is obtained by centrifugation methods. For example, marine sediments off southern California covering 15 million years span of sedimentation (EMERYand RITIZNBBRG,1952; R.~ENBERG et al., 1963) show no apparent change in the chemical composition of the interstitial solutions. The authors, therefore, concluded that changes in the chlorinity of the oceans since Middle Miocene are very minor. The same conclusion is suggested by the work of BRUEVICH (1957) and BRUWICH and ZAYTSFWA (1960) on interstitial solutions of some deep-water sediments in the northwestern part of the Pacific Ocean. ION-FfLTRATION BY CHARGED-=

CLAY MEMBRANES

Considering the fact that waters expelled at the initial stages of diagenesis at some pressures below approximately 100-200 p.s.i. are apparently only slightly enriched in salinity (10-20°A relative to the starting material, this mechanism cannot 80count for the high salt level up to values in the order of 30% salinity as found in some petroleum brine waters. Possibly, a more effective mechanism has to operate during diagenesis to produce the observed salinities. It has been proposed by ELLIS(1954), DAVIS(1955), WYLLIE (1955), McKm, VEY et al. (1957), VONENGELHARDT (1961), MCKELVEY and MILNB(1962), and BREDEHOEF~ et al. (1963) that buried waters may be subjected to ion-filtration by

493

DIAGENESIS OF SUBSURFACE WATERS

charged-net clay membranes. The atration of salt solutions through charged-net membranes has been suggested as a mechanism for producing fresh water from saline water. Shale beds in situ may be considered to be ideal membrane electrodes. The most suggestive argument for this assumption is the observed constancy of a “shale baseline” on “spontaneous potential” logs found in drill holes in every part of the world (WYLLIB, 1955). A quantitative theoretical treatment of the electrochemical properties of clays is given by the theory of membrane behavior of MeyerSevers-Teorell (DAVIS, 1955), which, according to calculations based on SP curve of e-logs, approximates the behavior of shales in situ in the earth. Considering a three-phase system consisting of two solutions of electrolytes separated by an intervening membrane, as represented by the following scheme: sol.

I

Is

membrane I

I I I

I

I

W’

sol.

I

I1

the Meyer-Sievers-Teorell theory is based on the following equation for the electrochemical potential E across the membrane: E=-

RT I1 t+-t-

s -t++t-

F I

d In c

where R is the gas content, T the absolute temperature, F the Faraday equivalent, and t the transference number which represents the relative amount of electricity carried by the ionic species, i, across a given plane. It is determined by the migration velocities, Ug,and concentrations, cz, of the ions at the given phase: tt

=

(2)

Usct/ZtUtct

so that: T1

RT ‘I U+c,- U-cE=d In c F I U+c++U-c-

s

(3)

~

Between the peripheral laminae of the membrane and the adjacent solution phases, i.e., between I and I’ and 11‘ and I1 two Donnan equilibria are presumed to exist, which may be expressed as: (fk2c2)I = cf+c+f-c-)I’

and (f.2c2)II

= cf+c+f-c-)II’

-

(4)

wheref represents activity coefficients. Inasmuch as f+ and f are not measurable, it is assumed that: (f*2)I

= cf+f-)I‘ and (f*2)II = (f+f-)II’

(5)

494 SO

E. T. DEGENS AND G. V. CHILINGAR

that: = (c+c-)I' and

(C2)'

(@)I1

=

(c+c-)II'

If A is the number of fixed unit charges on a negatively charged membrane, one may say that: C+

= C-

+A

(7)

where c+ and c- refer to the peripheral laminae of the membrane at I' and 11'. The integral of eq. 3. represents the sum of two external integrals, and an internal integral within and extending across the membrane, shown in the following scheme:

I

I'

CI

c+I' = c-I'+A

Liquid-junction potential

Donnan potential Donnan potential 1-1' =

Liquid potential 1'-11'

~

RT CI In F xl+A 2 ~

RT F

= __ u

In

xI+uA xII+uA

RT CII Donnan potential 11-11' = - -In ___ F xIIfA 2 where x

=

(4~2+A2)'fand u =

The total potential is thus:

u+- uu++u-

Donnan potential

DIAGENESIS OF SUBSURFACE WATERS

This equation has the mathematical properties that for A s CI or last term approaches zero, whereas the first term approaches In cI/cII. Thus, the equation for a “perfect” electrode is approached:

RT E=--lnF

CI

cn

495

6 1 ; the

(9)

On the other hand, if A d or cII, the first term approaches zero, whereas the last term approaches u In cI/cII. Consequently, eq. 8 is reduced to the ordinary liquid-junction potential:

These potentials, therefore, represent the upper and lower limits possible. From e-log calculations, one gets eq.9 and not eq.10; therefore, the conclusion is that shales function as perfect membranes. It follows that during compaction of clay containing sediments, the salt held back accumulates in the formation water retained in the strata. The process of salt removal or concentration depends on the large excess charge permanently attached to the clay membrane which prevents the passage of like-charged ions. In other words, the separation is effected because of the electrical properties rather than the size of the electrolyte. No such restrictions are placed on the water molecules. They move and therefore will pass the electrolytes. This process, therefore, should yield a lower salt level in the filtrate as compared to the original solution. Thus the salt is filtered by virtue of its electrolyticdissociation and the electricproperties of the clay membrane. ISOTOPE STUDIES OF INTERSTITIAL AQUEOUS SOLUTIONS

Aside from the electrolyte content, the stable isotopic composition of natural waters is another parameter to characterize a water. Deuterium and oxygen18 concentrations in meteoric surface waters vary by about 43 and 5.6 per cent, respectively, and are linearly related (FRIEDMAN, 1953; EPSTEIN and MAYEDA, 1953; DANSGAARD, 1953,1961; and CRAIG,1961). This comparison of the 180/160 and DH/H2 ratios shows that atmospheric precipitations normally follow a Raleigh process at liquid-vapor equilibrium. The atmospheric Raleigh process also explains why with higher altitudes and latitudes fresh waters become progressively lighter, whereas tropical samples show very small depletions relative to mean ocean water. Other factors which determine the isotope composition of meteoric waters have been elaborately discussed by EPSTEIN(1959), CRAJG(1961), and DANSGAARD (1961).

496

E. T. DEGENS AND G. V. CHLLINGAR

In contrast to meteoric waters, ocean waters are isotopically heavy and fall within a narrow range, i.e., 1 and 0.1% for deuterium and oxygen-18, respectively. Evaporation processes, however, strongly affect the 180/l60 and DH/H2 ratios of the water, because they cause a preferential depletion in the lighter isotopes H and 160, which become concentrated in the vapor phase. The remaining water consequently will be heavier, i.e., the D and 1 8 0 contents will show a relative increase. Compaction and filtration by charged-net clay membranes, on the other hand, should not noticeably influence at least the oxygen isotope ratios of waters. These processes are relatively slow and, inasmuch as only one phase is participating, yield no fractionation when equilibrium is reached. In this connection, it is noteworthy that hydrochemical studies, within the same aquifer (Nubian Series of the western Egyptian Desert), of the effect of long-range migration on the lsO/leO ratio of the water by KNETSCHet al. (1963) have shown that oxygen isotopes apparently are not fractionated during subsurface transportation. Over a distance of 700 miles, and at a depth of 500-2,OOO ft., the oxygen isotope ratios of water samples taken at intervals of 20-100 miles stayed within lx0,whereas the chemical composition fluctuated strongly in response to migration and diagenesis. In order to decide whether evaporation processes or compaction and ionfiltration are responsible for the concentration of electrolytes in petroleum brine waters, a study of the stable isotope distribution can be rather revealing. DEGENS et al. (1964) analyzed, for example, the oxygen isotope composition1 of a number of connate waters ranging in age from the Cambrian up to the Tertiary. Fig.6 shows that the 6 1 8 0 values of the highly saline petroleum brine waters do not deviate appreciably from the 6 1 8 0 of modern sea water2. Deviations from this mean value in some of the samples into the negative range of 6 1 8 0 are always well correlated to a decrease in salinity. This feature can be easily explained by effects of dilution with meteoric waters during migration of the brine, or perhaps by laterstage infiltrations as a result of a change in the geological settings by uplift, denudation, or other phenomena. The similarity between isotope characteristics of the brines and modern sea water leads immediately to the conclusion that the concentration of the inorganic salts has not been accomplished by syngenetic evaporation. Consequently, it probably occurred by processes of compaction and ion-atration by chargednet clay membranes. Slight deviations into the positive range of 8 1 8 0 values in some samples studied may have been caused by original evaporation in a surface environment, or by isotope equilibration with the surrounding mineral matter for millions of years (DEGENS and EPSTEIN,1962). Studies are presently under way to determine the deuterium content of the same brines used for the oxygen isotope work. The principal objective is to prove whether 180/160and DH/Hs are linearly related, as they are in most terrestrial surface waters. 6 = deviation relative to Chicago Belemnite standard (CRAIG, 1957).

x0

497

DIAGENJBIS OF SUBSURFACE WATJ3RS

351-

mean ocean water

g-

6'80 Fig.6. Comparison of salinity and oxygen isotope composition in waters associated with marine and fresh-water petroleum deposits. ( I ) Cambrian-ordovician (Oklahoma); (2) Devonian (Oklahoma); (3) Pennsylvanian (Oklahoma); (4)Tertiary (Texas); (5)Cretaceous (Colorado); et al., 1964.) and (6)Tertiary (Utah). (After DEGENS

Aside from the stable isotopes, the radioactive tritium and the carbon-14 content of the dissolved carbonate fraction produced by cosmic radiation, thermonuclear weapons, and nuclear industries is also of real usefulness for the elucidation of the origin and age of subsurface water samples (LIBBY,1959; BROECKER et al., 1960; EHHALT et al., 1963; MUNNICH, 1963).

ORGANIC MATTER IN OIL-FIELD BRINE5

A great variety of dissolved organic constituents are present in fossilized brines. They include humic acids, fatty and naphthenic acids, aromatic oxygen compounds amino acids, sugars, various heterocycles, and others. Naphthenic acids are more water soluble than are the calcium soaps; therefore, the maximum content of naphthenic acids occurs in sodium-bicarbonate (alkaline) type of waters (EREMENKO, 1960). The range in concentration is

498

E. T. DEGENS AND G. V. CHILINGAR

about 0.1-30mg-equiv./l. The oil-fieldwaters associated with heavy oils have higher content of naphthenic acids than the brines from light-oil deposits. Phenolic compounds are mostly found in waters associated with light oils, and may reach concentrations up to about 4 mg/l. The study of COOPER (1961) on the distribution of fatty acids in petroleum brine waters is of great significance, particularly in conjunction with the origin of oil. Whereas recent, and to a lesser degree ancient, sediments show a preference of fatty acids with even carbon numbers, in the case of the fatty acids in petroleum brine waters, concentration differences between neighboring odd- and even-numbered acids become very small. In addition, a nearly straight-linedecrease in relative abundance from 14C to 3OC acids is developed. Cooper tentatively suggested that the increase in abundance of odd-numbered acids with time apparently matches the generation of even-numbered paraffins (BRAY and EVANS,1961). This parallelism suggests related processes for the formation of these acids and paraffins. Of similar geochemical significance is the observation that the concentration et al., of amino acids in petroleum brine waters is a function of salinity (DEGENS +4

+2

I

A

mean

ocean water fO

to

-2 -

P

LQ

P

'

0 0%

+

0

0

0

I

-4-

-6

-

-8

-

-10

Tertiary iCarboniferous Devoniau marine A Ordovicion 0 Cambrian

o

0

I0 1

mean ocean water

60

80

100

120

140

160

180

200

Y / i amino acids

Fig.7. Comparison of amino acid distribution in petroleum brine waters of Paleozoic age and modern sea water. Q = Ocean water; 0 = petroleum brine water, adjusted to salinity of mean ocean water (3.5 %). Abbreviations:AL = alanine; ARG = arginine; ASP = aspartic acid; CYS = cystine; GLU = glutamic acid; GLY = glycine; HIS = histidine; LEU = leucine; LYS = lysine; MET = methionine; ORN = ornithine; PHE! = phenylalanine; PRO = proline; SER = serine; THR = threonine; TYR = tyrosine; VAL = valine. (After DEGENS et al., 1964.)

DIAGENESIS OF SUBSURFACE WATERS

499

1964). As the salts increase, so do the amino acids. Furthermore, if one adjusts the salinity of brine waters to that of today’s ocean (Fig.7) and applies the same calculation factors to the original amino acids values, the similarities between the amino acid spectra in the fossil brines and in recent sea water become striking. This may mean that amino acids were concentrated in the same fashion as the inorganic salts. Amino acids, however, do not occur in a free state, but are part of organic “heteropolycondensates” having quinones, phenols, and amino compounds as principal structural units. Such “heteropolycondensates~’are more commonly known as humie acids. ROLE OF WATER IN PETROLEUM FORMATION

The possible effect of subterranean waters on the formation of petroleum deposits is of considerable importance. It is generally agreed upon that water is the principal vehicle by which hydrocarbons and related substances are transported. Considerable disagreement, however, exists as to the state in which hydrocarbons are actually migrating. The degree of hydrocarbon solubility in water is different for various hydrocarbons. As a general rule, increase in molecular weight will decrease the solubility of hydrocarbons in aqueous solutions; and increase in salt concentration will increase their solubility. Marine-derived formation waters are known to produce salinities up to 30 %as a result of ion-filtration by charged-net clay membranes. The solubility of free hydrocarbons and allied substances will naturally be affected by these processes. An oil-droplet formation may be the consequence of chemical changes occurring during the migration of the formation water through the strata. It is also conceivable that during the transition from shales to more porous sediments (i.e., sandstones and limestones), or vice versa, formation waters release most of their hydrocarbons. There is no reason to assume that other organic constituents dissolved or emulsifiedin connate waters will not be affected by the outlined filtration processes. Perhaps kerogen-type material may be released when the solubility product of humic acids is exceeded. As a result, the organic matter content in sediments will increase. CONCLUSION

Subsurface waters undergo a considerable degree of chemical alteration during diagenesis. Most striking is an appreciable increase in salinity which may be a result of ion-filtration on charged-net clay membranes. The changes in the electrolyte spectrum may be caused by a variety of mechanisms such as dolomitization and chemisorption, pH-Eh relationships, or dissolution and precipitation phenomena. The oxygen-isotope distribution of the water is not noticeably affected by any of these processes operating during diagenesis. Increase in salinity is generally matched

500

E. T. DEGENS AND G. V. CHILINGAR

by an increase in some of the dissolved organic constituents, a feature which can possibly be linked to the same ion-filtration mechanism that produces the high salt concentrations. REFERENCES

AIEKIN,0. A.. 1953. Osnovy Gidrokhimii (Principles of Hydrochemistry). Gidrometeoizdat, Leningrad, 296 pp. (in Russian). as a clue to the recognition of BRAY,E. E. and EVANS, E. D., 1961. Distribution of n-parsource beds. Geochim. Cosmochim. Acta, 22: 2-15. BREDWOEIT,J. D., BLYTH,C. R., WHITE, W. A. and MAXEY, 0. B., 1963. Possible mechanism for concentration of brines in subsurfacx formations. Bull. Am. Assoc. Petrol. Geologists, 47: 257-269. BROEKER, W . S., GERARD, R., HEWING, M. and HEBZEN, B. C., 1960. Natural radiocarbon in the Atlantic Ocean. J. Geophys. Res., 65: 2903-2931. BRUEVICH, S. W., 1957. The salinity of the interstitialwaters (sediment solutions) of Okhotsk Sea. Dokl. Akad. Nauk S.S.S.R., 113: 387-390 (in Russian). BRUEVICH, S. W. and ~AYTSEVA, E. D., 1960. On the chemistry of the sediments of the northwestern part of the Pacific. Tr. Znst. Okeanol., Akad. Nauk S.S.S.R., 42:348(inRussian). CHAVE, K. E., 1960. Evidena on history of sea water from chemistry of deeper subsurface waters of ancient basins. Bull. Am. Assoc. Petrol. Geologists, 44: 357-370. CHEJKJTAREV, 1.1., 1955. Metamorphism of natural waters in the crust of weathering. 1-3. Geochim. Cosmochim. Acta, 8: 22-48; 137-170; 198-212. CHXLWGAR, G. V., 1957. Soviet methods of reporting and displaying results of chemical analyses of natural waters and methods of recognizing oil-field waters. Trans Am. Geophys. Union, 38: 219-221. CHILINGAR, G. V., 1958. Chemical composition of oil-field waters from Apsheron Peninsula, Azerbaidzhan. S.S.R: a summary. Geochim. Cosmochim.Acta, 14: 168-178. CHILMOAR, G. V. and DEGENS,E. T., 1963.Notes on chemistry of oil-field waters. Bol. Asoc. Mexicana Geol. Petrols., 15(7-8): 177-193. CooPBR, J. E., 1961. Fatty acids in recent and ancient sediments and petroleum reservoir waters. Nature, 193: 744-746. CRAIO.H., 1957. Isotopic standards for carbon and oxygen and correction factors for massspectometric analyses of carbon dioxide. Geochim. Cosrnochim. Acta, 12: 133-149. CRAIG,H., 1961. Isotopic variations in meteoric waters. Science, 133: 1702-1703. DANSGAAXD,W., 1953. The abundance of l W in atmospheric water and water vapor. Tellus, 5: 461-469. DANSGAARD, W., 1961. The isotopic composition of natural waters with special reference to the Greenland ice cap. Me&. Groenland, 165: 120 pp. DAVIS,L. E., 1955. Electrochemical properties of clays. Proc. Nutl. Conf. Clays C h y Technol. 1st-Calif., Div. Mines, Bull.. 169: 47-53. DEGENS, E. T. and E P S ~ I N S.,, 1962. Relationship between 180/160 ratios in coexistingcarbonates, cherts, and diatomitea. Bull. Am. Assoc. Petrol. Geologists,46: 534-542. DEOENS,E. T., HUNT,J. M.,REUTEX, J. H. and REED,W. E., 1964. Data on the distribution of amino acids and oxygen isotopes in petroleumbrine waters of various geologic ages. Sedimentology, 3: 199-225. DES I ~L. ,U., 1947. Diagenesis of oil-field brines. Bull. Am. Assoc. Petrol. Geologists, 31.20302040. Dmov, S. A., 1948. Classification of natural waters and graphical presentation of their composition. Dokl. Akad. Nauk S.S.S.R., 59(1): 87-90 (in Russian). EIIHALT, D., KNOTT,K., NAG% J. F. and VOGEL,J. C., 1963. Deuterium and oxygen-18 in rain water. J. Geophys. Res., 68: 3775-3780. ELLIS,C. B., 1955. Resh Water from ?he Ocean. Ronald Press Co.,New York, N.Y.

DIAGENESIS OF SUBSURPACB WATERS

501

EMERY,K. 0.and ENB BERG, S. C., 1952. Early diagenesis of California basin sediments in relation to origin of oil. Bull. Am. Assoc. Petrol. Geologists, 36: 735-806. EPSTEIN,S., 1959. The variations of the lsO/leO ratio in nature and some geologic implications. In: P. H. ABEUON (Editor), Researches in Geochemistry. Wiley, New York, N.Y., pp.217240. EPSTEIN,S. and MAYEDA,T., 1953. Variation of ‘*O content of waters from natural sources. Geochim. Cosmochim. Acta, 4: 213-224. EREMENKO,N. A. (Eiditor), 1960. Geology of Petroleum. 1. Principles of Geology and Petroleum. Gostoptekhizdat, Moscow, 592 pp. (in Russian). FRIEDMAN, I., 1953. Deuterium content of natural waters and other substances. Geochim. Cosmochim. Acta, 4: 89-103. GORRELL, H. A., 1958. Classification of formation waters based on sodium chloride content. Bull. Am. Assoc. Petrol. Geologists, 42: 2513. KNETSCH, G., SHATA, A., DEGENS, E. T., M~JNICH,K. O., VOGEL, J. C. und SHAZLY, M. M., 1963. Untersuchungenan Grundwhsern der Ost-Sahara. Geol. Rundschau, 1962,52(2): 587-610. KREJCI-GRAF, K., HECHT,F. and PALSER, W., 1957. uber ~ I f e l d w h des r Wiener Beckens. Geol. Jahrb., 74: 161-209. KRYUKOV, P. A., ZHUCHKOVA, A. A. and RENGARTEN, E. V., 1962. Changes in composition of solutions squeezed out of clays and ion-exchange resins. Dokl. Akad. Nauk S.S.S.R., 144: 1163-1165 (in Russian). V. D., 1954. About role of compaction processes of clayey deposits in the formation LOMTAJJZE, of underground waters. Dokl. Akad. Nauk S.S.S.R., 98: 451-454 (in Russian). LIBBY,W. F., 1959. Tritium in hydrology and meteorology. In: P. H. AEELSON(Editor), Researches in Geochemistry. Wiley, New York, N.Y., pp. 151-168. MCKELVEY Jr., J. G., and MILNE,1. H., 1962. The flow of salt solutions through compacted clay. Clays Clay Minerals, Proc. Natl. Conf. Clays Clay Minerals, 9(1962): 248-259. MCKELV~Y Jr., J. G., SPIEGLER, K.S. and WYLLIE,M. R. J., 1957. Salt filtering by ion-exchange grains and membranes. J. Phys. Chem., 61: 174-178. MEADE,R. H., 1963. Factors influencing the pore volume of he-grained sedimentsunderlow-tomoderate overburden loads. Sedimentology, 2: 235-242. M ~ m v S., F., 1956. Questions about Origin of Oil and Formation of Petroleum Deposits in Azerbaidzhan. Izd. Akad. Nauk Azerbaidzhanskoi S.S.R., Baku, 320 pp. (in Russian). M k m , K. O., 1963. Der Kreislauf des Radiokohlenstoffsin der Natur. Naturwissenschuften, 50: 211-218. RAINWATER, F. H. and W m , W. F., 1958. The solusphere-its inferences and study. Geochim. Cosmochim. Acta, 14: 244-249. RIEKE,H. H., CHILINOAR, G. V. and ROBERTSON, J. O., 1964. High-pressure (up to 500,000 psi) compaction studies on various clays. Intern. Geol. Congr., 2 2 4 New Delhi, 1964, Proc., in press. ~ N B E R GS ,. C., EMERY,K. O., HijLSEMA”, J., DEGENS, E. T.,FRAY, R. C., REIJTER, J. H., GRADY, J. R., RICHARDSON, S.H. and BRAY,E. E., 1963. Biogeochemistry of sediments in Experimental Mohole. J. Sediment. Petrol., 33: 140-172. SHL~HKINA, 0. V., 1959. Metamorphization of the chemical composition of muddy waters in the (Editor), Toward Knowle&e of Diagenesis of Sediments Black Sea. In: N. M. STRAKHOV (Symposium). Izd. Akad. Nauk S.S.S.R.,Moscow, pp.29-50 (in Russian). SIEVER,R., 1962. A squeezer for extracting interstitial water from modem sediments. J. Sediment. Petrol., 32: 329-331. S~EVER, R., GARRBLS, R. M., KANWISEIER, J. and BERNBR,R. A., 1961. Interstitial waters of recent marine muds off Cape Cod. Science, 134: 1071-1072. S m , V. A., 1948. Hydrogeology of Petroleum Deposits. Gostoptekhizdat, Leningrad. TOUTIKHIN, N. I., 1932. Toward question of graphical representation of analyses of waters. In: Sampling of Mineral Deposits. Gosgeolizdat (in Russian). (Editor), Companion of Field TOUTIKHIN,N. I. et al., 1954. Hydrogeology. In: N. B. VASSOEVICH Petroleum Geologist, ZI, 2 ed. Gostoptekhizdat, Leningrad, pp.101-145 (in Russian). VASS~EVICH, N. B. (Editor), 1954. Companionof FieldPetroleum Geologist, ZZ, 2 ed. Gostoptekhizdat, Leningrad, 564 pp. (in Russian).

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VONENGELHARDT, W.,1960.Der Porenraum der Sedimente. Springer, Berlin-Gottingen-Heidelberg, 207 pp. VONENGELHARDT, W. and GAIDA,K. H., 1963.Concentration changes of pore solutions during the compaction of clay sediments. J. Sediment. Petrol., 33: 919-930. W n m , M. R. J., 1955. Role of clays in well log interpretation. Proc. Natl. Conf. CIays Clay Technol., 1st-Calif., Div. Mines, Bull., 169: 282-305.

Chapter 11

INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS WOLF VON ENGELHARDT

Mineralogisch-PetrographischesInrtitut der Universitcit, Tiibingen (Germany)

SUMMARY

In this chapter the writer explains the diagenetic processes in sandstones and, starting from these observations, draws some conclusions concerning the processes occurring in the interstitial solutions. General conclusions about the concept of diagenesis and its distinction from metamorphism follow from this. INTRODUCTION

In the upper crustal zone, the sediments, which have been deposited as loose masses, undergo lithification and changes which are called diagenesis. In this zone the interspaces in the gradually consolidating sediments are, as a rule, filled with aqueous solutions which far exceed in quantity the local accumulations of natural gases or liquid hydrocarbons. These aqueous interstitial solutions, their chemical composition, and their movements, therefore, have particular bearing on the diagenetic processes. Their role in the diagenesis may be studied best in coarse-grained, highly porous, and highly permeable rocks. The extraction of interstitial solution is easiest from these rocks. New minerals, which grow from the solutions in such rocks, are most distinctly recognizable, and, as a rule, are clearly separated from the primary constituents. Also the dissolving effects of the interstitial solutions can be traced best in coarse-grained rocks. To illustrate general characteristics of diagenesis in sandstones, certain recent observations of the Keuper and Jurassic sandstones in Germany are presented here. GENERAL CHARACTEIUSTXCS OF DIAGENESIS IN S A ” E S

Table I shows the profile of the Middle Keuper Formation in the Stuttgart area. HELING (1963) recently examined the Stubensandstein (sandstone) using cores from a bore hole near Plochingen. Later ~ L I N G(1965) has, in a similar way, studied Schilfsandstein,a lower sandstone, using samples from northern and southern Germany.

W. VON ENGELHARDT

TABLE I LITHOLOQY AND TMCKNESSES OF MIDDLB AND LOWER KBLJPJ3R FORMATIONS NEAR STWTGART, GERMANY

Lithology

Formation

Average thickness (m)

Middle Keuper

Knollenmergel Stubensandstein Bunte Merge1 Schilfsandstein Gipskeuper

(marl) (sandstone with marls and clays) (marl with sandstone) (sandstone) (marl, gypsum)

Lower Keuper

Lettenkeuper

(marl with dolomite)

30 75 50 10 105 20

290

Total thickness

The Stubensandsteinin the Stuttgart area consists of a sequence of sandstones approximately 80 m thick, intercalated with clay and marly beds. Fig.1 shows the profile at Plochingen studied by HELING(1963). It consists of approximately 16 m of clays and marls and of about 57 m of sandstones. The sandstones, with the exception of a few clay-rich layers, are coarse-grained and very abundant in feldspar. The primary quartz/feldspar ratio is approximately 1/I. Most of the feldspar is orthoclase. Syngenetic layers and nodules of carbonates (calcite and dolomite) are to be found in the sands. In Fig.1 the distribution of the three principal sandstone types is indicated, as they have been formed by diagenesis: (I) consolidated sandstones having abundant kaolinite; (2) quartzitic sandstones; and (3) carbonate sandstones. These types belong to two principal diageneticperiods, which succeeded each other in time. During the earliest period, which must have started not too long after the deposition of the sandstone, decomposition of feldspar and garnet, as well as solution of quartz through pressure, took place on the one hand, and neoformation of kaolinite and quartz, on the other. The resulting sandstone, having a reduced content of feldspar and containing quartz grains with newly-grown rims, was filled with aggregates of kaolinite; where large quantities of quartz crystallized, quartzitic rocks have been formed. The decomposition of feldspar and the formation of kaolinite and of quartz must have taken place by way of the interstitial solutions. This process can be represented by the following formula:

+ +

+

+

+

2KAlSi30~ 16HzO + 2K+ 2M3+ 80H6H4Si04 + Al2(0H)4 Si205 4SiO2 2K+ 20H13H20

+

+

+

+

It is assumed, in this case, that Si dissolves in the form of orthosilicic acid, which

1NTERSTITLA.L SOLUTIONS AND DIAGENESIS IN SEDIMJBTS

505

CLAY

. . . .. .:::.. : . .. x) . . . . _..._ .._ ._..

SANDSTONE CONTAINING KAOLlNlTE AND CARBONATES

, -

SANDSTONE tONTAINING KAOLlNlTE

. ..-. .-.:.. SANDSTONE CONTAINI NG .... . . ...- - ------. . _ _

:

KAOLINlTE

CLAY

-....- *

ml

CONTAINING KAOLlNlTE QUARTZITE AND GRAVEL LAYERS

I

Fig.1. Lithologic section through the Stubensandstein, Plochingen, near Stuttgart, Germany.

does not dissociate easily; whereas Al and K are to be found mainly as ions. This would be in accordance with the conditions which exist when the pH values lie between 3 and 9, in which case the resultant solutions can only contain small concentrations of dissolved substances. The process, then, leads to an alkaline reaction, and can proceed only if there is a continuous removal of OH- and K+ions. This would be possible, for instance, if slightly acid solutions were introduced and the K+ion was extracted by formation of new minerals or as a result of adsorption. The observation that detrital grains of garnet dissolve, is compatible with slightly acid solutions acting during this first period, because garnet frequently has been observed to dissolve in slightly acid solutions in soils and rocks. With regard t o the decomposition of feldspar and the new formation of quartz, the calculation of a quantitative balance is impossible, because the decomposition and the neoformation do not have to take place in the same area; however,

506

W. VON ENGELHARDT

the quantity of newly-formed quartz seems to be too large to correspond with that of the feldspar decomposed. The newly-formed quartz is clearly recognizable as such; it is found in the form of pore-filling aggregates, and mainly as rims around detrital grains of quartz, which have often grown into small crystals with smooth facets. Some of the secondary quartz in the Stubensandstein most probably stems from pressure solution, because frequently stylolitic structures are to be found on grains of quartz which touch each other. Besides, additional SiOz may have been introduced into the sandstone from other sources. The processes of this first diagenetic period, as they have been observed in the Stubensandstein, are typical of diagenesis in many sandstones. Examples of the diagenetic dissolving of feldspars, of the neoformation of kaolinite, and especially of crystallization of quartz are known from numerous sedimentarybasins. At the same time, one always observes the decomposition of the so-called unstable heavy minerals, as for instance garnet, disthene, and staurolite. In this context, the writer would like to mention FUCHTBAUER’S (1961) fine observations on quartz growth in petroliferous Liassic and Dogger sandstones in northern Germany. The quantity of quartz grains with secondary rims found in the sandstones from oil fields in northern Germany is in distinct relation with the maximum depth of burial of the sandstone prior to the migration of the oil. This depth of burial apparently provides a measure of the intensity of the diagenetic change. Moreover, FUCHTBAUER (1961) was able to show that the formation of quartz and the simultaneousdissolving of unstable heavy minerals, like all diagenetic processes, comes to a standstill as soon as the interspaces are filled up with oil or with gases. At the same time the diagenesis continues in the water-filled rocks. In the oil-fdled sandstone, therefore, an early diagenetic stage is preserved. This accounts for the fact, known to all petroleum geologists, that oil-filled rocks within an oil pool usually are more porous and permeable than the same rocks from below the water table. Frequently this “diagenetic gap” does not coincide with the present boundary between the oil and water, which can be found at different levels, even in adjacent deposits. Quite a number of conclusions about the history of the oil migration and about the structural history of individual deposits can be drawn from these differences (see FUCHTBAUER, 1961; and PHILIPP et al., 1963). The first diagenetic period in the Stubensandstein was followed by a later one with different conditions prevailing in the interstitial solutions. Inasmuch as the products of the two diagenetic periods are distributed irregularly throughout the section, certain layers of sandstone being affected mainly by the processes of the first and others by those of the second period, one cannot determine whether the change in the chemical conditions occurred simultaneously throughout the entire sequence. It is quite possible that, during a period of transition, certain strata were subject to the conditions of the first period, and others already to those of the second. Yet, wherever mineral formations from both phases are to be observed i n the same bed, those of the second period always prove to be more recent.

INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS

507

During this later diagenetic stage, mainly calcite and dolomite were deposited in the pore spaces, in many instances replacing primary and secondary quartz and feldspar. The decomposition of garnet had ceased during this period, or, at least, it was taking place on a much smaller scale than during the first phase. Simultaneously with, or after, the carbonates, the Al-chlorite sudoite was formed in sandstones low in carbonate content. Finally, barite was formed, probably as the most recent product of the diagenesis. The late deposition of carbonate as the second diagenetic phase, after an initial formation of quartz and kaolinite associated with decomposition of feldspar, is typical in sandstones and is known through many examples. The successive order of these diagenetic processes, therefore, has to be explained from general principles. Whereas the Stubensandsteinthus represents a widespread form of the development of diagenesis in sandstones, HELING(1965) discovered an entirely different diagenetic process in the deeper-seated Schilfsandstein. Although the two sandstones are separated only by the marl sequence of the Bunte Mergel, 60-70 m thick, the chemical conditions existing in the interstitial solutions of the Schilfsandstein differ from those in the Stubensandstein. The Schilfsandstein consists throughout southern and northern Germany of a siltstone bed 0.5-4.0 m thick, containing a network of channel sandstones, which can reach thicknesses up to 50 m. The width of these channels varies between 100 m and several kilometers. Accordingly, one can distinguish between the so-called normal facies and the sandy channel facies. In the following discussion the writer is concerned mainly with the latter type of facies. The sandstones of the channel facies are rather homogeneous throughout Germany. The median diameter lies in the vicinity of 0.10 mm. Approximately 15 % of the clastic components are fragments of a fine-grained sericite schist, and 8-15 % are mica. There is also some quartz and feldspar, the quantitative relation of which varies between 2/1 and 1/1. Here, too, the orthoclase prevails over the plagioclase. During the most important diagenetic period in the Schilfsandstein, a new formation of orthoclase containing about 17% albite took place from the interstitial solution. This process must have started at a relatively early period when the sandstone was still highly porous. The newly-formed orthoclase crystallized as fringes around detrital grains of feldspar; during this process smooth crystal facets were formed wherever the spatial conditions would permit it. Fringes of orthoclase also formed around plagioclase. These fringes probably stand in a twin relation with the nucleus according to the Baveno law. The feldspar rims contain often zones which seem to be lower in Na content toward the exterior. The quantity of secondary feldspar is rather considerable, because approximately 60 % of all detrital grains of feldspar are surrounded by new rims. Chlorite (abundant in iron), which has been observed as a he-grained interspace-filling or as pseudo-

508

W. VON ENGELHARDT

morphs after primary biotite, is probably a somewhat more recent formation. Interstitial solutions which are oversaturated with alkali feldspar are, at any rate, slightly alkaline; for the formation of chlorite an alkaline reaction is likely to be necessary. Thus, the first diagenetic phase in the Schilfsandstein took place under slightly alkaline conditions. A second, later diagenetic phase was observed in samples of the Schilfsandstein from deep borings near Bruchsal in the Rhine Graben: anhydrite not only replaced feldspar, but also quartz and clay minerals. This happened to a very considerable extent in some instances, leaving a remainder of only 10% of the original clastic material. In the case of the feldspar, the inner parts, that is, the detrital grains, became more easily susceptibleto replacement than the newly-formed fringes. For this diagenetic process one has to suppose slightly acid solutions oversaturated with CaS04. Finally, as a third phase, the formation of carbonates (magnesite, dolomite, calcite) can also be observed in the Schilfsandstein. During this process, quartz, feldspar, and, in the Rhine Graben, anhydrite were replaced. Consequently, an oversaturation with carbonate and, probably, another slightly alkaline reaction must have prevailed during this period. In some samples of Schilfsandstein there has also been a formation of quartz as rims around detrital grains. In the whole, however, the formation of quartz took place on a much smaller scale in comparison with that in the Stubensandstein. For the processes observed in the Schilfsandstein, namely, the neoformation of orthoclase and of anhydrite, one could quote more examples from literature. A diagenetic development of this kind, however, seems to be less frequent in sandstones than the one observed in the Stubensandstein. The observationsmade in these specific Mesozoic sandstones were described elaborately in order to develop the principal and most general characteristics of all diageneticprocesses. All these observationscould be supplemented and confirmed by observations made in sandstones of different formations and sedimentary basins. In Germany, the thorough description by SCHERP (1963) of the diagenetic changes in sandstones of Carboniferous and Devonian ages at depths between 1,800 and 6,000 m, observed on examining cores of a recent deep bore hole near Miinster, deservesparticular mention in this context. There, again, the kaolinization of the feldspars, the formation of quartz and chlorite, and a late phase, during which carbonates were formed, can be observed. ROLE OF INTERSTITIAL SOLUTIONS IN DIAGENESIS

When trying now to determine the general characteristics of diagenesis, one has to point out first that the diagenetic changes and neoformations can not be considered as reactions in a closed system consisting of rock and interstitial solutions.

INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS

509

On the contrary, these processes take place in an open system and require, in each case, reactions with large quantities of solution. In specific cases the formation of quartz may be a reaction within a closed system, namely, whenever quartz is decomposed at grain contacts through pressure-solution, and precipitated again in places where the grains are not subject to pressure. If, however, as in most cases, the formation of quartz takes place elsewhere, where no dissolution is to be observed, considerable quantities of water have to be transported, as the following figures indicate. The solubility of quartz in HzO at room temperature is around 10 mg/l. If one supposes, at an increased temperature, a solubility of 30 mg/l, the interstitial solution present in 1 cm3 of sandstone with a porosity of 20% contains an approximate maximum of only 6 mg Si02 (= 2.3 * 10-6cm3 of quartz).Even if this entire Si02 content is precipitated in the pore spaces, the sandstone would have to be permeated by a quantity of solution 45 times the volume of its pores, before the porosity is reduced from 20 to 10%. For the solution of feldspar also large quantities of water are needed. In a closed system, an aqueous solution with potash feldspar reaches a state of equilibrium in which the solution becomes alkaline and contains relatively small amounts of dissolved feldspar. Larger quantities of feldspar can only be dissolved if, as in natural weathering in the soil or in artificial decomposition experiments in the laboratory, new solvent is continuously introduced. In experiments on the decomposition of potash feldspar, conducted by the writer in collaboration with CORRENS (CORRENS and VON ENGELHARDT, 1938), it was found that aqueous solutions with pH values between 6.5 and 11, which had been filtered through fine feldspar powder, contained between 20 and 40 mg/l of dissolved feldspar. These figures, certainly, do not represent equilibriumvalues, but they convey, nevertheless, an idea about the range of concentrations in solutions which are able to dissolve or to precipitate feldspar. The solubility of the unstable heavy minerals, such as garnet, staurolite, etc., is likely to be very low. Carbonates and anhydrite, it is true, have higher solubilities; but, here too, the quantities which were observed could not have formed through precipitation out of a stationary interstitial solution. This applies especially to barite. Thus, the decompositions and neoformations observed in sandstones show that the solution volumes involved in these reactions considerably exceeded the volume of the interspace. Large quantities of solution, therefore, must have flowed through the sandstones, bringing in, and also carrying out, various substances. Their chemical properties remained relatively unchanged for long periods of time; in the long run, however, their compositionchanged severaltimes. It can be assumed that the compaction of the argillaceous sediments is the principal cause for these flows which, in the course of geological periods, move large quantities of water in a subsiding sedimentarycolumn. The existence of such flows caused by compaction in recent sedimentary basins is evidenced, for instance, by the presence of superhydrostatic pressures in formation waters of sands surrounded by layers of clay in the

-

510

W. VON ENGELHARDT

Tertiary basin of the Gulf Coast and in the Tertiary basin of the Po River (see VONENGELHARDT, 1960, p.34). In older sedimentary basins, as, for instance, in northwestern Germany, the compaction apparently terminated, because everywhere in that area only hydrostatic pressures were observed. The quantities of water moved in the course of the compaction can be estimated, inasmuch as the porosity of the clay depends on the depth of subsidence. For the clays of the Lower Liassic in northwestern Germany, FUCHTBAUER (1961) determined the dependence of the porosity on the depth of burial (ROLL,1956; VONENGELHARDT, 1961), as shown in Fig.2. The curve plotted starts at a burial depth of 100 m and a correspondingfractional porosity of E = 0.35. At a depth T,1 cm3 of clay contains ET cm3 of pore space and (1 - E T ) cm3 of solid material. At a depth of To = 100 m, the same amount of clay material occupied a larger bulk volume, with the fractional porosity E~ = 0.35. The volume of solid clay material is:

Vf = (I-&!& and the initial pore volume corresponding to V f at To = 100 m was equal to:

v, =

Eo-Eo

ET

1- E o

Fig.2. Relation between porosity and maximum depth of burial for Liassic clays of northwestern Germany. (After F~~CHTBAUER, 1961.)

INTJ3RSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS

511

Thus the initial bulk volume ( V f + V p )cm3 of this clay was, in the course of the subsidence to the depth T compressed into 1 cm3. During this process a quantity of solution W T was squeezed out and, consequently, the following would hold true:

During the continuous sedimentation of clays in a subsiding basin a sequence of clay sedimentsis formed extendingfrom Toto the depth T. If the porosity decreases with the depth as shown in Fig.2, the quantity of interstitial solution yielded by the sequence per unit of surface area is found through integration of eq. 1: !r

T

The last integral in eq.2 is found by graphic integration of the curve of Fig.2; it is represented in Fig.3. If the depth T is measured in my WT is obtained in m3/m2 of surface, that is, in meters. The volume of solution yielded by any particular clay bed in the course of its burial, the top and bottom surfaces of which today lie at the depths TI and T2, respectively, is found according to the equation:

T1

This equation, too, may be found by means of the function presented on Fig.3. The following examples may illustrate these formulae. A succession of sediments 3,000 m in thickness, deposited from the Liassic time to the end of the Tertiary in the northern part of the Gifhorn trough in northern Germany, consists for the most part of clays. In the course of the subsidence they yielded 955 m3 of water/m2 of surface, provided that these clays reacted in the same way as the Liassic deposits. The compaction down to a depth of 100 m is neglected here. This large quantity of liquid was squeezed out in the course of a very long period of time; the average velocity of flow was, therefore, very low. On supposing a constant subsidence from the beginning of the Liassic until the end of the Tertiary, that is, during a period of approximately 180 million years, the flow of compaction must have penetrated a plane 100 m below the surface with an average velocity of only 5.3,ulyear. In reality, however, the subsidence was interrupted by periods of rest and of uplift. Yet, even on assigning only half of the time to the subsidence, an average velocity of flow of loplyear is still very slow. It will further be interesting to see what quantities of solution penetrated a particular section of sandstone during the compaction. The simplest supposition would be that the flow of compaction was directed upward, because it was caused

512

W. VON ENGELHARDT

Fig.3. Integration of data plotted in Fig.2.

by the force of gravitation which is directed downward vertically. In horizontal deposits the sandstone is, under these conditions, permeated by those solutions which are squeezed out of the subjacent clays. In the Liassic and Dogger sediments of the Gifiorn trough, the thickness relation of sandstone to clay lies between 1-10 and 1-20. Let us suppose, for example, that a layer of sandstone having a thickness of 30 m is underlain by a clay bed 600 m thick, which is situated, today, between 1,500 m and 2,100 m below the surface. According to eq.3 and Fig.3, one finds that, during the subsidence from a depth of 100 m to its present depth, this complex of clays emitted 213 m3 of solution/ m2. This quantity of solution traversed the overlying bed of sandstone, which, in a 30-m column with a 1-m2 base, contains 7.5 m3 of pore space (assuming porosity = 25 %). Under the above conditions, the interstitial solutions of the sandstone, therefore, have been renewed about 30 times by the flow of compaction. Considering the decompositions and neoformations, which were observed in sandstones, this is certainly less than what is required. With thinner clay beds, the quantities are even smaller. The Stubensandstein in southern Germany, for instance, is underlain by the Bunte Mergel (marl clay deposit) which is about 50 m thick today, and had probably a maximum depth of burial of 800-1,000 m. The Bunte Mergel contains about 40 m of marl and clay, and about 10 m of siliceous sandstone. If, during the compaction,the clay sediments responded the same way the Liassic clays in northern Germany did, 9.2 m3 of interstitial solution/m2 of horizontal bedding plane were squeezed out during the subsidence from 100 to 860 m (depth of top surface). The overlying Stubensandstein (with an approximate thickness of sandstones being 60 my and a porosity

+

INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS

513

of 25 %) contains 15 m3 of pore space/m2. Consequently, the quantity of solution squeezed out from the Bunte Merge1 would not have sufficed to renew even one time the interstitial solutions of the sandstone. Whereas the decomposition and neoformation of minerals lead to the conclusion that horizontally lying sandstones must have been permeated by large quantities of solutions compared with the pore volume, the calculations show that, in case the flow was directed upward vertically, the flow of compaction coming from subjacent clays frequently provided volumes which were definitely too small. The reason for this discrepancy is the fact that the assumption of vertical direction of the compaction flow through sandstones only applies to the ideal case and is practically never realized in nature (HUBBERT, 1940; VON ENGELHARDT, 1960). If one disregards the effect of directional permeability variations, the flow of compaction in clays has the same direction as the pressure gradient producing it, thus being directed vertically upward. If the solutions enter through a perfectly horizontal boundary plane from clay into rigid sandstone, in which the pore volume does not change through compaction any more, the vertical direction of migration is not altered. If, however, the boundary plane between clay and sandstone forms an angle with the horizontal, the flow lines will be refracted, as shown in Fig.4. If a1 and a 2 are the angles between flow directions in clay and sandstone and the line normal to the boundary plane, and kl and k2 are the permeabilities in clay and sandstone, respectively, then: tgal : tgaz = kl : k2

T

(4)

I I

I

\

Clay

a2

Fig.4. Refraction of flow at the boundaries between beds having different permeability.

514

W. VON ENGELHARDT

As shown in Fig.4, the refraction takes place in such a way that the direction of flow in the sandstone is deflected away from the normal and toward the lower boundary plane. Inasmuch as the permeabilities of clay and sandstone differ widely, the deflection of flow direction can become considerable even if the tilting is very slight. An example with figures may illustrate this: If the permeability of the clay kl = 10-6 darcy, that of the sandstone kz = 10-1 darcy, and the boundary plane is tilted by only lo', then: k2 = 10'; kl = 10-8; k2 = 10-1; tga2 = * tgal = 105 m i 10-3 = ki 2.91 102; a2 = 89'48'

-

~

-

The flow lines would then form an angle of only 12' with the boundary plane, thus running almost parallel with the bedding plane of the sandstone. If the thickness of the sandstone is only 10m, andifthe angle of incidence is constant, a single stream line would run within the sandstone for 3,000 m before reaching the top surface and entering into the overlying clay. Inasmuch as the boundary planes are never perfectly horizontal, the flows of compaction will collect in the sandstones, and travel great distances in them, despite the general direction of incidence. The flow of compaction is conducted up along the margins of basins from their lower parts mainly through sandstone strata. The quantities of liquid permeating a cross-section of sandstone, therefore will usually be much greater than the thickness of the clays underlying it would suggest. Sandstones in structurally high positions deserve particular attention. It is to be expected that the solution squeezed by compaction flow into anticlines and other structural highs from all directions, provided that the permeability of the overlying clays is constant. From the structural crests the solutions must enter the overlyingclay beds. There, a particularly intenseflow should be moving upward, and it ought to be possible to find its traces in the clay beds overlying the structure. If the solutions coming from underlying clays have carried small drops of oil or gas blebs besides various dissolved substances, the former are filtered out and remain in the sandstone, because they cannot enter into the small pores of the overlying clay from the large sandstone pores, if both are filled with water. This way, oil and gas deposits are formed which are products of diagenesis, as well as the decompositions and neoformations of difficultly soluble minerals in sandstones. In both cases they only could have formed through the flowing of great quantities of liquid. The picture of the permeation of sandstones by interstitial solutions, as it has, so far, been developed, is only valid if the permeability of the clay beds covering the sandstone is the same everywhere. If, by change of facies or because of joints and cracks, the permeability of these beds is higher in some parts than in adjacent areas, the solutions entering the sandstone might stream toward these places of higher clay permeability regardless of the dip. That interstitial solutions in sandstones sometimes flow from high positions toward structurally low ones

INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS

515

can be deduced from the tilting of the oil-water boundary in some oil deposits (HUBBERT, 1953; VONENGELHARDT, 1960, p.121). Such a boundary plane can only be tilted if the water is flowing. In these cases, the permeability of the strata overlying the sandstone must be lower in the area of the structural crest than in the area toward which the water is flowing. It is conceivable that in some cases the precipitation of dissolved substances from the solutions, flowing at an earlier time toward the crest, caused this effect. If fracture zones extend down into greater depths, it is possible that solutions which have been squeezed out of deeply-buried sediments will rise through them. Some thermal springs at the margins of recent sedimentary basins are probably fed entirely or partly through flows of compaction which owe their origin to a large volume of clay sediments, and which collect in the clefts leading to the surface. One might, for instance, examine the degree to which the hot springs at the margin of the Rhine Graben yield solutions, which owe their origin to the compaction of Tertiary sediments.

THEORIES OF DIAGENETIC REACTIONS

After explaining the probable origin of the great quantities of solution necessary for the diagenetic processes in sandstones, there remains a great deal to be answered about the theory of diagenetic reactions. Three groups of questions may be raised, which have to be answered. First of all, the origin of the substances causing the decompositions and the precipitations, and the whereabouts of other substances which have to be carried away continuously if the reactions are to proceed in the same direction, should be explained. Secondly, one has to explain why the concentration in the solutions permeating sandstones for a long time reached and passed saturation with respect to a particular solid phase, precipitating it in considerable quantities. Finally, it is necessary to explain how different chemical conditions could be maintained in different strata of sandstones belonging to the same sequence, not too far from each other, despite the fact that great quantities of solution flowed through the whole sequence. About the origin of the substances, one can only say, at present, that considerable amount of research remains to be done in this field. The solutions originate in the clays, and clays of different mineralogical composition and of different origin are likely to contain solutions having different composition. It was ascertained experimentally that the solution squeezed out from clay, soaked with salt solution, is more concentrated than that which remains in the clay (VONENGELHARDT and GAIDA, 1963). The question as to how various ions will behave if the interstitial solutions contain several kinds of ions has been only partially in-

516

W. VON ENGELHARDT

vestigated (Rmm et al., 1964; KRYUKOV, 1964). It may be assumed that disproportions would occur, that certain ions would be preferentially squeezed out initially and that the remaining liquid would be enriched in others. The slightly acid solutions necessary for the kaolinization of feldspars and the decomposition of heavy minerals, and which occur so frequently, may be caused by a production of H+ ions during the first stages of compaction in clays. The origin of the SiOz, which is found very frequently in sandstones as newly-formed quartz, is still uncertain. First of all, the formation of quartz may be due to the dependence of the solubility of SiOz on temperature. Because of the geothermal gradient, Si02 can dissolve at a greater depth and be precipitated closer to the surface. The transport can be effected either through the upward flowing solution, or, in cases of very slow subsidence and flow, by diffusion. The amorphous SiOz from the remains of organisms embedded in the sediments, and which has a much higher solubility than quartz, may be another source for quartz formation in sandstones. Secondary quartz can also owe its origin to the transformation of clay minerals abundant in SiOz into minerals having a lower Si02 content: for instance, as a result of transformation of illite-montmorillonite into illite and finally into muscovite. The writer already pointed out the pressure dissolving of quartz at grain contacts, according to the principle of Riecke. The origin of the substances necessary for the neoformation of feldspar is completely uncertain. Diagenetic processes in the clays, about which earth scientists hardly know anything definite, must have provided them. The thorough study of the clays of the Middle Keuper Formation, which is being conducted at the University of Tubingen, will, perhaps, produce some indications as to the reasons why the chemistry of solutions entering the Schilfsandstein was different from that of the solutions which moved through the Stubensandstein. One is also unable to say anything about the whereabouts of the potassium produced by the decomposition of feldspar in the Stubensandstein. Possibly, the increase of the illite content in the argillaceous fraction, which KROMER (1963) discovered in the lower parts of the Zanclodon marls, is caused by the migration of potassium from the Stubensandstein. The frequently observable supersession of an initial period of quartz formation and feldspar decomposition by a later phase, during which carbonate is formed and quartz and feldspar are replaced, must be ascribed to general causes, operating during the subsidence of clayey-sandy sediments. For the time being one can only state that the solutions, which are usually slightly acid at shallower depths, are replaced at a greater depth by slightly alkaline solutions, which have higher contents of Ca and Mg ions. The alkaline reaction, the higher Mg content, and probably also the higher temperature cause the formation of chlorite during this phase. In this chapter the author would like to present facts only, renouncing the formulation of dubious hypotheses. The formation of barite belongs to a still later diagenetic phase. It is likely

INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS

517

that, in some sandstones, it did not form during the subsidence, but during a later uplift. Sulfate is to be found only in those interstitial solutions which are relatively close to the surface. The lower formation waters usually do not contain sulfate anion (VON ENGELHARDT, 1961). Recently, the results of numerous analyses of mine waters in the Ruhr basin have been published by MICHEL and RULLER (1964) and by PUCHELT (1964). According to these authors, the highly concentrated waters from greater depths do not contain sulfate anion, but do contain 1,0003,000 mg/l of barium. In solutions which are low in salt content, and which come from shallower depths, sulfate anion is found; but as the solubility product of Bas04 indicates, very little barium is present. The presence in these solutions of sulfate-reducingBacteria, as shown by PUCHELT (1964), indicates that the sulfate is destroyed in the higher levels by bacterial action. There is, therefore, a lower boundary to the occurrence of sulfate. If the hydrothermal waters enriched in Ba content but devoid of sulfate ion come to the surface in the mines, oxidation of sulfide sulfur and consequent precipitation of barium sulfate takes place. A similar process probably causes a late diagenetic formation of barite, when deep-lying sandstones, the interstitial solutions of which have lost their sulfate through bacterial action and are enriched in barium content, are uplifted in the oxidation zone. There are various reasons for the frequent phenomenon of supersaturation in the pore solutions of sandstones with respect to certain kinds of minerals, which was mentioned in the second group of questions. The cooling of the solutions flowing in from deeper horizons with higher temperature may be one reason. But the author believes, above all, that the so-called filtering effect of clays with respect to electrolytic solutions plays an important part. At the present time it has been ascertained, not only by theoretical speculations, but also through experiments, that clays act as barriers against anions and are able to filter out to a large extent the electrolyte content of permeating solutions (VON ENGELHARDT and GAIDA,1963). If a solution containing electrolyte flows from sandstone into clay, the concentration within the sandstone must necessarily increase. Without entering into details about this mechanism which has been treated more thoroughly elsewhere (DEGENS and CHILINGAR, 1967, in this book, for example), the author would like to point out only that the formation of salt-rich interstitial solutions and the precipitation of minerals in sandstones permeated by such solutions can be explained in this way. Inasmuch as structurally high sandstone beds are characterized by a particularly intense flow of interstitial solutions permeating the overlying clay beds, a large increase in the concentration of dissolved substances in such sandstones is to be expected. In fact, all oil geologists are aware of the occurrence of cemented layers close to the top surface of oil-bearing sandstones, which seem to be formed more distinctly in structurally high positions. They were for the first time interpreted by FOTHERGILL (1955) as being a consequence of the filtering of ions. In Germany, such cemented layers are known in Dogger sands and Valanginian

518

W. VON ENGELHARDT

sands, among others. FUCHTBAUER (1961) has shown that in the area of the Gifhorn trough (northern Germany) the formation of secondary quartz is particularly extensive in structurally high layers of Dogger Sandstone immediately underlying the argillaceous cap rock. He has ascribed this phenomenon to the barrier action of the clays. In the Eldingen oil field, east of Celle (northern Germany), the gradual increase in the concentration of the interstitial solutions during the diagenesis can actually be proved (PHILIPP,1961). There, in a high position of sandstone at the western margin of the Gifhorn trough, the formation of oil deposit began already during the Liassic time. The oldest oil is preserved in the crest of the structure within a poorly cemented sandstone of an early diagenetic stage with a porosity of 32%; the associated interstitial water is low in salt content. Interstitial water in oil-filled sandstones forms thin layers or drops, which are isolated from one another. It cannot flow, usually, and also diffusion equilibrium does not exist here or in the water-filled part of the sandstone. The salt content at the time of oil migration is preserved, therefore, more or less unchanged. In structurally lower parts of the sandstone, filled with oil at a later period, the interstitial water contains more salt, in accordance with the growing concentration of formation water during diagenesis. The highest concentration, finally, is to be found in the free water below the oil-water boundary filling the pore spaces of the sandstone outside the oil pool. Table I1 shows the composition of water from wells producing from different structural depths, and also of the water below the oil-water boundary. Well no.9 was drilled on the crest of the structure, whereas wells no.6 and I I follow down dip in the direction toward the oil-water boundary. One recognizes the gradual increase in concentration, accompanied by a relative increase in Ca and a relative decrease in Mg ion content. For the third phenomenon, namely, the chemical isolation of sandstone beds not far from each other in the stratigraphic column, the barrier action of TABLE II CHEMICAL COMPOSITION OF ELDINGEN OIL-FIELDWATERS, NEAR CELLE, NORTHERN GERMANY

(After PHILIPP, 1961) Well Depth of top n p e of No. of sandstone water layer (m) 9 6 11 49

1,379 1,379 1,388 <1,388

Dissolved solids

Content of cations in % mMol

mgll

mMol/l

Na+

Caa+ Mgs+ K+

210 380 460 2,716

88.5

4.2 4.5 5.0 6.0

interstitial 11.8 interstitial 23.5 interstitial 28.5 free 169

94.3 93.7 92.5

7.0 0.9 1.0 1.4

0.3 0.3 0.3 0.1

S04a-ICImolar ratio

O.oo00 0.00308 0.00341 0.000112

INTERSTITIAL.SOLUTIONS AND DIAGENESIS IN SEDIMENTS

519

clays against anions is, again, likely to be the reason. The flow of compaction permeating a clay-sand sequence of sediments leaves most of the dissolved ions in the sandstones. Chemical conditions within the particular sandstone, therefore, will remain unchanged during long periods of time. They will be determined by the particular characteristics of the associated clay bed, the interstitial solutions of which migrate into the sandstone.

DISTINCTION BETWEEN DIAGENESIS AND METAMORPHISM

According to general usage, the term diagenesis comprises all processes of transformation through which, at moderate pressures and temperatures and at relatively shallow depths, compact rocks are formed from unconsolidated sediments. The diageneticprocesses, accordingly,start right after the depositionof the sediment;and diagenesis is followed by metamorphism, which begins to operate at higher temperatures and pressures. It has been attempted to define the boundary between diagenesis and metamorphism on the basis of the concept of mineral facies, through presence of critical minerals or characteristic mineral associations. Attempts of this kind lead to great difficulties, because the normal widespread kinds of sedimentary rocks contain only minerals which are altogether noncharacteristic: most of the minerals which form or dissolve as a result of diagenetic processes, such as quartz, feldspar, mica, chlorites, clay minerals, tourmaline, zircon, etc. are also to be found in unquestionably metamorphic facies. Minerals which may be critical, such as some of the zeolites (laumontite), are rarely to be found. Therefore, a definite distinction between metamorphism and diagenesis, which would apply to all practical cases, can not be made on using the facies concept. Consequently, the author would like to suggest that no sharp boundary line should be drawn between diagenesis and metamorphism, because it does not exist in nature. Instead, metamorphism and diagenesis should be identified by the types of processes prevailing during these two stages of transformation of primary sediments. This way it will be possible, in almost every practical case, to distinguish between diagenetic and metamorphic changes. In the diagenetic zone, the sediments contain interstitial fluids as continuous phases which can be moved by normal flow. These fluids (most commonly aqueous solutions) are, therefore, always involved in the diagenetic reactions, the transport of substances being effected through flow or normal diffusion. Because of the intercommunication and mobility of the interstitial fluids, reactions of the opensystem type prevail. Some sediments, as, for instance, fine-grained limestones, can already lose all of their porosity at a very shallow depth. Crystallizations and phase transformations in such dense rocks at moderate temperatures and pressures may be called diagenetic, if other closely associated sediments still have an intercommunicating porosity.

520

W. VON ENGELHARDT

The diagenesis stops at a depth where in all sedimentary rocks the intercommunicating pore spaces have been closed up by physical or chemical processes. In this zone the metamorphism begins, its reactions taking place in the solid state or by diffusion at grain boundariesl. There are, indeed, metamorphic reactions of the open-system type (metasomatosis). On the whole, however, in metamorphism, reactions of the closed-system type, and equilibria between solid phases, prevail. Solutions are to be found in the metamorphic zone also. They consist of remains of the interstitial .solutions, augmented by solutions released during the metamorphic transformation of hydrous minerals into unhydrous ones, and, perhaps, also by feedings of magmatic origin. They invariably collect in places, where, through tectonic movements, fractures and cavities with a lower pressure have opened. Such cavities are for instance formed by faults connectinglowerwith higher levels or by shearing, as clefts closed on all sides. In these cavities precipitation of mineral substances takes place similar to the diagenetic formations in the pore spaces of sandstones. Examples of these are mineral veins and the druses of the Alpine shearing clefts. During diagenesis, the intercommunicating porosity enables the interstitial solutions to penetrate and flow through the sedimentary rock, and thus participate in all diagenetic processes. A knowledge of the chemistry of these interstitial solutions or formation waters is very important for the comprehension of diagenesis. Unfortunately not enough general facts are known concerning the occurrence of the principal constituents in these waters; data pertaining to the frequency of the accessory constituents are still very scarce. Wherever solutions of this kind are found in mining and oil-drilling operations, they ought to be collected carefully and analyzed thoroughly.

REFERENCES

CORRENS, C. W. und VONENGELHARDT, W., 1938. Neue Untersuchungen iiber die Verwitterung des Kalifeldspates. Chem. Erde, 12: 1-22. E. T. and CHILINGAR, G. V., 1967. Diagenesis in subsurface waters. In: G. LARSENand DEGENS, G. V. CHILINQAR (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, pp. 477-502. C. A., 1955. The cementation of oil reservoir sands and its origin. World Petrol. FOTHERGILL, Congr., Proc., 4th, Rome, 1955, Sect. I, pp.301-314, H., 1961. Zur Quarzneubildung in Erdollagerstatten. Erd8l Kohle, 14: 169-173. FUCHTBAUER, HELLINQ, D., 1963. Zur Petrographie des Stubensandsteins. Beitr. Minerd. Petrog., 9: 251-284. D., 1965. Zur Petrographie des Schilfsandsteins. Beitr. Mineral. Petrog,, 11: 272-296. HELLING,

In accordance with this, SCHERP (1963) defined, for the sequence of sedimentary rocks of Devonian to Upper Carboniferous age from the deep bore hole Miinsterland 1, the boundary between diagenesis and metamorphism at a depth of approximately 3,000 m. At this depth the porosity of sandstones practically ceases to exist and reactions between minerals and interstitial solutions become insignificant.

INTERSTITIAL SOLUTIONS AND DIAGENESIS IN SEDIMENTS

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HUBBERT, K., 1940. The theory of groundwater motion. J. Geol., 48: 785-944. HUBBERT, K., 1953. Entrapment of petroleum under hydrodynamic conditions. Bull. Am. Assoc. Petrol. Geologists, 37: 1954-2026. KROMER, H., 1963. Untersuchungeniiber den Mineralbestand des Knollenmergel-Keupers in Wiirttemberg. Dissertation, Tiibingen, 71 pp. KRYUICOV, P. A., 1964. Soil, Mud and Rock Solutions. Summary. Thesis. Akad. Nauk S.S.S.R., Inst. Geokhim. Anal. Khim. V.I. Vernadskiy, Moscow, 56 pp. MICHEL,G. und RULLER,K. H., 1964. Hydrochemische Untersuchungen des Grubenwassers der Zechen der Hiittenwerke Oberhausen AG. Bergbau Arch., 25: 21-27. PHILIPP, W., 1961. Struktur und Lagerstiittengeschichte des Erdolfeldes Eldingen. Z. Deut. Geol. Ges., 112: 414-482. PHILPP,W., DRONG,H. J., FWCHTBAUER, H., HADDENHORST, H. G. und JANKOWSKY, W., 1963. Zur Migration im Gihorner Trog. Erdol Kohle, 16: 456-468. PUCHELT, H., 1964. Zur Geochemie des Grubenwassers im Ruhrgebiet. Z. Deut. Geol. Ges., 116: 167-203. RECKE,H. H. 111, CHILINCAR, G. V. and ROBERTSON JR., J.O., 1964. High pressure (upto500,000 psi.) compaction studies on various clays. Intern. Geol. Congr., 22nd, New Delhi, 1964, in press. ROLL,A., 1956. Zur Strukturgeschichte der Salzstocke von Wesendorf und Hohenhorn. In: F. L~TzE(Herausgeber), Geotektonisches Symposium zu Ehren von Hans Stille. Enke, Stuttgart, pp.228-245. SCHERP,A., 1963. Die Petrographie der palaeozoischen Sandsteine in der Bohrung Miinsterland 1 und ihre Diagenese in Abhhgigkeit von der Teufe. In: Die Aufschlussbohrung Miinsterland I-Fortschr. Geol. Rheinland Westfulen, 11 :251-282. VON ENGELHARDT, W., 1960. Der Porenraum der Sedimente. Springer, Berlin-Gottingen-Heidelberg, 207 pp. W., 1961. Zum Chemismus der Porenlosung der Sedimente. Bull. Geol. VONENGELHARDT, Inst. Univ. Upsala, 40: 189-204. VONENGELHARDT, W. and GAJDA,K. H., 1963. Concentration changes of pore solutions during the compaction of clay sediments. J. Sediment. Petrol., 33: 919-930.

Chapter 12

CONCLUDING REMARKS ON DIAGENESIS G. LARSEN AND G.V. CHILINGAR

Department of Geology, Universityof Aarhm. Aarhus (Denmark) University of Southern California,Los Angeles, Calif. (U.S.A.)

One of the main impressions gained from the very varied contents of this book i, that diagenesis is a field of geology in which research is undergoing a phase of veq rapid development; the reference lists appended to each chapter are, in themselves evidence of this. Another main impression, which is undoubtedly connected with the above, i! that there is not yet a universally accepted definition or delimitation of the tern “diagenesis”, let alone agreement on the specialized terminology which is used tc characterize the phenomena encountered in this field, despite the fact that the expression “diagenesis” has existed in the geological literature for about a hundred years (see AMSTUTZand BUBENICEK).~ Broadly speaking, two main points of view are found in this book regarding the delimitation of the term “diagenesis”. According to the one, which is advanced et al.), diagenesis is complete when the in the chapter on carbonates (CHILINGAR sediment has been converted to a sedimentary rock. This corresponds closely to the current practice in the U.S.S.R.where the distinctionis made between “syngenesis”, “diagenesis”, “epigenesis” and “metamorphism” (see RUKHIN, 1961,pp. 295331). In contrast with this, many other authors of this book, for exampleFAIRBRIDGE, DAPPLES, M ~ L E Rand , VON ENGELHARDT take the point of view that diagenesis includes all processes between deposition and metamorphism. PRAY and MURRAY (1965) express this as follows: “In its broadest sense diagenesis encompasses those natural changes which occur in sediments or sedimentary rocks between the time of initial deposition and the time-if ever-when the changes created by elevated temperature, or pressure, or by other conditions can be considered to have crossed the threshold into the realm of metamorphism.” The question of where to place the boundary between diagenesis in this broader sense and metamorphism has been touched on in the opening chapter (LARSEN and CHILINGAR). Furthermore, FAIRBRIDGE and VONENGELHARDT have discussed the question. According to the latter the boundary zone can be characterized as that “where in all sedimentary rocks the intercommunicating pore spaces have been closed up by physical or chemical processes”. ‘References that appear in the text without year of publication refer to authors of chapters in this book.

524

G. LARSEN AND G. V. CHILINGAR

With referenceto the terminological differences which arise, it can be mentioned that FAIRBRIDGE describes the late, “deep” diagenesis stage as the “anadiagenetic”; almost synonymous are DAPPLES’S “phyllomorphic stage” and M ~ L E R ’ S “deep burial stage”. It may be added to this that CHILINGAR et al. consider this stage to belong to epigenesis and not diagenesis. There are, therefore, as the above example shows, several different opinions. The editors of this book have not found it appropriate here to attempt to give one particular opinion more prominence than the others, but merely express the wish that before long some standardization may be established with regard to the definition of the term “diagenesis” and the terminology in t h i s field. Regardless of how the phenomena may be defined or named, diagenetic transformation may be characterized collectively as the reaction of the sediment or sedimentary rock to its physicochemical environment. The most important factors involved here are the pH and Eh of the environment, the concentration of various anions and cations, as well as pressure and temperature. As stated in several chapters in this book, many aspects of the development of the diagenetic processes are still inadequately explained. As stated earlier, however, investigation of the various subjects involved is proceeding at a rapid rate. This is undoubtedly due in part to the direct economic interest which is attached to certain aspects of diagenesis. An example of this is the recognition of the significance of diagenesis in the formation of certain minerals (AMSTUTZ and BUBFBICEK). In this connection the processes of coalification can be mentioned (TEICHMULLER and TEICHMULLER), as well as the origin of hydrocarbons (DEGENS). Within the realm of problems surrounding the formation of hydrocarbons are also certain other effects of diagenesis, namely the changes which occur in the composition of the formation water on the one hand (DEGENS and CHILINGAR), and on the other hand the cementing and dissolving processes which determine the permeability and other properties of the sedimentary formations (see CHILINGAR et al.). It is not really possible to distinguish precisely between those aspects of diagenesis which are of direct economic interest and those which are of purely scientific value. Because of the rapid developments in the investigations of diagenetic phenomena, a final, exhaustive account of diagenesis could scarcely be expected in the present book; it is, however, the editors’ opinion that the book represents an important step forward in the recognition of the nature of diagenesis.

REFERENCES

PRAY,L. C. and MURRAY, R. C. (Editors), 1965. Dolomitization and Limestone Diagenesis

( A Symposium)-Soc. &on. Paleontologists Mineralogists, Spec. Publ., 13 : 188 pp. RUKHM,L. B., 1961. Principles of Lithology. Gostoptekhizdat, Leningrad, 779 pp.

REFERENCES INDEX'

ABELSON,P. H., 35,79,350,353,371, 382,501 ABELSON,P. H. and HOERINQ, T. C., 375, 381, 382

ABRLSON, P. H., HOERINO,T. C. and PARKER, P. L., 361, 382 ADAMS, A. P., BARTHOLOMEW, W. V. and CLARK,F. E., 367, 383 ADAMS,J. E. and RHODRS,M. L., 68, 79 ABERHALDEN, E. and HEYNS, K., 361, 382 ABRAHAM,H., 379, 382, 383 ALDERMAN, A. R. and SKINNER, H. C. W., 298 ALDERMAN, A. R.and VON DER BORCH,c. c., 298

AULT,W. U., 35, 79 AULT,W. U. and KULP,J. L., 55, 79 AUSTEN,D. E. G. and INGRAM, 0. J. E., 403, 411

A m , J., 57, 59, 80

BAARS,D. L.,260, 270, 299 BAASBECKINQ, L. G. M., 26, 35, 54, 80, 448, 451,467

BAAS BECKING, L. G. M. and MACKAY, M., 467 BAASBECKINQ, L. G. M. and MOORE, D., 23, 25, 27,28, 29, 54, 80,467

BAASBECIUNQ,L. G. M. and WOOD, E. J. F.,

ALEKIN,0. A., 491, 500 467 BAASBECKINO, L. G. M., KAPLAN,I. R. and ALIMEN,H., 72, 79 ALLEN,E. T., CRENSHAW, J. L. and JOHN- MOORE,D., 80, 327, 340, 347, 383,467 STON, J. J., 54, 79 BAASBECKING, L. G. M., WOOD,E. 3. F. and ALLINO,H. L., 467 KAPLAN,I. R., 467 AMERICANGEOLOQICAL INSTITUTE, 161, 172, BADER,H., 453,468 BADER,R. G., 33, 35, 80 426,467 BADER,R. G., HOOD,D. W. and SMITH, J. B., Ammz, G. C., 181,298,421,422,423,426, 43 1,432,442,447,450,452,453,467 AMSTUTZ, G. C. and BUBENICEK, J., 1, 2, 16, 19, 79, 523, 524 AMSTUTZ, G. and PARK, w., 433 AMSTUTZ, G. C., RAMDOHR,P., EL BAZ, F. and PARK,W. C.,426, 442, 453, 467 ANDERS, E., 378, 383 ANDERSON, D. L. and BENSON,C. S., 21, 79 ANDERSON, F. W., 271, 298 ANDERSON, G. T., 367, 383 ANDERSON, G. T. and HAN,T. M., 452, 467 ANDRISE, R.,20,21,79,425,467 ANDREEV, P. F., BOGOMOLOVA, A. I., DOBRYANSKIY, A. F. and KARTSEV, A. A., 376, 383 ANDRICHUK,J. M., 270, 277, 298 ANONYMOUS, 69,19 AREND,J. P., 457, 467 ARRHEIWS, G., 36, 79, 134, 149, 150, 151, 172 ARRHENIUS, G., MERO,J. and KORKISH,J., 447,452,467 A u o u s ~ ~ l m S. s , S., 453, 467

c.

359, 383

BAERTSCHI,P., 379, 383 BAKER,E. G., 377, 383 BAKER,G., 445, 453, 468 BALUKHOVSKIY, N. F., 8 BANE-, A., 265, 266, 299 BANQHAM, D. H. and MAGGS,F. A. P., 403, 411

BARQHOORN, E. S., 353, 359, 383 BARQHOORN, E. S., MEJNSCHEIN, W. G. and SCHOPP,J. W., 299 BARQHOORN, E. S. and SPACKMAN, W., 359, 383

BARNES,I., 299 BARNES,I. and BACK,W., 299 BARON,G., 67, 80 BARTH, T. F. W., 24,28, 80 BARTH,T. F. W., C~RRENS, C. W. and ESKOLA, P., 70, 80, 301 BARTHOLOMI?, P., 435, 447, 452, 468 BASKIN,Y.,112, 122

1 The help extended by Herman H. Rieke, Ln in preparing the reference index is greatly appreciated by the editors.

526

REFERENCES INDEX

BRADLEY,W. H., 47, 80 BASTIN,E. S., 55,80,468 A. M., 421,424,468 O., 30,31,57,80,145,453,468 BATEMAN, BRAITSCH. BATES, R. L., 424,468 R. A. and POWERS, R. W., 300 BRAMKAMP, R. G. C., 210, 218, 219, 221, 227, BRAUN,H., 240, 300,452,457,468 BATHURST, F. E., 371,383 229, 233, 234, 263, 267, 270, 271, 273, 276, BRAUNS, E. D., 363, 375, 376, 279, 283, 299, 315, 317, 318, 319, 320, 321, BRAY,E. E. and EVANS, 322,446,468 377, 383, 498, 500 V. P., 41, 80 BATURIN, E. A., 433,445,468 BRECKE, BAUSCH, W. M., 299 J. D., BLYTH,C. R.,WHITE, W. A. BREDEHOEFT, BAUSCH, W. M. and WIONTZEK,H., 299 and M-Y, G. B., 65, 80, 492, 500 J. W., 299 BAXTER, BREGER,I. A., 21, 65, 80, 369, 377, 388 BEALES,F. W., 270, 271, 273,299, 311 J. M., 369, 384 BREMNER, H., 41 1 BECK,TH.and POSCHENRIBDER, J. M. and SHAW,K., 361, 384 BREMNER, H. and BUKATSCH, BRINKMANN, R., 134, 173 BECK,TH.,POSCHENRIEDER, F., 394,411 R., HEWING,M. BROEKER, W. S., GERARD, F. and BERG,G., 425,468 BEHREND, B. C., 497, 500 and HEEZEN, H. R., 260,276, 299 BELYBA, BRONGERSMA-SANDERS, M., 55, 80 BERG,G., 457,468 B. T., 376, 384 BROOKS, BERG,R. R., 64,80 C. W., 273, 300 BROWN, R. E. and TERRIERE, R. T., 299 BROWN, BERGENBACK, J. K. and HIRSCH,P. B., 403,411 N. and SCHEIN,G., 407,411 BERKOWITZ, J. S., 457,468 BROWN, BROWN,S. E. andTmYER, T. P., 119,120, I22 BERNARD,A., 433,442,452,468 S. W.,492, 500 BERNER,R. A., 35, 80 BRIJEVJCH, BERRY,F. A. F., 39, 80 S. V. and VINOGRADOVA, E. G., 6, BRIJEVJCH, J. and ANGOT,P., 457, 468 BICHELONNE, 9, 16, 300, 492, 500 D. E. andT~oMAs,W. H., BRYNER, BIEN,G. S., CONTOIS, L. C., BECK,J. V., DAVIS,I).B. and 36, 80, 328, 329, 338, 340 Wilson, D. G., 300 BIOOS,D. L., 34 BUBENICEK, L., 151, 173, 425, 450, 457, 459, BIRSE,D. J., 299 461,468 R. and WIESENEDER, BISCAYNE, P. E., 131, 132, 133, 134, 148, 149, BUCHTA,H., LEUTNER, 172 H., 300 BISSELL,H. J., 33, 56, 80, 89, 190, 222, 239, BURKHOLDER, P. R. and BURKHOLDER, L. M., 241, 299, 312, 313, 339, 340, 341 367, 384 B. H. and RILEY,CH. M., 96,122 BISSELL, H. J. and CHILINOAR, G. V., 284,287, BURMA, BURST,J. G., 147, 156, 173 293, 295, 297, 299 E. I., 457, 468 BUSHINSKIY, BLACK,W. W., 279,299 E., 40,80 BLACKWELDER, H. E., GIBSON,J. and RILEY,H. L., C A I L ~S., and KRAUT,F., 457, 468 BLAYDEN, 403,411 CALVIN,M., 357, 384 B ~ M ,L., EDELHAUSEN, L. and VON~ V E L E N , CMQBELL, C. V., 300 D. W., 398,411 CANAL, P., 300 BLUMER, M., 365,374,383 CAROZZI,A. V.,96,122,261,300,310,312,341 M. and OMENN,G. S., 365, 366,383 CAROZZI,A. V. and SODERMAN, BLUMER, J. G. W., 261, L. I. and BOWMOUIV,A. I., KHOTYNTSEVA, 300, 335, 341 PANINA, K. L.,376, 383 D., 99, 108, 122 CARROLL, BOQOMOLOV, A. I. and PANINA, K. I., 376,383 CARROLL, H. C., 142,143,173 D. and STARKEY, BOLT, G. H., 137, 165, 173 J. J. and GREENFIELD, L. J., 300 CARROLL, E., 47, 80, 135, 147, 148, 173 CASPERS,H., 33, 80 BONATTT, H., 240, 299, 448, 452, 457, 468 CASTENO,J. R. and GARRELS, R. M., 457,468 BORCHERT, BORCHERT,H. and M m , R. 0.. 30, 31,57, 80 CAWLEY, C. M. and KING,J. G., 361,384 BOSWELL,P. G. H., 64,80 CAYEUX, L.,59,61,62,64,75,77,80,285, 300, B ~ O M L E W. Y , B., 367, 383 457,468,469 BOTVINKINA, L. N., 245, 300 -A, s. K., 272, 300, 333, 341 V. S., 468 CHAVE,K. E., 38, 39, 65, 67, 80, 81, 152, 173, BA -, L. N. and YABLOKOV, BOWEN,H. J. M., 379, 383 279, 296, 300, 478, 500 M. A., 451,468 BRACONNIER, CHAVE,K. E., DEPPEYES,K. S., WEYL, P. K.,

527

REFBRENCES INDEX

GARRELS, R. M. and THOMPSON, M. E., 296, 300

MBOTAREV, I. I., 478, 500 ~HXLINGAR G. , V., 4, 16, 33,43,46,58,67,

74, 81,218,291,293,294,296,298,300,301,338, 341, 478, 481, 500 CHILINGAR, G. V. and BISSELL, H. J., 293,294, 301

CHILINGAR, G. V., BISSELL,H. J. and WOLP,K., 1, 523, 524

CHLINGAR, G. V. ~ ~ ~ D E GE. E T., N S478,500 , CHILINGAR,G. V. and KNIGHT,L., 8, 17, 164,

173 CIIILINGAR,G. V., RIEKB III, H. H. and ROBERTSON JR., J. O., 8, 17 CHILINGAR, G. V. and TERRY, R. D., 288,291, 301 CHOQUETTE, P. W., 337, 341 C H O Q U E ~P., W.and T u r n , J. D., 301 CISSARZ,A., 448, 452, 469 CLAYTON, R. N. and DEGENS, E. T., 301 CLAYTON, R. N. and EPSTEIN, S., 301 CLOUD JR., P. E., 45, 81, 185, 192, 193, 246, 298, 301, 319, 341 CONDON, M. A., 469 CONLEY, R. F. andBmY, W. M., 57, 81 CONWAY, D. and LIBBY, W. F., 350, 384 CONWAY, E. J., 46, 49, 52, 81 COOMBS, D. S., 3, 17 CooMBS, D. S., ELLIS,A. J., F m , W. S. and TAYLOR, A. M., 47, 81, 92, 122, 450, 469 COOPER, B. N., 68, 81 COOPER, J. E., 361,362,363,384,498,500 CQRB~ C., S., 453,469 CORNWALL,H. R., 450,469 CORRENS,C. W.,25, 45, 50, 70, 81, 118, 122, 262, 301, 327,425,457,469 CORRENS, c. and VON ENQELHARDT, 51, 81, 509 COU~SON,C.B., DAVIES, R. I. and KHAN,E.J. A.. 369, 384 COUTRY, G., 457,469 CRAIG, H., 297,301,379,380,384,495,496,500 CRICKMAY, G. W., 186, 189,217, 301 CRONOBLE, W. R. and MA", C. J., 281,301 CROOK, K. A. W.,112, 122,450,469 CROWLEY,M. S. and ROY,R., 166, 173 W G S , W. A., 112, 123 CURTIS, C.D. and KXUNGSL~Y, D., 301

w.

w.,

DAETWYL~R, C. C. and KIDWELL,A. L.,301 DAI~ME, A. and MACKOWSKY, M. TH.,401,411 DALY,R. A., 64, 81,90 DANCHEV, V. 1. and OL'KHA,V. V., 298, 301 D~AARD W., , 495, 500 DANGEARD, L. and RIOULT,M.,58,59,81

DAPPLES, E. C., 1,2, 32, 33,34, 36, 37, 61, 70,

71, 81, 97, 123, 161, 173,225,266, 301, 320, 329, 334, 337, 341, 453, 469, 523, 524 DAVIS,L. E., 492,493, 500 DAY,W. C. andEmMA~,J. G., 364, 375, 384 DEANS, T., 35, 81 DEBELMAS, J., 59,81 DEBYSER, J., 34, 81 DEER,W. A., Horn, R. A. and ZUSSMAN, J., 105, 123 DEFFEYES, K. S., 149, 173 DEFLANDRE, G., 59, 81 DEFORD, R. K., 260, 301, 311, 317 DEGENS, E. T., 2, 135, 173,217,224, 301, 302, 346, 368, 378, 384, 448, 451, 469, 524 DEOENS, E. T. and BAJOR,M., 353, 360, 384 DEGENS, E. T. and CHILINGAR, G. V., 2, 155, 173, 517, 524 DEGENS, E. T., CHILINGAR, G. V. and PIERCE, W. D., 384

DEGENS, E. T., EMERY, K. 0.and REVTER, J. H., 350, 351, 358, 371, 374, 384

DEGENS, E. T. and EPSTEIN, S., 61,67,81,296, 297, 302, 496, 500

DEGENS, E. T. and HUNT,J. M., 355,356,357, 370, 384

DEGENS, E. T., HUNT,J. M., REUTER, J. H. and REED,W. E., 353, 384, 496, 497, 498, 500 DEGENS, E. T., PRASEINOWSKY, A,, EMERY, K.0. and PIMENTA, J., 351, 352, 384 DEOENS,E. T., REUTER,J. H. and SHAW, K. N. F., 384 DELLWIG, L. F., 320, 341 DEMOLON, A. and BOISCHOT,P., 73, 81 DESL. U., 478,479, 500 DEVEIUN, L., 457,469 DIETRICH,R. V., HOBBS JR., C. R. R. and LOWRY, W. D., 302, 333, 341 DIXON,E. E. L. and VAUGHAN,A., 302, 318 DOBRYANSKIY, A. F., 376, 384 DONAHUE, J., 302 DOWNS, H. N. M., HUNTIENS, F. H. and VAN KREVELEN, D. W., 403,411 D o n JR., R. H., 336, 340, 341 DRAGUNOV, S. S., 384 DROOP,M. R., 367,385 DROSTE, J. B., 146, 173 DRYDEN, I. G. C., 403,407,411 DRYDEN, I. G. C.and G ~ m r mM., , 403,411 DULHUNTY, 3. A., 403,407,411 DUNHANI, K. C., 35, 81,457,469 DUNHANI, R. J., 219,302, 312, 315, 317 DUNNING, M. L. and HUNT,J. M., 365, 385 DUNNINGHAM, A. G., 407,411 DUNOYER DE SEOONZAC, G., 156, 173 Du~ov,S. A., 478, 500

528 EADES, J. L., and Gw, R. E., 105, 123 EARDLBY, A. J., 246, 271, 302 ECKELMAN, W. R., BROEKER, W. S., WHITLOCK, D. W. and ALLSUP, J. R., 379,380,385 ECKHARDT, F. J., 157, 158, 159, 173 EDIE,R. W., 261, 274, 302 EDWARDS, A. B., 407,411 EDWARDS. A. B. and BAKER,G., 40,54,81 EHHALT,D., KNW, K., NAGEL,J. F. and VOGEL,J. C.,497, 500 E m , W., 36,43,47, 82 EKIERT, F., 469 EL BAZ,F., 431,469 ELBAZ,F. and AMSTUTZ, G. C.,447,453,469 E m , G. K.,302 ELLIS,C. B , 492 EMERY, K.O., 302 EMERY, K.0. and HOGOAN, D., 445,453,469 EMERY,K. 0. and RIITBNBERo, S. C.,9, 17,

REFERENcm INDW

FORSMAN, J. P. and HUNT,J. M., 371,372,385 FO~RGIL C.LA., , 517 FOTIEV, A. V., 11 FOWLER, W. A., GREENSTEIN, J. L. and H o w , F., 378,385 Fox, D. L., 364, 385 Fox, D. L., UPDBGRAFF,D. M. and NOVELLI, D. G., 364, 385 Fox, W. T., 336, 341 FRANCIS,W., 363, 372, 374, 385 FRANKS, P. C. and SWINEFORD, A,, 337, 341 FREDMAN,G. M., 102,302,451,470,495,501 FROST, A. V., 385 FUCHTBAUER, H., 303, 430, 451,470,506,510, 518

FCTCHTBAUER, H. and GOLDSCHMIDT, H., 145,

153, 158, 160, 168, 169, 173 FUCHTBAUER,H. and -NECK, H. E., 138,173 FYI%,W. S. and BISCHOPP, J. L., 303 FYI%,W.S.,TURNER, P. J. and VERHOOGEN, J., 3, 17,47,82, 123

136, 151, 152, 163, 173, 302, 329, 347, 385, 492,501 EMERY, K. 0.. TRACEY JR., J. 1. and LADD, M. S., 189, 193, 216, 217, 302 GANSSER, A., 445, 453. 470 ENTWISTLE, L. P. and GOUIN, L. O., 452, 469 GARLICK,W. G., 452, 470 EPPRECHT, W., 452,469 GARRBLS, R M., 25, 26, 27,28, 34,67,78, 82, 99,123,450, 461,470 EPSIE~, S,. 495,501 EPSTEIN, S., GRAF,D. L. and D ~ E N S E., T., GARREIS, R M. and DRBYER, R. M., 303 GARREIS, R. M., DREYER,R. M. and HOW296,297, 302 EPSTEIN, S. and MAYEDA,T.,495, 501 LAND,A. L., 303 EREMBNKO, N. A., 478,481,497, 501 GARREIS,R. M., THOMPSON,M. E. and SEVER, R., 82 ERDMAN,J. G., 364, 374, 375, 377, 378, 385 EIUURT,H., 53, 77, 82 GASKELL, T.F., 288,289 ERICKSON, A. J., 302 GEDENK, R., 409,411 EUQSTER, H. P. and SMTH, G. I., 453,469 GEDENK, R., HEDEMANN,H. A. and RWL, W., EVANS, J. W., 302 409,412 GL?HMAN JR., H. M., 344,385 FAIRBAIRN, H. W., 72,82 GS~CIB, A., 82, 88,89 GEORGE, T.N., 271, 303 F A I R B ~R. EW , .,1,2,29,38,45, 51, 53,57, 58, 61, 67, 68, 69, 71, 78, 82, 88, 89, 147, GERMANOV, A. I., 448,451,470 Gms, R. J., 336, 341 151, 173,296, 302, 523, 524 GIDON,P.,59, 82 FAUST, G. T.,69, 82 FEDIUK, F., 435,451 GIGNOUX,M. and AVNIMELECH, M., 59, 82 FEELY, H. W. and KULP,J. L., 56,82 GILL, W. D. and KUENEN, PH. H., 445, 453 470 FERAY, D. E., HE=, E. and HEWAIT,W. G., GIMMELFARB, B. M., 73, 82 302 FERSMAN, A. E., 20,33. 82.89 GINSBURO, R.N., 180, 184,188,190,191,193. 211,266,283,284, 303, 311 FINCH,W.I., 470 G ~ T ~RRL., , 55,82 F~~CHER,R., 446 GLOSSARY OF GBOLOOY AND RELATED SCIENCES, FIS~~ER, W.,424,470 17 MRKE, 0. W., 337,341 FL~~E E.Land , FLUGE~KAEUX, E., 302 GLOVER, J. E., 38,82, 106, 123,303, 334, 341, 441,470 FOLK, R. L., 182, 185,218,219,220,221,222, 227, 233, 246, 256, 258, 259, 269, 272, 302, GOLDBERG, E. D., 345,385 GOLDBERG,E. D. and ARRHBNIUS, 0. 0. S., 311,312, 316, 317. 319 FOLH,R. L. and WEAVER,C. E.,103, 123 47, 52, 82 Gomrcx, S. S.,2,17,64,82,112,123 FORMOSOVA, L. H., 457,470

529

REFERENCES INDEX

V. M., 23,82 J. R. and GRAF,D. L., 67, 82,

GOLDSCHMIDT,

GQLDSWIX,

106, 123

GOLDSMITH,J. R.,G w , D. L. and WITTEW, 3.. 106, 123

GOLDSTEIN JR., A., 331, 341 GOLDSTEIN JR., A. and HENDRICKS, T. A., 332, 341

GORRELL, H. A., 478,501 G~RSLINE, D. S., 303 GQTHAN,W., 359, 385 GOTO,M., 188,303 GRABAU, A. W.,20,21,57,83,303,312,424,470 GRAF,D. L., 303 GRAF,D. L. and GOLDSMITH,J. R., 67,83 GRAF,D. L. and LAMAR, J. E., 215,303 GRAENICHER, H. and JONA, F., 453, 470 GRAHAM, E. R., 43, 83 GRAVELAND, D. N. and LYNCH,D. L., 361,385 GREENLAND,D. J., 359,385 GREENSMITH, J. T., 100,106,123,266,274,303 GREINER, H. R.,280, 303, 315 G m , G. M. and INGRAM, R.L., 49, 83 GRIPFIN, J. J. and GOLDBERG, E. D., 131, 132, 173

GRIFFIN,R. H., 303 GRIM, R. E., 48,49, 52, 83, 141, 158, 173 GRIM,R.E., BRADLEY, W. F. and WHITE,W. A., 170, 173

GRIM, R.E., BRAY,R.H. and BRADLBY, W. F., 83, 147, 174

G R ~R. , E.,DIETZ, R.S. and BRADLEY, W. F., 50, 51, 83, 145, 174

GRIM, R.E., DRWTB, J. B. andBRADui~,W. P., 146,174

HAMUTON, E. L., 130, 174 w.E., 364, 377, 385 HARBAUGH, J. W., 221,227,234, 304 HARDBR, H., 457,470 H-, P. E., 349, 350, 385 HARMS,J. E., WHITEHEAD, T. H. and HEATON, J. B., 470 HARRIS, L. D., 341 HARRINGTON, J. S., 353, 385 HARRINGToN, J. S. and CILLIERS,J. J. LER., HANSON,

353,385

F.H., RASTAU, R. H. and BLACK,M., 69, 75, 77, 83 HATHAWAY, J. c. a d SACHS, P. L., 146, 174 HAUSSOHL, S. and M ~ L E RG., , 153, 174 HAY,R.L., 47,83, 106, 123 HAYASHI,T. and NAGAX, T., 369,385 HAYES, A. O., 77, 83 IIAYES,J. B., 335,341 HEALD,M. T., 72, 83,94,96, 123 HATCH,

HEALD,M. T., THOMSON, A. and WILCOX, F. B., 341 HEcm, F., 34, 83 HEDBERO,H. D., 166,174 HEDEMANN, H.A., 410,412 HEDEMA",H. A. and T E I C ~ ~ L LRE .,R412 , HEER,I. C.,470 HEEZEN, B. C.,NESTEROFT, W. D. and SABATIER, G., 49, 50, 51, 83 HEW,A., 419,470 HEIM,D., 160,174 HELLINO, D., 503, 504,507 HEM,J. D., 326,328,341 HENDERSON,D. M. and RHODES,F. H. T., 304 HENDERSON, M. E. K., 371,385 HENDNCKS,S. B. and Ross, C. S., 147, 174 HENSON, F. R. S., 276,304 HERRMA",V. A. and RICJTTER-BERNBURG, G.,

GRIM,R.E. and JOHNS, W. D., 142, 145, 173 GRIM,R. E. and L~UGHNAN, F. C., 145, 173 GRIMM, W. D., 430,470 GROENNINGS, S,. 365,385 470 GRUN%R, J. W. and TIIIBL, G. A., 64,83 HFSS,K. and KOMARBWSKY, A., 359, 385 GRUSZCZYK,H., 448,452,470 IIINDMAN, J. c.,190 GRUSZCZYK,H. and OSTROWXCKI, B., 470 Hmm, L. F., 336, 341 GRUSZCZYK, H. and WAZBWSKA-RESEN-HODOSON, G. W. and BAKER, B. L., 365, 386 KADIPF, W., 470 HODQSON, G. W. and HITCHON, B., 366, 386 HODOSON,G. W.. IIITCHON, B., ELOFSON, HAMWA", E.,419,443,470 R. M., BAKER, B. L. and PEAKE, E., 365,366, HADDING, A., 216,277,303 386 HAQNI,R.D. and GRAWE, 0. R.,452 HOEHNE, K.,71, 83 HAIDINOJ3R, w., 67,83 HOFMANN, U. and HAUSDORF, A.. 139, 174 HAL.BACH, P., 303 H~I~LINO, H. J., 145,174 HAL.LA, F., C I ~ ~ O A G,R V.,and BISSBLL, H. J., HOHL.T,R.B., 304 293,294, 303 HOLMIS,A., 304,313 IIALLIMOND, A. F., 457,470 HOUER,W. T., 66,83 HAM,W. E.,303 H o r n , A. P. and JEPFRIES,C. D., 304 HAM,W. E. and PRAY,L. C., 257, 303 HONJO,S., FISCHER, A. G. and GARRISON, R., HAMBLBTON, A. W., 261,267,304,319 304,453

530

REFERENCES INDEX

KAZAKOV, A. V., 456,471 K~AKOV A., V., TIKHOMIROVA, M. M. and V. J., 471 PLOTNIKOVA, KEITH,M. L. andD~oENs,E. T., 366,386 KELLER,W. D., 42, 50, 83, 101, 123, 141, 142, 156, 174 KELLER,W. D. and REESMAN, A. L., 103, 123, 328, 341 KELLER, W. D., B u r n , W. D. and REESMAN, A. L.,84, 327, 328, 341 KELLEY,W.P., 43, 50, 84 KENDALL, D. L., 471 KESSLER, P., 41, 84 KHVOROVA, I. V., 290, 291, 292, 293, 304 KING,J. G. and WILKINS,B. T., 403, 412 KITANO, Y. and HOOD,D. W., 304 KLEBER, W., 304 K m , G. DEV., 99, 123 ILLING, L. V., 68, 83, 188, 270, 271, 274, 276, KLUBOVA, T. T., 7, 8, 17 304, 311, 317, 431, 453,471 KNETSCH,G., SHATA,A., DEGENS, E. T., INGELS, J. J. P., 336, 341 M ~ N I C HK., O., VOGEL,J. C. and SHAZLY, INGERSON, E., 304 M. M., 496, 501 INTERNATTONAL COMMITIXE FOR COALPETRO- Kom, H. W., 471 KOLBE, H., 452,457,471 LOGY, 392,400,412,414 IRVING, A., 53, 66, 83 KONONOVA, M. M., 368, 369, 371, 372, 386 IRVING, R. D. and VANHISE,C. R., 121, 123 KONSTANTINOV, M. M., 452,471 IRW, M. L., 304 KORNICKER, L. S. and PURDY,E. G., 275, 304 KWOVSKAYA, A. G., 156, 174 JAANUSSON, V., 61, 83, 186, 210, 245, 304 KOSSOVSKAYA, A. G., D m , V. A. and JACOB, H., 452, 471 ALBXANDROVA, V. A., 157, 174 JAMES, H. L., 452,455,471 KRAUSKOPF, K. B., 43, 70, 84, 116, 123, 325, JXNECKE, E., 31,83 327, 328, 329, 342, 366, 386 JANROWSKI, G. L. and =BELL, C., 304 KREJCI-GRAF,K., 65, 84, 381, 386 JEFFREY, P. M., COMPTON,W., GREENI~ALQH, KREJCI-GRAF, K. and WICKMAN,F. E., 379, D. and DELAETER, J., 379, 386 380, 386 A., 424,471 KREJCI-GRAP, K., HECHT,F. and PALSER, W., JOHANNSEN, JOHNS, W. D., 144, 174, 451, 453, 471 478, 501 JOHNS,W. D. and GRIM,R. E., 174 KROMER, H., 158, 174, 516 JOHNSON, J. H., 199, 203, 204 KROTOV,B. P., 46, 456, 457, 471 JONES,J. D. and VALLENTYNE, J. R., 350, 386 KRUMBEIN,W. C., 180, 181, 184,286, 304 JOULIA, F., BONIFAS, M., CAMEZ,T., MILLOT, KRUMBEIN, W. C. and GARRELS,R. M., 22, G. and W E I ~R., , 148, 174 41, 84, 187. 264, 304 KRYNINE, P. D., 105, 123 JUNG,C. G., 454,471 J~~TGE H.Nand , KARWEIL,J., 406, 409, 412 KRWROV,P. A., 516 KRYUKOV,P. A., ZHUCHKOVA, A. A. and KAARSBERO, E. A., 170, 174 RENOARTEN, E. V., 490, 501 KAHLE,C. F., 304 PH.H. and PERDOK, W. G., 116, 123 KUENEN, KUKAL,A., 304 KALKOWSKY, E., 21, 83 KAPLAN,J. R., DEGENS, E. T. and REUTER, KURBATOV, J. M., 395,412 J. H., 356, 357, 360, 373, 378,386 KWL, 0. S. and PATIJN,R. J. H., 407, 412 KARCZ, I., 304 K u z ~ ~ ~ sS. ov I.,, 451, 471 KARWEIL,J., 408,412 KAY,M., 38, 83 LALOU,C., 246,261,276, 304 LAMAR,J. E., 334, 342 &YE, C. A., 186, 188, 189, 190,200,210,246, LAMBE, T. W., 137, 174 304 KAYSER, H., 412, 396 LAMORT, C., 366, 386

HORSTMAN, E. L., 64, 83 Hoss, H., 58,83 HOUGH, J. L., 77, 83,452,471 HOWER, J., 456, 471 HSI, H. R. and CL~TON, D. F., 139, 174 HUBBERT, K., 513, 515 HUCK,G. and KARWEIL,J., 403,406,407,412 HUCK,G. and PA~EISKY, K., 406,412 HUCKENHOLZ, H. G., 122,123 HUMMEL, K., 41, 44,83, 89, 129, 174 HUNT,J. M., 344, 355, 364, 375, 377, 379, 386 HUNT,J. M. and JAMIESON, G. W., 375, 386 HURLEY, P. M., HEEZEN, B. C., PINSON, W. H., and FAIRBAIRN, H. W., 49, 83 HUTCHINSON, G. E., 367, 386 HUTCHINSON, G. E. and SETLOW, J. K., 364, 386

REFERENCES INDEX

LANDERGREN, S., 379, 386 LANDES,K. K., 304 LANE,A. C., 30, 37, 38, 84 LA PORT&L. F., 267,304 LARSEN,G. and CH~LINGAR, G. V., 284, 305, 523 LATIMER,W. M., 25, 84 LEES,A., 305 LEES, G. M., 60, 84 LEGATE, C.E. and JOHNS, W. D., 451,471 LEIGHTON, M. W. and PENDEXTER, C., 260,269, 305, 313, 317, 320 LEIIVZ,V., 133, 174 LEITH, C. K. and MEAD, W. J., 84, 89 LEMARRE, H., 59,84 LENSCH,G., 410,412 LEISTWEIN, F., 448,471 LEWIN,J. C.,58, 84 LIBBY, W. F., 497, 501 LICHTENBERG, K., 409, 412 LIBBENBERG, W. R., 435,452,453,471 LIESEGANG, R. E., 56 LD-LY,V. G. and LEONIAN, L. H., 367, 386 LINCK,G., 67,84 LINDBLOM, G., 350, 363, 386 LINDBLOM, G. and LUPTON,M. D., 347, 350, 363, 386 LINDGMN,W., 66,419,424,471 LINLJSTR~M, M., 210, 305 LINSCOTT, R. O., 305 LIPMAN,C. B., 347, 386 LIPPMANN, F., 154, 174 LISICYN,A. P., 471 LOCHTE,H. L., 367, 385 LOGAN, B. W., 245. 305 LOMBARD, A., 452,471 LOMBARD,J. and NICOLINI, A. P., 452, 471 L~MTADZE, V. D., 8, 130, 174,478, 501 Low, L. G., 54, 84, 152, 175,453,466,471 LOVERMG, T. G. and PATTEN, L. E., 330, 342 LOWENSTAM, H. A., 105, 123, 216, 217, 223, 260, 267, 271, 277, 279, 296, 305, 319, 320, 451, 452,471 H. A. and EPSTEIN, S., 260, 271, LOWENSTAM, 305 LUCAS,G., 68, 84 LUCIA,F. J., 215,233,268, 269,288, 305, 316, 319,322 LYELL,C., 84, 89,419 LYNCH, D. L., WRIGHT, L. M., HEARNS, E. E. and CQTNOIC,L. J., 359, 386 LYNCH,D. L., WRIGHT, L. M. and OLNEY, H. O., 369, 386 MACDONALD, G. J. F., 57, 84 MACFADYEN, W. A., 57,84

531 MACKOWSKY, M. T., 452,472 MACQUAR, J. C. and TREUIL,M., 452,472 MAG"IS, A., 419 MAGNUSSON, H. H., 424,472 MARCHER, M. V., 305 MARKEWCH, V. P.,448,452,472 MARSHALL, C.E., 452,472 MARSHALL, C. E. and UPCHURCH,W. J., 43,84 MA-, S. M., 440,453,472 MARTIN,R. T., 137, 175 MATSON, F. W., 423,472 MAUCHER, A., 442,452,472 AXWE WELL, W. G. H., 189, 305 MAXWELL, W. G. H., DAY,R. W. and FLEMING, P. J. G., 305 MAXWELL, W. G. H., JELL,J. S. and MCKELLAR, R. G., 305 MAYER,F. K., 216, 305 MCIVER,R. D., COONS, C. B., DENEKAS, M. O., and JAMIESON, G. W., 375,387 MCKEE,E. D., CHRONIC, J. and LEOPOLD, E. B. 305 MCKELWYJR., J. G. and MILNE,I. H., 492, 501 MCKELVEYJR., J. G., SPIEGLER, K. S. and W n m , M. R. J., 492, 501 MCKELVEY,V. E., SWANSON,R. W. and SHELDON,R. P., 75,84 MCKXNSTRY, H. E., 424,472 M c L m , J. F. and SCHLANGER, S. O., 305 MCNAB,J. G., SMITHJR., P. V. and BEITS,R. L., 376, 387 MEADE,R. H., 136, 137, 139, 164, 165, 166, 169, 170, 175, 478, 501 MEENTS,W. F., BELL, A. H., REES,0. W. and TILBURY, W. G., 116, 123 MEHTA,N. C.,DUBACH,P. and DEUEL,H., 359, 387 MEINSCKEIN, G. W., 364, 374, 375, 378, 387 MEINSCHEIN, G. W. and KENNY,G. S., 363, 364, 387 MEKHTIEV, S. F., 487, 501 MENNING,J. J. and VATAN,A., 68,84 MERO,J. L., 151, 175, 447, 448, 452, 453, 472 M I ~ LG., and R~~LLER, K. H., 517 MIDDLETON, G. V., 305 MILLERJR.,D. N., 100,124,190,283,305 MILLERJR., D. N. and FOLK, R. L., 101, 123 MILLER,L. J., 448,452,472 MIL-, S. L., 357, 387 MILLER,S. L. and UREY,H. C., 387 MIL.LOT,G., 21, 48, 51, 52, 84, 146, 147, 158, 175 MILNE,I. H. and EARLEY, J. W., 48, 84 MITCHELL,J. K., 137, 165, 169, 175 MOBERLEY, R., 134, 175

532

RBFERENCES INDEX

Mom, P. A., 452,472 MOLLAZAL, Y., 260,305 MONAOHAN, P. H. and LYTLE,M. L., 246,276, 305

MONOMAKHOFF, C.,407,412 MOORE,E. S. and MAYNARD, J. E., 75, 77, 84 MOORE,R.C.,269, 305 M o m , L., 305 MOREIT, F. J., 68, 84,264, 305 MORRIS, R. C.and DICKEY,P. A., 306 MORRISON,R. I., 371, 387 MORTEANI, G., 306 MOSEBACH, R.,55, 84 MOULTON,G. F., 53, 85 MULE, J. D. and ERDMAN, J. G., 364,375,387 MULLER,G., 1, 135, 136, 141, 145, 148, 149, 150, 153, 175, 523, 524

MULLER,W., 306 MUNOAN,N. and JESSEN,F. W., 139, 175 MIINKIcH,K. O., 497, 501 MURAWSKI, H., 424,472 MURRAY, G. E., 306, 313 MURRAY, H. H. and HARRISON, J. L., 153, 175 MURRAY, J. and RENARD, A. F., 45,47,52,85 MURRAY, R. C., 215,282, 306

OGNIBEN,L., 57, 85, 306 D. E.,339, 342 OJAKANGAS, R. W., 108,116, 124 OJAKANQAS, R. W. and KBLLER,W. D., 46,85 O K A M G., ~ , OKURA, T. and GOTO,K., 71,85 OPPENHELMER, C. H., 347, 387 ORME,G. R. andBRow, W. W. M., 219,221, WREN,

227,234, 306

0 ~ 6J., , 357, 387 0 ~ 6J., and KIMBALL, A. P., 357, 387 ORR,W. L., EMERY,K. 0.and GRADY, J. R., 365,366, 387

ORR,W. L. and GRADY, J. R.,365,387 ORR, A. P. and MOORHOUSE, F. W., 34, 85 OSMOND, J. C.,260,306 OSTROIJMOV, E. A. and SHILOV, V. M., 452,472 OTTE, C.and PARKS, J. M., 210, 306 OVERBEEK, J. TH.G., 137, 175

PACKHAM, G. H. and CROOK, K. A. W., 45,47, 85,92, 97,99, 124, 129, 175

PALACAS, J. G., 360, 387 PALM, C.and Calvin, M., 357, 387 PA^, H. M., 306 PARK, R. a n d D m m m , H. N., 379, 387 PARK,R. and E m m , S., 379, 380, 387 NAOY,B., 39, 65, 79, 85 PARK, W.C.,432,452,472 NAOY,B. and BITZ, M. C., 361, 364, 387 PARK,W. C.and AMsTuTz, G. C.,432,452,472 NAGY, B., MEINSWEIN,G. W. and HEN- PARKINSON,D., 278, 306,320 NBSSY, D. J., 378, 389 PATTEBKY, K., 409,410,412 NANZJR., R. H.,49, 85, 124 PATTEBKY,K. and "EI-LER, M., 392,400, NAUMANN, C.F., 66, 85 404,405,412 NBHER,J. and ROHRER, E., 347, 387 PATIERSON,M. N. A., BIEN,G. S.and BERNER, NELSON,H. F., BROWN,C. W. and BRINER. A., 67, 85 MAN, J. H., 308, 311, 319, 320 PAUL, E. A. and SCHMIDT,E. L., 352,387 NEWELL,N. D., 57, 85, 180,278,306,320 PAULSON,P. L., 448, 451, 472 NEWELL,N. D. and RIGBY,J. K., 275,306 PELTRO,C. R., 334, 342 NEWELL, N. D., R~GBY, J. K., FISCHER, A. G., PERKINS, R. D., 283, 306, 311 W ~ M A NA., J., HICKOX, J. E. and PETERSON, J. A. and OHLEN,H. R.,306 BRADLEY, J. S., 222,241,244,247,277,278, PETERSON, M. N. and GOLDBERG, E. D., 133, 285, 306, 330, 334, 335, 336, 342

NEWELL, N. D., RIQBY,J. K., WHITEMAN, A. J. and BRADLEY, J. S., 306 NBWHOUSB,W. H.,54, 85 NEWLANDS, D. R.and T Y R D. ~ S,., 435, 452,472

NICHOLLS,G. D., 448,452,472 NICOLINI, P., 472 NIOOLI, P., 426 NIOOLI,P. and NIQOLI,E., 187, 306 NITECKI,M. H., 263,264,306 NODDACK, I. and NODPACK, W., 448,451,472 NORIN, E.,47, 85, 148, 153, 175 NORTHROP, J. I., 59, 85 OpreDAHL,

c.,448,452,472

175

P E ~ K J., 452,451,472 , PETRASCHJXX, W.E., 407, 412, 424, 473 PETITJOHN, F. J., 20, 21, 22,41,44, 53, 64, 85,

88, 89,91, 111, 118, 124,215, 306, 312, 313, 316, 321,464,473 PPEIFPER, J. P and SAAL, R. N. J., 379, 388 PHILCOX, M. E., 306 PHILLIPe,W.. 506, 518, 521 PHILLIPP, W., DRONG, H. J., FUCWBAUER, H.,

HADDENHORST, H. G. and JANKOWSKY, W., 521

PIA, J.. 306 PIEKARSKI,K., 452,473 hovLewsKarA, L. v. and 396,412

RAKOWSKI,

v. E.,

REPBRBNCES

INDEX

PINSAK, A. P. and MURRAY, H. H., 145, 175 PI'ITMAN JR.,J. s., 342 POPOV,B. P.,457,473 PORTER, J. R., 347,388

533

352, 353, 355, 358, 361, 366, 388, 478,492, 501 &"ENHOUSE, G., 452,453,473 RMARE,A. and VERNHET,S.,307 -RE, A. and VISSE,L., 73,85 PoZIEpNk,F.,419,422,473 P ~ RP. E ,. and PBITUOHN, F. J., 435,452, ROLL,A.,510 Rom, I. I., 5, 17, 307 473 ROZENKRANTZ, A. and RASMUSSEN, H. W., P m , R. H., 141, 142,175 60,85 POURBAIX, M. J. N., 27, 85 I. TH.,138, 176 POUSTOVALOV(PUSTOWAUIFF), L. V., 239,448, ROSENQUIST, Ross, C.S., 50, 85 451,457,473 R m , C. S. and HENDRICKS, S. B., 159, 176 POWERS, M. C.,145, 155, 175 Ross, C.S. and KERR,P. F., 51, 85 POWERS, R. W.,306 PRASHNOWSKY, A., DEOENS, E. T., EMERY, Ross, C.S. and OANA,S.,224, 307 ROSSINI, F.D., 345, 374,388 K. 0. and PIMENTA, J. 358, 367, 388 ROULET, M. A. and SCHOPFER, W.H.,367,388 PRAY, L.C.,261,279,306 ROUTHIER, P., 424,450,452,453,473 PRAY, L.C.and MURRAY, R. C.,17, 523 RURHIN(Rum), L. B.,8, 17, 163, 176,238, PRAY, L. C.and WRAY,J. L., 306 239, 307,523 h O K O p o v I c H , N.,342 RUSNAK, G. A., 104, 124,276,307 PUCHELT, H., 517 RUSSELL, R.J., 189,307 P~CHEL.T, H.and M ~ L E RQ. , ,452,4731 RUSSELL,R. J. and RUSSELL, R. D., 58,85 PURDY, E.Q.,193, 306 RUTTEN, M. G.,55,60,61,62,86 RADCZEWSKI, 0.E., 133, 176 RAOUIN,E.,424,473 SAFIINS JR.,F. F., 307 RAINWATER, F.H.and WHITE,W.F., 478,501 SAC-, W.M., 307 RAKOWSKX, W.,BATURO,W. and PIGULEWS- SAOAN,C.,345,388 SALVAN, H.M., 75,86 KAJA, L., 395, 396, 412 SMONOWCZ, J., 58,86 RAMBERO, H., 463,473 SAMUELS, S. Q.,137, 176 RAMDOHR, P.,435,451,452,453,473' SANDER, B., 69, 86,427,237, 307,446,473 my H.H.and WATSON, J., 128, 176 SAW, D. D.,68,86,307 REED, R. D., 64,85 SAUER, R. W.,MELPOLDER, F. W.and BROWN, RENTZSCH,J., 433,473 R. A., 367,388 REU"ER, H., 411 S. M. and STEVENSON, F.J., 369,388 SAVAOE, REVI?LLE, R., 307 SAVEL'EV (SAWeLFJFJW), I. v., 163 REVEL- R. and EMERY,K. O., 62,85 D., 427,473 REVELLE,R. and FAIRBRIWE, R., 190, 193, SCHACHNBR-KORN, SCIIBIDEWER, A. E., 39, 86 194, 210, 211, 307, 318 SCHLANOER, S. O.,68, 86, 279, 280, 307, 310 REVELLE, R. and FLEMMINO, R. H., 307 REX,R. W.and GOLDBERO, E. D., 133, 176 SCHMACZ, R. F.,68,86,307 SCW~DT, E. L. and STARKIW, R. L., 367, 388 REYNOLDS, D. L.,64,85 S m , H.,38, 86, 450,473 &a,M.,307 SCANEIDER, H. J., 430,446,450,452,473 RICHTER-BBRNBURO, Q., 452,473 SCHNEIDER, S., 412 RIEDEL,W.R.,36, 58,85, 342 RIEDEL, W. R., LADD,H. S., TRACEY JR., SCANEIDERH~HN.H.,424, 443, 448, 452, 473 SCHERP,A.,410,412,508,520 J. 1. and BRAMLETTE, M. N., 353,388 M. andTRvRNIT, P.,452,453,473 RIEKE III, H. H., CHIUNGAR,G. V. and Scrimmws~~, ROBERTSON JR.,J. O., 17,478.488,489,490, SCBRBINER, 0.and LATHROP,E. C.,361, 365, 367,388 501, 516 ~CEIREINER,0.and ~HOREY,E. c., 367, 388 Rmm,C. M. and BYRNE,J. V., 57, 85 SCHtjLLw, A,, 391, 401, 413, 452, 473 RUWE,F., 31, 85 -ERG, S. c.,347,588 SCHULZ,O.,442,446,473 , 448, 551. 473 R ~ ~ ~ E N B ES.R C O., , EMERY,K. O., H ~ E -S ~ A R T B D., MA", J., Dmm, E. T., FAY,R. C., SCEIWARZACHW, W.,246, 280, 307, 321 REUTW,J. H., GRADY,J. R., RICHARDSON, SCHWENDMGER, R. B. and ERDMAN, J. G., S. H. and BRAY,E. E., 176, 295, 307, 347, 364, 388

534 SEIBOLD, E., 136, 147, 154, 155, 176, 296, 307, 464,473 SEIDL,K. 448,451,473 SELLERS, G., 355, 388 SELLEY,R. C., 446, 453,474 SHEARMAN, D. J., KHovru, J. and TAHA, S., 69, 78, 86 SHELTON, J. W., 51, 86 SHEPARD,F. P. and MOORE,D. G., 136, 176 SHISHKINA, 0.v., 448,451, 474,490,491, 501 SHOREY, E. c.,367, 388 SHORLAND, F. B., 388 SHROCK,B. R., 435,474 SHUMWAY, G., 176 SIFVETSOV, M. s., 17 SIDWELL, R. and WARN,G. F., 307 SIEGEL,F. R., 223,263, 307 SIEMENS, H., 427, 474 SEVER, R., 25, 32, 58, 70, 71, 86, 95, 116, 118, 124, 325, 326, 328, 329, 342, 478, 490, 491, 501 SIEVER,R., BECK,K. C. and BERNER, R. A., 151, 152, 176, 307 SIEVER,R., GARREU,R. M., KANWISHER, J. R. A., 490,491,501 and BERNER, SILLIKER, J. H. and RITTENBERG, S.C., 363,383 SILVERMAN, S. R., 376, 378, 379, 380, 383 SILVERMAN, S. R. and EPSTEIN,S., 380, 388 SIMPSON, G. G., 423,474 SINGEWALD, J. T., and MILTON,C., 307 SISLER, F. D., 347, 388 SKEATS, E. W., 307 SKEMPTON, A. W., 164, 176 SKINNER, H. C. W., 67, 86 SKINNER, H. C. W., SKINNER,B. J. and RUBIN,M., 307 SLOSS, L. L. and FERAY, D. E., 96, 124 SLOSS,L. L. and LAIRD, W. M.. 307, 315 SLOWEY, J. F., JEFFREY,L. M. and HOOD, D. W., 361, 389 SMIRNOW, L. P., 33, 86 S m H , H. M., 374, 389 S m , P. V., 374, 389 SMULIKOWSKI, K., 46,86 SOKOLOV, v. A., 375, 389 SORBY, H. C., 67, 86 SORENSEN, H., 359,360,389 SPENCER, E., 64,86 STACH,E., 401,413,474 STACH,H., 397,413 STADLER,G., 410,413 STANDARD, J., 196,230 STANTON, R. L., 447,452,474 STANTON, R. L. and BAASBECKING,L. G. M., 474 STARIKOVA, N. D., 448,451,474

REFERENCES INDEX

STAUDINGER, H. and JURISCHI., 359, 389 STAUFFER, K. W., 263, 284, 298, 308, 318 STEHLI,F. G. and HOWER,J., 217,222,223,308 STEINBRECHER, B., 452, 474 STERN,K. H., 56,86 STETSON, H. C., 56, 86 STEVENS, R. E. and CARRON, M. K., 42,86 STEVENSON, F. J., 36, 86, 350, 361, 389 STEWART, F. H., 70, 86,453,474 STOPES, M. C., 413 STORER, A., 166, 176 STRAKHOV(STRACHOV), N. M., 3, 4, 5, 7, 17, 67, 86, 131, 176, 239, 248, 284, 285, 291, 293, 294, 295, 296, 308, 435, 448, 451, 452, 474 STRAKHOV,N. M., BRODSKAYA, N. G., KNYAZEVA, L. M., RAZZHIVINA, A. N., RATEEV,M. A., SAPOZHNIKOV,D. G. and E. S., 4, 5, 6, 7, 17,285, 308 SHISHOVA, STRING-, B., 64,86,97, 124 SUGAWARA, K., 56, 86 SUGDEN, W., 308 SUIKOWSKI,ZB. L., 1, 3, 17, 37,42, 60,61, 86, 263, 285, 289, 308, 452, 453, 474 SULIN,V. A., 478, 479, 501 SUN, MING-SHANand ALLEN,J. E., 64, 87 SWAIN,F. M., 353, 389 SWAIN,F. M., BLUMENTALS, A. and PROKOPOWCH, N., 353, 389 SWINCHATT, J. P., 308 SWINEFORD, A., 99, 124 SWINEFORD, A. and FRANKS, P. C., 102, 124 T m , W. H., 217, 308 TALIAPERRO,N. L., 59, 87, 331, 337, 342 TAMM,O., 42,43, 87 TAN,T.K., 138, 176 TANAKA,J. and KUWATA, T., 361, 389 TARR,W. A., 54, 87, 308 TAIJPITZ, K. CH.,449,450, 474 TAYLOR, J., 93, 124 TAYLOR, J. H., 78,128, 155, 160,176,452,457, 474 TEICHM~LER, M., 397,409,413,452,474 TEICHM~~LER, M. and TEICHM~~LER, R., 1, 369, 389, 397,407,413,474, 524 TEIC-LER, R., 406, 407, 413 TEODOROVICH, G. I., 13, 16, 17, 21, 41, 45, 67, 74, 75, 76, 78, 87, 180, 289, 291, 293, 308, 435,451,457,474 TERMIER,H. and TERMER, G., 53, 58, 72, 87, 192, 216, 244, 246, 308 TERS,M., 59,87 TERZAGHI, R. D., 186, 308 TESTER, A. C. and ATWATER, G. I., 21,87.308 WANDER, O., 359, 389

REFERENCES INDEX

THIESSEN,G., 452,474 THODE, H. G., HARRISON, A. G. and MONSTER, J., 55, 87 THOMAS, D. W. and BLUMER, M., 453,474 THOMA& G.E., 308, 314 TXOMAS, G. E. and GLAISTER, R. P., 261,262, 276,282, 308, 321 THOMSON, A., 72,73,87,96,107,124, 342 THORP, E. M., 275, 308 TOCEIILIN,M. S., 457,474 TOLSTIKHIN, N. I., 478,479,480, 501 TOPKAYA, M., 435,451,474 TOWSE, D., 282, 308 TRASK, P. D., 1, 17 TRASK,P. D. and Wu, C. C., 364, 365, 389 TREIBS, A., 344, 365, 389 TUGARINOV, A. I and VINOGRADOV, A. P., 45, 87 T M , W., 409,413 TURNER,F. J. and VERHOOGEN, J., 47, 87 TWENHOFEL, W. H., 34, 53, 58, 75, 76, 87, 308 TZSCHORN, G., 448,452,475 UMBGROVE, J. H. F., 176 USDOWSKI, H. E., 186,218,220,223,245, 309 VALLENTYNE, J. R., 350, 364, 371, 389 VALLENTYNE, J. R. and CRASTON, D. F., 365, 366, 389 VALYASHXO, M. G., 294, 309 VANANDEL, T. and POSTMA, H., 50, 87, 147, 176 VAN DER MOLEN,w. H. and SMITS, H., 396, 413 VANDERSTAPPEN, R. and VERBEEK, T., 148,176 VANW S E , c. R., 20,34, 38,40,87,88,89 VANHOUTEN, F. B., 53, 87, 122, 124 VAN KREVELEN,D. W., 363,374,389,399,403, 413 VAN OLPHEN,H., 166, 176 VAN STRAATEN, L. M. J. U., 64, 87, 302, 307, 309 VANTUYL, F. M., 309, 342 VASSOEVICH, N. B., 7, 381, 389, 478, 480, 501 VEBER,V. V. and TURKELTAUB, N. N., 375,389 VINOGRADOV, A. P.,309 VISHNYAKOV, S. G., 296, 309 VISSE,L. D., 73, 87 VITAL’,D. A., 285, 309 VOIGT,E., 475 VONBURGER,K., 475 VON C O r r A , B., 419, 422, 475 VONENGELHARDT, w., 2, 144, 155, 161, 162, 166, 167, 168, 169, 176, 185, 309, 435, 475, 478, 479, 488, 492, 502, 510, 513, 515, 517, 523

535 VONENGELHARDT, W. and GAIDA, K. H., 99, 124, 140, 165, 169, 170, 478, 488, 502. 515, 517 VON ENGELHARDT, W., M~~LLER, G. and KROMER,H., 158, 177 VONGEHLEN, K., 427,475 VON GRODDECK, A., 419, 421, 422, 423, 475 VONG ~ E LC.,W., 1,17,19,87,88,419,424, 475 VON MORLOT,A., 67,69,87 WAKSMAN, S.A., 33, 87 WAKSMAN, S. A., CAREY, C. L. and REUSZER, H. W., 347,389 WALDSCHMIDT, W. A., 93,124 WAHLSTROM, E. E., 2, 17 WALKER,TH. E., 71, 87, 118, 119, 124, 309 WALLACE, R. C., 309 WALPOLE, B. P., 448,451,475 WALTHER, J., 19, 87, 88, 419, 422, 425, 475 WARDLAW, N. C., 218,262,282, 309 WEAVER,CH.E., 88, 142, 144, 158, 177 WEBER,J. N., BERGENBACK, R. E., WILLIAMS, E. G. and KEITH, M. L., 309 WEEKS,L. G., 344, 385, 389 Wmss, M. P., 61, 64,88, 160, 177 WELLER,J. M., 99, 124, 309,452,475 WELLS,A. J., 67, 88, 309 WELLS,J. W., 200, 309 WELTE,E., 395,413 WENGERD, S. A., 277, 309 WERMUND JR., E. G., 45, 88 WERNER,F., 185, 309,419 WESTOLL,T. S., 448,451,475 WETZEL,O., 59, 88 WETZEL,W., 21, 88 WEYL,P. K., 309 WEYNSCHENK, R., 309 WHITE, A., HANDLER, P., SMITH,E. L. and STETIEN, DEW., 348, 389 WHITE,D. E., 37, 38,88 WHITE,D. E., BRANNOCK,W. W. and MURATA, K. J., 325, 329, 338, 342 WHITE,W.A., 165, 177 WHITEHOUSE, U. 0. and MCCARTER, R. S., 145, 177 WICKMAN, F. E., 23,88, 389 WICKMAN,F. E., BLDC, R. and VONU a r s c ~H., , 309 WILK,H. B., 345,390 WILLIAMS, D., 35, 88 WILLIAMS, E., 407,413 WILLIAMS,H., TURNER,F. J. and GILBERT, C. M., 3, 17, 309, 313, 456, 475 WILLIAMS,P. M., 361, 390 WILLMAN, H. B., 64,88

536 W~ER, H. G. F., 172, 177 W-N, A. F.,452,475 WOOD,A., 271,279, 310 WRAY, J. L.,310 WRENSC. L. and DYBR, W. J., 367, 390 WRENSHALL, C. L and MCKIBBIN,R. R., 367, 390 W-, M.R.J., 492,493, 502 WOLF,K. H., 181,182,185,190,194,200,211, 215, 216, 219, 220, 221, 224, 225, 227, 228, 229, 233, 237, 241, 244, 246, 247, 248, 249, 250,251, 256, 257, 258, 267, 269, 271, 273, 275, 278. 279, 280, 283, 286, 298, 309, 310, 311,313,314,315,316,317,318,320,321 WOLF, K. H. and CoNou~, J., 310 WOLF, K.H.,CHLINC~AR, G. V. and BBALES, F.W.,310

REFERENCES INDEX

Worn, K. H.. EASTON, J. A. and WARNE,S., 310

YOUNO,R. B., 310

V. and WILK, L., 396.413 ZAPOROZRTSEVA, A. S., 46 ZAYISEVA (ZAITSEVA), E. D., 4, 10, 11, 12, 17, 310, 448, 451, 475 ZWBR, E. J. and WRAY.J., 188,263,310 ZEN, E-AN.,52, 68, 88,92,97, 124 zIMMEI1MA”, R. A, and AMSTUTZ,G. C., 437, 439,446,452,475 ZQBELL,C.E.,26, 34, 88, 191, 192, 310, 390, 448,451,475 ZijLLIo, H., 161. 162, 163, 177 ZAEZR,

SUBJECT INDEX

Abrasion pH, 42 Abrolhos Submergence, 61 Absaroka Range (Wyo.), 106 A D P (diphosphate), 72 Africa, 50, 51, 71, 72, 74, 75, 132, 133, 148, 149, 156,210,420,435 Agha Jari man), 289 qiibik Quartzite, 107 Alabama (U.S.A.), 158,327 Alabandine, 4, 13,27 Alberta (Canada), 113,260,270,271,277 Albian (Cretaceous), 159,160 Albite, 64, 92, 112, 120, 121, 122, 160, 172, 450,507 Albitization, 47 Albuquerque (N.M.), 326 Alcyonaria, 36, 58 Algae/algal, 26,29, 55, 67,68,71,73, 77, 134, 185, 186, 188, 190, 193-194, 195, 196, 197, 198, 199, 200, 201, 202, 203,204, 205, 206, 207, 208,209, 210, 211, 212, 213, 214, 215, 216, 219, 220, 225, 226, 229, 230,231, 232, 233,236,237, 241, 242, 243,244, 249,250, 251, 256,257, 258, 260, 261, 262, 266,267, 271, 272, 274, 275, 276, 277, 278, 279, 280, 291, 292, 296, 310, 312, 313, 315, 316, 317, 318, 319, 321, 325, 327, 336, 380, 432, 447 Alginite, 414,415 Allochthonous, 31 1 Allogenic, 311 Allophane, 13, 14,15, 101 Alpenrhein, 140 Alpine, 520 Alps, 69,141,442 AlzOs, 5, 64,70, 142 Al(OH)s,25,238 1.7AlsOs.0.6Mg0.8SiOa.2H~O (montmorillonite) 159 AlzOs.2SiOz.2HaO (kaolinite), 51 4AlPO4.2Al(OH)s.9HsO (wavellite), 74 Amazon, 51, 53 Amber, 364 Amino acids, 135,344,346, 348,364,368,369, 370, 371, 374, 375, 381, 382, 497, 498, 499 Anadiagenesislanadiagenetic, 19, 26, 32, 33,

3740, 41, 43, 44, 45,47,48,49, 55, 56, 58, 63, 64, 66, 68-69,70, 71,73, 74, 76, 77, 79, 88,524 Analcime, 77, 127, 148, 149, 172 Analcimolite, 148 Analcite, 47 Anamorphic/anamorphism,20,78,79,88,121 Anatase, 64,76 Andrews County (Texas), 288 Andros Island (Bahamas), 61,275 Anhydrite, 7, 13,21,27,42, 55, 56-57, 70, 76, 145, 215, 315, 324, 449, 452, 488, 508, 509 Ankente, 4,13, 14,68,76 Ankeritization, 410 Anorthite, 133 Antarctic, 132,133 Anthracite, 8,354,372,391,392,400404,405, 410 Apanepigenetic, 422 Apatite, 410 Aphanic, 259,260,261,217,311 Apo-epigenesis, 181,249, 311, 314,316 Appalachians, 53,479 Aquatolysis, 129, 130,140-141 Arafura Shelf (Australia), 45 Aragonite, 20,34,74,75,76,78,105,133, 183, 186, 187, 188, 189, 193, 216, 217, 218, 219, 220, 221, 222,223, 226, 228, 244,245, 246, 251, 260,263, 267, 271,273, 274, 275, 276, 279, 281, 283, 284, 295, 310, 311, 312, 317, 320, 332 Aransas Bay (off Texas), 144 Arctic Basin, 52 Ardennes, 69 Arizona (U.S.A.), 264,270,326,327 Arkansas (U.S.A.), 331, 339, 420, 421, 433, 437,439,479 Arsenopyrite, 438,439 Ascending water, 88 Asphalts, 346,350,378-379, 381, 382 Atlantic, 49, 50, 51, 131, 132, 133, 134, 135, 148, 150, 151,490,491 ATP (triphosphate), 72 Attapulgite, 146 Aua Reef (Samoa), 21 1

SUBJECT INDEX

Australia, 30, 45, 67, 71, 112, 145, 172, 182,

186, 190, 198, 199, 201, 202, 203, 205, 206, 207, 208, 209, 211, 213, 215, 224, 225, 226, 229, 230, 232, 233, 235, 236, 241, 242, 243, 244, 246, 249, 250, 251, 280, 321, 334, 442 Austria, 438,439 Authigenesis/authigenic,2,4,7, 19,21,22, 35, 41-78, 88,91,92,95,99, 100, 109, 110, 111, 112, 113, 114, 116, 119, 120, 121, 125, 142, 144, 146, 147, 148, 149, 150, 153, 158, 160, 184, 186, 244, 253, 269, 270, 271, 281, 287, 311, 325, 347, 363, 435, 441, 451, 465, 466 Autochthonous, 31 1 Azerbayjan (U.S.S.R.),481 Azores, 135 Azurite, 76 BaCO3 witherite, 77 Bacteria/bacterial, 5, 6, 9, 12, 23, 26, 27, 28, 33, 34, 35, 36, 37, 41, 46, 51, 52, 53, 54, 55,

57, 66,67, 150, 151, 152, 183, 184, 188, 189, 191-193, 210, 240, 242, 245, 246, 261, 265, 267, 275, 276, 284, 295, 347, 359, 361, 381, 394, 395,447,448,451,459 Bahamas, 61, 193,260,270,273,274,275,283, 311 Bahamite, 193, 257, 270, 271, 273, 280, 281, 311, 315, 318, 319 (Ba,H~O)Mndho(psilomelane), 77 Balakhany-Sabunchi-Romany (U.S.S.R.),487 Balkan, 433 Baltic Sea, 365 Bank, 311 Barite, 76, 420, 421, 433, 437, 439, 443, 452, 507,509, 516,517 Bas04 (barite), 76, 437,449, 517 Bathygenesis, 53 Bauxite, 13, 14, 15 Beach-rock, 57, 61, 179, 188, 189, 190, 191, 195,216,230,251,311 Bear’s Claw (Utah), 255 Beidellite, 4, 13 Belgium, 27,406 Bellefonte (Pa.), 335 Bentonite/bentonitic,50,64,115,156, 158 Bering Sea, 10, 11, 12, 74 Bibieybat (U.S.S.R.),487 Bighorn (Wyo.), 117 Bikini, 68,216 Bimini, 275 Biolithite, 311 Biorhexistasy, 53 Biotite, 64, 98, 100, 101, 102, 108-109, 110, 111, 119, 120, 122, 141, 147, 154, 157, 410, 456, 508 Birdseye, 185, 186, 199,235,245,283,311, 321

Bird Spring and Spring Mountains Formations, 222 Bituminous coal, 354, 372, 392, 397, 398, 399,

400404,405,407,408,410

- shale, 448

Black Sea, 33, 52, 130, 152, 163,448,490,491 Blairmore Formation, 113 Bleiberg (Germany), 422 Blind River, 420, 435 Blount Co. Venn.), 326 Bonneterre Formation (Upper Cambrian), 431 ,

440

Bornite, 35, 76 Boundstone, 312 Brachiopods, 75, 215,241,243,244,252,257, 267,281,291,335, 336 Brazil, 98, 133 Brazilian Shield, 71 Breccia, 312, 315, 318 British Guiana, 133 Broken Hill, 420 Brookite, 64, 76 Brown coal, 8, 391, 392, 394, 396-400, 401,

404,407,408,410

Bruchsal (Germany), 508 Brucite, 4, 67, 144, 145, 157, 158 Brushy Basin Member, Morrison Formation (Jurassic), 156 Bryalgal, 312 Bryozoa, 215, 237, 241, 242, 253, 257, 310,

312,335,336

Buchanan (va.), 112 Burke Co. (N.C.), 326 1zC/13Cratio, 56,297, 379, 381 SI3C, 224,297, 380 14C, 217,297, 359,360,497 CaAlzSi401z.4H~O (laumontite), 47, 172 (Ca, Ba, K,Na)eAl~(AI,Si)~Sii004015-20H~O (phillipsite), 76 CaCl2 (calcium chloride), 67, 165 CaC03, 5, 7, 13, 14, 20, 28, 34, 36, 44, 66, 67,

69, 70, 75, 76, 185, 186, 187, 188, 189, 190, 191, 192, 193, 194, 210, 211, 215, 216, 217, 218, 219, 222, 223, 226, 227, 234, 237, 238, 239, 240, 241, 244, 246, 250, 262, 263, 264, 273, 277, 278, 279, 284, 285, 286, 288, 290, 291, 293, 294, 296, 297, 311, 313, 316, 318, 321, 322,333,462 CaFz (fluorite), 430 Calcilutite, 312 Calcite, 13, 14, 15,20,22,24,27,30,34,44,52, 55, 56, 66,67,68,69,70,71,72, 76,78,93, 94,98,100,101,102,104,105,106,108,117,

118, 119, 120, 133, 143, 144, 154, 160, 172, 183, 186, 187, 188, 189, 190, 195, 202, 206,

SUBJECT INDEX

215, 216, 217, 218, 219, 220, 221, 222, 223, 224, 225, 226, 227, 228, 229, 233, 234, 236, 237, 241, 244, 245, 246, 247, 248, 249, 250, 251, 252, 253, 261, 263, 264, 265, 266, 267, 268, 269, 272, 273, 274, 277, 278, 279, 280, 281, 282, 283, 284, 285, 288, 290, 293, 295, 296, 297, 311, 312, 315, 316, 318, 319, 320, 321, 322, 332, 333, 335, 338, 394, 428, 429, 430, 458, 461, 462, 463, 465, 504, 507, 508 Calcitization, 69 Calclithite, 3 12 Calcrete, 24 Calhoun Co. (Ala.), 327 Caliche, 24, 189,258, 312, 322 California (U.S.A.), 9, 12, 47, 130, 136, 145, 146, 151, 152, 163, 170, 327, 350, 351, 358, 479 Cakisphaera, 236 Cambrian,73,107,108,109,111,112,113,217, 335, 361, 373, 431, 435, 440, 496, 497, 498 Cameroun, 156 CaMg(CO3)z (dolomite), 67, 69, 76, 288, 290, 294 Ca(Mg,Fe)(COs)a (ankerite), 76 (Ca,Na)(Al,Si)AlSiaOs (plagioclase), 77 (Ca,Naa)AlaSisOla * 5Hz0 (heulandite), 47 Canada, 113,260,261,262,270,271,273,274, 277,282, 321,431 Canaries, 135 Cannel coal, 415 Capitan Reef (Permian), 241,278,330,335,336 Caa(PO4)r . CaCOs 3Hz0 (dahllite), 74 CasPzOs HzO (collophane), 74, 76 Cap rock, 55, 56 Cape Verde, 133, 135, 149 Carbohydrate, 72,103,192,344,346,348,349, 358-361, 363, 368, 395, 396 Carbonification, 391 Carboniferous, 37, 71, 112, 157,261,262,278, 280, 290, 292, 293, 321, 334, 398, 405, 406, 407, 408, 409, 410, 498, 508 Carbonatization, 44 Cariaco Trench, 350 Carlsbad Cavern (New Mexico), 327 CaSO4, 21, 42, 56, 67, 69, 76, 238, 239, 240, 293, 508 CaS04 * 2Ha0,21, 34,42,56,69, 76,294, 318 Casper (Wyo.), 104 Caspian Sea, 6, 9, 67, 74, 486 Catagenesis, 3 Catskill delta, 53 Celestite, 76 Celle (Germany), 518 Cellulose, 103, 359, 360, 369, 378, 382, 393, 394, 395, 396, 397,400,414 &rnentation/cement, 3, 4, 6, 7, 13, 14, 15, 20,

-

539 21, 22, 37, 38, 40, 41, 46, 59, 61, 66, 71, 72, 89,91,92,93-96,98,101,102,104,106,118, 119, 124, 130, 150, 153, 160-161, 165, 179, 180, 181, 182, 184, 185, 186, 187, 188, 189, 190, 191, 192, 193-194, 208, 209, 215, 217, 218, 219, 220, 223, 225, 227, 229, 230, 232, 233, 234, 237. 242, 243, 244, 245, 246, 247, 248, 249, 250, 251, 256, 260, 263, 264, 265, 267, 268, 269, 270, 271, 272, 273, 215, 276, 277, 279, 281, 282, 283, 284, 285, 286, 288, 298, 311, 312, 313, 314, 315, 317, 322, 333, 337-338, 339, 422, 424, 429, 430, 432, 440, 457,464,465, 511, 518,524 Cenozoic, 4, 62, 94, 102, 104, 106, 149, 156, 160, 188, 189, 190, 248, 251, 257, 265, 279, 283, 337, 349, 350, 354, 361,408,492 Central Africa, 148 -America, 133, 135 - Asia, 481 - Congo Basin, 148 - France, 442 - Sahara, 148 Grtium Limestone (Danian, Denmark), 60 Cerussite, 76 CHI (methane), 54,60,404,409,410 Chabazite, 4, 77 Chalcedony, 40, 58, 72, 76, 94, 98, 102-104, 115, 149, 150, 171, 241, 243, 244, 247, 333, 334, 335, 337, 338, 339 Chalcocite, 13, 14 Chalcopyrite, 13, 35, 76, 153, 430, 431, 435, 438,439,441 Chamberlain Creek Syncline (Ark.), 437 Chamosite, 13, 14, 27, 46, 76, 77, 240 Charged-net clay membranes, 155, 377, 491, 492495,496,499 Charles Formation (Mississippian), 274 Chaves Co. (N.M.), 327 Chert, 13, 53, 56, 58,59,61, 71, 75, 78,94,95, 101, 103, 104, 113, 115, 116, 117, 118, 119, 146, 240,241, 286, 324, 325, 330, 331, 332, 333, 334, 335, 337, 338, 339, 340, 354 Chertification, 115 Chesapeake Bay (U.S.A.), 145 Chile, 74 Chitin, 36,210, 361 Chlorite, 14, 16, 46, 49, 52, 76, 99, 100, 101, 107, 108, 109, 110, 111, 113, 119, 120, 122, 127, 128, 131, 132, 133, 142, 144, 145, 151, 153, 154, 157, 158, 159, 160, 171, 172, 410, 450, 457,461,462, 463, 464, 465, 480, 507, 508, 516, 519 Chloritization, 410 Chlorophyll, 365, 366 CsHiaOe, 54 Chondrites, 345, 354, 356, 357, 364, 370, 373

540 Christobalite, 337 Chromatography, natural, 19, 39, 48, 65-66, 79,353,359 Chromite, 420 Cinnabar, 21 Clarain, 413,414 Clarite, 398, 414, 415 Clarodurite, 414 Clast, 313 Clay filtering, 39

Cleveland (England), 77 Clinoptilolite, 146,148,149, 172 Clinozoisite, 21 C/N ratio, 12, 36 COa, 7,9,23,24,25,28, 34, 36,40,43,44,46,

54, 56, 60, 67, 70, 72, 75, 78, 118, 119, 151, 172, 183, 187, 188, 189, 190, 191, 192, 193, 194, 210, 238, 240, 261, 279, 284, 285, 291, 293, 294, 295, 320, 321, 330, 359, 360, 380, 403,408,409,449,460,463 Coal, 1, 8, 30, 54, 71, 75, 102, 157, 338, 339, 344, 345, 347, 354, 359, 363, 364, 365, 369, 372, 314, 379. 380, 382, 391415, 424, 443, 446,447,452 Coamcation, 380, 391415, 524 CoWte, 413,414 Collophane (collophanik), 74,76 Colorado (U.S.A.), 98, 100, 110, 156, 270, 420,497 Columbia (South America), 133 Columbus (Ohio), 326 Colusa (Calif.), 327 Compaction, 7-8, 9, 20,21, 22, 31, 32, 37, 40, 56,60,65,66,69,70,71,89,96,97,99,115, 127, 130, 151, 154, 161, 162, 165, 166, 168, 170, 184-186, 215, 217, 222, 242, 262, 265, 266, 269, 274, 278,286,288, 320, 322, 334, 347, 353, 359, 377, 393, 397, 419, 421, 422, 424. 425, 430, 431, 432, 435, 445, 446, 459, 464,477, 478,488, 490, 491, 495, 496, 509, 510,511,512,513,514, 515,519 Concretions, 4, 5, 6, 7, 20, 22, 36, 37, 56, 59, 60, 62, 73, 75, 77, 102, 115, 128, 148, 153, 154-155, 160, 171, 245, 284-286, 295, 313, 324,354,422,462,463,464,466 ConaincOne, 22,186,223,245,313 Congelation, 20 Congo, 51, 148,435 Connate water, 20, 22, 26, 30, 32, 33, 34, 37, 39, 40,47, 48. 49. 65, 66, 68, 70, 71, 79, 88, 185, 262, 295, 347,431, 478, 487, 488, 491, 492,496,499 Coprolites, 73,147 Coquimbite, 56 Coquina, 313 Coquinite, 313

SUBJECT INDEX

Corals/codreef, 34, 202, 216, 217, 219, 223,

235, 241, 242, 258, 266, 267, 279, 291, 310, 335, 336 Cordilleran Geosyncline (U.S.A.), 185 Corocoro, 435 Corrasion, 179, 184,200,211, 321 Corrensite, 146, 159 Corrosion, 61, 179, 184, 186, 189, 193, 194, 200, 207, 210, 211, 212, 213, 215, 225, 229, 245,248,250,282,321,465 Cretaceous,46,59,60,62,64,73,77,104,113, 148, 154, 156, 159, 160, 170, 272, 354, 355, 356,437,479,485,497, 517 Crinoids, 204, 209, 211, 212, 213, 215, 218, 231, 232, 233, 237, 253, 263, 267, 268, 269, 276, 281, 283, 288, 291, 292, 313, 316, 319, 320,322,336 Criquina, 313 Criquinite, 253,254,255,281, 313 Crowsnest (Alberta), 113 Crude oil, 345, 347, 365, 366, 367, 374, 375, 376,377,378, 381,382 Cryptocrystalline, 313 Cuba, 133 CuaCOs(0H)a (malachite), 76 Cua(C0a)dOH)a (iyrite), 76 CuFeSa (chalcopynte), 76 Cu~FeSlr@ O d k ) , 76 Cutinite, 414, 415

Dahllite, 74,76 Danian, 60 Davis Shale, 444 Dead Horse Wash, White Pine County (Nev.), 252

Dead oil, 252,265 Dead Sea, 59 Decanturville (Mo.), 443,444 Decarbonatization, 44 Decementation, 22, 66,96, 179, 184,200 Dedolomitization, 66,69,314,320 Deep-burial (stage of diagenesis), 127, 128, 129,130,131,153,154,155-156,158,160,524

Defeldspathization, 97 Dehydration, 4, 6, 7,20,21,22, 39,40,42,43, 55,5644.66,70,89,166,353,371.377,407 Dehydrolysis, 4243.56 DelitMcation, 41

Denmark, 60,62

Density-gravity features, 427, 432, 433, 435442,447

Depocenter, 313 Depolymerization, 21,22 Desalinitication, 20 Desideritization, 78 Desilicification, 71.72

SUBJBCT INDEX

Desmoinesian (Utah), 254 DesMovibrio desulfriiccms,55 Deuteric, 450,451 Deuterium, 495,496 Devitrification, 47, 115, 135, 148, 160 Devonian, 53, 60, 68, 69, 112, 181, 190, 194, 197, 198, 199,200,202, 211, 212, 215, 225, 226, 268, 270, 271, 277, 280, 283. 288, 479, 485,497,498,508,520 DH/Ha ratio, 495 Diaclastic, 39,60,65 Diagenesis, 313 Diagenetic cycle, 39,44 - differentiation, 22 -evolution, 1, 3141 - fabric, 38 -rhythm, 37 - stages, 4-7 - -, 1st (depositional), 430 ,2nd (early burial), 430 - -, 3rd (pre-metamorphic), 430 Diastrophism/diastrophic, 31, 37, 38, 39, 40, 71,79 Diatoms, 58,134, 151,285 Disthene, 506 Dogger, 159,160, 506,512,517,518 Dololutite, 314 Dolomite/dolomitic, 4,5,13, 14,15,27,30,44, 51, 52, 66, 67, 68, 69, 70, 76, 78, 104. 105, 106, 133, 179,209,222,225,226,238, 239, 246, 247, 250, 255, 258,259, 260, 262, 263, 264,265, 268,270, 273, 274, 275, 279, 280, 282, 283, 284, 287-297, 310, 312, 314, 316, 317, 319, 320, 327, 330, 332, 333, 334, 338, 394, 437,443,444,504,507, 508 Dolomitization, 21,44, 61,6669,75, 78, 105, 151, 179, 181, 182, 186,217, 223, 240, 255, 258,260, 262,263, 264, 267,268, 269,270, 271,272, 273, 274, 276, 277, 279,280, 281, 282, 283, 287-294, 295, 296, 297, 314, 316, 319, 320, 322, 324, 332, 333, 334, 337, 338, 487,499 Dominican Republic, 133 Donetz Basin, 8 Drewite, 281 Ducktown, 420 Dunning (Neb.), 326 Durain, 413,414 D d t e , 398,414,415 Duroclarite, 414 Dutch coast, 161 Dvinelka, 291

--

Early burial (stage of diagenesis), 19, 32, 33, 34-37, 43, 57, 78, 97, 114, 338, 339 Eastern Great Bash (USA.), 265

541 Echinoids, 217,266 Egyptian Deaert, 496 Eh, 5, 7, 9, 16, 23, 25-28, 29, 31, 32, 34, 39, 45,46, 50, 51, 52, 54,55, 75, 78, 88, 89,91, 101, 147, 15€,152, 155, 183, 187, 192, 193, 238, 240, 245, 264, 366, 373, 425, 431,448, 449,450,459,460,461,499,524 Eifel (Germany), 69,422 Elberfeld (Germany), 408 Eldingen (Germany), 518 Elizabeth City (N.C.), 326 Ely Group (Nev.), 185 - Limestone (Nev.), 255 Emba (Munayli, U.S.S.R.), 479 Endogenous/endogem&c, 181, 186, 189, 239 314,418,419,437 England, 64,70,77,270,278,279 English White Chalk, 285 Eniwetok Atoll, 68,279 Enterolithic, 57 Eocene, 106,279,361 Eolianite, 250,251, 311 Epidiaeenesislepidiagenetic, 19,26, 32, 33, 37, 39, 40-41, 44, 51, 53, 56, 57, 58, 61,66,69, 71,72, 74, 76, 77,78,79, 88 Epidote, 21,450 Epigenesis/epigenetic, 3, 8, 19. 40,41, 42, 69, 75, 78, 88, 89, 92, 99, 179, 180, 181, 182, 200, 215, 217, 224, 235, 244, 246, 248, 249, 257, 290, 296, 298, 314-315, 316, 322, 339, 343, 391, 418, 419, 420,423,426,427, 432, 437, 446, 447, 448, 450, 451,453, 455,459. 466, 523,524 Erionite, 149 Erwin Quartzite (Va.), 112 Euxinic, 33, 36,52, 53, 78,240,245,264 Evaporation, 51,488,496 Evaporite, 2, 30,31, 55, 56, 68, 70, 156, 160, 265,280,283,297,324,330,338 Exinite, 393, 394, 400,401, 403,414, 415 Exodiagenesis, 4 Exogenous/exogenetic, 128,181,186,189,239, 315,418419,437,446 Exotic minerals, 21 Extraclast, 257.315 Faecal pellets, 46,185,222,270,271,272,275, 281,317,319 Falun, 420 Fat coal, 401 Fatty acids, 343, 344, 361, 362, 363, 364, 368, 374,375, 377, 381,382,497,498 FayetteviUe Shale (Mississippian), 339 FeCOa, 34,75,77,284,460 Feldspar, 42,43,45,47, 50,6344,76,77.91, 92. 96, 97, 99, 106, 109, 111-112, 113, 117,

SUBJECT INDEX

120, 121, 122, 125, 131, 133, 135, 141, 147, 150, 151, 155, 160, 170, 171, 172, 331, 410, 450, 504, 505, 506, 507, 508, 509, 516, 519

Feldspathization, 64 FeaOs, 5, 20, 28, 66, 76, 460 Fe304,28, 460 Fe(OH)a, 54, 238 Fe(OH)8, 238 FeO(OH), 54 FesPaOs.8HaO (vivianite), 74 Ferguson Mountain Formation (Wolfcampian), 255 Ferrihalloysite, 13 FeS, 34, 44, 45, 54,460 FeS n H ~ 0 (hydrotroilite), 152 FeSz, 13, 14, 15, 26, 28, 30, 34, 40, 44, 45, 54,

-

56, 76, 77, 245, 250, 430. 443, 444,448, 460 FeSi03, 460 FeSO4, 56 FeSO4 * 7Ha0 (melanterite), 56 Fe~(S04)s* 9Ha0 (coquimbite), 56 Fiji, 189 Finland, 55 Fish Haven Dolomite (Idaho), 336 Flint, 13, 37, 59, 60, 61, 62, 63, 285, 325

“Flint curtains”, 60 Florida (U.S.A.), 193,283 Florida Bay, 217, 348 Flour, 315 Fluorite (fluorspar), 420, 428, 430, 432, 433, 447,452

Fluorspar District (lll.), 439 Fontainebleau sandstone crystallization, 72 Foraminifera/foraminiferal, 34, 46, 147, 203, 216,217,241,255,282,291,315,316,336,456

Fort Union Formation (Wyo.), 104 Fossil charcoal, 394 France, 27,59,210,442,461,462,463,464,466 Fredericktown (Mo.), 431,437,440 Fredonia Limestone (Ill.), 428 French Guiana, 133 - Jura, 69 Fulton (Miss.), 326 Fulvic acids, 359 Funafuti Atoll, 68,296 Fusain, 414, 415 Fusinite, 394, 401,403, 413, 414,415 Fusite, 414 Gach Saran (Iran), 289 Galena, 28,35, 76, 153,422,430,431,432,433,

435,441,442,447 Ganister, 116, 117 Garnet, 504, 505, 506, 507, 509 Gas, 7, 26, 41, 185, 378, 408-410, 477, 478, 503, 506, 514

Gelification, 397 Geodes, 325,334-335 Geopetal, 437 Germany, 31, 69, 71, 140, 154, 157, 158, 159, 160, 410, 506, 518,

166, 167, 169, 405, 406, 407, 408, 409, 420, 422, 435, 451, 461, 503, 504, 505, 507, 508, 510, 511, 512, 515, 516, 517, 520 Gerster Formation (Permian, Utah), 253 Gibbsite, 131, 132, 134, 171 Gifhorn Trough (Germany), 511, 512, 518

Glacier National Park (Mont.), 64 Glarus Freiberg (Germany), 451 Glass, 113, 115, 134, 135, 149, 160 Glauconite, 5, 14, 15, 16, 27, 45, 73, 76, 101, 107, 108, 109, 111, 112, 119, 128, 146, 147, 171, 240, 245, 287, 293, 444, 455, 456 Glauconitization,45 Goethite, 46, 135 Grain growth, 216-224,234-237, 315 Grainstone, 315 Graphite, 343, 382, 391, 392, 399,401 Graphitization, 410 Graywacke, 72,92,112,115,120,121,122,431 Great Bahama Bank, 260 - Barrier Reef, 68, 189 - Basin (Nev., Utah), 190, 215, 265, 312 -Britain, 64,70, 77, 265,266,270, 214,278, 279, 321,461 Salt Lake (Utah), 146 Greenalite, 76 Greenlee Co. (Ariz.), 326 Grumous, 3 15 Guadalupe Delta, 144 - Mountains, 54 - River, 142, 144 Guano, 74

-

Guilmette Formation (Devonian, Utah, Nevada), 215 Guinea, 133, 147 Gulf Coast, 156,510 - of California, 145 -- Mexico, 50,142,144,145,146,152,153, 156, 329, 510

-- Naples (Italy), 135, 148, 150 -- Paria (Venezuela), 147

Gypsum, 7, 13, 21, 27, 34, 42, 51, 55, 56-57,

67, 69, 72, 73, 76, 101, 263, 294, 315, 449, 452,488

Haft Gel (Iran), 289 Haiti, 133 Halimeda, 216,217 Halite, 27, 30, 55, 76 Halloysite, 13, 14, 15, 101, 132, 142, 143 Hall Canyon Member (Morrowan, Utah), 253

543

SUBJECT INDEX

Halmyrolysis, 3,41,44-52, 78, 88-89, 97, 129, 130, 141-151, 159, 183, 316 Hancock (Mich.), 109, 110 Harney Co. (Ore.), 326 Hawaii (U.S.A.), 134, 149 Heat flow, 30, 31 Hematite, 7, 27, 55, 56, 76, 93, 99, 108, 110, 111, 114, 204, 205, 206, 207, 208, 241, 242, 243,247,249,250,287,435,459,460,461,463 Heron Island, 189 Heterocyclic compounds, 344, 346, 348, 365367, 371 Heulandite, 47, 77, 148, 172 HgS (cinnabar), 21 Hogan Formation (Desmoinesian, Utah), 254 Hoheneggelsen (Germany), 154 HsP04,29 HzS (hydrogen sulfide), 7, 16, 28, 29, 34, 35, 52, 53, 54, 55, 152, 191, 192, 194, 211, 240, 245,261,265,449 H~Si03,240 H4Si04, 70, 71, 325 HaS04, 56,211,449 Hudson (Wisc.), 109, 111 H d c acids, 34, 43, 189, 349, 354, 355, 359, 368, 369, 370, 371, 372, 373, 374, 381, 382, 395,400,497,499 - coal, 393,394 -matter, 343,394,395,396,397,400,414,415 Humification, 395 Humins, 359 Hydration, 21, 22, 28, 40, 42, 43, 45, 56-64, 66,96, 157, 166,294 Hydrocarbon, 21, 22, 33, 34, 55, 65, 343, 346, 355, 361, 364, 365, 366, 367, 374-378, 379, 381, 382,409,431,448,499, 503, 524 Hydrogoethite, 7, 15 Hydrohematite, 7, 15 Hydrolysis, 24,28, 31,42-43,44, 56,64,96-97, 151, 158, 294, 343, 344, 350, 353, 356, 357, 359, 360,361, 369, 371, 381 Hydromagnesite, 67, 76 Hydromica, 64,157 Hydrothermal, 3, 20, 35, 37, 69, 71, 78, 149, 296, 408, 423, 426, 427, 432, 437, 448, 450, 517 Hydrotroilite, 152, 265 Hypogene/hypogenic, 19, 37, 88, 89, 181, 316, 321,426,432 Idaho (U.S.A.), 326, 336 Igneous, 30, 50, 51, 140, 141,347,420,423 Illinois (U.S.A.), 49,95,428,430,433,439,447 Illite, 4,43,45,4849, 50,51,52,64,72,76,95, 100, 101, 105, 119, 127, 128, 131, 132, 133, 136, 137, 139, 141, 142, 143, 145, 147, 153,

154, 155-156, 157, 158, 164, 165, 169, 170, 171,410, 516 Ilmenite, 64 Impingement, 316 Inagua (Bahamas), 283 India, 64, 71,272 Indiana (U.S.A.), 261, 263, 283, 335, 336 Indian Ocean, 131,132,133,134,148,149,150 Inertinite, 394, 401, 414, 415 Internal sedimentation, 179, 184, 200, 205, 226226,237, 246, 248, 249, 250, 287, 316 Interstitial fluid/solution/water, 7, 8, 9, 10, 11, 20, 27, 33, 38, 51, 56, 65, 115, 151-152, 155, 158, 159, 183, 184, 185, 192, 223, 244, 245, 250, 284, 285, 295, 316, 329, 330, 340, 345, 477,488,490,491,492, 503-520 Intraclast, 316 lntrastratal solution, 22 Inversion, 20,68, 179, 184, 186, 188,216-224, 273,286, 316 Ion-filtration, 155, 377, 491, 492495, 496, 499, 500, 517 Iran, 288,289 Irelandwsh, 262,280, 321 Iron Mountain (Mo.), 107,108 Ishpeming (Mich.), 326 Isoprenoids, 346, 361-365 Itacolumite, 96,98 Italy, 47, 57, 130, 135, 148, 150, 161, 166, 167, 510 Jefferson City (Tern.), 327

-- Limestone, 445

Jerusalem, 59, 60 John Day Formation (Tertiary, Ore.), 160 Jordan-Dead Sea Rift system, 60 Judea, 59 Judean Desert, 60 Jurassic, 59, 72, 77, 78, 154, 156, 159, 160, 166, 167, 168, 169, 354, 461, 485, 503, 506, 510, 511, 512, 517, 518 Juvenile water, 65,478,487 Juxta-epigenesis, 181,248, 314, 316 Kaibab Limestone (Permian, Nev.), 241 KO.58(All.38 Fe3+0.37Fe*+o.o4 Ml30.34) (OH)~(Si3.41Alo.s0)010 (illite), 147 &%lSi308 (orthoclase), 76, 508 K A ~ ~ S ~ S O I ~ ((muscovite), OH)Z 76 Kansas (U.S.A.), 51, 94, 102, 337,485 Kaolinite, 13,14,15,42,49,50,51-52,76,100, 101, 105, 128, 131, 132, 133, 136, 137, 139, 142, 143, 145, 153, 157, 158, 159, 160, 164, 165, 169, 171, 172, 274, 410, 461, 504, 505, 506, 507 Kaolinization, 508, 516

544 Karachukhur (U.S.S.R.), 486 K/Ar isotope age, 46,49, 160 Kanvin (Upper Silesia), 410 Katagenesis, 20, 89 Katamorphism/katamorpbic, 20, 40, 89, 122 Katanga (Congo), 435 (KzCa)AlzSi401~* 4Hz0 (phillipsite), 47

SUBJECT INDEX

Liesegang rings, 56 Lignin, 35, 348, 359, 360, 368, 369, 371, 373, 378, 382, 395, 396, 397,400,414 Lignite, 354, 359, 361, 363, 372, 396 Limonite, 27, 55, 56, 76, 102, 108, 240, 274, 457,458,459,461,462,463,464,465 Lipids, 72, 344, 346, 348, 349, 361-365, 369, (%Cao.5,Na)o.s4(Alo.4~Fe3+~.g~Fe~+~.~gMg~.40) 377, 380 (OH)z(Sis.a~.Alo.a5)010(ghuconite), 147 Liptinite, 414,415 KCl, 139,239 Lisbon area (Paradox basin), 270 K2CO3, 72, 73 Lithification, 4, 6, 7,20, 21, 32, 37, 40, 41, 88, Keewenawan, 109, 110 89,91,92,93-97,98, 103,112,114, 116,128, Kentucky (USA.), 158 130, 131, 179, 180, 182, 185, 186-194, 248, Kerogen, 349, 354, 355, 356, 360, 368, 269, 264, 269, 278, 280, 283, 285, 296, 313, 314, 370, 371, 372, 373, 374, 375, 379, 380, 381, 316, 332, 333, 334, 338, 503 382,499 Lithoclastic, 269-272, 317 Kertschenite, 14, 16 Lithographic, 317 Keuper, 158,503,504,505,506, 507,508, 512, Locomorphism (or locomorphic stage of dia516 genesis), 32, 91, 98, 102-106, 112, 113, 115K2O * A1203 * 6SiOz (orthoclase), 64 119, 122, 125 KzO 3AlzOa * 6SiO2 * 2 H d (hydromica), 64 Loess, 47, 48 Krakatau, 134 Lone Tree or Blue Mesa (Colo.), 156 Krakow Leadbelt, 448 Loray Formation (Permian, Nev.), 252 Kuban-Black Sea, 479 Lord Howe Island (Australia), 230,251 Kupferschiefer (Permian, Germany), 420, 435 Lorraine (France), 461,462,463,464,466 Kuroko, 420 Los Angeles Basin (Calif.), 12 Lower Saxony Basin (Germany), 408 Labradorite, 133 Lump, 317 La Brea asphalt pit (Calif.), 350 Laguna Madre (Texas), 276 Maastrichtian, 60, 62 Lake Bonneville, 190 Macerals, 394, 400, 403, 413, 414, 415 -Constance, 136,138,140,141,153, 161,163 Madagascar, 132, 149 - Michigan, 325 Magmatic, 35, 38, 89, 407, 408, 411, 520 - Superior, 420,450,455 Maguesite, 13, 67, 76, 101, 420, 508 - Ziirich, 130, 162, 163 Magnetite, 460, 461, 463, 466 Lali (Iran), 289 Malachite, 76 Lameta Beds (Turonian, India), 272 Manganese nodules, 52,89,128,150-151,171, Lamotte Sandstone(Cambrian, Mo.), 107,108, 354, 447 431,435,440 Manganite, 15 Las Vegas Line (Nev., Ariz., Utah), 222, 264 Mansfeld (Germany), 435 Late burial (or pre-metamorphic stage of Marcasite, 4, 30, 36, 40, 54, 55, 56, 68, 71, 76, diagenesis), 32, 98, 102, 103 153,245,264,265,431,444,447 Laterite, 53, 71, 75, 146 Marine humus, 33 Laterization, 51 Marquette (Mich.), 107 Latosols, 51 Mascareignes, 149 Lau (Fiji), 189 Matrix, 317 Laumontite, 47, 77, 172, 519 Meadow Canyon Member (Utah), 254 Lead Belt (Mo.), 432, 439. 447 Mediterranean, 24, 50, 135, 148, 153 Ledikite, 127, 141, 154, 157, 158, 171 Melanterite, 56 Leduc Reef (Canada), 43 1 Melnikovite, 55 Leptochlorite, 4, 5, 7, 14, 15, 16 Memphis (Tenn.), 326 Leucoxene, 76 Merismitic, 426, 427 Leverrierite, 13 Mesopotamian Basin, 57 Liassic, 154, 159, 160, 166, 167, 168, 169, 506, Mesozoic, 46, 59, 60, 62, 64,70, 72, 73, 77, 78, 510, 511, 512, 518 104, 110, 111, 113, 148, 154, 156, 158, 159, Liesegang concretions, 286 160, 166, 167, 168, 169, 170, 272, 354, 355,

545

SUBJECT INDEX

356, 366, 433, 437, 461, 479, 485, 497, 503,

504, 505, 506, 507, 508, 510, 511, 512, 516,

517, 518 Mesquite Bay (Texas), 144 Metahannosis, 41, 42

Metamorphism/rnetamorphic,2, 3, 4, 8, 20,

21, 22, 30, 31, 32, 33, 38, 39, 47, 49, 57, 79, 88, 92, 97, 99, 112, 119, 121, 127, 128, 129, 130, 140, 141, 158, 170-171, 180, 217, 278, 280, 311, 314, 315, 317, 348, 356, 371, 373, 391, 401, 405, 411, 420, 422, 424, 427, 446, 451,461,464,487, 503, 519-520,523 Metasomatism/rnetasomatic, 21, 22, 31, 32, 44, 66-78, 263,274,295,297, 314,450, 520 Meteoric water, 19, 20, 32, 35, 38, 40, 41, 44, 79,88,188,222,464,477,478,487,495,496

Meteorites, 343, 344, 345, 347, 350, 354, 356, 357,360,361,369,373, 378 Meulerization, 72 MgClz, 67, 145, 238, 239, 240 MgCOs, 7, 44, 67, 68, 69, 76, 284, 293, 294, 296, 318 10.1MgO * 1.7f%os m6.4SiOa 8HzO (chlorite), 159 Mg(OH)a, 42, 44, 67,293,294 3Mg0 * 2SiOa 2Hz0 (serpentine), 42 MgzSiOe (olivine), 42 MgSOs, 67, 69,239,293,294 Miami Oolite (Pleistocene), 283 Mica, 21,32,43,45,48,64,76,91,95,98, 100, 101, 102, 106-111, 113, 114, 119, 120, 121, 122, 125, 127, 140, 141, 142, 147, 154, 157, 172, 331, 410, 456, 459, 507, 508, 516, 519 Michigan (U.S.A.), 107, 109, 110, 326, 450 Micrinite, 394, 401, 403, 413, 414, 415 Micrite/micritic, 179, 180, 185, 186, 195, 197, 198, 201, 202, 203, 204, 206, 207, 208, 209, 212, 213, 214, 215, 216, 219, 220, 222, 224, 227, 228, 229, 231, 232, 233, 235, 237, 242, 246, 251, 252, 256, 257-265, 266, 272, 281, 283, 298, 310, 313, 317, 318, 319, 321, 330 Microstylolitic, 96 Mid-Atlantic Ridge, 132, 146 Middle Kittaning Coal (Pennsylvanian), 117 Milwaukee Co. (Wisc.), 327 Mina Ragra, 420 Minneapolis (Minn.), 94 Minnesota (U.S.A.), 64, 94, 112, 160 Minturn Formation (Pennsylvanian), 98, 100, 110 Miocene, 354,492 M-i-S (Iran), 289 Mission Canyon (Mississippian), 274 Mississippi (U.S.A.), 136, 326, 328, 329, 420, 421,432,433,435, 439, Mississippian, 261, 262, 265, 269, 270, 274,

-

-

282, 321, 331, 335, 339, 370,437,485 Missouri (U.S.A.). 107, 108, 141, 431, 432, 435,437,439,440,443,444,447 Missourian age, 261,267 Mitterberg (Austria), 438, 439 MnCOs, 77,284 MnOa, 52, 53, 77 Mn(OH)z, 238 Moens Klint (Denmark), 60 Mohole, experimental, 151, 165, 295, 353, 358 Montana (U.S.A.), 64,274 MontmoriUonite,4,13,14,47,49,50-51,52,64, 65,76, 101,127,128, 131, 132, 133, 136, 137, 139, 141, 142, 143, 144, 145, 146, 147, 149, 153, 155, 156, 157, 158, 159, 164, 165, 166, 169, 170, 171,489,490, 516 Mordenite, 4 Morocco, 75 Morrison Formation (Jurassic), 156 Morrowan (Utah), 253 Moscow (U.S.S.R.), 408 Motion-disruption features, 427,44347 Mountain leather, 4 Mud aggregate, 317 Mudstone, 317 Mud volcano, 443,444,453 Muscovite, 43, 48, 64, 76, 95, 101, 102, 107, 108, 109, 110, 111, 113, 114, 120, 141, 172, 516 Miinster (Germany), 508 Miinsterland 1, borehole (Germany), 409,410, 520 Mytilus californianus, 350 15N/14N isotope. ratio, 381 NaCI, 67, 76, 139, 164, 165, 187,238,239,240 (Salt) NazCOs * loHz0 (natron), 76 Naft Khaneh (Iran), 289 - Sefid (Iran), 289 NaHCOs, 449 Namurian, 407 Naphthenic acids, 497,498 Na2S04, 187 Natrolite, 77 Natron, 76 Navajo Indian Reserve (Ariz.), 327 Navassa Island (West Indies), 74 Nebrasca (U.S.A.), 326 Neftechala (U.S.S.R.), 486 Nesquehonite, 76 49 Neuse River (N.C.), Nevada (U.S.A.), 185, 190,215,222,241,252, 255,264,265, 312, 336, New Caledonia, 57, 59 Newfoundland, 77, 132

SUBJECT INDEX

New Jersey (U.S.A.), 46, 325 (U.S.A.), 57, 241, 261, 267, 270, 278, 326, 330, 335, 336 - South Wales (Australia), 112, 182, 186, 198, 199, 201, 202, 203, 205, 206, 207,208, 209, 211, 213, 215, 224, 225, 226, 229, 232, 233, 235, 236, 241, 242, 243, 244, 246, 249, 251,280 - Zealand, 172 Niger (Africa), 51 (NH4)2co3, 187 NH40H. 29 (NH4)2Po4, 240 Nonesuch Shale (Keewenawan), 109, 110 Nontronite, 50, 132, 135, 149 Normandy (France), 59 North Atlantic, 49 - Carolina (U.S.A.), 49, 326 - Dakota (U.S.A.), 326 Northern Limestone Alps, 69 North Sea, 34, 138 Novaculite, 13, 324 Nubian Series (Egypt), 496 Nubrigyn (New South Wales), 182, 186, 198, 199, 201, 202, 203, 205, 206, 207, 208, 209, 211, 213, 215, 224, 225, 226, 229, 232, 233, 235, 236, 241, 242, 243, 244, 249, 251, 280 Nucleic acids, 72, 344, 348

- Mexico

lsO/lsO isotope ratio, 61, 297, 381, 495, 496 P O , 297,496,497

Ogallala Formation (Paleocene, Kansas), 94, 102,337 (OH)4Ky(A14,Fe4,Mga) (Sis-uAly)O~o(illite), 48 (OH)S(R~+,R%~~)S+~ . [Si~e-~(N,Fe~+)n] 3 040 nH20 (montrnorillonite), 50 Ohio (U.S.A.), 326, 327 Oil, 41, 55, 394, 506, 514, 515, 518 - accumulation, 153 - Shale Group limestones (Scotland), 274 Oklahoma (U.S.A.), 46, 326,331,339,485,497 Okmulgee Co. (Okla.), 326 Oligoclase, 92, 112, 133 Oligonite, 4, 14 Olivine, 42 Oiilite, 5, 15, 33, 64, 75, 77, 185,219,223,233, 234, 240, 245, 246, 257, 258, 259, 265, 270, 272, 273, 274, 275, 276, 277, 283, 284, 293, 313, 315, 316, 317, 318, 319, 321, 334, 335, 337, 418, 425, 428, 429, 430, 455, 456, 457, 458,459,463,464,465 Opal, 7, 14, 36, 39, 58, 59, 72, 76, 93, 94, 102-104, 115, 134, 147, 148, 149, 161, 171, 329, 337-338

-

Ophtalmite, 426,427 Oquirrh Formation (Utah), 253 - Mountains (Utah), 253 Ordovician, 58, 64, 77, 94, 158, 160, 336, 497, 498 Oregon (U.S.A.), 47, 160, 326 Organic matter, 2, 6, 7, 12, 13, 27, 33, 34, 35, 36,51,53,54,68,99,114,119,131, 135, 136, 147, 151, 152, 183, 187, 188, 190, 191, 192, 194, 210,222,245,261,265, 293, 295, 343382,414,448,449,459,480,497-499 Orinoco, 51, 53 Orthoclase, 42, 64, 76, 109, 111,120,133,160, 504, 507, 508 Orthomicrite, 228,229,316 Orthosparite, 225,228, 320 Oslo (Norway), 139 Ouachita Mountains (Ark., Okla.), 331, 339 Outukumpu, 420 Owyhee Co. (Idaho), 326 Oxidation/oxidizing, 4, 5, 7, 9, 12, 15. 16, 21, 22, 23, 26, 30, 32, 33, 34, 35, 36, 40,41, 43, 44, 45, 52-56, 66, 71, 78, 19. 91, 97, 98, 99, 100, 101, 102, 103, 108, 109, 110, 111, 114, 115, 116, 119, 120, 121, 125, 147, 150, 151, 184, 191, 192, 210, 211, 225, 242, 265, 267, 277, 293, 351, 352, 358, 359, 366, 371, 394, 395, 424, 448, 449, 456, 457, 458, 459, 517 Oxykertschenite, 15, 16 Ozark, 269,339 Pacific, 9, 12, 47, 49, 50, 51', 52, 131, 132, 133, 134, 135, 145, 147, 148, 149, 150, 151, 192, 348, 351, 358, 370,492 Palagonite, 47, 134, 135, 148, 149 Palygorskite, 127, 146, 148, 160, 171 Paleocene, 94, 104, 337 Paleozoic, 31, 37, 46, 49, 50, 51, 52, 53, 58, 60, 61, 64, 67, 68, 69, 70, 71, 73, 75, 77, 94, 95,98,100,102,107,108,109,110,111,112,

113, 117, 145, 147, 156, 158, 160, 181, 185, 190, 194, 197, 198, 199, 200, 202, 211, 212, 215, 217, 222, 224, 225, 226, 241, 252, 253, 261, 262, 264, 265, 266, 268, 269, 270, 271, 274, 277, 278, 280, 282, 283, 288, 290, 291, 292, 293, 321, 330, 331, 334, 335, 336, 339, 353, 354, 361, 367, 370, 373, 376, 398, 405, 406, 407, 408, 410, 431, 435, 437, 440. 443, 451, 458, 479, 485, 496, 497, 498, 508, 520 Palliser Formation (Devonian, Canada), 271 Paradox Basin (Ariz., Colo., N.M., Utah), 270 Paragenesis/paragenetic, 97,100, 102, 112, 125, 128, 158, 179, 180, 181, 182, 205, 226, 235, 243, 246-250, 318, 332, 333, 334, 338, 427435,441,447,448,455,459-466 Paris, 61

SUBJECT INDEX

Paris Basin, 64, 72 Parry Group (Devonian-Carboniferous, New South Wales), 112 PbCOa (cerussite), 76 PbS (galena), 76, 430, 443, 444, 448, 449 Peat, 14, 15, 354, 359, 361, 363, 367, 372, 382, 391, 392, 393, 394-396, 397,400 Peissenburger Molasse, 446 Pelagosite, 190, 193, 318 Pellet, 46,92, 109, 110, 185, 193,202,208,212, 222, 225, 226, 233, 234, 237, 241, 242, 243, 249, 254, 256, 257, 258, 259, 270, 273, 274, 275,276, 277, 317, 318, 319 Pennsylvania(U.S.A.), 110, 111, 113, 117, 326, 335 Pennsylvanian, 95, 98, 100, 102, 110, 111, 113, 117, 158, 185, 222, 264, 282, 336,354,376, 497 Peptides, 349, 353 Permeability/permeable, 9, 31, 36, 39, 40, 54, 65, 127, 140, 169, 180, 184, 186, 200, 210, 211, 215, 217, 222, 224, 269, 276, 286, 287, 297, 459, 464, 478, 503, 506, 513, 514, 515, 524 Permian, 31, 51, 70, 73, 75, 100, 145, 185, 222, 241, 252, 253, 264, 278, 282, 330, 335, 336, 451, 481 Persian Gulf, 57, 67 Perth (Australia), 30 Petroleum, 16,21, 33, 34, 55, 56,65, 75, 79,88, 270, 344, 345, 347, 354, 361, 362, 365, 367, 374, 375, 376, 377, 378, 379, 380, 381, 403, 448, 453, 477, 478, 480, 488, 492, 496, 498, 499 PH, 7, 9, 16, 23-25, 26, 27, 29, 31, 32, 34, 36, 40, 42, 43, 44, 45, 46, 48, 49, 50, 51, 52, 54, 55, 56, 58, 59, 64, 66, 67, 70, 71, 73, 74, 75, 78, 88, 89, 91, 99, 100, 101, 105, 106, 108, 115, 116, 118, 119, 120, 121, 143, 151, 153, 154, 155, 183, 188, 189, 190, 192, 193, 194, 210, 211, 215, 238, 240, 241, 244, 245, 261, 264, 285, 293, 327, 353, 425, 431, 448, 449, 450,459,460,461,499, 505,509, 524 pH-Eh ratio, 23,25,26,27,29,31,78,101,425, 431,448,450,460,461,499 Phenoles, 346, 355,368-374, 381,498,499 Phillipsite, 4, 47, 76, 77, 132, 135,147,148,149 Phosphatization, 66, 72-75 Phosphorites, 5, 14, 15, 27, 147, 448, 453 Photosynthesis/photosynthetic,23, 26, 33, 44, 67, 73, 75, 78, 187, 193, 278, 379, 380 Phyllomorphism (or phyllomorphic stage of diagenesis), 32, 91, 106-112, 113, 119-121, 125, 524 Pierce Canyon (Permian), 100,101 - Formation (Cretaceous), 355

547 Pierre Shale (Cretaceous), 356 Pinega River, 292 Pisolite, 194, 196, 216, 257, 258, 272, 273, 274, 275,276,277, 313, 318 Pittsburg (Pa.), 117 Plagioclase, 47, 64, 77, 92, 106, 111, 112, 120, 121, 122, 133, 160, 172,450, 507 Platteville Formation (Ordovician), 94 Playas, 93 Pleasantview Sandstone (Pennsylvanian), 95, 102 Pleistocene, 188, 189, 190, 248, 251, 251, 265, 283, 349, 350, 354 Plochingen (Germany), 503, 504, 505 Pneumatolytic, 20 Po (Italy), 130, 161, 166, 167, 510 Podol'skiy horizon, Onega (U.S.S.R.), 291 Podzolization, 24, 72 Poland, 58 Polypeptides, 344, 350 Polysaccharides, 344, 358, 359, 361, 381 Porcellanites, 324, 331-332 Porosity/porous,7,8,40,65,127,130, 135-138, 139, 140, 154, 155, 159, 160, 161-169, 171, 180, 184, 186, 200, 210, 211, 215, 217, 218, 224, 262, 269, 270, 273, 274, 276. 282, 283, 286, 287, 288, 289, 292, 295, 296, 297, 315, 316, 331, 391, 396, 398, 404, 407, 409, 503, 506, 507, 509, 510, 511, 512, 518, 519, 520 Portuguese Timor, 194, 195,236 Post-cementation (stage of diagenesis), 181, 182,208,209,230,244,248, 250, 314 Post-diastrophic (stage of diagenesis), 40 Postglacial, 62 Post-lithification (stage of diagenesis), 20 Pre-burial (stage of diagenesis), 127, 128, 129, 130, 1 6 1 5 1 , 153, 155, 158 Precambrian, 19, 31,46,49, 55, 64,67, 71, 77, 78, 107, 121, 265, 273, 291, 347, 353, 354, 360,369, 373,440,456,487 Pre-cementation (stage of diagenesis), 181,182, 208,209,243,244,248,250, 314 Pre-diastrophic (stage of diagenesis), 33 Pre-Hercynian, 455 Prehnite, 120, 172, 450 Pre-lithification (stage of diagenesis), 20 Pressure solution, 318 Proteins, 72, 192, 344, 346, 348, 349, 350, 352, 353, 357, 358, 361, 363, 369, 381, 382, 394, 395 Protobitumen, 400,414,415 Protodolomite, 67 Protoglauconite, 15 Proto-petroleum, 376, 379 Psilomelane, 15, 77 Puerto Rico, 188, 189, 191

548

SUBJECT INDEX

Purines, 343,344 Pyrimidines, 343, 344 Pyrite, 4, 13, 27, 28, 30, 35, 36, 40, 46, 51, 54, 55, 56, 68, 71, 73, 77, 91, 99, 109, 114, 153, 192, 211, 239, 240, 241, 243, 244, 247, 249, 250, 264, 265, 266, 287, 294, 394, 425, 430, 431, 433, 435, 438, 439, 444,453,460,461,463 Pyritization, 77 Pyrolusite, 15, 77 Pyrophyllite, 133, 145, 146, 172 Pyrrhotite, 28, 442,460

152, 245, 331, 441,

Quartz, 7, 15,24. 34, 36, 40,52, 58, 65, 66, 70,

71, 72, 73, 77,91, 93, 94,95, 96, 97, 98, 99, 102, 103, 104, 105, 106, 107, 108, 110, 111, 112, 113, 115, 116, 117, 118, 119, 120, 122, 131, 133, 135, 144, 147, 148, 149, 150, 160, 161, 166, 170, 171, 172, 189, 244, 272, 324, 325, 326, 327, 328, 329, 331, 333, 334, 335, 337, 338, 339, 394, 410, 422, 430,435,440, 441, 450, 453, 458, 461, 462, 463, 465, 466, 504, 505, 506, 507, 508, 509, 516, 518, 519 Quaternary, 51, 58, 62, 74, 148, 161,456 Quinones, 346,368-374

Radiolaria, 15, 36, 58, 134, 329, 331, 332 Rammelsberg, 420, 433, 434, 437 Ransom Co. (N.D.), 326 Ravenna (Italy), 161 Recrystallization, 3, 4, 6, 7, 20, 21, 22, 42, 44,

64, 104, 110, 179, 182, 184, 186, 188, 195, 196, 197, 198, 207, 216-224, 226, 227, 228, 229, 233, 235, 236, 237, 244, 245, 246, 262, 263, 265, 268, 269, 271, 272, 273, 275, 276, 277, 279, 281, 282, 283, 284, 286, 298, 312, 318, 319, 320, 322, 331, 332, 421, 428, 430, 435,450 Red clay/mud, 47, 52, 53, 133, 165 Red Sea, 191,240 Reduction/reducing, 4, 5, 7, 9, 12, 13, 16, 21, 22, 26, 30, 32, 34, 35, 36, 41, 43, 44, 45, 46, 52-56, 66, 68, 73, 74, 75, 78, 91, 97, 99, 100, 101, 102, 108, 109, 110, 111, 114, 115, 119, 120, 121, 125, 147, 152, 154, 184, 215, 244, 264, 265, 266, 293, 294, 295, 351, 352, 358, 359, 366, 393, 394, 395, 448, 449, 457, 459, 463,466,491

Redoxomorphism (or redoxomorphic stage of diagenesis), 32, 91, 97, 98, 99-102, 103, 108, 115, 119, 125

Reef, 319 Reefal limestone, 277-280 Rehydrolyzation, 56 Replacement, 13,21,22,43, 52, 59, 66, 68, 71, 77,91,92,94,97,98,102,103,105, 106,108,

110, 111, 112, 113, 114, 115, 161, 181, 183, 184, 186, 200, 226, 227, 233, 237-245, 247, 258, 259, 262, 266, 268, 276, 284, 287, 288, 289, 291, 292, 315, 316, 319, 322, 323, 324, 334, 335-337, 338, 339-340, 422, 426, 427, 430, 432, 433, 456,457, 508, 516 Resinite, 413, 414, 415 Rhexistasy, 53, 58 Rhine, 140 - Graben, 508, 515 Rhizoconcretion, 59 Rhizomorph, 59 Rhodesian Copper Belt, 435 Rhodochrosite, 4, 7, 14, 27, 77,

118, 211, 248, 281, 293, 330, 365, 435,

119, 215, 250, 282, 295, 332, 418, 446,

158, 225, 255, 283, 314, 333, 420, 447,

420

Rio San Antonio (New Mexico), 326

Rockport (Texas), 136, 142, 144 Rock salt, 13 Rowley Shelf (Australia), 45 Ruhr (Germany), 71, 157, 405, 406, 407, 517 Rushville (Ill.), 95 Rutile, 64, 77, 410 32S/s4S isotope ratio, 35, 55, 56 Saar-Nahe district (Germany), 160 Saccharoidal, 320 Sahara, 72, 133, 148 Sahul Shelf (Australia), 45 Sakhalin Island (U.S.S.R.), 479 Salt domes, 55 Salt Lake City (Utah), 246 Samarskaya Luka (U.S.S.R.), 290,292,293 Samoa, 211 San Antonio Bay (Gulf of Mexico), 142, 144 “Sand crystals’’, 118 San Diego Trough (Pacific), 348,351,358,370 Sanidine, 133 San Joaquin Valley (Calif.), 170 Santa Barbara Basin (Calif.), 130, 163, 351,

358

Saponite, 50, 159 Saprofication, 393, 394 Sapropelic coal, 393, 394,415 Saskatchewan (Canada), 261,274 Schilfsandstein (Triassic, Germany), 503, 504, 507, 508, 516 Schizophyta, 194 Sclerotinite, 414, 415 Scotland, 266,274 Scurry Co. (Texas), 376 Secondary, 320 Sedirnentogenesis, 3 Semifusinite, 414,415 Senonian, 479

SUBJECT INDEX

Sepiolite, 127, 146, 148, 160, 171 Septimus Limestone (Carboniferous, Australia), 334 Sencite, 128, 158, 171,410, 507 Sencitization, 410 Serpentine, 42, 133, 145 Shallow burial (stage of diagenesis), 108, 127, 128, 129, 130, 131, 145, 147, 151-155, 160, 165 Siberian Platform (U.S.S.R.),157 Sicily (Italy), 57 Siderite, 4, 7, 13, 14, 16, 34, 46, 75, 77, 78, 98, 100, 102, 104, 106, 154, 240, 285, 332, 394, 420, 455, 456, 457, 460, 461, 462, 463, 465 Siderite facies, 75 Sideritization, 75-78 Sideromelane, 135 Sidney (Ohio), 327 Siegenite, 431, 441 Silcrete, 24, 72 Silica gels, 57-63 Silicides, 13, 14, 15 Silicification, 66, 70-72, 179, 182, 241, 244, 247, 255, 272, 283, 332, 337, 338, 339,410 Silicoflagellates, 134 Silurian, 61, 77,95, 336 SiOa, 5, 6, 7, 20, 25, 36, 50, 58, 66, 70, 71, 72, 76,77,127,128,142,145, 146, 149, 150,151, 160, 161, 171, 238, 239, 240, 241, 244, 285, 318, 325, 326, 430, 457, 462, 463, 466, 504, 506, 509, 516 SiOa * nH~O,76 Siphoneae, 292 Skeletal, 265-269, 320 Socorro County (New Mexico), 261 Solution, 200-216 - transfer, 320 South America, 51, 53, 74, 130, 132, 133, 166, 167 Southeast Oxford Mine, 445 Sparite/sparry, 179, 185, 201, 202, 203, 205, 206, 207, 208, 209, 212, 218, 219, 220, 222, 224, 225, 226, 228, 229, 230, 231, 232, 234, 235, 236, 237, 241, 243, 246, 247, 248, 249, 251, 252, 255, 256, 257, 258, 259, 262, 265, 267, 271, 272, 273, 279, 280-284, 298, 311, 312, 313, 314, 317, 318, 320, 321, 334, 335 Sphalerite, 35, 77, 153,428,429,430,433,441, 442 Sphene, 21, 64 Spherulite, 321 Sponge, 15, 36, 58, 61,'134, 147, 241, 285, 310, 329, 331, 332, 336 Sporinite, 413,414,415 Springer Formation (Okla.), 46 Sr/Ca ratio, 223,224

549 SrCOs, 7, 77,223, 318 SrS04. 76 Stanley Shale (Mississippian,Ark., Okla.), 331 Stassfurt deposits (Permian, Germany), 31 Static metamorphism, 20 Staurolite, 506, 509 Steroids, 346, 361-365 Stems klint (Denmark), 60,62 Stockton Arkose (Triassic, Pa,), 110, 111, 113 Stromatactis, 179, 199,204,206,207,209,210, 245,250,278, 320, 321 Stromatitic, 426, 427 Stromatolite, 321 Stromatoporoids, 207,336 Strontianite, 77 Stubensandsteinflriassic, Germany), 503,504, 505,506, 507, 508, 512, 516 Stuttgart (Germany), 503, 504, 505 Stylolitelstylolitic,72, 201, 274, 428, 430, 432, 433,437-438, 506 Subalpidic foredeep, 408 Sub-graywacke, 102, 103, 104, 105, 106, 108, 113, 115, 122 Submarine weathering, 44,45, 89 Subvariscan foredeep, 406,408 Sudoite, 142, 145, 153, 158, 171, 507 Sugars, 343, 344, 355, 358, 359, 360, 361, 368, 369, 371, 381,394, 395,497 Suna Strait (Lndian Ocean), 134 Supergene/supergenic, 40, 88, 89, 181, 316, 321,432 Surinam (South America), 133 Sydney Harbor (Australia), 145 Syn-cementation (stage of diagenesis), 181, 182,208,209,230,235,243,248. 314 Syndiagenesislsyndiagenetic, 19, 20, 26, 32, 33-37, 39, 40, 41, 44,48, 49, 55, 56, 57, 59, 60,61,64,66,67,68, 70,71,73,75,77,78,89, 278, 314 Syneresis, 162, 262, 321, 333 Syngenesis/syngenetic, 13, 14, 15, 19, 33, 41, 42, 74, 77, 78, 89, 148, 179, 180, 181, 182, 184, 193, 200, 208, 209, 210, 230, 235, 243, 245, 246, 247, 248, 251, 256, 257, 264, 265, 266, 278,298, 314, 315, 321, 322, 340, 418419, 432, 435, 437, 447, 448, 455, 458, 459, 477,478,496, 504, 523 Syntaxial rim, 218,226,227,232,233,237,268, 272, 322 Tahiti, 194 Talc, 133, 144, 145, 160, 171 Tanners Creek Formation (Ordovician, Indiana), 336 Tannic acids, 34 Telinite, 413,414

550 Tennessee (USA.), 158, 326, 327, 439 Tensleep Sandstone (Pennsylvanian, Wyo.), 117

Terra rossa, 61 Tersko-Dagestan (U.S.S.R.), 479 TertiaIy, 46, 57, 73,94, 102, 104, 130, 156, 160,

166, 167, 236, 295, 337, 354, 359, 373, 398, 407, 466, 479, 496, 497, 498, 510, 511, 515 Tetrahedrite, 438 Texas (U.S.A.), 30, 50, 57, 136, 142, 144, 268, 276,278, 376,479,497 Thalassocratic stage, 58, 73 Thornton Reef (Silurian, Ind.), 336 TiOz, 76, 77 Tolga (Australia), 243, 244, 249, 250 Tornah Sandstone (Cambrian, Wisc.), 109, 111,112 Tooele County (Utah), 253,254,255 Tourmaline, 64, 77, 519 Trachytic-leucitic, 148 Transportation-weathering, 129, 130, 140-151, 171 Travertine, 322 Triassic, 70, 110, 111, 113, 366, 433, 503, 504, 505, 506, 507, 508, 512,516 Trinidad, 147 Tufa, 322 Turonian, 272 Tuscarora Sandstone (Silurian, Va.), 95 Tyrrhenian Sea, 148 Tyub-Karaganskiy Bay (U.S.S.R.), 9

Ukraine (U.S.S.R.), 479 Umatilla Co. (Ore.), 326 Upper Bavaria, 408 “Upper physicochemical zone”, 34 Upper Silesia, 410 UraI-Volga (U.S.S.R.), 479 U.S.A., 9, 12, 46,47, 49, 51, 53, 57, 64, 94, 95, 98, 100, 102, 104, 106, 107, 108, 109, 110, 111, 112, 113, 117, 130, 134, 136, 141, 145, 146, 149, 151, 152, 156, 158, 160, 163, 170, 185, 190, 193, 215, 222, 241, 246, 252, 253, 254, 255, 261, 263, 264, 265, 267, 274, 278, 283, 288, 297, 312, 326, 327, 328, 329, 330, 331, 335, 336, 337, 339, 350, 351, 358, 420, 421, 428, 430, 431, 432, 433, 435, 437, 439, 440,443,444,447,450,479,497 U.S.S.R., 8,9,46, 156, 157, 172,290,291,292, 293,479,487 Utah (U.S.A.), 146, 190, 215, 246, 253, 254, 255,264,265,270,297,312,420,497

Valanginian, 517 Vaterite, 187, 216, 217, 246, 263, 322 Venezuela, 130, 133, 166, 167

SUBJECT INDEX

Verkhoyano-Kolimskaya syncline (U.S.S.R.), 157

Vermiculite, 145, 157, 159 Vernadyte, 15 Vesuvius (Italy), 47 Virginia (U.S.A.), 95, 158 Vidan, 265 Vitrain, 413,414 Vitrinertite, 414 Vitrinite, 393,394, 398,400,401,403,404,405, 410,414,415

Vitrite, 397, 398, 400,403, 405, 414 Vivianite, 14, 16, 74 Void ratio, 164, 165, 166, 167, 168 Volcanism/volcanic, 4, 23, 29, 30, 47, 50, 58,

65, 74, 103, 115, 127, 132, 133, 134, 135, 146, 148, 149, 153, 156, 158, 171, 181, 183, 187, 194, 239, 314, 331, 332, 338, 340, 420, 437,439,448, 450,452, 487

Wadden Sea, 34, 138 Wales (Great Britain), 265, 270 Wamsutter (Wyo.), 104 Washington (U.S.A.), 326 “Waterbloom”, 55 Wavellite, 74 Waxes, 361, 363, 394 Weathering, 2, 3, 20, 32, 34, 39, 40, 41, 42, 43,

44, 48, 49, 50, 51, 53, 58, 64, 69, 71, 74, 79, 88,89,116,127,128,129,140, 146,157,181, 183, 200, 257, 274, 316, 317, 337, 339, 379, 422,456,466,477, 509 Weighscale (Pa.), 326 Welding, 322 Wendover (Utah), 254,255 West Cortez Basin, 376 Western Canada Basin, 261, 262, 321 - Turkmeniya, 479 Westfalian, 407, 408 White Pine County (Nevada), 185,255 Wilcox Formation (Eocene), 156 Winnowinglwinnow, 93, 114,256,315, 322 Wisconsin (U.S.A.), 109, 111, 112, 113, 327 Witherite, 77 Witwatersrand, 420,435 Wollastonite, 42 Woodbine Sand (Cretaceous, Texas), 479 Wurtzite, 153 Wutach-Schlucht (Germany), 154 Wyoming(U.S.A.), 104, 106, 117,274

Yakima Co. (Wash.), 326 Yangtze, 53 Zanclodon marl (Triassic, Germany), 516 Zechstein, 145

551

SUBJECT INDEX

Zeolite, 4, 5, 47, 77, 92, 99, 112, 120, 122, 127,

132, 147-148, 149, 171, 172,450, 519 - facies, 3, 47, 92, 172 Zeolitization, 47 Zircon, 64, 77, 519

ZnS (sphalerite), 77, 430, 443, 444, 448 Zoisite, 21 ZrSiO4 (zircon), 77 Zuger See, 161, 163 Zuider Zee, 136