Diagenesis in Sediments & Sedimentary Rocks (v. 2)

DEVELOPMENTS IN SEDIMENTOLOGY 25B

DIAGENESIS IN SEDIMENTS AND SEDIMENTARY ROCKS, 2

FURTHER TITLES IN THIS SERIES VOLUMES 1, 2, 3, 5, 8 and 9 are out of print 4 F.G. T I C K E L L THE TECHNIQUES OF SEDIMENTARY MINERALOGY 6 L. V A N D E R P L A S THE IDENTIFICATION O F DETRITAL FELDSPARS I S. D Z U L Y N S K I and E.K. W A L T O N SEDIMENTARY FEATURES O F FLYSCH AND GREYWACKES P.McL.D. DUFF, A . H A L L A M and E.K. W A L T O N 10 CYCLIC SEDIMENTATION 11 C.C. R E E V E S Jr. INTRODUCTION T O PALEOLIMNOLOGY 12 R.G.C. B A T H U R S T CARBONATE SEDIMENTS AND THEIR DIAGENESIS 13 A.A. M A N T E N SILURIAN REEFS O F GOTLAND 14 K.W. G L E N N I E DESERT SEDIMENTARY ENVIRONMENTS C.E. W E A V E R and L.D. P O L L A R D 15 THE CHEMISTRY O F CLAY MINERALS H.H. R I E K E III and G.V. C H I L I N G A R I A N 16 COMPACTION OF ARGILLACEOUS SEDIMENTS M.D. PICARD and L.R. HIGH Jr. 17 SEDIMENTARY STRUCTURES O F EPHEMERAL STREAMS 18 G. V. C H I L I N G A R I A N and K.H. W O L F COMPACTION O F COARSE-GRAINED SEDIMENTS 19 W. S C H W A R Z A C H E R SEDIMENTATION MODELS AND QUANTITATIVE STRATIGRAPHY 20 M.R. W A L T E R , Editor STROMATOLITES 21 B. V E L D E CLAYS AND CLAY MINERALS IN NATURAL AND SYNTHETIC SYSTEMS 22 C.E. W E A V E R and K.C. BECK MIOCENE OF THE SOUTHEASTERN UNITED STATES 23 B.C. H E E Z E N , Editor INFLUENCE O F ABYSSAL CIRCULATION ON SEDIMENTARY ACCUMULATIONS IN SPA.CE AND TIME 24 R.E. GRIM and N . G U V E N BENTONITES 25A G. L A R S E N and G . V . C H I L I N G A R I A N , Editors DIAGENESIS IN SEDIMENTS AND SEDIMENTARY ROCK, 1 26 T. SUDO and S . SHIMODA, Editors CLAYS AND CLAY MINERALS OF JAPAN 27 M.M. M O R T L A N D and V.C. F A R M E R INTERNATIONAL CLAY CONFERENCE 1978 28 A. N I S S E N B A U M , Editor HYPERSALINE BRINES AND EVAPORITIC ENVIRONMENTS 29 P.TURNER CONTINENTAL RED BEDS T. SUDO, S . S H I M O D A , H. Y O T S U M O T O and S. A I T A 31 ELECTRON MICROGRAPHS OF CLAY MINERALS 32 C.A. N I T T R O U E R , Editor SEDIMENTARY DYNAMICS OF CONTINENTAL SHELVES

DEVELOPMENTS IN SEDIMENTOLOGY 25B

DlAGENESlS IN SEDIMENTS AND SEDIMENTARY ROCKS, 2 EDITED BY

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

AND

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

ELSEVIER SCIENTIFIC PUBLISHING COMPANY 1983 AMSTERDAM -OXFORD -NEW YORK

ELSEVIER SCIENTIFIC PUBLISHING COMPANY Molenwerf 1, P.O. Box 211,1000 AE Amsterdam, The Netherlands Distributors f o r the United States and Canada: ELSEVIER/NORTH-HOLLAND INC. 52, Vanderbilt Avenue New York,N.Y. 10017

Librar? of ('ongrras ('ntaleging in tpUhlicptioti bata (Revised)

Main e n t r y under t i t l e :

Diagenesis i n sediments

and sedimentary rocks.

) (Developments i n s P d i a e n t o l o g y ; 2 5 A E d i t i o n of 1967 published under ti-tle: Diagenesis i n sediments. I n c l u d e s b i b l i o g r a p h i e s and i n d e x e s . 1. Diagenesis. I. Larsen, Gunnar, 192811. C h i l i n g a r i a n , George V . , 1929j oirit a u t h o r . 11. T i t l e . 111. S e r i e s . QE471.LsJ 1979 552' .5 78-23961 ISBN 0-444-41657-9 ( v . 1)

ISBN 0-444-42013-4 (Vol. 25B) ISBN 0-444-41238-7 (Series) Elsevier Scientific Puhlishing Company, 1983 All rights reserved. No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publisher, Elsevier Scientific Publishing Company, P.O. Box 330, 1000 AH Amsterdam, The Netherlands 0

Printed in The Netherlands

Dedicated to R.G.C. BATHURST, L. C A Y E U X K.O. EMERY, G.M. FRIEDMAN, A.P. LISITZYN, G,V. MIDDLETON, G. MILLOT, C. PRAY, S.G. SARKISYAN, H.C. SORBY and E. STEIDTMANN for their important contributions to the field of diagenesis

VI

CONTRIBUTORS

L. BUBENICEK

Societt Nationale Elf Aquitaine (Production), Pau (France)

G.V. CHILINGAR

University of Southern California, Los Angeles, Calif., U.S.A.

H.E. COOK

U.S. Department of the Interior, Geological Survey, Branch of Oil and Gas Reserves, Menlo Park, Calif., U.S.A.

R.M. EGBERT

U.S. Department of the Interior, Geological Survey, Branch of Oil and Gas Reserves, Menlo Park, Calif., U.S.A.

R.W. FAIRBRIDGE Department of Geological Sciences, Columbia University, New York, N.Y., U.S.A. H.J. KISCH

Department of Geology and Mineralogy, Ben Gurion University, University of the Negev, Beer Sheva, Israel

G . LARSEN

Department of Geology, University of Aarhus, Aarhus, Denmark

G . MULLER

Institut fur Sedimentforschung, Universitat Heidelberg, Heidelberg, Germany

A. SINGER

Department of Soils and Water, The Hebrew University of Jerusalem, Rehovot, Israel

VII

CONTENTS VI

Contributors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHAPTER 1. Introduction G.V. Chilingar and G. Larsen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHAPTER 2. Syndiagenesis- Anadiagenesis-Epidiagenesis: Phases in Lithogenesis R.W. Fairbridge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHAPTER 3. Diagenesis in Argillaceous Sediments A. Singer and G. Miiller . . . . . . . . . . . . . . . . .

1

17

...........

115

CHAPTER 4. Diagenesis of Deep-sea Carbonates H.E. Cook and R.M. Egbert . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

213

CHAPTER 5. Mineralogy and Petrology of Burial Diagenesis (Burial Metamorphism) and Incipient Metamorphism in Clastic Rocks H.J .Kisch . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

289

APPENDIX A. Diagenesis of Iron-Rich Ores (Illustrated by the Role of Diagenesis in Oolitic Iron Ores) L.Bubenicek . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

495

APPENDIX B. Mineralogy and Petrology of Burial Diagenesis and Incipient Metamorphism in Clastic Rocks-Literature Published since 1976 H.J .Kisch . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

513

References Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

543

Subject Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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1

Chapter 1 INTRODUCTION GEORGE V. CHILINGAR and GUNNAR LARSEN

THE EXPANDING INSIGHT

A revolutionary expansion of knowledge about sediments and sedimentary rocks has occurred during the last decade. This is largely due to the development of plate tectonics concepts, results of deep-sea drilling, and intensive studies by petroleum geologists. In the first edition of this work, the first one ever published on the subject of diagenesis, in 1967, it was pointed out that although the word “diagenesis” had been in the technical language for almost a century, it was not even listed in the general index of Encyclopedia Britannica: In the meantime, a new edition of the latter has appeared and the omission was corrected. Whereas in the first half of the present century. the subject could have been well-covered by a list of one hundred references, today we find well over that number appearing annually. According to Dr. Rhodes Fairbridge (personal communication, 198I ) , the revolutionary change in the knowledge and interest in diagenesis comes partly from need and partly from serendipitous discovery. The need emerged from the increasing emphasis being placed on the understanding of economic forces involved in the intensified search for petroleum, natural gas. coal, water, and other sedimentary raw materials. The discovery came hand-in-hand partly with the emergence of plate tectonics theory and, specifically, with the deep-sea drilling program. Rarely in the history of any discipline has its progress rested so squarely on the back of one single operation, the core-drilling of the “Glomar Challenger”. Repeatedly, from all parts of the world, this vessel has brought back long cores that not only contain the deep-sea stratigraphic history of the last hundred million years or so, but also disclose the actual development of diagenesis, an evolution that could previously only be deduced (R. Fairbridge, personal communication, 1981). It is also fortunate that when the cores get into the laboratory there are now a host of new techniques available for their study. In addition to the sieve-shaker, the binocular microscope, and the mineral stains of the earlier years, the big advance recently has been the development of the scanning electron microscope (SEM), which three-dimensionally supplements the information obtained by using the earlier-designed transmission electron mi-

2 croscope (TEM). There is also the luminoscope and the sophisticated equipment for isotopic analysis and trace-element identification. Arising in part from the surge of data, knowledge, and ideas, there has also appeared a number of textbooks (for example, Friedman and Sanders, 1978), important monographs, and the continuing series of Deep-sea Drilling Reports (DSDP). Also useful for those who need a quick capsule of information, alphabetically organized, is the Encyclopedia of Sedirnentoiogy of Fairbridge and Bourgeois (1978). A very convenient reprint collection of basic material of diagenesis compiled by Doris Curtis (1976) is entitled Sedimentary Processes: Diagenesis, and includes such classical works as those of ZoBell (1942), which brought out the role of bacteria, and Krumbein (1942), which established the multiplicity of physical and chemical changes. An important contribution is that of Davies and Supko ( 1973), which summarizes the results obtained during the first years of the DSDP-project. Another collection, volume 40 in the “Benchmark Series”, which was edited by Van der Lingen (1977), is devoted specifically to the hitherto little-known field of diagenesis of deep-sea biogenic sediments, now made accessible through the DSDP-drilling. The special type of diagenesis near sea-floor spreading centers is associated with hydrothermal systems, of particular interest to economic geologists. It has also been provided with a “Benchmark” treatment (Rona and Lowell, 1980).

ON THE DELIMITATION OF DIAGENESIS

An universally accepted definition of the term diagenesis does not exist. Many geologists define diagenesis as “all processes occurring between deposition and metamorphism”. This definition, however, is open for criticism, because in nature there are two distinct and different stages of alteration of rocks subsequent to deposition and prior to metamorphism: (1) conversion of sediments into sedimentary rocks, and (2) changes occurring in sedimentary rocks prior to metamorphism. One of the editors (G.V.C.) prefers to call the first stage “diagenesis” and the second stage “catagenesis”. These two stages differ not only in the character of physicochemical processes occurring in the materials composing sediments and sedimentary rocks, but also in the sources of energy required for the processes. The source of energy for diagenetic processes is the combination of various mineral components of the sediments and interstitial fluids. During the first (diagenetic) stage, an important role is played by the organic components (organic acids, etc.), microbial activity, solar energy, excess amounts of water present, etc. During the second (catagenetic) stage, on the other hand, the major role is played by temperature and pressure in

3 the alterations, which occur in the already formed sedimentary rocks. Briefly stated, diagenesis of sediments and catagenesis of rocks differ in sources of energy, physicochemical processes, and resulting products of alteration. As pointed out by Davies and Supko (1973), however, “many of the details of the vital link between the soft sediments and lithified sedimentary rocks remain a mystery.. .”. In the present book, the editors have found it necessary to leave the definition of diagenesis to individual contributors.

EFFECT OF SEDIMENTATION RATE ON DIAGENESIS

Studies on the influence of sedimentation rate on geological processes in contemporary sedimentary basins indicate that the degree of diagenetic alterations of sedimentary mineral components and associated organic matter in sedimentary deposits is a direct function of sedimentation rate and its duration (see Nazarkin, 1979). Sedimentation rate controls the degree of compaction of sediments and, consequently, determines direction, dynamics, and degree of organic matter alteration during diagenesis, which influence the intensity of hydrocarbon generation. As the sediments compact, dehydration occurs as a result of release of interlayer and adsorbed water. This dehydration intensifies hydrocarbon generation and migration. Associated liberation of heat supplements the petroleum-generation potential. Thus, in the case of low sedimentation rate one cannot expect presence of large oil and gas accumulations. Other factors being favorable, petroleum-generating potential of source rocks increases with increasing rate of sedimentation and its duration. Porosity and permeability of reservoir rocks would also be greater. Nazarkin ( 1979, p. 286) proposed the following three categories of petroleum-generating potential of paleobasins: (1) High petroleum-generating potential: average sedimentation rate ranges from 300 to 900 tons kmP2yr-I. (2) Average petroleum-generating potential: average sedimentation rate ranges from 160 to 300 tons km-2 yr-I. (3) Low petroleum-generating potential: average sedimentation rate ranges from 60 to 160 tons kmP2yr-’. The critical rate of sedimentation necessary for generation of sufficient quantities of hydrocarbons for accumulation of economical deposits is 60-100 tons kmP2 yr-I. The rate of sedimentation determines many diagenetic processes. In order to determine these relationships, however, they should be thoroughly studied.

4 SYNDIAGENESIS. ANADIAGENESIS A N D EPIDIAGENESlS

I n Chapter 2, Professor Rhodes W. Fairbridge recognizes three phases of diagenesis: syndiagenesis (penecontemporaneous), anadiagenesis (during burial and orogeny), and epidiagenesis (following emergence) as proposed by Fairbridge ( 1967). Slightly different (temporal) emphasis is expressed by the almost synonymous Choquette and Pray ( 1970) terms of eogeneric (early), mesogeneric (middle), and telogmetic (late) stage. The phase system expresses three conditions: time, place, and process. Initially, there is penecontemporaneous exchange with ambient media. Then there is a long-term exchange mainly with ascending and circulating connate and compaction waters, with an increasing role of temperature and pressure. Finally, there is the late-stage interplay with meteoric water, which often sinks to considerable depths with increased mixing at depth. Pervading each phase in the system is pore water, connate or meteoric, which is characterized by its dissolved components and by pH- and Eh-values. Besides precipitation and solution. there is reprecipitation (neoformation), authigenesis, and pseudomorphism. The first-phase waters start as normal lake or sea water, but become modified rapidly, either becoming supersaline as one goes down into the anadiagenetic realm, where temperature and pressure concurrently increase. or becoming fresher with depth due to compaction mechanism. According to Fairbridge, with uplift and re-exposure, the cycle reverses: temperature and pressure decrease and fresh water replaces or dilutes the saline fluids. In ancient formations, the cycle of anadiagenesis and epidiagenesis may be repeated again and again, but only one syndiagenesis cycle is possible. Cementation may, thus, be followed by leaching and then by recementation once more. In each phase, an equilibrium or steady state may be achieved. Events of geologic history, however, abruptly terminate the evolutionary development in any one of these phases, recycling or bypassing one or another. Carbonate diagenesis (treated in detail in Chapter 4 of this volume and Chapters 6 and 7 of Vol. I) is briefly discussed in Chapter 2. Carbonates are “peculiarly subject to many stages of diagenesis” (Wilson, 1975, p. 16) and are sensitive, with rapid response, to varying types of pore water. Remarkable advances have been made in this field during the last two decades, ranging from the actualistic settings and paleogeography to petrography and geochemistry. Bathurst ( 1 975, 2nd ed.) has provided the definitive study, whereas further treatments are available from Milliman (1974) and others. A contribution by Dr. L. Bubenicek on diagenesis of iron-rich ores is presented i n Appendix A.

5 DIAGENESIS IN ARGILLACEOUS SEDIMENTS

In Chapter 3, entitled “Diagenesis in Argillaceous Sediments”, Arieh Singer and German Miiller apply the term diagenesis to all changes which take place in a freshly deposited sediment until it reaches the stage of metamorphism. The following subjects are covered in that chapter: ( 1) initial composition of Recent argillaceous sediments, (2) initial (pre-burial) porosity and structure of argillaceous sediments, (3) changes in chemistry and mineralogy during diagenesis, (4)changes in structure during diagenesis, and ( 5 ) correlation of mechanical and chemical-mineralogical changes with depth of burial. The transitional zone between diagenesis and metamorphism is also discussed in that chapter. According to Singer and Miiller, argillaceous muds, with an initial porosity of 70-90%, are compacted during shallow-burial (burial depth down to 500m) diagenesis to mudstone (or shale if fissile) with a porosity of about 30%. Argillite with a porosity of < 10% forms on further compaction. The decreases in porosity and water content are rapid down to about 250-300 m burial depth and slow below that depth. Diagenesis passes into low-grade metamorphism when recrystallization of sheet-silicates occurs at about 200°C and about 10,000 m burial depth. Changes in porosity and water content are mainly a function of maximum overburden (burial depth) and of time. They are, however, also affected by geothermal gradients, tectonic stresses, lithology, depositional environments, overpressured fluid zones, and diagenesis. Under compaction, the individual particles or “domains”, making up the “card-house”-microstructure of dispersed muds in non-saline environments, attain a high degree of parallel orientation. The degree of orientation upon compaction of the much thicker domains (or oriented aggregates) making up the “book-house”-structure of flocculated muds, is less marked. During the pre-burial stage of marine diagenesis, ion exchange is the major chemical-mineralogical process. Exchangeable CaZi on the clays is exchanged for M g 2 + , K + and N a + from the sea water. Part of the adsorbed cations become non-exchangeable. Chlorite and illite-like minerals may thus be formed from degraded precursor clay minerals. Clay-mineral lattice transformations at this stage are minor. In the supersaline marine environment, transformations are more pronounced and include the transition of illites to chlorite through an intermediate stage of corrensite formation. In the supersaline lake environment, zeolites are the most prominent diagenetic products. Glauconite and manganese nodules are other products of pre-burial marine diagenesis. Zeolites, silica-minerals, palygorski te, and sepiolite form both during the pre-burial and shallow-burial diagenetic stages. Palagoni tc. smectites, and oxides form as a result of the submarine alteration of basalts and pyroclastics.

6 During deep-burial and compaction, the total electrolyte content of pore fluid decreases and its composition changes. With increasing depth, sediment pore water exhibits a strong depletion in Mg2+ and K + and a corresponding enrichment in C a 2 + . Dissolution of plagioclase and formation of Mg-rich smectites appear to be responsible for these changes. The deepest pore waters are enriched in dissolved SiO,. The major mineralogical process during deep-burial diagenesis is the progressive conversion of smectite into illite or into a mixed-layer illitesmectite, having a high proportion of illite layers, with increasing burial depth. This process is occurring in association with other minor changes, such as decomposition of coarser-grained mica and K-feldspar, decomposition of kaolinite, and formation of chlorite or chlorite interlayers. The fixation of K in newly-formed illite layers involves an increase in the net negative charge of the expandable layers, resulting either from the substitution of Al’+ for Si4+ in the tetrahedral layer or by the substitution of divalent for trivalent cations in the octahedral layer. The K cation necessary for the conversion is not supplied by pore solutions from a distant source, but is derived from K-feldspar and mica within the sediment. The K is thus being redistributed between the detrital and diagenetic phases. The conversion process appears to be directly related to temperature. Thus, reaction stages are controlled by burial depth and geothermal gradients. The magnitudes of the activation energy for the conversion process, determined experimentally, suggest that breaking of chemical bonds in the tetrahedral sheet is involved. Kaolinite is eliminated during deep-burial diagenesis, while chlorite and/or corrensite is formed, possibly by the interlayer adsorption of amorphous Fe20, and/or A1,0, onto expandable layer silicates. The deepest stage of diagenesis is characterized by a uniform clay-mineral association “illitechlorite”. During the transition to metamorphism it changes to a paragenesis quartz-albite-muscovite-chlorite. The burial diagenesis of silicate minerals is correlated with that of organic matter.

DIAGENESIS OF DEEP-SEA CARBONATES

As pointed out by H.E. Cook and R.M. Egbert in Chapter 4, four major processes modify physical properties, state of fossil preservation, texture, and geochemistry of deep-sea carbonates. These include: ( 1) gravitational compaction (mainly within the first 200m of burial); (2) dissolution (at the sediment-water interface, at shallow-burial depths, and some within the water column); (3) pressure-solution (mainly during the deep-burial phase); and (4) cementation (beginning within the first few centimeters of burial and

7 continuing with increasing burial depth and age). Pelagic carbonates are widespread and comprise about 70% of the carbonates deposited worldwide during the past 100 m.y. Deep-sea carbonates consist mainly of planktonic foraminifera, coccoliths, and discoasters whose tests are composed of stable low-magnesium calcite. This is in sharp contrast to the highly reactive aragonite and high-magnesium calcite that comprise shallow-marine carbonates. One of the most unique features of deep-sea carbonates is that they are generally not contacted by fresh water and, if they are, this does not occur until very late in their diagenetic history. In shallow-marine carbonates, early cementation, produced in part by contact with fresh water, preserves much of the original fabric and retards compaction in many cases. In deep-sea carbonates, on the other hand, because fresh water is absent and stable low-magnesium calcite is the main carbonate constituent, cementation is usually somewhat delayed and lithification by compaction is the rule within part of the sediment column. At shallow-burial depths of 50-200 m, gravitatiorial compaction is the dominant mechanism for reducing porosity from about 80 to about 60% with cementation being a subordinate pore-reducing process. Compaction in this shallow-burial stage takes place by closer packing of grains, some crushing of microfossil tests, and disintegration of delicate planktonic forams and coccoliths into abundant micron-sized crystals through dissolution along sutures. Discoasters appear to be the first organisms to develop secondary calcite overgrowth. This can occur within the first 35 m of burial before gravitational compaction produces a grain-supported texture, which suggests that dissolution-diffusion-reprecipitation processes occur early, while the grains are still dispersed. With increasing burial depth, a grain-supported fabric develops and pressure-solution becomes a major process in producing calcium carbonate for cementation, with gravitational compaction being a subordinate process. According to Cook and Egbert, the most advanced diagenetic state observed in Deep-sea Drilling Project cores is evidenced by the following features: (a) virtually all fossil grains are covered with subhedral to euhedral calcite cement, (b) the central areas of coccoliths are filled, (c) foram chambers are filled with granular calcite, (d) large amounts of interparticle cement are found, and (e) grain interpenetration and welding are common. Thus, at burial depths on the order of 1000- 1 200 m, porosities have been reduced to about 20%. Continued diagenesis produces an “ameboid mosaic” or “pavement mosaic” texture of completely interlocking grains. The degree to which a deep-sea carbonate becomes cemented, probably depends to a large extent upon the diagenetic potential the carbonate

8 sediment had at the time it was buried. Diagenetic potential is a measure of how much more diagenesis a sediment can undergo in the normal course of its history. There is a trend towards a decreasing Sr2+ content and increasing S " 0 values (negative) in pelagic carbonates with increasing burial depth. This is to be expected as most deep-sea carbonates are diagenetically altered after burial. With increasing burial depth, dissolution-reprecipitation processes reduce porosity. The Sr2+ originally present in the biogenic calcite is virtually excluded from the secondary cement derived from the dissolution and pressure-solution of calcite microfossils. The S " 0 values of this cement become increasingly more negative with burial depth, due to precipitation of the cement from pore fluids at increasingly elevated temperatures along the geothermal gradient. Acoustic reflectors, which are probably related to the degree of cementation in the sediments, may be predetermined in their broad characteristics by major oceanic events, which in turn largely control the nature (diagenetic potential) of the biogenic material reaching the sea floor. Deep-sea chalks form excellent petroleum reservoirs under certain circumstances: (a) when their porosity is retained at deep-burial depths if they are hydrostatically overpressured, which decreases grain-to-grain stress and, consequently, retards pressure-solution processes; (b) when early oil emplacement may retard dissolution-reprecipitation processes; (c) by lack of deep burial depths; and (d) when fracturing enhances the chalk's permeability.

BURIAL DIAGENESIS AND INCIPIENT METAMORPHISM

In Chapter 5 , Dr. H.J. Kisch reviews the mineralogical changes taking place during burial diagenesis and incipient metamorphism from the point of view of a metamorphic petrologist, stressing the attainment of equilibrium, the compositional constraints on the reactions, and the relation of the stages of burial during which the various changes take place. The similarities and largely gradual differences between diagenetic and metamorphic reactions are discussed, particularly regarding the degree of attainment of chemical equilibrium and the extent of mineralogical and textural reconstitution. Although persistence of metastable clastic relics and of clastic textures is more common at low temperatures due to kinetic factors (slow reaction rates), the processes operating in burial diagenesis are concluded to be essentially metamorphic in nature. Kisch gives a general introduction to the concept of diagenesis, the divergent uses of the term and its synonyms. He points out the absence of a

9 natural delimitation between late diagenesis and metamorphism and stresses the arbitrariness of such delimitations. Particular attention is paid to the various petrographic, textural, and mineral-facies criteria proposed for the distinction between diagenetically altered sedimentary rocks and metamorphic rocks. Following Coombs (1954, 1958, 1960, 1961, 1971), Kisch uses the degree and type of textural modification for a loosely defined subdivision of deep-burial phenomena into: ( 1) burial metamorphism, and (2) incipient regional metamorphism. Coombs’ term “burial metamorphism” is used for partial or complete reconstitution of clastic rocks on a regional scale, characteristically without development of a penetrative fabric. Some investigators consider burial metamorphism a synonym of burial, late, or epigenetic diagenesis (epigenesis or catagenesis of some authors I ) . “Initial metamorphism” is used for the more advanced stages of mineral modification, of which slaty cleavage is a common, though not an essential, attribute. It has very general mineral-facies connotations, including most of the prehnite-pumpellyite facies, includes terranes showing the attributes of Kubler’s (1967b) anchimetamorphism, and is considered by some authors to be approximately equivalent to the stage of “early metagenesis”. The second part of Chapter 5 deals with the modification of clay mineralogy of clastic sedimentary rocks upon deep burial. The following burialdiagenetic processes are reviewed: (a) The progressive replacement of smectite by illite through random illite-smectite mixed-layers and, commonly, through a subsequent stage of regular or partly ordered illite-smectite mixed-layers. Depth and temperature of the onset of mixed-layering, the non-continuous nature of the dehydration process, and the related question of the stability of illite-smectite mixed-layers are also discussed. (b) The progressive replacement of smectite by chlorite through chloritesmectite mixed-layers, including the regular mixed-layer corrensite. (c) Changes in the polytype of kaolinite-group minerals. (d)The replacement of kaolinite by illite and/or chlorite and, less commonly, by dioctahedral, aluminous chlorite. (e)Changes in the polytype of illite, and its dependence on the nature of its predecessor (illite-smectite mixed-layer; kaolinite) and on illite composition. ( f ) Changes in composition and polytype of trioctahedral chlorite. The

’

Possibly “early rnetagenesis” of Soviet authors (see discussion in the Introduction chapter of Vol. I of this book).

10 dependence of the various modification processes on compositional variables is stressed, including the effect of bulk composition, particularly the breakdown of clastic minerals such as biotite, K-feldspar, and plagioclase. The importance of the composition of interstitial solutions and, hence, of porosity and permeability is discussed. The following ongoing changes are reviewed during the incipient metamorphsm or anchimetamorphism, when random illite-smectite mixed-layers and kaolinite have disappeared and illite and chlorite have become the predominant layer-silicate minerals: (a) Composition and polytype of illite-muscovite. (b) Determination of the progressive increase in illite crystallinity (and its complicating factors), and the use of illite crystallinity in defining an anchimetamorphic zone. (c) Appearance of pyrophyllite and conditions of the kaolinite pyrophyllite-quartz equilibrium. (d) Appearance of Na-rich illite-smectite mixed-layers (rectorite or allevardite) and Na-rich, paragonitic micas. (e) Appearance of some other minerals locally associated with the anchimetamorphic zone, such as stilpnomelane and chloritoid. (f) The rank of coal associated with anchimetamorphism. Finally, tentative schemes of lowest-grade mineral facies for clastic rocks are presented. The third part of Chapter 5 deals with lowest-grade metamorphic facies in volcanic and lithic-feldspathic sedimentary rocks, as defined by the appearance of diagnostic zeolites such as laumontite and, at more advanced grades, by prehnite, pumpellyite, etc., without zeolites. Some well-documented cases of mineral zoning are described, including the classical terranes of South Island, N.Z., and Honshii, Japan. Particular attention was paid by Kisch to the various compositional controls on the zonal distribution of various burial-metamorphic zeolites and other Ca-Al-silicate minerals and telescoping of their distribution zones. He discusses: ( 1) selective replacement of different primary mineralogicpetrographic constituents, such as the preferential replacement of silicic pyroclastics by alkali-zeolites and analcirne; (2) persistence of laumontite into the prehnite-pumpellyite zone in silicic tuffs; and (3) dissimilar extent of reconstitution in rocks of contrasting mineralogic composition, including the notion of mineral facies. The various zeolite-facies equilibria are discussed, followed by a discussion of the constraints upon their application, which include: ( 1) reduced fluid pressure (osmotic conditions) and the consequent effect of porosity and permeability differences; (2) salinity and alkalinity of the fluid phase; (3) oversaturation in silica and ionic activity gradients controlled by hydrolysis

11 and solution of acid volcanic glass upon burial; (4) ionic equilibria and coupled solid-solution reactions, and ( 5 ) variation of the pco,/p H , O ratio (hydrous Ca-Al-silicates may be suppressed at high ratios). A number of distinctive types of progressive successions of lowest-grade metamorphic mineral assemblages or “facies series” are distinguished, based on the experimental evidence on the pressure dependence of diagnostic equilibria in the pumpellyitic facies. The final sections of Chapter 5 deal with the rank of coal and with associated clay minerals in the burial-metamorphic facies. Kisch concludes that the laumontite zone of the zeolite facies is associated with coals of not higher than low-volatile bituminous rank, and that it is entirely within the stage of “deep epigenesis” of Kassovskaya and Shutov (1963) or the zone of “diagenesis” of Kubler (1964, 1967a,b, 1968, 1970). The prehnitepumpellyite facies starts at a similar or somewhat lower rank than the stage of “early metagenesis” or the anchimetamorphic zone. Relations between illite crystallinity and coal rank and the onset of pumpellyitic facies are different in high-pressure, lawsonite-bearing terranes. Coordinated studies of the different kinds of parameters of burial diagenesis and incipient metamorphism are likely to contribute to the understanding of pressure- temperature gradients in burial diagenesis and incipient metamorphism, and to contribute to the recognition of thermal events and tectonic movements.

THE DOLOMITE PROBLEM

Although dolomitization was discussed in detail in Vol. I of this book, the following information sheds additional light upon the formation of dolomite, which is still one of the great puzzles of geology. Sea water is supersaturated with respect to dolomite and yet it does not appear to precipitate there. In addition, dolomite has not been synthesized in the laboratory under the present-day earth-surface conditions. Much of the dolomite forming at the present time and that of the geological record, at least the Phanerozoic, has formed by replacement. One particular dolomite problem concerns the dominance of dolostones over limestones in the Precambrian. Two basic hypotheses have been put forward to explain this: (1) age: being older, the original CaCO, sediments have had more chance to come into contact with dolomitizing fluids; or (2) different composition of sea water and of atmosphere: Precambrian sea water and atmosphere, unlike those of the Phanerozoic, permitted dolomite precipitation and/or extensive dolomitization of limestones. New information on the Precambrian dolomite problem has recently been presented by Dr. Maurice Tucker (University of Newcastle-upon-Tyne;

12

personal communication, 198l), based on a petrographic-isotopic study of the Late Proterozoic Beck Spring Dolomite of eastern California. Tucker’s results show that this Precambrian dolomite, and possibly many others too, are different from Phanerozoic dolomites in terms of their fabrics and isotopic ratios. The Beck Spring Dolomite (? 0.9 to 1.2 billion years old) is a typical platform carbonate with intertidal and subtidal facies of cryptalgal laminites, stromatolites, micrites, and grainstones (of peloids, ooids and pisolites). The most conspicuous feature of this dolomite is the preservation of fabric details. On a microscopic scale, all fabric details are preserved, even though the carbonate mineral is dolomite. There is no pervasive or selective dolomite replacement of calcite, as is usually the case with Phanerozoic dolomites (see figs. 7-41 to 7-50 in Vol. 25A in the Series “Developments in Sedimentology”). Ooids have a radial-fibrous and concentric structure identical to the calcitic ooids of the Phanerozoic. Diagenetic cements are directly comparable to those of limestones with the early isopachous fibrous dolomite crusts, which are present around grains and line cavities. There are also later sparry dolomite cavity fills, with delicate growth zones. The petrography shows that Phanerozoic-type dolomitization of an original limestone has not taken place. Carbon and oxygen stable isotope analyses of depositional grains (pisolites and micrite), early fibrous dolomite cements, and later sparry dolomites confirm that dolomitization has not taken place. The various components possess distinctive isotopic ratios. They exhibit a distinct trend towards lighter, more negative 6I8O and, to a lesser extent, towards lighter 6I3C, from the grains, through the fibrous dolomite, to the dolomite spar (Fig. 1-1). This isotopic trend is directly comparable to that obtained from the equivalent components in the Phanerozoic limestones (e.g., Walls et al., 1979). The 6I8O and 613C values of the depositional grains and micrite reflect the isotopic composition (and temperature) of sea water at the time. The ratios for the fibrous dolomite cements indicate a marine origin with some later recrystallization in lighter waters. The dolomite spar with more negative 6 ’ * 0 reflects precipitation at higher temperatures during burial. The decreasing S13C indicates a contribution of ” C from diagenetic decomposition of organic matter. Although many geologists believe that practically all Recent and ancient dolomites are replacive in origin, the isotopic and petrographic data of Maurice Tucker (personal communication, 1981) from the Beck Spring Dolomite of eastern California are best interpreted in terms of primary dolomite precipitation. Dolomitization is now believed to be by most geologists as being a wet process involving dissolution-precipitation, and in as much as water is involved, isotopic exchange will occur. Retention of

13

tl

+E

t6

15

+5

+l

6 1

t3

k 3

+2

,2

+1

11

0

0

-1

-1

-2

-2 I

-3 17

10

19

20

21

22

23

2L

&”o

25

26

27

28

29

30

31

-3

32

SMOW

Fig. 1-1. Carbon and oxygen isotope results from the Beck Spring Dolomite, late Proterozoic. eastern California. Symbols: 0 =pisolites; M =micrite; A =cryptalgal laminite (also micrite); =fibrous dolomite; V =sparry dolomite; I S =internal sediment. (Courtesy of Dr. Maurice Tucker, personal communication, 198 1.)

original isotopic ratios effectively rules out a major dolomitizing event, as does the fabric evidence for lack of replacement. An interpretation of original dolomite for this Late Proterozoic dolostone implies that the composition of the Precambrian sea water, which permitted primary dolomite precipitation, was somehow different. Higher Mg/Ca ratio, higher p C 0 , and lower SO:- content are all possibilities. Finally, it should be mentioned in this Introduction that the present Volume I1 of “Diagenesis in Sediments and Sedimentary Rocks” will be succeeded by a third volume, which will cover the following subjects: Compactional Diagenesis, Diagenesis of Ore Deposits, Diagenesis of Evaporites, and Diagenesis of Organic Matter and Generation of Oil.

REFERENCES A N D BIBLIOGRAPHY Ali, S.A. and Friedman, G.M. (Compilers), 1977. Diagenesrs of Sandstones. Am. Assoc. Petrol. Geol., Reprint Ser.. 20: 239 pp.

14 Aoyagi, K., 1979. Paleo-temperature analysis by authigenic minerals in sedimentary rocks. J. Jpn. Assoc. Petrol. Technol., 44(6): 1-5. Aoyagi, K. and Kazama, T., 1980. Sedimentary mineralogy of argillaceous sediments from deep-sea drilling project holes 436, 438A, and 439, Japan Trench. In: Honza et al. (Editors), Initial Reports of the Deep-sea Drilling Project, 56/57. U S . Gov. Print. Off., Washington, D.C., pp. 101 1-1017. Bathurst, R.G.C., 1975. Carbonate Sediments and Their Diagenesis. Elsevier, Amsterdam, 2nd ed., 658 pp. Choquette, P.W. and Pray, L.C., 1970. Geologic nomenclature and classification of porosity in sedimentary carbonates. Bull. Am. Assoc. Pet. Geoi., 54: 207-250. Coombs, D.S., 1954. The nature and alteration of some Triassic sediments from Southland, New Zealand. R. SOC.N. Z. Trans., 82( 1): 65- 109. Coombs, D.S., 1958. Zeolitized tuffs from the Kuttung Glacial Beds near Seaham, New South Wales. Aust. J. Sci., 21: 18-19. Coombs, D.S., 1960. Lower-grade mineral facies in New Zealand. Rep. Int, Geol. Congr., 21st Sess. (Norden), 13: 339-351. Coombs, D.S., 1961. Some recent work on the lower grades of metamorphism. Aust. J. Sci., 24(5): 203-215. Coombs, D.S., 1971. Present status of the zeolite facies. Ado. Chem. Ser., 101: 317-327. Coombs, D.S. and Whetten, J.T., 1967. Composition of analcime from sedimentary and burial metamorphic rocks. Geol. Sac. Am. Bull., 78: 269-282. Coombs, D.S., Ellis, A.J., Fyfe, W.S. and Taylor, A.M., 1959. The zeolite facies. with comments on the interpretation of hydrothermal syntheses. Geochim. Cosmochim. Acta, 17: 5 3 4 07. Coombs, D.S., Horodyski, R.J. and Naylor, R.S., 1970. Occurrence of prehnite-pumpellyite facies in northern Maine. Am. J . Sci., 268: 142-156. Curtis, D.M. (Compiler), 1976. Diagenesis. S.E.P.M. Reprint Ser. 1: 216 pp. Davies, T.A. and Supko, P.R., 1973. Oceanic sediments and their diagenesis: some examples from deep-sea drilling. J. Sediment. Petrol., 43(2): 38 1-390. Fairbridge, R.W., 1967. Phases of diagenesis and authigenesis. In: G. Larsen and G.V. Chilingar (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, pp. 19-89. Fairbridge, R.W. and Bourgeois, J., 1978. The Encyclopedia of Sedimentologv. Dowden, Hutchinson and Ross, Stroudsburg, 901 pp. Friedman, G.M. and Sanders, J.E., 1978. Principles of Sedirnentoiogy. Wiley, New York, N.Y., 792 pp. Kassovskaya, A.G. and Shutov, V.D., 1963. Facies of regional epigenesis and metagenesis. Izv. Akad. Nauk, S.S.S.R., Ser. Geoi., (7): 3-18. Krumbein, W.C., 1942. Physical and chemical changes in sediments after deposition. J. Sediment. Petrol., 12: 111-1 17. Kubler, B., 1964. Les argiles, indicateurs de metamorphisme. Reu. Inst. Fr. Pet., 19: 1093- 11 12. Kubler, B., 1967a. La cristallinite de I’illite et les zones tout a fait superieures du metamorphisme. In: Etages tectoniques. A la Baconnikre, Newhatel, pp. 105- 121. Kubler, B., 1967b. Anchimetamorphisme et schistosite. Bull. Cent. Rech. Pau, l(2): 259-278. Kubler, B., 1968. Evaluation quantitative du metamorphisme par la cristallinite de M i t e . Bull. Cent. Rech. Pau, 2(2): 385-397. Kubler, B., 1970. Crystallinity of illite. Detection of metamorphism in some frontal parts of the Alps (abstr.). Fortschr. Mineral., 47 (Beih. 1): 39-40. Kubler, B., Martini, J. and Vuagnat, M., 1974. Very low-grade metamorphism in the Western Alps. Schweiz. Mineral. Petrogr. Mitt., 54(2/3): 46 1-469.

15 Milliman, J.D., 1974. Marine Carbonates, 1 . Recent Sedimentary Carbonates. Springer, New York. N.Y., 375 pp. McBride, E.F. (Compiler), 1979. Diagenesis of Sandstone: Cement - Porosity Relationships S.E.P.M. Reprint Ser., 9: 233 pp. Nazarkin, L.A., 1979. Influence of Sedimentation Rate and Erosional Sections on Oil and Gas Potentials of Sedimentary Basins. Saratov Univ., 336 pp. Rona, P.A. and Lowell, R.P. (Editors), 1980. Seafloor Spreading Centers: I+drothermal Systems. Dowden, Hutchinson and Ross, Stroudsburg, Pa., 424 pp. Van der Lingen, G.J. (Editor), 1977. Diagenesis of Deep-sea Biogenic Sediments. Benchmark Papers in Geologv, 40. Dowden, Hutchinson and Ross, Stroudsburg, Pa., 385 pp. Walls, R.A., Mountjoy, E.W. and Fritz, P., 1979. Isotopic composition and diagenetic history of carbonate cements in Devonian Golden Spike reef, Alberta, Canada. Bull. Geol. SOC. Am., 90: 963-982. Wilson, J.L., 1975. Carbonate Facies in Geologic Histoty. Springer, Berlin, 471 pp. ZoBell, C.E., 1942. Changes produced by micro-organisms in sediments after deposition. J . Sediment. Petrol., 12: 127- 136.

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17 Chapter 2

SYNDIAGENESIS-ANADIAGENESIS-EPIDIAGENESIS: LITHOGENESIS

PHASES IN

RHODES W. FAIRBRIDGE

INTRODUCTION

Diagenesis of sediments is understood by the writer as those physical and chemical changes that the sediment undergoes during and after deposition and lithification, without introduction of great heat or pressure. Three phases of the overall rate process were defined (in the 1967-edition of this work) bn a temporal and spatial basis. In an ideal model, a dynamic evolution takes place within the constraints of increasing time, depth of burial or exposure, and changing hydrologic-geochemical systems. Syndiagenesis is the early phase, synchronous with deposition and early burial, being strongly influenced by biochemical agencies in a fluid regime dominated by the entrapped waters of the sea- or lake-floor. Anadiagenesis is the deep-burial phase marked by compaction and maturation, strongly influenced by increasing pressure and the upward expulsion or transit of connate waters that are often highly mineralized or saline. Epidiagenesis follows uplift or emergence that brings surficial, meteoric waters into circulation often displacing pre-existing fluids and reversing numbers of geochemical processes. Eustatic fluctuations or tectonism may accelerate or short-circuit the ideal triple-phase evolution. The term “diagenesis” has been in the geological literature for over a century, and it is hardly surprising, therefore, that it has evolved in its meaning. First introduced by Von Guembel (1868, p. 838), it comprised all post-sedimentary modifications of mineral constituents, ultimately to the level of metamorphism. This very broad definition was never adopted, and indeed Van Hise (1898, 1904), probably in ignorance of Von Guembel’s term, proposed that “metamorphism” for its part, should cover the identical field. By tacit rather than explicit agreement among geologists, metamorphism generally is taken to begin where diagenesis ends, i.e., at a temperature of around 200°C. There is clearly an area of overlap of about 10O-30O0C, depending upon the individual writer’s point of view (see, for example, Fyfe et al., 1958). In the new glossary of Bates and Jackson (1980), “diagenesis” is divided into a mineralogical usage and a sedimentological usage. In the former, it refers to any geochemical or crystal rearrangement of minerals (here they

18 specifically cite clay minerals, but this exclusivity is not necessary) that may affect the particles either before or after burial. In the sedimentological usage -the more usual approach-it means “all the chemical, physical and biological changes undergone by a sediment after its initial deposition, (both) during and after its lithification, exclusive of surficial alteration (weathering) and metamorphism”. Further.. . “it embraces those processes (such as compaction, cementation, reworking, authigenesis, replacement, crystallization, leaching, hydration, bacterial action, and formation of concretions) that occur under conditions of pressure (up to 1 kb) and temperature (maximum range of 100°C-300°C) that are normal to the surficial or outer part of the Earth’s crust; and it may include changes occurring after lithification under the same conditions of temperature and pressure”. In this sense the term was effectively utilized and popularized by Walther (1894, pp. 693-71 l), the “father of sedimentology”, as an actualistic science. In this meaning it was widely adopted in the English-speaking world (Twenhofel, 1932), in Germany (Correns, 1950), and France (Dunoyer de Segonzac, 1968). In the U.S.S.R., however, most geologists restrict the expression to cover only the initial, uncompacted phase (Fersman, 1922); after lithification, the term “epigenesis” (or more recently “catagenesis”) is employed (pers. comm. G.V. Chilingar). Some western writers have also adhered to this narrow approach, but the Bates and Jackson (1980) synthesis is preferred by the present author. 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), whereas 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.

LITHIFICATION

Lithification is the most striking aspect of diagenesis. I t is the “conversion of a newly deposited, unconsolidated sediment into a coherent, solid rock, involving processes such as cementation, compaction, desiccation, crystallization ... concurrent with, soon after, or long after deposition” (Bates and Jackson, 1980). Walther ( 1894) recognized that soft sediments did not become hard and “lithified” into rock merely by the action of time. Hutton ( 1788) believed that all lithification was due to heat and pressure that

19 involved partial fusion following deep burial. With his far greater experience of the younger formations and contemporary sedimentation, Walther was able to discern a continuum that led to his deduction that there were specific chemical and physical processes characteristic of diagenesis that took place under low temperature and pressure conditions. Lithification, according to Grabau (1913, p. 75 1) includes the following (with comments by the writer): (1) Congelation (e.g., the physical dehydration and hardening of silica or organic gels to form an amorphous or cryptocrystalline solid). (2) Crystallization (the primary reorganization of compounds, as in unstable pyroclastic sediments, but without participation of the interstitial water or of other 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 CaCO,, SiO, and Fe203, derived from circulating waters. Additional factors in diagenesis have been noted by Andree (191 1). They include: (a) formation of concretions (both in the pre- or post-lithification phases), and (b) desalinification (a post-lithification 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: ( 1) 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, CaSO, .2H,O, to anhydrite, CaSO,, 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 minerals). (4) Polymerization and depolymerization (e.g., natural catalytic “cracking” and other organic chemical reactions characteristic of the hydrocarbons). AUTHIGENESIS

After lithification, the next most striking aspect of diagenesis is the formation - of new minerals within an enclosing sediment or sedimentary

20 TABLE 2-1 More common authigenic minerals (excluding halides) I Mineral

Formula

Frequency

Usual development synanaepidiagen- diagen- diagenetic etic etic

Anatase Anhydrite Ankerite Aragonite Azurite Barite Bornite Brookite Calcite Celestite Cerussite Chalcedony Chalcopyrite Chamosite Chlorite group Collophane Dahllite Dolomite Galena Glauconite Greenalite Gypsum Halite Hematite Hydromagnesi te Illite

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

Kaolinite Leucoxene Limonite

X'

X

X Magnesi te Malachite Marcasite Mon tmorilloni te Muscovite Natron Nesquehonite Opal

X

X

X

x

X X X X

X X

X X

21 TABLE 2-1 (continued) Mineral

Formula

Frequency

Usual development synanaepidiagen- diagen- diagenetic etic etic

Orthoclase Phillipsite Plagioclase Psilomelane (Wad) Pyrite Pyrolusi te Quartz Rhodochrosi te Rutile Siderite Sphalerite Strontianite Sulphur Tourmaline

KAISi,O, (Ca,Ba,K,Na),Al x(Al,Si)2 SiloO,o. 15-20 H,O (Ca,Na)(Al,Si)AlSi ,O,

Wi therite Zeolites (Phillipsite, Heulandite, Laumontite, Chabazite, Natrolite, Analcime) Zircon ZrSiO,

C

X

R C

X

C C C C

X

R C R R R

X X X X

R R

X

C R

X

X

’ In part, after Twenhofel (1950, p. 288) and Teodorovich (1961). ’ R=rare; C=common. rock. It may be achieved by replacement (metasomatism), recrystallization, or by the filling of voids. Some authors would include secondary enlargement as in quartz overgrowths. As proposed by Kalkowsky (ISSO), this is authigenesis, the process of forming any new or secondary mineral, but now specifically in a sedimentary environment. It is also termed “neoformation” or “neogenesis” in the U.S.S.R. and France (Millot, 1970). We exclude from the term “authigenesis” such processes as solution and re-precipitation phenomena, simple cementation, decementation and intrastratal solution, including, for example, cone-in-cone, and overgrowths (addition of like ions). 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

22 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. Tester and Atwater (1934) emphasized that such minerals must be regarded as discrete crystallographic units, rather than rock-forming 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 of their origin has been provided by Teodorovich (1961). During diagenesis, new minerals often form as a result of a reaction between the ions of the interstitial water and the primary particles. For example, the reactions may involve the simple addition of some new and different ions (adsorption), the exchange of ions, or the replacement of certain ions by some new ions (metasomatism). The secondary overgrowth of some new ions onto an existing mineral of the same composition does not give rise to a new mineral. 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, whereas Wetzel ( 1955) has noted cinnabar (HgS) and barium minerals. Authigenic reactions are here interpreted as all those reactions leading to new mineral formation. They include: (a) metasomatism (ionic replacement); (b) ion exchange and adsorption (addition of new, and/or exchange of different ions, especially base exchange); (c) replacement (complete molecular substitution); (d) hydration and dehydration (addition to or release of H,O from the molecule or from solid solution); (e) oxidation and reduction (addition or release of 0, or hydroxyl ions); ( f ) polymerization and depolymerization (construction and breakup of hydrocarbon chains). Common authigenic minerals are indicated in Table 2-1. Inasmuch as any fortuitous mixture of minerals and ionic solutions, such as exists in the fresh sediment, is not 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.

BOUNDARY LIMITATIONS

There are limiting factors for diagenesis, but there are passage zones into metamorphism that vary according to the primary composition of the

23 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. Geochemical parameters In fresh, particulate sediment one deals with two components, namely, the solid sediment particles and the enclosing liquid. The latter is always present initially in marine deposits, but sometimes is 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); (b) oxidation-reduction potential (Eh); (c) ionic adsorption phenomena. A world-wide study by Baas Becking et 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 (Fig. 2-1). 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 CO, is the principal by-product of organic oxidation and is also the principal raw material of plant and bacterial photosynthesis, it is to be expected that it plays an all-pervading role. Thus, C 0 2 reactivity on the earth’s crust will be related to the rate of organic metabolism. 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. An advanced theoretical approach to the geochemistry of the early phase of diagenesis was published by Berner (1980). He derived a general “diagenetic equation”, with quantitative consideration of diffusion, compaction, pore-water flow, burial advection, bioturbation, adsorption, and radioactive decay. Special attention was given to decomposition of organic matter, cementation, and replacement. It is, therefore, unnecessary in this chapter to more than touch on a few environmental aspects. A good basic reference to solutions and equilibria is that of Garrels and Christ (1965). ( a ) Hydrogen ion potential. CO, dissolves freely in H,O, 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 upon saturation of water with CO, it rises to (pH = 5). CO, is thus involved with carbonic acid and the bicarbonate ion in sea

24 1,000

FI, Sp Oxidot ion

BOO

600

400

200

Eh mV 0

-200

-400

-600

\

I

I ACID

-800

I

0

pH

I

4

\

I

I

6

8

4*

%.\‘

I1 ALKALINE I

10

I

12

Fig. 2-1. Catenary diagram illustrating 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.) It is called a “catenary” diagram because the boundary is like a linked chain (“catena”).

water in equilibrium conditions as follows:

H 2 0 + C 0 2 = H , C 0 , = H C O ~ + H + = 2 H++CO;The corresponding equilibrium pH values in pure water (from left to.right) are 5, 6.3, and 10.3. In a closed system, the CO, reactions tend to move to the right as temperatures increase. Sea water in contact with COz will have a pH of 8.2-8.3. According to Sillen (1961), sea water is constantly buffered by the presence of clay-mineral particles in the ocean. Regardless of latitude and temperature, the oceans are thus always alkaline, whereas fresh waters tend to be acid.

25 Rainwater is normally saturated with CO, and, thus, has a pH of 5.5 or less, but it is usually buffered quite rapidly by soil carbonates except in acid situations. Rainwater p H is often much less (3-4) in the industrially polluted and volcanic areas, as well as downwind, for distances of up to 1000 km, where the “acid rain” is frequently toxic to lake and river life. Under organic control, however, very considerable modifications may be introduced, and the pH range in sediments may extend from about 2 to 12 (Baas Becking et al., 1960). It is also important to consider the ionic potentiul of the various components. The ionic potential is defined as the ratio of the ionic charge Z to the ionic radius r . According to Wickman (1944), these potentials fall into three categories : ( 1 ) Z / r = 0-3, soluble cations, which stay in true ionic solution even up to a very high pH, e.g., Na’, K + , M g 2 + , F e z + , M n 2 + , C a z t , S r 2 + , Ba2+; their hydroxides have ionic bonds and are, therefore, soluble. (2) Z / r = 3-12, elements of hydrolysates, which are precipitated by hydrolysis, e.g., A13+, Fe’+, S i 4 + , M n 4 + , etc.; these have hydroxyl bonds which makes them susceptible to hydrolyzation. ( 3 ) Z / r = > 12, soluble, complex ions, which form “complexes”, i.e., complex anions containing oxygen, and as a rule give true ionic solutions, e.g., B’+ , C 4 + , N 5 + , P 5 + , S 6 + ,Mn” ; they have hydrogen bonds, which also. like group ( l ) , lead to soluble compounds. These relationships have been clearly explained for geologists by Barth (1962). To quote from Barth (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; i t 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 of the earth’s crust. After oxygen, which occurs mostly in combination, there are only seven quantitatively important elements: Si4+, A13+, Fe2+ or F e 3 + , C a 2 + , Na’ , K + , and M g 2 + , in that order (see Table 2-11). Inasmuch as natural waters are everywhere subjected to organic inter-

26 ference (largely reflected by the concentration of CO,), one may observe in cool. humid climates with acid soils (high CO, content and 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 H,O and COz), the pH is high (8-9), so that calcium is precipitated, and results in the well-known lime “caliche” or Ycalcrete” 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 alteration 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 (Runnels, 1969; Fairbridge, 1975; Fairbridge and Bourgeois, 1978). Extreme swings, from one absolute p H boundary to the other are to be

TABLE 2-11 Common elements in the earth’s crust and their solubility at 25°C with respect to pure H,O (pH=7), or modified by solution of CO, ( p H = 5 ) or C0:- ion ( p H = 9 ) Element

Crustal abundance At p H 5 (at 25°C) (moles/l) (parts/1000, or g/kg)

Si Al Fe3+ Ca Na

277

K Mg

81 50

36

281

2 . lop3* 1.4. lo-’ 6 . lop9 very soluble

At p H 7 (moles/l)

At pH 9 (moles/l)

4.5. 1.4. l o - ”

6 . lo-’

6. at p H 12

1.4, 6 . lo-*’

3.2. lo-’

pH at which hydroxide begins to precipitate

**

2 4

2.5 12

very soluble (the hydroxide will not precipitate)

26

21

very soluble

1.1.10-‘

10.5

* Approximation, based on curve by Correns (1949, p. 210). Somewhat different according to Siever (1959). SiO, is in the form of the oxide, not hydroxide.

** This value may be too low. Due to the amphoteric nature of AI(OH),, i t begins to dissolve i n strongly alkaline solutions, forming the complex AI(OH,)+OH- =AI(OH), . At pH = 10. the solubility of AI(OH), increases sharply. Equally rapid (in the other direction) is the increase in solubility at ca. pH=4.

27 seen under exceptional conditions in some deserts. Normally, the dilution of all solutions by rain water (pure H 2 0 ) 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. ( b ) 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 reciprocal to the pH, but is influenced by certain other factors (Latimer, 1952: Blatt et al., 1980. p. 240). The pH-Eh relationship for natural environments, as established by Baas Becking et al. (1960), presents a boundary like a distorted shield, with small shoulders (or “ears”) in the low pH-high Eh corner and in the high pH-moderate Eh corner (see Fig. 2-1). The whole shield (i.e., all natural environments) fits between two parallel lines, the slope of which corresponds to -0.059 V/pH unit. These two parallel lines are absolute barriers (“fences”), representing the equilibrium limits of water at or near the earth’s surface. The upper diagonal bounds the upper limit ( H 2 0 / 0 , ) where the partial pressure of oxygen is equal to 1 atm, whereas the lower diagonal marks the lower limit ( H 2 0 / H 2 ) , where p H ,= 1. In the lower diagonal. the left-hand end corresponds to pH of 0.0 a n d E h of 0.0. (See also Fig. 2-2) The value of these relationships to an understanding of diagenetic reactions should hardly need emphasizing. Yet as brought out by Garrels (1960. 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 line-operated) 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, 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 it 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 geologist. In short, i t is absolutely basic to sedimentology. Whereas the pH readings are easily obtained and are closely repeatable, the Eh is more difficult to establish in natural media and must usually be regarded only as an approximation. It was through soil studies and bacteriology that pH-Eh relationships were introduced into sedimentology. ZoBell’s study ( 1946) brought to this writer his 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).

28 PH

-

7.0 I

-0.3

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

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, a sediment thickness of 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 p H of about 7, the Eh 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. After chloride, SO:- is the most important

29 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 CO, has long since passed into CaCO,, so that the only gas phase is H,S. 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). The stability fields of Fe,O,, Fe,O,, and FeS, have been illustrated by Garrels (1960, p. 145). Under favourable conditions, pyrrhotite, galena and other metallic sulfides will start to form. Baas Becking et al. (1960) have demonstrated the stability fields of some important naturally occurring non-metallic compounds in laboratory-controlled bacterial studies (Fig. 2-3). ( C ) 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

Fig. 2-3. 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 Baas Becking et al., 1960.)

30 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 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. ( a ) Temperature. The mean temperature at the earth’s surface through most of geological time for which there are identifiable indicators, say the last 3 . lo9 years (Fairbridge, 1967a), has remained at ca. 20 10°C. Soil temperatures today at depths of 0.1-1.0m fall generally within this range. Seasonal, latitudinal and altitudinal variables increase this range from ca. 60°C down to - 100°C, excluding volcanic phenomena, hot springs, and so forth. Certain algae are adapted to life in hot springs near 100°C, 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 early diagenesis, because in such ranges, crystallization is normally

31 slow and only simple compounds form. The complex, mixed crystals commonly involved at the temperatures of formation of metamorphic and igneous rocks (over 100°C) are encountered only in deep-burial diagenesis (“metagenesis” of Soviet geologists; “anadiagenesis” of the writer). Solubility and pH both change with rising temperature (see Fig. 2-4). 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 10’ years, temperatures below the surface are found to rise by ca. 0.5 to 1.5”C/100m, 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: (1) Abnormal concentrations of igneous or radioactive heat, as near volcanic vents and major fault lines, notably the celebrated “Mid-Ocean

5 0

1

20

1

1

1

40 60

80

I

5000

1

1

1

1

1

1

1

1

100 120 140 160 180 200 220 24(

I

I

I

I 1

IOQOO 15,000 25QOO 20,000 30,000

Feet

Fig. 2-4. Variation in the dissociation constant of water as a function of temperature and depth in a geosyncline. (After Blatt et al., 1972, fig. 7-1). One should note that neutrality varies with temperature, so that at an average burial depth of l0,000m, neutrality (equal concentration of H + and OH- ions) is indicated by pH equal to 6 .

32 Rift” along which the heat flow may rise to 8 . cal. cm-2 sec-’, in contrast with an average of ca. 1 elsewhere. (2) Sedimentary accumulations of minerals, which on oxidizing are exothermic, that is to say, that 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 basin sediments compact. One of the principal minerals involved is either marcasite or pyrite (FeS,) and these are commonly present as vast quantities of finely divided particles 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 a depth of only 300m comes from a pyritic shale-silt-sand sequence that brings the water temperature at the well head to 90°F (32°C). In deep oil wells on the Texas Gulf coast at depths down to 25,000 ft (8000m), the temperature through normal heat flow should be high, but may exceed 150”C mainly through this same exothermic mineral oxidation. ( b ) Pressure. In a gradually accumulating sedimentary basin, there is pro-

gressive increase of load pressure (Lane, 1922). This has sometimes been called load metamorphism or static metamorphism (Grabau, 1913, p. 750), 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 complete remobilization of ions and formation of minerals beyond even the limits of metasomatism. The results of this are seen in some ancient evaporites, in particular the Stassfurt deposits of the Permian in north Germany (Janecke, 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 the temperature of the alteration was probably not over 80°C. ( 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. I t 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.

33 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 were 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 lO7-1OXyears, 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 is involved, however. in a superficial “skin” that may not exceed 3000m, 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 . 10’ years.

DIAGENETIC EVOLUTION

Diagenesis begins at the moment a sedimentary particle comes to rest, for example, on the sea floor. It continues to a point in history when either deep burial and orogenic buckling cause the initiation of metamorphism, or when emergence leads to exposure and the initiation of weathering and erosion. Sever (1959) has called these phases, respectively, early, middle, and late diagenesis. It is an almost Davisian evolutionary cycle of youth, maturity, and old age, except that it may be complicated by rhythmic repetitions and “accidental” alternative courses, introduced by interaction with geotectonic, paleoclimatic and other cycles. For clarity, one may employ with “diagenesis” the classical prefixes “syn-” (together with, or synchronous with the sedimentation process), “ana-” (up, again, or thoroughly lithified), and “epi-” (outer or surficial, i.e., modified by surface phenomena). Thus, the three phases have been named by the writer (Fairbridge, 1967b, p. 32) as follows (Fig. 2-5): (a) syndiagenesis (the sedimentation phase, depositional and early burial);

34

RAINFALL

pH 7

+ + + + + + + + ’ + + +

Fig. 2-5. Idealized profile through a continental margin, showing the sites of contemporary marine sedimentation and the three phases of diagenesis. There is: ( 1 ) diffusion potential during syndiagenesis; (2) upward liquid motion in anadiagenesis (compaction fluids); and (3) downward motion during epidiagenesis.

(b) anadiagenesis (the compaction-maturation phase, with deep burial); and (c) epidiagenesis (the emergent-pre-erosion phase). This is a temporal and spatial classification designed to establish the environments of the mechanical and geochemical processes, i.e., time, place, and process. The three phases are evolutionary and dynamic, passing from one to the next across recognizable boundaries. Within each phase there are distinctive hydrologic systems associated with characteristic, though not unique geochemical systems. Thus, for example, lithification is the most common attribute of anadiagenesis, but in some specific environments (e.g., several types of carbonate sites), lithification is syndiagenetic. Dapples ( 1962) designated three geochemical stages, specifically for sandstones, and adopted for that purpose by several workers (e.g., Selley, 1976), as follows: ( 1) Redoxomorphic (mainly oxidation or reduction reactions; a metabolic control, that is 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; this is important during the lithification of the anadiagenetic phase). ( 3 ) Phyllomorphic (characterized by ion exchange such as is associated with clays and micas, to be observed during all phases).

35 The time factor is subject to great variation. Rapid subsidence of the basin can lead to accelerated transition to anadiagenesis. Slow subsidence may expose the fresh sediment to eustatic fluctuations and thus to epidiagenesis without going through the anadiagenetic stage. Choquette and Pray (1970, p. 219), on a basis of time and porosity types, also recognized three phases: eogenetic stage (time of brief evolution in proximity to sea floor), mesogenetic stage (long time period following burial and isolated from sea floor and preceding emergence), and tefogenetic stuge (time following emergence and influenced by processes leading to an unconformity). Essentially these terms, differently expressed, are synonymous with syn-, ana- and epidiagenetic, within each of which there is a trend towards a steady-state or equilibrium condition. In familiar terms one could simply restate the syn-, ana- and epidiagenetic phases as “mud”, “hard rock” and “weathering” stages. In many cases this would be a true designation, but it would also be misleading in some instances: hard, concretionary segregations may form during early burial stage, leaching often occurs during deep burial, and ha,rd-crust formation frequently results from exposure to weathering processes. Emphasis should include not only time and place, but also process. From a quite different point of view, the diagenetic processes involved in the evolution of sand to sandstone include six stages according to Pettijohn et al. (1973, p.387): five stages correspond to progressive depth of burial, 0-nearly 10,000m, with a pressure increase of about 1 bar/4 m; the sixth stage follows uplift and is equivalent to epidiagenesis (Fig. 2-6). 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 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”. This term, as originally defined by Fersman ( 1922), however,

Fig. 2-6. The stages of diagenesis of sandstones in relation to depth of burial and increase of pressure and temperature. (After Pettijohn et al.. 1972. fig. 10-1, p. 387; courtesy of Springer, New York.) Stages 1 and -7 are syndiagenetic. stages 3 to 5 are anadiagenetic, whereas stage 6 is epidiagenetic.

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): initial stage, controlled by the chemistry of the superjacent water, and early burial stage, controlled by the entrapped, connate water, chemically modified by the bacteria and other subsurface organisms.

( a ) 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 the organic matter in this zone of bioturbation 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, a 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

37 leads to diagenesis of petroleum hydrocarbons, but this is more a function of rapid accumulation than it is of euxinic conditions (Smirnow, 1958; Degens, 1967; Degens and Ross, 1974). 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 the connate water. The result will be a sharp rise in the pco, 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 Wadden Sea, 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 SO:- to H,S and its immediate reaction with the Ca2+ of sea water to form CaSO, a2 H,O. 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 (Orr and Moorhouse, 1933). In addition to C 0 2 , it is probable that humic and tannic acids are also liberated. Coral reefs of CaCO, are pocked by giant mud-filled pot-holes up to 5 m in diameter, wherever mangrove trees have been situated. ( b ) 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 Desuljovibrio desuljuricuns, liberate H,S. CaCO, precipitates freely at a pH of 8.5 and in this Eh range FeS is the stable iron compound (afterwards becoming FeS,; see section Oxidation and Reduction). Somewhat less commonly, siderite (FeCO,) is

38 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‘s 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 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 (Hunt, 1979). The transgressive phases of major eustatic fluctuations are particularly significant (Tissot, 1979; Arthur and Schlanger, 1979). After hydrogen, the most important active element in the early burial stage is sulfur. This element is present in sea water as the anion SO: - , which represents 7.68% of the total dissolved constituents of the ocean, and is the most important after Na+ and C 1 ~ “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 t 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 SO,‘- to S’- , the isotopes are fractionated, so that the sulfide ion is enriched in j’S, the more energetic isotope. The sulfur isotope ratio is, therefore, a valuable indicator of passage through the early burial stage (Holland, 1973; Holser, 1977). Thus, H,S and related authigenic minerals formed then show an 32S/34S ratio of 22.1 to 22.7 (with an everage of 22.49 for all sedimentary sulfides), in contrast to a constant 21.76 0.02 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. Baas Becking et al. (1960) have shown experimentally that, under certain conditions, marine bacteria can synthesize not only pyrite, but also the common ore sulfides. Deans ( 1950) reported

39 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). The principal sources of these thermal springs are sea-floor spreading centers, such as demonstrated near the Galapagos Islands in the Pacific (Spiess et al., 1980), or in the currently inactive trench of the Red Sea (Degens and Ross, 1969). 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 decreases sharply with depth of burial, the level of fixed ammonium remains rather constant. Thus, the C/N ratio can be 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 would occur at and above the sediment’s surface (Degens and Ross, 1974). If the oxidizing zone had been present, however, the acidizing experience of the sediments is the one that would 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, the soluble carbonates and other minerals are destroyed or modified under low pH, 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

40 material, regardless of grain size, is CaCO,) no amount of bacterial CO, production will cause the total solution of their substrate. The response of lime sands p proto-calcarenites") will be rather similar to that of quartz sands, but lime muds p 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 dissolved silica (Bien et al., 1958; Heath, 1974), as desert dust, or by solution from volcanic ash. It is biogenically concentrated as opaline silica from diatoms, radiolaria, sponge spicules, holothuria, and alcyonaria (Riedel, 1959). Penecon temporaneous resolution may occur, but is prevented in rapidly accumulating Globigerina oozes. The silica appears to accumulate in small globules of gel that dehydrate very slowly. Migration of silica often seems to occur while the sediments are still quite soft. Indeed much movement may be expected while sediments are permeable 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 SiO, stabilization, whereas under the high pH of deeper levels the reverse is true (see recent papers collected in Van der Lingen, 1977). 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. Additional discontinuities are also provided by “hardgrounds” which are often very widespread (Bathurst, 1975; Kennedy and Garrison, 1975). 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 determined 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

41 bacterial activity, which may range from a depth of ca. 1-100 m. In terms of organic metabolism, syndiagenesis may be taken to last as long as the food holds out. In terms of absolute duration, this may be for 1000-100,000 years, but considerably more research is needed on this aspect. It has been claimed that viable bacteria can be traced back to Carboniferous coal seams (nearly 3 lo6 years old), but possibilities of contamination are so great that it is very difficult to prove. +

A nadiagenesis

Anadiagenesis corresponds to deep burial, involving the compaction and maturation phase of diagenesis, during which the particulate sediment grains (or chemical ions) become once again (Greek: m a - ) 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 below). 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 indeed only since the introduction of the term has the importance of this phase of diagenesis been appreciated (see Lane, 1927; White, 1957; and Chave, 1960). The effect and weight of sediment-loading was also studied by Lane (1922). Earlier, Van Hise (1904), 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 (Pettijohn et al., 1975; Bathurst, 1975; Blatt et al., 1980). 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, begmning even at the surface. On mature plate boundaries, with a loading of up to = Z O O bars, i.e., in

42

orthogeosynclines, it is possible that the progressive downwarp and filling of the trough leads to burial in excess of 10,00Om, which is the approximate depth at which the geoisothermal level exceeds the normal operative limits of diagenesis (about 100°C). 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 may or may not be followed by vertical readjustment if subduction and excessive heating leads to granitization of the roots zone. Uplift then results, and further gravitational slides occur, this time externally 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 “Jura-type” displacement and “Alpino-type” nappe slides towards the exterior. These generally escape metamorphism, and should be associated only with lower temperature fields and more 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. On the other hand, in parageosynclines (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. 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 and volcanism. 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 surface 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 diaclastic revival (from “diaclase”, the classical term for joint or fracture).

43 Geochemically, anadiagenesis is often a de-watering stage. 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 subsidence to general epeirogenic uplifts 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” (Millot, 1970), inasmuch as it is essentially the clays that offer ionic adsorption potentials (see further discussion below). The deep waters may become progressively but irregularly more saline (Chave, 1960). Valuable compression experiments have been reported by Chilingarian et al. ( 1973).Whereas 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 upward 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. Sandstone porosity may also be progressively reduced by upward transport and deposition of clay minerals (Almon et al., 1976; Blatt et al., 1980). The depth-time limits of anadiagenesis may be broadly defined as extending from the lower limit of syndiagenesis (ca. 1-100 m) to ca. 10,000 m, and from say 1000-100,000 years to somewhere between lo7 and 10’ 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 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 104-10h years. Epidiagenesis Epidiagenesis is the emergent or post-diastrophic phase of diagenesis. The analogous stage in mineral genesis is associated with the terms epigene or

44 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 shortcircuit 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. Inasmuch as meteoric waters are normally saturated in oxygen and CO,, 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. Epidiagenesis of the old, stable cratons is a pervasive process in soil development and landscape control, inasmuch as it is constantly interrupted by transgressions and then reintroduced. A global eustatic rhythm of rise and fall of sea level tends to be superimposed on regional epeirogenic crustal warping. The alternation of stabilization and reactivation is the essence of the theory of biorhexistasy (Erhart, 1956; Termier and Termier, 1963). “Biostasy” is the condition marked by low relief, dense forest cover, and chemical leaching, with export of silica and calcium. “Rhexistasy” is the reactivation stage, when stream profiles are lowered and landscapes dissected, generating a large supply of clastics, while loss of vegetation leads to desiccation of plainlands. In tropical regions the clastics often include large supplies of hematite-rich debris from the mechanical breakup of laterites. The primary epidiagenetic effect of biostasy is leaching and that of rhexistasy is capillary concentration and deep cementation (Fig. 2-7). Ground water is divisible into two categories: (a) the uadose water that is associated with the well-aerated zone near the surface, and characterized by episodic saturation and seasonal capillarity; (b) the phreatic water that lies below the water table in the zone of more or less permanent saturation. The upper part of the vadose zone is usually marked by leaching. At the boundary between the two is the water table, and in the capillary fringe above it, especially in seasonally arid climates, there tends to be reprecipitation of carbonates (Bathurst, 1975; Milliman et a]., 1974), silica (Flach et al.,

45 Ca C O j Synqiagenesis lnselberg

Dense forest

/

/

Chemical

I

,

a

Warm, clear, wide

A. BlOSTATlC STATE 1, LOW land masses 2. High sea level (thalassocratic 1 3. Humid climate (high oceanicity) 4 . Rich vegetation 5. Chemical weathering

limited

€3.

RHEXISTATIC STATE 1. Emergent landscape, higher relief 2. Lowered base level (eustatic low) 3 . A r i d t o semiarid climate 4. Reduced vegetation, savanna 5. Mechanical erosion

reduced Shelf sea

Eustatic

leaching of fresh sediments)

RWF 1981

Fig. 2-7. Epidiagenetic effects of biostasy and rhexistasy.

1969; Stephens, 1971; Langford-Smith, 1978) and iron oxides (Turner, 1980; Allen, 1980), which become the “hard pans” of soil scientists and the “duricrusts” of geomorphologists. Rapid penetration of the bedrock by ground water is largely restricted to sedimentary rocks, although it also enters joints and fault zones in igneous rocks. In this way, given long periods of geologic time, weathering gradually penetrates into the most compact rocks, and even into mineral cleavage planes. One of the more spectacular processes of epidiagenesis is the rehydration of anhydrite to gypsum that produces enterolithic expansion with convolutions within thin-bedded layers, or surface geomorphologic effects, such as mounds, rolls, tepees, and chaotic blockfields, in semi-arid country. (Gypsum-anhydrite relationships are discussed further under “Hydration-dehydration”, below.) The rapid oxidation of such mineral compounds as pyrite (FeS,) will, as discussed earlier, tend to raise the temperature of deep ground waters far beyond the limits normally expected from the geothermal gradient, and further re-solution of certain salts will be facilitated. This general geochemical revival permits renewed cementation, and mineralization of fault and joint zones, thus effectively contributing to the final lithification of the rocks. An important aspect of the epidiagenetic phase is the state of permeuhility

46 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. Van Hise (1904) subdivided his “katamorphic zone” into two: a “belt of weathering”, and a “belt of cementation” (i.e., the anadiagenetic phase). The “weathering belt” was taken to include all of the zone affected by circulating water, which can be classed as epidiagenetic. Certainly there is a link as Blackwelder (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’5000 m or more, is a special @on-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 over 8000m. 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 10’ to lo9 years.

47 AUTHIGENESIS AND SOME DIAGENETIC PROCESSES

Other chapters in this book and its companion volumes deal more completely with almost all diagenetic processes, 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. 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, marked by authigenic mineralization; (b) late diagenesis, marked by epigenic mineralization. The suggested correspondence has not been widely accepted, however. Geochemical reactions The chemical reactions involved in authigenesis 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 ( 1 958) 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: hydration-dehydration, hydrolysis-dehydrolysis, ion adsorption, cation or base exchange, oxidation-reduction, carbonatization. ( a ) 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: CaSO,

+ 2 H 2 0+ CaSO, - 2 H 2 0

Hydration also occurs during hydrolysis, oxidation, and carbonatization (see, for example, Berner, 1971; Blatt et al., 1980; Garrels and Christ, 1965).

48 Hydration is most common during syn- and epi-diagenesis, whereas dehydration is prevalent during anadiagenesis. ( h ) 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 that yields hydrogen ions, which on reacting with H,O give H , O + , the hydronium ion. Most silicate minerals are susceptible to hydrolysis. Thus olivine hydrolyzes to serpentine:

2 Mg,SiO,

+ 3 H,O * 3 MgO ~2 SiO, . 2 H,O + Mg(OH),

Relative susceptibility is useful in identification of unknown minerals and can be demonstrated experimentally by means of the “abrasion pH” (Stevens and Carron, 1948; Keller et al., 1953). The mineral is ground up with distilled water and the p H is measured. Feldspars give a pH of about 10 and wollastonite gives a p H of 11. It is found that under these conditions metal cations are liberated, e.g., C a 2 + , Mg2+, N a + , 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 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 (Garrels and Mackenzie, 1974; Holland, 1972). Most detrital silicates (notably feldspars, micas, and clays) carried into the ocean are in an incomplete state of hydrolization. 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 hydrolization, or to the left, that is silicates become dehydrolyzed. Upon burial, the p H 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 hydrolization most of the silicate weathering products also have the capacity of base-exchange and adsorption. ( c ) Ion adsorption. 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: C u 2 + ,A13+, Zn2+, M g 2 + , C a 2 + , K + , and N a + ; and by the anions: S 2 - , C1- , and SO:-. 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

49 ion, unless the latter is present at a higher concentration (see Chilingarian and Vorabutr, 1981). After the adsorption of H + or O H - , 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 (Eitel, 1954, p. 458). ( d ) 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; Russell, 1970). An acid solution (rich in CO, 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 quite some time ago (Marshall and Upchurch, 1933), whereas attempts to carry out quantitative measurements in fresh marine sediments are only beginning. It is deduced that, for example, the 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. ( e ) 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 (see, for example, Friedman and Sanders, 1978; Berner, 1971 ; Krauskopf, 1967). Oxidation implies the loss of electrons. The substances that gain electrons 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 impossible to separate the concept of oxidation from hydrolysis. Furthermore oxidation is also involved when metallic iron is “oxidized” to FeS and then to FeS,. 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.

( f ) Carbonatization. The carbonate ion COj- (or the bicarbonate ion HCO, ) is often known to replace silicates during weathering, particularly

50 with Ca2+ and M g 2 + .Both being always present in the ocean, they are likely to be organically reprecipitated as CaCO,, “magnesian calcite”, or recrystallized later as dolomite. Decarbonatization normally occurs in the presence of solutions of CO, and then of HCO,. 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 be traced in the early syndiagenetic alteration of brucite Mg(OH), to MgCO, in dolomitization, which is discussed under “Metasomatism”. The concentration of the weak acid H,CO, in water that is in equilibrium with the atmosphere is independent of pH, but is always low. During syndiagenesis, however, the interstitial waters with pH around 8 develop far higher concentrations of carbonic acid than they do in rain water or in surface sea water. This is mainly due to the CO, of biogenic origin. 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 the nature of the reactions.

HALMYROLYSIS

Sediment particles reach the ocean: ( 1 ) in estuaries and deltas as rivertransported material (Fairbridge, 1980b); (2) from coasts subject to littoral erosion; (3) anywhere at the sea surface when the particles arrive wind-borne, as fall-out (pyroclastic ash) from volcanic eruptions, or by extraterrestrial transport. Reactions may begin immediately while the particles settle. Some may continue, as the sediments slowly drift along the bottom (Ewing’s “nepheloid layer”: see Jacobs, 1978). By virture of their fine grain size and exchangeable cations, the clays are particularly susceptible to electrolytic reactions. The alteration brought about is sometimes referred to as “predepositional diagenesis” or “submarine weathering”, but it is not really included in the original definition of diagenesis, nor is the analogy with subaerial weathering very precise. It seems that the best term for the process is halmyrolysis, an expression proposed by Hummel (1922), from the Greek roots hali- (sea) and myros (unguent), erroneously spelled “halmyrolosis” by Pettijohn (1957, pp. 649,662) in one place 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 is 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,

51 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 basins, submerged plateaus 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. Inasmuch as sea water takes up elements (e.g., Ca,Si, and P) at the expense of the sediment, the term “submarine weathering” is sometimes quite appropriate (Correns, 1950), but the whole picture is more complicated. Mackenzie and Garrels ( 1966) suggested a sort of “reverse-weathering reaction”, progressively, through geologic time, removing silica and bicarbonate from sea water: (amorphous aluminum silicate) HCO, H4Si04 cations + cation aluminum silicate (mainly clay) CO, H,O. Most marine silica, however, is removed biogenically (mainly by silica-shelled diatoms). Teodorovich (1954; see review by Chilingar, 1955; and discussion by Packham and Crook, 1960) recognized no less than thirteen diagenetic facies that are subject to halmyrolysis and/or processes of early diagenesis. They have been defined in terms of pH-Eh and mineral coinponents. Six fundamental geochemical environments are noted: (1) oxidizing (ferric oxides and hydroxides); (2) weakly oxidizing, Eh = 0 well below sediment surface; (3) neutral (iron chlorites), Eh = 0 slightly below sediment surface; (4) weakly reducing, Eh = 0 precisely at surface; ( 5 ) reducing (carbonates; scattered FeS and FeS,), Eh = 0 slightly above interface; (6) strongly reducing (sulfide zone proper), E h = O well above sediment interface and approaching sea surface.

+

+

+

+

+

Gfuuconite Probably the most characteristic of all halmyrolytic phenomena is that of glauconitization (Porrenga, 1967). The mineral gfuuconite (a hydrolyzate 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) believed that glauconitization was favored by cooler waters; this is not correct, for the most widespread occurrence of glauconite today is over the continental shelves north of Australia, as originally described by Murray and Renard (1891) and later extended b ~ Fairbridge r (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 inshore near river mouths. Although Murray and Renard found glauconites down to depths of 4000m in the ocean, it seems likely that they were transported there with

52 other sand-size particles, and that the optimum zone of glauconitization is from wave-base (say 15m) 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 at a p H of around 8. Reducing conditions can be established on the open shelf in a microenvironment, such as the rotting interior of a molluscan shell, and certain foraminifera, or associated with fecal pellets. Organic participation, even in micro-environments seems to be the rule (Ojakangas and Keller, 1964), in the case of existence of 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 10,000 years, so that the complete authigenesis of a particle 2 m m in diameter may require something of the order of 100- 1000 years. Although one lacks experimental evidence, the glauconite reaction probably lies in the activity field of the heterotrophic anaerobic bacteria ( p H = 7 - 8 and E h = O to -100 mV), and within a temperature range of 25-5°C. There is apparently a reduction in the amounts of glauconite in the pre-Mesozoic oceans (Conway, 1945), which may be related to the postulated higher pco, of the Precambrian and Paleozoic times. Tugarinov and Vinogradov-( 1961) reported Precambrian glauconites back to 1.5 . lo9 years. The older the glauconite, the lower is its Fe3+/Fe2+ ratio (Smulikowski, 1954). 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. Phillipsite and other zeolites Another product of halmyrolysis is the marine zeolite, phillipsite (its formula is sometimes given as (K,Ca)AI,Si,O,, - 4 H,O), which was first discovered in deep-sea environments by the ‘Challenger’ Expedition (Murray and Renard, 1891). These elegant little crystals, often twinned, make up an appreciable part of the “red-clay’’ sediments in parts of the Pacific, where they are not masked or inhibited by rapid deposition of other materials. In

53 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 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). Another more restricted deep-sea zeolite is clinoptilolite, which is most common in the Cretaceous-Eocene sediments (Venkatarathnam and Biscaye, 1973). Contemporary zeolitization following deep burial is hardly distinguishable from the first stage of metamorphism, the zeolite fucies (Fyfe et al., 1958; Coombs et al., 1959; Turner and Verhoogen, 1960, p. 532). This phenomenon may occur at burial depths down to about 8000 m (say 2000 bars pressure and 100-200°C), provided that there is suitably metastable parent material, in this case, fresh, rapidly accumulated volcanic ash. Tuffs, deposited from nuees 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, Na,)Al,Si,O,, . 5 H,O] and laumontite (CaAl,Si,O,, . 4 H,O), 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). Because of its volcanogenic derivation, zeolitization in the geologic record is an especially valuable indicator of plate tectonic activity (Sever, 1979). Clay minerals

A much-studied aspect of halmyrolysis is the authigenesis of clay minerals (reviewed, for example, by Drever, 1976). 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 the source region (“heritage” of Millot, 1953, 1970). 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.

54 Russell (1970) showed that some 11-47% of K-ions delivered in solution by rivers to the ocean is removed by clay minerals during a rapid halmyrolysis. This capturing process seems to apply also for Na, but is unimportant for Ca and Mg. The Mg-cation is only taken up in an anaerobic (syndiagenetic) setting, replacing Fe that later forms pyrite (Drever, 1971). Experimental work over a 6-month period showed that the most reactive species were the montmorillonitic clays, whereas kaolinite and illite appeared to be stable within this time frame (Whitehouse and McCarter, 1958). A “concept of dualism is essential to the understanding of clay genesis”, according to Weaver (1958); he refers here to the inherited characteristics on 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. Upon burial in a largely clayey facies, there is a natural tendency towards 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 l000m (Milne and 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 have been made by Biscaye (1965) and by Griffin et al. (1968). The broad pattern that is emerging seems to suggest that there are definite geographical, latitudinal provinces, which are dependent upon these two factors: (1) 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. Whereas a very large number of clay minerals are known, only three principal types will be discussed here: illite, montmorillonite, and kaolinite. 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% A1,0,). Crystallographically, they appear to have essentially the same lattice structure as muscovite, but are degraded with respect to K:

55 where y is around 1- 1.5. An ideal end-member could be:

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 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 (Drever, 1977). During syndiagenesis it may emerge from the halmyrolysis of montmorillonite or kaolinite (see below). Particularly K is adsorbed in these reactions, which are favored by the presence of C a 2 + ,M g 2 + , 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, there is considerable mixing with “chlorite”, tb 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, notably in the tropical latitudes, offshore from the deeply weathered cratons (Biscaye, 1965; Porrenga, 1967). 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 (Griffen et al., 1968). 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: montmorillonite: kaolinite ratio is approximately 1 : 1 : 1 (Biscaye, 1965), but 8 : 1 : 1 in the temperate latitudes. K/Ar isotopic age-dating of these North Atlantic illites give mean ages in the 300-400 m.y. brackets (Hurley et 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 “Paleozoic” 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 ages of

56 deep-sea illites decrease as one goes down the D.S.D.P. cores. Independent confirmation of the provenance rock age, transported dusts, and deep-sea sediments is provided by Rb/Sr ratio and Sr-isotope data (Biscaye and Dasch, 1971; Biscaye et al., 1974). It is noteworthy that the K values in clays rise on going back into the Precambrian (Nanz, 1953). Conway (1945) suggested earlier that there was a peak of K extraction in the oceans in Late Precambrian time. Whereas 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. Montmorillonite formula is generally given as A1,Si,O,o. (OH),nH,O. The structural formula of montmorillonite group of minerals is best expressed as:

where R3+-Al, F e 3 + ,Cr; R2+- F e Z f , Mn2+, Mg, Zn, Cu. In this “expanded” or three-layer lattice clay group, some of the aluminium may be replaced by Mg (suponire) or by Fe (nontronite). A general term for the montmorillonitic clays 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 Naf < K f < Mg2+< Ca2+< H + , i.e., that potassium replaces sodium, etc. Its cation exchange capacity is quoted as 60-150 mequiv./lOOg, as compared with 3-15 for kaolinite and 10-40 for illite. Rapid diagenetic reactions in the halmyrolytic stage are thus to be expected, even before the clay particle reaches the bottom; however, when buried and cut off from acid solutions, montmorillonite may even survive from Paleozoic times (although admittedly rather rarely). Burial in excess of 3000m clearly adds heat and pressure in addition to the chemical energy so that the end-product is normally illite (Keller, 1964). Montmorillonite seems to be the product of simultaneous weathering of feldspars and ferromagnesian minerals (Ross, 1943) from mafic igneous rocks, and especially fresh pyroclastics (ash), in waterlogged alkaline soils or newly-accumulated ash deposits of moderate pH but low Eh (poorly oxygenated). This is a feature of intermediate and “Mediterranean” latitudes, and formation of montmorillonite occurs particularly in semi-arid lakes (e.g., North Africa), swamps and littoral lagoons (e.g., Texas Gulf coast), as well as in volcanic regions (with alkali-pyroclastics). Experimental procedures have shown that in addition to pH, the flow rate of the interstitial water, the particle size, and the temperature are the key parameters in

57 such clay formation and decomposition (Correns, 1963). Weathering of freshly extruded basalts on the sea floor near spreading centers provides a K-rich smectite to the regional sedimentation (Hart, 1973). As regards distribution, montmorillonite is widespread in the Pacific (Griffen et al., 1968), no doubt because of the extensive occurrence of volcanic ash. In the tropical Atlantic it shares a 1 : 1 : 1 relationship with illite and kaolinite, as noted above, and may be related in part to wind-borne supplies from the semi-arid regions. It is also common (even dominant sometimes) in the Gulf of Mexico. In the North Pacific and North Atlantic, it seems to be related to the high-latitude volcanoes. The role of montmorillonite in the North Atlantic is very much masked by the high content of primary (Paleozoic) illite there. The absence of illite from the primary weathering cycle in most of the tropical soils, however, suggests perhaps that its presence in tropical oceans is the result of the breakdown of montmorillonite and kaolinite (see discussion by Van Andel and Postma, 1954). The fact that montmorillonite is widespread in the tropical Pacific without reacting rapidly with 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 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). Kuolinite is the simple two-layer lattice clay, characteristic of the so-called “high-alumina clays” (20-40% A1203). It does not appreciably expand upon increasing water content, and does not exchange iron or magnesium. The typical group formula is: A1,0, . 2Si0,. 2 H 2 0 , sometimes given as: A1,Si ,05(OH), or as A1,Si 40,,(OH)8 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.

58 Its formation is especially favored by aluminous igneous rocks such as granite (Ross and Kerr, 1931). In this process, C a 2 + , M g 2 + , and Fe2+ are leached out. In subtropical regions silica and iron rise to the surface by capillarity to form a ferruginous duricrust, whereas the Ca2+ and Mg2+ tend to pass into the ground water and drainage system and then to the sea. Kaolinite itself tends to be picked up in the arid lands by wind or subject to fluvial erosion only as a result of tectonic and eustatic oscillations of base level. Under such conditions it is transported to the ocean (as during the Quaternary). In the tropics it tends to dominate the interglacial clastics, being replaced in part by feldspar in the glacials (Damuth and Fairbridge, 1970). In the ocean today, kaolinite is, therefore, mainly a transported, nonauthigenic sediment, and is most widespread in the tropical Atlantic where the principal supply is eolian (Biscaye, 1965), though partly fluvial (fed by the Amazon, Orinoco, Congo, and Niger), according to Correns and Von Engelhardt (1938), Millot (1953), and Heezen et al. (1960). It is also widespread in the tropical Pacific (Griffen et al., 1968) and the Indian Ocean (Venkatarathnam et al., 1976). 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 waters are fresh (Shelton, 1964). 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 longterm geochemical changes in sea water, however, should not be forgotten (see Conway, 1945). Millot (1 953) 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 calcitedolomite-chlorite-kaolinite-quartz, 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 H 2 S lowers the Eh potential. To close this section, one may conclude that the limits of clay diagenesis are still rather poorly defined, especially the lower boundary of halmyrolysis.

59 Oxidation and reduction-processes and environments Inasmuch as normal marine basins are in constant circulation, oxygen saturation is 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 “sill depth” (Goldhaber, in: Fairbridge and Bourgeois, 1978, p. 297; Degens and Ross, 1974). In well-ventilated basins, the oxidation of sedimentary particles at the sediment-water interface 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 the red clay deposition of Murray and Renard (189l). 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 MnO, is a function of the time exposed at the sediinent-water contact. Divalent Mn2+ present in sea water is apparently oxidized to the tetravalent species Mn4+ (Goldberg and Arrhenius, 1958). Large boulders do not contain a coating of MnO, 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, but the present writer suggests about 5000- 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 MnO, accumulated is only 2 or 3mm, 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; Allen, 1980). 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 2000m, 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 SiO,. Such times (biostatic) are marked in the marine realm by limestones and cherts (Erhart, 1973a). A cyclic lowering of the water table which takes place eustatically during periods of cold climate and tectonically lowered marine basins (“bathygenesis” of Termier and Termier, 1963) leads to a mechanical break-up of old soils (rhexistasy), to stream dissection, and to transport of red, iron-rich

60 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 develop reducing environments, the red colors often persist, although green layers and mottled and bleached patches bear witness to reduction by H,S (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 H,S poisoning, this amount is very large. Inasmuch as open-sea floors are normally inhabited by both epi- and 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: ( 1 ) the size and vigor of the benthonic population and (2) the rat,e of deposition. As Twenhofel (1942) pointed out, a slow sedimentary rate permits thorough scavenging. He stated (p. 105): “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 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 deposition 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: (1) continental shelf and slope regions where the sedimentation and subsidence rates are low, and (2) abyssal plains and rises far removed from continental sediment sources. In contrast, shelf regions near deltas or other sediment source, and particularly 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 reducfion. In estuaries, Baas Becking and Moore (1959), have found that the aoeruge organic content is 12%, and may exceed 25%. In the same samples the iron, initially in the form of FeO(0H) is found to become completely reduced and anywhere that the organic content of the sediment is above 2% there is an excess of H,S.

61 The principal minerals so formed are the common ferrous sulfides. marcasite and pyrite, FeS,. These are produced by reduction of iron oxides and the various hydrates (first to FeS and later to FeS2) by the action of H,S liberated by sulfate-reducing 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 (Allen et al., 1912; Newhouse, 1927; Tarr, 1927). According to Berner (1970. 1971), the process occurs in steps, through the monosulfide. 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 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 p H 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 after saturation with H,S ( a weak acid), the p H 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 pyrirized 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 ( Brongersma-Sanders, 1957). Reducing environments are mainly syndiagenetic and also anadiagenetic. evidently produced during or soon after deposition (in the time range of about 1-1000 years), with Eh ranging from 0 to -400 mV and pH either slightly above or below 7, as appropriate either for marcasite or pyrite formation. Exceptionally, the iron sulfides are also formed at considerable depths (3000-5000 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/34Sisotope ratio (see introductory notes, in the section on Syndiagenesis). The similar isotope ratios in recent sulfides, and in a Precambrian

62 shale more than 1.8 . lo9 years old (Ault and 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 Holland (1973), however, that rapid fluctuations in the isotopic composition of marine gypsum and anhydrite disclose comparable variations in the sea water itself within time frames of lo7 yr or less, as dictated by global paleogeographic changes. 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 rnelnikovite. 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 sulfate-reducing bacteria ( Desuljovibrio desulfuricans). The H,S gas is then produced, which on oxidizing (assisted by thiobacteria) gives rise to enormous native sulfur deposits. The biologic nature of this reduction process is proven by the low 32S/34Sratio of the sulfur with respect to the sulfate (Feely and Kulp, 1957). Furthermore, the 12C/13Cisotope ratios in the calcite of the caprock are typical of petroleum and not of marine limestones, suggesting that the C0:- was derived likewise from the hydrocarbons. Under certain conditions of oxygenation, generally epidiagenetic, and under conditions of heavy tropical rainfall (high CO, 0, intake), there may even be a liberation of H,SO, from pyrite or marcasite. Some mine waters actually have a negative pH, which can be explained by the following reaction:

+

2 FeS,

+ 2 H,O + 7 0,

+

2 FeSO,

+ 2 H2S04

It may, however, be better represented (Baas Becking, 1959, p.61) by the equation: FeS,

+ 8 H,O

-+

FeSO,

+ HSOT + 15 H + + 14e-

Inasmuch as FeSO, is stable only in an anaerobic environment, it is likely to hydrate to melanterite, FeSO, . 7H,O, and eventually perhaps to

63 coquimbite, Fe,(SO,), 9H20, in the epidiagenetic stage. In the course of epidiagenesis, which is likely to be protracted and frequently episodic, 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 documented by R.E. Liesegang in 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-processes and environments

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 rehydrolization or revived hydration may be expected.

( a ) Gypsum-anhydrite. The most prominent reactions involve gypsum, CaSO, * 2 H,O, and anhydrite, CaSO, (Berner, 1971). The hydrated sulfate is produced as a primary evaporite deposit as soon as ordinary sea water is concentrated to somewhat less than 50% of its original volume, as in lagoons along the Texas coast or in the Persian Gulf. It has been stated that when the water temperature exceeds 42"C, anhydrite becomes the stable phase (Conley and 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). I t is also found in certain beachrocks (MacFadyen, 1950). In some deep-sea carbonate-ooze environments, authigenic gypsum is also found at 0.3-4 m below the surface, in water depths of 1000-5000 m (Briskin and Schreiber, 1978). Evidently early syndiagenetic, it is suggested that episodic influxes of C0,-rich Antarctic bottom water during interglacial climatic fluctuations cause carbonate solution, liberating calcium in the near-surface sediment layer. In these pore waters sulfate content is directly proportional to sedimentation rate (Berner, 1980). During glacial conditions,

64 the pack ice reduces the bottom conditions to near-stagnation and carbonate corrosion ceases. This euhedral gypsum is not to be confused with clastic, re-worked gypsum, sometimes carried into the deep sea as turbidites. Petrographically, two forms of this secondary gypsum are recognized (Holliday, 1970); a porphyroblastic and an alabastrine type. The first form may be syndiagenetic, but may recrystallize to the second form, whereas the latter is most common as an epidiagenetic product. Within the littoral sabkhas bf the Persian Gulf, North Africa, and Baja California, the present writer has indeed found anhydrite nodules a few centimeters below the surface and they are clearly of Holocene age. But this is not to demonstrate that they are direct precipitates. Quite the contrary, i t now seems clear that the nodules were formed as gypsum, but were dehydrated to anhydrite under high temperature, i.e., under insolation (West et al., 1979). Experimental growth of primary anhydrite at low temperatures and normal pressure, with salinities resembling those of coastal lagoons and sabkhas, has always failed in the past under abiogenic conditions. Recently, however, in the presence of macromolecular organic compounds, characteristic of the hot, semi-arid littoral belt, Cody and Hull (1980) have now succeeded in growing well-crystallized anhydrite by primary nucleation. A very remarkable process, operating in the Persian Gulf on a Kuwait sabkha has been described by Gunatilaka et al. ( 1 980). Gypsum forms a fine crust on the halophytic plants and rapidly dehydrates to anhydrite, which retains the pseudomorphic texture. Upon the eventual death of the plant, the anhydrite hydrates once more to gypsum, and back again to anhydrite forming a bed up to 2 m thick. From field evidence, the present writer is of the opinion that dehydration from gypsum to anhydrite occurs most commonly upon burial to depths of the order of 100m or so (confining pressures of ca. 50 kg/cm2). Thus, for example, anhydrite has never been observed forming today in the Persian Gulf, but i t is extremely widespread all over the Mesopotamian Basin in Tertiary formations that, near the basin margins, have been buried under little more than l00m (Fig. 2-8). The gross effect of dehydration is to reduce the volume of the formation; from l 0 0 m of gypsum only 62m 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) stated that even a very slight dynamic metamorphism (with shear stresses, produced by local faulting and uneven settling of the basin) would greatly reduce this figure. It also seems probable that at certain times in the geologic past under distinctly different paleogeographic conditions and somewhat higher mean

65 Growth of gypsum by displacement within pre-existing sediment Peposition of gypsum in standing body of water,

Meteoric water

Ill,,

replaces anhydrite E PIDIAG E,NES IS

SY NDlAGENESIS

lGypsuml ANADIAGENESIS

I

I 1000 m I

Growth of anhydrite by replacement and as

I

gypsum after anhydrite \

Fig. 2-8. Schematic diagram illustrating the gypsum-anhydrite diagenetic cycle. (Modified after Blatt et al., 1980.)

temperatures primary precipitation of anhydrite may have occurred, so that not all sedimentary anhydrite should be regarded as the result of diagenetic dehydration (Fairbridge, 1967a). Annual layers of gypsum in the Tertiary of Sicily are observed to pass upwards, with increasing grain size, into anhydrite, which apparently represents a high summer peak (Ogniben, 1955). Likewise, in Mesozoic anhydrites from Texas, primary structures, including graded bedding, have been recorded by Riley and 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 (Newel1 et al., 1953). ( b ) Silica gels. Even more important and certainly more complicated is the dehydration of silica gels, to form authigenic opal, chalcedony and yuurtz.

66 The opal is not completely dehydrated and is, thus, somewhat unstable. It will not normally pass through the anadiagenetic phase without dehydrating to chalcedony or quartz (Fig. 2-9). The formation of quartz from amorphous silica requires something in excess of 10 million years (Mizutani, 1970). 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: ( 1) colloidal form and solution, through river transportation; (2) dust, by eolian transport from deserts; (3) volcanic glass and ash, from eruptions; (4) organic fossils (radiolarians, diatoms, and sponge, alcyonarian and holothurian spicules). The annual increment of SiO, from rivers to the oceans is 4.3 . lOI4g or 0.31 mg/m3 of sea water (see Chilingar, 1 9 5 6 ~and Burton and Liss, 1973), 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) mainly due 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 Siever, 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). There are quantitative problems, however, with the contemporary silica cycle (Erhart, 1973; Drever, 1977). First, the supply from rivers may be doubled if one allows for littoral processes, Antarctic supply, and wind transport, to bring the total to 5-12 lOI4g SiO,/yr. There are strong, Thick- Age of Mineralogical composition of silica ness deposition

(Si02)

2

m.$.BP 0

Quartz

A

C

Fig. 2-9. An example of the amorphous silica (A)-cristobalite (C)-quartz transition from Japan. (After Mizutani, 1970.)

67 climatogenetic cycles, short- and long-term. Pelagic ooze removal at the present time may account for about 2 . lOI4g SiO,/yr and estuarine processes up to 1 * 1014 g/yr. The disparity lends strength to the Sillen’s (1961) idea of a silicon sink to the clay minerals (Mackenzie and Garrels, 1966). In the geologic record, silica concentration is found either as bedded chert or nodular flint (see recent collected papers: McBride, 1979). The origin of the silica and the nature of its petrogenesis is a classic problem. Non-actualistic mechanisms are often invoked. Ronov ( 1964) has demonstrated that while bedded cherts are dominant in the Precambrian, they become less frequent in the Phanerozoic, whereas colloform nodules become more and more common, especially in limestones and chalks. Non-biologic accumulation of silica gel today was first demonstrated in ephemeral lake basins in Australia (Peterson and Van der Borch, 1965), where the pH is raised to over 10 by algal photosynthesis, causing saturation by available silica. During desiccation, the pH falls and a colloform gel precipitates. At the present time, silica gels on the sea floos are distinctly rare (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 volcanism (Hoss, 1957). But this may not be the only factor. One of the major discoveries of the Deep-sea Drilling Project and, indeed, one of the major developments in sedimentology in recent decades has been the discovery of extensive abyssal cherts (Van der Lingen, 1977). A systematic plot of the horizons where cherts were encountered shows that most of the chert was concentrated in the Eocene, i.e., around 50 m.y. old, and almost none was less than 10 m.y. old. There is a rough correlation of the principal clusters with the seismologists’ “horizon A” (Mid-Eocene) and “horizon B” (Early Cretaceous). Herman (1972) proposed that these accumulation peaks corresponded to episodic sea-floor spreading which, in turn, triggered vigorous oceanic circulation favorable to planktonic life. At the same time submarine volcanic emanations provided abundant silica, augmenting further the oceanic fertility. The question of direct contribution of volcanic material to the chert bands is controversial, although certainly they are very rich in siliceous plankton. The presence in them of smectite and zeolite suggests a duality (see discussions in Van der Lingen, 1977). There is controversy also over the Tethyan radiolarites, now found on land (Folk and McBride, 1978). Some argue for an abyssal environment (e.g., Folk), whereas others favor shallow-water origin (e.g., McBride). Whereas many cherts and flint nodules are the result

68 of replacement, Folk and McBride (1978) offer compelling evidence for cyclic accumulation of radiolarian opal transported by rhythmic currents, with subsequent slow transformation through a 5-25 m.y. time-span. The supply of silica solutions from soil-forming processes may be greater than that from volcanism, 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: Termier and Termier, 1963, fig. 197, p. 333; see also Fig. 2-7). 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 would tend to be eclipsed by floods of detritals. Indications of silica gels at times other than the present are numerous. Many instances of evidence for Phanerozoic primary silica gels have been listed by Dangeard and Rioult (1961). In the lowermost Ordovician of Poland, Samsonowicz (1948) found chalcedony forming a cast of ripple marks. Several generations of ripples were similarly preserved in the Jurassic of Normandy (Lemaitre, 1960). A chert containing a worm burrow was noted by Debelmas (1959). A fossil Exogyra, 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 northern Europe and Deflandre (1936) reported in them delicate Flagellata and hystrichosphaeridae with pseudopodia perfectly preserved. From the perfect preservation of fossils, Cayeux (1897) long ago concluded that the dehydration of the silica gel began on the outside, and worked slowly inwards. There are numerous indications of the small-scale coagulation of silica gel around nuclei or centers of lowered p H during the initial stage of syndiagenesis. The roots of mangrove-type trees, for example, are permanently fixed in marine muds, where the pH-Eh conditions show remarkable fluctuations (tidal and diurnal). 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 in coastal eolianites and their paleosols (Fairbridge and Teichert, 1953). Discussing their form and occurrence in the Jurassic of Normandy (“Pierre de Caen”), Dangeard and h o u l t (1961) have made a helpful review of this entire field. Such concentrations were found forming today in the mangrove swamps of New Caledonia (Avias, 1956). The world’s oldest fossils, apparently bacteria and filamentous blue-green algae. contained in 3.0-3.6 billion year-old formations in Australia, South

69 Africa, and Canada are partly found in silicified limestones. Some of the cherts (e.g., Onverwacht, Fig-Tree, Soudan, and Gunflint formations), however, appear to be either primary or very early diagenetic, inasmuch as the extracted hydrocarbons contain curiously heavy ‘3C/12Cisotope ratios that seem to reflect an original condition (Oehler et al., 1972). The question of migration of gels towards nuclei presents further difficulty. Mutual attraction of like molecules in solutions are normally attributed to the so-called Van der Waals’ forces, assumed to exist particularly in liquids. In their absence, 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 to tridymite and cristobalite (Jones and Segnit, 1971). The slow crystallization to quartz changes the rock eventually to chert and leads to complex brecciation phenomena (Taliaferro, 1934, 1935). Referring to brecciated chert, CJayeux (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, 3-10m high. They are interstratified in soft white chalky limestones extending over several hundred km2 (see Figs. 2- 10 and 2-1 1). The adjacent limestones, however, 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 next 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 after burial. Penecontemporaneous slump structures, large and small, are features of almost all chert formations. Careful mapping of the Judean and Cretaceous cherts shows that the early

70 WEST

EAST

JUDEAN HILLS (Upper Cretaceous chaiks)

Fig. 2-10. 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.)

(slump) structures are oblique with reference to the principal folding (Steinitz and Kolodny, 1978). Chert-filled dikes transverse to the bedding point to the slow lithification of the silica (Steinitz, 1970). Microstructures of the syndi-

Fig. 2- I I . Slump structure in the Upper Cretaceous cherts, east of Jerusalem. Height of the scction is about 40 m. (Photograph: courtesy of G. Steinitz.)

71 agenetic phase vary according to horizon. Isotopic work discloses evidence of rapidly fluctuating, eustatic sea levels, introducing successively evaporitic and fresh-water conditions (Kolodny et al., 1980), i.e., the concept of “schizohalinity” (Folk and Siedlecka, 1974). Relict evaporite traces have been found to be widespread in the Senonian of Israel (Steinitz, 1977). One can turn now to a second example of a timing indicator. In eastern Denmark, south of Copenhagen at Stevns Klint and MQns Klint, the Upper Maastrichtian chalk is marked by parallel planes of scattered flints, i.e., parallel to the stratification (see Figs. 2-12 and 2-13). There are also vertical joints, and here and there the silica of the flints has flowed up along the joint

Fig. 2-12. Flints near the top of the Senonian chalk at Stevns Klint, Denmark, illustrate the slow dehydration of silica gel. The flints occur as concretions both parallel to bedding planes and merging (continuously) into vertical joint planes. Evidently, the gel was still mobile at the time when the chalk was sufficiently dehydrated to develop joints. (After Fairbridge, 1967b. fig. 6a, p. 62.)

72

Fig. 2- 13. Detail of Fig. 2- 12. Pencil= 15 cm. Apparently silica gel has migrated up and down along a joint plane in the chalk.

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 offshoots. One must conclude that the intrusion and dehydration of the silica gel occurred immediately subsequent to the diaclastic phase. Such jointing must result in loss of gas pressure, release of CO,, CH,, 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 Early Danian age (Rosenkrantz and Rasmussen, 1960). Evidently the jointing and syndiagenesis occurred ufter 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 . lo5 years. Complete dehydration of silica gels may be extremely slow, and Sujkowski

73 ( 1958) mentioned some “flints” encountered in deep bores that were still

rather soft. There was also a report of a vein of silica gel encountered during the construction of an Alpine tunnel. Nevertheless, the emplacement of the silica is usually an early phenomenon. Studies of the ‘80/160 isotope ratios in coexisting cherts and limestones suggest early diagenesis under similar marine environments (Degens and 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 and 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 2500 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 “hardgrounds”, “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 ; Bathurst, 1975; Wilson, 1975). In such cases, the epidiagenetic phase is temporarily applied, the mud becomes rapidly dehydrated, and sometimes even develops a karst crust with a trace of red soil (terra rossa). When

74 reburied, no further reactions are likely and the remaining syndiagenetic phase is locally short-circuited. Even briefer is the formation of tropical beachrock from calcareous sands of 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 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 eustatic fluctuations or the revival of an old anticline could cause the loss of 10-20m 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 10,000-100,000 years older than the contemporary sedimentation, and possibly as much as lo6 years (see discussion under Timing of Anadiagenesis). ( c ) Authigenic feldspar. The formation of authigenic feldspar is a phenome-

non that has long been known, but little understood, for it is easy to assume that the feldspars are high-temperature silicates. It now appears, however, that in some cases it may be little more than a leaching and dehydration event, probably in the anadiagenetic phase such as outlined by Kastner (1971). If one takes, for example, a hydromica (muscovite, or its degenerated form illite), it is not too difficu4t to forecast the loss of alumina and dehydration to form orthoclase. In simple empirical form it can be presented as follows:

K,O . 3 A1,0,

- 6 SiO, . 2 H,O + K,O . A1,0,

6 SiO,

+ 2 H,O + 2 A1,0,.

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

75 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 be provided. Attention was first drawn to authigenic feldspars in the Cretaceous chalk of the Pans Basin (Cayeux, 1897, 1916), and they have been very generally reported elsewhere: in England (Reynolds, 1929), India (Spencer, 1929, and North America (Reed, 1928; Goldich, 1934; Gruner and Thiel, 1937; Willman, 1942). An exceptional case was observed by Daly (1917) in the 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; Van Straaten, 1948). The most coherent paleogeographic model has been offered by Bryce and Friedman (1979, who showed that the authigenic K-feldspars of the Cambrian-Ordovician in the North Atlantic region were directly associated with the carbonate shelf regions landwards of the island arc, which then stretched from Spitsbergen and Norway through Scotland and Newfoundland through the Appalachians to Georgia. It is proposed that the authigenic feldspars here are all that is left of tephra that were remobilized to form overgrowths on pre-existing clastics. The transport of K-feldspars would have been favored by the high alkalinity of the weathering process in the Early Paleozoic pre-land plant environment. Zeolites were probably present, but, being unstable, could contribute to the feldspathization. The source of ions necessary for an independent authigenic growth of feldspar could come most readily from volcanic glass (Honess and Jeffries, 1940). There seems to be an inverse relationship to zeolites with time, i.e., as the latter decrease in frequency in older periods, the abundance of authigenic feldspars increases (Hay, 1966). Some almost pure K-feldspar beds in Cenozoic lake beds of western North America lack any trace of tuffaceous texture, although certainly they were derived from volcanic ash (Surdam and Parker, 1972; Sheppard and Gude, 1973). 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 unatase may be derived from biotite (also from ilmenite, etc.; Sun and Allen, 1957). Further,

76 tourmaline and zircon are known in the authigenic form (Boswell, 1933; Pettijohn, 1957), but little seems to have been done about investigating their origin. Other examples (sphene, etc.) were mentioned earlier. 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. 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 permeable 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 which 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 experiment to demonstrate natural chromatography through a quartz sand. Inasmuch as quartz carries very little charge, the separation is mostly mechanical, i.e., differential capillarity. In the case of clays-and most sediments carry at least some admixture of clays -the filtering can become mostly 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+, M g 2 + ,Na’ and K + ) will be trapped. Deep basin brines may thus achieve a salinity up to six times that of ordinary sea water. In studying these ancient brines, Chave (1960) came to the conclusion that the variability in the nature of the sedimentary membranes was so great that a tremendous variety of

77 brines would result, and that no deductions could be drawn about the salinity of the ancient oceans. Similarly, wide variety is experienced with the petroleum hydrocarbons, which are exposed 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 (Irving, 1892). The average sedimentary basin may accumulate over a period of 106-108 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 reactions are expressed by cementation and decementation phenomena. The cements are mainly CaCO, or S O , with Fe,O, 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 for the formation of Fe,O, are present in certain near-surfaGe bacterial environments in the initial stage of syndiagenesis. The acid oxygenated water can also be introduced in the epidiagenetic stage. Cementation, involving complete silicification, dolomitization, and dedolomitization introduces the special problem of sedimentary metasomatism. Diagenetic metasomatism The term metasomatism, coined by Naumann ( 1850), 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 Holser, 1947). This process can be defined as 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 and, thus, normally occurs during the anadiagenetic phase, though sometimes in 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”.

78 Some examples of diagenetic metasomatism may be considered briefly as follows. Dolomitization Even after countless research investigations, the understanding of the dolomitization phenomenon is still beset with problems (see Chapter 7 of Vol. I, entitled “Dolomites and Dolomitization”, by Chilingar et al.). Individual cases are often difficult to interpret because it occurs in all three phases of diagenesis. ( a ) Syndiagenesis. As pointed out by several authorities, for example, Carozzi (1975, p. 354), the history of the dolomite problem has gone through two stages of activity. The early studies, in the middle and late 19th century up till the 1920’s, disclosed the complexity of the problem (see review by Fairbridge, 1957). First, there was the non-actualistic development of dolomites: the “uniformitarian dilemma”, whereby their frequency increases in going back in time (Zenger, 1972a, b). Secondly, they have proved singularly difficult to synthesize in the laboratory under reasonable earthsurface conditions. The second phase of the dolomite problem’s history began in the mid-20th century with introduction of the X-ray technique and other sophisticated laboratory equipment, coupled with extensive field work on actualistic situations. The first critical discovery was that of “protodolomite”, a disordered calcium-rich dolomite, synthesized under “reasonable” earthsurface conditions by Graf and Goldsmith (1956). High-magnesium calcites of organogenic origin had been found in shallow-marine situations by Chave (1954); like aragonite, they are metastable over relatively brief periods of geologic time. Chave (1954) and Chilingar (1953, 1962) showed that the higher the environmental temperature, the higher the content of biogenic magnesium in the calcites. Here was the potential for syndiagenetic dolomitization. Then, in the 1960’s came a succession of spectacular discoveries in shallow-water lagoons and intertidal mudflats (“sabkhas” in Arabic) of the warm subtropics, as predicted by Chilingar (1956d). Wade (191 1) had reported dolomite rhombs in the Holocene lagoonal muds on the Red Sea coast of Egypt. Intertidal magnesium-enrichment in tropical seas (notably beachrock) was identified by Reuling ( 1934) as “Gezeiten-dolomi tisierung” (intertidal dolomitization). Sir Douglas Mawson (1929) found them in the Coorong lagoon of South Australia, where he later showed them to the present writer. Here the study was taken up again by Alderman and Skinner (1957) and further by Skinner (1963). A comparable lagoonal situation was shown to the writer by R.A. Bramkamp in 1959, in the Persian Gulf south of

79 Dhahran, where gypsum was observed actively precipitating with what proved to be protodolomite. Another lagoon setting, at Bonaire in the Dutch West Indies, was recognized by Deffeyes et al. (1965). On the supratidal mudflats (sabkhas) of the Persian Gulf, the seasonal reversal of wind systems alternately wets and dries the littoral zone. A much wider zone has been alternately flooded and dried by small eustatic oscillations (+/- 3m) during the Holocene in 500-1000-yr cycles (Fairbridge, 1961). A revolutionary study by Illing et a]. (1965) described the tiny rhombs of protodolomite that replace the original aragonite and gypsum when exposed to magnesium-rich brines. Under high temperature (> 40°C) and capillary rise, the pore water salinity rises to 6-8 times that of sea water and the Mg/Ca ratio rises to over 10, whereas the pH remains low (ca. 6.7) because of biogenic activity. Radiocarbon dating of sediments at depths of 20-30 cm gives 2000-3000 B.P. ages. Valuable reviews of the Persian Gulf carbonate environments have been presented by Bathurst (1975) and Wilson (1975); a special volume has been edited by Purser (1973), and the “Benchmark” papers have been collected by Grkland and Evans (1973). A somewhat comparable setting was established in the intertidal muds west of Andros Island in the Bahamas (Shinn et al., 1965). It seems that 1965 was the “vintage year” for protodolomite. Whereas most investigators believe that the Holocene protodolomites are syndiagenetic, the development of penecontemporaneous dolomites in ancient formations tends to be controversial. Sedimentational evidence often suggests that the dolomites are indeed penecontemporaneous, but the petrologic evidence often indicates complex subsequent histories of recrystallization or metasomatism. The geochemical setting whereby dolomite could develop from sea water during gypsum formation (i.e., in partial isolation) due to the rise in Mg/Ca ratio was recognized as a theoretical model already by Hunt (1859), but experiments with mixed carbonates required artificial heating to generate dolomite. With progressive concentration, the solubility curves of CaCO, and MgCO, intersect, and at this point CaSO, also reaches saturation (Teodorovich, 1955), but still this does not help to explain non-evaporative situations. In spite of much experimental work, however, it was not until Chave (1954) showed the biogenic origin and metastable nature of the high-Mg calcites that the dilemma was overcome. The principal Mg-calcite secreting organisms are the calcareous algae, which because of their photosynthetic requirements and preference for warm water, are characteristic of shallow tropical seas and tidal flats. As pointed out by the present writer (Fairbridge, 1957, p. 154), there is abundant evidence associating many ancient dolomites with warm, shallow seas. The diachronism of the Mississippian Greenbrier Dolomite (Rit-

80 TOP

BED N O I

EXPLANATION MgCO,

0CaCo3 INSOL

Fig. 2-14. Vertical section, with MgCO, and CaCO, analyses, reduced to beds of equal proportions, of a sequence of Lower Ordovician rocks west of Harrisburg, Pa. (After Fairbridge, 1957.) There is rhythmic alternation of dolomitized and nondolomitic limestones, the dolomitization coinciding always with the beds containing maximum terrigenous (insoluble) material, and therefore presumably closer to the original coastline. (Analytic data by Lesley, 1879.)

tenhouse, 1949) also demonstrated “a very precise relationship to the shoreline”. Most dramatic are the 1879 analyses by Leslie of an alternating series of Lower Ordovician limestones and dolomites from near Harrisburg, Pa., apparently a eustatic cycle, which show that the terrigenous components are always highest in the dolomitic beds (Fig. 2-14). Sarin (1962) has proposed a cyclic killing of algae in the limestone facies by supersaline magnesium-rich currents. In any case, the dolomite here is a near-shore indicator. An interesting, consequence of these conclusions was the reversal of some contemporary oil exploration strategy: previously, the dolomites were often assumed to be the deep-water indicators. Others who had earlier pointed out the warm, shallow-water characteristics of various dolomites included: Skeats (South Pacific, 1903; Tyrol, 1905), Dixon (Britain, 1907), Van Tuyl (Iowa, 1916, 1918), Steidtmann (Wisconsin, 1917), and Fondeur et al. (France, 1954). In the widespread epicontinental platform seas of the late Precambrian

81 and Paleozoic, there seem to have been times of almost universal dolomite production. It has been customary to refer to such rocks as “primary dolomites” on the assumption that a primary crystallization would be favored by a marine hydrology that was somewhat different from that of today. There seems little doubt that the atmospheric oxygen level was low, and the p C 0 , was appreciably higher than today (Chilingar, 1956~;Strakhov, 1967, 1969; Fairbridge, 1964, 1967a). Experimental work by Baron ( 1960) demonstrated the possibility of direct precipitation of dolomite. A forceful argument, however, against a primary dolomitization environment in the Late Precambrian is the fact that there were also widespread limestones that were quite low in magnesium. Clearly there were regional patterns. An important paleogeographic factor is that throughout the period in question the equator transected North America, where much of the evidence is found. Whether or not these ancient dolomites underwent a rapid syndiagenetic transition, is still an unsolved problem (see, for example, Carozzi, 1975, p. 357). What is certain, however, is that they are commonly associated with evidence of photosynthetic algae, so that vast amounts of high-magnesium calcites could be generated that would then be susceptible to penecontemporaneous dolomitization, i.e., syndiagenesis. Isotopic data appear to point to this interpretation (Degens and Epstein, 1962). In certain varieties of calcareous algae, the MgCO, in calcite reaches a maximum of 25-30%. Inasmuch as this is a metastable condition, within a period of time (10’-104 years) it is likely to invert to pure calcite, whereas in the presence of high Mg2+ concentration dolomite will develop. Probably this is true metasomatic replacement of Ca2+ by Mg2+ (see, for example, photographs of calcareous algae from an Eniwetok Atoll bore, with euhedral dolomite rhombs beginning to form in the middle, in Schlanger, 1957). Numerous syndiagenetic limestones and dolomites are so tightly crystallized that they still retain the fetid odor of decaying organic material, which 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 benthonic biota (Moretti, 1957), not likely to be found under evaporite conditions. Syndiagenetic dolomite, as isolated rhombs, occurs also in the deep-sea today, in depths of 3000-4000m (Boggild, 1912, 1916; Correns, 1939; Zen, 1960). It has been suggested by Weynschenk (195 la, 195Ib) that massive and extensive deep-sea dolomitization was possible, but the D.S.D.P. “Glomar Challenger”-drilling has not disclosed any examples. For the scattered dolomite rhombs it would seem that in some cases transported algal nuclei of

82

high-Mg calcite may provide the source. Bonatti (1966) envisioned an igneous association, whereas Davies and Supko ( 1973) discussed other suggestions. Dolomite is also forming today in the Atlantis Deep of the Red Sea, at a depth of over 2000m, but in this instance it is associated with hot metal-rich brine seepages which are properly classified as hydrothermal. ( b ) 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 that a refluxing of dense brines from the saline lagoons would lead to an ideal setting (Illing, 1959; Adams and 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 Ardennes; they spoke of “epigenesis”, but the present writer suspects that this,is merely a different use of the term. It seems likely that this metasomatism is not achieved without some elevation of pressure. Progressive dolomitization results in ca. 12% increase in porosity due to reduction in volume (see detailed discussion in Chapter 5 of VOI. I). Whereas it is observed on a massive scale below a depth of 196m in Funafuti Atoll corresponding to 20 atm (Schmalz, 1956; Fairbridge, 1957, p. 148), 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 a controlling factor, 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 (Fig. 2-15). 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 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 dolomitizing waters. Across the Paleozoic of the Midwest it is often noticed that dolomitization is more prominent as one approaches the tectonic belts. In the great evaporite basins of the world, there is often an interstratification of anhydrite-dolomite beds with the salts (halite, etc.). A clear distinction is usually made between these evaporite basins and the littoral evaporites such as those of the lagoons and sabkhas of the Persian Gulf and North Africa. With the progressive loading, with/or without some tectonic defor-

83 WATER PRESSURE DEPTH TEMP 28*C AM

s100-

4-

3200-

3002,

I I

l

I

l

I

Fig. 2-15. Diagram to illustrate, in an idealized atoll, the suspected relationship of primary Mg-rich calcite (largely fixed by algae) to contemporary dolomitization, which may take place beneath the lagoon floor at a depth of 600-700 ft, where there is a pressure of about 20 atm. a temperature of at least 10°C (possibly raised by organic activity), a free access of Mg2+ ions from sea water, but a high alkalinity (pH about 9-10) due to reducing conditions brought about by organic activity both in the atoll walls and in the sediments themselves. Below the level of contemporary dolomitization, the atoll core may have been dolomitized discontinuously, the process being affected by changes in depth (and pressure) due to eustatic oscillations of sea level and varying rates of atoll subsidence. Periods of stability permit complete dolomitization to be achieved, while rapid subsidence permits a quick buildup of sediment, partly sealing off the former zone of dolomitization, so that completion is not achieved (e.g., 5-20% MgCO, in some of the Funafuti analyses). Funafuti atoll stands as the prototype for this scheme; many other Tertiary dolomitized atolls in the Pacific and Indian oceans have been partially elevated, so that the critical top 600 ft have been removed by erosion. Nondolomitized atolls, such as Bikini and Eniwetok may, for some reason or other, have failed to provide the necessary physicochemical requirements, such as the reducing conditions (owing to freer circulation) or may have subsided too rapidly. (After Fairbridge, 1957.)

mation, there ensues within the halite-potash type salts what is known as a “geothermal metamorphism” (Borchert and Muir, 1964). This takes place in fact within the zone of normal anadiagenesis, but is somewhat more dramatic in its changes because of the low stability range of the various salt minerals. Whereas some of the dolomitization is recognizably syndiagenetic, resulting from downward percolation of brines from the lagoon floor, there is some replacement of dolomite by anhydrite (Stewart, 1965). DuriPg deeper burial, there is further dolomitization, now anadiagenetic and associated with ascending solutions, and with rising temperature and pressure. Lateral flux of saline solutions in massive anadiagenetic dolomitization has been mentioned by Hite (1970), in addition to the reflux seepage model of Adams and Rhodes ( 1960).

84 ( c ) Epidiugenetic 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; see Fig. 2-16). I t is thus sometimes known as “subsequent dolomitization” (Hatch et al., 1938, p. 193). Dolomitization by ground water (i.e., of meteoric origin) was considered by Steidtmann (1911) and some other early workers, but not taken very seriously. In most situations, insufficient volumes of magnesium solutions

Fig. 2-16. Four generations of dolomite in the Eifel district of West Germany (adapted by Fairbridge, 1957, from work of Udluft, Reuling, and others). Idealized section showing the Devonian sediments, consisting of noncalcareous Lower Devonian followed by the Middle Devonian Couvinian (low-magnesian) limestone and Givetian (penecontemporaneous) dolomite ( 1 ); these were broadly folded in Pennsylvanian-Permian times and peneplaned. During Permian soil formation, magnesium-saturated ground water (exposed to weathering of the Givetian dolomite) led to dolomitization of the bevelled surface layer of the Couvinian limestones, as a continental “hardpan” ( 2 ) . Uplift led to some dissection of the peneplain, and locally the hardpan was stripped to expose nondolomitized Couvinian limestones. With the progressive invasion of the Triassic sea, sands (Bunter sandstone) were provided with the Middle Devonian limestone boulders (now red from terra rossa soil formation), and Couvinian dolomite boulders ( S ) , thus by the two earlier generations of dolomite. Penecontemporaneous and subsequent faulting, associated with further uplift, allowed hydrothermal solutions in places to introduce a fourth generation of dolomite ( 4 ) . locally accompanied by ore formation.

85 would be available to cause large-scale dolomitization. This objection is overcome at the fresh-water-salt-water mixing zone in porous littoral formations, such as exist in Florida and Yucatan (Hanshaw et al., 1971), in Jamaica (Land, 1973), in Israel (Magaritz et al., 1980), and elsewhere. Although Holocene nodules and hardpans have been found in such settings, no massive epidiagenetic dolomitization has yet been identified. Nevertheless, an attempt to explain the Middle Ordovician dolomites of Wisconsin in terms of a fresh-water-salt-water mixing zone has been made by Badiozamani ( 1973). This so-called “Dorag” model postulates an early interruption of syndiagenesis by epidiagenesis, which can be instituted in a cratonic setting of this sort by eustatic fluctuations. Dedolomitization Dedolomitization is generally a metasomatic replacement of dolomite by calcite; for this, Smit and Swett (1969) prefer to use the term “calcitization”. I t may also result simply from leaching, with a molecule by molecule solution followed by a selective removal of the MgCO,.‘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(CO, ),

+ CaSO, . 2 H ,O * 2 CaCO, + MgSO, + 2 H 2O

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 described the dedolomitization textures in the French Jura. Hydrothermal solutions may also play a role in dedolomitization (Faust, 1949). A second reaction, possible where MgSO, dominates the brines, may be given as: CaMg(CO,),

+ MgSO, * CaSO, + 2 MgCO,

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). Multiple reversals are favored in the dolomitization-dedolomitization of a region subject to a complex geologic history. Walker (1962) pointed out the reversible nature of chert-carbonate replacements. In the Ordovician of Wisconsin he discovered chert nodules replaced by dolomite, but at a later stage the process was reversed as evidenced by chert pseudomorphs of dolomite rhombs.

86 Silicification Inasmuch as silica is relatively insoluble at normal temperatures in waters having pH of less than 9 (Correns, 1949; Krauskopf, 1959), it is rarely mobilized after dehydration of the primary gels in the syn- and early anadiagenetic phases (see Fig. 2-7). Even low concentrations of dissolved SiO, will be precipitated at the low pH (3-6) existing in some lakes, leading to silicification of wood, as CO, is liberated by bacterial action (Correns, 1949), but such conditions can rarely be achieved in oceanic sediments. With progressive compaction and concentration of the connate waters during advanced anadiagenesis, however, sharp contrasts in pH may be obtained. According to calculations by Siever (1957b), equilibrium with the ZaCO, is reached at pH of 9.8 at 25OC, particularly in a thin film of water containing originally atmospheric CO, between grains of quartz and calcite. At this point silica dissolves to the limit of the very restricted water film. Pitting (etching) of quartz grains is often 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 H,SiO, occurs at pH of 9.8 and, therefore, the often quoted solubility curve of C.W. Correns (in: Barth et al., 1939, p. 129) may be misleading. Siever (1959) pointed out, however, that at a pH of 9.8 partial dissociation of H,SiO, produces silicate ions and essentially doubles the quartz solubility at lower pH values (Fig. 2-17). An additional factor is the increased polymerization 12

r

Fe3++ I(FeOHIz4

PH

Fig. 2-17. Solubility of silica gel, aluminum hydroxide and ferric iron hydroxide in relationship to pH. (After Correns, 1939.) Solid line indicates a modified SiO, solubility, following the work of Krauskopf (1959). Broken line indicates original curve of Correns.

87 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 conditions which raise the pH. Hydrothermal waters may assist such temperature rise, and silicification is a marked feature of deep-seated faults. Silicification involves the cementation, or complete replacement of sedimentary particles or fossil remains, by silica, usually in the sequence opalchalcedony-quartz, ‘the end-point requiring lo7 yr or more. The process is often included in the term “induration”, which implies general hardening of the rock. During weathering under seasonally contrasting climates, the capillary transport of silica to the surface of exposed rocks results in reprecipitation, producing a siliceous rind, or “case-hardening”. Far-reaching silicification is even 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). 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 a 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). Quartz-calcite intergrowths occur often (Dapples, 1959; Walker, 1962). As epidiagenesis proceeds, there is a gradual tendency towards stabilization in the intermediate pH range, where calcite is still soluble and silica is completely stable. Thus, on old land masses, stable tectonically for a long period of time, there is a widespread silicification of limestones. Where, as in Africa, Australia, Peninsular India, and the Brazilian Shield, the climatic record of the last lo8 years or so has been tropical or subtropical, there have

88 been immense and almost continuous supplies of SiO, passing into the vadose water system by leaching down from lateritized 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 (stromatolites) of the Precambrian in those regions long went unrecognized, because they were found only in what appeared to be primary quartzites. A widespread phenomenon of epidiagenetic silicification of sands is the development of silcrete, the siliceous duricrust, or hardpan, of the semi-arid tropics and subtropics (Dury, 1974; Langford-Smith, 1978). A comparable argillaceous rock in this same way is converted to a porcellanite (also spelled “porcelanite”), which has the conchoidal fracture of unglazed porcelain. The environmental requirement for epidiagenetic silicification is a highly contrasting, seasonal climate fluctuation, whereby the silica is mobilized during the (warm) wet season and then reprecipitated during capillary lift during the hot dry season (Fairbridge, 1975). It is curious that, whereas the silcrete crusts of the present tropics and subtropics of the Southern Hemisphere are usually only 1-10 m thick, the massive silicification of many porous sandstone formations in the Northern Hemisphere may extend to a depth of 200-500m. Thus, for example, most of the Paleozoic quartz sandstone and conglomerate formations of the Appalachians are porous at depth (acting as oil and gas reservoirs); nevertheless, in outcrop, they are massively silicified, the induration having blocked all porosity. The explanation appears to lie in the fact that in plate tectonic history the Northern Hemisphere continents have shifted polewards since the Mesozoic, progressively traversing tropical to temperate climate belts and undergoing extensive differential uplifts. This amplified the entry of meteoric waters and promoted deep-working epidiagenesis. An interesting phenomenon that has received notice is what in the Paris Basin is called “meulerization” from the French word for millstone (Termier and Termier, 1963, p. 345). It is believed to be analogous to the formation of siliceous hardpan 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. The geomorphologic history of the Paris Basin discloses a Late Tertiary subtropical peneplanation, during which small ephemeral-lake depressions appear to have determined the sites of the “meulerization”. An opposite reaction is sometimes observed during the epidiagenetic phase in temperate latitudes. This has been called “Fontainebleau sandstone crystallization”, to describe the local formation of a calcite cement in a

89

Fig. 2-18. Desilicification of quartz sand grains. aided by the base exchange of intercalated clays (after E. Thomson. 1959). in four stages: A. Nature of quartz and clay shortly after deposition. Clay is present between some grains and not others. B. Base exchange of K by Ca2+ and Mg’+ begins along edges of sheets. C. Exchange proceeds rapidly along cleavage surfaces. K,CO, 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 p H and precipitates.

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 crystallographic continuity. 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) phases 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. 2-18). 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, 19551, 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 would liberate K,CO, if subjected to C0,-rich waters. K,CO, is a strong alkali, which would mobilize SiO, at the clay contact, only to be reprecipitated nearby in the generally acid solution.

90 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 limestones. The problems associated with phosphatization are quite as complex as those connected with dolomitization. Phosphorites (a collective name for phosphatic rocks) appear to be of both biogenic and inorganic origin. Most biogenic phosphorus 'is part of a food chain that begins with unicellular marine organisms, which pass to fish and then sea-birds. Droppings of the latter accumulate as guano. Leaching of this guano into coral or other porous carbonate substrate generates an epidiagenetic phosphorite. Phosphorus is usually introduced (in reactable form) into the sedimentary cycle at the present time by organic agencies. It is present in nucleic acids formed in all living matter. It is contained in many proteins (phosphoproteins), many lipids, and many carbohydrates. Its abundance is rather low in sea water (about 0.07 mg/l), so that the limiting factor in biogenic 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. The present-day inorganic phosphorites are mainly limited to nodular concretions of carbonate fluorapatite or a phosphormicrite, which show discontinuous but penecontemporaneous accumulations. Some are inherited (lag) deposits. Deposits are associated with shelf margins from 30- 1000m depth, off California and Mexico, Peru, Chile, northwestern and eastern Australia, northwest Africa, southwest Africa (Namibia), South Africa, Oman, and Somalia. The phosphatization of benthic foraminifera was reported by Manheim et al. (1975). From almost all periods of geological time, the sedimentary phosphates are of marine origin. Although the secondary terrestrial concentration often occurs, the initial segregation is marine. Certain periods seem to be more favorable than others, e.g., Cambrian, Permian, Upper Jurassic, Lower Cretaceous, Upper Cretaceous, and Tertiary (Gimmelfarb, 1956); these are essentially transgressive, 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; Riviere and Visse, 1954). In subsiding basins marked by a higher accumulation rate and probably by somewhat greater depths and poor circulation, the

91 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). Besides continental margins, some estuaries carry up to 1-ppm phosphate, so that a site parallel to an old, deeply eroded orogenic belt, such as the Appalachian Piedmont, could also be favorable. To explain the nodular and pelletal phosphorites of the geologic record, Cook (1976) suggests the following criteria: (a) influx of phosphate-rich water to a warm shallow marine shelf basin or margin relatively free from important terrigenous sedimentation; (b) development of a rich neritic biota; (c) existence of an anoxic benthic environment, which inhibits predation and bottom scavengers and, thus, permits the accumulation of a, phosphatic mud (during the early syndiagenesis phase there is a loss of C, N and H). In this way formation of an organogenic ooze is postulated; the possibility of purely inorganic precipitation, perhaps around organic nuclei, remains open. The next step is phosphatization under the low pH-high alkalinity conditions of early burial in the syndiagenetic phase. Leaching of the phosphaterich solutions leads to metasomatism of carbonate fossils, fecal pellets, bones, coprolites, shark teeth, and oolites by apatite. Even a siliceous ooze can be phosphatized. Bottom currents and wave action lead to periodic winnowing of the sediments, concentrating phosphatic nodules and pellets as lag deposits. The dilemma of phosphates, in general, is their apparently non-actualistic development in the geologic record. At certain times vast concentrations occurred, but the Present is not one of them. In the Late Cenozoic, sedimentary phosphates appear to be restricted to marginal-shelf regions, marked by strong upwelling, mainly in the trade-wind belts between 40°N and 40"s. This theory was proposed by the Soviet worker A.V. Kazakov in 1937 (see review by Blatt et al., 1980). For earlier periods, plate reconstructions and paleomagnetically established paleolatitudes suggest similar distributions (Cook and McElhinny, 1979). The saturation of PO:- in sea water is greatest near the shelf margin (0.3-0.8 ppm) in upwelling areas of the trade-wind belts. At shallower depths it is reduced by algal metabolism, whereas at greater depths it is reduced by the higher pC0,. Massive phosphate precipitation, however, would also require a reduction of the combined nitrogen level due to denitrification by bacteria, according to Piper and Codispoti (1975). This could have occurred during warmer times of the geologic past when there was expansion of the oxygen minimum layer (0, content of less than 0.1 ml/liter), which today is generally located below 200 m, at depths marked by higher pC0,. The Quaternary, in general, is an unfavorable period for this scenario.

92 Thus the stage for large-scale phosphatization is preset in geological time by certain geotectonic-paleogeographic events, that do not prevail at the present stage in history. The extreme and large-scaled oscillations of sea level during the Quaternary, however, favor a special epidiagenetic environment that is found in isolated atolls. These offer sites for sea-bird sanctuaries, free from predators until recently introduced by man. Deposits of phosphate-rich guano accumulated here. During the high sea-level stages, the islands were small and leachng of the soluble organic phosphates by rain water brought them quickly back into the sea. During the low sea-level episodes, however, the atolls became emerged and resembled mediaeval castles with limestone walls and floors. This was an 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), Ca,P20, . H 2 0 , or dahllite, Ca,(PO,), - CaCO, . $ H20. On small volcanic islands with a barrier or fringing reef, the incomplete atolls (e.g., Navassa Island, West Indies), or along semiarid mainland coasts (e.g., southwestern Africa, northern Chile), the same thing may happen. In such cases, however, the igneous weathering products may interact during the diagenesis, resulting in vivianite, Fe,P20, * 8 H 2 0 , or wavellite, 4 AIPO, 2 Al(OH), * 9 H 2 0 . The geochemistry of phosphatization of atolls suggests a low pH in a fresh-water-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. 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 phosphorus 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). 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 5000-50,000 years, which is the upper limit of the time required for phosphatization of quite thick formations. 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

93 majority show 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). 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 Sideritization is a fifth type of sedimentary metasomatic diagenesis. Primary precipitation as siderite, FeCO,, probably does not occur in open marine environments today. In swamps and other restricted basins, however, the ferric oxide hydrosols of river waters would be reduced to the ferrous state, and removal of CO, by photosynthesis of plants would dissociate the bicarbonate ions to cause the direct precipitation of FeCO,. Some sideritic iron ores are thus associated with some coal deposits. Teodorovich (1949, 1961) claimed also that there were at times distinct “siderite facies” of marine environments of the past, marked by reducing conditions and by strong fluctuations of pH and Eh. Much more usual, however, seems to be the condition where marine limestones are metasomatically replaced by siderite during syndiagenesis (Cayeux, 1916; Hatch et al., 1938, p. 135). The CaCO, 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, bryozoans 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 electrolytically flocculated and thus deposition is likely to occur on continental shelves rather than in the deep sea (Moore and Maynard, 1929, p. 507). On semi-tropical shelves one may thus expect the ideal environment. Iron is readily leached out of the laterite soils 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 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

94 preceding orogenic episodes was stressed by Cayeux, but the role of biorhexistatic soil cycles in concentrating the iron should not be forgotten (Erhart, 1956, 1973a, b). 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, Fe,O,, and with chamosite, 3 F e 0 . A1,0,. 2Si0, * H,O. 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 chamosite-oolite (Lower Jurassic of England), and the incorporation of penecontemporaneous pebbles of pyritized ironstone in the same formation (Hatch et al., 1938, p. 135). In the Middle Jurassic Northampton “ironstones” of England and in the contemporaneous Minette iron ores of Alsace-Lorraine, the iron was clearly derived from the tropical weathering of nearby land. The sedimentary environment was a shallow continental shelf where oolitic carbonates, mudstones, siltstones, and sandstones were accumulating in rhythmic sequences. The ores are now oolitic limonite-chamosite-siderite rocks, but the question of replacement is controversial (Kimberley, 1974). Others argue for a primary chamosite mud (Taylor, 1949; Knox, 1970). Eustatic fluctuations provided periodic exposure (Hallam and Bradshaw, 1979). It constitutes yet another non-actualistic problem area (Taylor, 1949). The situation in the case of the Silurian Clinton (Llandovery) iron ores of the eastern United States (New York State to Alabama) appears to be similar, but here the oolites and carbonate fossil debris have been replaced only by hematite (Hunter, 1970). Desideritization is observed in some instances in the Jurassic iron ores. Just as in the case of dedolomitization, siderites may be replaced by calcite forming very similar textures (Taylor, 1949). Certainly, the most remarkable of all sedimentary iron ore formations are the so-called “BIF”, the banded iron formations of the Precambrian (Trendall, 1968; James and Sims, 1973). The oldest are dated around 3.75 billion years, but the bulk are 2.6- 1.8 billion years old. Characteristically, they consist of finely interbedded chert and magnetite- or hematite-rich layers, in some cases with siderite, pyrite, or other minerals. In a few instances, there are oolitic textures. From the evidence of ancient stromatolites and other traces of photosynthetic life, it is likely that free oxygen was being generated at an early stage, so that ferrous iron liberated by weathering could be oxidized to hematite (Dimroth and Kimberley, 1976). Nevertheless, it is argued by Cloud (1973) that all, or almost all, of the oxygen was rapidly locked up in these massive oxides and so its seasonal supply may well have played a limiting boundary condition to the BIF-layering.

95 Moore 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 (Fe20,-SO,) 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 phenomenon in giant lakes, and it seems not improbable that the Precambrian sea had a salinity somewhat lower than the present (and a lower pH) so that the ocean of that time might be compared geochemically with a brackish lake (Fairbridge, 1967a). Siderites are also known in these deposits. It seems 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 CO, and elevation of pH. The deposits were modified in some cases by later hydrothermal action. Bauxitization Bauxite is the rock name for a mixture of alumina minerals, amorphous or crystalline hydrous aluminium oxides and aluminium hydroxides, mainly gibbsite, often with boehmite and diaspore. In many places it is an aluminous laterite, i.e., an indurated residual soil or paleosol. “Bauxitization” is believed to be usually the epidiagenetic alteration of the parent material, a kaolinite, by desilicification (Erhart, 1973b). In some cases, the parent material may be feldspars in the underlying bedrock or in volcanic ash carried in subaerially. The hydrolysis of kaolinite may be expressed as follows:

A120,. 2 SiO, . 2 H,O (kaolinite)

+ H 2 0+Al,O,.

+

n H 2 0 2 SiO,. 2 H,O (water) (bauxite) (silicic acid)

Contemporary large-scale bauxitization does not appear to have been demonstrated, so that yet another non-actualistic process must be assumed (Davidson, 1964). Bauxitization appears to be relatively restricted in time and space, occurring only in tropical latitudes in the littoral belt or at relatively low elevations. Quite exceptionally, a lacustrine site is known. The most likely sites include tropical mangrove swamps, because the most important deposits in South America (Surinam, Guyana) are associated with a Miocene shoreline and contain fossil mangrove roots (Valeton, 1972). Related deposits are colluvial spreads apparently derived from them, but distributed during the Pliocene and Quaternary phases of lowered base level. Other deposits do not fit this scenario and appear to be related to phases of lowered sea level in otherwise marine limestone sequences. Although it is widely assumed that the bauxites are derived from the clays liberated by

weathering of immense thicknesses of carbonate deposits, the volumetric problem is unsurmountable. Thus it seems likely that in most cases the clays are derived from volcanic ash weathering. Again, coastal swamps with warm temperatures and low p H may provide the necessary biogenically controlled epidiagenetic setting, to be followed by desiccation with partial dehydration.

GLOSSARY Anadiagenesis (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, 1967b). Anamorphism (adj. -ic): Metamorphism at depth, forming more complex minerals (Van Hise, 1898; modified to exclude low-temperature alteration, i.e., by diagenesis). Authigenesis (adj. -ic, -om): Formation of new sedimentary minerals in situ, within the enclosing sediment, during or after deposition (e.g., Pettijohn, 1957). Diagenesis (adj. -etic): Physical and chemical changes which a sediment undergoes after deposition and during lithification, without introduction of heat (6ver ca. 300°C) or great pressure (ca. 1000 bars) (Von Guembel, 1868; modified slightly by Walther, 1894; Fairbridge, 1967b). 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, 1967b). 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 (A.G.I. Glossary). Also applied to mineral deposits, as epigene. One may have both epigeneric supergene (with descending waters) and epigenetic hypogene (ascending waters). Halmyrolysis (adj. -1ytic): Geochemical modification of sediments during deposition, due to reactions with sea water (ionic transfer), originally called “submarine weathering” by Hummel (1922); but applies also to ionic rearrangement and replacement (Pettijohn, 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, involving magmatism and metamorphism (Lyell, 1833; Geikie, 1879). Ka;amorphism (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). Thus it includes both weathering (upper zone) and near-surface cementation (lower zone), but Leith and Mead (1915) excluded the latter. Lithification: The complex of processes that converts an unconsolidated sediment into a hard rock, including compaction, dehydration, cementation and induration (e.g., Pettijohn, 1957). Lithogenesis (adj. -etic): Synonymous with petrogenesis, relating to the origin of a rock (A.G.I. Glossary; Bates and Jackson, 1980). Supergene (adj.): Applied to mineral deposits or enrichment related to descending waters (A.G.I. Glossary). One may have both epigenetic supergene deposits and syngeneric supergene deposits (as in manganese nodules).

97 Syndiagenesis (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, 1967b). 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).

REFERENCES Abelson, P.H., 1959. Geochemistry of organic substances. In: P.H. Abelson (Editor). Researches in Geochemistty. Wiley, New York, N.Y., pp. 79-103. Adams, J.E. and Rhodes, M.L., 1960. Dolomitization by seepage refluxion. Bull. A m . Assoc. Pet. Geol., 44: 1912-1920. Alderman, A.R. and Skinner, H.C.W.. 1957. Dolomite sedimentation in the southeast of South Australia, A m . J . Sci., 255: 561-567. Alimen, H., 1958. Observations petrographiques sur les meulieres pliocenes. Bull. Soc. GPol. Fr., 8: 77-90. Allen, E.T., Crenshaw, J.L. and Johnston, J.J., 1912. The mineral sulfides of iron. A m . J . Sci.. 33: 169-236. Allen, J.R.L., 1980. Continental Red Beds. (Developments in Sedimenrolog~.29). Elsevier, Amsterdam, 562 pp. Almon, W.R., Fullerton, L.B. and Davies. D.K., 1976. Pore space reduction in Cretaceous sandstones through chemical precipitation of clay minerals. J . Sediment. Petrol.. 46: 89-96. Amstutz, G.C. and Bubenicek, L., 1967. Diagenesis in sedimentary mineral deposits. In: G. Larsen and G.V. Chilingar (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, pp. 4 17-475. Anderson, D.L. and Benson, C.S., 1963. The densification and diagenesis of snow. In: Ice clnd Snow. Mass. Inst. Technol., Cambridge, Mass., pp. 391-41 1. Andree, K., 1911. Die Diagenese der Sedimente. Geol. Rundsch., 2: 61-74. Arrhenius, G., 1950. Carbon and nitrogen in subaquatic sediments. Geochim. Cosmochim. Acts, 1: 15-21.

Arthur, M.A. and Schlanger, S.O., 1979. Cretaceous “oceanic anoxic events” as causal factors in development of reef-reservoired giant oil fields. A m . Assoc. Pet. Geol. BUN., 63: 870-885. Ault, W.U., 1959. Isotopic fractionation of sulfur in geochemical processes. In: P.H. Abelson (Editor), Researches in Geochemistry. Wiley, New York, N.Y., pp. 241-259. Ault, W.U. and Kulp, J.L., 1959. Isotopic geochemistry of sulphur. Geochim. Cosmochim. Acta, 16: 201-235. Avias, J., 1953. Sur la formation actuelle de gypse dans certains marais cBtiers de la Nouvelle-Caledonie. Congr. GPol. Int., 19e. Alger, I952, C. R .. 4: 7-9. Avias, J., 1956. La probleme des nodules petrifies des mangroves neocaledoniennes. Acres Congr. Assoc. int. Quaternaire, Rome- Pise, 1953, pp. 245-249. Baas Becking, L.G.M., 1959. Geology and microbiology. Contrih. Mar. Microhiol., N . Z . Oceanogr. inst., 22: 48-64. Baas Becking, L.G.M. and Moore, D., 1959. The relation between iron and organic matter in sediments. J . Sediment. Petrol., 29: 454-458.

Baas Becking, 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. Bader, R.G.. 1954. The role of organic matter in determining the distribution of pelecypods in marine sediments. J . Mar. Res., 13: 32-47. Bader, R.G., 1956. The lignin fraction of marine sediments. Deep-sea Res., 4: 15-22. Badiozamani, K., 1973. The Dorag dolomitization model-application to Middle Ordovician of Wisconsin. J . Sediment. Petrol., 43(4): 965-984. Baron, G., 1960. Sur la synthese de la dolomie; application au phenomene de dolomitisation. Reu. inst. Fr. Petrol. Ann. Combust. Liquides, 15: 3-68. Barth, T.F.W., 1962. Theoretical Petrology. Wiley, New York, N.Y., 2nd ed., 416 pp. Barth, T.F.W., Correns, C.W. and Eskola, P., 1939. Die Entstehung der Gesteine-ein Lehrbuch der Petrogenese. Springer, Berlin, 422 pp. Bastin, E.S., 1926. The presence of sulfate-reducing bacteria in oil waters. Science, 63: 21-24. Bates, R.L. and Jackson, J.A. (Editors), 1980. Glossary of Geology. Am. Geol. Inst., Falls Church, Va, 2nd ed., 749 pp. Bathurst, R.G.C., 1975. Carbonate Sediments and Their Diagenesis, (Developments in Sedimentology, 12). Elsevier, Amsterdam, 2nd ed., 658 pp. Baturin, V.P., 1937. Paleogeography on the Base of Terrigenous Components. Baku-Moscow, 292 pp. (in Russian with English summary). Berg, R.R., 1952. Feldspathized sandstone. J . Sediment. Petrol., 22: 221 -223. Berner, R.A., 1964. Distribution and diagenesis of sulfur in some sediments from the Gulf of California. Mar. Geol., 1: 1 17- 140. Berner, R.A., 1970. Sedimentary pyrite formation. A m . J . Sci., 268: 1-23. Berner, R.A., 1971. Principles of Chemical Sedimentology. McGraw-Hill, New York, N.Y., 240 PP. Berner, R.A., 1975. The role of magnesium in the crystal growth of calcite and aragonite from sea water. Geochim. Cosmochim. Acta, 39: 489-504. Berner. R.A., 1980. Early Diagenesis. Princeton University Press, Princeton, N.J., 250 pp. Berry, F.A.F., 1960. Geologic field evidence suggesting membrane properties of shales. Bull. A m . Assoc. Pet. Geol., 44: 953 (abstr.). Berry, R.W. and Johns, W.D., 1966. Mineralogy of the clay-sized fractions of some North Atlantic-Arctic Ocean bottom sediments. Bull. Geol. Soc. Am., 77(2): 183-195. Bien, G.S., Contois, D.E. and Thomas, W.H., 1958. The removal of soluble silica from fresh water entering the sea. Geochim. Cosmochim. Acta, 14: 35-54. Biscaye, P.E., 1964. Mineralogy and sedimentation of the deep-sea sediment fine fraction in the Atlantic Ocean and adjacent seas and oceans. Yale Uniu. Geochem. Tech. Rep., 8: 86 PP. Biscaye, P.E., 1965. Mineralogy and sedimentation of Recent deep-sea clay in the Atlantic Ocean and adjacent seas and oceans. Bull. Geol. SOC.Am., 76: 803-832. Biscaye, P.E. and Dasch, E.J., 197 1. The rubidium, strontium, strontium-isotope system in deep-sea sediments: Argentine Basin. J . Geophys. Res., 76(21): 5087-5096. Biscaye, P.E. et al., 1974. Rb-Sr, *'Sr,lE6Sr isotope system as an index of provenance of continental dusts in the open Atlantic Ocean. J . Rech. Atmosph., 8(3-4): 819-829. Bissell, H.J., 1959. Silica in sediments of the Upper Paleozoic of the Cordilleran area. In: Silica in Sediments-S.E.P.M., Spec. Publ., 7 : 150-185. Blackwelder. E., 1947. Diagenesis and weathering. Bull. A m . Assoc. Pet. Geol., 31: 500. Blatt. H.. Middleton, G. and Murray, R., 1980. Origin of Sedimentary Rocks. Prentice-Hall, Englewood Cliffs, N.J., 2nd ed., 782 pp. (1st ed. 1972.) Boggild, O.B., 1912. The deposits of the sea-bottom. Rep. Danish Oceanogr. Exped., 19081910. Medit.. l(3): 255-269.

99 Boggild, O.B., 1916. Meeresgrundproben der Siboga-Expedition. Siboga-Expeditie. Leiden. 60: 50 pp. Bonatti, E., 1963. Zeolites in Pacific pelagic sediments. Trans. N.Y. Acad. Sci., 25: 938-948. Bonatti, E., 1966. Deep-sea authigenic calcite and dolomite. Science, 153: 534-537. Borchert, H. and Muir, R.O., 1964. Salt Deposits. Van Nostrand, London, 300 pp. Boswell, P.G.H., 1933. On the Mineralogy of the Sedimentary Rocks. Murby. London, 393 pp. Bradley, W.H., 1929. The occurrence and origin of analcite and meerschaum beds in the Green River Formation of Utah, Colorado, and Wyoming. U.S., Geol. Suro., Proj. Pap., 158-A: 1-8. Braitsch, O., 1962. Entstehung und Stoffbestand der Salzlagerstatten. Mineral. Petrogr. Mitt., 3: 232. Bredehoeft, J.D., Blyth, C.R., White, W.A. and Maxey, G.B.. 1963. Possible mechanism for concentration of brines in subsurface formations. Bull. A m . Assoc. Pet. Geol., 47: 257-269. Breger, LA., 1960. Diagenesis of metabolites and a discussion of the origin of petroleum hydrocarbons. Geochim. Cosmochim. Acta, 19: 297-308. Briskin, M. and Schreiber, B.C., 1973. Authigenic gypsum in marine sediments. Mar. Geol.. 28: 37-49. Brongersma-Sanders, M., 1957. Mass mortality in the sea. In: J.W. Hedgpeth (Editor). Treatise on Marine Ecology and Paleoecology-Geol. SOC.A m . , Mem., 67( 1): 941 - I 110. Burton, J.D. and Liss, P.S., 1973. Processes of supply and dissolved silicon in the ocean. Geochim. Cosmochim. Acta, 37: 1761-1773. Bryce, M.R. and Friedman, G.H., 1975. Significance of authigenic K-feldspar in CambrianOrdovician carbonate rocks of the proto-Atlantic shelf in North America. J. Sediment. Petrol., 45(4): 808-821. Carozzi, A. (Editor), 1975. Sedimentary Rocks: Concepts and Histor,v (Benchmark Papers in Geology, 15.) Dowden, Hutchinson and Ross, Stroudsburg, PA, 468 pp. Carpenter, A.B., 1979. Discussion-Dorag dolomitization model by K. Badiozamani. J. Sediment. Petrol., 43: 965-984; 46: 254-256. (Reply, 46: 256-258). Caspers, H., 1957. Black Sea and the Sea of Azov. In: J.W. Hedgpeth (Editor), Treatise on Marine Ecology and Paleoecologv-Geol. SOC.A m . , Mem. 67( I): 803-889. Cayeux, L., 1897. Contribution a I’etude micrographique des terrains sedimentaires. MPm. SOC.Gkol. Nord (Lille). 24: 168- 187. Cayeux, L., 1916. Introduction a I’etude petrographique des roches sedimentaires. MPm. Carte Gkol. Fr., 2 vol. (2 td., 1931). Cayeux, L., 1929. Les Roches skdimentaires de France; Roches silicieuses. Masson, Paris. 774 PP. Cayeux, L., 1941. Causes anciennes et Causes actuelles en GPologie. Masson, Paris, 82 pp. Chave, K.E., 1954. Aspects of the biogeochemistry of magnesium, 1. Calcareous marine organisms; 2. Calcareous sediments and rocks. J. Geol., 62: 266-382; 587-599. Chave, K.E., 1960. Evidence on history of sea waters of ancient basins. Bull. A m . Assoc. Pet. Geol., 44(3): 357-370. Chilingar, G.V.. 1953. Use of Ca/Mg ratio in limestones as a geologic tool. Composs, 30: 203-209. Chilingar, G.V., 1955. Review of Soviet literature on petroleum source-rocks. Bull. A m . Assoc. Pet. Geol., 39: 764-768. Chilingar, G.V., 1956a. Black Sea and its sediments-a summary. Bull. A m . Assoc. Per. Geol.. 40: 2765-2769. Chilingar, G.V., 1956b. Joint occurrence of glauconite and chlorite in sedimentary rocks: a review. Bull. A m . Assoc. Pet. Geol.. 40: 394-398.

Chilingar, G.V., 1 9 5 6 ~ .Distribution and abundance of chert and flint as related to the Ca/Mg ratio of limestones. Bull. Geol. SOC.Am., 67: 1559-1561. Chilingar, G.V., 1956d. Relationship between Ca/Mg ratio and geologic age. Bull. A m . Assoc. Per. Geol., 40: 2256-2266. Chilingar. G.V., 1958. Some data on diagenesis obtained from Soviet literature: a summary. Geochim. Cosmochim. Acta. 13: 213-217. Chilingar, G.V., 1962. Dependence on temperature of Ca/Mg ratio of skeletal structures of organisms and direct chemical precipitates out of sea water. Bull. S . Cali/. Acad. Sci., 61: 45-60. Chilingarian, G.V. and Vorabutr, P., 1981. Drilling and Drilling Fluids. Elsevier, Amsterdam, 767 pp. Chilingarian, G.V.. Sawabini, C.T. and Rieke, H.H., 1973. Effect of compaction on chemistry of solutions expelled from montmorillonite clay saturated in sea water. Sedimentology, 20: 39 1-398. Choquette, P.W. and Pray, L.C., 1970. Geologic nomenclature and classification of porosity in sedimentary carbonates. Bull. A m . Assoc. Pet. Geol., 54: 207-250. Cloud Jr., P.E., 1955. Physical limits of glauconite formation. Bull. A m . Assoc. Pet. Geol., 39: 484-492. Cloud, P.E.. 1973. Paleoecological significance of banded-iron formation. Econ. Geol., 68: 1135-1143. Cody, R.D. and Hull, A.B., 1980. Experimental growth of primary anhydrite at low temperatures and water salinities. Geology, 8: 505-509. Conley, R.F. and Bundy, W.M., 1958. Mechanism of gypsification. Geochim. Cosmochim. Acra, 15: 57-72. Conway, E.J., 1945. Mean losses of Na, Ca, etc. in one weathering cycle potassium removal from the ocean. A m . J . Sci., 243: 583-605. Cook, P.J., 1976. Sedimentary phosphate deposits. In: K.H. Wolf (Editor), Handbook o/ Strata-bound and Stratiform Ore Deposits. Elsevier, Amsterdam, pp. 505-535. Cook, P.J. and McElhinny. M.W., 1979. A reevaluation of the spatial and temporal distribution of sedimentary phosphate deposits in the light of plate tectonics. Econ. Geol., 74: 3 15-330. Coombs, D.S., Ellis, A.J., Fyfe, W.S. and Taylor, A.M., 1959. The zeolite facies, with comments on the interpretation of hydrothermal syntheses. Geochim. Cosmochim. Acta, 17: 53-107. Cooper, B.N.. 1954. Fundamental problems of genesis of Appalachian dolomites. Virg. J . Sci., 5: 301-302 (abstr.). Correns. C.W., 1939. Pelagic sediments of the North Atlantic Ocean. In: D.W. Trask (Editor), Recent Marine Sediments. Am. Assoc. Pet. Geol., Spec. Publ., pp. 373-395. Correns, C.W., 1949. Ein/uhrung 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. Con/., Clays Ctay Miner., loth -Natl. Acad. Sci. Natl. Rex Coun., Publ., pp. 443-459. Correns. C.W. and Von Engelhardt, W., 1938. Neue Untersuchungen iiber der Verwitterung des Kalifeldspates. Chem. Erde, 12: 1-22. Curtis D.M. (Editor), 1976. Sedimentary Processes: Diagenesis- S.E.P.M. Reprint Ser., 1: 216 PP. Daly. R.A.. 1917. Low-temperature formation of alkaline feldspars in limestones. Proc. Natl. Al.ad. Sci. U.S.. 3: 659-665.

101 Damuth, J.E. and Fairbridge, R.W., 1970. Equatorial Atlantic deep-sea arkosic sands and ice-age aridity in tropical South America. Geol. Soc. A m . Bull., 81: 189-206. Dangeard, L. and Rioult, M., 1961. Observations nouvelles sur les accidents silicieux situes au sommet de la “Pierre de Caen”. Bull. Soc. Geol. Fr., 3: 329-337. Dapples, E.C., 1959. The behavior of silica in diagenesis. In: Silica in Sediments-S. E . P . M . , Spec. Publ., 7: 36-54. Dapples, E.C., 1962. Stages of diagenesis in the development of sandstones. Bull. Geol. Soc. A m . , 73: 913-934. Davidson, C.F., 1964. Uniformitarianism and ore genesis. Mining Mag., 110: 176-185; 244-253. Davies, T.A. and Supko, P.R., 1973. Oceanic sediments and their diagenesis: some examples from deep-sea drilling. J. Sediment. Petrol., 43: 381-390. Deans, T., 1950. The Kupferschiefer and the associated lead-zinc mineralization in the Permian of Silesia, Germany and England. int. Geol. Congr.. lath, London, 1948, Rept., 7: 340-352. Debelmas, J., 1959. Une curieuse contribution a I’etude de la genese des silex. Trav. Lab. GPol. Fac. Sci. Univ. Grenoble, 35: 135-136. Debyser, J., 1952. Variation du pH dans l’epaisseur vase fluvio-marine. C. R . Acad. Sci.. 234: 741-743. Deffeyes, K.S., Lucia, F.J. and Weyl, P.K., 1965. Dolomitization of Secent and Plio-Pleistocene sediments by marine evaporite waters on Bonaire, Netherlands Antilles. In: L.C. Pray and R.C. Murray (Editors), Dolomitization and Limestone Diagenesis-S. E . P . M . Spec. Publ., 13: 71-87. Deflandre, G., 1936. Les flagelk fossiles. Acta Sci. ind., 335: 98 pp. Degens, E.T., 1967. Diagenesis of organic matter. In: G. Larsen and G.V. Chilingar (Editors), Diagenesis in Sediments. Elsevier, Amsterdam. pp. 343-390. Degens, E.T. and Epstein, S., 1962. Relationship between ’ s O / ’ 6 0 ratios in coexisting carbonates, cherts, and diatomites. Bull. A m . Assoc. Pet. Geol., 46: 534-542. Degens, E.T. and Ross, D.A. (Editors), 1969. Hot Brines and Recent Heavy Metal Deposits in the Red Sea - A Geochemical and Geophysical Account. Springer, Berlin, 600 pp. Degens, E.T. and Ross, D.A. (Editors), 1974. The Black Sea-Geology, chemistry and biology. A m . Assoc. Pet. Geol. Mem., 20: 633 pp. Demolon, A. and Boischot, P., 1948. Observations sur le cycle du phosphore dans la biosphere. C. R . Acad. Sci., 227: 655-656. Dimroth, E. and Kimberley, M.M., 1976. Precambrian atmospheric oxygen: evidence in the sedimentary distributions of carbon, sulfur, uranium, and iron. Can. J. Earth Sci., 13(9): I161- 1 185. Dixon, E.E.L., 1907-21. Notes on the geology of the South Wales Coalfield. Mem. Geol. Sum., Engl. Wales (pts. 7, 8, 13). Drever, J.I., 1971. Early diagenesis of clay minerals, Rio Ameca basin, Mexico. J . Sediment. Petrol., 41: 982-994. Drever, J.I. (Editor), 1977. Sea Water: Cycles of the Major Elements. (Benchmark Papers in Geologv, 45). Dowden, Hutchinson and Ross, Stroudsburg, Pa, 344 pp. Dury, G.H., 1974. Duricrusts. In: Encyclopedia Brifannica, 15th ed., 5: 1088- 1093. Dunham, K.C., 1952. Age relations of the epigenetic mineral deposits of Britain. Trans. Geol. SOC. Glasgow, 21: 395. Dunoyer de Segonzac, C., 1968. The birth and development of the concept of diagenesis. Earth Sci. R m . , 4: 153-201. 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.

102 Eitel, W., 1954. The Physical Chemistry of the Silicates. Chicago Univ. Press, Chicago, Ill.. 1592 pp. Erhart. H., 1956. La GenGse des Sols -Esquisse d ‘une Thiorie geologique et geochimique: Biostasie et Rhexistasie. Masson, Paris, 90 pp. Erhart. H., 1963. Sur le cycle de la d i c e hydratke dans la biosphere. C . R . Acad. Sci.,256: 373 1-3734. Erhart. H., 1973a. Itineraires Giochimiques et Cycle Giologique du Silicium. Doin, Paris, 217 PP. Erhart. H., 1973b. Itineraires Geochimiques et Cycle Gkologique de I’Aluminium. Doin, Paris, 253 pp. Fairbairn. H.W., 1950. Synthetic quartzite. A m . Mineral., 35: 735-748. Fairbridge, R.W., 1950a. Recent and Pleistocene coral reefs of Australia. J. Geol.. 58: 330-40 1. Fairbridge, R.W., 1950b. Pre-Cambrian algal limestones in Western Australia. Geol. Mag., 87: 324-330. Fairbridge, R.W., 1953. The Sahul Shelf, northern Australia; its structure and geological relationships. J. R . SOC.W. Austr., 37: 1-33. Fairbridge, R.W., 1957. The dolomite question. In: R.J. Le Blanc and J.G. Breeding (Editors), Regional Aspects of Carbonate Deposition - S . E. P . M . , Spec. Publ., 5 : 125- 178. Fairbridge, R.W., 1959. Statistics of non-folded basins. Publ. Bur. Centr. Seismol. Int., Ser. A , 20: 419-440. Fairbridge, R.W., 1961. Eustatic changes in sea level. In: L.H. Ahrens, F. Press, S.K. Runcorn and H.C. Urey (Editors), Physics and Chemistry of the Earth. Pergamon, London, 4: 99- 185. Fairbridge, R.W., 1964. The importance of limestone and its Ca/Mg ratio to paleoclimatology. In: A.E.M. Nairn (Editor), Problems in Paleoclimatology. Wiley, New York, N.Y., 2: 43 1-478. Fairbridge, R.W., 1967a. Carbonate rocks and paleoclimatology in the biogeochemical history of the planet. In: G.V. Chilingar, H.J. Bissell and R.W. Fairbridge (Editors), Carbonate Rocks. Elsevier, Amsterdam, pp. 399-432. Fairbridge, R.W., 1967b. Phases of diagenesis and authigenesis. In: G . Larsen and G.V. Chilingar (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, pp. 19-89. Fairbridge, R.W., 1975. Epidiagenetic silicification. In: 9me Congr. Ini. Sedirnentol., Nice, pp. 49-54. Fairbridge, R.W., 1980a. Thresholds and energy transfer in geomorphology. In: D.R. Coates and J.D. Vitek (Editors), Thresholds in Geomorphology. Allen and Unwin, London, pp. 43-49. Fairbridge, R.W., 1980b. The estuary: its definition and geodynamic cycle. In: E. Olausson and I. Cat0 (Editors), Chemistry and Biochemistry of Estuaries. Wiley, Chichester, Chapter 1, pp. 1-35. Fairbridge, R.W. and Bourgeois, J. (Editors), 1978. The Encyclopedia of Sedimentology. Dowden, Hutchinson and Ross, Stroudsburg, Pa., 901 pp. Fairbridge, R.W. and Finkl, C.W., Jr. 1980. Cratonic erosional unconformities and peneplains. J. Geol., 88: 69-86. Fairbridge. R.W. and Teichert, C., 1953. Soil horizons and marine bands in the coastal limestones of Western Australia. J. Proc. R . SOC.,N . S . W., 86: 68-87. Faust. G.T.. 1949. Dedolomitization and its relation to possible derivation of a magnesium-rich hydrothermal solution. A m . Mineral., 34: 780-823. Feely, H.W. and Kulp. J.L., 1957. The origin of Gulf Coast salt dome sulfur deposits. Bull. A m . Assoc. Per. Geol.. 41: 1802-1853.

103 Fersman, A.E., 1922. The Geochemistty of Russia, 1. Goskhimizdat, Leningrad. Flach, K.W., Cady, J.G. and Nettleton, W.D., 1968. Pedologic alteration of highly weathered parent materials. Trans. Znt. Congr. Soil Sci., 9th, Adelaide, 4: 343-35 1. Folk, R.W. and McBride, E.F., 1978. Radiolarites and their relation to subjacent “oceanic crust” in Liguria, Italy. J. Sediment. Petrol., 48(4): 1069-1 102. Folk, R.L. and Siedlecka, A., 1974. The “schizohaline” environment: its sedimentary and diagenetic fabrics as exemplified by late Paleozoic rocks and Bear Island, Svalbard. Sediment. Geol., 11: 1-15. Fondeur, C. et al., 1954. Quelques aspects de la dolomitisation au jurassique en France. Int. Geol. Congr., I9th, Algeria, C.R., Sect. 13, 15: 47 1-49 1. Frankel, J.J. and Kent, L.E., 1937- 1938. Grahamstown surface quartzites (silcretes). Trans. Geol. SOC.S . Ajr., 40: 1-43. Freidman, G.M. and Sanders, J.E., 1978. Principles of Sedimentologv. Wiley. New York. N.Y.. 792 pp. Fyfe, W.S., Turner, F.J. and Verhoogen, J., 1958. Metamorphic reactions and metamorphic facies. Geol. SOC.A m . Mem., 73: 259 pp. Garrels, R.M., 1960. Mineral Equilibria, at Low Temperature and Pressure. Harper and Row, New York, N.Y., 254 pp. Garrels, R.M. and Christ, C.L., 1965. Solutions, Minerals and Equilibria. Harper and Row, New York, 450 pp. (re-issued by Freeman, Cooper and Co., 1975). Garrels, R.M. and Mackenzie, F.T., 1974. Chemical history of the oceans deduced from post-depositional changes in sedimentary rocks. In: Studies in Paleo-0ceanograph.v S . E . P . M . Spec. Publ., 20: 193-204. Garrels, R.M., Thompson, M.E. and Siever, R., 1961. Control of carbonate solubility by carbonate complexing. Am. J. Sci., 259: 24-45. Geikie, A,, 1879. Outlines of Geology. Stanford, London, 424 pp. Gidon, P., 1959. Sur les phenomenes de silicification dans I’Hauterivien du Jura meridional. Trav. Lab. GPol. Fac. Sci. Univ. Grenoble, 35: 133-134. Gignoux, M. and Avnimelech, M., 1937. Genese de roches sedimentaires brechoi’des par “intrusion et kclatement”. Bull. Soc. Geol. Fr., 7: 27-33. Gimmelfarb, B.M., 1956. Fundamental geological laws governing phosphorite deposits and their genetic classification. Int. Geol. Congr., 20th, Mexico, 1956, R e p . , 90 (abstr.). Ginter, R.L., 1938. Sulfate reduction in deep sub-surface waters. In: A.E. Dunstan, A.W. Nash, B.T. Brooks and H. Tizard (Editors), The Science of Petroleum. Oxford Univ. Press, London, 1: 908. Glover, J.E., 1963. Studies in the diagenesis of some Western Australian sedimentary rocks. J. R . SOC. W . A . , 46: 33-56. Goldberg, E.D. and Arrhenius, G.O.S., 1958. Chemistry of Pacific pelagic sediments. Geochim. Cosmochim. Acta, 13: 153-212. Goldich, S.S., 1934. Authigenic feldspar in sandstone of southeastern Minnesota. J. Sediment. Petrol., 4: 89-95. Goldschmidt, V.M., 1954. Geochemistry. Clarendon Press, Oxford, 730 pp. Goldsmith, J.R. and Graf, D.L., 1958. Structural and compositional variations in some natural dolomites. J . Geol., 66: 678-693. Grabau, A.W., 1913. Principles of Stratigraphy. Seiler, New York. N.Y., 1185 pp. Graf, D.L. and Goldsmith, J.R., 1956. Some hydrothermal syntheses of dolomite and protodolomite. J. Geol., 64: 173- 186. Graham, E.R., 1941. Colloidal organic acids as factors in the weathering of anorthite. Soil Sci., 52: 291-295.

104 Griffin, G.M. and Ingram, R.L., 1955. Clay minerals of the Neuse River estuary. J . Sediment. Petrol., 25: 194-200. Griffin, J.J., Windom, H. and Goldberg, E.D., 1968. The distribution of clay minerals in the world ocean. Deep-sea Res., 15: 433-459. Grim, R., 1951. The depositional environment of red and green shales. J . Sediment. Petrol., 21(4): 226-232. Grim, R.E., 1958. Concept of diagenesis in argillaceous sediments. BuN. A m . Assoc. Pel. Geol., 42: 246-253. Grim, R.E., Bray, R.H. and Bradley, W.F., 1937. The mica in argillaceous sediments. A m . Mineral., 22: 813-829. Grim, R.E., Dietz, R.S. and Bradley, W.F., 1949. Clay mineral composition of some sediments from the Pacific Ocean off the California coast and the Gulf of California. Bull. Geol. SOC.A m . , 60: 1785-1808. Gruner, J.W. and Thiel, G.A., 1937. The occurrence of fine-grained authigenic feldspar in shales and silts. A m . Mineral., 22: 842-846. Gunatilaka, A., Saleh, A. and Al-Temeeni, A., 1980. Plant-controlled supratidal anhydrite from AI-Khiran, Kuwait. Nature, 228: 257-260. Hallam, A. and Bradshaw, M.J., 1979. Bituminous shales and oolitic ironstones as indicators of transgressions and regressions. J . Geol. SOC.London, 136: 157-164. Hanshaw, B.B., Back, W. and Deike, R.G., 1971. A geochemical hypothesis for dolomitization by ground water. Econ. Geol., 66(5): 710-724. Hart, R.A., 1973. A model for chemical exchange in the basalt-seawater system of oceanic layer 11. Can. J . Earth Sci., 10: 799-816. Hart, S.R., Erlank, A.J. and Kable, E.J.D., 1974. Sea-floor basalt alteration: some chemical and Sr isotopic effects. Contrib. Mineral. Petrol., 44: 219-230. Hatch, F.H., Rastall, R.H. and Black, M., 1938. The Petrology ofthe Sedimentary Rocks. Allen and Unwin, London, 3rd ed., 383 pp. Hay, R.L., 1962. Stratigraphic and zeolitic diagenesis of the John Day formation of Oregon. Univ. Calif. (Berkeley) Pub[. Geol. Sci., 42(5): 199-261. Hayes, A.O., 1915. Wabana iron-ore of Newfoundland. Geol. Sum. Can., Mem., 78: 163 pp. Heald, M.T., 1955. Stylolites in sandstones. J . Geol., 63: 101-114. Heald, M.T., 1965. Lithification of sandstones in West Virginia. W. Va. Geol. Econ. Surv. Bull., 30: 28 pp. Heath, G.R., 1974. Dissolved silica and deep-sea sediments. In: W.W. Hay (Editor), Studies in Paleo-Oceanography-S. E. P . M . Spec. Publ., 20: 17-93. Hecht, F., 1933. Der Verbleib der organischen Substanz der Tiere bei meerischer Einbettung. Senckenbergiana, 15: 165-249. Heezen, B.C., Nesteroff, W.D. and Sabatier, G., 1960. Repartition des mineraux argilleux dans les sediments profonds de 1’Atlantique nord et equatorial. C . R Acad. Sci., 251: 4 10-41 2. Herman, Y., 1972. Origin of deep-sea cherts in the North Atlantic. Nature, 238: 392-393. Hite, R.J., 1970. Shelf carbonate sedimentation controlled by salinity in the Paradox Basin, southwest Utah. Third Symp. Salt, North. Ohio Geol. SOC.,1: 48-66. (Reprinted in D.W. Kirkland and R. Evans, Marine Evaporites, Benchmark Papers in Geology, 7). Hoehne, K., 1957. Zur Entstehungsgeschichte der Floezverkieselungen in den unterrotliegenden Steinkohlen von Stockheim in Oberfranken und Manebach in Thiiringen. Geologie (Bedin), 6: 806-836. Holland, H.D., 1972. The geologic history of sea water-an attempt to solve the problem. Geochim. Cosmochim. Acta, 36: 637-65 1.

105 Holland, H.D., 1973. Systematics of the isotopic composition of sulfur in the oceans during the Phanerozoic and its implications for atmospheric oxygen. Geochim. Cosmochim. Acta, 37: 2605-2616. Holliday, D.W., 1970. The petrology of secondary gypsum rocks: a review. J . Sediment. Petrol., 40(2): 734-744. Holser, W.T., 1947. Metasomatic processes. Econ. Geol., 42: 384-395. Holser, W.T., 1977. Catastrophic chemical events in the history of the ocean. Nature, 267: 403-408. Holser, W.T. and Kaplan, I.R., 1966. Isotope geochemistry of sedimentary sulfates. Chem. Geol., 1: 93-135. Honess, A.P. and Jeffries, C.D., 1940. Authigenic albite from the Lowville limestone at Bellefonte, Pa. J. Sediment. Petrol., lO(1): 12- 18. Horstman, E.L., 1957. The distribution of lithium, rubidium, and caesium in igneous and sedimentary rocks. Geochim. Cosmochim. Acta, 12: 1-28. Hoss, H., 1957. Untersuchungen uber die Petrographie kulmischer Kieselschiefer. Beitr. Mineral. Petrogr., 6: 59-88. Hough, J.L., 1958. Fresh-water environment of deposition of Precambrian banded iron formations. J . Sediment. Petrol., 28: 414-430. Houtz, R. and Swing, J., 1976. Upper crustal structure as a function of plate age. J . Geophys. Res., 81(14): 2490-2498. Hummel, K., 1922. Die Entstehung eisenreicher Gesteine durch Halmyrose. Geol. Rundsch., 13: 40-81. Hunt, J.M., 1979. Petroleum Geochemistty and Geology. Freeman, San Francisco, Calif., 617 PP. Hunt, T.S., 1859. On some reactions of the salts of lime and magnesium, and the formation of gypsum and magnesium rocks. Am. J. Sci., 2nd Ser., 28: 377-381. Hunter, R.E., 1970. Facies of iron sedimentation in the Clinton Group. In: G.W. Fisher, F.J. Pettijohn and J.C. Reed, Jr. (Editors), Studies in Appalachian Geology: Central and Southern. Wiley, New York, N.Y., pp. lQ1-121. Hurley, P.M., Heezen, B.C., Pinson, W.H. and Fairbairn, H.W., 1963. K-Ar age values in pelagic sediments of the North Atlantic. Geochim. Cosmochim. Acta, 27: 393-399. Hutton, J., 1788. Theory of the earth. R . SOC.Edinburgh, Trans., 1: 209-304. Illing, L.V., 1959. Deposition and diagenesis of some Upper Paleozoic carbonate sediments in western Canada. World Pet. Congr., Proc., Sth, N.Y., 1959, I : 23-52. Illing, L.V., Wells, A.J. and Taylor, J.C.M., 1965. Penecontemporaneous dolomite in the Persian Gulf. In: L.C. Pray and R.C. Murray (Editors), Dolomitization and Limestone Diagenesis-S. E . P . M . , Spec, Publ., 13: 89-1 1 1 . Irving, A., 1892. Organic matter as a geological agent. Proc. Geol. Assoc. (Engl.), 12: 227-238. Jaanusson, V., 1961. Discontinuity surfaces in limestones. Bull. Geol. Inst., Univ. Uppsala, 40: 22 1-241. Jacobs, M.B., 1979. Nepheloid sediments and nephelometry. In: R.W. Fairbridge and J. Bourgeois (Editors), The Encyclopedia of Sedimentology. Dowden, Hutchinson and Ross, Stroudsburg, PA, pp. 495-498. James, H.L. and Sims, P.K., (Editors), 1973. Precambrian iron-formations of the World. Econ. Geol., 6(7): 913-914. Janecke, E., 1915. Die Entstehung der deutschen Kalisalzlager. Wissenschuft, 59: 109 pp. Jones, J.B. and Segnit, E.R., 1971. The nature of opal, 1. Nomenclature and constituent phases. J . Geol. SOC.Austr., 18: 57-68. Kalkowsky, E., 1880. Uber die Erforschung der archaeischen Formationen. Neues Jahrb. Mineral.. 1: 1-28.

106 Kastner, M., 1971. Authigenic feldspars in carbonate rocks. A m . Mineral., 56: 1403-1442. Kay, M., 1951. North American geosynclines. Geol. SOC.A m . , Mem., 48: i43 pp. Keller, W.D., 1964. Diagenesis in clay minerals-a review. Clays Clay Miner., Proc. Natl. Conf. Clays Clay Miner., 11: 136-157. Keller, W.D., Balgord, W.D. and Reesman, A.L., 1963. Dissolved products of artificially pulverized silicate minerals and rocks. J. Sediment. Petrol., 33: 191-204, 426-437. Kelley, W.P., 1939. Base exchange in relation to sediments. In: P.D: Trask (Editor), Recent Marine Sediments. Am. Assoc. Pet. Geol.. Tulsa, Okla., pp. 454-465. Kennedy, W.J. and Garrison, R.E., 1975. Morphology and genesis of nodular cherts and hardgrounds in the Upper Cretaceous of southern England. Sedimentology, 22: 31 1-386. Kessler, P., 1922. Uber Lochvenvitterung und ihre Beziehungen zur Metaharmose (Umbildung) der Gesteine. Geol. Rundsch., 12: 237-270. Kimberley, M.M., 1974. Origin of iron ore by diagenetic replacement of calcareous oolite. Nature, 250: 319-320. Kirkland, D.W. and Evans, R. (Editors), 1973. Marine Evaporites: Origin, Diagenesis and Geochemistry. (Benchmark Papers in Geology, 7). Dowden, Hutchinson and Ross, Stroudsburg, Pa, 426 pp. Knox, R.W.O’B, 1970. Chamosite ooliths from the Winter Gill ironstone (Jurassic) of Yorkshire. J. Sediment. Petrol., 40: 1216-1225. Kolodny, J., Taraboulos, A. and Friedlander, U., 1980. Participation of fresh water in chert diagenesis: evidence from oxygen isotopes and boron alpha-track mapping. Sedimentology, 27: 305-316. Krauskopf, K.B., 1956. Factors controlling the concentrations of thirteen rare metals in sea water. Geochim. Cosmochim. Actu, 9: 1-32b. Krauskopf, K.B., 1959. The geochemistry of silica in sedimentary environments. In: Silica in Sediments-S.E.P.M., Spec. Publ., 7: 4-19. Krauskopf, K.B., 1967. Introduction to Geochemistry. McGraw-Hill, New York, N.Y., 72 1 pp. Krejci-Graf, K., 1963a. Diagnostik der Salinitatsfazies der Oelwasser. Fortschr. Geol. Rheinl. Wesrfalen, 10: 367-448. Krejci-Graf, K., 1963b. Origin of oil. Geophys. Prospect., 1 l(3): 32 pp. Krumbein, W.C., 1942. Criteria for subsurface recognition of unconformities. A m . Assoc. Pet. Geol. Bull., 26: 36-62. Krumbein, W.C., 1947. Analysis of sedimentation and diagenesis. Bull. A m . Assoc. Pet. Geol., 31: 168-174. Krumbein, W.C. and Garrels, R.M., 1952. Origin and classification of chemical sediments in terms of pH and oxidation-reduction potentials. J. Geol., 60: 1-33. Land, L.S., 1973. Holocene meteoric dolomitization of Pleistocene limestones, North Jamaica. Sedimentology, 20(3): 41 1-424. Lane, A.C., 1909. Mine waters and their field assay. BuN. Geol. SOC.A m . , 19: 501-512. Lane, A.C., 1922. Weight of sedimentary rocks per unit volume. Bull. Geol. SOC.A m . , 33: 353-370. Lane, A.C., 1927. Calcium chloride waters, connate and diagenetic. Bull. A m . Assoc. Pet. Geol., 11: 1283-1305. Langford-Smith, T., 1978. Silcrete in Australia. Armidale, N.S.W. Univ. New England (Geogr. Dep.), 304 pp. Latimer, W.M., 1952. Oxidation Potentials. Prentice-Hall, New York, N.Y., 2nd ed., 392 pp. Lees, G.M., 1928. The chert beds of Palestine. Proc. Geol. Assoc. (Engl.), 39: 445-462. Leith, C.K. and Mead, W.J., 1915. Metamorphic Geology. Holt, New York, N.Y., 337 pp. Lemaztre, H., 1960. Etude du banc de silex d’une poche de dissolution des calcaires bathoniens des Aucrais (Calvados). Bull. SOC.Linn. Norm., 10 (1959): 63-65.

107 Lewin, J.C., 1961. The dissolution of silica from diatom walls. Geochim. Cosmochim. Acta, 2 1 : 182- 198. Linck, G., 1909. Uber die Entstehung der Dolomite. Monatsber. Dtsch. Geol. Ges., 61: 230-241. Linck, G., 1937. Bildung der Dolomite und Dolomitisierung. Chem. Erde, 11: 278-286. Love, L.G., 1963. Pyrite spheres in sediments. In: M.L. Jensen (Editor), Biogeochemistry of Sulfur Isotopes, 193 pp. Lucas, G., 1952. Premiers rtsultats d’une etude sur les products odorants des calcaires fetides. C . R . Acad. Sci., 234: 121-123. Lyell, C., 1833. Principlt?s of Geology. Murray, London, 3 vol. MacDonald, G.J.F., 1953. Anhydrite-gypsum equilibrium relations. A m . J . Sci., 25 1 : 883-898. MacFadyen, W.A., 1950. Sandy gypsum crystals from Berbera, British Somaliland. Geol. Mag., 87: 409-420. MacKenzie, F.T. and Garrels, R.M., 1966. Silica-bicarbonate balance in the ocean and early diagenesis. J . Sediment. Petrol., 36: 1075- 1084. Magaritz, M., Goldenburg, L., Kafri, U., and Arad, A., 1980. Dolomite formation in the seawater-freshwater interface. Nature, 287: 622-624. Manheim, F., Rowe, G.T. and Jipa, D., 1975. Marine phosphorite formation off Peru. J . Sediment. Petrol., 45: 243-25 1 . Marshall, C.E. and Upchurch, W.J., 1933. Chemical factors in cation exchange between root surfaces and nutrient media. Proc. Soil Sci. SOC.Am., 17: 222-227. Mawson, D., 1929. South Australian algal limestones in the process of formation. Geol. SOC. London, Q . J., 85: 613-621. McBride, E.F. (Editor), 1979. Silica in Sediments: Nodular and Bedded Chert. S.E.P.M. Reprint Ser., 8, Tulsa, Okla. McKelvey, V.E., Swanson, R.W. and Sheldon, R.P., 1953. The Permian phosphorite deposits of the western U.S.A. Congr. Gtol. Int., 19e, Alger, 1952, C . R . , 11: 45-64. Mennig, J..J. and Vatan, A,, 1959. Repartition des dolomies dans le Dinantien des Ardennes. Rev. Inst. Fr. Pet. Ann. Combust. Liq., .14: 517-534. Milliman, J.D., Miiller, G. and Forstner, U., 1974-1975. Recent Sedimentary Carbonates, 1, 2. Springer, Berlin. Millot, G., 1953. Heritage et neoformation dans la sedimentation argileuse. Congr. Geol. lnt., 19e, Alger, 1952, C.R., 18: 163-175. Millot, G., 1957. Des cycles sedimentaires et de trois modes sedimentation argileuse. C. R . Acad. Sci., 244: 2536-2539. Millot, G., 1970. Geology of Clays: Weathering, Sedimentology, Geochemistry. Springer, New York, N.Y., 429 pp. Millot, G., Radier, H., Muller-Feuga, R., Defossez, M. and Wey, R., 1959. Sur la geochimie de la d i c e et les silicifications sahariennes. Bull. Serv. Carte Geol. Alsace-Lorraine. pp. 33-48. Milne, I.H. and Earley, J.W., 1958. Effect of source and environment on clay minerals. Bull. A m , Assoc. Pet. Geol., 42: 328-338. Mizutani, S . , 1970. Silica minerals in the early stage of diagenesis. Sedimentology, 15: 4 19-436. Moore, E.S. and Maynard, J.E., 1929. Solution, transportation and precipitation of iron and silica. Econ. Geol., 24: 272-303; 365-402; 506-527. Moretti, F.J., 1957. Observations on limestones. J . Sediment. Petrol., 27: 282-292. Mosebach, R,, 1952. Wasserige HIS-Losungen und das Verschwinden kalkige tierische Hartteile aus werdenden Sedimenten. Senckenbergiana, 33: 13-22. Moulton, G.F., 1926. Some features of redbed bleaching. Bull. A m . Assoc. Per. Geol.. 10: 304-31 1; 636-637.

108 Murray, J. and Renard, A.F., 1891. Report on Deep-sea Deposits, Based on the Specimens Collected During the Voyage of H . M . S . “Challenger” in the Years 1872 to 1876. H.M. Stationary Office, London, 525 pp. Nagy, B., 1960. Review of the chromatographic “plate” theory with reference to fluid flow in rocks and sediments. Geochim. Cosmochim. Acta, 19: 289-296. Nanz Jr., R.H., 1953. Chemical composition of pre-Cambrian slates with notes on the geochemical evolution of lutites. J . Geol., 61: 51-64. Naumann, C.F., 1850. Lehrbuch der Geognosie. Leipzig. Newell, N.D., Fischer, A.G., Whiteman, A.J., Hickox, J.E. and Bradley, J.S., 1953. The Permian Reef Complex of the Guadalupe Mountains Region, Texas and New Mexico-A Study in Paleoecology. Freeman, San Francisco, Calif., 236 pp. Newhouse, W.H., 1927. Some forms of iron sulphide occurring in coal and other sedimentary rocks. J . Geol., 35: 73-83. Norin, E., 1955. The mineral composition of the Napolitan Yellow Tuff. Geol. Rundsch., 43(2): 526-534. Northrop, J.I., 1890. Notes on the geology of the Bahamas. Trans. N . Y. Acad. Sci., 10: 4-22. Oehler, D.Z., Schopf, J.W. and Kvenvolden, K.A., 1972. Carbon isotopic studies of organic matter in Precambrian rocks. Science, 175: 1246-1248. Ogniben, L., 1955. Inverse graded bedding in primary gypsum of chemical deposition. J . Sediment. Petrol., 25: 273-281. Ojakangas, R.W. and Keller, W.D., 1964. Glauconitization of rhyolite sand grains. J . Sediment. Petrol., 34: 84-90. Okamotu, G., Okura, T. and Goto, K., 1957. Properties of silica in water. Geochim. Cosmochim. Acta, 12: 123-132. Orr, A.P. and Moorhouse, F.W., 1933. Physical and chemical conditions in mangrove swamps. Great Barrier Reef Exped., 1928-1929. Bull. Br. Mus. Nut. Hist., 2(4): 102-110. Packham, G.H. and Crook, A.W., 1960. The principle of diagenetic facies and some of its implications. J . Geol., 68: 392-407. Passarge, S., 1904. Die Kalahari. Dietrich Reimer, Berlin, 822 pp. Paterson, M.N.A., Bien, G.S. and Berner, R.A., 1963. Radiocarbon studies of Recent dolomite from Deep Spring Lake, California. J . Geophys. Res., 68: 6493-6505. Perel’man, A.I., 1967. Geochemistry of Epigenesis. Plenum, New York, N.Y., 266 pp. Peterson, M.N.A. and Von der Borch, C.C., 1965. Chert: modern inorganic deposition in a carbonate-precipitating locality. Science, 149: 1501- 1503. Pettijohn, F.J., 1957. Sedimentary Rocks. Harper and Row, New York, N.Y., 2nd ed., 718 pp. Pettijohn, F.J., Potter, P.E. and Siever, R., 1972. Sand and Sandstone. Springer, Berlin, 618 PP. Pingitore, N.E., Jr., 1976. Vadose and phreatic diagenesis: processes, products and their recognition in corals. J . Sediment. Petrol., 46: 985- 1006. Piper, D.Z. and Codispoti, L.A., 1975. Marine phosphorite deposits and the nitrogen cycle. Science, 188: 15-18. Porrenga, D.H., 1967. Clay Mineralogy and Geochemistry of Recent Sediments in Tropical Areas. Stolk-Dort, Dordrecht, 145 pp. Pourbaix, M.J.N., 1949. Thermodynamics of Dilute Solutions. Arnold, London, 136 pp. Purser, B.H. (Editor), 1973. The Persian Gulf- Holocene Carbonate Sedimentation and Diagenesis in a Shallow Epicontinental Sea. Springer, New York, N.Y., 471 pp. Pustowaloff, L.W., 1955. Uber sekundare Veranderungen der Sedimentgesteine. Geol. Rundsch., 43(2) 535-550. Reed, R.D., 1928. The occurrence of feldspar in California sandstones. Bull. A m . Assoc. Pet. Geol.. 12: 1023-1024.

Reuling, H.T., 1934. Der Sitz der Dolomitisierung: Versuch einer neuen Auswertung der Bohr-Ergebnisse von Funafuti. Abh. Senckenberg. Nut. Ges. (Frankfurt), 428: 44 pp. Revelle, R. and Emery, K.O., 1957. Chemical erosion of beach rock and exposed reef rock. (Bikini and nearby atolls, Marshall Islands). U.S. Geol. Suru., Prof. Pap., 260-T: 699-709. Reynolds, D.L., 1929. Some new occurrences of authigenic potash feldspar. Geol. Mag., 66: 390-399. Riedel, W.R., 1959. Siliceous organic remains in pelagic sediments. In: Silica in SedimentsS . E . P . M . , Spec. Publ., 7: 80-91. Riley, C.M. and Byrne, J.V., 1961. Genesis of primary structures in anhydrite. J. Sediment. Petrol., 31: 553-559.. Rinne, F., 1920. Die geothermische Metamorphosen und die Dislokationen der deutschen Kalisalzlagerstatten. Fortschr. Mineral., Kristallogr. Petrogr., 6: 101 - 136. Rittenhouse, G., 1949. Petrology and paleogeography of Greenbrier formation: dolomite and dolomitic limestone. Bull. Am. Assoc. Pet. Geol., 33( 10): 1724-1728. Riviere A. and Visse, L., 1954. L‘origine des mineraux des sediments marins. Bull. SOC.GPol. Fr., 4: 467-473. Rona, P.A. and Lowell, R.P. (Editors), 1980. Seafloor Spreading Centers: Hydrothermal Systems. (Benchmark Papers in Geology, 56). Dowden, Hutchinson and Ross, Stroudsburg, PA, 424 pp. Ronov, A.B., 1964. Common tendencies in the evolution of the Earth’s crust. Geochem. l n f . , 4: 713-737. Rosenkrantz, A. and Rasmussen, H.W., 1960. Southeastern Sjiilland and Mon, Denmark. Int. Geol. Congr., Zlst, Copenhagen, 1960, Guide Excursions, A42 and C37. Ross, C.S., 1943. Clays and soils in relation to geologic processes. J. Wash. Acad. Sci., 33: 225-235. Ross, C.S. and Kerr, P.F., 1931. The kaolin minerals. U.S. Geol. Suru. Prof. Pap., 165E: I5 1- 175. Runnells, D.D., 1969. Diagenesis, chemical sediments, and the mixing of natural waters. J. Sediment. Petrol., 39: 1188-1201. Russell, K.L., 1970. Geochemistry and halmyrolysis of clay minerals, Rio Ameca, Mexico. Geochim. Cosmochim. Acta, 34: 893-907. Russell, R.J. and Russell, R.D., 1936. Lower Mississippi delta. Louisiana, Dep. Conseru., Geol. BUN., 8: 454 pp. Rutten, M.G., 1953. Sur la genese des depBts a Ammonites pyriteuses. C . R . Somm. SPances SOC.GPol. Fr., 305-308. Rutten, M.G., 1957. Remarks on the genesis of flints. A m . J . Sci., 255: 432-439. Salvan, H.M., 1955. A propos des formations silicieuses des phosphates marocains. Geol. Rundsch., 43(2): 503-515. Samsonowicz, J., 1948. Les Graptolithes et quelques nouveaux groupes d’animaux du Tremadoc de la Pologne. Palaeontol. Pol., 3: 235 pp. Sander, B., 195 1. Contributions to the Study of Depositional Fabrics. Rhythmically deposited Triassic Limestone and Dolomites - A m . Assoc. Pet. Geol., Spec. Publ., 207 pp. Sarin, D.D., 1962. Cyclic sedimentation of primary dolomite and limestone. J. Sediment. Petrol., 32: 45 1-47 1. Scheidegger, A.E., 1957. The Physics of Flow through Porous Media. Toronto Univ. Press, Toronto, Ont., 236 pp. Schlanger, SO., 1957. Dolomite growth in coralline algae. f. Sediment. Petrol., 27: 181 - 186. Schmalz, R.F., 1956. The mineralogy of the Funafuti drill cores and its bearing on the physicochemistry of dolomite. J. Paleontol., 30: 1004- 1005. Schrnitt, H., 1950. The genetic classification of the bed rock hypogene mineral deposits. Econ. Geol., 45: 671-680.

110 Scholle, P.A., 1971. Diagenesis of deep-sea carbonate turbidites, Upper Cretaceous, Monte Antola Flysch, northern Apennines, Italy. J . Sediment. Petrol., 41: 233-250. Selley, R.C., 1976. An Introduction to Sedimentology. Acad. Press, London, 408 pp. Shearman, D.J., Khouri, J. and Taha, S., 1961. On the replacement of dolomite by calcite in some Mesozoic limestones from the French Jura. Proc. Geol. Assoc. (Engl.), 72: 1- 12. Shelton, J.W., 1964. Authigenic kaolinite in sandstone. J. Sediment. Petrol., 34: 102-1 11. Sheppard, R.A. and Gude, A.J., 111. 1973. Zeolites and associated authigenic silicate minerals in tuffaceous rocks of the Big Sandy Formation, Mohave County, Arizona. U.S. Geol. Sum., Prof. Pap., 830: 36 pp. Shinn, E.A., Ginsburg, R.N. and Lloyd, R.M., 1965. Recent supratidal dolomite from Andros Island, Bahamas. In: L.C. Pray and R.C. Murray (Editors), Dolomitiration and Limestone Diagenesis-S.E.P.M. Spec. Publ., 13: 112-123. Siever, R., 1957a. The silica budget in the sedimentary cycle. Am. Mineral., 42: 821-841. Siever, R., 1957b. Chemical factors in carbonate-quartz cementation. Bull. Geol. SOC.Am., 68: 1795-1796. Siever, R., 1959. Petrology and geochemistry of silica cementation in some Pennsylvanian sandstones. In: Silica in Sediments-S. E . P . M . , Spec. Publ., 7: 55-79. Siever. R., 1979. Plate-tectonic controls on diagenesis. J . Geol., 87: 127- 155. Sillen, L.G., 1961. The physical chemistry of sea water. In: M. Sears (Editor), Oceanography. Am. Assoc. Adv. Sci. Publ., 67: 549-581. Skeats, E.W., 1903. The chemical composition of limestones from upraised coral islands, with notes on their microscopical structures. Bull. Harvard Coll., Mus. Comp. Zool., 62: 53-126. Skinner, H.C.W., 1963. Precipitation of calcian dolomites and magnesian calcites in the southeast of South Australia. A m . J . Sci., 261: 449-472. Smirnow, L.P., 1958. Black Sea Basin. In: L.G. Weeks (Editor), Habitat of O i l - A m . Assoc. Pet. Geol. Symp., pp. 982-994. Smit, D.E. and Swett, K., 1969. Devaluation of “dedolomitization”. J. Sediment. Petrol., 39: 379-380. Smulikowski, K., 1954. The problem of glauconite. Pol. Akad. Nauk, Komit. Geol., Arch. Mineral., 18: 21-109. Sorby, H.C., 1879. Anniversary address. Q. J . Geol. SOC.(London), 35: 56-95. Spencer, E., 1925. Albite and other authigenic minerals in limestone from Bengal. Mineral. Mag., 20: 365-381. Spiess, F.N. et al., 1980. The East Pacific Rise: hot springs and geophysical experiments. Science, 208: 1421-1443. Steidtmann, E., 1911. Evolution of limestone and dolomite. J. Geoi., 19: 323-345; 393-428. Steidtmann, E., 1917. Origin of dolomite as disclosed by stains and other methods. Bull. Geol. SOC.Am., 28: 431-450. Steinitz, G., 1970. Chert “dike” structures in Senonian chert beds, southern Negev, Israel. J . Sediment. Petrol., 40(4): 1241-1254. Steinitz, G., 1977. Evaporite-chert associations in Senonian bedded cherts, Israel. Isr. J . Earth Sci., 26: 55-63. Steinitz, G. and Kolodny, Y., 1978. Chert-porcellanite-phosphorite-chalk association. In: Sedimentology in Isreal, Cyprus, and Turkey, Guidebook, 2, Jerusalem, 10th Int. Congr. Sedimentol., pp. 275-305. Stephens, C.G., 1971. Laterite and silcrete in Australia. Geoderma, 5 : 5-52. Stern, K.H., 1954. Liesegang phenomenon. Chem. Rev., 57: 79-99. Stetson, H.C., 1933. The bottom deposits. Sci. Results of the “Nautilus” Expedition, 1931(5): 17-37.

111 Stevens, R.E. and Carron, M.K., 1948. Simple field test for distinguishing minerals by abrasion pH. A m . Mineral., 33: 31-50. Stevenson, F.J., 1960. Some aspects of the distribution of biochemicals in geologic environments. Geochim. Cosmochim. Acta, 19(4): 261-272. Stewart, F.H., 1949. The petrology of the evaporites of the Eskdale No. 2 boring, east Yorkshire, 1. Mineral. Mag., 28: 621-675. Stewart, F.H., 1965. The mineralogy of the British Permian evaporites. Mineral. Mag., 34: 460-470. Strakhov, N.M., 1957. Mkthodes d 'Etude des Roches Skdimentaires. Bur. Rech. Geol., Paris, 1: 542 pp.; 2: 535 pp. Strakhov, N.M., 1959. Schema de la diagenese des dep6ts marins. Eclogue Geol. Helv., 51: 76 1-767. Strakhov, N.M., 1967-69. Principles of Lithogenesis, 1, 2. Consultants Bureau (transl. from Russian, 1962), 245 and 609 pp. Stringham, B., 1940. Occurrence of feldspar replacing fossils. A m . Mineral., 25: 139- 144. Sugawara, K., 1934. Liesegang's stratification developed in the diatomaceous gyttja from Lake Haruna, and problems related to it. BUN. Chem. SOC.Jpn., 9: 402-409. Sujkowski, Z.L., 1958. Diagenesis. Bull. A m . Assoc. Pet. Geol., 42: 2692-2717. Sun, Ming-Shan and Allen, J.E., 1957. Authigenic brookite in Cretaceous Gallup Sandstone, Gallup, New Mexico. J. Sediment. Petrol., 27: 265-270. Surdam, R.C. and Parker, R.D., 1972. Authigenic aluminosilicate minerals in the tuffaceous rocks of the Green River Formation, Wyoming. Geol. SOC.A m . Bull., 83: 689-700. Taliaferro, N.L., 1934. Contraction phenomena in cherts. Bull. Geol. SOC.Am., 45: 189-232. Taliaferro, N.L., 1935. Some properties of opal. A m . J. Sci., 30: 450-474. Tamm, 0.. 1925. Experimental studies on chemical processes in the formation of glacial clay. Sver. Geol. Llnders. Arsb., 18(5): 1-20. Tarr, W.A., 1927. Alternative deposition of pyrite, marcasite, and possibly melnikovite. A m . Mineral., 12: 417-422. Taylor, J.H., 1949. Petrology of the Northampton Sand Ironstone Formation. Geol. Surv. G. B. Mem., 6: 111 pp. Teodorovich, G.I., 1949. Siderite geochemical facies of seas and saline waters in general as oil-producing. 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. Pet. Inst., Acad. Sci. U . S . S .R . , 5. (English transl. in Int. Geol. Rev., l(3): 50-74). Teodorovich, G.I., 1961. Authigenic Minerals in Sedimentary Rocks. Consultants Bureau, New York, N.Y., 120 pp. Termier, H. and Termier, G., 1963. Erosion and Sedimentation. Van Nostrand, London, 433 PP. Ters, M., 1961. La Vendee Littorale. Thkse, 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. A m . Assoc. Pet. Geol., 44: 1809- 1817. Thomson, A., 1959. Pressure solution and porosity. In: Silica in Sediments-S. E. P . M., Spec. Publ., 7: 92-1 10. Tissot, B., 1979. Effects on prolific petroleum source rocks and major coal deposits caused by sea-level changes. Nature, 277: 463-465.

112 Trendall, A.F., 1968. Three great basins of Precambrian banded iron formation deposition: a systematic comparison. Geol. SOC.A m . Bull., 79: 1527- 1544. Tugarinov, A.I. and Vinogradov, A.P., 1961. Geochronology of the Precambrian. Geochemistry ( U . S . S . R . )(English Transl.), 1961 (9): 789-800. Turner, P., 1980. Continental Red Beds. Develop. Sedimentol., 29, Elsevier, Amsterdam, 562 PP. Turner, F.J. and Verhoogen, J., 1960. Igneous and Metamorphic Petrology. McGraw-Hill, New York, N.Y., 2nd ed., 694 pp. Twenhofel, 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. McGraw-Hill, New York, N.Y., 2nd ed., 673 pp. Valeton, I., 1972. Bauxites. Elsevier, Amsterdam, 206 pp. Van Andel, T. and Postma, H., 1954. Recent sediments of the Gulf of Paria. Verh. K. Ned. Akad. Wet., Afd. Natuurk., Reeks I , 20(5): 288 pp. Van der Lingen, G.J. (Editor), 1977. Diagenesis of Deep-sea Biogenic Sediments. (Benchmark Papers in Geology, 40.) Dowden, Hutchinson and Ross, Stroudsburg, Pa., 385 pp. Van Hise, C.R., 1898. Metamorphism of rocks and rock flowage. BuU. Geol. SOC.Am., 9: 269-328. Van Hise, C.R., 1904. A treatise on metamorphism. U.S. Geol. Surv. Monogr., 47: 1286 pp. Van Houten, F.B., 1961. Climatic significance of red-beds, In: A.E.M. Nairn (Editor), Descriptive Palaeoclimatolou. 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. A m . J. Sci., 246: 569-572. Van Tuyl, F.M., 1916. The origin of dolomite. Iowa Geol. Surv., Annu. Rep. (for 1914), 25: 25 1-422. Van Tuyl, F.M., 1918. Depth of dolomitization. Science, 48: 350-352. Venkatarathnam, K. and Biscaye, P., 1973. Deep-sea zeolites in the sediments of the Indian Ocean. Mar. Geol., 15: MI 1-M17. Venkatarathnam, K. et al., 1976. Clay mineralogy and sedimentation in the western Indian Ocean. Deep-sea Res., 23: 949-961. Veizer, J., 1977. Diagenesis of pre-Quaternary carbonates as indicated by tracer studies. J. Sediment. Petrol., 47: 565-581. Visse, L.D., 1953. Les facies phosphates. Rev. Znst. F r a y . Petrole Combust. Liquides, 8: 87-99. Von Guembel, C.W ., 1868. Geognostische Beschreibung des ostbayerischen Grenzgebirges. Gotha, 968 pp. Von Morlot, A., 1847. Uber die Dolomit und seine kiinstliche Darstellung aus Kalkstein. Haidinger Nat. Abt., 1: 305. Wade, A., 1911. Some observations on the Eastern Desert of Egypt. Q.J. Geol. SOC.London, 67: 238-261. Waksman, S.A., 1933. On the distribution of organic matter on the sea bottom and the chemical 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.A m . , 73: 237-242. Walther, J., 1894. Einleitung in die Geologie als historische Wissenschaft. Fischer, Jena, 1055 PP.

113 Weaver, C.E., 1958, Geologic interpretation of argillaceous sediments. BUN. A m . Assoc. Pet. Geol., 42: 272-309. Weaver, C.E., 1960. Possible uses of clay minerals in search for oil. Bull. A m . Assoc. Pet. Geol., 44: 1505-1518. 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. Wermund Jr., E.G., 1961. Glauconite in early Tertiary sediments of Gulf coastal province. Bull. A m . Assoc. Pet..Geol., 45: 1667-1696. West, I.A., Ali, Y.A. and Hilmy, M.E., 1979. Primary gypsum nodules in a modern sabkha on the Mediterranean coast of Egypt. Geology, 7: 354-358. Wetzel, O., 1933. Die in organischer Substanz erhaltenen Mikrofossilien des baltischen Kreidefeuersteins. Paleontographica, 47: 146- 188. Wetzel, W., 1955. Seltene Metallverbindungen in Sedimenten. Geol. Rundsch., 43(2): 464-469. Weynschenk, R., 1951a. The problem of dolomite formation considered in the light of research on dolomites in the Sonnwend Mountains (Tirol). J . Sediment. Petrol., 21: 28-31. Weynschenk, R., 1951b. A new sedimentary petrological interpretation of the results reported by the Swedish Albatross Deep-sea Expedition 1947-48. J . Seqiment. Petroi., 21: 82-84. White, D.E., 1957. Magmatic, connate and metamorphic waters. Bull. Geol. SOC.A m . , 68: 1659-1682. Whitehouse, U.G. and McCarter, R.W., 1958. Diagenetic modification of clay mineral types in artificial sea water. In: Clays and Cloy Minerals. Natl. Acad. Sci./NRC, Publ., 566: 81-119. Wickman, F.E., 1944. Some notes on the geochemistry of the elements in sedimentary rocks. Ark. Kemi, Mineral. Geol., 19(B): 1. Williams, D., 1960. Genesis of sulphide ores. Proc. Geol. Assoc. (Engl.), 71: 245-284. Willman, H.B., 1942. Feldspar in Illinois sands. Ill. State Geol. Sum., Rep. Inoest., 79: 87 pp. Wilson, J.L., 1975. Carbonate Facies in Geologic History. Springer, Berlin, 471 pp. Zen, E-An, 1960. Carbonate equilibria in the open ocean and their bearing on the interpretation of ancient carbonate rocks. Geochim. Cosmochim. Acta, 18: 57-71. Zenger, D.H., 1972a. Significance of supratidal dolomitization in the geologic record. Bull. Geol. SOC.Am., 83: 1-12. Zenger, D.H., 1972b. Dolomitization and uniformitarianism. J . Geol. Educ., 20: 107- 124. ZoBell, C.E., 1942. Changes produced by micro-organisms in sediments after deposition. J . Sediment. Petrol., 12: 127-136. ZoBell, C.E., 1946. Studies on redox potential of marine sediments. Bull. A m . Assoc. Pet. Geol., 30: 477-513.

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115 Chapter 3

DIAGENESIS IN ARGILLACEOUS SEDIMENTS ARIEH SINGER and GERMAN MULLER

INTRODUCTION AND DEFINITIONS

The writers prefer 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 Read and 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”. In physicochemical terms, diagenesis can be regarded as the process of equilibration of mineral (mostly detrital) phases with aquatic environments of differing salinities, under conditions of increasing temperature and pressure. According to Friedman and Sanders (1978), diagenesis involves, among other things: (1) compaction; (2) addition of new material; (3) removal of material and transformation of material by: (4) change of mineral phase, or (5) replacement of one mineral phase by another. The changes undergone by particles from the time of arrival into the aqueous basin up to their lithification under conditions of high pressure and temperature in the nearly total absence of liquid water can be seen as a continuum. The accurate delimitation of those processes to be designated as “diagenetic”, presents, therefore, some difficulties. According to most European authorities, such as, for example, Millot ( 1970), diagenesis starts only after deposition of the particles, that is, processes occurring during their transport or suspension are excluded. Diagenesis passes into iow-grade metamorphism when recrystallization of sheet-silicates occurs at about 200°C (Dunoyer de Segonzac, 1970). 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 processes (e.g., due to insolation, frost, water, atmosphere, and. organisms) leading to the formation of soils. Whereas weathering commonly involves the loss of metallic cations, compensated by the gain of water, the inverse is the case with diagenesis. 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.”

116 The largest part of the products of weathering, together with a certain amount of unweathered rock materials, 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 alteration. 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, they can remain floating in the marine sedimentary basins for centuries because of the low settling velocity. 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. Inasmuch as many physicochemical and chemical processes are similar or even identical during the subaqueous transportation-alteration 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) environment is the term “halmyrolysis” introduced by Hummel (1922) (halmyros = salty; lysein = to dissolve). Thus, halmyrolysis can be defined as: all chemical and physicochemical processes which occur during the marine transportation-alteration 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 fresh-water environment is lacking, Muller ( 1967) proposed the term “ a q ~ a t o l y s i s(aqua ~ ~ = water; lysein = to dissolve) for all processes which take place in fresh water during transportation-alteration and in the earliest stage of diagenesis. Thus, aquatolysis may be defined as: all chemical and physicochemical processes which occur during transportation-alteration and pre-burial stage of diagenesis in fresh-water environment (Table 3-1). The various stages of diagenesis can be divided into: ( 1 ) pre-burial stage, (2) shallow-burial stage, and (3) deep-burial stage. The pre-burial and shallow-burial stages correspond to the “early diagenesis” of Dunoyer de Segonzac (1970), whereas the deep-burial stage is equivalent to his “middle” and “late” diagenetic stages. 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

117 TABLE 3-1

Relationships between parent rock, weathering, diagenesis and metamorphism SOIL

SEDIMENT

Fresh - water 0000000 Marine

Submorine basement

~0000000000000

1

' /

AQUATOLYSIS

1

+++++++

P A R E N T ROCK

HALMYROLYSIS

1

Fresh-water P r e - b u r i o l M a r i n e DIAGE N E SIS

1

METAMORPHISM

those from the Santa Barbara Basin and the Black Sea (Table 3-11). Although the sediments in the Santa Barbara Basin and the Black Sea, covered by about 5 m of younger deposits, are at least 10 to 100 times older 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: (1) 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

+

118 TABLE 3-11 Decrease of porosity with depth in subaqueous argillaceous sediments (After Rukhin, 1958. Emery and Rittenberg, 1952, and Ziillig, 1956, cited by Miiller, 1967, p. 163.) Depth (m)

Porosity (%) Black Sea

0.00 0.20 0.50 1.oo

2.00 3.OO 4.00 5.00 6.00 7.00 8.00 I

' Santa Barbara Basin, Calif. (U.S.A.)

Lake Zurich (Switzerland) 88

-

79 73

72 71 70 65

82 81 79 77 75

74 73

78 77 75 73 71 68 66 64 62 60

Porosities in percent calculated from water content

about 30%. At this porosity, the initial bulk volume of the clay mud decreases by about 50%, with loss of at least 50% of the connate water. Hamilton ( 1959), who studied the relationship between overburden pressure and porosity, observed that no matter how porous the original material, when the pressure reaches about 100 kg/cm2, the porosity usually decreases 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. For subaqueous argillaceous sediments, Strakhov ( 1956) 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 above 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. During the third stage of

119 diagenesis it can be assumed that the pore solutions in most cases have already become similar in composition. Inasmuch as detrital clays have commonly been formed during continental weathering under the impact of meteoric, non-saline water, diagenetic effects are much more pronounced in the highly saline marine environment. Diagenetic effects, however, can also be obtained in non-marine sedimentation basins of comparable solute concentration, such as saline lakes. In this chapter, “argillaceous sediments” are defined as fine-grained sediments (with an average grain size in the range of about 1-10 pm) 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) feldspar, (4) carbonates, ( 5 ) amorphous silica and alumina, (6) pyroclastic material, and (7) organic matter. In addition, biogenic carbonates, biogenic amorphous silica, and organic matter are formed in the basin itself. Clay minerals and quartz are most stable against changes in the physicochemical environment, whereas pyroclastic materials are the least stable. The important allogenic and biogenic components of modern clay sediments are described below. Clay minerals (including gibbsite)

The distribution of clay minerals in Recent ocean sediments has been shown by Biscaye (1969, Griffin et .al. (1968) and Rateev et al. ( 1969) to be chiefly detrital and related to general source areas on adjacent continents. In the tropics, where leaching and chemical weathering are intense, there is a conspicuous abundance of the kaolin group minerals and gibbsite near continental masses (Fig. 3-1). The kaolinite content in the sediments varies from 40 to 60% of the total clay mineral content. Areas of its maximum content occur in the equatorial parts of the Indian and Pacific oceans, with a gradual decrease in abundance towards the poles. Gibbsite is localized in the tropical and subtropical latitudinal belts. The zone of maximum concentration of smectite (40-60%) coincides in general with the zone of high kaolinite content, although it is somewhat wider. This general distribution pattern is modified by superposition of

120

121

patches of azonal concentrations associated with volcanogenic ash material. Thus, in the Atlantic, the occurrence of southern and northern maxima is due to the development of areas of Recent and Quaternary volcanism, such as near Iceland (Rateev et a]., 1969). In these areas, the clay mineral distribution appears to be determined by the ratio of detrital input to in-situ neoformation. Smectite concentrations are greater in the Southern Hemisphere oceans, where there is a larger input of volcanic material relative to detrital clays, because the former lends itself to alteration to smectite. The highest concentrations of smectite occur in the South Pacific, associated with phillipsite and volcanic glass shards. Maximum concentrations of chlorite (30%)are located in northern-southern zones, whereas equatorial zones are nearly devoid of this mineral. The distribution pattern of chlorite is thus distinctly zonal. For that reason, in recent deep-sea sediments of the North and South Atlantic the distribution of chlorite has been observed to be nearly reciprocal to that of kaolinite (Biscaye, 1965; Zimmermann, 1977). More illite occurs in the Northern Hemisphere ocean sediments than in the southern oceans, reflecting the impact of continental input in the former. Mite concentrations in marine sediments bear close relation to river-borne sediments, whereby the highest illite concentrations are found in the North Atlantic, in association with chlorite. The clearest latitudinal zonation in the distribution of illite is observed in the sediments of the Indian, and to a lesser extent, of the Pacific oceans. While maximum contents (60-80%) i n the Atlantic Ocean were observed in its northern parts, a near-equatorial zone of minimal contents is missing. Small areas of very high illite contents occur at the southern end of the African continent, probably representing the clay weathering products of the arid climates prevailing in parts of this continental mass. This review of clay mineral distribution in Recent marine argillaceous sediments has been limited to the most common occurrence patterns. Although local accumulations of relatively rare layer-silicates do exist (palygorskite, talc, pyrophyllite, serpentine, etc.), in such cases their occurrence Fig. 3-1. Map showing the distribution of kaolinite in the fine-pelitic fraction of the bottom sediments of the world oceans. I , 2 , 3 and 4 =kaolinite contents expressed as percentages of the total clay mineral contents. Stations where observations have been made, are marked with solid circles ( 5 ) in each ocean basin. On the continents, the distribution of the main weathering zones is shown: I =zone of moderately humid climate; I I =zone of tropical humid climate; I I I = tectonically active areas lacking well-formed weathering crusts; I V = for each continent, the total river load discharged, given in million of tons per year; V=thc amount of river load (in the same units) and the direction of discharge, at each of the main discharge points. The kaolinite percentages are as follows: I =40-60; 2 =20-40; 3 = 10-20: 4 = 10; 5=stations. (After Rateev et al., 1969, p. 28, fig. I.)

122 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 (Biscaye, 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 illi te-montmorillonite. Quurtz and feldspars

Quartz and feldspars (particularly plagioclase) are the most important allogenic minor constituents of argillaceous sediments. The abundance and distribution of quartz in Recent ocean sediments is directed primarily by aeolian input from adjacent desert regions (Rex and Goldberg, 1958; Windom, 1975; Windom and Chamberlain, 1978). Quartz in the Holocene sediments of the Indian Ocean is derived mainly from continental sources and has been transported in the form of atmospheric dusts from Arabian and Australian deserts into the adjacent ocean (Kolla and Biscaye, 1977). It appears that during the last glacial period the aeolian influx of quartz into the Indian Ocean was higher. An aeolian origin for feldspar in North Atlantic sediments was shown by Windom and Chamberlain (1978). The increase in quartz content near Antarctica in the Northern Atlantic probably stems from glacial outwash (Biscaye, 1964); a similar tendency for feldspars in these areas has also been generally noted and explained in the study of Biscaye (1964). 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. Curhonutes

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 (Gfobigerina ooze covers enormous areas in the Atlantic, Pacific and Indian oceans) all transitions are possible.

123 Amorphous silica and alumina

In 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% (Biscaye, 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 Biscaye (1964) 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 basalt. 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 Age, 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 two thirds of all active volcanoes are situated here. According to Sapper (cf. Brinkmann, 1961), between 1500 and 1914 about 18 km’ of lava and 330 km7 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 km7 of pyroclastic material were hurled up as high as 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 magnitude (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 definite 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

124 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. 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. In the Pacific, the palagonitk 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 should 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). Orgunic mutter

The contribution of Degens and Mopper to the companion Volume of this book (Chapter 4) covers the subject of organic matter composition and diagenesis in argillaceous sediments.

INITIAL (PRE-BURIAL) POROSITY A N D STRUCTURE OF ARGILLACEOIJS SEDIMENTS

Initial porosity

The initial porosity and thus the water content' 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 70-9056, corresponding to a water content of about 50-80%. In sands, porosity is only 30-50% which corresponds to 20-30% water content. Figure 3-2 shows the dependence of the water content of Recent clay muds on the amount of the clay fraction (< 2 pm) 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

'

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

125

90

80-

-r 60--

80

s

;

-m U

Y

c

C

’O

:40.

60

al

s;

c

v)

U L

$

.

I11

20.

V

0

/

50

/

al c

a

/

J 20

40 % Fraction

c

60 2pm

80

Fig. 3-2. Correlation between water content and percentage of clay fraction ( Mg > K > Na. Influence of exchangeable cations Experiments carried out by Samuels ( 1950) 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 A1 have larger water contents than Ca- and Na-saturated kaolinites, 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 Ca-montmorillonites and these, in turn, contain more water than Al-montmorillonites. For further discussions on this subject, see Meade (1 964). Rate of sedimentation Up to now no comprehensive studies were made on the dependence of initial porosity on the annual rate of sedimentation. Observations made by Fuchtbauer and 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 low 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 3-IV) 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.

128 TABLE 3-IV Water content of young lake sediments (after Ziillig, 1956, cited by Miiller, 1967, p. 161, table IV) Zuger See

Lake Constance

depth of burial (cm)

water content

calculated porosity

(48)

(48)

0 1.2 2.4 3.6

83.6 74.0 74.2 70.6

92 87 87 85

.

depth of burial fcm)

water content

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

calculated porosity

(%I 87 86 84 80 79

The combined effects of different factors in natural sediments

In natural argillaceous sediments, several clay minerals &cur 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. 3-2, 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 behavior 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 2 p m was determined, without making a further subdivision.

-=

initial structure

The dependence of primary porosity on particle size, type of clay mineral, electrolyte 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. Van Olphen (1963b, 1964) suggested that the interaction between plate-like clay-particles may result in three different modes of association: (1) adhesion between the flat oxygen faces of two parallel particles (face-to-face, FF); (2) adhesion between broken bond surfaces of neighboring particles (edge-to-edge, EE); (3) adhesion of broken bond

129 surfaces to flat oxygen faces (edge-to-face, EF). The EE and EF types of association lead to three dimensional voluminous ”card-house’’ structures and to aggregates that can be classed as flocs. The FF type of association, which is the slowest, leads to thicker and larger flocs, to which the terms “oriented aggregates”, “books”, or “domains” have been applied. Figure 3-3 shows this fabric in electronmicrographs made by Rosenquist (1962) of several marine clays. In the case of sedimentation in non-saline river or lake environments, if the electrolyte concentration is below the critical coagulation concentration of the FF mode of particle association, EE and EF modes are obtained, forming the open arrangement referred to as “card-house’’ (Yariv and Cross, 1979). This structure may later, under compaction, give rise to an oriented type of structure. In fresh-water clays, Rosenquist (1962) observed a greater degree of parallel orientation between clay particles. At electrolyte concentrations above the critical coagulation concentration,

Fig. 3-3. Electronmicrograph showing mutual arrangement OF minerals in the blue Oslo clay. (After Rosenquist, 1962, p. 18, fig. 5.)

130

Fig. 3-4. Two-dimensional model of the aggregate structure. (After Von Engelhardt and Gaida, 1963, p. 927, fig. 1 1 . )

face-to-face (FF) coagulation becomes dominant. Therefore, sedimentation in a marine environment, where salt concentration is above the critical coagulation level and the FF mode of particle association becomes more stable than either the EE or EF mode, individual platelets aggregate together in “books” or “domains” and the “books” then form an open arrangement referred to as “book-house” (Yariv and Cross, 1979). From the behavior of permeability during the compression of clays with different electrolyte contents in their pore solutions, Von Engelhardt and Gaida (1963) concluded that clays which settled from solutions rich in electrolytes have a heterogeneous or aggregate structure (Fig. 3-4). 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 and pre-burial stage (aquatolysis and halmyrolysis)

A quatolysis

The clay minerals carried by rivers and deposited in lakes and oceans are

131 derived from soils, or directly from the outcropping sedimentary, igneous and metamorphic rocks in the hinterland (see Table 3-1). 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, which are suspended in the Alpenrhein (Rhine river before it enters Lake Constance) and deposited in 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. Muller, 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 (1) illite (dioctahedral) from muscovite, and (2) ledikite (trioctahedral) from biotite. I t is probable that after deposition these processes also c,ontinue during the pre-burial stage. Possibly, aquatolysis also occurs in the case of K-rich feldspars. Pedro et al. ( 1978) described diagenetic peloidal nontronite in surface sediments of Lake Chad. They showed that nontronite is obtained by the transformation of oolitic sand-sized Fe-oxide grains, in a process that begins in the interior of the grain and proceeds outward. They suggested that low silica activity in solution, having near-neutral pH and near-zero Eh, favor the process that is analogous to the ane responsible for chamosite and glauconite formation from peloidal material. The coprecipitation of silica from lake waters with hydrated oxides of iron and manganese has been demonstrated by Kato (1969), thus explaining the correlation of some lacustrine silica levels with the redox cycle in the lakes. Nontronite neoformation in Recent sediments of Lake Malawi has been explained by Muller and Forstner (1973) as resulting from the percolation of geothermal solutions rich in SiO, through the sedimentary fill of the basin, and their reaction with ferrous iron. Under reducing conditions and a pH lower than 7, iron and manganese are leached into the lake water. In more highly aerated areas of the lake, iron hydroxides and opal precipitate simultanously. Experimental synthesis of iron layer-silicates led Harder ( 1978) to suggest that low-temperature formation of iron layer-silicates is feasible under reducing conditions within relatively short spans of time in the presence of soluble iron and silica. In the clay minerals already derived from soils or sediments, or other layer-silicates of the metamorphic rocks, no considerable change is to be expected to occur during the subaqueous transportation.

132 Halmyrolysis When terrigenous clay minerals are transported to the sea, probably the first change to result from the fresh-water-saline-water transition is ion exchange. Russell (1970) suggested that ion exchange is the only rapid (time scale of a year or less) reaction occurring when land-derived clays encounter sea water. Exchangeable Ca2+ on the clays is exchanged for M g 2 + , K + and Naf from the sea water. Equilibration of the clay adsorption complex with the sea water appears to take place fairly rapidly. Experimental data obtained by Roberson (1974) show that after a few hours, and at most a few days of immersion in sea water, four expanding clays had exchanged most of their original interlayer cation Ca2+ for M g 2 + , K + and Na+ . Thereafter, there appears to be very little, if any, additional reaction. The ratio of adsorbed Mg2+ to adsorbed K varies between about 1 : 1 and 5 : 1 for these samples (Fig. 3-5). A large proportion (up to 40%) of the interlayer cations become non-exchangeable after prolonged contact with sea water. Results

a Poitras Sornple

401

30 20

L

-0>

~

2

m

, , , ;, 4 6 8 1 0 1 2 Time, w k s

%

2

4 6 8 Time, wks

1012

rn

0

0

0

0

i ,aP \

\

m

E c

Y)

; 70 D 0

9

Mg

0

0

V

60

V

D

50-

b Oquoga Somple

V

c:

d Loidig Somple

zziy,;o

40-

10

2

4 6 8 Time, wks

1012 Time, wks

Fig. 3-5. Cation adsorption in sea water as a function of time. (After Roberson, 1974, p. 443, fig. I . )

133 obtained by Russell (1970) indicate that over two thirds of the loss in exchange capacity can be accounted for by the fixation of K + . Also, some of the Na’ and Ca2+ cations become fixed. The decrease in exchange capacity is associated by Roberson (1974) with the formation of a poorly crystalline mixed-layer phase, consisting of illite-like expansible layers. Possibly. this phase can be regarded as a precursor for discrete illite layers. The decrease in expansibility of clays upon contact with sea water may be more pronounced with “degraded illite”, that is a potassium-depleted variety of illite, in which a large amount of K + had been replaced during weathering by exchangeable cations. Cation-exchange processes continue after deposition. A recent study of Atlantic Ocean sediments indicates that the reactions controlling the fluxes of most components across the water-sediment interface occur almost exclusively in the upper 100 cm of sediment (Sayles, 1979). Diagenesis has led to the uptake of Mg2+ and K + and the release of Ca2+, HCO, and Na+ by the solid phases. Contributions of M g 2 + , Ca2+, K + and HCO; from below 100 cm amount to less than 15% of the calculated fluxes across the interface. Reactions in the upper 30 cm account for 70-90% of the fluxes of the components across the interface. Only reactions involving Na may have a deeper source. Contrary to earlier views, it appears highly unlikely that much dissolved SiO, is sorbed on clay-mineral surfaces in the shallow areas of the continental shelf (Siever and Woodford, 1973). In some of the deeper waters of the oceans, however, or in the interstitial water of deposited sediments in areas of diatom ooze formation, dissolved silica levels might be high enough for sorption on clay surfaces. Can the newly acquired cation penetrate the silicate lattice of the clay mineral? Drever (1971a) has shown that in the Recent sediments of h o Ameca Basin, Mexico, the non-exchangeable Mg-content of the clay fraction is higher and the Fe-content is lower in sediments from strongly reducing environments than in similar sediments from less reducing environments. The mineralogy of the clay fraction did not show any parallel changes. According to Drever, in a strongly reducing environment, Fe leaves the smectite structure, to form a sulfide, and Mg enters the same sites from the water, so that gross clay-mineralogy remains unchanged. Commenting on this proposed process, Perry et al. (1976) have shown that addition of Mg by ion exchange can account for at most 50% of the dissolved Mg flux from rivers and ground water. They, therefore, proposed an additional mechanism of Mg removal from solution as a result of basalt alteration to smectite (see below). In an additional comment, Heller-Kallai and Rozenson ( 1978) indicated that direct Mg + Fe exchange in clays seems improbable, but that depletion of Mg in interstitial waters of anoxic sediments may be due to reaction of Mg-containing solutions with partially disintegrated clay.

134 Dunoyer de Segonzac (1970), following the concept of aggradation developed by Millot (1964), described the process whereby clay minerals that have been deprived of some of their cationic components during continental weathering (“degradation”) take up the same or similar cations on contact with sea water. Occurrences reviewed by Millot include mainly the formation of illite and chlorite from disordered mixed-layer structures. The major process involved is the adsorption of K + and Mg2+ into the interlayers. As early as 1956, Whitehouse and McCarter have shown by experimental studies that smectites exposed to artificial sea water for prolonged periods, were altered to yield illite-like and chlorite-like clay minerals. N o alterations whatsoever were produced in the original kaolinitic and illitic clay structures. From a consideration of the large amounts of Al, Fe and Si carried in solutions by fresh water into brackish or marine water, Jeans (1971) postulated neoformation or diagenesis of clay minerals in the very earliest stages of deposition and even before burial. Pre-burial diagenesis of clay minerals is also suggested by the comparison of adjacent deltaic and marine environments of deposition. The formation of metabentonites (K-bentonite) is relegated by Dunoyer de Segonzac (1970), at least partially, to the pre-burial stage of diagenesis. He concedes, however, that a considerable amount of evidence indicates burial diagenesis also for this type of clay mineral. The absence of extensive structural changes in clay minerals during halmyrolysis is possibly due to inhibition by dissolved organics which block interlayer sites otherwise available for uptake of cations (Berner, 1971, p. 184). Summing up, the relative importance of pre-burial diagenesis in the sense of clay-mineral lattice transformations, resulting from contact with sea water of normal salinity, appears to be only very minor. On the other hand, many significant clay-mineral transformations that can be attributed to early diagenesis have very probably resulted from the effect of hypersaline (and possibly warm) marine solutions on detrital clay minerals. Halmyrolysis in the supersaline marine environment During evaporation of sea water, the already high salt content is increased further. Increases in the salt content of sea water can also occur as a result of other phenomena such as hydrothermal activity. When a mineral passes from fresh into supersaline water, the radical change in environment could be expected to result in more significant mineral alterations than the freshwater-sea-water transition. The frequent occurrences of several clay-mineral types in salt clays appear to confirm this hypothesis. Corrensite, a regular mixed-layer chlorite-smectite is most common in the salt-rich Permo-Triassic deposits and is generally believed to be an intermediate step in the evolution of illites towards chlorite by fixation of magnesium. While Lucas and Ataman ( 1968) suggested that this evolution

135 occurred in the Triassic of the French Jura mainly during transport of the particles and before their deposition and burial, evidence summarized by Dunoyer de Segonzac (1970) indicates that this aggradation process could most probably be attributed to the action of highly saline solutions obtained by the gradual confinement of the Triassic basin. The aggradation process may have started during deposition and continued during shallow burial. From its appearance in sediments associated with hard beds, chlorite in Triassic sediments from southern Israel was proposed to have formed diagenetically through the sustained contact of detrital illite with constantly replenished saline water, that may have been of a higher salinity, percolating through the permeable rocks, perhaps in the inter- or supratidal zone (Heller-Kallai et al., 1973). Following diagenetic stages for this occurrence are proposed: dioctahedral illite dioctahedral illite-smectite dioctahedral smectite-di- and/or trioctahedral vermiculite-chlorite di- and/or trioctahedral chlorite. Chlorite, which is relatively abundant in some Messinian sediments from the Mediterranean area is considered by Chamley et al. (1978) to be of diagenetic origin for the following reasons: (a) proximity of the chlorite to sub-regular mixed-layer chlorite-smectite (corrensite) suggesting the initial stages of a “chloritization” of detrital minerals: (b) association with Mg-rich sediments such as authigenic dolomite: and (c) association with authigenic palygorskite. Chlorite and illite are the most abundant clay minerals in Paleozoic evaporite rocks (Droste, 1963). The origin of this association is controversial. It could represent not only primarily detrital accumulations, but also transformation products of detrital clay assemblages. Chlorite and illite are also the major clay minerals in silicate assemblages from an Upper Silurian rock salt bed in New York, U.S.A. (Bodine and Standaert, 1977). Textural features and Br-content of the salt indicate precipitation from shallow-marine brine. A relatively uniform chemical composition of the chlorite, that is distinctly different from that of normal chlorites, suggests an authigenic origin in the marine-evaporite environment (Fig. 3-6). The illite is clearly less degraded than normal shale illite, suggesting that some recrystallization occurred in the hypersaline environment. Postdepositional (and/or postburial) diagenesis resulted in improved crystallinity that involved isochemical recrystallization of the bulk silicate assemblage. The term “hyperhalmyrolysis” is introduced by Bodine and Standaert to denote mineral reactions which occur in the marine hypersaline environment. +

- -

Halmyrolysis in the supersaline lake environment Zeolites are the best known diagenetic products in saline lakes. Their occurrences have recently been summarized by Eugster and Hardie (1978). and Surdam and Sheppard (1978). They form through reaction of volcanic

136 1.0

++ N

2 0.4

N

d

0.2

00 20

22

24

26

20

30

32

3.4

36

Si (atoms per four tetrahedral sites)

Fig. 3-6. Compositional classification of the chlorite minerals with heavy boundaries outlining the probable limits of most “shale chlorite” compositions as determined ftrom X-ray diffraction data. o= stoichiometric end-member Mg-clinochlore. Chlorites associated with marine of four clinochlore samples from Retsof, N.Y.; =chlorites from evaporites: .=average other marine evaporite localities: I = clinochlore from the German Zechstein, KonigshallHindenburg; 2 Zpenninite from the German Zechstein, (?) Werra; 3 =clinochlore from an altered melaphyre, Austrian Haselgebirge, Hallstatt. (After Bodine and Standaert, 1977, p. 64, fig. 4.)

glass, deposited within the lake basins, with the lake brine. A zonal arrangement of the various diagenetic products reflects lateral salinity gradient. The peripheral areas are characterized by unaltered glass. They border on a zeolite zone, to be followed by analcime, and, in the very center, by potassium feldspar (Eugster and Hardie, 1978). This zonation can be observed in the Pleistocene Lake Tecopa, California (Fig. 3-7). In Lake Magadi, studied by Surdam and Eugster (1976), the Na-zeolite erionite is the most common diagenetic product. Depending on the nature of the dissolved salts, other zeolites, such as clinoptilolite or phillipsite have been observed to form in other lake sediments. The conversion of erionite into analcime has been represented by Surdam and Eugster as follows:

3 H,O Nao.5Ko,5Si3.509.

+ 0.5 Na+

+

Magadi-erionite

NaAlSi,O,. H,O

+ 0.5 K t + 1.5 SiO, + 2 H,O

molcimr

Analcime, that represents an environment of higher salinity, may also

137 116' 20'

116"lO'

Fig. 3-7. Mineral zonation in the tuffs of Pleistocene Lake Tecopa: dots ( I )=fresh glass; circles (2)=zeolites; lines (j)-K-feldspar. (After Eugster and Hardie, 1978, p. 254, fig. 10.)

form directly from volcanic glass, or from Na-Al-silicate gels. In the most saline environments, K-feldspars replace analcime as the common diagenetic product. Other common diagenetic products of saline alkaline lakes are bedded cherts. These appear to have formed from a sodium-silicate precursor, magadiite [NaSi,O,,(OH), - 3 H,O], first described by Eugster (1967) from Lake Magadi, Kenya. Some evidence appears to suggest that claymineral diagenesis may also take place in saline lake environments. Oolites composed principally of stevensite in the Eocene Green River Formation from Central Utah are believed to have formed diagenetically by precipitation from solution in the ancient Green River Lake (Tettenhorst and Moore, 1978). Earlier, diagenetic trioctahedral smectite has been reported to occur in the same formation (Dyni, 1976). The Mg-rich clay minerals such as Mg-montmorillonite, hectorite, vermiculite, and mixedlayers ( 14M-14c, 14,,,-14$,, and 14v-14c) are attributed to diagenetic neoformation in clayey Neogene sediments deposited in saline lacustrine environments in Turkey, in association with borates, silicates, and carbonates (Ataman and Baysal, 1978). The possibility of a smectite + illite conversion in the Tilton Shale Member of the Eocene Green Rtver Formation has been considered by Tank (1969). The vertical clay mineral distribution pattern in two sediment cores taken from two East African lakes distinctly suggests that illite may have formed diagenetically by the reaction of a smectite precursor with highly saline lake brines (Stoffers and Holdship, 1975; Stoffers and Singer, 1979; Singer and Stoffers, 1980). Existence of saline paleobrines is evidenced by

138 the presence of zeolites and protodolomite in sediments. In one lake sediment, illitization of smectite appears to have been favored by an unusually high K/N ratio in the lake water. The diagenetic illitization of smectite in the second lake sediment may have resulted from the parallel process of analcime formation, with the necessary K supplied by the volcanic glass + K, Na-zeolite + analcime conversion process (Fig. 3-8). Diagenetic palygorskite and sepiolite Whereas many palygorskite and sepiolite occurrences in marine and lake sediments can be traced back to continental origins, some are beyond doubt of diagenetic origin. The diagenetic formation of palygorskite in marine sediments has been documented extensively by: (1) Weaver and Beck (1977) for the Miocene deposits of the southeastern United States; (2) Couture ( 1977) and Church and Velde (1979) for Pacific Ocean occurrences; and (3) Timofeev et al. (1977) for the Atlantic Ocean sediments. The subject has been reviewed by Singer ( 1979). Smectite and volcanic glass are frequently mentioned as precursor minerals for palygorskite and sepiolite formation. In other occurrences the minerals appear to have formed by precipitation from solution. Requirements for their diagenetic formation appear to include high (alkaline) pH, high Si and Mg activity and low A1 activity. These requirements possibly are met in specific situations, such as: (a) near sites of hydrothermal activity; (b) in peri-marine, shallow-water environments, close to land-masses undergoing intensive desilicification by weathering; or (c) in response to fluctuations in ocean-water temperature that affect solubility levels of limiting chemicals, such as Si. The formation of palygorskite and sepiolite probably extends from the pre-burial stage well into the shallow-burial stage. The formation of glauconite “La glauconie est caracteristique du milieu marin: ceci appartient a SMECTITE GLASS

SMECTITE INTERLAYERS ILLITE

ILLITE

SOLUTION CHABAZITE ERlONlTE

ANALCIME

b

INCREASING SALINITY

AND

ALKALINITY

Fig. 3-8. Diagrammatic representation of the processes leading to the diagenetic illitization of srnectite and the formation of analcime from volcanic glass, with increase in water paleosalinity and paleoalkalinity of Lake Manyare, East Africa. (After Singer and Stoffers, 1980.)

139 l’alphabet de tout geologue” (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 of Trinidad. According to Seibold (1964), the green muds off Guinea at a depth of about l00m can contain up to 50% glauconite. Since the Paleozoic period, glauconite is found in almost all formations. The major mineral in glauconite pellets is an iron-rich, mixed-layer illite-smectite, analogous to aluminous illite-smectite. The nature of the interlayering varies with the properties of the layer types, i.e., from randomly interstratified to highly ordered. Glauconite differs from aluminous illitesmectite in that glauconite contains considerably less potassium per illite layer than does aluminous illite-smectite with the same proportion of illite layers, except near the pure illite composition (Thompson and Hower, 1975). Glauconite pellets frequently contain two textural forms of glauconite (aggregate and oriented), which differ in both mineralogy and chemistry. The aggregate glauconite composes the bulk of most pellets and consists of crystals in a nearly random arrangement. The oriented glauconite occurs primarily as rims on the periphery of pellets and has a honeycomb-like structure (Odom, 1976). Glauconite is formed diagenetically in marine sediments in reducing environments, primarily in fossil-rich carbonate sediments. Many processes have been proposed for the formation of the various mineralogical forms of glauconite. The most widely accepted is the “layer-lattice” theory developed by Hower (1961) after Burst (1958). The process of glauconite formation involves the absorption of K and Fe by any degraded detrital layer-silicate structure under suitable chemical and physical conditions. During the “glauconitization” process, the number of expandable layers is reduced and after considerable time the mineral glauconite is produced. According to Birch (1979) and Birch et al. (1976), illite clay-minerals are the precursors proto-glauconite") for glauconite pellets from the continental margins off the west coast of South Africa. Their data suggest that Fe is emplaced into the clay structure very early in the glauconitization process, possibly by a mechanism which is independent of, and prior to, the fixation of K . Odom (1976) suggested that many forms of glauconite pellets grow by the development of smectite in successive stages or layers, which then become oriented glauconite and later aggregate glauconite. The initial growth of smectite is probably related to decaying organic material in an environment with favourable pH, Eh, and physical conditions. The chemical composition

140 of the crystallized smectite and the rate of its development during successive stages of growth might be somewhat variable in different environments. Kohler and Koster (1976) cited the following proposed precursors for diagenetically formed glauconite: (a) amorphous gels, (b) clastic rock and mineral relics, (c) biotite and other micaceous minerals, and (d) montmorillonite. The detailed chemical and mineralogical investigation of glauconite from ten Cretaceous sediments leads the authors to conclude that the most likely precursors for glauconite formation are metal hydroxides in the form of gels, in addition to amorphous silica of primarily organic origin. Formation of glauconite from clastic material is not likely, because of the very minor amounts of these materials commonly associated with glauconite. Micaceous minerals or montmorilloni te would have to undergo drastic crystallographic transformations in order to change into glauconite and are, therefore, also ruled out as glauconite precursors. Possibly, the nature of the initial precursor materials, either iron-rich or iron-poor, are not of great importance for the “glauconitization” process. According to Velde and Odin (1975) and Odin (1978), pelletal glauconites are the result of the imposition of a chemical gradient upon sediments in a distinct spatial localization. The evolution of sedimentary environment towards that which is favorable for the formation of glauconite necessitates high iron contents and increasing K contents as more illite-like phases are produced. As a result, physicochemical properties of the pelletal aggregate, such as Eh and porosity, may constitute important factors. On the other hand, no distinct “parental” relations exist between the detrital precursor minerals and the diagenetic glauconite mineral. Formation of marine zeolites Bonatti’s (1963) studies in the Pacific show that phillipsite “is one of the most abundant mineral species of the upper layers of the earth’s crust”. In extensive areas of the Pacific Ocean, phillipsite concentrations are greater than 50%. Zeolites are among the most important diagenetic minerals in fine-grained pelagic sediments. Phillipsite and clinoptilolite are two common zeolites in deep-sea sediments. Another zeolite which occurs in significant amounts in deep-sea sediments, is analcime. Phillipsite is associated with argillaceous, volcanic, and siliceous sediments. It occurs in areas and sediments having slow sedimentation rates, at shallow depths in the sediment, and in very young sediments; it is most commonly associated with smectite (Stonecipher, 1976; Kastner and Stonecipher, 1978; Houghton et al., 1979). Deep-sea phillipsite from several Pacific core sediments appears to form near the sediment-water interface, incorporating elements from adjacent

141 mineral particles and interstitial sea water. Concentrations of Na, K, Rb, Ca and Sr that are quite uniform with sediment depth, suggest a common origin from the deep-sea alteration of volcanic material by pore water that contains excess amounts of these soluble cations (Bernat and Church, 1978). Phillipsite represents one of the principal products of the halmyrolytic alteration of basaltic glass, accompanying smectite (Honnorez, 1978). The initial stage of palagonitization of basaltic glass is characterized by the crystallization of an intergranular Na, K-phillipsite with saponite low in Ca and, probably, K-rich. During the mature stage of palagonitization the palagonitized glass granules are replaced in situ by intergranular phillipsite and, possibly, very minor amounts of Fe-rich saponite or Mg-bearing nontronite. At the final stage of palagonitization, the hyaloclastite has been completely replaced by an intimate mixture of authigenic K, Na-phillipsite, with almost no Ca, smectite, and Fe-Mn oxides. While phllipsite is dominant in younger deep-sea sediments, clinoptilolite is most abundant in Eocene and Cretaceous samples (Boles and Wise, 1978). Over 80% of the clinoptilolite occurrences are reported in brown clays and microfossil-rich sediments rather than in the ash beds, suggesting that volcanic ash is not a prerequisite for its deep-sea formation. Possibly, phillipsite forms as a metastable, silica-deficient phase in marine pore fluids and is eventually replaced by clinoptilolite. Time appears to have a more important control on this reaction than burial depth of silica concentration in pore fluids. Some clinoptilolite may also have formed from the dissolution of siliceous microfossils and clay minerals in the absence of volcanic glass. Houghton et al. (1979) noted the frequent association of clinoptilolite with alkali feldspar in the western North Atlantic. Nathan and Flexer (1977) compiled DSDP data on clinoptilolite occurrences and came to the conclusion that the mineral is more abundant in certain stratigraphic periods, namely Late Cretaceous, Eocene, and Miocene, that it is rarer in earlier and later periods, and that half of all recorded occurrences are concentrated between Late Cretaceous and Eocene (Fig. 3-9). This distribution is worldwide. Clinoptilolite occurs in deep-sea sediments as well as in shallow-water sediments and reflects, according to these authors, a warmer climate during this time period, which raises the soluble Si levels in the sea water. Clinoptilolite appears to be able to form from both volcanic and nonvolcanic precursors. In South Atlantic deep-sea sediments (cores), clinoptilolite occurs in association with three dominant sediment components, i.e., volcanic deposits, biosiliceous deposits, and turbidi te clays, each of which may provide a source of silica for diagenetic formation (McCoy et al., 1977). Volcanic material as precursor for clinoptilolite formation is dominant in the Cretaceous deposits. Diagenesis of volcanic material proceeds through de-

142

n

5

z

40

30 20 10

n

Fig. 3-9. Age distribution of clinoptilolite in D.S.D.P. Legs I to XXXIII (excluding Leg XXXII). (After Nathan and Flexer, 1977, p. 851, fig. 5.)

vitrification of volcanic glass. In Eocene time, biosiliceous material gained in importance. Biosiliceous materials act as precipitation cores and are being partially altered into euhedral zeolite crystals. Although the terrigenous association became more important during Late Tertiary time, it does not appear to be very significant. The Si for the formation of zeolites is probably supplied through degradation of clays. According to Petzing and Chester ( I979), the dominant factor controlling zeolite formation is the Si/A1 ratio of the precursor. The low-silica zeolite, phillipsite, is formed by the rapid breakdown of basic alkaline glasses and, therefore, is more abundant in young ocean sediments, whereas the high-silica zeolite, clinoptilolite, is formed by the slow breakdown of acidic glasses and is, therefore, more abundant in older sediments. There is a good correlation between the spatial and temporal distribution pattern of clinoptilolite and phillipsite and those of subaereal acidic and basic volcanicity. In addition, the dissolution of siliceous organisms in sediments may result in local silica enrichments in sea water, which will increase the Si/AI ratios and, thus, promote the formation of clinoptilolite. In the Recent sediments of the Gulf of Naples, Italy, analcime was found by Miiller (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

143 of the Pacific pyroclasts, these glasses have a trachytic-leucitic chemistry. This could possibly be the reason for the formation of analcime rather than phillipsite. A frequently encountered vertical zonation of zeolites in marine sediments is explained by Iijima (1978) by burial diagenesis. The zones are distinguished mineralogically on the basis of the reaction series of silicic glass + alkali zeolites albite (Fig. 3-10). Zone I is characterized by the presence of silicic glass which is partly altered to montmorillonite and opal-A or opal-CT. Zone I1 is characterized by the reactidn of silicic glass with interstitial water to form alkali clinoptilolite, alkali mordenite, opal-CT, and montmorillonite. The transformation of clinoptilolite and mordenite into analcime characterizes zone 111. In zone IV analcime transforms into albite. This downward succession of authigenic minerals in general represents decreasing hydration with depth. The precursor zeolites commonly persist as relics, especially in younger Tertiary sediments. +

Alteration of basalt and pyroclastics Iron-rich saponite, nontronite, montmorillonite, celadonite, mixed-layered

ILLITE PREHNITE

I

PWPELLY I TE

I

Fig. 3- 10. Schematic diagram showing zonal distribution of authigenic zeolites and silicates in silicic volcanic sediments in a thick column of marine deposits due to burial diagenesis. The zoning is based on the reaction series: silicic glass- alkali clinoptilolite and mordeniteanalcime-albite. Zone I11 is subdivided into two subzones on the basis of the reaction: heulandite- laumontite. Prehnite and pumpellyite may be in the metamorphic regime. (After Iijima, 1978, p. 179, fig. 3.)

144

chlorite-smectite and chlorite are among the more common alteration minerals produced during the halmyrolytic alteration of volcanic material. They are often localized in veins and vesicles. In addition to clays, alteration products often include zeolites (clinoptilolite and phillipsite), iron and manganese oxides, calcite, and minor amounts of K-feldspar. One of the common but ill-defined products of the halmyrolytic alteration of basaltic glass is palagonite, which essentially consists of a devitrified, hydrated glass in incipient stages of alteration into layer-silicate structures and zeolites. Phillipsite and smectite are the most common diagenetic minerals of the palagonitization process (Honnorez, 1978). From the study of DSDP Leg 34 basalts, Bass (1976) suggested two distinct diagenetic environments: ( 1) Non-oxidative diagenesis is characterized by relatively low oxidation rates, which produce assemblages of saponite, chlorite-smectite, talc and minor celadonite. Most of the Fe3+ is incorporated into silicate lattices and only a small amount is available for discrete oxide phases. (2) Oxidative diagenesis is characterized by the abundance of ferric oxides (goethite, limonite, and hematite) and celadonite; smectite and chlorite are less important and are frequently stained by iron oxides. Oxidative zones commonly occur in the upper few centimeters of each cooling unit, apparently due to direct contact with superjacent bottom waters. Oxidized phases can also be found, however, in lower portions of individual flows, filling veins and vesicles. An iron-rich saponite and accompanying celadonite are the dominant sheet-silicates in veins within the basalt cored during DSDP in the Nazca plate (Seyfried et al., 1978). Oxygen isotopic analyses of these clays and associated calcite indicate a formation temperature of 25 "C. Celadonite is commonly associated with goethite and hematite, suggesting that this phase was formed by precipitation within a dominantly oxygenated environment of components leached from basalt and also those provided by sea water. In contrast, iron-rich saponite, containing significant A1203, appears to have precipitated from a non-oxidizing, distinctly alkaline fluid containing a high Na/K ratio relative to unmodified sea water. Sea-water-basalt interaction at low temperatures, resulting in the formation of celadonite and smectite, may explain chemical gradients observed in the interstitial waters of sediments overlying basalts. On the basis of observed oxygen-isotope, C a z + , M g 2 + , and K + gradients in interstitial waters as a function of depth, Kastner (1976) proposed the following alteration sequence for basalt from DSDP Site 322: (1) At shallow burial depth, basalt releases C a 2 + , Mg2+ and iron. Iron precipitates locally as iron hydroxide and a residual amorphous phase enriched in silica develops. (2) In a second phase, smectite starts to form and consumes the Mg2+ that is released during pyroxene weathering and, possibly, some additional Mg2+ from sea water. Thus, depletion in Mg2+

145 and Ca2+-enrichment in the interstitial water results. Phillipsite formation is most probably responsible for the observed K depletion in the interstitial waters. A downhole decrease in I8O, Mg2+ and K + , an increase in Ca2+ content, and a low 87Sr/86Srratio of 0.7067 in the pore fluids of DSDP Site 323 were interpreted by Lawrence et al. (1979) as having been caused principally by the alteration of volcanic material into smectite, potassium feldspar, clinoptilolite, and calcite. A significant portion of the alteration of ash in the basal sequence must have'occurred before the deposition of the thick sequence of upper sediments. Mass balance considerations and the low 6 l 8 0 values of most of the alteration products suggest that much of the later alteration occurred progressively over the last 13 m.y. Iron- and manganese-enriched clays and even discrete metal oxide phases have probably been produced during the submarine alteration of volcanics without the intervention of hydrothermal processes (Natland, 1973; Jenkyns and Hardy, 1976). The most common halmyrolytic alteration products of pyroclastics are smectites, frequently accompanied by zeolites. Late Cenozoic ash deposits cored in Leg 19 of the Deep Sea Drilling Project in the far northwest Pacific and in the Bering Sea have altered to bentonite beds (Hein and Scholl, 1978). The bentonite beds are composed of authigenic smectite and minor amounts of clinoptilolite. A significant part of the Neogene volcanic history of land areas adjacent to the North Pacific is represented by these diagenetic deposits. Inasmuch as the ash beds become more highly altered with increasing depth of burial, formation of bentonite reduces the number of ash layers in the older deep-sea sediments. The amount of illite layers in the illitesmectite mixed-layer phase is used to distinguish diagenetic from detrital smectites. Diagenetic smectite has less than 15% interlayered illite, whereas detrital smectite contains 45%-55% illite interlayers. Apparently, silica is conserved when silicic ash transforms into smectite and minor amounts of clinoptilolite. The chemistry of bentonite beds reflects the chemistry of their parent ash. The diagenesis of silica in argillaceous sediments The alteration of biogenic opaline silica or opal A (primarily diatoms, radiolaria, sponge spicules and silica-flagellata) and the formation of opal-CT and quartz deposits are among the major diagenetic processes occurring in marine sediments. The studies of Jones and Segnit (1975) on opaline silica have revealed at least three distinct categories of opal, among which opal-CT is the most abundant mineral in common opal and in deep-sea chert. Studies of the structure of opal-CT and its relationship to the other silica minerals show that opal-CT has low cristobalite content and is unidimensionally disordered by interstratified stacking of tridymite layers.

146 According to the “maturation hypothesis” applied to marine cherts in pelagic sediments by Wise and Weaver (1974), disordered cristobalite is the first form of silica which is inorganically precipitated. Within periods of up to 100 m.y., this silica is then converted to a-quartz. Alternatively, the “quartz precipitation” hypothesis proposed by Lancelot ( 1973), states that the mineralogy and porosity of the host sediment is critical for the evolutionary path. A high initial content of impurities, such as exchangeable cations, will tend to retard the internal ordering of disordered cristobalite and hence delay the final timing of conversion to quartz. Consequently, the conversion to quartz will take place first in porous, impurity-free calcareous or siliceous oozes and only later in densely packed fine-grained argillaceous sediments. The presence of carbonates may also serve to chelate inhibiting cations such as iron and magnesium and, thus, to enhance the rate of solution of siliceous tests. Evidence from deep-sea sediments supports the following diagenetic maturation sequence: opal-A (siliceous ooze) + opal-CT (porcellanite) chalcedony or cryptocrystalline quartz (chert). The transformation of opal-A to opal-CT and then to quartz is primarily controlled by temperature, time, specific surface area of sediment, pore water chemistry (pH, types and concentrations of ions in solution and degree of saturation with respect to silica), sediment permeability, and the nature of the host sediment, whether primarily calcareous, argillaceous, or siliceous. Whereas most investigators agree that the opal-A to opal-CT transformation is a solution-reprecipitation reaction, opinions are divided as to whether the opal-CT to quartz transformation is also due to solution-reprecipitation or to solid inversion (Hein et al., 1978). Experiments carried out by Kastner et al. (1977) show that the transformation rate of opal-A to opal-CT is much higher in carbonate than in clay-rich sediments and that opal-CT lepisphere formation is aided by the precipitation of nuclei with magnesium hydroxide as important component. In carbonate-rich sediments, the dissolution of carbonate provides the necessary alkalinity, whereas the sea water supplies the magnesium cation necessary for the precipitation of magnesium hydroxide in the nuclei. In contrast, in clay-rich sediments the clay minerals impede the opal-CT formation by competing for the available alkalinity from sea water. As a result, the clays are enriched in magnesium and the rate of opal-CT formation is strongly reduced. This may serve as an explanation why in Mesozoic clayey sediments opal-CT predominates, whereas in carbonate sediments quartz is most common. The Troodos Massif of Cyprus is overlain by a variety of cherts in pelagic chalks, volcanogenic sediments, radiolarites, and radiolarian mudstones, all of Campanian to Upper Eocene age (Robertson, 1977). Most of the +

147 chalcedonic quartz is derived by recrystallization of previously inorganically precipitated disordered cristobalite rather than by direct precipitation. With time, internal solid-state reorganization of the disordered cristobalite is accompanied by gradual expulsion of impurities until its dissolution which is followed by quartz precipitation. Complete conversion of disordered cristobalite to quartz takes place first in porous calcareous sediments free of impurities, as in Cyprus calciturbidites. In fine-grained clay-rich sediments, such as Cyprus radiolarian mudstone, disordered cristobalite persists much longer. In continental-margin sediments off northwest Africa, no siliceous skeletons are preserved as opal-A in the pre-middle Miocene sediments studied by Von Rad et al. (1977). Siliceous organisms have been either dissolved or converted into opal-CT or quartz (Fig. 3-1 1). Eocene opal-A radiolarian skeletons have been replaced in situ by opal-CT crystallites, whereas diatoms and sponge spicules are always replaced by quartz. Calcite tests of foraminifera and mosaic cement in their chambers are replaced by quartz and not by opal-CT. These transformations represent the early diagenetic stage. During intermediate and late diagenesis, silicification fronts develop, involving the replacement of opal-CT by quartz. A progressive diagenetic maturation of the opal-CT structure within the porcellanite stage is suggested by the positive correlation of the opal-CT 4.1 A spacing with burial depth. This strongly suggests that maturation with time and increasing temperature is the major driving force for the chertification in the clayey and carbonate environments. As indicated by porcellanite relicts, all typical quartz cherts were developed by gradual maturation of porcellanites and not by direct precipitation, without the opal-CT precursor. The rate of opal-CT quartz transformation is much slower in the clayey sediments than in a carbonate environment. In the southern part of the Bering Sea, diagenesis of diatom frustules follows a series of changes that are related primarily to temperature, which is controlled by the depth of burial and local geothermal gradient (Hein et al., 1978). During the first 300-400 m of burial, frustules are fragmented and undergo mild dissolution. Upon reaching a depth of 600m, dissolution of opal-A (biogenic silica) becomes widespread. Silica reprecipitates abundantly as inorganic opal-A at depths between 600 and 700m, which is rapidly transformed by crystal growth to opal-CT. This results in the formation of silica-cemented mudstone and porcellanite beds. Opal-A is transformed to opal-CT at temperatures ranging from 35°C to 50°C. Deposition of at least 500m of diatomaceous sediment was required before the temperature at the base of this section was appropriate for silica diagenesis to occur. Reexamination of the results obtained experimentally by Mizutani ( 1977) showed that amorphous silica changes to quartz through an intermediate +

Silica Sources

.

I

I

I

Weakly

miqrotion of solutions

silicif.

Sediments

replacement of caIci1ic tests 1oIso in pores?

precipitation of dense op-CT inmatrix and of l e p i s p h e r e s in pores

in-situ replacement

[Qtzl I time -dependant ordering of opal-CT l a t t i c e

Porce -

2

lanites

I

I

calcar envir

clayey

envir op-CT latticeof ' metostoble l e g by cotions E low permeab 1

replaced op-CT lepisph

T

q t z transformation by dissol Iprecipil opal-CTin microenvironment

Opal- CT q t z nuclei

in situ re-

re1 fast

Cherts

Fig. 3- 1 1. Schematic diagram showing transformations and diagenetic processes for opal-A, opal-CT, and quartz. The vertical position of the silica phases within the diagram is correlated with their diagenetic evolution (right-hand column). A direct formation of quartz cherts without a porcellanite precursor was not presented. (After Von Rad et al., 1977, p. 894, fig. 6.)

phase of opal-CT, and that the (101) spacing of cristobalite progressively decreases from 4.10 A to 4.05 A. The rate of spacing-decrease increases with increasing reaction temperature. This spacing change represents ordering of opal-CT with the passage of time. Accordingly, it is suggested that stratigraphic boundaries should be parallel to isopleths of d (101) spacings, except in cases where folding had occurred. 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

149 quickly, all the silica liberated cannot be dissolved in the sea water. Consequently, there is at least a partial fixation of SiO, in the sediment. The sediments in the Gulf of Naples (Muller, 1961) can be mentioned as an example of recent halmyrolitic authigenesis of SiO, -minerals from finegrained 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 newly formed 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 continuously from the coast to the middle of the Gulf of Naples (Muller, 1961, 1967); on the other hand, the quartz and chalcedony contents increase (Fig. 3-12). Inasmuch as large quantities of SiO, are also released during the transformation of feldspars into clay-minerals, SiO, -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 diagenesis and is important for cementation.

. . .. .. .. . . . ..'.... . .:.. ..." ::'

.:. :. . .: . :. . . : ... .. .I ','

. ..

'

'

Fig. 3-12. Quartz content in Recent sediments in the Gulf of Naples, Italy. (After Muller, 196 1 , p. 7, fig. 2.)

150

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. All the facts indicate that these are not concretions which have been uncovered by water action from older strata, but are 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 even locally. The major mechanisms for their formation that have been proposed, include: ( 1) inorganic precipitation from normal sea water, (2) precipitation from volcanic hydrothermal solutions, and (3) diagenetic concentration at the sea water-sediment interface. Enrichment of manganese in surfaces of marginal sea floors is caused by the upward migration of Mn’+, due to a diffusive gradient, from anoxic sediment layers, and its precipitation in the oxic surface sediments (Elderfield, 1977a). A comprehensive treatment of the subject appeared in a book edited by Glasby (1977). Changes during the shallow-burial stage Changes in the chemistry of interstitial water 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 influenced. Emery and Rittenberg (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 with pH values ranging u p to 8.5 at a depth of about 8 m; most of the values were between 7.5 and 8.0. This downward increase of pH was not observed by Siever et al. (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 p H values of core samples were uniformly lower than that of surface sea water, ranging from 7.00 to 7.85, with most values clustering between 7.2 and 7.7. Siever et al. (1965, p.65) state: “The most likely explanation for the p H being lower than sea water is an increase in CO, 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 CO, pressure in the sediments is most likely the result of bacterial

151 oxidation of organic matter. The indications are that this effect is still operative at a depth of 10 m below the ocean bottom. Siever et al. did not determine the sulfate content in their cores. In the sediments of the California basins, Emery and httenberg (1952, p. 789) found a decrease in the sulfate content with increasing depth. In core 5 the sulfate/chloride ratio decreased from 0.13 1 at 0-6 inches to 0.030 at about 70 inches below the surface. 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 not only in the topmost layer of the sediment but also in the water layers over the sediment. Within the sequence of sediment layers, anaerobic zones can be intercalated among aerobic ones depending on the content of organic matter. Some of the changes in the chemistry of interstitial water can be regarded as a continuation of the processes occurring during halmyrolysis and, therefore, have been discussed in that context. Other changes can be better understood when relating to changes due to deep burial and are discussed later. 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 sulfides During the activity of sulfate-reducing bacteria, H,S is produced which in the presence of dissolved iron or Fe-hydroxides is transformed into black hydrotroilite (FeS . n H,O). 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).

152 The formation of hydrotroilite is not limited to marine sediments; many sediments in the fresh-water lakes (for instance, Lake Constance) locally contain several percents of this sulfide. In the Jurassic Oxford clay of central England, calcite-rich concretions include pyrite (Hudson, 1978). Concretion growth in pelleted anaerobic mud proceeded concurrently with bacterial sulfate reduction and pyrite precipitation. During post-compactional growth, 6I3C increased and pyrite content decreased, indicating a decrease in organic influence. Framboidal pyrite in clayey Holocene sediments and in Jurassic black shale in the northwestern part of the Barents Sea had formed diagenetically (Elverhafi, 1977). Pyrite formation may also be associated with clay-mineral diagenesis. The close textural association between chlorite and pyrite in two representative shales suggests a common diagenetic origin (Sever and Kastner, 1972). The geochemistry of this diagenetic pair indicates a Fe-Mg exchange in the clay mineral associated with sulfate reduction. In addition to pyrite and marcasite, the formation of which is favored by lower pH, very small quantities of other heavy-metal sulfides (mainly sphalerite, galena, and chalcopyrite) occur in clayey sediments of older formations. In pyrite-rich, bituminous, dark clay-marls of different Mesozoic strata, Miiller ( 1955) and Haussiihl and Miiller ( 1963) observed numerous idiomorphic wurtzite crystals which were formed during the early diagenetic stage. The conditions necessary for the formation of such heavy-metal sulfides have not yet been completely clarified. In all probability, however, 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 (marine enuironment) It is to be expected that most of the processes starting in the pre-burial stage also continue during 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 to occur in the Mediterranean Sea by Norin (1953) and Miiller (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. Changes in clay-mineral composition (continental enuironment) Within the continental environment, argillaceous sediments appear to

153 undergo diagenetic changes due to burial by organic material, such as peat. Among the more significant changes is the relative enrichment in kaolinite. Many underclays and tonsteins are rich in kaolinite. It seems that at least part of that kaolinite had formed diagenetically due to leaching by organic acid solutions originating from the organic matter overlying the argillaceous sediments. Thus, Staub and Cohen (1978) found that the kaolinite/smectite ratio in clay strata covered by some peat deposits in South Carolina is more than 2: 1, whereas in clays below this zone this ratio is below 1. Similar processes may have contributed to the formation of underclays below some Paleozoic coals. Due to occurrence beneath acidic marsh, illite clays in Pennsylvanian strata from Iowa-Missouri, U.S.A., were degraded rapidly to mixed-layer clays with maximum content of expansible layers (Brown et al., 1977). On the other hand, poor ordering of kaolinite in the underclays of Carboniferous sediments in Scotland is attributed to degradation of detrital kaolinite in the acidic swamp environment (Wilson et al., 1972). The enrichment in kaolinite of continental argillaceous sediments has also been ascribed to upward leaching by groundwater (dialysis), as in the case of the Kirkwood Miocene Formation in the New Jersey, U.S.A. (Isphording, 1970) or the Triassic flint clay in the Sydney Basin, Australia (Loughnan, 1970).

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, which is indicative of an earlier stage of compaction. Inasmuch as relics of organisms are frequently found in the center of the concretions, it seems likely that organic matter played a role in the formation of concretions. Lippmann (1955) explained the genesis of concretions as follows: Ammonia resulting from the decomposition of organisms or amines gives rise to strongly alkaline environments in the vicinity of the animal (or plant) embedded in the sediment. 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). As the carbonate concentration of the pore solution decreases in comparison to the surrounding environment, and because of the difference

154 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. Petrographic, chemical and carbon isotope data indicate a dominantly organic source for the carbon in calcite-rich concretions from the Jurassic Oxford clay of Central England (Hudson, 1978). Concretion growth in pelleted, anaerobic mud proceeded concurrently with bacterial sulphate reduction and pyrite precipitation. During postcompactional growth, 8 I3C increased and pyrite content decreased, showing waning organic influence. The mineralogical and chemical changes are explained by a replacement of originally marine pore waters, modified by bacterial activity, by water of ultimately meteoric origin. 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 (1964) made such calculations on Lower Cretaceous and Liassic concretions of argillaceous sediments from Hoheneggelsen (northwestern Germany) and WutachSchlucht (southwestern Germany). According to Lippmann, the calculated water content is about 55% corresponding to 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. Chunges during the deep-burial stage of diugenesis According to Taylor (1964), with increasing depth of burial hydrogen ion concentration and oxidation-reduction potential decrease in importance in controlling the diagenesis. Effects of 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 during diagenesis. Changes in the chemistry of interstitial water Large amounts of pore fluids are expelled during the compaction and dewatering process of continental and marine sediments. The chemistry of these pore fluids undergoes significant changes during this compaction

155

process. It seems that most of the salts present in the waters, which are trapped during sedimentation, are squeezed out during the initial stages of compaction. According to many investigations, reviewed by Rieke and Chilingarian ( 1974), the salinity of squeezed-out solutions progressively decreases with increasing overburden pressure. Results obtained by Chilingarian et al. (1973) indicate that the concentrations of solutions squeezed out of montmorillonite saturated with sea water go through a maximum, or at least remain constant, before starting to decrease with increasing overburden pressure (Fig. 3-13). The conclusion of Chilingarian et al. (1973) that the concentration of expelled solutions during the initial stages of compaction is slightly higher than the initial interstitial fluid was not confirmed by Rosenbaum (1976) who examined the variation in chemistry of pore fluid expelled during compaction of a montmorillonite material saturated with distilled water. The initial concentration of all the analyzed ions (K+ , Na' , Ca2+, Mg2+, SO:-, Cl-) decreased rapidly with increasing stress during

38,400 0

I

1

20 40 Amount of extruded solution ( m l )

3

Fig. 3- 13. Variation in the total dissolved solids content with increasing compaction pressure in subsequent fractions of extruded solutions from montmorillonite clay saturated in sea water. (After Chilingarian et al., 1973, p. 397, fig. 4.)

156 the initial loading and thereafter the rate of decrease declined markedly (Figs. 3-14a, b). This indicates that the decrease in concentration of dissolved solids in the pore fluid is related exponentially to the overburden pressure. The observed concentration decreases can be accounted for on the basis of the double-layer theory for highly charged clays, such as smectite. For these clays, the double-layer theory predicts that the electrolyte content of the expelled interstitial waters should decrease when the interaction between the diffuse layers begins to occur. This negative “adsorption”, taking place already during the initial stages of compaction, is caused by the preferential expulsion of ions having the same sign of charge as the colloids, when compaction brings the latter closer to each other. A refinement of this theory, using a model based on Donnan equilibrium, was recently offered by Appelo ( 1977). Not only the total electrolyte content of pore fluids undergoes changes during compaction and burial, but also the electrolyte composition. DSDP samples of interstitial solutions of deeply buried marine sediments throughout the world oceans have shown that in all but the most slowly deposited sediments, pore fluids exhibit changes in composition upon burial. In pelagic clays and slowly deposited (1 cm/103 yr) biogenic sediments these changes are relatively less pronounced than in the carbonate and biogenic sediments deposited at somewhat greater rates. Compositional changes become significant only at great depth. The major elements, except for SO:-, show little variation down to a depth of 300m. The solids and solutions of these sediments must represent a close approach to equilibrium. In partly consolidated sediment sections approaching igneous basement contact, on the other hand, interstitial Ca2+-enrichment, accompanied by depletion in N a f , Si4+ and CO, were observed (Sayles and Manheim, 1975). These changes are attributed to exchanges of Na for Ca in silicate minerals forming from submarine weathering of igneous rocks such as basalts. Water is also consumed in these regions, accounting for minor increases in total interstitial water salinity. It has been shown that the depth to which diffusional communications with the overlying sea water is possible, is small in rapidly accumulating sediments (50-500 cm/m.y.) on continental shelves, but can be large in slowly accumulating sediments ( < 2 cm/m.y.) (Gieskes, 1975). Any concentration-depth profiles, particularly when they are smooth and continuous, may originate at levels that are deeper than that which was sampled. Gieskes ( 1975, p. 450) concluded: “Therefore, in slowly accumulating sediments ( < 2 cm/m.y.) concentration-depth profiles of Ca, Mg, Na, SO, and other major sea-water constituents can be understood only from information derived from drillholes, preferably ones that have been drilled to the oceanic basement. Then, using information on possible changes with depth of the

157

..

350-

300

Potassium V Sodium A Calcium 8

8

-

.

A

-

5 E

'

A

c

v v

-2500

U

n.

200

A

150

.L

I

I

10

1

100

1

1000

Log e f f e c t i v e a x i a l pressure, kg/cm2

Fig. 3-14a. Variation of K, Na and Ca concentrations in expelled pore fluid with effective axial pressure at 40°C. (After Rosenbaum, 1976, p. 120, fig. 3.)

Magnesium 0 Chloride SulDhote 0

*Ol

8 0

0

l150.

O

15.

0

0

OO 650 51

@,@

e:.'o;sr@

808

400-

0

Fig. 3-14b. Variation of Mg”, C 1 and ~ SO,‘+ concentrations in expelled pore fluid with effective axial pressures at 40°C. (After Rosenbaum, 1976, p. 120, fig. 4.)

158 diffusional transport properties of the sediments, an analysis can be made of the probable reaction sites in the sediments and the elements involved in such reactions. Authigenic material in the sediments does not necessarily result from the reactions suggested by observed gradients and, conversely, suggested reactions must be verified by actual identifications in the sediments. Little information is available on the actual equilibrium state between interstitial waters and the solid sediments. Smooth concentration gradients, however, suggest that such an equilibrium state may not exist and that even the alteration of deeply buried volcanic rocks makes a contribution of dissolved material to the ocean, at least for Ca.” This alteration process may also form a sink for Mg2+ and Na+ , and possibly K + and CO,. On the basis of discrepancies between predicted and observed interstitial water profiles, Kastner and Gieskes (1976) identified two major sites of reactions at DSDP site 323 in the Bellinghausen Abyssal Plain of the southern Pacific Ocean (Fig. 3-15): (a) In the silicification zone, dissolution of plagioclase and a few coccoliths and formation of Mg-rich smectites and K-feldspars are responsible for the observed increases in Ca2+ and decreases in K + and Mg” contents. (b) Below the silicification zone, the observed interstitial water concentration gradients are to be accounted for by the alteration of the basal basalt minerals (pyroxene and plagioclase) and formation of celadonite, goethite, calcite, and smectite. DH

1 ’

meqll 2 3 4 “\

\f

f”

4.’

.

- , 10

u

600 il moles/l NHq 30 m m o Il e s i l S O 4

400 20 ,

t

rnoiesll

I

.

400

1

600

t

Alkalinity

800

Sulfate

I

Ammonia

8001

Silica

400

400

600

200

C

5

600

600

Fig. 3-15. Interstitial water chemistry, site 323, DSDP, Bellinghausen Abyssal Plain. I =clay, silt, and sand; I1 =diatomite, claystone, and porcellanite below a depth of 400 m; I I I claystone: I V = calcareous plankton and Fe-Mn sediments; V =zeolitic clays; W =basalt. (After Kastner and Gieskes, 1976, p. 14, fig. 3.)

159 The differences in the extent of diagenesis associated with basalt weathering at different sites could be due to various factors such as differences in age, the thickness of sediment cover, sedimentation rate, and possible interference of hydrothermal activity. Somewhat similar conclusions were reached by Perry et al. (1976) for DSDP Site 149 in the Caribbean Sea. With increasing depth, the sediment pore-water exhibits a strong depletion in Mg2+ and a corresponding enrichment in Ca2+ content, whereas the alkalinity remains relatively constant. Dissolved SiO, content is nearly constant (6 ppm) in the upper 100 m of the sediment, but is much higher (60 ppm) in the deepest pore waters. A depletion in K content and a decrease in 180/160 ratio with increasing depth were also recorded. The submarine alteration of volcanics to a smectite clay could account for the pore-water gradients (Fig. 3-16). By the alteration process, Ca2+ and Mg2+ are released to the pore water, but Mg2+ is retained by smectite formation. The net reaction allows for the observed relationship between Ca2+ and Mg2+ gradients of interstitial water, with little net change in alkalinity. Another process proposed for the removal of Mg2+ from the interstitial water, is the reaction with clay minerals in anoxic environments according to the following scheme:

3 Mg(S0l"tlo") + 2 Fe(clay) + 4 S

+

3 Mg(clay)+ 2 FeS~(pyrltc)

where Fe(clay) refers to the structural Fe3+ in the clay and Mg(c,av)to the Mg2+ added to the clay structure in the course of the reaction. According to Drever (1971a, b), Mg is derived from interstitial waters supplemented by diffusion from the overlying water body and sulfide is formed from dissolved sulphates by the action of reducing bacteria. From experimental data, Heller-Kallai and Rozenson (1978) recently concluded that direct Mg e Fe exchange in clays, as proposed by Drever and more recently by Elderfield (1977b), seems improbable, but that depletion of Mg in interstitial waters of anoxic sediments may be due to reaction of Mg-containing solutions with partially disintegrated clay. Also another mechanism, proposed by Sholkovitz (1973), which postulates addition of Mg to the exchange sites of clay minerals previously blocked by precipitated ferric hydroxides, appears improbable in view of the findings of Sayles and Mangelsdorf ( 1977) that the proportion of adsorbed Mg on marine clay had been grossly overestimated in the past. Changes in cia-v mineral composition-illite diagenesis 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

160 SlOZ p M O L E S / L I T E R

K

mMOLES/LITER

rnMOLES/LITER

Fig. 3-16. Vertical concentration profiles of (a) dissolved SiO,, (b) K + , and (c) Ca2+ and Mg2+ in the sediment pore-waters of Hole 149, DSDP. Crosses represent cold sediment squeezes, dots: warm squeezes. (After Perry et al., 1976, p. 415, figs. 3a, b, 4.)

161 with material from the sedimentation basin. In older sediments at greater depths, .this comparison is not possible and, thus, the unknown factors correspondingly increase.

Mineralogy and chemistry. The major mineralogical process occurring during the deep burial of argillaceous sediments is the progressive conversion with increasing burial depth of smectite into illite or into a mixed-layer illitesmectite, with a high proportion of illite layers. This major mineralogical process has been suggested as occurring in association with and being related to other minor changes, such as decomposition of mica and K-feldspar, decomposition of kaolin, and formation of chlorite or chlorite interlayers. In the Eocene to Pleistocene shales of the U.S. Gulf Coast the effects of burial on argillaceous sediments have been studied most extensively (Burst, 1969; Perry and Hower, 1970; Hower et al., 1976). When drawing conclusions from these studies, it must be kept in mind that mineralogical variation may simply reflect changes through time of the mineralogical detritus supplied to the sedimentary basin: For the interpretation of data obtained, this factor must, therefore, be carefully considered. A detailed mineralogical and chemical investigation was made by Hower et al. (1976) of shales from a well in Oligocene-Miocene sediments of the U.S.A. Gulf Coast. Major mineralogical changes with depth take place over the interval of 2000-3500m, after which no significant changes are detectable. The most abundant mineral, illite-smectite, undergoes a conversion from less than 20% to about 80% illite layers over this interval, after which the proportion of illite layers remains constant (Fig. 3-17). Over the same interval, calcite decreases from about 20% (bulk sample) to almost zero, disappearing from progressively larger size fractions with increasing depth: potassium feldspar (but not albite) content decreases to zero, whereas chlorite content appears to increase (Figs. 3-18 and 3-19). Variations in the bulk chemical composition of the shale with depth show only minor changes, except for a marked decrease in CaO, concomitant with the decrease in the calcite content. By contrast, the < 2 p m fraction, consisting of pure illitesmectite, shows a large increase in K,O and A1,0, and a decrease in SiO, content. The atomic proportions closely approximate the reaction: smectite +A13+ K+ = illite Si4+. A similar increase in illite layers with increasing burial depth was observed by Perry (1974) in the Gulf Coast sediments of Louisiana (Fig. 3-20). The measured K-Ar ages of the mixed-layer illite-smectite become progressively younger with increasing depth of burial because of the diagenetic addition of potassium, caused by conversion of smectite layers to illite layers. Smectite decreased in abundance and content of mixed-layer clays increased with increasing depth of shales in the Rhinegraben, southwestern Germany (Heling, 1974).

+

+

162

Percent illite layers

Fig. 3-17. Proportion of illite layers in illite-smectite as a function of depth in the finest and coarsest clay fractions of shales from a well in Oligocene-Miocene sediments of the U.S.A. Gulf Coast. (After Hower et al., 1976, p. 730, fig. 3 . )

The conversion of smectite to illite-smectite during burial metamorphism is reported from Middle Cambrian to Middle Ordovician sedimentary rocks of western Newfoundland (Suchecki et al., 1977). Potassium was derived from K-feldspar and/or K-mica in the detrital silt of the shale. Expandable chlorite and corrensite occurring in association with the illite-smectite incorporate Mg probably released during the smectite -+ illite conversion. Clay minerals from the Lower Cretaceous shale outcrops of northeastern British Columbia were investigated by Foscolos and Kodama ( 1974) in order to assess their degree of diagenesis and their oil-generating potential. Crystallinity index (the width of the illite 001 peak at half-height expressed in mm), sharpness ratio (the ratio of the height of the illite 001 peak at lO.0A to the height of the base line at 10.5 A), content (in W ) of 2M illite, polymorph, and presence of discrete minerals were used as indices of diagenesis. Data on the < 2-pm fraction show that the crystallinity index decreases, whereas the sharpness ratio and the 2M illite polymorph content increase with burial depth (Table 3-V). Results on the (0.08-pm size fraction reveal the ex-

163

Percent potassium feldspar

Fig. 3-18. Potassium feldspar content of > 10-pmand 2-10-pm fraction of shales from a well in the Oligocene-Miocene sediments of the U.S.A. Gulf Coast. (After Hower el al., 1976, p. 73 1, fig. 4.)

istence of a three-component interstratified clay mineral. In addition, Fourier transform calculations and chemical and physicochemical analysis indicate that both the ratio of the amounts of nonhydrated clays (illite) to hydrated clays and the K ,O content of clays increase with burial depth; cation exchange capacity and specific surface area decrease with burial depth.

Conversion process. Burst (1969) viewed the conversion process of smectite into illite as essentially a simple dehydration of smectite under the higher pressure and temperature conditions of deep burial. He distinguished three stages : (1) Expulsion of water from the pores at a depth of around 2500 m; the remaining water (about 30%) is present in between the layers where it forms a bimolecular layer. (2) At a depth range of 2500-4500 m, expulsion of one of the water layers occurs from the interlayer. (3) Below a burial depth of 4500m, there is an expulsion of the last layer

1

I

I

9

Percent

chlorite

Fig. 3-19. Chlorite content of the >2-pm and 0.1-2-pm fractions of shales from a well in the Oligocene-Miocene sediments of the U S A . Gulf Coast. (After Hower et al., 1976, p. 731, fig. 5.)

68000 I-

* c. 8,000

0 f0n) 10,000

12.000

14,000

*J

m 16,000

*J 1

I

I

I

20

40

60

80

% illite layers in illite/smectite

Fig. 3-20. Relation between sample burial depth and percentage of illite layers in mixed-layer illite-smectite in about 3 kbar), and (4) illite with imperfect crystallinities (Kubler’s values between about 10 and 4 mm, the latter value designating the high-grade limit of the anchimetamorphic zone). Other minerals that are considered to appear in parts of the “very-low-stage” of metamorphism, but to persist to different extents into the higher-grade “low-stage’’ facies, include pyrophyllite, glaucophane, jadeitic pyroxene, and stilpnomelane (Winkler, 1970, pp. 204, 225-226). Winkler (1970, p. 204) also claims appearance of chloritoid in the “verylow-stage’’ of metamorphism. The present writer considers this typical greenschist-facies mineral not to occur at lower grades of metamorphism, with the possible exception of rocks extremely rich in Mn. Unfortunately, with the exception of illite the distinctive minerals listed above appear only in rocks rich in intermediate or basic volcanic material. Strict application of Winkler’s mineral-facies criteria would assign many of the associated pelitic rocks (which, for instance, in the laumontite zone may still abound in sedimentary clay minerals such as montmorillonite) to diagenesis. Such rocks cannot be regarded as metamorphic on the basis of their own mineral facies, but can only be postulated to be such on the grounds of their association with some other rock types containing a diagnostic “metamorphic” mineral assemblage such as laumontite-quartz or pumpellyite-quartz. Winkler’s ( 1970) criteria for the onset of “metamorphism”, therefore, are directly applicable to a very restricted range of rock compositions and only on the basis of association with other rocks, at least in the lower grade of this group of metamorphic facies. Kubler (1967a, 1967b, 1968), followed by other French and Swiss authors (Dunoyer de Segonzac et al., 1968; Artru et al., 1969; Dunoyer de Segonzac, 1969, 1970; Frey, 1970) have introduced the concept of “anchimetamorphic zone” or “anchizone” intervening between the zone of late diagenesis and the epizone of regional metamophism (including the greenschist facies), reviving a term coined by Harrassowitz in 1927. Kubler’s approach makes use of the increase in “crystallinity” of illite during diagenesis and lowestgrade metamorphism, as monitored by the concomitant decrease of the width of the illite (10 A) X-ray diffraction peak at half-peak height. This anchimetamorphic zone is defined primarily by limiting values of the illite “crystallinity” index (Kubler, 1967a, p. 1 1 1; 1968, pp. 392-395). Pyrophyllite may appear, whereas kaolinite-group minerals and irregular illitemontmorillonite mixed-layers are usually absent.

307 Inasmuch as the low-grade limit of the anchimetamorphic zone corresponds closely to the appearance of pyrophyllite at the expense of kaolinite and its polymorph dickite in aluminous sedimentary rocks, which takes place at a more advanced stage of burial than does the alteration of kaolinite and montmorillonite to illite and chlorite (cf. Starke, 1968; K m h , 1969), Kubler (1967a, pp. 110- 11 1; 1968, pp. 394-395) correlates the zone with the “early metagenesis stage” of Kossovskaya and Shutov ( 1961, 1963, 1970). The illite-“crystallinity” method has the advantage of being applicable to a much wider range of clastic sediments than is the low-grade mineral facies method in volcanic sediments. Low potassium contents and abundance of clastic muscovite, however, may cause anomalously high and low crystallinity indices, respectively. The limitations of this method are discussed in greater detail in a later section of this chapter. Inasmuch as the “stage of early metagenesis” and the lower-grade “stage of deep epigenesis” are believed by Kossovskaya and Shutov (1961 and later papers) to encompass Coombs’ prehnite-pumpellyite-metagraywacke facies and laumontite zone of the zeolite facies, respectively, the suggested equivalenke of the anchimetamorphic zone with the “early metagenesis” stage would provide a correlation between the illite-“crystallinity” zones and the low-grade facies in volcanic sediments. The acceptance of such a correlation has apparently led Winkler ( 1970) to include a specified range of illite-“crystallinity indices” as one of the diagnostic criteria for his “very-low-stage metamorphism”. Although Kisch (1974) has shown that the anchimetamorphic zone as based on illite “crystallinities” and the appearance of pyrophyllite, on one hand, and the prehnite-pumpellyite-metagraywacke facies, on the other, are associated with a similar range of high coal ranks, direct correlation studies of diagnostic low-grade zones in the same sequences are lacking (but see Kisch, 1980b). Such studies will be required before this correlation can be considered to have general validity. In fact, unless the diagnostic mineral equilibria are affected in the same way by temperature and pressure, i.e., have the same d T/d P relationships, they may well have different relationships under various geothermal gradients. In the writer’s opinion, until well-defined diagnostic “metamorphic” mineral-transformation stages are firmly established to occur at approximately similar stages of deep burial or incipient metamorphism in various areas (at least in clay-rich clastic sedimentary rocks and volcanic sediments), low-grade mineral facies (though invaluable for the establishment of diagenetic and lowest-grade metamorphic zoning) should not be used as the criterion to label rocks as “metamorphic” or “non-metamorphic”, or to establish a boundary between the diagenetic and the metamorphic realms. In collectively attributing the lowest-grade metamorphic facies to “burial metamorphism”, Winkler ( 1965, 1967) has appreciably modified the meaning of this useful term as proposed by Coombs (1961); he defined it on the

basis of a temperature range and assigned the absence of schistosity, which is a major factor in Coombs’ definition, a second place: “In general, the absence of penetrative movement precludes a schistose structure. Although the original fabric of the rocks may be largely preserved, the mineralogical composition has been changed. Metamorphic changes are hardly visible in hand specimens; only in thin sections can they be clearly recognized” (Winkler, 1967, pp. 4-5). Arguing that in many occurrences the lawsonite-glaucophane-schist facies rocks show no schistosity, Winkler included this facies, as well as the somewhat lower-pressure lawsonite-albite facies, in the “burial metamorphism” stage. This inclusion would also apply to the many occurrences common schistosity of which originally gave rise to the name glaucophaneschist facies, whereas according to Coombs’ concept schistose gluacophaneschists should be regarded as regional-metamorphic and not as burial-metamorphic. Although rocks such as sandstones that are associated with glaucophane-schists may hardly show any textural imprint of metamorphism, it is very doubtful if any mineral assemblage containing the characteristic high-pressure minerals glaucophane, lawsonite, or jadeite, can form as a result of mere burial without the action of tectonically-controlled pressure (cf. De Roever, 1972). In the writer’s opinion, Winkler’s extension of “burial metamorphism” to cover series of predominantly schistose rocks of the glaucophanelawsonite-schist facies is at variance with Coombs’ concept, divests it of most of its usefulness, and should not be used. The proposed distinction between diagenesis and burial metamorphism is arbitrary and applicable directly only to a very restricted range of rock compositions. It is also undesirable in that it reduces the usefullness of the term “burial metamorphism” to cover all late-diagenetic and lowest-grade metamorphic processes up to the development of schistosity. Kubler (1967b) has shown that correlation between schistosity and illite-“crystallinity” indices were unsuccessful: cases are known of anchimetamorphism without development of schistosity and presence of schistosity without anchimetamorphism. Apparently, the dynamic factor which plays an important role in the development of schistosity, has hardly any effect on the development of illite crystallinity. The writer believes that there is no close parallelism between the development of metamorphic fabric and diagnostic mineral modifications. All attempts to attach a mineral-facies interpretation to a texturally delimited stage, such as burial metamorphism, therefore, must be rejected. Terminology

It may be useful, in summary, to present the terminology of burial diagenesis and lowest-grade metamorphism as used in this chapter.

309 The author uses the terms referring to specific stages, mineral zones, or mineral facies only for sequences containing the mineralogic criteria defining these stages, etc., and not for sequences that have been extrapolated from the presence of other criteria to have reached a similar degree of lowest-grade a1teration. Thus, the lowest-grade metamorphc facies are used only for sequences containing the. diagnostic zeolite and hydrous Ca-Al-silicate minerals or mineral assemblages. “Anchimetamorphic” or “anchizone” is recognized only on the basis of illite-“crystallinity” data and the appearance of pyrophyllite or rectorite-allevardite. The “stages of regional epigenesis and metagenesis” are defined on the basis of Kossovskaya and Shutov’s criteria. Usage of general terms in a more specific sense, or one departing from common practice, will be marked by parentheses. Such terms include “diagenetic zone” and “epizone”, as based on illite-“crystallinity” ranges, and the “facies” of Kossovskaya and Shutov. Burial diagenesis and burial metamorphism are used according to Coombs (1961) definition for partial or complete reconstitution, commonly on a regional scale. Characteristically there is no development of a penetrative fabric such as schistosity. There is also no mineral-facies connotations. These terms are virtual synonyms of late diagenesis and epigenetic diagenesis. Incipient metamorphism is used as a general term for the more advanced stages of mineral modification as characterized by the appearance of the attributes of anchimetamorphism. Thus it is approximately equivalent to the “stage of early metagenesis”. This term has very generalized mineral-facies connotations: Of Winkler’s ( 1970) “very-low-stage’’ metamorphism, it includes at least part of the prehnite-pumpellyite-metagraywacke facies and, probably, all of the lawsonite-albite-schist and glaucophane-lawsonite-schist facies. Slaty cleavage or schistosity, at least in pelitic rocks, is a common though not an essential attribute of incipient metamorphism. Lowest-grade metamorphism, finally is used as a general term covering all the above stages and metamorphic facies, up to the onset of the greenschist facies.

MODIFICATION OF CLAY MINERALOGY OF CLASTIC SEDIMENTARY ROCKS UPON BURIAL DIAGENESIS (BURIAL METAMORPHISM) AND INCIPIENT METAMORPHISM

By the appearance of the greenschist facies, which is regarded as signalling the onset of traditionally-recognized regional metamorphism, clastic sedimentary rocks have undergone major changes in mineralogy. Minerals such

3 10

as smectite, kaolinite, irregular illite-smectite mixed-layers, as well as clastic biotite, calcic plagioclase and K-feldspar, have entirely disappeared. Muscovite and chlorite have become the predominant layer-silicates in most clastic metasediments. Although minerals such as pyrophyllite and paragonite may be less abundant, they are highly characteristic constituents. Most of these changes d o not take place suddenly: many take place over wide range of burial conditions, and are in fact largely completed during burial diagenesis (burial metamorphism); others mainly occur in that ill-defined stage between diagenesis and conventional regional metamorphism variously referred to as incipient metamorphism, very-low-grade metamorphism, anchimetamorphism, metagenesis, etc. In this section, the writer discusses these various changes. Disappearance of primary clay minerals and some clastic constituents during burial diagenesis, as amply documented by mineralogic studies in deep oil wells, are discussed first, followed by the mineralogic characteristics of incipient metamorphism. Modification of primary mineralogy of clastic sedimentary rocks during burial diagenesis

The present sections deal with mineral modifications which occur during burial diagenesis for each one of the major primary, early-diagenetic, or clastic silicate mineral. The major alteration processes taking place during this stage include: ( 1) Replacement of smectite through progressive mixed-layering by illite and chlorite. (2) Replacement of kaolinite group minerals by illite and chlorite. (3) Increasing three-dimensional order and modification of the polytype of kaolinite group minerals. (4) Progressive change in the composition and polytypes of chlorite. ( 5 ) Progressive alteration of clastic biotite through vermiculite to chlorite. (6) Progressive albitization of clastic plagioclase. The formation of zeolites and other hydrous Ca-Al-silicate minerals at the expense of calcic plagioclase are discussed separately. Replacement of srnectite upon burial Illite-smectite mixed-layering During burial diagenesis, smectite tends to disappear gradually with increasing depth through formation of irregular illite-smectite mixed-layers with a decreasing percentage of expandable layers. The replacement can be

L O G 6 A B A WELLS ( D O U A L A BASIN, CAMEROON 1 haolinite. dickitr

chlorite

-

!EIGHT LO55 t JOOO C' ( i n % i f initial weiaht I

I

20

10

CARBON RATIO" non-volatile carbon olal organic c a r b o 0.2 0.1, 0.6 0

0

D 0

\

\ i

I

i \

d

\

i

\

. a o

0

0

t

,000

.. .. .......,

u 40 30 20

6

-

30

20

1

0

0.2

0.4

0.6

Fig. 5-1. Mineralogical and geochemical analyses of the Upper Cretaceous sequence in the wells of Logbaba, Douala Basin, southwestern Cameroon. (After Dunoyer de Segonzac, 1969. fig. 10.)

312 monitored by the pro8ressive shift of the glycolated (001) peak, from 17 A in smectite towards 10A in illite, with the latter without expandable layers being the final product. Smectite or irregular illite-smectite mixed-layers do not occur in either the greenschist facies or, as discussed later, in the anchimetamorphic zone or equivalent zones of incipient metamorphism. Progressive increase in illite-smectite mixed-layering with depth was first demonstrated in wells in the Tertiary of the Gulf Coast (e.g., Weaver, 1959; Burst, 1959; Powers, 1959) and in the Carboniferous deposits of Oklahoma (Weaver, 1959). Subsequently this was also demonstrated in the Upper Cretaceous deposits of the Douala Basin, Cameroun (Dunoyer de Segonzac, 1964, 1969, see Fig.5-1); a well in the Tertiary and Cretaceous deposits of the Camargue, southern France (Dunoyer de Segonzac, 1969); wells in the Jurassic “Terres Noires” of southeastern France (Artru and Gauthier, 1968); the Cenozoic of Azerbaydzhan (Teodorovich et a]., 1967); the Jurassic deposits of eastern Ciscaucasia (Teodorovich and Konyukhov, 1970); the Pliocene and Quaternary deposits of the geothermal areas of the Salton Sea, California (Muffler and White, 1969); and the Taupo Volcanic Zone of New Zealand (Steiner, 1968; Eslinger and Savin, 1973a). On a regional scale, the progressive illitization of montomorillonite with increasing degree of incipient metamorphism has been described from the Jurassic “Terres Noires” of the subalpine chains of the westernmost French Alps (Dunoyer de Segonzac et al., 1966; Artru et al., 1969; Dunoyer de Segonzac, 1969). The absence of montmorillonite from deeply-buried sedimentary sections which grade laterally into less deeply buried sections which contain montmorillonite or illite-montmorillonite mixed-layers has been noted by Carrigy and Mellon (1964) in Alberta, by Karpova (1969) in the Greater Donbas area of the U.S.S.R., by Karpova and Timofeeva (1971) in the northern Caucasus, and others.

’

Depth and temperature of illite-smectite mixed-layering. Although the progressive mixed-layering appears to be largely a function of maximum depth of burial, the depth ranges over which the process takes place differ widely between various regions. This is true in the case of those sequences that underwent no significant postdepositional uplift and, thus, may be considered to be at their maximum depth of burial at present. Table 5-111 gives the depth ranges over which the proportion of montmorillonite layers in illite-montmorillonite mixed-layers progressively decreases from 75 to about 20% in a number of wells for which detailed

’

The name “montmorillonite” is used following the original articles, although in many of these papers no clear distinction is made between montmorillonite sensu strict0 and the smectite (montmorillonite group) in general.

TABLE 5-111 Depth and temperature intervals of reduction in the percentage of expandable layers in illite-montmorillonite mixed-layers from 75 to about 20% in six wells Locality or area and age of formation

Depth interval

Near Galveston, Texas (well E) Oligocene

Geothermal gradient

Reference

about 1300 m 35" (approx. 2100-3400 m depth) (85°-900 to 12Oo-125"C)

3I0C/km

Perry and Hower (1970, 1972)

Coastal Louisiana (well C) Miocene

>3OOO m 75 " (approx. 2600-5800 m depth) (80"-155"C)

24"C/km

Perry and Hower (1970, 1972)

Chambers County (Gulf Coast), Texas Tertiary

about 1200 m 35" (approx. 2600-3800 m depth) ( 100°-1350C)

29"C/km

Burst (1959; 1969, pp. 77-79)

32"C/km

Dunoyer de Segonzac (1964, 1969)

60°C/km

Muffler and White (1969)

180°C/km

Muffler and White (1969)

Logbaba, Douala Basin, Cameroun about 2100 m Upper Cretaceous (Approx. 1500-3600 m depth)

Temperature interval

'

50"

(70"-120°C

I)

to 20Oo-210"C ')

Wilson No. 1 well, Salton Sea geothermal field, southeastern California Plio-Pleis tocene

about 2000 m (approx. 1300 m 2-3300 m depth)

110°C (95"-100"

I.I.D. No. 1 well, Salton Sea geothermal field, southeastern California Plio-Pleistocene

about 500 m (approx. 45 m 2 - 5 0 0 m depth)

not known (up to 2OO0-21O0C ')

'

I 35% expandanble layers in illite-montmorillonite mixed-layers; * disappearance of discrete montmorillonite: completion of conversion of illite-montmorillonite mixed-layers to illite; mean geothermal gradient to 5000 f t depth. calculated from Muffler and White (1969. fig. 3).

w c. w

3 14 clay-mineralogic data and well temperatures (geothermal gradients) are available. These depth ranges vary from 500 to more than 3000m, with differences much larger than can be accounted for by differences in the geothermal gradients. The temperature intervals over which the progressive mixed-layering takes place range from 35" to more than 100°C. although Perry and Hower (1972, p. 2017) note a "much closer coincidence of the temperature intervals over which the dehydration takes place (in the two Gulf Coast wells E and C) than there is for the depth intervals". The temperature of the onset of illite-montmorillonite mixed layering at the expense of montmorillonite or of mixed-layers with minor proportions of illite layers varies within a relatively narrow range, commonly 70" to 115°C (cf., Dunoyer de Segonzac, 1970, fig.5; Burst, 1969, fig. 10). Both the temperature interval of progressive mixed-layering and the temperatures at which the mixed-layers retain only about 20% expandable layers show a far greater range of variations. These temperature differences could in part be related to differences in geothermal gradients. The higher water pressures prevailing at greater depths at which a given temperature is attained under a low geothermal gradient may be expected to necessitate higher dehydration temperatures. Such an explanation has been invoked to account for differences in dehydration temperatures in the Gulf Coast wells by Burst (1969) and might also explain the differences in temperature at which 20% montmorillonite content is reached in the mixed-layers of the four wells of the Gulf Coast and the Douala Basin listed in Table 5-111. Such a conclusion, however, is at variance with the high temperatures at which the conversion of illite-niontmorillonite mixed-layers is completed in areas of extremely high geothermal gradients: ( 1 ) 200°C for the Salton Sea geothermal field (Muffler and White, 1969), (2) above 230°C for the OhakiBroadlands, New Zealand, geothermal area (Eslinger and Savin, 1973a), and ( 3 ) about 230°C for the Wairakei thermal area of New Zealand (Steiner, 1968). The lower temperatures found for the advanced stages of mixed-layering in some of the Gulf Coast wells are in accordance with the experimental data of Khitarov and Pugin (1966), who found that the loss of interlayer water from montmorillonite was completed at temperatures of about 130°C at 1000 bars pressure (equivalent to 3.5 km depth at a geothermal gradient of 35"C/km) and 170°C at 5000 bars pressure (equivalent to 18.5 km depth at a geothermal gradient of about 10"C/km). In the presence of potassium, the dehydrated montmorillonite is converted into a 10-A "hydromica". G. Sabatier (unpublished. results, quoted by Dunoyer de Segonzac, 1969, pp. 71-72) found that in a mixture of montmorillonite and kaolinite in an artificial seawater with 0.60 g KCI/I, illitization of the montmorillonite took

315 place at a temperature of 200°C and 15 bars pressure. Experimental results on Louisiana Gulf Coast argillaceous sediment in artificial seawater at 100°C and 7 bars and 200°C and 21 bars by Hiltabrand et al. (1973). showed formation of chamosite or iron-rich chlorite and illite. whereas feldspar, kaolinite, and mixed-layered montmorillonite were destroyed or decreased in relative amount. Montmorillonite showed a shift of the basal peak to 14.7A, probably due to the formation of chlorite-like sheets. The effect of the concentration of cations, particularly of potassium. o n mixed-layering is discussed in a later section. Development of regular illite-smectite mixed-la.vers. In recent years. a tendency towards partial stacking order has been noted in many illite-sniectite mixed-layers from bentonites (e.g., McEwan, 1956: Maiklem and Campbell. 1968), presumably formed at the expense of montmorillonite, and subsequently also in other clastic rocks (see references in Hamilton, 1968. p. 13). The ordering commonly is apparent from the presence of sequences of non-integral(O0 1) spacings associated with a low-angle reflection (at 26-28 A in the air-dried or 29-32 A in the glycolated samples), which disappears upon heating above 4OO0C, with a concomitant reduction of the first basal spacing to 10.0-10.2A (see Fig. 5-2). In some deep wells, for instance those in the Miocene of Sicily and the Triassic of Tunisia studied by Long and Neglia (1968), the partial ordering of the mixed-layers characterize a deeper zone than does irregular illitemontmorillonite mixed-layering and, thus, obviously represents an advanced stage in the illitizaiion of montmorillonite through K-fixation upon deep burial. The partial ordering in these cells occurs in clays that still contain ample kaolinite, but no discrete montmorillonite. The tendency towards stacking order generally is manifested when the proportion of montmorillonite mixed-layers has been reduced to less than about 40% during progressive burial. Mixed-layer illite-montmorillonites with 40-30% expandable layers from the sandstones and coarse siltstones of the Permian succession of the Lower Hunter Valley, N.S.W., show extensive, non-integral 001-spacing sequences (see Fig. 5-3). Fourier transform investigation of these sequences by Hamilton (1968, pp. 12-13) indicated a structure of alternating mica and montmorillonite layers in approximately 2 : 1 proportions, in a layer sequence showing a striking tendency towards 1 : 1 stacking regularity, somewhat similar to that of mica-montmorillonite mixed-layers from Kinnekulle (Sweden) and Woodbury (England) analyzed by McEwan ( 1956). Some silty claystones and siltstones from the Tomago Coal Measures of this area contain 50 : 50 mica-montmorillonite mixed-layers showing peaks at -25A (air-dried) and 28A upon glycolation, with peaks at integral

3 16 Kinnekulle A - 2

1-L

0

I

,

I

I -

02 e

i

Two Medicine

13 1

660

3025

463

2662 2 4 7 0

4

L ,

.____-

/-A

0 b

“2 8

Fig. 5-2. Calculated profiles (upper curves) and X-ray diffraction patterns (lower curves) of ethylene-glycol-treated illite-montmorillonite with maximum 1 M ordering. a. Kinnekulle A-2, Sweden: 32% montmorillonite layers. b. Two Medicine Formation, Mont. ( U S A . ) : 35% montmorillonite layers. (After Reynolds and Hower, 1970, fig. 1 . )

submultiples of this value, indicating regular 1 : 1 stacking (Hamilton, 1968, p. 12; Fig. 5-3, upper XRD trace). The 50: 50 mica-montmorillonite mixedlayers found in a few claystones are less regular. The most common mixed-layer type in the Permian sediments of the Lower Hunt Valley has mica-montmorillonite layer ratios of 60 :40 to 70 : 30. Samples from claystone and clayey siltstone show basically irregular stacking, usually with a weak low-angle peak indicating a slight tendency towards stacking order, particularly in the more silt-rich claystones and

317

Y

?

m

AIRDRIED

-I I

Fig. 5-3. (001) diffraction spacing sequences of ordered and partly ordered illitemontmorillonite mixed-layers in X-ray diffraction patterns of air-dried and ethylene-glycolsolvated oriented-aggregate ,

F

5

CLAY MINERALS

CARBONATE AND CRlSTALLlNlTY OF ILLIT€ QUARTZ CONTENT 53 Distribution inthe fraction Wciqht-% of t h e index after Kubler s tz t GO

I

I

\

5.00

7.0 ~

I I

t

I

I 14 O

II

t\ 23 27

A

Fig. 5-20. X-ray diffractograms of ( 2 p m fractions [untreated ( N ) , glycolated (C), and heated ( C h ) ] of an anchimetamorphic Middle Devonian slate from the Basque Pyrenees (France). A =allevardite; C =chlorite; M u =muscovite (illite); Py =pyrophyllite; Ch = chloritoid; Pa =paragonite. (After Dunoyer de Segonzac and Heddebaut, 1971, fig. 2.)

Rectorite (or allevardite) as an anchizonal index mineral In natural occurrences, rectorite (or allevardite) is commonly associated with pyrophyllite. These two minerals are persistently associated in (1) the Siluro-Devonian argillites and siltstones of the Ougarta, and of the Tindouf, McMahon and Ahnet Basins in the western Sahara (Beuf et al., 1966, p. 375; Dunoyer de Segonzac, 1969, pp. 115-1 19; Chennaux et al., 1970); (2)

368 carbonaceous shales of north-central Utah (Henderson, 1970), (3) the anchizonal Devonian deposits (but not the Ordovician and Silurian) of the Basque Pyrenees (Dunoyer de Segonzac and Heddebaut, 1971; see Fig. 5-20), and (4) the earlier-mentioned Helvetic terrane of the Glarus Alps. There is also ample evidence for the formation of rectorite (or allevardite) under conditions of incipient metamorphism even in the absence of pyrophyllite, which can usually be related to the high alkali contents. Logvinenko ( 1964) regards the occurrence of allevardite and allevardite-like minerals in phyllitic shales, from which kaolinite and irregular illitemontmorillonite mixed-layers have already disappeared, as due to the effects of incipient metamorphism. In the roof of the Bramscher Massiv, Westphalia, allevardite occurs in the Liassic shales of Vehrte, northeast of Osnabruck (“Vehrter schwarzen Kreide”), in association with corrensite (Eckhardt, 1958). The lean-coal to anthracite rank of the organic matter (Stadler, 1971b, p.488) attests to the advanced stage of diagenesis. The main occurrence of Na-rich micas in the anchizone thus appears to be in the form of regular mixed-layers, either with montmorillonite, rather than as a discrete paragonite phase. Origin of rectorite (or allevardite) The above-mentioned rectorite (or allevardite) occurrences, with the possible exception of the last one, can unequivocally be ascribed to incipient metamorphism. In the widely-studied Japanese occurrences (Sudo et al., 1962; Brindley and Sandalaki, 1963) and other areas listed by Dunoyer de Segonzac (1970, p. 309), however, rectopite is of hydrothermal origin, also usually associated with pyrophyllite. In the geothermal area of Broadlands, New Zealand, rectorite-like minerals occur at higher temperatures than the more common illite-montmorillonite mixed-layers (Eslinger and Savin, 1973a). These relationships point towards some similarity between conditions of formation during incipient metamorphism and in the hydrothermal zone (Dunoyer de Segonzac, 1970, p. 312). According to Henderson ( 1970), rectorite in the pyrophyllite-bearing carbonaceous Manning Canyon shales of north central Utah formed from paragonitic mica during late stages of diagenesis as an intermediate metastable phase, which subsequently was altered to pyrophyllite during prograde low-grade metamorphism. The mineral assemblage pyrophyllite-rectoritemuscovite found by Henderson suggests - presumably metastable - compatibility of rectorite under these conditions in alkali-deficient environments not only with muscovite (low Naf content) or pyrophyllite (K deficiency), but with both of these minerals in the presence of excess water. Rectorite (or allevardite) admittedly appears at a more advanced stage of +

369 incipient metamorphism than the earlier-mentioned regular K-micamontmorillonite mixed-layers. Nevertheless, the phase relationships found by Henderson could be analogous to Velde's ( 1969) experimental evidence that in K-poor compositions in the muscovite-pyrophyllite compositional join, pyrophyllite coexists with an ordered K-mica-montmorillonite mixedlayer between 310" and 420" at 1000 barspH20,and between 310" and 400" at 2000 bars p H Z OAbove . 300°C, the composition of Py,,, which corresponds to an ordered 30% K-montmorillonite phase, represents the maximum expandability present in a single phase in this join. At lower potassium contents (Py > 50), this phase was invariably accompanied by pyrophyllite. The regular paragonite-Na-montmorillonite mixed-layer phases (including the 1 : 1 synthetic analogue of rectorite), however, which were widely encountered in a subsequent experimental study of the reaction: 3 Na-montmorillonite

+ 2 albite

+

3 paragonite

+ 8 quartz,

were shown in the rate studies to be all transient and metastable products obtained during the transformation of the albite-Na;montmorillonite assemblage to paragonite-quartz (Chatterjee, 1973). Rectorite is apparently always less stable relative to the assemblage Na-montmorillonite-paragonite. The above reaction took place around 335°C at 2kbar p H l o , i.e., at approximately the same temperature as the reaction: kaolinite

+ quartz

+

pyrophyllite

+ H,O,

at the same p H z O which , is in good agreement with the field data. When p H , O > Ptota,, however, pyrophyllite should appear before paragonite-quartz in a prograde sequence. Appearance of miscellaneous metamorphic minerals during incipient metamorphism of pelitic rocks Appearance of chloritoid Another greenschist-facies index mineral that has been claimed to appear in anchimetamorphic roofing slates, is chloritoid (Kubler, 1967a, p. 11 1; see also Winkler, 1970, p.204). It had also been reported from the "stage of early metagenesis" of the eastern Yenisey Mountains by Kasanskiy (1967). Subsequently, Kubler (1968, p. 395) has argued that the apparent anchimetamorphic illite 10 A peak widths in obviously higher-grade metamorphic chloritoid schists may merely reflect peak broadening due to insufficient resolution of the (001) diffraction peak of muscovite at lOA from those of paragonite at 9.7 A or of pyrophyllite at 9.2 A (cf. p. 352 of this chapter), and that chloritoid only appears in the greenschist facies proper.

370 In the Quartenschiefer of eastern Switzerland, chloritoid only appears in the Urseren-Garvera zone in the quartz-albite-epidote-biotite subfacies of the greenschist facies (Frey, 1969a, pp. 101, 113-1 17; 1970, table I). Chloritoid has admittedly been reported from supposedly anchimetamorphic Paleozoic slates of the Armorican Massif, western France: from the Ch2teaulin Basin by Sagon (1965) and from Ordovician slates of other parts of the Massif by Boudier and Nicolas (1968) and Le Corre (1969, 1975). It is present, however, only in horizons very rich in both aluminium and iron. Moreover, Sagon and Dunoyer de Segonzac (1972) have argued that the 10A peak widths of most of the micas from the Chiiteaulin Basin are “epizonal”. They ascribe the low anchizonal “crystallinities” of some of the chloritoidbearing samples to peak-broadening due to the contribution of the 9.7 A peak of margarite. The lower stability limit of chloritoid is dependent on the partial pressure of oxygen (Hoschek, 1969; Ganguly, 1969; Kramm, 1973). Ganguly (1968) has argued that the curve defining the lower stability limit of chloritoid has a positive slope on a fo2-T plot. Kramm (1973) found that the chloritoid, occurring in the very low-grade greenschist facies (Mn-, Fe- and Al-rich Paleozoic slates of the southern Venn-Stavelot Massif, Ardennes), is unusually rich in Mn. Its appearance near the lower stability limit is restricted to host rocks with oxidation ratios 2 Fe,O, X 100/2 Fe,O, FeO of less than 85-90. The Mn and Mg contents of the chloritoids increase with increasing oxidation ratio. At oxidation ratios of 90 to 100, kaolinite, which is being replaced by pyrophyllite, and andalusite may be present instead of chloritoid in isometamorphic rocks. The appearance of chloritoid in the anchimetamorphic Middle Devonian black slates of the Basque Pyrenees is considered by Dunoyer de Segonzac and Heddebaut (1971) either (1) to constitute an additional marker of the anchizone under reducing conditions, or (2) to result from a thermal metamorphic event subsequent to the Hercynian anchimetamorphism. At the present-day state of knowledge, it is impossible to exclude the possibility of the local appearance of chloritoid at a metamorphic stage lower than the beginning of the greenschist facies. Although the reported occurrences of chloritoid in the anchizone appear to be restricted to rather extreme chemical rock compositions, such as very high Fe/Mg ratios, high Mn contents, and low oxidation ratios in aluminous rocks, chloritoid cannot unequivocally be regarded as an indicator of the beginning of the epizone.

+

Occurrence of stilpnornelane The occurrence of stilpnomelane in the anchizone is better established than that of chloritoid. This mineral is common not only in the glaucophane - lawsonite - schist facies, but also in the prehnite - pumpellyite -

37 1 metagraywacke facies (see references in IOsch, 1974, pp. 100-101, and p. 468, this chapter). Glauconitic horizons appear to be particularly favorable to the formation of stilpnomelane during incipient metamorphism. In Haute-Savoie, stilpnomelane appears in the Cretaceous glauconitic sandstones associated with “Gres de Taveyanne” metamorphosed in the lowestgrade part of the prehnite-pumpellyite-metagraywacke facies (Martini and Vuagnat, 1970, p. 58; Martini, 1972, p. 262). In the Helvetic zone of the Glarus Alps of eastern Switzerland, in which Frey (1969a, 1969b, 1970) earlier described the anchimetamorphism of Upper-Triassic and Lower-Liassic shales, Frey et al. (1973) have studied the effects of the progressive increase in grade of the low-grade Alpine metamorphism in the Cretaceous and lower Tertiary glauconitic horizons. The initial replacement of glauconite by stilpnomelane in the central Glarus Alps marks the beginning of zone 11, which covers the middle- and high-grade part of the anchizone (see Fig. 5-2 1). Appearance of stilpnomelane in the Glarnisch area is associated with the illite l O A peak widths of about 6.0 mm in the accompanying marly shales (Frey et al., 1973, fig. 6 ahd p. 212), i.e., well beyond the beginning of the anchizone, which is marked by illite 10A peak widths of 7.5 mm under the experimental conditions adopted. In the southern Glarus Alps in the southern part of zone 11, stilpnomelane has entirely replaced the glauconite. Further south, in the autochthonous cover of the eastern Aar Massif (e.g., in the Kistenpass area), stilpnomelane is accompanied by brown biotite and the illite 10A peak widths are largely narrower than 4 mm, i.e., “epizonal”. Similarly, Durney ( 1974) reported a mid-anchizonal illite “crystallinity” for a slate associated with the first appearance of stilpnomelane in the parautochthonous Mesozoic deposits of the Morcles area (southwestern Switzerland). The occurrence of stilpnomelane extends well beyond the high-grade limit of the anchizone. Occurrence of stilpnomelane has also been reported from quartzite-like sandstones of the Russian Platform that have reached the “stage of initial metagenesis”. Thus, in most areas studied stilpnomelane appears within the anchizone; however, locally it has been reported in lower - grade zones within the prehnite - pumpellyite metagraywacke facies (p. 468). Other minerals appearing during incipient metamorphism In the glauconite-bearing horizons of the Helvetic zone of eastern Switzerland mentioned above, the alkali - amphiboles magnesioriebeckite (“rhodusite”) and riebeckite locally appear in the anchizone (e.g., Frey, 1970). At the Urnersee and in the Glarnisch these amphiboles appear already during diagenesis, before the beginning of the anchizone. According to Frey et al. (1973, pp. 195 and 212), virtually all of these occurrences are

372

0 Glauconite 8 Glauconite

+ Stilpnomelane Stilpnomelane 0 Stilpnomelane + Biotite A Alkali-amphibole( Riebeckite) 70

Fig. 5-2 1. Distribution of lowest-grade metamorphic index minerals in glauconitic horizons of Cretaceous and Tertiary age in the Helvetic zone of the Glarus Alps, Switzerland. Zone I corresponds to part of the “diagenetic” zone and the low-grade part of the anchizone, zone I1 to the middle- and high-grade part of the anchizone, and zone 111 to the low-grade part of the greenschist facies on the basis of illite “crystallinity” (Frey, 1970). Distribution of some Alpine metamorphic minerals in other rocks in the same area is shown in Fig. 5-18. The numbers indicate K-Ar ages of glauconites in m.y. (After Frey et al., 1973, fig. 3.)

373 restricted to one particularly Na-rich horizon and are related to slickensides and joints. Coal ranks associated with anchimetamorphism The coal ranks associated with various stages of late diagenesis and incipient metamorphism have formed the subject of a number of papers by the writer. Only the main conclusions of a recent study of coal ranks in the anchizone (Kisch, 1974), however, are presented here. Exclusively anthracitic and low-rank meta-anthracitic coal and coaly matter (i.e., with less than 8% volatile matter (d.a.f.) and more than 2.5% have been found in association with anchimetamorphic illite “crystallinities”, pyrophyllite, and rectorite-allevardi te of the following anchizonal occurrences : (1) The Jurassic “Terres Noires” of the Gap-Embrun and Barcelonnette areas in the subalpine belt of the French Alps (Dunoyer de Segonzac et al., 1966; Artru et a]., 1969; see Fig. 5-17). (2)The Silurian and Devonian deposits in the well Assejmi 1, Tindouf Basin, western Sahara (Dunoyer de Segonzac, 1969, pp. 113- 119). (3) Underclays of Upper-Carboniferous coals of the Pennsylvania anthracite regions (Hosterman et al., 1960; Spackman and Moses, 1961). (4)Shales and sandstones from the roof of the Cretaceous “Bramscher Massiv” pluton near Osnabruck, Westphalia, Western Germany (Stadler, Teichmuller, and Teichmuller, unpublished, in: Teichmiiller and Teichmiiller, 1966a,b; Schreyer, 1969). ( 5 ) Tonsteins in Upper-Carboniferous coal seams in the southern Belledonne area near Grenoble, subalpine belt of the French Alps (Scherp et al., 1968; Stadler, 1971a). In some cases, anthracitic rank is reached in the higher-grade part of the “diagenetic” zone (as delimited by illite “crystallinities”, in which irregular illite-smectite mixed-layers and kaolinite are absent. Examples of such occurrences include (for details see Kisch, 1974): (1) The Jurassic “Terres Noires” in the deepest intervals of the wells Grand Luberon and Montagne de Lure (Dunoyer de Segonzac, 1969, pp. 167-170), and in the Gap area in the subalpine belt of the French Alps (Dunoyer de Segonzac et al., 1966; Artru et al., 1969). (2) The Upper Triassic Quartenschiefer at Quarten, Helvetic zone of the northernmost Glarus Alps (Frey, 1969a, 1970).

’

’

References given pertain to the descriptions of the anchizonal mineralogy. For references to the coal-rank determinations see Kisch (1974, particularly table 11).

374 (3)The Paleozoic deposits of the northeastern part of the Rheinische Schiefergebirge where the coaly matter, which is associated with anchimetamorphic illite “crystallinities” and pyrophyllite (Weber, 1972b; also Scherp et al., 1968), is exclusively high-rank anthracite, with Ro,,(vlt,,nlte) of 4%, and meta-anthracite (Wolf, 1972). (4) Jurassic cover of the southern termination of the “zone houillere Briangonnaise”, the Carboniferous coals of which are even higher-rank anthracites (e.g., Robert, 1971); both anchizonal illite “crystallinities” (Abbas, 1974) and high-rank anthracite (Chateauneuf et al., 1973) have been reported here. Thus, one can conclude that the appearance of anchimetamorphic illite “crystallinities”, pyrophyllite, and rectorite (or allevardite) takes place within the anthracite range of coal ranks. Conversely, the attainment of anthracite rank is associated with lower than anchimetamorphic illite “crystallinity indices” and, in many cases, with the occurrence of kaolinite ( e g , in underclays and tonsteins). Identification of the anchizone: a summary Although the anchizone is mainly defined by a range of illite “crystallinity indices”, it shows a number of additional characteristics that allow its recognition, even where i t might otherwise be overlooked in the absence of illite lOA peak widths or in the case of anomalous “crystallinity” data. Some distinctive minerals, including pyrophyllite, paragonite and, probably, stilpnomelane, are common to both the anchizone and the “epizone”. Their occurrence only indicates that the low-grade boundary of the anchizone has been exceeded. The former two minerals are particularly important in view of the tendency towards abnormally low illite “crystallinities” (i.e., 10 A peak broadening) in rocks containing these minerals. The high-grade limit of the anchizone in pelitic rocks is somewhat more difficult to establish in the absence of illite “crystallinity” data. I t can be considered to have been exceeded upon appearance of typical greenschistfacies minerals, such as biotite, andalusite, and, probably, also margarite and Mg-rich chloritoid. Negative characteristics for attainment of the anchizone must be used with caution. Inasmuch as absence of irregular illite-smectite mixed-layers and of kaolinite could be due to sedimentary control, it can be considered to be due to postdepositional processes only if it can be argued that the rocks were originally rich in smectite or kaolinite content (e.g., bentonites, underclays, and tonsteins). Anthracite and low meta-anthracite rank may also be regarded as primafacie indication for anchizonal environment, although there is an overlap

375 with both the “diagenetic” and the “epizonal” realms sensu Kubler. Finally, appearance of slaty cleavage may also serve as an indication that the anchizone has been reached, despite the possibility of its local development at lower-grade conditions. Towards a scheme of lowest-grade mineral facies for clastic rocks

As described in this chapter sequential changes in the mineralogy of clastic rocks upon deep burial can be used to define grades or stages of burial and incipient metamorphism in clastic sedimentary sequences. Establishment of generally applicable stages of this kind is likely to contribute greatly to various aspects of the study and understanding of sedimentary basins (e.g., in the exploration for fossil fuels and the reconstruction of burial history and tectonic development) and of marginal orogenic zones. Consequently, correlations must be established between the several sequential changes in mineralogy discussed above: ( 1) the decreasing content of swelling layers and the appearance of stacking order in illite-smectite mixed-layers, (2) the polytypes and ultimate disappearance of kaolinite group minerals, and (3) coal rank. For deeper levels and more advanced stages correlations must be established between illite “crystallinity”, polytypes of illite and chlorite, and the appearance of pyrophyllite and rectoriteallevardite. The changes in a given sedimentary mineral after burial, however, are controlled not only by temperature and pressure, but also to a major extent, by the chemical environment. The latter is a function of lithologic variables, such as the mineralogy of the clastic constituents, the composition of pore solutions, and the ability of pore solutions to circulate (i.e., permeability). The chemical factors affecting the stage of burial at which a particular discontinuous mineralogic change begins or is completed, or a continuous change attains a certain defined stage, are so numerous that no uniform relationships among various processes can be expected except by specifying reductions in the number of lithological-chemical variables. Thus, studies of the relationships among specific mineral-replacement processes should be restricted to a small number of well-defined ranges of lithology, e.g., siltstones rich in unstable K-rich detritus and minimal carbonate cement content, which could adversely affect permeability, or K-, Fe- and Mg-poor claystones having low permeability. An attempt in this direction was made by Kossovskaya and Shutov (1963) by developing a scheme for “regional epigenesis and metagenesis”, which allows for major differences in primary lithology by stating diagnostic mineral assemblages (“facies”) in each “stage” for each of five compositional “rock families”. An amplified and modified version of their scheme is

w

TABLE 5-V

4 o\

Schematic mineral distribution during the "stages" of regional epigenesis and metagenesis in clastic sedimentary rocks after Kossovskaya and Shutov (1961, 1963) and the approximately equivalent illite-"crystallinity" zones of Kubler (l967a, 1968) and their correlation with mineral distribution and metamorphic mineral facies (after Coombs et al., 1959 and Coombs, 1960, 1961) in intermediate volcanic graywackes and tuffs Schematic. mineral distribution'.

"stages" Of

cplgenesls and metagenesis

and "facies"

of regional eptgenesis and metagenests in d i f f e r e n t "families " of clastic rocks2 rock54

-

Quartz kaolinltc

Mineral zones of KUBLER (1967. 1968 i 3 major CharacteriSticS

Schematic mineral d i s t r i b ~ t i o n ~ and . metamorphic mineral facies after COOMBS (1960,1961)5, In Intermedmte volcanic graywackes and tufts

"L~thoclastlc sandstones" and "acid arkoses"

HEULANDITE-ANALCIME ZONE of t h e ZEOLITE FACIES (Coombs et al ,1959)

INITIAL EPlGENESlS

"heulandtte -analcimc l a c ~ 5 " (Kossovskaya and Shutov 1961.1963) "DIAGENETIC" ZONE i l l i t e crystallinities '7.5 [>4 irreaular illite/montmortllonite m-l kaoltnitc-group minerals

01

DEEP EPlGENESlS

EARLY METAGENESI!

Absent' pyrophyllite, rectorite, paragontte

ANCHMETAMORPHIC ZONE illite c r y s t a l l i n i t y index 7.5- 4.0 [ 4 0 - 2 . 5 1 pyrophyllite, r e c t o r i t e paragonite. phengitc

LAUMONTITE ZONE of the ZEOLITE FACIES (Coombs ct a1 ,1959) "laurnontlte facles" ( KOSSoVskaya and Shutov ,1961,1963)

PREHNITE-PUMPELLYITTE-METAGRAYWACKE FACIES (Coombs et al , 1 9 6 0 ) "prehnitc-muscovite-chlorite tacies" (Kossovskaya and Shutov. 1963)

Absent. irregular Illlte/

LATE METAGENESI!

Low-grade part of the GREENSCHIST FACIES "

E PI ZONE "

illite c r y s t a l l m t y index ZMi Mite (and polytype) --- Phengite Kaolinite Pyrophyllite Irregular illitejrnontmorillonite m-1 ? Regular illite /montmorillonite m-l (allev. ?Regulor paragonite jphengite rn-l Parogonite . _ . Regular chlorite/montmorlllonlte rn - I Chlorite Chloritoid Biotite >?omn 7-13 mrn lllite crystallinity**

_ _ _ _ - - -2 M 1 --

2M 1

, 7

---

__

LindaL Frick

Localities studied in Qtail by Frey ngures showing typical textures (Frey. 1969a)

Ouarten

32 B

32A

4.5-7.5rnm

Dia3enesis

Anchimetarnorphism

“Epimetornorphism”

Diagenesis I

I

I

I

8. Mesozoic and Eocene glauconitic hOriZOnS Glauconite Stilpnornelane Riebeckite Biotite 01.

< 4mm

< 4 5rnrn

Lirnrnernbmen Garvera Wnixerpass. Kistenpass 320

~

~

Metamorphic zones (based on illite crystallinity and presence of pyrophyllite and paragon i t e )

Zones of Frey et

4-53mm

Guggenegg. W. of Kobelruns,Wiss Linthal rnilen, Laui furkel 32C

1

I

(1973)

C. Volanic graywackes (“Gres de Taveyanne”) and associated slates Laumontite PurnDellvlte Prehnite Illite crystallinity’* (after Frey and Niggli, 1971)

, I

I

1

Fig. 5-43.Ranges of occurrences of some minerals in the Helvetic zone of eastern Switzerland, compiled after Frey (1969a, 1969b, 1970), Frey and Niggli (1971) and Frey et al. (1973). * Including the southernmost Glarus Alps. ** Low-grade limit of anchizone: 74 mm; of “epizone”: 4 mrn. *** In the Wageten, west of Lake Walen (Frey and Niggli, 1971, p. 232).

somewhat higher-grade than the boundary between that facies and the higher-grade pumpellyite-actinolite facies. Illite “crystallinity and paragonite associated with lawsonite-bearing pumpelbite facies. Data on these crystallinities can be obtained from a comparison of the lowest-grade Alpine metamorphic assemblages of the “zone houillere brianGonnaise” in the Pennine realm of the French Alps with the illite “crystallinites” in its Mesozoic sedimentary cover. Prehnite and pumpellyite are common in the external centra1 massif of the Pelvoux and its sedimentary cover and in the external, westernmost margin I’

47 1 of the “zone houillere brianqonnaise” (Saliot, 1973), accompanied in the Gres de Champsaur (the equivalent of the Taveyanne Graywacke of HauteSavoie and Switzerland) of the southern cover of the Pelvoux by laumontite (Martini and Vuagnat, 1965; Bocquet, 1971) and corrensite (Kubler, 1970; Aprahamian, 1974, p. 10). The illite lOA peak widths associated with the laumontite-bearing southwestern cover are still “diagenetic” (Aprahamian, 1974, p. 10). However, unequivocally anchimetamorphic (in part high-grade anchimetamorphic) illite 10 A peak widths predominate in the autochthonous and para-autochthonous Lower Tertiary flysch cover to the southeast and the northeast of the Pelvoux, whereas those in the Liassic autochthonous and parautochthonous Lower Tertiary flysch cover to the southlimit, and are accompanied by phengite-paragonite mixed-layer or paragonite (Kubler, 1970; Aprahamian, 1974, pp. 10-1 1). In the more internal, eastern part of the “zone houillere”, lawsonite has been reported from microdiorites west of Modane, with prehnite according to Fabre (1961, p.64; cf. Bocquet, 1971, p.98), but according to Bearth (1962) and Guitard and Saliot (1973, pp. 508-5 10) usually with pumpellyite without prehnite or actinolite. Guitard and Saliot (1971) and Saliot (1973) report several additional occurrences of lawsonite and/or pumpellyite from the “zone houillere brianqonnaise” further to the south. Predominance of low-grade “epimetamorphic” illite 10 A peak widths (2.5-3.4 mm) were reported in the Upper-Cretaceous calc-schists (“marbres en plaquettes”) of the sedimentary cover of the central part of the Brianqonnais zone (from around Brianqon in the north to the Guillestre area in the south) by Dunoyer de Segonzac (1969, pp. 176-180); some of the samples contain paragonite. In the more internal eastern margin of the Brianqonnais zone (Col d’Izoard area and Cristillan Valley), the lOA peak widths (2.2-2.7 mm at the conditions used) are higher-grade “epimetamorphic”, and several of the samples contain paragonite. Abbas ( 1974) has similarly found predominance of low-grade “epimetamorphic” lOA peak widths (means of 3.3 and 3.1 mm, respectively) and occurrence of paragonite in two sections of Rhaetian shales and dolomites at Charvie and Haut-Cristillan in the discontinuous “pre-piemontaise” zone which has a structural position intermediate between the internal Brianqonnais zone and the more internal, easterly Piemont zone (“schistes lustres” nappel. Predominance of low-grade anchimetamorphic 10 A peak widths (mean value = 5.1 mm), on the other hand, has been found by Abbas (1974) in the Rhaetian shales and dolomites of the Peyre-Haute nappe (part of the Brianqonnais zone, and overlying the “zone houillere”) at Lac d’Ascension, some 12 km south of Brianqon and 8 km southwest of the Col d’Izoard. One may note that the anthracitic coal ranks associated with the pumpel-

472 lyite- and lawsonite-bearing “zone houillere” (see p. 456) are quite similar to those found associated with prehnite-pumpellyite facies of Taveyanne Graywacke’in central and eastern Switzerland (Kisch, 1980b; cf. p. 456). These relationships appear to indicate that in the Briangonnais zone the pumpellyite- and lawsonite-bearing assemblages are largely associated with “epimetamorphic” illite 10 A peak widths and anthracite coal ranks, and that such “crystallinities” are thus attained well before the appearance of abundant Alpine glaucophane and jadeite to the east. These minerals appear immediately east of the Briangonnais zone sensu stricto, in the Permo-Trias of the “ecailles intermediaires” of Rio Secco and Chaberton (near the Montgenevre Pass, NE of Briangon), underlying the above-mentioned “pre-piemontaise” zone, and the Aceglio zone to the south (Bocquet, 1971, pp. 94-95, 99; Bocquet, 1974, pp. 165, 168; Saliot, 1973, Fig.5-30). The illite peak widths of associated rocks from the upper Valle Varai ta (northwestern Aceglio zone) reported by Dunoyer de Segonzac and Hickel ( 1972) are unequivocally “epimetamorphic”. Unequivocally “epimetamorphic” mica 10 A peak widths and presence of minor paragonite have also been found by this writer (Kisch, in preparation) to be associated with the lawsonite-pumpellyite (without actinolite or prehnite) zone in the Caples Group of the Humboldt Mountains in the upper Wakatipu district, western Otago, N.Z. (zone I1 of Kawachi, 1974, 1975). These very scanty data seem to suggest that at the metamorphic gradients giving rise to the appearance of lawsonite in pumpellyite-facies metamorphic terranes- presumably at low thermal gradients during metamorphism-“epimetamorphic” illite “crystallinities” are attained at a comparatively low degree of incipient metamorphism. This may be an expression of the existence of a pressure effect in enhancing the development of mica crystallinity at a given metamorphic temperature. Relationship of coal rank and authigenic layer silicates to burial-metamorphic facies in volcanic and lithic-feldspathic sediments in different lowest-grade metamorphic ‘yacies series” In the preceding sections the writer has noted the incorporation of the laumontite zone of the zeolite facies in Kossovskaya and Shutov’s “stage of deep epigenesis”. This incorporation, originally suggested by Kossovskaya and Shutov (1961), has been confirmed and refined by comparisons between the maximum coal ranks associated (either directly or as obtained by downward extrapolation) with laumontite-bearing mineral assemblages in various areas (low-volatile to semi-anthracite), and those associated with the “stage of deep epigenesis” or Kubler’s zone of “diagenesis” (or “non-metamorphic” zone) in other areas (Kisch, 1969, 1974). The writer has also arrived at a tentative correlation of the boundary

473

between the laumontite zone and the (zeolite-free) prehnite-pumpellyite facies with the high-grade part of “deep epigenesis”, or the transition between this stage and the higher-grade “early metagenesis” (this transition is equivalent to the “diagenesis”-anchimetamorphism boundary of Kubler). This correlation is based (1) on the above comparisons of ranks associated with laumontite-bearing assemblages, and (2) on direct correlations of zeolite-facies-prehnite-pumpellyite-facies successions with the associated modifications in layer-silicate mineralogy in very few areas (with Kossovskaya and Shutov’s zones in the Permian and Jurassic of the western Verkhoyanye, Siberia, and with illite crystallinity in the Lower Tertiary of the Helvetic zone of Switzerland and Haute-Savoie, and in the Permian of Otago, N.Z.). The correlation requires further refinement. Even then, it is valid only for terranes like the above, where a laumontite zone is succeeded by a zeolite-free prehnite-pumpellyite zone and then by a pumpellyite-actinolite zone, i.e., the intermediate- and high-intermediatepressure type “facies series” of Table 5-VII. Such a relationship cannot automatically be extrapolated to incipientmetamorphic terranes with different temperature-pressure gradients. I t remains to be seen, for instance, if the onset of pumpellyitic facies is associated with similar coal ranks and stages of clay-mineral alteration in such different temperature-pressure gradients-“facies series”-of lowest-grade metamorphism. For instance, incomplete evidence in terranes in which a lawsonite-bearing pumpellyite facies occurs (as a lower-grade predecessor to blueschist facies) indicates that the onset of such pumpellyitic facies could well be associated with different coal ranks and illite “crystallinities”. In the BrianGonnais zone of the French Alps. for instance, the coal ranks associated with the lawsonite-bearing pumpellyite facies appear to be anthracites, but the illite 10A peak widths are already characteristic for “epimetamorphism”. No illite “crystallinity” data are available for the presumably even higher-pressure (and lower-temperature) lawsonite- and pumpellyite-bearing metagraywackes of the Franciscan in the Diablo Range, California, but the associated coal ranks are appreciably lower: high- to low-volatile bituminous. These differences may appear at first sight disconcerting. However, they probably reflect fundamental differences in the way in which the various parameters of burial metamorphism react to changes in temperature and pressure. It may, for instance, reflect the major influence of temperature and duration of heating on the coalification process, pressure having little or no effect, whereas pressure may provide a major contribution to the improved crystallinity of illite and the disappearance of expandable layers from illite-smectite mixed-layers.

474

The differences between the relationship of the various parameters in various incipient-metamorphic “facies series” may, therefore, point the way to further studies, in which these differences will actually be used and evaluated in evaluating temperature-pressure-time gradients of burial metamorphism.

ACKNOWLEDGEMENT

The author is greatly indebted to his colleague Dr. Reg Shagam (BeerSheva) for improving the style of parts of the manuscript, and for his continuous encouragement. This chapter has benefited substantially from the editing by Dr. G.V. Chilingar, who suggested various improvements. REFERENCES * Abbas, M.. 1974. Mktamorphisme des argiles dans le Rhetien des Alpes sud-occidentales. Etude mineralogique et geochimique. These docteur 3eme cycle, Universite Louis Pasteur, Strasbourg, 73 pp. Albee. A.L. and Zen, E-An, 1969. Dependence of the zeolitic facies on the chemical potentials of CO, and H,O. In: V.C. Zharikov (Editor), Contributions to Physico-chemical Petrology (Korzhinskii Vol.), pp. 249-260. Alexandersson, T., 1970. The sedimentary xenoliths from Surtsey: marine sediments lithified on the sea-floor. Surtsey Res. Progr. Rep., V: 83-89. Althaus, E., 1969. Das System Al ,O, -SO, -H,O. Experimentelle Untersuchungen und Folgerungen fur die Petrogenese der metamorphen Gesteine. Teil I: Apparative und experimentelle Grundlagen; die Stabilitatsbedingungen der hydroxylhaltigen Aluminiumsilikate. Neues Jahrb. Mineral., Abh., 111 (1): 74-1 10. Aprahamian, J., 1974. La cristallinite de I’illite et les mineraux argileux en bordure des massifs cristallins externes de Belledonne et du Pelvoux. Geol. Alpine, 50: 5-15. Aronson, J.L. and Hower, J., 1976. Mechanism of burial metamorphism of argillaceous sediment, 2. Radiogenic argon evidence. Geol. SOC.Am. Bull., 87 (5): 738-744. Artru, Ph. and Gauthier, J., 1968. Evolution geometrique et diagenetique d’une serie miogeosynclinale (Lias inferieur a Berriasien) d’apres I’etude du sondage de Valvigneres (France sud-est). Bull. Centre Rech. Pau-SNPA, 2 ( I ) : 101-1 16. Artru, Ph.. Dunoyer de Segonzac, G., Combaz, A. and Giraud, A., 1969. Variations d’origine sedimentaire et evolution diagenetique des caracteres palynologiques et geochimiques des Terres Noires jurassiques en direction de I’Arc Alpin (France, sud-est). Bull. Centre Rech. Pau-SNPA, 3(2): 357-376. Bailey, S.W. and Brown, B.E., 1962. Chlorite polytypism: I. Regular and semi-random one-layer structures. Am. Mineral., 47: 819-850. Banno. S., 1964. Petrologic studies on Sanbagawa crystalline schists in the Bessi-Ino district, central Sikoku, Japan. Unio. Tokyo J . Fac. Sci., 15: 203-319.

* References up to 1976; reference list finalized in February, 1977. For literature since 1976, see Appendix B.

475 Barlier, J., 1974. Recherches Paleothermometriques dans le Domaine des Terms Noires Suhulpines Meridionafes. These docteur 3eme cycle. Univ. Paris-Sud (Centre d’Orsay). 96 pp. Barlier. J., Ragot. J.-P. and Touray, J.-C.. 1974. L’evolution des Terres Noires subalpines meridionales d’apres I’analyse mineralogique des argiles et la reflectomktrie des particules carbonees. Bull. B.R.G.M. (2eme Serie), 5: 533-548. Beall. A.O. Jr. and Ojakangas, R.W., 1967. Mineralogy of an Upper Cambrian K-bentonite from Missouri. J . Sediment. Petrol.. 37(3): 952-976. Bearth, P., 1962. Versuch einer Gliederung alpin-metamorpher Serien der Westalpen. Sdnt.ei:. Mineral. Petrogr. Mitt., 42: 127- 137. Beuf. S., Biju-Duval, B.. Stevaux, J. and Kulbicki. G.. 1966. Ampleur des glaciations “siluriennes” au Sahara: leurs influences et leurs consequences s u r la sedimentation. Rev. Inst. Fr. Petrol., 21: 363-381. Beugnies, A.. Godfriaux, L. and Robaszynski. F.. 1969. Contribution a l’etude des phengites. Bull. SOC.Beige GeoI. Paleontol. Hvdrol.. 77: 95- 146. Biljon, W.J. van and Bensch, J.J.. 1970. The “crystallinity” of illite as a measure of contact metamorphism in mudstone of the Karroo System, South Africa. In: Second Goriht.urur Symposium, Proceedings and Papers. C.S.I.R.. Pretoria. South Africa. pp. 45 1-453. Bishop, D.G., 1972a. Authigenic pumpellyite and other metamorphic effects in the Kveburn Formation, central Otago. N . Z . J . Geol. Geoph.vs.. 15(2): 243-250. Bishop, D.G.. 1972b. Progressive metamorphism from prehnite-pumpellyite to greenschist facies in the Dansey Pass area, Otago. New Zealand. Geol. Soc. Am. Bull.. 83: 3177-3198. Bocquet. J., 1971. Cartes de repartition de quelques mineraux du metamorphisnie alpin dans les Alpes franco-italiennes. Eclogue Geol. Helu.. 64( I ) : 71- 103. Bocquet. J.. 1974. Etudes Mineralogiques cr Petrologiques sur les MPtamorphisnies cl‘A^ge Alpiri dans les Alpes Franqaises. These. Univ. Grenoble. 489 pp. Boles, J.R. and Coombs. D.S., 1975. Mineral reactions in zeolitic Triassic tuff. Hokonui Hills. New Zealand. Geol. Soc. Am. Bull.. 86(2): 163-173. Bostick. N.H., 1971. Thermal alteration of clastic organic particles as an indicator of contact and burial metamorphism in sedimentary rocks. Geosci. Man. 3: 83-92. Bostick. N.H.. 1974. Phytoclasts as indicators of thermal metamorphism. Franciscan assemblage and Great Valley sequence (Upper Mesozoic). California. In: R.R. Dutcher. P.A. Hacquebard. J.M. Schopf and J.A. Simon (Editors). Carbonaceous Materials as Indi(.ator.T of Metamorphism. Geol. Soc. A m . Spec. Pap.. 163: 1-17. Boudier, F. and Nicolas. A,. 1968. Decouverte de chloritoide dans les schistes ardoisiers d’Angers. Bull. Soc. Fr. Mineral. Cristallogr.. 91 : 92-94. Bradley, W.F., 1950. The alternating layer sequence of rectorite. Am. Mineral.. 35: 590-595. Brindley, G.W.. 1956. Allevardite, a swelling double-layer mica mineral. Am. Minerd.. 41: 91-103. Brindley, G.W. and Sandalaki. Z.. 1963. Structure, composition and genesis of some longspacing, mica-like minerals. A m . Mineral., 48: 138- 149. Brown, C.E., 1961. Prehnite-pumpellyite metagraywacke facies of Upper Triassic rocks. Aldrich Mountains. Oregon. U.S. Geol. Surv., Prof: Pap.. 424-C: 146- 147. Brown, C.E. and Thayer, T.P.. 1963. Low-grade mineral facies in Upper Triassic and Lower Jurassic rocks of the Aldrich Mountains, Oregon. J . Sediment. Perrol., 33(2): 41 1-425. Brown, E.H.. 1967. The greenschist facies in part of Eastern Otago, New Zealand. Conrrih. Mineral. Petrol., 14: 259-292. Brown, G. and Weir. A.H.. 1963. The identity of rectorite and allevardite. In!. C1u.v. Cant 1963, Stockholm Proc.. 1. 27-35. Brown, G., Catt, J.A. and Weir, A.H.. 1969. Zeolites of the clinoptilolite-heulandite type in sediments of south-east England. Mineral. Mag., 37(288): 480-488. I

476 Browne, P.R.L. and Ellis, A.J.. 1970. The Ohaki-Broadlands hydrothermal area, New Zealand: mineralogy and related geochemistry. Am. J . Sci., 269: 97-131. Burst, J.F., Jr., 1959. Postdiagenetic clay mineral environmental relationships in the Gulf Coast Eocene. Clays Clqv Miner., Proc. 6th Nail. Confi, pp. 327-341. Burst, J.F., 1969. Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. Am. Assoc. Pet. Geol. Bull., 53( I ) : 73-93. Buryanova, Ye.Z. and Bogdanov. V.V., 1967. Distribution of the authigenic zeolites laumontite and heulandite in the sedimentary rocks of the Tarbagatai coal deposits. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 1967 (2): 195-202. Bystrom, A.M., 1956. Mineralogy of the Ordovician bentonite beds at Kinnekulle, Sweden. Suer. Geol. Unders. (Ser. C), 540: 62 pp. Caillere, S., Mathieu-Sicaud, A. and Henin, S., 1950. Nouvel essai d'identification du mineral de La Table pres d'Allevard, I'allevardite. Bull. Soc. Fr. MinPral. Cristallogr., 73: 193-201. Cameron, L.B. and Sabine, P.A., 1969. The Tertiary welded-tuff vent agglomerate and associated rocks at Sandy Braes, Co. Antrim. Inst. Geol. Sci., Rep. No. 69/6: 15 pp. Campbell, A.S. and Fyfe, W.S., 1965. Analcime-albite equilibria. Am. J . Sci., 263: 807-816. Carman, M.F., 1965. Nature of chlorite in some low-grade metamorphic rock in South Island, New Zealand. Geol. Soc. Am., Program 1965 Ann. Meet., Kansas City, p. 27. Carrigy, M.A. and Mellon, G.B., 1964. Authigenic clay mineral cements iQ Cretaceous and Tertiary sandstones of Alberta. J . Sediment. Petrol., 34(3): 461 -472. Cassan. J.P. and Lucas, J., 1966. La diagenese des gres argileux d'Hassi-Messaoud (Sahara): silicification et dickitisation. Bull. Seru. Carte GPO/.Alsace Lorraine, 19: 241-253. Castaiio, J.R. and Sparks, D.M., 1974. Interpretation of vitrine reflectance measurements in sedimentary rocks and determination of burial history, using vitrinite reflectance and authigenic minerals. In: R.R. Dutcher, P.A. Hacquebard, J.M. Schopf and J.A. Simon (Editors), Carbonaceous Materials as Indicators of Metamorphism. Geol. Soc. Am. Spec. PUP., 153: 31-52. Chamley, H., 1967. Possibilite d'utilisation de la cristallinite d'un mineral argileux (illite) comme temoin climatique dans les sediments. C.R. Acad. Sci., Ser. D, 265: 184-187. Chamley, H., 1968. Sur le r d e de la fraction sedimentaire issue du continent comme indicateur climatique durant le Quaternaire. C.R. Acad. Sci., Ser. D,267: 1262- 1265. Chateauneuf. J.-J., Debelmas, J., Feys, R., Lemoine, M. and Ragot, J.-P., 1973. Premiers resultats d'une etude des charbons jurassiques de la zone brianqonnaise. C.R. Acad. Sci. Poris, Ser. D,276: 1649-1652. Chatterjee, N.D.. 1973. Low-temperature compatibility relations of the assembalge quartzparagonite and the thermodynamic status of the phase rectorite. Contrib. Mineral. Petrol., 42: 259-271. Chennaux. G. and Dunoyer de Segonzac, G., 1967. Etude petrographique de la pyrophyllite du Silurien et du Devonien au Sahara. Repartition et origine. Bull. Serv. Carre GkoI. Alsace Lorraine, 20(4): 195-210. Chennaux, G., Dunoyer de Segonzac, G. and Petracco, F., 1970. Genese de la pyrophyllite dans le Paleozoique du Sahara occidental. C.R. Acad. Sci., Ser. D, 270: 2405-2408. Clauer, N. and Lucas, J., 1970. Mineralogie de la fraction fine des schistes de Steige-Vosges septentrionales. Bull. Groupe Fr. Argiles, 22: 223-235. Coombs. D.S.. 1954. The nature and alteration of some Triassic sediments from Southland, New Zealand. R. SOC.N.Z. Trans.,82(1):65-109. Coombs, D.S., 1958. Zeolitized tuffs from the Kuttung Glacial Beds near Seaham, New South Wales. Aust. J . Sci., 21: 18-19. Coonibs. D.S.. 1960. Lower grade mineral facies in New Zealand. Rep. In/. Geol. Congr., 21st Sess.. Norden, 13: 339-351.

477 Coombs, D.S., 1961. Some recent work on the lower grades of metamorphism. Ausr. J . Scr.. 24(5): 203-215. Coombs, D.S., 1971. Present status of the zeolite facies. Advanc. Cheni.. Ser. 101: 317-327. Coombs, D.S. and Whetten. J.T.. 1967. Composition of analcime from sedimentary and burial metamorphic rocks. Geol. SOC.A m . Bull., 78: 269-282. Coombs, D.S.. Ellis. A.J.. Fyfe. W.S. and Taylor. A.M.. 1959. The zeolite facies, with comments o n the interpretation of hydrothermal syntheses. Geochin7. Cosniochim.A u u , 17: 53- 107. Coombs, D.S., Horodyski, R.J. and Naylor, R.S.. 1970. Occurrence of prehiiite-punipell~iIe facies in northern Maine. Am. J . Sci., 268: 142-156. Coombs, D.S.. Nakamura, Y. and Vaugnat. M.. 1976. Pumpellyite-actinolite facies schists of the Taveyanne Formation near Loeche. Valois. Switzerland. J . Petrol.. 17(4): 440-47 I . Correns, C.W., 1950. Zur Geochemie der Diagenese. I. Das Verhalten von CaCO, und SiO?. Geochim. Cosmochim. Acta. 1 : 49-54. De Roever. E. W.F., 1972. Lawsonite-albite-facies Metamorphism Near F u s d d o . Culahriu (Southern Italy), its Geological Significance and Petrological Aspecrs. Thesis Univ. Anisterdam, 171 pp. De Roever, W.P., 1947. Igneous and metamorphic rocks in eastern central Celebes. I n : Geological Explorations in the Island of Cer'ehes Under Leadqrship of H .A . Brou,t.er. North-Holland, Amsterdam, pp. 65-173. De Roever, W.P.. 1950. Preliminary notes on glaucophane-bearing and other crystalline schists from South East Celebes, and on the origin of glaucophane-bearing rocks. K. Ned. Akad. Wet. Proc., 53(9): 1455-1465. De Roever, W.P., 1955. Some remarks regarding the origin of glaucophane in the North Berkeley Hills, California. Am. J . Sci., 253: 240-244. De Roever, W.P.. 1972. Glaucophane problems. Tschermaks Mineral. Petrogr. Miri.. 3. Folge. 18(1): 64-75. Devereux, I., 1968. Oxygen isotope ratios of minerals from regionally metamorphosed schists of Otago. New Zealand. N . Z . J . Sci., 1 1 : 526-548. Deverin, L., 1924. L'etude lithologique des roches sedimentaires. Schuei;. Mineral. Petrogr. Mitt., 4: 29-50. Dickinson. W.R., 1962a. Metasomatic quartz keratophyre in central Oregon. Am. J . Sci., 260: 249-266. Dickinson, W.R., 1962b. Petrology and diagenesis of Jurassic andesitic strata in central Oregon. A m . J . Sci.. 260: 481-500. Dickinson. W.R.. Ojakangas, R.W. and Stewart, R.J.. 1969. Burial metamorphism of the late Mesozoic Great Valley Sequence. Cache Creek. California. Geol. Soc: Am. Bull.. 80: 5 19-526. Drits. V.A. and Shutov, V.D.. 1966. New variety of mixed-layer mineral of the sudoite group. Lithol. Miner. Resour. (translated from Litol. Pole:n. Iskop.), 1966(4): 130- 133. Dunoyer de Segonzac. G.. 1964. Les argiles du Cretace superieur dans le bassin de Douala (Cameroun): Problemes de diagenese. Bull. Serv. Carte Geol. Alsace Lorruine, 17(4): 287-3 10. Dunoyer de Segonzac. G.. 1968. The birth and development of the concept of diagenesis. Earth-Sci. Rev., 4: 153-201. Dunoyer de Segonzac, G.. 1969. Les mineraux argileux dans la diagenese; passage au metamorphisme. MPm. Serv. Carte Geol. Alsace Lorraine. 29: 320 pp. Dunoyer de Segonzac, G.. 1970. The transformation of clay minerals during diagenesis and low-grade metamorphism: a review. Sedimentologv. 15: 28 1-346.

47 8 Dunoyer de Segonzac, G. and Chamley, H., 1968. Sur le r d e joue par la pyrophyllite comme marqueur dans les cycles sedimentaires. C.R. Acad. Sci., Ser. D, 267: 247-277. Dunoyer de Segonzac, G . and Heddebaut, C., 1971. Paleozoique anchi-metamorphique a illite, chlorite, pyrophyllite, allevardite, et paragonite dans les Pyrenees Basques. Bull. Serv. Carte Geol. Alsace Lorraine, 24(4): 277-290. Dunoyer de Segonzac, G . and Hickel, D.. 1972. Cristallochimie des phengites dans les quartzites micaces metamorphiques du Permo-Trias des Alpes piemontaises. Sci. Geol. Btdl., 25(4): 20 1-229. Dunoyer de Segonzac, G . and Millot, G., 1962. Pyrophyllite de diagenese dans le Devonien inferieur du synclinal de Lava1 (massif armoricain). C.R. Acad. Sci., 255: 3438-3440. Dunoyer de Segonzac, G., Artru, P. and Ferrero, J., 1966. Sur une transformation des mineraux argileux dans les “terres noires” du bassin de la Durance: influence de l’orogenie alpine. C.R. Acad. Sci., Sir. D, 262: 2401-2404. Dunoyer de Segonzac, G., Ferrero, J. and Kubler, B., 1968. Sur la cristallinite de I’illite dans la diagenese et I’anchimetamorphisme. Sedimentology, 10: 137- 143. Durney, D.. 1974. Relations entre les temperature d’homogeneisation d’inclusions fluides et les minkraux metamorphiques dans les nappes helvetiques du Valais. Bull. Soc. GPol. Fr., (7). 16(3): 269-272. Eckhardt, F.-J., 1958. Uber Chlorite in Sedimenten. Geol. Jahrb., 75: 437;474. Eckhardt. F.-J.. 1965. Uber den Einfluss der Temperatur auf den kristallographischen Ordnungsgrad von Kaolinit. Proc. Ist. Int. Clay Con&, 1963, Stockholm, 2: 137-145. Eckhardt, F.-J. and Von Gaertner, H.R., 1962. Zur Entstehung und Umbildung der KaolinKohlentonsteine. Fortschr. Geol. Rheinl. WesrJ, 3(2): 623-640. Ernst, W.G., 1963. Significance of phengitic micas from low-grade schists. Am. Mineral., 48: 1357-1373. Ernst, W.G., 1973. Interpretative synthesis of metamorphism in the Alps. Geol. Soc. Am. Bull., 84: 2053-2078. Ernst, W.G. and Seki, Y . , 1967. Petrologic comparison of the Franciscan and Sanbagawa metamorphic terranes. Tectonophysics, 4(4-6): 463-478. Ernst, W.G., Seki, Y . , Onuki, H. and Gilbert, M.C., 1970. Comparative study of low-grade metamorphism in the California Coast Ranges and the outer metamorphic belt of Japan. Geol. Soc. Am. Mem. 124: 276 pp. Esch, H., 1966. Vergleichende Diagenese-Studien an Sandsteinen und Schiefertonen des Oberkarbons in Nordwestdeutschland und den East Midlands in England. Fortschr. Geol. Rheinl. WesrJ, 13(2): 1013-1084. Eskola, P.. 1915. Om sambandet mellan kemisk och mineralogisk sammansattning hos Orijarvitraktens metamorfa bergarter (with Eng. summary: On the relations between the chemical and mineralogical composition in the metamorphic rocks of the Orijarvo region). Bull. Comm. GPO[.Finl.,No. 44: 145 pp. Eskola, P., 1920. The mineral facies of rocks. Norsk. Geol. Tidsskr., 6: 143- 194. Eslinger, E.V. and Savin, S.M., 1973a. Mineralogy and oxygen isotope geochemistry of the hydrothermally altered rocks of the Ohaki-Broadlands, New Zealand geothermal area. Am. J . Sci.,273: 240-267. Eslinger, E.V. and Savin. S.M., 1973b. Oxygen isotope geothermometry of the burial metamorphic rocks of the Precambrian Belt Supergroup, Glacier National Park, Montana. Geol. Soc. Am. Bull., 84: 2549-2560. Esquevin, J.. 1969. Influence de la composition des illites sur leur cristallinite. Bull. Centre Rech. Pau-SNPA, 3(1): 147-153. Fabre. J., 1961. Contribution a I’etude de la Zone Houillere en Maurienne et en Tarentaise (Alpes de Savoie). MPm. Bur. Rech. GPol. Min., 2: 315 pp.

479 Ferrero, J. and Kubler, B., 1964. Presence de dickite et de kaolinite dans les gres cambriens d'Hassi Messaoud. Bull. Serv. Carte Ghol. Alsace Lorraine, 17: 247-26 1. Feys, R., 1957. Etude geologique du Carbonifere briaqonnais (Hautes-Alpes). Mem. Bur. Rech. Geol. Min., 6: 387 pp. Flehmig, W., 1973. Kristallinitat und Infrarotspektroskopie naturlicher dioktaedrischer Illite. Neues Jahrb. Mineral. Monatsh., 1973 (7-8): 35 1-36 I . Frank-Kamenetsky, V.A., Logvinenko, N.V. and Drits. V.A.. 1965. Tosudite-a new mineral forming the mixed-layer phase in alushtite. Proc. 1st Int. Cla,v Con/.. 1963, Stockholm. 2: 181-186. Foscolos, A.E. and Kodama, H., 1974. Diagenesis of clay minerals from Lower Cretaceous shales from northeastern British Columbia. Clays C/ay Miner., 22: 319-335. Foscolos, A.E. and Stott, D.F., 1975. Degree of diagenesis, stratigraphic correlations and potential sediment sources of Lower Cretaceous shale of northeastern British Columbia. Geol. Sum. Can., Bull., 250: 46 pp. Foscolos. A.E., Powell, T.G. and Gunther, P.R.. 1976. The use of clay minerals and organic geochemical indicators for evaluating the degree of diagenesis and oil generating potential of shales. Geochim. Cosmochim. Acta, 40: 953-966. Frey, M., 1969a. Die Metamorphose des Keupers vom Tafeljura bis zum Lukmaniergebiet. Beitr. Geol. Karte Schweiz., 137: 161 pp. Frey, M., 1969b. A mixed-layer paragonite-phengite of low-grade metamorphic origin. Contrib. Mineral. Petrol., 24: 63-65. Frey, M., 1970. The step from diagenesis to metamorphism in pelitic rocks during Alpine orogenesis. Sedimentology, 15: 261 -279. Frey, M., 1974. Alpine metamorphism of pelitic and marly rocks of the Central Alps. Schweiz. Mineral. Petrogr. Mitt., 54(2/3): 489-506. Frey, M. and Niggli, E., 1971. Illit-Kristallinitat, Mineralfazien und Inkohlungsgrad. Schwei:. Mineral. Petrogr. Mitt., 5 I( I): 229-234. Frey, M., Hunziker. J.C., Roggwiller, P. and Schindler. C., 1973. Progressive niedriggradige Metamorphose glaukonitfuhrender Horizonte in den helvetischen Alpen der Ostschweiz. Contrib. Mineral. Petrol., 39: 185-218. Fuchtbauer, H. and Goldschmidt, H., 1959. Die Tonminerale der Zechsteinformation. Beitr. Mineral. Petrol., 6: 320-345. Fuchtbauer, H. and Goldschmidt, H., 1963. Beobachtungen zur Tonmineral-Diagenese. Proc. 1st Int. Clay Con/., 1963, Stockholm, 1 : pp. 99- 1 1 1. Furbish, W.J., 1975. Corrensite of deuteric origin. Am. Mineral., 60: 928-930. Fyfe, W.S., 1974. Low-grade metamorphism: some thoughts on the present situation. Cut?. Mineral., 12: 439-444. Fyfe, W.S. and Turner, F.J., 1966. Reappraisal of the metamorphic facies concept. Contrih. Mineral. Petrol., 12: 354-364. Fyfe, W.S.. Turner, F.J. and Verhoogen, J.. 1958. Metamorphic reactions and metamorphic facies. Mem. Geol. SOC.Am., 73: 259 pp. Ganguly, J., 1968. Analysis of the stabilities of chloritoid and staurolite. and some equilibria Am. J . Sci.. 266(4): 277-298. in the system FeO-AI,O,-SO,-H,O-0,. Ganguly, J., 1969. Chloritoid stability and related parageneses: theory, experiments, and applications. Am. J. Sci., 267(8): 910-944. Carrels, R.M. and Christ, C.L., 1965. Solutions, Minerals and Equilibriu. Harper and Row, New York, N.Y., 435 pp. Gavish, E. and Reynolds, R.C., 1970. Structural changes and isomorphic substitution in illites from limestones of variable degrees of metamorphism. Isr. J . Chem., 8: 477-485.

480 Gavrilov. A.A. and Alexandrova. V.A.. 1968. Postsedimentation argillitization in Paleozoic clastic deposits of the southern Urals and northern Mugodzhars. Dokl. Acud. Sci. U.S.S.R., Eurth G i . Sect. (transl. from Dokl. Akad. Nuuk S . S . S . R . ) . 182(5): 178-180. Gill. E.D.. 1956. Fossil wood replaced by laumontite near Cape Paterson. Victoria. Proc. R. SOC. Vict.. 69: 33-35. Glassley. W.. 1974. A model for phase equilibria in the prehnite-pumpellyite facies. Confrib. Minerd. Petrol., 43: 317-322. Goodwin. J.H. and Surdam. R.L.. 1967. Zeolitization of tuffaceous rocks of the Green River Formation. Wyoming. Science, 157(3786): 307-308. Gradusov. B.P., I97 1. Dioctahedral chlorites. Lithol. Miner. Resour. (transl. from Lrtol. Pole-n. Iskop.), 1971(4): 471-478. Guitard. G . and Saliot, P., 1971. S u r les paragenese a lawsonite et a pumpellyite des Alpes de Savoie. Bull. Soc. Fr. Mineral. Crvstullogr., 94: 507-523. Gurewitsch, A.B. and Toporez. G.A.. 1968. Uber die Epigenese der Sandsteine der Tunguska-Serie im Norden des Tunguska-Reviers (in Russian). Abstr. in Zeittrulbl. Geul. Paluonrol.. l(3): 42 I , Haas. H. and Holdaway. M.J., 1973. Equilibria in the system AI,O-SO,-H,O involving the stability limits of pyrophyllite, and thermodynamic data of pyrophyllite. Am. J . Sci., 273: 449-464. Hallam, A. and Snellwood, B.W., 1970. Montmorillonite and zeolites in Mesozoic and Tertiary beds of southern England. Mineral. Mug., 37(292): 950-952. Hamilton, J.D., 1968. Triphormic clay minerals from the Lower Hunter Permian succession of New South Wales. Geol. Soc. Ausr., 15(1): 9-24. Harrassowitz, H.. 1927. Anchimetamorphose, das Gebiet zwischen Oberflachen- und Tiefenurnwandlung der Erdrinde. Ber. Oberhess. Ges. Nut. Heilk. Giessen, Narurwiss. Aht., 12: 9-15. Hashimoto. M., 1966. On the prehnite-pumpellyite metagraywacke facies (in Japanese: English abstr.). J . Geol. Soc. Jpn., 72: 253-265. Hashimoto, M..1968. Glaucophanitic metamorphism of the Katsuyama district, Okayama Prefecture, Japan. J . Fuc. Scr. Uniu. Tokyo, Sect. 11, 17: 99-162. Hashimoto. M., 1972. Reactions producing actinolite in basic metamorphic rocks. Lirhos, 5: 19-31. Hashimoto. M. and Kashima, N., 1970. Metamorphism of Paleozoic greenstones in the Chichibu belt of western Shikoku (in Japanese; English abstr.). J . Geol. Soc. Jpn., 76: 199-204. Hawkins, J. W.. 1967. Prehnite-pumpellyite metagraywacke facies metamorphism of a graywacke-shale series. Mount Olympus, Washington. Am. J . Sci., 265(9): 798-8 18. Hay. R.L.. 1966. Zeolites and zeolitic reactions in sedimentary rocks. Geol. Soc. Am., Spec. Pup., 8 5 : 130 pp. Hay. R.L. and Iijima. A.. 1958. Petrology of palagonite tuffs of Koko Craters. Oahu. Hawaii. Conrrih. Mineral. Perrol., 17: 141- 154. Hayes. J.B.. 1967. Dickite im Lansing Group (Pennsylvanian) limestones, Wilson and Montgomery counties. Kansas. Am. M i n e r d , 52: 890-896 Hayes. J.B.. 1970. Polytypism of chlorite in sedimentary rocks. C1u.v.~Clay Miner., 18(5): 285 - 308. Heniley. J.J. and Jones, W.R., 1964. Chemical aspects of hydrothermal alteration with emphasis on hydrogen metasomatism. Econ. Geol., 59: 538-539. Henderson. G.V.. 1970. The origin of pyrophyllite and rectorite in shales of north central Utah. Clu,vs C1u.v Miner., 18: 239-246.

High, L.R. and Picard. M.D., 1965. Sedimentary petrology and origin of analcime rich Popo Agie Member, Chugwater (Triassic) Formation, west-central Wyoming. J . Sedimenr. Petrol., 35( I ) : 49-70. Hiltabrand, R.R.. Farrell. B.E. and Billings. G.K.. 1973. Experimental diagenesis of Gulf Coast argillaceous sediment. A m . Assoc. Pet. Geol. Bull.. 57(2): 338-348. Hinrichsen, Th. and Schiirmann. K., 1969. Untersuchungen zur Stabilitat von Punipellyit. Neues Jahrb. Mineral. Monatsh., 1969( 10): 441-445. Hinrichsen, Th. and Schiirmann. K.. 1972. Mineral reactions in burial metamorphism. N e w s Jahrb. Mineral. Monatsh., 1972(I ) : 35-48. Hoare, J.M., Condon, W.H. and Patton. W.W.. 1964. Occurrence and origin of laumontite in Cretaceous sedimentary rocks in western Alaska. U.S. Geol. Sura., Prof: Pap.. 501-C: 74-78. Hoschek, G., 1969. The stability of staurolite and chloritoid and their significance in metamorphism of pelitic rocks. Contrih. Mineral. Petrol.. 22(3): 208-232. Hosterman, J.W., Wood, G.H., Jr. and Bergin. M.G., 1970. Mineralogy of underclays in the Pennsylvania anthracite region. U.S. Geol. S u m , Prof: Pap., 700-C: 89-97. Hower, J. and Mowatt, T.C., 1966. The mineralogy of illites and mixed-layer illite/montmorillonites. A m . Mineral.. 5 1 : 825-854. Hower, J.. Eslinger, E.V.. Hower. M.E..and Perry, E.A.. 1976. Mechanism of burial metaniorphism of argillaceous sediment: 1. Mineralogical and chemical 'evidence. Geol. Soc,. Ani. Bull., 87(5): 725-737. Hoyer, P., Clausen, C.-D., Leuteritz, K., Teichmiiller. R. and Thome. K.N.. 1974. Ein Inkohlungsprofil zwischen dem Gelsenkirchener Sattel des Ruhrkohlenbeckens und dem Ostsauerlander Hauptsattel des Rheinischen Schiefergebirges. Forrschr. Geol. Rheinl. Wt>s/f:. 24: 161-172. Hutton, C.O.. 1937. An occurrence of the mineral pumpellyite in the Lake Wakatipu region. western Otago. New Zealand. Mineral. Mag., 24: 529-533. Hutton, C.O., 1940. Metamorphism in the Lake Wakatipu region. western Otago. New Zealand. N . Z . Dep. Sci. Ind. Res. Geol. Mem., 5 : 90 pp. Hutton, C.O. and Turner. F.J., 1936. Metamorphic z o n s in northwest Otago. Trans. R. Soc. N . Z . , 65(4): 405-406. Iijima, A,. 1961. Diagenetic alteration of some acidic tuffs in the Kushiro Coal Basin (clinoptilolite-montmorillonite-low-cristobaliterelation). Jpn. J . Geol. Geogr., 32(3-4): 507-522. lijima, A,, 1974. Clay and zeolitic alteration zones surrounding Kuroko deposits in Hokuroku district, northern Akita, as submarine-diagenetic alteration products. In: S. Ishihara et al. (Editors), Geology of Kuroko Deposits. Min. Geol. (SOC.Min. Geol. Jap.. Tokyo). Spec. Iss. No. 6: 267-289. Iijima, A. and Hay. R.L., 1968. Analcime composition in tuffs of the Green River Formation of Wyoming. A m . Mineral., 53: 184-200. Iijima, A. and Utada. M.. 1966. Zeolites in sedimentary rocks, with reference to the depositional environments and zonal distribution. Sedimentologv, 7: 327-357. Iijima, A. and Utada. M., 1971. Present-day zeolitic diagenesis of the Neogene geosynclinal deposits in the Niigata oil field, Japan. Aduan. Chem. Ser. 101: 342-349. lijima, A. and Utada, M., 1972. A critical review on the occurrence of zeolites in sedimentary rocks in Japan. Jpn. J . Geol. Geogr., 42( 1-4): 61-84. Jaron, M.G., 1967. A method for determining post-depositional alteration in underclays and associated lithologies. Penn. Geol. Suru. Inf: Circ. (4th Ser.) 57: 14 pp. Johns, W.D. and Shimoyama. A,. 1972. Clay minerals and petroleum-forming reactions during burial and diagenesis. Am. Assoc. Pet. Geol. Bull., 56(11): 2160-2167.

482 Jolly, W.T., 1970. Zeolite and prehnite-purnpellyite facies in south central Puerto Rico. Contrib. Mineral. Petrol., 27: 204-224, Kaplan, M.Ye., 1965. Diagenesis and epigenesis of Triassic sediments of the South Maritime region. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.), 163(4): 166-168. Kaplan, M.Ye., 1974. Catagenic zoning of the Mesozoic clastic complex of northeastern Siberia. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.), 2 18(2): 171 - 173. Karpova, G.V., 1966. On paragonite hydromicas in terrigenous rocks of the Great Donets Basin (in Russian). Dokl. Akad. Nauk S.S.S.R., 171(2): 443-445. Karpova, G.V., 1967. Muscovitic hydromicas in coal-bearing polyfacies deposits. Lithol. Miner. Resourc. (transl. from Litol. Polezn. Iskop.), 1967(6): 688-697. Karpova, G.V., 1969. Clay mineral post-sedimentary ranks in terrigenous rocks. Sedimentol00,13: 5-20. Karpova, G.V. and Shevyakova, E.P., 1965. Epigenetic features specific to the transitional region of the Greater Donets downwarp. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.), 160(4): 40-43. Karpova, G.V. and Timofeeva, Z.V., 1971. Post-sedimentational alteration of Aalenian rocks in the structural-facies complexes of the northern Caucasus. Lithol. Miner. Resour. (transl. from Litol. Polezn. Zskop.), 5 : 600-608. Karpova, G.V., Lukin, A.E. and Shevyakova, E.P., 1969. Catagenesis of Carboniferous deposits of the Dniepr-Donets basin. Lithol. Miner. Resour. (transl. from Litol. Polezn. Zskop.), 1969(1): 15-27. Kasanskiy, J.P., 1967. Besonderheiten sekundarer Umwandlungen in klastischen ober prakambrischen Karbonat-Gesteinen im Siidosten des Jenissey-Gebirges (Ost-Sibirien). In: Postsediment. Preobraz. Osadoch. Porod. Sibiri. Moscow, pp. 180-205 (abstr. in Zentralbl. Mineral., 1969(1/2): 181). Kawachi, Y., 1974. Geology and petrochemistry of weakly metamorphosed rocks in the upper Wakatipu district, southern New Zealand. N.Z. J . Geol. Geophys., 17(1): 169-208. Kawachi, Y., 1975. Purnpellyite-actinolite and contiguous facies metamorphism in part of upper Wakatipu district, South Island, New Zealand. N.Z. J. Geol. Geophys., 18(3): 40 1-44 1. Keller, W.D., 1963. Diagenesis in clay minerals-a review. Proc. 11th Natl. Conf: Clays Clay Miner. (Ottawa, 1962). Pergamon Press, pp. 136-157. Keller. W.D., 1970. Environmental aspects of clay minerals. J. Sediment. Petrol., 40(3): 788-813. Kerrick, D.M., 1974. Review of metamorphic mixed-volatile (H,O-CO,) equilibria. Am. Mineral., 59(7-8): 729-762. Kessler. P., 1922. Uber Lochverwitterung und ihre Beziehungen zur Metharmose (Urnbildung) der Gesteine. Geol. Rundsch., 12: 237-270. Kienast, J.R. and Velde, B., 1970. Le metamorphisme dans les Alpes franco-italiennes: mise en evidence d’un gradient de temperature et de pression. C.R. Acad. Sci. Paris, Ser. D., 27 1 : 637-640. Khitarov, N.I. and Pugin, V.A., 1966. Behaviour of montmorillonite under elevated temperatures and pressures. Geochem. Znt., 3: 621-626 (transl. from Geokhimiya, 7: 790-795). Kisch, H.J., 1966a. Chlorite-illite tonstein in high-rank coal from Queensland, Australia: notes on regional epigenetic grade and coal rank. Am. J . Sci., 264(5): 386-397. Kisch, H.J., 1966b. Zeolite facies and regional rank of bituminous coals. Geol. Mag., 102(5): 414-422.

483 Kisch, H.J., 1969. Coal rank and burial-metamorphic mineral facies. In: P.A. Schenk and I Havenaar (Editors), Advances in Organic Geochemistry, 1968. Pergamon Press, Oxford. pp. 407-424. Kisch, H.J., 1974. Anthracite and meta-anthracite coal ranks associated with “anchimetamorphism” and “very-low-stage” metamorphism. I, 11, 111. Proc. K. Ned. Akad. Wet., Amsterdam, Ser. B, 77(2): 81-1 18. Kisch, H.J., 1978. Incipient metamorphism of Cambro-Silurian clastic rocks in the Caledonides of central Jamtland, western Sweden: illite crystallinity and vitrinite reflectance. Isr. Geol. Soc. Annu. Meet. Zikhron Ya’akov, March 1978, and I.G.C. P. Project Caledonide Orogen Conj “The Caledonides of the British Isles Reviewed, and Deformation and Metamorphism in the Caledonide Orogen’: Dublin, August 1978, Abstracts, pp. 29-30 (abstracts of a research report of June 1978, 49 pp.). Kisch, H.J., 1980a. Incipient metamorphism of Cambro-Silurian clastic rocks from the Jamtland Supergroup, central Scandinavian Caledonides. western Sweden: illite crystallinity and “vitrinite” reflectance. In: W.E.A. Phillips and M.R.W. Johnson (Editors). Deformation and Metamorphism in the Caledonide Orogen. J . Geol. SOC.London (in press). Kisch, H.J.. 1980b. Illite crystallinity and coal rank associated with lowest-grade metamorphism of the Taveyanne Greywacke in the Helvetic zone of the Swiss Alps. Eclogue Geol. Helv. (in press). Kodama, H.. 1966. The nature of the component layers of rectorite. Am. Mineral., 51: 1035-1055. Konysheva, R.A., 1968. Gumbelite, an indicator of a high degree of sedimentary-rock alteration. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.), 181(3): 119-121. Kopeliovich, A.V., Kossovskaya, A.G. and Shutov. V.D.. 1961. On some aspects of the epigenesis of the terrigenous deposits of platform and geosynclinal regions. 1:o. Akud. Nauk. S.S.S.R., Geol. Ser., 1961(6) (in transl., 1962: 13-23). Koporulin. V.I.. 1961. The origin of zeolite cement in the sandstones and gravelstones of the coal-bearing stratum of the southeastern part of the Irkutsk basin. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.). 137(1 ): 467-470. Koporulin, V.I., 1962. Types of secondary alterations in sands and gravels of Irkutsk basin coal measures and their possible relationship with underground water. Irv. Akad. Nuuk S.S.S.R., Ser. Geol., 1962(3): 72-87. (Transl. in Int. Geol. Rev., 6(3): 531-540). Koporulin, V.I.. 1966. Composition, facies. and formation conditions of coal strata in the central part of the Irkutsk basin (in Russian). Tr. Geol. Inst. Akad. Nauk S.S.S.R., 160: 2- 164. (Abstr. in Chem. Abstr., 66(7): 57885j). Koporulin, V.I.. 1967. Comparative description of postdepositional changes in terrigenous rocks of the Irkutsk basin. Lithol. Miner. Resour. (Transl. from Litol. Po1e:n. hkop.), 1967(6): 679-687. Koporulin, V.I.. 1968. Epigenetic formation of crustification chlorite in sand-granule rocks. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.). 1968(5): 606-6 1 1. Koporulin. V.I.. 1972. Catagenetic changes in sandy-pebbly Upper Cretaceous rocks, Penzhinsk Bay area. Kamchatka. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop. ), 7(2): 23 1-238. Kossovskaya. A.G., 1960. Specific nature of epigenetic alteration of terrigenous rocks in platform and geosynclinal regions. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.), 130(1): 123-125. Kossovskaya, A.G., 1969. Specific features of the alteration of clay minerals under different facies-climatic conditions. Proc. Int. Clay Conk. 1969 Tokyo, I . Isr. Univ. Press, Jerusalem, pp. 339-347.

484 Kossovskaya. A.G. and Drits. V.A., 1970. The variability of micaceous minerals in sedimentary rocks. Sedinientologv, 15: 81-101. Kossovskaya. A.G. and Shutov. V.D., 1958. Zonality in the structure of terrigene deposits in platform and geosynclinal regions. Eclogue Geol. Helo., 5 l(3): 656-666. Kossovskaya. A.G. and Shutov. V.D.. I96 I . The correlation of zones of regional epigenesis and metagenesis in terrigenous and volcanic rocks. Dokl. Acud. Sci. U . S . S . R . , Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S . S . S . R . ) , 139(3): 732-736. Kossovskaya. A.G. and Shutov, V.D., 1963. Facies of regional epigenesis and metagenesis. I x . Akud. Nauk S.S.S.R., Ser. Ceol., 1963(7): 3-18 (transl. in I n t . Geol. Rev.. 7(7): 1157-1 167). Kossovskaya. A.G. and Shutov. V.D.. 1970. Main aspects of the epigenesis problem. Sedimen to log^, 15: 11-40, Kossovskaya. A.G.. Logvinenko. N.V. and Shutov, V.D., 1957. Stages of formation and alteration in terrigenous rocks (in Russian). Dokl. Akad. Nauk S.S.S.R., 116(2): 293-296. Kossovskaya, A.G.. Shutov, V.D. and Drits, V.A., 1963. Clay minerals as indicators of deep-seated alteration of terrigenous rocks (in Russian). In: Geokhymiya, Petrografya i Mineralogi.ya Osadochnykh Obrarovaniy. Akad. Nauk. S.S.S.R., Moscow, pp. 120- 130. Kossovskaya. A.G., Shutov, V.D. and Alexandrova, V.A.. 1964. Dependence of the mineral composition of clays in the coal-bearing formations on the sedimentation conditions. 5me Congr. Int. Strutigr. Geol. Carbong, Paris, 1963, C.R., 2: 519-529. Kossovskaya, A.G., Drits, V.A. and Aleksandrova, V.A.. 1965. On trioctahedral micas in sedimentary rocks. Proc. 1st I n t . Clay Conf., 1963, Stockholm, 2: 147-169. Kramm. U.. 1973. Chloritoid stability in manganese-rich low-grade metamorphic rocks, Venn-Stavelot Massif, Ardennes. Contrib. Mineral. Petrol., 41: 179- 196. Krumbein, W.C. and Sloss. L.L., 1963. Strutigruphy and Sedimentation. 2nd ed. Freeman, San Francisco. Calif., 660 pp. Kubler, B., 1964. Les argiles, indicateurs de metamorphisme. Rev. Inst. Fr. Petrol., 19: 1093-1 112. Kubler. B., 1967a. La cristallinite de l’illite et les zones tout a fait superieures du metamorphisme. In: Etuges Tectoniques. A la Baconniere, Neuchltel, Suisse, pp. 105- 121. Kubler. B.. 1967b. Anchimetamorphisme et schistosite. Bull. Centre Rech. Pau-SNPA, l(2): 259-278. Kubler. B., 1968. Evaluation quantitative du metamorphisme par la cristallinite de l’illite. Bull. Centre Rech. Pau-SNPA, 2(2): 385-397. Kubler. B., 1970. Crystallinity of illite. Detection of metamorphism in some frontal parts of the Alps (abstr.). Fortschr. Mineral., 47 (Beih. I ) : 39-40. Kubler, B., 1973. La corrensite, indicateur possible de milieux de sedimentation et du degree de transformation d h n sediment. Bull. Centre Rech. Pau-SNPA, 7(2): 543-556. Kulbicki. G. and Millot, G., 1963. Diagenesis of clays in sedimentary and petroliferous series. Cla.vs Clav Miner., Proc. 10th Natl. Conf., pp. 329-330. Kubler. B., Martini, J. and Vuagnat, M.. 1974. Very low-grade metamorphism in the Western Alps. Schweir. Mineral. Petrogr. Mitt., 54(2/3): 461 -469. Kulke, H.. 1969. Petrographie und Diagenese des Stubensandsteins (mittlerer Keuper) aus Tiefbohrungen im Raum Memmingen (Bayern). Contrib. Mineral. Petrol., 20: 135- 163. Landis. C.A., 1974. Stratigraphy, lithology, structure, and metamorphism of Permian, Triassic. and Tertiary rocks between the Mararoa River and Mount Snowdon, western Southland. New Zealand. J . R . SOC.N . Z . , 4(3): 229-251. Landis. C.A. and Bishop, D.G.. 1972. Plate tectonics and regional stratigraphic metamorphic relations in the southern part of the New Zealand geosyncline. Geol. Soc. Am. Bull., 83: 2267-2284.

485 Landis, C.A. and Coombs, D.S., 1967. Metamorphic belts and orogenesis in southern New Zealand. Tectonophysics, 4(4-6): 50 1-5 18. Landis, C.A. and Rogers. J., 1968. Some experimental data on the stability of pumpellyite. Am. Mineral., 53(5-6): 1038- 1041. Le Corre, C., 1969. Sur une paragenkse a chloritoyde dans les schistes de I’Ordovicien moyen des synclinaux du Sud de Rennes (Massif Armoricain). Bull. Soc. Geol. Mineral. Brr/agne (C). 1: 33-44. Le Corre, C., 1975. Analyse comparee de la cristallinite des micas dans le Brioverien et le Paleozoi’que centre-armoricains: zoneographie et structure d’un domaine epizonal. Bull. SOC.Geol. Fr. (7). 17(4): 547-553. Levi, B., 1970. Burial metamorphic episodes in the Andean geosyncline. Central Chile. Geol. Rundsch., 59(3): 994-1013. Liborio, G. and Mottana, A., 1975. White micas with 3T polymorph from the Calcescisti of the Alps. Neues Jahrb. Mineral. Monatsh., l975( 12): 546-555. Liou, J.G., 1970. Synthesis and stability relations of wairakite. CaAI,Si,O,,-2 H 2 0 . Contrih. Mineral. Petrol., 27: 259-282. Liou, J.G., 1971a. P-T stabilities of laumontite. wairakite. lawsonite, and related minerals in J . Petrol., 12(2): 379-41 1. the system CaA1,Si,O8-SiO,-H,O. Liou, J.G., I971 b. Analcime equilibria. Lithos, 4: 389-402. Lippmann, F., 1954. Uber einen Keuperton von Zaisersweiher bei Heilbronn. Heidelh. Beitr. Mineral. Petrol.. 4: 130- 134. Lisitsin, A.K., Kondrat’eva, LA. and Komarova. G.V., 1969. Methods of genetic interpretation of epigenetic alterations of sedimentary rocks. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 1969(3): 253-264. Logvinenko, N.V., 1964. Mixed-layer phase in Silurian phyllitized shale of Nura-Tau. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.). 157(4). 1 99-101. Logvinenko, N.V. and Karpova, G.V., 1968. Stages of bostdiagenetic alterations in rocks of coal-bearing formations. Int. Geol. Congr., Rep. 23rd Sess., Czechoslovakia 1968, Ahsrr., pp. 238-239. Long, G . and Neglia, S., 1968. Composition de I’eau interstitielle des argiles et diagenkse des mineraux argileux. Rev. Ins/. Fr. Pet., 23: 53-69. Lounsbury, R.W. and Melhorn, W.N., 1964. Clay mineralogy of Paleozoic K-bentonites of the eastern United States (Part 1). Clays Clay Miner., Proc. 12th Natl. Cant. pp. 557-566. Lucas, J., 1962. La transformation des mineraux argileux dans la sedimentation. Etudes sur les argiles du Trias. M6m. Seru. Carte Geol. Alsace Lorraine, 23: 202 pp. (English transl. publ. by Israel Program for Scientific Translations, Jerusalem, 203 pp.) Lucas, J. and Ataman. G.. 1968. Mineralogical and geochemical study of clay mineral transformations in the sedimentary Triassic Jura basin (France). C1a.p Cia-v Miner., 16(5): 365-372. Ludwig, V.. 1972a. Die Paragenese Muscovit, Pyrophyllit, 7 A-Chlorit und Kaolinit im Silur des Frankenwaldes (NE-Bayern). Neues Jahrb. Geol. Palaontol., Monarsh.. 1972(5): 303305. Ludwig, V., 1972b. Die Paragenese Chlorit, Muscovit. Paragonit und Margarit im “Griffelschiefer” des Ordoviziums in NE-Bayern (mit einem Beitrag zum Problem der IllitKristallinitat). Neues Jahrb. Geol. Palaonrol., Monatsh., 1972(9): 546-560. Ludwig, V., 1973. Zum Ubergang eines Tonschiefers in die Metamorphose: “Griffelschiefer” im Ordovizium von NE-Bayern. Neues Jahrh. Geol. Paliiontol., Abh., l44( I ) : 50- 103. Mahon. W.A.J. and Finlayson, J.B., 1972. The chemistry of the Broadlands geothermal area, New Zealand. Am. J . Sci., 272( I): 48-68.

Maiklem. W.R. and Campbell, F.A., 1968. A study of the clays from Upper Cretaceous bentonites and shales in Alberta. Can. Mineral., 8: 354-371. Martini. J.. 1972. La metamorphisme dans les chaines alpines externes et ses implications dans I’orogenese. Bull. Suisse Mineral. Petrogr., 52(2): 257-275. Martini. J. and Vuagnat. M., 1965. Presence du facies a zeolite dans la formation des “gres” de Taveyanne (Alpes franco-suisses). Bull. Suisse Mineral. Petrogr.. 45( I ) : 28 1-293. Martini. J. and Vuagnat, M.. 1970. Metamorphose niedrigst temperierten Grades in den Westalpen. Fortschr. Mineral., 47( 1 ): 52-64. Matsuda. T. and Kuriyagawa, S., 1965. Lower-grade metamorphism in the eastern Akaishi Mountains. central Japan (in Japanese; English abstr.). Bull. Earthquake Res. Inst. (Jpn.), 43: 209-235. Maxwell, D.T. and Hower, J., 1967. High-grade diagenesis and low-grade metamorphism of illite in the Precambrian Belt series. A m . Mineral., 52(5/6): 843-857. McEwan, D.M.C.. 1956. A study of an interstratified illite-montmorillonite clay from Worcestershire, England. Clays Clay Miner., Proc. 4th Natl. Conk, pp. 166- 172. McNamara. M.. 1965. The lower greenschist facies in the Scottish Highlands. Geol. Foren. Srockholni Forb.- 87: 347-389. Meisl. S., 1970. Petrologische Studien im Grenzbereich Diagenese-Metamorphose. Abh. Hess. Landesanzt Bodenforsch., 57: 93 pp. Mellon. G.B.. 1967. Stratigraphy and petrology of the Lower Cretaceous Blairmore and Mannville Groups, Alberta foothills and plains. Res. Counc. Alberta, Bull. 32: 270 pp. Mentser. M., O’Donnell. H.J. and Ergun, S., 1962. X-ray scattering intensities of anthracites and meta-anthracites. Fuel, 41: 153-161. Mentser, M.. O’Donnell. H.J. and Ergun. S., 1963. Development of three-dimensional crystallinity in natural graphitic materials. In: Proc. 5th Conk Carbon. Pergamon, London, pp. 493-497. Millot, G., 1964. Geologie des Argiles. Masson, Paris, 499 pp. Miyashiro. A,, 1961. Evolution of metamorphic belts. J. Petrol., 2(3): 277-31 1. Miyashiro. A. and Shido, F.. 1970. Progressive metamorphism in zeolite assemblages. Lithos, 3: 251-260. Mizutani. S.. 1967. Kinetic aspects of diagenesis of silica in sediments. J . Earth Sci., Nagoya Univ., 15(2): 99-1 11. Mizutani. S., 1970. Silica minerals in the early stage of diagenesis. Sedirnentology, 15(3/4): 419-436. Moiola, R.J., 1970. Authigenic zeolites and K-feldspar in the Esmeralda Formation, Nevada. A m . Mineral., 55(9-10): 1681-1691. Mossler, J.H. and Hayes, J.B., 1966. Ordovician potassium bentonites of Iowa. J . Sediment. Petrol.. 36: 414-427. Muffler. L.J.P. and White, D.E.. 1969. Active metamorphism of Upper Cenozoic sediments in the Salton Sea geothermal field and the Salton trough, southeastern California. Geol. SOC. A m . Bull., 80: 157-182. Muller. G.. 1963. Zur Kenntnis di-oktaedrischer Vierschicht-Phyllosilikate (Sudoit-Reihe der Sudoit-Chlorit-Gruppe). Int. Clay Conk, 1963, Stockholm, Proc., I . Pergamon Press, Oxford. pp. 121- 130. Muller. G.. 1967a. Sudoit (“dioktaedrischer Chlorit”. “Al-Chlorit”) im Cornberger Sandstein von Cornberg/Hessen. Contrih. Mineral. Petrol., 14: 176- 189. Muller. G.. 1967b. Diagenesis in argillaceous sediments. In: G. Larsen and G.V. Chilingar (Editors). Diagenesis in Sediments. Elsevier, Amsterdam, pp. 127- 177. Murata. K.J. and Whiteley, K.R., 1973. Zeolites in the Miocene Briones Sandstone and related formations of the central Coast Ranges, California. J. Res. U.S. Geol. Suru., l(3): 255-265.

487 Muravjew, W.I. and Salyn. A.L., 1969. Epigenetische Umwandlungen der Schichtsilikate in einem Perm-Triasprofil in Mangyschlak (Kasachstan). Proc. In?. Clay Conb, IY69, Tokw. 1. Isr. Univ. Press, Jerusalem, pp. 325-333. Nakajima, W. and Koizumi. M.. 1966. The chemical composition of analcite from the low-grade metamorphic rocks in Japan. J . Geol. SOC.Jpn.. 72(1 I ) : 517-521. Nakajima, W. and Tanaka, K., 1967. Zeolite-bearing tuffs from the Izumi mountain range. southwest Japan, with reference to mordenite-bearing tuffs and laumontite tuffs (in Japanese; English transl.). J . Geol. SOC.Jpn., 73(5): 237-245. Nitsch, K.H., 1968. Die Stabilitat von Lawsonit. Naturwissenschaften. 55(8): 388. Nitsch. K.H., 1971. Stabilitatsbeziehungen von Prehnit- und Pumpellyit-haltigen Paragenesen. Contrib. Mineral. Petrol., 30: 240-260. Otalora, G., 1964. Zeolites and related minerals in Cretaceous rocks of east-central Puerto Rico. A m . J . Sci.,262: 726-734. Packham, G.H. and Crook, K.A.W., 1960. The principle of diagenetic facies and some of its implications. J . Geol., 68: 392-407. Perry, E.A., 1974. diagenesis and the K-Ar dating of shales and clay minerals. Geol. So(,.Am. Bull., 85: 827-830. Perry, E. and Hower, J., 1970. Burial diagenesis in Gulf Coast pelitic sediments. Clays C/o,. Miner., 18: 165-177. Perry, E.A. and Hower. J., 1972. Late-stage dehydration in deeply’ buried pelitic sediments. A m . Assoc. Pet. Geol. Bull., 56(10): 2013-2021. Pettijohn, F.J., 1957. Sedimentary Rocks. 2nd ed. Harper and Row. New York. N.Y.. 718 pp. PiquC, A., 1975. RCpartition des zones d’anchimetamorphisme dans les terrains dinantiens du Nord-Ouest du Plateau central (Meseta marocaine). Bull. SOC.GPol. Fr., 17 (3): 417-420. Portugal V. Ferreira, M.R., 1967. Nota previa sobre uma facies zeolitica em Cabo Ledo. Angola. Mem. Not. (Coimhra), 63: 17 pp. Portugal V. Ferreira. M.R.. Pires. C.A.C. and Sousa. M.B.. 1971. Sobre a facies zeolitica de Cab0 Led0 (Angola). Mineralogia de sondagem Puaqa 1. Mem. Not. (Coimhra). no. 70: 36 PP. Powers, M.C., 1959. Adjustment of clays to chemical change and the concept of the equivalence level. Clays Clay Miner., Proc. 6th Natl. Conk. pp. 309-326. Prozorovich, G.E., 1967. Distribution of clay minerals and trace elements in clay fractions of Neocomian rocks of the West Surgut oil field (western Siberia). Dokl. Acud. S c i U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S . S . S . R . ) , 174(1): 948-951. Pustovalov, L.V., 1955. Uber sekundare Veranderungen der Sedimentgesteine. Geol. Rundsch., 43: 535-550. Raam, A., 1968. Petrology and diagenesis of Broughton Sandstone (Permian). Kiama district. New South Wales. J . Sediment. Petrol., 38(2): 319-331. Rateev, M.A. and Murav’ev. V.I.. 1971. International conference on the study of clays. Tokyo, 1969. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 3: 386-390. Read, P.B., 1968. Correction to the reported occurrence of iron-rich saponite in New Zealand. N . Z . J . Geol. Geoph.ys., 11: 1236-1237. Read, P.B. and Eisbacher, G.H.. 1974. Regional zeolite alteration of the Sustut Group, north-central British Columbia. Can. Mineral., 12: 527-541. Reed, B.L. and Hemley, J.J., 1966. Occurrence of pyrophyllite in the Kekiktuk Conglomerate, Brooks Range, northern Alaska. U.S. Geol. Sum., Prof. Pup., 550-C: 162-166. Reynolds, R.C., Jr. and Anderson, D.M.. 1967. Cristobalite and clinoptilolite in bentonite beds of the Colville Group, northern Alaska. J . Sediment. Petrol., 37: 966-969. Reynolds, R.C., Jr. and Hower, J., 1970. The nature of interlayering in mixed-layer illitemontmorillonites. Clays Clay Miner., 18: 25-36.

Reynolds, W.R., 1970. Mineralogy and stratigraphy of Lower Tertiary clays and claystones of Alabama. J. Sediment. Petrol., 40(3): 829-838. Robert, P., 1971. Etude petrographique des matieres organiques insolubles par la mesure du pouvoir reflecteur. Contribution a I’exploration petroliere et a la connaissance des bassins sedimentaires. Reu. Inst. Fr. Pet., 26(2): 105-135. Robie, R.A., Bethke, P.M. and Beardsley, K.M., 1967. Selected X-ray crystallographic data, molar volumes, and densities of minerals and related substances. U.S. Geol. Suru. Bull., 1248: 87 pp. Robinson, P.T., 1966. Zeolitic diagenesis of Mio-Pliocene rocks of the Silver Peak Range, Esmeralda County, Nevada. J . Sediment. Petrol., 36(4): 1007- 10 15. Robinson, P.T., McKee, E.H. and Moiola. R.J., 1968. Cenozoic volcanism and sedimentation, Silver Peak region, western Nevada and adjacent California. Geol. Soc. A m . Mem., 116: 577-6 1 1. Roehler, H.W., 1972. Zonal distribution of montmorillonite and zeolites in the Laney Shale Member of the Green River Formation in the Washakie Basin, Wyoming. U.S. Geol. Suru. Pro/.. Paper 800-B: 21-24. Ross, C.S., 1958. Welded tuff from deep-well cores from Clinch County, Georgia. Am. Mineral., 43: 537-545. Ross, C.S., 1960. Review of the relationships in the montmorillonite group of clay minerals. Clays Clay Miner., Proc. 7th Natl. Congr., pp. 225-230. Rukhin, L.B., 1961. Bases de la Lithologie. Etude des Formations Sedirnentaires. Bur. Rech. Geol. Minieres, Paris, 880 pp. (transl. from Russian). Sabine, P.A., 1963. Volcanic and associated rocks in the coal measures of Colston Bassett (South) borehole, Nottinghamshire. Geol. Mag., lOO(6): 551-555. Sagon, J.-P., 1965. A propos du chloritoide dans les schistes devoniens du bassin de ChAteaulin (region d’Uzel, Saint-Gilles-du-Vieux-Marche; C6tes-du-Nord). C .R. Somm. Soc. Geol. Fr., 8: 269-270. Sagon, J.-P. and Dunoyer de Segonzac, G., 1972. La cristallinite des micas dans les schistes paleozoiques et brioveriens du Bassin de Ch%teaulin(Massif armoricain). C.R. Acad. Sci., Ser. D, 275: 1023-1026. Sakhibgareev, R.S. and Galikeev, K.Kh., 1971. Effect of fracture disturbances on the epigenesis of argillaceous minerals in oil-bearing deposits of the Neocome of the WestSiberian depression. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), I97 l(5): 609-617. Saliot. P., 1973. Les principales zones de metamorphisme dans les Alpes franqaises. Repartition et signification. C.R. Acad Sci., Ser. D., 276: 3081-3084. Schermerhorn. L.J.G., 1975. Pumpellyite-facies metamorphism in the Spanish Pyrite Belt. Pitrologie, l ( 1 ) : 71-86. Scherp, A., 1963. Die Petrographie der palaozoischen Sandsteine in der Bohrung Munsterland 1 und ihre Diagenese in Abhangigkeit von der Teufe. Fortschr. Geol. Rheinl. West/., 1 1 : 25 1-282. Scherp. A., Stadler, G . and Schmidt, W., 1968. Die Pyrophyllit-fuhrenden Tonschiefer des Ordoviziums im Ebbesattel und ihre Genese. Neues Jahrh. Mineral. Abh., 108(2): 142- 165. Schreyer. E.D.. 1969. Zum Vorkommen von Pyrophyllit, Gumbelit und Quartziten in der Kontaktaureole des Bramscher Massivs. Geol. Rundsch., 58(3): 983-997. Schroeder. R.J. and Hayes, J.B., 1968. Dickite and kaolinite in Pennsylvanian limestones of southeastern Kansas. Clays Clay Miner., 16: 41-49. Seki. Y.. 1958. Glaucophanitic regional metamorphism in the Kanto Mountains, central Japan. Jpn. J . Geol. Geogr., 29(4): 233-258. Seki. Y . . 1961. Pumpellyite in low-grade metamorphism. J . Petrol., 2(3): 407-423.

489 Seki, Y., 1966. Wairakite in Japan. Jpn. Assoc. Mineral. Petrol. Econ. Geol., 56(1): 254-261. Seki, Y., 1969. Facies series in low-grade metamorphism. J . Geol. SOC.Jpn., 75(5): 255-266. Seki, Y., 1971. Wairakite-analcime solid solution as an indicator of water pressures in low-grade metamorphism. J . Geol. SOC.Jpn., 77( 10): 667-674. Seki, Y., 1972. Lower-grade stability limit of epidote in the light of natural occurrences. J . Geol. SOC.Jpn., 78(8): 405-413. Seki, Y., 1973a. Distribution and mode of occurrence of wairakites in the Japanese island arc. J. Geol. SOC.Jpn., 79(8): 521-527. Seki, Y., 1973b. Ionic substitution and stability of mordenite. J . Geol. SOC.Jpn., 79(10): 669-676. . and pressure scale of low-grade metamorphism (in Japanese; Seki, Y., 1 9 7 3 ~ Temperature English abstr.). J . Geol. Soc. Jpn., 79( 1 I): 735-743. Seki, Y., Ernst, W.G. and Onuki, H., 1969. Phase Proportions and Physical Properties of Minerals and Rocks from the Franciscan and Sanbagawa Metamorphic Terranes, a Supplement to Geological Society of America Memoir No. 124. Jpn. SOC.Promot. Sci., Tokyo, 85 PP. Seki, Y., Oki, Y., Matsuda, T., Mikami, K. and Okumura, K., 1969a. Metamorphism in the Tanzawa Mountains, central Japan. J . Jpn. Assoc. Mineral. Petrol. Econ. Geol., 61( I ) : 1-75. Seki, Y., Onuki, H., Okumura, K. and Takashima, I., 1969b. Z,eolite distribution in the Katayama geothermal area, Onikobe, Japan. Jpn. J . Geol. Geogr., 40(2-4): 63-79. Seki, Y., Oki, Y., Odaka, S. and Ozawa, K., 1972. Stability of mordenite in zeolite facies metamorphism of the Oyama-Isehara district, east Tanzawa Mountains, central Japan. J . Geol. SOC.Jpn., 78(3): 145-160. Seki, Y., Onuki, H., Oba, T. and Mori, R., 1971. Sanbagawa metamorphism in the central Kii Peninsula, Japan. Jpn. J. Geol. Geogr., 41(2): 65-78. Senderov, E.E., 1968. Experimental study of crystallization of sodium zeolites under hydrothermal conditions. Geochem. Int., 5( I): 1-12 (transl. from Geokhimiya, 1968(1): 3-16). Sheppard, R.A. and Gude, A.J.. 1968. Distribution and genesis of authigenic silicate minerals in tuffs of Pleistocene Lake Tecopa, Inyo County, California. U.S. Geol. Suro., Prof: Pap., 597: 38 pp. Sheppard. R.A. and Gude, A.J., 1969. Diagenesis in tuffs of the Barstow Formation, Mud Hills, San Bernardino County, California. U.S. Geol. Surv., Prof: Pap., 634: 35 pp. Shimazu, M., Tabuchi, A. and Kusuda, T., 1971. Metamorphism in the northeastern part of the Tanzawa Mountainland (in Japanese; English abstr.). J . Geol. SOC. Jpn., 77( 1 1 ) : 701-722. Shutov, V.D. and Dolmatova, T.V., 1961. The character of kaolinite alteration in terrigenous rocks during deep-seated epigenesis (in Russian). Acta Univ. Carolinae. Geol. Suppl., I : 393-415. Shutov, V.D., Drits, V.A. and Sakharov, B.A., 1969. On the mechanism of a postsedimentary transformation of montmorillonite into hydromica. Proc. Int. Cluy Con&., Tok-vo. l Y 6 Y . Isr. Univ. Press, Jerusalem, pp. 523-533. Shutov, V.D., Aleksandrova, A.V. and Losievskaya, S.A., 1970. Genetic interpretation of the polymorphism of the kaolinite group in sedimentary rocks. Sedimentologv, 15: 69-82. Smith, J.V. and Yoder, H.S., Jr., 1956. Experimental and theoretical studies of the mica polymorphs. Mineral. Mag.. 31(234): 209-235. Smith, R.E., 1969. Zones of progressive regional burial metamorphism in part of the Tasman geosyncline, eastern Australia. J . Petrol., 10: 144- 163. Smykatz-Kloss, W. and Althaus, E., 1974. Experimental investigation of the temperature dependence of the “crystallinity” of illites and glauconites. Bull. Groupe Fr. Argiles, 26(2): 319-325.

490 Spackman, W. and Moses, R.G., 1961. The nature and occurrence of ash-forming minerals in anthracite. Miner. ind. Exp. Stn. Pennsyluania State Univ., Bull. 75 ("Proceedings of the Anthracite Conference"): 1- 15. Stadler, G., 1963. Petrographie und Diagenese der oberkarbonischen Tonsteine in der Bohrung Miinsterland, 1. Fortschr. Geol. Rheinl. West&, 1 1: 283-292. Stadler, G., 1971a. Die Kaolin-Kohlentonsteine aus dem Westfal C und B der Untertagebohrung 150 der Steinkohlenbergwerke Ibbenbiiren und ihre Bedeutung fur die Karbonstratigraphie Nordwestdeutschlands. Fortschr. Geol. Rheinl. West&, 18: 79- 100. Stadler, G., 1971b. Die Vererzung im Bereich des Bramscher Massivs und seiner Umgebung. Fortschr. Geol. Rheinl. West&, 18: 439-500. Starke, R., 1968. Tonmineralparagenesen der Sedimentgesteine. Freiberg. Forschungsh.. C23 1 : 2 1 3-22 1. Steiner, A., 1968. Clay minerals in hydrothermally altered rocks at Wairakei, New Zealand. Clays Clay Miner., 16: 193-213. Steinike, K., 1966. Die Charakterisierung der Diagenese im Bereich der tonhaltigen klastischen Sedimente durch Faziesreaktionen im Sinne von Eskola. Chem. Erde, 25(2): 153- 168. Stewart, R.J., 1970. Petrolop, Metamorphism and Structural Relations of Graywackes in the Western Olympic Peninsula, Washington. Ph.D. Thesis, Stanford Univ., 74 pp. Stewart, R.J., 1974. Zeolite facies metamorphism of sandstone in the western' Olympic Peninsula, Washington. Geol. SOC.Am. Bull., 85(7): 1139-1 142. Stewart, R.J. and Page, R.J., 1974. Zeolite facies metamorphism of the Late Cretaceous Nanaimo Group, Vancouver Island and Gulf Islands, British Columbia. Can. J . Earth Sci., 11: 280-284. Strakhov, N.M., 1958. Schema de la diagentse des depBts marins. Eclogue Geol. Helu., 51: 76 1-767. Strens, R.G.J., 1965. Stability and relations of the AI-Fe epidotes. Mineral. Mag., 35(271): 464-475. Sudo, T., Hatashi, H. and Shimoda, S., 1962. Mineralogical problems of intermediate clay minerals. Clays Clay Miner., Proc. 9th Natl. Conf., pp. 378-392. Surdam, R.C., 1967. Low-grade metamorphism of the Karmutsen Group, Butte Lake area, Vancouver island, B.C. Ph.D. thesis, Univ. California, Los Angeles, 313 pp. Surdam, R.C., 1973. Low-grade metamorphism of tuffaceous rocks in the Karmutsen Group, Vancouver Island, British Columbia. Geol. SOC.Am. Bull., 84(6): 191 1- 1922. Surdam, R.C. and Parker, R.B., 1972. Authigenic aluminosilicate minerals in the tuffaceous rocks of the Green River Formation, Wyoming. Geol. SOC.Am. Bull., 83(3): 689-700. Szadecky-Kardoss, E., Bardossy, Gy., Fiirst, I., Pesty, L., Kliburszky, B., Tomor, E. and 1968. On the montmorillonite facies. Acta Geol. Acad. Sci. Hung., l2(1-4): Tomschey, 0.. 61-65. ratios of coexisting minerals in Taylor, H.P. and Coleman, R.G., 1968. 0'8/0'6 glaucophane-bearing metamorphic rocks. Geol. Soc. Am. Bull., 79: 1727- 1756. Teichmiiller, M. and Teichmiiller, R., 1966a. Geological causes of coalification. Aduan. Chem. Ser., 5 5 : 133-155. Teichmiiller, M. and Teichmiiller, R., 1966b. Inkohlungsuntersuchungen im Dienst der angewandten Geologie. Freiberg. Forschungsh., C219: 155- 195. Teodorovich, G.I., 1961. Authigenic Minerals in Sedimentary Rocks. Consultants Bureau, New York, N.Y., 120 pp. Teodorovich, G.I. and Konyukhov, A.I., 1970. Mixed-layer minerals in sedimentary rocks as indicators of the depth of their catagenetic alteration. Dokl. Acad. Sci. U . S . S . R . ,Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S . S . S . R . ) , 191(5): 174-176.

Teodorovich, G.I., Chernov, A.A. and Kotel’nikov, D.D., 1967. Postsedimentation alteration of clayey material in sediments of the lower division of the Azerbaydzhan productive series. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R . ) , 182(2): 152-155. Thompson, A.B., 1970a. A note on the kaolinite-pyrophyllite equilibrium. Am. J. Sci.. 268: 454-458. Thompson, A.B., 1970b. Laumontite equilibria and the zeolite facies. Am. J. Sci., 269: 267-275. Thompson, A.B., 1971a. Analcite-albite equilibria at low temperatures. Am. J . Sci., 271 : 79-92. Thompson, A.B., I97 1 b. pco, in low-grade metamorphism; zeolite, carbonate, clay mineral, prehnite relations in the system CaO-AI,O, -SO, -CO, -H,O. Contrib. Mineral. Petrol., 33: 145-161. Tomita, K. and Sudo, T., 1968. Conversion of mica into an interstratified mineral. Rep. Fuc. Sci. Kagoshima Univ. 1: 89-1 10. Triplehorn, D.M., 1967. Occurrence of pure, well-crystallized 1 M illite in Cambro-Ordovician sandstone from Rhourde el Baguel field, Algeria. J. Sediment. Petrol., 37(3): 879-884. Triplehorn, D.M., 1970. Clay mineral diagenesis in Atoka (Pennsylvanian) sandstones, Crawford County, Arkansas. J . Sediment. Petrol., 40(3): 838-847. Turner, F.J., 1968. Metamorphic Petrology: Mineralogical and Field Aspects. McGraw-Hill. New York, N.Y., 403 pp. Turner, F.J. and Verhoogen, J., 1960. Igneous and Metamorphic Petrology. 2nd ed. McGrawHill, New York, N.Y.. 694 pp. Twenhofel, W.H., 1926. Treatise on Sedimentation. Dover, New York, N.Y., 926 pp. Twenhofel, W.H., 1950. Principles of Sedimentation. 2nd ed. McGraw-Hill, New York, N.Y.. 673 pp. Umegaki, Y. and Ogawa, T., 1965. A note on occurrence of zeolites in the Miocene formation in Shimane Prefecture, Japan. J. Sci. Hiroshima Univ., Sect. C, Geol. Mineral., 4(4): 479-497. Utada, M., 1965. Zonal distribution of authigenic zeolites in the Tertiary pyroclastic rocks in Mogami district, Yamagata Prefecture. Tokyo Univ., Coll. Gen. Educ., Sci. Pap., 15: 173-2 16. Utada, M. and Minato, H., 1972. Zeolitic zoning of the Neogene pyroclastic rocks in the western area of Shimane Prefecture (Nima-Yunotsu district) (in Japanese; English abstr.). J . Geol. SOC.Jpn., 78(7): 329-340. Velde, B., 1965a. Experimental determination of muscovite polymorph stabilities. Am. Mineral., 50: 436-449. Velde, B., 196513. Phengite micas: synthesis, stability and natural occurrence. Am. J . Sci., 263: 886-913. Velde, B., 1966. Mixed-layer mineral associations in muscovite-celadonite and muscovitechlorite joins. Clays Clay Miner., Proc. 13th Natl. Cont, 1964, pp. 29-32. Velde, B., 1967. Si4+ content of natural phengites. Contrib. Mineral. Petrol., 14: 250-258. Velde, B., 1968. The effect of chemical reduction on the stability of pyrophyllite and kaolinite in pelitic rocks. J . Sediment. Petrol., 38: 13-16. Velde, B., 1969. The compositional join muscovite-pyrophyllite at moderate pressures and temperatures. Bull. Soc. Fr. Mineral. Cristallogr., 92: 360-368. Velde, B. and Hower, J., 1963. Petrographical significance of illite polymorphism in Paleozoic sedimentary rocks. Am. Mineral., 48: 1239- 1254. Velde, B. and Kornprobst, J., 1969. Stabilite des silicates d’alumine hydrates. Contrib. Mineral. Petrol., 2 1 : 63-74.

Veselovskaya, M.M., 1967. Characteristics of argillaceous strata in the stage of catagenesis and initial metagenesis as illustrated by a study of the Wendian and Riphean of the Russian Platform. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Doki. Akad Nauk S.S.S.R.), 176(1): 165-166. Vlasov, V.V., 1966. An example of the dependence of epigenetic changes in clay minerals from lithological characteristics of the rocks. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akud. Nauk S.S.S.R.), 170(5): 184-186. Vlodarskaya, V.R. and Nosov, G.I., 1973. Species stability of clay minerals in the zone of catagenesis. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.),210(5): 212-214. Volkova, A.N., Ivanova, N.V. and Rekshinskaya, L.G., 1974. Giimbelite, a satellite of high-grade metamorphic coals. Lithol. Miner. Resour. (transl. from t i t o / . Polezn. Iskop.), 9(6): 708-718. Von Engelhardt, W., Miiller, G . and Kromer, H., 1962. Dioktaedrischer Chlorit (“Sudoit”) in Sedimenten des Mittlereu Keupers von Plochingen (Wiirtt.). Nuturwissenschaften. 49(9): 205-206. Walther, J., 1893- 1894. Einleitung in die Geologie als historische Wissenschaft. Beobachtungen uber die Bildung der Gesteine und ihrer organischen Einschliisse. Fischer, Jena, 1055 pp. Weaver, C.E., 1958. The effects and geologic significance of potassium “fixation” by expandable clay minerals derived from muscovite, biotite, chlorite, and volcanic material. Am. Mineral., 43: 839-861. Weaver, C.E., 1959. The clay petrology of sediments. Clays Clay Miner., Proc. 6th Natl. Conj, pp. 154- 187. Weaver, C.E., 1961. Clay minerals of the Ouachita structural belt and adjacent foreland. In: P.T. Flawn, A. Goldstein, Jr., P.B. King, and C.E. Weaver, The Ouachita Belt. Univ. Texas Publ. 6120: 147-160. Weaver, C.E., 1967. The significance of clay minerals in sediments. In: B. Nagy and U. Colombo (Editors), Fundamental Aspects of Petroleum Geochemistry. Elsevier, Amsterdam, pp. 37-75. Weaver, C.E. and Wampler, J.M., 1970. K, Ar, illite burial. Geol. Soc. Am. BUN., 81: 3423-3430. Weber, K., 1972a. Notes on determination of illite crystallinity. Neues Jahrh. Mineral. Monatsh., 1972(6): 267-276. Weber, K., 1972b. Kristallinitat des Illits in Tonschiefern und andere Kriterien schwacher Metamorphose in nordostlichen Rheinischen Schiefergebirge. Neues Jahrb. Geol. Palaeontol. Abh., 141(3): 333-363. Whetten, J.T., 1965. Carboniferous glacial rocks from the Werrie Basin, New South Wales, Australia. Geol. SOC.Am. Bull., 76: 43-56. Wilkinson, J.F.T. and Whetten, J.T., 1964. Some analcime-bearing pyroclastic and sedimentary rocks from New South Wales. J . Sediment. Petrol., 34: 543-553. Williams, H., Turner, F.J. and Gilbert, C., 1954. Petrography. Freeman, San Francisco, Ca., 406 pp. Winkler, H.G.F., 1964. Das T-P-Feld der Diagenese und niedrigtemperierten Metamorphose auf Grund von Mineralreaktionen. Beitr. Mineral. Petrol., 10: 70-93. Winkler, H.G.F., 1965. Die Genese der metamorphen Gesteine. Springer, Berlin, 218 pp. Winkler, H.G.F.. 1967. Die Genese der metamorphen Gesteine. 2. Auflage. Springer, Berlin, 231 PP. Winkler, H.G.F., 1968. Wandel auf dem Gebiet der Gesteinsmetamorphose. Ceoi. Rundsch., 57: 1002-1019.

493 Winkler. H.G.F.. 1970. Abolition of metamorphic facies, introduction of the four divisions of metamorphic stage, and of a classification based on isograds in common rocks. Neues Jahrb. Mineral., Monatsh., (1970(5): 189-248. Wise, W.S. and Eugster, H.P., 1964. Celadonite: synthesis, thermal stability and occurrence. Am. Mineral., 49: 1031-1083. Wolf, M., 1972. Beziehungen zwischen Inkohlung und Geotektonik im nordlichen Rheinischen Schiefergebirge. Neues Jahrb. Geol, Palaeontol. Abh., 141(2): 222-257. Wolf, M., 1975. Uber die Beziehungen zwischen Illit-Kristallinitat und Einkohlung. Neues Jahrb. Geol. Palaeontol., Monatsh., 7: 437-447. Wyart, J. and Sabatier, G., 1966. Synthese hydrothermale de la corrensite. Bull. Groupe Fr. Argiles, 18(13): 33-40. Yoshitani, A,, 1965. Zeolites in the Neogene pyroclastic rocks in the eastern part of Tanzawa mountainland, central Japan-Studies on the alteration of the Green Tuff formation, 1. Mem. Coll. Sci., Univ. Kyoto, Ser. B, Geol. Mineral., 31(4): 199-213. Yoder, H.S. and Eugster, H.P., 1955. Synthetic and natural muscovites. Geochim. Cosmochim. Act4 8: 225-280. Zaporozhtseva, A.S., 1960. On the regional development of laumontite in Cretaceous deposits of the Lena coal basin. Izv. Akad. Nauk S.S.S.R., Ser. Geol., 1960(9): 61-69 (in transl.: 1961: 52-59). Zaporozhtseva, AS., 1963. Relation of buried clastic material associations to the sedimentation medium. Dokl. Acad. Sci. U.S.S.R., Earth Sci. Sect. (Dokl. Akad. Nauk S.S.S.R . ) , 151(2): 139-141. Zaporozhtseva, A.S., Vishnevskaya, T.N. and Dubar, G.P., I96 1. Successive change in calcium zeolites through a vertical section of sedimentary strata. Dokl. Acad. Sci. U.S.S.R . , Earth Sci. Sect. (transl. from Dokl. Akad. Nauk S.S.S.R.), 141(2): 1264-1266. Zaporozhtseva, AS., Vishnevskaya, T.N. and Glushinskiy, P.I., 1963. Zeolites from Cretaceous formations in northern Yakutia (in Russian). Litol. Polezn. Iskop., 1963(2): 161-177. Zen, E-an, 1961. The zeolite facies: an interpretation. Am. J . Sci., 259: 401-409. Zen, E-an, 1963. Components, phases, and criteria of chemical equilibrium in rocks. Am. J. Sci., 26 I ; 929-942. Zen, E-an, 1967. Mixed-layer minerals as one-dimensional crystals. Am. Mineral., 52: 635-660. Zen, E-an, 1974a. Prehnite- and pumpellyite-bearing mineral assemblages, west side of the Appalachian metamorphic belt, Pennsylvania to Newfoundland. J . Petrol., 15(2): 197-242. Zen, E-an, 1974b. Burial metamorphism. Can. Mineral., 12: 445-455. Zen, E-an and Thompson, A.B., 1974. Low-grade regional metamorphism: mineral equilibrium relations. Annu.,Rev. Earth Planer. Sci., 2: 179-212.

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495 Appendix A DIAGENESIS OF IRON-RICH ROCKS (Illustrated by the Role of Diagenesis in Oolitic Iron Ores) L. BUBENICEK

INTRODUCTION- DEFINITIONS

Because of its abundance in the earth’ crust, iron is one of the major constituents of rocks and is always present, in varying quantities, in sedimentary formations. The diagenetic evolution, however, may be related to iron-phase changes only in iron-rich rocks. This is particularly true in sedimentary iron ores. It should be kept in mind, however, that the term “ore” is restricted here to iron formations that are economically minable, which implies meeting requirements such as location, thickness, and grade. Iron-rich rocks have more widespread distribution through geological scale and on the earth’ surface than ores. A major difficulty arises when an attempt is made to assign a definite sedimentary origin to ores. It is not sufficient for an iron formation to be interbedded in sedimentary rocks in order to assign a sedimentary origin to iron, because it may be introduced either by solution through replacement of host rocks (substitution type of iron deposits) or through magmatic injection (Kiruna type). The difficulty in assigning an unquestionable sedimentary origin in such cases is the fact that mineralogical changes that occur in iron-rich rocks are so destructive of pre-existent textures and structures that the various stages of evolution cannot be recognized. Iron ores of sedimentary origin have many different physical and chemical properties and have been found in all types of rocks and from all stages of sedimentary evolution. Classically, four fundamental categories, which are presented in order of importance below, may easily be distinguished: (1) The Lake Superior ore type, including low-grade banded ores, B.H.Q, taconites; etc., i.e., different iron-bearing formations (James, 1955); (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

496 particular the role of diagenesis, however, it is necessary to study all the stages of transformation. Many authors working mainly on considerably modified pre-Hercynian ores 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. I t is an extremely important process, particulary in the case of iron, an element the chemical behavior of which varies greatly according to its ionic state. (See discussions in Grubb, 1971; Dimroth and Chauvel, 1973; Tsu-Ming Han, 1968.) At present, after tremendous amount of studies, the ores of Lake Superior type, which are generally called “iron-formations”, have still not been clearly related to the original state of deposition that remains conjectural. (See discussions and references in James, 1955; Gross, 1965.) The main contribution is in relative mineralogical changes that may be observed through morphological transformations and justified by mineral equilibria. To date, however, no complete balance from depositional to post-diagenetic stage has been determined for any one of those ores. The importance of sphaerosiderites 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 exploited 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: (1) texture, (2) nature of associated minerals, and (3) depositional environment. The great peculiarity of these ores is that 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. Some Soviet authors assigned a diagenetic origin to this mineral in a slightly oxidizing environment (Krotov, 1952; Kazakov, 1957; also see Hower, 1961). New categories probably will be added to this list. For example, personal observations convinced the author (Bubenicek, 1968) that in Kiruna ores many residual features may be explained by sedimentary processes, and that the actual state reached by this iron ore may be explained through metamorphic phenomena applied to a sedimentary formation. In the same way, the substitution type of iron ore, which is well illustrated by the Ouenza deposit in Algeria, leads one to believe in replacement of calcareous formation by

497 ferruginous brines at early stages of diagenesis. The brief general survey presented here 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 and through various stages of evolution between syngenesis and metamorphism. These stages have been studied in detail by Bubenicek ( I 968).

HISTORICAL REVIEW OF LITERATURE ON DIAGENESIS IN OOLITIC IRON ORES

In earlier publications on oolitic iron deposits, emphasis was placed on the process of diagenesis in the formation of the deposits, i.e., the concentration of the iron. Thus, numerous hypotheses have been proposed, involving mineralizing solutions and the replacement of pre-existing rocks. As a result of progressive accumulation of new data, however, it wag 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 by various authors varies a great deal and depends largely on which observational scheme is regarded as decisive. The different theories may be classified as follows: (I) 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, compactional fluids, etc.), with replacement of oolitic and calcareous detritus by iron solutions containing SiO, and Fe (Castano and Garrels, 1950). ( 1) Without reworking phenomena: the different minerals were directly formed during the replacement processes (Arend, 1933). (2) With reworking phenomena: the different mineralogical types were formed by oxidation during the reworking process (Cayeux, 1909, 1922; Deverin, 1945). (11) Iron ore deposits resulted from a sedimentary differentiation process predating the deposit, i.e., predating sedimentation. (1) With direct precipitation of iron along with other components at the bottom of the sedimentary basin. (a) Direct formation of different iron ore minerals during precipitation. Formation of oolites in situ as a result of diagenesis. The mineralogical distribution depended on the distance from the shore line (Poustovalov, 1940; Caillere and Kraut, 1954, 1956; Braun, 1963; Petranek, 1964). (b) Formation of oolites in situ in minerals having low iron content

498 (e.g., 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, 1833). (2) Formation of oolites in dynamic conditions before deposition. (a) The oolites were formed directly from various mineralogical types of iron minerals prior to deposition: (i) as part of primary cement (Borchert, 1952); (ii) alone: cement or a few minerals only being of diagenetic origin (Popov, 1955; Tochilin, 1956; Courty, 1959, 1961 ; Formosova, 1959; Dunham, 1960; Teodorovich, 1961). (b)The oolites are of an oxidized nature; different reduced iron minerals present are the result of diagenetic processes (Brown, 1943; Harder, 1951, 1957; Correns, 1952; Kolbe, 1958; Bubenicek, 1961, 1963). These various hypotheses reflect the treading of a new path of research, as new data were obtained. Every genetic theory, however, should not only give an explanation of the facts, but also allow practical conclusiohs and inferences to be drawn. A brief review of the present knowledge on this subject is presented below.

PRESENT KNOWLEDGE OF THE DIAGENESIS OF OOLITIC FORMS OF ORESCONDITIONS OF DEPOSITION

The data available at present leads one to believe that the sedimentary differentiation of 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 place in the basin itself, where iron is precipitated and oolites are formed. It appears that there is no direct relation between the influx of iron and the evolution of the basin. Because of the oolitic form, iron is deposited together with various detrital grains, mainly quartz grains or fragments of various shells. At this stage the last possible concentration o f t h e iron takes place by variation of the relative proportions of detrital particles. Consequently, there are deposits exhibiting primary structures of currentbedding and cross-bedding, which reflect current effects. 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 most oxidized form of iron (Fe”) 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

499 iron deposits subsequent changes did not lead to the subtraction or addition of major components, allow a preliminary interpretation of the chemical composition of these ores. Some conclusions can be drawn on the basis of the almost constant amounts of A1,0, and P, along with the dominant iron, especially in the original limonite: (1) In the case of non-clay ores, A1,0, and phosphorus contents’ are proportional to the amount of iron (the more common Al,O,/Fe ratio is 0.10-0.12 in Recent deposits and 0.05 in Paleozoic deposits). This is valid for all the iron contents, which depend on the degree of original dilution with quartz grains or calcite from shell debris; (2) An excessive amount of A1*03 indicates clay ores.

DIAGENETIC EVOLUTION

Limits of diagnesis

When discussing diagenesis, it is always necessary to define the limits of this process, because of lack of agreement among various authors. The present author assigns to diagenesis all the processes which act on the sediment after the end of mechanical movements of the particles 2 , whether during 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 water-sediment interface can preserve syngenetic conditions over a long period of time. Generally, 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 slow circulation which depends on the permeability of the sediment and on possible ionic composition and charge differences. Circulation occurs on the beaches where the water of waves returns to the sea through the deposited sand3. When the exchange of matter with exterior is low, the

’ ’

For the whole ore body, taking into consideration the whole thickness of a layer, and not only sections where segregation phenomena could be involved. It should be kept in mind, however, that the “compactional diagenesis” inuolutng movement of particles is very important. (Editorial comment.) Cementation and/or solution occurring as a result of this process is considered as part of diagenesis by many authors. (Editorial comment.)

500 “closed milieu” conditions of evolution are approached. The end of diagenesis is more difficult to establish. The present writer places it at the beginning of epigenesis, when hydrated minerals are transformed into less hydrated forms (e.g., limonite to hematite and 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. Other transformations, however, may occur beyond this point as has been shown for metamorphosed ores, e g , for hematite and magnetite type ores (see Bubenicek, 1965). Role of diagenesis in the paragenesis of iron-bearing minerals Nature of the diagenetic environment. The diagenetic environment mainly consists of two phases: (1) solid phase: the detrital sedimentary phase where the dominant iron-bearing components are present in the mdst oxidized forms; (2) liquid phase: 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 (e.g., 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 (see Bubenicek, 1964). The diagenetic environment has a certain oxidation-reduction potential and capacity, which is a function of the quantity of the confined organic matter or 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 sulfate-reducing bacteria. Furthermore, interstitial solutions contain an important reserve of a variety of different ions in variable concentrations. A new diagenetic environment is established more or less rapidly at various depths 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: (1)If the level at which Eh=O is within the sediment, its depth will depend on the conditions given above. In this case the changes in the sediment will be of a diagenetic nature. (2) If the level at which Eh = 0 is above the top of the water-sediment interface, deposition of products by direct precipitation from solution could occur. In t h s case one may only apply the term syngenesis, as defined by sedimen tologists.

50 1 Nature of the transformations. The passage from one environment to another, having different physicochemical characteristics, results in a reorganization of the chemical elements in order to establish a new and more stable equilibrium. The previously-formed minerals become unstable and new minerals are precipitated in a stable form, accompanied by exchange of chemical elements (addition or subtraction) with the interstitial solutions. The formation possibilities of the principal iron minerals were discussed by Garrels (1961), who defined the theoretical conditions of their stability (Fig. A-1). Although, theoretical, these diagrams show that all pH-Eh variations involve crossing of stability boundaries which occurs during diagenesis. The following should be noted, however: (1) These diagrams should be expanded by adding areas where the mineral species can exist in a metastable state. (2) The stability diagrams only give an imperfect picture of the reactions, because on assuming the existence of a closed environment the concentrations would vary in accordance with the displacement ,of the reactions. Consequently, threshold to other reactions would appear (Fig. A-2). Main transformations. The main reactions which can occur, as already suspected on studying many deposits, have been confirmed by studies of the Lorraine iron ore deposits. The first fundamental reaction for the non-clay ores is:

limonite

+ quartz

-

siderite

+ chlorite.

The determination of the relative proportions of siderite and chlorite is made possible by the invariance of A1,0, and iron contents and by the knowledge of the Al,O,/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 composition of the phases given in Table A - I 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 A1,0,, Fe, SiO,, Mn, and P; MgO and S are added, whereas H,O is subtracted. The origin of the CO, is open to discussion; however, pyrite forms in the siliceous sediments where calcite is not abundant. This would suggest that at least this CO, 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, the first threshold would

0 10

N

3

Fig. A-1. Stability diagrams of the main iron minerals under various experimental conditions; T = 25°C; p = 1 atm. (After GARRELS, 1961.) A. Absence of silica and LO,, ZS = 10 '. B. Absence of silica, ZCO, = ZS = C. Presence of silica, CCO, = zs =

503 100

A 50 c .-

c

c 0

0 3

0

A

100

0

Fig. A-2. Diagrams of the evolution of limonite-quartz-calcite ores during diagenetic reduction. These ores have not undergone metamorphism. Lorraine, France; central Great Britain; Jurassic deposits of Germany, etc. (After Bubenicek, 1963.) A. Evolution when quartz is in excess. B. Evolution when limonite is in excess.

TABLE A-I Chemical composition of the iron-bearing phases of the minette of Lorraine

Limonite Chlorite Siderite

52.0 31.9 31.1

4.0 29.1 -

0.5 1.5 5.3

1.2 6.5 5.6

6.0 9.9

0.7 0.2

-

-

11.5 10.9 39.2

504

appear when one of these components had been completely consumed in the reaction forming chlorite and siderite. New transformations then intervene: ( 1 ) In the case of excess of the limonite, there is formation of hematite, siderite and then magnetite. (2) When quartz is present 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 becomes equal to 6.05. The various parageneses and their evolutionary relationships are presented diagrammatically in Fig. A-2. The limonite and the quartz transformations as a result of fundamental reduction reaction can cease in the case of exhaustion of 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 limonite would be the result. The boundaries beyond which new parageneses occur may be determined through calculations.

Fig. A-3. Calcite concretions. There is a continuity of the laminae 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 CaCO, displacement; X 0.25.

505 Role of diagenesis in modifying structures As noted earlier, the primary and fundamental structures within the oolitic iron ores are current-bedding (most common) and contorted structures. Diagenesis can modify these structures in two different ways: (1) By the general volume shrinkage due to the 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. A-3). usually by segregation of some primary component (e.g., calcite) or of components produced during diagenesis (e.g., siderite and 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 energy, this phenomenon is frequent for minor components (Seibold, 1955; Pettijohn, 11956). Differences in structure or texture can control the arrangement of the concretions, which may be scattered in beds rich in concretions or may be perfectly localized and rounded. The presence of center 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.

ROLE OF DIAGENESIS IN THE FORMATION OF TEXTURES

The final texture of the rock reflects all of 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 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

506 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 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 common case for chlorite; Fig. A-4). (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. A-5). (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. Siderite shows this very often against clay minerals and

Fig. A-4. Development of chlorite from the limonite of oolites. Chlorite appears as a pellicular cement around the limonitic oolites. The advanced transformation of limonite into chlorite appears clearly on some oolites. Lorraine ore; grey bed=chlorite paragenesis; X 150, natural transmitted light.

507

Fig. A-5. Corrosion of a quartz grain by siderite. The two residual areas of quartz show intended 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.

sometimes against calcite. In ores containing very small amounts of quartz, magnetite (euhedral grains) formed as a result of diagenetic reduction. 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, A1,0,, P and Mn in solid phase 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 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, H,O, CO,, 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

508 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 Fe"/Fe,,,,, ratio 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 Fetora,,Fe" , S O z , A1,0,, 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. All deposits which have been deeply buried or affected by an orogeny, and especially the older deposits, have been subjected to important modifications which frequently mask the earlier textures. It would appear, nevertheless, that the chemical characteristics established during diagenesis (particularly the reduction index FeZt/Fe,,,,,) are preserved for a long time if the deposit has not been subjected to weathering, recent or ancient. REFERENCES A N D BIBLIOGRAPHY Ailing, H.L., 1947. Diagenesis of the Clinton hematite ores of New York. Bull. Geol. Soc. A n ] . . 58: 99-1018. Amstutz. G.C.. 1962. L'origine des gites mineraux concordants dans les roches sedimentaires. Chron. Mines Outre-Mer Rech. Miniere, 308: 1 15- 126. Anderson. G.T. and Han, T.M., 1957. The relationship of Diagenesis, Melamorphisrn and Secondary Oxidation to the Concentrating Characteristics of the Negaunee Iron-Formation of /he Marquette Range. Unpubl. Rep., 4 pp. Arend. J.P.. 1933. Les particularites genetiques du bassin de Briey et leurs rapports avec la repartition, la constitution et les proprietes metallurgiques des minerais oolithiques. Reo. MPt. (Paris). 31: 43-53, 142-151, 188-199, 227-231. Augusthitis, S.S., 1962. Mineralogical and geochemical changes in the diagenetic and postdiagenetic phases of the Ni-Cr-ion oolitic deposit of Larymna/Locris, Greece. Chem. Erde, 2: 1-17. Baas Becking. 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. Berg. G.. 1924. Die Entstehung der sedimentaren Eisenerze. Geol. Rundsch., 15: 96-1 10. Bichrlonne. J. et Angot. P., 1939. Le Bassin Ferrifkre Lorraine, 1. Berger-Levrault, Nancy-Strasbourg. 464 pp.

509 Blondel, F.. 1955. Les types de gisements de fer. Chron. Mines Outre-Mer Re&. Mirii6re. 23 1: 226-246. Borchert. H., 1952. Die Bildungsbedingungen mariner Eisenerzlagerstatten. Chem. Erde. 60: 49-73. Braconnier, M.A., 1883. Description Geologique et Agronomique des Terrains de Meurthe-etMoselle. Berger-Levrault, Nancy-Paris, 436 pp. Braun, H., 1963. Zur Entstehung der marin-sedimentiiren Eisenerze. Borntrager. Berlin. 133 pp. Brown. J.S.. 1943. Supergene magnetite. Econ. Geol.. 38: 137- 148. Bubenicek, L., 1961. Recherches sur la constitution et la repartition des minerais de fer dans I'Aalenien de Lorraine. Sci. Terre, 8( 1-2): 5-204. Bubenicek, L., 1963. Les parageneses des minerais lorraines. R e v . Ind. Miner.. 45(7): 503-506. Bubenicek. L., 1964. L'oxydo-reduction en sedimentologie. Revue synthetique et critique. Bull. Bur. Rech. Geol. Miniere. 4: 36 pp. Bubenicek, L.. 1965. Les parageneses des minerais de fer oolithiques des terrains primaires de I'Ouest de la France. R e v . Ind. Miner.. 47: 1-23. Bubenicek, L., 1966. Geologie des gisements de fer: metallogenie ou geochimie? Minerul. Deposita. I( I ) : 43-55. Bubenicek, L.. 1968. Geologie des minerais de fer oolithiques. Mineral. Depositu. 3: 89- IOX. Bubenicek, L., 1971. Geologie du gisement de fer de Lorraine. Bull. Cent. Rech. Ptru-SNPA. 5(2): 223-320. Bushinskij, I., 1956. Sur la diagenese en rapport avec la genese des argiles refractaires. des minerais de fer sedimentaires et des bauxites. 1:v. A k a d . Nauk S . S . S . R . . Ser. Geol.. 1 1 : 3-115. Cailkre. S. et Kraut, F.. 1954. Les gisements de fer du bassin Lorraine. Meni. Mus. Nuti. Hist. Nut., Ser. C , 4( I ) : 175 pp. Caillere, S. et Kraut, F., 1956. Quelques remarques sur la genese du fer Ordovicien de Segre (Loire inferieure). C . R . Acad. Sci.. 238: 1499-1501. Castano, J.R. and Carrels. R.M.. 1950. Experiments on the deposition of iron with special reference to the Clinton iron ore deposit. Econ. Geo!.. 45: 755-770. Cayeux, L., 1909. Les Minerais de Fer Oolithique de France. I . Minerair de Fers Priniaires. Imprimerie Nationale, Paris, 294 pp. Cayeux, L., 1922. Les Minerais de Fer Oolithiques de France. E. Minerais de Fers Sec,onduire.r. Imprimerie Nationale, Paris, 105 1 pp. Correns. C.W.. 1947. Uber die Bildung des sedimentaren Eisenerze. Forsch. Fortschr.. 2 1 -23(4,5.6): 59-60. Correns, C.W., 1952. Zur Geochemie des Eisens. Congr. Geol. Int. lYme, Al,giers, 1952. C.R.. 2: 23-27. Coutry, G., 1959. Contribution a I'etude du mineral de fer de May-sur-Orne (Calvados), I . Bull. Soc. Geol. Fr.. 5(7): 500-510. Country, G.. 1961. Sur la sideritisation d'oolithes chloriteuses au sommet de la couche de mineral de fer de St. Remy (Calvados). C . R . Acad. Sci.. 252: 301-303. Deverin. L., 1945. Etude petrographique des minerais de fer oolithiques du Dogger des Alpes suisses. Mater. Geol. Suisse, Ser. GPotech., 13(2): 1-1 15. Dimroht, E. and Chauvel, J.J., 1973. Petrography of the Sokoman iron formation in part of the central Labrador Trough, Quebec, Canada. Geol. SOC.A m . Bull., 84: 1 11- 134. Dunham, K.C.. 1960. Syngenetic and diagenetic mineralization in Yorkshire. Proc. Yorks. Geol. Soc., 32(1 I): 229-284. Epprecht, W., 1946. Die Eisen- und Manganerze am Gonzen. Beitr. Geol. Schweiz.. Geotech. Ser.. 24: 1-128.

5 10 Formosova, L.H., 1959. Les Minerais de Fer du PreQrafNord, 1. Geol. Inst. Acad. Sci. USSR., Moscow, 444 pp. Garrels, R.M., 1960. Mineral Equilibria. Harper and Row, New York, N.Y., 254 pp. Gross, G.A., 1965/1967. Geology of Iron Deposits in Canada. 1. General Geology and Evaluation of Iron deposits, 181 pp.; 2. Iron deposits in the Appletrian and Greenwille regions of Canada, 11 pp. Geol. Sum. Can. Econ. Geol. Rep., 22. Grubb, P.L.C., 1971. Silicates and their paragenesis in the Brockman iron formation of Wittenoom Gorge, Western Australia. Econ. Geol., 66: 282-292. Hallimond, A.F., 1925. Iron ores, bedded ores of England and Wales. Geol. Sum. G . B., Mem. Geol. Surv. Spec. Rep. Miner. Resour., G.B., 29: 26-27. Harder, H., 1951. Uber den Mineralbestand und die Entstehung einiger sedirnentarer Eisenerze des Lias y . Heidelb. Beitr. Mineral. Petrogr., 2: 455-476. Harder, H., 1957, Zum Chemismus der Bildung einiger sedimentarer Eisenerze. Z . Dtsch. Geol. Ges., 109(1): 69-72. Harms, J.E., Whitehead, T.H. and Heaton, J.B., 1961. Syngenesis in some Australian iron formations. In: Syngenesis in Ore Deposition-Symp. A . N . Z . A . A . S . , Brisbane, pp. 1-10 (unpubl.). Hough, J.L.. 1958. Fresh-water environment of deposition of Precambrian banded iron formations. J . Sediment. Petrol., 28: 414-430. Hower, J., 1961. Some factors concerning the nature and origin of glauconite. A'm. Mineral., 46: 313-334. James, H.L., 1955. Sedimentary facies of iron formation. Econ. Geol., 49: 235-293. Kazakov, A.V., 1957. La glauconite. Tr. Inst. Geol. Nauk Akad. Nauk S . S . S . R . , Geol. Ser., 152(64): 39- 142. Kazakov, A.V.. Tikhomirova, M.M. et Plotnikova, V.J., 1957. Systerne FeO-C0,-H,O et conclusions sur la paragenese des siderites et phosphorites. Tr. Inst. Geol. Nauk S.S.S.R., Geol. Ser., 152(64): 59-71. Kolbe. H., 1958. Die Erzablagerungen im Salzgittergebiet. Geogr. Rundsch., lO(3): 92-99. Lemoalle, J. and Dupont, B., 1973. Iron bearing oolites and the present conditions of iron sedimentation in Lake Tchad (Africa). In: G.C. Amstutz and A. Bernard (Editors), Ores in Sediments. Springer, Berlin, pp. 167-178. Lindgren, W., 1933. Mineral Deposits. McGraw-Hill, New York, NY, 930 pp. Love, L.G., 1964. Early diagenetic pyrite in fine-grained sediments and the genesis of sulphide ores. In: G.C. Amstutz (Editor), Sedimentologv and Ore Genesis. Elsevier, Amsterdam, pp. 1 1- 19. Krotov, B.P., 1952. Sur le probleme d e la diagenese des sediments: diagenese et teleogenese. Dokl. Akad. Nauk. S . S . S . R . , 82(6): 973-976. McKinstry. H.E.. 1949. Mining Geology. Prentice-Hall, New York. NY, 680 pp. Mohr, P.A.. 1963. Geochemistry of authigenic magnetite from a sedimentary carbonate rock. Univ. CON.Addis Ababa, Fac. Sci., Contrib. Geophys. Obs. Ser. A , 3: 1 I . Oftedahl, C., 1958. A theory of exhalative-sedimentary ores. Geol. Foren. Stockholm Forh., 80(1): 19 pp. Ostroumov. E.A. and Shilov, V.M., 1956. Distribution of sulfides of iron and hydrogen in deep sediments in the northwestern Pacific Ocean. Geochemistry ( U .S. S . R . ) (English transl.). 1960(7): 669-683. Petranek, J.. 1964. Shallow-water origin of Early Paleozoic oolitic iron ores. In: L.M.J.U. van Straaten (Editor), Deltaic and Shallow Marine Deposits. Elsevier, Amsterdam, pp. 3 19-322. Pettijohn. F.J.. 1956. Sedimentary Rocks. Harper, New York, NY, 2nd ed., 718 pp. Popov, B.P.. 1955. Au sujet des carbonates et des silicates dans les minerais de fer de la

511 presqu'ile de Kertch. Tr. lnst. Geol. Nauk, Akud. Nuuk Ukr. R . S . R . . Ser. Prtrogr.. Mineral., Geokhim., 61: 97- 100. Poustovalov, L.V., 1940. Petrographie des Roches Sedimenraires. 1-2. Gostoptekhirdat. Moscow. Ramberg, H., 1952. The origin of Metamorphic and Merasomaric Rocks. 1. Univ. Chicago Press, Chicago, Ill.. 317 pp. Ramdohr, P., 1960. Die Erzmineralien und ihre Veru~uchsungen.Akademie Verlag. Berlin. 1089 PP. Routhier, P., 1963. Les Gisements Metalliferes. Geologie et Priricipes de Recher.che.s. Masson. Pans, 1282 pp. Schnelderhohn, H., 1962. Erzlagerstatten, 4. Fischer, Stuttgart, 37 1 pp. Seibold, E., 1955. Zum Phosphat. Eisen- und Kalkgehalt einiger Horizonte des suddeutschen Jura. Geol. Jahrh., 70( 1955): 577-610. Stanton, R.L., 1964. Textures of stratiform ores. Nature, 202(4928): 173- 174. Strachov, N.M., 1953. La diagenese des sediments et son importance pour la metallogenese sedimentaire. I n . Akad. Nauk S.S.S. R., Ser. Geol.. 5 : 12- 19. Taylor, J.H., 1949. Petrology of the Northampton Sand-Ironstone Formation. Geol. Surv. G. B., M e m . Geol. Suru. G .B., Engl., Wales. 1949: 11 1 pp. Taylor, J.H., 1955. Concentration in sediments. In: Natural Processes of Mineral Concentration. Inter-Univ. Geol. Congr., 3rd, Durham, Proc., pp. 15-20. Teodorovich, G.I., 1958 (1961). Authigenic Minerals i n Sedimentuq Rocks. Consultants Bur.. New York, NY, 120 pp. Tochilin, M.S., 1956. Geochemistry of authigenic siderites. Vopr. Mineral. Osud. O h : . , 3-4: 203-2 1 1. Tsu-Ming Han, 1968. Ore mineral relations in the Cuyuma sulfide deposit. Minnesota. Mineral. Depositu. (3): 109- 134. Williams, H., Turner, F.J. and Gilbert, C.M., 1954. Perrogruphx. Freeman, San Francisco. CA, 406 pp. 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. A m . Assoc. Pet. Geol., 30: 477-513.

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513 Appendix B MINERALOGY A N D PETROLOGY OF BURIAL DIAGENESIS (BURIAL METAMORPHISM) A N D INCIPIENT METAMORPHISM IN CLASTIC ROCKS (Chapter 5)

LITERATURE PUBLISHED SINCE 1976 * (Arranged by subject) HANAN J. KISCH

Changes in clay-mineral assemblages during burial diagenesis and incipient metamorphism: general reoiew

Hoffman, J. and Hower, J.. 1979. Clay mineral assemblages as low grade metamorphic geothermometers: application to the thrust faulted disturbed belt of Montana, U.S.A. In: P.A. Scholle and F.R. Schluger (Editors), Aspects of diagenesis. Soc. Econ. Paleonrol. Mineral., Spec. Publ., 26: 55-79. [Mineral assemblages (principally mixed-layer clays and zeolites) indicate low-grade metamorphism at 100"-2OO0C; this heating is ascribed to burial between thrust plates.] Lippmann, F., 1977. Diagenese und beginnende Metamorphose bei Sedimenten. Bull. Acad. Serbe Sci. Arts, CI. Sci. Nat. Math., 56(15): 49-67. Timofeev, P.P., Kossovskaya, A.G., Shutov, V.D., Bogolyubova. L.I. and Drits, V.A., 1974. New aspects of the study of stages of sedimentary rock development. Lirhol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 9(3): 3 18-336. [Review of diagenesis of clay minerals and of organic matter; very little concerning their correlation.] Experimental and thermodynamical studies and structural transformations of tlarious groups of clqr~minerals (smectite and kaolinite to illite or chlorite)

Lippmann, F., 1979. Stabilitatsbeziehungen der Tonminerale. Neues Jahrb. Mineral. Abh., 136(3): 297-309. Velde, B., 1977. Clays and Clay Minerals in Natural and Synthetic Systems. Elsevier. Amsterdam, 218 pp. [Extensive treatment of all major clay-mineral groups and zeolites.] Frank-Kamenetskij. V.A., Kotov, N.V., Gojlo, Eh. A. and Tomashenko. A.N., 1976. Strukturelle und genetische Wechselbeziehungen zwischen Schichtsilikaten (Phyllosilikaten) und einige Problemen der Tonmineralogie. Z . Angew. Geol., 22: 85-92 [Stepwise transformation of kaolinite and montmorillonite in aqueous and other solutions. Discussion of thermodynamical factors, stability-metastability of the phases formed, and controls by the structure of the starting materials.] Kotov, N.V., Frank-Kamenetsky, V.A. and Goilo, E.A., 1975. Crystal chemistry and thermodynamics of structural transformations of some layer silicates under hydrothermal conditions. Mineral. Pol., 6( 10): 3-27. [Experimental transformation of kaolinite and montmorillonite in the presence of Na, K, Ca and Mg chlorides, sulphates and carbonates under hydrothermal conditions. Inheritance of structure of initial minerals and stability range of the mixed-layer clays are discussed. A dehydration-ionic model is preferred.]

*

The references in this list (completed June 7, 1982, and added in proof) are not inserted in the index of this book.

5 14 Experrmentoi und thermodvnamical studies and structural transformations of dioctahedral three-layer clay minerals

Eberl. D.. 1978. Reaction series for dioctahedral smectites. Clays Clay Miner.. 26: 327-340. [Hydrothermal production of the eight reaction series relating to various dioctahedral clay minerals from beidellite and montmorillonite by making simple changes in interlayer and solution chemistry; significance of assumptions of stability or metastability of the mixed-layer phases for the significance of paragenesis in a reaction series.] Eberl. D.. 1979. Reaction series for dioctahedral smectite: the synthesis of mixed-layer pyrophyllite,’ smectite. In: M.M. Mortland, and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier, Amsterdam, pp. 375-383. [Synthesis from Ca- and Na-saturated montmorillonite in AI3+ solution between 320’ and 400°C.] Eberl, D. and Hower, J., 1976. Kinetics of illite formation. Geol. Soc. Am. Bull., 87: 1326-1330. [Large activation energies suggest that the alteration of smectite to illite involves breaking of chemical bonds in the 2 : 1 layers. The rate of the formation of illite from smectite on the Ocean floor is very slow.] Eberl. D. and Hower, J., 1977. The hydrothermal transformation of sodium and potassium smectite into mixed-layer clay. Clays Clay Miner., 25: 215-227. [Course of the reactions and the appearance of ordered interlayering in the mixed-layer phases are strongly affected by interlayer chemistry ( K versus Na); the difference in hydration energy may account for the fixation of K rather than Na in illite during burial diagenesis.] , Eberl, D., Whitney, G . and Khoury, H., 1978. Hydrothermal reactivity of srnectite. Am. Mineral., 63: 40 1-409. Robertson. H.E. and Lahann, R.W.. 1981. Smectite to illite conversion rates: effects of solution chemistry. Clays Clay Miner., 29(2): 129-135. [Reaction rate and the rate of ordering of mixed-layers were retarded by the addition of N a + , Ca2+ and M g 2 + . ] Velde, B. and Odin, G.S., 1975. Further information related to the origin of glauconite. Clays Clay Miner., 23(5): 376-381. [There is no mineralogical or chemical continuity between illite and glauconite when the K-content is 6 wt,% or greater. Low-grade metamorphism of illitic and glauconitic mica-smectite mixed-layers produces different mica phases.] at 300°C Velde. B. and Weir, A.H., 1979. Synthetic illite in the chemical system K,O-AI,O,-Si0,-H,O and 2 kb. In: M.M. Mortland and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier. Amsterdam. pp. 395-404. [Limited solid solution in a partially-ordered 1 M rnicaceous mineral; compositions more pyrophylliterich than 80% muscovite contain illite-beidellite interstratification.] Also: -

Velde, 1977, Chap. 4 “Montmorillonites” (ref. p. 513).

.YRD methods for determination of percentage expandables

in

illite-smectite mixed-layers

Johns. W.D. and Kurzweil. H., 1979. Quantitative estimation of illite-smectite mixed phases formed during burial diagenesis. TMPM Tschermaks Mineral. Petrogr. Mitt., 26: 203-215. [Modifications to the methods of Reynolds and Hower (1970) and Perry and Hower (1970).] Rettke. R.C.. 1981. Probable burial diagenetic and provenance effects on Dakota Group clay mineralogy. J . Sediment. Petrol.. 51(2): 541-551. [Use of “saddle”/l7 A intensity ratio and of the I/S composite peak position to determine expandability of I/S mixed-layers.] Schultz. L.G.. 1978. Mixed-layer clay in the Pierre Shale and equivalent rocks, northern Great Plains region. U.S. Geol. Surv.. Prof. Pap., 1064-A: 28 pp, reflection for [Determination of the proportion of illite- and smectite-type layers using the S,,/I,, glycolated material; distinction between beidellite and montmorillonite using the expanding behaviour after the Li+ -200°C-glycerol treatment.]

5 15 Srodon, J., 1980. Precise identification of illite-smectite interstratifications by X-ray powder diffractlon. Clays Clay Miner., 28(6): 401-411. [Methods that take layer-spacing variability of the EG complex of dioctahedral smectites into account. as well as techniques fob quantifying the degree of layer ordering and minimizing the error due to the influence of domain size on the positions of the reflections.] Burial-diageneiic transformation of smectite to K-illite (including regular mixed-layers)

Blank, P. and Seifert, U.. 1976. Zur Untersuchung diagenetischer Tonmineralbildungen und deren experimentelle Modellierung. Z. Angew. Geol., 22( 12): 560-564. [Profiles in sedimentary sections compared with experimental data: importance of cation concentrations is stressed.] Boles, J.R. and Franks, S.G., 1979. Clay diagenesis in the Wilcox sandstones of southwest Texas: implications of smectite diagenesis on sandstone cementation. J . Sedrment. Peirol.. 49: 55-70. [Samples from depths of 975 to 4650 m (representing the temperature range 55°-2100C). Temperatures given for ( I ) disappearance of discrete smectite: (2) disappearance of kaolinite: and ( 3 ) replacement of calcite cement by ankerite. Smectites with high ( F e Mg)/AI ratios appear to resist conversion to illite until temperatures high enough to produce ordering are attained.] Eslinger, E. and Sellars, B., 1981. Evidence for the formation of illite from smectite during burial metamorphism in the Belt Supergroup. Clark Fork. Idaho. J . Sedrment. Petrol., 51( I ) : 203-216. [The ratio K-feldspar/plagioclase decreases, and the illite/quartz ratio inSreases downward in a 8500 m section. K-rich 1Md illite formed from K-poor smectite rather than from direct weathering of feldspar.] Eslinger, E., Highsmith, P.. Albers, D. and DeMayo, B., 1979. Role of iron reduction in the conversion of smectite to illite in bentonites in the Disturbed Belt, Montana. Clqvs Clur Miner.. 27(5): 327-338. [Increase in Fe2+/Fe3+ with increasing percentage of illite layers is tentatively attributed to a redox reaction involving the oxidation of organic matter.] Heling, D., 1974. Diagenetic alteration of smectite in argillaceous sediments of the Rhinegraben (SW Germany). Sedimentologv, 2 1 : 463-472. [Different temperatures of disappearance of smectite in different formations related to differences in permeability, and hence of availability of potassium ions.] Heling, D.. 1978. Diagenesis of illite in argillaceous sediments of the Rhinegraben. Clay Miner.. 13: 21 1 . [Alteration of smectite to illite depends primarily on temperature X time. Potassium is supplied by decomposition of feldspar, rather than from distant sources. Mite diagenesis is affected by the inherited layer charge of the initial smectite.] Jonas, E.C., 1975. Crystal chemistry of diagenesis in 2 : 1 clay minerals. In: S.W. Bailey (Editor). Proceedings of ihe International Clyv Conference 1975. Applied Publishing Ltd.. Wilmette. Ill.. pp. 3-13. Lahann, R.W.. 1980. Smectite diagenesis and sandstone cement: the effect of reaction temperature. J . Sediment. Petrol., 50(3): 755-760. [The temperature at which illitization proceeds may control the spatial distribution zones of cementation through the effect of temperature upon the distance of solution transport of silica. Percentage expandable layers versus T are given for four wells in the Gulf Coast.] Nadeau. P.H. and Reynolds, R.C., 1981. Burial and contact metamorphism in the Mancos Shale. Clays Clay Miner.. 29(4): 249-259. [Correlation of percentage expandable layers in illite-smectite mixed-layer with coal ranks and Laramide tectonic activity. Presence of carbonate inhibits illitization. The use of mixed-layered illite-smectite compositions to infer thermal regimes is misleading unless allowance is made for local chemical controls.] McDowell. S.D. and Elders. W.A., 1980. Authigenic layer silicate minerals in borehole Elmore 1. Salton Sea Geothermal Field, California. U.S.A. Contrrh. Mineral. Petrol.. 74: 293-3 10. [Decrease in percentage expandable layers in illite-smectite mixed-layer phase decreases from 10- 15% at 185'C (41 1.5 m depth) to 6% at 210' (494 m); no expandable layers below 725 m. The recrystallized white mica below 850m depth tends to a progressively more muscovitic composition.]

+

5 16 Rettke, R.C., 1981. Probable burial diagenetic and provenance effects on Dakota Group clay mineralogy. J. Sediment. Petrol., 51(2): 541-551. [Ordered illite-smectites are attributed to provenance differences, which gave lower-expandability illite-smectite a “head start” of diagenetic trends.] Schultz, L.G., 1978. Mixed-layer clay in the Pierre Shale and equivalent rocks, northern Great Plains region. U.S. Geol. Surv., Pro6 Pap., 1064-A: 28 pp. [Beidellite- and montmorillonite-type layers are distinguished; generally 20-60% illite layers. In the Montana disturbed belt much of the clay has been altered to regularly interlayered IS-ordered mixed layer with 60-85% illite layers.] Thorez, J. and Pirlet, H., 1979. Petrology of K-bentonite beds in the carbonate series of the Visean and Tournaisian stages of Belgium. In: M.M. Mortland and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier, Amsterdam, pp. 323-332. [Mainly illite-smectite and illite-vermiculite mixed-layers; volcanic origin.] Yeh, H.-W., 1980. D/H ratios and late-stage dehydration of shales during burial. Geochim. Cosmochirn. A d a , 44 (2): 341-352. [Significant D/H fractionation between residual and expelled pore waters. The montmorillonite to illite transformation during burial diagenesis is considered to be the most important mechanism of late-stage dehydration.] Also: -

Aoyagi and Kazama, 1980 (ref. p. 529).

- Foscolos and Powell, 1979 (ref. p. 516). - Hoffman and Hower, 1976 (ref. p. 513). -

Kubler, 1980 (ref. p. 517).

- Powell et al., 1978 (ref. p. 517). - Velde, 1977, Chap. 5, “Illite, montmorillonite and mixed layered minerals in sequences of buried rocks ( P - T space)” (ref. p. 513). Kaolinite-group minerals: polyrypism and transformation to illite

Heling, D., 1980. Tonmineraldiagenese und Palaotemperaturen im gebleichten mittleren Buntsandstein am Westrand des Rheingrabens. Neues Jahrb. Mineral. Monatsh., 1980(1): 1- 10. [Near the western fault of the Rhinegraben the “Bunter” is bleached, and the clay fraction is almost entirely 2M illite (the Bunter normally has IM illite with some 50% kaolinite). This is ascribed to high paleothermal gradients and heat transfer by ascending waters along deep faults during periods of rapid subsidence (e.g., the Oligocene).] Rodionova, A.E. and Koval’skaya, M.S., 1974. Dickite distribution in coal-bearing formations of the Donets Basin. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 9(4): 75 1-755. [Two dickite-forming stages are recognized: (1) in late epigenesis, and (2) in the post-inversion stage (“regressive epigenesis”).] Relationship between burial-diagenetic modification of clay mineralogy (except illite crystallinity) and of organic matter, including hydrocarbon generation

Gindorf, L. and Paetz, H., 1979. Wechselbeziehungen zwischen Organiten und Anorganiten als Indikatoren geologischer Prozesse wahrend der superkrustalen Gesteinsbildung. Z. Geol. Wiss. (DDR), 7(2): 235-240. Foscolos, A.E. and Powell, T.G., 1979. Mineralogical and geochemical transformation of clays during burial-diagenesis (catagenesis): relation to oil generation. In: M.M. Mortland and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier, Amsterdam, pp. 26 1-270. [First dehydration of the interstratified clays coincides with 0.5% R, vitrinite reflectance and occurs several thousand feet above the main phase of oil generation, whereas the second dehydration step takes place below the oil generating zone.] Heroux, Y.,Chagnon, A. and Bertrand, R., 1979. Compilation and correlation of major thermal maturation indicators. Am. Assoc. Pet. Geol. Bull., 63(12): 2128-2144. [Correlation chart of the most commonly used organic and mineral thermal maturation indicators.]

517 Kubler, B., 1980. Les premiers stades de la diagenese organique et de la diagenese minerale. Deuxieme partie: Zoneographie par les transformations mineralogiques, comparaison avec la reflectance de la vitrinite, les extraits organiques et les gaz adsorbes. Bull. Ver. Schweiz. Pet.-Geol. -Ing., 46( 110): 1-22. Powell, T.G., Foscolos, A.E., Gunther, P.R. and Snowdon, L.R.. 1978. Diagenesis of organic matter and fine clay minerals: a comparative study. Geochim. Cosmochim. Acta, 42: 1181-1 197. [Relationships in subsurface samples from Canadian Northwest Territories. Mixed-layer clays comprising smectite-vermiculite-illite are transformed during thermal diagenesis to smectitevermiculite-illite-chlorite. The first clay dehydration occurs at 0.5% R prior to hydrocarbon generation. Vermiculite is an intermediary in the smectite-illite transformation. and in the (common) presence of Ca2+ ions delays the second dehydration step to the zone where cracking of liquid hydrocarbons to gas occurs (between 1.O- 1.2 and 1.4% R o,l ).] Zhelinskii, V.M.. 1980. Catagenesis of terrigenous rocks and metamorphism of coals in south Yakutia. Lirhoi. Miner. Resour. (transl. from Lifol. Polezn. Iskop.), 15(2): 187- 199. [Lateral variation of relationship between alteration of minerals and coal rank is correlated to effects of magmatic bodies or hydrothermal processes.] Burial-diagenetic transformation of smectite and disappearance of kaolinite fwithouf rllire cn.sral1init.v) in sections with coal-rank data

Chudaev, O.V., 1978. Occurrence of clay minerals in flyschoid sediments of eastern Kamchatka. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.). 13(I): 89-97. [Upper part of volcano-sedimentary sequence, with montmorillonite and randdm chlorite-montmorillonite mixed-layers, and middle part, with corrensite-like minerals and “defective chlorite”, have respectively long-flame coal and gas coal, and are assigned to initial epigenesis; the lower part. with chlorite and hydromica, and coking coal, to the “zone of hypogenesis”.] Hutcheon, I., Oldershaw, A. and Ghent, E.D., 1980. Diagenesis of Cretaceous sandstones of Kootenay Formation at Elk Valley (southeastern British Columbia) and Mt Allan (southwestern Alberta). Geochim. Cosmochim. Acta, 44: 1425- 1435. [At Elk Valley the authigenic kaolinite-dolomite assemblage is associated with vitrinite reflectances from 0.8 to at least 1.6% R,,, oil; at Mt Allan kaolinite disappears at 265 m depth (at approx. 1.4% R o,l), and chlorite calcite appear. Reaction is considered to involve a CO, fluid which is immiscible with water under the extrapolated diagenetic conditions.] Ivanova, N.V., Volkova, A.N., Rekshinskaya, L.G. and Konysheva, R.A., 1980, Pyroclastic material in coal measures of the Donets Basin and its diagnosis. Lithol. Miner. Resour. (transl. from Litol. Polern. Iskop.). 1980(6): 709-718. [In the zone of coking and lean coals almost exclusively illite-smectite mixed-layers with ca. 25-30% expandable layers, which tend to regular mixed-layering. In association with anthracites well-crystallized 2M, mica. “Lag” in transformation of pyroclastic material.] Kisch, H.J.. 1981. Burial diagenesis in Tertiary “flysch” of the external zones of the Hellenides in central Greece and the Olympos region, and its regional significance. Ecologae Geol. Helu., 74(3): 603-624. [Wide-spread occurrence of smectite and illite-smectite mixed-layers in the external zones is associated vitrinite reflectance.] with 0.44 to 0.65% R , Kisch. H.J., 1982. Coal rank and illite crystallinity associated with the zeolite facies of Southland and the pumpellyite-bearing facies of Otago, southern New Zealand. N . Z . J . Geol. Geoph-vs.,24 (3): 349-360. [The zeolite facies of the North Range Group of Southland is associated with smectite and vitrinite reflectance.] illite/smectite mixed-layers, and with 0.60 to 1.33 R Pevear, D.R., Williams. V.E. and Mustoe, G.E., 1980. Kaolinite, smectite. and K-rectorite in bentonite: relation to coal rank at Tulameen, British Columbia. Clavs Clus Miner.. 28: 241 -254. [Smectite cristobalite clinoptilolite and smectite kaolinite associated with 0.60-0.708 R , l ; regular illite-smectite with 55% illite layers and rectorite-type (IS) superlattice with up to 0.86% R,,.]

+

+

+

Also:

- Gill et al.. 1977 (ref. p. 520).

+

518 Polvtppes of illire and degree of incipient metamorphism

Cameron. T.D.J. and Anderson, T.B.. 1980. Silurian metabentonites in County Down. Northern Ireland. Geol. J . . 15: 29-15 (Mineral. Abstr.. 32: 81-0164). [Two polymorphs coexist. The clay-size 1M is mainly derived from montmorillonite; the coarser 2M may be an anchimetamorphic alteration of the former. Collapse of mixed-layer clay to illite during Caledonian anchimetamorphism.] Also; - Brime and Perez-Estaun. 1980 (ref. p. 519). - Schramm. 1982b (ref. p. 520) Illite cystallinity-methods

McConchie. D.M.. Ward, J.B., McCann, V.H. and Lewis, D.W.. 1979. A Mossbauer investigation of glauconite and its geological significance. Clays Clay Miner., 27: 339-348. [Defines “disorder coefficient” for glauconites on the basis of the shape of the 10 A X-ray diffraction peak for the heated oriented sample.] Weber, F.. Dunoyer de Segonzac, G. and Economou, C.. 1976. Une nouvelle expression de la “cristallinitt“ de I’illite et des micas. Notion d”‘epaisseur apparente” des cristallites. C.R. Somm. Soc. GCoI. Fr.. 1976(5): 225-221. [“Apparent thickness” based on the Scherrer formula for diffraction by very small crystals.] Illire crystallinity-complications accompanying minerals)

in assessment of the peak width (effect of chemical treatments, grain size,

Clauer. N. and Kroner, A,, 1979. Strontium and argon isotopic homogenization of pelitic sediments during low-grade regional metamorphism: the Pan-African upper Damara sequence of northern Namibia (South West Africa). Earth Planet. Sci. Lett.. 43: 117-131. [Difference in crystallinity index and I,,/I,, ratio between illites in parageneses with or without stilpnomelane and/or microcline.] For comparison of 10 A peak widths of untreated and EG-solvated samples see also: Ahrendt et al., 1977 (ref. p. 527). - Brime and Perez-Estaun. 1980 (ref. p. 519) - Frey et al., 1980 (ref. p. 523), - Kisch. 1980a (ref. p. 520). - Kisch. 1980b (ref. p. 523, 539). - Kisch. 1981 (ref. p. 517). For comparison of 10 A peak widths in different grain-size fractions, see also: - Kisch. 1980a (ref. p. 520). - Kisch. 1980b (ref. p. 523, 539). - Stalder, 1979 (ref. p. 523). - Teichmiiller et al., 1979 (ref. p. 519, 521). For effects of the presence of biotite see also: - Bril and Thiry. 1976 (ref. p. 539). ~

Illite crvstallinit-v- and coal rank-

in relation to intrusive bodies

Rohde. Agnes, 1980. Clay minerals and illite crystallinity of the Almeshkra Group. Geol. Foren. Stockholm Farh.. 102: 26. [Crystallinity values range from “diagenetic” to high-grade anchimetamorphic; samples from the vicinity of diorite tend to have high crystallinities.] Deutloff. 0.. Teichmiiller, M.. Teichmuller, R. and Wolf, M.. 1980. Inkohlungsuntersuchungen im Mesozoikum des Massifs von Vlotho (Niedersachsisches Tektogen). Neues. Jahrb. Geol. Paluontol. M ~ t ~ t s h1980(6):32 ., 1-341.

519 [Both coal rank and chlorite crystallinity increase towards the Vlotho Massif. The retardation of the improvement in illite- but not in chlorite-crystallinity towards the intrusive body is due to high content of organic matter.] Teichmiiller, M., Teichmiiller, R. and Weber, K.. 1979. Inkohlung und Illit-Kristallinitat-vergleichende Untersuchungen, im Mesozoikum und Palaozoikum von Westfalen. Fortschr. Geol. Rheinl. Westf.. 27: 201-276. [Around the Upper Cretaceous “Bramsche Massif” intrusive coalification is more sensitive to heating than illite crystallinity (IC); also in deeper parts of the Miinsterland 1 well. Due to post-kinematic annealing, the R , , in the Lippstadt dome (well Soest-Envitte l / l a ) is higher (up to 10%) than for similar high-grade anchimetamorphic IC values in the Ostsauerland anticline.] IIItte crystallinity-and

coal rank-

in burial-diagenetic sequences

Hutcheon, I., Oldershaw, A. and Ghent, E.H., 1980. Diagenesis of Cretaceous sandstones of Kootenay Formation at Elk Valley (southeastern British Columbia) and Mt. Allan (southwestern Alberta). Geochim. Cosmochim. Acta, 44: 1425- 1435. [Illite crystallinities (after Weber) associated with R,, D,I of up to 2.18 are within the “diagenetic” range. Their increase with depth is irregular: illites at higher depth tend to be somewhat more poorly crystalline than those from samples hgher in the section; this reverse trend of illite crystallinity with depth is ascribed to degradation of detrital illite. However, SEM indicates a regular improvement in crystallite size and morphology with depth.] Kiihn, L., 1979. Untersuchungen am Illit der flozfiihrenden Schichten im Ruhrkarbon (abstr.). Fortschr. Mineral., Beih. (Jahrestag. Dtsch. Mineral. Ges. 1979, Darmstadt), 57( I): 76. [Kaolinite and only very subordinate illite in the coals, and well crystallized illite and chlorite, as well as kaolinite, in the shales: effect of the coaly matter during diagenesis. No relation found between the half height width of the illite 10 A peak and the coal rank.] Also: - Blank and Seifert, 1976 (ref. p. 515). - Teichmiiller et al., 1979 (ref. p. 519. 521). Illite ctystallinity-regional

studies in pre-Alpine belts without relation to coal rank

Aparicio, A. and Galan, E., 1978. El metamorfismo de bajo grado en el area central del Macizo Hesptrico Espaiiol (Sistema Central-Toledo). Bol. Geol. Min. (Spain), 89: 475-486. [Epizonal illite crystallinities. Study of the Na/(Na + I() ratios of the muscovites and the chemical composition of the chlorites allows distinction between the low-pressure type metamorphism of the Lower Cambrian and the intermediate-pressure type of the Lower Ordovician.] Aparicio, A. and Galin, E., 1980. Las caracteristicas del metamorfismo hercinico de bajo y de muy bajo grado en el sector oriental del Sistema Central (Provincia de Guadalajara). Esiud. Geol.. 36: 75-84. [Kubler and Weaver indices and intensity ratios given. Local pyrophyllite with allevardite. or paragonite; some chloritoid in the Silurian. Degree of metamorphism increases from predominantly highest-grade anchimetamorphic in the Carboniferous and Devonian, to “epizonal” in the Silurian and Ordovician.] Brime, C. and Perez-Estaun, A,, 1980. La transicion diagenesis-metamorfismo en la region del Cabo Peiias. Cuad. Lab. Geol. Laxe (Publ. Semin. Estud. Galegos). 1: 85-96. [From E to W along the coast: “diagenetic” Lower Devonian. “diagenetic” to anchimetamorphic Silurian (with pyrophyllite, paragonite. and minor kaolinite). and anchizonal to “epizonal” Middle Ordovician. Illite crystallinity and Im2/Iml ratios on both untreated and EG-solvated samples.] Cailleux, Y., 1979. Les contrBles de la cristallinite des illites dans la partie W du Massif Central Marocain (abstr.). Reun. Annu. Sci. Terre, 7/1979, Lyon, p. 97. Galan, E., Aparicio, A. and Villegas, F., 1978. El metamorfismo de muy bajo grado (anquimetamorfiamo) de la cuenca carbonifera Ciiiera-Matallama (provincia de Leon). Estud. Geol.. 34: 505-5 10. [Kubler and Weaver indices determined. Anchimetamorphism with local paragonite, hut rather common kaolinite.]

520 Also: -

Arkai. 1977 (ref. p. 525). Bevins et al.. 1981 (ref. p. 539). Bril and Thiry, 1976 (ref. p. 539). Hartnady et al., 1978 (ref. p. 525). Lecolle and Roger, 1976 (ref. p. 540). Leitch, 1975 (ref. p. 540). Padan et al., 1982 (ref. p. 526).

Illire crystallinit.~-regional studies in pre-Alpine belts in relation to coal rank

Gill. V.D.. Khalaf. F.I. and Massoud. M.S., 1977. Clay minerals as an index of the degree of metamorphism of the carbonate and terrigenous rocks in the South Wales coalfield. Sedimentology, 24: 675. [The illite sharpness (Weaver) and intensity ratios, and expandabilities of illite/smectite mixed-layers allow distinction of three lateral zones, correlated with coal rank ranges. The highest-grade anchimetamorphic zone (with local pyrophyllite and allevardite) corresponds to the anthracite area in the northwestern region.] Jackson, T.A.. 1977. A relationship between crystallographic properties of illite and chemical properties of extractable organic matter in pre-Phanerozoic and Phanerozoic sediments. Clays Clay Miner., 25: 187-195. (Correlation between illite peak width and degree of lumification. Variations ark mainly primary of predetermined by the primary character of the material: original nature of organic matter could influence post-depositional clay modification.] Kisch. H.J.. 1980. Incipient metamorphism of Cambro-Silurian clastic rocks from the Jamtland Supergroup, central Scandinavian Caledonides, western Sweden: illite crystallinity and “vitrinite” reflectance. In: W.E.A. Phillips and M.R.W. Johnson (Editors) Deformation and Metamorphism in the Caledonide Orogen. J . Geol. Soc. London, 137(3): 271-288. [Illite-crystallinity ranges from “diagenetic” in the E to “epizonal” in the W. Four zones are distinguished, the three higher-grade zones showing a similar reflectance range of 3.7 to 4.3% R max The incipient metamorphism was at least in part due to the overthrust metamorphic allochthon.] Rehmer, J., Hepburn, J.C. and Schulman, J., 1978. The diagenetic to metamorphic transition in an Appalachian coal basin. Geol. SOC.Am. Abstr., lO(7): 477-478. [The < 2 p fraction is non-detrital, as indicated by its more ferromagnesian composition than the coarser. more aluminous fraction. The trends in the illite crystallinity index correlate closely with coal rank data (see following abstract).] Rehmer. J.. Hepburn, J.C. and Ostrowski, M., 1979. lllite crystallinity in sub-greenschist argillaceous rocks and coal, Narrangansett and Norfolk Basins, USA. 9th I n f . Congr. Carbonif. Stratigr. Geol., Urhona (Ill.), 1979, Abstr., p. 176. [Most of the northern part of the Narrangansett Basin has been subjected to anchimetamorphic conditions. The crystallinity indices of anchizone-lower greenschist illites correlate well with anthracite to meta-anthracite coal’ranks; however, in the diagenetic zone the coal rank is higher than normally found with illites of this crystallinity.] Robinson. D.. Nicholls, R.A.. and Thomas, L.J., 1980. Clay mineral evidence for low-grade Caledonian and Variscan metamorphism in south-western Dyfed. south Wales. Mineral. Mag.. 43: 857-863. [Largely high-grade anchizonal illite crystallinities in pelites are associated with prehnite-pumpellyite facies in Lower Paleozoic basic igneous rocks. The Lower and Upper Paleozoic rocks south of the Variscan front show predominantly low-grade anchimetamorphic illite crystallinities, locally with subordinate pyrophyllite: the Pembroke coalfield, in this southern area, has largely high-rank semi-anthracitic and low-rank anthracitic coals.] Rowsell. D.M. and de Swardt. A.M.J., 1976. Diagenesis in Cape and Karroo sediments, South Africa, and its bearing on their hydrocarbon potential. Trans. Geol. Soc. S. Afr., 79( 1): 81-129. [Includes maps of illite-crystallinity values (Kubler index). clay-mineral distribution, porosity/permeability. hulk density, and of various parameters of organic maturity (CR/CT. etc.). There is a general

52 1 decrease in diagenesis from S to N. I n the southern fold belt and some distance to the N. and in the central Karroo basin, the argillites are in the state of “incipient or very early metamorphism” (IargeIy anchimetamorphism-HJK); only in the N part of the Karroo basin is the degree of diagenesis much lower.] Saupe, F.. Dunoyer de Segonzac. G. and Teichmiiller, M., 1977. Etude du metamorphisme regional dans la zone dAlmaden (Province de Cuidad Real, Espagne) par la cristallinite de l’illite et par le pouvoir reflecteur de la matiere organique. Scr. Terre (Nancy). 21(3): 251-269. [The Ordovician-Devonian has undergone a much weaker metamorphism (deep diagenesis and low-grade anchizone) than the Precambrian (anchizone-epizone); however. the reflectance of the organic matter in the Paleozoic suggests a somewhat stronger, high-grade anchizonal metamorphism.] Teichmiiller, M. and Teichmiiller, R., 1979. Ein Einkohlungsprofil entlang der linksrheinischen Geotraverse von Schleiden nach Aachen und die Inkohlung in der Nord-Sud-Zone der Eifel. Fortschr. Geol. Rheinl. WestJ, 27: 323-355. [The maximum values of both illite crystallinity and coal rank (with semigraphite) are reached in the Ordovician and Lower Devonian of the southern limb of the Venn anticline rather than in the Cambrian core, indicating the metamorphism continued after the pre-Asturian folding of the anticline.] Teichmiiller, M., Teichmiiller, R. and Weber, K., 1979. Inkohlung und Illit-Kristallinitat-vergleichende Untersuchungen im Mesozoikum und Palaozoikum von Westfalen. Fortschr. Geol. Rheinl. Westf.. 27: 201-276. [Different coal rank-illite crystallinity relationships depending on whether the recrystallization of the illite is pre- or synkinematic, and whether a post-kinematic re-heating has increased the coal rank (but not the illite crystallinity). It is proposed to define the onset of the anchizohe in terms of coal rank rather than of illite crystallinity.] Vinchon, C., 1977. Contribution a IYtude petrogruphique du Silurien des Pyrenees centrales espagnoles (Region du Rio Eseru, Province de Huesca et Region de Llavorsi, Province de Leridu). Mem. Diplome Etud. approfondies, Univ. Sci. Tech. Lille, 86 pp. [XRD of organic matter, and use of H/L ratio of 002 diffraction peak. The organic matter is perfectly ordered graphite, graphite-d, and graphite-d,, in Landis’ (1971) classification. with d,,3.36-3.39 The muscovite shows epimetamorphic crystallinities; the presence of subordinate kaolinite is ascribed to secondary alteration.] Also: - Ahrendt et al., 1977 (ref. p. 527). - Teichmiiller et al., 1979 (ref. p. 519, 521).

A.

Illite crystallinity-regional studies in Alpine belts, without relation to coal rank

Aprahamian, J. and Pains, J.-L., 1981. Very low grade metamorphism with a reverse gradient induced by an overthrust in Haute-Savoie (France). In: Thrust and Nappe Tectonics. Geol. SOC.London, Spec. Publ., pp. 159-165. [Illite crystallinity of five composite sections in the Plate Massif, corresponding to laumontite zone to prehnite-pumpellyite facies (in the most internal parts) in the Taveyanne sandstones. The crystallinity of illite in the upper part of the internal sections shows a reverse gradient (from anchimetamorphic to “diagenetic”) superimposed upon an earlier gradient, and ascribed to heat produced by friction along the thrust plane of the overlying pre-Alpine nappes.] Aprahamian, J., Pains, B. and Pairis, J.-L. 1975. Nature des mineraux argileux et cristallinite des illites dans le massif de Plate et le revers occidental des Aiguilles Rouges-implications possibles d u n point de w e sedimentaire, structural et metamorphique. Ann. Cent. Uniu. Saooie. I1 (Sci. Nat.): 95- 119. [Two stages of metamorphism recognized (see preceding abstract). The appearance of the assemblage prehnite-pumpellyite in the internal parts of the Plate massif is associated with low-grade anchimetamorphc illite crystallinities. Corrensite locally persists into the anchizone.] Blanc, P. and Obert, D., 1979. Le metamorphisme lie a la phase technique antecenomanienne du domaine tellien septentrional (Babors, Algerie). Bull. Soc. Giol. Fr., 21(2): 189- 193. [Cumulative illite-crystallinity curves for different tectonic units and geologic periods given. The N-S gradient of the late Albian metamorphism can be distinguished from that of the weaker post-Senonian metamorphism: a metamorphic discontinuity at the level of the Cenomanian.]

522 Dumont, J.-F. and Desprairies, A,, 1977. RCsultats preliminaires d’une etude du mttamorphisme dans I’autochtone du Taurus occidental (Coupole d e Karacasihar, Turquie). C.R. Acad Sci. Paris, SPr.D, 284: 1017-1020. [Anchizonal Triassic and low-grade “epizonal” lower Paleozoic is ascribed to successive post-Cambrian and post-Triassic metamorphic phases.] Dunoyer de Segonzac, G. and Abbas, M., 1976. Metamorphisme des argiles dans le Rhetien des Alpes sud-occidentales. Sci. Geol., Bull. (Strasbourg), 29: 3-20. [Sampled at six points from W to E. The anchimetamorphic zone is reached in the subalpine and Brianqonnais domains, the upper “epizone” in the pre-Pikmontais (with minor paragonite). The carbonate rocks of the anchi- and low-grade epizone contain an aluminous montmorillonite, which is considered a metamorphic mineral.] Dunoyer de Segonzac, G . and Bernoulli, D., 1976. Diagenbe et metamorphime des argiles dans le Rhetien Sub-alpin et Austro-alpin (Lombardie et Grisons). Bull. Soc. Geol. Fr., (7), 18(5): 1283- 1293. [Diagenesis to anchimetamorphism in the higher or more external Austro-alpine nappes (Silvretta, Tschirpen); epizone is reached in the more deeply buried lower Austro-alpine Bernina and Err nappes.] Frey, M. and Wieland, B., 1975. Chloritoid in autochthon-parautochthonen Sedimenten des Aarmassivs. Schweiz. Mineral. Petrogr. Mitt., 5 5 : 407-418. [Several occurrences of fine-grained colourless chloritoid. The assemblage pyrophyllite + chlorite does not persist into the chloritoid zone. The chloritoid isograd, where defined, is associated with “epizonal” illite crystallinities, but is external of the pumpellyite-actinolite facies assemblage of Leuk.] Frey, M., Jager, E. and Niggli, E., 1976. Gesteinsmetamorphose im Bereich der Geotraverse Basel-Chiasso. Schweir. Mineral. Petrogr. Mitt., 56: 649-659. [Onset of anchizone, as well as localities of kaolinite, stilpnomelane, pyrophyllite, and pumpellyite are indicated.] Kleberger, J. and Schramm, J.-M., 1980. Ein Metamorphosehiatus an der Salzach-L~gsstorung?Osterr. Akad. Wiss., Anr. Math.-Natunviss. KI., 1980(5): 1-6. [The similarity in the “epizonal” illite crystallinities on both sides of this fault-between the N margin of the penninic schist cover and the S margin of the Graywacke Zone-does not support the assumption of a break in degree of metamorphism.] Schramm, J.M., 1978. Anchimetamorphes Permoskyth an der Basis der Kaisergebirges (Siidrand der nordlichen Kalkalpen zwischen Worgl und St. Johann in Tirol, Osterreich). Geo/. Paliiontol. Mitt. (Innsbruck), 8: 101-1 11. Schramm, J.-M., 1982a. Anchmetamorphose im klastischen Permoskyth der Schuppenzone von Gostling (Nordliche Kalkalpen, N.O.). Verh. Geol. Bundesanst. (Wien), 1982 (2): 53-62. [Better crystallinities in the vicinity of this fault are ascribed to post-metamorphic upward drag of deep elements of the Northern Calcareous Alps.] Schramm, J.-M., 1982b. Uberlegungen zur Metamorphose des klastischen Permoskyth der Nordlichen Kalkalpen vom Alpenostrand bis zum Ratikon (Osterreich). Verh. Geol. Bundesanst. (Wien), 1982(5): 73-83. [The Graywacke Zone is “epimetamorphic”. There is no hiatus between the metamorphic overprint of this zone and the adjoining southern margin of the Northern Calcareous Alps. The anchimetamorphism in the latter decreases northwards, and ceases 5-10 km S of the northern margin of the Calcareous Alps.] Venturelli, G. and Frey, M., 1977. Anchizone metamorphism in sedimentary sequences of the northern Apennines. Rend. Soc. Ital. Mineral. Petrol., 33( I): 109-123. [The shales of all except one tectonic unit (Monte Caio) show illite crystallinities characteristic of deep diagenesis and anchimetamorphism, while nearby ophiolitic rocks show prehnite-pumpellyite facies (see Cortesogno and Venturelli, 1978). Some data on do,, of illite.] Wieland. B.. 1979. Zur Diagenese und schwachen Metamorphose eozaener siderolithischer Gesteine des Helvetikums. Schweiz. Mineral. Petrogr. Mitt., 59: 41-66. [Extremely Fe-rich illites, tending towards tri-octahedral; local pyrophyllite. rectorite, paragonitephengite. and montmorillonite-illite mixed-layers. In E and central area illite crystallinities of limit anchi-epizone, and pyrophyllite; in W area medium- and low-grade anchizone, and presence of kaolinite.]

523 Also: - Arkai, 1973 (ref. p. 539). - Bonhomme et al., 1980 (ref. p. 527). Illite crystallinity and lowest-grade metamorphic zoning in the Alps-in

conjunction with coal-rank studies

Frey, M., Teichmiiller, M., Teichmiiller, R., Mullis, J., Kiinzi, B., Breitschmid, A,, Gruner, U. and Schwizer, B., 1980. Very low-grade metamorphism in the external parts of the Central Alps: Illite crystallinity, coal rank and fluid inclusion data. Eclogue Geol. Helo., 73(1): 173-203. [Four cross-sections were studied. Illite crystallinity and coal reflectance generally increase from tectonically higher to lower units, and from external to internal parts in the same tectonic unit. General evolution of fluid composition in inclusions with metamorphic grade. Two cases of thrusting of higher-grade upon lower-grade units are mentioned.] Kisch, H.J., 1980. Illite crystallinity and coal rank associated with lowest-grade metamorphism of the Taveyanne greywacke in the Helvetic zone of the Swiss Alps. Eclogue Geol. Helv., 73(3): 753-777. [Illite crystallinities associated with laumontite-bearing and laumontite-free, prehnite- and pumpellyite-bearing Taveyanne greywackes are respectively “diagenetic” and middle- to high-grade anchimetamorphic; the associated mean vitrinite reflectances are respectively 0.85- 1.3% and 3.3-4.2% R max ”,,. The onset of anchimetamorphism seems to be approximately in the coal-rank range 2.3-3.358 R,,,,,,, (semi-anthracite to anthracite).] Kiibler, B., Pittion, J.-L., Heroux, Y . , Charollais, J. and Weidmann, M., 1979. Sur le pouvoir reflecteur de la vitrinite dans quelques roches du Jura, de la molasse et des nappes prealpines. helvetiques et penniniques (Suisse occidentale et Haute-Savoie). Eclogae Geol. Helv., 72(2): 347-373. [Vitrinite reflectance data of various tectonic units in NW-SE sections in the southern part of Haute-Savoie (including the ThBnes syncline and the Plate massif-cf. Aprahamian et al.. 1975, ref. p. 521), the Lake of Geneva, the Chablais and Romande Prealps, and the Valais RhBne. In part of these sections information is available on illite crystallinity or clay-mineral diagenesis.] Stalder, P.J., 1979. Organic and inorganic metamorphism in the Taveyannaz Sandstone of the Swiss Alps and equivalent sandstones in France and Italy. J. Sediment. Petrol., 49(2): 463-482. [Diagnostic zeolite-facies and prehnite-pumpellyite facies assemblages, and the associated coal ranks and illite crystallinities show a systematic temperature relationship. The onset of prehnite-pumpellyite facies correlates with high-grade “diagenetic” crystallinities and anthracite rank. Temperatures based on coal rank given (assuming some 5-10 m y . effective heating time). The lowest-grade metamorphism is correlated with the Lepontine event (approx. 38 m.y.).] Teichmiiller, M. and Teichmiiller, R., 1978. Coalification studies in the Alps. In: H. Closs, D. Roeder and K. Schmidt (Editors), Alps, Appennines, Hellenides. Schweizerbart, Stuttgart, pp. 49-55. [Coalification in the Helvetic nappe complex postdates its internal structure but predates the emplacement of the complex upon the molasse (similar to time relations based on illite crystallinity studies at the Glarnisch-Frey et al., 1973). (cf. Eggert, P., Grebe H., Teichmiiller, M. and Teichmiiller, R., 1976. Inkohlungsuntersuchungen an Treibholz aus den Unteren Junghansen-Serie (Unterkreide) der Feuerstatter Decke (Nordpenninikum) westlich Oberstdorf/Allgau. Neues Jahrb. Geol. Palaontol. Abh., 152( 1): 112-1 36).] Zingg, A., Hunziker, J.C., Frey, M. and Ahrendt, H., 1976. Age and degree of metamorphism of the Cavanese Zone and of the sedimentary cover of the Sesia Zone. Schweiz Mineral. Petrogr. Mitt., 56: 361-375. [The occurrence of meta-anthracite to meta-graphite (Stadler, Teichmiiller, and Teichmiiller. 1976) in association with illite at the anchizone-“diagenesis” boundary in the Tertiary cover of the Sesia zone could be due to contact effects of the overlying andesite flow.] Illite crystallinity and chemical composition of dioctahedral mica in relation to development of metamorphic fabric and of second (crenulation) cleavage

Gray, D.R., 1977. Differentiation associated with discrete crenulation cleavages. Lithos, 10: 89- 101. [Analyses of muscovites in coarse-grained muscovite schist from Broken Hill show little or no

5 24 chemical variation between those in the original schistosity and those along the discrete crenulation cleavages.] Knipe, R.J., 1981. The interaction of deformation and metamorphism in slates. Tectonoph.jsics, 78: 249-272. [Differences in composition of phyllosilicates in the oriented phyllosilicate-rich and the disoriented quartz-rich domains in developing crenulation cleavage: the former domains contain a more phengitic mica and Fe-poor chlorite.] Liewig, N., Caron, J.-M., and Clauer, N., 1981. Geochemical and K-Ar isotopic behaviour of Alpine sheet silicates during polyphased deformation. Tectonophysics, 78: 273-290. [Si-tetrahedral substitution in the phengites depends on their microstructural position: it is about 3.30 in the S , schistosity, and 3.37-3.46 in the deformed pre-existing mica lamellae from crenulation zones and neoformed phengites. Scatter of the apparent K-Ar ages of the phengites (38-65 my.) could be due to deformation-dependent isotopic behaviour at temperatures close to the blocking temperature.] Stephens, M.B., Glasson, M.J. and Keays, R.R., 1979. Structural and chemical aspects of metamorphic layering development in metasediments from Clunes, Australia. Am. J . Sci., 279(2): 129- 160. [Metasediments suffered only one deformation event and low-grade metamorphism. New phengites in the white mica chlorite ( P ) and quartz ( Q ) layers defining the metamorphc layering are similar in composition. They are richer in Si, Fe, and Mg compared to the detrital micas, and grew in equilibrium with the ambient pore fluid through an orientation-dependent growth mechanism.] Weber, K., 1976. Gefiigeuntersuchungen an transversalgeschieferten Gesteinen aus dem ostlichen Rheinischen Schiefergebirge (Ein Beitrag zur Genese der transversalen Schieferung). Geol. Jahrb. (Hannover), Reihe 0,H. 15: 3-98. [In anchimetamorphic range. At lowest grades no nucleation of phyllosilicates on the first cleavage: purely mechanical alignment. Increasing nucleation of phyllosilicates on the cleavage planes and recrystallization within the uncleaved “cleavage lamellae” with increasing metamorphic grade. The second (crenulation) cleavage is entirely post-crystalline.]

+

Pyrophyllite, rectorite, paragonite in anchimetamorphism (including experimental studies )

Day, H.W., 1976. A working model of some equilibria in the system alumina-silica-water. Am. J . Scr., 276(10): 1254-1284. [Stability of pyrophyllite in diagenesis and very-low-grade metamorphism of sediments, with application to natural assemblages.] Eberl, D., 1979. Synthesis of pyrophyllite polytypes and mixed-layers. Am. Mineral., 64(9- 10): 1091- 1096. Frey, M., 1978. Progressive low-grade metamorphism of a black shale formatoin, central Swiss Alps, with special reference to pyrophyllite and margarite bearing assemblages. J . Petrol., 19(1): 95- 135. [Distinction between muscovite, paragonite, and margarite in the 45’-48’ 2 8 range of diffractometer traces. In the anchizone, pyrophyllite formed at the expense of kaolinite; mixed-layer paragonitemuscovite presumably from mixed-layer illite-montmorillonite. The assemblages are treated in the High XcH, and low two subsystems MgO (or Fe0)-Na,0-Ca0-AI,0,-’(KA1305-Si0,-H,0-C0,). in the anchizone.] Gomez-Pugnaire, M., Sassi, F.P. and Visona, D., 1978. Sobre la presencia de paragonite y pyrofilita en las filitas del complejo Nevado-Filabride en la Sierra de Baza (Cordilleras Beticas, Espana). Bol. Geol. Min. (Spain), 89(5): 468-474. Tomita, K., 1977. Experimental transformation of 2M sericite into a rectorite-type mixed-layer mineral by treatment with various salts. Clays Clay Miner., 25: 302-308. [Rectorite-like mixed-layer formed when dehydroxylated 2M sericite treated with solutions of Na. Ca or Mg salts. Random mica-montmorillonite mixed-layer is formed from 2M sericite.] Also: - Aparicio and G a l h , 1980 (ref. p. 5 19). - Brime and Perez-Estaun, 1980 (ref. p. 519). - Frey et al., 1976 (ref. p. 522), - Frey and Wieland, 1975 (ref. p. 522). - Galan et al., 1978 (ref. p. 519).

525

- Gill et al.,

1977 (ref. p. 520).

- Kisch. 1980b (ref. p. 539) Lecolle and Roger, 1976 (ref. p. 540). Robinson et al., 1980 (ref. p. 520). - Schramm, 1978 (ref. p. 522). - Schramm, 1982b (ref. p. 522). - Wieland, 1979 (ref. p. 522). -

-

Change in chemical composition of potassic white micas with grade in lowest-grade metamorphism

McDowell, S.D. and Elders, W.A., 1980. Authigenic layer silicate minerals in borehole Elmore I , Salton Sea Geothermal Field, California, USA. Contrib. Mineral. Petrol., 74: 293-3 10. [“Illite” ( = textural sericite) is free of expandable layers below 725 m (275°C). Change towards recrystallized phengitic white mica below 850 m (ca. 290°C) involves more muscovitic compositions with increasing temperature. In the same interval chlorite shows an increase in total Mg + Fe.] Also: - Dunoyer de Segonzac and Abbas, 1976 (ref. p. 522). - Frey, 1978 (ref. p. 524). - Timofeev et al., 1974 (ref. p. 513). Use of the cell parameter b(]of potassic white micas as a porameter of P / T gradients!of metamorphism

Arkai, P., 1977. Low-grade metamorphism of Paleozoic sedimentary formations of the Szendro Mountains (NE-Hungary). Actu. Geol. Acad. Sci. Hung., 21(1-3): 53-80. [Use of b,-do,, diagrams for rocks of greenschist facies and adjoining zone of low and very-low-grade metamorphism (“anchi-epi-zones”); low- to medium-pressure metamorphism ( b , 9.003 A).] Fettes, D.J., Graham, C.M., Sassi, F.P. and Scolari, A,, 1976. The lateral spacing of potassic white micas and facies variation across the Caledonides. Scott. J . Geol., 12(3): 227-236. [Almost 200 samples from five areas in the lowest-temperature zone of the Scottish Caledonides. A decrease in the b, values demonstrates a gradual transition in metamorphic facies series from the SW-Highlands ( b , - 9.017 A), across the area of “Barrovian” metamorphism in the central Highlands, into the area of “Buchan” metamorphism ( 6 , 8.992 A).] Guidotti. C.V. and Sassi, F.P.. 1976. Muscovite as a petrogenetic indicator mineral in pelitic schists. Neues Jahrb. Mineral. Abh., 127(2): 97- 142. K) ratio and celadonite content of muscovite using d,, and b,. [Determination diagram of Na/(Na Exhaustive discussion of control of muscovite composition by temperature. pressure, and H ,O activity. Importance of considering bulk composition and mineral assemblage. Diagrams for variation of muscovite composition with increasing temperatures at different pressures.] Hartnady, C.J., Antrobus, B. and Spector, D., 1978. Reconnaissance studies of regional metamorphism in the Malmesbury Group and the Name Group of southern Namaqualand. Unio. Cupetown, Dep. Geol., Precambrian Res. Unit Annu. Rep., 14-15: 204-207. [Different b, values are correlated with different ages of porphyroblastesis. Lower-than-greenschist facies low to intermediate pressure (b, 9.004 or 9.007 A) Malmesbury metamorphism appears to overprint an earlier Barrovian intermediate-high pressure metamorphism ( h , 9.033 A).] Kisch. H.J. and Padan, A,, 1981. Use of the lattice parameter h, of dioctahedral illite/muscovite for the characterization of the P-T gradient of incipient metamorphism in the Caledonides of Jamtland. western Sweden. Terra Cognitu. I ( I ) ; 54-55. [The mean b, values for the illite crystallinity zones established earlier (Kisch, 1980a) are lower for the partly “diagenetic” zone A than for the three higher-grade zones (mean h, = 9.032 A). It is shown that the method can be used in the anchizone: the P / T gradient found agrees with the intermediate-pressure Barrovian series found further W.] KrButner, H.G.. Sassi. F.P., Zirpoli, G . and Zulian, T.. 1976. Barrovian-type Hercynian metamorphism from the Poiana Rusca Massif (South Carpathians). N e w s Jahrb. Mineral. Monatsh., 1976( 10): 446-45 5.

-

-

+

-

-

526 [More than half the samples are from the chlorite zone. The intermediate-pressure (“Barrovian”) type metamorphism indicated by the b, values (mean b, = 9.021 A) contrasts with the general low-pressure character of Hercynian metamorphism in Europe.] Padan, A,, Kisch, H1. and Shagam, R., 1982. Use of the lattice parameter b, of dioctahedral illite/muscovite for the characterization of P / T gradients of incipient metamorphism. Contrib. Mineral. Petrol., 79( 1): 85-95. [ b, curves for different incipient-metamorphic zones in marginal zones of the Swedish Caledonides (see Kisch and Padan, 1981, above), Swiss Alps, and Venezoelan Andes. b, tends to increase with grade during incipient metamorphism. Distinct differences between the P / T gradients for the for the Venozoelan Andes indicates a much lower different terranes are found: mean b, = 9.005 P / T gradient than for the other two terranes.] Robinson, D., 1981. Metamorphic rocks of an intermediate facies series juxtaposed at the Start boundary, southwest England. Geol. Mug., 118(3): 297-301. [Devonian phyllites of low-intermediate pressure facies series (mean b, = 9.002 A) have been juxtaposed, N of the Start boundary, against the high-intermediate pressure type (mean b, = 9.032 A without paragonite) of the Start schists.] Sassi, F.P., Krautner, H.G. and Zirpoli, G., 1976. Recognition of the pressure character in greenschist facies metamorphism. Schweir. Mineral. Petrogr. Mitt., 56: 427-433. [On the basis of approx. 2000 b, values of white micas from the low-grade part of the greenschist facies the baric type of several metamorphic terranes is given; a range of baric types is recognized within the field of intermediate-pressure (‘Barrovian’) metamorphism.] Seidel, E., 1977. Lawsonite-bearing metasediments in the phyllite-quartzite series of S ~ - C r e t e(Greece). Neues Jahrb. Mineral. Abh., 130(1-2): 134-144. [The comparatively low celadonite contents of the white micas from lawsonite schists (about 20%) compared to other HP/LT rocks is related to the Ca-rich and Fe-poor bulk composition.] Zhang, Q., Zhang, Z., and Li, S., 1980. Muscovite of Ppetamorphic’rocks in east and south Xizang and its petrological significance (Chinese with English abstr.). Sci. Geol. Sinicu, 340-347 (Mineral. Abstr., 32: 81-3128). [b,, d,,,, and MgO content of nine muscovites vary with metamorphic pressure. The high-P area that can be distinguished on diagrams of MgO vs RM and Si vs Mg of the muscovite-phengite series of the greenschist-blueschist facies, can be subdibided in a glaucophane-bearing and a glaucophane-free field.] Zhang Zhaozhong, Zhang Bingliang, Feng Jinjiang, and Li Songbin, 1981. b, values of muscovites and metamorphic belts of Dabie Mouhtains metamorphic terrains. Kexue Tongbao Sci. Bull. (English transl. from K’o-Hsueh T’ung-Pao), 26(4): 341-345. [The metamorphic facies series of three metamorphic belts were determined on the basis of b, values. Under similar conditions b, values of 3T-phengites tend to be higher than those of 2M-phengites.]

A

Corrensite and other chloritic mixed layers in incipient metamorphism

Lippman, F. and Rothfuss, H., 1980. Tonminerale in Taveyannaz-Sandsteinen. Schweiz Mineral. Petrogr. Mitt.,60: 1-29. [In the presence of laumontite, corrensite is a major clay mineral; some of the samples contain “para-corrensite” with reduced expandability.] Suchecki, R.K., P e w , E.A., and Hubert, J.F., 1977. Clay petrology of Cambro-Ordovician continental margin, Cow Head Khppe, western Newfoundland. Clays Clay Miner., 25: 163-170. [Mg-rich volcanic detritus and its alteration products in the Lower and Middle Ordovician reacted during burial metamorphism to form illite-smectite with 5- 10% expandable layers plus corrensite or expandable chlorite.] Velde, B., 1977. A proposed phase diagram for illite, expanding chlorite, corrensite and illite-montmorillonite mixed-layered minerals. Clays Clay Miner.. 25: 264-270. [Based on experimental data on hydrothermally treated natural clay minerals. lmportance of R3’ content of the assemblage and of P-T-X variables as controls for appearance of expanding chlorite or corrensite in earliest metamorphism.]

527 Velde, B.. 1977. Clays and Clay Minerals in Natural and Synthetic Systems. Elsevier, Amsterdam, 218 pp. Chapter 6-Chlorites, and 7-Corrensite. Velde, B., Proust, D. and Meunier, A,, 1979. Chlorite compositions during sedimentation. Sci. Geol., Mem., No. 53: 71-73. Zingg, A,, Hunziker, J.C., Frey, M. and Ahrendt, H., 1976. Age and degree of metamorphism of the Cavanese Zone and of the sedimentary cover of the Sesia Zone. Schweiz. Mineral. Petrogr. Mitt.. 56: 361-375. [Occurrence of regular chlorite-montmorillonite mixed-layer in a shear zone and in a massive talc-bearing metadolomite in the “epizone” of the Cavanese Zone between Biella and Valle d’Ossola.] Also: - Aprahamian et al., 1975 (ref. p. 521). - Chudaev, 1978 (ref. p. 517). - Dunoyer de Segonzac and Bernoulli, 1976 (ref. p. 522). - Kisch, 1981 (ref. p. 517). Change in chemical composition and crystallinity of trioctahedral chlorites with grade of lowest-grade metamorphism

Deutloff et al., 1980 (ref. p. 518). McDowell and Elders, 1980 (ref. p. 524). Resetting of K-Ar, Rb-Sr, and U-Pb ages in incipient metamorphism

Ahrendt, H., Hunziker, J.C. and Weber, K., 1977. Age and degree of metamorphism and time of nappe emplacement along the southern margin of the Damara Orogen/Namibia (SW-Africa). Geol. Rundsch., 66(2): 719-742. [K/Ar ages of white mica from the basement are around 1160 my.; those from the anchimetamorphic Naukluft nappes and the underlying Nama beds adjoining to the SE define two isochrons with ages 495 and 530 m.y. The latter represents the peak of the anchimetamorphism, the former indicates the emplacement of the Naukluft nappes.] Ahrendt, H., Hunziker, J.C. and Weber, K., 1978. K/Ar-Alterbestimmungen an schwachmetamorphen Gesteinen des Rheinischen Schiefergebirges. Z . Dtsch. Geol. Ges., 129: 229-247. [The K/Ar ages of anchimetamorphic micas are not cooling ages, but date the peak of metamorphism; they range from 300 m y . in the N E Rheinisches Schiefergebirge to 315 m y . in the S, and up to 330 m y . in the Taunus. The somewhat younger ages of around 310 m y . from the somewhat higher grade rocks of the “Taunus Pre-Devonian” to the S are interpreted as cooling ages.] Bonhomme, M.G., Saliot, P. and Pinault, Y . , 1980. Interpretation of potassium-argon isotopic data related to metamorphic events in south-western Alps. Schweiz. Mineral. Petrogr. Mitt., 60: 81 -98. [The non-metamorphic Rhetian suffered a late diagenesis at 155 m y . The apparent K-Ar ages of the fine fractions decrease towards the E with increasing metamorphic grade as shown by the increasing crystallinity of illites. Samples from the anchizone show mixed ages between 108 and 72 my.; those from the low-grade epizone between 95 and 38 m.y. Importance of mica chemistry is discussed.] Clauer, N. and Kroner, A., 1979. Strontium and argon isotopic homogenization of pelitic sediments during low-grade regional metamorphism: the Pan-African upper Damara sequence of northern Namibia (South West Africa). Earth Planet. Sci. Lett., 43: 117-131. [Two successive regional events of anchizonal intensity dated at about 535 and 455 my., respectively. Anomalously high K-Ar ages from some stratigraphic horizons can be related to open system behaviour and K migration during formation of stilpnomelane from ferromagnesian illites.] Gebauer, D. and Griinenfelder, M., 1977. U-Pb systematics of detrital zircon from some unmetamorphosed to slightly metamorphosed sediments of Central Europe. Contrib. Mineral. Petrol.. 65: 29-37. [Strong discordancy of pre-Assyntic zircon population from the Algonkian of Bohemia is tentatively ascribed to recrystallization and lead loss during the Assyntic very-low-grade metamorphism (“zeolite facies”) at temperatures as low as 300°C.]

528 Hoffman, A.W., Mahoney, J.W. and Giletti, B.J., 1974. K-Ar and Rb-Sr data on detrital and postdepositional history of Pennsylvanian clay from Ohio and Pennsylvania. Geol. Soc. Am. Bull. 85: 639-644. [No systematic relation between coal rank and the fairly uniform whole-rock K-Ar ages (355-383 m y . in five samples) is apparent. The decreasing K-Ar and Rb-Sr ages (510 to 320 m.y.) with decreasing grain size fraction is ascribed to crystallization or reconstitution of IMd-type illite.] Hoffman, J., Hower, J. and Aronson, J.L., 1976. Radiometric dating of time of thrusting in the disturbed belt of Montana. Geology, 4( I): 16-20. [Burial metamorphism below thrust plates. Bentonite has been transformed to K-bentonite giving a K-Ar age of 72 to 56 m y . ; associated with laumontite-bearing volcanic sand and tuff (see also Hoffman and Hower, 1979).] Kroner, A. and Clauer. N.. 1979. Isotopic dating of low-grade shale in northern Namibia (South West Africa) and implications for the orogenic evolution of the Pan-African Damara Belt. Precambrian Res., 10: 59-72. (2 p m fractions of higher-grade (illite-chlorite and stilpnomelane-bearing) assemblages show younger Rb/Sr ages (about 457 my.) than those of a lower-grade (smectite-bearing) assemblage (about 537 my.). These ages are related to two separate low-grade regional tectono-thermal events (see also Clauer and Kroner, 1979, above).] Leitch, E.C. and McDougall, I., 1979. The age of orogenesis in the Nambucca slate belt: a K-Ar study of low-grade regional metamorphic rocks. J . Geol. Soc. Aust.. 26: 1 1 1-1 19. [Prehnite-pumpellyite to greenschist facies metasediments yield a range of ages (some comparable with depositional ages). The more coherent group of K/Ar ages from actinolite-pFmpellyite and greenschist facies rocks is considered to represent orogenesis at 250-255 m.y.1 Odin, G.S., Velde, B. and Bonhomme, M.. 1977. Radiogenic argon in glauconites as a function of mineral recrystallization. Earth Planet. Sci. Lett., 37: 154- 158. [The extent to which incipient metamorphism affects the apparent radiogenic age of glauconites depends on their composition and the temperature at metamorphism; experiments in the range 200-414°C at 2 kbars.] Perry, E.A. and Turekian. K.K., 1974. The effect of diagenesis on the redistribution of strontium isotopes in shales. Geochim. Cosmochim. Acta., 38: 929-935. [Attending the diagenetic changes with depth there is a trend towards the homogenization of the 87Sr/86Sr ratios of the size fractions of the shale, but diagenesis and homogenization are not complete in the deepest part (5523 m) of the Miocene shale section studied. Rb/Sr shale dates in many cases probably represent the time of the major diagenetic construction of new phases.] Alterarion and dissolution of clastic feldspar during burial diagenesis; material rransfer of mineral-mineral and water-rock equilibrium

Land, L.S. and Milliken, K.L., 1981. Feldspar diagenesis in the Frio Formation, Brazoria County. Texas Gulf Coast. Geology, 9(7): 314-318. [The material transfer involved in the dissolution of albitization of detrital feldspars below about 4000m affects at least 15% of the rock volume. Important implications for several other diagenetic processes such as precipitation of cements, and evolution of formation waters.] Milliken, K.L., Land, L.S. and Loucks, R.G., 1981. History of burial diagenesis determined from isotopic geochemistry, Frio Formation, Brazoria County, Texas. Am. Assoc. Pet. Geol. Bull.. 65(8): 1397- 1413. [C and 0 isotopic data. Temperatures of formation of quartz cement. kaolinite. albitization. Extensive shift towards water-rock equilibrium. Organic maturation, albitization and the smectite to illite transformation contribute most of the constituents required for the precipitation of the cements. Yeh, H.-W. and Savin, S.M., 1977. Mechanism of burial metamorphism of argillaceous sediments: 3, 0-isotope evidence. Geol. Soc. Am. Bull.. 88: 1321- 1330. [0-isotope disequilibrium among clay fractions less marked as burial T increases, but persists even at burial to 170°C. Temperatures calculated from 0-isotope fractionations between fine-grained quartz and clay approach measured well temperatures as depth of burial and temperature increased. but agreement was found absent at measured temperatures as high as 120'C.J

529 Also: - Eslinger and Sellars, 1981 (ref. p. 515). - Heling, 1978 (ref. p. 515). C-isotope ratios in carbonaceous matter in incipient metamorphism

Hoefs, J. and Frey, M., 1976. The isotopic composition of carbonaceous matter in a metamorphic profile from the Swiss Alps. Geochim. Cosmochim. A d a , 40: 945-951. [ SI3C values are around -25%0 in the unmetamorphosed and anchimetamorphic (Glarus Alps) sediments, but shift to higher I3C content with increasing grade of metamorphism above the chloritoid isograd. 6I3C values of around - 11%0were measured in rocks of the highest metamorphic grade (staurolite schists).] Progressive ordering of silica during burial metamorphism

Mizutani, S., 1977. Progressive ordering of cristobalite silica in the early stage of diagenesis. Contrib. Mineral. Petrol,, 61: 129-140. [The ordering of opal-CT crystals with time is reflected by an decrease in the d,,, spacing of cristobalite. The isopleths of d , , , spacings should usually be parallel with stratigraphic boundaries, but should be discordant where the strata have been folded. This discordancy is chiefly controlled by thermal history during burial and folding.] Murata, K.J., Friedman, L. and Gleason, J.D., 1977. Oxygen isotope relations bktween diagenetic silica minerals in Monterey Shale, Temblor Range, California. Am. J. Sci., 277(3): 259-272. [The isotopic temperatures in the silica phases increase from the diagenetic opal (1) through cristobalite (2) to the microquartz (3) zone, hut remain fairly constant within each zone. The progressive structural ordering of cristabalite of virtually constant 0-isotopic composition within zone 2 seems to he a solid-state reaction.] Murata, K.J. and Norman, M.B., 1976. An index of crystallinity for quartz. Am. J . Sci., 276: 1120- 1130. [Is largely a function of crystallite size (up to 1 pm), but may be affected by lattice distortions due to mechanical stress.] Pisciotto, K.A., 1981. Diagenetic trends in the siliceous facies of the Monterey Shale in the Santa Maria region, California. Sedimentology, 28: 547-571. [Three zones as in Murata et al. (1977; see above). Ranges in temperatures for top and base of the opal-CT zone from present geothermal gradients and reconstructed burial depths, and from 0-isotopic compositions of opal-CT and quartz.] Zeolite facies in general (including relations with diagenetic clay mineralogy and coal rank)

Aoyagi, K. and Kazama, T., 1980. Transformational changes of zeolites and clay minerals during diagenesis. Sedimentology, 27: 179- 188. [Temperatures for montmorillonite -,montmorillonite-illite mixed-layer (- 100°C), montmorilloniteillite mixed-layer illite ( 140OC). clinoptilolite heulandite and/or analcime (120'). heulandite and/or analcime -, laumontite and/or albite (140'C). Based on these transformations, seven mineral zones are recognized in argillaceous sediments. Relation to compaction stages.] Boles, J.R., 1977. Zeolites in low-grade metamorphic rocks. In: F.A. Mumpton (Editor), Mineralogy and Geology of Natural Zeolites (Mineral. SOC.Am., Short Course Notes, Vol. 4). Southern Printing Comp., Blacksburg, Va., pp. 103-135. Boles, J.R., 1977. Zeolites in deep-sea sediments. In: F.A. Mumpton (Editor), Mineralogy and Geology of Natural Zeolites (Mineral. SOC.Am., Short Course Notes, Vol. 4). Southern Printing Comp., Blacksburg, Va., pp. 137-163. Ghent, E.D., 1979. Problems in zeolite facies geothermometry, geobarometry and fluid composition. In: P.A. Scholle and P.R. Schluger (Editors), Aspects of Diagenesis. Soc. Econ. Paleontol. Mineral., Spec. Publ., no. 26: 81-87. [Complications of estimating P9, T and fluid compositions from correlation of mineral assemblages

-.

-

5 30 from experimental and computed phase equilibria (including P H t o < P,, aSio,. porosity and permeability). Correlation with coal rank and clay mineral assemblages in any one area will lead to the best estimates of P,. T and fluid composition.] Hay, R.L., 1977. Geology of zeolites in sedimentary rocks. In: F.A. Mumpton (Editor). Mineralogy and Geologv of Natural Zeolites (Mineral. SOC.Am., Short Course Notes, Vol. 4). Southern Printing Comp., Blacksburg, Va., pp. 53-64. Hay, R.L., 1978. Geologic occurrence of zeolites, In: L.B. Sand and F.A. Mumpton (Editors), Natural Zeolites-Occurrence, Properties, Use. Pergamon, Oxford, pp. 135- 143. Iijima, A,, 1978. Geological occurrences of zeolite in marine environments. In: L.B. Sand, and F.A. Mumpton (Editors), Natural Zeolites-Occurrence, Properties, Use. Pergamon, Oxford, pp. 175- 198. Kisch, H.J., 1982. Coal rank and illite crystallinity associated with the zeolite facies of Southland and the pumpellyite-bearing facies of Otago, southern New Zealand. N . Z . J . Geol. Geophys., 24(3): 349-360. [High- and medium-volatile bituminous ranks (0.60- 1.33% R ol, ) and expandable mixed-layers are associated with the zeolite facies of the North Range Group, Southland Syncline. Only illite is associated with the laumontite-bearing area in the Torlesse terrane.] McCulloh, T.H., Cashman, S.M. and Stewart, R.J., 1979. Diagenetic baselines for interpretive reconstructions of maximum burial depths and paleotemperatues in clastic sedimentary rocks. In: D.F. Oltz (Editor), A Symposium in Geochemisty: Low Temperature Metamorphism of Kerogen and 0la.v Minerals. SOC.Econ. Paleont. Mineral., Pac. Sec., Los Angeles, Calif., pp. 65-96. [Relationship between laumontite zone and coal rank in some California sedimentary basins.] Shimoyama, T. and Iijima, A., 1976. Influence of temperature on coalification of Teriary coal in Japan-Summary. In: Circum-Pacific Energy and Mineral Resources. Am. Assoc. Pet. Gdol., Mem., 25: 98- 103. [Zoning of zeolites replacing felsic glass in vitric tuffs in the coal measures. Lignite and subbituminous coal in mordenite-clinoptililite zone; coking bituminous coal ( R,,,> 0.6%) exclusively in analcime zone. Bottomhole temperatures at the base of these zones are 85"-9OoC and 120°-1250C, respectively.] Stability of zeolites: experimental (exclusive of high-grade boundary of zeolite facies)

Arima, M. and Edgar, A.D., 1980. Importance of time and H,O contents on the analcime-H,O system at 465°C and 1 kbar P H Z 0 .Neues Jahrb. Mineral. Monatsh., 1980(5): 543-554. [In runs of up to 50 days continuing increase in the amount of albite and progressively less siliceous analcime, when no excess water is present; no change after 20 days when 10, 20 or 252 H,O are present. Most previous studies have been done with excess H,O: stability relations should be used with caution.] DeKimpe, C., 1976. Formation of phyllosilicates and zeolites from pure silica-aluminium gels. Clays Clay Miner., 24: 200-207. [Formation in presence of NaOH solution. Zeolites formed at low gel/solution ratios: kaolinite produced at less alkaline pH and large gel/solution ratios.] Goto, Y . , 1977. Synthesis of clinoptilolite. Am. Mineral., 62: 330-332. [At 200°C at pH 7.9, in presence of K as well as Na in starting materials.] Hawkins, D.B., Sheppard, R.A. and Gude, A.J.. 1978. Hydrothermal synthesis of clinoptilolite and comments on the assemblage phillipsite-clinoptilolite-mordenite. In: L.B. Sand and F.A. Mumpton (Editors), Natural Zeolites-Occurrence, Properties. Use. Pergamon, Oxford, pp. 145- 174. Kim, M.-T. and Burley, B.J., 1980. A further study of analcime solid solutions in the system NaAISi,O,NaAISi0,-H,O, with particular note of an analcime phase transformation. Mineral. Mag.. 43(332): 1035-1045. [Investigation of the variations of the room-temperature cell parameters of analcime as a function of the temperature of synthesis and of composition; most solid solutions encountered are equilibrium compositions.] Velde. B., 1977. Clays and Clay Minerals in Natural and Synthetic Systems. Elsevier, Amsterdam. 218 pp. Chapter 8-Zeolites (pp. 116- 140).

53 1 Also: - Ghent, 1979 (ref. p. 529). - Iijima, 1975 (ref. p. 532). Formation of alkali zeoliies ai low temperatures and shallow depth in silicic volcanic rocks

Boles, J.R. and Surdam, R.C., 1979. Diagenesis of volcanogenic sediments in a Tertiary saline lake; Wagon Bed Formation, Wyoming. Am. J. Scr., 279(7): 832-853. [Three diagenetic facies formed after burial as a result of different pore fluid compositions inherited from the different depositional environments. Diagenetic reactions took place in moderately saline, but not highly alkaline pore fluids.] Dibble, W.E. and Tiller, W.A., 1981. Kinetic model of zeolite paragenesis in tuffaceous sediments. Clays Clay Miner., 29(5): 323-330. [Kinetic factors may determine the specific authigenic phases. Sequence of assemblages formed during series of metastable reactions resembling Oswald step rule. Explanation for occurrence of metastable reactions.] Ratterman, N.G. and Surdam, R.C., 1981. Zeolite mineral reactions in a tuff in the Laney Member of the Green River Formation, Wyoming. Clays Clay Miner., 29(5): 365-377. [Two successive diagenetic stages. The second produces analcime from early zeolites + Na-carbonate brine, and involves significant mass transfer.] Surdam, R.C., 1977. Zeolites in closed hydrologic systems. In: F.A. Mumpton (Editor), Mineralogy and Geology of Natural Zeolites (Mineral. Soc. Am., Short Notes, Vol. 4). Southern Printing Comp., Blacksburg, Va., pp. 65-91. Surdam, R.C. and Sheppard, R.A., 1978. Zeolites in saline, alkaline-lake deposits. In: L.B. Sand and F.A. Mumpton (Editors), Natural Zeolites-Occurrence, Properties, Use. Pergamon, Oxford, pp. 145- 174. Taylor, M. and Surdam, R.C., 1981. Zeolite reactions in the tuffaceous sediments at Teels Marsh, Nevada. Clays Clay Miner., 29(5): 341-352. [Hydratation of Holocene rhyolitic glass, mostly to phillipsite; also analcime and clinoptilolite. Si concentration is controlled by authigenic reactions at less than 100 p.p.m.1 Van, A.V. and Kolodezhnikov, K.E., 1979. Mineralogical types of tuff in Middle Paleozoic deposits in the west of the Vilyui Syneclise. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), l4( 1): 79-89. [Deeper analcime zone and shallower heulandite zone are ascribed respectively to a lagoonal-saline and a fresh-water environment; the formation of analcime in the deeper zones was enhanced by subsequent "regional epigenesis".] Walton, W.A.. 1975. Zeolitic diagenesis in Oligocene volcanic sediments, Trans-Pecos Texas. Geol. Soc. Am. Bull., 86: 615-624. [Montmorillonite, clinoptilolite, and analcime formed during diagenesis in an open hydrologic system at depths of not more than a few hundred meters. Distribution of clinoptilolite was controlled locally by permeability of the hos! rocks.] Low-temperature formation of laumontiie

McCulloh, T.H., Frizzel, V.A., Stewart, R.J. and Barnes, I., 1981. Precipitation of laumontite with quartz, thenardite, and gypsum at Sespe Hot Springs, western Transverse Ranges, California. Clays C l q Miner., 29(5): 353-364. [Laumontite precipitates at 89" to 43"; no other zeolites were observed. Little or no carbonate minerals. The subsurface water source is thought to have a temperature of 125°-1350C.] Sands, C.D. and Drever, J.I., 1978. Authigentic laumontite in deep-sea sediments. In: L.B. Sand and F.A. Mumpton (Editors), Natural Zeolites-Occurrence, Properties, Use. Pergamon. Oxford-New York, pp. 269-279. [Associated with major clinoptilolite. 0-isotope data indicate maximum temperature of 60"C.l Also: - Barnes et al., 1978 (ref. p. 534).

532 Effect of pore-water chemistry and of AI/Si and C a / N a ratios of parent material on depth zoning of diagnostic zeolites

Boles, J.R. and Coombs, D.S., 1977. Zeolite facies alteration of sandstones in the Southland Syncline, New Zealand. Am. J. Sci., 277: 982-1012. [Individual mineral ranges in the 10.4 km sequence overlap even more than previously described. Effect of host rocks. Evidence of mass transfer on macroscopic and sometimes larger scale. Complexity of mineral distribution patterns is attributed to the effects of parent materials. permeability, ionic activity ratios in stratal waters, relationship of Pnutdto P,,,,,.] Davies, D.K., Almon, W.R., Bonis, S.B. and Hunter, B.E., 1979. Deposition and diagenesis of TertiaryHolocene volcaniclastics, Guatemala. In: P.A. Scholle and P.R. Schluger (Editors), Aspects of Diagenesis. Soc. Econ. Paleontol. Mineral., Spec. Publ., 26: 281 -306. [Boundaries between different diagenetic assemblages (montmorillonite-goethite; montmorillonite plus hematite; montmorillonite plus heulandite) are determined more by groundwater chemistry than by T o r P . ] Hay, R.L. and Sheppard, R.A., 1977. Zeolites in open hydrologic systems. In: F.A. Mumpton (Editor), Mineralogy and Geology of Natural Zeolites (Mineral. SOC.Am., Short Course Notes, Vol. 4). Southern Printing Comp., Blacksburg, Va., pp. 93- 102. Iijima, A,, 1975. Effect of pore water on clinoptilolite-analcime-albite reaction series. J . Far. Sci.. Unio. Tokyo, See. II, 19(2): 133-147. [The concentration of Na+ in pore water plays an important role in lowerin5 the equilibrium temperatures of the reaction series.] Moncure, G.K., Surdam, R.C. and McKague, H.L., 1981. Zeolite diagenesis below Pahute Mesa, Nevada test site. Clays Clay Miner, 29(5): 385-396. [Three vertical zones, caused by ( I ) changing pore-water chemistry in an essentially closed hydrologic system; (2) disequilibrium or kinetic precipitation of metastable phases; and (3) a higher thermal gradient than now present.] Surdam, R.C. and Boles, J.R., 1979. Diagenesis of volcanic sandstones. In: P.A. Scholle and P.R. Schluger (Editors), Aspects of Diagenesis. Soc. Econ. Paleontol. Mineral., Spec. Publ.. 26: 227-272. [Temperature effects have been overestimated. Broad overlap of individual mineral ranges cannot be explained by differences in geothermal regime between areas. Importance of fluid phase and ionic species in the fluid phase in diagenetic reactions: chemical or ionic stability. Significance of fluid flow and composition in controlling distribution of diagenetic mineral phases.] Wirsching, U., 1981. Experiments on the hydrothermal formation of calcium zeolites. Cla.ys Clay Miner., 29(3): 171-183. [From basaltic and rhyolitic glass, nepheline. and oligoclase, and CaCI, and CaCl + NaOH solutions at 10O"-25O0C. Importance Si/AI and Ca/alkali ratio of the starting materials, of the Ca activity of the reacting solution, presence of an open alteration system, and T , for the zeolite formed.] Phase equilibria of prehnite, pumpellyite, Iawsonite, epidote: experimental and thermodynamical

Brown, E.H., 1977. Phase equilibria among pumpellyite. lawsonite, epidote and associated minerals in low grade metamorphic rocks. Contrib. Mineral. Petrol., 64: 123- 136. [Phase relations analysed on Al-Ca-Fe'+ diagram in which all minerals are projected from quartz. albite or jadeite, chlorite and fluid. This procedure reveals several reactions relating rocks formed at different P-T conditions in the blueschist, greenschist and pumpellyite-actinolite facies.] Frost, B.R., 1980. Observations on the boundary between zeolite facies and prehnite-pumpellyite facies. Contrib. Mineral. Petrol., 73: 365-373. [Graphical analysis of system CaO-AI,O,-SiO,-H ,O-CO,. First appearance of the assemblage epidote-chlorite-quartz ('-albite) should mark the upper boundary of the zeolite facies. This assemblage forms at the expense of laumontite-bearing assemblages. Monitoring composition of minerals from low-variance assemblages may provide a sensitive indicator of metamorphic grade.] Glassley, W., Whetten, J.T., Cowan, D.S. and Vance, J.A., 1976. Significance of coexisting lawsonite, prehnite, and aragonite in the San Juan Islands, Washington. Geology, 4(5): 301-302. [Stability of prehnite could extend to pressures above those of the calcite-aragonite transition.]

533 Nakajima, T., Banno, S. and Suzuki, T., 1977. Reactions leading to the disappearance of pumpellyite in low-grade metamorphic rocks of the Sanbagawa metamorphic belt in central Shikoku, Japan. J . Petrol., 18: 263-284. [The minimum Fe3+ content of epidote can be used to define the metamorphic grade. The temperature range in which the assemblage pumpellyite + epidote chlorite actinolite is stable and is about 90°C in metabasite. The higher temperature limit of the pumpellyite-actinolite facies corresponds to coexistence of epidote with Fe3+/(Fe3+ +Al) = 0.10 0.15 with pumpellyite, actinolite, and chlorite; the lower temperature limit with about 0.33.1 Schiffman, P. and Liou, J.G., 1977. Synthesis and stability relations of Mg-pumpellyite. In: Proc. 2nd Inr. Symp. Water-Rock Interaction, Strasbourg. Cent. Natl. Rech. Sci. Inst. Geol., Strasbourg, pp. 157-164. Schiffman, P. and Liou, J.G., 1980. Synthesis and stability relations of Mg-AI pumpellyite, Ca,Al,MgSi,O,,(OH),. J . Petrol., 21: 44-474. [Stability relations determined using subequal mixtures of synthetic Mg-A1 pumpellyite and its high-temperature assemblage. Schreinemakers' relations for pumpellyite and associated minerals constructed in pseudo-ternary system Ca0-A1,0,-MgO(Si0,-H20). The invariant point ITR( was located at approx. 5.7 Ib Pnuidand 375"C.I Thompson, A.B., 1976. Investigation of lawsonite and prehnite stabilites in natural andesitic rock compositions. In: G.M. Biggar (Editor), Progress in Experimental Petrology; Third Progress Report. Natl. Environment Res. Counc. Publ. Ser. D , No. 6: 1 1 12. [Test of suggestion that lawsonite forms from breakdown of dense Ca-Alihydrosilicates such as prehnite rather than from Ca-zeolites, using the weighed-crystal method.] Also: - Glassley, 1975 (ref. p. 533). - Kuniyoshi and Liou, 1976 (ref. p. 535). - Pluysnina and Ivanov, 1981 (ref. p. 534).

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Effect of substitution of Fe for Mg and A1 on the stability of minerals in the pumpellyite facies

Bird, D.K. and Helgeson, H.C., I98 1. Chemical interaction of aequous solutions with epidote-feldspar mineral assemblages in geologic systems. 11. Equilibrium constraints in metamorphic/geothermal processes. Am. J. Sci., 281(5): 576-614. [Thermodynamic analysis of the system Na,O-K ,O-CaO- FeO- Fe,O,-AI ,O,-SiO,-H ,0-H01CO, at P and T u p to 5 kb and 600OC. Data on low-temperature stability of clinozoisite + quartz.] Coombs, D.S., Kawachi, Y.,Houghton, B.F., Hyden, G. and Pringle, I.J.. 1977. Andradite and andraditegrossular solid solutions in very low-grade regionally metamorphosed rocks in southern New Zealand. Contrib. Mineral. Petrol., 63: 229-246. [May form over a wide range of fo,. but pco, in fluid must be low. Occurs with prehnite and pumpellyite, but not with epidote or Ca-zeolites.] Glassley, W.E., 1975. Low variance phase relationshps in a prehnite-pumpellyite facies terrain. Lithos, 8: 69-76. [Assemblages observed in a basalt member in the Olympic Peninsula are believed to represent stable reaction relationships. Evaluation of P-T-XCo2 conditions is difficult: high iron content of natural phases requires modification of the equilibrium conditions defined in the Fe-free system.] Liou, J.G., 1979. Zeolite facies metamorphism of basaltic rocks from the East Taiwan ophiolite. Am. M i n e d , 64: 1-14. [Zeolite- and pumpellyite-bearing assemblages. Pumpellyites are Fe-rich (up to 25 wt% total Fe as FeO). Substitution of Fe3+ for A1 in pumpellyite enlarges its P-T stability field relative to the zeolite facies assemblages under oxidizing conditions.] Offler, R., Baker, C.K. and Gamble, J., 1981. Pumpellyites in two low-grade metamorphic terranes north of Newcastle, NSW, Australia. Contrib. Mineral. Petrol., 76: 171- 176. [Extreme range in composition. Bulk chemical composition of host rock is not the controlling factor in determining pumpellyite composition. Intensity of alteration (in part of opaque minerals), fluid chemistry, and variation of oxidation potential are more important variables.]

5 34 Pluysnina, L.P. and Ivanov, 1.P.. 1981. Thermodynamic regime of greenstone metamorphism of basic volcanic rocks after experimental data. Can. J . Earth Sci.. 18(8): 1303-1309. [Stability fields of laumontite. prehnite. pumpellyite, zoisite and tremolite bearing assemblages investigated in system Ca0-Mg0-A1,0,-Si0,-C02. Influence of Fe-content on the shift of the upper stability boundary towards both lower T and Xcoz equilibrium values is shown for pumpellyite.] Shimazu, M. and Kusuda, T., 1977. Pumpellyite and prehnite in low-grade metamorphic rocks. Sci. Rep. Niigata Univ., Ser. E . No. 4: 67-81 (Mineral. Abstr., 80-4767). [Fe,03 contents of pumpellyites from the Tanzawa Mountains and the Kita-Akita area seem to be independent of metamorphic grade; they are higher than in most pumpellyites in various other metamorphic terranes.] Tulloch, A.J., 1979. Secondary Ca-AI silicates as low-grade alteration products of granitoid biotite. Conrrib. Mineral. Petrol., 69: 105- 117. (Andradite-grossular, epidote, pumpellyite. and prehnite are extremely common. Correlation of prehnite Fe3+ with host biotite Fe3+ and oxidation state support evidence of prehnite replacing biotite. Plagioclase is the chief source of Ca.] Also: - Brown, 1977 (ref. p. 532). - Frost. 1980 (ref. p. 532). - Nakajima et al., 1977 (ref. p. 533). Ca-Ai-hydrosilicates and C 0 2 in fluid phase

Barnes, I., Downes, C.T. and Hulston, J.R., 1978. Warm springs. South Island, New Zealand, and their potential to yield laumontite. Am. J . Sci., 278(10): 1412-1427. [All the fluids are supersaturated with laumontite, and are generally either in equilibrium with or are unsaturated with calcite and albite. Wide range of CO, partial pressures; these are dependent variables and do not control the chemical reactions. Analcime and prehnite are not necessarily lower-grade and higher-grade facies indicators.] Ghent, E.D. and Miller, B.E., 1974. Zeolite and clay-carbonate assemblages in the Blairmore Group (Cretaceous), southern Alberta Foothills, Canada. Contnb. Mineral. Petrol.. 44: 3 13-329. [Laumontite and barian-strontian heulandite in plagioclase-rich sandstones without kaolinite. Alternative assemblages calcite-kaolinite-quartz and laumontite suggest gradients in /coz/fH20.Computed ionic equilibria suggest among other things that late-formed calcite may not have equilibrated with laumontite. Carbonaceous material corresponds to d , of Landis ( 1971).] Giggenbach, W.F., 198 1. Geothermal mineral equilibria. Geochim. Cosmochirn. Acta, 45: 393-4 10. Ivanov, I.P. and Gurevich, L.P.. 1975. Experimental study of T-XC02 boundaries of metamorphic zeolite facies. Contrib. Mineral. Petrol., 53: 55-60. [Experimental study of the T-Xco2 conditions of the reactions lau = pr mont qz H,O and lau H,O + CO, = cal mont + qz at P,= 1000 bars. Boundaries of the zeolite facies are 200'270'C up to 2500 bars P,, up to 40-60 bars Pco,.] Pearce, T.H. and Birkett, T.C.. 1974. Archean metavolcanic rocks from Thackeray Township, Ontario. Can. Mineraf., 12: 509-519. , ~ pumpellyite-bearing assemblages. Both quartz-chlorite [Effects of p H 2 0 , / p c o 2 and P , / P ~ upon epidote-actinolite-pumpellyite-magnetite-calcite and quartz-prehnite-pumpellyite-epidoteactinolite-chlorite-magnetite are isothermally univariant in pc0,-pH20-Pso,,dsspace. Pluysnina, L.P. and Ivanov, I.P., 198 I . Thermodynamic regime of greenstone metamorphism of basic volcanic rocks after experimental data. Can. J . Earth Sci., 18(8): 1303- 1309.1 [Plot of limits of zeolite, prehnite-pumpellyite, and greenschist facies plotted on schematic T-Xco, diagram; possible Pn limits are discussed. The Xco, equilibrium value for some dehydration-decarbonation reactions decreases for even low salt content of the fluid.] Seki, Y . , 1974. Comparison of CO, and 0, in fluids attending the prehnite-pumpellyite facies metamorphism of the Central Kii Peninsula and the Tanzawa Mountains. Japan. Proc. I n f . Symp. Water-Rock Interaction, Praha. Geol. Survey, Prague, pp. 230-235. [Calcite and calcite-bearing veins are more common in the prehnite-pumpellyite than in the pumpel-

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535 lyite-actinolite facies areas of the Central Kii Peninsula; Xco2 in fluid phases generally increases with decreasing grades of regional metamorphism. In mafic rocks of the same metamorphic rocks in the Tanzawa Mountains calcite is rare: much lower Xco2 in the fluid phase.] Senderov, E.E., 1974. Effect of pH and dissolved carbon dioxide on the replacement of zeolites by clay minerals. Lithol. Miner. Resour. (transl. from Litof. Polern. Iskop.), 9(5): 575-580. [Calculations of the equilibrium constants of analc-kaol-qz, lau-kaol-qz, and lau-kaol-calc-qz. For replacement of laumontite by kaolinite a very low pH and considerable dissolved CO, are required. Analcime is replaced at higher pH values, and is apparently unstable in contact with sea water.] Thompson, A.B., 1976. Investigation of laumontite-calcite-quartz relations at low X co,. In: G. Biggar (Editor), Progress in Experimental Petrology; Third Progress Report. Natl. Environment Res. Counc. Publ. Ser. D. No. 6: 7-9. [Investigation of lau calc = pr qz H,O CO, in the range 17O0-35O0C along the H,O liquidvapour curve, using the weighed-crystal (calcite) method. Calculated values of Xco,and mc02 are in agreement with those found at Broadlands (Browne and Ellis, 1970). These values are an order of magnitude greater for lau CO, = calc kaol SiO, H,O.] Also: - Brown, 1977 (ref. p. 532).

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Metasomatism and change in bulk chemical composition of volcanic rocks during zeolite facies and prehnite-pumpellyite facies metamorphism

Barrows, K.J., 1980. Zeolitization of Miocene volcaniclastic rocks, southern Desatoya Mountains. Nevada. Geol. SOC.Am. Bull., 91(4): 199-210. [Heulandite, clinoptilolite, mordenite, analcime, thomsonite (?), erionite, and chabazite (?). Formed mainly through breakdown of glass; paragenetic sequence is given. Chemical comparisons indicate that Si, Ca, Na, K and H,O were mobile; the same inferred for Fe and Mg from petrography. Na and possibly some K were lost during diagenesis.] Hashimoto, M., Kashima, N., Kato, A,, Katto, J., Kuwano, Y . , Matsubara, S., Saito, Y . , Suyari, K. and Tiba, T., 1976. Acid volcanic rocks of the Okanaro Group in the Kurosegawa Tectonic Zone, Shikoku. Mem. Natl. Sci. Mus. (Tokyo), No. 9: 9-16 (in Japanese, English summary). [Na enrichment of rhyolites and K enrichment of tuffs during prehnite-pumpellyite facies metamorphism (see also Hashimoto, M., 1977. Low-grade metamorphism of the Okanaro Group of the Kurosegawa belt, Shikoku. Bull. Natl. Sci. Mus. (Tokyo), Ser. C (Geol.), 3(3): 147-149).] Kuniyoshi, S . and Liou, J.G., 1976. Burial metamorphism of the Karmutsen volcanic rocks, northeastern Vancouver Island, British Columbia. Am. J . Sci., 276: 1096- 1 1 19. [Depletion of Na, Si, Ca, and A1 from aquagene tuffs and pillow rims during prehnite-pumpellyite facies metamorphism. Local equilibrium was approached in most mineral assemblages under high p H 2 0and pco2. The spilitic features of the volcanic rocks are metamorphic and not metasomatic or deuteric).] Schau, M., 1974. Low-grade metamorphism and metasomatism in the Nicola Group, B.C. Can. Mineral., 12: 543.

[Absence of minerals from assemblages indicate that pNazO and pcaovaried in flows that had more or less the same initial composition.] Smith, R.E., 1980. Recognizing the paths of metamorphic/metasomatic fluids in a basic volcanic pile, Hamersley Basin, Western Australia. 26me Congr. Geol. Int., Paris, 1980, Abstr., 1: 93. [Relict domains allow assessment of departures from original compositions during prehnite-pumpellyite to greenschist facies metamorphism. Extent of mass transport by metamorphic fluid-rock interaction calculated.] Strong, D.F., Dickson, W.L. and Pickerill, R.K., 1979. Chemistry and prehnite-pumpellyite facies metamorphism of calc-alkaline Carboniferous volcanic rocks of southeastern New Brunswick. Can. J . Earth Sci., 16: 1071-1085. [Locally significant silicification and variable chloritization of most samples. Ti, P, Zr, Rb, Nd. Ga. and Y are relatively immobile (assuming constant AI).] Wood, D.A., Gibson, I.L. and Thompson, R.N., 1976. Elemental mobility during zeolite facies metamor-

phism of the Tertiary basalts of eastern Iceland. Contrrb. Mineral. Petrol., 55: 241-254. [Significant mobilization of Si. Mg, K, Rb, Sr and light REE. Values for Ti, P. Zr, Y. Nb, Ta. Hf are relatively unaffected by metasomatic transport.] Also: -

Boles and Coombs, 1977 (ref. p. 532).

Inversions and repetitions in zeolite- to pumpellyrte-facies zoning

Aguirre, L., Levi, B. and Offler, R., 1978. Unconformities as mineralogical breaks in the burial metamorphism of the Andes. Contrib. Mineral. Petrol., 66: 361-366. [Each of several stratigraphical-structural units in the Andes of Peru and Chile shows a facies series covering part or all of the range between the zeolite and the greenschist facies: mineralogical breaks coincide with the regional unconformities. Cases of high-grade assemblages overlying lower-grade assemblages. Cf. Levi (1970).] Offler, R., Aguirre, L., Levi, B. and Child, S., 1980. Burial metamorphism in rocks of the Western Andes of Peru. Lithos, 13(1): 31-42. [See preceding abstract. The presence of wairakite and the development of a wide range of metamorphic facies in thin sequences suggest high geothermal gradients.] Tzeng, S.-Y. and Lidiak, E.D., 1976. Low-grade metamorphism in east-central Puerto Rico. Geol. Soc. Am., Annu. Meet. 1976, Abstr. Progr., 8(6): 1150. [Grade is generally related to depth of burial, but locally actinolite qone overlies zeolite and prehnite-pumpellyite zones. This is ascribed to either ( I ) two cycles of sedimentation, or (2) high thermal regime in the vicinity of a fault zone.] Zeolite zoning in geothermal and other high-temperature areas

Besse, D., Desprairies, A,, Jehanno, C. and Kolla, V., 1981. Les parageneses de smectites et de zeolites dans une serie pyroclastique d’lge eocene moyen de I’Ocean lndien (D.S.D.P.. leg 26. site 253). Bull. Mineral., 104: 56-63. [Three successive zeolite zones (phillipsite; clinoptilolite-mordenite; analcime-clinoptilolite) in a 550 m thick hyaloclastic sequence are attributed t o hydrothermal alteration in a high-temperature geothermal area (- 20O0C/km).] Jefferis, R.G. and Voight, B., 1981. Fracture analysis near the mid-ocean plate boundary, ReykjavikHvalfjordur area, Iceland. Tectonophvsics, 76(3/4): 171-236. [Fluid inclusion temperature data from fracture and vug minerals: temperatures and depth for the seven zeolite zones distinguished in different areas. The geothermal gradient was approx. 80°C/km during the secondary mineralization.] Kristmansdottir, H., 1979. Alteration of basaltic rocks by hydrothermal activity at 1O0-30O0C. In: M.M. Mortland and V.C. Farmer (Editors), International Clay Conference, 1978. Elsevier, Amsterdam. pp. 359-367. [Deep wells in six high-temperature geothermal areas (> 200’ at I km depth). Four mineralogical alteration zones are distinguished. Smectites have transformed into mixed-layer clay minerals and swelling chlorites at 200’-230OC. Zeolites and Ca-silicates (except wairakite) disappear by about 200OC. Epidote and prehnite are formed slightly above 240°C; actinolite appears near to 300°C.] Kristmansdottir, H. and Tomasson, J.. 1978. Zeolite zones in geothermal areas in Iceland. In: L.B. Sand and F.A. Mumpton (Editors), Natural Zeolites, Occurrence, Properties. Use. Pergamon. Oxford. pp. 269-275. [Four zones in “low-temperature” areas: chabasite; mesolite-scolecite; stilbite; and laumontite. In high-temperature areas (> 200’ at 1 km depth) the sequence includes mordenite. heulandite. laumontite, and analcite. At even higher temperatures analcime and wairakite are formed.] Leitch, E.C., 1978. Hydrothermal metamorphism of the Whangakea Basalt, New Zealand. N.Z. J. Geophys., 21(3): 287-291. [The second of two metamorphic episodes produced several zeolites: i t is due either to a late stage of the earlier hydrothermal metamorphism, or to burial during the Miocene.]

537 Sheridan, M.F. and Maisano, M.D., 1975. Zeolite and sheet silicate zonation in a LateTertiary geothermal basin near Hassayampa, central Arizona. Proc. 2nd U . N . Symp. on Development and Use of Geothermal Resources, 1 : 597-607. [For zones of zeolites and associated phyllosilicates-ranging from (I) mordenite, epistilbite, kaolinite, 1Md muscovite, to (IV) heulandite, chabazite. thomsonite, chlorite, 2M muscovite-formed in a geothermal system that has cooled since.] Pumpellyite in geothermal, oceanic or other high-temperature-low-pressure terranes

Franks, S.G., 1974. Prehnite-pumpellyite metamorphism of the New Bay Formation, Exploits zone, Newfoundland. Can. Mineral., 12: 456-462. [The andesitic sandstones contain prehnite-epidote, but pumpellyite is uncommon. The high metamorphic temperatures (estimated at 300'-400OC at Plead < l kb) are ascribed to high heat flow due to dykes and sills.] Mevel, C., 1981. Occurrence of pumpellyite in hydrothermally altered basalts from the Vema Fracture Zone (Mid-Atlantic Ridge). Contrib. Mineral. Petrol., 76(4): 386-396. [Occurrence discussed in terms of temperature, p H Z 0 fO,. . Recrystallization by hydrothermal circulation of sea water at very low pressures (< 1 kb). Strong modification of bulk composition of the rocks during hydrothermal metamorphism. Smewing. J.D.. Simonian. K.O. and Gass, I.G.. 1975. Metabasalts from the Troodos Massif, Cyprus: genetic implication deduced from petrography and trace element geochemistry. Contrib. Mineral. Petrol., 5 I ; 49-64. [Zeolite to greenschist facies metamorphism of lower pillow lavas and the sheeted dykes ( = axial sequence) during sea-floor geothermal cycle adjacent to the axis. The zeolite facies of the upper pillow lavas is not related to this cycle (these were erupted further from the ridge).] Also: - Kuniyoshi and Liou, 1976 (ref. p. 535). Pumpellyite facies with data on phase petrology and mineral compositions (incl. relation to hulk composition and extrapolated P - T relations and geothermal gradients)

Jolly, W.T., 1980. Development and degradation of Archean lavas. Abitibi area, Canada, in light of major element geochemistry. J . Petrol., 21(2): 323-363. [Absence of stilpnomelane, relatively little coexistence of pumpellyite and actinolite. and wide range in Mg-Fe content of pumpellyites indicates relatively low-pressure metamorphism.] Katagas, D. and Panagos, A.G., 1979. Pumpellyite-actinolite and greenschist facies metamorphism in Lesvos Island (Greece). TMPM Tschermaks Mineral. Petrogr. Mitt., 26: 235-254. LThe extensive distribution of chlorite-calcite instead of the Ca-AI-silicate-bearing assemblages is ascribed to local variations in pLcO,. Phase relations suggest metamorphism at 27O"-36O0C and pressures little lower than 5 kb.] Kirchner. E.Ch., 1979. Pumpellyitefiihrende Kissenlavabreccien in der Gips-Anhydrit-Lagerstatte von Wienern am Grundlsee, Steiermark. TMPM Tschermaks Mineral. Petrogr. Mitt., 26: 149- 162. [Pumpellyite along glass matrix is associated with carbonate, gypsum, and anhydrite in vugs. Metamorphism at very low P and T i n a gas phase of very unusual composition due to the vicinity of the sulfate deposits.] Schreyer, W. and Abraham, K., 1978. Prehnite/chlorite and actinolite/epidote bearing mineral assemblages in the metamorphic igneous rocks of La Helle and Chalks, Venn-Stavelot-Massif. Annu. Soc. Geol. Belg.. 101: 227-241. [Prehnite + chlorite in metatonalites; actinolite + epidote in metabasalts. Prehnite may have formed at the expense of pre-existing pumpellyite.] A Iso: -

Brand, 1980 (ref. p. 539).

Pumpellyite facies in orogenic terranes-without cophanitic terranes)

phase petrology (exclusive of pumpellyite :ones in glau-

Bevins, R.E., 1978. Pumpellyite-bearing basic igneous rocks from the Lower Ordovician of North Pembrokeshire, Wales. Mineral. Mag., 42: 81-83. [Caledonian prehnite-pumpellyite facies metamorphism with prehnite, pumpellyite, stilpnomelane, and actinolite.] Davies, H.L., 1980. Folded thrust fault and associated metamorphics in the Suckling-Dayman massif, Papua New Guinea. Am. J . Sci., 280-A (Jackson Volume), part 1: 171-191. [Metamorphic grade from unmetamorrjhosed basalt through prehnite-pumpellyite and pumpellyite-actinolite (-blueschist) to greenschist facies towards the overlying ultramafic thrust sheet. Nystrom, J.O. and Levi, B., 1980. Pumpellyite-bearing Precambrian rocks and post-Svecokarelian regional metamorphism in central Sweden. Geol. Foren. Stockholm Forh., 102( I): 37-39.] [Prehnite-pumpellyite facies in Jotnian, and pumpellyite-actinolite facies in sub-Jotnian rocks. Common overprinting of “retrograde” minerals on amphibolite facies assemblages in Svecokarelian rocks could in part be ascribed to these lowest-grade metamorphic episodes.] Oliver, G.J.H., 1978. Prehnite-pumpellyite facies metamorphism in County Cavan, Ireland. Nature, 274: 242-243. [A regional distribution of prehnite-pumpellyite facies metamorphism is believed to occur in Ordovician rocks of the paratectonic and the southern orthotectonic Caledonides of the British Isles.] Oliver, G.J.H. and Leggett, J.K., 1980. Metamorphism in an accretionary prism: prebnite-pumpellyite facies metamorphism of the Southern Uplands of Scotland. Trans. R. Soc. Eainburgh, Earrh Sci., 71(4): 235-246. Papezik, V.S., 1974. Prehnite-pumpellyite facies metamorphism of Late Precambrian rocks of the Avalon Peninsula, Newfoundland. Can. Mineral., 12: 463-468. [Broad prehnite zone-narrow prehnite-pumpellyite zone-actinolite zone (without pumpellyite). The metamorphic grade increases in the direction of tighter folding and increasing penetrative deformation of Acadian (Devonian) age.] Richter, D.A. and Roy, D.C., 1974. Sub-greenschist metamorphic assemblages in northern Maine. Can. Mineral., 12: 469-474. [Successive prehnite-analcime (this assemblage only in quartz-free rocks), prehnite-pumpellyite. and pumpellyite-epidote-actinolite (with prehnite) zones within the prehnite-pumellyite facies. Acadian metamorphism did not exceed the prehnite-pumpellyite facies; the possibility of preceding Taconic metamorphism to the same grade remains open.] Richter, D.A. and Roy, D.C., 1976. Prehnite-pumpellyite facies metamorphism in central Aroostook County, Maine. Geol. Soc. Am. Mem., 146: 239-261. Roberts, B., 1981. Low grade and very low grade regional metabasic Ordovician rocks of Llyn and Snowdonia, Gwyneld, north Wales. Geol. Mag., 1 18(2): 189-200. [The isograds pumpellyite-in, pumpellyite-out-clinozoisite-in, and biotite-in have been mapped. The metamorphism was syn- and immediately post-tectonic (end-Silurian to Devonian).] Ryan, P.D., Floyd, P.A. and Archer, J.B., 1980. The stratigraphy and petrochemistry-of the Lough Nafooey Group (Tremadocian), western Ireland. J . Geol. SOC.London, 137(4): 443-458. [Assemblages of zeolite, pumpellyite. and greenschist facies (with increasing stratigraphic depth) in basic volcanics of the South Mayo Trough formed prior to Llandovery sedimentation. and are probably related to the Grampian orogeny.] Williams, H. and Einarson. G .W.. 1976. Discussion of “Prehnite- and pumpellyite-bearing mineral assemblages, west side of the Applachian metamorphic belt, Pennsylvania to Newfoundland” by E-an Zen. J . Petrol.. 17(1): 135-136 (reply by E-an Zen. p. 137). [Suggest that the prehnite-pumpellyite assemblages relate to pre-transport burial metamorphism not requiring exceptionally high pressures, and have no genetic affinity with the blueschist facies metamorphism related to Taconic orogeny and possible subduction.] Also: - Kisch, 1980 (ref. p. 523. 539). Stalder, 1979 (ref. p. 523). - Frey et al., 1976 (ref. p. 522). ~

539 Relationships between pumpellyite facies and illite crystallinity in orogenic belts

Aprahamian, J. and Pairis, J.L., 1981. Very low grade metamorphism with a reverse gradient induced by an overthrust in Haute-Savoie (France). In: Thrust and Nappe Tectonics. Geol. SOC.London, London, pp. 159-165. [Local reverse gradients, as expressed by (1) the metamorphic assemblages in Taveyanne sandstones. and (2) mineralogical composition and illite crystallinity in the argillaceous fraction of the associated shales, are ascribed to heat production along a thrust plane. This reverse gradient is superimposed on the earlier known gradient (decrease in grade towards the top of the series and towards the external part of the chain).] Arkai, P., 1973. Pumpellyite-prehnite-quartz facies Alpine metamorphism in the Middle Triassic volcanogenic-sedimentary sequence of the Biikk Mountains, northeast Hungary. Acta Geol. Acad. Sci. Hung., 17(1-3), 67-83. [The anchimetamorphism (“illite-chlorite facies”) and textures (“initial metagenesis”) in the slates is correlated with the prehnite-pumpellyite facies in the associated volcanics, which is ascribed to Alpine (Cretaceous) metamorphism. A table gives the correlation of zeolite and prehnite-pumpellyite facies with textural and mineralogical criteria in rocks without critical Ca-Al-silicate minerals.] Arkai, P., 1980. Metamorphic evolution of the Paleozoic and Mesozoic formations in one of the Alpine mobile belts of the Pannonian Basin. 26e Congr. Gkol. Int., Paris, 1980, Abstr., 1: 12. [On the basis of mineral assemblages, illite-sericite crystallinity, b,, and vitrinite reflectance values, the metamorphic grade in a Devonian to Upper Triassic profile decreases upwaTds from low-temperature-low-intermediate pressure greenschist facies to prehnite-pumpellyite-quartz facies; no unconformities or sudden changes in grade that would indicate a Hercynian orogenic phase or metamorphism.] Bevins, R.E., Robinson, D., Rowbotham, G. and Dunkley, P.N., 1981. Low-grade metamorphism in the Welsh Caledonides (abstr.). In: Metamorphic Studies: Research in Progress. J . Geol. Soc. London, 138(5): 634. [Zoning from prehnite-pumpellyite through pumpellyite-actinolite to greenschist facies. “Epimetamorphic” illite 10 A peak widths are associated with greenschist facies rocks; anchizonal values with the prehnite-pumpellyite zone.] Brand, R., 1980. Die niedriggradige Metamorphose einer Diabas-Assoziation in Gebiet Berg/Frankenwald. Neues Jahrb. Mineral Abh., 137: 82-101. [Changes in the composition of chlorite (Al/Si ratio) and clinozoisite-epidote group minerals (Fe/(AI Fe) ratio) with progressive zoning from prehnite-pumpellyite though pumpellyite-actinolite to low greenschist facies. Medium-pressure metamorphsm. Facies is compared to the illite and chlorite crystallinities of Ludwig (1973).] Bril, H. and Thiry, M., 1976. Le metamorphisme de basse pression anchi- a mesozonal de la region de Bodennec (Finistere): essai methodologique. C.R. Acad. Sci. Paris, Skr. 0,283(3): 227-230. [Hercynian prehnite-pumpellyite, pumpellyite, and actinolite zones (from S to N). “Epizonal” illite crystallinities are within the actinolite zone. Anchizone is poorly defined, with many “diagenetic” 10 A peak widths. The imperfect correlation with the Ca-Al-silicate assemblages is due to errors in the crystallinity method (e.g., anchizonal values within the “epizone” are in part due to presence of biotite). The results are complementary to those of Sagon (1970).] Cortesogno, L. and Venturelli, G., 1978. Metamorphic evolution of the ophiolite sequences and associated sediments in the northern Apennines-Voltri Group, Italy. In: H. Closs, D. Roeder, and K. Schmidt (Editors), Alps, Apennines, Hellenides. Schweizerbart, Stuttgart, pp. 253-260. [Prehnite-pumpellyite and subordinate prehnite-zeolite facies in the Sestri-Voltaggio Zone. In the Northern Apennines to the east prehnite-pumpellyite and subordinate prehnite-zeolite facies; the equivalent anchimetamorphism is found in the sedimentary sequences (Venturelli and Frey, 1977).] Kisch, H.J., 1980. Illite crystallinity and coal rank associated with lowest-grade metamorphism of the Taveyanne greywacke in the Helvetic zone of the Swiss Alps. Ecologae geol. Helv., 73(3): 753-777. [Illite crystallinities associated with laumontite-bearing, and with laumontite-free, prehnite- and pumpellyite-bearing Taveyanne greywackes are respectively “diagenetic” and middle- to high-grade anchimetamorphic (see also Stalder, 1979).] Kisch, H.J., 1982. Coal rank and illite crystallinity associated with the zeolite facies of Southland and the

+

5 40 pumpellyite-bearing facies of Otago, southern New Zealand. N . Z . J. Geol. Geophys., 24(3): 349-360. [Highest-grade “diagenetic’ illite 10 A peak widths are found in the laumontite-bearing Torlesse terrane. Predominantly low-grade anchimetamorphic values associated with the prehnite-pumpellyite facies of the Caples-Pelorus terrane. Unequivocal “epizonal” values appear somewhat before (in South Otago), or already beyond (in North Otago) the pumpellyite-actinolite isograd, and predominate in a lawsonite-bearing zone of western Otago.] Lecolle, M. and Roger, G., 1976.Metamorphisme regional hercynien de “faible degrk” dans la province pyrito-cuprifkre de Huelva (Espagne). Consequences petrologiques. Bull. Soc. Giol. Fr., Sir. 7 , 18(6): 1687-1698. [The metamorphic zones explained by Schermerhorn (1975)as the result of one metamorphic phase, and separated by a pumpellyite-prehnite isograd, are now ascribed to two successive phases of low-grade and very-low-grade (prehnite-pumpellyite facies) metamorphism. Illite crystallinities largely “faible degrk“ (scale of Sagon and Dunoyer de Segonzac, 1972); paragonite, pyrophyllite.] Leitch, E.C., 1975.Zonation of low grade regional metamorphic rocks, Nambucca slate belt, northeastern New South Wales, J. Geol. Soc. Aust., 22(4): 413-422. [Three isograds mapped across pumpellyite facies in metabasics: (1) stilpnomelane and pumpellyite in; (2) prehnite out, and pumpellyite-actinolite in; (3) pumpellyite out. Parallel changes in textures of metaclastic rocks. Mica crystallinity zones show overlapping peak width values. Changes in chlorite composition with grade.] Also: - Robinson et al., 1980 (ref. p. 520). - Stalder, 1979 (ref. p. 523). Pumpellyite zones in blueschist terranes

Hashimoto, M. and Kanehira, K., 1979.Preliminary study on mineral paragenesis of quartz schists of the Iimori district, Sambagawa terrane, Japan. Mem. Natl. Sci. Mus., Tokyo, 12: 23-27. [Pumpellyite-actinolite zone (A) is succeeded by glaucophane-actinolite zone (B). Pumpellyite is absent from zone B; glaucophane appears in the middle of zone A,] Hen+., F., 1975. Petrography of the Kamuikotan metamorphic belt at the Ubun-Orowen cross section, central Hokkaido, Japan. J. Far. Sci., Hokkaido Uniu., Ser. IV. 16(4): 453-470. [Zone I: pumpellyite, chlorite, jadeitic pyroxene. Zone 11: crossite, jadeitic pyroxene, pumpellyite, lawsonite, and actinolite. Zone 111: actinolite, chlorite, and epidote (no Na-amphiboles, Na-pyroxenes, lawsonite, or pumpellyite). do,, of phengite and d,, of chlorites given. The facies series is intermediate between Seki’s (1969)type I (Franciscan) and I1 (Sambagawa).] Watanabe, T., 1977. Metamorphism of the Sambagawa and Chichibu belts in the Oshika district, Nagano prefecture, central Japan. J. Far. Sci., Hokkaido Uniu.,Ser. IV. 17(4): 629-694. [Appearance of Na-amphibole is controlled by Fe,O,/FeO ratio and MgO, and that of pumpellyite depends on MgO/Ca and Fe,O,/FeO ratios: these minerals cannot be used as index minerals without considering bulk composition. Changes in composition of actinolite and epidote with metamorphic grade.] Surveys of the occurrence of miscellaneous minerals in the pumpellyite facies

Cortesogno, L. and Lucchetti, G., 1976. Carfolite nei diaspri della Val Graveglia: caratteristiche mineralogiche e considerazioni genetiche. Ofioliti (Bologna), l(3): 373-382. [Carpholite in manganiferous layers in radiolarian cherts has formed during prehnite-pumpellyite metamorphism.] Hashimoto, M. and Kanehira, K., 1975. Some petrological aspects on stilpnomelane in glaucophanitic metamorphic rocks. J. Jpn. Assoc. Mineral. Petrol. Econ. Geol., 70(1 1): 377-387. [Stilpnomelane appears in pumpellyite zones (without glaucophane) of many glaucophanitic terranes. Stilpnomelane is commonly associated with quartz, chlorite, and calcite, and rarely with epidote, pumpellyite, and actinolite.] Pringle, I.J. and Kawachi, Y.,1980.Axinite mineral group in low-grade regionally metamorphosed rocks

541 in southern New Zealand. Am. Mineral., 65( 11- 12): 11 19- 1129. [In quartz-bearing vein assemblages in prehnite-pumpellyite. pumpellyite-actinolite. and chlorite zone of greenschst facies. Fe-rich axinite in spilitized volcanite and graywacke, sometimes with prehnite. pumpellyite, Fe-rich epidote, and chlorite. Mn-rich axinite in ferruginous and manganiferous cherts.] Furies series of lowest-grade metamorphism and their geologic and plate-tectonic enarronmeni

Oliver, G., 1980. Metamorphism of the paratectonic Caledonides of the British Isles. 26e Congr. GCol. Int.. Paris, 1980, Abstr., 1 : 69. [Four metamorphic environments are recognized: ( I ) sedimentary burial (Mayo. N and S Wales); (2) subduction zone (Southern Uplands, Longford-Down, Clare); (3) obduction zone (Ballantrae-Girvan): (4) interior ocean-plate during spreading (Bail Hill, Southern Uplands). HP/LT facies series in paratectonic Scotland and Ireland is paired with a contemporaneous moderate-high P/HT facies series in the northern orthotectonic Caledonides. Mite-crystallinity and vitrinite-reflectance methods also used.]

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543

REFERENCES INDEX Abbas, M., 348, 374, 395, 471, 474 Abelson, P.H., 38, 97 Adams, J.E. and Rhodes, M.L., 82, 83, 97 Adelsech, C.G. and Berger, W.H.. 221, 225. 236, 237, 238, 282 Adelsech, C.G., Geehan. G.W. and Roth, P.H., 241, 249, 268, 282 Albee, A.L. and Zen, E-An., 432, 433, 439, 4 74 Albers, D., see Eslinger, E. et al. Albrecht, P., see Durand, B. et al. Alderman, A.R. and Skinner, H.C.W., 78, 97 Aleksandrova, V.A., see Gavrilov. A.A. and Aleksandrova. V.A. Alexandrova, V.A., see Kossovskaya, A.G. et al. Aleksandrova, V.A., see Shutov, V.D. et al. Alexandersson, T.. 213, 282, 431. 474 Ali, S.A. and Friedman, G.M.. 13 Ah, Y.A., see West, LA. et al. Ahmen, H., 88, 97 Allen, E.T., Crenshaw. J.L. and Johnston, J.J., 61. 97 Allen, J.E., see Sun, Ming-Shan and Allen. J.E. Allen, J.R.L., 45. 59, 97 Alling, H.L., 508 Almon, W.R., Fullerton, L.B. and Davies. D.K., 43, 97 Al-Termeeni. A., see Gunatilaka, A. et al. Althaus, E., 362, 363, 474 Althaus, E., see Smykatz-Kloss, W. and Althaus, E., American Geological Institute. 185, 203 Amstutz, G.C., 508 Amstutz, G.C. and Bubenicek. L., 97 Anderson, D.L. and Benson, C.S.. 18. 97 Anderson, D.M.. see Reynolds, R.C. and Anderson, D.M Anderson, G.T. and Han, T.M., 508 Anderson, N.R. and Malahoft. A.. 214, 282 Anderson, T.F. and Schneiderman. N.. 265, 282 Anderson. T.F., see Lawrence, J.R. et al. Andree, K., 19, 97 Angot. P., see Bichelonne. J. and Angot, P. Aoyagi, K.. 14 Aoyagi, K. and Kazama, T., 14

Appelo. C.A.J., 156. 203 Aprahamian. J., 324. 348. 356, 364. 366. 395. 461, 466, 471. 474 Arad. A,. see Magaritz.. M. et al. Arend, J.P., 497. 508 Aronson. J.L. and Hower. J.. 328. 474 Arrhenius, G.. 39. 97. 123. 203 Arrhenius. G.O.S.. see Goldberg. E.D. and Arrhenius. G.O.S. Arthur. M.A. and Schlanger. S.O.. 38. 97 Artru. Ph.. Dunoyer de Segonzac. G.. Combaz, A. and Giraud. A,. 306, 312. 330. 343. 348, 354. 355. 356. 357, 358. 359. 366. 373. 377, 474 Artru. Ph. and Gauihier. J.. 312. 319. 474 Artru. P.. see Dunoyer de Segonzac. G . et al. Ataman, G. and Baysol. 0.. 137. 203 Ataman. G.. see Lucas. J. and Ataman. G . Athy, L.F.. 190 Atwater. G.I., see Tester, A.C. and Atwater. G.I. Augusthtis. S.S.. 508 Ault. W.U.. 38, 97 Ault. W.U. and Kulp, J.L.. 62. 97 Avias. J.. 63, 68. 97 Avnimelech. M., see Gignoux. M. and Avnimelech. M. Aylmore, L.A.G. and Quirk. J.P.. 199, 203 Baas Becking, L.G.M.. 23, 28. 97 Baas Becking, L.G.M.. Kapland. I.R. and Moore, D.. 25, 27. 29, 38. 98, 508 Baas Becking. L.G.M. and Moore. D.. 60, 97 Back. W., see Hanshaw, B.B. et al. Bader. R.G.. 36. 38, 98 Badiozamani, K., 85, 98 Bailey, S.W. and Brown. B.E., 339, 474 Bailey. S.W., see Brown, L.F. et al. Bain, D.C., see Wilson. M.J. et al. Baker. G., see Edwards. A.B. and Baker, G. Balgord, W.D., see Keller, W.D. et al. Banno, S.. 450, 474 Barash, M.S., 216, 282 Bardossy, Gy., see Szadecky-Kardoss, E. et al. Barlier, J.. 348, 356, 465. 475 Barlier, J., Ragot. J.-P. and Touray, J.-C., 356, 465, 475 Baron, G., 81, 98

544 Barth, T.F.W., 25, 29, 98 Barth, T.F.W., Correns, C.W. and Eskola, P., 86, 98 Bartlett, G.A. and Greggs, R.G., 213, 282 Bass, M.N., 144, 203 Bastin, E X , 61, 98 Bates, R.L. and Jackson, J.A., 17, 18, 96, 98 Bathurst, R.G.C., 4, 1 4 , 40, 41, 44, 73, 79, 98, 213, 214, 216. 219, 220, 226, 242, 243, 254, 255, 279, 282 Baturin, V.P., 47, 98 Baysol, 0.. see Ataman, G . and Baysol, 0. Be, A.W.H., Morse, J.W. and Harrison, S.M., 225, 282 BC. A.W.H., see Sliter, W.V. et al. Beall, A.O. and Fischer. A.G., 255, 282 Beall, A.O. and Ojakangas, R.W., 327, 475 Beardsley, K.M., see Robie, R.A. et al. Bearth, P., 471, 475 Beck, K.C., see Siever, R. et al. Beck, K.C., see Weaver, C.E. and Beck, K.C. Bensch, J.J. see Biljon, W.J. and Bensch, J.J. Benson, C.S., see Anderson, D.L. and Benson, C.S. Bentley, S.P. and Smalley, F.J., 180, 200, 203 Berg, G., 498, 508 Berg, R.R., 75, 98 Berger, W.H., 18, 214, 216, 218, 220, 221, 223, 224, 225, 226, 235, 236, 238, 251, 282 Berger, W.H., Ekdale, A.A. and Bryant, P.P., 239, 282 Berger, W.H. and Von Rad, U., 25 1, 283 Berger, W.H. and Winteren, E.L., 218, 283 Berger, W.H., see Adelseck, C.G. and Berger, W.H. Berger, W.H., see Ekdale, A.A. and Berger, W.H. Berger. W.H., see Johnson, T.C. et al. Berger, W.H., see Parker, F.L. and Berger, W.H. Berger, W.H., see Roth, P.H. and Berger, W.H. Berger, W.H., see Sliter, W.V. et al. Bergin, M.G., see Hosterman. J.W. et al. Bernat, M. and Church, T.M., 141, 203 Berner, R.A., 23, 38, 47, 49, 61, 63, 98, 134, 203, 222. 238, 254, 283 Berner. R.A., see Morse, J.W. and Berner, R.A. Berner, R.A.. see Paterson, M.N.A. et al. Berner, R.A., see Siever, R. et al. Berrow, M.L., see Wilson, M.J. et al. Berry. F.A.F., 43, 98 Berry, R.W. and Johns, W.D., 98 Bethke. P.M.. see Robie. R.A. et al.

Bettardy. W.J., see Wilson, M.J. et al. Beuf, S., Biju-Duval, B., Stevaux, J. and Kulhicki, G., 367, 475 Beugniet, A,, Godfriaux, L. and Rohaszynski, F., 343, 475 Bichelonne, J. and Angot, P., 498, 508 Bien, G.S., Contois, D.E. and Thomas, W.H., 40, 98 Bien, G.S., see Paterson, M.N.A. et al. Biju-Duval, B., see Beuf, S. et al. Biljon, W.J. and Bensch, J.J., 351, 475 Billings, G.K., see Hiltabrand, R.R. et al. Birch, G.F., 139, 203 Birch, G.F., Willis, J.P. and Richard, R.S., 139, 204 Biscaye, P.E., 54, 55, 56, 58, 98, 119, 121, 122, 123, 204 Biscaye, P.E. and Dasch, E.J., 56, 98 Biscaye, P.E., see Kolla, V. and Biscaye, P.E. Biscaye, P., see Venkatarathman, K. and Biscaye, P. Bishop, D.G., 338, 386, 388, 390, 391, 403, 430, 443, 451, 453. 464, 468, 475 Bishop, D.G., see Landis, C.A. and Bishop, D.G. Bissell, H.J., 35, 63, 97, 98 Bissell, H.J., see Chilingar, G.V. et al. Black, M., 216, 283 Black, M., see Hatch, F.H. et al. Blackwelder, E., 46, 98 Blatt, H., Middleton, G., and Murray, R., 27, 31, 41, 43, 47, 65, 91, 98 Blondel, F., 509 Blyth, C.R., see Bredehoeft, J.D. et al. Boardman, M.P., see Mullins, H.T. et al. Bocquet, J., 393, 395, 457, 471, 472, 475 Bodine, M.W. and Standaert, R.R., 135, 204 Bogdanov, V.V., see Buryanova, Ye.2. and Bogdanov, V.V. Boggild, O.B., 81, 98, 99 Boischot, P., see Demolon, A. and Boischot, P. Boles, J.R. and Coombs, D.S., 383, 400, 403, 408, 426, 431, 461, 475 Boles, J.R. and Wise, W.S., 141, 204 Bolt, G.H., 126, 197, 204 Bonatti, E., 53, 82, 99, 140, 204 Borchert, H., 498, 509 Borchert, H. and Muir, R.O., 31, 32, 64, 83, 99 Bostick, N.H., 455, 456, 475 Boswell, P.G.H., 76, 99 Boudier, F. and Nicolas, A., 370, 475 Bourgeois, J., see Fairbridge, R.W. and Bourgeois, J. Boyce, R.E., see Pimm, A.C. et al.

545 Braconnier, M.A., 498, 509 Bradley, J.S.. see Newell, N.O. et al. Bradley, W.F.. 366, 475 Bradley, W.F. and Serratosa, J.M.. 195. 204 Bradley, W.F., see Grim. R.E. et al. Bradley, W.H., 53, 99 Bradshow, M.J.. see Hallam. A. and Bradshow. M.J. Braitsch, 0..31, 32. 64, 99 Bramkamp, R.A., 78 Bramlette. M.N., 216, 283 Braun, H.. 497. 509 Bray, R.H.. see Grim, R.E. et al. Bredehoeft, J.D., Blyth. C.R., White. W.A. and Maxey, G.B.. 76, 99 Breger, I.A.. 77, 99 Bricker, O.P., 213, 283 Brindley. G.W., 366, 475 Brindley, G.W. and Sandalaki, 2.. 366. 368. 4 75 Brinkmann, R., 123, 204 Briskin, M. and Schreiber, B.C., 63, 99 Broecker. W.S., see Takahashi, T. and Broecker. W.S Bromley, R.G., 226. 283 Brongersma-Sanders. M.. 61, 99 Brown. B.E., see Bailey, S.W. and Brown, B.E. Brown. C.E.. 388, 400, 462, 475 Brown, C.E. and Thayer, T.P.. 388. 400. 401, 436, 462, 475 Brown, E.H.. 386, 475 Brown, G.. Catt, J.A. and Weir, A.H.. 420, 458, 475 Brown, G. and Weir, A.H., 366, 475 Brown. J.S., 498. 509 Brown, L.F.. Bailey. S.W., Cline. L.M. and Lister. J.S., 153. 204 Browne. P.R.L. and Ellis, A.J., 440, 476 Brueckner, H.K.. see Lawrence. J.R. et al. Bryant, P.P., see Berger, W.H. et al. Bryce, M.R. and Friedman. G.H., 75, 99 Bubenicek, L., 4, 496. 497, 498, 500, 503, 508. 509 Bubenicek, L., see Amstutz. G.C. and Bubenicek, L. Budd, D.A. and Perkins, R.D.. 240. 283 Bukry, D.. 236, 283 Bundy, W.M.. see Conley. R.F. and Bundy, W.M. Burst, J.F.. 139, 161, 163, 190, 195, 196. 198, 204, 312. 313. 314, 320, 376 Burton. J.D. and Liss, P.S.. 66, 99 Buryanova, Ye.Z. and Bogdanov, V.V., 476 Bushinskij, I., 498, 509

Byrne. J.V.. see Riley. C.M. and Byrne. J.V Bystrom. A.M.. 322. 476 Cady, J.G.. see Flach. K.W. et al. Caillere. S. and Kraut. F., 497. 509 Caillere, S.. Mathieu-Sicand. A . and Henin. S.. 366, 476 Califet-Debyser. Y..see Tissot. B. et al. Cameron, L.B. and Sabine, P.A.. 420. 4 76 Campbell, A S . and Fyfe. W.S.. 407. 408. 409. 418. 421, 424, 448. 476 Campbell. F.A.. see Maiklem. W.R. and Campbell, F.A. Carman, M.F.. 338. 476 Carmouze. J.P.. see Pedro. G . et al. Carozzi. A.. 78. 81. 99 Carpenter. A.B.. 99 Carrigy, M.A. and Mellon, G.B.. 312. 462. 476 Carron. M.K.. see Stevens. R.E. and Carron. M.K. Caspers. H.. 36, 99 Cassan. J.P. and Lucas. J., 337. 476 Castano. J.R. and Garrels. R.M.. 497. 509 Castoiio. J.R. and Sparks. D.M.. 455. 476 Catt. J.A.. see Brown, G . et al. Cayeux. L.. 68. 69. 73. 74, 75, 93. 99. 497. 509 Chamberlain, C.F.. see Windom. H.L. and Chamberlain. C.F. Chamley. H.. 351. 476 Chamley. H.. Dunoyer de Segonzac. G . and Melieres. F., 135. 204 Chamley. H.. see Dunoyer de Segonzac, G . and Chamley. H. Chateauneuf, J.-J.. Debelmas. J.. Feys. R., Lemoine, M. and Ragot. J.-P.. 374, 395. 4 76 Chatterjee, N.D., 369. 476 Chauvel, J.J.. see Dimroth, E. and Chauvel. J.J. Chave, K.E.. 41, 43. 76, 78, 79, 99. 151. 204, 224, 283 f h e n , C., 238, 283 Chennaux, G. and Dunoyer d e Segonzac. G.. 360. 476 Chennaux. G., Dunoyer de Segonzac. G. and Petracco, F., 348. 360, 367. 476 Chernov, A.A., see Teodorovich, G.I. and Chernov, A.A. Chernov. A.A.. see Teodorovich. G.I. et al. Chilingar, C.V., 18, 36. 51. 66, 78, 81, 99, 100 Chilingar, G.V., Bissell. H.J. and Wolf, K.H., 213, 283 Chilingar, G.V. and Knight, L., 193. 194. 197, 204

Chilingarian, G.V. and Rieke, H.H., 193, 204 Chilingarian, G.V., Sawabini, C.T. and Rieke, H.H., 43, 100, 155. 204 Chilingarian, G.V. and Vorabutr, P., 49, 100 Chilingarian, G.V., see Rieke, H.H. and Chilingarian, G.V. Choquette, P.W. and Pray, L.C., 4, 14, 35, 100 Christ, C.L., see Carrels, R.M. and Christ, C.L. Church, T.M. and Velde, B., 138, 204 Church, T.M., see Bernat, M. and Church, T.M. Clauer, N. and Lucas, J., 336, 354, 476 Clausen, C.-D., see Hoyer, P. et al. Cline, L.M. see Brown, L.F. et al. Cloud, P.E., 51, 94, 100 Codispoti, L.A., see Piper, D.Z. and Codispoti, L.A. Cody, R.D. and Hull, A.B., 64, 100 Cohen, A,, see Staub, J. and Cohen, A. Coleman, R.G., see Taylor, H.P. and Coleman, R.G. Combaz, A., see Artru, Ph. et al. Condon, W.H., see Hoare, J.M. et al. Conley, R.F. and Bundy, W.M., 63, I00 Contois, D.E., see Bien, G.S. et al. Conway, E.J., 52, 56, 58, 100 Cook, F.M. and Cook, H.E., 226, 230, 241, 243, 283 Cook, H.E., 276,283 Cook, H.E. and Enos, P., 213, 283 Cook, H.E., see Cook, F.M. and Cook, H.E. Cook, H.E., see Piper, D.Z. et al. Cook, P.J., 91, 100 Cook, P.J. and McElhinny, M.W., 91. 100 Coombs. D.S., 9, 14, 296, 297, 300, 303, 306, 307, 308, 309, 376, 379, 381, 385, 386, 388, 389, 390, 402, 403, 408, 409, 410, 412, 414, 417, 418, 420, 428, 430, 436, 441, 447, 448, 454, 461, 464, 466, 468, 476, 477 Coombs, D.S., Ellis, A.J., Fyfe, W.S. and Taylor, A.M., 14, 53, 100, 296, 302, 376, . 379, 380, 381, 382, 385, 386, 388,402,403, 407, 408, 411, 412, 414, 415, 416, 417, 420, 424, 427, 440, 477 Coombs, D.S., Horodyski, R.J. and Naylor, R.S., 14, 390, 436, 439, 443, 448, 450, 454, 462. 464, 477 Coombs. D.S., Nakamura, Y. and Vaugnat, M.. 430, 469. 477 Coombs. D.S. and Whetten, J.T., 14, 424, 425, 477 Cnomhs. D.S., see Boles, J.R. and Coombs, D.S.

Coombs, D.S., see Landis, C.A. and Coombs, D.S. Cooper, B.N., 81, 100 Coplen, T.B. and Schlanger, S.O., 265, 283 Correns, C.W., 18, 26, 51, 57, 81, 86, 100, 294, 477, 498, 509 Correns, C.W. and Von Engelhardt, W., 58, I00 Correns, C.W., see Barth, T.F.W. et al. Country, G., 498, 509 Couture, R.A., 138, 204 Crenshaw, J.L., see Allen, E.T. et al. Crook, A.W., see Packham, G.H. and Crook, A.W. Cross, H., see Yariv, S. and Cross, H. Crowley, M.S. and Roy, R., 194, 204 Curtis, C.D. and Spears, D.A., 176,204 Curtis, D.M., 2, 14, 100 Daly, R.A., 75, 100 Damuth, J.E. and Fairbridge, R.W., 58, 101 Dangeard, L. and Rioult, M., 68, 101 Dapples, E.C., 36, 37, 39, 41, 73, 86, 87, 101, 180, 204 Dasch, E.J., see Biscaye, P.E. and Dasch, E.J. Davidson, C.F., 95, 101 Davies, D.K., see Almon, W.R. et al. Davies, T.A. and Supko, P.R., 2, 3, 14, 82, 101, 214, 241, 283 Deans, T., 38, 101 Debelmas, J., 68, 101 Debelmas, J., see Chateauneuf, J.-J. et al. Debyser, J., 37, 101 Deffeyes, K.S., Lucia, F.J. and Weyl, P.K., 79, I01 Deflandre, G., 68, 101 Defossez, M., see Millot, G. et al. Degens, E.T., 37, 101 Degens, E.T. and Epstein, S., 73, 81, 101 Degens, E.T. and Mopper, K., 124. Degens, E.T. and Ross, D.A., 37, 39, 59, 101 Deike, R.G., see Hanshaw, B.B. et al. DeMayo, B., see Eslinger, E. et al. Demolon, A. and Boischot, P., 90, 101 De Roever, E.W.F., 308, 395, 477 De Roever, W.P., 390, 391, 392, 477 Derov, G., see Tissot, B. et al. Devereux, I., 45 I , 477 Deverin, L., 295, 477, 497, 509 Dibbel, W.E., see Seyfried, W.E. et al. Dickinson, W.R., 403, 404, 431, 461, 477 Dickinson, W.R., Ojakangas, R.W. and Stewart, R.J., 403, 436, 455, 463, 477 Dimroth, E. and Chauvel, J.J., 496, 509

547 Dimroth, E. and Kmberley, M.M., 94, 101 Dixon, E.E.L., 80, 101 Dolmatova, T.V., see Shutov, V.D. and Dolmatova, T.V. Douglas, R.G. and Savin, S.M., 259, 283 Douglas, R.G., see Matter, S.M. and Douglas, R.G. Douglas, R.G., see Savin, S.M. and Douglas, R.G. Douglas, R.G., see Schlanger, S.O. et al. Douglas, R.G., see Schlanger, S.O. and Douglas, R.G. Drever, J.I., 53, 54, 5 5 , 66, 101, 133, 159, 205 Drever, J.I., see Lawrence, J.R. et al. Drits, V.A. and Shutov, V.D., 338, 477 Drits, V.A., see Frank-Kamenetsky, V.A. et al. Drits, V.A., see Kossovskaya, A.G. and Drits, V.A. Drits, V.A., see Kossovskaya, A.G. et al. Drits, V.A., see Shutov, V.D. et al. Droste, J., 135, 205 Dubar, G.P., see Zaporozhtseva, A.S. et al. Dunharn, K.C., 39, 101,498, 509 Dunharn, R.J., 213, 243, 284 Dunnington, H.V., 255, 284 Dunoyer de Segonzac, G., 18, 101, 115, 116. 134, 135, 170, 171, 174, 176, 177, 205, 294, 306, 311, 312, 313, 314. 319. 324, 325, 327. 330, 334, 335, 337, 340, 343. 347, 348, 350. 35 I , 354, 355, 356. 357. 358, 359, 360, 364. 366, 367, 368, 373, 395, 468, 471, 477 Dunoyer de Segonzac, G., Artru, P. and Ferrero, J., 312, 356, 373, 478 Dunoyer de Segonzac, G. and Chamley, H., 360, 478 Dunoyer de Segonzac, G., Ferrero, J. and Kubler, B., 306, 346, 348, 355, 356, 377. 478 Dunoyer de Segonzac, G. and Heddebaut, C.. 348, 352, 354, 357, 358, 359, 360, 364. 367. 368. 370, 478 Dunoyer de Segonzac, G. and Hickel, D., 343. 344, 354. 468, 472, 478 Dunoyer de Segonzac, G. and Millot, G., 340, 359. 478 Dunoyer de Segonzac, G., see Artru, Ph. et al. Dunoyer de Segonzac, G., see Chennaux, G. et al. Dunoyer de Segonzac, G., see Chamley, H. et al. Dunoyer de Segonzac. G., see Chennaux. G. and Dunoyer de Segonzac, G. Dunoyer de Segonzac, G., see Durand, B. et al.

Dunoyer de Segonzac, G., see Sagon, J-P. and Dunoyer de Segonzac, G. Dupont, B., see Lemoalle, J. and Dupont. B. Durand, B., Dunoyer de Segonzac. G., Albrecht, P. and Van den Broueke, M., 182, 205 Durney, D., 371. 466, 468, 478 Dury, G.H., 88, 101 Dyni, J.R., 137. 205

.

Earley, J.W., see Milner, I.H. and Earley, J.W. Eberl, D., 174, 180, 205 Eberl, D. and Hower, J., 172. 173. 174, 205 Eckhardt, F.-J. 177, 178, 205, 324. 336, 338. 368, 478 Eckhardt, F.-J. and Von Gaertner, H.R., 336, 4 78 Edmunds, W.M., Lovelock, P.E.R. and Gray, D.A., 275, 284 Edwards, A.B. and Baker, G., 44. 61. 101 Eisbacher, G.H., see Read, P.B. and Eisbacher. G.H. Eitel. W., 49. 53, 102 Ekdale, A.A. and Berger. W.H.. 239, 240, 284 Ekdale, A.A., see Berger, W.H. et al. Elderfield, H., 150, 159, 205 Ellis, A.J., see Browne, P.R.L. and Ellis, A.J. Ellis, A.J., see Coombs, D.S. et a]. ElverhB, H., 152, 205 Emery, K.O. and Rittenberg. S.C.. 118. 150. 151, 205 Emery, K.O., see Revelle, R. and Emery, K.O. Enos, P., see Cook, H.E. and Enos, P. Epprecht, W., 509 Epstein, S., see Degens, E.T. and Epstein, S. Eremeev, V.V., see Timofeev, P.P. et al. Ergun. S.. see Mentser, M. et al. Erhart, H., 44. 59, 66. 68, 93, 95, 102 Erlank, A.J.. see Hart, S.R. et al. Ernst. W.G., 343, 457. 478 Ernst, W.G. and Seki. Y . , 457, 478 Ernst, W.G., Seki, Y., Onuki, H. and Gilbert, M.C., 343, 457. 467, 468, 478 Ernst, W.G., see Seki, Y. et al. Esch, H., 334, 478 Eskola, P., 302, 303, 380, 402, 478 Eskola, P., see Barth, T.F.W. et al. Eslinger, E., Highsmith, P., Albers, D. and DeMayo, B., 166, 205 Eslinger, E.V. and Savin, S.M., 321, 314, 329, 368, 478 Eslinger, E.V., see Hecht, A.D. et al. Eslinger, E.V., see Hower, J. et al. Esquevin, J., 343, 352, 354, 355, 478

Eugster. H.P., 137. 205 Eugster. H.P. and Hardie. L.A.. 135. 136. 137, 205

Eugster. H.P.. see Surdam. R.C. and Eugster, H.P. Eugster. H.P.. see Wise, S.W. and Eugster, H.P. Eugster. H.P., see Yoder, H.S. and Eugster. H.P. Evans. R., see Kirkland, D.W. and Evans. R. Fabre, J.. 471. 478 Fairbairn. H.W., 89. 102 Fairbairn, H.W.. see Hurley. P.M. et al. Fairbridge, R.W.. 1. 4, 14, 26, 30. 33. 42, 50, 51, 57, 59, 65. 68, 71, 73, 78, 79, 80, 81, 82, 83, 84. 88, 95. 96, 97, 102 Fairbridge. R.W. and Bourgeois. J.. 2, 14. 26, 59. 102 Fairbridge. R.W. and Finkel. C.R.. 102 Fairbridge, R.W. and Teichert, C., 68. 102 Fairbridge, R.W., see Damuth, J.E. and Fairbridge, R.W. Farrell, B.E., see Hiltabrand, R.R. et al. Faust, G.T.. 85. 102 Faust. L.Y., see Weatherby, B.B. and Faust, L.Y. Feely, H.W. and Kulp. J.L., 62, 102 Ferrero. J. and Kubler. B., 337, 479 Ferrero. J., see Dunoyer d e Segonzac G. et al. Fersman. A.E.. 18, 35, 97, 103 Feys. R.. 456, 479 Feys. R.. see Chateauneuf, J.-J. et al. Finch, J.W., see Ramsay, A.T.S. et al. Finkel, C.W.. see Fairbridge, R.W. and Finkel, C.W. Finlayson, J.B., see Mahon. W.A.J. and Finlayson, J.B. Fisher. A.G. and Garrison. R.E., 213, 284 Fischer. A.G.. Honjo, S. and Garrison. R.E.. 223. 256. 257. 268, 271, 274, 280, 284 Fischer. A.G.. see Beall, A.O. and Fischer, A.G. Fischer, A.G.. see Newell. N.D. et al. Flach, K.W.. Cady. J.G. and Nettkton. W.D., 44. 103 Flehmig. W.. 347. 350. 479 Flexer. A,. see Nathan, Y. and Flexer, A. Folk. R.L. and Siedlecka. A,. 71, 103 Folk, R.W. and McBride. E.F.. 67. 68, 103 Fondeur. C. et al.. 80, 103 Forniosova. L.B.. 498. 510 Fdrstnrr. U.. see Milliman, J.D. et al. Fiirstner. U.. see Miiller. G . and Forstner. U.

Foscolos, A.E. and Kodama. H.. 162. 165, 166. 167, 175, 205, 465, 479 Foscolos, A.E., Powell, T.G. and Gunther, P.P.. 182, 183, 205, 465, 479 Foscolos, A.E.. see Powell, T.G. et al. Foster, S.S.D.. 275, 284 Frankel, J.J. and Kent, L.E., 103 Frank-Kamenetsky. V.A., Logvinenko, N.V. and Drits, V.A., 337, 479 Frey. M., 177, 201, 205, 306, 343, 344, 348. 352, 358, 359, 360, 361, 364, 365, 370, 371. 372, 373, 468, 469. 470. 479 Frey, M., Hunziker, J.C. Roggwiller, P. and Schindler, C., 359, 371. 372, 455, 468. 469, 470, 479 Frey, M. and Niggli. E., 466, 469, 470, 479 Friedlander, U.. see Kolodny. J. et al. Friedman, G.M.. 213, 284 Friedman. G.M. and Sanders, J.E., 2, 14. 49, 103, 115, 205 Friedman, G.M., see Ali, S.A. and Friedman. G.M. Friedman, G.H., see Bryce, M.R. and Friedman, G.H. Friedman, G.M., see Gervitz, J.H. and Friedman, G.M. Fritz, P., see Walls, R.W. et a]. Fuchtbauer, H., 192. Fuchtbauer, H. and Goldschmidt, H.. 179, 205, 324, 334, 337, 479 Fiichtbauer, H. and Reineck, H.E., 127, 206 Fullerton, L.B., see Alrnon, W.R. et al. Funnel, B.M. and Rxdel, W.R.. 214, 284 Furbish, W.J., 461, 479 Fiirst, I., see Szadecky-Kardoss, E. et al. Fyfe, W.S., 406, 407, 414, 479 Fyfe, W.S. and Turner, F.J., 303, 380, 392, 479 Fyfe, W.S., Turner, F.J. and Verhoogen, J., 17, 53, 103, 299, 302, 379, 381, 386, 415. 479 Fyfe, W.S., see Campbell, A.S. and Fyfe, W.S. Fyfe, W.S., see Coombs, D.S. et al.

'

Gaarder, K.R., 219, 284 Gaida. K.H., see Von Engelhardt, W. and Gaida, K.H. Galehouse, J.S., see Houghton, R.L. et al. Galikeev, K.Kh., see Sakhibgareev, R.S. and Galikeev, K.Kh. Ganguly. J.. 370, 479 Gardner. J.V.. 235. 238, 284 Gardner, J.V.. see Piper, D.Z. et al. Garmon, L.B., see Hecht, A.D. et al. Carrels, R.M., 27, 29, 38, 103, 501, 502, 510 Carrels, R.M. and Christ, C.L., 23, 47, 103, 334, 479

549 Carrels, R.M. and Mackenzie, F.T., 48, 103 Carrels, R.M., Thompson, M.E. and Siever, R., 103 Carrels, R.M., see Castaiio, J.R. and Garrels. R.M. Garrels, R.M.. see Krumbein. W.C. and Garrels, R.M. Garrels, R.M., see Mackenzie, F.T. and Garrels, R.M. Garrison, R.E., see Fischer, A.G. and Garrison, R.E. Garrison, R.E., see Fischer, A.G. et al. Garrison, R.E., see Kennedy, W.J. and Garrison, R.E. Garrison, R.E., see Pimm, A.C. el al. Gauthier, J., see Artru. Ph. and Gauthier. J. Gavish, E. and Reynolds, R.C.. 330, 479 Gavrilov, A.A. and Aleksandrova, V.A., 327. 343. 480 Gealy, E.L., 241, 253, 254, 284 Geehan, G.W., see Adelseck, C.G. et al. Geikie, A,, 96, 103 Gervirtz, J.L. and Friedman. G.M., 213. 284 Gidon, P.. 68, 103 Giesker, J.M., 156, 206 Giesker, J.M.. see Kastner. M. and Giesker, J.M. Giesker, J.M., see Kastner. M. et al. Giesker, J.M., see Perry, E. et al. Gignoux, M. and Avnimelech, M.. 69. 103 Gilbert, C.M., see Williams, H. et al. Gilbert, M.C., see Ernst, W.G. et al. Gill, E.D., 455, 480 Ginsburg, R.N., 213, 284 Ginsburg, R.N., Shinn, E.A. and Schroder, J.H., 213, 284 Ginsburg, R.N., see Shinn, E.A. el al. Ginter, R.L., 61, 103 Giraud, A,, see Artru, Ph. et al. Glasby, G.P., 150, 206 Glassley. W., 439. 480 Glover, J.E., 41, 103 Glushinskiy, P.I., see Zaporozhtseva, A.S. et al. Godfriaux, L., see Beugniet, A. et al. Goldberg, E.D. and Arrhenius, G.O.S.. 59. I03 Goldberg, E.D., see Griffin, J.J. et al. Goldberg, E.D., see Peterson. M.N. and Goldberg, E.D. Goldberg, E.D., see Rex. R.W. and Goldberg, E.D. Goldenburg, L., see Magaritz, M. et al. Goldich, S.S., 75, 103

Goldschmidt. H.. see Fuchtbauer. H. and Goldschmidt. H. Goldschmidt, V.M.. 23. 103 Goldsmith, J.R. and Graf. D.L.. I03 Goldsmith. J.R.. see Graf. D.L. and Goldsmith, J.R. Goodwin. J.H. and Surdam. R.L.. 419. 480 Gorbunova. Z.N.. see Rateev. M.A. et al. Goto. K.. see Okamotu. G. et al. Grabau. A.W., 19. 32. 65. 103 Gradusov. B.P., 337. 480 Graf. D.L. and Goldsmith. J.R.. 78. 103 Graf, D.L. see Goldsmith. J.R. and Graf, D.L. Graham, E.R.. 49. 103 Gray. D.A.. see Edmunds. W.M. et al. Greggs. R.G.. see Bartlett. G.A. and Greggs. R.G. Griffin, G.M. and Ingram. R.L.. 55. 104 Griffin. J.J., Windom. H. and Goldberg. E.D.. 54, 55. 57. 58, 104. 119. 206 Grim, R.E.. 54, 58. l a 4 Grim. R.E.. Bray, R.H. and Bradley. W.F.. 54. I04 Grimmelfarb. B.M., 90, 103 Gross. G.A., 496. 510 Grubb. P.L.C.. 496. 510 Gruner. J.W. and Thiel. G.A.. 75, 104 Gude. A.J.. see Sheppard. R.A. and Gude, A.J. Guitard. G. and Saliot. P., 393. 395, 471. 480 Gunatilaka. A,, Saleh. A. and Al-Termeeni. A,. 64. 104 Gunther, P.R., see Foscolos. A.E. et al. Gunther, P.R.. see Powell, T.G. et al. Gurewitsch. A.B. and Toporez. G.A., 335. 480 Haas. H. and Holdaway, M.J.. 362. 480 Hallam. A. and Bradshaw, M.J., 94. 104 Hallam, A. and Snellwood. B.W.. 420, 458, 480 Halley. R.B.. see Shinn, E.A. et al. Hallimond. A.F., 498. 510 Hamilton, E.L.. 118, 125, 206, 214, 226 232. 241, 251, 253, 284 Hamilton. E.L., Shumway. G., Menard, 1.W. and Shipek, C.J.. 235, 241, 284 Hamilton, E.L., see Johnson, T.C. et al. Hamilton, J.D.. 315. 316. 317. 318, 326, 327, 480 Han, T.M., see Anderson, G.T. and Han, T.M. Hancock, J.M., 216, 284 Hancock, J.M. and Scholle, P.A., 266, 277, 284 Hanshaw, B.B.. Back, W. and Deike, R.G., 85, I 04

Harder, H., 131, 206, 498, 510 Hardie, L.A., see Eugster, H.P. and Hardie, L.A. Hardy, R.G., see Jenkyns, H.J. and Hardy, R.G. Harms. J.E., Whitehead, T.H. and Heaton, J.B., 510 Harrassowitz, H., 306, 347, 480 Harrison, A.G., see Thode, H.G. et al. Harrison, S.M., see Be, A.W.H. et al. Harrison, W.E., 182, 206 Hart, R.A., 57, 104 Hart, S.R., Erlank, A.J. and Kable, E.J.D., I04 Hashimoto, M., 377, 379, 388, 390, 391, 393, 394, 403, 429, 430, 443, 447, 449, 450, 454, 464, 480 Hashimoto, M. and Kashima, N., 393, 394, 464, 480 Hatashi, H., see Sudo, T. et al. Hatch, F.H., Rastau, R.H. and Black, M., 84, 93, 94, 104 Haussiihl, S. and Miiller, G., 152, 206 Hawkins, J.W., 403, 468, 480 Hay, R.L., 53, 75, 104, 417,419, 420,421,458,

480 Hay, R.L. and lijima, A,, 430, 480 Hay, R.L. see Iijima, A. and Hay, R.L. Hay, W.W., Southain, J.R. and Noel, M.R., 216, 284 Hayes, A.O., 93, 104 Hayes, J.B., 337, 339, 480 Hayes, J.B., see Mossler, J.H. and Hayes, J.B. Hayes, J.B., see Schroeder, R.J. and Hayes, J.B. Heald, M.T., 89, 104 Heath, G.R., 40, 104 Heath, G.R., see Van Andel, T.H. et al. Heaton, J.B., see Harms, J.E. et al. Hecht, A.D., Eslinger, E.V. and Garmon, L.B., 225, 238, 285 Hecht, F., 37, 104 Hedberg, H.D., 190, 193, 206 Heddebaut, C., see Dunoyer de Segonzac, G. and Heddebaut, C. Heezen, B.C., Nesteroff, W.D. and Sabatier, G., 58. 104 Heezen, B.C., see Hurley, P.M. et al. Hein, J.R. and Scholl, D.W., 145, 206 Hein, J.R., Scholl, D.W., Barron, J.A., Jones, M.G. and Miller, J., 146, 147, 206 Heling. D.. 161. 170, 172, 174, 189, 191, 200, 206 Heling, D. and Teichmiiller, M., 182, 206

Heller-Kallai, L., 166, 206 Heller-Kallai, L., Nathan, Y. and Zak, I., 135, 206 Heller-Kallai, L. and Rozensen, I., 133, 159, 206 Hemley, J.J. and Jones, W.R., 362, 480 Hemley, J.J., see Reed, B.L. and Hemley, J.J. Henderson, G.V., 318, 352, 366, 368, 369, 480 Henin, S., see Caillere, S. et al. Herman, Y., 67, 104 Hickel, D., see Dunoyer de Segonzac, G. and Hickel, D. Hickox, J.E., see Newell, N.D. et al. High, L.R. and Picard, M.D., 458, 481 Highsmith, P.. see Eslinger, E. et al. Hilmy, M.E., see West, I.A. et al. Hiltabrand, R.R., Farrell, B.E. and Billings, G.K., 315, 481 Hinrichsen, Th. and Schiirmann, K., 429, 450, 45 1, 481

Hite, R.J., 83, 104 Hoare, J.M., Condon, W.H: and Patton, W.W., 417, 436, 463, 481

Hoehne, K., 86, 104 Holdaway, M.J., see Haas, H. and Holdaway, M.J. Holdship, S.,see Stoffers, P. and Holdship, S. Holland, H.D., 38, 48, 62, 104 Holliday, D.W., 64, 105 Holser, W.T., 38, 77, 105 Holser, W.T. and Kaplan, J.R., 105 Honess, A.P. and Jeffries, C.D. 75, 105 Honjo, S., 214, 225, 235, 236, 285 Honjo, S., see Fischer, A.G. et al. Honjo, S., see Okada, H. and Honjo, S. Honnorez, J., 141, 144, 206 Horodyski, R.J., see Coombs, D.S. et al. Horstman, E.L., 75, 105 Hoschek, G., 370, 481 Hoss, H., 67, 105 Hosterman, J.W., Wood, G.H. and Bergin, M.G., 343, 361, 364, 373, 481 Hough, J.L., 95, 105,510 Houghton, R.L., Rothe, P. and Galehouse, J.S., 140, 141, 206 Houtz, R. and Swing, J., 105 Hower, J., 139, 207, 496, 510 Hower, J., Eslinger, E.V., Hower, M.E. and Perry, E.A., 161, 162, 163, 164, 168, 169, 175, 177, 178,207, 320, 321, 328, 481

Hower, J. and Mowatt, T.C., 170, 207, 329, 332, 333, 481 Hower, J., see Aronson, J.L. and Hower, J. Hower, J., see Eberl, D. and Hower, J.

55 1 Hower, J., see Maxwell, D.T. and Hower, J. Hower, J., see Perry, E. and Hower, J. Hower, J., see Reynolds, R.C. and Hower, J. Hower, J., see Thompson, G.R. and Hower, J. Hower, J., see Velde, B. and Hower, J. Hower, M.E., see Hower, J. et al. Hoyer, P., Clausen, C.-D., Leuteritz, K., Teichmiiller, R. and Thome, K.N., 465, 481 Hsii, K.J. and Jenkyns, H.C., 214, 285 Hsii, K.J., see Lloyd, R.M. and Hsii, K.J. Hubert, J.F.,see Suchecki, R.K. et al. Hudson, J.D., 152, 154, 207 Hudson, J.H., see Shinn, E.A. et al. Hull, A.B., see Cody, R.D. and Hull, A.B. Hummel, K., 50, 96, 105, 116, 207 Hunt, J.M., 38, 105 Hunt, T.S., 79, 105 Hunter, R.E., 94, 105 Hunziker, J.C., see Frey, M. et al. Hurley, P.M., Heezen, B.C., Pinson. W.H. and Fairbairn, H.W., 5 5 , 105 Hutton, C.O., 385, 386, 481 Hutton, C.O. and Turner, F.J., 379, 386, 464, 481 Hutton, J., 18, 105 Iijima, A., 143, 207, 420, 458, 481 Iijima, A. and Hay, R.L., 419, 425, 481 Iijima, A. and Utada, M., 324, 383, 391, 397, 398, 399, 401, 421, 458, 459, 481 Iijima, A., see Hay, R.L. and Iijima, A. Illing, L.V., 82, 105 Illing, L.V., Wells, A.J. and Taylor, J.C.M., 79, 105 Ingram, R.L., see Griffin, G.M. and Ingram, R.L. Irving, A.. 60, 77, 105. Isphording, W.C., 153, 207 Ivanova, N.V., see Volkova, A.N. et al. Jaanusson, V., 73, 105 Jackson, J.A., see Bates, R.L. and Jackson, J.A. Jackson, T.A., 183. 207 Jacobs. M.B., 50, 105 James, H.L., 495, 496, 510 James, H.L. and Sims, P.K.. 94, I05 James, N.P., see Schlager, W. and James, N.P. Janecke, E., 32, 105 Jaron, M.G., 340. 481 Jeans, C.V., 134, 207 Jeffries, C.D., see Honess. A.P. and Jeffries. C.D.

Jenkyns, H.C., see Hsii, K.J. and Jenkyns, H.C. Jenkyns, H.J. and Hardy, R.G., 145, 207 Jipa, D., see Manheim, F. et al. Johns, W.D. and Shimoyama, A., 320, 481 Johns, W.D., see Berry, R.W. and Johns, W.D. Johnson, T.C., Hamilton, E.L. and Berger. W.H., 226, 232, 233, 234, 235, 285 Johnston, J.J., see Allen, E.T. et al. Jolly, W.T., 383, 388, 403. 410. 417, 482 Jonas, E.C., 174, 175, 207 Jones, J.B. and Segmit, E.R., 69, 105, 145, 207 Jones, M.G., see Hein, J.R. et al. Jones, W.R., see Hemley, J.J. and Jones, W.R. Kable, E.J.D., see Hart, S.R. et al. Kafri, U., see Magaritz, M. et al. Kalkdwsky, E., 21, 105 Kapland, I.R., see Baas Becking, L.G.M. et al. Kaplan, J.R., see Holser, W.T. and Kaplan, J.R. Kaplan, M.Ye., 455, 458, 462. 482 Karpova, G.V., 321, 326, 331. 332, 333, 337, 339, 342, 343, 364. 377. 482 Karpova. G.V.. Lunkin, A.E. and Shevyakova. E.P., 335, 483 Karpova, G.V. and Shevyakova. E.P.. 334, 335, 336, 482 Karpova, G.V. and Timofeeva, Z.V., 3 12, 336, 482 Karpova. G.V., see Logvinenko. N.V. and Karpova, G.V. Kasanskiy, J.P., 360, 369, 482 Kashima. N.. see Hashimoto, M. and Kashima, N. Kastner, M., 74, 105, 144, 180, 207 Kastner. M. and Giesker, J.M., 158, 180, 207 Kastner, M., Keene, J.B. and Giesker, J.M., 146. 207 Kastner. M. and Stonecipher, S.A., 140, 207 Kastner, M., see Siever, R. and Kastner. M. Kato, K., 131, 207 Kawachi, Y.. 394, 447, 464. 468, 472, 482 Kay, M., 42, 105 Kazakov, A.V.. 91, 496, 510 Kazakov, A.V., Tikhomirova, M.M. and Plotnikova, V.J., 510 Kazama, T., see Aoyagi, K. and Kazama, T. Keene, J.B., see Kastner, M. et al. Keller, W D., 56, 106. 200, 207. 325, 331, 482 Keller, W.D., Balgord, W.D. and Reesman. A.L.. 48, 106 Keller, W.D., see Ojakangas, R.W. and Keller, W.D.

552 Kelley. W.P.. 49. 56. 106 Kelts. K.R., see Wise, S.W. and Kelts. K.R. Kennedy, W.J. and Garrison. R.E., 40. 106. 226. 285

Kennedy, W.J., see Scholle. P.A. and Kennedy, W.J. Kent, L.E., see Frankel, J.J. and Kent, L.E. Kerr. P.F., see Ross. C.S. and Kerr. P.F. Kerrick. D.M., 434. 435, 482 Kessel. P.. 106. 294, 482 Khitarov. N.1. and Pugin. V.A.. 314. 482 Khouri, J.. see Shearman. D.J. et al. Kienast. J.R. and Velde, B., 457. 482 Kimberley. M.M., 94, 106 Kimberley, M.M., see Dimroth. E. and Kimberley, M.M. Kinsman, D.J.J., 258? 285 Kirkland, D.W. and Evans, R., 79, 106 Kisch. H.J.. 8. 9, 10, I I , 307, 330, 335. 348. 351, 357. 359, 360, 371. 373, 377. 378,454. 455. 456. 458, 459, 461, 465, 466, 467. 468, 469. 472, 482, 483 Kliburszky. B.. see Szadecky-Kardoss, E. et al. Knight, L.. see Chilingar, G.V. and Knight, L. Knox. R.W.O’B., 94. 106 Kodama. H., 366. 483 Kodama, H., see Foscolos, A.E. and Kodama. H. Kohler, E.E. and Koster, M., 140, 207 Koizumi. M., see Nakajima, W. and Koizumi, M. Kolbe, H.. 498. 510 Kolla. V. and Biscaye, P.E., 122, 207 Kolodny. J., Taraboulos, A. and Friedlander. U.. 71. 106 Kolodny, Y.,see Steinitz, G. and Kolodny. Y. Komarova, G.V.. see Lisitsyn. A.K. et al. Kondrat’eva. J.A., see Lisitsyn, A.K. et al. Konyukhov. A.I., see Teodorovich, G.I. and Konyukhov. A.I. Konysheva, R.A., 343. 483 Kopeliovich. A.V.. Kossovskaya, A.G. and Shutov, V.D., 339, 483 Koporulin. V.I.. 324, 326. 339, 340, 459. 462, 4x3

Kornprobst, J., see Velde. B. and Kornprobst, 1.

Kossovskaya. A.G.. 321, 340. 483 Kossovskaya. A.G. and Drits, V.A., 33 1, 332, 343. 484 Kossovskaya, A.G., Drits, V.A. a n d Alexandrova. V.A.. 179.208, 334, 336,340, 4x4

Kossovskaya. A.G.. Logvinenko. N.V. and Shutov. V.D.. 301. 302. 340. 484

Kossovskaya, A.G. and Shutov, V.D., 1 I. 14, 294, 296. 301. 302. 303, 304, 307, 309. 336, 337, 340, 342, 360, 375, 316, 377. 378, 458. 464. 465, 466, 467, 474, 484 Kossovskaya, A.G., Shutov, V.D. and Alexandrova, V.A., 319, 340, 484 Kossovskaya, A.G., Shutov, V.D. and Drits. V.A., 334, 340, 484 Kossovskaya, A.G., see Kopeliovich, A.V. et al. Koster, M.. see Kohler, E.E. and Kuster, M. Kotel’nikov. D.D., see Teodorovich. G.l. et al. Kramm, U., 362, 370, 484 Krauskopf, K.B., 49, 86, 106 Kraut. F.. see Caillere, S. and Kraut, F. Krejci-Graf, K., 77, 106 Krinsley, D., see McCoy, F. et al. Kromer, H., 178, 208 Kromer, H., see Von Engelhardt, W. et al. Krotov, B.P., 496, 498, 510 Krumbein, W.C.. 2, 14. 46. 106 Krurnbein, W.C. and Carrel’s, R.M., 28, 106 Krumbein, W.C. and Sloss, L.L., 294, 484 Kubler. B., 9, 1 I , 14, 306, 307, 308, 324, 325, 342, 345, 346, 347, 348, 350, 351, 352, 356, 358, 359, 360, 369, 375, 316, 377, 395, 451, 465, 466, 461, 469, 470, 472, 413, 484 Kubler, B., Martini, J. and Vaugnat, M., 14, 324, 431, 459, 461, 466, 469, 484 Kubler, B., see Dunoyer d e Segonzac G. et al. Kubler, B., see Ferrero, J. and Kubler, B. Kulbicki, G. and Millot, G., 334, 484 Kulbicki, G., see Beuf. S. et al. Kulke. H., 338, 484 Kulp, J.L., see Auk, W.U. and Kulp, J.L. Kutp, J.L., see Feely, H.W. and Kulp. J.L. Kuriyagawa, S., see Matsuda, T . and Kuriyagawa, S. Kusuda, T., see Shimazu, M.et al. Kvenvalden, K.A., see Oehler, L. et al. Lambe, T.W., 198, 208 Lancelot, Y., 146, 208 Lancelot, Y.,see Schlanger, S.O. et al. Land, L.S.. 85. 106 Landis. C.A., 388, 394, 447, 484 Landis, C.A. and Bishop, D.G., 386, 387, 443, 484

Landis, C.A. and Coombs, D.S., 386. 390,443, 485

Landis, C.A. and Rogers, J., 451. 485 Lane, A.C., 32, 41, 106 Langford-Smith, T., 45, 88, 106 Latimer, W.M., 27, 106 Laughton, A.S., 253, 285

553 Lawrence, J.R., Drever, J.I., Anderson, T.F. and Brueckner, H.K., 145, 208 Lawrence, J.R., see Perry, E. et al. Le Corre, C., 370, 485 Lees, G.M., 69, 106 Leith, C.K. and Mead, W.J., 96, 106 Lemaitre, H., 68, 106 Lemoalle, J. and Dupont, B., 510 Lemoine, M., see Chateauneuf, J.-J. et al. Lesley, 80 Leuteritz, K., see Hoyer, P. et al. Levi, B., 405, 485 Lewin, J.C., 66, 106 Leyell, C., 96, 107 Liborio, G. and Mottana, A,, 344, 485 Lidz, B.H., see Shinn, E.A. et al. Linck, G., 107 Lindgren, W., 77, 510 Liou, J.G., 382, 385, 405, 408, 410, 411. 413. 414, 416, 419,424, 425, 427, 485 Lippman, F., 153, 154, 208, 324, 485 Lisitzin, A.P.. 214, 216, 217, 218, 223, 285 Lisitzin, A.P., Kondrat’eva, J.A. a n d Komarova, G.V.. 334. 485 Lisitzyn, A.P., see Rateev, M.A. et al. Liss, P.S., see Burton, J.D. and Liss. P.S. Lister, J.S., see Brown, L.S. et al. Lloyd, R.M. and Hsii, K.J.. 259. 285 Lloyd. R.M., see Shinn, E.A. et al. Logvinenko, N.V., 327, 368, 485 Logvinenko, N.V. and Karpova, G.V., 332, 485

Logvinenko, N.V., see Frank-Kamenetsky, V.A. et al. Logvinenko, N.V., see Kossovskaya. A.G. et al. Lomtadze, V.D., 117. 208 Long, G. and Neglia, S., 315, 325, 334, 485 Losievskaya, S.A., see Shutov, V.D. et al. Loughnan, F.C., 153, 208 Lounsbury, R.W. and Melhorn. W.N., 327, 485

Love. L.G., 61, 107. 151,208. 510 Lovelock, P.E.R., see Edmunds, W.M. et al. Lowell, R.P., see Rona, P.A. and Lowell, R.P. Lucas, J., 81, 107, 459, 485 Lucas, J. and Ataman, G., 134, 208. 324, 485 Lucas. J., see Cassan, J.P. and Lucas, J. Lucas, J., see Clauer, N. and Lucas. J. Lucia, F.J., see Deffeyes, K.S. et al. Ludwig, V., 343, 350, 351, 352, 360, 365, 485 Lukin. A.E.. see Karpova. G.V. et al. MacDonald. G.J.F., 64, 107

MacFadyen, W.A., 63. 107 Mackenzie, F.T. and Carrels, R.M., 51, 67. I07 Mackenzie, F.T., see Carrels, R.M. and Mackenzie, F.T. Magara, K., 195, 208 Magaritz, M., Goldenburg. L.. Kafri. U. and Arad, A,. 85, 107 Mahon, W.A.J. and Finlayson, J.B.. 440, 485 Maiklem, W.R. and Campbell, F.A., 315. 318, 486

Malahoff. A., see Anderson, N.R. and Malahoff, A. Mangelsdorf, P.C. see Sayles. F.C. and Mangelsdorf, P.C. Manghnani. M.H., Schlanger, S.O. and Milholland, P.D., 232, 254, 258. 259, 262. 263, 275, 285 Manghnani, M.H., see Milholland, P.D. et al. Manheim, F., Rowe, G.T. and Jipa, D.. 90. 107

Manheim, F.T.. see Sayles. F.C. and Manheim, F.T. Mapstone, N.B., 276, 285 Marshall, C.E. and Upchurch, W.J., 49, 107 Martini, J.. 371, 403, 437, 455, 459, 463. 464. 468, 486 Martini, J. and Vaugnat, M.. 371, 388, 406, 414, 430, 468, 469. 471, 486 Martini, J., see Kubler. B. et al. Mathieu-Sicaud. A,, see Caillere, S. et al. Matsuda, T. and Kuriyagawa, S., 388, 391, 430, 449. 460. 46 1, 486 Matsuda, T., see Seki. Y.et al. Matter, A., 172, 208. 226, 243. 247, 255, 256, 258, 285 Matter, A,, Douglas, R.G. and Perch-Nielsen, K., 213, 214. 241, 242, 244, 245, 246, 247, 248, 251, 254. 255, 257, 258, 259. 260, 265, 277, 285 Mawson, D., 78, 107 Maxey, G.B., see Bredehoeft, J.D. et al. Maxwell, D.T. and Hower, J., 329, 343, 486 Mayer, L.A., 232, 243. 276, 278. 285 Maynard, J.E., see Moore, E.S. and Maynard. J.E. McBride, E.F., 67, 107 McBride, E.F., see Folk, R.W. and McBride, E.F. McCarter, R.W.. see Whitehouse, U.G. and McCarter, R.W. McCoy, F., Zimmerman, H. and Krinsley. D., 141, 208 McElhinny, M.W., see Cook, P.J. and McElhinny. M.W.

554 McEwan. D.M.C., 315, 486 McIntyre, A. and McIntyre, R., 216, 225, 235, 236, 285 McIntyre. R. see McIntyre, A. and McIntyre, R. McKee. E.H.. see Robinson, P.T. et al. McKelvey. V.E.. Swanson. R.W. and Sheldon, R.P., 93 107 McKinstry, H.E., 510 McNamara. M., 439. 486 Mead, R.H.. 126, 127, 194, 197, 198, 199. 208 Mead. W.J.. see Leith, C.K. and Mead, W.J. Meisl, S.. 468. 486 Melhorn. W.N., see Lounsbury, R.W. and Melhorn, W.N. Melieres. F.. see Chamley, H. et al. Mellon, G.B.. 462, 486 Mellon. G.B.. see Carrigy. M.A. and Mellon, G.B. Menard. H.W.. see Hamilton, E.L. et al. Mennig. J.J. and Vatan. A,. 82, 107 Mentser. M., O’Donnell, H.J. and Ergun, S., 456, 486 Meyer. W.J.. 285 Middleton. G., see Blatt, H. et al. Mikami, K.. see Seki, Y.et al. Milholland. P.D., M a n g h n a n i , M.H., Schlanger. S.O. and Sutton, G.H.. 235, 251. 252, 275, 278. 285 Milholland. P.D.. see Manghnani, M.H. et al. Miller. J., see Hein. J.R. et al. Milliman. J.D., 4, 15, 213, 214. 216, 218, 219. 220, 221, 222, 285, 286 Milliman. J.D. and Muller, J.. 214. 286 Milliman. J.D.. Miiller, G. and Forstner, U., 44, 107 Millot. G.. 21, 43, 53, 57, 58, 107, 115, 134, 139, 177. 208, 324, 486 Millot. G., Radier, H., Muller-Feuga, R., Defossez. M. and wey, R., 107 Millot. G., see Dunoyer de Segonzac, G. and Millot, G. Millot. G.. see Kulbicki. G.and Millot, G. Milner. I.H. and Earley, J.W.. 54. 107 Minato, H.. see Utada. M. and Minato, H. Mitchell. J.K.. 126. 197, 208 Miyashiro. A,. 441, 486 Miyashiro. A. and Shido. F.. 383, 486 Mizutani. S.. 66. 107. 147. 208, 383, 421, 486 Moberley, R.. 123, 208 Mohr. P.A.. 510 Moiola. R.J.. 383. 423. 458. 486 Moiola. R.J.. see Robinson, P.T. et al. Monster. J.. see Thode. H.G. et al.

Moon, C.F.. 198, 199, 209 Moore, D., see Baas Becking, L.G.M. and Moore, D. Moore, D., see Baas Becking, L.G.M. et al. Moore, D.G., 268, 286 Moore, E.S. and Maynard, J.E., 93, 95, 107 Moore, G., see Tettenhorst, R. and Moore, G. Moore, T.C.. see Schlanger, S.O. et al. Moore. T.C., see Van Andel, T.H. et al. Moorhouse, F.W., see Orr, A.P. and Moorhouse, F.W. Mopper, K., see Degens, E.T. and Mopper, K. Moretti, F.J., 81, 107 Mori. R., see Seki, Y.et al. Morse, J.W. and Berner, R.A., 225. 286 Morse, J.W. see, Be, A.W.H. et al. Mosebach. R., 61, 107 Moser, R.G., see Spackman, W. and Moser, R.G. Mossler, J.H. and Hayes, J.B.. 327, 331, 486 Mottana, A., see Liborio, G.and Mottana, A. Moulton, G.F., 60, 107 Mountjoy, E.W., see Walls, R.W. et al. Mowatt, T.C., see Hower, J. and Mowatt, T.C. Muffler, L.J.P. and White, D.E., 312, 313, 314, 439, 440, 486 Muir, R.O.. see Borchert, H. and Muir, R.O. Miiller, G., 116, 118, 124, 125, 128, 131, 142, 149, 152, 190, 197, 202, 209, 330, 337, 338, 486 Miiller, G. and Forster, U.,131, 209 Miiller. O., see Haussiihl, S. and Miiller, G. Miiller, G., see Milliman, J.D. et al. Miiller, G., see Von Engelhardt, W. et al. Muller, J., see Milliman, J.D. and Muller, J. Muller-Feuga, R.S., Millot, G. et al. Mullins, H.T., Boardman, M.R. and Neumann. A.C., 213, 214. 226, 235. 243, 253, 286 Murata, K.J. and Whiteley, K.R., 458, 486 Murav’ev, V.I., see Rateev, M.A. and Murav’ev, V.I. Muravjew, W.I. and Salyn, A.L., 338, 487 Murray, J. and Renard, A.F., 51, 52, 59, 108 Murray, R.. see Blatt, H. et al. Murray, R.C., see Pray, L.C. and Murray, R.C. Nagy. B.. 43, 76, 108 Nakajima, W. and Koizumi, M., 424, 425, 487 Nakajima, W. and Tanaka, K.,401. 414, 418, 426, 487 Nakamura, Y,see Coombs, D.S. et al. Nanz, R.H.. 56, 108

555 Nathan. Y. and Flexer, A,, 141, 412, 209 Nathan, Y, see Heller-Kallai, L. et al. Natland. J.H., 145. 209 Naurnann, C.F.. 77, 108 Naylor. R.S., see Coornbs. D.S. et al. Nazarkin, L.A.. 3, 15 Neglia, S., see Long, G. and Neglia, S. Nesteroff, W.D., see Heezen. B.C. et al. Nettkon, W.D., see Flach, K.W. et al. Neugebauer, J., 254, 286 Neurnann, A.C., see Mullins, H.T. et al. Newell, N.D., Fischer, A.G., Whiternan, A.J., Hickox, J.E. and Bradley. J.S.. 65, 108 Newhouse, W.H., 61, 108 Nicolas, A,, see Boudier, F. and Nicolas. A. Niggli, E., see Frey. M. and Niggli. E. Nitsch, K.H.. 408, 410, 429, 430, 450, 451, 452, 454, 487 Noel, M.R., see Hay, W.W. et al. Norin, E., 53, 108, 142. 152, 209 Northrop, J.I.. 68. 108 Norton, M.F.. 57 Nosov, G.L., see Rateev, M.A. et al. Oba, T., see Seki. Y . et al. O’Brian, N.R., 200, 209 Odaka, S.. see Seki, Y . et al. Odin, G.S., 140, 209 Odin, G.S., see Velde, B. and Odin. G.S. Odorn, E.I., 139, 209 O’Donnell, H.J., see Mentser. M. et al. Oehler, D.Z., Schopf, J.W. and Kvenvolden, K.A.. 69, 108 Oftedahl, C., 510 Ogawa, T.. see Umegaki. Y . and Ogawa. T. Ogniben, L.. 65. 108 Ojakangas, R.W. and Keller, W.D., 52. 108 Ojakangas, R.W., see Beall, A.O. and Ojakangas. R.W. Ojakangas, R.W., see Dickinson. W.R. et al. Okada, H. and Honjo, S., 219, 286 Okamotu, G., Okura, T. and Goto, K., 87,108 Oki. Y . , see Seki, Y. et al. Okumura, K., see Seki, Y. et al. Okura, T., see Okarnotu, G. et al. Onuki, H., see Ernst, W.G. et a]. Onuki, H., see Seki, Y . et al. Orr, A.P. and Moorhouse, F.W., 37. 108 Orr, W., 238, 286 Ostrournov, E.A. and Shilov, V.M., 510 Otalora, G., 382, 383, 388, 400, 410, 417. 431, 487 Oudin, J.L., see Tissot. B. et al. Ozawa, K., see Seki, Y. et al.

Packharn. G.H. and Crook. A.W.. 51. 53. 108. 116. 209. 294, 296. 299. 300. 382. 383. 386, 391. 403, 430, 447. 449. 454, 487 Packham. G.H.. see Van der Lingen. G.J. and Packham. G.H. Page. R.J.. see Stewart. R.J. and Page. R.J. Parharn, W.E.. 176. 209 Parker. F.L.. 216. 238, 249. 286 Parker. F.L. and Berger. W.H.. 221. 222. 238. 286 Parker. R.B.. see Surdarn. R.C. and Parker. R.B. Passarge. S.. I08 Paterson. M.N.A.. Bien. G.S. and Berner. R.A., I08 Patton, W.W.. see Hoare. J.M. et al. Paul. A.Z.. 240. 286 Pedro. G.. Carrnouze. J.P. and Velde. B.. 131. 209 Perch-Nielsen. K., see Matter. A. et al. Perel’rnan. A.1.. 108 Perkins, R.D.. see Budd. D.A. and Perkins. R.D. Perkins. R.D.. see Zeff. M.L. and Perkins, R.D. Perry, E.. 161. 164. 168. 209. 328, 487 Perry. E.. Gisher. J.M. and Lawrence. J.R.. 133, 159. 160, 209 Perry. E. and Hower, J.. 161. 166. 168. 170. 171, 195. 209. 313. 314. 318. 319, 320, 322, 327. 340. 487 Perry, E.A.. see Hower. J. et al. Perry, E.A.. see Suchecki, R.K. et al. Pesty, L., see Szadecky-Kardoss. E. et al. Peterson. M.N.A.. 224. 225, 284 Peterson, M.N. and Goldberg. E.D.. 122, 209 Peterson, M.N.A. and Van der Borch. C.C. 67. I08 Petracco. F.. see Chennaux. G . et al. Petranek, J.. 497. 510 Pettijohn, F.J.. 21, 47, 50, 59. 76. 96. 108. 294, 295, 487. 505, 510 Pettijohn. F.J., Potter. P.E. and Siever. R., 35, 36, 41, 108 Petzing, J. and Chester, R., 142, 209 Picard, M.D., see High. L.R. and Picard. M.D. Pirnrn, A.C.. Garrison, R.F. and Boyce. R.E. 241, 251, 286 Pingitore, N.E., 108 Pinson, W.H., see Hurley, P.M. et al. Piper, D.Z. and Codispoti, L.A., 91, I08 Piper, D.Z., Cook, H.E. and Gardner, J.V., 276, 286 Pique, A,. 356, 358, 487 I

556 Pires, C.A.S., see Portugal Ferreira, M.R. et al. Plotnikova, V.J., see Kazakhov, A.V. et al. Pollard, C.O., 166, 209 Popov, B.P., 498, 510 Porrenga, D.H., 51, 55, 108 Portugal Ferreira, M.R., 487 Portugal Ferreira, M.R., Pires, C.A.S. and Sousa, M.B., 383, 458, 461, 487 Postma, H., see Van Andel, T.H. and Postma, H. Potter, P.E., see Pettijohn, F.J. et al. Pourbaix, M.J.N., 27, 108 Poustovalov, L.V., 497, 511 Powell, T.G., Foscolos, A.E., Gunther, P.R. and Snowdon, L.R., 177, 183, 184, 185, 209 Powell, T.G., see Foscolos, A.E. et al. Powers, M.C., 166, 169, 195, 209, 210, 312, 325, 487 Pray, L.C., 243, 286 Pray, L.C. and Murray, R.C., 213, 286 Pray, L.C., see Choquette, P.W. and Pray, L.C. Preiss, K., 186, 187, 210 Prozorovich, G.E., 334, 487 Pugin, V.A., see Khitarov, N.I. and Pugin, V.A. Purdy, E.G., 213, 286 Purser, B.H., 19, 108 Pusch, R., 198, 200, 210 Pustovalov, L.V., 294, 487 Pustowaloff, L.W., 22, 108 Quirk, J.P., see Aylmore, L.A.G. and Quirk, J.P. Raam, A., 388, 400, 403, 462, 463, 464, 487 Radier, H., see Millot, G. et al. Ragot, J.-P., see Barlier, J. et al. Ragot, J.-P., see Chateauneuf, J.-J. et al. Ramberg, H., 505, 511 Ramdohr, P. 511 Ramsay, A.T.S., 216, 287 Ramsay, A.T.S., Schneidermann, N. and Finch, J.W., 222, 238, 287 Rasmussen, H.W., see Rosenkrantz, A. and Rasmussen, H.W. Rastall, R.H., see Hatch, F.H. et al. Rateev, M.A., Gorbunova, Z.N., Lisitzyn, A.P. and Nosov, G.L., 119, 121, 210 Rateev, M.A. and Murav’ev, V.I., 319, 487 Rateev, M.A., see Tirnofeev, P.P. er al. Read. H.H. and Watson, J., 155, 210 Read. P.B., 461. 487 Read, P.B. and Eisbacher, G.H., 383, 400, 422. 461. 462, 463, 487

Reed, B.L. and Hemley, J.J., 362, 487 Reed, R.D., 75, I08 Reesman, A.L., see Keller, W.D. et al. Reineck, H.E., see Fiichtbauer, H. and Reineck, H.E. Rekshinskaya, L.G., see Volkova, A.N. et al. Renard, A.F., see Murray, J. and Renard, A.F. Reuling, H.T., 78, 109 Revelle, R. and Emery, K.O., 74, 109 Rex, R.W. and Goldberg, E.D., 122, 210 Reynolds, D.L., 75, 109 Reynolds, R.C. and Anderson, D.M., 420,458, 487 Reynolds, R.C. and Hower, J., 316, 318, 487 Reynolds, R.C., see Gavish, E. and Reynolds, R.C. Reynolds, W.R., 420, 458, 488 Rhodes, M.L., see Adams, J.E. and Rhodes, M.L. hchard, R.S., see Birch, G.F. et al. Riech, V., see Von Rad, U. et, al. Riedel, W.R., 40,66, 109, 214, 215, 216, 220, 221, 222, 235,287 Riedel, W.R., see Funnel, B.M. and Riedel, W.R. Rieke, H.H. and Chilingarian, G.V., 155, 185, 186, 187, 188, 193, 194, 198, 210 Rieke, H.H., see Chilingarian, G.V. and Rieke, H.H. Rieke, H.H., see Chilingarian, G.V. et al. Riley, C.M. and Byrne, J.V., 65, 109 Rinne, F., 32, 109 Rioult, M., see Dangeard, L. and Rioult, M. Rittenberg, S.C., see Emery, K.O. and Rittenberg, S.C. Rittenhouse, G., 79, 80, 109 Rivibre, A. and Visse, L., 90, 109 Robaszynski, F., see Beugnief, A. et al. Roberson, H.E., 132, 210 Robert, P., 374, 395, 456, 488 Robertson, A.H.F., 146, 210 Robie, R.A., Bethke, P.M. and Beardsley, K.M., 407, 488 Robinson, P.T., 423, 458, 488 Robinson, P.T., McKee, E.H. and Moiola, R.J., 423, 458, 488 Roehler, H.W., 419, 458, 488 Rogers, J., see Landis, C.A. and Rogers, J. Roggwiller, P., see Frey, M. et al. Rona, P.A. and Lowell, R.P., 2, 15, 109 Ronov, A.B., 67, 109 Rosenbaurn, M.S., 155, 157,210 Rosenquist, I.T., 129, 198, 200, 210 Rosenkrantz, A. and Rasmussen, H.W., 12, I09

557 Ross, C.S., 56, 109, 461, 488 Ross, C.S. and Kerr, P.F., 58, I09 Ross, D.A., see Degens, E.T. and Ross, D.A Roth, P.H. and Berger, W.H., 225, 236, 287 Roth, P.H., see Adelseck, C.G. et al. Roth, P.H., see Schlanger, S.O. et al. Rothe, P., see Houghton, R.L. et al. Routhier, P., 511 Rowe, G.T., see Manheim, F. et al. Roy, R., see Crowley, M.S. and Roy, R. Rozenson, I., see Heller-Kallai, L. and Rozenson, I. Rukhin, L.B., 118, 294, 488 Runnels, D.D., 26, 109 Russell, K.L., 49, 54, 109, 132, 133, 210 Russell, R.D., see Russell, R.J. and Russel, R.D. Russell, R.J. and Russell, R.D., 61, 109 Rutten, M.G., 12, 73, 14, 109 Rosch, H., see Von Rad, U. et al. Sabatier, G., 325, Sabatier, G., see Heezen, B.C. et al. Sabatier, G., see Wyart, J. and Sabatier, G. Sabine, P.A., 458, 488 Sabine, P.A., see Cameron, L.B. and Sabine, P.A. Sagon, J.-P., 370, 488 Sagon, J.-P. and Dunoyer de Segonzac, G., 348, 352, 359, 310, 488 Sakharov, B.A., see Shutov, V.D. et al. Sakhibgareev, R.S. and Galikeev, K.Kh., 321, 488

Saleh, A,, see Gunatilaka, A. et al. Saliot, P., 393, 395, 411, 412, 488 Saliot, P., see Guitard, G. and Saliot, P. Salvan, H.M., 93, 109 Salyn, A.L., see Murajew, W.I. and Salyn, A.L. Samsonowicz, J., 68, 109 Samuels, S.G., 127, 210 Sandalaki, Z., see Brindley, G.W. and Sandalaki, Z. Sander, B., 85, 109 Sann, D.D., 80, 109 Savin, S.M. and Douglas, R.G., 238, 287 Savin, S.M., see Douglas, R.G. and Savin, S.M. Savin, S.M., see Eslinger, E.V. and Savin, S.M. Sawabini, C.T., see Chilinganan, G.V. et al. Sayles, F.C., 133, 210 Sayles, F.C. and Mangelsdorf, P.C., 159, 210 Sayles, F.C. and Manheim, F.T., 156, 210, 258, 287

Scheidegger, A.E., 43, 109 Schermerhorn, L.J.G., 390, 488 Scherp, A,, 334, 340, 488 Scherp, A,, Stadler, G. and Schmidt, W., 360, 361, 363, 313, 314, 488 Schindler, C., see Frey, M. et al. Schlager, W. and James, N.P., 214, 287 Schlanger, S.O., 81, 109 Schlanger, S.O. and Douglas, R.G., 214, 215, 223,224, 226, 232, 235, 241, 243, 244, 241, 249, 250, 251, 253, 254, 251, 259, 261, 268, 269, 215, 287 Schlanger, S.O., Douglas, R.G., Lancelot, Y., Moore, T.C. and Roth, P.H., 214,226,241, 243, 241, 250, 251, 256, 268, 287 Schlanger, S.O., see Arthur, M.A. and Schlanger, S.O. Schlanger, S.O., see Coplen, T.B. and Schlanger, S.O. Schlanger, KO., see Manghnani, M.H. et al. Schlanger, SO., see Milholland, P.D. et al. Schmalz, R.F., 82, 109 Schmidt, W., see Scherp, A. et al. Schmitt, H., 41, 109 Schneiderhohn, H., 511 Schneidermann, N., 236, 287 Schneiderman, N., see Anderson, T.F. and Schneiderman, N. Schneidermann, N., see Ramsay, A.T.S. et al. Scholl, D.W., see Hein, J.R. and Scholl, D.W. Scholl, D.W., see Hein, J.R. et al. Scholle, P.A., 110, 214, 232, 243, 244, 251, 252, 254, 263, 264, 265, 215, 216, 271. 287 Scholle, P.A. and Kennedy, W.J., 265, 287 Scholle, P.A., see Hancock, J.M. and Scholle, P.A. Schopf, J.W., see Oehler, L. et al. Schott, W., 249, 287 Schreiber, B.C., see Briskin, M. and Schreiber, B.C. Schreyer, E.D., 343, 313, 488 Schroeder, R.J. and Hayes, J.B., 337, 488 SchrMer, J.H., see Ginsburg, R.N. et al. Schurmann, K., see Hinrichsen, Th. and Schiirmann, K. Segnit, E.R., see Jones, J.B. and Segnit, E.R. Seibold, E., 139, 154, 210, 505, 511 Seki, Y., 383, 385, 388, 390, 391, 392. 416, 426, 427, 430, 440, 443, 447, 448, 450,453, 454, 456, 488, 489 Seki, Y., Ernst, W.G. and Onuki, H., 344,382, 443, 467, 489 Seki, Y., Oki, Y., Odaka, S. and Ozawa, K., 401, 418, 422, 489

Seki, Y., Oki, Y., Matsuda. T., Mikami, K. and Okumura, K. 383, 385, 388, 391, 403, 410, 411, 427, 430, 450, 459, 460, 464, 489 Seki, Y., Onuki, H., Oba, T. and Mori, R., 388, 390, 430, 443, 468, 489 Seki, Y . , Onuki, H.. Okumura, K. and Takashima, I., 384, 385, 41 1, 416,427, 440, 489 Seki, Y., see Ernst, W.G. and Seki, Y. Seki, Y.,see Ernst, W.G. et al. Selley, R.C.. 34, 110 Senderov, E.E., 419, 489 Serratosa, J.M., see Bradley, W.F. and Serratosa, J.M. Seyfried, W.E., Shanks, W.C. and Dibbel, W.E., 144, 210 Shanks, W.C., see Seyfried, W.E. et al. Shearman, D.J., Khouri, J. and Taha, S., 85, I10 Sheldon, R.P., see McKelvey, V.E. et al. Shelton, J.W., 58, 110 Sheppard, R.A. and Gude, A.J., 110,419,422. 458, 489 Sheppard, R.A., see Surdam, R.C. and Sheppard, R.A. Shevyakova, E.P., see Karpova, G.V. and Shevyakova, E.P. Shevyakova, E.P., see Karpova, G.V. et al. Shilov, V.M., see Ostroumov, E.A. and Shilov, V.M. Shimazu, M., Tabuchi, A. and Kusuda, T., 383, 385, 388, 410, 41 I , 427,447, 458, 459, 460, 461, 489 Shimoda, S., see Sudo, T.et al. Shimoyama, A., see Johns, W.D. and Shimoyama, A. Shinn, E.A., Ginsburg, R.N. and Lloyd, R.M., 79,110 Shinn, E.A., Halley, R.B., Hudson, J.H. and Lidz, B.H., 244, 279, 287 Shinn, E.A., see Ginsburg, R.N. et al. Shipek, C.J., see Hamilton, E.L. et al. Sholkovitz, E., 159, 210 Shumway, G., see Hamilton, E.L. et al. Shutov, V.D., Aleksandrova, A.V. and Losievskaya, S.A., 336, 489 Shutov, V.D. and Dolmatova, T.V., 334, 336, 360, 489 Shutov, V.D., Drits, V.A. and Sakharov, B.A.. 319, 338, 489 Shutov, V.D., see Drits, V.A. and Shutov, V.D. Shutov, V.D., see Kopeliovich, A.V. et al. Shutov, V.D., see Kossovskaya, A.G. and Shutov, V.D.

Shutov, V.D., see Kossovskaya, A.G. et al. Siedlecka, A,, see Folk, R.L. and Siedlecka, A. Siever, R., 26, 33, 53, 66, 86, 87, 110 Siever, R., Beck, K.C. and Berner, R.A.. 150. 151, 210

Siever, R. and Kastner, M., 152, 210 Siever, R. and Woodford, N., 133, 210 Siever, R., see Garrels, R.M. et al. Siever, R., see Pettijohn, F.J. et al. Sillen, L.G., 24, 67, I10 Sims, P.K., see James, H.C. and Sims, P.K. Singer, A,, 138, 210 Singer, A. and Stoffers, P., 137, 138, 211 Singer, A,, see Stoffers, P. and Singer, A. Skeats, E.W., 80, 110 Skinner, H.C.W., 78, I10 Skinner, H.C.W., see Alderman, A.R. and Skinner, H.C.W. Sliter, W.V., Be, A.W.H. and Berger, W.H., 214, 287 Sloss, L.L., see Krumbein. W.C. and Sloss, L.L. Smalley, F.J., see Bentley, S.R. and Smalley, F.J. Smirnow, L.P., 37, I 1 0 Smit, D.E. and Swett, K., 85, 110 Smith, J.V. and Yoder, H.S., 329, 489 Smith, R.E., 391, 430, 447, 449, 454, 464, 489 Smulikowski, K., 52, 110 Smykatz-Kloss, W. and Althaus, E., 350, 489 Snellwood, B.W., see Hallam, A. and Snellwood, B.W. Snowdon, L.R., see Powell, T.G. et al. Sorby, H.C., 110 Sousa, M.B., see Portugal Ferreira, M.R. et al. Southain, J.R., see Hay, W.W. et al. Spackman, W. and Moser, R.G., 361, 373,490 Sparks, D.M., see Castaiio, J.R. and Sparks, D.M. Spears, D.A., 200, 211 Spears, D.A., see Curtis, C.D. and Spears, D.A. Spencer, E., 75, 110 Spiess, F.N. et al., 39, 110 Stadler, G., 334, 336, 359, 361, 364, 368, 373, 490 Stadler, G., see Scherp, A. et al. Standaert, R.R., see Bodine. M.W. and Standaert, R.R. Stanton, R.L., 511 Starke, R., 303, 307, 340, 490 Staub, J. and Cohen, A,, 153, 211 Steidtmann, E., 80, 84, 110 Steiner, A,, 312, 314, 325, 385, 411, 416, 427, 440, 490

559 Steinike, K.. 302, 490 Steinitz, G., 70, 71, 110 Steinitz, G. and Kolodny, Y . , 70, 110 Stephens, C.G., 45, 110 Stern, K.H., 63. 110 Stetson, H.C., 63, I10 Stevaux, J., see Beuf, S. et al. Stevens, R.E. and Carron, M.K., 48, 111 Stevenson, F.J. 39, 111 Stewart, F.H., 83, 85, 111 Stewart, R.J., 403, 417, 436, 490 Stewart, R.J. and Page, R.J., 383, 403, 417, 437, 463, 490 Stewart, R.J., see Dickinson, W.R. et al. Stoffers, P. and Holdship, S., 137, 211 Stoffers, P. and Singer, A., 137, 211 Stoffers. P., see Singer, A. and Stoffers, P. Stonecipher, S.A., 140, 211 Stonecipher, S.A., see Kastner, M. and Stonecipher, S.A. Strakhov, N.M., 81, 1 1 1 , 118, 211, 294, 490, 511

Stringham, B., 75, 111 Suchecki, R.K., Perry, E.A. and Hubert, J.F., 162, 177, 211 Sudo, T., Hatashi, H. and Shimoda, S., 366, 368, 490 Sudo, T., see Tomita, K. and Sudo, T. Sugawara, K., 63, I I I Sujkowski, Z.L., 40, 47, 72, 111 Sun, Ming-Shan and Allen, J.E., 75, 111 Supko, P.R., see Davies, T.A. and Supko, P.R. Surdam, R.C.. 382, 385, 388, 400, 403, 410, 426, 427, 428, 429, 430, 490 Surdam, R.C. and Eugster, H.P., 136, 211 Surdam, R.C. and Parker, R.B., 75, 111, 419, 458, 490 Surdam, R.C. and Sheppard, R.A., 135, 136, 211 Surdam, R.L., see Goodwin, J.H. and Surdam, R.L. Sutton, G.H., see Milholland, P.D. et al. Swanson, R.W., see McKelvey, V.E. et al. Swett, K., see Smit, D.E. and Swett, K. Swing, J., see Houtz, R. and Swing, J. Szadecky-Kardoss, E., BBrdossy, Gy., Fiirst, I., Pesty, L., Kliburszky, B., Tomor, E. and Tomschey, O., 325, 490 Tabuchi, A., see Shimazu, M. et al. Taha, S., see Shearman, D.J. et al. Takahashi, T., 225, 288 Takahashi. T. and Broecker, W.S.. 235, 288 Takashima. I., see Seki. Y . et al.

Taliaferro, N.L., 69, 111 Tamm, O., 49, 111 Tan, T.K., 198, 211 Tanaka, K., see Nakajima, W. and Tanaka. K. Tank. R., 137,211 Taraboulos, A,, see Kolodny, J. et a]. Tarr, W.A., 61, 111 Taylor, A.M., see Coombs, D.S. et al. Taylor, H.P. and Coleman, R.G., 45 I , 490 Taylor, J.C.M., see Illing, L.V. et al. Taylor, J.H., 94, 111, 115, 154, 179, 211, 498, 51 I

Teichert, C., see Fairbridge, R.W. and Teichert, C. Teichrniiller, M. and Teichmuller, R., 373. 490 Teichmuller, M., see Heling, D. and Teichmiiller, M. Teichmiiller, R., see Hoyer, P. et al. Teichmiiller, R., see Teichmiiller. M. and Teichmiiller, R. Teodorovich, G.I., 21, -22, 47, 51, 79, 92, 93, 111, 294, 490, 498, 5511 Teodorovich, G.I. and Chernov, A.A., 191, 211

Teodorovich, G.I., Chernov, A.A. and Kotel’nikov, D.D., 312, 491 Teodorovich, G.I. and Konyukhov, A.I., 312, 327, 490 Termier, G., see Termier, H. and Termier, G. Termier, H. and Termier, G., 44, 59, 68, 88, 111

Ters, M., 68, 111 Tester, A.C. and Atwater, G.I., 22, 111 Tettenhorst, R. and Moore, G., 137, 211 Thiede, J., 249, 288 Thiel, G.A., see Gruner, J.W. and Thiel, G.A. Thode, H.G., Harrison, A.G., and Monster, J., 111

Thomas, W.H., see Bien, G.S. et al. Thome, K.N., see Hoyer, P. et al. Thompson, A.B., 362, 408, 409, 410, 41 1, 412, 417, 429, 430, 437, 438, 440, 441, 450, 453, 491

Thompson, G.R. and Hower, J., 139,211 Thompson, M.E., see Garrels, R.M. et al. Thomson, A,, 89, 111 Tikhomirova, M.M., see Kazakhov, A.V. et al. Timofeev, P.P., Eremeev, V.V. and Rateev, M.A., 138, 211 Timofeeva, Z.V., see Karpova, G.V. and Timofeeva, Z.V. Tiratsoo, E.N., 276, 288 Tissot, B., 38, 111 Tissot, B., Califet-Debyser, Y . , Derov, G. and Oudin, J.L., 182, 211

Tochilin, M.S., 498, 511 Tomita, K. and Sudo, T., 318, 491 Tomor, E., see Szadecky-Kardoss, E. et al. see Szadecky-Kardoss, E. et al. Tomschey, 0.. Toporez, G.A., see Gurewitsch, A.B. and Toporez, G.A. Touray, J.-C., see Barlier, J. et al. Tourtelot, H.A., 182, 183, 211 Tracey, J.I., 226, 241, 243, 276, 288 Trendall, A.F., 94, 112 Triplehorn, D.M., 330, 334, 491 Triimpy, R., 201 Tsu-Ming Han, 496, 51 I Tucker, M., 11, 12, 13 Tugarinov, A.I. and Vinogradov, A.P., 52, 112 Tunn, W.L.M., see Von Engelhardt, W. and Tunn, W.L.M. Turner, F.J., 295, 299, 300, 306, 379, 386, 407, 425, 426. 491 Turner, F.J. and Verhoogen, J., 53, 112, 415, 491 Turner, F.J., see Fyfe, W.S. and Turner, F.J. Turner, F.J., see Fyfe, W.S. et al. Turner, F.J., see Hutton, C.O. and Turner, F.J. Turner, F.J., see Williams, H. et al. Turner, P., 45, 112 Twenhofel, W.H., 18, 21, 37, 59, 60, 67, 112, 294, 491 Umegaki, Y. and Ogawa, T., 397, 491 Upchurch, W.J., see Marshall, C.E. and Upchurch, W.J. Uschakova, M.G., 219, 288 Utada, M., 383, 398, 399, 421, 458, 459, 491 Utada, M. and Minato, H., 383, 491 Utada, M., see Iijima, A. and Utada, M. Valeton, I., 95, 112 Van Andel, T.H. and Postma, H., 57, 112 Van Andel, T.H., Heath, G.R. and Moore, T.C., 216, 288 Van den Broueke, M., see Durand, B. et al. Van der Borch, C.C., see Peterson, M.N.A. and Van der Borch, C.C. Van der Lingen, G.J., 2, 15, 40, 67, 112 Van der Lingen, G.J. and Packham, G.H., 214, 226, 228, 230, 232, 241, 244, 251, 270, 272, 273, 279,288 Van Hise, C.R., 17, 38, 41, 46, 96, 112 Van Houten, F.B., 59, 112 Van Olphen, H., 127, 128, 194, 199, 211 Van Straaten, L.M.J.U., 75, 112 Van Tuyl, F.M., 80. 112

Vatan, A,, see Mennig, J.J. and Vatan, A. Vaugnat, M., see Coombs, D.S. et al. Vaugnat, M., see Kubler, B. et al. Vaugnat, M., see Martini, J. and Vaugnat, M. Veizer, J., 112 Velde, B., 181, 211, 322, 323, 329, 330, 343, 369, 467, 491 Velde, B. and Hower, J., 329, 491 Velde, B. and Kornprobst, J., 362, 491 Velde, B. and Odin, G.S., 140, 212 Velde, B., see Church, T.M. and Velde, B. Velde, B., see Kienast, J.R. and Velde, B. Velde, B., see Pedro, G. et al. Venkatarathnam, K. and Biscaye, P., 53, 58, 112 Verhoogen, J.. see Fyfe, W.S. et al. Verhoogen, J., see Turner, F.J. and Verhoogen, J. Veselovskaya, M.M., 468, 492 Vinogradov, A.P., see Tugarinov, A.I. and Vinogradov, A.P. Vishnevskaya, T.N., see Zaporozhtseva, A.S. et al. Visse, L.D., 90, 112 Visse, L., see Riviere, A. and Visse, L. Vlasov, V.V.. 326, 333, 492 Vlodarskaya, V.R. and Nosov, G.J., 328, 492 Volkova, A.N., Ivanova, N.V. and Rekshinskaya, L.G., 343, 492 Von Engelhardt, W., 185, 190, 192, 212 Von Engelhardt, W. and Gaida. K.H., 130, 198,212 Von Engelhardt, W., Miiller, G . and Kromer, H., 178. 212, 337. 338, 492 Von Engelhardt, W. and Tunn, W.L.M., 195, 212 Von Engelhardt. W., see Correns, C.W. and Von Engelhardt, W. Von Gaertner, H.R., see Eckhardt, F.-J. and Von Gaertner, H.R. Von Guembel. C.W., 17, 96, 112 Von Morlot, A., 85, 112 Von Rad, U., Riech, V. and Rosch, H., 147. 148, 212 Von Rad, U., see Berger, W.H. and Von Rad, U. Vorabutr, P., see Chilingarian, G.V. and Vorabutr, P. Wade, A,, 78, 112 Waksman, S.A., 36. 112 Walker, T.E.. 85, 87, 112 Walls, R.W., Mountjoy, E.W. and Fritz, P., 12, I5

Walther. J., 18, 96, 112, 294, 492 Wampler, J.M., see Weaver, C.E. and Wampler, J.M. Watson. J., see Read, H.H. and Watson, J. Weatherby, B.B. and Faust, L.Y., 253, 288 Weaver, C.E., 54, 112, 113, 166, 169, 175. 176, 178,212. 312,319, 329, 331, 344, 345. 346, 347, 348, 350, 492 Weaver, C.E. and Beck, K.C., 138, 169. 212 Weaver, C.E. and Wampler, J.M., 327, 492 Weaver, F.M., see Wise, S.W. and Weaver. F.M. Weber, K., 347, 348, 350, 35 I , 357, 358, 359, 360, 374, 465, 492 Weir, A.H., see Brown, G. et al. Weiss, M.P., 73, 74, I13 Weller, J.M., 243, 288 Wells, A.J., 113 Wells, A.J., see Illing, L.V. et al. Wermund, E.G., 51, 113 West, I.A., Ali, Y.A. and Hilmy, M.E., 64, 113 Wetzel, O., 68, 113 Wetzel, W., 22, 113 Wey, R., see Millot, G. et al. Weyl, P.K., see Deffeyes. K.S. et al. Weynschenk, R., 81, 113 Whetten, J.T., 383, 398, 462, 492 Whetten, J.T., see Coombs, D.S. and Whetten, J.T. Whetten, J.T., see Wilkinson, J.F.T. and Whetten, J.T. White, D.E., 41, 113 White, D.E., see Muffler, L.J.P. and White, D.E. White, W.A., see Bredehoeft, J.D. et al. Whitehead, T.H., see Harms, J.E. et al. Whitehouse, U.G. and McCarter, R.W., 54, 113, 134, 212 Whiteley, K.R., see Murata, K.J. and Whiteley, K.R. Whiteman, A.J., see Newell, N.D. et al. Wickman, F.E., 25, 113 Wilber, R.J., 226, 288 Wilkinson, J.F.T. and Whetten, J.T., 398, 424, 492 Williams, D., 39, 113 Williams, H., Turner, F.J. and Gilbert, C., 294. 295, 296, 291, 492. 496, 511 Willman, H.B., 75, 113

Wilson. J.L., 4, 15, 73, 79, 113, 213, 288 Wilson, M.J., Bain, D.C., Bettardy, W.J. and Berrow, M.L., 153, 212 Windom. H.L.. 122, 212 Windom, H.L. and Chamberlain. C.F., 122, 212 Windom, H., see Griffin, J.J. et al. Winkler. H.G.F., 201, 203, 212, 296, 303, 306, 307, 309, 359, 364. 369, 392, 393, 409, 412, 430, 451, 492, 493 Winterer, E.L., Ewing, J.I. and co-workers, 215, 226, 275, 288 Winterer, E.L., see Berger, W.H. and Winterer, E.L. Willis. J.P., see Birch, G.F. et al. Wise, S.W., 214, 226, 245, 247. 251. 267, 288 Wise, S.W. and Eugster, H.P., 331, 461, 493 Wise, S.W. and Kelts, K.R., 226, 242, 288 Wise, S.W. and Weaver, F.M., 146, 212 Wise, W.S., see Boles, J.R. and Wise, W.S. Wolf, K.H., see Ch/ingar, G.V. et al. Wolf, M., 374, 465, 493 Wood, G.H., see Hosterman, J.W. et al. Woodford, N., see Siever, R. and Woodford. N. Wyart, J. and Sabatier, G., 324, 493 Yariv, D. and Cross, H., 127, 128, 130, 212 Yoder, H.S. and Eugster, H.P., 329, 330, 493 Yoder, H.S., see Smith, J.V. and Yoder, H.S. Yoshitant, A., 383, 398, 400, 403, 493. Zak, I., see Heller-Kallai, L. et al. Zankl, H., 243, 288 Zaporozhtseva, A.S., 340, 462, 464, 493 Zaporozhtseva, A.S., Vishnevskaya, T.N. and Dubar, G.P., 383, 455, 493 Zaporozhtseva, A.S., Vishnevskaya, T.N. and Glushinskiy, P.I., 383, 493 Zeff, M.L. and Perkins, R.D., 240, 288 Zen, E-An, 58, 81, 113, 322, 390, 405, 432, 440, 443, 468, 493 Zen, E-An and Thompson, A.B., 440, 493 Zen, E-An, see Albee, A.L. and Zen, E-An Zenger, D.H., 78, 113 Zimmermann, H.B., 121, 212 Zimmermann, H., see McCoy, F. et al. ZoBell, C.E., 2, 15, 27. 37, 113, 510, 511 Ziillig, H., 118, 128, 187, 189, 212

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SUBJECT INDEX Aar Massif, 371 Abrasion pH, 48 Abrolhos Submergence, 73 Acoustic stratigraphy, 275-276 Actinolite, 283, 305, 338, 376, 379, 389, 390, 391, 392, 394, 395, 396, 412, 429, 430, 435, 439, 442, 443, 447, 448, 449, 450, 451, 452, 453, 454, 460,464, 468,469, 473 Adularia, 299, 399 Aerobic bacteria, 28, 61 Amazon, 58, 59 Ameboid mosaic, 7 Amphibole, 293, 301, 391, 392, 393 Anadiagenesis, 4, 17, 34, 35, 41-43, 61, 62, 65, 74, 77, 82-83, 84, 86, 87, 91 Anaerobic bacteria, 28, 37 Analcime, 10, 21, 53, 136, 138, 140, 142, 143, 203, 291, 299, 303, 304, 376, 378, 381, 382, 383, 384, 385, 396, 397, 398, 399, 400, 402, 404, 406, 407, 408, 409, 410, 419, 421, 423, 424, 426, 428, 432, 433, 448, 454, 458, 460 Anatase, 20, 75, 301 Anchimetamorphism, definition, 347 Andalusite, 362, 370, 374 Andesine, 437 Andros Island, 73, 79 Angola, 383, 458, 461 Anhydrite, 20, 28, 45, 62, 63-65, 82, 83, 293, 324 Ankerite, 20, 81, 439 Anorthite, 409, 41 1, 414, 437 Anthracite, 373, 374, 456, 466, 471, 472 Alabama, 94 Alabandite, 28 Alaska, 436 Alberta, 312, 462 Albite, 6, 74, 143, 291, 299, 303, 304, 305, 341, 359, 369, 370, 376, 38 1, 382, 386, 390, 391, 393, 399, 403, 405, 407, 408, 409, 410, 412, 417, 419, 421, 422, 424, 428, 431, 433, 448, 452, 456, 464 Albitization of plagioclase, 299, 402, 403, 431 Algeria, 496 Algerian Sahara, 359, 360 Allevardite, 10, 291, 309, 319, 337, 342, 352, 357, 359, 363, 365, 366, 367, 368, 373, 374, 375, 378, 465,467 Alsace-Lorraine. 94

Apatite, 91 Appalachians, 60, 75, 88, 91, 390, 443, 445, 468 Apsheron, Azerbayjan, 193 Aquatolysis, 116, 117, 130-131 Arabian desert, 122 Arabian Sea, 172 Aragonite, 7, 20, 78, 122, 222, 238, 456 Arctic Basin, 59 Ardennes, 82, 370 Armorican Massif, 359, 370 Atlantic Ocean, 5 5 , 57, 58, 121, 122, 123, 124, 133, 138, 150, 218 Augite, 404, 405 Australia, 67, 68, 78: 87 Australian desert, 122 Austria, 256 Authigenesis, 19-22, 47-50 Authigenic feldspar, 74-76, 179- 180 Azerbaydzhan, 312, 327, 343 Azores, 124 Azurite. 20 Bacterial action, 58, 61 Bahamas, 73, 79, 255 Baja California, 64 Barents Sea, 152 Barite, 20 Bassin de Douala, Cameroun, I82 Bauxite, 95 Bauxitization, 95-96 Beachrock, 74 Bedded chert, 67 Beidellite, 327 Belgium, 358, 362 Bellinghausem Abyssal Plain, Pacific Ocean, 158, 180 Belt Series of Idaho, 342 Bentonite, 74, 145, 178, 315 Bering Sea, 145, 147 Bikini-Eniwetok, 82, 83 Biorhexistasy, 59 Biostasy, 44, 45, 68 Biotite, 10, 75, 131, 140, 179, 290, 301, 304, 310, 326, 330, 359, 370, 371, 372, 374, 376, 379, 388, 391, 460, 462, 464, 465, 470 -, alteration during burial diagenesis, 340 -, decomposition, 463-464

Black Forest, Germany, 337 Black Sea, 36, 59, 117, 118, 151. 188 Bonaire, Dutch West Indies, 79 Bornite, 20, 39 Bowen Basin, Queensland, 335 Brazil, 122 Brazilian Shield, 87 Brecciated chert, 69 BrianGonnais zone, 395, 456. 470, 471, 472 British Columbia, 162, 165, 167, 382, 383, 385, 388, 399, 403, 461, 462, 465 British Guiana, 122 Brookite, 20, 75, 301 Brunsvigite, 338 Burial diagenesis and metamorphism, equilibrium processes, 292-293 Burial-metamorphic zeolites, 396-405 Calabria, 393 Calcite, 7, 20, 28, 58, 78, 79, 81, 82, 86, 87, 88, 94, 122, 144, 145, 147, 158, 180, 203, 220, 236, 238, 242, 247, 248, 250, 254, 255, 256, 258, 267, 268, 270, 271, 272, 273, 275, 279, 280, 341, 381, 389, 390, 399, 412, 418, 432, 433, 434, 435, 436, 437, 439, 440, 441, 449, 451, 452, 503, 504, 505 Calcitization, 85 Caliche crusts, 275 California, 12, 13, 90, 393, 403, 419, 436, 444, 456, 457, 458, 463, 478 California Basin, 125, 1SO, 15 1 Cambrian, 75, 90, 162 Cambro-Silurian, 357, 465 Campanian, 146 Canada, 69 Canaries, 124 Cape Verde Islands, 124 Carbon isotope data, 12, 13, 261 Carbonate concretions, 153-154 Carbonatization, 49-50 Carboniferous, 87, 153, 312, 319, 328, 331, 332, 334, 335, 336, 337, 339, 342, 361, 364, 373, 374, 386, 455 Caribbean, 159, 239, 265 Case-hardening, 87 Catagenesis, 2, 18, 289, 290, 294, 296, 331, 332, 333 Caucasus, 3 12 Celadonite, 143, 144, 158, 331, 332, 333, 404, 405, 458, 460, 467 -, in burial-metamorphic facies, 461 -462 Celebes, 390, 391 Celestite, 20 Cementation, 7, 8, 19, 41, 77, 87

- of argillaceous sediments, 180 Cenozoic, 75, 91, 145, 176, 312, 327, 383, 419, 458 Central America, adjacent basins, 122 Central Pacific, 238, 243, 265 Cerithium limestone, 72 Cerussite, 20 Chabazite, 21, 138, 406 Chalcedony, 20, 65, 66, 68, 87, 88, 146, 147. 149, 381, 399 Chalcopyrite, 20, 39, 152 Chalk, definition, 215 -, 8, 227, 228, 230, 250-252, 256, 258, 264, 269, 270, 271, 272, 273, 276, 277, 281 Chalk stage, 269 Chamosite, 20, 28, 94, 131, 315 Changes in microstructure of clay, 198-200 Chlteaulin Basin, 370 Chert, 40, 69, 70, 87, 94 Chile, 90, 405 Chlorite, 5, 6, 9, 20, 5 5 , 58, 121, 122, 134, 135, 143, 144, 152, 161, 164, 179, 181, 182, 201, 291, 298, 300, 303, 304, 305, 307, 310, 31 1, 324, 325, 326, 331, 334, 335, 336, 337, 340, 341, 353, 357, 359, 362, 363, 375, 316, 381, 386, 389, 390, 391, 392, 399, 404, 405, 429, 430, 434, 435, 439, 448, 449, 450, 451, 452, 453, 454, 459, 460, 461, 462, 463, 467 -, composition during burial diagenesis, 338-339 Chloritoid, 10, 306, 359, 361, 362, 367, 374, 379, 470 -, appearance during incipient metamorphism 369-370 Chlorophyceae, 90 Cinnabar, 22 Ciscaucasia, 3 12, 326 Clay, change in structure during diagenesis, 185-200 Clay filtering, 43 Clay minerals, deep-burial diagenesis, 159- 178 _ _ , distribution in ocean sediments, 119122 - -, halrnyrolytic authigenesis, 53-58 - _ , shallow-burial diagenesis, 152-153 Cleveland, 94 Clinochlore, 449, 450 Clinoptilolite, 53, 136, 140, 141, 142, 143, 144, 145, 291, 303, 378, 381, 383, 396, 398, 399, 401, 406, 414, 418, 420, 421, 422, 423, 425, 426, 427, 441, 454, 458, 460 Clinopyroxene, 464 Clinozoisite, 22, 391, 406, 450 Clinton iron ores, 94

565 C/N ratio, 39 Coal rank, 395, 465-466, 472-474 Coal ranks associated with anchimetamorphism. 373-374 Coast Ranges, California, 343 Coccoliths, 7, 219, 220, 225, 233, 236, 242, 243, 245, 246, 247, 248, 249, 266, 267 Coccolithophorids, 218-220, 235 Collenia, 88 Collophane, 20, 92 Colorado. 276 Columbia, 122 Compaction, 3, 5, 7, 19, 41, 77, 86, 87, 155, 156, 185, 193, 194, 195, 196, 198-200, 243, 244, 251, 253, 270, 276, 278, 279, 280 Concretions, 180 Congelation, 19 Congo, 58 Connate water, 41, 43, 44, 54, 55, 76, 86, 87 Coorong lagoon, 78 Coquimbite, 63 Corrensite, 6, 9, 134, 135, 177, 181, 182, 324, 325, 368, 378, 395, 459, 461, 471 Cow Head Klippe, Newfoundland, 177 Cretaceous, 67, 68, 69, 70, 71, 74, 75, 90, 93, 140, 141, 154, 162, 165, 167, 171, 216, 221, 250, 260, 264, 265, 274, 3 1 1, 3 12, 3 13, 334, 355, 356, 358, 361, 364, 371, 373, 375, 382, 383, 403, 455, 462, 465, 469 Crimea, U.S.S.R.. 337 Cristobalite. 145, 146, 147, 399, 420, 421, 423 Crystallinity, illite, 342-351, 353, 355, 356, 358, 359, 364, 365, 371, 372, 374, 375, 378, 465, 470, 472, 473 Cuba, 122 Cyprus, 147 Dagestan, U.S.S.R., 337, 343 Dahllite, 20, 92 Danian, 72 Dead Sea, 69 Decementation, 77 Dedolomitization, 77, 85, 94 Deep-burial diagenesis, 154-182, 25 1-258 Deep-sea carbonates, distribution, 216-218 Definition,. anadiagenesis, 96 -, anamorphism, 96 -, authigenesis, 96 -, cementation, 2 15 -, chalk, 215 -, diagenesis. 2, 17, 18, 96, 1 15-1 18, 222 -, dissolution, 215 -, epidiagenesis, 96 -, epigenesis, 96

-, halmyrolysis, 96 -, hypogene, 96 -, katamorphism, 96 -, limestone, 215 -, lithification, 96, 215 -, lithogenesis, 96 -, ooze, 215 -, planktonic, 215 -, pressure-solution, 2 15 -, supergene, 96 -, syndiagenesis, 97 -, syngenesis, 97 Dehydration, 47 Dehydrolysis, 48 Denmark, 71 Density, 185-187, 226, 229, 231, 232, 252, 253, 268, 271 Desilicification, 87, 89 Desulfooibrio desulfuricans, 37, 62 Devonian, 82, 84, 328, 333, 359, 360, 362, 367, 370, 373 Devonian Catskill delta, 60 Dewatering stage, 43 Dhahran, 79 Diabanite, 338 Diaclastic revival, 42, 43 Diagenesis-metamorphism transition, 201 203 Diagenetic equation, 23 Diagenetic evolution. 33-46, 499, 505 Diagenetic fabric, 41 Diagenetic potential, 266-268, 275, 280, 28 1 -, realms 223, 224 -, rhythm, 40 -, versus metamorphic, 299-308 Diaspore, 303 Dickite, 194, 304, 306, 31 I , 336, 337, 338, 341, 359, 376 Discoaster, 7, 248, 249, 258, 266, 267 Dissolution, 245, 246, 247, 250, 254, 255, 258, 259, 264, 273, 275, 277, 278, 280 Distinction between diagenesis and incipient metamorphism, 295-297 Disturbed Bed, Montana, 166 Dniepr-Donetz Depression, U.S.S.R.,33 I , 332, 333, 335, 337, 339 Dogger coal, 456 Dolomite, 11-13, 20, 28. 58, 122, 135, 439, 47 1 -, origin, 11-13, 79-85 Dolomitization, 19, 77, 78-85, 92, 93 Dominican Republic, 122 Douala Basin, Cameroun, 3 1 1, 3 12. 3 13, 3 14, 319, 324, 334, 355

566 Donbas area, U.S.S.R., 331, 332, 333, 335, 339, 342, 364 Dutch West Indies, 79 Dynamic metamorphism, 32

French Guiana, 122 French Jura, 85, 135, 324 French-Italian Alps, 393, 395, 468 Funafuti Atoll, 82, 83

Early burial diagenesis, 37-41 East Africa, 137 Echnoid trails, 239 Egypt, 78 Eifel district, Germany, 84 Ekofisk field, 276 England, 75, 85, 94, 264, 275, 315 Eniwetok Atoll, 81 Enterolithic structures, 65 Eocene, 67, 137, 141, 142, 146, 147, 260, 265, 361, 458, 469, 470 Eocene to Pleistocene shales, U.S. Gulf Coast, 161, 162, 163, 164, 168, 170, 171, 173, 175, 177 Eodiagenesis, 183, 184 Eogenetic stage, 4, 35 Eolianites, 68 Epidiagenesis, 4, 17, 34, 35, 42, 43-46, 58, 64, 65, 73, 77, 84-85, 87, 88, 90, 92, 95, 96 Epigenesis, 18, 93, 289, 290, 294, 296. 301, 304, 331, 332, 333, 336, 340, 376 Epidote, 22, 291, 303, 305, 341, 370, 379, 381, 386, 391, 392, 393, 394, 395, 396, 399, 400, 403, 406, 409, 412, 414, 427, 429, 430, 431, 435, 439, 440, 447, 448, 449, 451, 452, 453, 454 Equatorial Atlantic, 238 Equatorial Pacific, 236, 237, 238, 257 Erionite, 136, 138, 422, 423 Euxinic conditions, 36, 37, 59, 60 Exogyra, 68 Expandable clays, in burial-metamorphic facies. 458-46 1

Galapagos Islands, 39 Galena, 20, 29, 39, 152 Garnet, 305, 340, 359, 391, 392, 395 Gibbsite, 119, 123 Geochemical changes with depth, carbonates, 258-266 Geochemical parameters, 23-30 Geophysical parameters, 30-33 Georgia, 75 Geothermal gradients, 384, 385 Geothermal metamorphism, 83 Germany, 154, 161, 172, 178, 190, 192, 324, 334, 335, 336, 337, 338, 360, 362, 368, 373, 503 Glarus Alps, 343, 344, 352, 353, 360, 361, 364, 365, 366, 368, 371, 372, 373, 405, 470 Glauconite,5, 20,28,51-52,90, 131, 138-140, 177, 303, 33 1, 332, 340, 37 1, 372, 469, 470, 495, 496 Glaucophane, 306, 308, 309, 343, 370, 379, 430, 443, 456, 467, 468, 472 Glaucophanitic terranes, 391-396 Globigerina ooze, 40, 122, 253 Globoquadrina hexagona, 22 I Goethite, 144, 158 Graphite, 182 Great Barrier Reef, 82 Great Britain, 503 Greenalite, 20 Green River Formation, Utah, 137 Grossular, 436, 450 Guadaloupe Mountains, New Mexico, 65 Guano, 90 Guinea, 139 Gulf Coast, 276, 277, 312, 314. 318, 319, 320, 321, 328 Gulf of Mexico, 57, 139, 151 Gulf of Naples, 124, 142, 149 Gulf of Trinidad, 139 Giimbelite, 343 Gunflint Formation, 69 Guyana, 95 Gypsum, 20.28, 37,45,62,63-65.79. 85.88, 91. 293

Facies series, lowest-grade metamorphism, 441 -454 Fig-Tree Formation, 69 Flagellata, 68 Flint, 40, 67. 69, 71, 73, 74 Flocculation intensity, 127 Florida, 85 Fluorapatite, 90 Fontainebleau sandstone crystallization, 88 Foraminifera, 220-22 1 France, 312, 319, 335, 336, 337, 354, 357, 358, 364, 370, 437, 461, 503 Franciscan Formation, California, 274 French Alps, 334, 335, 343, 355, 356, 358, 361, 364, 366, 373, 393, 456, 457, 459, 465, 470, 473

Haiti, 122 Halite, 20, 28, 62, 82, 83 Halmyrolysis, 50-58, 116, 117, 132-138. 139, 141, 142, 144, 145

567 Hardgrounds, 73, 265, 266 Harrisburg, Pennsylvania, U.S.A., 80 Hassi Messaoud, Sahara, 337 Hawaii, 123, 430 Hectorite, 137 Helvetic zone, Switzerland, 343, 344, 360, 371, 372, 373, 388, 414, 455, 456, 465, 466,468, 469, 473 Hematite, 20, 28, 62, 63, 94, 144, 304, 500, 502, 503, 504 Hercynian, 356, 358, 370 Heulandite, 21, 53, 143, 203, 291, 299, 303, 304, 341, 376, 378, 381, 382, 383, 384, 389, 396, 397, 398, 399, 400, 401, 402, 403, 404, 405, 406, 407, 408, 409, 410, 417, 421, 422, 425, 426, 427, 431, 432, 434, 436, 441, 448, 454, 458, 460 461 Holocene, 64, 78, 79, 85, 122 Honshu, Japan, 10, 324, 383, 384, 385, 388, 390, 391, 392, 393, 394, 397, 398, 399, 401, 403, 416, 418, 421, 422, 426, 427, 429, 443, 444, 445, 446, 447, 448, 456, 458, 459, 460, 461, 468 Hornblende, 390, 404, 405, 460, 464 Hydration-dehydration, 63-76 Hydrobiotite, 363 Hydrocarbons, 69, 182, 183 Hydrolysis, 29, 48, 74, 75, 95 Hydromagnesite, 20 Hydrotroilite, 151, 152 Hystnchosphaeridae, 68 Iceland, 121, 431 Illite, 5 , 6, 9, 10, 20, 54-56, 58, 74, 75, 89, 121, 122, 126, 131, 133, 134, 135, 137, 138, 139, 143, 145, 152, 153, 176, 177, 178 181, 182, 183, 195, 199, 201, 203, 290, 291, 293, 298, 300, 303, 306, 307, 308, 309, 310, 31 1, 312, 313, 314, 315, 316, 317, 318, 319, 320, 321, 323, 324, 325, 327, 332, 333, 334, 335, 336, 338, 341, 343, 357, 359, 360, 363, 364, 366, 367, 368, 369, 371, 372, 373, 374, 375, 376, 378, 395, 458, 459, 461, 462, 464, 465, 466, 468,469,470,472, 473 -, composition during burial diagenesis, 328, 333 -, “crystallinity” in the anchimetamorphic zone, 344-358 -, diagenesis, 159-175 Illite-muscovite in incipient metamorphism, 342 Ilmenite, 75 India, 87 Indian Ocean, 119, 121, 122, 126, 150

Induration, 228, 230 Initial porosity of argillaceous sediments, 124-127 Initial stage of syndiagenesis, 36-37 Initial structure of argillaceous sediments, 128-130 Interstitial water, chemistry, 150-151, 154159 Ionic adsorption phenomena, 29-30,48 Iowa-Missouri, 153 Ireland, 264 Irkutsk Coal Basin, Kamchatka, Siberia, 459, 462 Israel, 71, 135, 275 Italy, 142, 149 Jadeite, 392, 396, 409, 452, 456, 472 Jamaica, 85 Japan, 66, 337, 368, 382, 383, 385, 435, 462 Jerusalem, 69, 70 John Day Formation, Tertiary, Oregon, 179 Jordan-Dead Sea Rift, 69 Judea, 69, 70 Jurassic, 68, 90, 93, 94, 152, 154, 216, 221, 256, 312, 319, 326, 334. 335, 343, 344, 354, 355, 356, 357, 361, 366, 373, 374, 400, 404, 455, 458, 465, 473 Juvenile water. 76 Kansas, 337 Kaolinite, 6, 9, 10, 20, 55, 56, 57-58, 95, 119, 122, 126, 127, 134, 153, 161, 173, 179, 181, 182, 194, 198, 199, 201, 290, 293, 301, 303, 304, 306, 307, 310, 311, 314, 315, 317, 319, 321, 322, 323, 325, 326, 330, 331, 338, 340, 341, 345, 357, 359, 360, 362, 363, 369, 374, 375, 376, 389, 412, 432, 436, 437, 439,440, 465, 467, 470, 503 Kaolinite, alteration during burial diagenesis, 334-338 - and chlorite diagenesis, 175-178 -, in burial-metamorphic facies, 462, 463 Karaganda Basin, Kazakstan, 3 19 Karagana Basin, U.S.S.R., 338 Karst, 73, 275 Katamorphic zone, 46 Keuper formation, Germany, 178 K-feldspar, 6, 10, 328, 329, 341, 376, 391, 419, 422, 423, 440, 465 -, decomposition, 463-464 Kinnekulle, Sweden, 315, 316, 319 Kirkwood Miocene Formation, 153 Kiruna type, iron ore, 495,496 Krakatau, 123

568 Kuwait, 64 Kyanite, 305 Lake Chad, 131 Lake Constance, 125. 128, 131, 152, 187 Lake Magadi, Kenya, 136, 137 Lake Malawi, 131 Lake Manyare, East Africa, 138 Lake Superior ore type, 495, 496 Lake Tecopa, California, 136, 137 LakeZiirich, 116, 117, 118. 187, 188, 189 Laterized soils, 88 Laumontite, 10, 21, 53, 143, 203, 291, 299, 303, 304, 306, 307, 324, 341, 376, 378, 381, 382, 383. 384, 385, 388, 389, 391, 396, 397, 398, 399, 400, 401, 402, 403, 404, 405, 406, 407, 408, 409, 410, 411, 412, 413, 414, 416, 417, 418, 419, 422, 425, 426, 427, 428, 429, 431, 432, 434, 435, 436, 437, 439, 440, 441, 444,448, 453,455, 458, 459,460, 461, 463, 464, 465, 466, 467, 471, 472, 473 Laumontite zone, 464-466 Lawsonite, 303, 306, 308, 309, 343, 370, 392, 393, 395, 396, 405, 406, 409, 410, 411, 443, 447, 448, 453, 456, 467, 468, 471, 472, 473 Ledikite, 131, 179 Lena Coal Basin, Yakutia, Siberia, 383, 455, 462 Leucoxene, 20, 301 Liassic, 154, 190, 192, 353, 357, 360, 361, 364, 368, 370, 371 Liesegang rings, 63 Limonite, 20, 28, 62, 63, 94, 144, 501, 503, 504, 506 Limestone stage, 269 Lithification, 18- 19 Load metamorphism, 32,42 Locomorphic stage, 34 Louisiana, 161, 313, 315 Lower Hunter Valley, New South Wales, 317, 326, 327 Low-grade metamorphism, 5, 115 Lowest-grade metamorphism, recognition, 379-380 Lowest-grade mineral facies, 375-378 Magadiite, 137 Magnesioriebeckite, 37 1 Magnesite, 20 Magnetite, 94, 500, 502, 503, 507 Maine, U.S.A., 390, 462 Malachite, 20 Manganese nodules, 5, 59, 150 Marcasite, 20, 32, 40, 44,61, 62, 81, 152

Margarite, 374 Mediterranean, 124, 135, 152 Melnikovite, 62 Mesodiagenesis, 183, 184 Mesogenetic stage, 4, 35 Mesopotamian Basin, 64 Mesozoic, 55, 65, 88, 146, 152, 176, 334, 371, 470 Metabentonite, 134 Meta-anthracite, 456 Metagenesis, 31, 289, 294, 296, 301, 305. 322, 331, 332, 336, 338, 376 Metamorphic versus sedimentary rocks, 297299 Metasomatism, 19, 21, 32, 77-96 Metharmosis, 294 Methane, 182 Meulerization, 88 Mexico, 90 Micarb, 244, 270, 271, 273 Mineralogical changes in burial diagenesis, 290-29 1 Minnesota, U.S.A., 74 Miocene, 95, 141, 253, 260, 265, 313, 315, 327. 328, 382, 397, 459 Miocene-Pliocene, 423 Mississippian. 327 Mississippian Greenbrier Dolomite, 79 Mississippi Delta, 125, 328 Model of progressive diagenesis, 268-275 Moisture content, argillaceous sediments. 193-198 Montana, U.S.A., 75 Montmorillonite, 20. 55, 56-57, 58, 74, 76, 122, 126, 127, 137, 140, 143, 152, 155. 171, 174, 178, 179, 181, 182, 193, 194, 195, 201, 203, 290, 291, 293, 301, 303, 304, 306, 307, 311, 312, 313, 314, 315, 316, 317, 318, 319, 320, 321, 322, 323. 324, 325. 326, 327, 328, 330, 331, 332, 333, 334, 335, 337, 338, 340, 341, 342, 352, 357, 359, 363. 365, 366, 368, 376, 389, 399, 412, 422, 423, 432, 433, 439, 458, 459, 460, 461, 462, 465, 470 Morcles area, Switzerland, 37 I Mordenite, 143, 291, 379, 383, 384, 396, 398, 399, 401, 406, 414, 418, 420, 421, 422, 423, 426, 441, 458, 460 Moroccan Meseta, 357, 358 Morocco, 93 Muscovite, 6, 10, 20, 74, 131, 181. 291, 305, 3 10. 3 18, 322, 323, 326, 327, 329, 330, 33 I , 332, 333, 344, 351, 363, 364, 367, 369, 376. 386, 390, 391, 467, 468 -, alteration during burial diagenesis, 340 M6ns Klint, Denmark, 71

569 Natrolite, 21 Natron, 20 Natural chromatography, 43, 54, 76-77 Navassa Island, West Indies, 92 Nazca plate, 144 Neogene, 137, 145, 383, 398 Neogenesis, 2 1 Nesquehonite, 20 Neuse River, North Carolina, 55 New England, 403 Newfoundland, 75, 162 New Jersey, U.S.A., 153 New South Wales, Australia, 382, 383, 388, 391, 398, 400, 403, 445, 446, 447, 449, 454, 455, 462 Nevada, U.S.A., 383, 423 New York, U.S.A., 94 New Zealand, 10, 314, 325, 338, 368, 382, 385, 386, 387, 388, 390, 391, 394, 400, 403, 410, 416, 417, 425, 426, 427, 428, 430, 431, 435, 440, 442, 445, 446, 447, 448, 455, 458, 459, 461, 462, 464, 466, 468, 469, 472, 473 Niger river, 58 Nontronite, 56, 131, 143 Normandy, 68 North Africa, 8, 56, 64 North America, 75, 81 North Atlantic, 55, 57, 75, 141, 264 North Pacific, 145 North Sea, 37, 127, 264, 276, 277, 364 Northwest Territories, Canada, 177 Norway, 75 Oklahoma, U.S.A., 312 Oligocene, 171, 253, 260, 265, 313 Oligoclase, 376, 388, 403, 437 Oman, 90 Ontong-Java Plateau, Equatorial Pacific, 228, 230, 232, 233, 234, 238, 240, 251 Onvenvacht Formation, 69 Oolitic iron ore, 497-499 Ooze stage, 269 Opal, 20, 65, 66, 87, 142, 143, 145, 146, 147, 148, 180, 423, 431 Ordovician, 68, 74, 75, 80, 85, 93, 162, 177, 178, 351, 360, 362, 370 Oregon, U.S.A., 388, 400, 403, 404, 436, 461 Organic and mineral diagenesis, correlation, 182- 185 Organic coatings, 235, 236 Organic metabolism, 18, 23 Orinoco, 58, 59 Orthoclase, 21, 75, 122, 390, 398, 464 Orthogeosynclines, 42

Oslo, 129 Ouachita Belt, Texas and Oklahoma, 345 Overpressured zones, 276 Oxford clay, England, 152, 154 Oxidation-reduction, 49, 59-63 Oxygen isotope data, 12, 13, 259-266, 280, 28 1

Pacific Ocean, 5 5 , 57, 58, 59, 119, 121, 122, 123, 126, 128, 138, 140, 143, 145, 150, 218, 245, 253, 263, 275, 276 Palagonite, 5, 123, 124, 144 Paleocene, 260, 265 Paleosols, 68 Paleozoic, 5 5 , 56, 57, 58, 75, 81, 82, 88, 135, 139, 153, 176, 326, 327, 329, 345, 350, 355, 356, 358, 370, 374, 382, 383, 465 Palygorskite, 5 , 121, 135, 138, 179 Paragenesis of iron-bearing minerals, 500-504 Parageosynclines, 42 Paragonite, 181, 203: 291, 303, 310, 335, 342, 352, 359, 361, 362, 374, 376, 465, 467, 470, 47 1, 472 Paragonitic mica, appearance during incipient metamorphism, 364-369 Paris, 73 Pans Basin, 75, 88, 89, 334, 355 Pavement mosaic, 7 Pelvoux Massif, French Alps, 324, 364 Pennsylvania, U.S.A., 364, 373 Pennsylvanian, 153, 343, 361 Periglacial, 89 Permeability, 43, 61, 251, 252, 277, 281, 417, 418 Permian, 57, 90, 315, 316, 317, 320, 326, 327, 335, 337, 383, 405, 455, 473 Permian Phosphorite Formation, 93 Permo-Triassic, 84, 134, 334, 355, 472 Persian Gulf, 63, 64, 78, 79, 82 Perth, Western Australia, 32 Peru, 90 Petchora Basin, U.S.S.R., 321 Petroleum, evolution of, 76 Petroleum-generating potential, 3 Petroleum hydrocarbons, 18, 61, 62 Petroleum potential, 276-277 Phanerozoic, 67, 68, 183 Phengite, 201, 203, 332, 344, 352, 353, 359, 362, 365, 370, 371, 376, 467, 468, 469, 470, 47 1 Phillipsite, 21, 52-53, 136, 140, 141, 142, 144, 145, 422, 423 Phosphatization, 77, 90-93 Phosphorite, 28

570 Photosynthesis, 79, 81, 90, 94, 95 Phreatic water, 44 Phyllomorphic stage, 34 Piemontite, 392 Placoliths, 272 Pistacite, 303 Plagioclase, 10, 21, 291, 299, 301, 310, 376, 379, 381, 382, 391, 398, 399, 401, 402, 404, 405, 418, 428, 430, 431, 437, 439, 441, 464 alteration during burial diagenesis, 341 Pleistocene, 136, 137, 238, 419 Pliocene, 95, 171, 260, 265, 312 Plio-Pleistocene, 3 13 Po Basin, 118, 190, 192 Podzolization, 89 Poland, 68 Polignac Basin, Algerian Sahara, 355 Porcellanite, 88, 146, 147, 148, 180 Porosity, 118, 187-193, 226, 227, 229, 231, 232, 241, 243, 244, 251, 252, 253, 254, 261, 264, 268, 269, 270, 271, 272, 273, 276, 277, 279, 281, 301, 417, 418 Precambrian, 11, 12, 13, 56, 61, 67, 75, 80, 81, 88, 94, 95, 360 Precambrian Belt Series, Idaho, 329 Prehnite, 9, 10, 143, 289, 303, 305, 306, 307, 309, 338, 341, 359, 376, 378, 379, 381, 382, 383, 385, 386, 388, 389, 390, 391, 393, 394, 395, 396, 400, 402, 403, 404, 405, 406, 412, 414, 427, 428, 429, 430,431,432, 433, 434, 435, 436, 437, 439, 442,443,444, 447, 448, 449, 450, 451, 452, 453, 454, 456, 457, 459, 460, 461, 464, 466, 467, 468, 469, 472, 473 Prehnite-pumpellyite-metagreywacke facies. 385-388 Pressure-solution, 254, 255, 256, 271, 273, 278, 280 Proterozoic, I76 Proto-calcarenites, 40 Proto-calcilutite, 40 Protodolomite, 138 Proto-glauconite, 139 Psilomelane, 2 1 Pteropods, 221-222, 238 Ptilolite, 383 Puerto Rico, 382, 383, 388, 400, 403, 431 Pumpellyite, 9, 10, 11, 143, 289, 291, 299, 303, 306, 307, 309, 338, 341, 359, 361, 370, 371, 376, 378, 379, 381, 382, 383, 385, 386, 388, 389, 390, 391-396, 400,403,404,405,406, 412, 414, 427, 428, 429, 430, 431, 434, 435, 436, 439, 441, 442, 443, 444,447, 448, 449, 450, 45 1, 453, 454, 456, 457, 458, 459, 460, 461, 463, 464, 466, 467, 468, 469, 471, 472, 473

-.

Pumpellyite-actinolite-schist facies, 388-391 Pumpellyite facies, 385-396, 466-472 Pynochlorite, 338 Pyrenees, 325, 352, 356, 357, 358, 360, 364, 367, 368, 370 Pyrite, 21, 28, 32, 38, 40, 44, 45, 46, 61, 62, 81, 91, 94, 151, 152, 154, 159, 502 Pyritized fossils, 6 1 Pyroclastic material, 123- 124 Pyrolusite, 2 I Pyrophyllite, 10, 121, 122, 173, 201, 291, 303, 305, 306, 307, 309, 310, 322, 323, 332, 333, 335, 336, 340, 342, 353, 365, 366, 367. 368, 369, 370, 373, 374, 375, 376, 436, 465, 467. 470 -, formation of, 359-363 Pyroxene, 293, 301, 306, 390, 457 Pyrrhotite, 29, 502

Quartz, 6, 10, 21, 58, 65, 66, 76, 86, 87, 88, 89, 122, 142, 143, 146, 147, 148, 149, 168. 169, 173, 180, 201, 203, 299, 303, 304, 305, 306, 323, 336, 360, 362, 363, 369, 370, 376, 38 1, 386, 390, 391, 399, 404, 407, 408, 409, 410, 41 1, 412, 414, 417, 419, 420, 421, 424, 425, 432, 433, 434, 437, 439, 440, 448, 449, 450, 451, 452, 453, 454, 501, 503, 504, 505, 506, 507 Quaternary, 58, 68, 91, 92, 95, 121, 171, 260, 312 Rank of coal, associated with mineral facies, 454-457 Reaction series, 180- 182 Recrystallization, 19, 21 Rectorite, 10, 181, 291, 309, 319, 337, 338, 342, 352, 359, 363, 365, 366, 367, 368, 369, 373, 374, 375, 376, 378, 465, 467, 469 Red beds, 59 Red clay, 59 Red Sea, 39, 78, 82 Replacement, 19, 21, 67, 75, 77, 87 Redoxomorphic stage, 34 Rhabdolith, 255 Rhaetian, 47 1 Rheinische Schiefergebirge, 357, 358, 360, 374, 465, 468 “Rhexistasy”, 44, 45, 68 Rhine, 131 Rhinegraben, 161, 170, 172, 174, 189, 191,200 Rhizoconcretions, 68 Rhizomorphs, 68 Rhirophora, 37 Rhodochrosite, 21, 28

57 1 Rhodusite, 361, 371 Riebeckite, 371, 372, 470 Rio Ameca Basin, Mexico, 133 Rockport area, Texas, 125 Ruhr area, Germany, 87, 336 Russian Platform, 333, 359, 360, 371, 468 Rutile, 21, 75 Sahara, 88, 351, 367, 373 Salton Sea, California, 171, 312, 313, 314 Sanbagawa, Japan, 343, 344 Sanidine, 122 Saponite, 56, 141, 143, 144, 460, 461 Santa Barbara Basin, 117, 118, 188 Saussurite, 341 Schizohalinity, 7 1 Scolecite, 341, 406 Scotian Shelf of Canada, 276 Scotland, 75, 153, 439 Sedimentation rate, 3, 261 Sediment-water interface, 226-241 Seismic reflectors, 228, 230 Sepiolite, 5, 138, 179 Sericite, 178, 301, 303, 332, 340, 341, 392, 460 Serpentine, 121 Shallow-burial diagenesis, 150- 154, 241-25 1 Shikoku, Japan, 394 Shinjo Basin, Honshii, Japan, 383 Siberia, 326, 327, 462 Siberian platform, 179 Sicily, 65, 315 Siderite, 21, 28, 37, 93, 94, 95, 154, 501, 502, 503, 504, 506, 507 Siderite facies, 93 Sideritization, 93-95 Sideromelane, 430, 43 1 Sidney Basin, Australia, 153 Silcrete, 88 Silica gel, 65-74 Silica in argillaceous sediments, 145- 149 Silicification, 77, 86-89 Silurian, 73, 93, 94, 135, 359, 360, 367, 373 Slump structures, 69, 70 Smectite, 5, 6, 9, 10, 67, 119, 121, 135, 137, 138, 139, 140, 141, 144, 145, 153, 156, 158, 161, 162, 164, 166, 168, 170, 171, 172, 173, 177, 180, 195, 198, 310, 334, 338, 341, 351, 374, 375, 378, 381, 461. 473 -, replacement upon burial, 310-328 Somalia, 90 Sonic velocity, 229, 231, 234, 235 Soudan formation, 69 South Africa, 68, 69, 88, 90, 139 South America, 95

South Atlantic, 141 South Carolina, U.S.A., 153 Spessartine, 301, 362 Sphaerosiderite, 495, 496 Sphalerite, 21, 39, 152 Sphene, 22, 76, 381, 390, 400, 464 Sphenolithm moriformis, 27 1 Spiroraphe, 240 Spitsbergen, 75 Stable oxygen isotopes, 259-266 Stassfurt deposits (Permian), evaporites, north Germany, 32 Static metamorphism, 32 Staurolite, 293 Stevensite. I37 Stevns Klint. Denmark, 71 Stilbite, 382, 406, 460 Stilpnomelane, 303, 305, 306, 359, 361, 372, 374, 376, 388, 391, 392, 394, 395, 443, 467, 468, 470 -, occurrence during incipient metamorphism, 370-37 I Stromatolites, 88, 94 Strontianite, 21 Strontium, 258-259, 262 Stylolite, 89, 255 Sudoite, 152, 178, 337, 338 Sulphur, 2 1 Sulfides, formation of, 151- 152 Sulfur bacteria, 61 Sunda Straits, 123 Supersaline lake environment, 135- I38 Supersaline marine environment, 134- 135 Surface weathering, 42 Surinam, 95, 122 Sverdrup Basin, Northwest Territories, Canada, 184 Sweden, 322, 357, 465 Switzerland, 357, 370, 371, 456, 459, 468, 469, 472, 473 Syndiagenesis, 4, 17, 33, 34, 35-41, 45, 61, 63, 64, 65, 72, 74, 77, 78-82, 83, 84, 86, 87, 91, 93 Talc, 121, 144, 179, 450 Tanzawa Mountains, Honshii, Japan, 383 Taringatura, Southland, New Zealand, 381, 383, 410 Taupo Volcanic Zone, New Zealand, 3 12 Taveyanne Graywacke, 371, 388, 403, 437, 455, 456, 459, 464, 466, 469 Telodiagenesis, 183, 184, 185 Telogenetic stage, 4, 35 Terra rossa, 73

572 Tertiary, 64, 65, 88, 90, 118, 142, 143, 177, 190, 192, 195, 200, 250, 264, 312, 318, 338, 361, 371, 383, 403, 455, 456, 471 Texas, 65, 275, 276, 313, 320 Texas Gulf Coast, 32, 56 Thalassocratic stage, 68 Thomsonite, 34, 460, 462 Tilton Shale Member, 137 Tosudite, 325, 337, 338 Tourmaline, 2 I , 76 Travertine, 68 Triassic, 135, 153, 324, 343, 344, 353, 357, 360, 371, 373, 381, 382, 383, 400, 431. 455, 462, 465, 470 Tridymite, 145 Troodos Massif of Cyprus, 146 Tuarnotu Archipelago, 222 Tuff, 53 Tunguska area, Siberia, 335 Turbidites, 64 Turkey, 137, 393 Two Medicine Formation, Montana, 316 Tyrrhenian Sea, 142 Ural, 327, 328 Utah, 368 Vadose water, 44 Vancouver Island, British Columbia, 403, 410, 427, 428, 431, 435, 436, Velocity, acoustic, 241, 251, 252, 253, 273, 275 Venezuela, 122, 190, 192 Verkhoyano-Kolimskaya syncline, 179 Verkhoyanye, Siberia, 301 Vermiculite, 135, 137, 177, 179, 290, 310, 311, 324, 338, 339, 340, 363, 366, 459, 460 Victoria, 455 Vivianite, 92 Volcanic and lithic-feldspathic sediments, 378-474

Wabana deposits, Newfoundland, 94 Wairakite, 341, 382, 384, 385, 391, 397, 406, 407, 41 1. 412, 413, 414, 416, 419, 426, 427, 428, 429, 431, 440,448, 449, 454, 460 Wakatipu metamorphic belt, New Zealand, 387 Washington, 436 Waterbloom, 61 Water column, changes, 224-226 Wavellite, 92 Weathering, 115, 116, 117, 119, 134, 144, 159, 179, 302 Wisconsin, U.S.A., 85 Witherite, 21 Wollastonite, 437 Wurtzite, 152 Wyoming, U.S.A., 419, 425,458 Yagawralite, 384 Yangtze river, 59 Yenisey range, Siberia, 360, 369 Yucatan, Mexico, 85 Zechstein, 136, 324 Zeolites, 5 , 21, 67, 75, 135, 136, 137, 138, 140-143, 203, 289, 303 Zeolite facies, 53, 291 _ _ , recognition, 381-385 _ _ , progressive mineral facies, 406-418 - zonation, controlled by composition of fluid phase, 418-441 Zeolitization of plagioclase, 402 Ziron, 21, 76 Zoisite, 22, 395, 406, 432, 433, 436, 439, 450 Zuger See, 128, 187 Zuiderzee, 125