Green Marine Clays (Developments in Sedimentology)

DEVELOPMENTS IN SEDIMENTOLOGY 45 Green Marine Clays Oolitic Ironstone Facies, Verdine Facies, Glaucony Facies and Cela...

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DEVELOPMENTS IN SEDIMENTOLOGY 45

Green Marine Clays Oolitic Ironstone Facies, Verdine Facies, Glaucony Facies and Celadonite-Bearing Facies - A Comparative Study

FURTHER TITLES IN THIS SERIES VOLUMES 1-1 1, 13-1 5 and 2 1-24 are out of print 12 R.C. G. BATHURST CARBONATE SEDIMENTS AND THEIR DIAGENESIS 13 H.H. RlEKE Illand G. V. CHlLlNGARlAN COMPACTION OF ARGILLACEOUS SEDIMENTS 17 M .0. PICARD and L . R. HIGH Jr. SEDIMENTARY STRUCTURES OF EPHEMERAL STREAMS 18 G. V. CHlLlNGARlAN and K.H. WOLF, Editors COMPACTION OF COARSE-GRAINED SEDIMENTS 19 W . SCHWARZACHER SEDIMENTATION MODELS AND QUANTITATIVE STRATIGRAPHY 20 M.R. WALTER, Editor STROMATOLITES 25 G. LARSENand G. V. CHILINGAR, ,Editors DIAGENESIS IN SEDIMENTS AND SEDIMENTARY ROCKS 26 T. SUDOand S. SHIMODA, Editors CLAYS AND CLAY MINERALS OF JAPAN 27 M . M . MORTLANDand V.C. FARMER, Editors INTERNATIONAL CLAY CONFERENCE 1978 28 A . NISSENBAUM, Editor HYPERSALINE BRINES AND EVAPORlTlC ENVIRONMENTS 2 9 P. TURNER CONTINENTAL RED BEDS 3 0 J.R.L. ALLEN SEDIMENTARY STRUCTURES 3 1 T. SUDO,S. SHIMODA, H. YOTSUMOTO and S. AlTA ELECTRON MICROGRAPHS OF CLAY MINERALS 3 2 C.A. NITTROUER, Editor SEDIMENTARY DYNAMICS OF CONTINENTAL SHELVES 33 G.N. BATURIN PHOSPHORITES ON THE SEA FLOOR 3 4 J.J. FRIPIAT, Editor ADVANCED TECHNIQUES FOR CLAY MINERAL ANALYSIS 35 H. VAN OLPHENand F. VENIALE, Editors INTERNATIONAL CLAY CONFERENCE 1981 36 A . IIJIMA, J.R. HElN and R. SIEVER, Editors SILICEOUS DEPOSITS IN THE PACIFIC REGION 37 A . SlNGERand E. GALAN, Editors PALYGORSKITE-SEPIOLITE: OCCURRENCES, GENESIS AND USES 3 8 M.E. BROOKFlELDand T.S. AHLBRANDT, Editors EOLIAN SEDIMENTS AND PROCESSES 3 9 B. GREENWOODand R.A. DAVIS Jr., Editors HYDRODYNAMICS AND SEDIMENTATION IN WAVE-DOMINATED COASTAL ENVIRONMENTS 4 0 B. VELDE CLAY MINERALS - A PHYSICO-CHEMICAL EXPLANATION OF THEIR OCCURRENCE 4 1 G. V. CHlLlNGARlAN and K.H. WOLF, Editors DIAGENESIS, I 42 L.J. DOY’ E a i d H.H RORERTS, Editors CARBONATE-CLASTIC TRANSITIONS 43 G. V. CHlLlNGA RIA N and K. H. WOLF, Editors DIAGENISIS, II 4 4 C.E. WEAVER CLAYS, MUDS AND SHALES

DEVELOPMENTS IN SEDIMENTOLOGY 45

Green Marine Clays Oolitic Ironstone Facies, Verdine Facies, Glaucony Facies and Celadonite-Bearing Facies - A Comparative Study

Edited by G.S. ODlN Mairre de Recherche - CNRS Universire Pierre er Marie Curie, Paris

ELSEVIER Amsterdam - Oxford - N e w York - Tokyo

1988

ELSEVIER SCIENCE PUBLISHERSB.V. Sara Burgerhartstraat 25 P.O. Box 21 1, 1000 AE Amsterdam, The Netherlands Distributors for the United States and Canada: ELSEVIER SCIENCE PUBLISHINGCOMPANY INC. 655, Avenue of the Americas New York, NY 10017, U.S.A.

Library of Congress Cataloging-in-Publication Data

Oolitic ironstone facies, verdine facies, glaucony facies, and celadonite-bearing facies. (Developments in sedirnentology ; 4 5 ) Bibliography: p. Includes index. 1. Clay minerals. 2. Facies (Geology) 3 . Submarine geology. I. Odin, Gilles S . 11. Series. QE389.625.055 1988 552l.5 88-3 1042 ISBN 0-444-87120-9

ISBN 0-444-87 120-9 (Vol. 45) ISBN 0-444-4 1238-7 (Series) 0 Elsevier Science Publishers B.V., 1988

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 Science Publishers B.V./ Physical Sciences & Engineering Division, P.O. Box 330, 1000 AH Amsterdam, The Netherlands. Special regulations for readers in the U.S.A. -This publication has been registered with the Copyright Clearance Center Inc. (CCC), Salem, Massachusetts. Information can be obtained from the CCC about conditions under which photocopies of parts of this publication may be made in the USA. All other copyright questions, including photocopying outside of the USA, should be referred to the publisher. No responsibility is assumed by the Publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Printed in The Netherlands

V

CONTENTS

Contributors Detailed contents of the volume Foreword (in French) by Georges Millot, de 1' AcadCmie des Sciences English translation

VIT IX XVII XXI

Introduction to the study of green marine clays

1

Part A THE OOLITIC IRONSTONE FACES

Introduction and contents of Part A *A1 Chamosite, the green marine clay from Chamoson by M.F. Delaloye and G.S. Odin *A2 Green Marine Clays from the oolitic ironstone facies by G.S. Odin, R.W.0.B' Knox, S . Guerrak and R.A. Gygi

5 7 29

Part B THE VERDINE FACES

Introduction and contents of Part B *B1 The verdine facies from the lagoon off New Caledonia by G.S. Odin *B2 The verdine facies from the Senegalese continental shelf by G.S. Odin and J.P. Masse *B3 The verdine facies off French Guiana by G.S. Odin, I.D.R. Mackinnon, and M. Pujos *B4 The verdine facies deposits identified in 1988

53

by G.S. Odin, J.P. Debenay and J.P. Masse *B5 Mineralogy of the verdine facies by G.S. Odin, S.W. Bailey, M. Amouric, F. Frohlich and G.A. Waychunas *B6 Geological significance of the verdine facies by G.S. Odin and B.K. Sen Gupta

131

57

83 105

159 205

Part C THE GLAUCONY FACES Introduction and contents of Part C *C1 Glaucony from the Gulf of Guinea

22 1

by G.S. Odin *C2 Glaucony from the margin off northwestern Spain by G.S. Odin and M. Lamboy *C3 Glaucony from the Kerguelen Plateau by G.S. Odin and F. Frohlich *C4 Geological significance of the glaucony facies by G.S. Odin and P.D. Fullagar

225 249

277 295

VI

Part D THE CELADONITE-BEARING FACES

333 Introduction and contents of Part D *D Nature and geological significance of celadonite by G.S. Odin, A. Desprairies, P.D. Fullagar, H. Bellon, A. Decarreau, 337 F. Frohlich, and M. Zelvelder Conclusion to the study of green marine clays Acknowledgements

399 405

Glossary

407

References

419

Index of collaborators

44 1

Geographical index

443

VII

CONTRIBUTORS

M. Amouric

Centre de Recherche sur les mCcanismes de la croissance cristalline, Case 913, F13288 Marseille Cedex 9, (France) S.W. Bailey Department of Geology and Geophysics, University of Wisconsin 1215W, Dayton St, Madison, WI 53706, U.S.A. H. Bellon Laboratoire de GCochimie, G.I.S. OcCanologie , 6 Avenue Le Gorgeu, F29287 Brest Cedex, France. J.P. Debenay Laboratoire de GCologie, UniversitC du Maine, F72017 Le Mans Cedex, France. A. Decarreau Laboratoire de PCtrologie de la Surface, UniversitC, 40 Avenue du Recteur Pineau, F86022 Poitiers Cedex, France. M. Delaloye DCpartement de MinCralogie, UniversitC de Genkve, 13, rue des Maraichers, CH 1211 Genkve, Switzerland. A. Desprairies Laboratoire de GCochimie sedimentaire, UniversitC de Pans Sud, B2timent 504, F91405 Orsay Cedex, France. F. Frohlich Laboratoire de GCologie, MusCum National d'Histoire Naturelle, 43 rue de Buffon, F75005 Paris, France. P.D. Fullagar Department of Geology, The University of North Carolina, Mitchell Hall CB 3315, Chapel Hill, NC 27599, U.S.A. S . Guerrak Institut de GCologie, UniversitC de Rennes I, Campus de Beaulieu, F35042 Rennes Cedex, France. R.A. Gygi Naturhistorisches Museum, Augustinergasse 2, CH 405 1 Basel, Switzerland. R.W.O'B. Knox British Geological Survey, Keyworth, Nottingham, NG 125GG, England, U.K. M. Lamboy Laboratoire de GCologie, UniversitC de Rouen, BP 118, F76134 Mont-Saint-Aignan Cedex, France. I.D.R. Mackinnon Department of Geology, The University of New Mexico, Northrop Hall, Albuquerque, NM 87131, U.S.A. J.P. Masse Laboratoire de Stratigraphie,UniversitC de Provence, Centre St-Charles, F13331 Marseille Cedex 3, France. G. Millot Institut de GCologie, UniversitC Louis Pasteur, 1, rue Blessig, F 67084 Strasbourg Cedex, France. G.S. Odin DCpartement de GCologie dynamique, UniversitC P. et M. Curie, 4, Place Jussieu, F75252 Paris Cedex 05, France. M. Pujos DCpartement de GCologie, UniversitC de Bordeaux I, Avenue des FacultCs, F33405 Talence Cedex, France. B.K. Sen Gupta Department of Geology and Geophysics, Louisiana State University, Baton Rouge, LA 70803, U.S.A. G.A. Waychunas Center for Materials Research, 105 McCullough Building, Stanford University, Stanford, CA 94305, U.S.A. M. Zelvelder DCpartement de GCologie dynamique, UniversitC P. et M. Curie, 4, Place Jussieu, F75252 Paris Cedex 05, France.

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IX DETAILED CONTENTS OF PARTS A, B, C, and D:

CONTENTS of PART A *Introduction to the oolitic ironstone facies *Contents of Part A *Chapter A1 Chamosite the green marine clay from Chamoson by M.F. Delaloye and G.S. Odin 1. Introduction 2. Petrography of the Swiss oolitic ironstone 2.1. Elements of the sediment 2.2. The groundmass of the oolitic sediment 3. Mineralogical data on the green clay 3.1. X-ray diffraction study 3.2. Chemical study 3.3.Thermal study 4. Discussion and conclusions on the genesis of the facies 4.1. Environment of deposition of the oolitic ironstone 4.2. Mineralogy of the green clay 4.3.Discussion on the genesis of the green clay from Chamoson 5. summary

5 6

7 9 9 14 15 15 19 22 24 24 25 26 28

*Chapter A2 Green marine clays from the oolitic ironstone facies: habit, mineralogy, environment by G.S. Odin, R.W. O'B. Knox, R.A. Gygi, and S. Guerrak 1. Introduction 29 30 2. The problem of the ooids 30 2.1. Interpretations of the ooids 31 2.2. The extra-sedimentary accretion hypothesis 32 2.3. The intra-sedimentary concretion hypothesis 34 2.4. The mid replacement hypothesis 2.5. The non-oolitic hypothesis 38 39 3. The green clay in the groundmass 4. Mineralogy of the ironstone green clays 40 40 4.1. Green clays of the ooids 42 4.2. Presumed initial green clays 5. Discussion of the environmental factors 44 44 5.1. General environment of oolitic ironstone formations 47 5.2. Microenvironment of green clay genesis in ooids 6. Summary 51

CONTENTS of PART B * Introduction to the verdine facies * Contents of part B

53 54

X

*Chapter B1 The verdine facies from the lagoon off New Caledonia by G.S. Odin 1. Geological setting 57 2. Sampling and sample treatment 59 3. Distribution and habit of the green pigment 62 3.1. Abundance and distribution 62 63 3.2. Habit of the green pigment 4. Mineralogy of the green pigment 66 4.1. X-ray diffraction study 66 69 4.2. Acid treatments of green pigment 4.3. Chemical study 73 5. Mineralogical nature and geological significance 75 75 5.1. Ease of alteration of the green pigment 77 5.2. Mineralogical nature of the green pigment 77 5.3. Environment of genesis of the phyllite V 79 5.4. Geological significance of the phyllite V 6. Conclusions 80 *Chapter B2 The verdine facies from the Senegalese continental shelf by G.S. Odin and J.P. Masse 1. Introduction 83 2. Sedimentology 85 85 2.1. The southern shelf (Baie de Rufisque) 2.2. The northern shelf 86 3. Study of the green pigment 89 89 3.1. Separation, proportion, and habit 93 3.2. Mineralogy of the green pigments 4. Discussion and conclusions 100 100 4.1. Sedimentology of the verdine facies off Senegal 101 4.2. Mineralogy of the verdine facies off Senegal 102 4.3. Microenvironment and genesis of the 'phyllite V' mineral 4.4. Conclusions 103 *Chapter B3 The verdine facies off French Guiana by G.S. Odin, I.D.R. Mackinnon and M. Pujos 1. Introduction 2. Geological setting of the shelf off French Guiana 3. Authigenic green grains 3.1. The glaucony facies 3.2. The verdine facies 4. Mineralogy of the verdine facies 4.1 .X-ray diffraction study 4.2. Analytical electron microscopy 4.3. Chemical study 5. Geological significance of the verdine facies 5.1. Age of the neoformation

105 106 109 109 110 112 112 115 123 124 124

XI

5.2. Recent history of the shelf 6. Summary and conclusions

125 129

*Chapter B4 The verdine facies deposits identified in 1988 by G.S. Odin, J.P. Debenay and J.P. Masse 1. Introduction 131 2. The Ogooue River mouth (Gabon) 131 3. The Orinoco River mouth and eastern extension (northern South America) 134 4. The Niger delta (Nigeria) 136 5. The Konkoure River mouth (Guinea) 138 6. The verdine off Sarawak (N. Borneo, Malaysia) 142 7. The Congo River mouth 145 8. The shelf off Ivory Coast (Comoe River and other rivers) 149 9. The verdine facies from Mayotte (Comoro Islands) 152 10 The Casamance estuary (Senegal) 153 156 11 Conclusions 6

*Chaper B5 Mineralogy of the verdine facies by G.S. Odin, S.W. Bailey, M. Amouric, F. Frohlich and G. Waychunas 1. Introduction 159 2. Mineralogical data on the grains of verdine by G.S. Odin 159 3. X-ray diffraction study of the verdine facies by G.S. Odin 162 162 3.1. The green clay with a dominant peak near 7.2 A: phyllite V 169 3.2.The green clay with a dominant peak near 14 A: phyllite C 4. High resolution transmission electron microscopy on verdine by M. Amouric and G.S. Odin 171 4.1. Verdine from New Caledonia 172 4.2. Verdine from Guinea 172 4.3. Verdine from Senegal (phyllite V) 174 4.4. Verdine from Senegal (phyllite C) 176 4.5. Conclusions on the TEM studies 177 5 . Chemical study of the minerals from the verdine facies by G.S. Odin 178 5.1. Chemical composition of phyllite V 178 5.2. Chemical composition of phyllite C 182 6. Physico-chemical study of the verdine minerals by G.S. Odin, S.W. Bailey, F. Frohlich and G.A. Waychunas 183 6.1. Thermal study of phyllite V 183 6.2. Infra-red study of phyllite V 185 6.3. Mossbauer study of phyllite V. 188 7. The clay minerals from the verdine facies by S.W. Bailey and G.S.Odin 189 7.1. The diffraction peak near 14 %, for phyllite V 189 7.2. Mineralogical interpretation of phyllite V 191

XI1

7.3. Mineralogical variation of phyllite V as a function of time 7.4. Mineralogical interpretation of phyllite C 7.5. Comparison between phyllite V and chlorites 8. Conclusions

198 200 20 1 202

*Chapter B6 Geological significance of the verdine facies by G.S. Odin and B. Sen Gupta 1. Broad features of the environment 2. Localsettings 3. Substrates and microenvironment 4. Recent verdine facies versus ancient ironstone facies 5. Synopsis

205 205 205 212 215 217

CONTENTS OF PART C *Introduction to the glaucony facies *Contents of part C

22 1 223

*Chapter C1 Glaucony from the Gulf of Guinea by G.S. Odin 1. Introduction 2. Geological setting 2.1. Hydrology 2.2. General data on glaucony 3. Habit and nanostructure of glaucony 3.1. Morphological data 3.2.Nanostructure of the faecal pellets from the Congolese shelf 3.3. Morphological evolution of the faecal pellets 4. Mineralogy of the green grains 4.1. X-ray diffraction study 4.2. Chemical study 4.3. Stable isotope study 4.4. Radioactive isotope study 5. Geology of the glaucony from the Gulf of Guinea 5.1. Age of the glauconitization process 5.2. The glauconitization process 5.3. Environment of glauconitization 6. Summary and conclusions

225 225 225 227 229 229 232 232 233 233 236 237 238 240 240 242 245 246

*Chapter C2 Glaucony from the margin off northwestern Spain by G.S. Odin and M. Lamboy 1. Introduction 2. Distribution of glaucony 3. Morphological features 3.1. The verdissement of echinoderm fragments 3.2. The verdissement of bored shell fragments

249 249 252 252 254

XI11

3.3. The verdissement within the foraminifera1 tests 3.4. The verdissement of detrital mica flakes 3.5. The verdissement of quartz grains 3.6. The verdissement of non-bored shell fragments 3.7. Other morphological features 4. Mineralogical study 4.1. X-ray diffraction sudy 4.2. Chemical study 4.3. Isotopic study 5. Discussion 5.1. Morphological features and glauconitization process 5.2. Mineralogical features and glauconitization process 5.3. Age of the glauconitization process 5.4. Mechanism of glauconitization 5.5. Environment for glauconitization 6. Summary

256 257 260 262 264 266 266 268 269 27 1 27 1 27 1 272 272 273 274

*Chapter C3 Glaucony from the Kerguelen Plateau by G.S. Odin and F. Frohlich 1. Pi-esentation 1.1.Glauconies from high latitude deposits 1.2. The Kerguelen Plateau 2. Sedimentology 2.1. The sedimentary cover 2.2. Glaucony 3. Nature and origin of the green grains 3.1. Detailed data on the northeastern shelf sediments 3.2. Substrates of verdissement 3.3. Electron microscopy of evolved grains 3.4. Mineralogy of the green grains 4. Discussion 4.1. History of the glaucony from the Kerguelen Plateau 4.2. Factors of the glauconitization process 5. summary

277 277 27 8 279 279 280 282 282 283 289 29 1 29 1 29 1 292 294

*Chapter C4 Geological significance of the glaucony facies by G.S. Odin and P.D. Fullagar 1. Introduction 2. Habits of glaucony 2.1. Classification of habits 2.2.The granular habits 2.3.The film habits 2.4. Discussion of the habits of glaucony 3. Mineralogy of glaucony 3.1. Substrate components 3.2. Authigenic marine clays

295 296 296 297 300 301 305 305 309

XIV

3.3. Post-genesis components of green grains 4 . Genesis of glaucony: the verdissement process 4.1. The layer lattice theory 4.2. The mechanism of verdissement 5 . Environment of glauconitization 5.1. Microenvironment 5 . 2 . General environment 6 . Summary

316 318 318 3 19 323 323 324 33 1

CONTENTS of PART D *Introduction to the celadonite-bearing facies *Contents of part D

333 334

*Nature and geological significance of celadonite by G.S. Odin, A. Desprairies, P.D. Fullagar, H. Bellon, A. Decaneau, F. Frohlich and M. Zelvelder 1 . Presentation of celadonite by G.S. Odin

337

2. Occurrence and geological setting of celadonite by G.S. Odin and A. Desprairies 2.1. Examples of celadonite-bearing outcrops 2.2. Morphological features of celadonite 2.3. Petrographic environment of celadonite

337 337 340 342

3. Mineralogical properties of celadonite 3.1. Physical properties of celadonite by G.S. Odin, A. Desprairies and M. Zelvelder 3.2. Physico-chemical properties of celadonite 3.2.1. Infra-red absorption spectra of celadonite and related minerals by F. Frohlich and M. Zelvelder 3.2.2. Mossbauer spectra of celadonite and related minerals by A. Decaneau 3.3. Chemical properties of celadonite by G.S. Odin and A. Desprairies 3.3.1. Major element analyses on large samples 3.3.2. Discussion of chemical analyses on macro-samples 3.3.3. Micro-chemical analyses for major elements 3.3.4. Trace and rare earth elements 4. Environment of formation of celadonite 4.1. Environment of formation of celadonite based on petrography by G.S. Odin and A. Desprairies

345 353 362 365 365 368 369 374

375

xv 4.2. Environment of formation of celadonite based on mineralogy by G.S. Odin 4.3. '*O isotopic study on celadonite by A. Desprairies and G.S. Odin 4.4. Time of formation of celadonite after basalt emplacement by G. S . Odin, H. Bellon, P. D. Fullagar, A. Desprairies and M. Zelvelder 4.4.1. Radiometric dating applied to celadonite 4.4.2. Formation of celadonite in young oceanic basalts

376 379 382 382 392

5. Nature and geological significance of celadonite by G.S. Odin and A. Desprairies 5.1. General characteristics of celadonite 5.2. Comparison between celadonite and glauconitic minerals

393 393 395

6. Acknowledgements

398

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XVII

PREFACE

par Georges Millot, de 1'Academic des Sciences

Nous devons toujours &re reconnaissants aux scientifiques qui prennent le temps d'Ccrire un livre. En fait, le prCsent livre n'est ni un manuel ni une fresque pour le grand public, mais un truite'; ce qui nkcessite un travail beaucoup plus considCrable. Et ce trait6 fait dflCchir. J'ai retenu quelques thkmes de rkflexion. a a .

La couleur Ce livre ttudie les argiles marines vertes. La couleur est leur apparence, mais quelle est leur nature? Voici un demi-sikcle, les gCologues ignoraient la nature des roches argileuses: ils distinguaient les argiles grises, bleues, noires, rouges, bariolCes... 11s commenckrent alors l'inventaire min6ralogique de leurs constituants: seuls ou en melanges, une vingtaine desp5ces furent inventorides, ainsi que leurs interstratifiks. I1 en fut de meme pour les glauconies en grains, connues depuis longtemps. h i s on dktermina que le constituant principal Ctait la glauconite, variCtC fine de mica ou illite verte, riche en fer ferrique; la cCladonite des altkrations de basaltes fut confondue avec la glauconite. Peu A peu, une systdmatique rigoureuse de ces minCraux verts s'est enrichie. Le present trait6 nous donne 1'Ctat actuel de cet inventaire. Et nous voyons venir au moins cinq espkces, la smectite glauconitique, la glauconite et la cCladonite, aujourdhui distingukes, la phyllite V A 7 %, et la phyllite V B 14 8, du faciks verdine. A ces pales s'ajoute la population des interstratifiks. Enfin, sont traitkes avec prdcision deux espkces diagCnCtiques et ferreuses: la berthikrine et la chamosite.

Pour un mindral, la couleur est une apparence. Sa structure rdvdle sa nature. Les couleurs peuvent "converger" comme l'oiseau et la chauve-souris convergent dans leur apparence. Au contraire, la structure est propre d une espke: elle est une signature gdndtique. Faciks et mineral Ce livre insiste sur une distinction trks importante: il sCpare categoriquement faciks et minCral. On avait montrC voici 30 ans que les "glauconitic pellets", bien reconnaissables comme faciks, pouvaient contenir toute une variCtC de minCraux. C'est pourquoi, il a CtC proposC de distinguer le faciks sous le nom de gluuconie et le mineral type sous le nom de glauconite. C'est la meme exigence qui impose de distinguer dolomie et dolomite, calcaire et calcite. Ce trait6 dkmontre la grande importance gkologique de cette distinction, mais aussi son importance pratique en

XVIII

gkochronologie. En effet, plus l'kchantillon, analysC en spectromktrie de masse, sera riche en un minkral authigknique et trks potassique, la glauconite, plus les dkterminations isotopiques seront significativeset fidkles, et non pas erratiques. La mEme distinction "catkgorique" est menke pour le facibs verdine. Et ce faciks prksente deux espkces principales singulikres: une phyllite V B 7 qui est une kaolinite ou serpentine trks ferrique, et une phyllite V B 14 A, qui est une chlorite trks ferrique. Leurs caractkres cristallographiques s'accumulent en vue d'une reconnaissance par le ComitC international de nomenclature. Faciks gkologiques et espkces miniralogiques sont, tous deux et Cgalement, importants B connaitre pour reconstituer les conditions de genkse. Le faciks indique l'environnement avec ses particules en mouvement, les conditions de dkpbt, les traces de la vie. La paragenkse minkrale dkpend de la composition des solutions qui ont cristallisk ou recristallisk.

La dife'rence entre f a c i b et mine'ral est celle de deux kchelles dbbservation. De plus, la premikre est bioge'odynamique,la seconde est gkochimique.

Les miniiraux argileux interstratifiiis Les minCraux argileux ont "inventd" un type dorganisation particulier, qui est celui des minkraux interstratifiks. Ce sont des Cdifices oh alternent dans un meme cristal des feuillets ou parties de feuillets de natures diffkrentes. Cette invention est prkcieuse. En effet, le caractkre propre du monde minCral est discontinu. Au contraire, les solutions et les magmas fondus prisentent la suite continue des melanges en toutes proportions. C'est pourquoi, quand les magmas et les solutions cristallisent, ils donnent un mClange de minCraux diffkrents: ce sont les roches et les minerais. A la suite continue des magmas et des solutions, correspond une r6ponse discontinue par le mklange de minkraux diffdrents en proportions diffirentes. I1 en est de meme pour les argiles, mais elles sont plus souples pour deux raisons. D'abord, elles sont plus tolkrantes aux substitutions stoechiomktriques. Ensuite, elles peuvent donner des cristaux mixtes, interstratifiks rCguliers ou irrkguliers, o i ~ alternent des feuillets de minQaux argileux diffkrents. Le rksent trait6 illustre cette propriktk des argiles: on peut citer l'interstratifik (7A-14 ) ou le composk nommk phyllite C (interstratifid smectite-chlorite) tous deux propres au facibs verdine et l'interstratifik (smectite-glauconite) dans le facibs glauconie. Cette capacitk de s'organiser en edifices mixtes (mixed-layers) ou interstratifiks est trks intkressante. Elle donne une grande souplesse dans la cristallisation des microcristallites argileux, en fonction de la variation infinie des solutions naturelles. De meme, dans les recristallisations successives d'un materiel andrieur, des ktapes intermkdiaires sont reconnaissables.

K

A la varie'te' continue des milieux de genbse peuvent re'pondre en me'lange ou bien les mine'raux argileux types, ou bien leurs mine'raux interstrat@e's.Ces derniers jouent le r6le des diphtongues dans une langue, ou des di2zes et

be'mols dans une me'lodie.

XIX

Microenvironnement et ntioformations Les premiers travaux sur les relations entre milieux de genkse et argiles sddimentaires n'avaient pas un grand "pouvoir skparateur". On Cvoquait les milieux marins, sursalks, lacustres... Aujourdhui, et ce trait6 en donne un be1 exemple, le progrks de toutes les microscopies (optique, Clectroniques...) coup16 avec les progrks de la systkmatique des minCraux argileux permet de distinguer des microenvironnements. On comprend que le micromilieu qui se trouve au coeur d'un grain ou dun pellet prisente des caractkres physicochimiques et biochimiques diffkrents de ceux de l'eau de mer oh circulent les crabes et les turbots. Et ceci est de grande consQuence. En effet, la plupart des modkles thermodynamiques disponibles nous donnent les conditions dCquilibre "minCraux-solutions", en solution diluCe, oh 1'activitC de l'eau est considCrCe comme Cgale h l'unitC. Or, les travaux rkcents nous montrent que dans les petits pores, une partie de l'eau est like aux parois et l'activid de l'eau baisse: la prkcipitation d u n grand nombre despkces minCrales en est facilitde et leurs domaines de stabilid changCs. Ces precipitations diminuent encore le diamktre des pores et I'activitC de l'eau baisse encore. Ainsi les milieux microporeux, par rapport aux milieux environnants oh l'eau est libre, vont fonctionner comme "puits" et favoriser les minCralisations. Ceci s'effectuera avec ou sans dissolution des parois, ce qui correspond 2 deux mCcanismes distinguCs dans ce livre: cristallisation dans les microvides ou cristallisation avec contribution des minQaux de la roche. Aux extremes, on se trouve devant la nCocristallisation ou devant la recristallisation; on comprend les cas interm6diaires.

La notion de microenvironnement,largement utiliske dans le prksent trait&,est tr2s importante: elle pennet d'envisager des microsites avec leurs micromilieux difkrents. Et ces microsites sont autant de "pi2ges"pour les cations environnants alimentant la croissance de cristaux difkrents. VarititC des mintiraux et milieux de gen&se G r k e B une systdmatique raffinCe des "Argiles marines vertes", griice B l'observation distincte des faciks et des minCraux, griice h l'examen minutieux des microenvironnements, ce trait6 prisente une discussion serrCe sur les relations entre les minCraux argileux verts et les milieux de genkse. De cette faqon, ce livre est un livre de giologie aussi bien que de minkralogie. En effet, son but ultime est de comprendre l'origine et l'histoire de ces minCraux, histoire qui s'kclaire aujourdhui de faqon decisive, parce que 1'Ctude gCologique des faciks et 1'Ctude minCralogique des constituants sont menCes toutes les deux de faCon approfondie. Et puis, aprks la sedimentation et la diagenkse prCcoce qui se produit dans les vases, vient la diagenbse d'enfouissement. Lh, les contraintes thermodynamiques augmentent, et les minCraux recristallisent. Les trois termes de ma rkgle d o r augmentent en m6me temps: ordre, puretC, taille. L'ordre est caractCrisC par la cristallinid. Parce que chaque mintral tend vers un minimum dCnergie interne, les ions Ctrangers sont chassis dans le mineral voisin, et "la puretd", comme l'ordre, augmente. Enfin, parce que l'ordre et la puretC augmentent, la taille des minCraux

xx augmente. On parvient aux minCraux de la diagenkse, puis du mCtamorphisme. L'Ctude comparke des "ironstone facies" B chamosite et berthikrine et des "verdine facies" est trks suggestive dune diagenkse denfouissement. Quund l'alpiniste quitte la vallke et s'klhe sur les versants puis sur les hauteurs de la montagne, il voit la vkgktation changer: changement des caracttres dune esptce, passage d des sow-esplces puis des esptces diflkrentes, puis trts diflkrentes. La gore change avec l'environnement. I1 en est de m&meici pour les milieux skdimentaires. 0 0

.

Depuis un demi-sikcle, j'ai CtC le tCmoin de l'aurore, puis des progrks de la "GCologie des Argiles". Plus gCnCralement, au cours de ce long voyage, j'ai CtC admiratif des renouveaux des connaissances en "GCochimie de la Surface de la Terre". C'est un honneur pour moi de prksenter ce livre sur les "Argiles marines vertes". J'en remercie les auteurs et le principal dentre eux, Monsieur G.S. Odin et ceci pour deux raisons. D'abord parce que ce livre nous Cclaire de lumikres nouvelles sur un vieux problkme: ceci est trks utile pour sortir de la confusion. Mais aussi et surtout parce qu'il nous fait rCflCchir sur les Cquilibres dClicats qui, dans les micromilieux des sCdiments marins, nous conduisent B des nCoformations minCrales par croissance des cristaux. Faciks et minCral: une nouvelle fois, il existe une relation entre la composition et le milieu de genkse des roches argleuses. Ce trait6 est un exemple significatif de la mCthode B suivre, en Sciences de la Terre: combiner les arguments gCologiques et gComCtriques et les arguments mindralogiques et gbochimiques. Ces deux familles d'arguments doivent etre combinks de facon entrelacee, come serpenti in m o r e . Strasbourg, au jour du printemps le 21 Mars 1988 Georges Millot

XXI

FOREWORD

by Georges Millot de 1' Academic des sciences

We should always be thankful to scientists who take the time to write a book. In fact, this volume is neither a manual nor a vulgarization, but a treatise; and that entails a considerably greater amount of work. And this treatise is thoughtprovoking. I have taken note of some of the themes upon which to reflect. 0 0

0

The colour This book is about green marine clays. Such is their appearance; but what is their nature? Half a century ago, geologists knew little of the nature of clayey rocks: to them they were grey, blue, black, red, smped ... A mineralogical list of their components was then begun, and twenty or so kinds were recognized, alone or combined, as well as their interlayerings. The same went for glaucony grains, which had long been known. It was then found that the principal component was glauconitic mica, a fine variety of mica or green illite, rich in f e m c iron; the celadonite produced by alteration of basalts was confused with glauconitic mica. Little by little, the rigorous inventory of these green minerals was added to. The actual state of this inventory today is presented in this treatise. We note the arrival of at least five types, glauconitic smectite, glauconitic mica and celadonite whose differences are recognized today, and the 7 A and 14 A phyllite V of the verdine facies. Between these extreme falls a sequence of mixed-layers. Finally, two diagenetic and ferrous types: berthierine and chamosite, are discussed with precision.

With minerals, the colour is superficial. I t is their structure which reveals their nature. Colours may "converge" in the manner of the shape of a bird and a bat. In contrast, the structure remains proper to each species: it is a genetic signature. Facies and mineral This book dwells on a very important distinction: it categorically separates the mineral from the facies. Thirty years ago, it was shown that glauconitic pellets, while recognizable as a facies, could contain a wide variety of minerals. It was for this reason that the suggestion was made to distinguish the facies with the name glaucony (in French, glauconie) from the mineral, called glauconitic mica (in French, glauconi&& The same difficulty obliges us to distinguish the magnesian limestone (in French, dolomie) from the mineral dolomite (in French, dolomitg), and limestone from calcite. This treatise demonstrates the great geological impor-

XXI I

tance of the terminological distinction, as well as its practical contribution in geochronology. Indeed, more a sample destined for mass spectrometry analysis is found to be rich in the highly potassic authigenic mineral, glauconitic mica, more the isotopic determinations will be reliable and less erratic. The same "categorical" distinction is made for the verdine fucies. Now this facies presents two singular princi al types: a 7 8, phyllite V which is a highly femc serpentine or kaolinite, and a 1 4 phyllite V which is a highly ferric chlorite. Their increasingly well-understood crystallographic characters make them likely to be recognized by the ad-hoc International Committee on mineralogical nomenclature. The nature of both geological facies and mineralogical types are each equally important in determining the conditions in which the rock was formed. The facies indicates the environment with its free-moving particles, depositional conditions, and traces of life. The mineral paragenesis depends on the composition of the crystallizing or recrystallizing solutions.

f

The difSerence between facies and mineral is that of two scales of observation. Furthermore, the first is biogeodynamic, the second is geochemical. The mixed-layered clay minerals Clay minerals have "invented" a particular type of organization, that of mixed-layered minerals. These are constructions where sheets in whole or part of different natures, alternate within the same crystal. This invention is invaluable. In fact, the very character of the mineral world is discontinuous. In contrast, solutions and molten magmas present a continuous suite of mixtures in varying proportions. And so, when these magmas and solutions crystallize, they give a mixture of different minerals: these are the rocks and the ores. To the continuous suite of magmas and solutions corresponds a response made discontinuous by the mixture of minerals in different quantities. The same goes for the clays, but they are more flexible for two reasons. First, they are more tolerant towards stoechiometrical substitutions. Secondly, they may give mixed crystals, with regular or irregular mixed-layering, where different clay mineral layers alternate. An illustration of this property in clays is given by the present treatise: we may cite the 7 A-14 8, mixed-layer, or the compound called phyllite C (smectite-chlorite mixed-layer) both of which are found in the verdine f a c i e s , and the smectite-glauconite mixed-layers of the gluucony fucies. This capacity for mixed-layer organization is most interesting. It lends a great flexibility to the crystallization of the clay microcrystallites, as a function of the infinite variation of natural solutions. In this way, the intermediate stages in the successive recrystallizations of earlier material may be shown to exist.

From the continuum of varying milieux may come a mixture either of clay mineral types, or of mixed-layering of these minerals. These last serve in the same way as diphtongs in language, or sharps andjlats in a melody.

XXIII

Microenvironment and neoformations The early work done on the relationship between sedimentary milieux and sedimentary clays was of rather low "resolution". Terms were invoked such as marine, brackish, lacustrine ... Nowadays - and this treatise is a good example - progress in all the microscopies (eg. optical, electronic ...) associated with the extended inventory of clay minerals, allows different microenvironments to be distinguished. We understand that the micromilieu found at the centre of a grain or pellet presents physicochemical and biochemical characteristics different from those of sea-water where crabs and turbots swim. Now this is of great consequence. For most available thermodynamic models provide for mineral-solution conditions in equilibrium, in diluted solution, where the activity of the water is taken as unity. Now, recent work has shown that in small pores, part of the water is attached to the wall, and the waterk activity is diminished; the precipitation of a great number of mineral species is thereby made easier and their stability domains are changed. Such precipitations further diminish the diameter of the pores and the activity of the water diminishes even more. Thus microporous milieux, as against the surrounding regions of free water, will act as "sinks" favourable to mineralizations. This will occur with or without dissolution of the walls, which corresponds to two mechanisms distinguished in this book: crystallization in microcavities or crystallization wherein the minerals of the host rock contribute. In extreme cases, we thus observe neocrystallization or recrystallization. We understand the cases that lie between.

The notion of the microenvironment, widely used in the present treatise, is very important: it allows the conception of microsites each with their different micromilieu. And these microsites will be so many traps for neighbouring cations which feed the growth of different crystals. Variety of minerals and formation milieux Thanks to a greatly perfected inventory of the "Green Marine Clays", thanks to the separate observation of mineral and facies, thanks to the examination in minute detail of microenvironments, this treatise presents a close discussion of the relations between the green clay minerals and the formation milieux. The book thus is as much about geology as it is about mineralogy. Its ultimate aims, indeed, is to understand the origin and history of these minerals, a history which today becomes decisively clearer, because the geological study of the facies and the mineralogical study of the components are both undertaken in depth. Then, after the sedimentation and the early diagenesis produced in the ooze, comes the burial diagenesis. Here, the thermodynamic constraints increase, and the minerals recrystallize. The three terms of my golden rule: order, purity, size increase at the same time. The order is characterized by the crystallinity. Since each crystal tends towards a minimum internal energy, external ions are chased away to neighbouring minerals, and the "purity" increases with the order. Finally,

because order and purity increase, the size of the minerals increases. We reach the diagenetic minerals, then those of metamorphism. The comparative study between the ironstone facies (which include chamosite and berthierine) and the verdine facies, is very suggestive of a burial diagenesis.

When the alpinist leaves the valley to climb the mountain-sides and then the mountain-tops, he sees the vegetation change: change in the characters of a species, passing to sub-species, to diferent species and then to very digerent species. The flora changes with the environment. It is the same here with sedimentary milieux.

* * Over the last half century, I have witnessed the beginnings and then progress of "Geology of Clays". In more general terms, I have, during this long journey, admired the increase in knowledge of the "Geochemistry of the Earth's surface". It is for me an honour to present this book on the "Green Marine Clays". For it I thank the authors, and the major contributor, Dr. G.S. Odin and I thank them for two reasons. First, because this book sheds new light for us on an old problem: this is very useful for getting out of the confusion. But also and above all because it causes us to reflect upon the delicate balances which, in the micromilieux of marine sediments, brings us to mineral neoformation by crystal growth. Facies and minerals: once again there exists a relation between the composition and the formation milieux of clays. This treatise is a significant example of working methods in the Earth Sciences: to combine geological and geometric arguments with mineralogical and geochemical arguments. These two families of arguments must be combined as though they were interwoven, come serpenti in amore. Strasbourg, the first day of Spring March 2lSt,1988 Georges Millot

1

INTRODUCTION TO THE STUDY OF GREEN MARINE CLAYS by G.S. Odin About twenty years ago, the studies of the present editor were mostly devoted to glauconitic sediments. To obtain a better understanding of the conditions and process of genesis of the facies before diagenetical alteration (glauconitization),particular attention was paid to the unburied sediments lying on the continental shelves of present oceans (Odin, 1975a). In the course of that study, it became clear that some green grains were formed of authigenic clay minerals, with a dominant X-ray diffraction peak at 7 8, (Odin, 1985a). This sort of 7 8, clay mineral is known from preQuaternary series e.g., oolitic ironstones (Orcel et al., 1949). However, substancially different environmental or mineralogical characteristics have been observed for the following three facies namely 1) glaucony, the most frequent ancient and recent marine facies characterized by the formation of green clay minerals at comparatively high depth; 2) verdine, the facies encountered in recent sediments at shallow depth; and 3 ) oolitic ironstone, the facies described from shallow depth pre-Quaternary sediments. The three facies correspond to three specific combinations of different green clay minerals with particular habits, and therefore, must be identified using specific names. Incidentally, amongst the diverse aspects of the study of the glaucony facies, one was the use of crushed green grains for paint pigments during Roman times and later (Odin and Delamare, 1986). However, detailed studies of the green pigment of wall paintings showed that the main natural pigment used was a celadonitic clay mineral (a green clay found in volcanic rocks). The search for the possible geographical source of the green pigment led to a study of the celadonite-bearing facies, which showed sufficient points in common with the glaucony facies to allow a fruitful comparison from an environmental point of view. The four facies quoted above have sometimes been confused in the past literature i.e., ironstone green clays for verdine green clays, verdine facies for glaucony facies, and some glauconitic minerals for the celadonitic mica. The common point is that green clays are involved, and all of them are iron-rich. Partly for that reason, the order of presentation chosen in this volume follows the geochemical path of iron from continent to ocean (Odin, 1975b): ancient ironstone facies is known to form at the boundary between land and ocean; recent verdine facies appears to develop at a similar location but is essentially marine; glaucony facies occurs farther from continental water input in an open marine environment; in the celadonite-bearing rocks, iron appears dominantly "oceanic" in the sense that many of these rocks are found in volcanic series formed in deep waters where oceans are opening. The general organization of each of the four parts of the volume (detailed

2

content at the beginning of each part) has been varied to avoid monotony in the form and to take into account the present state of knowledge as discussed below. The oolitic ironstone facies is mostly discussed here for the purpose of comparison and as it is widely described in the literature, this part is short comprising two chapters. The first chapter (Al) concerns a selected example of outcrop which is interesting both because it shows many of the main characters of the facies and because it is little known to English speaking readers (most literature on these Swiss sediments from Chamoson area is in other languages). The second chapter (A2) discusses the diverse hypotheses on the origin and process of formation. A new interpretation of the green marine clays of the oolitic ironstone facies is proposed in the light of knowledge of recent verdine facies: the ironstone clays would be essentially diagenetic in origin. The verdine facies is presented at length because previously little known; many new data are gathered. This part gives what is intended to be an exhaustive view. Four chapters describe the presently known deposits for this facies, three deposits are considered in details from the most recent deposit (Chapter B 1, New Caledonia) to the presumably oldest and relict (Chapter B3, French Guiana) deposit. Others are shortly described in Chapter B4.A detailed mineralogic study constitutes a fifth chapter where a variety of green marine clays are precisely identified for the first time. A sixth chapter summarizes the geological significance of that facies in comparison to the formerly discussed one (oolitic ironstone facies) as well as to the following one (glaucony facies). The glaucony facies is discussed in part C; new data have been obtained from the present sea-bottom. Three chapters give examples of unburied glauconitic sedimentary deposits from different latitudes and different ages and illustrate the variety of the initial substrates favourable to glauconitization. A fourth chapter summarizes the general characteristics of the facies; this chapter is not intended to be exhaustive since many data have been published in the literature especially from the mineralogical point of view; but the geological significance is considered in detail and new data obtained since ten years from isotopic analyses are particularly fruitful for this discussion. The fourth part on celadonite-bearing rocks, is arranged as a multiauthored single monograph which reviews the data recently gathered on a subject which is extensively discussed in the literature of the last decade because celadonite occurs frequently in deep-sea basalts. New unpublished results are added in this review; for example, results of isotopic studies which allow both a better knowledge of the environment of formation to be obtained and the geological significance of celadonite to be understood. Following a general conclusion on green marine clays, a glossary is proposed; it might be useful to consult this glossary before reading the main text because a certain number of new or uncommonly used words have been used in order to designate some of the concepts frequently referred to in the volume. This glossary may serve as a subject index. In short, the present volume concerns clay formation in more or less close contact with sea-water i.e., synsedimentary processes; the central part of this

3 volume is the discussion of two facies, verdine and glaucony, observable in the present oceans; each of these two facies has been compared to a brother facies, mineralogically different and environmentally apparently similar for ironstone facies versus verdine facies; or environmentally apparently different and mineralogically close for celadonite-bearing facies versus glaucony facies. The study of these four facies in a single volume was a good opportunity to derive general ideas on the relationship between the marine environment and clay minerals. Two leading ideas were constantly present when writing analytical discussions regarding examples of sedimentary deposits as well as chapters of general discussion. The first leading idea is that authigenic marine clays are widespread and diverse in oceans. Although quantitatively more restricted than inherited clays, the authigenic clay minerals are a more precise reflection of their environment of sedimentation and deserve more study. Moreover, the geological significance of authigenic clay minerals can be precisely discussed thanks to the study of sediments of the present-day sea-bottom. The variety of authigenic clay and other minerals formed at the sea-bottom still needs to be investigated, but appears wider than commonly known. In particular, all clay mineral groups are present. The present volume only gives a limited view on the question with authigenic serpentinic (7 A), smectitic to illitic (glauconitic minerals), intermediate smectitic-chloritic, and purely chloritic clay minerals; this first set of data should be supplemented in the future. The second leading idea results from the observation that the diverse clay minerals observed are all formed following a similar fundamental process which is crystal growth. Therefore, the fashionable hypothesis developed during the sixties and seventies, and according to which the sea was essentially the site for a transformation process (moderate modification in a permanent inherited crystal structure) of inherited clays, is considered obsolete here. The marine environment entirely creates its own suite of clay minerals. It is one of the most interesting aspects of the study of these green marine clays, to have been, in some aspects, the first instance in which this revision of a fundamental idea of the sedimentology of clay was made necessary.

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5

Part A

THE OOLITIC IRONSTONE FACIES INTRODUCTION TO THE OOLITIC IRONSTONE FACIES The oolitic ironstone facies has been described from sedimentary formations of Cambrian to Pliocene age (James, 1966). It can be regarded as absent from modern sea-bottom sediments, with one possible exception reported from Scotland (Rohrlich et al., 1969). The present part intends to summarize what is known (or suspected) about the environment of formation of the ancient oolitic ironstones and the general mineralogy of the authigenic clay minerals present in this facies. These clay minerals are green, iron-rich, and occur in a shallow marine facies; therefore, the data summarized will allow a comparison with the modern, apparently equivalent facies, i.e. the verdine facies described in the following Part B of this volume. Part A comprises two chapters. The first describes a single case study initially considered by M.F. Delaloye in his thesis. This example has been selected because the concerned outcrop is 1) historically important, 2) comparatively little quoted in the English literature, 3) well preserved, and 4)really representative of the facies. This outcrop is located near Chamoson and was the source from where chamosite was defined by Berthier (1820). Based on the information provided by this and other related outcrops in Switzerland, the second chapter discusses the diverse hypotheses proposed in the literature about the process and environment of formation of the oolitic ironstone facies. Central to these hypotheses is the question of the puzzling original oxido-reduction factor. A new hypothesis concerning the original nature of the green marine clays will be proposed in order to suggest a new approach at the Eh problem. This hypothesis proposes that the iron cations in the original synsedimentary green marine clays were initially ferric in contrast to their present ferrous state. This implies that all oolitic ironstone clay minerals analysed today result from diagenetic reactions. This fundamental, previously non-considered evolution of clay minerals formed during sediment deposition, would have taken place during early burial diagenesis; the hypothesis is supported by several observations including the comparatively good crystallinity observed today for these green clays, which is unusual for marine authigenic clay minerals, and the common occurrence of authigenic ferric marine clay minerals in recent unburied sediments.

6

CONTENTS of PART A *Introduction to the oolitic ironstone facies *Contents of Part A

5 6

*Chapter A1 Chamosite, the green marine clay from Chamoson; a study of Swiss oolitic ironstones by M.F. Delaloye and G.S. Odin 1. Introduction 7 9 2. Petrography of the Swiss oolitic ironstone 9 2.1 .Elements of the sediment 14 2.2.The groundmass of the oolitic sediment 15 3. Mineralogical data on the green clay 3.1 .X-ray diffraction study 15 3.2.Chemical study 19 3.3.Themal study 22 24 4. Discussion and conclusions on the genesis of the facies 24 4.1 .Environment of deposition of the oolitic ironstone 25 4.2.Mineralogy of the green clay 26 4.3.Discussion on the genesis of the green clay from Chamoson 5 . Summary 28 *Chapter A2 Green marine clays from the oolitic ironstone facies: habit, mineralogy, environment. by G.S. Odin, R.W. O'B. Knox, R.A. Gygi, and S. Guerrak 1. Introduction 29 30 2. The problem of the ooids 2.1.Interpretations of the ooids 30 31 2.2.The extra-sedimentary accretion hypothesis 2.3.The intra-sedimentary concretion hypothesis 32 2.4.The ooid replacement hypothesis 34 2.5.The non-oolitic hypothesis 38 39 3. The green clay in the groundmass 40 4. Mineralogy of the ironstone green clays 40 4.1.Green clays of the ooids 4.2.Presumed initial green clays 42 44 5. Discussion of the environmental factors 44 5.1.General environment of oolitic ironstone formations 47 5.2.Microenvironment of green clay genesis in ooids 6. Summary 51

7

Chapter A1 CHAMOSITE, THE GREEN MARINE CLAY FROM CHAMOSON; A STUDY OF SWISS OOLITIC IRONSTONES by M.F. Delaloye and G.S. Odin INTRODUCTION

Previous studies of Swiss oolitic ironstones The sedimentary oolitic ironstone near Chamoson (Western Swiss Alps) allows the presentation of the historical, sedimentological and mineralogical significance of the term "chamosite". Berthier (1820) proposed the name chamoisite after the name of the village erroneously written "Chamoison" in his study. The correct spelling of "Chamoson" for the locality led to the term of charnosite which now designates the mineralogical component first identified by Berthier (Lacroix, 1895). Berthier undertook an optical and chemical study. He correctly identified the green component as a new (nearly pure) mineral: a dominantly ferrous and slightly aluminous silicate which he ascribed to the chlorite family. DCverin (1945) published a detailed petrographic study on the oolitic ironstones of Switzerland. In the sixties, this problem was resubmitted to one of us for his Ph. D. thesis (Delaloye, 1966).

Geological setting The geological setting of the iron ore of Chamoson, and of equivalent outcrops in Switzerland, is shown in Figure 1. The sediments have been locally submitted to a light metamorphism which provoked an induration of the formations. Therefore, the sediments appear generally more cemented than the iron ores of northern France or Great Britain which are also of Jurassic age. The historical outcrop is located 6 km NE of Chamoson; DCverin (1945) quotes this outcrop after the local name of Chamosentze nearby; the sediments are regarded as Callovian in age, like the Erzegg and Planplatte wid-bearing iron-rich sediments and those in the northernmost outcrop of the Glarnisch Mount. In Urbachtal, the age is upper Bathonian; and in Windgale, the oolitic ironstone formation seems to cross the Bathonian-Callovian boundary. The lithological sequence is mainly schistose and limy with ammonites as common stratigraphic fossils. Near Chamoson, the oolitic facies has the form of a restricted lens, 250 m wide, and a few metres thick. Elsewhere, the facies appears as irregular, 1 cm to 10 cm thick layers, or as series of oolitic lenses gradually passing into schists or limestones. Often, the limy sediments underlying the iron-rich oolitic lenses are composed of bioclastic debris of crinoids

8 (calcaire 3 Entroques) typical of a shelf environment. At the top of the iron-rich sediments, DCverin has reported localized concentrations of belemnite and ammonite remnants. Due to its influence on the mineral assemblage, it is of interest to note here that the area near Chamoson has been more affected by tectonics and related low-grade metamorphism than the well preserved area near Erzegg and Planplatte, but less affected than the area of Windgalle, and similarly affected compared to the area of Urbachtal.

Figure 1. Geological setting and location of the Jurassic oolitic ironstones in Switzerland. 1) Chamoson; 2) Urbachtal; 3) Planplatte; 4) Erzegg; 5) Windgillle; 6) Glmish Mount. The outcrops are located in the Helvetic Nappes and were more or less slightly metamorphosed during Alpine events.

Mineralogical nomenclature Since the original proposal by Berthier (1820), the term "chamosite" has been assigned to the silicate of Chamoson. Because it has been strictly defined only on optical and chemical grounds, the name has been used for other more or less green clays with an oolitic facies. This is a source of confusion because it is now known (Orcel et al., 1949) that crystallographically different minerals have been designated by this name. In this sense "chamosite" took the meaning of a facies: any green ferrous clay in an oolitic ironstone formation. The situation was considered inadequate by several authors including Orcel et al. (1949). Millot (1964, p. 246) suggested to define chamosite sensu strict0 as a 14 8, mineral, i.e. a chlorite. This suggestion was later approved by the Nom-

9 enclature Committee of the Clay Mineral Society (Brindley et al., 1968). Therefore, chamosite designates a ferrous chlorite like the one found in the sedimentary iron ore outcropping near Chamoson. The 7 A clay mineral that has been called "chamosite" by English writing authors (for example Yerskova et al., 1976) does not have priority for this name and the term berthierine should be prefered. The mineral name berthierine was first proposed by Beudant (1832, in Lacroix, 1895) for the green oolitic clay collected from the Jurassic ironstone of Hayange (Lorraine, France). These ooids are purely made of a 7 8, iron-rich clay mineral. Berthierine is therefore a highly femferous clay of the serpentine family. However, the correct use of these two terms is not easy due to the difficulty in distinguishing whether the mineral has a periodicity at 14 A or not. In the natural feniferous chlorites, the first diffraction peaks of odd order in the 001 series, are indeed frequently of very low intensity (Orcel et al., 1949). Specific treatments may however reveal the 14 8, periodicity poorly expressed on untreated samples (Tomita and Takahashi, 1986). If a detailed mineralogical study is not available, we will quote ironstone green clay for the clay minerals of the iron-rich oolitic sediments. PETROGRAPHY OF THE SWISS OOLITIC IRONSTONES

The main characteristic of the studied sediments is their high iron content; these formations are therefore easy to observe in the field where they appear as brownish lenses due to meteoric oxidation. In thin section as well as in the field, the most typical criterion is the presence of ooids locally reaching 1 mm or more in diameter. These ooids are irregularly distributed and usually draw clouds or irregular bands of more or less dense spheroidal or ellipsoidal elements in a fine-grained, cemented groundmass. This distribution as well as the general sedimentary structure (lens) clearly results from deposition in presence of bottom currents.

Elements of the sediment

Carbonate remnants Excluding the cement and the ooids, the only important constituents of the iron ore lens near Chamoson are bioclastic carbonate grains, in addition to very little detrital quartz. Ninety percent of the biodetrital fraction is made of crino'id fragments. Most are strongly eroded and can be observed in all thin sections either as free grains or as nuclei in the ooids. Spines or other skeletal parts of echinids are not rare; like the crino'id fragments, they can be identified by their characteristic structure delineating a regular lattice: the stereom. The rest of the carbonate fraction consists of pieces of molluscs (bivalve shells or belemnite rostra), sponge spines, tests of brachiopods or ostracodes. According to Dkverin (1945), the small (pelagic) foraminifers were probably abundant in the initial sediment; their present abundance, when preserved by silicification in pyrite or in a few oolitic nuclei, supports this view. Planktonic fossils are abun-

10 dant in the ironstone of the G l b i s c h Mount where DCverin (1945) observed the presence of foraminifers, ostracodes, microcephalopodes and microgastropodes. These fossils indicate an open marine environment of deposition. In thin sections, the relationship between the carbonate grains and the green clays is of three types in the outcrop near Chamoson: the clays 1) fill the voids, 2) impregnate and replace, or 3) surround and coat the previously deposited clasts of carbonate. The latter case leads to the oolitic structure. DCverin (1945) illustrated green clay in the pores of the echinoderm stereom as well as fillings in ostracode tests. The partial or complete replacement of carbonate by green clay minerals includes echinoderm and mollusc fragments, belemnite rostra and brachiopod tests in the outcrop near Chamoson. It is very frequent in Erzegg or Planplatte regions. From the relationships observed in thin section, it is difficult to conclude whether the replacement process took place before burial or later on during diagenesis (Bradshaw et al., 1980).

Well preserved ooids Oolitic structures are illustrated by DCverin (1945). A discussion of the formation of the green material is based on those oolitic structures. The problem in interpreting the thin sections is the difficulty to locate the time at which the present aspect of the rock has been reached because the presently observed paragenesis is a combination of cumulated geochemical reactions which occurred before or at the time of deposition, during early burial, diagenesis and dynamic metamorphism. Noneless, the sediments show distinct and unambiguous oolitic structures and the material forming their layers is a light-green clay resembling chlorite (Fig. 2). The size of the ooids is frequently homogeneous in a given thin section. The mean diameter varies from 0.2 to 1.0 mm. Usually of regular, ellipsoidal form when correctly preserved, the ooids show a restricted variation of the maximum versus minimum diameter ratio between 1.5 and 1.3. These characters strongly suggest a dynamic deposition and reject a strictly in situ intra-sedimentary genesis of the ooids. The nuclei of many ooids are comparatively small, and numerous green

----3-+----3

on the form and size of the initial nucleus. Photomicrograph 2 shows a large triangular calcitic nucleus (shell fi-agment) surrounded first by black and green layers followed by purely green layers. The nucleus of the darker mid is made of an angular layered fragment of broken ooid (broken ooids are present on Photo.1). The light-green layers are underlined by black sheets representing interlaminae of diagenetic iron oxide additions to the initial grain. Photomicrographs 3 and 4 ( crossed and parallel nicols, respectively) show an mid of which the nucleus is an already rounded piece of echinoderm. The carbonate stereom voids are filled with green clay and the stereom itself has begun to be replaced by the green clay. Only a few of the calcitic nuclei have partly turned to green before the coating process. Photomicrographs 5 and 6 (parallel and crossed nicols respectively) show another green mid of which the nucleus is a fragment of the layered coating of a previously formed and broken mid. Photo. 6 and 3 show that the layered coating is formed of concentrically oriented clay particles because the optical figure of a black-cross appears under crossed nicols.

11

Figure 2. Thin sections of pieces of iron ore from Chamoson. The large majority of the elements are coated nuclei, i.e. ooids. Photomicrograph 1 shows abundant ooids cemented by calcite. The cement is locally green. The general form of the ooids is ellipsoidal but appears much less regular than it is usual for carbonate ooids. The form of the ooids depends... (Figure caption follows p. 10)

12 clay layers surround it. A moderate proportion of grains shows a large nucleus and a thin envelope which presents only 3 or 4 clay layers. The maximum number of more or less distinct laminae was counted for an ooid illustrated by DCverin; the ninety laminae have a constant thickness of about 4 to 5 pm for that ooid of about 0.8 mm in diameter. The colour of the clay varies from a very light bluish-green to a bright-green. The pleochroism is more or less obvious between the above quoted colours and a more yellow one. According to us, the evolution from a piece of crinoid toward an mid comprises three steps: 1) the pores of the stereom fragment are filled with green clay; 2) the clay replaces the carbonate skeleton which sheltered the first crystallization of green clay; 3) this element acts as a nucleus around which the authigenic clay concentrically crystallizes forming an mid growing with time (Fig. 3). This interpretation differs from that suggested by DCverin (1945,p. 34,45 and 47) who considers that the fragment of substituted carbonate is progressively and structurally modified (digested) from the exterior to the interior. The total volume of the grain being time-constant, the thickness of the oolitic aspect increases by centripetal rearrangement and recrystallization of the already present green clay. This was called by Dtverin "metamorphose". In other words, according to DCverin, the oolitic structure would develop itself at the expanse of the preliminary substrate, previously substituted for green clay (Fig. 3); we believe that the oolitic structure results from the accretion of green clay layers around a nucleus and this phenomenon increases the size of the initial nucleus (Fig. 3). We reject the hypothesis of DCverin mostly because the substituted fragments of echinodermal stereom have no reason to have a regular ellipsoidal shape before the "metamorphose". However, the mechanism proposed by DCverin would be possible if the initial carbonate substrate had already an oolitic structure (Fig. 3); in other words, if the green clay would form by substitution of carbonate ooids. This is the heart of the problem and this third hypothesis is difficult to reject definitely. However, we feel that it is difficult to accept a hypothesis implying primary carbonate ooids in a sediment where no carbonate ooids have never been quoted nor any trace of carbonate remnants is present in the now observed oolitic structure. A large majority of the nuclei is bioclastic carbonate fragments but other possibilities have been described. The most frequent is the re-use of pieces of brocken ooids (Fig. 2). This indicates that already hard ooids were present and were used again as nuclei for the growth of new ooids. In other words, favourable conditions for growing green clay ooids are also compatible with an in situ reworking of previously deposited hard ooids of similar composition. Therefore, the genesis location of some ooids is similar to that of deposition of others and consequently, the two environments are not far from each other. In summary, in the less modified ironstone sediments, the ooids are green; limy ooids have never been reported; the nucleus of the ooids is preferably made of biodetrital carbonate or broken pieces of previous argillaceous ooids. Pieces of already cemented sediment have also been observed. The green clay

13

presently observed in the ooids seems to result from three different processes: 1) filling the voids of a preliminarily deposited carbonate bioclastic grain; 2)replacing the carbonate fragment itself; and 3) oolitic accretion of clay around a nucleus. The last process appears predominant and common to most of the green grains. This point makes the initial oolitic clay mineral a result of geochemical reactions at or near the sea-bottom and, therefore, directly concerns the subject of the present volume.

nCARBONATE

G E E N CLAY

SEDIMENT

Figure 3. Three mechanisms could explain the internal structure observed in green ellipsoidal grains of Swiss oolitic ironstones. 1) the mechanism proposed here; 2) the mechanism suggested by Dtverin (1945); 3) the mechanism of intra-sedimentaryreplacement of previous carbonate ooids. In 1, the green clay successively fills, replaces and surrounds the carbonate fragments by an accretion process. In 2, the last stage is replaced by a centripetal recrystallization of the previously formed clay and this might occur either above or, preferably (according to Deverin), below the sea-water/sedimentinterface (bold line).

Evolution of the o o i h The evolution of the green clay ooids after burial may be of diagenetic or tectonic origin. According to DCverin (1945), the green oolitic material can be

14

substituted for siderite or feldspar; ooids can be silicified and pyrite or magnetite may grow on their margin. Carbonate is sometimes introduced between the argillaceous laminae. The tectonic effect is a deformation of the ooids which become more and more elongated despite their toughness. This deformation presumably favours the mineral substitutions quoted above (these descriptions are interpretative however -i.e. subjective-; in fact the sediments show different structures, and the observer tries to arrange them in a manner which presumably reflects the chronologic evolution). Thus three conclusions have been obtained from the Swiss ironstone study. The first conclusion accepts the fact that the formation of ferric oxides and hydroxides occurs after the genesis of the green clay which will be shown later to be ferrous. This is of great importance because several authors have reversally suggested that the first process of iron accummulation during ironstone genesis was the formation of iron oxide ooids (see Chapter A2, p. 36). The latter suggestion is not supported by observations made in the Swiss Alps. A second conclusion still concerns the chronological succession of the paragenesis. Here again, when carbonate is present in the ooids, its formation follows the accretion of green clay i.e., no carbonate ooids can be suspected before the formation of the green clay (hypothesis 3 in Fig. 3 is not supported). A third conclusion is related to the habit of the ooids. When the present form of the ooid is not ellipsoidal, this is not due to an initially soft nature; the ooids are assumed to be hard very early in the sedimentogenesis. But their deformation (flattening, lengthening) is linked to the numerous laminae allowing the clay layers to shift one onto the other under an important dynamic pressure.

The groundmass of the oolitic sediment Between the grains studied above, there is always a certain proportion of calcite which is locally interpreted as the fine crushing of bioclasts where and when ooids have been formed. For the ironstone near Chamoson, the presence of siderite makes the rock difficult to break and the oolitic grains difficult or impossible to separate and purify. A further interesting component of the groundmass is a green clay. This is a fourth habit of this component in the rock (besides the former three described from within the ooids). This diffuse green pigment is frequently darker green than the ooids themselves. It is very difficult to identify the time of genesis: it could be synchronous with the deposition or early burial time, or the result of a late diagenesis. We did not find strong arguments supporting the formation of the green clay together, before or shortly after the genesis of the green ooids. However, E v e r i n (1945) quotes the presence of pieces of cemented green clay reworked in the groundmass or used as a nucleus for some ooids. If this is correct, the green clay of the groundmass is penecontemporaneous with the formation of ooids. Elsewhere, the same author states that the diffuse green clay really seems to originate from an exudation out of the ooids. This would imply a diagenetical origin.

15 Finally, the outcrop near Chamoson is known to show numerous fissures filled with a mixture of calcite and green clay. In that case, the green clay is diagenetic and no longer indicative of the environment of deposition. In our opinion, a large proportion of green clay is formed during the time of deposition or during very early burial; this clay forms the angular, hard pieces of clay in the groundmass. But the later diagenetical processes frequently remobilize the green clay of the groundmass as well as the one from ooids. This material either recrystallizes in situ or fills the voids of the sediment. The resulting clay minerals are diagenetic products probably not related to sea-water fluids from the environmental point of view. MINERALOGICAL DATA ON THE GREEN CLAY

Only hard sediments are available in the Swiss outcrops. The ooids have been mechanically concentrated by gentle crushing of the whole-rock, by sieving, and by magnetic separation. After removal of most of the carbonate by this treatment, decinormal acetic or hydrochloric acids have been used by adding them drop by drop in a Becher containing the rock and constantly stimng. The ooids have been then concentrated by rolling the dried grains on an inclined plane and finally purified by hand picking. After crushing the selected material in an agath mortar, moderate acid leaching followed by several (up to ten) distilled water cleanings and centrifugation, glass slides were used to obtain oriented clay preparations.

X-ray diffraction study The green clay from the ooi& A particular attention was paid to ooids which represent an early step of genesis: a primary component. The purified ooids from the well-preserved sediments from Chamoson were selected, crushed, liberated of their calcite, and cleaned with distilled water. Oriented slides were prepared; Figure 4 shows two of many X-ray diffraction patterns. At first sight, the material appears somewhat complex, especially considering the number of peaks present between 7 A and 20 A. The X-ray diffraction peak at 12 A is due to stilpnomelane a mineral previously observed in thin section. However, the most developed peaks indicate the dominant presence of a chlorite: 14 A, 7.05 A, 4.68 A, 3.50 A. The 001 and 003 reflections are clearly smaller than the 002 and 004 ones, which shows that the chlorite is very rich in iron. The precise location on the diagrams of the 001 peaks for I equal to or above 5 is difficult. The nature of the peaks between 7 A and 20 A has been understood after several conventional treatments of the oriented slides. Ethylene glycol treatment shows that a mixed-layer chlorite-smectite is probably present. This smectite has been interpreted as a nontronite by Delaloye (1966).

16 Si

UNTREATED OOLITHS

30'

100

KaCu

Figure 4. X-ray diffraction patterns of well preserved ooids from Chamoson. The green clay appears to be a mixture of chlorite with other clay minerals. Foreign minerals as stilpnomelane (St), siderite (Si) and quartz (Q) are also present. (Redrawn from Delaloye, 1966)

Progressive heat treatments have been undertaken on several series of oriented clay specimens. An interesting observation is that the peak at 7.05 8, evolves and yields a doubZe peak: one portion remains at 7.01 A and the other shifts toward 7.68 A when the temperature goes up. Delaloye (1966) considers that the shifting peak characterizes the mineral berthierine ("ferrokaolin" of this author). Therefore, Delaloye (1966) considers that chlorite and berthierine are present in the green clay of the well preserved ooids; a mixed-layer chloritesmectite would complement the composition of that clay.

The clay from veins Near Chamoson the presence of 2 cm thick veins filled with pure clay minerals makes this sediment useful to easily prepare the material for analysis. The clay sampled there results from a secondary process of genesis. Oriented slides have been prepared and Figure 5 shows one of the diagrams obtained. It shows the characteristic 001 reflections for a nearly pure chlorite. The compara-

17

tively high 001 diffraction peak remains smaller than the 002 and 004 reflections. The 001 reflections have been carefully measured both in position and relative intensity in order to calculate the structural factors (F). Table 1 gives the results for the diagram of Figure 5 and for an other one not shown here. The position and intensity of the peaks are not reproducible from a diagram to another. Furthermore, the real "recognition" of the 005 reflection and higher orders is very tentative. CHLORITE from a fissure (sample 8 0 )

r-

(0

2

0

I

0 0 0

0

0

50.

40'

0

Ir)

N

0

0

0

0

300

0

-

0 0

20"

Figure 5. X-ray diffraction pattern of a clay from a vein collected near Chamoson. (Redrawn from Delaloye, 1966)

Table 1. Positions and relative intensities for the series of 001 reflections of two "chamosites"from veins collected near Chamoson. (Data from Delaloye, 1966) d

(A)

1

(%I

d

(A)

1 (%I

001

14.25

40

14.13

29

002

7.10

100

7.02

100

003

4.68

10

4.69

11

004

3.537

50

3.528

46

005

2.931 ( ? )

5

2.818

7

006

2.272 ( ? )

12

2.334

1

007

2.008

15

2.008

10

008

-

0

009

-

0

0

0

00.10

-

0

1.408

2

1.691 ( ? )

(?) Question marks indicate inaccurate estimates for the location of the reflections.

1

18

: 2 e T k.

R

T

$4 CHLORITE from a f i s s u r e

Figure 6. Electronic density of the green clay along the C, axis computed according to the 001 diffraction peaks. The curve obtained corresponds perfectly to the chlorite structure shown on the left hand side.

In order to verify the purity of the analysed material and its identity as a chlorite, oriented slides have been heated to different temperatures between 400" and 650°C before X-ray diffraction. Heating leads to a progressive decrease of all the 001 reflections but two of them; the 004 reflection remains stable until 550°C and disappears later; the 001 reflection remains similar or increases slightly after heating, its top becomes wider at 600°C. The chamosite of the veins appears to be pure, but the absence of a 7 A mineral, still remains to be proved at that stage.of the study. In order to test again the purity of the material, the calculation of the electronic density was undertaken. Projection of the electronic density of a set of planes separated by increments of 0.025/C0 on the C, axis of the elementary cell undertaken by Delaloye (1966) have been improved using computer facilities. It permitted a slightly different reconstruction of the electronic density projection on the C, axis. It is presented in Fig. 6 along with a normal chlorite structure. Very good matching indicates that the secondary iron silicate is a pure chlorite. In summary, the green clay from the ooids (primary clay) appears to be composite and the green clay of the fissures (secondary clay) is pure and strictly chloritic.

19

Chemical study Well preserved pure ooids were obtained from four samples and analysed using the classic wet chemical methods. Table 2 summarizes the data. If we except the presence of calcite and possibly iron carbonates (the two known possible impurities) the chemical results could well be used for characterization of a chlorite and a structural formula of ferrous chlorite would result. However, this exercise would not be correct since the X-ray diffraction study has proved the presence of some nontronite. This may be presumed from the chemical data because, compared to other pure chlorites (Foster, 1962), the silica contents appear rather high. Chemical analyses are also available for the clay from the veins sampled near Chamoson and at Windgtille (wet chemistry). They are given in Table 3. The two analyses published by Everin (1945) were widely quoted in the literature amongst the very few chemical data available for "chamosite". They were obtained after acid treatment which explains the absence of CaO. Amongst the three new analyses published by Delaloye (1966) for the locality of Chamoson, the one for sample 8a showed about 10% of CaO and the presence of about 7% of CO,. Therefore, the values measured were recalculated in order to eliminate calcite; the recalculated values alone are given in Table 3. Compared to the data available for the ooids, it is evident that the clay from the veins contains less silica and more alumina, FeO and MgO. This corresponds better to a ferrous chlorite composition. Amongst the clay from the veins themselves, the green clay from Chamoson is richer in silica and the Fe 0, content is 3 or 4 times less compared to the other green clays from Windghle. This may result from an oxidation process; however, the ferrous iron content does not decrease concurrently. Table 2. Chemical analyses of the green clay from purified ooids of the ironstone near Chamoson. The composition of recent green clays referred to phyllite V (Part B) is given for comparison. (G. Krummenacher analyst; data reproduced from Delaloye, 1966) ~

001 i t hs

Si02

93

44

94

60d

35.29

34.22

34.41

38.67 16.00

6

5 -

*lZ03 FeO

16.96 32.87

14.67

16.00

33.56

33.49

32.94

Fe203

0.00

3.97

0.00

CaO

1.12 0.79

8.22

0.63

0.07

MgO

2.98

2.71

2.39

2.45

H20'

9.05

6.00

8.89

8.69

99.06

99.38

99.78

98.82

Phyllite V 34 - 38

- 12 7

18 - 22 8.5 - 13.5

20 Table 3. Wet chemical analyses of the clay from Jurassic ironstone veins of Switzerland. (1) J. Jakob and (2) G. Krummenacher analysts; after (1) DCverin (1945) and (2) Delaloye (1966). Sample b is from Windgiillele;other samples from Chamoson.

Si02

28.86

26.20

27.50

30.07

22.43

*l203

18.49

20.12

19.14

17.26

14.27

Fe203 FeO

1.44

2.39

5.9

3.66

11.98

36.51

35.36

33.20

37.61

37.63

4.73

4.93

<4.5

1.69

3.42

9.95

10.33

10.00

9.01

9.70

0.21

-

0.0

0.0

0.16

0.04

0.44

0.22

0.12

0.19

100.02

99.88

MgO H20' CaO

(0)

-

K20 Na20 Ti02

Total

99.98

99.54

The calculation recommended by Foster (1962) applied to the pure chlorite sarnple 8a leads to the following formula:

(si3.12 A' 0.88)

1.47 Fe3+~.17Fe2+3.29 Mg 0.76) '10.7

OH73

Very similar structural formulae are obtained for the three other clay samples from Chumoson; therefore chamosite appears to be a somewhat homogeneous component (at Chamoson) of the chlorite family. It is worth noting that one of the samples from Chamoson in the collection of the MusCum (MusCum National dHistoire Naturelle, Paris) was labelled as bavalite. We sampled this dark-grey material called bavalite from Chamoson for further study (sample MusCum 53.300). The rock was essentially made of a cement without any apparent oolitic structure. X-ray diffraction patterns of this sample indicate the presence of a single clay mineral: an iron-rich chlorite with a comparatively very low 001 reflection. Admixtures of siderite, calcite and quartz were also present in restricted proportions. According to Lacroix (1895) and Kerforne (1908) the name bavalite was proposed first by Huot in 1841 for the green clay of the oolitic ironstone collected from the locality of Bas Vallon in Normandy. A later chemical analysis given by Orcel(l927) for this Devonian bavalite, seems to represent a typical chlorite for which we have calculated the following formula:

(si2.54 A' 1.46)

1.42 Fe3+0.06 Fe2+4.11 Mg 0.40) O10 OH,

21

Figure 7. Location of the Swiss chamosites in the SiO, - R2+- R3+ triangle. K indicates the position of kaolinite, and V the position of phyllite V (see Part B). The open triangle distinguishesthe sample from Windgalle from those of the ironstone at Chamoson.

This trioctahedral chlorite (5.99 cations in the octahedral sites) is characterized by a large Fe2+content;it has a comparatively large A1 for Si substitution. This chlorite is therefore a thuringite in the classification of Foster (1962; see Fig. 8, below). Due to the tetrahedral substitution, this mineral is different from the chlorite found in veins near Chamoson. However, the bavalite from Bas Vallon is comparatively similar to the mineral analysed by Dkverin from Windgalle. For the latter, we have also calculated a formula showing real similarities in Al-for-Si substitutions and ferrous character:

(si2.61 A1 139) (A10.57 Fe3+1.~5Fe2+3.66 Mg 0.59) OlO OH, In order to obtain more chemical details, our colleague F. Arbey submitted a small piece of sample Museum 53.300 to the micro-chemical analyzer of a scanning electron microscope. On about ten areas selected for their presumed absence of recognizable carbonate nanostructures, the data obtained were as follows: SiO, = 31% ; A1 0, = 18% ;FeO = 47OA ; MgO =3% ; (TiO, + P,O, + CaO + N 0 +K 0)= &ss than 1%. If these ata re ect the composition of the clay mineral from the cement of the Chamoson iron ore, then the concerned clay is a chlorite which looks closer to the clay from the veins than to the bavalite of Bas Vallon or the similar chlorite from Windgale. This proves that using specific names for different chlorites collected from different areas was not always done in the past with the necessary rigour from the mineralogical point of view. The data obtained for the chemical composition of "chamosite" in Table 3 have been used to plot these components in the Si0,- R2+ - R3+ triangle; Figure 7 plots four samples of the Chamoson ironstone slightly nearer to the

2

d

22 SiO, pole than the clay from Windgalle (open triangle). All five points are in the area where chlorites normally lie. The kaolinite (K), the chlorite from Chamoson, as well as the compositional field of our phyllite V mineral (Chapter B5 of this volume) lie on the same line projecting from the FeO+MgO pole. The analytical data already obtained indicate that the mineralogical component from Chamoson has the composition of a comparatively highly ferrous chlorite. This mineral may be located in the classification of Foster (1962) which combines the substitutions in tetrahedral layers to the substitutions in octahedral layers (Fig. 8). The green clays from Chamoson plot at the extreme corner (just outside, in fact) of maximal substitution of Fe2+ and minimal substitution of A1 for Si compared to the chamosite field. However, only two samples amongst her 154 chlorites fall in the domain suggested for chamosite. In that sense, the definition of chamosite by Foster (1962) lacked the foundation from a geological and chemical point of view; today, her classification could be improved according to our supplementary data.

RlPlDOLlTE

0 4-

I I

BRUNSVIGITE

'

I

I I

I

I

DIABANTITE

I I -------- + ---02-

1-- - - - --

I

I I

I

I -r

I 1

Thermal study Delaloye (1966) has undertaken some thermal analyses on the purest green clay from the veins of the Chamoson iron ore. After X-ray diffraction analysis, the material prepared has shown the presence of some calcite; it has been purified again with dilute acid. Figure 9 gives the cumulative weight loss curve and the differential thermal analysis curve. Four endothermal reactions have been recorded. The first small endothermal reaction between 60°C and 200°C

23 is broad and corresponds to the expulsion of water mainly adsorbed by the sample during acid leaching (it is nearly absent from a curve undertaken on the unleached sample: Delaloye, 1966). The largest endothermal reaction occurs at about 550°C and corresponds to a first departure of the OH groups. These OH groups are from the less stable brucitic sheet in the interlayer, which is richer in OH than the second octahedral sheet in the talc layer. The second large endothermal reaction at 770°C corresponds to the departure of the OH groups of this second octahedral sheet. The rapid increase of the curve after the endothermal peak possibly marks the beginning of the high-temperature exothermal reaction usually present in chlorites (Caillkre et al., 1982, p. 94). But this usual reaction is not well-shaped on the curve shown. Caillkre et al. (1982, p. 41) report DTA curves obtained for natural berthierines; these usually show a single endothermal reaction which could be confused with the 550°C reaction of a chlorite. If we consider that both chlorite and berthierine have an endothermal reaction near 550°C, and if we compare the surfaces of the first and second large endothermal reactions between them on Figure 9, their ratio is nearly similar to the calculated ratio of the OH groups in the brucitic layer versus the OH groups of the mica sheet of the chlorite alone. Therefore, the berthierine-type structure does not contribute to the endothermal reaction at 550°C. We may consider that the chamosite from the vein at Chamoson is purely made of chlorite.

N

CHLORITE from a fissure

I

200'

400"

600"

800"

Ipo 0°C

Figure 9. Thermal curves for the green clay from the veins collected near Chamoson.

24 DISCUSSION AND CONCLUSIONS ON THE GENESIS OF THE OOLITIC IRONSTONE FACIES

The study of the Swiss Jurassic ironstones, and especially the disused iron ore quarry near Chamoson, allowed us to gather information on three different points of interest: the genesis environment of the facies, the mineralogical nature of the green clays related to this facies, and the possible genetical process of these clays.

Environment of deposition of the oolitic ironstone

General environment The iron-rich lens near Chamoson is a marine sediment overlying a marine biodetrital formation and underlying a limy ammonite-rich deposit. The lens includes the same marine fossils as the surrounding formations and therefore, at the time of deposition, the locality was probably covered by one or several decametres of sea-water. The abundance of crinoids, which are passive suspension feeders (Gygi, 1981) indicate that bottom currents able to bring clay-size particles to that fauna were present. There is no rough change of marine regime indicated by the nature of the fossil record or by the sediments, and no modification of the sea-level is obvious before, during, or after the deposition of the oolitic facies. Limestones or schists are laterally present; the absence of sand-size detrital input indicates that continental erosion was mainly of chemical nature (flat coast). This environment exists in restricted to very restricted areas scattered on a distance reaching 170 km in Callovian time (Fig. 1). The low economic interest of the iron-rich layers is essentially due to this restricted development of the local thickness and the extent of the ores.

Local environment The precise depth at which the green ooids were formed is constrained by the fact that several indications clearly translate a sensible roughness of the sea-water. The oolitic structure itself indicates that the sea-water movements needed to be episodically strong enough to maintain the nucleus and growing ooids mobile during the oolitization process. The sea-water roughness is also indicated by the figures drawn in the sediments and underlined by the ooids: clouds, lenses of ooids in a clayey or limy groundmass clearly indicate a deltaic-like emplacement, and suggest powerful tide or wave currents. Finally, the movements must have sometimes been very strong in order to break already indurated groundmass or ooids found today as nuclei in other ooids. Therefore, green ooids appear to have been formed in a very stirred environment usually regarded as necessarily linked with highly oxidizing conditions in sea-water. However, the presently unsolved cmcial paradox of the genesis of green ooids is precisely here. The minerals presently met in the sedimentary ironstones indicate fundamentally reducing formation conditions: the major clay minerals of the sediment are ferrous; siderite seems to precipitate

25

during or soon after the deposition of the sediment, and pyrite is also frequently present. Due to this paradoxal situation many authors have suggested that, during a first step of sediment genesis, oxides or hydroxides were deposited, and during a second step, all material was reduced, sometimes implying a long preliminary transportation toward deeper sea-bottom where more reducing conditions occurred. The study of the Chamoson outcrop seems to indicate that the genetical environment was close to the area of deposition because mid nuclei can be pieces of indurated ooids or pieces of the already cemented groundmass of the sediment (reworked in situ). With regard to transport requirements, this two-step hypothesis implicitely considers that the oolitic structure of the clayey ooids observed today is the result of an epigenesis of carbonate or goethite ooids by a diagenetic clay. Neither the assumed preliminary presence of carbonate ooids nor that of iron oxide ooids has been confirmed by field or petrographic observations. Therefore, there is no room for a two step genesis of the green clay in the Swiss oolitic ironstones. The local environment for this original clay genesis needs to conjugate an obvious reducing factor and a stirred character of the environment at the sea-waterhediment interface, where the green clay minerals (or their precursor) should have been formed. The combination process of oolitic structure and ferrous nature of the green clay is certainly, although difficult to imagine, the key factor for the formation environment of the Swiss and other sedimentary ironstones.

Mineralogy of the green clay It has been shown that the green clay from Chamoson could be split in three sedimentologically different categories: 1) in the ooids, 2) in the groundmass and 3) in the veins. The chlorite is the main clay mineral as revealed by X-ray diffraction studies and by DTA curves. The detailed analyses of the ooids show that they contain a mixture of berthierine, of ferrous chlorite, and of nontronite. Nontronite is probably under the form of a mixed-layer in association with chlorite; the proportion of two smectitic layers for one chloritic layer has been suggested. The veins show the presence of a single mineral: a ferrous chlorite which is probably at the origin of the first chemical analysis of Berthier and to which the name chamosite was therefore appropriately given at that time. The green clay from the groundmass appears to have a similar chloritic nature; its composition is difficult to define due to its intimate mixture with siderite and quartz. The chemical study of several chlorite samples from the veins indicate that they do not reveal a similar content in silica. However, a dominant ferrous character is common to all these green clays. The green clay of the ooids is the most interesting for the aim of the present volume because it could have been similar to the one formed in contact with sea-water. This clay was probably not monomineralic. The clay mineral mixture is considered as resulting from possible modifications of the environment during the time when the laminae were added to the nucleus. Possibly, each lamina had initially its own chemical

26 and mineralogical composition presumably depending on variable local Eh, pH, and ion concentrations in the fluids surrounding the growing mid. This heterogeneity, contrasting with the homogeneity of the later diagenetic clay of the fissures, suggest that the initial composite nature of the ooids has been preserved.

Discussion on the genesis of the green clay from Chamoson Five different relations have been observed in thin sections between the green clay and the other elements of the rock. These relationships are: 1) the chlorite fills the holes of a biodetrital structure, 2) the green clay presents itself as a minutely detailed pseudomorph of a biogenic skeleton, 3) the green clay accretes itself to a nucleus, 4) the pigment impregnates or constitutes the groundmass of the rock, and 5) post-depositional fissures are filled with nearly pure green clay. The single mechanism able to give these five different habits is the de novo crystal growth of neoformed clay minerals using the available cations at the sea-water/sediment interface and later in the interstitial liquids. From a mineralogical or crystallographical point of view, we do not see variations between the chlorites in these different habits within a given sediment. The only difference concerns the degree of purity of the clay. This homogeneity raises a question because the clay mineral formed within the veins (in comparatively "deep" conditions of temperature, pressure, and cation availability) cannot be similar to the original clay formed in close contact with sea-water. Moreover, in several opportunities, Mverin (1945) indicates that the green clay from the ooids seems to be exuded toward the groundmass and, similarly, that the clay from the groundmass is exuded toward the fissures. By this process, the clay becomes purer than in the previous substrate and we obtain a nearly pure chlorite in the fissure. Therefore, we suggest that, because all the chlorites appear today more or less similar in the different portions of the sediment, most probably all the green clays analysed in the Callovian sedimentary ironstone of Chamoson are, to a certain degree, diagenetic clays. It will be difficult to know what kind of clay formed at the bottom of the Jurassic sea. Three questions are interesting to answer in this context: - Was there a clay, at the time of its origin, in the mid ? - If yes, was it a 7 A, 10 A or 14 A clay ... or a mixture ? - Was it a ferrous or a ferric clay ? To the first question, our answer is yes because of the total absence of carbonate remnants as well as of carbonate ooids; the total absence of non transformed oolitic initial substrate, would be an extraordinary phenomenon if these carbonate ooids were the exclusive source of the green ooids. Moreover, it is a common observation for all Swiss outcrops -as well as for many areas where Jurassic oolitic ironstones are present- that, where the green oolitic clay is present, the carbonate oolitic formation is absent in the vicinity and vice versa. It seems that the two facies exclude themselves. Therefore, two distinct

27

environments for these two facies exist, and we believe that the oolitic clay does not result from a diagenesis (replacement) of carbonate ooids. After studying the outcrop near Chamoson, we are unable to answer definitely and precisely the second question. Delaloye (1966) considers that the original clay from the ooids was probably a 7 A mineral, i.e. a berthierine. The overall presence of chlorite in Switzerland would be the result of the general dynamic metamorphism of the nappes where the iron-rich sediments presently lie. In Lorraine, where the iron ore age is similar but the tectonization much less important, the green clay is either a pure berthierine or a chlorite; finally, even more to the north, the tectonically very quiet Jurassic ironstones of Great Britain show mainly a 7 h y p e clay. This fact seems to be reinforced by the observation that the berthierine appears exclusively in the youngest sedimentary iron ores: all Palaeozoic oolitic ironstones show the presence of chlorite (Gale, 1980). Therefore, the chlorite would be the result of a stronger diagenesis of the original sediment compared to the presumably less modified berthierine. This hypothesis needs to be checked in favourable basins. Our scepticism is linked to the fact that the two clays considered here -ferrous chlorite and berthierine- are both resulting from a diagenetical process, and to the question of the clay type present in the original sediment. The only opinion presently supported is that the green oolitic clay was initially composite. Obviously, the answer to the third question needs to solve the same problem as the one discussed above, i.e. the homology or not between the initial component of the sediment and the clay presently analysed in the rock. In this chapter, our preliminary opinion is that the systematically ferrous nature of the clay present today in the Swiss sedimentary ironstones possibly reflects the initial environment of deposition of the sediment. Finally, the iron-rich formation in Switzerland does not differ from the equivalent formations of different age in different areas of the world; therefore, the general ferrous character of the green clays presumably reflects a fundamental character of the sediment surface where these ooids were deposited. A clay which is believed to be formed by oolitic accretion, and not modified diagenetically, integrates the physico-chemical conditions gathered in the surrounding fluids. If this clay is very rich in iron, the immediate environment also has to be iron-rich. If this iron is combined in its reduced form, the interface with the sediment had a low Eh. In summary, the present ferrous character of the green clay of the oolitic ironstones seems to result from the ferrous character of the clay initially deposited near the sea-bottom. This initial clay, in turn, was ferrous because the liquid layer where the ooids lay had a reducing character. Obviously, the same character makes the formation of a green clay directly at or just below the sea-bottom very easy. This green clay will form the future groundmass. As usually at continental stable margins, iron originates from the fluvial input and there is no doubt that the localized oolitic ironstone lenses are linked to localized fluvial influxes to the sea.

28 SUMMARY

Although not studied as extensively as the outcrops in Lorraine, Great Britain, or North Africa, the Swiss oolitic ironstones, and particularly the iron ore lens near Chamoson, constitute a representative case study. The green clay found there was at the origin of the general term charnosite given to most ironstone green clays. The specific mineral name is restricted to a chlorite-type characterized by iron as the major octahedral cation. This iron is mainly in its reduced state. The results of the mineralogical study confirm that the chamosite of Chamoson is a pure chlorite when the clay is extracted from the fissures or the groundmass of the rock. The green clay from the ooids is composite; the presence of chamosite (14 A), berthierine (7 A) and mixed-layer chloritesmectite can be suspected. The green clay occurs under three main habits which have different sedimentological significance: the ooids represent the time of deposition; the green clay of the groundmass may represent a period close to the deposition near the sea-water/sediment interface; the clay of the fissures is diagenetic in origin and might be posterior to the tectonic events. The similar nature of the chlorite from the three parageneses, as well as the thin section observations indicating exsolution processes from the ooids to the groundmass and from the groundmass to the fissures, suggest that the presently analysed chamosites, to a certain degree, all resulted from a diagenetic process. However, the mineralogical modification from the initial green clays to the present ones is probably not very important since the composite nature of the clay from the ooids has been preserved. The facies studied in Switzerland is typically marine with a restricted detrital contribution; the proximity of an iron-rich fluvial input appears necessary. Initially, the clay formed the ooids layer by layer. These became rapidly hard in situ. Although marks of bottom current, stirred environment, and redeposition processes are present, there is no apparent difference between the location of genesis of the ooids and the location of final deposition in the Swiss outcrops. There is a paradoxical situation between the reducing character shown by the authigenic marine minerals found in these ironstones and the apparent shallow and stirred characters of the same sediment. The genetical palaeoenvironment of the green clays of this outcrop will only be clearly understood when an acceptable compromise will be proposed taking into account these two factors. ACKNOWLEDGEMENTS

The analyses and discussion by F. Arbey and the comments by S. Guerrak made on a previous draft of this chapter are appreciated. We wish to thank Alex Waibel and Prof. R. Ledoux for critically reading the manuscript and improving the English version. Photo processing by 0. Fay is appreciated.

29

Chapter A2 GREEN MARINE CLAYS FROM THE OOLITIC IRONSTONE FACIES: HABIT, MINERALOGY, ENVIRONMENT by G.S. Odin, R.W.O'B. b o x , R.A. Gygi and S. Guerrak INTRODUCTION

The case study of the Alpine Swiss ironstones discussed in Chapter A1 allowed us to present the morphological, mineralogical, and environmental factors and characters of a comparatively young (Mesozoic) deposit. These young deposits are usually petrographically less modified than those of the Palaeozoic and are, therefore, more favourable to the reconstruction of the conditions of formation of the green marine clays. Based on this case study, the present chapter will propose a summarized review of the data and interpretations relating to the study of other deposits. The matter is complex but we are mainly interested in the factors that can be compared with those observed in the modern verdine facies considered in Part B of this volume. Central to the understanding of the green clays of the ironstones is a knowledge of their habit. Six different habits have been identified for the green clays: 1) void filling in bioclasts 2) replacement of carbonate bioclasts 3) oolitic envelopes 4)groundmass of the sediment 5) nucleus of oolitic grains 6) filling of tectonic fissures The first three habits represent synsedimentary processes; the fourth is also presumed by many authors (Taylor, 1949, pp. 7-81) to be synsedimentary; the fifth represents inherited particles of the first four habits; the last habit is unrelated to the time of sedimentation. The question to be discussed is the precise mechanism and location of green clay genesis by a crystal growth process. Four phenomena have been considered by previous authors. 1) The extra-sedimentary accretion process is a free growth around a nucleus located above the sediment surface (Hallimond, 1925, p. 92; Bichelonne and Angot, 1939). The increase in size may result from external addition either by adhesion of previously deposited material or by direct precipitation at the surface of the growing particles. 2) The intra-sedimentary focused cementation process is the formation of a hard mass by precipitation from solution about a centre (Schellmann, 1969). The resulting intra-sedimentary "concretion" may be less than 1 mm in size or larger and finally constitutes a specific particle with sharp edges.

30

3) The replacement process i.e., a substitution by green clays of an already deposited grain (Sorby, 1856; Cayeux, 1909; Kimberley, 1979). 4) The diffuse precipitation process, i.e. the formation of secondary minerals by "secretion" from solution. This occurs in the intergranular space of the soft sediment or within cavities of porous indurated substrates; it does not necessarily imply a hardening of the sediment. This precipitation is also possible at the su~aceof the sediment. Table 1. Possible processes of genesis of the various habits of the green clays of ironstones.

1. lnfilling of

Secretion

Replacement

+ (early)

0

Accretion Concretion 0

0

0

0

0

3. Oolitic structure

0

?

?

4 . Ground-mass

?

0

0

5. Green nuclei

depends on the structure : 1-2-3 or 4

skeletal cavities 2. Replacement of substrate

6 . Fissure filling

+

(late)

?

0

0

0

Table 1 indicates the possible mechanisms which may be suggested for the different habits quoted above. The main problem appears to be the interpretation of the oolitic habit for which three processes of genesis (three question marks in Table 1) are possible. THE PROBLEM OF THE OOIDS

Interpretations of the ooids The interpretation of the oolitic structure is still much debated. Although ferriferous ooids rarely constitute the major portion of the rock, they are highly conspicuous. They may consist of a variety of minerals: iron oxides, iron hydroxides, various clay minerals, or mixtures of these giving a wide range of colour and textures. Geologists have usually characterized the iron-rich sediment by this very specific constituent, sometimes forgetting the presence of other less noticeable forms of green clay. However, the selective study of the oolitic grains is justified because the understanding of the environment of genesis of the whole formation largely depends on the interpretation of these concentric structures. The various interpretations of the oolitic green grains are generally concerned with two main themes: the site of formation and the mechanism of formation of the envelope. The former is usually discussed in

31 terms of development above or below the sea-water/sediment interface, the latter in terms of successive additions of layers around a nucleus, or focused cementation, or replacement without volume increase of a previous substrate. Four main interpretations of the ooids exist as summarized in Table 2. Table 2. Four possible interpretations of the ooids that occur in ironstones; the corresponding genetic interpretationof the ellipsoidal pellets is given in parentheses. I

Mec hanisrn Location

I

Growing around a nucleus

Replacement or filling of a previous substrate

II

I

Concentric structure Accretion process Above sediment surface

(true ooids)

I Below sediment surface

1l1

Concentric structure Concretion process (intrasedimentary concretion)

N o true concentric

structure Rep 1ace ment or secretion process (faecal pellet or ellipsoidal carbonate structure

IV Concentric structure Replacement process (inherited ferric or carbonate ooids 1

These diverse interpretations are important not only in terms of the actual mechanism of formation of the ooids, but also in that they imply different relationships between the femferous clays and the source of iron. The following paragraphs will discuss four possibilities i.e., the extra-sedimentary accretion hypothesis (section 2.2, p. 31); the intra-sedimentary "concretion" hypothesis (section 2.3, p. 32); and the replacement hypothesis for which two different sorts of substrate are considered in the literature: true ooids (section 2.4, p. 34), or non-oolitic structures (section 2.5, p. 38).

The extra-sedimentary accretion hypothesis Hypothesis I in Table 2 was supported by Bichelonne and Angot (1939) and appears to be well examplified in the case of the Jurassic Alpine ironstones of Switzerland (Chapter A l , and Fig. 1 below). The ooids are of the concentric type (Knox, 1970) with a nucleus surrounded by a regular concentric cortex, which implies a progressive addition of material layer by layer as in a typical marine carbonate mid. The size, form, and granulometry of the green grains are compatible with an accretion of oolitic crusts around particles that lay near the surface of unconsolidated muddy sediment, possibly accompanied by remobilization of the grains.

32 Gehring (1983, among others, is of the opinion that the iron ooids (goethite) from the northernmost Jurassic outcrop of the Swiss Jura also formed by a true oolitic accretion process at the sediment surface with alternating accretion of phosphate and iron hydroxide. For this author, the formation of green and red ooids close together or of alternating green and red cortices within one mid, suggest an ooid accretion mechanism that was periodically influenced by more oxidizing conditions at the sediment surface. Knox (1970) indicates that the ooid envelopes are not always as regularly concentric as illustrated in the Swiss outcrops. He explains the eccentric nucleus in some ooids from a Bajocian ironstone of Yorkshire as due to an accretion process on nuclei which were stationary at the top of the sediment. This interpretation is compatible with hypothesis I in Table 2, but requires that the bottom current, if present, is too slack or intermittent for traction and saltation of the nucleus to result in the development of a concentric envelope; these are the 'quiet water ooliths' of Freeman (1963). This observation is fundamental and introduces the possibility that oolitic structures can be formed without movement of the accreted grains. Hypothesis I, when applied to green clay ooids, carries with it two possible implications. The first is that the marine bottom water is reducing, in order to allow the formation of a ferrous green clay, a theory supported by Taylor (1949), or Borchert (1952, in Harder, 1964). Alternatively, the sea-bottom is not necessarily strongly reducing during formation of the ooids, because it is possible that the green clay is formed before ooid formation and is physically accreted, not chemically precipitated, onto the mid (Knox, 1970). It is also possible that the material initially accreted is a ferric clay or that the clay is a replacement of a primary non-clay substrate (see p. 34-38). These last alternatives could appear to solve the paradox of ferrous clay being associated with shallow water sediment, as emphasized in Chapter Al.

The intra-sedimentary concretion hypothesis Due to the prerequisite for the ferrous ironstone green clays to form in a reducing environment, various authors have followed Schellmann (1 969) in postulating that the concentric internal structure of the ooids resulted not from true oolitic accretion but from an intra-sedimentary concretion process. The presence of oxygenated water above the sediment surface presents no problem to the formation of ferrous ooids in this way. During a study of Palaeozoic ironstones Chauvel and Guerrak (1986) observed that the oolitic ironstones are often associated with argillaceous sandstone

-4-+-+

with carbonate. Photo. 2: detail of the upper left comer of Photo.1. Photo. 3: detail of Photo.2 with a composite mid probably indicating a formation at the surface of the sediment in a calm environment. Photo. 4: thin section of an mid with an m i d fragment as nucleus, crossed nicols; note the black cross due to the tangential arrangement of the clay microcrystals. This specific arrangement supports an accretion process in free water or within a very soft water-rich mud. Photo. 5: another example of composite mid.

33

Figure 1. Thin section of the oolitic ironstone of Chamoson. The white bar represents 5 mm (Photo.1); 0.5 mm (Photo. 2); or 0.2 mm (Photo. 3, 4, and 5). Photo. 1: general view showing the abundance of grains with oolitic envelopes around variously shaped nuclei. Most ooids are composed of green clay; a few are black (probably hematite); fissures are filled

34 composed of abundant demtal gains in a chamosite-rich matrix. The groundmass often shows concentric organization around the detrital grains. The latter are sometimes surrounded by concentric chlorite laminae, with the boundary between this concentric structure and the groundmass being gradual. A second feature of the Sahara ironstones is that the ooids often enclose several nuclei (Chauvel and Massa, 1981; Guerrak and Chauvel, 1985; see Fig. 2). According to Guerrak (1987), this habit is not compatible with the presence of ooids in agitated water. Such composite ooids, with one or several oolitic structures surrounded by a common envelope, are numerous in the Saharan formations. They are interpreted by Guerrak (op. cit.) as having originated through the fragmentation of oolitic sediment with the oolitic fragments subsequently being redeposited and then submitted to a second intra-sedimentary concretion process. However, the same fragments could also be used as nuclei for an accretion process on the surface of the sediment. A strictly concretionary hypothesis could not explain an oolitic structure with alternation of ferric oxide and ferrous silicate laminae, as was pointed out by Gygi (1981). This might be explained by considering that the concretion (future mid-like structure) formed within a silica- and iron-rich mud but near the sediment surface. Episodic winnowing of the sediment from around the concretion would expose its surface to oxidation, with the formation of an external layer of iron oxi-hydroxide. Repeated burial and winnowing would allow the concretion to develop alternating layers of iron oxide and chlorite (Joseph and Beaudoin, 1983; Guerrak, 1987).

The ooid replacement hypothesis The hypothesis IV in Table 2 implies that the green clay ooids found in many ironstones are replacements of primary ooids of a different mineralogy. It solves the incompatibility between a presumably oxidizing environment, where carbonate or hydroxide ooids form, and the ferrous character of the green clay in the oolitic cortices presently analysed. The replacement could take place within the sediment or above the sediment surface.

-+++

Photo.2: Ahnet area (Famennian, Central Sahara, Algeria). Polarized reflected light. Double mid composed of two hematite dominating ooids enveloped by chamositic layers. Photo.3: Ahnet area (Fmennian, Central Sahara, Algeria). Polarized reflected light. Three chamositic-hematiticooids surrounding an hematitic intraclast (white), are enveloped by thin alternating chamositic (grey) and hematitic (white) layers. Photo.4: Mecheri Abdelaziz deposit (Famennian, Western Sahara, Algeria). Polarized transmitted light. Composite ooid made of seven chamositic (white) and hematitic (grey) ooids and one broken mid. Photo.5: Tassilis NAjjer (Lochkovian, Eastern Sahara, Algeria). Polarized reflected light. Example of envelopes around quartz grains (grey). The primary chamosite has been completely transformed to hematite and goethite (white). Photo.6: Ougarta (Arenig, northwestern Sahara, Algeria). Polarized transmitted light. Development of chamositic-hematiticooids around detrital grains (white).

35

Figure 2. Sections of the oolitic ironstones from the Sahara showing various types of oolitic envelopes. All scale bars show 100 pm. Photo.1: Ahnet area (Famennian, Central Sahara, Algeria). Polarized reflected light. Composite ooid formed of two chamositic (grey) and hematitic (white) ooids surrounded by a first chamositic cortex (grey) and by f i e chamositic (grey) and hematitic (white) cortices.

36 Replacement of Carbonate ooiak The hypothesis that the primary substance of ferriferous ooids was carbonate was put forward by Sorby (1856); it was supported by Cayeux (1909) and in recent years by Kimberley (1979), Kearsley (1987), and Dreesen (1987). The last author concluded that ironstones most probably 'originated through early diagenetic replacement, coating or impregnation (by Fe-hydroxides and chlorite) of originally calcareous allochems'. The hypothesis of carbonate mid replacement, (hypothesis IV, Table 2) during early diagenesis or later, appears to be supported by the observation of replacement of various non-oolitic carbonate substrates by green clay as quoted by many authors (see p. 38-39). Kimberley (1974, 1978) has put forward the hypothesis that primary marine carbonate (aragonite) ooids are replaced by ferrous green clays in a continental environment after regression, with iron and aluminium being eluviated from surrounding clay minerals. The replacement hypothesis would be more acceptable if relicts of the initial carbonate ooids were present in the oolitic ironstones, or if the ironstone and carbonate ooids were horizontally or laterally often associated. On the contrary, however, this facies association has not been found in nature and some sort of incompatibility seems to exist between ferriferous oolitic and carbonate oolitic formations (Chapter Al). Replacement offerric iron-richooidr Following Caill&reand Kraut (1954) or Bubenicek (1969), and after studying the Oxfordian (Jurassic) ironstones of southern England, Talbot (1974) proposed the hypothesis of a primary marine accretion of ferric ooids -above (or very near) the sediment surface- because he noted an absence of irregular or interference growth patterns in chamosite ooids. The authors quoted above support the view that the coating material accreted onto the mid is a gelatinous, ferric iron-rich aggregate which, after reworking and burial, transforms to a ferrous ironstone green clay (hypothesis IV in Table 2). Although marine ferric hydroxide precursors are favoured by most authors, Siehl and Thein (1987) proposed that the Aalenian (Jurassic) green clay ooids of the minette of Lorraine resulted from the post-depositional marine diagenetic modification of soil-derived ooids initially made of ferric oxi-hydroxides. The experimental replacement of a ferric iron-rich precursor by a ferrous green clay mineral has been carried out by Harder (1964; 1978; 1987). This author emphasizes the necessity of a negative Eh for quick crystallization of ironstone-like ferrous green clay. The initial mixture from which the clay was suspected to be formed was precipitated in an oxidizing solution with silica and ferric hydroxides in adequate proportions. These theoretical precursors are compatible with the environment where natural ironstone green clays seems to form and the hypothesis in this paragraph is a good candidate for explaining the origin of some green ooids. However, as shown in the Swiss Alps the ferric state of the iron element in some ironstones can be explained as a meteoric alteration product rather than as a precursor of the ferrous clay. A complete replacement of oolitic goethite by

37 green chamosite or berthierine is difficult to imagine without modification of the well-defined concentric structure initially present; a molecule by molecule replacement can be postulated, however. But a greater problem is the observation that green clay envelopes are commonly more delicately structured than their hydroxide counterpart. Finally, the presence of alternating coatings of goethite and chamosite around a nucleus illustrated by Gygi (1981) in ooids from a Jurassic ironstone in the Swiss Jura, cannot be explained by the above discussed hypothesis which only considers the argillization of an entire mid following its formation.

The accretion-replacement hypothesis The two processes of the extra-sedimentary accretion and the replacement have been combined by several authors as a means of explaining both accretion in an oxidizing environment and ferrous clay mineral genesis in a reducing environment, this being the paradox that remains the crucial problem for most authors. This combined hypothesis was proposed by Gygi (1981) after the study of an ironstone in the Swiss Jura. Gygi (1981) postulated that each green clay lamina as defined by Knox (1970) is the result of a two-step genesis. During the first step, the nucleus, together with its ferriferous envelope, is rolled by a water current at the surface of muddy sediments and is coated by iron hydroxides and silica, probably in the form of a gel, in an environment with a positive Eh. During the second step, the mid is buried in mud that contains clay and organic matter. Decay of organic matter within a fresh sediment creates a reducing environment favourable for the crystallization of ferrous clay from the gel coating the mid. This cycle is repeated as long as episodic currents are able to remove the thin layer of mud covering the ooids between the accretion of successive laminae. Alternating accretion and replacement, repeated many times, would then produce such mids. This mechanism allows us to understand the alternation of femc goethite and ferrous green clay laminae in some ooids, if we assume that sometimes a goethite lamina accreted during step 1 remains unaltered or is only partly altered to green clay during step 2. This could occur either because of short-lived burial or because burial is insufficient for the environment to become sufficiently reducing. In that case, the goethite lamina remains as it is within the oolitic envelope. We have to explain how the oxide layer may become reduced during burial, but also why the previously formed reduced clay lamina may remain unoxidized on re-exposure. This suggests that the reducing environment in the sediment is more aggressive to iron oxides than the oxidizing extra-sedimentary environment is aggressive to the ferrous clay, or that clay laminae are less permeable than goethitic ones. The environmental factors involved and able to concur with the above duality remain to be precisely defined and examplified in the sedimentary record; we here discuss only those models in the literature for which more or less convincing scenarios were proposed. In summary, episodic currents between long intervals of quiescence may allow the entrapment of ions or other components around a nucleus followed by mineral conversion of this material in a more reducing environment, below

38 the sediment surface. Re-exposure does not necessarily alter the previously formed minerals. The process ends when the current is no longer able to remobilize an mid, either because too heavy or because too deeply buried.

The non-oolitic hypothesis Some authors have come to the conclusion that a number of ellipsoidal structures observed in ironstone did not result from a physico-chemical addition of material around a nucleus (Champetier et al., 1979, 1987a, 1987b). This is hypothesis I1 in Table 2. These authors do not reject the occurrence of concentric mid-like structures but emphasize the fact that one could confound several sorts of bioclasts with ooids. The non-oolitic possible substrates for green clay formation suspected by these authors in two ironstone formations (Aalenian of Lorraine, France, and Devonian of the Gara Djebilet, Algeria) are of three types: 1) Oncoid particles, which are algal balls formed by a concentric accretion of algal or cyanobacterial layers. 2 ) Faecal pellets, of anomuran crustaceans (Favreina sp.) or other pellets, that show two characteristic sections. The transverse one is clearly recognizable by a pattern of spany calcite dots which are cross-sections of canals formed by elongate appendices from the intestinal walls. 3) Nubecularia, an incrusting foraminifer of the Miliolidea. Initially, the wall of this foraminifer is cryptocrystalline calcite associated with iron and organic matter. The test is plectogyroid with pear-shaped chambers in longitudinal sections. The streamlined parts of the chambers are opposite from one pole to the other. Champetier et al. (1979) claim that they have discovered Nubecularia in the Devohian of Gara Djebilet; these forms would then be older than their vertical range as known to date (Carboniferous to Recent). The common characteristics of all of these constituents are their biogenic origin, their non-ferriferous nature and their organic matter content. They are often associated with crusts formed by microbial mats (Dahanayake and Krumbein, 1986). If we suppose that these biogenic structures can be filled or replaced by green clays, the primary precipitation theory (the accretion theory of Bichelonne and Angot, 1939; Caillkre and Kraut, 1954; Bubenicek, 1961) could not be applied. However, in spite of the presence of abundant green clay in the groundmass of the. sediments studied from Lorraine and Algeria, the ellipsoidal constituents considered above have never been observed to be filled with, or replaced by green clay. Moreover, although oncoids or Nubecularia may have a concentric structure, this is not the case in faecal pellets, where confusion with ooids must result from an inadequate observation. Y. Champetier is convinced that the predominant portion of the "ooids" described from oolitic ironstones do not possess true oolitic structure but are the result of replacement processes (personal communication). Examples of infilling or of replacement by green clay of more or less ellipsoidal bioclasts (without concentric structure) have been observed in thin sections. Chapter A1 has illustrated the role of crinoid remains. Among others,

39 Kearsley (1987) reports that Hettangian (Jurassic) ironstones of Skye (Scotland) contain evidence of fragments of brachiopod shells, crinoid ossicles, foraminifers or gastropodes being replaced by chlorite in which the original shell structure is very well preserved. In addition to the replacement of carbonate bioclasts, we may quote the replacement of detrital mineral fragments. For example, Bubenicek (1961) has illustrated quartz grains being partly replaced by a green clay in the Aalenian (Jurassic) iron ore of Lorraine. THE GREEN CLAY OF THE GROUNDMASS

Green clay is common and even abundant in the groundmass of sedimentary oolitic ironstones. The question here is to know whether this green clay can be interpreted as a synsedimentary product formed by a precipitation process within soft sediment (or above it), or by early or late diagenetic precipitation in more or less lithified sediment (see Table 1, p. 30, the two question marks in columns precipitation and replacement). A scrutiny of the literature reveals that a variety of interpretations exists but that the majority of them propose that the green clay groundmass of the ironstone has formed following a synsedimentary process (Taylor, 1949). For example, Knox (1970) is of the opinion that his curiously structured ooids from the Bajocian (Jurassic) ironstone of Yorkshire have formed at the sediment surface of green clay mud. More recently, Knox and Fletcher (1987) indicated that the green clay groundmass of the Sinemurian (Jurassic) Frodingham ironstone of eastern England can be interpreted as the primary muddy constituent of the sediment in the site of final sediment accumulation. The formation of original mud with green clay very early in the sedimentation, is also evidenced by the interpretation of DCverin (1945) who thought that some of the nuclei of ooids in Jurassic ironstones of Switzerland are fragments of previously hardened chamosite mud. According to this theory, the ferrous green clay of the groundmass must have crystallized very close to the sediment surface. This indicates that the zero Eh plane was located near the top of the sediment during deposition. These observations contradict the hypothesis of an oxidizing and agitated water environment above the sediment. However, an oxidizing environment at the sediment surface is indicated by the presence of an abundant benthic fauna in the Schellenbriicke Bed described by Gygi (1981). Supplementing the synsedimentary hypothesis, we must remember that some chlorite was generated during a later diagenetic episode. For example, Bubenicek (1961) has illustrated the crystallization of chlorite coating the walls of pores and welding grains of the rock. This relatively early diagenetic formation of chlorite has also been described and illustrated from many sedimentary formations other than ironstones (Hayes, 1970). Probably a similar (but much later) process has been responsible for the formation of chamosite as fissure filling in the site near Chamoson. In summary, the major portion of the green clays in the groundmass is usually interpreted as a synsedimentary product, but other authors consider it as

40 diagenetic. Continued diagenetic modification of this material is probable. It is thus not known to what extent our analysed green clay from the groundmass is representative of the clay originally formed, or is the result of later chemical or crystallographic modifications, (recrystallization) during early diagenesis. MINERALOGY OF THE IRONSTONE GREEN CLAYS

It is important to bear in mind that the oolitic ironstones have generally been submitted to various burial, tectonic, or meteoric evolutions since their deposition at or near the sediment surface. They are comparatively unstable because of their high iron content, because this element can easily change its state of oxidation. Therefore, we begin in this section,with an account of the present mineralogy of the green clays, followed by a short discussion of the possible original composition of these clays.

Green clays of the ooids Berthierine and chamosite are two different terms that were proposed for the fermginous clays 150 years ago (Berthier, 1820; Beudant, 1832). However, the precise mineralogical distinction of these two minerals was not made until Orcel et al. (1949) distinguished on the one hand a mineral from the chlorite group and on the other a mineral from the serpentine group. X-ray diffraction patterns of the chlorite mineral show a small 14 8, reflection which increases when the mineral is heated to 530°C and 600°C; the differential thermal analysis diagram of the mineral shows two endothermal reactions, representing two sorts of OH groups, one in the interlayer (hydroxide sheet) and one in the 2:l layer (talc layer). X-ray diffraction patterns of the serpentine mineral show a 7 8, reflection which is destroyed after heating to 530"C, and the differential thermal analysis indicates the presence of a single endothermal reaction with only one sort of OH group in the mineral. Figure 3 gives X-ray diffraction patterns of two of the samples probably analysed by Orcel et al. (1949). These samples are from the collection of the Mustum National d' Histoire Naturelle (Paris), and are labelled MusCum 53-300 (the chlorite from Chamoson) and MusCum 95-334 (the berthierine from Hayange, Lorraine). The whole rock was analysed for sample 53-300 and the obtained powder was greyish; separated ooids were analysed from sample 95-334, the obtained powder being light-green. For comparison with the procedures of treatment considered in part B for clay minerals of the verdine facies, the two samples were leached with acids using normal solutions heated to 70°C for one hour. Neither the hydrochloric nor the acetic acid leaching modified the chamosite X-ray diagram, but the HC1 slightly lightened the colour of the powder. The berthierine from Hayange was not modified after leaching with acetic acid but its structure was completely destroyed by the hydrochloric acid leaching under the conditions quoted above: the corresponding X-ray diffraction pattern was flat and the initially green material became white.

41

r

001

, 002

1

I .

340

32'

280

24O

209

1W

120

80

__

..

4'

Figure 3. X-ray diffraction patterns of chamosite from Chamoson: 53-300 and the berthierine from Hayange: 95-334. The bulk sample from Chamoson is a mixture of chamosite, siderite, calcite, and quartz. The green ooids from Hayange (Lorraine) are pure berthierine. X-ray power: 780 W, sensibility: lo3.

In addition to chamosite and berthierine, a third component has been reported from a few ironstones under the name swelling-chlorite. This clay typically shows a 14 8, diffraction peak; but, treatment with ethylene glycol produces a slight swelling, whereas heating does not collapse the layers toward 10 A. However, the precise behaviour of the swelling-chlorites is not reproducible; Brindley (1961a) and Carroll (1966) remark that some of these clays partly collapse to 10 A when being heated. The chemistry of chamosites has been discussed in the preceding chapter and that of berthierine will be further discussed in the second part of this volume. Natural chamosites and berthierines (Brindley, 1982) have a similar general chemical composition. This composition is characterized by a very high iron content, the majority of which is reduced to ferrous state. Therefore, with their large proportion of magnesium and ferrous iron, the green clays from the ironstones are all trioctahedral clays. This character is found similar in all the sedimentary chlorites analysed by Hayes (1970). There is, however, one paper which is frequently quoted in the literature (as usual when an exceptional material is mentioned) in which an example of ferric chamosite is documented (Brindley and Youell, 1953). These authors studied Jurassic berthierines from England and showed that the natural ferrous iron of

42 the mineral can turn to ferric when the substance is heated to 400°Cin air. These conditions hardly occurred in nature, however. Moreover, heating of the berthierine produced a red colour similar to the one shown by all of the ferrous clays when heated in air above 200-300°C. Finally, the initial H20+ content of about 10% for berthierine was lowered to less than 3%. It is our opinion that the "mineral" synthesized by the authors has nothing in common with natural clays and that therefore the mythical "ferric chamosite" still remains to be discovered in ancient ironstones. More will be said below on femc green clays (see Chapter B5).

Presumed initial green clays A number of synsedimentary green clay minerals of many Recent sediments from most oceans and seas have been analysed (Giresse and Odin, 1973; Lamboy and Odin, 1975; Odin, 1972, 1973, 1975; Odin, Debenay et al., 1987 in press; Odin and Dodson, 1982; Pujos and Odin, 1986). There is a considerable difference between these minerals collected from the present sea-bottom and those from the ironstones. Disregarding the question of the chemistry, this difference is also of crystallographic nature: the shape of the X-ray diffraction peaks obtained from ancient ironstone green clays is much too "perfect" to be that of a clay mineral formed in the marine environment (Fig. 4). This point has rarely been considered in the past. The "crystallinity" of berthierine and chamosite is much better than that usually observed in minerals from the verdine and recent or ancient glaucony facies, with the X-ray diffraction peaks being much higher and narrower in the ironstone minerals. They are almost certainly minerals which have been recrystallized to a better organized structure in a diagenetic environment without contact with the open sea-water. Such recrystallization is a common feature of authigenesis during burial. It does not necessarily result in substantial modification of the components concerned and the basic crystal structure (mineral family) is usually preserved. Therefore, our present opinion is that the synsedimentary green clays were in many respects comparable to those of the present ironstone, but differred in possessing a much poorer crystallinity . The question of the precise primary mineralogy i.e., serpentine or chlorite structure, has been debated. The general view, as expressed by Millot (1964, p. 246) and discussed by various authors, including Delaloye (1966) in his work on the outcrop near Chamoson, is that the initial clay was a serpentine-like mineral and that the development of a chlorite structure is related to the age and burial history of the rock. It is a common feature that Mesozoic ironstones contain mainly serpentine-like minerals whereas Palaeozoic ironstones contain mainly chlorite minerals. The general opinion, repeatedly quoted in the literature (Iijuma and Matsumoto, 1982; Van Houten and Purucker, 1984; Maynard, 1986) is that a 7 8, phase precedes a 14 8, phase. However, the question is confused by the incorrect use of the mineralogical term chamosite, when berthierine is concerned, by many dilettante mineralogists (including some of us) although the two terms have been officially distinct for 20 years

43 (Brindley et al., 1968; Bailey, 1980). In absence of a detailed mineralogical study, a general term like green clay would properly replace the incorrect use of a mineralogical term for the designation of a facies.

C

001

Ka 28"

24O

200

16O

120

Cu

ao

Figure 4. X-ray diffraction patterns of randomly oriented powders of the chamosite from Chamoson (Museum 44-45) and of a phyllite V from Senegal (sample 601-95). The latter shows the highest diffraction peaks for a Recent green clay of that sort. It is probably a mixture of 14 A and 7 8, clay minerals. Emission power: 910 W, sensibility: 4.102 for the two X-ray diagrams. C = carbonate; Q = quartz; S = siderite.

Based on the present-day models of oceans, different primary green clay minerals might have been generated in the very specific and diverse environments that occur at the boundary between oceans and continents. Three arguments may be put forward for this interpretation: 1) The postulated recrystallization of a 7 A clay mineral to a 14 8, clay mineral is a very probable possibility. However, if that process was universal for all ironstone formations, the occurrence of intermediate minerals, i.e., interstratified structures (7 A-14 A), should be common and probably more common than the end members. Neither of these conditions appears to be met in the ironstones studied so far.

44

2) The ooids in ancient oolitic ironstones may be composed of several clay minerals: 7 A, 14 A, smectite (Chapter Al). This could be interpreted as indicating that the primary clay was probably not monomineralic. Furthermore, in addition to the presence of berthierine and chamosite, swelling-chlorites have also been observed in ancient formations. These observations indicate that the general environment in which oolitic iron-rich sediments accumulated might have been favourable for the formation of diverse green clay minerals in different localities, or at different times. 3) Recent sediments contain neither true berthierine nor true chlorite. However, the study of Recent nearshore sediments indicates that different parageneses exist where diferent green clay minerals are formed in comparable environments. It is thus possible that a chlorite-like, a berthierine-like, and a swelling-chlorite mineral might have formed concurrently at the bottom of ancient seas. DISCUSSION OF THE ENVIRONMENTAL FACTORS

General environment of oolitic ironstone formations There are a very few points on which everybody agrees with respect to the general environment of oolitic ironstones genesis. One such point is that the ooids are formed somewhere near the boundary between the continent and the ocean (see Fig. 5). However, brackish or even evaporitic environments are sometimes considered (Dreesen, 1987); sheltered sea-water lagoons or embayments are also suggested (Talbot, 1971). Deposition in shallow water on a flat delta, near sandwaves or barrier islands, was proposed by Siehl and Thein (1978; 1987) for the minette of Lorraine; the goethite ooids are thought to have been derived from the fluvio-marine reworking of pedogenic goethite ooids, but the ironstone green clays are thought to have been formed in the shallow marine environment. That shallow marine environment is frequently said to be much calmer than necessary for carbonate mid genesis (Knox, 1970; Talbot, 1974). Gygi (1981) concluded that the green ooids might well have formed at water depths of up to 100 m. In contrast to these marine hypotheses, Kimberley (1979) proposed a model in which the aragonitic ooids, originally formed at the bottom of a shallow inland sea, are replaced by "chamosite"on land following a regression. Many authors have stressed the problem of the source of iron; this aspect was central to the proposal by Kimberley of a continental genesis of iron ooids because he pointed out that there is too little iron in solution in sea-water to account for a marine genesis. All authors agree that oolitic ironstones accumulated at a time of reduced terrigenous influx or of non-deposition (Dreesen, 1987; Knox and Fletcher, 1987). Gygi (1981, Fig. 2) documented this with ammonite successions. Van Houten and Purucker (1984) assumed a condensed facies where iron can be concentrated. This iron is presumed to be derived from rivers draining the adjacent landmasses (Talbot, 1974). The proximity of a substantial terrestrial run-off is therefore a prerequisite for a large oolitic iron-

45

stone formation to develop; this aspect will be elaborated further in Chapter B6. A zone of iron concentration at the mouth of present-day rivers has long been suspected (Odin, 1975b; Harder, 1978, between others) and was recently discussed in the case of tropical rivers (Pujos and Odin, 1986), where the heavy load of iron is accompanied by efficient processes of accumulation by precipitation. Oolitic or sub-oolitic ferruginous structures result sometimes from the precipitation of this iron in Recent tropical areas (Mahakam delta in S Borneo: Allen et al., 1979; Ogooue delta in Gabon: Giresse, 1965b; off Ivory Coast: Leneuf, 1962). Part B of this volume will examplify that point.

Figure 5. The regional environment of formation of the oolitic ironstone facies according to various authors: 1. Dreesen, (1987); 2. Siehl and Thein, (1978, 1987); Nahon et al., (1980); 3. Talbot, (1971); 4. Kimberley, (1979); 5. Guerrak and Chauvel, (1985); Guerrak, (1987); 6. Gygi, (1981).

The relationship between the landmass, the pro-delta facies and the marine facies has been examplified in the Early Oxfordian sediments of the Swiss Jura (Gygi and McDowell, 1970; Gygi, 1981; see Fig. 6 below). From the land (north-west)to the open sea (east), three palaeogeographical zones are present. They are represented by 1) a 15 m thick pro-deltaic facies of predominantly argillaceous mud (Terrain B Chailles) which grades toward the deeper sea into 2) an intermediate facies: a 0.25 m thick oolitic ironstone bed and then into 3) an outer facies: a thinner glauconitic bed. This sedimentary array clearly indicates where the source of terrigenous mud was, and that glaucony was formed even farther from the fluvial input in a fully open epicontinental sea.

46 I

Figure 6. Lower Oxfordian oolitic sediments in northern Switzerland. The map gives an indication of the geographic extent of the three facies concerned: 1) pro-deltaic "Terrain B Chailles", 2) oolitic ironstone (Schellenbrucke Bed) and 3) glauconitic horizon. (Modified from Gygi, 1981)

Many authors agree (Bubenicek, 1964; Hallam and Bradshaw, 1979) that oolitic ironstones were formed at the end of a regressive cycle of sedimentation, for example at the top of coarsening-upward sequences in the Devonian of the Sahara (Guerrak, 1987) or shoaling-upward cycles in the Devonian of Ardenne (Dreesen, 1987). Fifteen such sequences were identified in the Jurassic rninette of Lorraine (Siehl and Thein, 1987). Finally, large ironstone deposits appear to coincide with regional regression, according to Guerrak (1987). From a climatic point of view, Siehl and Thein (1987) indicated a palaeolatitude of about 35"N. The Oxfordian sediments of northern Switzerland lay at a similar palaeolatitude and the facies indicates that the climate was at least subtropically warm and seasonal (Gygi, 1986, p. 146). A tropical climate is assumed by many authors who postulate an abundant vegetation on land to account for the leaching of ions, especially iron from the soils without significant

47

mechanical erosion. However, a temperate climate is sometimes considered for the Devonian of Algeria (Guerrak, 1987).

Microenvironment of green clay genesis in ooids From the microenvironmental point of view, the observation of ooid formation is informative. All authors agree that the dominant factor for the genesis of the presently analysed green clays is a low Eh, allowing ferrous iron to be incorporated in green clay minerals. Moreover, the green clays are generally accepted to be formed near the sediment surface. The debate is: how and when were these reducing conditions established? An apparent paradox must be resolved between the presence of both: (1) ooids c=> shallow + agitated oxidizing environment versus ( 2 )berthierine-chamosite c=> ferrous clays formed in a reducing environment. The solutions proposed so far for this paradox are many and may be broadly grouped in two categories: Cutegory I : The two terms of the above paradox are not interrelated in time and/or space. Category 2 : The oolitic grains are not formed by extra-sedimentary accretion in free shallow, agitated and oxidizing water. Let us briefly remember the discussion of solution 1. A two step model has been preferred by many authors who have proposed different solutions. The intra-sedimentary "verdissement" (replacement) of previous carbonate ooids is one of the oldest hypotheses (Sorby, 1856). However, remains of initial carbonate ooids have never been found. The replacement of ooids initially made of iron oxi-hydroxides following (or without) lateral transport from an area with an oxidizing water to a deeper and more reducing site, would only be acceptable in the absence of nuclei that are fragments of green ooids, or of alternating laminae of ferric and ferrous minerals. Cyclic reworking of ooids is preferred by Gygi (1981). This model combines, for each lamina, a physical accretion process in an oxidizing environment and a replacement process in a reducing environment. According to Gygi (1981), iron ooids are coated with a layer made of 'Fe (OH), and clay minerals' in well-aerated sea-water while being rolled by episodic, oscillating water currents at the surface of muddy sediment. The same episodic, storm-driven currents that rolled the ooids, stirred up muddy sediment into suspension. After the storms, many of the ooids would be partially or completely buried by the settling mud. Provided this mud was rich in organic matter, a negative Eh would develop during the long quiet intervals between storms, within the interstitial water of the mud. Under these conditions, a ferrous green clay could be formed 'out of clay minerals and Fe(OH),' (i.e., by replacement) in the outermost layer of deeply buried ooids. In the external laminae of the ooids that remained at the sediment surface, the iron would remain in the ferric state. Winnowing and burial of ooids would be repeated many times.

48 The above quoted category 2 also comprises several models. The first proposal is that ooids are the result of a concretionary process below the sediment surface. In the scenario proposed by Guerrak and Chauvel(l985) or Guerrak (1987), there is in fact a combination of a two step model with a concretion process. This proposal differs from the above suggestion by Gygi (1981) in that the material is added to the previous grain within the sediment (not above it) in a suspectedly reducing environment. An intermittent reworking process is also postulated but its effect is indeed to modify the initial Fe2+to Fe3+in contrast to the proposal by Gygi who postulates a modification from the original Fe3+to Fe2+. However, the very regular shape of the laminae of most ooids of the Mesozoic ironstones perfectly "imitate" true oolitic structures freely grown at the sediment surface; but the crystallite arrangement should also be considered. The iron- stone green clay ooids show a tangential arrangement; but the concretions found in many rocks like continental laterites and marine sands, where chlorite develops by precipitation in the intergranular space during diagenesis, as well as the very similar case of the cortex added to glauconitic grains after early burial (Zumpe, 1971; Lamboy and Odin, 1974), all show a radial structure. However, Richter (1983) indicates that the orientation of crystallites in the carbonate ooid envelopes cannot be used as a general criterion for the depositional environment. Radially oriented crystals exist in marine carbonate ooids (Gygi, 1969, P1. 9); moreover, Calvet and Julia (1983) describe tangential structures in soil pisoids. Richter (1983) reminds us that radial carbonate ooids appear specific for quiet sea-water and tangential ooids for agitated sea-water; this was confirmed by Strasser (1986, Fig. 6). Moreover, the tangential arrangement might be more apparent than real and depends on the scale of observation. Scanning electron microscope studies have shown that the microcrystals within individual laminae are randomly oriented, whereas the optic microscope shows the black cross phenomenon (Fig. 1 above). The poorly defined ooid margins that where interpreted by Guerrak (1987) as incompletely formed laminae during a concretionary process provide evidence of a diagenetic recrystallization rather than ooid accretion at the sediment surface according to this author. However, they have not been reported from Mesozoic or younger ironstones. The presence of polynucleated ooids and composite ooids is considered to be a serious argument against an extra-sedimentary accretion process in agitated water. Such ooids were also observed in Mesozoic formations. However, these habits (polynucleated ooids and composite ooids) do not exclude the accretion process itself but only the presence of an agitated environment above the sediment surface. This would support the many observations indicating that green ooids are formed in calm water in contrast to the Recent aragonite ooids of the Bahamas. In fact, it is known that many oolitic carbonate grains (or oncolitic structures) result from accretion around a nucleus, without movement of the coated grain and in still water; for example, the cave pearls or the ooids from the Persian Gulf described by Loreau and Purser, (1973) or the 'quiet water ooliths' of Laguna Madre (Freeman, 1963). In these two examples, the carbonate microcrystals

49 are radially oriented (Richter, 1983). The accretion in calm water is a process which is compatible with the presence of ooids showing several nuclei as well as with re-oolitized ooids. In summary, the concretion process could explain some situations observed in nature. This process requires complex cyclic environmental changes to be effective. A second hypothesis has to be considered when assuming that ferriferous ooids are not formed by accretion when rolled at the sediment surface. This hypothesis takes into account the observation that mid-like grains are not ooids but ellipsoidal faecal pellets or foraminifera1 infillings. Table 2 locates the process discussed here above the sediment surface but the filling of microtests of foraminifers, gastropodes etc., of micropores in echinoderm bioclasts, or the mineralization of faecal pellets may occur below the sediment surface as well. What is significant here, is that the favourable microenvironment of chlorite crystallization is located inside biogenic remains which were probably rich in organic matter. Above or below the sediment surface i.e., in a generally slightly oxidizing or reducing environment, the respective microenvironment will be partly isolated, possibly reducing, and therefore, favourable to the growth of ferrous minerals sheltered from the diluting and oxidizing action of the open sea-water. This environment is physico-chemically similar to the semi-confined microenvironment defined by Odin and Matter (1981) for the genesis of glauconitic minerals (see Chapter B5). Non-oolitic green grains do exist in the ironstones, but they do not help in solving the problem of the undoubtedly oolitic habit of the vast majority of green ooids found in different Mesozoic or younger formations. If these two hypotheses (concretion and cavity-fill processes) are not fully satisfying, then the accretion process of ooids above the sediment surface has to be re-examined. The equation:

ooid c=> shallow + agitated + oxidizing environment has to be modified according to what is discussed above i.e., oolitic structures do not necessarily indicate a constantly agitated environment. This point is important because it provides a simple solution to many problems in which the assumption of a concretion process would require postulation of complex situations. Among other problems, ooid accretion in calm water can easily explain the alternating of ferric and ferrous laminae or that of ferrous and phosphate laminae depending on the chemistry of the water present at the time of formation of a given lamina. The welding of several nuclei or ooids is also possible in a quiet environment. If there is a little agitation in the sea-water, the sheltered lagoons or embayments could have had a general low Eh explained by a large amount of organic matter present both as suspension and at the seabottom. High algal productivity in phosphate- and nitrate-rich waters could be postulated. This is substantiated by the presence of phosphate-goethite ooids (Gehring, 1985). Therefore, it appears that the accretion process at the sediment surface could easily explain the oolitic habit of most of the ironstones as was suggested by Gygi (1981).

50 In addition to the two categories of hypotheses discussed above to explain the apparent paradox: ooid versus ferrous clays, there is a third possibility for which many arguments are presented in this volume. This new theory considers that the green clay minerals in ancient ironstones are secondary minerals and do not represent the initial component(s) formed at the time of sedimentation. We have discussed above the problem of the unusually good crystallinity of berthierine and chamosite as compared with other green clay minerals formed in the domain of the marine surface sediments. Together with the probable recrystallization process inferred in order to explain this particular property of the ironstone green clay minerals, we suggest that the initially formed clays were not dominantly ferrous but ferric. The femc state of iron in these (postulated) clay minerals (as well as the iron hydroxides postulated by Gygi, 1981, and others) would be more compatible with shallow deposition conditions very near to a general marine environment with high Eh. This suggestion can easily be combined with the previously discussed hypotheses that invoke an accretion or concretion or replacement process. Our proposal is not only a theoretical possibility and can be supported by three observations. 1) Formation of green clay goes along, very often, with that of goethite, and there is a smaller environmental gap between the genesis of a Fe3+-rich clay and goethite than between a ferrous clay and goethite. 2) In the light of this first observation, most of the incompatible hypotheses previously discussed could be reconciled on the assumption that a positive Eh could well be present both in the free sea-water above the sediment, and in the interstitial water circulating within the uppermost centimetres of the sediment. Therefore, depending on the outcrops, we could accept an accretion, concretion, replacement, or diffuse precipitation process without real opposition between them. All are crystal growths and might occur in a generally neutral or slightly oxidizing environment. 3) The study of Recent marine authigenic minerals has shown that the postulated dominantly ferric green clay minerals exist and are common in nature. Moreover, possible equivalents of the three sorts of minerals from ancient ironstones have also been discovered; berthierine, chamosite, and swelling-chlorite would correspond to what is called later (PartB): 7 A phyllite V (not berthierine), 14 A phyllite V (not femc chamosite), and phyllite C (some sort of intermediate between swelling-chlorite and smectite). We do not claim that we.have found a perfectly equivalent series of minerals in a similar facies in Recent and ancient sediments. It will be shown later that the ancient ironstone facies is sedimentologically different from the Recent verdine facies. For example, two very specific aspects of the ironstone facies i.e., diffuse green clay minerals and concentric structures, are not present in Recent sediments. But the natural occurrence of the shallow femc green clays in Recent sediments is a convincing model for reconstruction of the real nature of initial clays formed in the probably comparable ancient ironstone facies. All such metastable (femc) initial clays would become recrystallized due to the early diagenetic processes to form more stable, better crystallized minerals, so that the initial clays cannot be found in ancient series.

51 SUMMARY

We have tried to review the conflicting interpretations of ironstone formation in order to understand the palaeosedimentological history of the facies and, more specifically, the nature and conditions of green clay genesis. It has been shown that green clays were characterized by two specific habits: ooids and diffuse clay in the groundmass. These two particular habits are dominant but not exclusive. Replaced carbonate bioclasts, replaced mineral debris, infillings of microfossil tests and possibly mineralized faecal pellets do exist. The four last habits are common to green grains collected from other green marine clay facies (glaucony; verdine). The mechanisms invoked for green clay genesis in the ironstone facies are diverse and include extra-sedimentary accretion, intra-sedimentary concretion, replacement, and precipitation either in voids or diffuse in the sediment. It is suggested here that the variety of interpretations might reflect a real variety of habits of which the reciprocal importance can hardly be estimated in this short study. All habits seem to correspond to an authigenic crystal growth of green clay minerals. From the mineralogical point of view, it appears that the green clay minerals found in the ironstone facies are diverse, and include three distinct phases, namely: berthierine (serpentine group), chamosite (chlorite group), and a swelling-chlorite which is the least common in this facies. All of these minerals are trioctahedral and contain dominantly ferrous iron. From the environmental point of view, the most controversial problem is the apparent incompatibility between criteria indicating an agitated oxidizing environment (the presence of ooids and ferric iron, often present in oxi-hydroxides) and a reducing environment (the presence of ferrous clays). This paradoxical situation has been explained by a variety of hypotheses which have been discussed by many authors. It is difficult to reconcile to all of these hypotheses. However, we presume that a new theory could probably explain most of the apparently incomptabile observations found in the literature on ironstones. This proposal should take into account the characteristics of green clay minerals from Recent shallow marine sediments which will be discussed in the next part of this volume. In the final analysis, we can reasonably postulate that the ferrous green clays now found in the ironstones are diagenetically modified clays which originate from similar precursor minerals in which the iron was dominantly in a ferric state. It is reasonable to suppose that such dominantly ferric clays would be similar to those which are systematically found in Recent sediments all over the world. We do not believe, however, that the ironstone facies has an equivalent or even direct precursor in the Recent continental shelves. Our proposal only concerns probably equivalent minerals, though the habits are different. We suggest that this new point of view could well be considered in the future as sufficiently valid either to complement or to replace some previously accepted hypotheses.

52 ACKNOWLEDGEMENTS

The authors are grateful to F. Arbey (Orsay), Y.Champetier (Nancy), U. Gehring (Zurich). and A.F. Person (Paris) for their useful comments on the subjeccwe alsothank B. Mahabaleswar (Rennes) for his review of the English in the preliminary version of this paper. The curators of the mineralogical collection of the Mustum National d Histoire Naturelle in Paris and Professor Delaloye (Geneva) provided us with some of the samples studied here; their help is greatly appreciated. Photo processing was realized by 0.Fay.

53

Part B

THE VERDINE FACIES INTRODUCTION TO THE VERDINE FACIES The verdine facies is known since about twenty years from the sea-bottom (Giresse, 1965a; Porrenga, 1965; Von Gaertner and Schellmann, 1965); but it has been specifically distinguished recently (Odin, 1985a) because of its many particular features illustrated in details for the first time in this volume. Among other particularities, the mineralogical composition is the most original. The verdine facies is characterized by a variety of green clay minerals. All of them are different from those present in the ironstone facies as well as from those in the glaucony facies. Two of these clay minerals have been found pure and abundant enough for a detailed study; one is a 7 A clay mineral, the other is an "hybrid smectite-vermiculite-chlorite" clay mineral. All of these clay minerals are iron-rich and iron is dominantly ferric. As a result, the newly described 7 A clay mineral: some sort of "ferric kaolinite" is a new mineral species. Because the habit of the verdine facies is also different from the one of the ironstone facies, the two facies are distinct and their geological significance are different. The following six chapters will emphasize these differences. The verdine facies is presented in some details below and three case studies are proposed illustrating different regional environments and sources of iron. In the lagoon of New Caledonia (Chapter B l), the facies is very young; it characterizes an oceanic island from a tectonically active area. The case study off Senegal (Chapter B2) shows an older facies developed on a stable continental shelf where iron has probably two different sources: the river input to the north, the volcanic rocks to the south of Cap Vert. The third case study off South America (French Guiana, Chapter B3) appears to be a relict Holocene facies on a continental shelf fed with fluvial water from the largest river in the world, Amazon River. Follows a review of all known deposits which can be related to that facies (Chapter B4); most of them were investigated or reinvestigated by the present editor and many unpublished observations are gathered. Follows a detailed mineralogical study (Chapter B5) where numerous collaborators have tried to gather a first data base utilizing a variety of techniques of investigation in order to characterize for the first time the very particular clay minerals discovered in this facies. Finally, a synthetical chapter on the environment and geological significance of the verdine facies is proposed (Chapter B6). By comparison, this synthesis on the verdine facies allows the palaeosedimentological environment of formation of the ancient ironstone facies to be better understood. The hope of the editor is to invite sedimentologists to pay attention to this specific facies with its specific green marine clay minerals.

54

CONTENTS of PART B

* Introduction to the verdine facies

* Contents of part B * B1 The verdine facies from the lagoon off New Caledonia by G.S. Odin 1. Geological setting 2. Sampling and sample treatment 3. Distribution and habit of the green pigment 3.1. Abundance and distribution 3.2. Habit of the green pigment 4. Mineralogy of the green pigment 4.1. X-ray diffraction study 4.2. Acid treatments of green pigment 4.3. Chemical study 5. Mineralogical nature and geological significance 5.1. Ease of alteration of the green pigment 5.2. Mineralogical nature of the green pigment 5.3. Environment of genesis of the phyllite V 5.4. Geological sigmficance of phyllite V 6. Conclusions

53 54

57 59 62 62 63 66 66 69 73 75 75 77 77 79 80

* B2 The verdine facies from the Senegalese continental shelf by G.S. Odin and J.P. Masse 1. Introduction 2. Sedimentology 2.1. The southern shelf (Baie de Rufisque) 2.2. The northern shelf 3. Study of the green pigment 3.1. Separation, proportion, and habit 3.2. Mineralogy of the green pigments 4. Discussion and conclusions 4.1. Sedimentology of the verdine facies off Senegal 4.2. Mineralogy of the verdine facies off Senegal 4.3. Microenvironment and genesis of the phyllite V mineral 4.4. Conclusions

83 85 85 86 89 89 93 100 100 101 102 103

* B3 The verdine facies off French Guiana by G.S. Odin, I.D.R. Mackinnon and M. Pujos 1. Introduction 2. Geological setting of the shelf off French Guiana 3. Authigenic green grains 3.1. The glaucony facies 3.2. The verdine facies

105 106 109 109 110

55 4. Mineralogy of the verdine facies 4.1. X-ray diffraction study 4.2. Analytical electron microscopy 4.3. Chemical study 5. Geological significance of the verdine facies 5.1. Age of the neofonnation 5.2. Recent history of the shelf 6. Summary and conclusions

112 112 115 123 124 124 125 129

* B4 The verdine facies deposits identified in 1988 by G.S. Odin, J.P. Debenay and J.P. Masse 1. Introduction 2. The Ogooue River mouth (Gabon) 3. The Orinoco River mouth and eastern extension (nortern South America) 4. The Niger delta (Nigeria) 5. The Konkoure River mouth (Guinea) 6. The verdine off Sarawak (N. Borneo, Malaysia) 7. The Congo River mouth 8. The shelf off Ivory Coast (Comoe River and other rivers) 9. The verdine facies from Mayotte (Comoro Islands) 10 The Casamance estuary (Senegal) 11 Conclusions

131 131 134 136 138 142 145 149 152 153 156

* B5 Mineralogy of the verdine facies by G.S. Odin, S.W. Bailey, M. Amouric, F. Frohlich and G. Waychunas 1. Introduction 159 2. Mineralogical data on the grains of verdine by G.S. Odin 159 3. X-ray diffraction study of the verdine facies by G.S. Odin 162 3.1. The green clay with a dominant peak near 7.2 A: phyllite V 162 169 3.2. The green clay with a dominant peak near 14 A: phyllite C 4. High resolution transmission electron microscopy on verdine 171 by M. Amouric and G.S. Odin 4.1. Verdine from New Caledonia 172 4.2. Verdine from Guinea 172 174 4.3. Verdine from Senegal (phyllite V) 176 4.4. Verdine from Senegal (phyllite C) 4.5. Conclusions on the TEM studies 177 5. Chemical study of the minerals from the verdine facies by G.S. Odin 178 178 5.1. Chemical composition of phyllite V 182 5.2. Chemical composition of phyllite C 6. Physico-chemical study of the verdine minerals 183 by G.S. Odin, S.W. Bailey, F. Frohlich and G.A. Waychunas 183 6.1. Thermal study of phyllite V

56 6.2. Infra-red study of phyllite V 6.3. Mossbauer study of phyllite V 7. The clay minerals from the verdine facies by S.W. Bailey and G.S.Odin 7.1. The diffraction peak near 14 8, for phyllite V 7.2. Mineralogical interpretation of phyllite V 7.3. Mineralogical variation of phyllite V as a function of time 7.4. Mineralogical interpretation of phyllite C 7.5. Comparison between phyllite V and chlorites 8. Conclusions

* B6 Geological significance of the verdine facies 1. 2. 3. 4. 5.

by G.S. Odin and B. Sen Gupta Broad features of the environment Local settings Substrates and microenvironment Recent verdine facies versus ancient ironstone facies Synopsis

185 188 189 189 191 198 200 20 1 202 205 205 205 212 215 217

57

Chapter B1 THE VERDINE FACIES FROM THE LAGOON OFF NEW CALEDONIA by G.S. Odin GEOLOGICAL SETTING

New Caledonia is an elongate island, 400 km long and less than 80 km wide. It is located about 1500 km East of Australia between 20" and 22"30' latitude South. The climate is equatorial. The island shows an asymmetric relief with the highest altitude reaching 1600 m on the NE border. The SW border is flatter and mangrove-like vegetation develops at sea-level (Balker, 1969). Bays have been created on this SW coast following a Recent subsidence (Debenay, 1985). Rainfalls show a gradient from less than 1 m on the SW coast to 3 m on the NE coast with as much as 10 m of water on the summits (Moniod, 1966). The island is surrounded by a lagoon 10 to 65 km wide with a barrier reef toward the ocean (Fig. 1). The barrier reef is cut by channels sometimes more than 80 m deep. These channels are submazinc palaeo-canyons and correspond to the ancient course of the coastal rivers 15,000 to 20,000 years ago when sea-level was located 120 m below the present (Launay, 1972).

Figure 1. Location of New Caledonia and Loyalty Islands in the SW Pacific Ocean. A barrier reef delimits a lagoon around the main Island.

58

The mean depth in the southwestern lagoon is 20 to 25 m but reaches 40 m to the south and on the northeastern coast. A deeper basin with a bottom at about -80 m, exists to the south in the widest part of the lagoon (see Fig. 4). The topographical relations between the different morphological structures of the lagoon (Fig. 2) were proposed by Coudray (1976) and published again by Debenay (1985). The tide is about 1.5 m high and partly renews the lagoonal water twice a day. When the tide goes in, the sea-water mainly enters through the channels located at the SE end of the island to create a general stream from SE to N W in the SW lagoon (Jarrige et al., 1975; Rougerie, 1986). The current is rapid at the surface and slower at the bottom (5 c d s ) ; it is very rapid in the channels (25 to 80 c d s ) and prevents sediment deposition locally. When the tide goes out, the currents are reversed (see m o w s in Fig. 3). The mean residence time of the waters in the lagoon is about ten days. The temperature of the water is always above 20°C with an annual mean of 23.5'C (Launay, 1972), and with monthly fluctuations lower than 2.5"C around that mean. The salinity is constant at about 35.5%0. Very slightly lower than the sea-water salinity, it allows coral to grow up to the river mouths in the bays in spite of the fresh-water input there (Launay, 1972). The numerous coastal rivers on the southwestern coast have a very variable flow of torrential nature. However, the charge is very low and usually not measurable. The mangrove, swamps, and other vegetation trap the already low

Figure 2. Schematic section in the lagoon of New Caledonia. (Partly after Coudray, 1976) The known green clay-bearing sediments are located in the deepest portion of the lagoon not far from a river mouth opening into a bay.

59 detrital output (Baltzer, 1969; Dugas, 1974). Exceptional demtal outputs occur during and shortly after the cyclones; this leads to the deposition of clay or sandy clay on a restricted submarine delta (Baltzer and Trescases, 1971). For these authors, the continental detritus feeding the rivers is essentially due to the chemical alteration of mainly ultra-basic rocks. The ions reaching the lagoon are mainly silica and magnesium (10 to 20 m a ) ; iron is also present (0.1 mg/l) in the rivers. Sedimentological studies were published for areas immediately linked with river mouths (Launay, 1972; Dugas, 1974) as well as for the SW lagoon (Dugas and Debenay, 1978; 1980; 1981; 1982; and Debenay, 1985) or the southernmost lagoon (Chevillon, 1986). The sediments in the lagoon itself are biodetrital: sands and muds with molluscs, foraminifers and, locally, reefal debris. However, the present sedimentation rate appears low in spite of the abundant bioclast-building fauna (Debenay, 1985). According to the few cores available, we may consider that during the last 7 to 8,000 years about 1-2 m of these sediments were deposited (Launay, 1972; Dugas, 1974). The interstitial water of these cores, collected from the lagoon, shows pH between 7.5 and 8.5 while the Eh is always positive and located around +200 mV, which indicates a relatively open environment, basic and oxidizing. However, when collected, the shelly green pigment-bearing sediments are light-grey in colour; after a few months of storage, the external part of the sediment becomes yellow-ochre, indicating that the sediment was not initially in very oxidizing conditions. Finally, although permeable, the sediment is comparatively compact and allows only slow water circulation between the deposited bioclasts, which show no trace of rolling or reworking. In summary, the green pigment-bearing sediments are from a typical tropical marine lagoon in which the general environment is that of a normal, slightly oxidizing and basic sea-water with a comparatively hot temperature and a complementary continental ionic input. The sediment is formed of nearly 100% biogenic carbonate, and the accumulation rate is moderate. Above the sediments, the water is constantly renewed by alternating currents. But these currents are not strong enough to dissociate the already deposited sediment and to roll the bioclasts. SAMPLING AND SAMPLE TREATMENT

Sampling Since 1984, about fifty sediments were analysed. They were collected either by dredging during sedimentological research undertaken by members of the "Institut Francais de Recherche Scientifique pour le ddveloppement en Coo@ration" (O.R.S.T.O.M.): J.P. Debenay, B. Richer de Forges, C. Chevillon, P. Rigolot, or by direct sampling by scientific divers (1984 to 1986): B. Gout, P. Rigolot, and B. Thomassin. The last sampling using dredging realized in 1986 in the Lagon Nord basin, as well as the last two collections of samples by diving in the Lagon Nord and Lagon Sud basins were undertaken for obtaining

60 VEROINE

.

c SlOOOpprn 1000b>100

0 A

100l-blO < 10

o absent

3coast C=Y reef -canyon curren 3other tide curreni

Figure 3. The south west lagoon of New Caledonia. The green clay was searched for in about fifty samples located on the map. The proportion of green clay extracted is given; two areas appear comparativelyrich facing "Baie de la D u m W and "Baie du hony".

large quantities of material for chemical analyses, distribution to interested colleagues, and preparation of a reference material. In summary, the samples were obtained from five areas in the SW lagoon: 1) about 20 km N W of "Baie de Saint Vincent"; 2) lagoon off "Baie de Saint Vincent"; 3) lagoon off "Baie de la DumbCa", west of Noumea (Lagon Nord); 4) lagoon east of Noumea (Lagon Sud); 5 ) the area SE of "Baie du Prony" where the lagoon is very wide and widely open to the Ocean (Fig. 3).

Sample treatment The green pigment in the present study is nearly invisible even for people studying the sediment for bio-sedimentary purposes. J.P. Debenay (personal communication) indicated that he saw something green in 1977 when mapping the sedimentary cover of the lagoon. Later, C. Froget and J.P. Masse indicated that a green pigment was probably present in the lagoon. The usual technique of concentration of green grains was applied by the present author to a first series of samples for obtaining enough material for X-ray diffraction analysis. In this

61

GRAIN

SIZE

MAGNETIC FRACTION

n25 L

IF53

3%

-

->soopm 500>>50

- esoprn

--

acid insoluble portion of the magnetic fraction

GS0.M

Figure 4. Granulometric composition and green clay content for samples of the two areas where green clay is comparatively abundant (sample location in Fig. 5 and 6). All of the sizefractions are mostly composed of carbonate. The "magnetic"grains are mixture of composite nonmagnetic carbonate with magnetic green clay inside, and of demtal magnetic pyroclastic minerals.

aim, the sediments were washed on a 50 pm or 64 pm sieve and then split into restricted granulomeuic fractions, usually with three sieves: 500 pm, 200 pm and 100 pm. The three size-fractions: 500-200 pm, 200-100 pm and 100-50 pm are then separately treated with a magnetic separator at high intensity with a lateral slope of 15' and then at low intensity to remove the very magnetic fraction. Following this treatment, a small portion of grey to black bioclasts was obtained from a certain number of fractions. These bioclasts mainly show a diffractopm of magnesian calcite

62

Figure 5. Details of the southeastern part of the Neo-Caledonian lagoon. Circled asterisks indicate sediments with more than 17- of green clay; simple asterisks indicate an absence of green clay; intermediateproportions were found in two sediments (squares). Note that verdine is present in the deepest part of the area immediately facing the fluvial output feeding "Baie du Prony".

of no interest for our purpose. The first magnetic fractions obtained were then treated with acetic or hydrochloric acids (about 1 mole/litre). Only at that time does the green pigment become visible in a very small number of samples. The later study showed that, for preservation of the green mineral, only N/10 acetic acid could be used at ambient temperature. After acid leaching, a separation with bromoform followed by magnetic fractionation was necessary to obtain better concentrates,which were nearly pure in the most favourablecases. DISTRIBUTION AND HABIT OF THE GREEN PIGMENT

Abundance and distribution The green pigment is never abundant in the investigated sediments. The magnetically attracted fraction rarely reaches 2% of the whole sediment although in some size-fractions (500-200 pm or 200-100 pm) up to 10% is attracted. This magnetic fraction is mainly formed of carbonate bioclasts grey to grey-black in colour (Fig. 4).The abundance of green pigment can only be estimated after acid leaching, which eliminates 60 to 90% of that magnetically selected fraction. It is interesting to note that, when looking at the residue of the

63 acid leaching, before washing of the solution, one may observe a ghost of the initial carbonate grain composed of a transparent jelly. This is possibly due to remnants of organic material. Finally, the abundance of the green pigment rarely exceeded 1%0and never reached 1% in the sediments studied in the lagoon off New Caledonia. The distribution summarized in Figure 3 shows that we know two main zones where the green pigment is abundant: the Lagon Nord, to the west of Noumea and the area to the SE of Baie du Prony. The green pigment was also found sporadically in one sediment from the Lagon Sud (E of Noumea) as well as in three localities linked with the vicinity of a channel crossing the barrier reef (canyons shown in Fig. 3). It is of interest to look more carefully at the areas where the green pigment is the most abundant. Figure 5 shows the area facing Baie du Prony. Fourteen asterisks indicate the locations where dredging samples were obtained by B. Richer de Forges. Two of them on the northeastern side of the lagoon are located in channels (see Fig. 3) and correspond to hard rocks. From the other localities, only those located in the centre of the figure (circled asterisks) gave an abundant green pigment. All are located at about -80 m of depth and form a restricted zone of about 20 km2 in area. On the one hand, this zone is in the middle of a basin widely open to the SE and partly open to the NE through the two channels quoted above. On the other hand, this basin is fed with an ionic supply by rivers reaching Baie du Prony. Moreover, this basin is also fed with clays from the continent (C. Chevillon, personal communication). Figure 6 shows the distribution of the samples studied in Lagon Nord (their origin is indicated -abbreviation of the collector- because similar numbers were used by the different collectors). Dredgings were done by B. Richer de Forges (RF) and the vessel Vauban in 1986 (V); other samples were collected by diving by B. Thomassin (NC), B. Gout (B.) and B. Gout and P. Rigolot (PR). The sediments richest in green pigment are underlined; except for RF 277 (depth -33 m) they are located slightly deeper than -21 m near the bottom of a basin which is not deeper than about -25 m. This basin is open to the south through the barrier reef on the course of the canyon originating from the rivers feeding the Baie de la DumbCa and cutting the lagoon in a general N/S way. Sample RF 277 is also on this course; it seems that we have there a good example of the two kinds of environment where the green pigment mainly develops: 1) basins fed by rivers in the lagoon and 2) in the vicinity of canyons. The green pigment is less abundant in the shallower samples: (B5, B6) and is also present in some of the other samples (Fig. 3).

Habit of the green pigment As indicated above, there are no grains made of entirely free green pigment in the sediment; this explains why it is difficult to see indications of the presence of the pigment before magnetic separation and acid leaching. After these treatments, the green pigment presents itself in two main forms. The first one consists of particles, green in colour, spongy and easily crushed when one

64

Figure 6. Location and contents of verdine in sediments collected from the lagoon near Noumea. Five different figures are used to designate samples collected by Richer de Forges (RF); B. Thomassin, in 1983 (NC); B. Gout, in 1984 (B); B. Gout and P. Rigolot (PR)and the vessel Vauban, in 1986 (V). Sediments with more than loA of green clay are underlined. Verdine is mostly abundant in the deepest part of the area immediately facing the fluvial output feeding the Baie de la D u m k .

tries to hand-pick them with a needle or with pliers. There is no recognizable form or structure. After acid leaching one observes that these green elements come from initially grey to white bioclasts. There is a suite of intermediate cases between this kind of habit and the second one which consists of remarkably detailed infillings of tests of foramini-

65 fers, ostracodes, fragments of bryozoans or of echinoderms, and microtests of gastropodes. These remnants are usually grey in colour, but sometimes, they are still white. In other cases, they are yellow to reddish. In the area west of Noumea, 40 to 75% of the green pigment extracted consisted of bioclast replacements; but in the area SE of Baie du Prony infillings are dominant. In the two areas, a very restricted number of the green grains extracted were replacements of ellipsoidal faecal pellets. Finally, off Baie du Prony a number of mica flakes have a green colour and possibly host the genesis of the green pigment. The pigment usually has a silky-green colour, especially for carbonate replacements. However, a low proportion of the green infillings, especially for miliolids, has a dark-green colour, nearly black. These infillings are statistically more magnetic than the lightest ones. In the area west of Noumea (Fig. 6) it was observed that the samples at depth around 15-18 m (B5-B6) gave a pigment with a series of colours from green to ochreous-yellow and ochreous. It seems there that, after the green pigment was formed, the oxidizing character of the water altered the green mineral. Another related observation concerns the sediments off Baie du Prony. Among the numerous foraminiferal tests extracted from these sediments using magnetic separation, a low proportion is made of miliolids. Some of them have a test, white in colour, with thin black bands between the chambers; other tests are ochreous-yellow to ochreous-red although their surface remains bright and porcelanous. These tests of miliolids show the only trace of oxidation observed in these comparatively deep lying green pigment-bearing sediments. What is of interest for us is that only a small proportion (about 1/10) of the black and white tests contain the green pigment, but all those with an ochreous colour which were peeled of with pliers showed a dark-green infilling. From the observations made above, we can deduce that the interior of the tests of miliolids, a foraminifer living at low depth, constitutes a very favourable microenvironment for genesis of the green pigment. We suggest a multi-step evolution as follows: 1) death of the miliolid; 2) the test traps iron and becomes grey to grey-black (the X-ray diffraction of acid dissolved grey tests shows that the dark colour is due to a mixture of poorly crystallized pyrite and goethite); 3) oxidation of the dark pigment which is transformed to iron oxi-hydroxides of red colour; 4) genesis of the green pigment favoured by the presence of these poorly crystallized iron-rich minerals; and 5 ) transportation of the low depth tests to the present depth or subsidence of the sea-floor. In summary, the green pigment is not very abundant in the Neo-Caledonian lagoon; two distinct restricted areas showed the pigment in all sediments collected and especially in the 500 to 100 pm size fractions, but it may well be present sporadically in other localities still to be determined. A minimum depth of about -20 m seems necessary for preservation of the green pigment although it has been found in sediments at about 15 m of depth. According to its habit, the green pigment is certainly authigenic because it is formed inside microcavities of bioclasts lying at the bottom of the lagoon.

66 MINERALOGY OF THE GREEN PIGMENT

X-ray diffraction study The green pigment was obtained in about twenty fractions of the fifty two samples selected for this research and for which one to four size fractions were available. X-ray diffraction was applied systematically to all of the acceptedly green magnetic fractions obtained. Randomly oriented powders were prepared. All diffractograms obtained have in common a diffraction peak near 7.2 d which dominates all other peaks. The better shaped diffraction peaks were obtained from the most magnetic fractions of the richest sediments. These fractions were dark-green in colour and gave grey-green, densely coloured powders like sample B5 (Fig. 7). The base of the peak near 7.2 A usually tails off to lower 28 angles sometimes up to 10 A. A smaller peak at about 3.6 A is correlated with the peak near 7 A. Furthermore, a peak at about 4.6 A and two peaks, usually poorly defined, exist near 2.65 8, and 2.5 A. These peaks have a relatively constant intensity ratio between themselves, and indicate that the green pigment is a clay. 5 lJ

p”

i”

T”

loa

148

20x

T T T j

co3

I5 Normal

EG

Figure 7. X-ray diffraction patterns of the green clay from sample B5. Untreated sample (Normal), treated with ethylene glycol (EG), treated with hydrazine (Hz), heated to 490’C for one hour (490°C 1 h). Dotted areas indicate the presence of diffracting material.

67 In addition to these constant diffraction peaks and neglecting the possible traces of quartz or remaining carbonate, a small wide diffraction peak is often present around 10 A (Fig. 7). Finally, a hump between 14 A and 20 A, was the most variable feature of the preliminary series of diagrams obtained from the green pigment extracted using 1N hydrochloric or acetic acid reagents. The treatments with ethylene glycol or hydrazine do not sensibly modify the diagrams although a small proportion of the peak near 7 A seems to shift to the lower angles. Heating to 350°Cclearly lowers the intensity of this peak which nearly disappears after heating for one hour to 490°C;the reflection is totally destroyed in all cases after two hours (Fig. 7).

I

3558, t

I

INHERITED CHLORITE/

+

14

a

T

Figure 8. X-ray diffraction patterns of magnetic green grains of the lagoon of New Caledonia. These grains are made of more or less altered (goethitized) inherited chlorite. Dotted peaks show goethite.

68 In summary, the mineral of the green pigment of New Caledonia is distinct from the glauconitic minerals with which it could have been confused, and shows a main peak at 7.2 A. It is a clay, poorly crystallized, and easily destroyed by heating. It is now time to return to two approximate assertions mentioned above: the f i s t is that no free green grains were present and the second is that all of our diagrams showed no sharp diffraction peak at 7 A. There are a few free dark-green grains, and one of the first diagrams obtained from them in 1984, showed a very sharp peak at 7 A. These two observations were made from the same area in the lagoon off Baie de Saint Vincent (see Fig. 3). The four sediments studied in that area contain magnetic dark-green grains floating in bromoform. A number of them are reddish-brown but others are black-green. A comparatively high proportion of detrital minerals (pyroxene) is present in the magnetic fraction sinking in bromofoxm. However, the free dark-green grains never show remains or structures, nor the form of a bioclast. Several series of these grains were hand-picked and submitted to X-ray diffraction analysis. Figure 8 shows three diagrams that are typical of a chlorite. The peak near 7 A is not precisely at the same place for the different fractions of different samples indicating that different chlorites are probably present. In any case, that peak is always much sharper than the one obtained from the green clay growing in the bioclasts. Moreover, a very high peak at 14 A is always present and always falls abruptly at 14 A on the high 20 angle side; on the other side of the peak, it tails off toward the smallest angles. From diagram 1 to 3 (Fig. 8) there appears to be an alteration of the originally well-crys tallized structure. This alteration is also recognizable from the presence of goethite (dotted peaks in Fig. 8) and the corresponding reddish colour of the most altered of these grains. Therefore, the environment is presently aggressive with regard to these grains of free green pigment. It is suggested that the minerals from which diagrams of Figure 8 were obtained are detrital chlorites for four reasons: 1) the peaks are those from a sheet silicate much too well-crystallized to be of recent sedimentary origin; 2) the sediments from which the dark-green grains were hand-picked are on the submarine course of the rivers feeding Baie de Saint Vincent and contain a noticeable proportion of detrital minerals; 3) none of the dark-green handpicked grains show morphological characters indicating a marine authigenic origin; 4) the environment appears much too oxidizing to allow crystallization of an authigenic ferriferous chlorite that is not protected in some manner against alteration (i.e., inside a test). Finally, one may note that at least two of the four sediments where the free dark-green grains were observed also contain the authigenic green pigment protected inside bioclasts; the latter shows a much lighter colour and a normal X-ray diffraction pattern like in Figure 7. The presence of a variable diffraction peak between 14 A and 20 A in many concentrates of green pigment could be explained by the presence of detrital components that cannot be removed by any physical process. However, the possibility exists that the process of concentration of the green clay and the use of acid treatment could also modify the structure of the authigenic marine mine-

69 ral and thereby affect the diffraction peaks above 14 A. Acid treatments of the green pigment

Concentrated acid treatments In a first series of experiments, hydrochloric and acetic acids were used with concentrations of 1 to 6 moles per litre. Different temperatures from ambient to 75°C were selected. The leachings were undertaken for 1 to 10 hours. A selection of the most significant results is shown in Figures 9 and 10. The green pigment extracted from sample B1 (see Fig. 6) has been used for all experiments. Diagram 1 in Figure 9 gives the reference X-ray diffraction pattern of that sample before acid leaching. As a first approximation, it is considered here to be obtained from a nearly unaltered mineral. One may note the usual main peak at 7.2 8, to which are added, in this case, two broad humps near 10 A and between 14 A and 20 A; the probable background of the diagram is shown by a dashed line in order to identify these domes. At ambient temperature (22°C in September 1986 in Paris) the hydrochloric acid noticeably alters the carbonate-free green pigment; for example, after leaching with a 1N solution for one hour (diagram 5, Fig. 9) the 7 A peak is already very small; when a 6N solution is used, the 7 A peak disappears after about 1 hour (diagram 7, Fig. 9). A series of ex eriments has shown that, when using a 1N HC1 solution, a dome between 14 and 20 A develops during the first hour of treatment and then slowly disappears with the top of the dome shifting to 20 A with time. The same behaviour exists when a 6N HCl solution is used, but the reaction is much quicker than with a more dilute solution. Concurrently, the green pigment loses its colour and becomes creamy, then white. If we use a hot solution (50-75°C) of hydrochloric acid (lN), the reaction is even quicker and the carbonate-free green pigment loses its colour after a few minutes. Acetic acid solutions show a much weaker aggressiveness with regard to the green pigment. However, a 1N solution at ambient temperature diminishes the peak near 7 A after one hour (diagram 2, Fig. 9). At 7 5 T , the alteration is slightly quicker, but we need to wait for more than three hours at 75°Cto reach the stage of alteration obtained after only one hour at ambient temperature with a similar concentration of hydrochloric acid (compare diagrams 4 and 5 in Fig. 9). The acetic acid must therefore be preferred for the dissolution of the carbonate when we want to preserve the clay; this was previously shown for other iron-rich clay minerals like glauconitic ones (Pomerol and Odin, 1974). Concerning the hump between 14 A and 20 A , there is something remarkably clear in Figure 9 (diagrams 1-2-3-4). Moderate in size at first on the "untreated" sample, the hump develops itself progressively (while the 7 A peak is altered concurrently) and reaches an intensity much higher than the initial intensity of the 7 A diffraction peak. It is now presumed that the 14-20 A dome initially present in the reference diagram is mostly due to the sample purification treatment. This is a second possible explanation, in addition to the possible presence of detrital chlorite mentioned above, for that very variable hump in the preliminary series of diffractometric analyses undertaken on the green pigment

81

70

Bj

i"h

A-

Reference

14a

lo^

T --

20i I!

T

76

~

14%

r

208

7 1-

- _-

I

&

6N-lh 16O

120

80

K a 40C u

Figure 9. X-ray diffraction patterns of the green clay from sample B1 following acid leaching treatments; acetic acid: left hand column, hydrochloric acid: right hand column. Different temperatures, different concentrations (normality), and different durations (hours) were tested. When the leaching is more efficient, the peak at 7 A is lowered and disappears, and a dome at (14-20) 8, is created first and destroyed later during a second stage.

from the lagoon of New Caledonia. It was of interest to reproduce that hump under different conditions and then to try to identify the nature of the phyllite "synthesized" by the experiments. Another portion of the same sample of carbonate-free green pigment was extracted from B1; it gives the X-ray diffraction pattern in Figure 10 (diagram 1) similar to that in Figure 9. Leached with a 5 N acetic acid solution for one hour at ambient temperature, the green pigment shows a very slightly lowered 7 A peak, but the dome at (14-20 A) is higher and wider (diagram 2, Fig. 10) as was anticipated. Ethylene glycol treatment of the previous preparation provokes a slight but distinct shift of the dome toward lower angles (diagram 3, Fig. 10). The synthesized phase therefore has a swelling behaviour. However, the diverse heating treatments undertaken on the acid leached pigment obtained

71 have shown that the layers never totally collapse like smectitic layers would. After heating to 350°C,there is a shift toward high angles in such a manner that a nearly continuous spectrum of diffraction exists between 20 8, and 7 8, (diagram 4, Fig. 10). When the tem erature of heating increases, the higher angle end of the spectrum (near 7 ) disappears but the spectrum remains hinged at a high intensity on the 20 8, distance at the other end.

w

148

f

2 0 ~

T T

loi 1 4

T

T T

""c\

120

K a Cu 40

EG SO861F"

I

I

_--'

k 1"

CH3COOH-5N-lh-22"

16O

~Z 0 a

I

K a Cu ,

120

80

40

Figure 10. X-ray diffraction patterns of the green clay showing the behaviour of the dome at (14-20) A. The dome develops after acid leaching using 5N acetic acid (diagram 2); it partly shifts to lower 2 9 angles after ethylene glycol treatment (EG diagram 3) and partly to higher angles after various heating treatments.

The clay formed during acid leaching of the green pigment has a complex and irregular structure. A portion of the layers swell with ethylene glycol another, or the same, irregularly shifts to the smaller distances; another is stable and diffracts between 14 8, and 20 A.

Dilute acid treatments In order to determine satisfactory conditions of preparation of the green pigment before analysis, decinomal solutions of hydrochloric and acetic acids have been used at various temperatures for different durations. Figure 11 shows significant diagrams related to these experiments. The reference diagram (1 in Fig. 11) was still obtained from the most magnetic carbonate-free green

72

pigment of sample B1. Note that a dome between 14 A and 20 A is already present due to the process of purification. Leaching using the N/10 acetic acid for about two days does not modify noticeably the initial diagram if ambient temperature is retained; however, at 60°C, long leaching alters the peak at 7 A and the dome at 14-20 A is intensified (diagram 3, Fig. 11). At a similar temperature, the decinormal hydrochloric acid alters the carbonate-free green pigment after less than one hour and the peak near 7 8, is destroyed after only three hours (diagrams 5-6, Fig. 11). Even when ambient temperature is maintained, a long leaching with decinormal hydrochloric acid alters the peak at 7 A of the carbonate-free green pigment of the lagoon of New Caledonia (diagram 4, Fig. 11). 3 l J

f

f Ta

f

F'

/

I

yJ_/:2*j

Reference

i

/

22"-45h

I

I

Figure 11.X-ray diffraction patterns of the green clay submitted to leaching using decinormal acids: acetic (to the right), hydrochloric (to the left). The reference diagram obtained from sample B1 already shows a dome at (14-20) A, probably due to the preliminary treatment (purification of the green pigment). Diagram 2 obtained after acetic acid leaching at ambient temperature is the only one which shows no modification of the reference diagram.

73

Therefore, hydrochloric acid cannot be used without careful mineralogical control for extraction of the green pigment from the bioclasts. But the acetic acid seems to be usable, without noticeably modifying the mineral of the green pigment, if a low temperature is maintained. The alteration of the green pigment can be seen by the shape of the 7 8, peak (the intensity of this peak decreases when the alteration occurs) as well as by the development of a dome between 14 A and 20 A. Concurrently with the modification of the X-ray diffraction patterns, the colour of the pigment changes from green to yellow and white. This probably explains why, in some of the preliminary extractions undertaken on the samples from the lagoon of New Caledonia, the presence of green infillings observed after magnetic concentration was not confirmed after later acid leaching: the too strong acid leaching destroyed the green pigment.

Chemical study

Figure 12. X-ray diffraction patterns of the green clay separated from sample B3. (1) magnetic size-fraction 500 to 160 pm treated with 1 normal acetic acid; (2) magnetic sizefraction 160-100 pm treated with 1 normal hydrochloric acid; the peak at 7 A is slightly lowered in the latter case. In the absence of obvious alteration and in view of the very low content of impurities, the two size-fractions were combined in order to have enough material for chemical analysis (see Table 1).

Only one sediment: B3 from the Lagon Nord of Noumea (see Fig. 6) gave enough material (=4 g) to obtain a chemical analysis using X-ray fluorescence

74 and wet chemical techniques. In this sample, the green pigment has a comparatively light-green colour like many samples from this basin. The carbonates were dissolved using a 1N acetic acid solution at ambient temperature for the size-fraction 500-200 pm (A) and a 1N hydrochloric acid solution for the sizefraction 200- 100 pm (B) (at the time of preparation of these green pigments, the results of the experiments described above were not yet known). In spite of these conditions, shown later to be aggressive for the green mineral (see paragraph above), the X-ray diffraction patterns were not noticeably disturbed (Fig. 12). This may well be understood if one considers that our experiments of acid leaching were done on a carbonate-free green pigment. The initial presence of carbonate obviously diminishes the aggressiveness of the acid solutions, which are neutralized by the dissolution of that carbonate. The necessary caution is therefore to avoid the replacement of the acid solution as soon as the carbonate remains are mostly dissolved. Table 1. Chemical analysis of the green clay from sample B3. (X-ray fluorescence by Miss M. Lenoble, DCpartement de PCtrographie,Paris) SiO2 83 A

+

B

Herthierine

34.0 24.0

A1203

Fez03

6.13 22.0

22.0 3.5

FeO

RlgO

CaO

K20

Ti02

P2O5

6.95

0.16

0.14

0.3

H2O+CO2

Total

13.7

0.7

15.2

99.28

35.5

3.5

-

11.5

100.00

Chlorite M g

29.7

19.8

0.3

6.6

30.8

-

12.7

99.90

C h l o r i t e Fe

22.2

20.0

7.4

35.2

3.8

-

9.2

97.80

Berthierine: the data are mean values calculated from analyses given by Brindley (1982). Typical magnesian chlorite (Chlorite Mg) and femferous chlorite (Chlorite Fe) according to Foster (1962): samples 38 and 141 of that author.

We performed chemical analysis of the mixture of the two size-fractions above (A+B) which a pear little altered after the acid treatment (no dome between 14 8, and 20 , Fig. 12). The purity of the pigment was improved using bromoform, magnetic fractionation, and hand-picking of the few black and reddish grains still present after the other physical processes of concentration were used. Table 1 shows the results (line B3 A+B). The roportions of the elements measured are compatible with those of a 7 A or 14 clay mineral. The purity is verified by the low proportions of CaO and K,O. Table 1 also gives the proportions of the elements usually found for minerals to which the green clay can be referred: a berthierine (defined according to Brindley, 1982), a magnesian chlorite and a ferroan chlorite, (chlorite Mg and chlorite Fe respectively) according to Foster (1962). One may note the remarkably low content ofalumina in the green clay of the lagoon compared to the other analyses given. However, what appears to be the most extraordinary characteristic of the green clay of the lagoon concerns the

i

1

75

high iron content; this iron is mainly in the ferric f o m . The calculated Fe203 content is higher than the total: FeO+MgO when oxides of the chemical analysis are considered. This gives a very specific formula to the octahedral sheet of that clay since the previously known natural pure iron-rich 7 A and 14 A cla minerals are all dominantlyferrous. Compared to the other known iron-rich 7 or 14 A phyllosilicates, the clay from the Neo-Caledonian lagoon is: - always richer in silica, especially compared to berthierines; - always much poorer in alumina; - always much richer in ferric iron (Fe3+is a dominant ion of the octahedra); - compared to berthierines of Brindley (1982) it is also richer in magnesium. In summary, none of the five major cations of the crystallographic architecture of the Neo-Caledonian green clay has a content compatible with the range of variation of natural berthierines of sedimentary ironstones. It is not a berthierine. One might also consider the Neo-Caledonian clay mineral as a poorly crystallized iron-rich chlorite. In these circumstances, the peak at 14 A might be so small that it could not be seen on the randomly oriented powder diagrams shown here: a very small first order peak is a property of the iron-rich chlorites in which iron is concentrated in the interlayer sheet. However, a problem remains in that we do not know of any dominantly ferric natural chlorite.

x

MINERALOGICAL NATURE AND GEOLOGICAL SIGNIFICANCE

Ease of alteration of the green pigment The experiments above have shown that acetic and hydrochloric acids may alter the diffractograms of the original green pigment formed in the carbonate bioclasts. The alteration is much less visible when some carbonate remains present in the treated sediments (Fig. 12). In this case, acid leaching with 1N solutions appears to be acceptable. However, acetic acid is preferred because the 7 8, peak of the hydrochloric acid treated fraction is slightly diminished. Furthermore, for better preservation, decinormal solutions at ambient temperature have to be used. This solution replaced twice a day has been used for 200 hours in some cases in order to remove the carbonate; the diagrams obtained then showed a normal peak at about 7 A and a small dome of diffraction between 14 A and 20 8, (Fig. 13). The ease of alteration with regard to the acids discussed above has already been quoted for other clay minerals. According to Brindley (1961b) hydrochloric acid (concentration 2N; temperature 1W C ) destroys berthierines, chlorites, and serpentines after 1.5 hours i.e., decomposes all the 7 A and 14 A iron-rich clay minerals. Similarly, Carroll (1966) indicates that a warm 1N solution of hydrochloric acid slowly destroys the well-crystallized chlorites and more quickly the poorly crystallized ones (see also p. 40 above). With regard to the heating treatment, well-c stallized chlorites are not destroyed after one hour to 6OO0C,the peak at 7 disappears but the one at 14 A is increased.

T

76

I

CH3COOH

v2

N/IO

P

9

Figure 13. X-ray diffraction patterns of phyllite V purified with decinormal acetic acid. The carbonate was removed by leaching at ambient temperature for 450 hours (diagram above) and for only 1 1 hours (diagram below). There is no significant difference in the diagrams obtained. This shows that decinormal acetic acid at ambient temperature may be used without major problems during long periods when necessary.

The mineral of the green pigment of New Caledonia is already destroyed in less severe conditions using acids; heating to 490" or 550°C does not noticeably reinforce the diffraction peak at 14 A. In short, the Neo-Caledonian mineral is much more alterable than the other minerals quoted above. Although no reinforcement of a peak at 14 A is visible in the randomly oriented powder diagrams of the Neo-Caledonian clay mineral after heating, it is probably relevant to remember that a dome is created (or markedly enhanced) in some cases using acid solutions. The literature does not quote such a behaviour for chlorites, berthierines or serpentines. It would be useful to consider that point specifically for these minerals in the future. In the meantime, because that behaviour has repeatedly been observed in the Neo-Caledonian green clay mineral, it is emphasized that this can be regarded as a characteristic behaviour. This property might help in identifying clay minerals referable to the Neo-Caledonian one. The question is now: what is the modification introduced so easily into the crystalline architecture which causes that (relatively diffuse) periodicity? In other words: what is the significance of the dome between 14 A and 20 A: a neoformed mineral synthesized or simply a rearrangement of an octahedral sheet, and what kind of rearrangement? This question cannot be answered until the mineral structure that produces the band of diffraction is precisely known.

77 However, the band of diffraction enhanced by the acid leaching is difficult to interpret. It is not simply centred on a 14 A distance but clearly indicates a periodicity higher than 14 A. It could also be interpreted as a double band with tops at 14 A and 18 A. According to the behaviours of this band following ethylene glycol and heating treatments, one may speculate on the existence of a chlorite-type structure for some of the layers in which incomplete brucitic interlayers would allow ethylene glycol to enter the structure but where the occupied brucitic sites would prevent collapsing to 10 A with heating treatments. This is what has been termed a swelling-chlorite.

Mineralogical nature of the green pigment The green pigment is therefore made of a clay phase with a dominant diffraction at 7.2 A. This phase is green and can be concentrated by magnetic separation; it is therefore iron-rich. It is very susceptible to alteration using acid solutions and is destroyed by heating. All of these criteria are referrable to a berthierine-like clay. However, the creation of a broad spacing above 14 A with some acid solutions (although not by heating) could also lead to the presumption that this clay is partly made of a chlorite-like phase. If one looks in details at the X-ray diagrams obtained from the green pigments separated using very dilute acids (and consequently considered to have been little modified) one suspects that a very small dome is already present. The diagrams in Figure 13 are representative of a number of others obtained from similarly obtained materials: most of time, the 7.2 A peak is preceded by larger d-value eaks which can tentatively be interpreted as a weak double peak at (18-20) and (14-15) A. Within this study, the problem is not yet solved as to the real nature (7 A or 14 8, or mixture) of the green marine clay of New Caledonia; more data will be proposed in Chapter B5. However, the chemical analysis is very informative and confirms the specificity of the Neo-Caledonian phyllite. In this structure, the iron is dominantlyferric and this is not yet known from the classical 7 A or 14 8, clay minerals. Moreover, it is very poor in alumina, poorer than any chlorite; finally, even the silica and magnesia contents are incompatible with the green clays presently described from sedimentary ironstones. The green clay mineral from the lagoon off New Caledonia is therefore a presently undescribed clay. As suggested previously from the study of similar material from different deposits of the present sea-bottom, this clay justifies the use of a specific designation (Odin, 1985a). Leaving open a definitive name, the preliminary term phyZlite V , proposed in 1985, will be used in this work (see Glossary).

x

Environment of genesis of the phyllite V At the present stage of this work, it is difficult to define precisely the factors which are necessary to favour the genesis .of phyllite V. However, the Neo-Caledonian deposit is sufficiently well-characterized to give an interesting general view of the macro-environment. Moreover, the mineral is clearly con-

78 temporaneous with conditions of deposition similar to the ones observed today in the lagoon, because the sea-water has only been present for about 7-8,000 years in the lagoon, following the Holocene transgression, and the facies has not been deeply modified during that period. The physico-chemical properties of the lagoonal environment, or some of them, would therefore be favourable for growth of phyllite V. They are: -a marine environment i.e., a normal salinity (=35.5%0) and a pH of 7.5 to 8.5, -a temperature higher than 20°C and rather near 25"C, -an abundance of oxygen and a positive Eh above the sediment, -a minimal depth of about -15 m, especially at the bottom of basins, -an absence of detrital deposition and a low rate of accumulation of bioclasts (less than 10 cm/lOOO years), -a mainly sandy sediment from the granulometric point of view, with a noticeable proportion of particles above 500 pm, -a sea-water circulation (tide currents four times a day), -a continental ionic input with abundant Si, Mg, and Fe appears useful. Phyllite V does not grow in this general environment however. The clay only grows in microcavities of bioclasts. This implies several factors. First of all, the microenvironment, surrounded by carbonate and basic waters, is not acid. Secondly, these microcavities were initially filled with organic material which, possibly, implies a redox potential more reducing than the general environment. The grey to black colour of many bioclasts is a good criterion of these local conditions, as previously quoted by Debenay (1985). Secondly, the grey-black colour linked with the presence of a mixture of pyrite and goethite indicates that, at least temporarily, the environment was reducing during the evolution of the carbonate fragments. Thirdly, these microcavities isolate a particular microenvironment with regard to the general environment described above. The fact that phyllite V only crystallizes in this isolated microenvironment (and not between or at the surface of the bioclasts) allows us to presume that a sort of protection against the general environment is an important factor favouring the mineral genesis. It has been emphasized above that phyllite V was susceptible to quick alteration by oxidation at depths shallower than -20 m. Obviously, phyllite V itself does not form in a very oxidizing environment that would destroy it. A fourth series of observations concerns the grain-size of the sediment where phyllite V grows and the role of the currents. These two factors favour the renewing of ions around the carbonate particles. This, in turn, indicates that in spite of the isolation created by the internal microcavities of the bioclasts, there is a possibility of exchange of cations with the general environment through the filter of the external pores of the bioclasts. In these conditions, the microenvironment where the clay minerals grow is a restricted volume isolated from the aggressive factors of the general environment. The ions present in this volume may combine in order to crystallize a clay. But the reaction would immediately stop if a controlled opening was not available to the exterior allowing a re-equilibration of the ionic concentration useful for favourably feeding the crystallization.

79

In summary, the microenvironment where phyllite V grows is a semiconfined one. It is present in carbonate substrates, especially those which naturally have microcavities. This explains why phyllite V often presents as infillings of microtests or other organic carbonate structures.

Geological significance of phyllite V The geological significance of phyllite V can be deduced from its mineralogical composition as well as from its environment of formation; one therefore discusses a facies. This facies is a sedimentary one; it is linked (in the above example) to a reefal environment; it is developed during the earliest stage of evolution of a deposited sediment; it is characterized by the development of clay minerals near the sea-water/sediment interface. A good opportunity to precise the nature of this facies is to emphasize its specificity with regard to other apparently similar facies. For example, it can easely be confounded with the glaucony facies looking only at the appearance. Obviously, the main difference concerns the mineralogical nature of the green clay. This nature is an efSect (the result of the geochemical reaction); but the effect being specific, one may try to analyse what is particular as to the cause (the elements, the environment, and the speed of the geochemical reaction). Concerning chemical elements, there is little difference between glaucony and phyllite V-bearing facies since silica, alumina, dominantly femc iron, and magnesium are concerned in the two facies. The relative richness in magnesium of the continental water input possibly influences the comparatively higher Mg content in the phyllite V versus that in the glauconitic minerals. Concerning the environment, the comparatively low depth and proximity of the continental water input seem specific of the phyllite V-bearing facies. The water temperature is clearly higher in the investigated lagoon (25°C) compared to the usually quoted temperature of 10 to 15°C for the Recent glaucony-bearing sediments. Finally, the speed of the reaction appears very high for phyllite V growth (1000 years as an order of magnitude) compared to the growth of glauconitic minerals which needs 10 to 100 times more (Odin and Matter, 1981). This probably results from the temperature of reaction. A second facies from which Recent phyllite V-bearing sediments have not been clearly distinguished until recently (Von Gaertner and Schellmann, 1965; Porrenga, 1967; 0din.and Giresse, 1972; Martin, 1973) is the ancient chamositeberthierine-bearing facies of oolitic sedimentary ironstones. In this latter case, the result of the geochemical reactions is either a 7 8, or a 14 8, wellcrystallized mineral or a mixture of both with frequent admixtures of iron oxi-hydroxides. The general architecture of the sedimentary ironstone green clays (forgetting their nice crystallinity) is comparable to that of phyllite V mineral but the chemical composition (elements and oxidation state) is different. In a certain manner, phyllite V mineral and the ironstone clay minerals are more comparable in terms of their crystallographic architecture; but phyllite V and glauconitic minerals are more similar as to the chemical composition (Fe3+, Fe2+, A1 cations).

80

Concerning the environment, the noticeable specificity for the ironstone clay minerals is given by their oolitic habit. This habit has never been observed for the Neo-Caledonian phyllite V mineral. The significance of this oolitic habit is fundamental: it is the general environment which is favourable for the ironstone clay mineral genesis as opposed to the microenvironment localized in the carbonate grains for phyllite V genesis (see Chapter A2). This is confmed by the fact that chamosite and berthierine also form between the ooids; in comparison, phyllite V had never been observed out of the bioclasts. Therefore, the phyllite V-bearing sediments constitute a facies that is different from the one where berthierine and chamosite form as well as from the one where glaucony forms. The environment is different, the mineral resulting from this environment is different; therefore the facies is original. This facies has never been distinguished nor described in details until now, a specific name is therefore needed for its designation and it has been suggested to call it the verdine facies (Odin, 1985a). CONCLUSIONS

The biodetrital sands of the lagoon off New Caledonia shelter the genesis of a green marine clay. This mineral phase can be concentrated thanks to its magnetic properties, green colour, and particular habit; - it forms in microcavities of bioclasts and can be separated using a moderate acetic acid leaching; - the X-ray diffraction study indicates that the mineral phase is a clay with a dominant peak at 7.2 A; it is poorly-crystallized however; - dilute acid leaching causes the formation of a clay-type diffuse peak in the area 14-20 8, but quickly destroys the mineral when more aggressive solutions are used; - chemical analysis indicates that the iron content of this clay is high and that iron is in the trivalent form; - this iron-rich ferric clay is presently undescribed and is called phyllite V in this study waiting for more precise mineralogical results; - the general environment of genesis of phyllite V is marine, lagoonal, and oxidizing, the temperature is comparatively high and the present depth of occurrence can be as small as -15 m; - the microenvironment of genesis is not very oxidizing however, and it is semi-confined with regard to the general environment; - the mineralogical nature and environment of genesis of phyllite V are different from what is known from two comparable facies: glaucony and pre-Quaternary sedimentary oolitic ironstones; - the name verdine is suggested for designation of this newly described sedimentological facies; - the present environmental characters in the lagoon off New Caledonia are similar and probably identical to those favourable for the formation of the verdine facies because the green marine clay studied in this chapter formed less than 7-8,000 years ago and is likely to form today.

81 ACKNOWLEDGEMENTS

This work was undertaken following preliminary observations by C. Froget and J.P. Masse. The geological department of the ORSTOM in Noumea headed by Mr. RCcy was very efficient for sample collecting. Sediments were collected or provided by C. Chevillon, J.P. Debenay, B. Gout, B. Richer de Forges, P. Rigolot, and B. Thomassin. The samples collected using aqualung by B. Gout, P. Rigolot, and B. Thomassin were especially appreciated. Previous manuscripts on this work have been discussed and improved by J.P. Debenay, C. Froget and P. Rigolot. The help of these colleagues is greatly appreciated. Comments on the mineralogy of the Neo-Caledonian phyllite V by Professor S.W. Bailey were very helpful. Thanks are also due to S.W. Bailey who was kind enough to improve the English of the present chapter. Figure 2 was partly prepared by M. Petzold.

82

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83 Chapter B2 THE VERDINE FACIES FROM THE SENEGALESE CONTINENTAL SHELF by G.S. Odin and J.P. Masse INTRODUCTION

General environment The continental shelf of Senegal is located on the eastern border of the North Atlantic Ocean between about 14" and 16"North latitude. The northern limit is the delta of the Senegal River, which is 200 km north of Cap Vert. The latter, where the city of Dakar is situated, is a tongue-shaped cape which divides the study region into two areas (Fig. 1). The coast is generally low and flat; but the Tertiary sedimentary rocks and Late Cenozoic volcanic extrusions of the cape have been eroded by the sea to form cliffs. The continental shelf, less than 10 km wide at the latitude of Cap Vert, widens to more than 50 km northward and southward.

Figure 1. Location of the study area off Senegal.

The Senegal River today brings a limited quantity of detrital material to the sea. This is transported southward by longshore currects generated by the oblique oceanic swell moving towards the SE. Thus the mouth of the river has shifted southwards, flanked by an active spit and offshore bar system.

84 Although no other permanent river feeds the studied platform, a coastal water influx is observed during the rainy season. The mean annual rainfall is 500 to 800 mm and the continental climate is tropical. Locally, mangroves with Rhizophoru develop on the estuaries, margins of the seasonal rivers, and in corresponding coastal swamps. The sea-water temperature varies according to season and the corresponding oceanic current patterns. Temperature and salinity characteristics are as follows: 28°C and 35.7%0 for the two warm surface streams which are active during the humid season (monsoon); mean 18°C and 35.5%0 during the dry season (trade winds) when upwelling is active. Coral-reefs are absent on the inner shelf.

Sampling The sedimentology of the southern area was deduced from the collection of about 120 superficial sediments for an area of 60 km x 40 km. The detailed study of the green pigment was undertaken on three series of samples located radially from the coast to about 200 m depth (Fig. 2).

Figure 2. Location of the samples collected from the area to the south of Cap Vert (Baie de Rufisque). 1) individual samples; 2) continuous series of samples, some with their numbers above; the other numbers may be interpolated; 3) percentage of components of the sediment showing traces of green pigment. (Counts after Masse, 1967)

85 The sedimentology of the northern area was established from about 140 superficial sediment samples and a series of cores which were collected or drilled on the shelf in the years 1970 to 1975 by several French organizations including the ORSTOM, the Bureau de Recherche GCologique et Minikre (B.R.G.M.) and the Institut de GCologie du Bassin d’Aquitaine (I.G.B.A., University of Bordeaux). A small number of samples were provided to the present editor by Y. Monteillet (Dakar) and J. Pinson (Dakar and Bordeaux) as well as by the I.G.B.A. for the study of the green material. SEDIMENTOLOGY

The southern shelf (Baie de Rufisque)

Main sedimentological characteristics The sedimentological data presented in figure 2 have been published by Masse, (1968; 1970). According to these works, five sedimentological zones roughly parallel to the coast may be identified. 1) Between 0 m and about -35 my pre-littoral siliciclastic sands with molluscs (4 in Fig.2) and/or barnacle fragments (5 in Fig. 2) of Recent age are found; the quartz content decreases rapidly with depth. 2) Between about -35 m and -55 m, reddish brown bioclastic sands (6 in Fig. 2) with bryozoans indicate Recent oxidation. A small submarine cliff is frequently present at a depth of about -50 my and is probably the result of a palaeo-strand line. 3) Between -55 m and -90 to -100 m depth, are f i e greenish sands with in situ mollusc debris, abundant foraminifers and quartz (7in Fig. 2). Foraminifers are more abundant than in the shallowest zones and include an increasing variety of benthic genera: Elphidium, Cibicides, Discorbis, and Nonion as well as pelagic groups: Globigerinids and Globorotalids. However, the proportion of the more coastal Miliolids decreases. In this zone, traces of oxidation occur at about -80 m depth; the first specimens of Amphistegina gibbosa, (a relict Quaternary foraminifer linked to a low sea-level phase) is recorded with some corals and faecal pellets in the deeper parts. 4) Between -100 m and -180 m depth, biogenic coarse to medium green sands are present (8 in Fig. 2). They contain dominant mollusc debris (40%), abundant bryozoan fragments (20%): Porella, Retepora, Membranipora, calcareous algae, foraminifers including the relatively abundant relict A. gibbosa while the large foraminifer: Schizammina, is locally abundant, associated with some quartz grains. This material is mainly relict (thanatocenosic) and belongs to the “Amphisteginafauna” (sensu Lagaaij, 1973) which has been dated from 11,000 years to 13,000 years B.P. (Barusseau et al., 1987). 5) In the deeper part of the shelf (9 in Fig. 2) the fine biogenic sands are characterized by the foraminifera1 genus Cyclammina. This facies, of bathyal character, is dominated by mollusc shell fragments, foraminifers, bryozoans (more than lo%), pieces of coral, very corroded pieces of calcareous algae, and a little quartz. As above in 4 relict material of Quaternary age predominates.

86 The quartz is less abundant than in the above zone. To the south of Cap Vert the fine fraction (less than 50 pm) is almost absent. When present (with a maximum proportion of 3% in our samples), it is mostly composed of quartz silt and calcilutite. On an X-ray diagram, a low 7 8, reflection and an irregular low hump between 10 A and 15 A exists.

Green pigment Special attention was paid to the green pigment during the study of the sediments of the Baie de Rufisque (the southern area of our platform). Green or partly green grains were observed in the main part of the sediments collected from -8 m down to -200 m depth. The green material is mainly abundant in the fine sandy fraction (between 200 pm and 100 pm). In sediments deeper than the reddish brown zone, the high proportion of green material gives a generally green colour to the sediment. Counting of the bioclastic components of the sediment was undertaken on a selection of samples (Masse, 1968). A specific class was drawn for elements showing the presence of green pigment. These counts c o n f i i that the green material is an ubiquitous component of the sediments. In the zone between -100 m and -180 m depth, brilliant black biogenic debris are associated with the partly green grains; the nature of the mineralization giving that black colour has not been identified here. It was noticed by Masse (1968) that the green material was only exceptionnaly under the form of free and fully green grains. Rather, the green pigment usually a) diffusely impregnates carbonate bioclasts, b) fills the perforations of bioclasts (perforations of about 10 pm in diameter) made by boring algae, sponges, and polychaetes; small borings may also originate from fungi and Cyanophyta, c) fills the chambers of foraminifers, bryozoans, ostracodes or microgastropodes. When they occur, free grains are very small and correspond to the internal moulds of one of the biogenic structures described above. Other grains which can be presumed to be entirely formed with the green pigment are ellipsoidal faecal pellets, 100 pm to 200 pm in length; however, they usually still contain recognizable carbonate silty debris or minute quartz particles. The northern shelf

Main sedimentological churacteristics The continental shelf to the north of Cap Vert is cut by Cayar canyon; the sedimentary cover has been studied by Pinson (1980) from whom are taken most of the following data. From a granulometric point of view, the surface sediments may be divided into sands, muddy sands (for which more than 30% of the sediment are larger than 63 pm) and muds (Fig. 3). The sands and muddy sands are located deeper than -75 m to -80 m as well as shallower than -30 m to 50 m. Pure sands are restricted to the North of Cayar Canyon, along the coast as well as deeper than -90 m. These sands contain quartz grains, ochre in colour, and similar to those found today in the continental red dunes. The muddy sands are widespread to

87

0

1

0

B

2

6 3 2

B

3

r

ZOtm, GSO ffi

.

1

3

3

ZOLm, GSO 86

Figure 3 (left). Sedimentology of the area to the north of Cap Vert. Note the presence of Cayar Canyon; many other canyons are shown on the -200 m isobath. 1) sands; 2) muddy sands; 3) mud. Note that the mud seems to originate from the north. (Distribution according to Pinson, 1980) Figure 4 (right). Nature of the macrofaunal fragments on the area north of Cap Vert. 1) White, Recent in age; 2) Beige to yellow ochre, relict and probably partly oxidized; 3) Grey, relict and sub-reefal facies. (Distribution according to Pinson 1980)

the south of Cayar Canyon; they are restricted to the margins of the pure sands. Elsewhere, mud is present. Cores show that this mud overlies the previous sandy facies. The clay fraction of this mud shows a constant composition: mixed-layer (illite-smectite) + smectite (40%), kaolinite (30%), illite, and chlorite (about 15% each). This clay seems to originate from the Senegal River to the north. Deeper, clayey sediments are located on the slope (deeper than -200 m). The latter clay shows 50% of kaolinite, 30% of smectite and 10% of both illite and chlorite, respectively. A series of cores was drilled along the platform from St Louis on the Senegal River to Cayar at the head of the Canyon (Fig. 3). From these cores and complementary seismic studies, it has been established that the thickness of the soft sedimentary cover (mud and sand) decreases from 40 m to 10 m southward, clearly indicating the northern origin of the sediments the deposi-

88

0

Figure 5. Presence of faecal pellets on the continental shelf north of Cap Vert. 1) Samples associated with the relict thanatocaenosis, with grey remains showing green pigment filling ostracode tests between others; 2) less than 10% of faecal pellets in the sediment; 3) more than 10%.(Distribution according to Pinson, 1980)

tion of which can be considered younger than 20,000 years (Pinson, 1980). In many samples, at depth of -100 m to -1 10 m, the sedimentary components of different sizes are covered by three or four aragonite layers indicative of carbonate precipitation. Study of the macrofauna allowed Pinson (1980) to distinguish three major zones (biofacies) in the surface sediments. A first zone is shallower than -75 m, the macrofauna is restricted to the southern part of an area where the sediment is not too muddy (Fig. 4). Here, the fauna is Recent in age and white coloured. A second narrow zone is discontinuous with a relict macrofauna, yellow to beige in colour, and located between - 100 m and -75 m depth. Ooids are often associated with this facies interpreted as resulting from a standstill of sea-level. A third large zone between -100 m and -200 m depth (at the top of the slope) is

89 characterized by a thanatoccenosis (indicating a comparatively hot temperature) with an abundant macrofauna, grey in colour (Fig. 4). At a depth of about -1 10 m, more than 75% of the sediment is composed of grey faunal remains, usually very fragmented. The fauna includes corals, bryozoans, molluscs, and benthic foraminifers, all marine and fossil. Locally, reef like structures (corals or bryozoans) were postulated by Pinson (1980). This interpretation agrees with the suspected reefal structures ('pitons assimilables ii d'Cventuelles structures rkcifales') quoted by Ruffman et al. (1977) according to seismic records. However, the reefal nature of these structures and sediments is doubtful (Masse, unpublished).

Green pigment The green pigment in the area north of Cap Vert shows two aspects of special interest: green ellipsoidal faecal pellets and infillings of ostracode tests. The faecal pellets, frequently dark-green in colour, are widespread and sometimes abundant (Fig. 5) especially on the median to deep part of the platform. This distribution is not strictly related to the boundaries between the sedimentary zones defined above. Near Cayar Canyon, there are dark-green to black glossy ellipsoidal faecal pellets with an admixture of cracked green grains the initial substrate of which is unknown. Infillings of ostracodes appear more depth dependent; they are clearly related to the outer shelf thanatocaenosis. STUDY OF THE GREEN PIGMENT

Separation, proportion, and habit

Separation and proportion of the green pigment The green pigment was magnetically separated from the water washed sizefractions: 500 pm to 200 pm and 200 pn to 100 pm. The coarser size-fraction concentrated the faecal pellets, and the finer-size fraction, the infillings. The magnetic fraction was observed under the binocular microscope, the main substrate identified and then, the carbonate removed using a half normal acid solution (acetic for a best preservation of the green material). The obtained material was ultrasonically cleaned washed and sieved. The lower size-fractions were usually enriched after those treatments due to the dissociation of the foraminifer chamber infillings. A magnetic purification was done after drying. Figure 6 shows the very high proportion of nearly pure green pigment (without carbonate) extracted from the sediments from the south of Cap Vert. The proportions indicated in Figure 6 are lower than in Figure 2 due to the removal of carbonate. However, in two cases, it is higher. This may be due to the fact that, before acid leaching, it is not easy to recognize whether the chambers of a foraminifer are fiiled wih green pigment or not.

Aspects of the green pigment I n f i l l i n g s of microtests are the most ubiquitous substrates, e.g., foraminifers, ostracodes, microgastropodes, and scaphopodes (Fig. 7AB).

90

-

I

10 km

-

1

-

2

a

3

I

T N%

Figure 6. Nature and percentage of the green pigment identified after acid leaching (Baie de Rufisque). The proportion of each of the three categories of substrate identified, is a function of the area shown in the ellipses drawn above each sample. 1) infillings of bioclast pores; 2) faecal pellets; 3) infillings of microfaunal chambers. I, I1 or I11 above the ellipses indicate the mineralogical type discussed below.

Bioclastic debris of various origins frequently shelter the formation of the green pigment: fragments of tests of echinoderms, including spines, chambers of bryozoans, perforated pieces of molluscs. These two first categories of bioclasts are really equivalent substrates since, in both cases, the green pigment forms in minute cavities surrounded by carbonate. Faecal pellets (Fig. 7CD) locally constitute more than 90% of the green particles magnetically separated from a sediment. The acid leaching and ultrasonic cleaning frequently shows that an appreciable proportion of the particle is still made of carbonate or minute quartz debris. We have not observed mineral debris really replaced by green pigment; however, some sediments have shown the presence of detrital quartz grains superficially corroded or internally fissured with the green pigment seeming to line or penetrate these different kinds of cavities.

91

Figure 7. Pictures of magnetic green grains. A-B samples 95 and 96, size-fraction 100200 pm: dominant infillings in deep sediments south of Cap Vert. C-D dominant faecal pellets from sediments north of Cap Vert; C = sample 552; D = sample 486 (see Fig. 5).

92 Figure 6 gives a summary of our observations on the southern area. The proportion of green pigment in the sediment is not related to the depth; but the aspect of the green pigment is depth dependent. The green bioclasts are dominant in the sediments shallower than -30 m; the infillings found in these shallow sediments mainly occur in microgastropodes and scaphopodes. The infilled foraminifers are mostly Miliolids. Infilled ostracodes and bryozoans are also present. In the sediments deeper than -1 10 m, infilled foraminifers are very dominant but Miliolids are absent. Faecal pellets impregnated or made of green pigment are abundant or dominant between -80 m and -100 m depth. The green faecal pellets are widespread on the shelf north of Cap Vert (Fig. 5 and 7CD). In the same area, the outer zone of the shelf is also characterized by infillings of microtests, particularly ostracodes. The absence of green material in large portions of the platform to the north of Cayar canyon must be related to the rapid deposition of fine detrital sediments coming from the Senegal River. These detrital sediments either prevented the mineralization or avoided the formation of favourable substrates or even buried the mineralized substrates. The depth dependent zonal distribution of the habit of the green grains in the southern area precludes the interpretation of the deeper green material as a reworking and downward transport of shallower grains by currents perpendicular to the coast; in other words, the mineralized grains are mainly in situ in the different zones of the shelf. Particular attention is drawn to the mineralization in the shallowest sediments. At about 9 m depth, (601-4 in Fig. 2) the green pigment-bearing substrates are diverse: a minority of glossy red to ochre (oxidized) tests of Miliolids contain a dark-green pigment. This aspect is similar to what was observed in the lagoon off New Caledonia (Chapter Bl). Many bioclasts, and microtests infilled with light to dark-green material appear very well-preserved: white and not friable. This shows that the mineralization is recent and therefore, may occur in shallow environments. In this shallow sediment, about 80% of the green pigment is not dark-green in contrast to the general colour of the material in the deepest sediments (601-95). However, this shallow sediment shows traces of oxidation in a small proportion of grains which indicates that the green pigment is very sensitive to the presence and abundance of oxygen. This oxidation is absent in the slightly deeper sediment collected at -15 m (601-41 in Fig. 2). Similar traces of oxidation and reddish brown Miliolids are present in sediments collected at -70 m to -80 m depth in an area of probable sea-level still stand. This interpretation is also supported by the presence of abraded grains with a glossy appearance. In the zone from -35 m to -55 m depth, the sediments are entirely oxidized but ochre to red bioclasts show perforations infilled with dark-green material indicating that the green mineralization occured after the oxidation phase of the bioclasts. The dominant faecal pellets separated from the sediments collected from between -80 m to -100 m depth in the Baie de Rufisque, generally have a particular earthy-green slightly olive colour different from the green material elsewhere.

93

Mineralogy of the green pigments X-ray diffraction patterns The green material separated from the sediments collected from the Senegalese continental shelf has been analysed using X-ray diffraction technique applied to randomly oriented powders; two to three hundreds hand-picked green grains were crushed for each analysis. The grains were selected from different size-fractions, or different magnetic fractions obtained after acid leaching and ultrasonic treatment. Fourteen samples from the southern area and four samples from the northern area were selected. Most of the fractions were successively analysed before and after heating or ethylene glycol treatments. From more than 100 X-ray diffraction patterns it has been possible to identify three categories of green pigment: I, 11, III.

l

?

32O

,

28'

24'

200

Figure 8. X-ray diffractograms of purified highly magnetic green pigment; randomly oriented powders. The only impurity is quartz (Q) included in the pellets. Note that for the sample 601-93 we have analysed a) the dominant olive-green earthy faecal pellets and b) the clear-green gloss pellets. Copper tube, voltage: 35 KV; intensity: 26 mA; speed: l'/minute; sensibility: 4.10 ; inertia: 10 s; entrance slit: lo,exit slit 0.2. Diagram 3 is an exception for which we have used a power of 700 W instead of the usual 910 W.

?

94 @

601.108b

J

I

I

490"-3h

;SO86

p+

i"

,

,

,

12" , ,

,

,

8O I

40 I

I

I

,

550"-3h

KE Cu 12'

80

40

Figure 9. X-ray diffractograms of green pigment of type I; randomly oriented powders. The fraction selected from sample 601-108 is different from the one used in Figure 8 and was selected in order to identify its comparatively high hump at about 14 A (diagram 1). It is suggested that this hump represents an admixture of type I11 green piment with the type I pigment examplified by sample 601-65 (diagram 7). With the exception of the main diffraction at about 7 A, the limits of the surfaces probably representing the presence of diffracting mineral structures have been suggested using dashed lines. These diagrams suggest the presence of a mineral structure diffracting between 8 A and 10 A as well as at about 20 A (diagram 7). The diffraction between 8 A and 10 A is even more apparent after the ethylene glycol treatment (diagram 6). Note that heating to 490'C for 1 hour destroys the diffraction peak at about 7 A for sample 601-108 (diagram 4) but only lowers the equivalent one for sample 601-65 (diagram 8). The hump between 14 A and 20 A (untreated sample 601-108) progressively shifts toward the smaller distances as the temperature and duration of the heating treatment increase (diagrams 2 to 5).

The first observation is that the different green pigments present in the studied areas are all clay minerals. These clay minerals are illustrated by their X-ray diffraction patterns obtained without treatment in Figure 8. The first and predominant green clay minerals display a main peak at about 7.15 A with a smaller one at about 3.55 A. Many samples also show a flat and sometimes composite dome at 13 to 18 A. This kind of diagram was obtained for 10 of the 14 samples analysed from the Baie de Rufisque and 3 of the 4 samples analysed in the area north of Cap Vert (Fig. 9).

95 A second category of X-ray diffraction pattern was obtained from a single sample from the southern area (sample 91, -75 m depth, see Fig. 2). The diagram is characterized by a nearly similar size for two first-order peaks at about 7 A and about 14 A. The peak at 7 A always remains smaller than the highest corresponding peaks of category I diagrams. The peak culminating near 14 A is broad and possibly represents diverse layer thicknesses. A third category of X-ray diffraction patterns was obtained from fractions originating from four samples, i.e. three samples: 601-60 (the ultrasonically disaggregated green grain fraction from a sediment at -55 m depth), 601-92 (faecal pellets, -90 m depth), and 601-93 (olive-green faecal pellets, -100 m depth) from the southern area (see Fig. 6); and one sample: 603-552 (olivegreen faecal pellets, -75 m depth) from the northern area (Fig. 5). These diagrams show a peak located between 14 A and 15 A, and which is so high that conditions of recording different from those used reviously had to be selected (see Fig. 8, diagram 3). A peak at about 7 is always present; however, its height never reaches the one shown when analysing samples of category I or 11using similar conditions of diffraction and recording. Diagrams of category 111 were obtained from samples where the dominant substrate was faecal pellets (Fig. 10). Nevertheless, diagrams of categories I and I1 have also been obtained from faecal pellets (samples 603-486 and 603-578). The faecal pellet-rich sample 601-93 draws our attention to the fact that most of the ellipsoidal pellets have an earthy brightness and olive-green colour, but a few elements have a glossy brightness and are clear-green to dark-green in colour. About 150 of these clear-green pellets have been analysed separately on one side and the olive-green ones on the other side. The two diagrams obtained are given in Figure 8. The olive-green grains show a type 111 diagram and the clear-green a type I diagram. This indicates that: 1) the mineralogy of the authigenic green pigment is not unequivocally linked with a particular sort of substrate, 2 ) it is possible to differentiate the different categories of green pigments on morphological grounds, and therefore, 3) two separate green pigments coexist on the continental shelf, and they are locally found mixed in a single sediment but in different grains. The clays of the three categories have been independently characterized with regard to their behaviour following heating and ethylene glycol treatments. Within the samples sho.wing a composition of category I the results of the treatments are acceptably reproducible. Figure 9 gives diagrams from the green pigments of a shallow sample (601-108, green to dark-green in colour, -30 m depth) and of a deep sample (601-65, very dark-green nearly black in colour, -1 15 m depth). Due to the black colour of the authigenic clay and the much more corroded state of the carbonate substrates, sample 601-65 may be considered more evolved than 601-108. The diagrams obtained show the presence of a low proportion of diffracting material between 7 A and 10 A as was observed for verdines separated in other regions (see Chapter B 1). Heating treatment for one hour does not severely modify the peak at 7 A for temperatures of 150°, 250" or 350°C. After heating to 490°C for one hour, the

8:

96 lOA

14 6m'92b

T I

T I

601.92a

I I

/ ,

,

,

.

,

.

.

.

,

,

,

Figure 10. X-ray diffractograms of green pigment of type 111; comparison with a Recent glauconitic smectite; randomly oriented powders. Diagram 1 obtained after ethylene glycol treatment shows a clear and reproducible shift to about 16 8, compared with the untreated sample used in diagram 2. Note also the shift to high distances of a portion of the comparatively small 7 A peak, following the ethylene glycol treatment. The diagram obtained after heating to 150'C for 1 hour is similar to the one obtained after heating to 250°C for 1 hour: diagram 3. The powder successively heated to 150'C, 250"C, 350'C and 490'C for 1 hour each shows a dome at about 10 A: diagram 4. Note also a possible diffractionat about 20 8,on diagrams 2 and 4. Diagram 5: another fraction heated to 490'C for 3 hours; the same plus 550°Cfor 1 hour: diagram 6; the same fraction heated to 4 W C for 3 hours and to 550°C for 3 hours: diagram 7. Note (diagrams 4 to 7) that the shift toward 10 8, is progressive and incomplete. Diagrams 8 and 9 were obtained from a Recent glauconitic smectite from the Makassar Strait using untreated powder (normal) and a powder heated to 490'C for 4 hours (49O0C-4h).The peak, initially at 14 A, entirely and clearly shifts to 10 A.

97 peak at 7 A is considerably lowered or even disappears in the fractions selected from the less evolved sample 601-108; it remains partly visible for the more evolved black-green fractions as in sample 601-65. For all samples, the 7 8, peak disappears after heating to 490°C for two hours. The grain fractions separated from sample 601-108 initially showed a low portion of diffracting clays between 18 A and 14 A. After a one hour treatment, the dome was slightly lowered at 150°C and 250"C, and partly shifted toward 10 8, at the temperature of 350°C. This shift progressively increased after heating to 490" and 550°C; but the resulting dome remained between 10 8, and 14 A and never showed the appearance of a peak at 10 A as would be the case for a smectite. Furthermore, there was no significant modification of the X-ray diffraction pattern after ethylene glycol treatment of randomly oriented powders made of clays of category I. The analyses of green pigments of category I11 are also consistent with regard to their behaviour during the heating and ethylene glycol treatments. Figure 10 gathers some representative diagrams obtained from faecal pellets separated from sample 601-92 (-90 m depth). The weak 7 A peak is still present but lowered after a one hour treatment at a temperature of 150", 250", or 350°C but disappears at 490°C. The dominant, strong and broad peak at 14 A is lowered but does not shift at 150" or 250°C. It shifts to the lower distances at 490°C (right hand side in Fig. 10).For some fractions of the same sample, the dome obtained after heating for one hour to 490°C is very broad and progressively shifts to 10 A after treatments at 550°C for 1 hour and 3 hours. This behaviour is not similar to that shown by glauconitic minerals, an example of which is shown for comparison in Figure 10. The diffraction patterns are from a Recent glaucony collected from the Makassar Strait (South Borneo) and are similar to what is usually obtained for that sort of green pigment. Following the ethylene glycol treatment the peak at 14 A clearly shifts to about 16 A; the peak at 7 A is lowered and a portion of it seems to shift towards 8 A which is probably the 2nd order peak of the 16 A peak (Fig. 10).

Interpretation of X-raydifSraction patterns The diagrams characteristic of the dark-green pigment showing a main peak at 7 A in our conditions of recording (category I) have been observed in different regions such as New Caledonia (Chapter Bl). The problematic green clay was called phyllite V.and we have here a new example of this material. An interpretation of the diagrams for the clay of category III is not simple. The peak at 14 A is not the usual chlorite peak because: 1) it is very much lowered at 150°Cas for a smectite; 2) it shifts to low angles after the ethylene glycol treatment; 3) it shifts towards 10 A after heating to 490°C. The shift obtained after ethylene glycol treatment is not fully comparable with that shown by some smectites: the clay structure does not collapse to 10 A as would that of a true smectite or a glauconitic mineral. This behaviour can be interpreted as due to a mineral halfway between a smectite and a chlorite. It could be called a swelling-chlorite:a group of hybrid clay minerals in which are

98 601.91~

?

?

I

I

148

T

'rs" /

,'

Y Norma

Figure 11. X-ray diffraction patterns of the green pigment of type I1 collected from the Senegalese shelf. The untreated sample (Normal) of the fraction 601-91b (diagram 1) shows a lower peak at 14 A than the fraction 601-91c (diagram 6). The fraction 601-91b has been cumulatively heated to 150"C,250'C (diagram 2), 350'C (diagram 3), 490'C (diagram 4) and 550°C (diagram 5 ) for 1 hour each. Note the progressive shift of the dome from about 14 A to about 10 A. The ethylene glycol treatment gives a shift at about 16 A for the fraction 601-91b (diagram 7). There is not yet any shift for the same fraction heated to 490'C for 3 hours (diagram 8).

gathered minerals showing a variety of behaviours; however, these clays do not usually collapse to a distance as small as that shown on our diagrams; moreover, the 7 A peak is very low on the untreated samples and is not characteristic of a chlorite. One may also imagine a mixed-layer structure (smectite-chlorite). We propose in the first instance to call the green clay of the olive-green faecal pellets of the Senegalese shelf phyllite C (C being for its partly chloritic behaviour). The behaviour of the green components of category I1 is similar to that shown by category I11 for the peak at 14 A; this peak shifts to 16 A following

99 ethylene glycol treatment: it is very much lowered after heating to 150" and 250°C. Following heating to 350" or 490°C a dome remains present at about 14 8, for some preparations; but two small eaks at 10 8, and 14 8, seem to be present for others (Fig. 11). The peak at 7 (unchanged after heating to 150" or 250°C, lowered at 350"C, and destroyed at 490°C) has a behaviour similar to the main peak of the components of category I (Fig. 11). In conclusion, we consider that the X-ray diffraction patterns of category I1 denote the presence of a mixture of green pigments of the categories I and III.

8)

Chemical study Chemical analyses of green pigments collected off Senegal was only possible €or "deep" sediments of the southern area. There, the magnetically separated pigment consists of predominant dark-green grains; pale-green or slightly olive-green grains are also present. Substrates are made of about 70% infillings, 5 to 10% faecal pellets, and about 20% bioclast replacements. The X-ray diffraction patterns were of type I. Samples 95 and 96 were mixed in order to have enough material to perform a precise chemical analysis. Although ultrasonically cleaned and leached three times with N/10 acetic acid, carbonate was still present after the purification process as the only visible impurity. The chemical results (conventional wet chemistry) shown on Table I indicate a high content in CaO and low total percentage. This prevents the direct use of these results for a mineralogical interpretation; however, the two analyses are similar for a material presumed to be homogeneous. Using the mean percentages found for the elements which can be assumed to be present in clay layers, we have recalculated "corrected" percentages (assuming that CaO was not in the clay and that the total had to be about 100%). From this exercise, it is possible to draw the following conclusions. Table 1. Chemical composition of two verdine separates (type I) from the Senegalese shelf (Baie de Rufisque). Recalculated results obtained after elimination of the polluting carbonate.

601-95/96AB

SiO2

A1203

27.8

8.08

Fe2O3

FeO

CaO

MgO

NaaO

K20

17.2

5.99

7.26

11.8

0.2

0.41

P2O5 H 2 0 0.4

2.29

HzO' 11.1

28.9 601-99 AB

29.9

dntn

34.8

92.53 93.63

.

7.38

19.1

5.17

8.48

10.4

0.2

0.36

0.4

2.40

7.1

29.6 recalculated

Total

90.89 90.59

9.3

21.7

6.6

(0)

13.2

0.2

0.5

0.4

2.8

10.8

99.9

The silica content of this mineral is compatible with a clay structure at 7 8, or at 14 A. The mineral is very rich in iron; this iron is dominantly ferric; the magnesium content is higher than the aluminium content. All these properties as well as the individual contents in each element are very similar to the corresponding properties of the phyllite V mineral found in New Caledonia.

100 DISCUSSION AND CONCLUSIONS

The mineralogical study of the green pigment collected from the Senegalese shelf shows that the investigated material cannot be referred to the glaucony facies. We have here a new example of the verdine facies.

Sedimentology of the verdine facies off Senegal First of all, the facies was observed everywhere in the sampled area off Senegal; the platform is located between the tropics like the formerly described Neo-Caledonian area. The environment is marine with respect to the salinity parameter. There is a large river to the north of the studied platform the influx of which moves southward where the verdine is present; the green pigments are located in areas where no detrital sediment deposition occurs. Northward, where the muddy tongue derived from the Senegal River has been formed prior to the present conditions, no green pigments are found. It seems that the green pigment develops deeper i.e., offthe depositional area of the detrital sediments. The green pigment is present at depths (locally less than -10 m) shallower than is usually accepted for the glaucony facies. Traces of oxidation are always present there but less abundant in sediments deeper than -15 m. We have quoted traces of oxidation in elements formed before as well as after verdine genesis at different depths. This indicates that the environment needed for verdine genesis is usually very near the zone where oxidation of iron occurs: a shallow depth in the sea as well as a location near the sea-waterhediment interface. This, a contrario, implies that the verdine formation occurs in a microenvironment protected against the oxidation processes. Nevertheless, abundant green pigment has also been observed in sediments 200 m below present sea-level; the question then arises as to when the green pigment has formed on the continental shelf. Several hypothesis are to be discussed. Hypothesis I assumes that the green pigment formed at a slight depth (less than about 60 m). According to the sea-level curve evolution (Fig. 12) and the fact that about 10 m of water are needed to avoid a systematic oxidation process, the green pigment shallower than about -35 m is about 5,000 years old. In this context, the oxidized green pigment collected between 50 m and 80 m could represent a shallow alteration process older than this age; in this case, the genesis of verdine should be considered diachronous. We suggest that the verdine located presently between about -140 m and -10 m depth has formed diachronously from about -18,000 years ago to the present time. Hypothesis 2 would be that verdine can form deeper than -60 m. This would explain the presence of abundant verdine at about 180 m depth or more. In that case, the verdinization process could have been realized only once at various depths (and recently) and the oxidation processes observed in different areas would be explained by comparatively deep oxidizing currents created by the submarine topography. However, a short distance transport downward could explain the occurrence deeper than -180 m as well.

101

Figure 12. Sea-level changes off Africa during the last 70,000 years. (Scheme according to data by Faure and Elouard, 1967 and H. Faure, personal communication; redrawn from Odin and Dodson, 1982)

Another indication of interest concerns the temperature at which the verdine forms. Today, we know that during the humid season the temperature of the sea-water is relatively high. Although the sub-reefal character of some of the relict sediments quoted above is questionable, a temperature of 20°C or slightly above, occurs today several month a year on the shelf, and probably during the main part of the year about 12,000 years ago in connection with the decreasing importance of Trade Winds and the corresponding upwelling system (Masse,

1967). Mineralogy of the verdine facies off Senegal The mineralogical study has shown that the green pigment collected on the Senegalese shelf was always clayey. In contrast to the Neo-Caledonian deposit, the Senegalese continental shelf shows the presence of two mineralogically distinct green clays. The most frequent green clay appears very similar to the one idenMied in the Neo-Caledonian lagoon. The dominant peak is at about 7.2 A; the intensity and width indicates a weak crystallinity, and the base of the diffraction peak reaches 10 A. The chemical data are very specific: abundance of ferric iron, high SiO content compared with berthierine of Brindley (1982), relatively high Mg, an4 very low ferrous iron contents. This is our preliminarily called phyllite V (Odin, 1985), the presence of which, in marine sediments, is the main character of the verdine facies. The study of the Senegalese shelf allowed us to obtain a green marine clay

102 mineral showing X-ray diffraction patterns very different both from those obtained from glauconitic minerals and from phyllite V. This clay mineral is characterized by a main peak near 14 A, very strong compared with that shown by phyllite V; the behaviour of this mineral is comparable with that which would be shown by a swelling-chlorite or a mixed-layer structure chloritesmectite. We suggest to give it the provisional name phyZZite C.We think it useful to identify that green phyllite by a specific name because its mineralogy is not usual in sediments and it was found in a very particular environment: in faecal pellets and near to the sea-waterhediment interface; it might well correspond to a specific undescribed clay mineral phase. Finally, a few green pigments showed both a large peak at 14 A and a reasonably well-shaped peak at 7 A. From the series of diagrams available, this kind of green pigment must be interpreted as a mixture of the two minerals identified above because: 1) The two peaks at 14 8, and 7 A have independent behaviours after treatments by heating or with ethylene glycol (sample 601-91); this argument might be rejected since, for many chlorites, the peak at 7 A disappears when the 14 A one is unchanged or increased; however, we have shown that, after heating to low temperature, the peak at 14 A was modified and decreased in the mixture while the peak at 7 A remained unmodified. 2) The ethylene glycol treatment does not modify the peak at 7 A of phyllite V but the phyllite C portion of the mixture swells; in this sense, even the sample 601-92 with its high 14 A peak, seems to contain a small proportion of phyllite V and sample 601-65, with its very dominant peak at 7 A, possibly contains a proportion of phyllite C. 3) If a continuous series of minerals (from phyllite V to phyllite C) is postulated, the two end-members of the series would be much less frequent than the mean intermediate component in contrast to what we have observed. 4) In several samples of our end-members, we may observe a small portion of the other mineral; this observation has been shown to be in connection (as for sample 601-93) with the presence of a small proportion of grains slightly different in colour from the dominant pigment.

Microenvironment and genesis of the phyllite V mineral The phyllite V mineral fills many kinds of pre-existing microcavities in different biogenic substrates. These cavities are a structural property of the carbonate bioclasts (chambers of microfaunal remains; stereom of echinids; loculi of bryozoans ...) or a newly formed porosity linked with biogenic boring. These cavities create large surfaces, and allow sea-water to circulate within the grains; they also delimit volumes relatively isolated from the open marine surrounding environment. It is only in such microenvironments that phyllite V develops while the clay-size fraction of the sediment, when present, has a composition and a colour different from those of the green pigment. We may therefore conclude: 1) the microenvironment corresponding to these pores is a necessary prerequisite for the green clay crystallization; 2) this crystallization

103 occurs mainly using ions from the sea-water while the substrate is usually free of the necessary ions (Mg from magnesian calcite could be utilized however); 3) exchanges of ions with the exterior of the substrate are needed; 4) the relative local confinement created within the micropores is necessary. In these conditions, a favourable environment of clay crystallization may occur at the top of the deposited sediments.

Conclusions On the tropical shelf off Senegal, where minor detrital sedimentation occurs, we have found authigenic green pigments. According to sedimentological observations, these green pigments formed in shallow sediments. However, the genesis of the pigment itself is not compatible with oxidizing conditions which alter the neoformed material. The genesis is probably diachronous and older in the deepest zone of the continental shelf than in the shallowest one. The age of the green pigment found deeper than about -140 m deph cannot be estimated with precision but that of the sediments shallower than this depth are younger than 20,000 years the main part of which is about -12,000 years BP; some material is certainly younger than about 5,000 years or even 3,000 years. The detailed mineralogical study indicates that two different green clay minerals compose these pigments. The most widespread clay is similar to the phyllite V first named by Odin (1985a) and already met and described in the lagoon of New Caledonia. The second green clay shows X-ray diffraction characteristics partly similar to those of a swelling-chlorite and could also be considered as a mixed-layer chlorite-smectite; this mineral was not described previously from Recent marine sediments and it is proposed to identify that newly described marine sedimentary green clay as the phyllite C. Phyllite V forms in a particular microenvironment limited to microcavities in sedimentary grains deposited near the sea-waterhediment interface. It forms mostly in sands or muddy sands under the influence, at least temporarily, of a relatively high temperature sea-water. Phyllite V and phyllite C seem to form in very similar conditions and may be interpreted as two different expressions of the verdine facies. The first mineral shows a dominant peak near 7 A, the second mineral shows a dominant peak near 14 A. It is.very tempting to Compare this association: phyllite V-phyllite C with the one observed in ancient ironstones: berthierine and chamosite. However, this idea needs further analyses to test and evaluate in what sense it might be developed. AKNOWLEDGEMENTS

Comments by J.P. Debenay on the earlier versions of that chapter are greatly appreciated. We are deeply indebted to Peter Skelton from the Open for his improvement of the English of the University, Milton Keynes (U.K.), final version of this chapter.

104

The Makassar Strait glauconitic sediment (p. 96) was provided to the editor by Dr Gayet (Univ. Bordeaux). Chemical analyses (p. 99) were performed by M. Lenoble (Univ. Curie, Paris). Photo processing was realized by 0. Fay.

105

Chapter B3 THE VERDINE FACIES OFF FRENCH GUIANA by G.S. Odin, I.D.R. Mackinnon and M. Pujos INTRODUCTION

Sedimentological investigations were undertaken on the northern continental shelf of South America before 1970 by Porrenga (1965) and Nota (1969). Green grains have been observed and analysed off the Orinoco delta (Porrenga, 1967a) and on the continental shelf of Surinam (Hardjosoesastro, 1971). These green grains were identified as "chamosite" and "glauconite". Later, green grains were collected more to the east, off French Guiana and have been described by Moguedet (1973) and Bouysse et al. (1977) following three series of samplings by 1) the Institut Scientifique et Technique des Pkches Maritimes (I.S.T.P.M.) in summer 1971, 2) the Bureau de Recherches GCologiques et Minihres (B.R.G.M.) in 1975 and 3) a joint effort between B.R.G.M. and Institut de GCologie du Bassin d'Aquitaine (I.G.B.A.) in 1976. Samples from these sediments were provided to the present editor by the I.G.B.A. and by G. Moguedet for a detailed study of the nature and palaeosedimentological significance of the green grain-bearing sediments. Data were especially gathered in order to supplement the theses of Jeantet (1982), RCniC (1983), and Chagnaud (1984). A sedimentologic synthesis for the continental shelf has been proposed recently by Pujos and Odin (1986). This chapter will essentially focus on green grains from the shelf of French Guiana as a regional example of sedimentologic process on the whole stable continental margin from the Amazon to the Orinoco River. The location of the areas considered here is shown in Figure 1. Local variations exist on this vast portion of the continental margin which is about 1800 km in length. However, the general relationship between the different kinds of sediments discussed here remains very similar from the Amazon delta to the Orinoco delta. These deltas are located between the equator and 12" N latitude. The continental shelfis over 200 km wide off the Amazon mouth and about 80 km wide westward compared to a mean width of about 40 km for all oceans. Another feature of the present conditions of sedimentation is a general superficial stream along the coast which pushes the continental waters from the Amazon River mouth towards the N W .The Amazon River really influences the sedimentology of all the continental margin due to its considerable discharge of 175.103 m3/s. This discharge results in about 500 to 900.106T/year of material being brought to the sea. The finer sized material is trapped along the coast in the form of a wide tongue of mud off the Amazon, and which progressively narrows until the Orinoco mouth (Gibbs, 1973, 1976). Everywhere, coarser sediments are initially deposited on the middle part of the continental shelf.

106

Figure 1. The continental margin between the Orinoco and Amazon River deltas. The area presently studied is located off French Guiana (Guyane). The dotted areas show the locations where the verdine facies is known to occur according to our study and a reinterpretation of previously published data.

These conditions only concern most Recent time as characterized by a high sea-level. During the last sea-level regression, about 18,000 years ago, the Amazonian material was deposited on its deep delta facing the river mouth through a canyon cutting the slope. GEOLOGICAL SETTING OF THE SHELF OFF FRENCH GUIANA

Continental factors The muddy coast of French Guiana is generally very flat and occupied by wet swamps and mangrove as a result of the equatorial climate. Mangroves are also present along the banks of the very wide rivers like the Oyapock and the Maroni which define the eastern and western borders of the country respectively. Complementary to the dominant influence of the Amazon River, a number of coastal rivers add water to the sea today as in the Recent past. These rivers usually bring mainly ions to the sea. Iron is precipitated in the estuary as soon as the influence of marine salinity appears: 10 to 20 km up river from the coast (Roche, 1977; Chapter B6). However, large quantities of ions and suspensions

107

are brought to the sea during the humid season (Blancaneaux, 1981). These rivers erode the hilly pre-Palaeozoic Guiana Shield (Fig. 1) where the tropical wet forest grows. Sediments on the continental shelf

Figure 2.Relief and granulometric data on the continental shelf off French Guiana. 1) terrace boundary; 2) muds and silty muds; 3) fine to very fine sands; 4) mid to coarse sands. (Simplified from h j o s and Odin, 1986)

In general, the shelf off French Guiana presents a sediment distribution parallel to the coast. This distribution is underlined by a series of topographic breaks, more or less continuous at about -20 m and -110 m (Fig. 2). Muds are deposited in two zones: the coastal zone is linked to the Recent Amazon input as noted above; the external zone at the shelf edge, is relict and the mud has the same origin as the coastal mud as shown by their nearly similar natures (illite, 40%; illite-smectite, 40%; and kaolinite, 20%). However, the chlorite content is usually higher in these muds (2 to 8% of the clay-size fraction) compared to the clay analysed from the Amazon waters themselves (0% to traces) by Eisma and Van der Mare1 (1971). This difference might be explained by the complementary clay input of local rivers, as suggested by Bouysse et al. (1977), because chlorite contents of 13 to 20% have been measured in clays from the Oyapock and Maroni Rivers. Presently however, the input of these two rivers, estimated at 0.03.lo6 T/year and 0.6. lo6 T/year respectively (Jouanneau and Pujos, 1986), is practically negligible compared to the Amazon input (500.106 T/year according to Gibbs, 1967) and can hardly influence the composition of

108 these shallow sedimented clays. Between the two muddy zones, the main portion of the shelf is covered with fine to very fine sands. Facing the Maroni, coarser oxidized sands mark the palaeo-delta and suspected palaeo-course of that river oriented eastward (Fig. 2).

Figure 3. Carbonate proportion on the continental shelf off French Guiana. 1) 10 to 20% of the sediment; 2) more than 20% of the sediment; 3) presence of carbonate ooids; 4) abundant mi&. (After Bouysse et al., 1977)

Carbonates are more or less present in the fine to very fine sands and abundant in the external zone between -80 m and -150 m depth. These carbonates are formed from corals, Porifera, benthic foraminifers (Amphistegina), and bryozoans (Fig. 3). This assemblage represents a palaeo-thanatoccenosis of sub-reefal character. Reefal structures are mainly present along a line located at a present depth of about -100 m (Moguedet, 1973; Bouysse et al, 1977). On the southeastern half of the external shelf, abundant ooids were described first by Moguedet (1973) and have been observed locally by later authors. These ooids are calcitic and usually spherical; they have been illustrated by Moguedet (1973) and Chagnaud (1984). Ooids are sometimes white to cream in colour, but also become grey or ochre to reddish-brown due to an increasing ro portion of iron oxide (about 2 wt% of Fe 0,). Moguedet reports a radiometric dating of 17,000 f 400 years for dese ooids and concludes that the deeper ooid-rich sediment, collected at -130 m depth, clearly marks the time and location of the last regressive phase because the oolitic facies is known to form at a few metres of depth.

P46

109 However, the apparent 14C age of 17,000 years must be discussed because the dated ooids are not completely made of neofonned carbonate. The heart of these ooids sometimes consists of "dead" carbonate i.e., contains reworked carbon for which the 14C isotope was initially absent and, consequently, the carbonate nucleus can be considered an inherited old component. Therefore, the apparent age calculated from the presently remaining radioactivity of the ooids must be considered a maximum age (H.A. Geyh, personal communication, December, 1984). According to thin section pictures published by Moguedet (1973), the proportion of initially "dead" carbonate nuclei in the dated ooids is low and can be estimated at about 10%. The age of the ooids can therefore be corrected to 15,000 years with an error bar of about 2,000 years. Finally, patches or lines of hard rocks have been observed on the shelf using seismic detectors. They are interpreted as a palaeo-barrier reef and occur between -80 m and -150 m depth (Moguedet, 1973; Bouysse et al., 1977). A similar sedimentological scheme has been described from Brazil to Venezuela. The reefal structures off North Brazil have been considered in detail by Santos (1972), and off Surinam and Guyana by Nota (1958; 1969) and Morelock (1972). Nota (1969) noted fossil reefs at relatively low depths (-26 m, -70 m and -90 m) according to the presence of submarine height and he suggested ages of 12,000 to 17,000 years for the deeper reefs. AUTHIGENIC GREEN GRAINS

The glaucony facies In his preliminary work, Moguedet (1973) describes magnetic green grains (glauconie between quotation marks) everywhere on the continental shelf off French Guiana, but discusses only the richer green sands at depths greater than -200 m. Bouysse et al. (1977) distinguishes 1) the 'glauconie verte' for which they identify two facies: the 'glauconie du large' (deep glaucony) and the 'glauconie verte de la plateforme' (glaucony from the shelf) and 2) the 'glauconie brune' (brown glaucony). The green grains deeper than -150 m, actually represent true glaucony. Mineralogical study shows that the clay has a mean periodicity near 11 A and can be related to a slightly-evolved glauconitic facies. The proportion of green grains in the sediment is higher than 10% and can reach 50%. The facies has only been unambiguously identified from a portion of the outer shelf and slope between -150 m and -800 m depth, but its lateral extent is very probable (Fig. 4). The green grains are usually coarse and represent the major portion of the sediment coarser than 0.125 111111. These grains are still abundant in the sizefraction 0.6-0.8 mm (Moguedet, 1973). The substrates of verdissement are essentially microfaunal test chambers and faecal pellets. Green grains made entirely of the glauconitic clay are deeply cracked. Finally, on the outer part of the shelf and on top of the slope (-150 m to -200 m depth), these green grains are generally more or less oxidized and modified to goethite. This form is the brown glaucony of Bouysse et al. (1977).

110

Figure 4. Green grains on the continental shelf of French Guiana. The samples submitted to a mineralogical study are shown. 1) verdine facies; 2) glaucony facies; 3) very oxidized grains. When recognizable the dominant substrates are identified ; C = faecal pellets (Coprolites); I = Infillings; M = more or less chloritized biotite Mica.

T h e verdine facies Most green grains on the continental shelf represent the verdine facies. Figure 4 shows the sites where this green pigment has been mineralogically analysed. However, the same type of green grain has been observed in the whole area between about -20 m, the border of the coastal mud zone, and about -150 m. The proportion of green material is between 0.5% and 12% of the total sediment. The fine to very fine sands which contain these green grains were considered to be a very thin .blanket of about 1 dm in thickness by Bouysse et al. (1977). However, six short cores analysed by Chagnaud (1984) off French Guiana as well as two additional cores obtained off the Amazon River mouth at depths greater than 20 m show that these verdine-bearing fine sands may be 2 m thick. These sands overly relict indurated oxidized clays presumed to be 30,000 years old or more (Bouysse et al., 1977). In contrast to the observations on glaucony, green grains from the shelf are predominantly smaller than 160 p.m. Figure 4 also gives the identification of the verdissement substrates when it was possible to recognize its nature. Although the dominant recognizable substrates are shown separately in Figure 4, a mixture of substrates can often

111 be observed. Infillings of benthic foraminifers sometimes dominate and are particularly abundant on the outer shelf (I on Fig. 4). Faecal pellets are abundant locally. These pellets probably deposited at the mouth of a river when the sea-level was low. However, the main substrate on the Guianese shelf is a mineral substrate: biotite mica flakes (M in Fig. 4). These mica flakes were already very altered before undergoing the verdissement process. Thin sections and mineralogical study show that the micas were more or less deeply chloritized. However, the accordion-like green grains are usually moderately opened in that the thicknesses of the grains are smaller than the initial diameter of the flakes (Fig. 5). Therefore, the vermicular facies so characteristic of some glauconitized biotites (Odin, 1972; and Chapter C2, p. 259) was not obtained. The distribution of the mica flake substrate shows two trends. Micas are less frequent from the coast toward the deep shelf and from the SE to the NW. The lateral trend gives an indication of the origin of the flakes which, most probably, were brought from the Amazon River.

Figure 5. The substrates of the verdissement processes. Left hand side photo.: verdine from sample 29, surface sediment of a core at -95 m depth, 500-200 pm size-fraction showing mature black-green grains including infillings of foraminifera1chambers (F). Central photo.: verdine from sample 508-175A, sediment collected at -50 m depth, 500-200 pm size-fraction showing an expanded mica (M) and a faecal pellet (C) both green. Right hand side photo.: glaucony from sample 508-217 collected at -205 m, size-fraction coarser than 500 pm, the substrates are mostly more or less modified ellipsoidal faecal pellets. (Photomicrographscourtesy of 0. R6ni6.1983 and M. Chagnaud, 1984)

112 Compared to the Neo-Caledonian and the Senegalese deposits described in the two previous chapters, the carbonate residues initially surrounding infillings (the tests) have mostly been dissolved. This feature indicates a more mature sediment and, therefore, a more evolved verdine facies. Moreover, a number of samples show that they underwent an oxidizing process especially at depth between -80 m and -150 m, and shallower off the Maroni and the Oyapock Rivers. Figure 4 only indicates the sites (circles) where the oxidation process was visible on the majority of the grains; but the phenomenon is very extensively developed everywhere on the shelf. MINERALOGY OF THE VERDINE FACIES

X-ray diffraction study

4

108

14H 20A

T T T

Figure 6. X-ray diffraction patterns of green grains from the shelf (relevant samples are located on Fig. 4). The intensity at which grains were attracted with the magnetic separator is given (F) in amperes. Admixture of quartz (Q) is frequent in these separates. All samples are ultrasonically cleaned with N/2 acetic acid for six minutes. The purest grains (1) are mainly foraminiferal infillings and are composed of phyllite V. The grains of the fraction (4) are mainly green biotite thin flakes. Thicker green biotite flakes containing the authigenic phyllite V gave diagrams (2) and (3).

lp

113

i" 1y

508-202 F0.43

r Q

r

100" 2 h

1

K a CU

Figure 7. X-ray diffraction patterns of green grains of verdine after biotite (see location of the samples on Fig. 4).The untreated sample (1) shows a small diffuse peak at 14 A. After heating for 2 hours to 300'C (2), the trace is not fundamentally modified. After heating to 490°Cfor 4 hours (3), the peaks at about 7.2 8, and 3.6 A (about 25' 28) disappear and the diffraction at 14 A is increased. The grains probably contain inherited chlorite together with phyllite V.

The interpretation of the X-ray diagrams obtained from purified green grains from the Guianese shelf was problematic as long as it was not understood that the nature of the'original substrate had an "influence" on the apparent mineralogy. However, it was clear, as soon as the first X-ray diagrams were obtained, that most of the continental shelf was covered with grains characterized by a main peak at 7.2 A . Figure 6 gives a series of diagrams which illustrates the situation. Diagram 1 is from green grains which are dominantly infillings (see Fig. 4 for location). The diagram is dominantly that of phyllite V with a small admixture of quartz. This admixture of quartz is very difficult to avoid because the impurity is located inside the grains. When green grains from an area with a dominant biotite substrate (samples 106 or 171, see Fig. 4) are analysed, then a new peak

114 develops at 14 A; but it is nearly always associated with a diffraction peak at 10 A. At first sight, this peak is not that of a clay mineral (sharp shape) but represents mica remains. Moreover, when a series of diagrams is undertaken (on different magnetic fractions of a single sample), an evolution can be recognized (see diagrams 3 and 4, Fig. 6). The less magnetic the green grains, the less they look like rounded grains (they are flakes) and the more the 14 A peak develops while the peak near 7 8, becomes sharp and nearer to 7 A. For the less magnetic fraction, one obtains an X-ray trace of chloritized mica (very high and sharp peaks at 14 A, 10 A, and 7 A) without trace of marine authigenic mineral. In this case, the powdered flakes have a grey colour rather than a green colour, and thus these flakes cannot be mistaken for verdine so that the corresponding X-ray diagram is not shown here. Due to this situation, it was necessary to systematically obtain three or four diagrams on different magnetic fractions for each verdine sample in order to be sure that the peak observed near 7 8, was that of the authigenic marine mineral: phyllite V. In some cases, the situation was not very clear but the heat treatments allowed us to obtain a correct interpretation of the diagrams. Figure 7 shows an example (see location of sample 202 in Fig. 4). A heat treatment at 300°C for 2 hours does not sensibly modify the diagram, but a temperature of 490°C for four hours destroys the peak at 7.2 8, and reinforces the 14 A peak. One may also note the presence of a small peak at 10 A. These green grains can be interpreted as a mixture of dominant authigenic phyllite V with a small proportion of very chloritized mica in which the iron-rich chlorite itself has been altered and therefore gives comparatively broad peaks. In summary, a new interpretation for the presence of a 14 8, peak in the diagrams obtained from presumed verdine is as follows: the 14 A peak is not due to an admixture of foreign detrital chlorite grains as in some samples from New Caledonia but is due to the favourable substrate itself in which verdine forms. It must be em hasized that the top of the diffraction peak near 14 8, remains clearly at 14 in these samples i.e., there is not a hump between 14 8, and 20 8, as in the samples from Senegal or New Caledonia described in the previous chapters. On some occasions, the diffraction peak near 14 A is too low for a clear interpretation (e.g. diagram 1, Fig. 6). In many samples, an admixture of goethite has also been observed when a portion of the grains was ochre or red (diagram 2, Fig. 6). In spite of these impurities such as 1) admixture of quartz within the grains, 2) remains of chloritized mica from the substrate, and 3) goethite due to oxidation, nearly pure fractions of green marine clay have been separated from a number of samples using granulometry, density, form, and paramagnetism. The systematic control of these fractions by X-ray diffraction shows that, everywhere, the authigenic clay from the shelf is characterized by a main peak at 7.2 A. As quoted above, the outer part of the shelf and the slope is rich in green grains which can be related to the glaucony facies. A particular case has to be quoted at the boundary between these two sedimentologic domains. The dark-green gains of sample 205 (location in Fig. 4) shows an X-ray diffraction

i

115

pattern with a strong hump at 14 A and a lower intensity hump at 7 A. This peak disappears after heating to 490°C for four hours (Fig. 8), while the 14 A peak shifts to I0 A . These X-ray diagrams are interpreted as a mixture of glauconitic minerals with phyllite V. The green grains concerned have an homogeneous aspect; therefore, the same substrates have successively sheltered the crystallization of the two kinds of minerals. This example shows a second new possible reason for the presence of a peak near 14 A in the X-ray diagrams obtained from green grains of the verdine facies. 508-205

I'

0

SO-86.

32'

28'

24"

200

IW

120

80

Kd CU

40

Figure 8. X-ray diffraction patterns of green grains collected from about -100 m. The diffraction peak near 7 A is low, the one at 14 A is high. The latter is due to a glauconitic smectite which collapses after heat treatment to 490'C. The former is due to a poorly crystallized phyllite V. Quartz (Q) is still present.

Analytical electron microscopy (AEM) In order to obtain additional data on the actual nature of the green clay in the verdine facies, a sample from the French Guiana shelf was investigated using The sample (508-175A; see location in analytical electron microscopy (AEM). Fig. 4) is a coarse grained fraction, between 150 pm and 300 pm in diameter. These black-green grains were ultrasonically cleaned in a dilute acid bath. The X-ray diffraction pattern from this sample is shown in Figure 9 and is typical for phyllite V, with a predominant peak at 7.2 A. Selected green grains from

116 this coarse fraction were then prepared for AEM by embedding in an epoxy and sectioned to less than 100 nm thickness with an ultramicrotome. The embedding procedure is a relatively gentle process and is known to expand smectite clays only to 15 A basal spacing. Other non-expandable clays do not show expansion of the basal spacing using this technique (Vali and Korster, 1986). However, the precise behaviour of minerals from the verdine facies remains to be determined. Each slice was then mounted on a carbon substrate attached to a 150 mesh copper grid and placed in a low-background, double-tilt sample holder. Samples were analysed with a JEOL 2000FX AEM and attached Tracor-Northern TN5500 energy dispersive spectrometer (EDS) for secondary X-ray collection. Analysis procedures for clay minerals using this instrumentation are given by Mackinnon and Kaser (1987).

I 508 - I75A

108

148

1

K a Cu 36

32

28

24

20

16

12

8

Figure 9. X-ray diffraction pattern of the verdine from sample 508-175A This sample was chemically analysed (Table I) and submitted to the AEM analysis.

Observations Low magnification photomicrographs of sample 508- 175A from a number of different areas show a wide range of grain-size and morphology. Representative images from the same section and from within an area of 10 pm2are given in Figure 10. Figure 10 shows (A) large, 1 pm long euhedral grains, (B) very fine-grained (less than 0.1 pm) laths, (C) more equant, subhedral grains between 0.2 and 0.4 p.m size and @) intermediate size 0.2 to 1 pm laths. These low magnification micrographs suggest that at least three different phases occur

-+-+-+

Figure 10. Low magnification (x 25,000) transmission electron microscope images of a microtomed section of sample 508-175A from French Guiana. Four different types of grainsize and morphology are apparent in each image. Scale bars in A and D are one micrometre. (A) Predominant euhedral quartz (Fig. 13a); (B) non analysed small particles; (C) numerous particles made of a 10 A spacing clay (see Fig. 11) with Si, Al (Fig. 13b); (D) predominant 7 A spacings (Fig. 12) with Si, Fe, Al, Mg (see analyses in Fig. 14C. 14D).

117

118 in this sample of verdine. This observation is generally confirmed by high resolution transmission electron microscopy (HRTEM) and elemental analyses (AEM) discussed below. HRTEM images from regions shown in Figure 1OC and 10D mostly indicate the presence of three layer silicate minerals. These layer silicates have basal spacings of about 7 A, 10 A, and 14 A. The mineral with =lo A spacing is related to those morphologies shown in Figure lOC, while the silicate with =7 A spacing is more closely associated with the image shown in Figure 10D. Crystal thicknesses in the basal direction for 10 A particles are between 150 8, and 300 A and typical examples of these clays are shown in Figure 11. This 10 A clay mineral usually shows well-defined crystals which appear to be structurally well-ordered. The 7 8, clay is much smaller in the basal direction, usually between 50 A and
Figure 11. High resolution image of particles similar to those shown in Figure 1OC. Both examples show 10 A basal spacings for crystals ranging in basal dimension from 150 8, to 300 A. These crystals are well-ordered and are primarily alumino-silicates.

119

Figure 12. Examples of 7 A (A) and 14 A (B and C) phases from areas as shown in Figure 10D.These crystals are smaller in the basal dimension and do not appear to be as well-ordered as the clays shown in Figure 11. Clays shown in this figure display the presence of Si, Fe, Al, and Mg. Magnification in Figures 11 and 1 2 is approximately 2 million times.

120 275

cu

%

t .-

u)

c

0) t

-c 0 1020

0 2

4

6

Energy

8

10

(KeV)

Figure 13. Elemental analyses from particles shown in Figures 10 to 12. (a) EDS spectrum from the large euhedral particles in Figure 10A and (b) typical spectrum from particles showing 10 8, spacings as in Figure 11.

Figure 13 shows typical elemental analyses (for atomic number above 11) of individual particles from the regions shown in Figure 10. These analyses were obtained from areas smaller than 300 8, in diameter from many particles and so represent the approximate element abundances for at least three different types of minerals in this sample. A spectrum from the large euhedral particles shown in Figure 10A is given in Figure 13a. The small copper peaks in all spectra arise from rescattered electrons interacting with the sample grid and is not of significance in these'analyses. The spectrum in Figure 13a shows the presence of only silicon in these grains and preliminary analysis of electron diffraction data (not shown) indicates that these grains are quartz. The relative abundance of this material in the sample is estimated at about 15%. The micro-chemistry of particles shown in Figure 1OC and Figure 11 is given in Figure 13b. These particles show predominantly aluminium and silicon with minor amounts (less than 5 element YO)of iron. When considered in conjunction with a 10 8, basal spacing, this clay can be identified as a pyrophyllite- or mica-type. Elemental analyses from the particles in Figures 10D and 12 show that the

121 material is predominantly a Mg, Al, Fe-silicate. Examples of fine-probe (<30 nm) analyses from the regions shown in Figure 10D are presented in Figure 14. Spectra shown in Figures 14B, C, and D are from the individual 7 A (or 10 A) particles denoted by corresponding symbols (b, c, d) in Figure 10D. However, since the particle size is smaller than the region which generates X-rays (due to beam broadening effects), a weak signal from other overlapping particles may be present. Thus, the presence of minor elements (e.g., potassium) may not be directly attributed to the predominant mineral present.

I I

00

40

80

120

keV

160

00

Fe

20

40

6.0 keV 00

134E

D

B

Fa

(

00

20

40

60

00

00

20

40

6.0 keV 00

Figure 14. Elemental analyses from particles shown in Figure 10D. (A): composite spectrum; (B) spectrum obtained from a dominantly mica-type mineral. (C and D): typical spectra for particles showing 7 A spacings.

The analysis shown in Figure 14A contains abundant Ca, Sr, Ag, and lanthanide elements and indicates that there is some mineralogical heterogeneity even in sections of the sample which are predominantly 7 A material. Sample analysis location b in Figure 1OD corresponds to the spectrum shown in Figure 14B. This spectrum shows a small potassium peak and relatively high Al, Si

122 peaks (compared to Mg and Fe) probably due to the overlap of both 7 8, and 10 A particles within the area of analysis. A HRTEM image from part of this analysis region is shown in Figure 12A. Nevertheless, the predominant chemistry for individual 7 A particles is represented by the spectra in Figures 14C and D. Thus, the 7 8, material in this verdine sample is similar to a serpentine-type clay. In order to provide the reader with a quantitative estimate of the element abundance for typical particles, these thin-film analyses are calculated as weight per cent oxides in Table 1. These data have been obtained by standard quantitative thin-film techniques as outlined by Lorimer (1987) and have a relative error of about 7%. Analysis A in Table 1 is from the 10 A material and shows a typical low Mg content (<1 wt YO),while analyses B, C, and D are from the corresponding analysis points (b, c, d) in Figure 10D (and spectra in Figures 14B, 14C, and 14D). As mentioned previously, analysis B is probably due to overlap of the two phases, but suggests that potassium may be present in the 10 A material at about the 1 weight YOlevel. Analyses C and D are from the predominantly 7 A material and are consistent with the bulk chemical analyses given in Table 2, p. 123. Table 1. AEM analyses of typical phases in verdine from French Guiana. Oxide

A

Weight percent* B C

D

SiOg

48.71

58.81

49.12

47.73

MgO

0.72

3.63

12.05

12.37

47.70

29.62

11.99

13.26

FeO

2.87

7.19

26.78

25.49

K20

0.0

0.75

0.05

1.14

A1203

* Normalized to 100%. Interpretation of AEM results Detailed AEM analyses ,of the verdine sample 508-175A show that three layer silicates as well as quartz are present in various proportions. The predominant species are a serpentine-type Mg, Al, Fe-rich clay (about 50% abundance) and a mica-type alumino-silicate (about 30% abundance). Smaller proportions of quartz (about 15%) and chlorite-type clay (about 5%) are also present. These different phases are not only distinguished by their fine-scale (sub-micrometre) chemistry and structure, but also by their distinctive grain-size and morphology. In general, these AEM analyses agree with the spatially less precise mineral identifications made using X-ray diffraction techniques and bulk chemical analyses. Both X-ray diffraction and electron microscopy show that chlorite is present only in very small amounts. AEM

123 emphasizes the presence and apparent abundance of a well-ordered mica-type phase which is remarkably poor in iron. For this reason, we suspect that this phase is probably different from the inherited altered biotite and should represent an authigenic phase.

Chemical study The chemical analyses undertaken in 1980 on the Guianese verdine were the fiist detailed analyses for Recent material of this type. Four purified samples of green grains were analysed in the Laboratoire de PCtrographie, Paris, by Miss M. Lenoble (sample numbers: 102, 169, 172, 175 in Table 2). At that time, a composition of berthierine or chamosite was expected. However, the internally very consistent results were in disagreement with typical analyses for berthierine. One sample (508-175A) and two fractions of another sample collected near sample 102 (sample 103, Table 1) were analysed in another laboratory (Ecole Normale SupCrieure, Paris) by Miss E. Lebrun. The results perfectly confirmed the earlier results by Miss M. Lenoble and these chemical data can be considered representative of the Guianese verdine. Compared with the data reported in the two preceding chapters, these analyses generally confirm that the green pigment of the verdine facies has a remarkably homogeneous chemical composition in spite of collection from three widely dispersed locations each with somewhat different sedimentary conditions (depth and maturity of the sediment). In addition, these locations have different types of substrate of verdissement, contributions of impurities, and processes of alteration; therefore, the analytical results precisely characterize the authigenic phase itself. Table 2. Chemical analyses of green grains purified from the continental shelf off French Guiana. An important proportion of the substrate of verdissement is made of chloritized biotite flakes; the potassium contents of the grains probably result partly from the presence of mica remains. Samples 103 and 175 were analysed twice in two independant laboratories for crosschecking. 102B

169B

1728

175A

175A

Si02

36.9

39.1

37.5

38.4

A1203

11.0

11.8

10.8

11.0

Fe203

17.9

'18.3

19.5

FeO

6.5

5.7

Ti02

0.4

0.5

103B'

hlean

38.0

37.0

36.7

37.7

10.6

11.7

12.3

11.3

18.0

19.6

18.3

18.2

18.5

0.6

6.1

6.3

5.6

4.9

5.1

5.8

0.6

0.5

0.5

CaO

0.6

0.6

0.5

0.7

0.7

hlgO

11 .o

8.3

8.9

9.9

11.3

10.4

11.0

Na2O

0.2

0.2

0.2

0.3

0.3

0.2

0.2

0.9

K20

1.4

1.3

1.3

1.5

H2O-

3.0

3.5

3.8

3.1

10.2

10.1

10.3

9.3

H2O'

(+C02) ~

Total

99.1

*

103B

12.6

-

-

-

-

99.4

99.4

99.0

99.6

0.26

lo

0.9 _+

0.6

0.25 9.9

f

1.2

1.15

1.13

1.3 2 0.1

1.9

2.1

3.1

12.2

11.7

98.0

98.7

0.7

1 0 . 4 2 1.0

124 In details, a slightly higher alumina content of 11-12% (A1,O ) in Guiana compares with 6-9% for verdine from New Caledonia and 3enegal. In contrast, the iron content is slightly lower in Guiana: between 18% and 19% (Fe203) and between 5.0% and 6.5% (FeO) compared to contents of 20-23% (Fe203) and 6.5-7.0% (FeO), respectively, for the samples from New Caledonia and Senegal. The magnesium content of the Guianese phyllite V (9 to 11%) is also lower than in Senegal and New Caledonia where it reaches 13% (MgO). Another difference in chemistry concerns the comparatively high potassium content (1 to 1.5% K20) in the Guianese green grains. This potassium content possibly results from an admixture of micaceous substrate, and may represent a minimum of 10% of the green grains analysed; the admixture of mica could also account for the higher alumina content. In addition, the observation of the 10 A phase using AEM suggests that the samples chemically analysed are not made of pure phyllite V but include a still problematic 10 A clay mineral. GEOLOGICAL SIGNIFICANCE OF THE VERDINE FACIES

The mineralogical study undertaken on the green grains from the continental shelf of French Guiana has shown that the largest portion of the studied area is the site for a neoformation of a green clay related to the verdine facies, as defined in the Neo-Caledonian and Senegalese deposits.

Age of the neoformation Due to the large palaeoclimatic and eustatic changes occumng during the most Recent history of the shelf, the geological significance of the verdine facies of the present deposit cannot be completely understood as long as the phenomenon has not been accurately dated. The following observations are relevant: 1) the verdine facies appears characteristic of high temperature sea-water at a shallow depth; 2) off French Guiana, the verdine is present between -20 m and -150 m depth or slightly more; 3) the presence of a palaeobarrier reef seems to be well established at about -100 m depth; 4) the green clay, locally initially formed in carbonate bioclasts, now forms free green grains which means that the carbonate has had time to dissolve; 5 ) traces of oxidation processes are frequent and are proof of a long history of the green grains after their genesis. In summary, the green grains from the Guianese shelf appear to be relict on the one hand. On the other hand, due to the ease with which the material is altered, it is difficult to imagine that the green clay could survive to the pre-Holocene regression. Moreover, the green grain-bearing fine sands are considered to be Holocene in age by Bouysse et al. (1977). Therefore, the time at which verdine could have been formed is bracketed between -18,OOO years and "some-time ago", possibly a few thousand years. The presence of green grains at depths of -20 m to -25 m indicates that these green grains must be younger than 6-8 thousand of years according to the known evolution of the sea-level on the shelf.

125 If we consider processes at greater depths, one has to remember and combine on the one hand, the presence of a reefal complex dated at about -15,000 years and, on the other hand, the case study described off New Caledonia. About 15,000 years ago, the situation off French Guiana appears very similar to the present one off New Caledonia. In these circumstances, verdine from the middle part of the shelf (mostly foraminifera1 infillings) could be interpreted as having formed in a lagoon of sub-reefal character. If this is assumed, the interpretation of the verdine lying deeper than -100 m (i.e., off the barrier-reef) is problematic. There is no argument against the possibility that verdine itself forms in front of a barrier-reef complex or deeper than where reef patches are present (see, Chapter B1, the area located to the SE of Baie du Prony off New Caledonia). In addition, the case study off Senegal shows that verdine may form without the protection of a barrier-reef. The sedimentological data obtained from the shelf off French Guiana as well as from other portions more to the west of this South American margin (off Columbia) indicate that there is a tendency for the shallower sediments, including carbonate remnants of palaeo-reefal origin and benthic foraminifers, to slip along the shelf edge (Pujos and Win, 1986). This phenomenon may account for the presence of verdine deeper than the palaeoreef. The above discussion indicates that the verdine might be of different ages on different portions of the shelf: about 15,000 years old in the deepest area, and younger than about 7,000 years in the shallowest area. Thus, the verdine would have formed diachronously following the transgression between about 15,000 and 6,000 years before the present. Another possibility is that the verdinization could have developed on the shelf during a shorter period of time between -10,000 and -7,000 years. This hypothesis would require the possibility that verdine forms at depth between a few metres and about -100 m (or shallower if we accept the transportation to deeper levels) and would account for the tenuous differences in the present aspect of the green grains as a function of depth. In fact, there are small differences: the shallower green grains appear "less evolved" with a chloritized micaceous substrate less altered than in deeper sediments; as a result, the first hypothesis (diachronous formation) appears the most convincing presently.

Recent history of the shelf The above study of green grains was very helpful for understanding the palaeogeographical and sedimentological history of the area off French Guiana. It is discussed here according to the most recent unpublished results. Three periods can be selected in the continuum of evolution: 1) the time when the regression was at its maximum 19 to 18,000 years ago, 2) the time when reefs were present 15,000 years ago, and 3) after the first reefs were killed about 10,000 years ago. Eighteen to nineteen thousand years ago, the sea-level was located near -100 m. The Amazon River ejected sediments into the deep ocean at the head

126 of a canyon cutting the outer part of the shelf (Milliman and Barretto, 1975). The narrow shelf, only about 10 km wide, was covered by sea-water with a temperature similar to the present sea surface, despite the colder climate. Figure 15 indicates the palaeo-valleys for the two main rivers of the area. The most recent sedimentological studies indicate that other rivers were also present. A general NNW course is suspected for all of these rivers and is similar to the present-day direction of the estuaries. The Maroni River, and also probably the other rivers, primarily bring coarse sands to the shelf. The NNW course of the easternmost river (Oyapock) in the area makes the eastern half of the shelf off French Guiana free of detrital local input and allows reef and ooids to develop possibly during the first time period and certainly during the following period. The coastal plain was occupied by an herbaceous vegetation (as shown by pollens) indicating a drier climate in Surinam (Roeleveld, 1969) or in NE Brazil (Caratini et al., 1975). Swamps and oversalted lagoons were present on the flat plain. Coarse sands were probably brought to that coastal plain due to the more efficient erosion. To the east of the palaeo-course of the Maroni River drawn in Figure 15, coarse sands are still visible today (Fig. 2) following a W-E direction. According to their heavy mineral contents they originate from the Maroni River. It is suggested that these sands were deposited in fresh-water lagoons parallel to the coast. No verdine was present, but glaucony was probably forming deeper than

Figure 15. Palaeosedimentology of the area near the shelf off French Guiana between -19,000 YWS and -18,000 years.

127

SURINAM

._

coast reef barrier

ooo

abundant glaucony possible verdine

... ooliths GSO 86

Figure 16. Palaeosedimentology of the area near the shelf off French Guiana about 15,000 years ago.

the present isobath -180 m due to the fluvial iron output and open marine environment. The very first reefal or sub-reefal facies (as well as the first oolitic facies) might well have been present during this fiist period, very near the coast as suggested by 14C radiometric apparent ages between - 14,000 years and -21,000 years measured on ooids off Brazil (Milliman and Barretto, 1975). It was suggested above that these apparent ages partly overestimate the true age of the sediment deposition and cannot be used as definitive evidence. Therefore, it might be considered that reefs and ooids were emplaced slightly later, after the beginning of the transgression, less than 17,000 years ago. Fifteen thousand years ago, the sea-level was near -80 m (Fig. 16). The Amazon River complex still ejected most of its waters and detrital sediments directly to the deep slope because it is assumed that the inversion of the coastal hydrologic regime occurred when the sea-level was higher at about -60 m (Bouysse et al., 1977). In spite of the sea-level rise, the shelf was only a little wider than in the previous situation: 10 to 20 km but not more. A narrow reefal-type lagoon was present. The barrier-reef was about 5 to 10 km off the coast and allowed growth of benthic foraminifers (with Amphistegina). The palaeo-valleys of the Maroni and Oyapock Rivers have been interpreted as far as the -70 m isobath from the heavy minerals study. The

128 palaeo-valleys show a general direction toward the north (15" E and about 20" W respectively) as shown in Figure 16. The southeastern side of the lagoon off French Guiana, allows the formation of an oolitic sediment. The main rivers brought ions to the western part of this lagoon which, however, showed a typical marine environment with relatively high temperature waters. This environment can be considered locally similar to the present one off New Caledonia. However, where they exist, the presence of ooids shows a slightly different situation with probably a lower depth, a higher temperature, and a higher cation concentration. Elsewhere, if the Neo-Caledonian model is applied, it is possible that the verdine facies was locally present at that time in the lagoon. Glaucony was forming contemporaneously on the shelf edge at depth higher than the present -150 m isobath (deeper than about 70 m at that time) but not immediately off the river output. 10,000 BPI

-

coast

+++ living reef

___

;SO 86

\

,.

..

dead reef glaucony forming verdine

Figure 17. Palaeosedimentology of the area near the shelf off French Guiana about 10,OOO years ago.

Ten thousand years ago the sea-level was about 40 m below the present one (Fig. 17). For about two thousand years the turbidic water originating from the Amazon River progressively invaded the Guianese shelf. The deeper reefs were initially killed by this fluvial output and the rapid sea-level rise. However, more recent reefs have possibly grown farther from the Amazonian output particularly off Surinam (Fig. 17). The shelf width was roughly enlarged during this period, the number of coastal rivers increased and the detrital sediment output

129 was very low. It is at this time that all favourable substrates were present at the sea-bottom: relict faecal pellets from the palaeo-course of the rivers, benthic foraminifers from the palaeo-reefal area, and mica flakes brought from the Amazon River. These substrates were necessary for the genesis of the phyllite V mineral. This situation developed until about 7,000 to 5,000 years ago at low depths when the surface water temperature was high enough at the sea-bottom. Later on, the environment was modified and good growth conditions for phyllite V were not available. On the contrary, green material was frequently submitted to an oxidation process due to the lack of burial of the fragile green clay formed. This general oxidation process may be linked with a too large winnowing occurring on the continental shelves where the open sea freely circulates. The same general scheme may be applied from the Amazon River mouth until the present Orinoco estuary because the same authigenic minerals have been described off Surinam by Hardjosoesastro (1971) and off Guyana and Venezuela by Porrenga (1967a) and Morelock (1972). While the situation off the northern coast of South America 15,000 years ago is probably similar to the present Neo-Caledonian model (a comparatively narrow reefal lagoon), the situation between -10,000 and -5,000 years is comparable to the Senegalese model (a tropical wide shelf where the sea-water is sometimes too cold or too turbidic for reef growth). As a final remark, it must be emphasized that the size of the phenomenon is very different in these three deposits. Off New Caledonia, the localized basins where verdine is known to form are small patches about 500 km2 in area. Off Senegal, the presently known verdine-bearing basin is an order of magnitude larger in area. Off South America, the area probably covered by the verdinebearing sediments is another order of magnitude higher being probably between 50,000 and 100,000 km2. This deposit is most probably the largest in the world for the facies considered in this section. SUMMARY AND CONCLUSIONS

Green grains are ubiquitous on the shelf and the top of the slope off French Guiana. Two sedimentological facies exist: glaucony deeper than -150 m, verdine at shallower depths. The verdine facies.has mainly developed from mineral debris and especially chloritized biotite. Carbonate bioclasts and faecal pellets are also utilized. The mica flakes were never wholly replaced by authigenic clay and the phenomenon leads to mixed grains where authigenic and substrate remains are recognizable. Carbonate substrates lead to mainly clay pure green grains since the initial carbonate has been dissolved. The authigenic mineral found on the shelf is the typical phyllite V as shown by X-ray diffraction (main peak at 7.2 A and low or dubious 14 A peak) and by the chemical composition showing a comparatively high silica and ferric iron and low alumina contents. The AEM studies have shown the composite nature of the green grains of the verdine facies. The 7 8, phyllite V is predominant; a

130 noteworthy proportion of the grains is made of well-ordered clay minerals with a 10 A spacing. In contrast to other minerals from the verdine facies, the 10 A clay is iron-poor. The irregular presence of a 14 A peak on X-ray powder diffraction patterns has been shown to be mostly related either to the presence of remnants of the inherited chloritized mica (the substrate), or to the presence of glauconitic smectite. However, 14 A spacings are also present as authigenic marine phase in few mixed-layered structures and as more abundant pure 14 A chlorite-type clay minerals as observed during the AEM study. The study of the green grains allows us to draw conclusions on the main changes in the palaeogeography of the shelf for the Holocene time. The verdine is relict today and has been formed diachronously between -15,000 years and -5,000 years and perhaps during a shorter period between -10,OOO years and -7,000 years. The formation of verdine can be located in a general marine environment, at a comparatively warm sea-water temperature, and at a depth probably shallower than -60 m. The intertropical latitude is concerned as with all the other known deposits of this facies (see Chapter B4). The verdine genesis seems to be fundamentally influenced by the variation of the Amazon River output both in terms of timing and environment. On the whole, the presence, nature and geological significance of the phenomenon appear to be similar for the whole northern shelf off South America from the Amazon River mouth to the Orinoco River mouth. This makes this verdine-bearing deposit the largest of the Recent continental margins. ACKNOWLEDGEMENTS

We are grateful to Dr Odile RCniC and Dr Michel Chagnaud, UniversitC de Bordeaux, for their help in the sedimentological study. We thank Professor H. A. Geyh for his valuable 14Cdating of the Guianese ooids and comments, and Miss E. Lebrun, Ecole Normale SupCrieure, Paris, and Miss M. Lenoble, DCpartement de PCtrographie, UniversitC Pierre et Marie Curie, Paris, for their accurate chemical data.

131

Chapter B4

THE VERDINE FACIES DEPOSITS IDENTIFIED IN 1988 by G.S. Odin, J.P. Debenay and J.P. Masse INTRODUCTION

The history of the knowledge of the facies here identified as the verdine facies is relatively short. It began twenty years ago when, the same year, three different authors, in three different areas, showed fairly clearly that what would otherwise have been called glaucony (green grains composed of a smectitic to illitic ferric marine clay) was in fact composed, at least partly, of a different authigenic clay mineral. The first reaction of all the authors was to identify the clay minerals in these green grains with those of an already known facies: ancient sedimentary ironstone. The terms "berthierine" (French) or "chamosite" (English) were used to name the green mineral of those newly discovered, Recent, green grainbearing sediments. The use of these terms in the literature allows us to presume that the authors were describing the presence of verdine facies in spite of the fact that, at that time, the specificity of the facies was not recognized. It was somewhat unfortunate that the ancient ironstone minerals were already well known because it was tempting to use the terms "berthierine" or "chamosite". This, in turn, obscured the fact that not only was the mineral in question distinctively different, but also, in consequence, that the correspondingfacies had a different geological significance. The present chapter summarizes the main characteristics of the verdine facies deposits which show 1) that the facies is common and extensive on earth, 2) that several characteristics of the facies very similar to those examplified in the three preceding chapters can be found in the other regions, and 3) that the geological significance of the facies can be defined from the factors common to all deposits. The order of presentation of the regions below will follow the chronology of the first description in the literature. THE OGOOUE RIVER MOUTH

(Gabon)

Literature Comparison of the dates of publication of the three papers published in 1965, shows that the first Recent green grain-bearing deposit for which the word "berthierine" has been used is located in the immediate proximity of the mouth of the Ogooue River (Gabon). Giresse (1965) quotes shallow green grains for which he writes 'la "glauconie"... n'est pas une glauconite au sens illite fenifere, mais un melange de berthikrine (50%) dinterstratifit5vermiculitechlorite et de quartz'. Caillkre and Giresse (1966) confirm that the green sands

132 off the Ogooue River mouth are mainly made of green clay showing a 7 8, X-ray diffraction peak. Further data were gathered by Giresse (1969) but all the mineralogical results obtained at that time concerned mixtures of various demtal and authigenic minerals frequently oxidized. More precise results on carefully purified and selected grains, have been accumulated since 1971 by the editor and partly published by Odin and Giresse (1972) and Giresse and Odin (1973).

Morphological data The main substrate of verdissement observed on the Ogooue delta is ellipsoidal faecal pellets (see Fig. 7). Usually, a sediment contains fair sorted pellets, locally small (about 100 pm in length) but elsewhere larger (around 500 pm in length). Frequently, the dark-green colour of the glossy grains grades from brownish to dark-brown, at the surface of the pellets or throughout. This aspect helped to increase the confusion with the ancient, oolitic sedimentary ironstone facies. The question was even more obscured by the presence of granular, iron oxi-hydroxides, which externally resemble ooids (Giresse, 1969). These superficial ooids are considered as faecal pellets (mineralized to a green femc clay), oxidized later, and to which one or two layers of goethite have been added by chemical precipitation on in situ reworked grains. In addition to the faecal pellets, about 10% of the magnetically separated fraction is composed of green infillings, sometimes with remains of their carbonate tests (see Fig. 7). Finally, a small proportion of the quartz grains of the sediment is encrusted with a green pigment, which sometimes penetrates into the grain along cracks. The same phenomenon was observed in other areas (off French Guiana, Chapter B3).

Mineralogical data The faecal pellets were initially formed of a grey clay mixture with kaolinite dominating (Giresse and Odin, 1973). This fact, allied to the frequent oxidation of the magnetic pellets, makes the mineralogical characterization of the green component difficult. However, the X-ray diffraction study of series of pellets: more or less magnetic, more or less green, or more or less dense, allows the diffraction peaks of each component to be distinguished i.e., substrate, authigenic marine mineralization, and alteration product. Four representative X-ray diffraction patterns of the purest green grains available were published by Giresse and Odin (1973). They show the usual diagram of phyllite V with a diffraction peak at 7.2 A dominating (Fig. 4, diagram 1). Admixtures of quartz and goethite are visible; remains of kaolinite are also probably present, in spite of the selection of highly magnetic fractions for X-ray analysis. The chemical analyses of pellets published by Giresse (1969) are varied and can hardly be used for mineral characterization because of the complexity of the mixtures analysed. Mineralogical compositions calculated by this author assume the presence of chlorite, vermiculite, 'berthierine', goethite, 'glauconie',

133 quartz, and calcite. Moreover, because the presumed authigenic 7 8, phase was erroneously accepted to be ferrous while it is probably ferric, the interpretation is generally very speculative.

Distribution of the suspected verdine facies The studied area is located between latitudes 0" and 1" S . The precise distribution of the facies is difficult to assert since, at the time of the study, ten years ago, the authors were much less familiar with the subject than at the present time. However, Figure 1 proposes a general map based on X-ray diagrams from half a dozen samples.

t N

Figure 1. Distribution of green grains off the Ogooue River mouth. (1) verdine facies; (2) glaucony. The curved arrows indicate the main fluvial inputs into the sea.

134 Verdine is present near the mouths of the river at depths as little as -5 m. The deepest verdine was collected from a depth of about -40 m where magnetic grains were available in the very restricted area towards the north. Towards the west of Mandji Island, formed of Recent sands brought down by the river and deviated northwards by a longshore current, the verdine facies has been identified in a single sample. W e lack the data needed to know whether the verdine facies extends on both sides of that area or is restricted to the immediate proximity of the local river inputs into the sea. In samples from deeper than about 50-60 m, the magnetic green grains contain a mineral referred to as glauconitic smectite. For that reason, Giresse, (1969) wrote that "berthierine" was a preliminary stage of the glauconitization process. This hypothesis, however, was not based on detailed analyses and is now rejected. Looking at the present depth distribution of verdine, the age of the mineralization is probably between -7,000 years and -3,000? years. Giresse (1969) insists on the fact that the area where verdine is present today is covered by hot temperature waters similar to those present in the southern and the northem areas where verdine is absent; however, where verdine is present, the pH appears lower (7.5) than elsewhere on the shelf, because of river water input. Finally, verdine seems to be related to areas where no terrigenous deposition occurs. The deposit off the Ogooue River mouth would need to be submitted to further research for a better knowledge of the verdine environment. THE ORINOCO RIVER MOUTH AND EASTERN EXTENSION (northern

South America) Literature Porrenga, (1965) recognized the presence of "non-glauconitic" green grains off the Orinoco River mouth and in western Guyana. These grains were known ten years before his own study was made but, in the absence of detailed analyses, the green grains were confused with glaucony. A little more data were discussed by Porrenga (1967a).

Morphological data Most green grains occur as faecal pellets; however, foraminifer and ostracode infillings are also common. There is a transition from dark-green grains and externally brownish grains, to totally brown grains due to oxidation. The green grains usually appear free of the material of the initial substrate and are therefore evolved and probably relict.

Mineralogy No mineralogical analysis has been published, but Porrenga (1967a) writes that the green grains are similar to the grains collected from the Niger delta. He called them "chamosite" (see section Niger delta, p. 136 below).

135 Distribution The distribution of Porrenga's "chamosite" is shown in Figure 2 simplified from Porrenga (1967a). The map shows the facies to cover an area over 600 km in length and 20 lcm to 60 km in width. The 7 A mineral mainly occurs at depths from -60 m to -150 m. Porrenga notes that the absence of shallower "chamosite" is probably related to rapid burial by sediments coming from the Orinoco River. Deeper sediments contain true glaucony locally. The coasts of eastern Venezuela and western Guyana are very flat and swamps are present nearly everywhere. In our study of French Guiana, it was shown that the pattern of distribution observed off the Orinoco delta was similar to distribution off Surinam, French Guiana and Brazil. Most of this area eastward of the Orinoco delta is sedimentologically dependant on the Amazon River influx. But it is clear that the Orinoco River itself helps, or helped, to create the favourable conditions for verdine growth at the western end of the area under discussion, which is the largest verdine-bearing deposit to be identified in contemporary oceans.

Figure 2. Green grains off the Orinoco River mouth. (1) swamps; (2) area where the verdine facies is known to exist; (3) glaucony; (4) brown grains inteqreted as oxidized verdine. Note that this deposit is at the westemmost end of the area discussed in Chapter B3.

136 THE NIGER DELTA (Nigeria)

Literature Porrenga, (1965; 1967a) analysed and identified "charnosite" in Recent sediments collected from the Niger delta and detailed results were obtained.

Morphological data Most of the green grains collected from the area derive from the mineralization of faecal pellets. Scarce infillings of foraminifers and ostracodes exist. The green grains are locally more or less deeply oxidized.

Mineralogy Porrenga (1967a) gives an X-ray diffraction pattern of a Recent "chamosite" from the Niger delta. The figure completely resembles the usual diagrams shown in the three preceding chapters (Bl, B2, B3), with a dominant peak at about 7.2 A. A very small, broad diffraction peak exists at 10 A and nearly nothing exists near 14 A; there are admixtures of quartz and goethite. It is interesting to note that, because this author assumes that he observes a '7 A chamosite', he seeks, and finds, a reflection at 1.54 A, which is really not very clear on his diagram. For him, this reflection is linked to the assumption that the green mineral must have a trioctahedral structure (magnesium and ferrous iron in octahedral sites). Curiously, Porrenga compares the X-ray diagrams and chemical data of the Nigerian green grains with ancient Algerian chamosites, and he systematically underlines the similarities without insisting on the differences clearly visible in his figures. Two chemical analyses are given by Porrenga (1966), but the quantity of material available did not allow the analyst to distinguish ferrous from femc iron. However, the author assumes the iron to be in the ferrous state. The very high silica content (Table 1, below), which corresponds to the presence of quartz should be noted. Finally, the total iron content appears low compared to our own analyses, probably due to the quartz admixture. It is not possible to calculate a chemical formula from these results unless a quartz admixture of about 20% is assumed. Table 1. Chemical analyses of "chamosite" from the Niger delta (Porrenga, 1966). Samples

heated to 700'C before spectro-chemicalanalyses. SiO2 R ecen t Miocene

52 46

A1203 F e z 0 3 8

13

20 23

MgO

CaO

Na20

K20

H20'

8.3

0.5

0.3

<0.5

11

0.4

<0.5

11

4.7

0.5

137

Figure 3. Green grains off the Niger River mouth. The verdine facies is restricted to depths shallower than -40m (thick dashed line). Within the area where verdine is present, the densest dots indicate a proportion greater than 2% of the total sediment.

Verdine distribution off the Niger delta A simplified version of the distribution of the verdine facies is shown in Figure 3 based on the detailed map published by Porrenga (1966). Verdine is present at depths generally shallower than 30-40 m. Below -50 m to -100 m and down to -200 m, glaucony is present, and is developed in substrates similar to. those favourable for verdine formation. At depths shallower than - 10 m all pellets are brown. This demonstrates generally good depth-dependent zonation and leads to the deduction that verdine is less than about 10,OOO years old in the surface sediments. However, Porrenga also writes that similar green grains ("chamosite") are found in Miocene sediments, drilled 2,300 m below the sediment surface; he concludes that, if there is a genetic link between the ironstone-type chamosite and the facies found today on the sea-bottom, the diagenetic process needed to modify the latter into the former better crystallized form happens either at greater depth or at higher temperatures. From an environmental point of view, Porrenga notes that the distribution of verdine facies is related to the thermocline present on the shelf. Shallower than -40 m, there is a water mass at a temperature of 25"-27"C; deeper, the water mass is 10" to 15°C colder. Therefore, Porrenga concludes that this temperature control is the fundamental factor which distinguishes favourable conditions for the 7 A mineral growth (temperature above 25°C) from conditions favourable to glauconitization (temperature at about 11 f 4°C).

138 THE KOUKOURE RIVER MOUTH (Guinea)

Literature and new results The third study published in 1965 in which the new formation of Recent "chamosite" is assumed, is by Von Gaertner and Schellmann. This study was often quoted later as an example of Recent "chamosite" formation. In fact, until very recently, that interpretation was considered dubious by the present authors mainly because the area is in the close vicinity of volcanic and plutonic outcrops including ultra-basic rocks. In a similar situation "berthierine" was identified off northwestern Spain following X-ray diffraction and chemical studies (Caillhre and Lamboy, 1970a). However, our more detailed later studies, including thin sections and X-ray diffraction, showed that this Spanish "berthierine" was the result of an alteration of the material derived from near ultra-basic rocks; thus, it was re-interpreted as a serpentine (Odin and Matter, 1981) because of its aspect resembling a stained-glass window and the very sharp peaks shown by the X-ray diffraction patterns. Fortunately, two samples from an area less than 10 km away from the one where the single highly oxidized sample analysed by Von Gaertner and Schellmann was collected, were provided to the authors (end 1986) by C. Moreau and B. Robineau (University of Dakar). They allowed us to c o n f m the Recent, marine, authigenic nature of the material and to identify it with the other materials considered in this part of the volume.

Morphological data Von Gaertner and Schellmann (1965) consider that their "chamosite" develops from substrates formed of goethite, which indicates that what they observed was deeply oxidized "green grains". A picture of a thin section in their paper shows that they studied faecal pellets magnetically separated from fine sand. Our two samples are also fine to very fine sands; about 10% is made up of magnetic green grains. Eighty to ninety percent of this fraction is ellipsoidal dark-green faecal pellets, about 100 pm in length. Infillings of microgastropodes constitute most of the rest of the green grains. Thin, white, carbonate test remains, although partly dissolved, are still present. A small proportion of green flat flakes, is present: probably biotite used as a substrate for the development of the green marine mineral. No alteration into goethite was observed in these two samples, allowing the collecting of significant mineralogical data.

Mineralogical data

-+++

(5) sample 601-95 from the Senegalese continental shelf (Chapter 2) for comparison;

(6) sample from the Niger delta redrawn after Porrenga (1967a); (7)sample from Sarawak redrawn after Porrenga (1967a). Note the great similarity of all of these diagrams obtained from green grains collected from areas far apart; they represent a specific widespread facies.

139

Figure 4. X-ray diffraction patterns of verdine from different areas. (1) sample 292 from the Ogooue delta; (2) sample 699 from the Koukoure River mouth (locality 2 in Fig. 5); that sample was chemically analysed (Table 2); (3) sample 498 off the NDogo fresh water lagoon (see Fig. 8); (4) sample 467 off the Comoe River (Ivory Coast); (caption follows p. 138)

....

140 Von Gaertner and Schellmann (1965) write that their magnetic fraction is a mixture of goethite and clay. The latter has a kaolinite structure according to X-ray diffraction. Our own X-ray diffraction study shows that the pellets are composed of a phyllite V similar to the one analysed in the Neo-Caledonian reefal lagoon. The base of the peak at 7.2 A tails off slightly to lower angles up to about 10 A. A low hump between 14 A and 20 8, is present; no obvious goethite peaks are visible (Fig. 4). The chemical results published by Von Gaertner and Schellmann (1965) are shown in Table 2. The analysis (1) gives the rough results. These authors recalculated the composition (2) by eliminating what they presumed to be goethite according to the high Fe,O, content assuming that the authigenic component was "chamosite" i.e., a 7 A clay with Mg and ferrous iron as the main octahedral ions. This recalculation was biased by that hypothesis, and they obtained an SiO, content much too high for that kind of mineral i.e., for a berthierine (see, for example, analyses collected by Brindley, 1982, where the SiO, content is always clearly lower than 30%). This recalculation (2) is quoted alone by Porrenga (1967a) in order to confirm that the recent green mineral from the Niger delta is actually a chamosite. It can be seen that, by this circular reasoning, a hypothesis of recalculation is taken as a fact which leads to postponing the real understanding of the significance of the Recent, shallow, green marine clay. A sufficient quantity of relatively pure material was prepared from one of our samples. The purification was improved using bromoform, N/10 cold acetic acid and magnetic fractionation. The selected fraction was a mixture of variously coloured grains (dark-green to light-green) with an earthy aspect. Traces of greenish mica were also present, but no oxides were observed. The chemical analytical results are given in Table 2 (analysis 3). All cation contents are very similar to those observed in green grains collected from other areas like New Caledonia, Senegal and French Guiana. Therefore, phyllite V appears to be a chemically homogeneous component indicating a remarkably similar genesis environment in these four geographical areas scattered from the Pacific to the Atlantic Ocean. Table 2. Chemical analyses of the verdine from Guinea. Si02 ( 1 ) 27.4

A1203 Fez03 9.82

31.8

( 2 ) 42.4

16.4

6.7

( 3 ) 35.96

12.18

19.48

FeO 5.96 10.0 6.21

.

MgO

CaO

Na2O

K20

Ti02

P2O5

MnO

HzO-

H20t

7.22

0.86

0.18

0.59

0.32

-

0.36

3.55

11.7

-

-

0.13

tr

4.10

10.96

12.0 9.72

0.35

0.37

0.17

0.33

(1) magnetic fraction analysed by Von Gaertner and Schellmann (1965) and (2) their recalculation assuming that the fraction was a mixture of "7 A chamosite" and goethite; (3) pure green grains, Los Islands; analysis by Madeleine Lenoble (X-ray fluorescence and wet chemistry, unpublished).

141

Distribution and environment of the Guinean verdine Von Gaertner and Schellmann (1965) indicate that their analysed sample comes from 600 m off "Plage PCronne" (Conakry). The surface sediment is oxidized at shallow depth. Our two samples were collected from inside the Los Islands, to the north of the central island called Roume. This complex of islands is of volcanic origin; its petrography, mainly a nepheline syenite, was recently studied by Moreau et al. (1986). The area is under the influence of a local stream (Fig. 5) deviating the Koukoure River input southward.

N

f i \- t

C

o

r 9'30"

1'

loKm 'GS0.87

&\ 0

'

o

a

n

1

Figure 5. The verdine facies off the Koukoure River mouth. (1)location of the sample analysed by Von Gaertner and Schellmann (1965); (2) location of the two samples recently collected and analysed here; (3) ultra-basic rocks and gabbro. The small map of Africa indicates the location of the present area (F5) as well as those of Figure 8 (F8) and 9 (F9). The Los Islands are those islands surrounding site number 2. The whole area shown is shallower than -20 m depth.

From the environmental point of view it is important to note that the verdine-bearing samples were collected by hand in less than one metre of water at low tide. This makes this locality the shallowest known to date (see also the

142 Casamance Estuary, below), and the absence of oxidation very surprising. This shallow depth implies that the verdine formation is less than 5,000 years to 7,000 years old because, before that time, the sea was not present there. The good preservation of these grains in this tropical area, where the syenites of the nearby islands are altered to bauxitic laterite, probably implies that the verdine is either forming today, or was formed a few hundreds of years ago. The shelf off Guinea is very flat and shallow, and the shore is bordered by recently formed tidal marshes with mangrove swamps on the bank of numerous tidal channels. The whole area shown in Figure 5 is shallower than -20 m and can be considered as a large estuary due to the numerous coastal rivers. The temperature of the superficial sea-water is nearly constant (about 27°C all the year round). Salinity reaches its maximum (37.8%0, as in the open sea) at the end of the dry season, from April to May; during the rainy season, the coastal rivers lower the salinity of the superficial sea-water to an average minimum of 21%,, and exceptionally down to 9%0, and bring turbidic waters to the Los Islands (Uschakov, 1970; Debenay and Konad, in press). The topographic organization of the volcanic Los Islands looks very similar to a reefal atoll. The precise distribution of the verdine-bearing sediments off Guinea is unknown. Because the green clay forms in a shallow area where sample collection is easy, the Los Islands would be a good possibility for further research. However, it is already known that sedimentologically and mineralogically, the verdine has a geological significance similar to that of the other known deposits: recent, neoformation of a green clay mineral at shallow depth, in the sea not far from a fresh-water input in a tropical area. THE VERDINE OFF SARAWAK

(N. Borneo, Malaysia)

Literature The first and only study devoted to the green grains of the surface sediments of the very wide shelf off Sarawak was published by Porrenga (1967a); the observations below are taken from that precious study. There, again, Porrenga called the green grains 'chamosite' and assumed that their composition and significance were similar to those of ancient sedimentary ironstones.

Morphological data According to Porrenga (1967a), the main substrate of verdissement consists of infillings of microtests in most of the eighty sea-bottom samples submitted to the study; faecal pellets are also present. The colour is pale yellowish-green, and the aspect is earthy in the shallowest sediments; deeper than -20 m, the grains become blackish-green. At the same time, the carbonate tests disappear, but the form of the infillings still allows recognition of the initial substrates. This shows that the deeper grains are older and more evolved than the shallower ones. Brownish grains are present in the deepest sediments as a result of an oxidation process.

143

Mineralogical data X-ray diffraction patterns of the green grains of the shelf off Sarawak are given by Porrenga (1967a). They are similar to our own diagrams, with a dominant peak at 7.2 8,. The occurrences of a weak reflection at 10 8, can be noted as well as a diffuse dome at about 14 8, (see Fig. 4). The sample was submitted to heat treatment; the corresponding X-ray diffraction patterns show a slight lowering of the 7.2 8, peak at 350"C, a more pronounced lowering at 450°C, and the disappearance of the 7.2 8, peak at 550°C;there is no noticeable difference for lower angle reflections after heat treatment. This behaviour indicates that the mineral from the shelf off Sarawak is generally similar to our phyllite V and especially to the one from New Caledonia.

7 CHINA

Figure 6. Verdine from off Sarawak (Borneo Island). (1) samples from depths exceeding -60m without green grains; (2) swamps; (3) area sampled at shallow depth, all samples with verdine; (4) area where verdine develops on chloritized biotite flakes.

144 As a complement to these analyses, Porrenga (1967a) also gives X-ray diagrams showing a phenomenon similar to that observed off French Guiana i.e., the growth of phyllite V within chloritized biotites. This phenomenon is particular to a limited area of the shelf (see Fig. 6). Chemical results are available (Table 3). Compared to our data for verdine, a high silica content may be noted, probably related to insufficient purification of the grains with which quartz is admixed. But the main surprising result concerns the iron, which is given as mainly ferrous. If correct, this analysis would show the only example of ferrous phyllite V; nevertheless, further results must be waited for checking this hypothesis. Table 3. Approximate chemical composition of the grains collected from the shelf off Sarawak. Sample 8782: green grains with quartz admixture; sample 8878: brown grains.

(According to Porrenga, 1967a) Sample

SiOz

A1203

8782

49

9

8878

36

11

FeO

CaO

16.94

0.4

31 ( t o t a l )

0.9

Fez03 4

MgO 10 3.4

NaZO

KzO

Ti02

RlnO

0.3

0.5

0.3

0.2

0.6

0.5

0.4

0.2

HZOi 9.3 16

Distribution and environment of the verdine off Sarawak Verdine from the area concerned has been found in all samples collected at depths ranging from less than -5 m to -60 m. Between -60 m and -170 m depth verdine is rare or absent; deeper, no green pigment is visible. The aspect of the green material is depth dependent, and indicates that the genesis of the material has probably immediately followed the last transgression. The genesis appears therefore, as diachronous and generally younger than about 15,OOO years. The area where verdine is found is restricted to between 3" N and 4" N (Fig. 6). The nearby continent has a very high rainfall, and abundant continental waters are delivered to the shelf. The occurrence of a flat coastal margin with swamps probably avoids the presence of detrical clay and sand material in the waters reaching the South China Sea. As a result, corals are present on the northern coast of Borneo; this also indicates a constant, relatively high temperature of the sea. That situation is generally similar to what is observed in the other areas described above. It is not now known whether the verdine facies extends northward or southward of the intensively sampled area studied by Porrenga in 1967. But southward of Borneo, between that large island and the Celebes (Sulawesi), the Makassar Strait has been investigated using several series of samples from the shelf (samples provided to the editor by Dr. Gayet, UniversitC de Bordeaux). However, most of these samples were collected from depths of -100 m or more. A number of these samples contain a green clay, usually filling tests of foraminifers; all this material must be related to glaucony. Finally, Porrenga quotes a personal communication by G.H. Keller who might have observed "chamosite" in the Strait of Malacca, south of the Malacca Peninsula and north of Sumatra (Fig. 6).

145 THE CONGO RIVER MOUTH

Literature The shelf between the Congo River and the Ogooue River was studied by Giresse and his co-workers. Few analyses are available concerning verdine because the largest portion of that area is essentially spectacular for its glaucony-bearing sediments (Odin and Giresse, 1972; Giresse and Odin 1973; Giresse, 1976a and 1976b). From the abundant material collected, few samples have shown the presence of verdine. Some results were published in Odin (1975) and Odin and Matter (1981). These results are summarized below together with new results obtained in the course of the present study.

Figure 7. Green grains mostly made of phyllite V. To the left (sample 293 from the area north of the Ogooue delta) the green clay impregnates faecal pellets initially made of kaolinite, but also carbonate bioclasts: bivalve fragments, and fills ostracode tests. To the right (sample 449 near 448 in Fig. lo), mineralized faecal pellets. Note that no cracks are present at the surface of the variously evolved grains. Black bar is 1 mm long.

146

Morphological data The green grains from the whole shelf are very similar to those initially observed from the Ogooue submarine delta: they are dominantlyfuecal pellets. Usually, they have a glossy appearance, and the colour varies from dark-green nearly black, to brown. In fact, most of the grains appear oxidized. However, the less oxidized magnetic grains may usually be purified by using bromoform (oxidized grains sink). Figure 7 shows two pictures illustrating green grains from the Gulf of Guinea off the Ogooue River mouth (to the left) as well as off the Congo River mouth (to the right).

Mineralogical data

I -jx

CONGO River Mouth -0

d

-0 Y

GSO 87

KdCo

Figure 8. X-raydiffraction patterns of more or less magnetic green grains collected from off the Congo River mouth (sample 449-near 448, Fig. 10). The current intensity of the elecuo-magnet is increased from the top diagram (0.35 A) to the bottom diagram (0.75 A) allowing first the most magnetic (verdine-rich)grains and finally the less magnetic (verdinepoor, kaolinite-rich)grains to be separated.

The mineralogical properties of the verdine observed on the shelf presently being studied are difficult to observe. One reason is that the substrate of verdissement is itself mostly clayey and the dominant clay is kaolinite. Therefore, the X-ray diffraction patterns show a more or less sharp 7 A peak following the less or more evolved stage of the grains as shown by Odin and

147 Matter (198 1, Fig. 9); the more green and magnetic the grains, the lower and wider the 7 A peak. Figure 8 shows a series of X-ray diffractograms from low magnetic grains (grey coloured) to very magnetic grains (green coloured). The difference in the X-ray diffraction patterns obtained is small; in fact, all faecal pellets are mixtures of authigenic green clay with demtal kaolinite. Concurrently with the decrease of the 7 A peak height, there is a widening toward the low angles. In the absence of good criteria by which to estimate whether kaolinite was present or not in the most magnetic fraction, a study using infra-red spectrometry was undertaken by F. Frohlich. This author showed that the mineral of which verdine is made and kaolinite have very different absorption spectra (Chapter B5) which makes this technique ideal for judging the purity of the verdine from the Gulf of Guinea deposits. Figure 9 shows the IR spectrum of the purest fraction of verdine selected from sample 1150 collected from off the Congo River mouth. The quantitative method developed by Frohlich (personal communication) allowed him to suggest from this spectrum, and by comparison with a spectrum of verdine entirely composed of the phyllite V mineral, that the analysed green grains contain 83 f 5% of phyllite V, 1 1 f 1% of kaolinite, and very approximately 3% of quartz. CONGO 1150

3400

3000

1500

1000

500

cm-1

Figure 9. Infra-red spectrum of a verdine collected from off the Congo River mouth. The bands, or portions of bands, probably due to kaolinite are shown in black; quartz (Q) is also present. (Spectrum courtesy of F. Frohlich)

A second reason why the precise mineralogical properties of the verdine from the area under study is difficult to define is mainly puzzling when chemical analyses are undertaken, because it is linked to the very general oxidation process which increases the Fe203 content and lowers the content for all other cations (Table 4). If we simply lower the Fe20 content of the analysis from about 30% to 20% (the usual content for phylhte ) and split the missing

zr

148 10% by increasing the other cation contents proportionally to total 100%, then, the contents obtained are remarkably similar to the chemical data available from other areas such as New Caledonia, French Guiana, Guinea or Senegal. The mineral phase involved is therefore our phyllite V. It may be noted, however, that some admixtures of mica or illite (presence of K,O) and of carbonate are present. Table 4. Chemical composition of the faecal pellets from the shelf off the Congo River mouth (see sample location on Fig. 10). The dark-green glossy pellets are partly oxidized. SiOz

A 1 2 0 3 FepO3

FeO

MgO

COO

NO20

K20

Ti02

HzO-

H20'

G448

32.7

11.7

30.2

4.1

4.1

1.2

0.2

1.1

0.4

3.8

10.5

PG 1150b

31.6

9.5

30.7

4.75

7.0

1.1

0.1

1.0

0.45

3.2

10.3

corrected*

37.0

12.2

(20)

5.1

6.3

1.3

0.2

1.2

0.45

4.0

12.0

* corrected mean values assuming that F%03 must be 20%. Distribution and environment of verdine between 0" and 6" S We have undertaken X-ray diffraction analyses on about fourty of the numerous samples available. The resulting distribution is shown in Figure 9. For most of the shelf, glauconitic minerals are identified or suspected in sediments deeper than -80 m even to -1000 m (Bongo Passi, 1984). But, immediately off the Congo River mouth, phyllite V is present or suspected to the north in a small area about 70 km long and 25 km wide. This must be connected with the fact that a general surface current moves northwards the ions and particles brought into the sea by the Congo River. This fact is well evidenced by the distribution of mainly kaolinitic clays transported by the river, which form a large tongue of ancient muds from -10 m to -80 m depth on the shelf, to the north of the river mouth, extending to about 4" S (Giresse and Kouyoumontzakis, 1973). Obviously, there is a link between the distribution of the verdine facies and the Congo River input to the sea, which makes this area dependant of that river, and therefore sedimentologically and genetically distinct from the Ogooue delta. Between these two areas (Ogooue and Congo River mouths) about 550 km apart, magnetic faecal pellets shallower than -80 m are either absent or do not show the morphological and mineralogical characteristics of the verdine facies. There is a noticeable exception to this picture in the area facing the N'Dogo lagoon. There, another series of samples has shown the presence of verdine in the surface sediments. This third and newly identified area could indicate the Occurrence of a palaeo-delta corresponding to the presently dried up river which fed the N'Dogo lagoon. For this third area, as well as for those of the Ogooue and the Congo deltas, the verdine is certainly younger than 10,000 to 15,000 years. But the general oxidation also indicates that conditions favourable for

149 verdine genesis ended some time ago. It is possible that other small patches of verdine-bearing sediments are present on that shelf, which would need detailed study from that point of view.

Figure 10. Green grains distribution between the Congo and Ogooue River mouths (partly after Giresse, 1976a). (1) area where more than 5% of the sediment is made of unidentified green grains, or of identified verdine (2) and glaucony (3) facies. Green grains deeper than -80 m can be considered to be glaucony. Verdine is presently known from three distinct areas: off the Congo River mouth, off the NDogo Lagoon and off the Ogooue River mouth. THE SHELF OFF IVORY COAST

(Comoe River and other rivers)

Literature "Berthierine" is quoted from the shelf off Ivory Coast by Martin (1970; 1973). In his sedimentological mapping of the shelf, this author has tried to compare his area of study with the previously described shelf off Nigeria. However, some difficulties occur in this comparison. For example, the identification of "berthierine" was again problematic because of the presence of kaolinite as the main substrate for the shallow verdissement. Louis Martin was kind enough to provide the senior author with some material from which it has been possible to obtain results comparable to the ones obtained from other verdine-bearing sites.

150 Morphological data Like everywhere in the Gulf of Guinea, the very dominant substrate of verdissement is faecal pellets. Most of the relatively shallow, submarine areas contain oxidized grains, which leads Martin to write that, compared to the Niger delta, 'La zone B berthikrine pure n'existe pas' (there is no zone with pure berthierine). Instead of that, there is only a shallow zone of oxidized pellets directly followed, on the deeper side, by glauconitic pellets.

Mineralogical data Martin (1973)has tried to observe in his shallow green (to brown) faecal pellets the mineralogical characteristics of what he thought to be true berthierine according to the characteristics of that mineral as known from literature dealing with sedimentary ironstones. Above all, he tried to distinguish the supposed diffraction peaks of true berthierine from those of kaolinite by using the precise position of the peaks (002) at 3.51 A (true berthierine) or 3.57 A (kaolinite) and (060) at 1.55 A (true berthierine) or 1.49 A (kaolinite). The X-ray diffraction patterns are very difficult to interpret from that point of view. It is suggested here that this author has probably underestimated the extent of the phyllite V distribution. For us, it seems sufficient to observe green grains, with high magnetic response and a dominant peak at about 7.2 A to be sure that the analysed grains are mainly formed of phyllite V (see Fig. 4)because kaolinite is neither green nor magnetic. Martin has undertaken a systematic treatment of the pellets using 2 N boiling hydrochloric acid on all the samples; he has shown that a peak at 7 A remained present after the treatment. This indicates either that detrital kaolinite was always present within the green pellets separated or, possibly, that impure fractions of grains were analysed. New analyses were undertaken on pellets ultrasonically cleaned and purified using a series of magnetic and heavy liquid fractionations. It was possible to separate nearly pure faecal pellets showing all characteristics of phyllite V. The X-ray diagrams are similar to those obtained from other areas (Fig. 4). Moreover, the mineral is destroyed after heating to 490°Cfor two hours.

Distribution and significance of the verdine facies According to Martin (1973),mineralized faecal pellets, green or brown coloured, have been observed on large portions of the shelf between the Cavally River mouth to the west and 50 km to the east of the Comoe River mouth. This area is located between 4"N and 5" N (Fig. 11). However, the verdine facies seems to be present only as restricted patches at a depth of less than -50m. Owing to the small number of convincing crystallographical analyses, there is only one area where the facies is certainly present: off the Comoe River mouth. Similar concentrations of pellets are present locally off the Bandama, Sassandra, and probably the Cavally River mouths. These sediments

151

can presently only be suspected of being verdine-bearing. The criteria indicating the probable presence of the verdine facies are (1) the Occurrence of green (and brown) pellets showing a distribution very similar to the one off the Comoe River mouth, as well as (2) the abundance of these pellets. It seems, therefore, that each relatively important river (about 600lun in length) is linked with its own verdine-bearing site off Ivory Coast.

Figure 11. Distribution of green grains off Ivory Coast. (Modified from Martin, 1973) (1)Occurrence of green or brown magnetic grains; (2) verdine mineralogically identified; (3) glaucony mineralogically analysed; (4) swamps. Note that the presence of shallow green grains (presumably verdine) is linked to the mouths of the four main rivers of the area. The figure at the bottom follows the western side of the figure at the top.

Below -60 m, the ma netic green pellets give X-ray diffraction patterns showing both a peak at 7 (probably kaolinite remains of the initial substrate), and a large and broad diffraction peak at 14 A or slightly less. This peak behaves like that of a smectite and characterizes the presence of a glauconitic smectite; this mineral is particularly abundant at depths of between -90 m and - 120 m. The description of cores by Martin indicates that the green pellets are essentially present in the first upper metre of sediment; deeper in the cores, the pellets are grey and soft.

1

152 The coast of Ivory Coast is flat and swampy with lagoons between 10 km and 30 km wide. The rivers presently bring fresh-water without sand or clay particles into the sea. Martin notes that the Ivory Coast shelf is presently submitted to low temperature waters (less than 17°C) at depths as shallow as -10 m, for several months a year. Together with the large proportion of very oxidized mineralized pellets, this suggests that the verdine facies is relict. THE VERDINE FACIES FROM MAYOTTE

(Comoro Islands)

Figure 12. Map of Mayotte Island and its coral reef lagoon. The verdine facies has been identified in the eastern lagoon (stars). The main passages through which sea-water is renewed are pointed out (arrows); the island is located between Africa and Madagascar (small arrow).

153 The presence of verdine in the sediments of the Comoro Islands has not yet been reported. At the beginning of 1987, one of us (J.P.M.) suggested that green material, similar to that studied off Senegal, was probably present according to the sedimentological research undertaken by one of his students. Three samples were selected for detailed analyses. The three samples showed a comparatively large propomon of black to grey grains (about 10% to 20% of the sand-size fractions) magnetically attracted. The green material was not visible at that stage of the treatment, but became recognizable after acid leaching using 0.1N acetic acid. However, the really green material was a very small fraction of the magnetic grey to black particles mostly composed of volcanic rock debris more or less magnetite-rich. Moreover, in all three samples, the initially green material was oxidized to ochre or red-brown. In spite of this, the authigenic nature of this material was easily recognizable because it was entirely formed of infillings of microtests (such as foraminifers, adorned ostracodes, echinoderm fragments, gastropodes) minute infillings of perforated bioclasts, and infillings of bryozoan chambers. The deeper sediment collected at -24 m showed a slightly less oxidized and more abundant green pigment (about 0.3%0 of authigenic clay). The X-ray diffraction pattern obtained from this untreated green pigment showed the usual aspect of phyllite V with a dominant peak at 7.2 A and a very low dome between 15 A and 20 A. Mayotte Island can be considered a representative locality for the verdine facies; although few samples are available, the general environment appears similar to the one observed in other localities. The three samples where green infillings were observed came from shallow depths estimated at -3 m; -15 m and -24 m; all these three samples were collected from Mayotte's eastern lagoon (stars on Fig. 12), which, as in New Caledonia, is reefal. This lagoon, a few kilometres wide, encircles a small volcanic island with marked relief summits of 500 m to 600m rise close to the narrow, flat, mangrove bordered coast. Offshore reefs indicate a low continental sedimentary input and the wholly marine character of the waters of the lagoon. At the bottom of the lagoon, mostly located at depths around -40m, dominantly carbonate bioclastic sediments are deposited with very local admixtures of volcanic debris. In short, the local and the general environments are very similar to the ones described for the New Caledonia coral reef lagoon. THE CASAMANCE ESTUARY

(Senegal)

The Casamance estuary is located between the already described areas of the Baie de Rufisque (south of Dakar) 250 km to the north, and the Koukoure River mouth 450 km to the south-east. About 300 km long, the Casamance River is a very particular hydrographic system compared to a "normal" river. One to ten km wide in its lower course, the river is already nearly at the sea-level 230 km inland. Because of the persistency of particularly dry conditions since 1968, the current is weak, even during the rainy season, and a sea-water inflow occurs during most of the year (Pagks et al., 1987). As a result, the salinity of this "inverse estuary" increases inland up to a maximum

154

which, presently, reaches more than 150”/00at the end of the dry season (170”/00 in June 1986). In short, the Casamance estuary acts as an oversalted lagoon, but this is only the present situation and in a recent past, the river brought fresh-water to the sea. Figure 13 below shows the Casamance estuary and its location on the NW African Coast. The area where mangrove is present is shown in order to indicate the very flat and muddy character of the banks of the river.

I

1OOkm

I

,

GS0.87

Figure 13. The Casamance estuary. Today, this estuary effectively constitutes an oversalted lagoon open to the sea with a fresh-water input inferior to evaporation. The three samples studied here (1, 2,4) -with green faecal pellets- were located at very shallow depth at the mouth of the estuary. The dashed area indicates the presence of mangroves and flat muddy banks.

Our main interest has been the region located at the mouth of the estuary. The presence of green grains had never been reported from that area. The preliminary study of these fine quartz sands indicated that magnetic grains with a proportion reaching 2.5% of the whole sediment, were present. Their darkgreen colour and aspect similar to that seen in some places of the shelf north of Dakar make them of interest for the present study.

155

Three sediments were studied in details; they are located in Figure 13. Samples 1 and 4 are fine sands; sample 2 is similar, but has been noted as rich in organic matter. Data are available for the salinity, temperature and pH of the water at the time of sample collection for samples 1 and 2. We have nothing for sample 4, but measurements are available for a several-months-period for two points: 4- 1 and 4-2 located near the collection site of that sample. These results are presented in Table 5 . It must be noted that these data concern present-day water conditions, which could be slightly different from those prevailing at the time when the authigenic clays of the faecal pellets formed. However, according to Table 5, it is reasonable to keep in mind that the waters favourable to green clays were probably hot, with salinity approaching that of sea-water, but with a pH possibly lower than usual for the open sea. Finally, the two fine sand samples (1 and 2) were collected by hand in about 0.50 m of water at low tide. Sample 4 was collected at a depth of about -4 m. Table 5. Physico-chemical data for the water overlying the green grain-bearing fine sands of the Casamance estuary. Seasonal variations are quite wide, but the salinity is roughly similar to that of sea-water and the pH is more acid. Period

Temperature

("C)

Salinity

PH

(%o)

Sample 707.1

J a n . 1987

20.0

33.4

7.34

Sample 707.2

J a n . 1987

21.4

33.2

7.76

Site 4 . 1

1985-1986

Site 4 . 2

1984-1985

20-46 22 .O-29.6

34-40

7.5-8.1

The green grains separated (707-1, 707-2, 707-4) are mostly dark-green ellipso'idal faecal pellets about 250 pm long and 100 pm wide or smaller. In some fractions the green grains were attracted at low intensity by the magnetic separator; in others, higher intensities were required. A variable proportion of the magnetic grains were superficially or wholly ochre to brown, but ultrasonic treatment preferentially disintegrates these brown grains; the green ones, more compact, can therefore be purified. X-ray diffraction study showed that the different minerals described from the area off Cap Vert (Senegal) are also present in the Casamance River estuary (Senegal). Figure 14 gives some examples of the X-ray diffraction patterns obtained where phyllite V (707-4), phyllite C (707-2) and a mixture of the two (707-1) could be identified. X-ray diffraction patterns obtained for sample 707-4 vary from a magnetic fraction to another with a more or less important proportion of phyllite C admixed to the phyllite V. The X-ray diffraction study has shown that the more or less magnetic character of the green grains was linked to the mineralogical nature of the clay composing the grains: phyllite C is clearly less magnetic than phyllite V.

156

35

32

28

24

20

16

I2

KUCU

Figure 14. X-ray diffraction patterns of various green faecal pellets from the Casamance River estuary. The green pellets of 707-4 were separated between 0.37 A and 0.39 A, green pellets 707- 1 between 0.37 A and 0.42 A; and 707-2 between 0.45 A and 0.48 A.

The restricted number of analyses presently available does not allow us to distinguish the preferential environment favouring phyllite C as compared to phyllite V in the Casamance estuary. However, this site does allow the first description of our authigenic (presumed) marine clays apparently inside a river mouth complex. The environmental conditions of this "river mouth" are very particular, however, and could be better considered as a sort of creek or fjord. The presence of authigenic minerals in this environment is therefore very significant and reinforces the marine facies significance of the analysed minerals. The problem of the particular conditions of formation of phyllite V versus phyllite C could probably be solved by a more detailed study of that deposit. CONCLUSIONS

About 20 years ago, the presence of a particular facies characterized by green grains different from glaucony was identified from five areas off the Ogooue, Orinoco, Niger, and Koukoure River mouths in the Atlantic, and off N Borneo in the South China Sea. Although the facies was not correctly interpreted at that time (Odin, 1985a), the recognition of its specificity versus

157 glaucony was a real discovery in the knowledge of the sedimentology of authigenic clays from Recent margins. Since that time, we have obtained complementary results from two of these five areas (Ogooue and Koukoure) and the Orinoco site has been shown to extend (probably continuously) as far as the Amazon River mouth. In collaboration with other French sedimentologists including C. Froget, P. Giresse, L. Martin, G . Moguedet and M. Pujos, six more areas have been studied, more or less in details, by the present authors: the Congo, Comoe, Casamance, and Senegal River mouths on the East Atlantic margin, the lagoon off New Caledonia in the West Pacific, and the lagoon off Mayotte between Africa and Madagascar. According to this list, it would appear that the facies is relatively common in tropical areas between 23"S and 16"N (Fig. 15). The extent of the distribution of the facies varies greatly from about 10 km2 to about 105 km2. This will be further commented in Chapter B6. In view of this world wide distribution and the frequency of the facies, it is surprising that most of these verdine-bearing sediments are only known from less than 15,000 years old deposits. The only ancient site probably referrable to that facies concerns the Miocene sediments cored in the Niger delta.

Figure 15. General distribution of the verdine facies in the present oceans. (1) Ogooue; (2) Amazon to Orinoco; (3) Niger; (4) Koukoure; (5) Sarawak; (6) Congo + NDogo; (7) Ivory Coast; (8) New Caledonia; (9) Senegal; (10) Mayotte; (11) Casamance.

Most of the presently known localities (7 out of 11) are from the eastern Atlantic margin; an eighth one is located on the western margin of the same ocean, and is the largest discovered to date. It is almost sure that other sites exist on the African coast; probably restricted in surface, they have to be searched for opposite the various river mouths still undocumented from that sedimentological point of view.

158

Similarly, there are many possibilities for the occurrence of small deposits of verdine in warm water shelves or atolls in the West Atlantic, Pacific, and Indian Oceans (eastern Australia, Malaysia or Indonesia are good candidates as well as the eastern margin of Africa). Consequently, distribution as known at present -somehow concentrated in the Atlantic Ocean- is mainly a reflection of the area from where present authors have been able to obtain samples. As a result of this review, the main question to be asked concerning the distribution of verdine facies in the present sea is: are there shelves in tropical areas near a continental water input to the sea where green clay formation does not occur in the surface sediment, and why? ACKNOWLEDGEMENTS

We are very grateful to P. Giresse and L. Martin who were kind enough to provide us with sediments collected from their favourite shelf off GabonCongo and Ivory Coast respectively. Others useful samples were provided by G. Moguedet (Congolese shelf'), J. Gayet (Makassar Strait), C. Moreau and B. Robineau (Los Islands, off Guinea). NOTES ADDED IN PROOFS

1- A new example of the tropical verdine facies has recently been identified. It is located off Martinique Island, Lesser Antilles, 61' E - 14' 50 N. Samples provided (March 1988) to the editor by Ch. Henocq and M. h j o s (UniversitC de Bordeaux) have shown the presence of abundant green grains (up to 15% of the sediment) at depth of about -50 m. These samples were collected in 1985 (CaracolanteI1 cruise) from the shallow shelf to the NE off Ste Marie, and within the Baie de Fort de France to the SW of Martinique Island. The green grains bearing sediments are presently located below 1 or 2 m of ashes probably Recent in age. All green grain samples analysed by the editor produce X-ray diffraction patterns characteristic of the minerals from the verdine facies. The abundance of the 14 8, phyllite V together with the 7 8, phyllite V in the same grains, indicates a relict facies. The present facies may be compared with the Guianese verdine; however it is mostly buried below volcanic products; an early burial diagenesis may have disturbed the initially marine green clays and this will be studied in a near future. 2- Dr Ch. Henocq informed the editor (April 1988) that, according to the literature, a facies similar to the one quoted above off Martinique Island was also known from several areas about 700 km to the NW, off Puerto Rico (Seiglie, 1970, Rev. espafi. Micropal., 2: 183208). Seiglie describes a shallow facies which appears similar to our verdine facies: abundant green clay within essentially relict foraminiferal shells or substituting shell material; many other observations by this author correspond to what was seen off New Caledonia, but the green clay proportion is higher, like it was reported from off Senegal. However, Seiglie (1970) uses the word "glauconite" to designate the green clay located at depths between -5 m and -10 m, between a typically fluvial-marine facies and a typically reefal facies. Although confirmatory analyses remain to be undertaken, the shallow depth and general environment of these green grains suggest better the presence of the verdine facies.

159

Chapter B5 MINERALOGY OF THE VERDINE FACIES by G.S. Odin, S.W. Bailey, M. Amouric, F. Frohlich and G. Waychunas INTRODUCTION

Studies of the verdine-bearing deposits were presented in detail in Chapters B 1 to B4.They allowed us to obtain and discuss some particular mineralogical results for each area studied. The present chapter will 1) review the preceding results; 2) try to show whether or not each observation on a given deposit is valid for the others; 3) supplement the previous results; and 4)propose our present view on the minerals concerned. The first remark is that we have to distinguish clearly the precise nature of the mineral(s) composing the green pigment from its general aspect and environment, i.e., to separate the data characterizing the facies from the crystallographical data characterizing the mineral(s). Specifically, we now mention again that the term verdine designates a facies and that this facies, among other characteristics, is composed of minerals which obviously must have a designation different from that of the facies. This might appear trivial, but the literature on the green grains is full of such confusion (refer to "glauconite" versus glaucony or "chamosite" versus berthierine). The second remark concerning the mineralogy of the verdine facies is that the green clay minerals can be purified by a combination of magnetic and heavy liquid separation. This allows observations and analyses to be made much more easily than for most clays which are rarely available as monospecific separates. MINERALOGICAL DATA ON THE GRAINS OF VERDINE

Colour The green colour of the grains considered here can be compared to that of the better known glauconitic grains. No real difference in colour is obvious between grains from the two facies. The colour of the green grains that were later classified as verdine may vary from a very light white-green to a very dark black-green. The darkest grains can only be identified as green after crushing. The powder obtained has a blue-green colour, often a little grey for the darkest grains. The dark-green material has been found to be nearly exclusively from the deepest (i.e., probably oldest) deposits. The light-green material is the one usually found in the shallowest sediments (i.e., probably more recent). But dark-green clays are also present in shallow sediments such as those from faecal pellets of the Islands of Los (Guinea). Similarly, a number of foraminifer infillings (in Miliolids for example) are dark-green in the lagoon of New Caledonia in sediments at depth around -20 m.

160 A common factor for the light-green clays is that they have always been found surrounded by carbonate remnants; in contrast, the grains totally free of the initial substrate within which the clays have been crystallized are always dark-green. There is therefore a general indication of evolution or maturity given by the colour.

Paramagnetism The green clays of the verdine facies can be concentrated mechanically using a magnetic separator allowing concentration of large quantities of pure material in favourable cases. Like colour, this property also applies to glauconitic grains. It was shown (Chapter B 1) that this property also applies to the detrital chlorite grains of New Caledonia. This could lead to a concentration of mixtures of grains, all green in colour, that would give misleading analytical results not representative of the authigenic clays studied. Compared to glauconitic grains, the green grains of the verdine facies are attracted at lower values of the magnetic field (i.e., they are more magnetic). In practice, a "Frantz isodynamic" separator with a longitudinal slope of about 25" and a lateral slope of 17' allows verdine to be attracted for a 0.35 to 0.45 ampere current to the magnet, while the glauconitic grains are mainly attracted between 0.4 and 0.5 ampere current. This observation was first made for the Senegalese samples in which a very curious circumstance occurred. Most samples first appeared "normally" attractable, but two or three were surprisingly much less attractable even though the general colour of the grains and quality of the separates appeared generally similar. Particular attention was paid to these differently attractable grains, which were compared to glauconitic grains, and the histograms of Figure 1 were obtained. The most attractable grains were identified later as the phyllite V clay mineral and the clearly less attractable grains corresponded to the phyllite C clay mineral which just begins to be attracted at 0.45 A. (Fig. 1).

I

.'

I

1

.

I

I

0.>5

030

035

040

' 045

0.50

055

06OA

Figure 1. Histograms showing the paramagnetic properties of various green grains. (1) ooids of berthierine from Hayange (near Thionville, Lorraine, France); (2) verdine grains made of phyllite V; (3) glaucony; (4) verdine grains made of phyllite C. The x axis indicates the current intensity (amperes) of the electric magnet of a "Frantz Isodynamic" separator with lateral and longitudinal inclinations of 17' and 25', respectively.

161 In general no important difference in magnetic attraction was found for light-green or dark-green clays when present in a given deposit. As a result, it has not been easy to mineralogically characterize these two morphologically different sorts of clay separately in a single sediment.

Density and hardness Heavy liquids (bromoform or bromoform-acetone mixtures) were often used in the course of the purification process of the green grains in order to eliminate inherited pyroclastic minerals denser than the bromoform, as well as oxidized grains frequently heavier than the other green grains. The density of the grains made of green clays was always clearly lighter than the bromoform and very near that of quartz which, either trapped in the green clay or trapping a small inclusion of magnetite, could not be entirely separated mechanically from the rest of the green grains. As a result, many of our X-ray diagrams show a small proportion of quartz. The difference in density between light-green and dark-green clays, if any, was not great enough to allow concentration using heavy liquids. Concerning hardness, light-green clays are always very soft and a powder is obtained easily. In contrast, the dark-green grains are hard and need a few minutes of hand crushing in an agate mortar to obtain a powder adequate for X-ray diffraction analysis. Related to hardness, the resistance to ultrasonic treatment is higher for the dark-green grains; the latter are not dispersed by several minutes immersion, whereas the light-green grains are dissociated in a few seconds.

Selection of samples for further mineralogical studies Following these preliminary mineralogical results different samples of verdine were selected for more detailed studies. A light-green sample was selected from the Neo-Caledonian deposit where this material can be obtained pure enough. A large quantity (ten grammes) was separated from twenty kilogrammes of sediment; neither concentrated acid nor long acid leaching were used for the purification of the green grains of that sample (704-2). An older verdine (presently lying deeper than any Neo-Caledonian verdine-bearing sediment) was selected from the Baie de Rufisque (south of Cap Vert, Senegal); the very magnetic material from sample 601-95 is mostly dark-green. In the same area, a sample of the less magnetic green grains (601-92) has been cautiously purified and is mainly composed of mineralized faecal pellets. Finally, one of the probably oldest known verdine (green grains entirely free of carbonate remnants from the initial substrate even without treatment, dark-green in colour, and comparatively hard to crush) was selected from the sediments lying off French Guiana (508-103). These four sorts of grains have been analysed in similar conditions in order to determine the common properties of the components and to what extent they are homogeneous.

162 X-RAY DIFFRACTION DATA ON THE VERDINE FACIES

The study has c o n f i e d that the X-ray diffraction patterns obtained from the green grains referable to the verdine facies are of two kinds; the first kind shows a dominant peak near 7.2 A; it is present everywhere in all known deposits. The second kind shows a very strong peak near 14 A.

The green clay with a dominant peak near 7.2 A:phyZZife V

X-raydigraction of untreated samples Figure 2 gives the X-ray diffraction patterns of three of the purest materials obtained mechanically in the course of the present work. With the exception of

-.----

36"

32"

28"

24"

200

16"

12"

4"

Figure 2. X-ray diffraction patterns of three phyllite V-type components. The green grains were purified mechanically for sediments from French Guiana (508), Senegal (601). and New Caledonia (704). The oldest, hardest, and darkest material is the one from French Guiana, which is considered more evolved than the shallower light-green material from New Caledonia. The former does not give sharper peaks than the latter; the reverse seems correct. The three diagrams are the reference patterns for comparison with diagmms of acid-leached and heated materials below.

163 a small proportion of quartz, all the diffiaction patterns shown can be considered as due to the authigenic clay minerals formed near the sea-water/sediment interface. Patterns using randomly oriented powders have a similar position for the dominant peak at about 7.2 A whether the grains are light or dark, very recent or older. The least evolved light-green material from New Caledonia does not show a main peak that is broader or smaller than that for the most evolved dark-green grains. The reverse appears sometimes correct.

1.51A

1.56A

7 - r

704.2

699

.-. ....

601.92

GSO87

66"

62"

58"

Figure 3. X-ray diffraction patterns of minerals of the verdine facies in the vicinity of the 060 diffraction peak. Examples are taken from light-green grains from New Caledonia (700 and 704) mid- to dark-green grains from Guinea (699), Senegal (601), and dark-green grains from French Guiana (508). The diagrams obtained are very homogeneous for phyllite V; phyllite C (sample 601-92) shows a diffraction peak clearly different at 1.52 A.

164 On the low 28 angle side of the 7 A peak, it is always difficult to identify any discrete diffraction peak (when no foreign material is present as will be discussed later). However, a large majority of the diagrams obtained show that there is diffracting material which clearly increases the normal background in the area corresponding to between 20 A and 10 A and sometimes up to the base of the peak at 7.2 A. This indistinct area of diffraction is underlined in Figure 2. It must be emphasized that, in spite of very diverse stages of evolution, depth of deposition, or area of sampling, the diagrams obtained appear remarkably homogeneous (see also Fig. 4 in Chapter B4,p. 139). Another point of interest concerning the X-ray diffraction patterns of untreated verdine is the location of the 060 diffraction peak known to be connected with the dioctahedral or trioctahedral nature of the analysed mineral. Figure 3 indicates that the clearest diagrams obtained from materials collected in the different deposits show a rather constant 060 peak at about 1.54 A. This peak is wide and asymmetric; it slowly decreases towards 1.50 A. This probably indicates non coincidence of the 060, and 331- peaks as a result of distortion of the structure, plus the appearance of 061,330,and 332-peaks. Finally a large d-060 value is to be expected as a result of the high iron content regardless of the dioctahedral or trioctahedral nature. X-ray diffraction of thermally-treated phyllite V No clear change of the diffraction patterns were obtained on the natural samples after ethylene glycol or hydrazine treatments, but heat treatments are more interesting. As soon as the temperature of 350°Cis reached, there is a marked decrease of the diffraction peak near 7.2 A, which underlines the real fragility of the mineral. After one hour at 490°C,most of the 7.2 8, peak has disappeared. The darkest material has not kept its 7.2 A diffraction peak longer than the lightest material. In the area corresponding to 10 to 15 A, after heating to 490°Cfor one hour, the two samples 95 and 103 show a clear development of a dome. Figure 4 gives the diagrams obtained after heating to 490°C for two hours. The increase of a diffraction peak around 14 A is spectacular for the two dark-green samples. The Neo-Caledonian sample seems to keep its previous diffraction in that area. A complementary experiment at 600°Cfor two hours does not sensibly modify the situation. When the components are heated to 600°C for 4 hours a portion of the diffraction peak shifts toward 10 8, (Fig. 4). In summary, heating treatment shows that the peak at 7.2 A is quickly destroyed at less than 500°C.The most evolved green material also shows the noticeable development of a wide peak around 14 A, a part of which shifts toward 10 8, when heating to 600°Cis undertaken for longer than two hours. The less evolved light-green material does not develop a good diffraction peak at 14 A, but the previous low diffraction in that area remains present and partly shifts to 10 A after heating to 600°C for four hours. This behaviour generally c o n f m s the previous experiments (Chapter B 1, B2) and appears reproducible. A few of the layers do not have stable interlayers and collapse to 10 A on prolonged heating.

165 ,

1490% 2hourr1

1

1600O C 4 hours]

Figure 4. X-ray diagrams of phyllite V after heating treatments. 1 and 4: sample 704-2, light-green grains from New Caledonia; 2 and 5: sample 601-95 mid- to dark-green grains from Senegal; 3 and 6: dark-green grains from French Guiana. Note the absence of the peak near 7.2 8, and the development of the wide peak near 14 8, after heating to 490'C for two hours (left hand side). The development of the peak near 14 8, is more pronounced for the oldest and darkest material. After heating to 6OO'C (right hand side) a limited portion of the sheets may collapse to 10 8,. Compare to the diagrams in Figure 2 and other diagrams in Chapter B2.

In order to characterize the material considered here, it appears useful to recommend three different heating treatment i.e., at 350°Cfor one hour, at 490°Cfor two hours, and at 600°C for four hours. Three different diagrams will be obtained, all three different from that realized on the untreated material. The most interesting figure is the one obtained after heating to 600°Cfor four hours which indicates the presence of a diffracting structure partly stable at 14 8, and partly collapsing with difficulty to 10 8, after a long duration treatment. The light-green and the dark-green clays do not show a perfectly similar be haviour.

X-ray difraction of phyllite V following acid leaching The normal diffraction patterns of phyllite V are altered by dilute acid solutions. For example, a 1N solution of hydrochloric acid at ambient temperature for two hours, leads to a large decrease in intensity of the peak near 7.2 8,. The light-green clay is more quickly altered than the dark-green clay. The origi-

166 HCI 6N 22OC lhour

HCI I N 7OoC lhour

8 /

/'

I

'

J

1,1) hje

0 I

0

+%

\

209

'

I &

16'

89

49

Figure 5. X-ray diffraction patterns of phyllite V after hydrochloric acid treatment. A 6N solution has been used at ambient temperature for one hour (right hand side side) by comparison to a 1N solution at 70'C for one hour (left hand side side). Note a better resistance of the dark-green clay of French Guiana (3 and 6) compared to the material from the dark-green grains of Senegal (2 and 5). The light-green material of New Caledonia (1and 4) is quickly destroyed and becomes whitish.

nal green colour is noticeably modified: a very soft whitish material is obtained; the acid solution itself becomes brilliant yellow. Using a 1N hydrochloric solution heated to 70'C for less than one hour the light-green clay no longer diffracts X-rays coherently. For similar conditions, some diffraction peaks still exist for the dark-green clay and, curiously, a sharp peak at 7 A remains for the Guianese verdine (possibly chlorite remains from the substrate). The use of 6N hydrochloric solutions at ambient temperature leads to a similar alteration (Fig. 5). Therefore, the phyllite V mineral is easy to destroy, even at low temperature with a concentrated solution of hydrochloric acid. When moderate acid treatments are undertaken, it is possible to have a diffraction peak around 14 A. Several reagents have been tested on the three different sorts of green clay selected here. Figure 6 shows the results of experiments undertaken at 70°C for one hour using N/10 hydrochloric acid or 1N acetic acid solutions. There is a clear development of a broad diffraction peak around 14 A on the six diagrams obtained. This is very interesting since,

167 D

20'

16'

12'

KuCu

8"

HCI N/10 70°C l h o u r

4'

Figure 6. X-ray diagrams of acid-leached phyllite V components; acetic acid (left side) and hydrochloric acid (right side). Compare to the untreated samples (seeFig. 2). 1and 4: sample 704-2; 2 and 5 : sample 601-95; 3 and 6: sample 508-103. The decinormal solution of hydrochloric acid quickly destroys the lightest, least evolved mineral of the verdine facies. The development of the peak around 14 A is larger for the more evolved, darker material from French Guiana than for the lightest one from New Caledonia when hydrochloric acid is used. In contrast, the 1 normal acetic acid solution does not preferentially modify any sample.

like the heating treatment, it seems that the acid leaching reveals a periodicity that was not clearly expressed in the diagrams obtained on untreated samples. Therefore, a specific treatment had to be chosen in order to obtain the most efficient and reproducible behaviour. Figure 6 indicates that the dark-green clay (diagrams 3 and 6) reacts more clearly to the hydrochloric acid treatment than the light-green clay (diagrams 1 and 4), confirming the tendency observed with the heating treatments. However, acetic acid solution appears more efficient and reveals better the 14 8, periodicity of the light-green clay; this reagent was preferred and further tested at different temperatures and for different durations. Finally, it was found that the most efficient treatment was to use a 1N acetic acid solution heated to 70°C for five hours. In practice, the material to be tested (20 to 30 mg necessary for realization of a randomly oriented preparation) is transferred into a 20 cm3 glass tube, the solution is added, the tube is closed and deposited in an oven at the desired temperature (70°C). After reaction for the duration selected, the acid solution is removed, the material is cleaned with distilled water and dried.

168

I

.

.

Figure 7.X-ray diagrams of acid leached phyllite V-type components. 1N acetic acid for six hours (left hand side side); similar reagent for three hours, and complementary treatment with ethylene glycol (right hand side side). After three hours of acid leaching the peak around 14 A is more developed than after one hour (Fig. 5 ) and less developed than after six hours. The intensity of the peak near 7.2 A slowly and progressively decreases during this leaching. Ethylene glycol treatment distinctly leads to a slight swelling of the layers generating the peak around 14 A, for which the top shifts to 16 A. 1 and 4: light-green Neo-Caledonian sample 704-2; 2 and 5: dark-green Senegalese sample 601-95; 3 and 6: sample 508-103 from French Guiana.

Figure 7 gives the diagrams obtained after leaching for six hours with 1N acetic acid solutions heated to 70°C. The development of the peak at about 14 A is spectacular, and it is suggested that this sort of behaviour is a very specific one for the material under study here. The colour of the previously light-green clay has been modified to yellowish and the dark-green clay also becomes lighter than before. The nature of the mineral giving a well-shaped diffraction peak near 14 A has been further evaluated using complementary treatments following moderate acid leaching. For example, the use of ethylene glycol leads to a small, but systematic, swelling of the sheets toward about 16 A (Fig. 7). Heating to 490°C for two hours after acid leaching modifies the 14 A peak so that a portion shifts toward 10 A and the remaining portion stays at 14 A or even higher. This

169 is true for the two dark-green clays, but the Neo-Caledonian light-green clay seems to be destroyed by this heat treatment following the acid leaching. The X-ray diffraction patterns obtained are roughly similar to the ones obtained by heating untreated samples but the shift toward 10 A is more pronounced after acid leaching. In summary, acid treatment of the phyllite V destroys the clay structure when moderately concentrated solutions of hydrochloric acid are used at high temperature. The light-green and the dark-green clays have distinctly different behaviours. A specific treatment using 1N acetic acid at 70°C for five hours is recommended here in order to emphasize the very noticeable development of a diffraction peak near 14 A. The 7.2 A diffraction peak is also diminished by this treatment, especially for the dark-green component. The 14 A diffraction peak generated by the moderate acid leaching swells slightly with ethylene glycol and partly collapses to 10 A at 600°C.

The green clay with a dominant peak near 14 A:phyZZite C The phyllite C-type mineral has only been analysed from two independent occurrences: the area to the south of Dakar (Baie de Rufisque) and the Casamance estuary, 250 km more to the south. The material of these peculiar green grains is very similar in the two areas as well as very different from phyllite V as far as X-ray diffraction characteristics are considered. The main characteristics have been discussed in the study of the Senegalese deposit. The untreated grains are made of a single nearly pure mineral showing a very strong peak corresponding to a d-value slightly above 14 A. The top of the peak may be between 14 A and 15 A. The other peaks, and especially the peak at 7 A, are always very weak. Another specific character of phyllite C is the position and shape of the 060 diffraction peak, which is illustrated in Figure 3 (p. 163). The top of the diffraction peak corresponds to about 1.52 A, which possibly reflects a dioctahedral structure more dominant than for the phyllite V mineral. It has already been shown on one hand that the ethylene glycol treatment leads to a distinct, although not very large, swelling of the layers. On the other hand, the very strong 14 A peak is easily decreased by heating. At 250°C for one hour, the intensity (surface) of the peak is decreased by a factor of two or more. If the temperature is increased to 600'C, the peak height and surface slowly decrease; this peak is replaced by a dome still beginning at 14 A but reaching 10 8, on the high angle side. Therefore, the mineral looks like a smectite to which some characters of chlorite would be admixed. The strong, sharp peak at 14 8, means that all layers have the same thickness i.e., a regular crystal architecture. The layers variously modify their spacing when a moderate physico-chemical treatment is undertaken. In order to verify this fragility further, a sample was treated with hydrochloric acid solutions. Figure 8 gives the diagrams obtained for two of these treatments. The green clay reacts very similarly to the most evolved dark-green phyllite V of the Guianese samples. However, a peak near 7 A is

170 surprisingly preserved while the peak near 14 8, is mostly destroyed. An inherited chlorite structure was suspected in French Guiana samples; this explanation is unlikely in the Senegalese sediments where chlorite has not been observed in the clay-size fraction. During the leaching, the green grains have modified their colour, which is now white to cream to ochre depending on the grains. In summary, according to the X-ray study, there are clear differences between the most common green clay with a dominant badly shaped peak near 7.2 A and a clay mineral rarely encountered in this study, that shows a strong and dominant 14 A peak. This diffraction peak is easily modified by ethylene glycol, heating, and acid treatments. Phyllite C is not purely a smectite-type mineral, since the swelling as well as the collapsing are far from complete.

/A

~~~

601-92 Acid treated

601-92 Untreated

A HCI 1N 70% (hour

~

J

&d-%&, st.

.f"

36"

32'

,

28'

24"

20"

.

16'

,

28'

24"

20"

.

16'

12'

.

8'

40

Figure 8. X-ray diffraction patterns of the phyllite C mineral. The untreated sample gives a strong 14 8, peak which could not be kept on scale using the conditions of emission, diffraction,and recording usually selected for the analysis of verdine. Therefore, the emission power has been diminished from 900 W to 700 W for diagram(1) compared to the other figures in this chapter. In diagram (2) the peak near 14 A is greatly diminished and possibly partly shifted to the high angle side. In diagram (3) there is no diffraction peak at 14 A but a well-shaped peak at 7 A is still present; there is also an indication for a peak at 10 A.

171 Concerning a possible genetic link (or progressive passage) between phyllite V and phyllite C, diagrams intermediate between the two characteristic patterns (i.e., diagrams with both a mean peak near 14 8, and near 7 A), were obtained; however, it has been possible to separate from the bulk samples giving these diagrams, green grains of slightly different aspects which, analysed separately, gave two distinct diagrams the one characteristic of phyllite V, the other of phyllite C. This is the case in the Baie de Rufisque where a probably important proportion of the sediments collected contain the two components; the same is true in the Casamance deposit where all magnetic separates are mixtures of mineralized faecal pellets made of phyllite V and pellets made of phyllite C. The distinction may be made in treating the green grains. Heating to 150°C or 250°C lowers the 14 A peak when ph llite C is concerned. In contrast, phyllite V tends to develop a peak near 14 when heated. The location of phyllite C at a depth around - 100 m off Senegal is bracketed between deeper (i.e., older) and shallower (i.e., younger) green grains which mostly show the presence of phyllite V. In this context, it would not be easy to draw an evolution from one component to the other as a function of time because the age of the green material is likely to be a function of depth. It is suggested here that there are two different components representing two distinctly different geochemical answers to two slightly different environments that are both located at the shallow sea-water/sediment interface.

K

HIGH RESOLUTION TRANSMISSION (HRTEM) ON VERDINE

ELECTRON

MICROSCOPY

The sedimentological and X-ray diffraction studies allowed us to select a small number of samples able to represent the diverse mineralogical compositions of the green grains met in the verdine facies. Three presumably very recent neoformed green clays were prepared from sediments collected from New Caledonia (B1 and V4, Fig. 6 in Chapter B1, p. 64)and from Guinea (699, Fig. 4 and 5 in Chapter B4, p. 139 and 141). Two older green clays referable to the phyllite V clay mineral (601-58 and 601-108) and one green clay referable to phyllite C (601-92) were selected from the Baie de Rufisque. All samples were embedded in araldite resin before sectioning with an ultramicrotome equipped with a diamond knife. A E M lOOC (100 KV) and a JEOL 2000 Fx (200 KV) electron microscope were successively used for the HRTEM study. Images were selected from experimental through-focus series recorded in the 800-1200 A range of underfocusing. These optimal imaging conditions for phyllosilicates were defined referring to previous image simulations in micas (Amouric et al., 1981). A loss of contrast with time, without apparent phase transformation, was observed under the beam due to the sensitivity of clays to the radiation damages. As a result, quickly recorded one dimensional images were mainly analysed here. Most of these images showed characteristic 001 lattice fringes of the dominant mineral phases present in the samples.

172

Verdine from New Caledonia The infillings of foraminiferal or other tests and epigenetic replacements of carbonate bioclasts of New Caledonia were initially purified using 1N hydrochloric acid. The microtomed sections of these green grains (sample B1) show highly predominant spacings near 7 A (Fig. 9), hotomicrograph 2925 1 shows an actual mean spacing between 7.3 8, and 7.6 which is slightly higher than shown on the X-ray diffraction patterns. The sections of the microcrystals are usually formed of 15 to 20 layers and they are more than 0.1 pm in length. The spindle shaped general section is very common. The detailed observation of a number of sections allowed us to discover that some sheets are sometimes thicker or thinner and may reach 9 to 10 A or 6 to 5 A. In photomicrograph 29249 such a pair (9+5) A can be seen in the middle of several groups of usual 7.5 A layers. This pair is structurally significant of a chlorite layer. That situation has been met in other views; however, the proportion of (9+5) A layers remains low (less than 5-10%). It must be concluded that, when such a mixture of spacings exists, the apparent thickness of the layers appears to become irregular and suggests a possible local evolution from one layer-type to another i.e., from a 7 A mineral to a 14 A mineral. Less commonly, the microstructure may show thick 7.5 8, stacking sequences and, at the periphery, 3 or 4 layers of which the mean thickness may reach 11 or 11.5 A (Photo. 29237). This might correspond to small foreign crystals attached to the main structure.

i

Verdine from Guinea The small black-green faecal pellets of sample 699 were collected from between the Los Islands at low tide and at a depth shallower than I rn. They were purified from their carbonate using a less aggressive treatment than for sample B1 above (N/10 acetic acid). The HRTEM study confirmed that more than 90% of the existing spacings were near 7 A. Photomicrograph 30995 allowed us to measure a mean spacing at 7.3 A. The microcrystals appear to be formed of ten to twenty layers each. Laterally, a few 7 A spacings seem to become wider and reach 10 A or more; the continuity of the layers is not obvious, however. Most of crystals apparently are pure 7 A clay minerals i.e., the green grains of the verdine facies are sometimes monomineralic. In one or two cases (30993) a spacing at 21 A is observed as a grouping of three individual 7 8, layers.

-+++

Figure 9. HRTEM pictures of "young"verdine. The 7 8, phyllite V from New Caledonia (sample B1: 29251, 29249,29237) and Guinea (30993,30995). Pictures, by M. Amouric, show a l,OOO,OOO times enlargement. Note the very dominant, often exclusive, 7 8, spacings (29251,30995). Rare (9+5) A pairs indicate the presence of individual chlorite layers (arrows on 29249). A few 10 A spacings are also present out of the 7 A groups (29237). Finally, rare groups of three 7 A layers give apparent spacings at 21 A (arrows on 30993).

174

Figure 10. HRTEM pictures of the Senegalese verdine minerals from sample 601-58. From the top to the bottom 7 A, (7x2) A, and 10 A spacings are illustrated. The actual mean thicknesses are about: 7.5 A, 15 A and 10.7 A, respectively. Note that all microcrystals are monomineralic. (Pictures by M. Amouric, enlargement: x 1,OOO,OOO)

Verdine from Senegal (phyllite V) The two samples selected for characterization of phyllite V from the Senegalese shelf are sedimentologically and mineralogically equivalent, both being collected at about -30 m depth. Sample 601-58 (Fig. 10) was cleaned using 1N hydrochloric acid in order to remove the carbonate of the substrates; sample 601-108 (Fig. 11) was prepared later using a less aggressive N/10 solu-

175

Figure 11. HRTEM pictures of the Senegalese verdine sample 601-108. Photo. 30905 shows two microcrystals; one is purely made of 7 A spacings the other of irregular 14 8, spacings. Photo. 30888 shows a microcrystal externally showing 7 A spacings but the interior is made of a chloritic mineral characterized by (9+5) A pairs. (Pictures by M. Amouric, enlargement: x 1 ,OOO,OOO)

tion of acetic acid. According to the X-ray diffraction study these verdines are formed of minerals showing a predominant peak near 7.2 A; but a diffraction hump is also frequently present between 14 A and 20 A. Three sorts of spacings were observed in the ultra-thin sections of sample 601-58: 7.5 A, 15 A and 10.7 A. The 7.5 A spacing is not uncommon and several particles have been found purely formed of fifteen to twenty five times 7.5 A thick layers (see Photo. 29347). These 7.5 A sheets have been observed once systematically grou ed two by two in a thick microcrystal containing more than twenty of these 15 thick pairs (see Photo. 29368). However, in view of the equal thickness of the two units of the combined structure, and assuming that we used the best imaging conditions, that organization is representative of a (2 x 7.5) A polytype rather than of a true chlorite which must show regular (9+5)A alternating spacings.

51

176 The surprisingly predominant spacing observed in the sections of sample 601-58 is 10.7 A thick as a mean. Several precise measurements were done on the numerous well-ordered particles showing this structure; extreme mean values of 11.4 and 10.4 8, were obtained. Frequently, these particles contain twenty five to fourty layers: they are comparatively thick (see Photo. 29355). Sample 601-108 gave even more composite views since four sorts of particles were observed with spacings at about 7 A, 10 A, 14 A, and mixedlayer (7-14) A. The latter mixed-layer and the crystallites comprising pure 7 A spacings are predominant in some views. Photomicrograph 30905 is a view of two particles near each other; the first shows a well-defined spacing near 7 A, and the other poorly defined layers of 14 A thickness. Some sections of 250 to 300 A thick particles made wholly of 10 A thick layers were observed. They are similar to the equivalent particles of sample 601-58: in these particles, there is usually a good periodicity without thinner or thicker sheets. The most interesting sections in sample 601-108 show irregular mixtures of 7 A and 14 A layers. Photomicrograph 30888 shows a more than 300 A thick particle which is schematically made of 7 8, layers surrounding 14 A layers. From the bottom to the top of the section, one may observe seven or eight layers of equivalent thicknesses at about 7 A, then five to seven layers of which the thickness is more or less alternatively thicker and thinner than 7 A, then we reach the central part of the crystal where the 14 A periodicity is typical of a chloritic clay mineral. In contrast to what was observed in photo. 29368 above, the 14 A layers are clearly composed of similarly defined sheets at about 9 A and 5 A. The two samples (58 and 108) that were treated differently with acid solutions give roughly similar pictures.

Verdine from Senegal (phyllite C) The X-ray diffraction patterns of the earthy-green faecal pellets of sample 601-92 are characterized by a strong peak at 14 A and a very small 7 8, peak. The corresponding mineral was called phyffiteC. The preliminary study was disappointing because the above quoted 14 A peak did not correspond to well defined spacings in the HRTEM views. The microtomed sections showed microcrystals purely made either of 7.5 A or 10.2 A or 14 A type spacings (Fig. 12). The latter were poorly defined and resembled the structure shown in photomicrograph 30905. A second series of photomicrographs has shown that the 14 A spacings are predominant, many are not subdivided and irregular: they show small lateral changes in thickness; this observation is common when smectites are observed; this confirms the smectitic behaviour observed when heating during the X-ray diffraction study. The (9+5) A pair indicative of a chlorite structure was rarely observed, but the (7+7) A pair was slightly more common. The 10 A spacings, sometimes observed in 250 A thick crystallites, were often equally subdivided by a line as it is commonly seen with these minerals. In photomicrograph 600, the 10 A spacings are inequally subdivided but this is a technical artifact linked to non-optimal imaging conditions.

177

Figure 12. HRTEM pictures of the Senegalese verdine sample 601-92. According to X-ray study, this sample is composed of phyllite C. However, small and rare microcrystals show 7 A spacings (Photo. 587) which sometimes seem to be modified to (9+5) A pairs (arrows on Photo. 587). Rare but nicely defined and strictly monomineralic 10 A microcrystals occur too (Photo. 600). The majority of the crystallites, often larger than those precedingly quoted, show poorly defined 14 A spacings (Photo. 611) which are similar to what would be seen for smectite or vermiculite better than chlorite. Photo. 586 shows a composite crystallite with dominant 14 A spacings in the centre(a) one or two 10 to 11 A spacings on the borders (b) and laterally, apparent 7 A spacings (c). (Pictures M. Amouric x l,OOO,OOO)

Conclusions of the TEM studies Taking into account the results obtained from French Guiana, (Chapter B3) the high resolution transmission electron microscopy allowed us to identify two sorts of verdine: one mostly monomineralic, the other composite. The first concerns shallow verdines known to be Recent in age and possibly in the process of genesis at the time of sampling in the sediment; the "young" verdines are aggregates of particles fundamentally made of 7 A thick layers.

The second concerns probably more ancient verdines of which the grain sections show mixtures of about 7 A, 10 A, and about 14 A microcrystals. The suspected less mature verdine dominantly shows layers which are never thinner than 7.3 A as a mean. This corresponds to the predominant thickness observed in the X-ray diffraction patterns. The general shape of the 7 A peaks observed on the X-ray diffraction patterns can be understood thanks to this HRTEM study. The top of the peak, at about 7.2-7.3 A, represents the largest proportion of the layers present. There are very few thinner layers and the X-ray diffraction peak decreases sharply on the high 28 angle side. On the other side of the top, the peak decreases more slowly because a number of layers are locally thicker, commonly reaching 8 or 9 A. When we measure the mean thickness of a dozen adjacent la ers on a photomicrograph, the obtained mean is frequently about 7.4 to 7.5 because a few thicker layers (and no thinner ones) are added to the most common 7.2 A thick typical layers. The more "mature" verdines contain both 7 A and 14 A type layers. These layers sometimes are met separately in distinct particles but sometimes are combined in particles where there is a geometrical passage between pairs of 7 A layers (one of which becomes wider, the other thinner) and the pair (9+5) 8, which finally forms the 14 A chlorite-type layer. In some samples (especially the one identified as phyllite C) the 14 A layers are not subdivided and their specific nature cannot be resolved using the TEM observation. Groups of two, and even three, 7 A layers were also observed. This supposes a polytypic behaviour of the 7 A phase. In summary, the presence of 10 to 11A thick layers forming comparatively thick particles is a problem discovered during the E M study. The X-ray diffraction patterns did not give the impression that these layers were so common and sometimes predominant in the most evolved green grains. The particles made of 10 A type layers are exclusively formed of this single sort of layer. The main problem solved using the TEM observations is the identification of a mineral purely made of 7 A type layers in some samples of what was previously identified as phyllite V. This mineral is predominant in the less mature verdines. In complement to that, the presence of rare combinations of 7 A and 14 A spacings in more evolved verdines might be interpreted as the possible evolution of the 7 A initial clay toward a more stable chlorite-type 14 A mineral. This "chlorite" may entirely form specific crystallites. The knowledge of the phyllite C has not made any progress during this TEM study compared to the X-ray diffraction study, and the research has to be pursued further.

f

CHEMICAL STUDY OF THE MINERALS FROM THE VERDINE FACIES

Chemical composition of phyllite V There has been some difficulty in obtaining enough pure material for a detailed and precise chemical analysis. At the beginning of this research analy-

179 ses were undertaken on green pigment for which it was known from X-ray diffraction analysis that a certain amount of impurity was present. For example, the two samples from off the Congo River mouth are partly oxidized (excess Fe,O,), contain a small proportion of kaolinite, and probably a potassic mineral. Concerning French Guiana, there is a possible small admixture of chloritized biotite in most samples. Therefore, an excess of silica coming from that mica is to be assumed if a chlorite or serpentine formula is to be calculated. HRTEM photomicrography of ancient verdine confirms the presence of a micaceous component which will give an excess of silica if a chlorite or serpentine formula is calculated. Finally, the first samples analysed from the Senegalese margin were not treated until totally free of their carbonate substrate in order to avoid a possible alteration due to the acid leaching.

Analytical results The chemical data discussed below will strictly concern analyses which were obtained and verified in the course of the present study. For most of the Table 1. Chemical analyses of verdine grains of various Quaternary ages, and showing a main peak at about 7.2 A (phyllite V). (3)

(3)

(3)

700.3AB

704.4

601.108

601

34.0

31.2

32.6

34.8

5.2

5.9

9.3

23.1

25.1

SiO2 A1203 Fez03

6.13 22.0

(2)*

(1)+(2)**

(2)**

(1)

(1)

508.102

508.103

508.169

508.172

36.9

36.9

39.1

37.5

38.2

36

36.0

11.0

12.0

11.8

10.8

10.8

12

12.2

21.7

17.9

18.3

19.5

19.5

18.8

20

19.5

5.0

5.7

6.1

5.4

5

0.5

0.5

0.5

6.95

9.4

6.9

6.6

6.5

Ti02

0.14

0.11

0.12

-

0.4

CaO

0.7

1.1

0.4

(0)

0.6

0.3

0.6

0.5

0.7

MgO

13.7

10.8

12.0

13.2

11.0

10.7

8.3

8.9

10.6

Nap0 K20 H2OH20+ (+C02)

Total

(3)

508.175 Congo 699

FeO

MnO ~205

***

(1)

-

6.21 0.4 0.13

6.5

9.7

0.4

-

-

-

tr.

0.2

0.2

0.2

0.2

0.2

0.3

0.17

0.5

1.4

1.1

1.3

1.3

1.2

0.35

3.8

3.0

2.8

3.0

2.0

3.5

3.8

3.2

4.10

13.4

13.3

i0.8

10.2

12.0

10.1

10.3

9.4

10.96

98.7

98.73

99.9

99.1

98.5

99.4

99.4

100.1

100.05

0.02 0.30

0.03 0.39

0.02 11.22

tr.

0.03

0.16

0.18

-

-

0.33 0.17

tr.

15.2 99.30

(1) wet chemistry by M. Lenoble, Dbpartement de Pbtrographie, Paris; (2) wet chemistry by E. Lebrun, Ecole Normale Supkrieure, Paris; (3) X-ray fluorescenceby M. Lenoble. (*) Mean of two analyses after correction for CaC03 impurities; (**) Mean of two analyses for two sediments collected at the same site; (***) Mean of two analyses after correction assuming that Fe203 must be 20 %. Samples 700 and 704 are from New Caledonia; 601 from Senegal; 508 from French Guiana: 699 from Koukoure River mouth off Guinea.

180 Table 2. Comparison between chemical results for phyllite V, berthierine (Brindley, 1982).

and chlorites (Fosicr;1962). "ratioFez03"represen& value of Fe203/Fe203+FeO+Mg0, in percent. Phyllite V

Berthierine

Chlorites

range

mea n

range

mean

33-39

35.8

19-27

23.3

21-34 ( 1 )

*l2O3

5.5-12

9.8

18-28

22.1

12-26 ( 2 )

Fe203

18-25

21.2

3.2

0-13 ( 3 )

FeO

5-9

6.4

30-37

34.8

0-40

MgO

8-14

10.7

1-8

3.5

1-36

7.7

0-31 ( 4 )

SO2

r a ti o F e z 0 3 51-57

0-5.5

(1) two chlorites from 154 showed extreme values at 36.1 and 36.4% Si02; (2) two chlorites from 154 showed extreme values of 10.1% and 10.6% A1203;(3) three chlorites from 154 showed extreme values of 15.3%. 17.6%, and 20.2% Fe203; (4) one chlorite from 154 showed an extreme value of 44% of F%03.

older results little data are available concerning the valence of iron and the purity of the material analysed. Table 1 gives the results obtained for phyllite V from five different areas in the world. These results can be interpreted by comparison with data available for other iron-rich minerals. The contents measured for each major element are remarkably homogeneous; compared to the analyses given by Brindley (1982) for berthierines or by Foster (1962) for the various sorts of chlorites (Table 2), the ranges of variation of the oxide contents for phyllite V are much narrower. The silica content is extremely high compared to the usual berthierine as well as the large majority of chlorites. In contrast to that, the alumina content of phyllite V is constantly lower than for any berthierine and the large majority of chlorites. There is always much more Fe203,much less FeO, and more MgO in phyllite V compared to berthierines. Chlorites are too variable for comparison of FeO and MgO contents, but their Fe,O, contents are usually much lower than in phyllite V (Table 2). An important criterion for comparison is the ratio of Fe2O3 relative to the sum of Fe203+FeO+Mg0. From this point of view, berthienne shows a ratio about eight times lower than that for phyllite V; all the varieties of chlorite are also very different from the phyllite V mineral: the former always have a ratio clearly lower than 50% and usually lower than 25% while the latter has a nearly constant ratio slightly above 50%. In summary, none of the main cations of the phyllite V clay mineral phase has a range of variation similar to the commonly known values for natural berthierine; from a strictly chemical point of view, phyllite V can no longer be

181 considered as a berthierine similar to that of the sedimentary ironstones. Concerning chlorites, there is no presently described natural species or variety of chlorite with a Fe 0, content as high as that of the least ferric phyllite V. Concerning silica antalumina contents, the silica/alumina ratio in phyllite V is unusually high compared to most chlorites, but Ib (8 = 97")chlorites may have similar ratios. It is therefore suggested that phyllite V is a chemical component which has never been described in the mineralogical literature on natural clays, neither with the berthierines nor with the chlorites, from which it differs in high Fe3+content (Fe3+is supposed to be primary and not the result of an oxidation following sample collection).

Triangular diagramsfor phyllite V A simple way to compare mineral components is to use the rough composition in oxides given by the chemical analysis on bi- or tri-axis diagrams. Three triangles are proposed in Figure 13; the compositions shown in Table 1 are used for phyllite V (squares in the diagrams); the data for 14 berthierines by Brindley (1982) were added (stars); finally, the fields drawn by the 154 chlorite analyses gathered by Foster (1962) were calculated. These fields could be increased in order to include the low-alumina chlorites found later by Bailey and Brown (1962) and others. In diagram SiO /R3+O/R2+0,there is no common range between phyllite V on the one hand ana chlorites and berthierines on the other hand. The latter two are nearer to the MgO+FeO pole than phyllite V which is in between these and kaolinite (K in Fig. 13). Phyllite V has more silica than any berthierine and most chlorites, and essentially less FeO+MgO than any of these minerals. The diagram A1 0 /Fe O,/FeO+MgO also indicates complete lack of a common range of c?ie&cdcomposition between phyllite V on the one hand and chlorites and berthierines on the other hand. It is clear that a small proportion among the chlorites goes in the direction of the field of the phyllite V; but the proportion of Fe20 always remains far lower in these Fe203-rich chlorites than in any phyllite $mineral. Like in the preceding diagram, it may be noted that not all berthierines are in the compositional field drawn by chlorites and that all these berthierines are farther from the compositional field of phyllite V than are chlorites in general. Similar conclusions can be obtained from the diagram FeO/MgO/Fe203. The berthierines define a field restricted to the FeO comer; the chlorites are mostly widespread along the MgO-FeO axis; some are very rich in MgO, they are then very poor in FeO and Fe 0 , some others are poor in MgO. Among these, a small proportion contain bo& 8eO and Fe203in quantity and the compositional field obtained goes in the direction of the points characterizing phyllite V but, again, even the most extremely ironrich chlorites do not meet the compositional field characteristic of the phyllite V mineral of the verdine facies. These three diagrams clearly indicate that the green clay separated from shallow marine sediments of the present sea-bottom in the tropical area is different from berthierine as well as from any natural mineral species of the chlorite family described until now.

182

Figure 13. Comparison of the rough chemical composition of phyllite V with berthierines and some chlorites from the literature. The data for phyllite V (black squares) are taken from Table 1; berthierine (stars) from Brindley (1982); the chlorite fields were calculated from the 154 analyses gathered by Foster (1962). K indicates the position of kaolinite.

Chemical composition of phyllite C Phyllite C was obtained pure and in a quantity sufficient for chemical analysis from a single Occurrence in Baie de Rufisque (off Senegal). According to the X-ray diffraction pattern shown in Figure 8 (p. 170) and thin sections, the single crystallized impurity present within the mineralized faecal pellets is a low proportion of quartz and calcium carbonate. The chemical results shown in Table 3 contain a measure of both surprise and confirmation. The surprise consists in the generally comparable chemical composition of the phyllite C component compared to that of phyllite V (see for comparison sample 508-169 in Table 1, p. 179). If we calculate the points representing the chemical composition of 601-92 in the three triangles of Figure 13, they all fall at the contact of the array of points drawn by the phyllite V. This presumably indicates that the chemical compositions of the distinct environments where the two sorts of components form, are similar. The silica content, recalculated for the presence of extraneous calcium carbonate, is higher than for any phyllite V mineral. In the latter, it was already very high for a chlorite-type component. As a result, there is too much silica in the analysed mineral to give an acceptable structural formula for a true chlorite (Table 3). The iron content is smaller in phyllite C than in phyllite V; this is true for both the femc and ferrous iron. This is presumably the reason why the mineralized faecal pellets of 601-92 needed a higher intensity of the current feeding the electromagnet for an efficient separation from the non-magnetic sediment.

183 Table 3. Chemical composition of the phyllite C mineral, sample 601-92 purified from a Senegalese sediment. Rough

Recalculated

analysis Si02

Cations per f o r m u l a unit if c h l o r i t e

if s m e c t i t e 3.22

39.1

40.8

4.16

*l203

9.2

9.6

1.15

0.78 + 0.11

Fe2°3 FeO

16.2

16.9

1.30

1.oo

3.8

4.0

0.34

0.27

10.9

1.66

1.28

MnO

0.01

MgO CaO

10.4

Na20

traces

K20 Ti02

0.47

p205 H20+( +C02)

0.17 9.4

8.1

1120-

8.4

8.8

Total

99.56

2.15

(0) X+ 0.35

0.26

Rough analysis: Analysis by X-ray fluorescence supplemented using other techniques (M. Lenoble analyst). Recalculated: due to the presence of calcium carbonate (inclusions in the mineralized faecal pellets), corrected values have been calculated for the main cations. Cation per formula unit: contents have been calculated assuming a chlorite structure and a smectite structure.

In summary, the chemical composition of phyllite C is generally similar to that of phyllite V; all the considerations developed concerning the latter are therefore valid for the mineral discussed here: there is no chemical similarity with the berthierine (nor chamosite) minerals analysed from the oolitic ironstone facies. PHYSICO-CHEMICAL STUDY OF THE VERDINE FACIES

Thermal study of phyllite V The differential thermal analyses (DTA) undertaken on verdines from the Gulf of Guinea (Ogooue delta and Ivory Coast) gave reproducible curves. Figure 14 mainly shows two endothermal reactions at 125°Cand 550°C.Other

184 diagrams were made and the temperatures at 120 f 5°C and 550 f 5°C can be considered specific for the phyllite V mineral of the verdines from the Gulf of Guinea. No exothermal reaction was observed for OUT material in contrast to what is usually shown for ferriferous chlorites or serpentines. This is probably due to the low proportion of ferrous iron, which is not abundant enough to provoke a measurable exothermal reaction: Fe2+ => Fe3+.

cn cn

0

I

2

N

ul

Figure 14. DTA curve for the purified phyllite V from Ivory Coast. The temperatures for the reactions are shown in 'C. Table 4. Three interpretations of the two endothermal reactions shown by DTA curves of phyllite V. 125'

C

550'

C

1 chlorite-type mineral

adsorbed water

brucitic OH + t a l c layer OH

2 chlorite-type mineral

adsorbed water + brucitic OH adsorbed water

t a l c layer OH

3 serpentine-type mineral

octahedral O H

The two endothermal reactions are a problem because the first one occurs at too low a temperature for a chlorite. The curve is more characteristic of a 2: 1 mineral (adsorbed water extracted at 125"C,octahedral OH groups extracted at 550°C). Several solutions can be suggested to interpret these two endothermal reactions which can be due to three sorts of water: adsorbed water (chemical analyses give 3-4%); OH groups from an interlayer sheet; OH groups from a 2:l layer (Table 4).

185 It is suggested here that the first endothermal peak indicates a reaction which is much too large (if we consider the low quantity of adsorbed water shown by the chemical analyses) to be interpreted as resulting from the extraction of adsorbed water alone. Consequently, this large first reaction presumably represents the extraction of OH groups from an interlayer sheet. The second endothermal peak represents the departure of the octahedral OH groups of the 2:l layer of the chlorite. This second temperature is nearly normal (although low by about 100°C;see p. 196-197 below) for a chlorite; but the temperature of the first endothermal reaction, needs to be discussed further. The dehydroxylation temperature of the interlayer sheets of chlorites is known to depend on their main cation (Caillhre and HCnin 1960), Mg interlayers dehydroxylate at 640"C, Fe2+at 450-500°C and Fe3+at 250°C. The latter temperature is still higher than the one observed for phyllite V but low degree of crystallinity and small crystallite size should presumably decrease this temperature. This low temperature would suggests that the interlayer cation is mostly ferric iron.

Infra-red study of phyllite V Our technique is presented briefly in part D of this volume (p. 353-354). Samples of the phyllite V mineral purified from sediments collected from New Caledonia, Senegal, French Guiana, and Congo delta were analysed. The proportion of 0.75 mg of green pigment for 300 mg of mixture with KBr allowed us to obtain reproducible spectra defining comparable intensities of absorption i.e., the specific absorbance for the respective bands of vibration.

3Mx)

32W

2600

200

Figure 15. Infra-red spectrum of the Neo-Caledonian phyllite V. Double beam spectrophotometer with optical compensation, reference beam with pure KBr (300 mg), analyser beam with 0.75 mg of clay heated to 105'C. Note the comparatively small number of wide bands (interpreted in Table 5). The y axis shows the relative transmission (T) in % (linear scale) which is linked to the specific absorbance (A) following the relation: A = log l/T.

186 Table 5. Interpretation of the spectrum of the Neo-caledonian phyllite V. The specific absorbance is computed for our conditions of analysis. Interpretation

Frequency v c m - l

Specific absorbance

v Fe-OH

3560

0.1740

H-0-H

3420

6 H-0-H

1630

0.0535

v Si-0

1010

1.4540

V

6 Fe-OH

v Si-0

680 660

450

-

0.1055

0.7230

Concerning the major bands, the spectra obtained were comparable for all samples. The spectrum shown in Figure 15 (phyllite V from New Caledonia) represents one of the purest materials analysed. Some small bands were specific to one sample or another; they have been identified as resulting from the presence of traces of impurities such as quartz, kaolinite (see Chapter B4, Congo River deposit), or organic matter. The main absorption bands of phyllite V are first characterized by their small number (six) and are linked to three chemical groups: H,O, OH, and SiO (Table 5). The molecules of water give two medium bands i.e. stretching mode: vH-0-H at 3420 cm-' and bending mode: 6H-0-H at 1630 cm-l. These two bands remain present after a heat treatment at 110°C for 48 hours. The OH groups give two bands one for the stretching mode: vR-OH (3560 cm-') and the other for the bending mode: 6R-OH. The latter shows a primary maximum at 660 cm-' and a secondary one barely defined at 680 cm-'. The two largest bands are those indicating that the material obviously is a silicate. The stretching mode band: vSi-0 at 1010 cm-' also shows a small complementary shoulder near 990 cm-'. The bending mode band: 6Si-0 at 455 cm-' sometimes shows a secondary maximum at 470 cm-l. The analysed green pigment therefore shows a spectrum of a silicate which can be compared to the spectra known for other layer silicates. The wide stretching R-OH band at 3560 cm-' with a secondary maximum at 3420 cm-' linked with the interlayer bonded water can be compared with the spectrum of a nontronite (Farmer, 1974). According to this author, the comparatively low frequency of this band (3560 cm-') results from the abundance of Fe3+in the octahedral sites of the layer silicates. The remarkable width of this vR-OH band can be computed as the width (Av in cm-l) at mid absorbance; this width is equal to or higher than 120 cm-' in our samples. This is a criterion of low homogeneity of the octahedral site occupancy. Due to this high width, it is not possible to differentiate absorptions due to cations other than Fe3+.

187

c I " " 1-1

3500

1

3ooo

'

3

1

'

00

1

1

I

I

l

l

I

I

I 500

1wO

I

I

cm

Figure 16. Infra-red spectra of phyllite V after different treatments. U: untreated; AA: 1N acetic acid at 70'C for six hours; HA: 0.1N hydrochloric acid at 70'C for one hour; and H: heated to 490°C for one hour. Note the decrease of the band at 3560 cm-' in frequency when alteration goes on. Phyllite V from Senegal and French Guiana showed similar spectra.

The comparative weakness of the band near 3420 cm-' does not permit comparison of our mineral with chlorites for which the spectra show two bands of similar intensity for these frequencies (Farmer, 1974; Van der Mare1 and Beutelspacher, 1976). Below 1200 cm-', the frequency of the stretchin mode Si-0-Si is very low (1010 cm-'), lower than for a nontronite (1020 cm- ). All samples also show a similar R-OH band with two maxima at 660 and 680 cm-'. This band appears common to a number of &octahedral ferro-magnesian micas which show a specific band at 680 cm-l (see part D, this volume). Sensible modifications of the structure of the phyllite V were observed using X-ray diffraction after various aggressive treatments. Figure 16 illustrates the modification of the infra-red spectra when the green pigment is treated before analysis. The R-OH bands are the most affected; the absorbance progressively decreases from the untreated sample (U) -to the sample pre-treated with 1N acetic acid (AA) -to the sample pre-treated with N/10 hydrochloric acid (HA) and -to the sample pre-heated to 490°C for 1 hour (H).

f

188 The computation of the ratios between the absorbance of the stretching vibration mode bands of Si-0 and those for the two R-OH bands is very indicative of the modification of the spectra; the results are shown in Table 6. After the two acid leachings, small bands linked with the resence of organic matter become visible at frequencies above 1380 cm-?and below 3000 cm-' (Fig. 16). This is due to the fact that acid leaching dissolves a portion of the mineral analysed (and not the organic matter), therefore the 0.75 mg of material weighed for analysis contain a larger proportion of organic matter which becomes visible on the spectrum. The stretching mode band for Si-0 initially located at 1010 cm-' progressively shifts to 1025 cm-' after acid leaching and 1035 cm-' after heating. Concurrently, the increase of the absorbance ratio Si-Om-OH (Table 6) shows that the evolution is due to a relative increase of the Si-0 bonds versus the Fe-OH bonds. The acid leaching would thus partly destroy the octahedral sheets with a solubilisation of iron. This departure of material appears to be linked with a broadening of the Si-0 band initially at 450 cm-'. The single possible trace of reorganization of structure shown on the spectra obtained after pre-treatment (Fig. 16) is the appearance of a small band near 520-530 crn-'. Table 6. Mean values computed for the relative absorbance between the band at 1010 cm-l and the two R-OHbands at 660 cm-l (GFe-OH) and 3560 cm-l (vFe-OH). U, AA, HA and H as for Figure 16. Treatment

w Si-O/

6 Fe-OH

v Si-O/

v Fe-OH

U

13.2

8.0

AA

17.0

13.3

HA

32.0

12.5

H

39.4

34.5

In summary, the infra-red analysis has shown that little differences were visible from one sample to another when the material was pure. The comparatively low number of clearly drawn bands can be linked both to the abundance of Fe3+in the structure and to a very irregular occupancy of the octahedral sites by cations. The acid and heat treatments do not seem to provoke new bands in contrast to what was shown with X-ray diffraction analyses, which indicated new well-defined periodicities previously not clearly visible with routine techniques.

Mossbauer study of phyllite V The Mossbauer technique is very informative for minerals that contain iron. A preliminary study was undertaken at Stanford University on the Senegalese sample 601-58. Figure 17 below gives the spectrum obtained, and a simplified

189 104

535 .

cn t-

f

s 530. -I

a

I-

2 525 -

520

.

-2

-1

0

1

2 3 4 WPPLER VELOCITY (mm/s)

Figure 17. MUssbauer spectrum for the green grains from the Senegalese sediment 601-58. The spectrum has been interpreted here as the result of the sum of two doublets: one for Fe3' and one for Fe2'. A more complex nature of the spectrum is likely however.

interpretation assuming a single doublet for Fe2' and a single doublet for Fe3+. The peaks are rather broad, this could be due either to structural inhomogeneity or to multiple sites for the iron atoms. From the relative areas of the doublets, the Fe3+/Fe2+ratiocan be calculated as 73.6/26.4.This gives for Fe2O3: 21.4% and for FeO: 6.9%. These results compare favourably with the chemcal results of 21.7% and 6.6%, respectively (see Table 1, p. 179). The spectrum confirms that there is no iron in tetrahedral sites. The doublet deduced for Fe2+is symmetrical but that for Fe3+is asymmetric. The latter can be split into two doublets for Fe3' in two different locations. The intensity ratio between these is 0.895/0.765with the larger amount most likely in a less distorted and smaller octahedron. If we consider that the analysed clay mineral is part of the chlorite family, the smaller octahedra are apt to be the two cis M 2 sites in the 2:l layer and the symmetrical M4 site in the interlayer sheet. THE CLAY MINERALS FROM THE VERDINE FACIES

The diffraction peak near 14 A for phyllite V The diffraction peak near 14 A for the green clay mineral with a dominant peak near 7 A, is the most variable characteristic encountered in this study. Five different interpretations have been mentioned previously.

190

The first interpretation was discussed in the study of the Neo-caledonian green clay. In the course of that study, it was shown that the use of acid solutions, for making the clay free of its carbonate substrate, leads to an artifact (?). When too concentrated an acid is used, the ori inal peak at 7.2 A is partly or totally destroyed, and a wide dome between 14 and 20 A is created. This diffraction must be compared to the one which is developed during a heating treatment. Very similar characterististics have been encountered in the two cases i.e., destruction of the 7 A diffraction, development of a 14 A diffraction which, when the treatment becomes more severe, partly shifts toward 10 A. This %tifact" could be considered as a characteristic equivalent to that of the ferriferous chlorites which show a similar I; .haviour when heating treatments are applied. The second interpretation of the presence of a 14 A peak was also discussed in the case of the Neo-Caledonian deposit. It was due to the presence of foreign inherited chlorite. The nature of this second sort of peak at 14 A can be suspected from two criteria. The first is the fact that the peaks obtained at 14 A (and at 7 A) are much narrower than those obtained for the marine clay mineral, and their positions are different. The second is the fact that, under the binocular microscope, the grains giving these narrow peaks do not show any trace of the preliminary biodetrital substrate in which the marine green clay usually develops. The third possible interpretation of the presence of a peak at 14 A was suggested during the study of the Guianese sediments. The marine green clay frequently develops there on substrates which already are partly composed of a chlorite. The chloritized biotites, inherited from the geological formation crossed by the river feeding the shelf, cannot be separated from the authigenic marine clay because the same grains contain both components. Obviously, this possibility ma be recognized from the X-ray diffraction patterns (sharp peaks at 7 A and 14 and presence of a complementary peak at 10 A) as well as from the observation of the grains (presence of green accordion-like grains). The study of the Guianese deposit also permitted to propose a fourth interpretation: the presence of a double verdissement process with both the genesis of a 7 8, clay mineral and of a glauconitic smectite successively on the same substrate. This case may be suspected in the numerous deposits where the area of genesis of phyllite V at shallow depth is directly followed into the deep sediments by an area where glaucony forms. A moderate heating treatment permits collapse of the glauconitic clay to 10 A and eliminates the diffraction peak initially at 14 8, or less. The study of the Senegalese deposit has proved that there exists afifth possibility, i.e. an authigenic clay mineral distinct both from the glauconitic smectite and the phyllite V. This clay, still of marine origin, is characterized by a strong peak near 14 A on the X-ray diffraction patterns. The admixture of this phyllite C together with phyllite V has been observed in several Senegalese sediments. A hand-picking selection may help in identifying a mixture. The five interpretations outlined above from examples met in nature do not reject the possibility that the hump observed on the X-ray powder diffraction

f

x

191 patterns between 20 A and 14 A are organically linked to the crystallographical architecture diffracting more clearly at 7.2to 7.4A. The heating and acid treatments indicate that a concealed chlorite structure can reasonably be suspected for many of our phyllite V samples. The wellcrystallized iron-rich true chamosites, with a 7 A peak about ten times greater than the one usually seen for phyllite V, usually show a very small diffraction peak at 14 A (see Chapter A l , p. 17);even if chloritic, it would be abnormal that a peak at 14 A would be prominent in the X-ray diagrams of a poorly crystallized mineral like phyllite V.

Mineralogical interpretation of phyllite V Interpretation of the X-ray difiaction patterns (untreated samples) Film powder patterns have been used to supplement the view given by the diffractograms above in order to discuss the precise nature of the mineral from the verdine facies. An example of a film powder pattern is shown in Figure 18 which was obtained in Madison using monochromatic Fe X-radiation. The Senegalese verdine submitted to X-ray analysis was dark-green in colour and showed a small admixture of quartz. The film patterns were taken of individual pellets or infillings mounted on glass fibers without grinding. The resulting films show the 001: 14.4 A, 003:4.85 A and 005:2.90 A lines of a chlorite structure as well as the 002 and 004 lines.

Figure 18. X-ray film powder pattern of a phyllite V pellet from Senegal. The position of the diffraction lines are interpreted in A below the pattern. The two sharper lines are due to a trace of quartz impurity (Q) which was used as an internal standard for measuring the other lines.

The interpretation of the film is summarized in Table 7 which gives an indexing of the lines obtained. The indices listed c o n f i i a Ib chlorite of the monoclinic-shaped cell type, although the broadness of the lines suggests a large number of stacking faults tending toward the Ib orthorhombic-shaped cell. The pattern illustrated in Figure 18 is representative of all those that were taken from this sample. The intensity and broadness of the lines are identified in Table 7 as follows: very strong (VS), strong (S), medium (M), medium weak (MW), weak (W) or very weak (VW); broad (B) or very broad (VB).

192 Table 7. Interpretation of the film powder patterns obtained for phyllite V of the Senegalese sample 601-58. Distance

(A)

Intensity ( s h a p e )

h k l indices

M

001 002

4.85

vs vw

4.62

S

02. ; 11.

3.60

S

004

2.90

vw

005

2.66

MS

201

14.4 7.2

23 . 54 26

003

I

i

201 205 20!1 ; 007

2.05

W(B)

1.74

MW

205

1.550

S

060

1.518

M

062

1.440 1.420 1.320 1.310 1.28

I

vw

I I

0,0,10 064 0 , 0066 ,11 209

Figure 18 and Table 7 show that the individual particles that these data represent are probably chloritic in nature. Although not well crystallized, the structure of this Senegalese phyllite V is well enough organized that the interlayer sheet has a fixed relationship to the 2:1 layers above and below. That relationship is what ident.ifies the presumed chlorite as having the I b (B = 97") structure (terminology of Bailey and Brown, 1962). Imperfections in the form of 1) b /3 layer shifts are indicated by the presence of two-dimensional bands for reflections where k f 3n and 2) shifts of some layers from b-type superpositions to a-type superpositions relative to the interlayer sheet. The latter superposition types differ in whether the interlayer cation do (a-type) or do not (b-type) project normally onto tetrahedral cations of the 2:l layers below and above, and some disruption of the regular baba sequence along 2 of the ideal I b (B = 97")structure is indicated by the breadth of some of the 201 reflections. The 001 chlorite reflection at approximately 14.45 8, is large compared to that of most chlorites, but in fact is quite the usual case for the I b (13 = 97")

*

193 structure. Bailey and Brown (1962) state that the latter structure forms at the lowest temperature and has the least substitution of A1 for Si in tetrahedral positions. Thus, the layers are held together less tightly than in most chlorites, and the d-001 values are correspondingly larger ( ~ 1 4 . 4A). Application of this d-value to the graph of Shirozu (1958) indicates a tetrahedral composition of about Si, 5Ab.5. Mossbauer spectra indicate no tetrahedral iron present. The observed d-060 value of 1.55 A is also large and consistent with a high iron content. Shirozu's 1958 graph cannot be used to evaluate the octahedral iron content precisely from this d-value, because the graph is calibrated for Fe2+-richspecies rather than Fe3+.The 1.55 A spacing indicates about 2.0 Fe atoms for a ferrous chlorite from the graph, however, and the presence of a large proportion of the smaller Fe3+ ion should increase that value. The indicated value of 2.0 Fe atoms is in agreement with the analysed Fe2+and Fe3+contents of phyllite V listed in Table 9 (p. 198), however, it must be kept in mind that these verdine samples (except 601-108, 704-4 and 699) have undergone chemical treatment that may have leached the interlayer cations of the minerals preferentially; two of the unleached samples show the highest Fe atom contents: near 2.9. Further structural interpretation hinges on the origin of the Fe3+ now present. Primary ferrous smectites dredged from the Red Sea proved to be unstable on exposure to the atmosphere and oxidized to a ferric form with concomitant dehydroxylation of the interlayer to maintain charge balance (Badaut et al., 1985). Ferric smectites from altered continental basalts in some cases also are dehydroxylated and thus may have originated in the ferrous form (Daynyak et al., 1983). For an ideal chlorite OH content of 8.0 atoms per formula unit, the H,O+ content by analysis should range from about 10 to 13%, with the water value decreasing with increasing iron content (Foster, 1964). Table 1 (p. 179) shows that H O+ analyses for different phyllite V samples range from 9.4 to 13.4%. %he minerals are not appreaciably dehydroxylated, therefore, and most of the Fe3+must be considered primary. Furthermore, the environmental factors are relevant. In New Caledonia (samples 700 and 704), the Eh of the phyllite V-bearing sediment was measured at +0.2 V; in the Islands of Los (off Guinea, sample 699) the phyllite V-bearing sediments are presently shallower than 1 m at low tide; in these oxidizing conditions, it is dificult to imagine that a dominantly ferrous clay could be formed and.preserved. An Fe3+-richchlorite must have an appreciable &octahedral character for charge balance. This is born out by Table 9 in which the total number of octahedral cations is given as 5.0 or less per formula unit, instead of the 6.0 total expected for a fully trioctahedral chlorite. The values lower than 5.0 may well be due to preferential leaching of octahedral cations relative to tetrahedral cations during acid treatment. This would have the effect of decreasing the number of octahedral cations and increasing the Si cations, which is the trend that is observed in Table 9 for these specimens. An admixture of a 2:l clay mineral would have a similar result, however, and must be considered according to the HRTEM analyses.

194 For other chlorites in the literature with similar octahedral cation totals near 5.0, the octahedral sheet within the 2:l layer is usually dioctahedral and the interlayer is usually trioctahedral, but with enough trivalent ions present in the latter to give it the positive charge necessary to balance the negative charges of the tetrahedral sheets. Only two di-trioctahedral chlorite species with dioctahedral2:l layers and trioctahedral interlayers are known to date, and both are Al-rich. Cookeite is Al- and Li-rich and sudoite is Al-rich but Li-poor. The chlorite(?) assumed to be represented by some samples of phyllite V is a new species, therefore, if it can be confirmed to be homogeneous enough for these samples (see TEM study above) to warrant species status. The intensities of the 001 reflections are sensitive to the distribution of heavy (Fe) and light (Mg, Al) cations between the two octahedral sheets of the chlorite structure. Because the two sheets are separated along 2 by c /2, the diffracted contributions from the two sheets are exactly in phase for the even orders of 001 so that these intensities vary as a function of total heavy atom content only, irrespective of their distribution between sheets. If the heavy atoms are distributed symmetrically between the two octahedral sheets, the odd-order 001 intensities are independent of octahedral composition due to exact out-of-phase cancellation of the diffracted contributions from the two sheets. The odd-order intensities then are due to the diffracted conmbutions from the tetrahedral sheets alone. For an asymmetric distribution of heavy atoms between the two octahedral sheets, on the other hand, the odd-order 001 intensities have an added contribution that is a function of the difference in scattering power between the cations in the two sheets. Calculations using the composition of phyllite V from Table 9 having about 5.0 octahedral cations give even-order 001 intensities in good agreement with those observed in Table 7. This indicates the bulk chemical composition is approximately correct. Fourteen different distributions of heavy and light atoms between the two octahedral sheets were then used as models to calculate the odd-order 001 intensities. The best fit was for a model with a dioctahedral sheet in the 2: 1 layer of composition p 2 + 0 . 5 7 , Fe3+i.43)2.00-0-57and a trioctahedral But the fit interlayer of composition (Fe +0.25, A10.72. Mg2.03)3.00+~.~~. between calculated and observed intensities is not good enough to be considered correct. In particular, the calculated intensities of the 001 and 003 reflections are considerably larger than those observed, and it is not possible to decrease one without increasing the other. For example, adding heavy atoms to the interlayer decreases the 001 peak intensity but increases that of 003, and vice-versa. One possible explanation for the lack of fit between the observed and calculated odd-order 001 intensities is interstratification with the chlorite host of some 7 A 1:l type layers, or the admixture of an independent 7 A iron-rich mineral. The latter layers contribute zero intensities at the 28 angles of the odd-order 001 chlorite reflections, and thus would decrease these resultant intensities from the interstratified assemblage without appreciably affecting the even-order 001 intensities. Ahn and Peacor (1985) used high resolution transmission electron

195 Table 8. Interpretation of the film powder pattern obtained from the young verdine particles of the Islands of Los (Guinea). The clay mineral is a better than 90% pure 7A phyllite V. d (observed)

Intensity

Joint

1M

W

c h 1orite

7.15

VS

001

4.62

MS

020

4.50

MW

4.25

ww

3.58

S

3.34

MW

quartz

2.84

ww ni s

022

2.67 2.56

W

112

2.51

ww

2.42

M B

2.34

W

2.29 2.14

ww ww ww

2.02

W B

203, 132

1.90

202, 1 3 5

1.82

ww ww

quartz

1.74

W B

1 5 0 , 240

1.67

W B

203, 1 3 3

1.62

ww

1 5 2 , 241

1.550

MS

1.492

MW

1.43

WW B

1.38

ww

1.32

M B

14.4

2.24

1T

110

quartz 002

zoi,

130

200 201

202, 131 040

quartz quartz 202

060 204 06 2

quartz 204, 1 3 5 4 0 0 , 402

Specimen 699-AC-6 (light-greengrains from the Islands of Los) 1M a = 5.371, b = 9.316, c = 7.361 A, I3 = 103.87' 1T a = 5.369, b = 9.307, c = 7.172 A, I3 = 90' (indexed on orthogonal axes)

196 microscopy (HRTEM) for argillaceous sediments of the U.S. Gulf Coast region to prove intercalation of berthierine layers within a trioctahedral iron-rich chlorite host. Electron diffraction of these regions showed a reduced intensity of the odd-order 001 reflections of the chlorite as a result of this interstratification, similar to that observed for phyllite V of the Senegalese sample. Although in small proportion, a similar interstratification (7-14) 8, was evidenced by HRTEM study in the case of some samples of hyllite V (Senegalese samples); however microcrystals showing both the 7 and 14 8, spacings are always rare. Other samples (New Caledonia, Guinea) did not show any mixture of 7 A and 14 8, spacings, the former being present alone or the two structures being in different microcrystals (with little 14 8, specimens). X-ray film atterns have been obtained from phyllite V samples showing predominant 7 spacing on HRTEM photomicrographs. Individual light-green grains from the Islands of Los and New Caledonia have shown sufficiently pure 7 A phase f o r definition of a new mineral species. This species is characterized by the presence of both a monoclinic and a trigonal form; the monoclinic form is much more abundant than the trigonal form in sample 699 from Guinea (Table 8). In the Neo-Caledonian sample 700-5A, only the monoclinic form of the 7 8, phase could be identified, possibly due to Mgcalcite impurity. A common "impurity" in all these young phyllite V is the presence of poorly crystallized Ib (13=97")-typeof chlorite as described above. But this chlorite is less perfectly organized than in older verdine and is much less abundant allowing good X-ray diffraction patterns to be obtained of the 7 8, phase. An example of record is given in Table 8; it was repeatedly obtained in patterns of nine individual grains from the same bottle as the one used for chemical analysis (Table 1, p. 179).

B1

1

Intensity changes upon heating and chemical treatment of phyllite V With the preceding paragraphs as background i.e., the hypothesis that the phase called: phyllite V, is primarily a mixture of serpentine- and chlorite-like minerals with a progressive enrichment in chlorite for older verdine, it is convenient to discuss the changes in 001 intensities observed on heating and chemical treatment of phyllite V. Brindley and Ali (1950) showed that the great increase in intensity of the 14 8,001 peak observed on heating well-crystallized chlorite at 500°C to 550°C is due to reducing the scattering power of the atoms at the mid-plane of the. interlayer sheet. This is accomplished by dehydroxylation of the interlayer in which two-thirds of the interlayer OH groups leave the structure in the form of water at that temperature. The interlayer cations then migrate into the positions vacated by the OH groups in order to satisfy their coordination needs. This leaves only a small amount of scattering power (cations) at the middle of the interlayer, which is an asymmetry of distribution between the two octahedral sheets that causes an enhancement of the 001 peak intensity at the expense of the other 001 peaks. Poorly-crystallized iron-rich chlorites encountered in sedimentary rocks do not show a comparable enhancement of the 14 8, peak. Study at several petroleum company research laboratories shows that these poorly-crystallized

197 chlorites become amorphous on heating near 5OO0C, instead of simp1 dehydroxylating. The failure of these clay minerals to develop a visible 14 peak on heating, therefore, cannot be used as definitive evidence that a chlorite structure is not present. Mention has been made in previous sections that phyllite V after chemical treatment shows an X-ray diffractogram with a hump or dome of diffracted intensity extending between 14 A and 20 A. With continued acid treatment the hump may sharpen and move closer to 14 A. The diffraction hump indicates that a range of spacings is present. This could be due to primary interstratification within the chlorite host of other structural types or of expandable layers, disruption of the structure by acid to precipitate iron oxides or oxihydroxides in the interlayer, partial removal of the interlayer to provide additional expandable layers, etc. The observed enhancement of the 14 A peak upon continued acid treatment is suggested as due to preferential leaching of the interlayer cations. This reduces the scattering power at the mid-plane of the interlayer and thereby increases the 0 0 1 intensity, as in the case of heating well-crystallized chlorite at 500 to 550°C.

1

Interpretation of the chemical analysis From the above discussion, it can reasonably be assumed that some phyllite V samples have primarily a serpentine- or a chlorite-type structure. Therefore, a rough calculation of "structural formulae" can be undertaken. This exercise has two main limitations here. The f i s t concerns the presence of impurities in the analysed green grains of Table 1. Therefore, it is useful to remember here that all data for French Guiana (508) are apt to be slightly in excess for silica. In addition, it is probable that the Senegalese samples, and especially 601-108 according to the X-ray diffraction patterns, are a mixture of the minerals phyllite V and phyllite C, the former being predominant. The second limitation obvious1 concerns the real nature of phyllite V which is a mixture of chlorite-type (14 ) with a serpentine-type (7 A) clay mineral ("mixture" implies actual mixtures of separate phases as well as interlayered structures). However, the very similar chemical composition of young (i.e., dominant 7 A clay mineral) and old (i.e., 14 A-rich) verdine indicates that the two phases have very similar chemical composition. It is also useful to remember that the calculations done here (according to the method proposed by Foster, 1962) like any other method of calculation, do not allow us to distinguish the actual nature of the two octahedral sheets of chlorite which are fundamentally different in their anionic composition and could be different in their cationic composition as well; the data for the octahedral cations are therefore mean values for these two sorts of octahedra. Table 9 gathers the results of the formula calculations undertaken according to the data of Table 1. These calculations emphasize that phyZZite V has a small tetrahedral charge; this fits the large 001 distance (14.4 A observed for chlorite and 7.3 8, for serpentine-type mineral), on the X-ray diffraction patterns and might be expected for a very "young" material just being formed. However, the Neo-Caledonian mineral, the one for which a very recent origin is

x

198 Table 9. Calculated composition for phyllite V (cations per formula unit assuming a chlorite structure). Si

A l . IV

A l . VI

Fe3+

Mg

Fez+

H3+

22+

ZOCta

508.169

3.99

0.01

1.41

1.43

1.26

0.49

2.84

1.75

4.59

508.172

3.89

0.11

1.21

1.55

1.38

0.53

2.76

1.91

4.67

507.175

3.89

0.11

1.19

1.47

1.61

0.46

2.66

2.07

4.73

508.102

3.81

0.19

1.15

1.41

1.69

0.56

2.56

2.25

4.81

508.103

3.80

0.20

1.25

1.44

1.64

0.43

2.69

2.07

4.76

699

3.72

0.28

1.20

1.54

1.50

0.53

2.74

2.03

4.77

601

3.59

0.41

0.72

1.68

2.03

0.57

2.40

2.60

5.00

601 . l o 8

3.67

0.33

0.36

1.96

1.81

0.88

2.32

2.69

5.01

700.3

3.63

0.36

0.40

1.80

2.18

0.70

2.20

2.88

5.08

700.4

3.60

0.40

0.30

1.99

1.86

0.90

2.29

2.76

5.05

likely is not the one with the lowest charge. The very low tetrahedral charge obtained for sample 508-169 may well be considered as an artifact due both to the presence of a small proportion of quartz and to admixtures of 10 A phase making the silica content too high. In any case, there is no chlorite listed by Foster (1962) with a layer charge as small as the highest charge calculated for the presently discussed components. The formula calculations allow us to emphasize interesting relationships between the proportions of octahedral cations. The most abundant octahedral cation is either Mg or Fe3+and the least abundant is usually Fe2+; Fe3+ is always more abundant than Al. According to the comparison between the total of the trivalent cations (Z3+) and the total of the bivalent cations (Z2+)there is no systematic dominance of a dioctahedral or trioctahedral character. The formula calculations in Table 9 have been arranged as a function of an increasing number of octahedral cations (Z octa) from the top to the bottom. As a result, the other cations are quite well ordered too: Si, A1 VI, and R3+ decrease, the other cations increase from the top to the bottom. However, it seems that we may propose a cut in this series between samples 699 and 601. The cut is visible for the sum of octahedral cations, the Mg content or the sum of Fe atoms.

Mineralogical variation of phyllite V as a function of time The degree of evolution of the facies is difficult to observe from the mineralogical data. No modification of the X-ray diffraction patterns has been recognized for untreated samples, i.e. the 7 A peak has a similar shape for light-green Recent and shallow material compared to the presumably older dark-green and deep lying material. However, it seems that the resistance to alteration processes (acid leaching and heating treatments) is better for the more evolved green material compared to the more recent light-green clay. The following Table 10 summarizes some of the criteria which may be considered.

199 Table 10. Comparative characters of phyllite V following aging without burial. Criterion

Young material

O l d material

~~

Habit

lnside substrate

Free grains

Colour

Light-green

Dark-green

Hardness

Soft

Compact

X-ray behaviour after N / 1 0 HCl (leaching a t 7OoC for 1 hour)

S m a l l diftuse hump near 14A

Large hump around 141i

X-ray behaviour after heating a t 49OOC for 2 hours Colour after leaching with 6 N HCl a t 22'C

Diffuse hu-mp near 14A

Large peak a t 14A

White

Greenish t o white

Spacings observed during HRTEM study

Predominant 7 A

Various spacings including 7 A or l 0 . i or 14A or (7/14)A

It is possibly useful to remember here that one of the problems with which we are faced in the present study, is the lack of detailed data dealing with a component resembling phyllite V in the literature. This may well be due to the absence of adequate study: the sub-reefal facies are usually more studied for their carbonate component than for their clay component. However, this may also be due to the absence of a similar material in ancient series (sediments older than 20,000 years ago). This absence, in turn, could be due either to exceptional conditions of environment linked to the Recent time or, more likely, to the systematic destruction of the phyllite V during burial diagenesis or later. The only study which rejects the latter hypothesis is the description by Porrenga (1966) of a "chamosite" similar to the superficial material in Miocene sediments cored at depth as great as 2300 m below the sea-watedsediment interface off the Niger delta. However, a confirmatory study would be needed. The evolution of the most frequent clay mineral of the verdine facies during burial remains an open question; our experiments have shown that it is fragile and should be modified during early burial diagenesis. The HRTEM study allowed us to suspect the main probable evolution of phyllite V with time. Two differences have been observed: the Neo-Caledonian and Guinean "young" verdines are monomineralic and the 7 8, spacings are alone or widely predominant; the Senegalese and Guianese "old" verdines are multimineralic and the 14 A spacings are more abundant. It is of interest to compare these observations with the chemical data interpreted in Table 9. On the one hand, the Guinean phyllite V resembles the Guianese phyllite V; on the other hand, the Neo-Caledonian phyllite V compositions resemble more the Senegalese ones. Therefore, the apparent crystallographical difference between young and old phyllite V observed during the HRTEM study does not appear to be concurrent with a chemical change. Our present interpretation is that the primary clay formed would show 7 A spacings, but this "unstable" iron-rich clay would rapidly recrystallize still near the sea-bottom and without chemical changes, into one or several other clay

200 minerals showing either a true 14 8, chlorite-type spacing or a 10 A spacing, or intermediate mixed-layer (7- 14) A structures. There are diverse problems with the well-shaped 10 A spacin s observed in ultra-thin sections, 1) there is no corresponding well-shaped 10 peak visible on X-ray diffraction patterns (or at least the 10 A peak is not really important), and 2) the small potassium contents measured in the "authigenic" green clay from French Guiana appear smaller than what they should be according to the relative abundance of the 10 A spacings shown by the HRTEM photomicrographs.

i

Mineralogical interpretation of phyllite C The X-ray diffraction pattern of phyllite C is characterized by an extremely intense 001 reflection with a d-value between 14 A and 15 A. Such an intense reflection can only occur if there is a minimum of scattering power in the interlayer region of a 2:1-type layer. The observations that the layers expand upon solvation with ethylene glycol and contract toward 10 A upon heating indicate that phyllite C partly has a smectitic behaviour. Phyllite C does not solvate completely, however, as the expansion is only to about 16 A rather than the 17 A expected with ethylene glycol. Likewise, the collapse upon heating is not complete as there is a band of intensity extending from 14 A to 10 A. With acid treatment at different temperatures (Fig. 8, p. 170), the 14 A peak largely disa pears into a continuous background extending from about 10 8, to over 20 whereas the 7 A peak remains intact. These observations suggest an inhomogeneous structure having variously expandable layers that are interstratified with non-expandable 2:l and possibly 1:l layers. Some of the 14 A thick layers definitely do not expand with solvation nor collapse with prolonged heating and may be considered to have chloritic interlayers, some of which may be incomplete. Therefore, the presence of five different families of structures: chlorite, chlorite with incomplete interlayer sheets, illite, smectite, and serpentine is suspected above. However, the thicknesses of three of these five layers are approximately 14 to 15 A in the untreated state, and this gives rise to a sharp 001 diffraction peak. Solvation, heating, and acid treatments change the thicknesses of the different kinds of layers in different ways, however, and the range of resultant spacings in the treated samples gives rise to broad diffraction peaks at intermediate spacings. In summary, according to X-ray diffraction, phyllite C could be primarily a smectite. Some characters are intermediate between smectite and chlorite however, and the mineral could also be considered as of a swelling-chloritetype mineral; but, usually, the minerals referred to swelling-chlorites swell and collapse less than phyllite C. From a chemical point of view, phyllite C cannot be a chlorite (too much silica). The HRTEM study confirms that 14 8, spacings are predominant. No mixed-layer crystals were observed. A few crystals show either a (7+7) A or an unclear (9+5) A subdivision. The obtained images resemble smectites more than chlorites.

8,

201

Comparison between phyllite V and chlorites The literature dealing with chlorite is abundant because this clay mineral family comprises a variety of minerals which can be met in very diverse environments. From a chemical point of view, two main poles can be distinguished: the magnesian chlorites for which the ratio Fe2+/R2+is lower than 0.2 and the ferroan chlorites for which the same ratio is around 0.8 or more (Fig. 19). Between these two, the intermediate chlorites constitute a large portion of the minerals referred to that family for which the Fe,O, content is usually smaller than 14%. If we compare these data with the ones available for phyllite V (Fig. 19) very interesting conclusions emerge. The phyllite V components have a constant Fe2+/R2+ratio and, from a strictly chemical point of view, can be considered as magnesian "chlorites" at the boundary between magnesian and intermediate chlorites and not as ferroan chlorites. For these specific minerals (magnesian true chlorites) all the FeZO contents are below 9% and 73 of the 75 samples reviewed by Foster contam fess than 4.5% of Fe203.In contrast to that, all of the phyllite V samples have a Fe203content of about 18% or more. In summary, although they are comparatively rich in iron, the phyllite V components have also an important proportion of magnesium in the octahedra and this places them in a compositional field which is again entirely void of other chlorites in Figure 19. 0

0

0

0.8

-

a

>0.4-.

k

0.2 .

0

Ferroan chiorlte

0

0

-

0

0

0

..

. .. . . .

. ... . .....

0

0

0

A---

.

0

.

O

.

. lntermedbte

chlorite

Magnesiol chlorite

.

.=. ..

.

10

Fez03

%

m .

20

Figure 19. Diagram showing the specific chemical composition of phyllite V compared to chlorites for the relation Fe2+/R2+ratio versus F 9 0 3 content. (Redrawn from Foster, 1962 and supplemented)

In order to propose a normalized classification of the chlorites, Foster suggested compositional fields in the diagram Fe2+/R2+ratio versus tetrahedral Si. This diagram emphasized the substitutions in octahedra versus those in tetrahedra. In this diagram, reproduced in Figure 20, the chlorites where Fe,O, is higher than 4% have been distinguished (open circles) from the ones where the Fe20:, is low (dots); and the chlorites very rich in Fe203(10% or more) are also dstinguished (stars). One may note that the points representative of the

202

phyllite V (black squares) are not especially near the area where the open circles nor the stars are abundant. Once again, the compositional field of phyllite V is outside the fields attributed to the previously described natural chlorites. Especially, the area where Foster suggested the name chamosite is far from the area where phyllite V lies. Moreover, the Fe203-richest chlorites are mostly characterized by a comparatively high tetrahedral substitution (Si = less than 2.70 per formula unit) in contrast to the phyllite V samples which all show a low tetrahedral substitution. Other more recent classifications could be considered. But it seems that the conclusions will remain invariably the same i.e., the compositional field drawn by the components here identified as phyllite V is outside the fields usually attributed to any chlorite in general and to chamosites more specifically.

Figure 20. Location of phyllite V in the classification of chlorites proposed by Foster (1962); redrawn from that author. Note that phyllite V is far from other F%03-richchlorites as well as from the field attributed to chamosites. The three "chamosites"from Chamoson quoted in Delaloye and Odin (this volume, Chapter Al) are shown together with the similar chlorites from the other Swiss ironstones.

CONCLUSIONS

-The verdine facies comprises two separate series of clay minerals with a common green colour, general habit, and shallow marine environment of formation. These two series have been called phyllite V and phyllite C . Initially thought to be monomineralic it is now known that phyllite V comprises several mineral species. As a group, phyllite V is sensibly more magnetic than phyllite C (Table 11). -X-ray diffraction patterns allow an easy distinction between phyllite V, characterized by a dominant, although comparatively poorly shaped peak near

203 7.2 A, and phyllite C characterized by a very strong peak near 14.5 A. -HRTEM study allowed us to subdivide phyllite V into different sorts of minerals, the variety of which is presumably linked to the evolution of the green grains. The youngest grains show crystals with a dominant monomineralic 7 A spacing. The oldest grains reveal the presence of a variety of structures, sometimes monomineralic and including the still abundant 7 A phase, but sometimes interlayered (Table 11). Table 11. Mineralogy of the verdine facies. Phyllite V family Young sediment

. Magnetism . X-ray 7 14

.

Phyllite C family

O l d sediment

high

high

medium

diffraction

A A

peak

dominant

dominant

comparatively s m a l l

peak

diffuse to s m a l l

s m a l l to high

dominant and well-shaped ( p a r t l y s m e c t i t i c behaviour)

dominant to exclusive

common

HRTEM study

pure 7 pure 14

A spacings A spacings

r a r e t o absent (14

pure 10

A

spacings

composite ( 7 + 1 4 )

. Chemistry

. Nomenclature

A

A

common = (9+5) A )

rare dominant to exclusive ("smectite")

r a r e t o absent

common

rare

present

present

absent

no difference whatever young or o l d "7 A p h y l l i t e V " unnamed

mixture with "chlorite-type p h y l l i t e V"

s l i g h t l y more s i l i c a ; more adsorbed water ; less F e z 0 3 and FeO Smectite to s w e l ling-chlorite

-The most important result from the HRTEM study is that a pure ubiquitous clay mineral phase with a unique 7 8, spacing has been separated from some samples. This confirms that a specific previously unrecognized clay mineral of the serpentine group exists and can be defined from the verdine facies: this mineral is sufficiently dominant in certain samples for the bulk chemical analyses to be representative of its real structural formula. -Similarly, the HRTEM study confirms that a chlorite-type mineral exists. Finally, the HRTEM study reveals the presence of a mineral with 10 A spacings which was not clearly visible on X-ray diagrams either using the randomly oriented powder or using the single grain-camera technique. -X-ray powder film patterns have confirmed the dominant presence of 7 A structures in light-green grains from young deposits. -The chemical study indicates that both phyllite V and phyllite C mineral families are characterized by an abundance of iron which is dominantly in the ferric state. The chemical data do not allow us to distinguish between the various sorts of phyllite V.

204 -The chemical composition of the phyllite V or phyllite C minerals of the verdine facies does not correspond to that shown by the green clay minerals from oolitic ironstones. Therefore, it is emphasized here that neither the term berthierine nor the term chamosite are appropriate to designate the green clay minerals from the Recent shallow marine sediments. -For the present time, it appears useful to suggest the following designations. The young 7 8, mineral could be called the 7 A phyZZite V . The 14 A clay mineral present in the more evolved facies could be called the 14 A phyllite V the chlorite structure has been clearly identified; however, the latter material has not yet been obtained pure. -The 7 %, phyllite V has been obtained nearly pure and will eventually need a more specific name. This 7 A mineral would fill a vacant niche in the clay mineralogy classification because it is half dioctahedral and half trioctahedral as well as being rich in ferric iron in its natural deposits. -The phyllite C green clay mineral shows a behaviour which is intermediate between a smectite and a swelling-chlorite. The HRTEM stud has not been able to identify precisely the clay mineral(s) which shows 14 spacings and has a partly smectitic behaviour. Presently, we are led to interpret phyllite C as a possible mixed-layer (14 8, smectite-14 A chlorite) although no interlayered (9+5)pairs have been observed. -The study will have to be pursued further; however, important progress has been realized during the last four years since the time when the shallow marine green grains were considered to be composed of berthierine or even glauconitic minerals by most authors.

K

AKNOWLEDGEMENTS

During the last four years, the problematic nature of phyllite V has been discussed from diverse points of view with G.W. Brindley, J.J. Fripiat, G. Millot, and R. Prost. The editor deeply aknowledges the comments by these experts. Miss E. Lebrun and Miss M. Lenoble are thanked for their chemical analyses and 0. Fay for photographic processing.

Note added in proof The 7A phyllite V mineral has been approved as a new mineral species by the Commission o n new minerals and mineral names of the International Mineralogical Association following presentation by Professor Bailey. It will be published in the paper by S. W. Bailey : "Odinite, a new dioctahedral-trioctahedralFe3+rich 1: 1 clay mineral", Clay Minerals, 1988,23, 3, (available in November 1988). The present editor acknowledges this honour to become the 'root' for a mineralogical species and thanks Professor Bailey.

205

Chapter B6 GEOLOGICAL SIGNIFICANCE OF THE VERDINE FACIES by G. S. Odin and B. K. Sen Gupta In the context of the descriptions of the verdine facies in previous chapters, we now discuss the mode of origin and the geological significance of verdine. BROAD FEATURES OF THE ENVIRONMENT

The known verdine deposits, with one exception, occur in environments that are fully marine. In the Casamance River estuary, the Occurrence is inland, but the salinity here is that of the open sea. In general, the present evidence indicates a marine origin for the verdine facies. Verdine has been collected from diverse depths, In the Casamance estuary, the Baie de Rufisque (south of Cap Vert), the Islands of Los (Guinea), and the Ogooue estuary, the observed depths are sometimes less than 5 m. In French Guiana and in areas north and south of Cap Vert, verdine has been locally found in water depths of 100-200 m; these deeper occurrences are possibly relict or related to current transportation perpendicular to the coast. Overall, a depth range of 20-60 m seems to be most favourable for the formation and preservation of the green clays specific to the verdine facies. Porrenga (1967b) suggests that temperature is the critical factor for the deposition of verdine off the Niger delta. For the Ogooue delta deposit, however, the influence of lower pH fluvial water may be more significant (Giresse, 1976b). As discussed later in this chapter, an adequate supply of iron also seems to be a critical factor. The typical verdine facies is shallow marine, and is frequently replaced by the glaucony facies at greater depths; a reverse situation has never been encountered. Thus, the verdine-glaucony facies transition may be used for the palaeobathymetry of ancient sediments. On the basis of present information, this transition would represent a depth between 50 m and 80 m. As reported in Chapter B4 (p. 157), the known occurrences of the verdine facies are from the tropical belt (Porrenga, 1976a; Odin and Matter, 1981). Specifically, the eleven deposits listed (Table 1) are located between 22"30' S (New Caledonia) and 16"N (Senegal). Apparently, this tropical distribution of verdine goes back to at least the last 20,000 years. The estimated ambient sea-water temperatures are above 2WC, probably close to 25°C. LOCAL SETTINGS

Sediment Verdine-bearing sediments may be sub-reefal, e.g., off New Caledonia,

206 Mayotte, and French Guiana (Pujos and Odin, 1986). Off Senegal and Los Islands, however, the development of verdine is not in the proximity of a reefal facies. Thus, the verdine facies may be linked to a high water temperature, but not necessarily to a reefal environment (Chapter B2). Verdine occurs in biodetrital carbonate sediments (e.g., in reef lagoons) and in silicic sands that contain a small amount of carbonate, usually shell fragments. These diverse sediments of various mineralogical compositions generally consist of sand-size grains, with less than 10% silt or clay. This suggests an unrestricted circulation of sea-water in the outer sediment where verdine forms. The accumulation rate of verdine-bearing sediments is slow. The green clays occur mainly in front of the marine delta where the bulk of the detrital components accumulate. Where verdine forms in shallow coastal waters, the detrital influx is usually restricted because of sediment trapping in swamp and mangrove environments. Thus, the verdine facies may be present over deltaic sediments, but there is no significant deposition of terrigenous particles during its development. However, examples of slow, conconiitant, biodetrital accumulation are seen in lagoonal carbonates and faecal pellets constructed mainly with clay-size particles.

Hydrology and physiography

Figure 1. Locations of the verdine facies and major river influxes along the western coast of Africa. Detailed information on the verdine deposits is given in Chapters B1 to B4. Arrows indicate the dominant surface currents along the coast. B = Bandama River; Ca = Casamance estuary; Co = Comoe River; K = Konkoure River; S = Sassandra River.

207 Table 1. Locations and estimated extents of verdine-bearing basins compared to 1) lengths of associated rivers (1-8, for passive margin continental shelves) and 2) adjacent continental drainage arms (9-11, for islands with young relief). RiverfCoun try

Ocean

Latitude

1. Koukoure

12' 30 N

3 . Cornoe

E. Atlantic E. Atlantic E. Atlantic

4 . Ogooue

E . Atlantic

5 . Senegal

8 . Amazon

E. E. E. W.

9 . Mayotte

2 . Casamance

6 . Niger 7 . Congo

9' 30 N

Verdine-bearing Basin ( k m 2 )

River length or basin area

-10

250 km

>

?

200 km

5' N

102

s

102

970 krn

104

1700 krn

10

700 km

Atlantic

14-16'

Atlantic

4-5'

N

2.104

4200 krn

Atlantic

5-6'

S

103

4640 krn

Atlantic

0-10'

105

7025 krn

N

N

W . lndian

12-50 S

1 0 . New Caledonia

W . Pacific

22-23'

1 1 . Sarawak

S . China Sea

3-4'

-10

S N

2 . 1 0 2 km2

102

>

104

>

10

krn2

10

krn2

Bottom waters are found to be relatively rich in oxygen in areas where verdine is presently forming or has recently formed. This is particularly true for the shallowest deposits. Moreover, the overall sandy character of the sediment ensures the movement of this oxygenated water within the sediment. Available analyses (Chapter B1, p. 59) c o n f m that this pore-water has a positive Eh, particularly in deposits where goethite is associated with verdine. The relation between fluvial influxes and verdine deposits is well illustrated along the Atlantic coast of Africa (Fig. 1). In the tropical part of this coast, verdine deposits have been identified in the immediate vicinity of most large rivers. Furthermore, the precise location of the deposit on either side of a river mouth is related to the general direction of the locally dominant surface current. Off Senegal, the main current is southward and verdine is known only south of the river mouth. In contrast, verdine is known only north of the Congo River mouth where the dominant surface current is northward. Data on verdine occurrences at river mouths are given in Table 1. The first eight areas are parts of nearly flat continental shelves, listed in order of increasing length of the main river feeding the basin. The last three areas in Table 1 (9-1 1) are oceanic islands with comparatively young relief, and are listed in order of increasing drainage area. These criteria have been selected because 1) data are available, 2) the present length of the main river of an eroded area generally reflects the volume of the continental water supplied during the last 20,000 years, and 3) the measure of a continental drainage area with young relief (and significant erosion by many rivers) is, overall, a good representation of the continental water supply. In essence, Table 1 compares the amount of fresh-water draining into a marine basin with the surficial extent of the related verdine deposit. In each of the two groups of occurrences in Table 1 (1-8 and

208 Table 2. Composition of the water in the Kourou River estuary. Two representative analyses are given for each site; the values, given in m a , are influenced by the tide and the season. (After Roche, 1977) D i s t a n c e up t o t h e coast l i n e

Si02

M g ++

Fe”’

Salinity

PH

28 krn

10.1

0.5

0.8

14.0

7.1

28 krn

10.0

0.8

0.9

16.0

6.8

16 k m

8.6

59

0.02

2,800

16 k m

7.6

420

0.02

12,000

7.6

1 . 5 krn

4.4

1170

0.08

31,000

7.8

1 . 5 km

3.5

1040

0.05

27,400

7.5

Conductivity

7.2

mhoskm 259:

Figure 2. Iron and silica in the Kourou River estuary. The conductivity (x-axis) is related to the salinity of the water resulting from a mixture of sea-water and ion-poor tropical fluvial water. Iron in solution precipitates in the estuary well before reaching the coast line. Each dot represents a measurement. Asterisks represent theoretical values calculated on the basis of a simple mixing of fluvial and marine waters. (After Roche, 1977)

9-1 l), the sizes of the verdine deposits generally reflect the amounts of the continental water input. The deposit at the mouth of the Congo River does not fit this pattern. It is possible that, unlike the other rivers in Table 1, its water input is not proportional to its length. On the other hand, the areal extent of

209

I

-5m FLAT LAND

Om FLAT

-60m CONTINENTAL SHELF

Figure 3. Sedimentological model for the verdine facies at a tropical river mouth. The extent of the facies varies by two orders of magnitude, depending on the amount of the fluvial input. The Occurrence is common today at the boundary between a large peneplain and a flat continental shelf. 1: zone of iron precipitation; 2: mangrove; 3: swamps; 4: sandy, deltaic sediment; 5: demtal clay; 6: zone of verdine formation.

these verdine-bearing sediments may have been underestimated because of insufficient sedimentological and mineralogical data. Overall, however, the distribution of data in Table 1 leads to the conclusion that the formation and size of verdine deposits in Holocene sediments are directly related to the volume of local fresh-water discharge onto the continental shelves. The presence of iron in verdine deposits is also related to the influx of tropical fluvial waters. The dissolved iron of these waters is often removed by precipitation at river mouths, with an increase in salinity and pH. For instance, large amounts of Fe and Si are delivered to the sea by the coastal Kourou River of French Guiana, which drains a dense tropical forest for more than 100 km. Table 2 gives the SO,, Mg, and Fe concentrations at three different sites of the Kourou estuary. The tidal effect is felt as far as 80 km from the coastline, but the salinity is significant only over a much shorter distance. Figure 2 shows the relation between dissolved iron in the Kourou estuary, and the conductivity of the water which reflects the salinity and the sea-water contribution. The iron content abruptly decreases within a zone where little sea-water is admixed with the fluvial water. A simple mixing model of iron-rich fluvial water and iron-poor marine water (Fig. 2) also suggests rapid iron preci-

210 ITROPICALCLIMATE

I

WARM

SEA

WATER

-30m LAND

F L A T

CONTiNENTAL

Land Vcrdlne tacler

0

Volcanics

D

[71

GSO 8;

SHELF

Figure 4. Verdine facies related to volcanic outcrops. Except for the fluvial input, the situation is similar to that of Figure 3. A moderately rainy climate is necessary for the supply of weathered iron into the basin. This iron is precipitated on the sea-bottom in the immediate vicinity of the volcanoes.

pitation when the conductivity of the water increases from about 0.2 to about 0.4 mho/cm. The observed silica concentration in this estuarine water (Fig. 2) is clearly the result of a simple mixing of the fluvial and marine waters; values calculated on this basis (asterisks) agree with the observed values. A large quantity of iron is flocculated at the bottom of the river, near the river mouth. This precipitate is pushed by the river toward its mouth at low tide or during the rainy season. The iron thus concentrates in the sediment in the immediate vicinity of the river mouth, although the sea-water is iron-poor. Because it is not in a stable, crystalline form, the iron is available for the formation of synsedimentary minerals at the river mouth. Locally, goethite will crystallize in this zone of primary iron deposition; elsewhere, a verdine facies will develop (Odin, 1975b). A schematic representation of the process is given in Figure 3 (see also Fig. 1, p. 403). The availability of iron within adjacent fluvial basins also influences the formation of verdine-bearing sediments. Among the eleven deposits listed in Table 1, at least four (New Caledonia, Mayotte, Los Islands, and the Baie de Rufisque) are adjacent to voZcanic rocks subject to vigorous chemical alteration. We suggest that the weathered volcanic iron-rich rocks favour the development

21 1

_ _ _ _ _ _

- - -

i

l0h YOUNG RELIEF

*5m

0

-30/-60m LAGOON

I

Om ]OPEN SEA

Figure 5. Verdine facies on an active margin, particularly around a volcanic tropical island. Terrigenous sedimentation is absent or rare (coastal trapping by swamps or mangrove), but a slow bioclastic deposition occurs. Where the sea-floor is steep, a barrier reef is necessary in order to avoid the dispersion of iron.

of the verdine facies in adjacent marine basins. It is uncertain whether the current that flows near the Koukoure River mouth (and transports some detrital material and fluvial iron southward) really influences the composition of the very shallow water-mass off Los Islands; more probably, the nearly closed basin within the islands receives its supply of iron from the surrounding volcanic rocks. Baie de Rufisque, south of Dakar, is another example where the possibility is remote that a river (Senegal) is the source of iron in the basin, because the sediment transport is blocked by Cap Vert (Chapter B2, p. 83). The weathering of the .surrounding volcanic rocks would explain the presence of iron favouring the formation of verdine facies in this bay. The role of weathered volcanic rocks in the formation of verdine is illustrated in the model proposed in Figure 4. Finally, the growth of the verdine facies may be favoured when an active margin, with a young relief and a steep submarine bottom, is coupled with an offshore barrier. The resulting marine lagoon will trap iron weathered from the land. The precipitated iron will preferentially accumulate in the quieter, deeper zones of the lagoon. The situation would be even more fzvourable when the weathered terrigenous material is volcanic, and this situation is common in oceanic islands; the corresponding model is proposed in Figure 5.

212 SUBSTRATES

and MICROENVIRONMENT

Necessity of substrates The presence of an initial granular substrate is critical for the development of the verdine facies; the known association of the facies is restricted to sand-size particles on the sea-floor (Porrenga, 1967a; Odin and Giresse, 1972; Pujos and Odin, 1986; Chapters B1 to B4). Diffuse authigenic green clays have not been reported, but the sample preparation process used might have hindered the detection of small amounts of diffuse, intergranular, synsedimentary, authigenic green clays. Such clays, in small amounts, would remain indistinguishable from inherited kaolinite with which they would be admixed. As in the case of the glaucony facies (Odin and Matter, 198l), verdine is known to develop in four settings: as infillings of microfaunal shells, by replacement of biodetrital carbonate remains, impregnation within faecal pellets, and along cracks, fissures, or cleavage surfaces of mineral grains. Infillings of microfaunal shells constitute the usual mode of Occurrence in marine lagoons and off Senegal (Fig. 6, and Chapter B2). Partial replacement or incrustation of bioclasts such as shell fragments (Chapter B4) is sometimes the dominant habit, and is frequently coupled with infillings of micropores of biogenic carbonate remains. Abundant faecal pellets were observed in numerous deposits off the western coast of Africa (Fig.6, and Chapter B4). Off the northern coast of South America (Chapter B3), the growth of phyllite V is protected mostly by mineral debris (biotite flakes). The mineralogy of these substrates is diverse, and no link is observed between substrates of particular mineralogical compositions (e.g, kaolinite) and specific clay minerals (e.g., 7 A phyllite V) of the verdine facies. Consequently , the hypothesis of transformation implying that kaolinite or femferous kaolinite is a necessary or useful precursor for the genesis of phyllite V, as has been suggested for ancient berthierine (Bhattacharyya, 1983; Van Houten and Purucker, 1984), is not valid. This conclusion is sup orted by the fact that 7 A inherited clays (kaolinite-rich faecal pellets) and 10 mica (chloritized biotite) are both suitable substrates for verdine. Furthermore, carbonate substrates are particularly common, although calcite can hardly be considered a "precursor" for the development of a 7 A (for phyllite V) or smectitic (for phyllite C) clay. The implication is that the green clay minerals of the verdine facies arefomed by crystal growth; the substrates simply provide a physical environment that promotes mineral growth. Two other observations are relevant in the context of the relationship between substrates of verdissement and mechanisms of clay mineral genesis. First, in contrast to the case of glaucony, verdine minerals rarely occur as films; there are no green boulders or hard-grounds known from this facies. Second, an oolitic texture has never been reported from a verdine facies or from carbonate sediments in its proximity. The only verdine-bearing continental shelf from which carbonate ooids have been collected is off French Guiana. In this area one of the rare verdine-free localities is in the vicinity of the oolitic facies.

1

213

Figure 6. Magnetic grains of the Senegalese verdine facies. Bars = 0.5 mm. Photo. 1, sample 601-92; photo. 2 to 6, sample 601-95. All photo. except n'2 are of araldite-embedded thin sections; n'2 is of grains leached with dilute acetic acid. 1) faecal pellets (dark-green) constitute 99% of the sample, carbonate remains are white. 2) shells of foraminifers, echinoid spine fragments, and bryozoans with filled chambers. 3 and 4) views in parallel and crossed nicols, respectively, showing green clay (dark-grey in picture) filing voids within a carbonate substrate (white in 4). 5 and 6) parallel- and crossed-nicol views, respectively, of one set of grains. The green clay in foraminiferal chambers is pure, whereas that of the faecal pellet is mixed with carbonate (lighter in 6). Note the absence of oolitic and similar structures.

214

Locations of favourable substrates The precise distribution of the verdine facies in surface sediments off New Caledonia is well documented; samples have been repeatedly dredged or handcollected by divers on our request. The verdine facies is present in the very first decimetre of sediment. This sediment is a compact carbonate sand with molluscan and foraminiferal shells and a small proportion of organic material, but there is no H,S or decomposition odour. The colour is generally grey; it is stable in fresh-water but changes to light or dark ochre in air. Overall, the circulation of sea-water is unrestricted within all known verdine-bearing sediments; but this circulation is slow and the particles are not moved. On the basis of sampling of diverse surface and near-surface sediments associated with the verdine facies, it is well established that the green clays are commonly formed in the uppermost metre of sediment in a fluid whose properties are very similar to those of sea-water. When the sampling is extended to cores (e.g., off Ivory Coast, Gabon, and French Guiana), we conclude that the mineralization takes place within a few metres of the sediment-water interface, mainly in the first metre. Therefore, all green clay minerals discussed in the preceding chapter (B5)can be related to synsedimentary geochemical processes.

Role of substrates The role of a suitable substrate is to provide the necessary microenvironment for geochemical reactions and crystal growth. The fluid within the substrate is slow-moving and will support crystallization, and hinder dissolution, of the verdine clay minerals. The environment beyond the substrate is comparatively poor in ions, mobile, and oxidizing. Moreover, the granular substrates also shelter the green clays from short-lived conditions of very high oxidation that occasionally develop in shallow marine environments. The necessity of a sheltered microenvironment in the substrate is demonstrated by the fact that when carbonate substrates, such as foraminiferal shells, are dissolved, green clays, previously protected within shells, are rapidly oxidized to goethite. The initiation of this process is seen in some miliolid shells in the New Caledonia verdine. The external surfaces are red, because of the oxidation of iron-rich compounds that impregnate the shells, whereas the chambers remain filled with well-preserved (green) phyllite V. The presence of organic matter in substrates, such as faecal pellets or biogenic carbonates, may facilitate the genesis of green clays by allowing the transformation of insoluble ferric iron into soluble ferrous iron (Giresse, 1969). This, however, is not a prerequisite, because chloritized biotites and other substrates free of organic matter are also known to be favourable substrates for verdine formation. In summary, the substrates create a subtle semi-confined microenvironment (Odin, 1975a) which allows the slow entry (and combination by crystal growth) of external cations and protects the growing minerals from the diluting and oxidizing actions of the general marine lagoonal environment.

215 RECENT VERDINE FACIES

versus ANCIENT IRONSTONE FACIES

Comparison of environmental factors It has been suggested that the Recent verdine facies is a modern equivalent of the ironstones (Bhattacharyya, 1983; Van Houten and Purucker, 1984). This needs to be carefully examined. In terms of general environment, both facies are deposited in nearshore, shallow, tropical, marine waters. Oolitic ironstones are never associated with a reefal facies, whereas verdine may locally develop in the vicinity of reefs. The accumulation rate is low in both cases. Variable bottom currents have been reported for the verdine facies; they were probably present but episodic in the case of the ironstone facies. The presence of abundant iron is a prerequisite for the development of the characteristic green clays in both facies. Overall, the two facies appear to have formed in approximately similar general environment. The two mineralogical characters common to the verdine and oolitic ironstone facies: presence of the same two serpentine- and chlorite-type clay mineral families (distinct from glauconitic minerals), and abundance of iron appear to be related to regional environments, and not to microenvironments. In contrast, a difference between microenvironments of genesis may be inferred from the different oxidation states of iron and a systematic difference in the relative proportion of chemical elements in the green clay minerals, i.e., berthierine and chamosite of the ironstone facies, and phyllite V and phyllite C of the verdine facies. Of course, as discussed in Chapter A2, part of the mineralogical differences between the two facies can be explained by the probability that the minerals present today in the oolitic ironstones are different from those initially deposited. The original clays should have been poorly crystallized and could have incorporated femc iron better than ferrous iron. This could then be followed by a strong diagenetic reaction resulting in the reduction of the femc iron to ferrous iron, and a recrystallization and modification of the general chemical composition of the initial synsedimentary green clay minerals. Even in this scenario, however, the problems of the different chemical compositions of chamosite and phyllite V and the lower Mg and Si contents of berthierine and chamosite would remain. Thus, the mineralogical data do not demonstrate a complete sedimentological similarity between ancient oolitic ironstones and Recent verdines. The comparison is difficult, however, because the former facies has been affected by diagenesis, but the latter has not. The habits of marine green clays in the two facies are clearly different. The clays of the ironstone facies are mainly oolitic or diffuse, whereas the oolitic habit is absent and the diffuse habit is rare in the verdine facies. The verdine facies occurs in sandy sediments, whereas the ironstone facies is in muddy sediments. Overall, these are indications that microenvironmental factors supporting the developments of verdine and oolitic ironstone are quite different. Therefore, specific names should be used for designations of the two facies, as specific mineral terms are recommended for designations of the two series of minerals (Chapter B5).

216

Interpretation of ironstones in the context of Recent verdines There are no modem oolitic ironstones, but some of the similarities and differences between the verdine and the ironstone facies offer clues on the depositional environment of the latter. 1) Original oxidation state of iron in the green clays. The study of verdine indicates that synsedimentary iron-rich clays are more likely to be ferric than ferrous. Modem marine sediments, both shallow and deep, include a number of Occurrences of ferric (but not ferrous) iron-rich clay. A number of features in ironstones, including oolitic texture, broken ooids, abundance of goethite, and overall shallow water sites of deposition suggest that iron-rich clays of ironstones were also initially ferric, but were modified to ferrous clays during diagenesis. 2) Significance of berthierine and chamosite. Our present understanding of the causes of the mineralogical diversity in verdine is incomplete, and cannot be used to explain fully the mineralogy of ironstones. The similarity between the mineralogy of verdine ( 7 A phyllite V, 14 A phyllite V, and phyllite C) and that of ironstones (7 A berthierine, 14 A chamosite and swelling-chlorite) has been examined in the context of the depositional significance of berthierine versus chamosite (Chapters A l , A2). Do these two latter oolitic ironstone minerals represent two successive stages of diagnetic mineralogenesis or two geochemical products in different depositional environments? It has been suggested (Chapter B5)that the 14 A phyllite V is the result (or one of the products) of the "aging" of the 7 A phyllite V in close contact with sea-water. A similar process may be invoked for ironstone facies minerals. The presence of different minerals in pre-diagenetic Recent surface sediments of the verdine facies suggests that berthierine and chamosite in ironstones might represent two different initial or very early parageneses. The existence of 150 Ma old or even Palaeozoic 7 8, berthierine also indicates that the 7 A =>14 A reaction does not always take place at an early stage. 3 ) Source of iron. The availability of iron is a limiting factor for the growth of both facies. In this context, modern depositional environments of verdine may serve as analogues for ironstone palaeoenvironments. For example, the presence of a major ironstone unit in a sedimentary sequence may imply the proximity of a large river mouth. A smaller ironstone deposit may imply a smaller river or a volcanic outcrop in the vicinity. Thus, using models based on verdine, aspects of the hydrographic system, climate, and continental erosion can be deduced for oolitic ironstones (Fig. 3 , 4 , 5 above). 4) Habit of the green clays. Substrates supporting the growth of verdine provide a domain that is relatively protected from an oxidizing marine environment. The difference in habit between green clays of verdine (infilling and replacement and those of ironstones (ooids and diffuse mud) indicate that berthierine and chamosite crystallized in comparatively open environments; they were not sheltered from sedimentary fluids. Thus, the fluids in which berthierine and chamosite (or their synsedimentary precursors) formed were considerably less oxidizing than the fluids in which phyllite V and phyllite C develop

217 today. Figure 7 illustrates the relationships: zero Eh planehediment surface for both facies. If we agree with previous authors that the clays formed in the oolitic ironstone facies were initially ferrous, the zero Eh plane must be at or above the sediment surface. If, on the other hand, we consider it more probable (as we have done earlier in this chapter) that the initial clay minerals were femc, the scheme can be modified by slightly lowering the zero Eh plane. In both schemes (femc versus ferrous clays), the general Eh will be more reducing for the ironstone facies than for the verdine facies. A higher proportion (possibly a thin bed) of organic matter at the sediment surface could be the explanation for this lower Eh.

I

p

OOLITIC

VERDINE FACIES

IRONSTONE

FACES

*

,:&::,: .,.. a d .

+ '.

.:.@@

. . . .. . .. . . . 0 ..'.-.-.... . . . .-..-.-c..,- . . . . . . .. . . . . . . , . _. . . . _. . . . .

I@

lnf i//ing s

@

Accretion

+

0Concretion +

Secretion

Figure 7. Locations of the zero Eh plane in relation to the sediment surface. In Recent verdine facies (A), it is located well below the sediment surface, and a sheltered microenvironment is necessary for the growth of femferous green clays in the first decimetre of sediment. In the oolitic ironstone facies, the zero Eh plane has been postulated to be above the sediment surface if the mids form by accretion (B), and at the sediment surface if the ooids grow by an intra-sedimentaryconcretionaryprocess or derive from an accretion-replacementprocess (C).

SYNOPSIS

The newly identified verdine facies is characterized by its habit, mineralogical composition, environment of formation, process of genesis, and differences from related facies. Habit is characterized by the fact that a green pigment is formed within deposited granular substrates which may be 1) filled or 2) replaced carbonate bioclasts, 3) impregnated faecal pellets, and 4) fissured mineral debris. X-ray diffraction studies demonstrate that the iron precipitated in the substrates at the

218 continent-ocean boundary, is utilized by two main kinds of shallow marine clay minerals. Phyllite V and phyllite C show their dominant peaks at about 7 8, and 14 A, respectively, on the X-ray diffraction patterns. A more detailed study, using both X-ray diffraction and HRTEM techniques, indicates that phyllite V may be monomineralic, with dominant 7 A spacings on HRTEM photomicrographs, or plurimineralic, with four species of microcrystals: 7 A, 10 A, 14 A, and 7+14 A; the last three species appear to have recrystallized from the first. The characters of phyllite C are intermediate between those of smectite and chlorite, the similarity with smectite being more pronounced. The HRTEM study indicates that it is essentially monomineralic. Chemically, the difference between phases called phyllite V and phyllite C is that the latter has more silica and less iron. Overall, however, a high proportion of ferric iron is common to all of these clay minerals but the 10 A still problematic mineral. This dominant character and the different relative proportions of Si-Al-Fe and Mg distinguish the green clay minerals of the verdine facies from those of the oolitic ironstone facies; some of the verdine minerals (a "ferric kaolinite" and a "ferric chlorite") are not presently described in the literature. The verdine facies is marine. It is locally present in waters less than 5 m deep, and is probably authigenic in tropical waters as deep as 60 m. In situ formation of verdine in deeper waters is uncertain. Generally, the facies develops in comparatively sheltered areas, but bottom currents have also been reported. The presence of iron appears to be the main geological factor related to the genesis of the verdine facies. The supply of iron originates from river mouths or from volcanic outcrops near the depositional basin. Regional conditions of deposition postulated for ancient oolitic ironstone facies are similar to those observed for the verdine facies. For example, ferriferous ooids in ironstones have been interpreted to have formed on shallow continental or volcanic platforms and within inland seas (Kimberley, 1978). This agrees with known regional environments of verdine. The verdine facies develops at and immediately below the sea-water/ sediment interface during sedimentogenesis. The host sediment is a quartz or carbonate sand and allows sea-water circulation. The Eh of this interstitial sea-water is definitely positive. The green clay minerals of the verdine facies, however, do not form in direct contact with the oxidizing water, but in a sheltered microenvironment within granular substrates where the diluting and oxidizing effects of the surrounding fluids are reduced. The clay minerals form by a crystal growth process, and all ions used in this process are extracted from the intergranular fluid; the mineralogy of the substrates themselves is of no significance in this context. In contrast to regional environments, the respective microenvironments of the verdine and ironstone facies are fundamentally different. Verdine minerals form in a neutral to oxidizing environment, ironstone minerals in a putative reducing environment. Verdine minerals form only within sheltered, granular microenvironments, ironstone minerals in open microenvironments at the sea-watedsediment interface. The overall characters and the mineralogy of clay

219

minerals of the verdine and oolitic ironstone facies are distinct, although both facies are iron-rich and have formed in tropical nearshore environments. The important question why verdine is known only from Holocene sediments whereas sedimentary oolitic ironstones are known only from Pliocene or older sediments remains unanswered at this time. ACKNOWLEDGEMENT

We appreciate the help of D.R. Lowe and P. Aharon (Louisiana State University), who reviewed an early version of this chapter. Araldite mounts were prepared by G. Rouget; photo processing was performed by 0.Fay.

220

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22 1

Part C

THE GLAUCONY FACIES INTRODUCTION TO THE GLAUCONY FACIES The third part of this volume gives a modem view of the glaucony facies; glaucony has been widely studied and discussed in the past thirty years; each year it is still the major subject of several papers and the incidental subject of several dozens of papers. We do not present an extensive new review of this abundant literature, often more mineralogical than geological because several recent reviews are available (McRae, 1972; Odin, 1975a; Buckley et al., 1978; Odin and Matter, 1981; Odin and Morton, 1988); also, we prefer to present new data and new views. Moreover, interest in such a review would not be great since many different and sometimes incompatible interpretations and opinions on the nature, genesis, and environment have been published. This leads to some confusion for readers who simply need to apply knowledge of the facies to their own sedimentological investigations. Instead of a review, we present a series of three case studies (Chapters C1, C2, and C3) selected amongst the numerous occurrences investigated by the editor in the past twenty years. A synthesis chapter based on the data gathered in these case studies and supplemented with significant data from other sources follows (Chapter C4). The case studies were selected from investigations of deposits occumng in modem oceans. These sediments have not been submitted to burial, allowing us to describe a sedimentological phenomenon entirely developed by contact with the marine environment. The marine facies has not been modified by any early diagenetical or later changes which can constitute a source of misinterpretation of sedimentological and mineralogical observations; such changes are often neglected in the literature dealing with older glaucony-bearing formations. The first case study concerns a Pleistocene to Holocene deposit of the eastern equatorial Atlantic continental shelf characterized by faecal pellets as the marine substrate of glauconitization. This west African deposit is suitable for the investigation of the beginning of the verdissement process. The second case study concerns a Plio-Quaternary sediment of mid latitude in the northern hemisphere. The continental shelf off northwestern Spain is characterized by a variety of substrates, most of them carbonate bioclasts. The earliest phases are present but the deposit also shows the latest phases of the glauconitization process still in contact with sea-water. The continuous morphological evolution is illustrated for six substrates using macrophotography, microphotography, and scanning electron microscopy. The third case study deals with a Quaternary deposit at high latitude in the Indian Ocean. These sediments located at the top of an oceanic high were never studied for authigenic clay minerals. Glauconitization occurred mainly on rock

222 fragments and mineral debris; however, green bioclasts are not rare. Both mineral debris and bioclasts are characterized by their siliceous nature. Morphological, mineralogical, genetical, and environmental data are given for these three examples. These data are more complete than would have been the case for older formations; however, data from older formations have sometimes been considered in the above three examples to show that the phenomena observed were not specific of young sediments. These data are discussed in the last chapter along with supplementary observations made on older sedimentary series by the authors or taken from the literature. These show that glauconitization (sensu largo) seems to include a previously non-recognized phase of evolution: recrystallization during early burial at a time when the environment is formed of interstitial waters. Our knowledge of the geology of glaucony has made progress during these last thirty years. From a morphological point of view, our present understanding, mostly based on our own observations, allows elaboration of a genetical classification instead of the early strictly descriptive "lists" of habits. This classification is partly based on the nature of the substrate sheltering the formation of the marine green clay genesis and partly on the degree of evolution of these substrates. From the mineralogical point of view, it is important to distinguish the two fundamental components of green grains: the substrate, mineralogically very diverse and generally disaggregated, and the authigenic minerals. The latter also are diverse and need specific mineral names. The duality of composition of the grains has been especially emphasized using the results of isotopic studies which show that the contribution of the substrate elements and isotopes is much greater than could have been presumed from routine diffractometric or chemical analyses; this is of importance for geochronological application and has often been neglected in the past. Our most original contribution deals with the genesis of glauconitic minerals. At the beginning of our study (late 1960'), glauconitic minerals were believed to be formed by a progressive transformation of an inherited illitic degraded layer. This theory, supported by most authors at that time, was generally accepted for many clay mineral processes of formation. Observations gathered in the four chapters below prove that this theory is incorrect. Glauconitic minerals result from a neoformation, they form by crystal growth and recrystallization processes. This was one of the main ideas presented more than ten years ago (Odin, 1975a). This idea now has begun to be accepted by clay specialists. Finally, from the environmental point of view, the editor believes that the determining factor is the presence of a specific microenvironment at the boundary between the open sea-water and the buried sediment. This semi-confined microenvironment is linked to the physical nature of the substrate of verdissement. The role and influence of this substrate is one of the key factors for all aspects of glaucony. The contents of this section on the glaucony facies follow:

223

CONTENTS of PART C *Introduction to the glaucony facies *Contents of part C

22 1 223

*Chapter C1 Glaucony from the Gulf of Guinea by G.S. Odin 1. Introduction 2. Geological setting 2.1. Hydrology 2.2. General data on glaucony 3. Habit and nanosmcture of glaucony 3.1. Morphological data 3.2. Nanostructure of the faecal pellets from the Congolese shelf 3.3. Morphological evolution of the faecal pellets 4. Mineralogy of the green grains 4.1. X-ray diffraction study 4.2. Chemical study 4.3. Stable isotope study 4.4. Radioactive isotope study 5. Geology of the glaucony from the Gulf of Guinea 5.1. Age of the glauconitization process 5.2. The glauconitization process 5.3. Environment of glauconitization 6. Summary and conclusions

225 225 225 227 229 229 232 232 233 233 236 237 238 240 240 242 245 246

*Chapter C2 Glaucony from the margin off northwestern Spain by G.S. Odin and M. Lamboy 1. Introduction 2. Distribution of glaucony 3. Morphological features 3.1. The verdissement of echinoderm fragments 3.2. The verdissement of bored shell fragments 3.3. The verdissement within the foraminiferal tests 3.4. The verdissement of detrital mica flakes 3.5. The verdissement of quartz grains 3.6. The verdissement of non-bored shell fragments 3.7. Other morphological features 4. Mineralogical study 4.1. X-ray diffraction sudy 4.2. Chemical study 4.3. Isotopic study 5. Discussion 5.1. Morphological features and glauconitization process 5.2. Mineralogical features and glauconitization process 5.3. Age of the glauconitization process

249 249 252 252 254 256 257 260 262 264 266 266 268 269 27 1 27 1 27 1 272

224 5.4. Mechanism of glauconitization 5.5. Environment for glauconitization 6. Summary

272 273 274

*Chapter C3 Glaucony from the Kerguelen Plateau by G.S. Odin and F. Frohlich 1. Presentation 1.1. Glauconies from high latitude deposits 1.2. The Kerguelen Plateau 2. Sedimentology 2.1. The sedimentary cover 2.2. Glaucony 3. Nature and origin of the green grains 3.1. Detailed data on the northeastern shelf sediments 3.2. Substrates of verdissement 3.3. Electron microscopy of evolved grains 3.4. Mineralogy of the green grains 4. Discussion 4.1. History of the glaucony from the Kerguelen Plateau 4.2. Factors of the glauconitization process 5. Summary

277 277 27 8 279 279 280 282 282 283 289 29 1 29 1 29 1 292 294

*Chapter C4 Geological significance of the glaucony facies by G.S. Odin and P.D. Fullagar 1. Introduction 2. Habits of glaucony 2.1. Classification of habits 2.2. The granular habits 2.3. The film habits 2.4. Discussion of the habits of glaucony 3. Mineralogy of glaucony 3.1. Substrate components 3.2. Authigenic marine clays 3.3. Post-genesis components of green grains 4. Genesis of glaucony: the verdissement process 4.1. The layer lattice theory 4.2. The mechanism of verdissement 5. Environment of glauconitization 5.1. Microenvironment 5.2. General environment 6. Summary

295 296 296 297 300 30 1 305 305 309 3 16 318 318 319 323 323 324 33 1

225

Chapter C1

GLAUCONY FROM THE GULF OF GUINEA by G.S. Odin INTRODUCTION

The present chapter considers the continental shelf located between Ivory Coast and the Congo River mouth (between 6" N and 6" S). This area called the Gulf of Guinea, is part of the East Atlantic margin where green grains of the glaucony facies were described in surface sediments, or in cores penetrating the Quaternary sediments from off Morocco to the extreme south of Africa. More to the north of the Gulf of Guinea, (Fig. 1) the Moroccan glauconies were observed by Bell and Goodell (1967), Mathieu (1968), and Emelyanov (1970). Moroccan sediments provided to the editor by R. Mathieu mainly show glaucony filling foraminifer chambers in sediments collected from the outer margin of the shelf, in agreement with Tooms et al., (1970). Southwards, Correns (1939) quotes glaucony from between the Cap Blanc and the Cap Vert, off Mauritania. The detailed mineralogy of these green grains is unknown. Off Guinea, glaucony has been quoted by Mc Master et al. (1970). Samples from short cores in Quaternary sediments from the outer portion of the shelf and the slope were provided to us by J.C. Faugkres (Casamance region) and J.P. Masse (off Guinea). Light-green to dark-green glaucony is abundant in these sediments (unpublished results; Chapter C4: Fig. 14, p. 325) and sometimes mixed with phosphate. Further to the south of the Gulf of Guinea, the presence of glaucony has been reported (from north to south respectively) from off the Cumene River, Walvis Bay, Orange River and southern coast by Collet (1908), Bezrukov and Senin (1970), Emelyanov (1970) or Simpson (1970), and more recently by Birch et al., (1976) or Odin (1985b). The continental shelf in the Gulf of Guinea shows a nearly continuous deposit of green sands which have common characters everywhere including morphology, depth, location on the shelf, abundance, age, local and general environment. In this context, the region between Ogooue and Congo river ill be considered more in detail as a local example of the case study. mouths w GEOLOGICAL SETTING

Hydrology This region is characterized by a peneplaned continent covered with tropical forest. Two very large rivers are present: Niger and Congo; the Niger River has a large delta indicating a major detrital output. Many other coastal rivers are present, from west to east, they are the Cavally, Sassandra, Bandama, and Co-

226 *fMOROCCO

I000krn SSO 87

Figure 1. Glaucony in Quaternary sediments of the East Atlantic border. Starred deposits were studied by the present author. (Revised after Odin, 1973)

moe rivers in Ivory Coast; the Volta River in Ghana, the Sanaga River in Cameroun, and the Ogooue River in Gabon (Fig. 2). The Congo River brings 45,000 m3/s of water to the sea and follows the biggest river of the world (the Amazon River) from this point of view. The iron content of the Congo River water is comparatively high (255 m a ) and higher than that of the Amazon (34 m a ) according to Figukres et al. (1978). The major portion of the fluvial iron usually flocculates in sea-water as shown by Sholkovitz et al. (1978). However, the water output is so rapid at the Congo River mouth that fluvial water does not mix immediately with sea-water as it does in the estuary of the Kourou River (Chapter B5). As a result, the light fluvial water mass moves far into the sea at the surface of the salted water mass; by this process, iron is transported into the ocean far from the river mouth. The continental shelf is a flat platform, 40 to 80 km wide from the coast to the break located at about -120 m depth where the slope begins. The history of most recent sea-level changes is well documented on this stable margin and is illustrated above (Fig. 12 in Chapter B2, p. 101; and Fig. 8 below, p. 241). The most important feature is that the sea-level was located around 110 m below the present one 18,OOO years ago; that period was followed by a rapid transgression reaching the present sea-level 6,000 to 7,000 years ago. The Quaternary sedimentary cover is thin especially for the most recent Holocene

227

Known I V V Glaucony

Figure 2. Glaucony from the continental shelf of the Gulf of Guinea.

period; facing the main river mouths submarine deltas exist.

General data on glaucony The presence of glaucony from the continental shelf of the Gulf of Guinea has long been known (Collet, 1908); but the mineralogy was mainly considered 60 to 70 years later. Three areas are more precisely known i.e., off Ivory Coast (Martin, 1970; 1973); off the Niger delta (Porrenga, 1967a); and off Congo and Gabon. The precise mineralogy of the latter Congolese green grains was examined initially by Odin and Giresse (1972). Glauconyfrom the shelf offIvory Coast Of the three areas quoted above, that off Ivory Coast is the least interesting for glaucony genesis because few samples were collected from deeper than -100 m and because a number of sediments were submitted to an oxidation process after the genesis of the green material. Oxidation is visible in shallow sediments (Chapter B4) and widespread between -100 m and -120 m where all grains are rich in goethite. The green pigment is mainly present in the form of faecal pellets, green in colour. These pellets are restricted to the f i t metre of

228 the sedimentary cover, deeper in the sediment, faecal pellets exist but are soft and grey coloured. The mineralogical study allowed the glaucony facies to be distinguished from the shallow verdine facies. X-ray diffraction patterns of powdered glauconitic pellets show the presence of poorly crystallized clay minerals; the 001 diffraction peak is located between 14 A and 12.6 A. In addition, a wellshaped 7 8, diffraction peak is always present and indicates the presence of kaolinite (Martin, 1970). From the chemical point of view, pellets collected from depth between -20 m and -70 m contain 0.4 to 0.8 wt% K 0; between -80 m and -250 m depth the content increases from 1.2 to 2 . 3 % . h e iron contents of many green pellets was measured between 15% and 32% (Fez03) plus 1 to 6% (FeO). The highest ferric iron contents correspond to a beginning of oxidation process. However, the iron content of the magnetically separated pellets are always high and distinctly superior to those measured in the clayey matrix: 5.5 to 7.0 total wt% Fe203 (Martin, 1973). This indicates that the pellets have selectively undergone a mineralogical change mainly characterized by an important trapping of iron (concomitant with a loss of aluminium).

Glauconyfrom the Niger delta Green grains related to the glauconitic facies have been identified by Porrenga (1967a) from the continental shelf between Lagos and Fernando Poo (Fig. 2). The facies is generally better preserved than off Ivory Coast at depths between -90 m and -400 m. Green faecal pellets are more abundant between -125 m and -250 m depth, at the top of the slope, than elsewhere on the shelf. Other green pellets are present shallower than -80 m; they were related to the verdine facies (Chapter B4).Deeper than -350 m the pellets are soft and did not undergo the mineralization process. According to Porrenga (1967a), X-ray diffraction study shows that the light-green clay is mainly of smectitic nature with 70% expandable layers and a low potassium content (2.7 wt% K,O). Porrenga emphasizes the striking difference in chemical composition between the green pellets and the clay matrix, the former have much higher iron, magnesium, and potassium and lower aluminium content compared to the latter. Porrenga concludes that these green pellets do not consist of merely agglutinated detrital clay-size material but do represent a newly-formed mineral'... in an early stage of development compared to the glaucony from older and buried sediments. ...I

Glauconyfrom the Congolese shelf The area located between the Ogooue and the Congo river mouths is better known than the two previously quoted and shows a wider variety of mineralogy and habits. Green grains are probably present everywhere on this continental shelf; they are essentially abundant at the deepest extremity of the shelf and the top of the slope (Giresse and Odin, 1973). Their known distribution is given above (Chapter B4, p. 149). Figure 3 gives a synthetic view for the sedimentary nature of the deposits along a profile, it is generally

229

I

I$

100

T

%om

T

100

cracked and broken green large faecal pellets shelly dork-green carbonaie mud greensmid

T

T

120

shelly sand

50

smooth faecal pellets ochreous to iight-green

1

0

-

DEPTH

GLAUCON'

sand t o mud

Figure 3. Magnetic grain contents in a profile cutting the Congolese continental shelf. The grains shallower than about -110 m are younger than -18,000 years; deeper grains are older. This distribution is representative of the margin in the Gulf of Guinea; variations may occur locally (Giresse, 1976a) depending on burial or winnowing processes.

representative of the sediments from the shelf of the whole Gulf of Guinea. On the continental shelf itself, and from the coast seaward, glauconitic sands and muds are present at depths of about -50 m; the green grain contents are most important at depth between -100 m and -250 m forming a band about 10 km wide and parallel to the coast. Deeper, magnetic grains have been identified down to -1000 m by Bongo Passi (1984). Glaucony is not always in surface sediments because the sedimentation is active in these deep sediments; this allows a biostratigraphic relative dating of the time of deposition to be done for these probably perigenic green pellets coming from the top of the slope. HABIT AND NANOSTRUCTURE OF GLAUCONY

Morphological data The green grains are mainly formed from faecal pellets in the three regions discussed above. The present study will consider the phenomena related to these faecal pellets. These pellets were originally ma& of the clay-size fraction of the sediment. Pellets probably similar to the original pellets may still be observed either at great depths or in many places in the shallow deltaic sediments as already emphasized by Porrenga (1967a) in the Nigerian area. Facing the Congo and Gabon coasts, there is a band of glauconitic material mainly formed of carbonate bioclasts in the zone of relict shelly facies at about -1 10 m i.e., where the palaeo-coast was located 18,000 years ago. It seems that the verdissement does not preferentially select some grains but indifferently occurs within the substrates fortuitously present in the deposit. The distribution

230

Figure 4. Faecal pellets from the Congolese shelf. 1) Sample 313 collected from -95 m 2) Sample 319 collected from -300 m depth; coarse size-fraction for both samples; the deeper grains are darker and more cracked that the shallower well recognizable faecal pellets. The bars on the pictures are 1 mm in size. (Modified from Odin and Dodson, 1982)

of the different habits as a function of depth (Fig. 3) indicates that there is no general transportation perpendicular to the coast. Therefore, each kind of green grain characterizes the in situ mineralogenesis along a restricted band parallel to the coast on the continental shelf. This is also supported by the different grainsizes of the faecal pellets as a function of depth; shallow green grains are smooth and 250 pm to 400 pm in diameter; deep grains are cracked, dark-green in colour and may reach .1 mm in diameter. In the latter case, smaller grains occur in the same sediment and appear to originate from larger grains which have been broken as a result of deep cracking (Fig. 4). The deep dark-green grains appear therefore as faecal pellets more evolved than the shallow smooth light-green pellets.

-+++

Figure 5. Scanning electron microscopy of light-green and dark-green faecal pellets from the Congolese shelf. Light-green faecal pellets collected from -90m depth are illustrated in photomicrographs 1 to 6; dark-green faecal pellets collected from -250 m depth are illustrated in photomicrographs 7 to 12. (Scale bars are 100pm for photomicrographs 1,2,7,8 and 10 pm for others)

232

Nanostructure of the faecal pellets from the Congolese shelf Scanning electron microscopy has been used for characterization of the faecal pellets obtained from the Congolese shelf. Two sorts of grains have been selected; ochreous to light-green ellipsoidal shallow pellets are illustrated in Figure 5: photomicrograph 1 to 6; dark-green deep cracked pellets are illustrated in photomicrographs 7 to 12. The light-green faecal pellets are smooth (Fig. 5, Photo.1). When broken, they show many pores in an heterogeneous interior (Fig. 5, Photo.2). The pores appear to result from the dissolution of minute bioclasts such as foraminifers (Fig. 5, Photo.4). Some structures are possibly still made of carbonate (Fig. 5, Photo.5); but others show that the initial carbonate has been replaced by green clay (Fig. 5 , Photo.6). The green clay is mainly formed of ill-defined globules frequently associated in caterpillar (Fig. 5, Photo.3 and 5). This structure is typical for the early evolution of the green gains. The dark-green pellets show deep cracks modifying the initial ellipsoidal form of the particle (Fig. 5, Photo.7). When broken, these pellets show an homogeneous interior without large pores (Fig. 5, Photo.8). The green clay cristallites are better shaped and larger than in shallow pellets (Fig. 5, Photo.9). The elements of the nanostructure are sometimes associated to form boxwork and rosette nanostructures (Fig. 5, Photo. 10). When present, caterpillars may reach 2 or 3 pm in lengh (Fig. 5, Photo.11). These caterpillars sometimes represent the external border of small flakes, 2 to 3 pm in diameter (Fig. 5, Photo. 12).

Morphological evolution of the faecal pellets The striking feature in the general distribution of the faecal pellets is the morphological evolution as a function of depth. Shallow sediments between -60m and -90 m depth give generally ochreous faecal pellets which become light-green in the deeper area. In contrast, the pellets collected from between -150 m and -300 m depth are darker. Together with the colour, the form and size are depth dependent. Smooth ellipsoidal small grains evolve to deeply cracked, large grains. The pellets increase in size as a function of depth while they progressively lose their general form due to cracks. In a single sample, the same sequence of morphologies can be found in comparatively deep green sands. This observation, and other similar ones in other circumstances, rejects the idea that cracks could be due to dehydration figures correlative with volume decrease during the evolution. The internal structure evolves from heterogeneous, pore-rich pellets toward homogeneous dense pellets in the oldest and deepest sediments. The pores present in the early stage of evolution are demonstrably due to carbonate remnants dissolved very soon after the pellet formation. Therefore, from a series of morphological criteria we may distinguish less evolved pellets from more evolved pellets. This evolution appears related to the duration of the presence of sea-water above the sediment at its present site.

233 Local conditions of deposition also influence the mineralization process i.e., glauconitization and its distribution on the shelf. For example, local bottom currents may provoke transportation and mixtures; local detrital input may bury the pellets; as a result, the general distribution described above shows many local disorders. But the important observation is the general trend as a function of depth, and its presumed cause which is the time available for glauconitization to proceed within the faecal pellets initially formed in a shallow environment. The progressive evolution may easily be observed just by collecting samples at various depths on the Congolese margin, a unique opportunity. MINERALOGY OF THE GREEN GRAINS

X-ray diffraction study Evolution of the mineralogical composition with depth Taking into account the results of the morphological study, X-ray diffraction study was undertaken as a function of depth of collection of the glauconybearing sediments. For each sediment, faecal pellets have been separated from different size fractions using a magnetic separator, and analysed after ultrasonic treatment. Randomly oriented powder diagrams have been preferred in order to avoid the size selection of particles and resulting mineralogical bias which occurs when oriented preparations are made from heterogeneous material. Figure 6 gives a few X-ray diffraction patterns obtained using cobalt radiation. Glaucony samples (G) are designated with their depth of occurrence. The evolution of the mineralogical composition is obvious. The ochreous-green faecal pellets collected from shallow depth (-85 m) show a hump at 14 A. The dark-green cracked faecal pellets collected from the slope (-300 m) show a main peak at about 11 A. Intermediate phases are observed at -95 m and -120 m. The comparatively broad peaks obtained are typical for unburied authigenic marine clays analysed using randomly oriented powder mounts. In addition to this dominant peak between 10 8, and 14 A, two other peaks are present at about 4.5 A and 3.3 A. These two peaks correspond to the 020 and 003 diffractions of a 2:1-type clay mineral. The height of these peaks is increased in deeper pellets compared to shallower ones; but the presence of quartz does not favour this observation. Finally, two other peaks at 7 A and 3.5 A are interpreted as traces of the initial composition of the faecal pellets i.e., kaolinite which is the clay dominantly present in the less than 50 pm size-fraction of the sediments of the shelf. An X-ray diffraction pattern of this fine fraction is drawn in Figure 6 and called "clay". This fine fraction probably formed the material ingested by mud-eaters and it is reasonable to find traces of it in the less evolved faecal pellets (G312). The proportion of kaolinite decreases from the shallow to the deeper faecal pellets and this traces again the general mineralogical evolution of the green grains. On the one hand, an authigenic phase at 14 A grows in the originally kaolinitic faecal pellets, this was similarly described from the Niger delta; off

234

Figure 6. X-ray diffraction patterns of green faecal pellets and clay-size fraction from the Congolese continental shelf.

Congo, the 14 A peak progressively evolves further toward 10 to 11 A. On the other hand, the detrital phase (mostly kaolinite but also quartz and calcite as shown in Fig. 6) is progressively eliminated from the grains now green coloured. This double mineralogical evolution is well illustrated in the depth dependent progressively evolved grains of the Congolese margin.

235

Mineralogical nature of the green marine clay GLAUCONY 312

7

I

10

14

17A

I

E Glycol

Figure 7. X-ray diffraction patterns of the green clay from the shallow faecal pellets.

The authigenic phase of the faecal pellets is characterized by its 001 diffraction peak between 14 8, and 10 A. This clay has been submitted to usual treatments such as heating or ethylene glycol vapour. Figure 7 gives the X-ray diffraction patterns obtained from randomly oriented powders of a glaucony collected from -85 m depth. The untreated peak at 14 8, shifts to 17 A after ethylene glycol treatment indicating a swelling behaviour (absence of cations in the interlayers of the 2:l clay). Heat treatment at 350°C for two hours easily collapses the clay layers to 10 A (this indicates the absence of brucite in the interlayer). Therefore, the authigenic phase of the shallow green pellets is a smectite. One may also note that the small peak at 7 A is not destroyed at 350°C and is destroyed at 490°C indicating that this peak acmally represents a kaolinite similar to that present in the fine size-fraction of the sediment (Fig. 7). A similar study has been undertaken for p e n grain fractions from the Congolese and Ivory Coast continental shelves. The results are similar everywhere concerning the heat treatment which collapses the layers to 10 8, at comparatively low temperature. However, the ethylene glycol vapour does not

236 lead to a swelling of all samples. More precisely, as soon as the untreated green clay mineral shows a main peak at 13 A or less the swelling behaviour becomes more difficult to observe in the absence of pre-treatment eliminating the interlayer potassium.

Chemical study The chemical composition of green grains already studied for morphology and X-ray diffraction is shown in Table 1. Results have been obtained by three different analysts using two different main techniques (X-ray fluorescence and wet chemical technique) in order to have reliable results. Three to four grammes of magnetically separated and ultrasonically cleaned (using N/10 dilute acid) faecal pellets have been used. Special attention was paid in trying to avoid the analysis of oxidized grains which are common. Table 1. Chemical composition of the green clays of variously evolved faecal pellets from the Congolese continental shelf; comparison with co-occurring demtal fine fraction. Sample

Depth

Si02 A1203 Fez03 FeO

T i 0 2 MnO C a O hlgO NaZO K 2 0 P 2 O 5 H2O- H2O'

(1) G.312.A

(-85m)

46.7

9.75 (20.85 t o t a l )

0.7 0 . 0 3

<1.4

3.05

0.2

2.3

0.15

5.0

10.0

( 1 ) G.313.A

(-95m)

45.7

9.8

21.6

1.1

0.8

0

<1.4

2.95

0.5

3.0

0.25

4.8

8.6

(1) G.490.hA

(-120m)46.9

6.9

21.2

0.85

0.5

1.4 3.6

0.3

3.41

-

6.55

8.35

( 2 ) G.49O.cA

(-120m)46.3

5.6

23.5

1.2

0.3

1.4

4.3

0.2

4.2

-

6.5

4.2

( 3 ) G.49O.cA

(-120m)46.9

6.5

24.1

1.1

-

-

1.4

4.0

0.2

3.4

-

( 1 ) G.318.A

(-225m147.6

6.6

22.9

1.1

0.8

0

<1.4

2.8

0.4

5.1

0.2

4.65

( 1 ) (3.319.0

(-300m)47.5

6.4

(23.8 total)

0.4

0

<1.4

2.8

0.15 6 . 6

0.15

3.4

6.9

( 1 ) (3.319.0

(-300m)45.9

6.4

(23.8 t o t a l )

0.4

0

<1.4

3.5

0.1

6.6

0.05

2.3

9.1

4 0 . 8 21.8

(11.7 t o t a l )

1.1

-

2.6

1.7

0.2

1.52

3.6

14.1

( 2 ) A.448

-

(11.7 t o t a l ) 6.55

1) G . Richebois, Suatigraphie, Paris (wet chemistry); 2) M. Lenoble, PCuographie, Paris (X-ray fluorescence);3) E. Lebrun, GCologie, Ecole Normale, Paris (wet chemistry).

The analysis of eight samples collected from between -85 m and -300 m depth points out a remarkable homogeneity for most of the cations analysed including silica: 46% to 4.89'0, iron oxides: 21% to 24% or MgO: 3% to 4%. Alumina contents are comparatively less homogeneous although always low; but the two higher values are presumably due to detrital contamination within the green grains. Potassium shows wide variation, its content being a factor of three greater in the deepest samples compared to the shallowest ones. This is consistent with the X-ray diffraction results which show thinner layers for deeper green grains. The potassium is the cation which binds together the negatively charged layers of the green clay mineral. The correlation between increasing potassium content, thinning of the layers as seen in X-ray diffraction patterns, and depth of occurrence, is good.

237 The data discussed here usually concern the most evolved green grains from each sediment because, in order to obtain the best purity and homogeneity, the most paramagnetic, darkest, and densest grains are always selected for analysis. A given sediment may be composed of diverse green grain fractions with analytically distinct potassium contents; for example, glaucony G363 separated from a sediment collected from -180 m depth was split in two magnetic fractions where the potassium and iron oxide (F903) contents were measured at 6.3% and 20.1% on the one hand and 5.0% and 23.2% respectively on the other.

The comparison between the chemical composition of the magnetic green faecal pellets and that of the fine fraction of the Congolese sediments confirms the observations published from other areas of the Gulf of Guinea by Porrenga (1967a) and Martin (1973). These authors have emphasized the striking chemical difference between the green clay in the pellets and the clay in the matrix, the former contains much more iron, more silica, potassium and magnesium and much less alumina than the latter. Table 1 gives the composition of a shallow fine fraction (A448). Others have been analysed with similarly low iron oxide contents between 8% and 12%. In conclusion, the chemical study of the major elements c o n f i i s that there is an evolution of the faecal pellets initially made of dominant kaolinite. As soon as they are green and paramagnetic, these faecal pellets are made of a smectite with a chemical composition very different from the initial kaolinitic clay. Later on, the evolution mainly consists of an increase in potassium content which both traces the closure of the authigenic clay minerals evolving from a smectite toward a mica-like structure (but never reaching this end-member in the Gulf of Guinea) and the elimination of the detrital components out of the faecal pellet. In order to follow that evolution in more detail, more sophisticated analytical techniques were used as described below.

Stable isotope study The isotopic composition of oxygen may help in following the general evolution of the faecal pellets because the isotopic composition of oxygen is not similar in detrital clay minerals and in marine authigenic clay. Detrital clays may show comparatively low 6l80values (about 15%0).In contrast to that, a green clay formed in equilibrium with the present sea-water oxygen isotopic composition (6l80= 0) tends to show a higher 6 l 8 0 value of 23-24%, (see Table 13, p. 379 in Part D). Table 2 below gives the oxygen isotopic composition of four samples from the Congolese shelf already analysed above for X-ray diffraction and major element composition. The approximate age of deposition of the faecal pellets is also indicated. The 6l80values indicate that the potassium poor light-green pellets still contain a large proportion of detrital material, much higher than that shown on X-ray diffraction patterns (Fig. 6) because the measured value of 18.2%0 is still very far from the equilibrium value for an authigenic material (about 23%0). This means that the oxygen isotopes initially present in tetrahedra or octahedra of the detrital minerals have not been exchanged with those of sea-

238 water. Taking into account the X-ray diffraction patterns (kaolinite structure mostly destroyed) this may be interpreted in three ways: 1) the tetrahedra (and octahedra) of the detrital material, with their oxygen, have been dissociated but not eliminated from the pellets and do not diffract X-rays although still present; 2) the tetrahedra and octahedra of the detrital material, or group of them, have been re-used as elementary pieces for growth of the neofonned smectite; and 3) tetrahedra and octahedra of the detrital material have been completely dissociated as ions but immediately re-used for growth of authigenic minerals in an environment too confined to allow exchange with the general sea-water environment. It is not yet possible to favour one or another hypothesis. But the important conclusion is that although X-ray diffraction study seems to indicate that the major proportion of detrital material is eliminated in the green grains analysed here, this is not correct. The next section will show that hypothesis 3 above is unlikely and that still combined demtal material remains present but has lost its crystallographical organization. Table 2. Oxygen isotopic composition of some green grains from the Gulf of Guinea.

(unpublished analytical results by E. Keppens and J. ONeil on our samples) Sample

Depth

Deposition age

K20

6 180

G 313 A

-95 m

16, 000 years

3. 0

18.23

G 490h A

-120 m

20,000 years

3.4

20.80

G 490e A

-120 m

20,000 years

4.2

21.48

G 490a A

-120 m

20,000 years

4.25

21.10

Radioactive isotope study The K and Ar isotopes may be used here to trace the evolution (elimination) of the detrital material from which faecal pellets are initially formed. Table 3 gives the potassium and radiogenic argon contents of five fine fractions (clays) and five variously evolved samples of green faecal pellets from a mostly smectitic (G313) to a dominantly (but not entirely) micaceous green mineral (G319). The fine fraction composition is presumed to represent the initial K-Ar isotopic composition of the faecal pellets because that fraction is similar to the material ingested by mud-eaters. In this fine fraction, the radiogenic argon content is high (about 25 nanolitre per gramme: nl/g) and the potassium content is low. The resulting potassium-argon apparent age indicates that the fine fraction contains demtal components inherited from the old Congolese shield. The less evolved green faecal pellets still contain more than halfof the initial radiogenic argon content shown by the fine fraction. This means that the same proportion of detrital material is still present (because evolving faecal pellets are not able to selectively trap radiogenic argon from the exterior). Moreover, the

239 argon atoms have no architectural role in any crystal structure and must naturally be eliminated as soon as a crystal structure is altered; as a result, the above estimate of more than 50% is only a minimal one if we assume that some argon from the substrate has diffused out of the incompletely dissociated inherited constituants. When the potassium content reaches its maximum value on the Congolese continental shelf (sample G319: K,O = 6.6%), the faecal pellets still contain about 10% of the initial radiogenic argon content. This is surprising since no trace of inheritance has been identified using the X-ray diffraction analysis nor the scanning electron microscope. The nature of the mineral structure which is able to keep this argon in spite of the deep alteration is problematic. An approach to this problem has been undertaken using leaching experiments (Odin and Rex, 1982). This study showed the argon trapped in sample G490 to be in a mineral structure different from the authigenic phase. In effect, Odin and Rex showed that leaching with concentrated NaCl hot solution leads to an easy extraction of radiogenic argon from the pure authigenic glauconitic minerals but not out of the Congolese green pellets. The presence of inherited mineral fragments (unaltered by acid solution) is demonstrated by these experiments. Table 3. Potassium-argon data for green grains and clay-size fractions (less than 50 pm) from the Congolese shelf. (Editor's analyses in Berne, Hannover, Strasbourg, and Leeds; after Odin and Dodson, 1982)

Sample

Potassium (wt % K 2 0 )

Radiogenic argon (nl/g)

Apparent age Ma (+ 2 a

5 clays

1 . 1 to 1 . 7

20 to 30

4 7 5 to 5 2 0 1 5 0 2 10

G . 313

3.0

15.0

G . 490 h

3.4

8.9

80

G . 490 e+a ( 2 samples)

4.2

7.1

50 2 3

G . 319

6.6

2.6

12 2 1

7

A similar Rb-Sr isotopic study was undertaken in 1984 in collaboration with E. Keppens. The preliminary analytical results indicate a similar tendency with traces of detrital isotopes in all analysed samples from the Congolese shelf (unpublished results supplemented by P.D. Fullagar in Chapter C4, p. 307). In summary, a predominant proportion of the nascent to slightly-evolved Congolese pellets, showing an authigenic green clay mostly of smectitic nature, is made of detrital components; these components are still partly present in the most evolved green pellets of the same shelf. These conclusions agree with the ones previously proposed following the study of the stable isotopes.

240 GEOLOGY OF THE GLAUCONY FROM THE GULF OF GUINEA

Age of the glauconitization process The most important point in estimating the age of glauconitization of green grains presently dredged from the sea-floor is to recognize whether or not they are reworked from older sediments or in situ. Reworking is very frequent because glaucony is a very stable component in the sea (with which it is thermodynamically in equilibrium). For example, glauconitic chalk or limestones outcropping at the sea-bottom are easily altered by dissolution of carbonate and then glaucony is concentrated ( O h , 1985b). This is not the case in the Gulf of Guinea where glauconitic formations are absent immediately below the Recent sediments except for Ivory Coast where Palaeogene glauconitic clays occur in coastal cliffs; but the general distribution of glaucony characteristics (such as the presence, proportion in the sediment habit, and detailed mineralogy) indicate that glauconitization occurred in situ everywhere. The precise time at which the in situ glauconitization process occurred in the Gulf of Guinea may be discussed considering the sea-level changes curve during the last 100 ka. From this mean curve (Fig. 8) and the postulation that glaucony forms at the sea-bottom and would be destroyed during emersion, two possibilities occur for dating the glauconitization process. At depths shallower than -1 10 m, glaucony must be younger than 18,000 years; deeper, it could be older. Concerning glaucony presently shallower than -110 m depth, the sea is present for a larger time at -1 10 m compared to -60 m; therefore, the beginning of the glauconitization process can be older at -110 m depth compared to -60 m. This is likely to be correct since deeper glaucony is more evolved than the shallower one. The question arises of when after the immersion, does the glauconitization actually begin. Glaucony is only present and abundant deeper than -80 m off Ivory Coast (Martin, 1973) and Nigeria (Porrenga, 1967a). According to Figure 3 above, there are glauconitic green grains present off Congo only at depth greater than -60 m. Therefore, it is likely that glaucony forms only when a minimum of 60 m of sea-water is present above the sediment. Consequently, the faecal pellets now collected from -85 m began their glauconitization process when the sea-level was at about -25 m i.e., about 8,000 to 9,000 years ago (Fig. 8). This is the maximum duration needed for obtaining a very early (smectitic) stage of evolution of the faecal pellets. Deeper than -1 10 m, the question is more complex. Sea-water was always present there during the last 100 ka. However, because glauconitization seems to need about 60 m of sea-water above the sediment to occur, the possibility for glauconitization to occur in sediments presently lying at -110 m will exist by the time when the sea-level was located at 110-60 = 50 m below the present one. Figure 8 allows us to conclude that glauconitization was not possible at - 110 m in the two periods -1O,OOO/-30,000 years and -45,OO0/-75,000 years. We also know that glaucony was already present on the shelf at the time of minimum sea-level, about 18,000 years ago. This is because a nearly continuous band of

24 1

Figure 8. Age estimate of the glauconitization process as a function of depth of deposition in the Gulf of Guinea. The sea-level curve accepted here (see p. 101) was established after observations in sediments off Senegal (Faure and Elouard, 1967; H. Faure, personal communication); off Ivory Coast (Martin, 1973, p. 292); and off Congo (Delibrias et al., 1973).

oxidized glaucony is known at about -1 10 m depth in the different areas of the Gulf of Guinea: this previously developed glaucony was altered during the regression time. Therefore, at -1 10 m and deeper,'the glaucony present in the sediment began its evolution either between about 30,000 years ago and about 45,000 years ago (Fig. 8) or before 75,000 years B.P. The depth at which glauconitization was possible during the whole period of time considered can be calculated as the maximum depth of regression (1 10 m) plus 60 m i.e., 170 m. The question is to know whether glaucony presently observed deeper than 170 m is older than 45,000 years or not. This question was debated in the past (Odin and Giresse, 1976), the latter author supported the view that glaucony is younger than 24,000 years B.P. everywhere in the region. However, we have seen above that glaucony must be older than 30,000 years deeper than -1 10 m if our hypothesis of in situ genesis deeper than -60 m is correct. Moreover, the study of cores collected from -1,000 m by Bongo-Passi (1984) showed that sediments biostratigraphically confidently dated at 75,000 years B.P. already contained probably perigenic glaucony. In conclusion, there are proofs of a glauconitization process older than 75,000 years on the Congolese continental slope. At the top of the slope and on the deep shelf, glaucony seems to be older than 30,000 years; on the continental shelf, glaucony is younger than about 12,000 years and the smectitic stage

242 of evolution of glaucony observed at -80 m -100 m depth may be obtained following a period of time of less than 8 to 9,000 years of immersion.

The glauconi tizat ion process The study of the glaucony facies in the sediments of the Gulf of Guinea is helpful in order to reconstitute the history of the glauconitization process because various stages of a continuous evolution are present, and because no burial process has modified the purely surface evolution.

Morphological evolution In stage 1 (Fig. 9) the faecal pellet, deposited in a shallow environment, is made of the mud of the sediment, probably enriched in biogenic particles since the mud-eaters most probably select a mud rich in organic matter. The stage 2 in Figure 9 concerns an ochreous-green to light-green faecal pellet with a large proportion of the initial biogenic material dissolved. This leads to a high porosity in the pellets. The pellet appears to be made of small globules sometimes arranged in caterpillar-like structures. A few biogenic structures are preserved but most of them are due to a replacement process which substitutes authigenic clay for carbonate.

1.5-3.0 Yo

K20 = Q5-1.5 O/O

- 12

FezOa.

8

AI2O3-

15- 20%

O/O

21-24

O/o

24

Yo

7- I I

Yo

6- 7

O/o

s r n e c i i t e m globules

clay s i z e 0 kaolinite DETRITAL

AUTHIGENK

coarse

4.0-7.0 Yo

quartz-carbonate

118 clay af lakes-rosettes

GSO 87

Figure 9. Schematic evolution of the faecal pellets of the Gulf of Guinea during glauconitization.Three stages from a continuous evolution, are shown. Stage 1 is the initial faecal pellet. Note that there is a volume increase between stage 2 and stage 3, both green coloured. The chemical data are from micro-analysis estimates.

243 Stage 3 in Figure 9 is a dark-green cracked grain which has partly lost its initial form due to an increase of the volume provoking cracks in the grain surface. Well-shaped microcrystals (rosettes or flakes) are present in most of the grain but for the periphery. The latter still shows poorly organized globules. This indicates a more active crystal growth in the interior of the grain than in the outer part. The pores of the grain have been entirely filled during the evolution and the grain appears now generally homogeneous. This indicates both that the greatest portion of the initial detrital material has been removed and that the authigenic material initially crystallized has been re-arranged i.e., recrystallized. Off Ivory Coast and Nigeria it seems that only stage 2 was reached; off Congo, stage 3 is present and abundant. The following evolutionary stages are absent in the Gulf of Guinea and will be documented in next Chapters.

Mineralogical evolution The mineralogical evolution of a pellet may be subdivided in two questions: the first concerns evolution of preliminary components and the second considers the neoformed minerals. The evolution of the initial components of the grains depends on their nature; but the result is always similar: elimination; this needs a short duration for carbonate, a longer duration for kaolinite and quartz. Our detailed elementary and isotopic analyses have shown that detrital remnants may be present at a comparatively late stage of evolution of the green grains at a time when their initial morphology and mineralogy is difficult to recognize. X-ray diffraction may not be sufficient to detect such remnants at a time when they still greatly influence the isotopic composition of the bulk grains. The actual nature of the components which contain unstable isotopes is not known. But the fact that inherited isotopic compositions are preserved for stable and unstable isotopes shows that inherited very stable crystalline or sub-crystalline structures are preserved. This, in turn, leads to a problem concerning the elementary composition of the grains and more specifically to the iron content. The mud of the sediment contains about 10% iron oxides, the less evolved pellets already contain more than twice as much iron and half as much aluminium than the mud (Table 1). If only half of the pellet is made of authigenic iron-rich, aluminium-poor green clay as deduced from isotopic data above, then there is something wrong somewhere: either more than 50% of the pellet is made of authigenic material or the high iron (and low aluminium) contents must be explained. It is suggested here that an important proportion of iron analysed in the less evolved green faecal pellets is not in the form of a silicate or any well-crystallized structure and, therefore, that it is not observed on X-ray diffraction patterns. The low alumina content could be explained by the fact that the inherited alumino-silicate structures first make free the octahedral cations during alteration (eliminating the octahedral aluminium) while the tetrahedra remaining bonded still keep the oxygen and argon isotopic composition trapped. The discussion remains speculative however, and this problem would merit a specific study. Of more interest to the present study is the evolution of the authigenic green marine clay. The combination of X-ray diffraction and nanostructure studies

244 suggests that there are two steps in the development of authigenic minerals. The first step is the crystal growth of a smectite-type mineral which develops in the pores of the faecal pellets. This smectite makes globules or caterpillars in stage 2 of Figure 9. The authigenic clay is already rich in iron but poor in potassium. The second step of evolution is the re-arrangement of these smectitic globules which entirely recrystallize to form better shaped microcrystals. During this evolution, the crystal growth is very active and new authigenic material is added to that formed previously. The general mineralogy of the green grains is characterized by a progressively increased potassium content and a concomitant decrease of the interlayer thickness. In short, authigenic ferric smectites are formed initially and this material is recrystallized later to form potassium richer 2: 1 clay minerals which, also by crystal growth increase the general volume of the neoformed material.

Mechanism of glauconitization In the Gulf of Guinea, the mechanism of glauconitization was mainly considered for kaolinite faecal pellets. This mechanism has two effects. The destruction of the initial detrital and biogenic components progressively eliminated from the sediment particle is the first effect. The second effect is the appearance of a marine green clay which is made of 2:1-type ferric clay minerals. These clay minerals compose a continuous series of phases of which the initial end-member is a smectite. The most evolved member, in the Gulf of Guinea, has a 11 A thick layer as a mean and is found off Congo. The early theory concerning the mechanism of glauconitization (and that of a number of clay mineral evolutions, probably the large majority) was a transformation process which would modify previously altered detrital 2: 1 crystal structures by replacing iron for aluminium (Burst, 1958a and b; Hower, 1961, and many others). This mechanism was accepted until our own research showing, as above, that the initial material within which the glauconitic minerals form, was usually not composed of a 2: 1 precursor (Odin, 1975a). In the case of the glaucony from the Gulf of Guinea the transformation process would require acceptance of the possibility for kaolinite to be progressively modified by replacement of iron for aluminium in the octahedra and addition of a tetrahedral sheet to the 1:l structure in order to form a smectite-type layer. This mechanism is unlikely. In contrast to that mechanism, it is considered here that the glauconitic smectite (formed at the beginning of the glauconitization and leading to the verdissement of the initial particle of the sediment) results from a crystal growth of neoformed minerals using cations. Consequently, the initial mechanism of glauconitization is the neoformation of a ferric, slightly potassic, smectite. As for other clay mineral evolutions, the transformation process of early sedimentologists was also invoked in order to explain the evolution from a potassium-poor to a potassium-rich glauconitic (2: 1-type) clay mineral. This phenomenon of potassium enrichment, thought to be the fundamental part of the glauconitization process by previous authors, is only the second side of this evolution. Many authors until recently suggested that the potassium enrichment

245 was linked to a concomitant iron enrichment. In agreement with Foster (1968), the Congolese case study shows that iron is mostly present before potassium enrichment. However, a portion of this iron is possibly not combined in the silicate phase as discussed above. It is argued here that the potassium enrichment does not result from a transformation process for the following reasons. The above study has shown that the quantity of authigenic clay was clearly increased between stage 2 and stage 3 in Figure 9. Therefore, the evolution cannot be a simple transfornation of previously available material. Moreover, the nanostructure study has shown that the shape of the microcrystals was generally a function of the stage of evolution; this can only be understood if a recrystallization process occurred with re-use, in a potassium-richer mineral architecture, of cations previously gathered in a potassium-poorer architecture. In short, the two steps of formation and evolution of the glauconitic minerals during glauconitization both result from crystal growths.

Environment of glauconitization The glaucony of the Gulf of Guinea is mainly abundant at the top of the slope; glauconitized material is also present up to about -80 m of depth and probably slightly less (-60 m). At these depths, even in tropical areas, the temperature is not high (less than 15°C).The marine character of the glauconybearing sediments is obvious everywhere. Compared to the verdine facies discussed above in Part B, the glaucony facies is clearly deeper everywhere in the Gulf of Guinea and the east Atlantic border. The detailed study of the Congolese deposit indicates that glauconitization occurs in the superficial sediment, which it has been possible to dredge i.e., the uppermost decimetres of the deposited sediment. From observations made in cores with a length of up to 7 m, glaucony has been observed from all horizons but no mineralogical evolution has been detected as a function of depth. Therefore, all data from the Congolese shelf concern geochemical processes dominated by pure sea-water as the fluid surrounding the glauconitized material. In this sediment in close contact with sea-water, glauconitization occurs exclusively in deposited sand-size particles. Faecal pellets were mainly considered here, but foraminiferal tests, mollusc fragments, or nearly everything with a grain-size between 100 pm and 500 pm undergoes a verdissement in the Gulf of Guinea. The role of an initial granular substrate of verdissement appears determinant. It seems that the grains determine a specific microenvironment different from that at the exterior of the grain, and from that in the sea-water itself. The specificity of this microenvironment is discussed further at the end of this Part: Chapter B4. Within the grain itself, differences of microenvironment have been observed. These observations generally indicate that the internal portion of the initial substrate is more favourable for the glauconitization process than the periphery. Three sorts of observations are relevant. Nanostructural features are the easiest to interpret. The outermost part of the

246 green grains always remains formed of globules even in the most evolved deep glaucony. In addition, the small flakes and rosette-like idiomorphic figures are only present in the central portion of the evolved grains. The central part of the grain preferentially allows recrystallization to occur and glauconitization to reach its more evolved stage. Micro-chemical data were obtained from the most evolved pellets of the Congolese deposit. Table 4 indicates that the outermost portion of the grains (less than 10 pm in thickness) contains more alumina and less potassium and iron. The absolute values lack precision due to the difficulty in obtaining satisfying polished surfaces; moreover artifacts cannot be entirely ruled out on the border of the grains, but the differences are very significant and indicate that the central part of the grains is more evolved than the external layer: higher potassium content. Table 4. Micro-chemical analytical data for evolved faecal pellets of the Congolese

sediments. Periphery of t h e grains Fe2O3 ( t o t a l ) K20 A 203

Central portion of the grains

20 %

25 %

4 t o 5 %

6to796

1 0 to 1 2 %

7 t o 8 %

The third observation concerns morphological data. The most evolved grains are larger than the initial pellets and are cracked. The cracks result from a differential speed of crystal growth with quicker crystallization in the centre of the grains. This cracking of the external crust of the pellet occurs in a similar manner to the surface of a loaf when raising and baking. These cracks therefore result from a general volume increase of the initial grain. SUMMARY AND CONCLUSION

-Glaucony is present in the sedimentary cover dredged or cored in the Gulf of Guinea and other portions of the east Atlantic border mainly at depths between -60m and -300 m. -In the Gulf of Guinea, the main habit of the green marine clay consists of faecal pellets; diverse other sand-sized sedimentary particles also undergo the verdissement. -An interpretation of these glauconitic grains can only be correct following recognition of their fundamental dual nature: 1) the material of the initial grain (the substrate) and 2) the authigenic green marine clay. Like the mud of the sediment, the initial faecal pellets are composed of a major proportion of kaolinite; the rest comprises other detrital clays, silt sized mineral or organic debris, bioclasts, and probably other components which do not diffract X-rays.

247 This initial material progressively disappears during evolution. The second category of material, authigenic green marine clays, is formed within that substrate and gives to the particle its specific physical characters: the green colour, and the paramagnetism. The green marine clay of the glaucony facies is not a unique mineral, its composition varies from a smectitic potassium-poor initial member to a potassium richer 2:1-typeclay mineral. The elementary and isotopic compositions of these authigenic minerals are different from those of the muddy fraction of the sediment. They are characterized by a comparatively homogeneous elementary composition with a remarkable very high proportion of ferric iron and low aluminium content compared to other 2:1 clay minerals. The beginning of the glauconitization process is diachronous on the Congolese shelf allowing us easily to collect variously evolved samples as a function of depth of occurrence. The evolution phenomena can be understood and described independently of any burial or diagenetic alteration. The glauconitization process appears fundamentally to follow a two step model. The first is the de-novo crystal growth of a femc smectite (the glauconitic smectite). The second is the recrystallization of this smectite which leads to the formation of provessively K-richer 2:1 clay minerals. These K-richer clay minerals pursue their crystal growth as shown by the sizeable volume increase of the initial substrate. The detailed study of this glauconitization process therefore suggests that the fashionable theory of transformation, developed and accepted in the years 1960 to 1980 by most sedimentologists, as generally explaining the clay mineral evolutions in the sedimentary environment is not correct for glauconitization. The work summarized above was one of the fiist demonstrations of the inadequacy of that transformation theory, now rejected by many detailed studies of clay mineral evolutions in nature, and especially glauconi tization. Glauconitization is greatly favoured by sand-sized substrates, therefore the microenvironmental role of the granular habit is fundamental for crystal growth. This mechanism occurs in superficial sediments in close contact with sea-water far from any detrital continental input to the sea. AKNOWLEDGEMENTS

Sediments of the continental shelf off Congo and Gabon were provided by P. Giresse, those of the shelf off Ivory Coast by L. Martin. The discussion of the sedimentological results with these colleagues is aknowledged. I thank Eddy Keppens for his isotopic analyses and the related discussions during the meetings where he presented the data. Argon isotopes were measured by the author thanks to the kind welcome of M.G. Bonhomme in Strasbourg, J.C. Hunziker in Berne, H. Kreuzer in Hannover, and D.C.Rex in Leeds. E. Lebrun, M. Lenoble, and G. Richebois are thanked for their chemical data. Careful photographic processing by 0. Fay is much appreciated. This chapter has been revised for English by John Thompson.

This Page Intentionally Left Blank 248

249

Chapter C2 GLAUCONY FROM THE MARGIN OFF NORTHWESTERN SPAIN by G.S. Odin and M. Lamboy INTRODUCTION

Glaucony from northwest Spanish continental margin has been described by Lamboy (1967). This study of superficial sediments is of interest because detailed morphological investigation was carried out using high magnification. Evolutionary processes have affected several sorts of substrate; this is illustrated in the French literature (Lamboy, 1968; 1975; 1976; Lamboy and Odin, 1974; Odin, 1972a; 1975a; Odin and Lamboy, 1975). The mineralogy of the Spanish green grains was first studied by Caillkre and Lamboy (1970a, b) and revised by Lamboy and Odin (1975). The latter authors also published a review of the morphological features and a discussion of the geological significance of the facies in the area. The present chapter is mainly intended to summarize the large volume of available data which has allowed the formation mechanism and geological significance of glaucony facies to be understood. The study covers about 550 sediment samples collected off Galice Province (northwestern Spain) from depths ranging between -50 m and -1000 m. From a granulometric point of view, the dredged sediments are clayey sands with a clay-size fraction content generally lower than 25 percent (Fig. 1). The sandsize fraction at the sea-bottom is characterized by the presence of carbonate. Figure 1 shows the proportion of carbonate in the sand-size fraction. This carbonate is bioclastic and is particularly abundant at depths shallower than -100 m. The rest of the sand is composed of quartz and glaucony. The clay-size fraction is remarkably homogeneous from a mineralogical point of view; it is mainly formed of illite with a small proportion of chlorite and kaolinite. This composition is similar to that of the clay found in the coastal river sediments. Finally, an interesting characteristic of the sedimentary cover in the area is the presence of green grain-bearing nodular phosphate, briefly quoted by Collet (1908) and studied in detail by Lamboy (1976). DISTRIBUTION OF GLAUCONY

Figure 2 shows the distribution of magnetic grains separated from washed sediments. The measured contents attain 50% or more on the deepest portion of the shelf down to depths of around -200 m. To the north, the magnetic grains contain smooth, brown particles which were initially considered to be authigenic 'berthierine' (Caillkre and Lamboy, 1970b; Lamboy, 1976). Elsewhere, the magnetic grains represent typical glaucony with depth distribution similar to

250

CAR EON ATE

0

<20%

20-50% 50-80%

,

50km

> 80%

Figure 1. Distribution of the clay-size fraction (< 65 pm) and carbonate fraction (> 65 pm) off NW Spain.The area under study comprises the deeper half of the shelf and the slope. Dots in the bottom figure represent nodular phosphate. (Reproduced from Lamboy, 1976)

25 1

GLAUCONY

0 0

0-1096 10-20%

20-50x

> 50%

Figure 2. Magnetic grain (mostly glaucony) content in superficial sediments off northwestern Spain. All samples deeper than 50 m contain green or partly green grains in a proportion above 1%. (Reproduced from Lamboy, 1976)

that observed on the west African Atlantic margin. The distributions of glaucony and clay are unrelated (compare Fig. 1 top and Fig. 2). From a granulometric point of view, separate analyses have been undertaken for the quartz fraction and the glauconitic fraction. This study has shown that there is no relation between the grain-size and mode of glaucony (mode between 0.6 mm and 0.18 mm) and that of quartz. Glauconitic grains are often coarser than quartz grains; but the difference in size between glaucony and quartz varies greatly from one site to another (Lamboy, 1976). In short, quartz and clay represent a detrital component, and glaucony has an independent distribution showing its in situ authigenic nature. There is, however, a clear, inverse relationship between the abundance of carbonate and that of glaucony (compare Fig. 1 bottom to Fig. 2). This relation is very significant since the glauconitization process mostly occurs at the expense of carbonate bioclasts, as will be shown below. Where glaucony is abundant, carbonate bioclasts have disappeared. This dependance is corroborated by the distribution of the combination: shelly remnants/green grains with intermediate forms, on the western side of the Spanish margin. Figure 3 considers four habits: 1) white, shelly debris, more or less perforated, 2) partly green shelly remnants, 3) olive-green grains, 4) dark-green grains.

252 present sea-level

1 Om

~

100

1

I

2m

3

0

4

0

2a

30(

Figure 3. Relationship between the distribution of carbonate bioclasts and that of glauconitic grains off northwestern Spain. Cross section of the continental shelf. 1) darkgreen grains: 2) olive-green grains: 3) partly green bioclasts; 4) white to grey bioclasts many of them may show borings with green clay inside. (Modified from Lamboy, 1976)

MORPHOLOGICAL FEATURES

The verdissement of echinoderm fragments On the Spanish shelf, various sorts of substrate shelter the glauconitization process. The glauconitic grains resulting from the genesis of green marine clay within a fragment of echinoderm skeleton are easily identified by their distinctly reticulated structure nicely illustrated by Bignot (1976). Photomicrograph 1 in Figure 4 shows this typical habit locally observed in green grains frequently as much as 1.5 to 2.0 mm in diameter. A green grain slightly more evolved than the one in Photo.1 is pictured in Photo.2; cracks appear in locally swollen areas. An even more evolved grain (Photo.3, Fig. 4) shows a large portion of the particle with deep cracks; this grain has been broken in order to show that the reticulated structure has disappeared from the main portion of the grain except for the lower left-hand comer. A detail of a cracked zone of the grain from Photo.2 is enlarged in Photo.4 (Fig. 4). This picture clearly demonstrates that cracks characterize an area which is in reliefon the exterior of the grain. This verruca-like formation essentially represents a growth feature. These observations negate the shrinkage fissure interpretation, often quoted in past literature and more recently by Morton et al. (1984), supplemented by the hypothesis of a dessication (dehydration) process. The cracking of green grains of the glaucony facies results from slow or zero growth at the grain surface at a time when the inside is expanding (see Chapter C1, p. 232). Details of the internal arrangements of the first stage of evolution of the echinoderm skeleton fragments are shown in Photo.5 (Fig. 4). The reticulated carbonate stereom is clearly shown with its smooth surface. The enlarged view in Photo.6 (Fig. 4) shows that the voids in the stereom represent a volume very similar to that of the carbonate skeleton itself. However, photomicrographs 5 and 6 demonstrate that the green marine clay mineral first develops in the pores

253

Figure 4. Verdissement process of fragments of echinoderm skeleton (Scanning Electron Microscope). The marine green clay first fills the pores of the stereom fragments. The carbonate stereom later dissolves and the voids are filled with neoformed green clay. Finally, a recrystallization process destroys the original reticulated texture. (According to Odin and Lamboy, 1974) Scale bars: 100 pm (Photo. 1,2,3,4,5) or 10 pm (Photo. 6.7) or 2 pm (Photo. 8,9).

within the carbonate fragment and does not replace the stereom. During a second stage, the carbonate stereom is dissolved as shown in Photo.7 (Fig. 4),

254 where some carbonate remnants, not totally dissolved, are still visible. The nanostructure of the clay mineral is shown enlarged in Photo.8 (Fig. 4). The globules and caterpillar arrangement are characteristic of an early stage in the evolution of the green clay. At a later stage, as, for example, the one depicted in the upper right interior of the grain shown in Photo.3, better shaped automorphic figures, like the rosettes illustrated in Photo.9 (Fig. 4), may be observed. At this stage, the grains become homogeneous, and the initial porosity disappears due to filling of the voids with neoformed clay minerals. The reticulated structure disappears following a general recrystallization process in the interior of the grain, but the external reticulated habit partly remains between verrucous areas. In summary, all grains observed above are light-green or olive-green in colour. During the first stage of evolution, glauconitic minerals grow de novo in the pores of the carbonate substrate; secondly, the substrate is dissolved and glauconitic minerals can grow in the newly acquired porosity; thirdly, the glauconitic minerals undergo a general recrystallization process which obliterates the initial reticulated structure more or less completely.

The verdissement of bored shell fragments The most abundant initial substrate for glauconitization identified off northwestern Spain was shell fragments. These shell fragments, usually 0.2 to 2.0 mm in diameter, are generally not directly glauconitized. The first stage in the evolution of the carbonate particle is the formation of a porosity which is produced by boring micro-organisms. Photomicrograph 1 in Figure 5 is a view of a section of a 5 mm long shell fragment. From the interior to the exterior, the grey coloured centre is made of well-preserved carbonate; the whitish portion represents partly dissolved carbonate; and the external black filaments are filled with dark-green glaucony. Photo.2 in Figure 5 shows the external surface of the bioclast. Borings parallel or perpendicular to the surface are now filed with dark-green clay. Such perforated grains have been submitted to dilute acid solution to remove the carbonate. Photo.3 (Fig. 5) was obtained from a carbonate free grain after removal of the above portion in order to see the arrangement of the filaments within the grain. Note that the external zone of the 2 mm long decarbonated grain is much richer in clayey fiaments and forms a nearly continuous film compared to the interior where few filaments are present. An enlarged view of the external zone is shown in Photo.4 (Fig. 5). An enlarged view of a filament from the interior is pictured in Photo.5 (Fig. 5). The section of one filament is enlarged in photomicrograph 6 (Fig. 5); the contorted blades, 1 to 2 pm long, are typical of a clay mineral. The clayey filaments were found in all glauconitized carbonate substrates from the Spanish sediments. For example, a few echinodermal skeleton fragments were bored by

-+-+-+

Figure 5. Verdissement process of bored shell fragments. Optic microscopy (Photo.1, 2) and scanning electron microscopy (Photo3 to 13). Scale bars show 100 pm (Photo. 2,3,4, 10) or 10 pm (Photo. 5, 7,8,11) or 2 pm (Photo. 6 , 9 , 12).

256 micro-organisms (Photo.7 in Fig. 5). Borings are not always homogeneously filled with green clay, and some clayey filaments of the decarbonated grains show a composite section (Photo.8 in Fig. 5). An idiomorphic nanostructure resembling rosettes can be seen in the enlarged section in Photo.9 (Fig. 5). All pictures mentioned above represent comparatively early stages of evolution where the recrystallization processes have still not modified the organization of the primary green clay fillings. However, a highly-evolved stage of the same process has been identified in many grains after observation in thin section. These grains are black-green to blue-green in colour and comparatively dense; they display a zonation with a lighter and soft central part. Following ultrasonic treatment, this soft heart of the grain is removed (Photo.10 in Fig. 5). In an enlarged view of the interior of a dark-green grain submitted to ultrasonic treatment (Photo. 11 in Fig. 5), no trace of the filaments in the cortical zone (i.e. the previous densely bored part of the carbonate grain) can be seen. The nanostructure is lamellar and characteristic of a highly-evolved stage of glauconitization (Photo. 12 in Fig. 5). Proof that these dark blue-green grains result from the evolution of bored carbonate particles has been observed in some thin sections where filament ghosts are present in the soft centre of the grain. In summary, the densely bored shell fragments constitute a substrate physically equivalent to echinoderm fragments with similar porosity. Green clays first develop by filling the pores. Later on, the rest of the carbonate is dissolved and replaced by green clay. Finally, recrystallization occurs which mostly obliterates the initial filament arrangement. The phenomenon is more efficient on the outside than in the interior of the comparatively large grains, within which ghosts of filaments may be found in the heart of the particles at a very late stage of evolution. Off northwestern Spain, glauconitization of bored shell fragments is predominant.

The verdissement within the foraminiferal tests Glauconitization of tests of micro-organisms is widespread in the superficial sediments collected from northwestern Spain. Infillings of foraminiferal tests are the most abundant. They are generally similar to those described elsewhere by several authors (Murray and Renard, 1891, pp. 378-391; Collet, 1908; Caspari, 1910; Wermund, 1961; Ehlmann et al., 1963; Bjerkli and Ostmo-Saeter, 1973). The microscope allows numerous miliolids, of which the test is filled with yellowish to dark-green clay (Photo. 1 in Fig. 6) to be observed. It may be noted that the test is well preserved when the first clay minerals are formed within it; but when the clay minerals are dark-green, the carbonate test is dissolved. Photomicrograph 2 (Fig. 6) shows the infilling of an Elphidium; the clay entirely fills the test. One test has been opened with pliers (Photo.3 in Fig. 6) and the nanostructure of the infilling is pictured in Photo.4 (Fig. 6). Blades typical of a clay mineral are present. Cracks appear at the surface of the lightgreen coloured infilling (Photo.5 in Fig. 6 ) when the test has been dissolved in the sediment. Later on, the infilling becomes dark-green with cracks deeper

257

Figure 6. Verdissement within the tests of foraminifers. Optic microscopy (Photo.1); scanning electron microscopy (Photo.2 to 6) Scale bars show 100 pm (Photo.l,2,3,5) or 10 pm (PhOt0.4).

than previously. Photomicrograph 6 (scale similar to Photo.2, Fig. 6 ) shows a fragment of Elphidium infilling after natural dissolution of the test, cracks are present. By this process, the grain loses its original form. As a result, the initial substrates of the most evolved grains are difficult to identify especially when the cracks are deep enough for the grain to be broken in two. In short, verdissement within foraminifera1 tests mainly consists of filling chambers with authigenic green clay. The carbonate test itself is dissolved allowing the green grain to evolve freely into the common dark-green cracked grain. Therefore, the role of the test is to create the favourable physico-chemical microenvironment for starting the verdissement.

The verdissement of detrital mica flakes The evolution of mica flakes into 'vermicular pellets' (Triplehorn, 1966) or 'accordion-like grains' (Odin, 1972a) is one of the famous examples quoted in literature (see below) as supporting the hypothesis of a transformation process.

25 8 This hypothesis assumes that evolution from denital mica (or from another 2: 1 phyllosilicate) into glauconitic mica occurs following a restricted and progressive cation exchange process (Fe for A1 especially) within a perennial minera architecture. This hypothesis has been proposed by Galliher (1935) and was beleived to apply to all the glaucony collected from the Quaternary sediments in Monterey Bay (California). However, Hein et al. (1974) showed that, in Monterey Bay, a considerable proportion of the glauconitic green grains did not result from the evolution of detrital mica flakes. The transformation hypothesis was later generalized and improved by Burst (1958a, b), Hower (1961), Seed (1965), Triplehorn (1966), to explain the genesis of all glauconitic minerals as well as most other clay minerals evolving on the surface of the earth. Detailed scanning electron microscope study of the verdissement of mica flakes has shown the transformation hypothesis to be invalid for this example (Odin, 1972a), leading to some doubt on the general validity of this theory widely accepted in the sixties and seventies. Odin (1972a) identifies three preliminary reasons why the transformation process could not have taken place. The first one alone is conclusive and concerns the question of the mass of material involved. The densities of biotite flakes and glauconitic accordion-like pellets are similar (between 2.5 and 3 g/cm3). However, in a sediment, detrital mica flakes are usually thin compared to their diameter. In contrast to this, green vermicular pellets, which we prefer to call accordion-like grains in order to emphasize the presence of cleavages perpendicular to their longitudinal axis, are usually thicker than they are wide. In northwestern Spain sediments, the glauconitized biotite flakes present at depths shallower than -90 m are often wide (more than 1 mm in diameter) and thin. Deeper, true accordion-like green grains were observed, their diameter being between 0.1 mm and 0.4 mm and their thickness often more than 1 mm. Therefore, the wide flakes appear much less prone to glauconitization than the small flakes. Photomicrographs 1 and 2 (Fig. 7) show some examples of accordion-like grains found in the Spanish sediments. These grains are all several tens of times heavier than were the initial mica flakes from which they formed since the diameter has remained the same but they have grown in thickness. Galliher himself (1935) noticed that his final 'vermicular pellets' evolved from biotite flakes 10 to 20 times thinner. It is clear, therefore, that material from the biotite cannot produce all material from the glauconitic clay by a simple chemical exchange or re-arrangement of octahedral cations as implied in the transformation hypothesis. Secondly, transformation is unlikely to occur because the chemical composition and mineral nature of the initial substrate and that of the resulting authigenic clay are in no way connected. Off northwestern Spain and in other formations of various ages, both biotite and muscovite have been observed as the initial substrate of the accordion-like glauconitic grains. Consequently, the presence or absence of iron in the initial mica is not a fundamental need for the formation of glauconitic minerals. In addition, it has been shown previously

259

Figure 7. Verdissement of mica flakes. Optic microscopy by A. Matter (Photo.1); SEM by G.S. Odin in Dijon (Photo.2 to 7). All pictures from Spanish sediments collected by M. Lamboy. The accordion-like green grains are formed by de novo crystal growths of green marine clay within the cleavage planes. Scale bars show 100 pm (Photo. 1 , 2 , 3 ) or 5 pm (Photo. 4,5,6,7). (According to Odin, 1972a; 1974; Lamboy and Odin, 1974)

260 (Chapter B3) that, in a different environment, the resulting authigenic clay is not a glauconitic mineral, but phyllite V which does not have a 2: 1-type mineral architecture. Therefore, the nature of the resulting authigenic mineral is not influenced by the nature of the substrate; a simple octahedral cation exchange or re-arrangement cannot account for these observations. There is a third reason why the transformation process is an unlikely explanation for the verdissement of mica flakes. Even when evolution of the mica into an accordion-like grain is nearly complete, mica remains are preserved in the cleavages of the broken grain. Therefore, the cations of the initial micaceous substrate are not only quantitatively insufficient to form the final glauconitic material, but sometimes they are not used at all. The cleavage surfaces perpendicular to the longitudinal axis of accordionlike grains were observed using a scanning electron microscope. In Figure 3, photo.3 shows a broken mica flake. Locally, the surface is flat and represents the unaltered or little altered biotite mica (Photo.4 in Fig. 7). Small blades begin to develop on this surface by a crystal growth process. Photomicrograph 5 (Fig. 7) testifies to the multiplication of these blades from which a box-work fabric results. Upon this structure, rosettes develop unrelated to the micaceous substrate (Photo.6 in Fig. 7). The relationship between the mica substrate and the neoformed green clay is shown in Photo.7 (Fig. 7). In this picture, unaltered mica is visible in a fissure while in the open cleavage between two mica layers the green clay is growing into authigenic figures. The material grown in the cleavage leads to an opening of the accordion and the thickening of the initially thin flake. In summary, the transformation theory developped after observation of the evolution of biotite mica flakes into glauconitic, accordion-like grains is not c o n f i i e d by more detailed study. Mica flakes must only be considered as physical substrates within which cleavage planes delimit lamellar pores where green marine clay may grow as in the various types of bioclast pores studied above. The material from the detrital flakes, cations or groups of cations, does not significantly contribute to the formation of the authigenic clay and a fortiori cannot be suspected of being a prerequisite chemical or crystallographical precursor to the formation of glauconitic minerals.

The verdissement of quartz grains The verdissement phenomenon in quartz grains has long been known (Cayeux, 1916). A number of quartz grains collected from superficial sediments from the Spanish continental shelf and upper slope are green coloured. It has been difficult to obtain good pictures from them, but some are available for a similar phenomenon observed in Albian sandstones from Boulonnais, Paris Basin, France (Odin, 1975a; and Fig. 8). Photomicrograph 1 in Figure 8 was obtained from a thin section of a rounded quartz grain, 2 to 3 mm in diameter. A fissure in this quartz has been invaded by green clay, grey in the photomicrograph. A more highly fissured quartz grain is shown in Photo.2 (Fig. 8). Note that the fissures are more or

261

Figure 8. Verdissement of quartz grains. Optic microscopy of an Albian sandstone: Photo.1-3; scanning electron microscopy of green quartz collected from Spanish sediments: Photo.4-6 (according to Odin, 1975a). The glauconitic minerals are formed in the fissures and rarely replace the silica. Scale bars show 100 pm @hoto.l,2,3) or 5 pm @hOt0.4,5,6).

less parallel in the form of a net, which is green coloured in thin section. In Photo.3 (Fig. S), a broken, fissured, and dissolved quartz grain can be seen to the left, and, to the right, a dark-grey rounded particle of a very similar size and form. In thin section this grey grain is green coloured; this has been interpreted as the ghost of an initial quartz grain because numerous small remnants of quartz are present; those remnants are more or less parallel (NS in the photomicrograph) and this orientation most likely represents the orientation of the previous fissures along which glauconitization proceeded. Quartz grains similar to that in Photo.2 are not rare in the Spanish deposit; one of these was broken for scanning electron microscopic study. Photo.4 (Fig. 8) illustrates a

262 quartz (Q) fissure which has been opened, on which a film of green clay is present. In Photo.5 (Fig. 8) to the left, the green clay shows a xenomorphous surface following the form of the fragment of quartz which has been removed. To the right, more or less automorphic figures, also shown in Photo.6 (Fig. €9, are very similar to some of those observed in the opened accordion-like grains illustrated in Figure 7. In summary, fractures in a quartz grain may harbour the genesis of green marine clay, which seems to develop at the expense of the siliceous material. Quartz appears less favourable than biotite for green clay genesis because the green clay only replaces the progressively 'dissolved quartz grain, which retains its volume and shape. This process of replacement is not specific to quartz and was also observed for certain carbonate substrates as shown below.

The verdissement of non-bored shell fragments Mollusc shell fragments are sometimes characterized by a well-shaped structure showing prisms or crystallographically oriented layers. A similar arrangement has been observed in magnetic green grains from the Spanish shelf and in several ancient sediments such as the Lutetian glauconie grossi2re from the Paris Basin (Odin, 1969; 1971). The structure of these green grains obviously results from the delicate replacement of the carbonate structure as illustrated below. This mechanism appears different from the one illustrated in the four paragraphs above. The most typical green grains are flat and green coloured, and may reach large dimensions. In some cases, the original external ornamentation of the shell (such as ribs or striae) is still preserved (Photo.3 in Fig. 9). The enlarged view of the green grain in the upper left-hand corner of Photo.3 (Photo.4 in Fig. 9) shows a nice zebra-structure similar to that of a carbonate shell fragment (Bignot, 1974). A thin section of this sort of green grain (Photo.2 in Fig. 9) displays layers alternatively oriented in different directions (crossed nicols observation). At a later stage of evolution, a recrystallization process occurs concomitantly with crystal growth; cracks develop and the grain becomes thicker. Photo.1 (Fig. 9) illustrates this stage of evolution where one side of the grain exhibits the initial zebra-structure and the other a verrucous, cracked aspect. The cracked side is convex but the side showing the zebra-structure preserved is generally concave. This observation again proves that the cracked side actually represents later growth which followed the replacement phenomenon. The scanning electron microscope allowed more detailed pictures showing the relations between the layers to be obtained (Photo.5 in Fig. 9). Photo.6 (Fig. 9) illustrates the various alternative orientations of the successive layers of green material. An enlarged view (Photo.7 in Fig. 9) displays a very particular arrangement of the clay minerals which are actually oriented and do not form idiomorphic figures.

263

Figure 9. Verdissement of non-bored shell fragments. Optic views (Photo.l,3,4); thin section (Photo.2, courtesy of A. Matter); SEM (Photo. 5, 6,7). All pictures illustrate Lutetian green grains from the Paris Basin; similar green grains with a zebra-structure are present on the shelf off northwestern Spain. (Partly according to Odin, 1969; 1971) Scale bars show 500 pm (Photo. 1 , 2 , 3 , 4 ) or 10 pm (Photo. 5,6,7).

264 Lamboy (1976) has illustrated a similar process of replacement for biogenic carbonate fragments without initial zebra-structure and without borings. Again, dark-green cracked grains result from the complete evolution of the carbonate bioclasts. Such a replacement (without zebra-stucture) has also been observed in bored shell fragments and involves glauconitization of the areas between the borings. Therefore, the replacement process is common and constitutes a phenomenon apparently distinct from the filling process. In summary, magnetic green grains with a zebra-structure are easily recognizable (Triplehorn, 1966). They result from the verdissement of shell fragments and represent a replacement process. Following this process, later evolution consists of disordered crystal growth in the centre of the grain. The zebra-structure may be preserved on the surface but is usually destroyed when cracks develop.

Other morphological features

Formation of rounded grains The foregoing six examples of verdissement describe the formation of green marine clay within a granular substrate. The resulting green grains are initially formed either by replacement or by filling pre-existing pores. Later on, crystal growth modify the form of the substrate and a cracked grain results. However, a large proportion of the dark-green grains separated from the superficial sediments of the deepest portion of the Spanish shelf have a rounded form and a smooth surface. Lamboy (1976) has shown that these rounded grains result from the evolution of cracked grains, and represent a new evolutionary stage. Two processes are able to transform the cracked grains into rounded grains. The f i s t process consists of filling the cracks in already well evolved darkgreen grains (Lamboy, 1975; 1976 p.123). To observe the presence of one or several generations of filled cracks, special preparation in oils of different indices is needed. The green clay which fills these earlier cracks is usually lighter than the rest of the grain. An easy means of revealing these concealed cracks is to submit the rounded grains to ultrasonic treatment. This removes the later green clay which is softer than that formed earlier. The second process able to produce rounded grains is the addition of what has been described as fibro-radiated rims (Collet, 1908; Cayeux, 1916). According to observations by Bailey and Atherton (1969), Zumpe (1971), or Lamboy (1976), these cortices may be either clayey or clayey and phosphatic. Photomicrograph 1 in Figure 10 displays a dark-green rounded grain, about 600 pm in diameter. This grain has been superficially broken with pliers and the external envelope can be observed (Photo.2 in Fig. 10). The cortex is more than 30 pm thick. The green clay crystals are radially arranged, and this explains the fibro-radiated structure observed in thin sections. In some grains, the cortex is composed of several layers with the external one invariably thicker and radially oriented (Photo.3 in Fig. 10). In the majority of rounded grains it is difficult to recognize the original substrate. However, sometimes ghosts of structures have been observed from

265

Figure 10. Dark-green grains made round by addition of a cortex. The phenomenon illustrated follows the evolution shown in Figure 5; it possibly occurs within the sediment after moderate burial. Scanning electron microscopy. Scale bars show 100 pm (Photo.1) or 10 pm (Photo.2,3).

which it can confidently be presumed that the original substrates were bored shelly fragments. The above observations mainly prove that the rounded appearance of dark-green grains from northwestern Spain does not result from an in situ reworking and abrasion of previously irregular grains but correspond to an addition of a later generation of green clay around evolved green grains. A physical rounding of the green grain before this addition of green clay cannot be ruled out, however. T k verdissement of large lithoclasts The complete verdissement of comparatively small sediment particles has been described above. The diameter of these particles may vary from 0.1 mm for foraminifer chambers and quartz grains, to several millimetres for bored shelly fragments. But the sediments from northwestern Spain contain much larger substrates which have partly undergone verdissement. For example, numerous boulders or large fragments of sedimentary rocks are green coloured externally. Carbonate-rich blocks are more favourable to the development of a green film than other boulders.

266

As far as size is concerned, a complete range exists from fully glauconitized small grains (0.5 mm in diameter), through larger grains (more than 2-5 mm in size) where external portions are entirely green and indurated, whereas their centres, although green, have remained soft (see Fig. 5 , p. 255), and onto centimetric or decimetric boulders and blocks with a millimetric green film present at the surface. Detailed observation on thin sections or on broken grains indicated that the green clay had developped at the surface because that part of the particle had undergone physical boring and chemical alteration. This had allowed pores to form and green clay to accumulate in them. MINERALOGICAL STUDY

X-ray diffraction study

Figure 11. X-ray diffraction patterns of decarbonated green grains. Slightly-evolved grains contain less than 6% K 2 0 (ML 210: fillings of bored shell fragments; ML 348: green echinoderm fragments with reticulated structure and cracked olive-green grains). Highlyevolved grains contain about 8.8% K20, and are dark-green in colour (ML 212: dark-green grains).

267 The facies term glaucony has been used above to designate the green grains from northwestern Spain because preliminary mineralogical results have shown glauconitic minerals to be present throughout that area. X-ray diffraction patterns are simple and easy to interpret since the pure authigenic phase can be separated by dissolution of the carbonate substrate, in contrast to the previously discussed example from the Gulf of Guinea (Chapter C1) where the clayey nature of the initial substrate prevented pure authigenic phases, especially in relatively unevolved grains, from being identified. The diagrams obtained varied, but an obvious relationship exists between the morphological evolution and the X-ray diffraction patterns as shown in Figure 11. Bored carbonate fragments were dissolved in dilute acid to obtain pure green filaments; in sample ML 210, the powder was olive-green and the diagram indicates a large proportion of glauconitic smectite. The green clay filling the fragments of echinoderm stereom is also made of very open minerals (sample 348 in Fig. 11). The diagram for sample 349 was obtained from selected, olive-green grains with many cracks. Sample 212 produced a diagram with a comparatively well-shaped peak near 10 A; this sample is predominantly composed of dark-green rounded grains. It is worth noting that, in spite of a highly-evolved morphology and of a sharp 001 diffraction peak located at 10 A, the dark-green grains give an X-ray diffraction pattern with a restricted number of diffractions. This is characteristic of a disorder in the crystallographical structure (Bentor and Kastner, 1964). GLAUCONY ML253

Iq ITa ;

I

Hard middle

Soft centre

./=CO.R7:

Figure 12. X-ray diffraction patterns of various portions of dark-green grains resulting from the evolution of bored carbonate fragments later surrounded with a cortex. The hard nuclei are illustrated in Fig. 5, the cortices are illustrated in Fig. 10. (According to Lamboy and Odin, 1975)

268 The crystallographic characteristics of the green clay vary not only from sediment to sediment, but, within a given sample, from one grain to another. For example, the magnetic grains from sample ML 349 included bored green grains, olive-green grains with numerous cracks, and dark-green smooth grains. Powders prepared respectively from these different grains gave diagrams very similar to the three lower patterns shown in Figure 11. The different portions of the dark-green grains with a cortex were separated using dilute acid and ultrasonic treatment. Three diagrams were obtained for sample ML 253 (Fig. 12). In contrast to the rest of the grain, the cortex is made of glauconitic minerals poorly crystallized (open). The selected hard nuclei show better shaped peaks, and all hkl peaks of the glauconitic mica are visible. However, the 112 and 112- peaks remain much smaller than the 003 peak. This is related to their disordered structure, and possibly to the fact that the randomly oriented powder is easily re-oriented when deposited on the glass slide; this orientation increases all 001 diffraction peaks. This artifact is especially difficult to avoid for the apparently very fine-grained soft, blue-green clay removed from the centre of the grains. The resulting 001 peaks and probably the 003 ones are artificially increased for this fraction. This allows the mean thickness of the layers, which is about 10.2 A, to be observed.

Chemical study Table 1. Chemical analyses of dark-green grains collected off northwestern Spain. Potassium data are maximum values. SiOz

A1203

Fe2O3

FeO

CaO

MgO

Na20

K20

H20-

H2O'

Total

M L 212

46.4

0

27.9

3.2

1.0

4.1

0.2

8.75

( C 242A)

45.7

6.2

25.1

2.1

0.7

4.1

0.2

(8.8)

1.4

6.3

100.6

M L 253

46.9

6.5

23.1

1.6

0.8

3.8

0.2

(8.9)

1.4

6.1

99.3

47.6

4.3

24.6

1.7

0.7

3.9

0.3

(8.9)

1.4

6.1

99.5

45.4

5.8

6.4

1.6

0.7

3.5

0.3

(8.2)

1.6

6.2

99.7

48.7

7.8

19.9

2.6

0.7

3.8

0.2

(8.0)

1.8

6.2

99.8

8.05

99.6

( C 243A)

M L 276 ( G 245A)

M L 350 ( C 250A)

M L 279 (G 252 A )

The chemical composition of glaucony separated from the Spanish sediments was studied by Caillkre and Lamboy (1970b). The authors measured diverse potassium contents down to 3% in light-green grains and around 8% in dark-green grains. They also emphasized the exceptionally low alumina content (which may reach zero percent in a few samples) and the very high iron content. The purity of the separates analysed by these authors was not very good, however, especially for the light-green grains, and the range of values obtained in this preliminary study was much too high. More recent analyses

269 were undertaken on the same sediments after a better purification. From these results, it may be concluded 1) that some goethite may be present in dark-green grains; 2) that the samples thought to be aluminium-free in fact contained a measurable proportion of that element (Odin, 1975a). For comparison with previously published data, Table 1 indicates both the number of the sediment (ML,) and the specific number of the glauconitic fraction purified later (G) for which the first author of this chapter has usually selected the size-fraction -500 pm +160 pm (A). Note that the high potassium contents must be considered as maximum values because reference materials measured together were found systematically 5% too high and because sample G 245A measured at 8.9% in 1975 was remeasured recently at 8.1%.

In summary, the purified green clays have a fairly constant chemical composition, the main variation being the potassium content, which increases when the morphology of the grains indicates a more evolved stage.

Isotopic study Stable isotopes have been measured in various green grains from Spain (Keppens et al., in press). Figure 13 suggests a trend from low 6 l 8 0 values measured for K-poor green grains toward high 6l80 values for K-rich green grains. However, sample ML 253 deviates from the regular trend shown by the other four samples. Confirmatory experiments would be of interest; however, the particular nature of the initial substrate suggests some remarks. Because the purified green grains do not contain remnants of the initial substrate dissolved by acid leaching, the measured values concern authigenic minerals alone. In spite of this, the smectite-rich (K-poor) green grains show a 6l80 value clearly lower than the K-rich grains, and, therefore, mimic a light isotope inheritance. I

ML276

23

22

i

---I

ML212

I1

ML253 ML279

4

4

I

GSO 87

1

I

1

5

6

7

I

8

K20

9

*

W0/o

Figure 13. Oxygen isotopic composition of diversely evolved glauconitic grains from northwestern Spain.(According to Keppens et al., in press)

270 Precise and definitive conclusions cannot be drawn from analysis of these sediments because the moment of glauconitization may have widely influenced the presently measured composition due to the occurrence of large fluctuations in sea-water composition during Quaternary time. Two types of glauconitic grains were analysed for K and Ar: olive-green grains made of an open clay mineral separated from the sediment ML 210, and dark-green grains of a mica-like clay mineral. Table 2 shows rather imprecise results for potassium content, probably due to sample inhomogenity. Argon content was also measured with a large error bar due to atmospheric argon contamination. This was usual 15 years ago when measurements were made by routine analysis. However, the results clearly indicate that the radiogenic argon content of the green grains was not negligible. The apparent ages calculated taking into consideration the potassium content are analytically similar for both types of glauconitic grain. This is in contrast to what was shown for the green faecal pellets from the Gulf of Guinea (Chapter C l ) , where the apparent ages were dependent on the potassium content i.e., on the stage of evolution. This is easy to understand when it is remembered that, off northwestern Spain, the initial glauconitization substrates (mainly carbonate bioclasts) do not contain radiogenic argon. Therefore, the radiogenic argon measured today is totally authigenic in both types of grain, evolved or otherwise, and comes from the decay of the potassium present in the grains. The analysed glauconies were probably formed between 5 Ma and 6 Ma ago. A corollary of this assumption is that, since that time, the measured glauconies have been "dead" i.e., they have not exchanged cations with sea-water and the system is closed. In other words, the glaucony observed on the deep part of the northwestern Spanish continental shelf is relict, and no sediments have been deposited during the last 5 Ma (except for some foraminiferal tests and a small quantity of detrital clay). This conclusion agrees with our present knowledge of the shelf where Miocene sediments underlie the soft sedimentary cover. Table 2. K-Ar analysis of two types of green grains from the sedimentary cover off the NW Spanish shelf. Data by G.S.O. in the University of Berne. (According to Odin, 1975a; Lamboy and Odin, 1975) wt.

Sample

W,

Atmospheric A r

Rad. Ar (n l .g-l)

Apparent age (Ma+ 2 0 )

5.6 ? 0.3

87.4

1.017

5.7 ? 1.7

8.0

62.3

1.504

5.6 ? 0.6

K20

F.P.

ML 210 ( - 175 m )

5.3 2 0.3

h l L 279 ( - 190 m )

8.0

0.2

A.A.

0.3

The conclusion proposed above is similar to what can be deduced from the glauconies from Chatham Rise (New Zealand) analysed by Cullen (1967) or the major portion of the glauconies from the south African continental shelf (Odin, 1985b).

27 1 DISCUSSION

Morphological features and glauconitization process The sediments from northwestern Spain allowed a large variety of initial glauconitization substrates to be identified. Their diverse mineralogical and chemical nature has no particular relationship to the green clay formed within them. A number of substrates are carbonate bioclasts, which usually undergo very complete evolution leading to dark-green grains with little morphological resemblance to the initial clasts. Some substrates observed off northwestern Spain are mineral debris (quartz, biotite); in this case, evolution is not finished: mineral remnants and the initial structure of these substrates can be recognized. Carbonate substrates appear specifically favourable for harbouring glauconitization, as had long been noted by earlier authors (Cayeux, 1932). This unequivocal fact was forgotten by most authors supporting the mechanism of the transformation of inherited clay into glauconitic minerals as the major process in glauconitization. Lamboy (1976) notes that the ability of carbonate constituents to harbour glauconitization is mostly due to their alterability. This is supported by the need for the presence of pores in a substrate for glauconitic minerals to grow, as discussed below. There is a clear trend in all the evolution described for the various substrates from northwestern Spain; during an initial stage, the general structure of the substrate remains unchanged; later, the substrate is destroyed, (more or less depending on its alterability); ghosts of the initial structure may still be recognized during this second stage. A third stage consists of a general deformation of the grain, which becomes cracked whatever the initial substrate. Finally, these cracked grains may further evolve into dark-green grains with the external fissures and cracks filled in.

Mineralogical features and glauconitization process Glaucony from the sedimentary cover off northwestern Spain allowed a clear relationship between morphological and mineralogical evolution to be established. The latter is similar to that described for the Gulf of Guinea in its initial stages, but progresses further (like morphological evolution). The glauconitic minerals form a continuous series from smectitic nascent components up to micaceous highly-evolved components. As already observed for the sediments from the Gulf of Guinea, the potassium content is all what is needed to reliably estimate the degree of evolution. However, with some exceptions, the above general scheme needs developing, for example, because not all substrates in a given sediment are equally favourable, the degree of glauconitization will never be uniform. Furthermore, it has been noted that not all portions of a given substrate are equally favourable for glauconitization to develop; therefore, even a single grain will display various evolutionary stages, as has been shown by X-ray diffraction. This indicates that multiple analysis of various fractions of each grain and sediment

272 will need to be undertaken in connection with each mineralogical characteristic, if the history of a sediment is to be accurately documented.

Age of the glauconitization process off northwestern Spain The sedimentological significance of the glauconitic sedimentary cover off northwestern Spain is made more complex by its long history. Geochronological and sedimentological evidence indicates that a large proportion of the glaucony collected from the shelf was formed 5 or 6 Ma ago. No trace of a burial process is evident in the glauconitic grains; however, alteration has occurred since traces of goethite have been observed on X-ray diffraction patterns resulting in an abnormally high ferric oxide content. The major portion of the glaucony observed in the sediments dredged off northwestern Spain is relict and related either to the late Miocene regressive period or to the general transgression identified at the beginning of Pliocene time on this margin (Durand, 1974). However, on the one hand, we have observed in situ reworked nummulites and other Late Eocene microfossils of which the chambers were filled with marine green clay suggesting the possibility of an inheritance. On the other hand, the episodic presence of green clay within foraminifera1 tests of Holocene age, on the outer shelf and slope, indicates that glauconitization occurred at different periods.

Mechanism of glauconitization The study of the Spanish sediments has proved that glauconitic clays were initially formed by a crystal growth mechanism and not by a transformation mechanism because 1) in the majority of cases, the green grains were obtained from non-clayey substrates; 2) in cases where an earlier micaceous substrate was utilized (biotite mica flakes), the latter would not have been able to support a transformation mechanism from a simple quantitative point of view; and 3) the same substrate was able to give different authigenic green clays in different environments (mica flakes evolve to verdine off French Guiana -see Chapter B3- and into glaucony off northwestern Spain) whereas the same green clay could be obtained from different mineralogical substrates. Therefore, the study of the Spanish sediments allows us to dissociate the nature and composition. of the substrates from the nature and composition of the authigenic green-clay. This is in conflict with the theory of transformation which postulates a mineralogical link between the inherited material and the authigenic one (Burst, 1958a; Shutov et al., 1970) including for Recent sediments (Bell and Goodell, 1967). This conclusion concerns the first stage of evolution of the substrates. During the second evolutionary stage, the initial potassium-poor marine green clay becomes potassium-rich. It would be possible to invoke a transformation process by potassium absorption for this second stage. However, the morphological and nanostructural features indicate that the initial structure is always destroyed; this whole reorganization on a nanostructural scale can only

273

be understood by assuming a recrystallization process i.e., dissolution of the previously formed clay and use of the available cations to form new microcrystals, richer in potassium. Another point supporting the crystal growth mechanism has already been observed in the glaucony from the Gulf of Guinea and is well illustrated in the verdissement of echinoderm and mollusc skeleton fragments, foraminiferal casts and biotite mica flakes; it concerns the volume increase of the initial substrate and resulting cracks at the surface of the many glauconitic grains.

Environment for glauconi tizat ion The most important feature emphasized by the study of the Spanish glauconies is the role of pores in the genesis of the green clays. A variety of pores are involved: chambers of microfaunal tests (50 p m to 300 pm in diameter), pores of echinodermal stereom or borings in carbonate bioclasts (all are cylindrical pores, 10 pm to 15 pm in diameter), fissures along cleavage planes in mica flakes, or fissures in quartz grains (about 1 pm in thickness). All these pores are initially filled with glauconitic minerals which preserve the initial structure of the substrate. But glauconitic minerals which intimately replace the initial carbonate substrate have also been found. It is suggested that the fundamental process involved is similar in both cases. As far as the replacement is concerned, the "pore" filled is simply a dissolved microcrystal of carbonate immediately replaced by clayey material which, consequently, takes on the form of the earlier carbonate crystal and the morphology of the bioclast. This leads to an apparent orientation of the authigenic clay particles moulding the previous carbonate nanostructure. Therefore, pores constitute the favourable microenvironment for green clay crystal growth. The size of the porous substrates is also an interesting character influencing green clay formation. The most favourable substrate appears to be formed of deposited grains 100 pm to 500 pm in size. As a function of their alterability (ability to form pores), the favourable substrates and resulting green grains will be smaller or larger. Quartz or mica appear more favourable when their size approaches 100 pm,whereas 0.5 to 1 mm large carbonate particles are usually entirely glauconitized. Larger substrates will only be partly glauconitized at the surface. The size of the deposited substrate, the portion of the substrate submitted to glauconitization, and the presence of pores are all aspects of a single phenomenon which can be interpreted in terms of confinement of the microenvironment favourable for green-clays to grow. As was the case for the sediments from the Gulf of Guinea, there are no glauconitic minerals in the free clay-size fraction of the Spanish shelf. If green clay forms in pores of the deposited grains, it is because they find a more favourable microenvironment there, more confined than between sand-sized grains for example. Inside a large boulder, this confinement will be much too closed, and the cationic exchanges will not be rapid enough for the substrate to dissolve or for the green clay to gather the necessary cations for effective crystal growth. On the other hand, sand-sized particles of the sediment allow

274

an easier equilibrium between crystal growth, concomitant cationic feeding, and substrate alteration. A semi-confined microenvironment is, therefore, a prerequisite for the formation and evolution of glauconitic minerals. This notion is fundamental to glaucony genesis; a corollary is that, if glauconitic minerals are always more abundant with a granular habit, it is because the favourable initial substrate was itself in the form of grains. This conclusion differs from that suggested by papers which suppose the first step in glauconitization to be the presence of inherited nontronite (a smectite which resembles glauconitic smectite), and the second the formation of pellets by physical agglomeration or biologic agglutination (Trauth et al., 1969). In short, in glaucony genesis, granular habit precedes the formation of green clay and not vice versa. SUMMARY

The glauconitic sediments from the continental shelf off northwestern Spain allowed various stages of the glauconitization process for various types of substrates to be observed. Verdissemen t mostly occurred within biogenic carbonate fragments which appear very favourable for harbouring green clay growth. Microfaunal tests and mineral debris (biotite mica or quartz) have also been utilized in these Spanish sediments. Verdissement mostly took place at a time near the Miocene-Pliocene boundary. The obtained green grains are essentially authigenic and dominantly result from the evolution of shelly bioclasts. It can be seen that all green clays have formed by a crystal growth process, and the transformation process theory is contradicted by the observations. The mechanism comprises two stages; the preliminary stage is largely favoured by the presence of micropores in the substrates; it consists of filling pores with potassium-poor glauconitic minerals while preserving the substrate. The second stage is conditioned by the dissolution of the substrate remnants and consists of crystal growth of a second generation of green clay followed by a recrystallization process allowing potassium-rich glauconitic minerals to grow partly from the dissolution products of the potassium-poor glauconitic minerals formed earlier. The most abundant and usual product of glauconitization is green grains. This habit results from the fact that granular substrates are initially the most favourable for allowing green clay to crystallize i.e., the green clay genesis follows the granular habit. The pores within the grains delimitate a favourable microenvironment which is characterized by isolation from open sea-water. But pores also facilitate, to a certain degree, cationic exchanges with interstitial fluids. This microenvironment is defined as semi-confined and its influence is well illustrated off northwestern Spain. Apparently, the whole evolution, including the last stage leading to potassium-rich dark-green grains, is possible in close contact with the open sea-water environment without burial diagenesis.

275 ACKNOWLEDGEMENTS

Scanning electron microscope study was undertaken by the authors at the DCpartement de GCologie, UniversitC de Dijon (in 1970-1971), and at the Laboratoire de micropalkontologie, UniversitC Pierre et Marie Curie, Paris (in 1972-1979). The microfaunal tests filled of glauconitic minerals were identified by A. Blondeau and P.A. Dupeuble. Argon isotopes were measured by the editor thanks to the kind welcome and help of J.C. Hunziker in Berne. Chemical analyses, especially for potassium, were carried out by M. Lenoble and G. Richebois. High quality photographic processing by 0. Fay is greatly acknowledged. This chapter was improved for English by Dr. Pennington.

275

This Page Intentionally Left Blank ACKNOWLEDGEMENTS Scanning electron microscope study was undertaken by the authors at the DCpartement de GCologie, UniversitC de Dijon (in 1970-1971), and at the Laboratoire de micropalkontologie, UniversitC Pierre et Marie Curie, Paris (in 1972-1979). The microfaunal tests filled of glauconitic minerals were identified by A. Blondeau and P.A. Dupeuble. Argon isotopes were measured by the editor thanks to the kind welcome and help of J.C. Hunziker in Berne. Chemical analyses, especially for potassium, were carried out by M. Lenoble and G. Richebois. High quality photographic processing by 0. Fay is greatly acknowledged. This chapter was improved for English by Dr. Pennington.

277

Chapter C3 GLAUCONY FROM THE KERGUELEN PLATEAU (southern Indian Ocean) by G.S. Odin and F. Frohlich

PRESENTATION

Glauconies from high latitude deposits Few green grains were quoted and studied in detail (mineralogy, origin, age of formation) from the bottom of the Indian Ocean. Glaucony was described a long time ago by Collet (1908) from off southern Africa (Agulhas Bank) at the western boundary of the Indian Ocean. However, the grains were presumed to be mostly reworked from Cretaceous and Eocene glauconitic limestones (Birch et al., 1976) or to be partly relict and Miocene in age (Odin, 1985b). Collet also indicated the occurrence of green grains on the eastern continental margin off Africa but no detailed study was performed since then. In addition, this author quoted glaucony from off the southern and western coast of Australia. Glaucony was also reported from sites 330 (Leg 36; Islas Malvinas) and 281 (Leg 29; south of Tasmania). Finally, several Dutch authors have quoted the presence of presumed glauconies in some restricted deposits off the Indonesian Islands: south of Sumatra, of Java, and of Timor (Odin, 1969; and Fig. 14, p. 325 in Chapter C4). We do not know of any detailed study dealing with the nature and origin of these green grains. However, Recent glaucony has been identified in the Makassar Strait (Boichard et al., 1985; Fig. 10 in Chapter B2, and unpublished personal results). During the preliminary examination of smear slides from samples collected from the Kerguelen Plateau (Fig. l), Frohlich (1983) and Wicquart (1983) observed green grains which they called glaucony according to the infra-red spectrum of one sample showing the presence of a smectite-type mineral. These seem to be the southernmost records of green grains in the Indian Ocean; they are located very far from any continental shelf. However, glaucony is already known at high latitude in other oceans of the southern Hemisphere. Relict Plio-Quaternary green grains were described by Norris (1964) Cullen (1967) and Seed (1968) from the Pacific Ocean on top of the Chatham Rise (off eastern New Zealand). They are located between -200 m and -2000 m depth, 40" to 45" latitude south. In the Atlantic Ocean, Bell and Goodell (1967) described green grains observed in surface sediments from the Scotia Ridge (around and between the Islas Malvinas and South Georgia, South America). The glauconitic sediments are located between -200 m and -3000 m depth, at about 50" to 55" latitude south and they also occur mainly at the top of the tectonical submarine height.

27 8

\

I /

\/

\

CROZET BASIN \

\

50.

-

AUSTRALIA- ANTARCTIC BASIN

ENDERBY BASIN

Figure 1. Location of the Kerguelen Plateau in the Indian Ocean.

The Kerguelen Plateau

Location, morphology, and geology The Kerguelen Plateau is an aseismic ridge located in the southern Indian Ocean, near Antarctica. This plateau, oriented NO-SE is up to 3500 m high above the adjacent marine basins, about 2000 km long and 500 km wide (Fig. 1). The northern part of the ridge, namely the Kerguelen-Heard Plateau is the shallowest one, less than -700 m deep (Fig. 2). The only islands of the ridge are the Heard Island and the Kerguelen Archipelago. They are of volcanic origin and were built during late Eocene to Quaternary times (Giret, 1983). The northeastern edge of the Plateau displays a rift morphology with normal faulting and associated volcanic reliefs, related to the Kerguelen Plateau-Broken

279 Ridge separation during Late Eocene time (Wicquart and Frohlich, 1986). The Plateau is constituted by a 2000 m thick sedimentary sequence, early Cretaceous (Wicquart, 1983) to Holocene in age, lying on a still unknown acoustic basement. The major seismic feature of this sedimentary basin, uplifted in the Eocene time (Wicquart and Frohlich, 1986), is a transgressionlike pattern of the upper reflectors upon the summital reflector ("acoustic discordance") of the Palaeogene sequence, and towards the northeastern edge. Upper Cretaceous to Middle Eocene sediments are widely cropping out through the prominent scarps which form the northeastern margin of the Plateau (Frohlich et al., 1983).

Sample collection After the first geophysical researchs aboard the R.V. Gallieni and the R.V. Eltanin (Houtz et al., 1977; G o s h , 1981), the main geophysical and geological exploration of the Kerguelen-Heard Plateau was accomplished by the R.V. Marion Dufresne during years 1980 to 1983 (mainly multichannel seismic reflection, and geological sampling with piston cores and dredges). The geological purpose was to perform a stratigraphic sampling of the sequence and collect the oldest sediments. As a consequence, the geological information, concerning the sediment distribution on the present bottom, is restricted to the edge areas. The lithological and stratigraphical description of the cores is given in Frohlich (1983) and Wicquart (1983). These two authors mentioned glaucony occurrence based on smear slide observation. 42 cores have been sampled again for verification of the glaucony occurrence and a mineralogical study of the green grains (Fig. 2). For practical purpose, we have subdivided the Plateau into four sectors: 1) the northeastern shelf, 2) the northern border, 3) the Skiff Bank, 4) the southwestern shelf and the adjacent deep oceanic basin (Enderby Basin). Green magnetic grains were searched for in the soft sediments after washing on a 50 pm sieve using magnetic separation. SEDIMENTOLOGY

The sedimentary cover The sequence is composed of semi-indurated nanno-oozes and hard chalks with cherts and flints of upper Cretaceous to middle Eocene age. Middle Eocene chalks and hard-grounds constitute the eastern scarps. They are locally covered with a phosphatic layer (Frohlich, 1986). Volcanoclastic and sedimentary sands and pebbles overlie Eocene chalks on the flat top of the cliff edge, where the "acoustic discordance'' seems to crop out on the seismic records. Towards the centre of the Plateau, Neogene and Quaternary soft sediments lie upon this discordance, linked with emersion followed by aerial erosion of the ridge during the Late Palaeogene (Wicquart and Frohlich,l986). This young sequence, with a thickness ranging from 900 m in the centre to 0 m towards the northeastern border, is dominantly diatomaceous. The thickness of the Quaternary sediments (ice-rafted diatom oozes), is unknown: only the top (several

280 metres) of the sequence was cored in these areas. Quaternary deposits are missing in the eastern and northern edges of the Plateau. In the area near the Kerguelen Archipelago, the hard seismic reflector at the interface corresponds to a rhythmic sedimentation, dominantly volcanoclastic and diatomaceous, with an important ice-rafted, coarse fraction (sands and pebbles) and a typical polar low rate of sedimentation. In the western (Skiff Bank) and northern areas, bottom sediments (Miocene to Holocene, and locally Late Eocene in age), are mainly diatom-rich nanno-oozes deposited with a presumably high rate of accumulation (50 m/Ma) and without any coarse ice-rafted detritus. Such a sedimentary contrast (low rate versus high rate and diatomaceous versus calcareous deposits) is a consequence of the opposition between antarctic and sub-antarctic waters and is related to the outstanding salinity/ temperature discontinuity of the Polar Front, presently located near the Kerguelen Archipelago latitude. The Kerguelen Plateau strongly disturbs this hydrologic system and is an obstacle to the circulation of the Circum Antarctic Current. As a result, the sedimentation is different from the west to the east; eastwards, turbulences and upwellings result in the absence of significant Quaternary sedimentation. The probably very thin Quaternary deposits of the eastern area are enriched in ice-rafted and volcanoclastic components and in authigenic green grains.

Glaucony Green grains have been observed in a number of smear slides prepared from surface sediments in the different areas of the Plateau (Frohlich, 1983; Wicquart, 1983). A few greenish grains from the southwestern shelf and its related basin, encountered in biozone NN 21, probably represent chloritized detrital debris. Six cores collected from the southern margin of the Skiff Bank (Fig. 2) have shown soft sediments corresponding to NN 21. The smear slides show a predominance of microfaunal tests (radiolarians, foraminifers, and some ostracodes). In two cores, a small proportion of olive-green to yellowishochre grains were found in some restricted layers, one to five metres below the sea-watedsediment interface. The yellowish-ochre grains represent the beginning of a glauconitization process; they are very indented, excluding a reworking from the underlying sediments, but they possibly fell down from the top of the Bank onto their present depth, at - 1500 m and -3300 m. From the northern border of the Plateau, 500 km N-NW off Kerguelen, six cores were studied. The soft sediments are regarded as Pleistocene in age (equivalent to the NN 19 and NN 21 biozones). A small number of yellow-ochre grains were magnetically separated from the uppermost metres of sediment recovered. The grains are comparatively small (less than 100 pm) and may be perigenic. The northeastern shelf of the Plateau was sampled along a distance of 500 km. Glaucony was recognized from the soft sediments either at the surface or somewhat below in numerous cores presently located between -600 m and

28 1

Figure 2. The Kerguelen-Heard Plateau. Astenscs show the location of the studied cores. The dotted area indicates the presumed location of glaucony formation, the extent of which still remains to be investigated especially towards low depth.

-2500 m depth. Deeper, green grains are absent while, shallower, no cores were available. Biostratigraphic dating is difficult; however, in one of the deepest cores considered (sea-bottom at -2544 m) glaucony was identified in a subsurface layer bracketed between sediments biostratigraphically correlated, using radiolarians, with the NN 21 biozone. The green grains at that depth were probably carried downslope from the Plateau at a time close to their genesis. The richest glauconitic formation is found at the top of the shelf between -600 m and -800 m. To summarize, in three of the four sampled areas, the deep-sea sediments show few or no green grains. A reworking from older layers altered at the sea-bottom is unlikely; a short transport from shallower depth is possible on the slope. In the northeastern area and especially on the edge of the Plateau, the facies is very characteristic of an in situ glauconitization process occurring within a low accumulation rate sediment.

282 NATURE AND ORIGIN OF THE GREEN GRAINS

Detailed data on the northeastern shelf sediments On the northeastern shelf, samples were studied from three groups of cores: 1) 6 sites in the northern area, at the Kerguelen Archipelago latitude; 2) 13 sites in the central area; and 3) 5 sites more to the south (Fig. 2). In the northern area, 150 km east of Kerguelen, the cores collected on the border of the Plateau show glauconitized particles at their top (several decimetres) as well as deeper (about 5 m below the sea-water/sediment interface) which probably indicates several episodes of glauconitization. The layers at the top are the richest (about 8% of green grains) whereas the deeper layers contain no more than 2%. The green grains at the surface show all intermediate stages of morphological evolution from pumice (0.2 mm to 1 mm in size) to green grains entirely made of green material. However, the evolution was never completed since no dark-green grains were observed there; the more or less glauconitized pumice grains are frequently very indented and delicate which points to an in situ origin. These pumice grains constitute the main substrate for glauconitization in the sediment. In contrast to the green grains observed at the top of the cores, there are comparatively few intermediate forms between pumice and green grains in the sediments collected from deeper parts of the cores. Moreover, the green grains are darker in these deep sediments compared to the most evolved grains at the surface of the same cores. Thirteen sites were studied at the centre of the northeastern shelf; in seven cores, taken from the shallowest sites, green grains were encountered at the top of the cores, whereas from the deepest sites, they were found below the sediment-surface. Three cores collected from the Plateau itself at a depth between -650 m and -750 m are made of a sediment with about 50% of slightly-evolved to evolved green grains. The cracks of the dark grains are filled with green material somewhat lighter in colour. All grains here appear more evolved than in the northern area. This central area has been more favourable than elsewhere for the glauconitization process. 1 to 2 m below the sediment-surface, it is possible to observe intermediate evolutionary stages between spheroidal radiolarian infillings and free green grains. Finally, it can be noted that the dark-green grains are comparatively quite dense and hard to crush in the agate mortar, a factor generally met for relict green grains (i.e., formed long ago and never buried). The core collected from -1030 m depth also contains abundant green grains: 7% to 8% of the bulk sediment at levels -0.3 m, -0.8 m and -2.3 m below the sea-waterhediment interface and 25% at -4.9 m. The grains are deeply cracked, olive-green at -0.3 m, grey-green to ochre at -0.8 m, green to olive-green at -2.3 m and stronger green at -4.9 m depth. The colour is never dark-green in contrast to what happens on the Plateau. The main substrates were probably pumice grains. One may note that the volcanic debris are less abundant in this core and that agglomerates of pyrite are present at -4.9 m in the core.

283 Therefore, there is some sort of gradient of oxidation from the top of the core, where the green grains may be interpreted as oxidized (ochre colour), to the bottom where reducing conditions (pyrite) have preserved the original colour of the green grains. The deeper sediments of this central area where green grains have been observed, were cored at -2580 m depth. The sediment is mainly composed of more or less magnetic, volcanic debris, from which it is difficult to concentrate green grains. These green grains were present in minute proportion at -0.15 m, -3.8 m, and -4.2 m depth in the core. Their colour varies from very dark-green to yellow-green at the surface. At deeper levels, the grains are smooth and frequently reddish. This may be considered as the track of an oxidation and of a long transport. The green grain-bearing sediments were tentatively correlated with NN 21 biozone which overlies Pliocene deposits. The glauconitization cannot be considered in situ in these sediments; green grains are perigenic.

Substrates of verdissement Amongst the four possible kinds of substrates usually sheltering the verdissement process (Odin and Matter, 1981), two are virtually absent here: faecal pellets and carbonate bioclasts. The two others: mineral debris and microfaunal test infillings are both siliceous.

The verdissement of pumice debris (Fig. 3 and 4) Pumice debris were found in several areas of the Plateau. In the northern portion of the northeastern shelf, cores obtained from -650 m depth show a number of pumice debris 0.3 to 3.0 mm in size (Photo.1 in Fig. 3). All photomicrographs in this paragraph are from this core but similar habits have been observed elsewhere. At a higher magnification under the binocular, broken debris display small green patches inside; sometimes the whole fragment is green. Scanning electron microscope observation indicates that these green pumice debris are made of two components (Photo.2 in Fig. 3). One appears as a fibrous structure (Photo.3 in Fig. 3) similar to the unmodified pumice. The material is siliceous, but not crystallized according to X-ray diffraction analyses; it may include well-shaped and well-preserved volcanic crystals. The second component grows within this very light and porous structure (Photo.3 in Fig. 3) filling sometimes large cavities (Photo.4, Fig. 3). An enlarged view of this material, green under the binocular microscope, shows the nanostructure of the green clays (Photo.5 in Fig. 3) including a honeycomb pattern with sand rosette-like structures in larger micro-geodes. The clay mineral particles are smaller than two micrometres in length (Fig. 3, Photo.6). They are sometimes very irregularly arranged (Photo.2 in Fig. 4); the latter photomicrograph also illustrates the relationship between the siliceous substrate and the clay developed within it. On the one hand, we have shown (Photo. 1 in Fig. 3) clay minerals beginning to grow on the siliceous material on which they are simply attached; on the other hand, photomicrograph 2 (Fig. 4) shows that the smooth siliceous substrate is probably dissolved at its contact

284

Figure 3. Glauconitized pumice debris, first stage; scanning electron microscopy. White bars represent 100 pm (Photo. 1); 10 pm (Photo. 2,3,4,5) or 1 pm (Photo. 6).

285

Figure 4. Glauconitized pumice debris, late stage; scanning electron microscopy. White bars represent 100 pm (Photo.l,3) or 1 pm (Photo.2.4).

with the growing green clay mineral crystals. Therefore, the clayey material grows de novo (i.e., from ions) in the pores of the substrate; the assumption that ions from the substrate itself are utilized for growth of the authigenic phase cannot be excluded. In the darker green grains suspected to be originally pieces of pumice, the fibrous structure disappears (cf. Photo.1 in Fig. 4 -no more fibrous structureand Photo.1 in Fig. 3). At higher magnification, the clayey material appears nearly alone: the substrate has been dissolved (Photo.3 in Fig. 4:enlarged view of Photo.1 immediately above). The dissolution creates new pores in which new clay minerals grow with a better shape but with still a small size (Photo.4 in Fig. 4).

The verdissement of siliceous tests (Fig. 5 and 6 ) As explained above, the biogenic fraction of the surface sediments is siliceous being made of radiolarians, diatoms, or siliceous sponge remains. We

286

Figure 5. Glauconitized discoidal radiolarians; scanning electron microscopy; white bars represent 100 pm (Photo.l,3); 10 pm (Photo.2); or 1 pm (Photo.4).

have chosen sediments from the central zone of the northeastern shelf to illustrate the verdissement of these biogenic remains. All the pictured remains come from the border of the Plateau, between -700 m and -800 m depth. Amongst the radiolarians submitted to the verdissement, some are flat disks which have preserved their original form (Photo.1 in Fig. 5). However, an enlarged view shows that clay minerals are present between the meshes of the reticulated test which is in process of disaggregation (Photo.2 in Fig. 5). A similar origin may easily be attributed to a series of green elements of the sediment; they are more or less splitted with fissures and cracks (Photo.3). Inside the cracks, the characteristic pattern of the clay minerals may be observed (Photo.4 in Fig. 5). Other radiolarian tests are spherical (Photo.1 in Fig. 6). Often, they became substrates for the verdissement and numerous spherical green grains of a size

287

Figure 6. Gauconitized siliceous biogenic remains; scanning electron microscopy; white bars represent 100 pm (Photo. 1,2,3,5,6); 10 pm (Photo. 7); or 1 pm (Photo. 4).

288

Figure 7. Dark-green grains; scanning electron microscopy. White bars represent 100 pm (Photo.1) or 1 pm (Photo.2,3,4).

similar to that of the radiolarians were magnetically separated. Sometimes, the rounded green grains exhibit a nucleus and an external layer (Photo. 1 in Fig. 8) like the tests of the radiolarians which may comprise several spheres of silica one within the other. During its evolution, the green sphere develops cracks (Photo.2 and 5 in Fig. 6, both from dark-green grains), which, in a final step, are filled with a new generation of clays which are usually light-green in colour (Photo.3 in Fig. 6). The clay mineral structure (Photo.4 in Fig. 6) may be observed within the cracks of the grain shown on Photo.2 (Fig. 6). Finally, a spectacular aspect of the verdissement process is the presence of green material inside siliceous, transparent sponge spicules as shown in Photo.6. The spicule itself forms a tube which has been partly peeled here in order to show the filling. The limit between the tube with its conchoidal break, and the green clay infilling, is shown in Photo.7. The typical nanostructure of

289

Figure 8. Coccoliths in the cracks of the green grains; scanning electron microscopy. White bars represent 100 pm (Photo.l), 10 pm (Photo.3,4), or 1 pm (Photo 2).

the clay mineral can be observed at the surface of the infilling. The above observations were undertaken on grains for which the initial substrate was recognizable. In other words, the verdissement process still preserved traces of the initial substrate: pieces of the substrate itself or the general morphology of the grain which remind the probable substrate. However, in the most evolved sediments it is very difficult to imagine the initial substrate of a number of dark-green grains as documented below.

Electron microscopy of evolved grains (Fig. 7 and 8) In the samples where the siliceous biogenic debris were studied, the sediment contains sometimes abundant dark-green grains, very hard to crush,

290 generally with a smooth surface, and without any characteristic structure. Some of these dark-green grains were broken in order to observe their internal structure. The grain pictured in Figure 7 (Photo.1) does not show any special texture, being very homogeneous and compact (Photo.2 in Fig. 7). At high magnification the green clay shows a dense arrangement of small thick sheets (Photo.3 in Fig. 7). This structure is very common and homogeneous. However, honey comb pattern has been observed at the surface or inside grains which appear less dark, (Photo.4 in Fig. 7). It is worth mentioning that, very often, the cracks, fissures or other narrow invaginations in the grains were filled or covered with minute biogenic remains efficiently trapped and preserved there. For example, Photo.7 (in Fig. 6 above) shows a number of biogenic remains, including coccoliths located between the spicule and the green filling as well as in the fissure of the tubular spicule. Similarly, between the glauconitic nucleus and the glauconitic envelope of the radiolarian filling (Photo. 1 in Fig. 8), many coccoliths and diatom remains are trapped (Photo.2 in Fig. 8). The surface of the spherical green grains is frequently divided by cracks (Photo.3 in Fig. 8) in which calcareous nannofossil and diatom tests are trapped (Photo.4 in Fig. 8). The observed DARK-GREEN GLAUCONY

KERGUELEN

G702-2A

020

d

003

ki

490'C

for I hour

Figure 9. X-ray diffraction patterns of dark-grwn grains of the Kerguelen Plateau. The 001 diffraction peak is at about 10.8 8, on the natural randomly oriented powder sample and reaches 10 8, after heating. Note the poorly shaped hkl peaks except for the 020 one.

29 1 nannofossils are calcitic and were previously not described from the sediment. Since the modem sea-water above the Kerguelen Plateau is not favourable to the life of coccoliths this observation will have to be explained.

Mineralogy of the green grains The mineralogical study of the green grains of the Kerguelen Plateau has been restricted to the general recognition of the nature of the green clays first analysed using infra-red spectra. The latter technique showed the absorption bands characteristic of a disordered Fe-Mg dioctahedral mica which confirms that the green grains present on the Kerguelen Plateau constitute a newly known glaucony deposit lying on the present-day sea-bottom. The X-ray diffraction analysis indicates that the dark-green grains (the morphologically most evolved rains collected from the Plateau) display a strong 001 peak between 10.5 and 11 A which shifts to 10 A after heating for four hours to 490°C (Fig. 9). This corresponds to about 6.0% K,O (Odin and Dodson, 1982) and is less than could have been presumed from the dark colour and the induration of the grains but confirms the indication of the nanostructure of the grains which never displays the large sheets of a highly-evolved glauconitic mineral. An interesting indication is also given by the remarkably low size of the 112- and 112 peaks comparatively to the 001 peak, which indicates an absence of ordering of the crystallographical structure. This, in turn, may be an indication for the absence of burial of these grains (see Fig. 9, p. 313).

f

DISCUSSION

History of the glaucony from the Kerguelen Plateau Glaucony is a significant component of the surface sediments deposited on the northeastern border of the Kerguelen Plateau in an area where the general sedimentological observations indicate condensed series. The age estimate of the glauconitization was one of the questions to be solved. Green grains have been found mixed with siliceous fossils considered to be contemporaneous with late Pleistocene NN 21 biozone. This correlation was possible in the deeper sediments, where the series is more complete than on the Plateau. On the Plateau, the reworking from older series is unlikely considering a) the long distance from any continent and the strictly volcanic nature of the emerged islands; b) the presence of very brittle, glauconitized substrates which would have been destroyed by reworking processes; c) the coincidence between the substrates of glauconitization and the components of the sediment assumed to have been recently deposited. Moreover, the X-ray diffraction patterns exhibit characters similar to those of green grains which have never been significantly buried, e.g., southern Africa submarine continental margin (Odin, 1985b), northwestern Spain margin (Lamboy and Odin, 1975), French Guiana margin (Pujos and Odin, 1986). The glaucony is therefore authigenic and Quaternary in age on the Plateau. Deeper than -1000 m, glaucony can be

292 suspected to be perigenic. On the other hand, the glauconitization process is not really Recent nor contemporaneous. This can be assumed because many samples show traces of an oxidizing alteration of the grains whatever they appear evolved or nascent. This indicates that the environment was modified (becoming too oxidizing) following a period during which the general environment was favourable for glauconitization. The present environment is presumably such a too oxidizing one. Therefore glaucony is relict and reflects sedimentological conditions which preceded the present ones. A second criterion is the fact that the grains are relatively indurated and locally display some sort of glazing, two characters which were also observed for relict grains of the Atlantic continental shelf. We suggest that the duration of the favourable conditions was not very long (less than about 0.5 Ma) since the evolution of the most evolved grains was not completed i.e., the glauconitic minerals have not reached the micaceous, potassium rich, end-member of the family (see Chapter C4). A good idea of the environment which preceded the present one and was favourable for the glauconitization process may be obtained from the biogenic remains preserved in the cracks of the green grains. Coccoliths are numerous in these cracks; at present, living organisms producing these tests are nearly absent in the marine environment of the Kerguelen Plateau; they only develop at lower latitude. The temperature of the sea-water where they grew was higher than the one present today on the Kerguelen Plateau. Therefore, the glaucony found today on top of the Plateau was probably formed at a time when the sea-water was warmer than today; more precisely, between the time of glaucony genesis and the present one, fossils characteristic of comparatively hot temperatures have been trapped within the cracks of the already formed dark-green grains. These conditions occurred during a relatively short time, too short to allow a complete evolution of the glauconitic minerals; the comparatively warm sea-water was replaced later by colder waters. The environmental modifications are also known from the evolution of the planktonic assemblages (Denis-Clocchiatti, 1982); they are probably related to a shifting of the Polar Front southward from its present position during Plio-Pleistocene time (Wicquart and Frohlich, 1986).

Factors of the glauconitization process On the Kerguelen Plateau, the glauconitization occured in biogenic or volcanogenic siliceous substrates. This is an original situation compared to the occurrences reported in the literature (see Chapter C2) where dominantly calcareous bioclasts are involved. No layer silicates with degraded structure (illite or smectite) are involved in the genesis mechanism of the glauconitic clay. Accordingly, it appears that the theory first developed by Burst (1958a) and followed by a number of authors since that time i.e., a transformation of a prerequisite degraded layer silicate, is not adequate to describe the process of glauconitization on the Kerguelen Plateau. Instead, we have observed glauconitic minerals growing inside the voids of granular siliceous substrates. When

293 the pores are entirely filled, the substrate itself is dissolved and disappears allowing new green marine clays to crystallize. Narrow fissures or larger cracks develop on the green grains formed at that stage. During the final stage, the cracked grains may become smooth mainly because of the filling of the cracks by lately formed green clay minerals. All of these glauconitic minerals result from de novo crystal growths. The role of the substrates appears to be mostly physical. In fact, radiolarian tests and pumice debris are very similar from a physical point of view. They bear a lot of pores; for the pumice debris, the section of the elongated pores is about 2-5 pm wide whereas the subspherical voids inside the radiolarians are 5-10 pm in diameter. The marine clays grow only inside these substrates which demonstrates that they need to be sheltered. Moreover, the assumption that the pores play a significant role is supported by the fact that no green compact basalt debris nor volcanogenic minerals of a size similar to the glauconitized substrates, have been observed during our investigation although they occur in the sediment. We presume that this is because they were not porous. Some kind of relationship between the growing of glauconitic clay and the dissolution of the siliceous material of the pumice has also been observed. This can be interpreted as a possible utilization of silicic ions provided by the substrate for the crystal growth. Amongst the other ions of the glauconitic minerals, iron has an important role, it is likeky to originate from the alteration of volcanic debris. The abundance of volcanic debris appears as a favourable factor for glauconitization. Elsewhere in the Indian Ocean, the deep-sea environment allows neoformations of amorphous iron silicates or iron-rich smectite (Frohlich, 1980, 1981, 1982). In the Kerguelen-Heard area the glaucony formation observed at shallower depths is probably related to a "deepening" of the oceanic oxygen minimum level. During its tectonic history (uplift events and subsidence), the Kerguelen-Heard Plateau was probably on several occasions at the level of the oxygen minimum depth (Frohlich, 1986). Concerning the general conditions of glauconitization, the above study indicates that the process may occur between -500 m and -1000 m depth provided that the present depth of the Plateau has not been substantially modified since Pleistocene time. This depth is far greater than the one generally favourable for glauconitization process on the stable continental margins (100 m to 300 m depth). The morphology of the sea-bottom may play a larger role than the absolute depth to allow glaucony to form; in other words, the fact that we are either on the fop of a submarine high topographically similar to the Chatham Rise in the Pacific Ocean (Cullen, 1967) or the Scotia Ridge in the Atlantic Ocean (Bell and Goodell, 1967) for active margins, or on the top of the continental slope for stable margins could be the determinating factor. The presence of glaucony at a relatively high latitude does not mean that the present environment is favourable for glauconitization. Like on the Chatham Rise, the glaucony of the Kerguelen Plateau, known to be relict (Plio-Quaternary), indicates a marine environment different from the present one at the same latitude.

294 SUMMARY

Glaucony is present in the sedimentary cover of the Kerguelen-Heard Plateau and especially abundant on its northeastern margin. Its distribution is not known for the shallow portion of the Plateau. The facies is relict and may be assigned a Pleistocene age. Glaucony only develops as grains originating from siliceous, highly-porous substrates of two main types i.e., mineral pumice debris and biogenic siliceous tests. Glauconitic green clay minerals develop f i s t within the pores of these two physically equivalent substrates; after dissolution, the substrate itself is replaced by the same authigenic material and the initial structure of the grain may disappear. The final granular habit of glaucony is therefore due to the initial granular habit of the favourable substrates. The genesis mechanism of the glauconitic grains of the Kerguelen deposit only implies crystal growth utilizing ions from the immediate environment. Si and Fe probably originate from the volcanoclastics, abundant in the vicinity. The glauconitization probably occurred between -500 m and -1000 m depth on this oceanic height. Hydrographic conditions varied widely during Pleistocene time but it is suggested that glaucony formed during a period when sea-water was generally warmer than today above the Plateau. In spite of this limitation, the high latitude location of the deposit indicates that a temperature much cooler than the one usual for the verdine facies allows glauconitization to proceed. ACKNOWLEDGEMENTS

This chapter is illustrated with pictures obtained using facilities made available in 1986-1987 at the Department of Geology, Museum National d'Histoire Naturelle, Paris. In the course of this study we have been able to analyse samples from the Makassar Strait (south Borneo) provided by J. Gayet and J.C. Faugkres UniversitC de Bordeaux. We thank Aline Ehrlich from the Geological Survey of Israel, for her discussion and improvement of the English language of that chapter. Really good photogaphic processing by 0. Fay, DCpartement de Stratigraphie, UniversitC P. et M. Curie, Paris, is greatly acknowledged.

295

Chapter C4 GEOLOGICAL SIGNIFICANCE OF THE GLAUCONY FACES by G.S. Odin and P.D. Fullagar INTRODUCTION

The information presented in Chapters C1, C2, and C3 allows us to reach certain conclusions regarding the significance of the sedimentary glaucony facies. The use of a specific facies name, glaucony, is a condition for an adequate understanding of the subject. This name was first suggested by Odin and Utolle (1980) and more precisely defined by Odin and Matter (1981) in order to fill a need in the English terminology. It allows an appropriate distinction between the facies 'glauconite' (= green grains) which is generally used after superficial observation of the grains' habit, etc., and the mineral 'glauconite' (= glauconitic mica) which is one of the specific components of the authigenic phase of the green grains (Table 1). Distinction is commonly made in other languages (e.g., French, Spanish) and should clearly indicate the type of material actually considered; in some cases, such as the use of glaucony for radiometric dating, the interpretation of the analytical results depends on the type of "glauconite" which was analysed. To avoid further confusion, it is recommended that the term "glauconite" be discontinued, and instead the term glaucony be used for the facies and the term glauconitic mineral be used for the minerals (Table 1). Table 1. Nomenclature for the glaucony facies and its components.

French ter mino 1ogy (old) Facies (green grains) l a glauconite Minerals (components) l a glauconite Mica end member l a glauconite '

Eng 1ish terminology (old) (suggested new 1 gl auconite glauconite glauconite

g laucony glauconitic minerals glauconitic mica

The glaucony facies will be described according to habit, composition, process of formation and environment of formation, the four characters defining a geological facies. As has been done elsewhere in this volume, this chapter will focus on conditions prevailing at the time of deposition (i.e., geochemical processes occurring when sediments are in contact with sea-water). Fortunately, there is abundant glaucony at the present-day sea-bottom. We will preferentially consider observations and analyses from these Quaternary and, when possible, Holocene deposits. This means that we may disregard observa-

296 tions made on Palaeozoic, deeply buried (i.e., more than 1000 m), tectonized, or lithified glauconitic sediments. The same approach was chosen in our preceding syntheses (Odin, 1975a; Odin and Matter, 1981; Odin 1985a) and contrasts with procedures followed by some colleagues (McRae, 1972; Buckley et al., 1978) whose aim was different. We wish to describe the glaucony facies as it is at the time of formation, rather than after it may have undergone different types and episodes of alteration. Under these circumstances, some of the variable mineralogical nature of 'glauconite' (e.g., Ireland et al., 1983) is not relevant for our purpose as it probably results from local diagenetic modifications. Some of the chemical variations observed in Palaeozoic sediments have been interpreted as resulting from differences of initial environmental factors characterizing the original sea (Berg-Madsen, 1983). According to what we now know, most of these mineralogical or chemical variations are diagenetical changes rather than initial variations. This view is based on detailed studies of many Recent deposits, comparisons with ancient sediments of various Phanerozoic ages, and the fact that the widest variatiom described in the literature are demonstrably linked with Palaeozoic or deeply buried sediments of complex or very long geochemical history. It is also based on the fact, documented below, that glaucony results from interaction between oper Lza-water and deposited sediments; these two phases have had a relatively constmt physico-chemical nature during the portion of the Earth's history with whish we are concerned. Glaucony could have formed durkg early phases of sediment deposition (Odin, 1986) and is known from 1 o 2 Ga old sediments in the U.S.S.R. (Polevaya et al., 1961), Australia ('3ebb et al., 1963), and China. But here below, we consider glauconies younger than about 100 Ma. HABITS OF GLAUCONY

Classification of habits The morphological diversity of glaucony is great and was recognized by early workers who tried to classify this heterogeneity. Cayeux (1916), Millot (1964), and Triplehorn (1966, 1967) published the most widely quoted proposals. A combination of the proposal by Cayeux (1916), and revision by Millot (1964), is shown in Table 2 along with the present suggestion. These earlier classifications are strictly descriptive and could be expanded following the many observations available today. We prefer to suggest an interpretative genetic classification which is now possible thanks to the observations made in the three case studies discussed earlier. Four main factors influence the habit of glaucony. They are 1) the presence of a prerequisite material which has undergone verdissement; this material is here termed the substrate; 2) the presence, origin and size of the pores in which forms the authigenic phase characterizing the facies; 3) the size of the substrate and 4) the stage of evolution of the glauconitized material. Two main categories of habit can be distinguished: the better known granular habit, plus the film habit.

297 Table 2. Variety of glaucony habits as a function of the original substrate. According to Cayeux (1916) revised by Millot (1964)

Present classification

1.

Remplissage d e microfossiles (infillings of microfossils)

2.

Epigknie d e grains d e boue et coprolites (replacement of mud grains or faecal pellets)

1.1 Internal moulds

3.

Epigknie d e tests calciques (replacement of carbonate shells)

1.2 Faecal pellets

4.

Pseudomorphose d e spicules d'Eponge (pseudomorphism of sponge spicules)

1.3 Bioclasts

5.

Enduits sur rninkraux et kpigknie (coating and replacement of minerals)

1.4 Mineral and rock grains

6.

Transformation d e paillettes d e biotite (transformation of biotite flakes)

1.5 Green grains without recognizable origin

7.

Globules d e remplacement d'opale (replacement of opal globules)

8.

Grains v e r b quelconques (common green grains)

2.1 Large clasts

9.

Enduits et filons dans phosphate (coating and fissure filling in phosphate)

2.2 Hardground

10. Pigment diffus (diffuse pigment)

1. GRANULAR HABlT

2. FILM HABlT

2.3 Diffuse habit

Our classification intends to be interpretative rather than descriptive in that it mes to recognize and interpret the green material as a function of its origin. For example, it includes only one category for small or large microfossil infillings or replacements observed within foraminifer or radiolarian chambers, grains resulting from the disaggregation of these tests, as well rounded or cracked grains resulting from the evolution of these infillings or replacements. Similarly, categories 5,6, and 7 of Millot (Table 2) can be placed in our category 1.4, which also includes many other glauconitized non-biogenic particles. The sedimentological significance of the glaucony-bearing sample then can be considered as in section p. 327.

The granular habits Four main groups of granular habits have been proposed by Odin (1975a), and Odin and Matter (1981): Internal moulds are probably the most pervasive habits and were noted by many authors who observed calcareous microfossil tests filled with green clay (Murray and Renard, 1891; Collet, 1908; C a s p a , 1910; Cayeux, 1916; Wermund, 1961; Ehlmann et al., 1963; Bjerkli and Ostmo-Saeter, 1973 and many others). These substrates were predominant on the deepest part of the present continental shelf off Morocco, Norway, southeastern United States, and Makassar Strait, Borneo (Boichard et al., 1985; and unpublished personal observation). In these sediments, foraminifers are the most common fossils but ostracodes, plus small molluscs or bryozoans are also present. Conversely,

298

Figure 1. Granular habit of glaucony. 1) Sponge spicules replaced (photo. courtesy A. Matter); 2) thin section showing a green infilling of Miliolid (darker), Lutetian at Paris (France); 3) thin section of a light-green replaced test of Alveolinid, same sample; 4) darkgreen highly-evolved grains of the Bartonian at Cassel, France; 5) common dark-green, cracked grains, Cuisian, Paris Basin, the initial substrate is presumably carbonate bioclast. (Photo. 2 to 5 after Odin, 1967). Scale bars =1 mm (1,4,5) or 0.1 mm (2, 3).

299 internal moulds of siliceous microfossil tests seem to characterize high latitudes and tectonic submarine highs (Morton et al., 1984; Odin and Frohlich, Chapter C3). The chambers of these tests define comparatively large voids, and the question arises as to whether or not the green authigenic clay directly fills these voids or replaces a clay-size material which previously filled the chambers. Both infilling and replacement processes are observed (Fig. 1). Faecal pellets form the predominant substrate for glauconitization in many Recent sediments (Takahashi and Yagi, 1929; Moore, 1939; Bell and Goodell, 1967; Porrenga, 1967a; Tooms et al., 1970; Chapter C1 in this volume). Mud grains have sometimes been noted (Millot, 1964; Bornhold and Giresse, 1985). It is not clear how these mud grains could have formed by a strictly physical process. Our opinion is that these ellipsoidal clay pellets in which the green clay forms are all biogenic. Pryor (1975) has reported that considerable amounts of these faecal pellets can be produced by mud-eaters or filter feeding organisms usually living in the shallowest part of the continental shelf but sometimes at more than -100 m depth (Moore, 1939). Within these pellets, the verdissement seems to occur by filling the diffuse pores of the substrate with green clays. These pores result from the dissolution of previously aggregated material. Carbonate or silicate bioclasts constitute our third category of substrate. Although cited by a few authors (Dangeard, 1928; Cayeux, 1932; Houbolt, 1957; Millot, 1964), these substrates were long considered of little importance in the glauconitization process. Their notable significance was first pointed out by Lamboy (1968) off northwestern Spain. This sort of substrate is now known to be predominant in a number of glaucony deposits (Lamboy, 1976; Odin, 1969; 1975a; and Chapter C2 in this volume). These bioclasts appear to be glauconitized in two different ways which may occur in a single grain. On the one hand, the green authigenic clay may fill voids: either natural pores of the bioclasts or pores created by biogenic or chemical alteration. This filling process is generally similar to that observed in microfaunal chambers. On the other hand, the green clay may replace the carbonate or silicate material sometimes preserving the initial structure at the scale of microcrystals (e.g., zebra-structure of glauconitized bivalve shells). This replacement process is similar to pseudomorphism. Mineral and rock grains constitute another important group of granular substrates. Regardless of their chemical composition, a wide range of grains is susceptible to glauconitization. Examples have been given by Cayeux (1916); Galliher (1935); Wermund (1961); Ojakangas and Keller (1964); Odin (1972a); and Hein et al. (1974), among others. Glauconitized quartz, silica, chert, feldspar, mica (black or white), calcite, dolomite, phosphate, volcanic glass, and diverse rock fragments have all been observed. However, there are few occurrences of green grains dominantly formed from mineral grains. The case study of the Quaternary Kerguelen deposit (Chapter C3 in this volume) illustrates this point; this deposit resembles the Miocene Rockall Plateau sediments with both dominant volcanoclasts and some glauconitized siliceous bioclasts (Morton et al., 1984).

300 The verdissement of these mineral substrates usually occurs along fissures or cleavage planes. These planes are a peculiar form of pores which are filled and have significance similar to that of the voids found in pumice substrates. In a second stage, however, a replacement process may occur, green clays seem to invade the space between the previously filled pores. Most categories of grains identified by previous authors can be placed within the four genetical groups described above. But there is a practical difficulty because the substrates in which the glauconitic clay minerals form finally reach a stage of evolution where it is no longer easy to identify the original substrate. A fifth common category: grains without recognizable origin (Fig. 1, Photo. 4) will therefore be occupied mostly by grains representing a late stage of evolution of the substrate rather than a particular nature of the substrate.

The film habits Two varieties of thin layer of green clay have been observed in nature. A green film is often present on large carbonate bioclasts such as mollusc shells, rock boulders such as in the Miocene molasse, or various synsedimentary nodules such as phosphate in Quaternary sediments or flint in Cretaceous chalk. A thin section study of these films shows that the green clay mainly develops along fissures, biogenic borings, or any other void resulting from physico-chemical or biogenic alteration. The thickness of the green film is related to the extent of alteration (e.g., typically a few millimetres for siliceous or silicate boulders, and up to 1 cm or more for more reactive constituents, especially carbonate. A second variety of the film habit of the glaucony facies occurs as extensive surfaces, often more than one km2in area (Fig. 2 top). These surfaces are usually present in carbonate sediments such as Jurassic marine limestones or Late Cretaceous chalk. They have characteristics of hardgrounds, indurated at the top and preferentially bored by large organisms. The green colour which extends along the hard surface is due to a clay similar to that in green grains (Aubry and Odin, 1973); the facies has been used as valid regional lithologic marker to trace a specific time of deposition in a particular basin (Gosselet, 1901; Juignet, 1974; Odin, Renard et al., 1982b). The two varieties of the film facies (boulders or hardgrounds) are really equivalent and differ only in size. There is possibly a third variety of glauconitization which has sometimes been called diffuse habit (Table 2). It can be regarded as a vertical extension of the film habit. The glauconitic clay would develop in decimetre or metre thick layers giving a green colour to the entire formation. However, it is generally not possible, in sediments where inherited clay can be present, to efficiently separate the green clay presumed to be authigenic from the rest of the clay-size fraction. Therefore, the glauconitic facies cannot be identified with certainty since the mineralogical characteristics of the green pigment cannot be defined, a precondition for identification of the facies. In general, a good indication that the diffuse green clay or film is actually related to glauconitization would be the superposition with a typical granular facies (Fig. 2).

301

Figure 2. Film habit and granular habit in the field. 1) glauconitized bored hardground in the chalk of Maastrichtian age at Hallembaye, Limburg, The Netherlands; 2) thinly glauconitized hardground with glauconitized,oxidized, and phosphatized grains above; chalk of Maastrichtian age at Maastricht, Limburg, The Netherlands.

Discussion of the habits of glaucony From the variety of morphological habits described in the literature we have proposed eight groups (five granular and three film habits, Table 2). However, these groups merely represent diverse expressions of a unique phenomenon (Odin, 1975a). The common feature for all groups is a prerequisite substrate within which authigenic clay genesis is sheltered. The apparent variety is mostly in the nature, size, and stage of evolution of the substrate. The variety of the mineralogical nature of the substrates is now well established. In addition, similar substrates may host either glaucony (or verdine) genesis. Therefore, any theory about the process and mechanism of glaucony (or verdine) genesis must account for the independence between the chemical and

302 mineralogical nature of the initial substrate (i.e., the immediate environment of genesis) and the nature of the authigenic mineral formed and facies developed. However, all substrates are not equally favourable for glauconitization (i.e., the nature and size of the substrates appear to be somewhat interrelated with the observed stage of evolution in a given sediment).

Figure 3. Glauconitization of substrates differing by size at the sea-water/sedimentinterface. a) small grains; b) coarse grains; c) lithoclasts; d) hardground. Dots indicate the density of the glauconitization and arrows represent cation exchanges. (After Odin and Matter, 1981)

A first observation is that sand-sized grains smaller than about 100 pm in diameter are always less evolved than larger ones. A second observation is that when grains generally larger than 1 mm in diameter are submitted to glauconitization, the centre of the grains is the least evolved or is not green at all, the intermediate zone is dark-green and the outermost zone is lighter. These two observations lead to the conclusion that some sort of optimal conditions of glauconization exist a specific distance from the exterior of a glauconitized substrate. Finally, a third observation is that for a similar grain-size, dfferent minerals do not result in a similar stage of evolution in a given sediment: the optimal size depends on the nature of the mineral submitted to the verdissement

303 (see data p. 273). Figure 3 summarizes the situation. Four cases considered are a) small grains; b) large grains; c) lithoclasts (boulders) and d) hardground. The four different cases are part of a continuous single series of physical substrates differring only in size; we progressively pass from a granular (a and b) to a film habit (c and d). The optimum quality of the substrate (i.e., the substrate able to completely evolve to pure dark-green clay) is met for an intermediate size, usually 200 pm to 500 pm in diameter (Fig. 3, b). The optimal dimension for carbonate grains (500 pm) has been observed to be larger than for faecal pellets (200-500 pm) or than for mica flakes or quartz grains (less than 200 pm). This is illustrated in Figure 4. For faecal pellets, four size fractions have been considered with diameters of about 1500 pm, 300 pm, 150 pm, and 35 pm (y axis). The x axis shows the time in years as deduced from observations on present-day continental shelf, and identifies four stages: 1) nascent, 2) slightly-evolved, 3) evolved and 4) highly-evolved. Large faecal pellets (I) become green (dotted area), their size does not change through time and after a certain period of evolution (about lo4 years) all pores are filled and evolution stops. Faecal pellets with diameters of approximarely 200-400 pm (11) are the most favourable for glauconitization. After a period of verdissement, cracks appear (stage 2), and the mean size of the initial substrate is generally increased. However, some cracks are so deep that the grain is broken. Thus from one type of substrate in stage 1, we obtain two populations of green grains in stage 3: one larger and the other smaller than the initial stage. Evolution follows and dark-green grains are obtained. Small faecal pellets (111) become green and evolve until the end of stage 2 when evolution is stopped without change in size. These grains keep their initial form and remain light-green in colour. Finally, the very small pellets (IV) may host the formation of green clay at the beginning of evolution, but they are finally disaggregated by alteration and after a long period of contact with sea-water, these small pellets may be destroyed. A similar set of evolutionary changes occurs for carbonate bioclasts. However the size of each category (I, 11, IV on Fig. 4b) is larger than for faecal pellets; the dissolution of the more alterable small bioclasts occurs more quickly than for faecal pellets. Figure 4c indicates the largest quartz grains and mica flakes where authigenic green clay has been observed. For quartz, evolution usually stops very quickly and green fissures can be observed. Micas about 100 pm in diameter allow green authigenic clays to grow for long periods and the volume of the initial substrate may be greatly expanded (Fig. 4c, II). Figure 4d shows glauconitization patterns for infillings of test. The largest ones (I) have thick walls which are not destroyed by alteration; after perhaps lo4 years, the chambers are filled and evolution is stopped. In this case, the substrate remains easily recognizable. The most favourable size (Fig. 4d, 11) allows green clay to form, and after a certain time the walls of the test are disaggregated. Some infillings remain coherent, follow their evolution and become larger, cracked and dark-green (Elphidium, for example, see Fig. 6, p. 257). But others, like globigerinid infillings, are dissociated when the carbonate test is dissolved; each chamber infilling becomes a free grain which is

304

Figure 4. Diagrams showing the size evolution of various substrates submitted to the glauconitization process. I) size too coarse to be fully glauconitized; 11) the most favourable size; 111) size allowing glauconitization to proceed but not until the latest step; IV) size too small for creating the favourable confinement. 1,2,3 and 4 indicate the nascent, slightlyevolved, evolved and highly-evolved stages of glauconitization. Dots indicate active glauconitization. (Redrawn from Giresse et al.. 1980)

frequently "too small" to host the formation of new authigenic green clay; therefore, evolution is stopped. This situation is similar to that for the small tests (Fig. 4d, 111) for which evolution is stopped after the test dissolution; the resulting cham- ber infillings remain light-green in colour and keep the form of IV), the chambers. The very small tests, 30 pm or less in diameter (Fig. 4, generally do not host the formation of green clay. These observations on the size of the substrates of verdissement sometimes permit deduction of the nature of the original substrate even though no concrete

305 trace of the structure or form is still visible. For example, most of the very large green grains in sediments presumably result from the complete evolution of carbonate bioclasts, while small rounded green grains could be interpreted as resulting from glauconitization of foraminifers (chamber infiiling). MINERALOGY OF GLAUCONY

A complete mineralogical description of the green grains (glaucony) separated from sediments has to take into account the entire evolution of the facies (deposition of sediment, beginning and ending of the verdissement process, burial, and exposure as outcrops). Three components may be present in the green grains: preliminary substrate, authigenic green clay minerals, and minerals formed during early burial. We will not consider later modifications.

Substrate components n 500

-m

z-

0

glaucony

n

clay

W

u Q k

z W

m

2

a Q

150

79 -

52

12

-

I

1

2

3

4

5

K,O

(%)

6

7

Figure 5. Evolution of the K-Ar equilibrium in Recent faecal pellets from the Congolese continental shelf. The y axis shows both the radiogenic content inherited from the substrate and the approximate K-Ar apparent age calculated for each analysed phase.

306 X-ray diffraction studies as well as microscope observations allow us to identify the mineralogical composition of the substrate. Nearly all components of a sediment, provided that they are in a granular form, may become green. Particles made of detrital clay are not rare; the most frequently recognized inherited clay mineral is kaolinite (Ehlmann et al., 1963; and Chapter Cl). This is because the intertropical climate appears today especially favourable for the glauconitization process; under these conditions, kaolinite is usually the most abundant clay derived from the continent. In mud pellets from higher latitudes a mixture of illite, chlorite, and smectite is usually found. A 2:l sheet silicate is relatively uncommon as a substrate component. Therefore, in contrast with assumptions made in earlier theories, there is no "precursor" mineralogically similar to glauconitic mica which must be present in order for the glauconitization process to occur. The general destiny of the substrate components is to progressively disappear. We regard them, however, as an integral part of green grains since they are always present during part of the evolution of the facies and even very late during that process as shown by preliminary isotopic data in Chapter C1, plus the study of the verdissement of biotite and quartz in Chapter C2. We will give more details concerning this sort of inheritance of isotopes because this phenomenon is of fundamental importance for radiometric age interpretation when glaucony is used as a geochronometer in stratigraphy. Studies of the Congolese faecal pellets provide the best known isotopic behaviour of glaucony at the time of genesis. The K-Ar system was discussed by Odin and Dodson (1982); the Rb-Sr system was investigated by Keppens et al. (in press), and additional data are presented here. The principle of radiometric dating schematically consists of measuring the ratio radiogenic/radioactive isotope in a mineral or group of minerals considered to be geochemically closed systems. In a closed system, this ratio increases with time by accumulation of the radiogenic isotope i.e., 40Ar or 87Sr) resulting from the decay of the radioactive one (40Kor Rb). In the example reported here, an inheritance of radiogenic isotope will easily be identified since the radiometric age of the faecal pellet is negligible (zero age, given the analytical precision). Therefore, calculated apparent ages, isotopic ratios, or radiogenic isotope contents different from those characteristic of the present time will denote the presence of inherited components within the faecal pellets. The potassium-argon data presented above (Table 3, p. 239 in Chapter C l ) are plotted in Figure 5. The evolution of the isotopic equilibrium 40Ar versus K is shown from the presumed original clay-size agglomerated fraction to the K-rich faecal pellets. The most evolved grains have not yet reached the zero radiogenic argon content and zero apparent age; thus the presence of radiogenic isotope inheritance is demonstrated in this Recent deposit. The same trend has been observed for other Recent glauconitic grains. There is variation from one deposit to another in the amount of initial radiogenic argon, but the trend remains similar everywhere (Table 3, below). An absence of inherited 40Arcan only be suspected a riori for highly-evolved glaucony or for green grains known to result from Ar free substrate.

S

B

307 Table 3. Relationship between K and Ar contents of Recent glauconies from the seabottom. The remaining proportion of initial inherited argon is quoted in parentheses. (Data from Odin and Dodson, 1982; Odin and Stephan, 1981)

.

Radiogenic argon ( n l g - 1 ) Congo Senega 1 Mexico

K20

Sample

(wt

%)

Clay size

< 2

30 ( 1 0 0 % )

G laucony

3

15

(50%)

G 1aucony

3.5-3.7

9

(30%)

G 1aucony

4.1-4.7

7

(25%)

G laucony

2.6 (9%)

6.5

1 4 (100%)

3.7 (100%)

-

-

3.5 (25%)

-

0.5-0.9

(14-24%)

-

Table 4 gives Rb-Sr analytical results for faecal pellets from the Congolese continental shelf. The isotopic evolution was traced from the clay-size fraction presumed to represent the initial substrate (A 594) to the most evolved cracked dark-green faecal pellets available which, however, are far from having reached the final stage of possible evolution for glaucony. Samples are more and more evolved from top to bottom of Table 4. All data in the table are not comparable because acid treated samples (Keppens and Pasteels, 1982) are probably freer from common strontium contamination than the untreated ones. However, all results show the same pattern (Figure 6). Data from Table 4 are plotted such that the relative age for each sample is a function of the slope of the line drawn between the representative point for that sample and an initial strontium isotopic ratio (y axis for x = 0) which is assumed to be that of the environment in which the rubidium-bearing minerals grew. Table 4. Rb-Sr isotopic data for faecal pellets (samples G) and presumed initial substrate (A594) from the Gulf of Guinea. 87186Sr

Rb PPm

Sample

E x c e s s 87Sr ( ppb ) ~~

Apparent age Ma 2 2 c

A594**

1.58

65.1

290.5

0.65

0.713

108.8

~500

(3313 A*

3.00

103.5

23.7

12.7

0.738

62.2

165

(3490 h*

3.41

120.6

14.3

24.6

0.743

44.0

96

2*

(3490 e *

4.19

131.1

11.8

32.5

0.741

34.2

68

1*

G490 a *

4.25

132.1

10.8

35.6

0.742

32.3

65

1*

G362 0**

6.6

155.9

12.4

36.5

0.718

10.2

* mean of two sample aliquots differently treated with HCl (E. Keppens in Denver) ** without acid treatment (P.D. Fullagar in Chapel Hill)

_+

-17

*** excess 87Sr relative to 87Sr/86Sr = 0.7092 for common strontium in the present sea

3*

308 0 7 ~I ''Sr r 0 740

0 730

0 720

0.71C ;INITIAL RATIO I

10

20

30 87R b /e%r

*

Figure 6. Rb-Sr isotopic data obtained from Recent glauconitized faecal pellets (Table 4). The apparent ages decrease from about 500 Ma (a very imprecise age due to uncertainty of the initial strontium isotopic ratio) to 17 Ma for the cracked green grains with 6.6 wtoA K20.A zero Ma age would be represented by a horizontal line. Two differently treated sample aliquots were measured for samples 313,49Oa, 49Oe, and 490h; the two data points are at the end of the dashed lines and the mean is shown as a black mangle.

Rb-Sr apparent ages progressively decrease (Table 4, Fig. 6 ) as glauconitization proceeds; Rb content and the 87Rb/86Srratio increase along with potassium contents of the green grains. The 87Sr/86Srratios remain higher than 0.7092 which is the approximate value for Recent sea-water (Faure, 1982; De Paolo and Ingram, 1985); this indicates that all these samples have inherited an isotopic signature which has resulted in apparent Rb-Sr ages similar to the K-Arones although slightly older. In summary, in the most evolved grains, about 10% of the initial substrate with its inherited isotopic composition is still present in the Congolese samples.

309 Table 5. Examples of radiogenic argon inheritance in ancient glauconies. Apparent ages calculated are clearly inversely function of the K content and proves the inheritance. Stratigraphic age

Substrate

Potassium increase (wt % K 2 0 )

Concomitant age decrease (radiometric age in Ma)

( 1 ) Mid Eocene (45 Ma)

Chlorite + kaolinite

from 0.9 to 5.5

from 136

to 46.5

( 2 ) Early Miocene (18 Ma)

Biotite + chlorite

from 1.8 to 5.5

from 108

to 18

( 3 ) Late Oligocene (25 Ma)

unknown

from 3.3 to 8.0

from 140

to 25

( 4 ) Late Miocene ( 8 Ma)

unknown

from 3.0 to 7 . 4

from

36.9 to 8.1

These observations are not unique for Recent sediments. An inheritance of radiogenic Ar-bearing components in older glaucony suites has been observed (Table 5). This indicates that the substrate, although not always visible in the green grains, is an important component which will contribute to the mineralogy, chemistry, and isotopic composition of the analysed glaucony during a large portion of the verdissement process. From a geochronological point of view, the glaucony apparent ages obtained on K-poor grains may therefore, give numbers higher than the time of deposition of the sediments; the exclusive ill generally avoid this problem. selection of highly-evolved, K-rich grains w

Authigenic marine clays X-ray diffraction ( X R D ) study The mineralogical variety of the green clays found in the glaucony facies is well illustrated by X-ray diffraction techniques. A specific mineralogical family or series has been defined under the collective term of glauconitic minerals (Odin, 1975a; Velde and Odin, 1975; Odin and Matter, 1981). XRD traces provide specific information on the stage of morphological evolution of the green grains. Because oriented samples tend to bias results in favour of the thinnest crystals (which scanning electron microscopy study shows to be the least evolved), XRD analyses are best carried out on powder mounts. Moreover, in order to obtain peaks with the best possible shapes it is necessary to analyse homogeneously evolved populations of grains with similar physico-chemical properties (i.e., to select grains of similar size, colour, magnetism and density). Figure 7 gives a comprehensive set of XRD traces for glauconies of Upper Cretaceous and Palaeogene age. The variation in position of the first order basal reflexion 001 is the most useful criterion for identification. The variation between 14 A and 10 A corresponds to the mean thickness of the sheets of the clay mineral. Further details on the mineralogical nature of the initial (less evolved) thickest sheet silicate can be obtained by the study of its behaviour following glycolation or heat treatment. With simple glycolation, the 001 peak shifts only slightly toward lower 28 angles but expansion is enhanced if, prior to glycolation, potassium is removed by cation exchange. On heating to 4WC, the 001 peak is shifted to 10 8, (cf Fig. 10 in Chapter B2 and Fig. 7, p. 235, in

310 BURIED GLAUCONIES 112

003

112

020

002

001

Figure 7. X-ray diffraction patterns for Cretaceous and Palaeogene (buried) glauconies. The position of the 001 diffraction peak and the shape of hkl peaks (especially 112 and 1127 vary concomitantly with the potassium contents. In contrast to that, the position of the 020 diffraction peak is constant and easely located. Note that the 002 diffraction peak at about 5 A is very small indicating the high iron content of all these minerals.

311 9

highly evolved 0

Y

evolved

7

%

o 0

5

O O

0

0

slightly evolved 0

3

0 I

I 11.0

I

I

12.0

0

distance 001420

*

13.0 cm

Figure 8. Distance between diffraction peaks 001 and 020 on X-ray diffraction patterns as a function of potassium contents of glauconitic minerals. Distances measured on diagrams obtained with K,Cu emission with a goniometer running 1' per minute and recorder 1 cm per minute. Some data are indicated with error bars. (Redrawn from Odin, 1982b)

Chapter C l ) . These and other analyses indicate that all glauconitic minerals have a 2:l-type structure. This distinguishes the glaucony facies from the verdine and ironstone facies discussed in previous chapters. The most constant peak in both position and shape is the 020 one at 4.53 A. Because 020 is constant and the position of 001 varies as a function of potassium content (Fig. 7), the distance between the two peaks allows us to obtain an easy estimate of potassium content (Odin, 1982b). Figure 8 shows 001-020 distance versus K 2 0 percent for XRD traces obtained using the K,Cu radiation at 1"/minute/cm. Low potassium contents exhibit scatter due to variable contribution of the substrate to the total '(20content. In practice, we may refer to different glauconitic minerals in several ways. The position of the 001 peak is one possibility: we could refer to a 10.5 A-, or a 12 A-glauconitic mineral, etc. Another possibility is to emphasize the potassium content; we could refer for example, to a 4% or a 7% K20-rich glauconitic mineral. Finally, one may calculate the theoretical proportion of 14 A (smectite) layers which must be added to the 10 8, (mica) layers in order to shift the position of the 001 peak to the position actually measured on the X-ray difffac-

'

312 tion pattern; the glauconitic mineral would then be identified by its theoretical proportion of smectite, or open sheets, or expandable sheets (e.g., a glauconitic mineral with 10% smectite: 001 peak at 10.4 A; with 30% smectite: 001 peak measured at 11.2 A; or 50% smectite: 001 peak at 12 A). Several other peaks show distinct changes during the glauconitization process. As the clay mineral evolves from a smectite-type to a mica-type, the shape and size of the 112 and 112- peaks on either side of the 003 one are modified. According to Bentor and Kastner (1965) these two peaks give a reliable measure of the order-disorder of the layer silicate structure; however, rather than actually disappearing when the structure is disordered as Bentor and Kastner suggest the peaks mostly decrease in height. Figure 9 presents a series of diffractogrammes for glaucony samples collected from the present-day continental shelf; these samples apparently have not been reworked. Therefore, they are likely to have never been buried. Their age of formation is Late Miocene to Recent as supported by sedimentological and geochronological studies off Congo (samples 312, 313,319,490, Giresse and Odin, 1973; Odin and Dodson, 1982), off northwestern Spain (sample 210, Lamboy and Odin, 1975; and Chapter C2 in this volume), off California (sample 585, Odin and Dodson, 1982) and off South Africa (samples 640 and 642, Odin, 1985b). The respective potassium contents are given for comparison with Figure 7. Detailed examination of the two figures (9 and 7) indicates that there is a noticeable difference in the shape of the hkl peaks for a given potassium content. The 112 and 112- peaks (see also the small 021 and 111peaks) are generally much smaller for unburied samples. As indicated by Bentor and Kastner (1965), the poorly shaped hkl peaks would indicate less order in the unburied samples. In contrast to the hkl peaks, the 001 peaks are rather well-shaped in our diagrams, and sometimes better shaped than for glauconies separated from buried sediments. We suggest that during the burial and aging of the grains a further recrystallization process occurs which leads (without noticeable potassium enrichment) to a better crystallized and more stable phase. The whole glauconitization process from substrate to highly-evolved and even relict green grains may occur in direct contact with sea-water; however, when burial occurs at any time during this synsedimentary process, it causes further systematic recrystallization leading to more ordered glauconitic minerals.

Chemical study There is a considerable amount of published information on the chemical composition of green grains. Published data often include results for samples which have undergone late diagenetic evolution or are mixed with inherited material. Reliable original data obtained from large samples can be found in classic works by Hendricks and Ross (1941), Smulikovski (1954), Foster (1969), Cimbalnikova (1971a), and Odin and Matter (1981). Toda the tenden-

-4-+-+

Figure 9. X-raydiffraction patterns of unburied glauconies. The potassium contents increase from the bottom to the top. Diffraction peaks due to impurities are shown by dotted pattern.

hON BURIED GLAUCONG?

I

020

001

I

I

314 cy is to give results from micro-analyses. Although certainly useful for other purposes we disagree with the assumption that the spots analysed represent the mineral formed near the sea-bottom. Also, many recent results assume values for the state of oxidation of iron and water contents. Generally speaking, the precision of the analyses is not good enough and the data are not reproducible enough for consideration here (see discussion of micro-analyses, p. 369, in part D of this volume). Finally, from a sedimentological (environmental) point of view, data obtained from many grains appear more significant. Table 6. Cation contents of glauconitic green grains. Oxide Si02 A 203

Usual contents 47.5 to 50.0% 3.5 to 11%

Fe2°3

19 t o 25%

FeO

1 . 0 to 3.2%

MgO

2 t o 5%

K20

3 to 9%

Deviating results

a f e w higher v a l u e s indicate quartz admixture some v a l u e s up to 15% higher values when oxidized

lower v a l u e s when substrate abundant

Typical abundances for the major cations present in glaucony are summarized in Table 6. Figure 10 gives histograms of compositions for about 60 Cambrian to Holocene glauconies, most of them published in Odin and Matter (198 1). With the exception of potassium contents (scattering between 2.6 and 9.1% K20,and which are not shown in Fig. lo), the cations involved are present in the same general ranges as for the green clays from the ironstone and verdine facies. Both glauconitic minerals and minerals from the verdine facies have low A w e ratios, high Fe3+/Fe2+ratios and high Fe/Mg ratios compared to other clay minerals with the same structure. The influence of a common environment (sea-floor) is indicated by these similarities. The structural formula of the resulting glauconitic minerals may be written:

where : x varies from 0.2 to 0.6, and : y (the sum of the divalent octahedral cations) ranges from 0.4 to 0.6. Therefore, the layer charge of glauconitic minerals results from both tetrahedral and octahedral substitutions. The relationship between iron content in glauconitic and other 2:l clay minerals, potassium content, plus stage of evolution, are sometimes not clear in the literature. Figure 11 was composed (Odin, 1975a) in order to emphasize the lack of continuity of the chemical composition between illito-smectitic and glauconitic minerals at a time when the genesis of glauconitic minerals was con-

315 4

5-

2

50 %

10 %

57 0

0

MgO

1

10

n

20

Fen03

Figure 10. Chemical compositions of glauconies, plotted as histograms. Data from Odin (1975a), Odin and Matter, (1981), and supplemented with a few unpublished results. One discrepant value of 7.2% FeO for a Palaeozoic glaucony was omitted. All samples carefully purified before analysis; however the distribution of total iron (Fe203) indicates that iron oxide impurities might have been present in several samples resulting in excess iron between 27% and 29.5% Fe203.

316 sidered to result from the evolution of illitic (aluminium-rich) degraded clays. Figure 11 shows an obvious compositional break between 10% and 15% of total Fe203. The very few intermediate points are for material with low interlayer cation content which can be interpreted as resulting from mixtures of an illite-rich substrate with authigenic glauconitic minerals. Some other intermediate points (stars) all represent ferric iZZites which have been related to restricted hypersaline conditions (Kossovskaya and Drits, 1970). Thus, there are essentially no intermediate 2:l clay minerals with Fe 0 contents between 8% and 15%. This definitely indicates that the majority o?' glauconitic minerals do not result from the progressive transformation of illitic clays by Fe for A1 substitution in the octahedra, a theory still supported by most authors at the end of the 1970s. The persistency of that theory occurred in spite of pertinent observations by many authors such as Takahashi and Yagi (1929) who recognized from studies of faecal pellets that glaucony is characterized by a high iron content even in the earliest stages of evolution. Ehlmann et al. (1963) also observed the abundance of iron in the earliest infillings of foraminiferal tests, an observation subsequently confirmed by Pratt (1963) and Seed (1968). Finally, Foster (1969), Odin (1975a), and Birch et al. (1976) have all shown that the iron content is not related to the content of interlayer cations and that iron is fixed in the mineral structure prior to the incorporation of potassium. These studies contradict the hypothesis which considered that incorporation of iron and potassium were interrelated during the transformation process of an altered Al-rich 2: 1 layer structure to a glauconitic mineral. The abundance of potassium appears to be the best criterion for distinguishing one glaucony from another. For example, the potassium content appears to govern most physical properties of glaucony including XRD patterns as shown above, amount of expandable layers (Manghnani and Hower, 1964; Velde and Odin, 1975), density of the grains (Shutov et al., 1970), refractive index (Cimbalnikova, 1970),cation exchange capacity (Cimbalnikova, 1971b), and paramagnetism (Odin, 1982b). All these properties can be used to identify the degree of evolution and select homogeneous sets of grains for analyses.

Post-genesis components of green grains Following genesis in contact with sea-water, we have suggested that early burial induces recrystallization resulting in a better crystallographical ordering of the clay structure. This evolution can be considered as part of the general process of glauconitization sensu lato. W e do not intend to discuss in detail other alteration processes which characterize deep burial (Odin, 1982c), tectonization (Conard et al., 1982), and weathering or reworking (Odin and Rex, 1982). However, we will consider some alteration processes which take place at the present-day sea-bottom. One type of alteration results in the frequent presence of goethite (hydrated femc oxides) in green grains. This situation appears linked to a long hiatus of deposition during which occurs a regression, or any arrival of more oxidizing water concomitant with mobilization of green grains. Especially in a subtropical

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Figure 11. Position of the illitic and the glauconitic mineral families with respect to their iron and interlayer cation content. Data for glauconitic minerals (black circles) are from Cimbalnikova (1971a), Parry and Reeves (1966), and Odin (1975a). Black circles in mangles are from Hower (1%1) who interpreted data as supportinga continuous series from aluminous to iron-rich minerals. We suggest that they would be better considered as mixtures between authigenic (iron-rich) and substrate (aluminium-rich)clay minerals. Stars represent iron-rich illite from saline deposits (Kossovskaya and Drits, 1970). Open circles represent illitic minerals analysed by Hower and Mowatt (1966). (Figure redrawn after Odin and Matter, 1981)

climate, green grains quickly become brown at depths shallower than 30 m, and Fe,O increases. The same type of alteration is even more common with the verdine f3acies. For both the glaucony and verdine facies it is accepted that iron was already ferric before alteration, and verdissement appears to be a first step of iron enrichment of the substrate, and guethitization a second step. A second type of alteration common on the present-day sea-bottom and which also occurred in older sediments, results in the association phosphateglaucony. This association has been reviewed by Odin and Utolle (1980) who conclude, in agreement with Lamboy (1976), that although both the general environment and the microenvironment of the two parageneses are different there are many similarities. Glauconitized phosphate (i.e., phosphate substrate) and phosphatized glaucony (i.e., phosphate invading green grains or previously glauconitized hardgrounds) are common. Both facies form during a hiatus of sedimentation. When the hiatus is very long some slight modification of the environment (e.g., change in the chemical or physical properties of bottom currents) induces a change in the paragenesis. In summary, at first glance, the mineralogy of the glaucony facies appears

318 to be characterized by a wide variety of compositions. In fact, glauconitic grains are made of several components: the substrate (variable in composition) which tends to disappear during the evolution, and the authigenic clay which tends to replace the former substrate. When the purely authigenic unaltered phase can be analysed, this material is found to be restricted to a specific series of clay minerals which form a continuous family between smectitic and micaceous end-members. All members of the family are characterized by potassium as the interlayer cation and a high iron content most of which is in the ferric state. The glauconitic minerals are a unique series of clay minerals which is distinct from other sedimentary clays. Mineralogical study, as well as the previous study of habits, has shown that there is no particular link between the mineralogical composition of the substrate and that of the authigenic phase. GENESIS OF GLAUCONY: T H E VERDISSEMENT PROCESS

The layer lattice theory We will mainly summarize discussion by Odin and Matter (1981) and Odin and Morton (in press). The model for the genesis of glaucony proposed by Burst (1958a and b) and Hower (1961), namely the luyer lattice theory, became widely accepted in the 1960s. This model was suggested for the evolution of many clay minerals. This model proposes that the authigenic green clay results from the transformation of a degraded layer silicate, with the authigenic mineral retaining a memory of past structure. According to this model, a pre-condition for glauconitization is that a mineral with similar crystal structure to glauconitic minerals, such as illite or smectite must be present. Odin and Matter (1981) have listed seven observations which are incompatible with that theory: 1) Glauconitization of detrital mica is quoted as an example of the layer lattice theory. However, Odin (1972a; see also Chapter C2) has shown that verdissement of detrital mica takes place through growth of glauconitic minerals between the mica sheets, which remain unaltered for a considerable period. Therefore, glauconitization requires neither the crystal architecture nor the ions of the initial mica. 2) In many cases, verdissement proceeds on granular substrates which are wholly calcareous (Chapter C2), a situation which the layer lattice theory cannot explain. 3) Similarly, most glauconitized hardgrounds in ancient formations are limestones or chalks (Aubry and Odin, 1973; Juignet, 1974). 4) Two mineralogically different authigenic clay minerals can be generated from the same substrate. For example, biotite undergoes glauconitization off northwestern Spain (Chapter C2) and California, but undergoes verdinization off Sarawak and French Guiana (Chapter B3). Similarly, in different parts of the Gulf of Guinea, glaucony and verdine have both been generated from similarfuecalpellets, composed largely of kaolinite (Chapter B4 and Cl). 5) In areas where the sea-floor is muddy, verdissement only proceeds in the faecal pellets not in the diffuse clay of similar mineralogical nature. The process

319 is clearly governed by the physical nature of the substrate, rather than its chemistry and crystallography. 6) If illite or smectite were specifically favourable substrates for glauconitization, the authigenic clays would frequently display a continuum between aluminous 2:l and ferric 2:1 structures, and this is not the case (Fig. 11, p. 317). Similarly, there is no chemical composition continuum between kaolinite and phyllite V for the verdine facies. 7) The layer lattice theory postulates that the octahedral layer loses aluminium at the same time as it gains ferric iron. The similarity in geochemical behaviour of these two irons makes this difficult to achieve. The layer lattice theory, therefore, does not adequately explain the glauconitization process.

The mechanism of verdissement

Our understanding of the mechanism of verdissement starts with two critical observations: 1) Evolution begins and proceeds close to the water-sediment interface. Cores from present shelves invariably show that glaucony grains occur over a relatively thin zone, usually several metres and often few decimetres, immediately below the sea-bottom. 2) The mechanism proceeds more efficiently, and often exclusively, where the sediment is in granular form. This is the reason why glaucony is found mainly in granular form. As already discussed, the starting material is highly porous. Crystal growth begins in these pores, which may extend across an entire grain. The neoformed minerals grow as blades attached to internal surfaces or as minute lepispheres in the pores. The glauconitic minerals may closely mimic the shape of the pores. Clearly, the porosity of the particles is an important factor. The grains essentially act as a sponge and provide an environment where geochemical reactions can take place. The development of new minerals soon imparts a green colouration to the grains. Even at this early, or nascent stage, the main green clay is iron-rich and characteristic of the glauconitic mineral family: K20 contents are in the order of 2%-4% (stage 1, Fig. 12). The minerals of the substrate (in geochemical disequilibrium with sea-water) become progressively destroyed as verdissement proceeds. The more stable the substrate, the longer it takes to disappear. Calcareous substrates being least stable, are easily and rapidly replaced by authigenic clays (Lamboy, 1976). Conversely, micas and quartz are much more stable, and remnants of these materials are commonly found in ancient glauconies, even in evolved grains. As the substrate is destroyed, it leaves a new system of pores that, in turn, becomes filled with authigenic clays, developing as blades or rosettes. At this stage, an individual grain largely consists of glauconitic minerals with K20 contents between 4% and 6%; these minerals show globular, caterpillar-like or blade-like habits. The grain is said to be slightly-evolved (stage 2, Fig. 12).

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Figure 12. Glauconitization of a granular substrate: 1) nascent stage; 2) slightly-evolved stage; 3) evolved stage; 4) highly-evolved stage. Stars represent glauconitic minerals. The grain (about 5 mm in diameter) is shown at the boundary between sea-water and sediment. After this evolution, the glauconitic minerals do not evolve anymore at the sea-water/ sediment interface and the green grain is relict. (Redrawn from Odin and Dodson, 1982)

With continuing evolution, a series of recrystallization takes place, which tends to obscure the initial texture of the grains. Because the grain interiors are more favourable for crystal growth than the surfaces, the larger and betterorganized crystallites are found at the grain centres. This results in more rapid growth at the centre compared to the margin, and cracks appear at the surface. Earlier studies attributed these cracks to dehydration with reduction in volume (Ehlmann et al., 1963), but SEM analyses indicate that this is not correct. In glauconitized mica flakes the growth of authigenic clay between the individual sheets causes the flakes to open into accordion-like grains. At this stage, q 0 contents are between 6% and 8%, and the grains are said to be evolved. If environmental conditions remain suitable, the cracks created in the preceding stage are filled (Chapter C2) imparting a smooth aspect to the grains. This is typical of highly-evolved grains, in which K20contents exceed 8%. The minerals filling the cracks are less rich in potassium than the rest of the grain, again illustrating that the surface of the grain is less favourable for clay mineral authigenesis than the interior. The evolution process may be halted at any stage of the above continuum if the environment becomes unsuitable. Two main factors appear to be involved: sea-level change and burial. A regressive phase may introduce the grains to a more oxidizing environment; a transgressive phase may cause phosphatization. Although verdissement may still proceed at distances of a metre or more from open sea-water, burial below this level rapidly halts the process. Consequently,

32 1 a high rate of detrital infux will inhibit or entirely prevent glauconitization. The entire evolution process described in Figure 12 as well as possible subsequent alteration (goethite, phosphate) occurs in close contact with sea-water. However, if this synsedimentary evolution is halted at any stage, a further crystallographical reorganization occurs in the green grains during early burial. This reorganization results in additional crystallographical ordering of the already formed glauconitic minerals. This point is important because it suggests that the final environment of equilibrium of these minerals could be interstitial water which differs in composition from sea-water. This late recrystallization does not necessarily re-equilibrate the green grain with the isotopic or elemental composition of its interstitial water, however. Odin et al. (1977) have shown that even a crystallographically observed experimental phase change induced by hydrothermal pressure does not completely remove argon isotopes from glauconitic minerals. This indicates that, although recrystallized, the green clay phase may retain certain isotopic and chemical characteristics of the earlier phase. In some cases, a late intra-sedimentary addition of green clay to the already formed glauconitic grains can also be postulated from the observation of fibro-radiated r i m s on the grains (Fig. 10 in Chapter C2). In brief, the above observations and proposed mechanism suggest that the closure of the geochemical system for a green grain is a complex question. It probably occurs distinctly after the time of deposition of the sediment. This indicates that the fossil remnants present in the glauconitic horizon could have an age noticeably older than the time of closure which is dated when a radiometric measurement is undertaken. An idea of the duration of the evolutionary process is of interest to estimate the possible difference between time of deposition and time of final closure, and, concomitantly, the duration of the corresponding hiatus of deposition. Observations made on samples from the present-day sea-bottom are relevant. Recent glaucony from the Congolese continental shelf, from depths shallower than 110 m, are younger than 20,000 years and have maximum K20 of 5%. A similar situation was observed for samples from the Aegean Sea (Robert and Odin, 1975). Bornhold and Giresse (1985) estimate that less than 3,000 years was required to form glaucony with about 3% K20 off Vancouver Island. The relict highly-evolved glauconies off South Africa are Pliocene or older if we disregard the locally observed in situ reworking (Odin, 1985b). The situation appears generally similar off northwestern Spain and on the Chatham Rise (New Zealand; Cullen, 1964). We suggest that evolution to the highly-evolved stage may take lo5 to lo6 years. After that time and if the conditions remain similar (absence of both alteration and burial) green grains progressively become relict. In other words, after about lo6 years they do not recrystallize or grow anymore, and they keep their maximum K20 content of about 8.5 to 9.0%. A summary of the radiogenic argon behaviour in glaucony is presented in Figure 13. It summarizes both the contribution of the inherited substrate at the beginning (assuming it contains significant radiogenic argon) and that of authi-

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EVOLVED EVOLVED

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Figure 13. Evolution of the radiogenic argon content during glauconitization of faecal pellets. The thick line shows the evolution occurring in absence of burial. When burial occurs at time 1,2, or 4, the remaining (or for time 4, in situ accumulated) argon content tends to diminish first (early burial recrystallization) before accumulating due to the normal decay of potassium.

genic minerals at the end. If the inherited substrate is free of radiogenic argon, the evolution simply begins along the x axis. The thick line shows the radiogenic argon content when no burial occurs; when burial occurs at time 1,2,3, or 4, then radiogenic argon generally is presumed to decrease (cases 1 , 2 or 4) but probably not reach zero, as a result of early burial recrystallization. The evolution should be generally similar for strontium isotopes. However, effect of burial could be an increase in initial 87Sr/86Srratio because the interstitial fluids may have higher 87Sr/86Srratio than sea-water. In summary, the mechanism of verdissement in the glaucony facies is similar to that suggested for verdine; it is based on crystal growth and recrystallization. The first stage of the process is the crystal growth of glauconitic smectite which is a neoformed clay mineral using cations from sea-water above

323 the grains, and from interstitial water below. Later, recrystallization processes re-use these ions to grow minerals with high potassium contents within the green grain system. The speed of the geochemical reactions is initially typical for clay genesis although low compared to reactions on land (where soil including clay needs only a few centuries to form). As potassium content increases, the reactions become slower and slower, with the resulting green clay becoming more and more stable. ENVIRONMENT OF GLAUCONITIZATION

Microenvironment The presence of pores is essential for genesis of glaucony (as well as for verdine). The quality of the microenvironment determined by this porosity is an important factor in explaining many observations. First of all, the specific environment of formation results from sea-water introduced in the substrates. The sea-water becomes an intragranular interstitial fluid equilibrating intra- and extra-grain physico-chemical factors. This microenvironment is best defined in terms of confinement given the fact that although glauconitized substrates are in close contact with the infinite reservoir represented by sea-water, the degree of confinement or isolation of the grain system from sea-water determines the extent to which mineral-forming reactions occur. The necessity for confinement in the glauconitization process can be deduced from the following observations: 1) Growth of glauconitic minerals inside microfossil tests, pores, fissures in particles, burrows, or any void in hardgrounds and not at the exterior is an indication that verdissement requires a confined microenvironment different from that present in sea-water. 2) Carbonate grains with diameters smaller than 100 pm are less evolved than larger grains with diameters between 200 pm and 500 pm; this suggests that because the interior of small grains is relatively unconfined, significant interaction with sea-water inhibits evolution. 3) Crystal growth is most effective in the centre of a grain rather than at its less-confined periphery. However, too confined a microenvironment is not favourable either as shown by the following points 4 to 6. 4) Glauconitic green clays generally do not grow diffusely within layers of sediment presumably because the interior of most sedimentary units is too confined for cationic exchanges with sea-water to be effective. 5 ) The absence of ions necessary for crystal growth in many substrates in which glaucony forms indicates that exchange between substrate and surrounding fluids is necessary; similarly, the cations composing the substrate must be permitted to depart. The large volumetric increases which occur in the late stages of glauconitization testify to the introduction of ions from the exterior. 6) Coarse sedimentary particles are only glauconitized on their surfaces because their interiors are too confined to allow ions to enter and contribute to the growth of glauconitic minerals.

324 Therefore, a key factor in the glauconitization process is the presence of a semi-confined microenvironment which ions may enter and leave, but where exchange is not too rapid. Ions primarily come from sea-water, but also from interstitial fluids in the sediment, and presumably, in a few cases, from the substrate itself; the porosity of the substrate acts as a controlled passageway providing optimum conditions for interaction of the relevant ions. Within mica flakes, such favourable conditions of semi-confinement are found in the cleavages. In microfossil tests, semi-confinement is created by the wall of the test which acts as a semi-permeable barrier for migrating ions. In hardgrounds, semi-confinement is controlled by the porosity of the deposit: according to Juignet (1974), glauconitization of porous chalks occurs over a thickness of up to lcm, less porous rocks such as phosphates or cherts develop a glauconitized zone less than l m m thick. Grains which are mobile on the sea-floor are particularly susceptible to verdissement, because motion facilitates renewal of the ion supply. As soon as the substrate is buried, grains become isolated from sea-water, the main source of the cations, and exchanges become more restricted because water circulation diminishes. At this stage, the favourable zone of semi-confinement may transfer from the grain interiors to the pore spaces between grains, thus allowing the formation of a layer of green silicates around grains. This frequently-observed layer is known as the peripheric oriented rim or the fibro-radiated cortex (Collet, 1908; Zumpe, 1971; Odin, 1975a; Lamboy, 1976). Following this, verdissement halts altogether, no matter what stage was reached in the general scheme of evolution proposed above (Fig. 12). Some characteristics of the microenvironment can be deduced from the study of present-day deposits. For example, Boichard et al. (1985) indicate that the pH of the glauconitic sediment is usually 7.9 to 8.0 in the deposit studied in the Makassar Strait where glaucony is abundant and presumably forms today mostly in bioclasts. The results of the Eh measurements depend on the depth of the sample in the core; they change from 0 to 200 mV at the seahediment interface (an oxidizing environment) to about -100 mV five centimetres below. Glauconitization thus occurs at the boundary between oxidizing sea-water and reducing interstitial water; this explains the remarkable geochemical activity of iron which, at this boundary, changes ionic state. This observation agrees with earlier assumptions by several authors (Ireland et al., 1983).

General environment Climatic environment The effect of latitude on the glaucony formation can be discussed in terms of its present-day distribution at the sea-bottom. The map in Figure 14 gives an indication of that distribution. It is not exhaustive because many areas are yet to be investigated in detail. It is based on previous reviews, including Collet (1908), Murray and Chumley (1924), Bell and Goodell (1967), Odin (1973; 1975a), Odin and Matter (1981), Odin and Morton (in press), and supplemented with further studies made for this work ( e g , Makassar Strait, Kerguelen).

325

Figure 14. Presence of glaucony on present-day sea-bottom. Hachured areas indicate presumed glaucony for which no mineralogical details are available. Data reviewed by Odin and Matter (1981), supplemented by data from Murray and Chumley (1924), Keller and Richards (1967), Burnett (1980), Odin and Stephan (1981). The areas shown are generally larger than are the actual deposits in order to facilitate legibility.

Some areas are shown in hachured patterns because no detailed mineralogical studies are available and the identity of the green grains is not established; however, the presence of true glauconitic facies is presumed. According to Figure 14, it might appear that glauconitization occurs in all oceans and that the latitudinal effect excludes only the extreme coldest areas of the globe. However, the situation is more complex. For example, during Quaternary time, very wide climatic variations occurred. This, together with the fact that climatic changes occur more rapidly than glaucony evolution, makes the interpretation of glauconitic deposits possible only if one precisely knows when glauconitization O C C U K ~for the sediment collected from the present seabottom. Low latitude green clays can generally be considered still in the environment in which they formed (Gulf of Guinea, Makassar Strait, California, eastern United States). In high latitudes, many deposits are known to be relict: South Africa, Kerguelen, Chatham Rise, and Scotia Ridge. Bornhold and Giresse (1985) give a Holocene age to the glaucony from off Vancouver Island (50" north latitude). Bjerkli and Ostmo Saeter (1973) have not thoroughly discussed the age of their Norwegian glaucony (63.4" north latitude); moreover, the small quantity of nascent green clays available allowed them to obtain only semi-quantitative chemical microprobe analyses, and complex X-ray diffraction patterns. Therefore, the nature of that deposit would need to be confirmed. We do not need to discuss all known deposits; the important point is that each deposit needs to be categorized: Holocene, Pleistocene, relict, perigenic or in situ reworked. Low latitudes in general (with

326 the present tropical climate) appear to be favourable for glauconitization; however, the formation of in situ recent glaucony can reasonably be expected in much cooler regions. This situation is very different from the one favouring the ironstone or verdine facies described previously. In high latitudes off Norway, the temperature of glauconitic sediments was precisely reported at +7"C; Bornhold and Giresse (1985) are rather vague regarding the exact nature of the environment of deposition of their high latitude Holocene glaucony. In low latitudes, Boichard et al. (1975) published a temperature of 14.3"C for a glauconitic sample collected from -300 m depth in the Makassar Strait. This temperature is similar to that usually reported for the intertropical sea-bottom on the continental shelf; for example Porrenga (1967b) indicated that in such a setting, glaucony forms at temperatures below 15°C. In contrast, he also noted that a temperature of above 20°C seems to be necessary to allow "berthierine" (in fact our verdine facies) to form.

Depth of glauconitization The depth factor can be considered by evaluating the present distribution of glaucony. The facies occurs schematically in two distinctly different oceanographic environments: stable margins and tectonically affected areas. Glauconybearing stable margins are comparatively well-known from the Atlantic Ocean. The facies occurs mostly between -60 m and -550 m depth, sometimes forming up to 90 wt % of the sediments collected. The optimum depth is about -200 m, near the top of the slope (Odin, 1973). Deeper than -500 m, the green grains are usually interpreted as perigenic being transported from the top to the bottom of the slope. Shallower than -60 m, glaucony is either absent or reworked from ancient rocks and this character again emphasizes the incompatibility of glaucony and verdine. The problem is whether or not the observed optimal conditions between -150 m and -300 m depth are linked to a factor which is particular today at this depth. Isolated active tectonic highs lack a nearby continent to provide material formed at shallower depths to the deep-sea sediments. Such areas as the Chatham Rise (New Zealand), the Scotia Ridge (South America), or Kerguelen Heard Plateau (South Indian Ocean) show that depth of -2,400 m (Chatham Rise), -500 m to -2,000 m (Scotia Ridge), and -500 m to - 1,000 m (Kerguelen; see Chapter C3) were favourable for glaucony genesis. Glaucony is not known from the shallower sea-bottom in these areas except for the Kerguelen-Heard Plateau from where Logvinenko (1982) incidently reports glaucony at -60 m or -270 m depth but does not give more specific information. A common characteristic of the glaucony on stable margins and in tectonically active areas is the fact that in both cases it is located at or near the top of a topographic high with a deep-sea basin nearby. According to Morton et al. (1984) glaucony can also form in deep-sea basins because green grains were found in Miocene sediments which probably formed at depths of around -2,000 m to the southwest of the Rockall Plateau (North Atlantic Ocean). Finally, deep occurrences of glaucony have also been reported from active margins near large land masses. On both sides of Japan, glaucony from deeper

327 than -2,000 m (Karig et al., 1975) occurs with faunas indicating shallower water than the present depth. Off Los Angeles (California) to Panama, glaucony was recovered at depths in excess of -1,000 m (Odin and Stephan, 1981). In both areas, suggestions of recent rapid subsidence could explain the deep occurrences. In brief, glauconitization seems to occur at various depths between -60 m and -1,000 m and possibly deeper; thus the facies cannot be considered as a bathymemc criterion. However, all deposits appear typically in an open marine environment where bottom currents are present. The intensity of these currents can be related in part to the topographic highs considered above. The way in which currents contribute to glauconitization can easily be postulated. Sea-bottom currents probably lead to accumulation of glaucony on the outer part of the shelf. Close to the shore, detrital influx exceeds winnowing producing accumulation rates which prevent glaucony formation. Where water is deeper on the shelf (-60 m to -300 m) winnowing causes continual redistribution of sedimentary particles; grains are exposed at the sea-floor for long periods sufficient to allow glauconitization. Below -300 m, energy is less and sediments accumulate more rapidly again inhibiting in situ glauconitization while shallower glaucony is transported deeper forming perigenic deposits. Moving sea-water enhances the addition of ions to the evolving grains by causing an in situ reworking of the sediment itself allowing slightly buried grains to return to contact with sea-water, and by replenishing ions in the fluids adjacent to the substrate.

Glaucony and time relationship Additional remarks may be done on the estimates of the time needed for glaucony to reach successive morphological and mineralogical stages in nature. The proposed sheme (Fig. 12 and Fig. 13) indicates that initial stages are quite rapidly reached; the K-poor green grains may be formed in a few thousand years. Potassium enrichment becomes less with time. This must be due to the fact that the more and more stable phases obtained are increasingly difficult to recrystallize. The rate of glauconization presumably varies depending on internal factors such as the nature of the substrate, and on external factors such as the temperature of sea-water. Even if the general scheme we have presented above is correct at a given time, every point within a grain need not to have reached the same stage of evolution. On the contrary, a single grain often appears heterogeneous. The observation by Bornhold and Giresse (1985) of 'focus' of more advanced glauconitization within a grain is relevant. Within sediment, glaucony is heterogeneous, and an average should be considered for sedimentological interpretation. Figure 15 is a picture of a young glaucony collected off Casamance (Senegal). The mixture of different stages of evolution in a single sediment is indicated by the light- and dark-green colours. Moreover, within some grains, one may see that dark-green particles have been re-ingested by mud-eaters and thus incorporated in a new faecal pellet.

328

Figure 15. More or less glauconitized Recent faecal pellets from a surface sediment collected off Casamance (Senegal, sample J.C. Faughres, Univ. of Bordeaux). The black bar is 1 mm in length. Note that several very light-green faecal pellets agglomerate small white lithoclasts or bioclasts and dark-green pieces of previously glauconitized faecal pellets.

Because glauconitization occurs in contact with sea-water, the presence of glaucony in a sediment will identify a hiatus in the accumulation of sediment. This is especially evident when hardgrounds are involved. With grains, the stage of evolution reached will trace the length of the hiatus. Smectitic glaucony, in fact, can be found dispersed in a sediment even if there was not a significant hiatus. We considerer that glaucony with a mean potassium content of 7% K,O) indicates a geologically significant break in sedimentation ( lo4-1 years). A more lengthy interruption in sediment accumulation (about lo6 years) will be indicated by highly-evolved glaucony, dark-green in colour, rounded, and possibly either partly goethitized or phosphatized. This last type of relict glaucony corresponds to sediments found off northwestern Spain, off South Africa, or on the Chatham Rise where essentially nothing was deposited during the last 2 to 5 Ma.

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Sea-level

Figure 16. Distance between the zone of production of substrates and the zone of glauconitization. The larger the distance, the larger the transgression needed before glauconitization is possible. Glauconitized shell bioclasts will denote an important transgressive phase; glauconitized planktonic tests will not.

The relationship between glaucony formation and sea-level change is another aspect of the passage of time. In the sedimentological record, the base of a transgressive sequence is frequently marked by a highly glauconitic layer: e.g., world-wide Cenomanian transgression (Odin and Hunziker, 1982); Lutetian transgression in the Paris Basin (Odin et al., 1982b); Holocene transgression on the Congolese continental shelf (Chapter Cl). Transgression results in many factors favouring glauconitization: e.g., supply of substrates, time needed for glauconitization to proceed. Following transgression, suitable sedimentary particles, such as shell debris, mineral grains, and faecal pellets which were produced and deposited at too shallow depths for glauconitization to occur, are at depths greater than -60 m where glauconitization may proceed. Moreover, transgression causes a decrease of sediment supply where these potential substrates are deposited by encroaching onto the continental land mass; thus the potential substrates are not rapidly buried. However, transgression is not necessary for glauconitization because other circumstances may combine favourable factors such as: accumulation of pelagic or benthic foraminiferal tests at the outer part of the continental shelf, during a biostasic period. Consequently, the identification of the original substrate of glauconitization may be a way to

330 determine the absence or presence plus the scale of a transgression phenomenon (Fig. 16). In this figure, the extent of the transgression is indicated by the length of the arrow (lower part of the figure) drawn between the depth of the usual zone of production of the substrates and the depth of the usual zone of glauconitization.

saline deposits

2 @

fibrous clays

2

L,

carbonates ox idates - - -_ _ - - - - - -

hydr olysates

coarse residues

LU

3

@

smectites

% -- _ 4

P 8

E

glauconitic minerals phyllite V + C - - -- _ _ - - - - CONTINENTAL CLA Y!

NO C L A Y

Figure 17. Location of the glaucony facies in the normal evolutionary sequence of Goldschmidt. A parallel sequence of clay minerals is proposed in the right hand column. Phyllite V and phyllite C are the authigenic green marine clay minerals of the verdine facies which are identified and described in the preceding part B of this volume (see also Glossary). A number of iron-rich clay minerals are formed within about the same member of the sequence. (Modified from Odin and Utolle, 1980)

Finally, using an even longer time scale, glaucony can be considered in terms of the classical evolutionary sequence of Goldschmidt as revised by Millot (1964). In this model (Figure 17), the normal sequence of five members involves epicontinental sedimentation in a shallow marine basin. The first two members (bottom left) correspond to coarse, and then fine grained detritus of continental origin brought and deposited into the basin; the last two (top left) correspond to the deposition of chemical products, successively carbonate and saline deposits. Glaucony mainly develops at the boundary between these first two and last two members, as part of the oxidate member. In comparison, phosphates can be placed later in the sequence but before deposition of chemical saline deposits. The ironstone and verdine facies, together with other iron oxides, also would be grouped with the oxidate member of Goldschmidt. This member could also be designated as the authigenic ferric clay member. Considering clay minerals, this authigenic ferric clay member would follow a detrital member without many clay minerals, and a second member rich in detrital clays. It would be successively followed by 1) Al-rich 2:l clays, and 2) Mg-rich fibrous clays (saline deposits) as suggested by Odin (1975a).

33 1 SUMMARY

This chapter summarizes present knowledge on the geology of the glaucony facies. As for any facies, glaucony can be described by four characteristics: habit as seen in the field or through the microscope; composition as analysed in the laboratory; origin as deduced from the above observations; and sedimentary environment in which it forms. The ability to observe glaucony in various stages of formation in the present-day Oceans prior to early diagenetic alteration has proven to be very helpful in developing an understanding of this facies. The habit of glaucony is very diverse, but it essentially comprises two groups: a granular and a film habit. The most common habit is the granular one. It is due to the fact that the green clay develops within a substrate which is usually granular itself. There are four subgroups of granular substrate: microfaunal infillings, faecal pellets, biogenic carbonate debris, and mineral debris. Within these substrates, glaucony apparently develops either by filling pores or by replacement. In fact, all habits have been shown to result from physically similar substrates: they all are more or less porous material, which undergoes alteration, and within which authigenic material develops. In this sense, infilling or replacement appear to be equivalent phenomena differing only in the scale of the voids where authigenic clay forms. No particular mineralogical or chemical character appears necessary for these substrates to be glauconitized. The general composition of glaucony results from the combination of two different materials: the substrate which will eventually disappear and which is very variable in composition, and the authigenic phase which replaces the substrate and which may undergo further changes as a result of recrystallization. The elimination of the substrate is very gradual and traces have been observed even at a very late stage of the evolution. The authigenic phase has a limited range of compositions, and forms a mineralogical family between smectitic and micaceous end members. All members have potassium as the main interlayer cation, plus a high iron content, mainly in the femc state. Earlier suggestions for the formation of glaucony (i.e., transformation of degraded layer silicate), have been shown to be incorrect. Instead of transformation, it has been observed that the fundamental geochemical reaction is the de novo crystal growth of smectite. This smectite appears to be iron-rich even when the surrounding sediment is iron-poor. This reaction takes only a few thousand years. Several recrystallization processes then occur. These appear mainly to be characterized by an increase in potassium content. Recrystallization (which occurs from ions previously gathered by the neoformed smectite) is accompanied by a significant amount of crystal growth causing an increase in the volume of the initial granular substrate which is now mostly disaggregated. This volume increase, more pronounced in the interior than in the exterior of the grains leads to the formation of the very characteristic cracks at the surface of the grains. Although the entire evolution occurs as a continuous phenomenon, we found it useful to identify four stages: nascent, slightly-evolved, evolved, and

332 highly-evolved. These have different geological significance resulting from the varying amounts of time needed to reach each of them. A final stage can be identified as relict when the grain has achieved complete evolution and remains in an environment still apparently favourable for glauconitization. The general synsedimentary evolution of the grains may be stopped by burial at any time during this evolution. After burial, additional recrystallization occurs allowing disordered crystal structures to become more ordered. Therefore, an early burial stage of evolution will immediately follow any of the nascent to relict stage reached at the sea-floor because the green clays are in an environment which is different from the one where they formed as soon as burial occur. The microenvironment within which the glauconitic green clays form is mainly characterized by semi-confinement which shelters both initial crystal growth and recrystallization from the aggressive diluting and oxidizing character of the typical sea-water environment. The semi-confiied environment allows the necessary cation exchanges to occur for the glauconitic minerals to form. Conditions with temperatures below 15°C and the pH around 8 are favourable. The Eh conditions are at the oxidation-reduction boundary which allows iron to be mobilized as Fe2+ from the environment, and stabilized mainly as Fe3+in the crystal structure. The general environment appears to be open sea-water. Glauconitization occurs near the sea-bottom where substrates are in direct contact with sea-water, and at a time when little or no demtal deposition occurs. It is also favoured during a transgressive period but this is not necessary. Compared to verdine facies, the glaucony facies characterizes deeper water and cooler conditions. The proximity of an iron-rich fluvial input to the sea is not required for glaucony, and the latitudinal distribution is much wider. In many ways, the process of formation is very similar for glaucony and verdine. Initial substrates providing a similar microenvironment are involved in both cases. Glauconitization first forms smectite b crystal growth, and K-rich minerals develop later; verdinization forms 7 phyllite V first, which later recrystallizes to other phases very soon during the synsedimentary process itself. Knowledge of conditions of formation and evolution of glaucony during synsedimentary time and early burial diagenesis is necessary to correctly interpret radiometric ages obtained from this material. A phenomenon of systematic inheritance af radiogenic isotopes has been pointed out. Discussion of the closure time of the geochronometer and possibility of inheritance of isotopes both indicate that K-rich glauconitic minerals (only) are appropriate for evaluation of the time of deposition of the glauconitic formation provided that no excessive geochemical alteration occurs later on.

x

333 Part D

THE CELADONITE-BEARING FACIES INTRODUCTION TO THE CELADONITE-BEARING FACIES The fourth part of this volume discusses a green clay occurring in volcanic rocks. There has been a renewal of interest for celadonite in recent years because it is now recognized as common in deep-sea basalts. Published and unpublished data from volcanics cropping out both on land and in oceans will be presented and evaluated. The present study supports the view that although occurring in volcanic formations, celadonite is essentially marine in origin and forms when lavas are covered by the sea. The opinion of the editor is that celadonite formation is more closely linked to sea-water than has been suggested in previous investigations. This is the reason why the subject is discussed in this volume. Celadonite is usually considered hydrothermal in origin, but it will be shown below that 1) its temperature of genesis is frequently lower than is usually accepted for true hydrothermal alteration, and 2) the fluids from which it crystallizes are primarily marine (not juvenile), although they are later modified by passage through fissures in basalts. Celadonite is commonly compared to and grouped with glauconitic clay minerals; it is emphasized in this part of the volume that these two sorts of clay minerals are distinct from morphological, chemical, and crystallographical points of view. Therefore, geologists should distinguish between these minerals. In contrast to the variety shown by glauconitic minerals, celadonite seems to be mainly a single homogeneous, dioctahedral, mica-like, potassium-rich clay mineral. Celadonite is often mixed with smectitic green clay (saponite); however, this saponite is distinctly different since it is moctahedral. Numerous new data are tabulated and discussed concerning the time of genesis of celadonite. Celadonite may form up to 20 Ma after the emplacement of the host volcanic rocks. Consequently, the alteration of oceanic basalts seems to be a very long term process; from a geochronological point of view, the radiometric age of celadonite will be a minimum age of emplacement of the host volcanics. Celadonite was known to many ancient civilisations. The editor's interest in this material was initially stimulated by a problem posed by archaeologists working on green Roman wall-paintings. A detailed study showed that the main green pigment was celadonite (Odin and Delamare, 1986). This portion of the volume will be presented as a monograph. Subchapters or sections were written by one or more authors, the names of which are indicated in the contents below. The editor acknowledges the particular help of P.D. Fullagar for reviewing the content and English language of this last part. The contents of this part follows next page.

334

CONTENTS of PART D *Introduction to the celadonite-bearing facies *Contents of part D

333 334

*Nature and geological significance of celadonite by G.S. Odin, A. Desprairies, P.D. Fullagar, H. Bellon, A. Decarreau, F. Frohlich and M. Zelvelder 1. Presentation of celadonite by G.S. Odin

337

2. Occurrence and geological setting of celadonite by G.S. Odin and A. Desprairies 2.1. Examples of celadonite-bearing outcrops 2.2. Morphological features of celadonite 2.3. Petrographic environment of celadonite

337 337 340 342

3. Mineralogical properties of celadonite 3.1. Physical properties of celadonite by G.S. Odin, A. Desprairies and M. Zelvelder 3.2. Physico-chemical properties of celadonite 3.2.1. Infra-red absorption spectra of celadonite and related minerals by F. Frohlich and M. Zelvelder 3.2.2. Mossbauer spectra of celadonite and related minerals by A. Decarreau 3.3. Chemical properties of celadonite by G.S. Odin and A. Desprairies 3.3.1. Major element analyses on large samples 3.3.2. Discussion of chemical analyses on macro-samples 3.3.3. Micro-chemical analyses for major elements 3.3.4. Trace and rare earth elements 4. Environment of formation of celadonite 4.1. Environment of formation of celadonite based on petrography by G.S. Odin and A. Desprairies 4.2. Environment of formation of celadonite based on mineralogy by G.S. Odin 4.3. '*O isotopic study on celadonite by A. Desprairies and G.S. Odin 4.4. Time of formation of celadonite by G. S. Odin, H. Bellon, P. D. Fullagar, A. Desprairies and M. Zelvelder 4.4.1. Radiometric dating applied to celadonite 4.4.2. Formation of celadonite in young oceanic basalts

345 353 353 362 365 365 368 369 374

375 376 379 382 382 392

335 5. Nature and geological significance of celadonite by G.S. Odin and A. Desprairies 5.1. General characteristics of celadonite 5.2. Comparison between celadonite and glauconitic minerals

393 393 395

6. Acknowledgements

398

This Page Intentionally Left Blank

337 Chapter D NATURE AND GEOLOGICAL SIGNIFICANCE OF CELADONITE by G.S. Odin, A. Desprairies, P.D. Fullagar, H. Bellon, A. Decarreau, F. Frohlich and M. Zelvelder PRESENTATION OF CELADONITE

Celadonite is a clay mineral, green, and iron-rich such as other green marine clays described up to this point in this volume. Mineralogically, celadonites are very similar to some of the glauconitic minerals. According to Hendricks and Ross (1941) the name celadonite was proposed in 1847 by Glocker. This material was previously called terre verte or terre verte de Ve'rone by many authors beginning in the 16th century (de Brignoli de Brunnhoff, 1820). The word celadonite comes from Celadon which was the name of a warrior in "The metamorphoses'' by Ovide. The original meaning was: the noisy, the resounding. The name Celadon was popularized by Honor6 d'Urf6 in his famous 16th century pastoral novel "L'AstrCe". In this novel, Celadon was a neoplatonic lover whose clothes were adorned with green ribbons. At the time this novel was published, the first Chinese porcelains were imported into France. This porcelain was an unknown pale-green and the colour was immediately named "celadon" (like the porcelain itself) by the French court after the fashionable hero. Therefore, in French, celadon means a soft-green colour which is slightly bluish. The term celadonite refers to a mineral which is usually blue-green in colour after crushing, and is found in volcanic formations. It has the crystallographical structure of the micas. Due to similar colour, properties and occurrence, celadonite has long been confused with chlorite. However, during the last ten years, celadonite has become more and more familiar due to its frequent occurrence in volcanics found in deep-sea drill cores. OCCURRENCE AND GEOLOGICAL SETTING OF CELADONITE

Examples of celadonite-bearing outcrops The terre verte of the 19th century authors had been recognized by Vitruve in his volume "De architectura" (1st century B.C.). He noted the creta viridis from Verona and from Smyrna (presently Izmir, Turkey). The first location must refer to the celadonite found in the Late Eocene volcanics from the Monte Baldo area, 30 km north of Verona in northern Italy (Odin and Delamare, 1986; Odin, Hernandez et al., 1986). This celadonite was successively called Baldogke (earth from Baldo) by De Saussure and Vkronite by DelamCtherie (de Brignoli, 1820; Delesse, 1848). These specific references suggest that celado-

338

4A

Verona0

Smyrna

CELADONITE commercial abundant t localized

c,

A

iS0.87

Figure 1. Some celadonite-bearing volcanics in Europe. 1) Cyprus (Cretaceous); 2) Monte Baldo (Eocene); 3) Val di Fassa (Triassic); 4) Bohemia; 5) Faeroe Islands (Eocene); 6) Rockall Plateau (Eocene); 7) Taypon and Wormit Bay; 8) Skye and Island of Rum; 9) Giant's Causeway and Antrim Mounts; 10) Teruel (Toarcian, Sierra de Javalambre); 11) Zonguldak; 12) Icelandic basalts; 13) Sicily; 14) V6ring Plateau (Eocene). Note that these Occurrences are either related to the Alpine domain or to the ridge formed by the opening of the North Atlantic Ocean.

nite from the Monte Baldo area was one of the notable occurrences in the Roman World as well as the best known location to the 19th century European scientists. Not far from Monte Baldo, celadonite can be found in other portions of the southern Alps, and especially in the Val di Fassa in the Dolomites where it occurs in Triassic volcanics. Another area where celadonite has been observed in large quantity and intensively studied is the island of Cyprus. There, the green clay in the Massif of Troodos has been quarried until recently for paint pigment (Bear, 1963; Desprairies and Lapierre, 1973). Considering the very large quantity of celadonite available and the relative ease of obtaining it, there is little probability that the Romans ignored this material. It is suggested that Smyma was simply the port where the celadonite from Cyprus was sold under

339 Table 1. Celadonite from oceanic basalts. 1) Atlantic; 2) Pacific; 3) Indian Ocean. Leg Site

Locality

Depth

Stratigraphy

Mineralogy

37-332

W . Azores ( 1 )

1800 rn

Pliocene

Celadonite and

37-334

W . Azores

2600 m

Mid-Miocene

Mixtures ( a )

51-417

Bermuda Rise ( 1 )

5500 rn

Aptian

Ce 1adonite :

51-418

Bermuda Rise

5500 rn

Aptian

K20

81-553

SW Rockall ( 1 )

2200 rn

Top Palaeocene

Pure celadonite

81-554

SW Rockall

2600 rn

Early Eocene

Celadonite and saponite mixture (')

17-164 to

Central Pacific ( 2 )

2300 rn

Barremian

Celadonite and

17-171

Central Pacific

5800 rn

Barremian

srnectite ( d )

34-321

Peru Trench ( 2 )

4800 rn

Late Eocene

Pure celadonite

35-323

Bellingshausen Plain ( 2 )

5000 rn

Palaeocene

Celadonite ( e )

70-506 to

Ga 1apagos ( 2 )

2700 rn to

0 . 5 to 3 . 0 Ma

Dark-green clay ( f )

70-510

Galapagos ( 2 )

2800 rn

0 . 5 to 3 . 0 Ma

25-249

Mozambique Ridge

2100 rn

Neocomian

Blue-green clay

26-250

and Basin ( 3 )

5100 m

Coniacian

hiixture ( g )

26-256

Wharton Basin :

5300 rn

A 1bian

Celadonite ( g )

26-257

off Australia ( 3 )

5300 rn

Albian

Celadonite ( g )

7-8% ( b )

the name of Earth of Smyrna (Fig. 1). To the present authors' knowledge, there is no other location in the world where celadonite could be obtained in such large quantities. Between the years 1926 and 1958,15 tons of celadonite have been exported each year from Cyprus (Bear, 1963). In addition to these two exceptionally rich occurrences at Monte Baldo and Cyprus, celadonite can be found in small quantities in numerous volcanic outcrops. In Europe, celadonite occurs in Tyrol in northern Italy (Levi, 1914), and in Sicily (Scherillo, 1938). Celadonite of Maegdefrau and Hofmann (1937) has the characteristic colour but its abnormally high A w e ratio and its continental occurrence at Vesuvius would need verification. The Terre de BohEme from Bohemia, now in Czechoslovakia (Delesse, 1848; Beutelspacher and Van der Marel, 1968), may correspond to celadonite; the identification of celadonite from Turkey (Bayramgil et al., 1952; T. Hugi, personal communication, 1985) is more certain. In the Faeroe Islands and Iceland, celadonite has been studied recently by Delmont (1985).Heddle (1879) reported on celadonite in Scotland and Ireland. In Spain, a small quantity of celadonite has been found in the Toarcian volcanics of the Sierra Iberica south of Teruel (Gautier and Odin, 1985; see more data p. 390-391, below).

340 In North America, celadonite has been found in New Jersey (Delesse, 1848), Arizona, California, and Nevada (Hendricks and Ross, 1941), and Washington (Wise and Eugster, 1964). Celadonite occurs in South America in Brazil as indicated by samples which we have seen in mineralogical collections (British Museum, London and Ecole des Mines, Paris). Celadonite has been reported from New Zealand by Buckley et al. (1978). Foster (1969) reviewed data on possible celadonite from the U.S.S.R. but their low iron content suggests that these samples may not be celadonite. All of the occurrences listed above are from continental outcrops. The deepsea drilling project has led to the recognition of celadonite in many submarine volcanic rocks. A summary of the main known celadonite-bearing submarine basalts is given in Table 1. Detailed data on samples from Leg 92 (South Pacific) where celadonite was found in Oligocene basalts, and from Leg 104 (North Atlantic) where pure celadonite is abundant remain to be published. There is no doubt that celadonite is present in many places where basalts (tholeiitic and non-tholeiitic) are altered in a submarine environment.

Morphological features of celadonite Based on older publications and recent numerous descriptions of celadonite occurring in oceanic basalts, plus our own observations, it is possible to describe celadonite in terms of five morphological categories: 1) filled vesicles; 2) replaced phenocrysts; 3) diffuse; 4) films; 5) veins. There are certainly intermediate forms between these categories. Vesicles which become filled with celadonite are usually between 0.1 cm and 1.5 cm in diameter. These vesicles may be spherical, lenticular, ellipsoydal or irregular. This type of celadonite is probably the most frequently observed both in continental outcrops (e.g., Tertiary of the Faeroes, Eocene of the Monte Baldo, and Early Tertiary of Washington) and in submarine oceanic outcrops. Replacement of phenocrysts by celadonite has also been described from both continental and submarine outcrops. Triassic volcanics of the Val di Fassa contain pseudomorphs after augite. In Cyprus, replacement of olivine and augite is locally present. Replacement of pyroxene by celadonite has occurred in samples from Leg 35-site 323 in the South Pacific (Kastner and Gieskes, 1976). Scherillo (1938) reports that there are pseudomorphs after olivine in Sicily. In the Tertiary of Washington, celadonite replaces plagioclase phenocrysts (Wise and Eugster, 1964). Augite, olivine and plagioclase are the minerals most frequently replaced by celadonite. Hendricks and Ross stated that hypersthene or groundmass can also be replaced. The formation of celadonite does not seem to be dependent on the composition of the phenocrysts . The diffuse type of celadonite may be present locally in some tuffs or very altered lava flows which are green in colour. For example, Humphris et al. (1979) described a prominent green weathered zone in the groundmass of hole 417A (oceanic basalt of the Bermuda Rise). However, it is usually difficult to identify such fine-grained minerals and celadonite can be confused with green illite or smectite (Bellon et al., 1986).

341 Films of celadonite are ubiquitous; they can be observed in thin sections (e.g., Icelandic volcanics and many deep-sea basalts). Frequently, celadonite forms a film around vesicles or any voids filled with zeolite, smectite, or calcite. Following alteration of the rock, these fillings can be obtained pure; although externally entirely green, X-ray diffraction patterns obtained after crushing show that they only contain a small proportion of celadonite.

Figure 2. Section of a hydrothermal clad pipe (after Delmont 1985). 1) external part of the pipe (basalt); 2) celadonite film in vesicles and as first layer in the pipe; 3) zeolite layer (stilbite); 4) brown clay partly filling the pipe; 5) interior of the pipe (void).

The most important source of anciently industrial and analytically useful celadonite is from veins. We have observed in the field veinlets less than 0.5 cm in thickness and several decimetres in length in the Monte Baldo area and in the Toarcian lavas south of Teruel (Spain). The same kind of veinlets, which are blue-green in colour, have been observed in oceanic basalts (A. Desprairies' samples from North Atlantic). De Brignoli (1820) described vertical veins cross-cutting horizontal volcanics in the Monte Baldo area. These veins, 3 to 4 cm in thickness, may be followed for long distances. Locally, they show much thicker lenses, from which a piece weighing about five hundred kilogrammes was once collected. Celadonites in vertical fissures in sheet flows on Cyprus have been studied by Staudigel et al. (1986). Most of the celadonite in this ophiolitic complex occurs in these veins, some of which approach one metre in thickness; replacement of olivine and augite phenocrysts by celadonite also occurs in the same massif (Desprairies and Lapierre, 1973). Some deep-sea cores exhibit celadonite-filled fractures cutting the basaltic flows (e.g., samples from legs 35 and 81). A variety of these veins from the Faeroes has been described by Delmont (1985) as 'hydrothermal celadonite-clad pipes'. These pipes are not fully filled with the green clay (Fig. 2). Similarly, Bayramgil et al. (1952) describe the structures containing the green clay as tubes, a few decimetres long and 4 to 7 mm in diameter. These features raise questions about the relationship between the localization and form of the celadonite and the formation of the volcanics.

342

Petrographic environment of celadonite Celadonite and the texture of basalts To our knowledge, celadonite has been recognized only in contact with basaltic rocks which have aporous texture. Some type of opening such as vesicles, cracks, hydrothermal pipes, fissures, dissolved phenocrysts, scoriaceous character, spaces between pillow lavas seems to be necessary in order for celadonite to form. Accumulations of celadonite have been related to an abundance of vesicles. In the southern Alps (Monte Baldo area), vesicles are generally observed in the upper 5 to 10 m of the 50 to 100 m thick volcanic formations. Similarly, Desprairies et al. (1984) observed that the abundance of celadonite is linked with the abundance of vesicles in the oceanic basalts of the Rockall Plateau (Leg 81) and the Voring Plateau (Leg 104). The vesicles or fissures are essentially localized at the top of the flow and are partly filled with celadonite, as are the voids of the breccia lying on the top of the flow. Delmont (1985) also observed that amygdules and vesicles are more abundant at the top and bottom of lava flows studied in the Faeroe Islands. Andrews (1980), studying Mid-Atlantic oceanic volcanics, observed that celadonite forms mainly in voids, veins, vesicles, and more rarely as an alteration of plagioclases. The observation of fissures open to the top of the volcanic pile so that nannofossil-rich sediments were introduced down into the volcanics is very significant in terms of the probable origin of the fluids from which celadonite forms. Celadonite and other secondary minerals It is extremely rare that celadonite occurs alone as a secondary product in volcanic units. Most recent descriptions of deep-sea volcanics and of volcanic rocks now exposed on continents show that celadonite is only one of many secondary minerals which commonly occur together. Included in this suite are other clays, nontronite and saponite, as well as calcite, silica, zeolites, and sometimes chlorite. These minerals may be intimately mixed in vesicles or they may occur together as replacements of phenocrysts (e.g., association chloriteceladonite in Val di Fassa, or zeolite-celadonite which occurs locally in Cyprus). Frequently, vesicles or other voids are filled with several clay minerals in a sequence which may differ from one void to another. For example, saponite may have crystallized either at the earliest stage, at the latest, or both. Some examples of mineral association may give an idea of the nature of the fluids from which celadonite forms. An association with silica was the first to be emphasized. De Brignoli (1820) deschbed a quarry in the Monte Baldo area in which veins of flint are usually honey-coloured. Miners frequently neglected the relatively minor quantities of celadonite in green vesicles or kidneys scattered in the rock in order to trace the veins of flint because they knew that the flint provided the best guide to find the terre verte in appreciable quantity. Desprairies and Lapierre (1973) noted the presence of silicified borders around celadonite lenses in Cyprus. Finally, Humphris et al. (1979) reported silica in polyphasic veins of oceanic basalts.

343

Figure 3. Scanning electron microscope view of a rock section in a basalt from Leg 81: relationship between saponite and celadonite. A portion is drawn to the right and shows five components from left to right: basalt, massive Fe-Mg saponite (white), mixture saponiteceladonite, saponite (Mg), and void (black) of the vesicle. The white bar at the bottom of the picture is 1 mm in length. (Photomicrograph by Desprairies)

Calcite and celadonite can occur in vesicles in the same volcanic unit, but these minerals are found in separate outcrops, perhaps separated by only 10 m or 100 m. W e have observed this situation in the Monte Baldo and Teruel areas. Hendricks and Ross (1941) also noted this association in basalts from New Mexico. W e suggest that the same general phenomenon produces vesicular celadonite and calcite (either in different locations at the same time or at different times in a given site) as indicated by the common association of these two secondary minerals in many volcanics in the field as well as in oceanic basalts. The association of several clays in a single formation is a general feature of most celadonite-bearing outcrops. Saponite and celadonite were described by Hendricks and Ross (1941) as occurring together in New Mexico and Nevada. In the Faeroes, Delmont (1985) emphasized the ubiquitous association of celadonite with a black ferriferous saponite (a smectite where octahedral sites are dominantly filled with divalent Fe and Mg). This black or brownish smectite has also been found to fill vesicles in the Monte Baldo area. Among others, Seyfried et al. (1978), Andrews (1980), Desprairies et al. (1984) have described the association celadonite-saponite in deep-sea basalts (Fig. 3). The colour of saponite varies: black, brown, light-green (Desprairies et al., 1984), pale-

344 green to olive-green (Andrews, 1980), or blue-green (Seyfried et al., 1978); these authors have found that the saponite could also be chemically variable: either rich in Mg, or in Fe, or in Al. The questionable mixed-layer clay minerals nontronite-celadonite (Buckley et al., 1978; Alt and Honnorez, 1984) or saponite-celadonite (Andrews, 1980) have also been reported. Seyfried et al., (1978) and Desprairies et al., (1984) identify a chlorite-smectite mixed-layer. Finally, celadonite is associated with various zeolites as in Cyprus (Staudigel et al., 1986). A sequence of formation of these different secondary minerals in volcanic rocks has been observed. Following the general lines published by Laverne and Vivier (1983), Bohlke et al. (1983) and Alt and Honnorez (1984), we suggest a five step scheme: 1) alteration with palagonitization and the formation of black halos, followed by silicification which begins to reduce the porosity of the rock; 2) short period of clay formation under oxidizing conditions, usually leading to films of saponite (e.g., thin sections of Leg 81 or Leg 104 basalts, and samples from the Faeroe Islands, as reported in 1985 by Delmont); 3) celadonite forms and greatly reduces the porosity of the rock; 4) second period of clay formation under less oxidizing conditions, resulting in the formation of smectites which contain more Mg and less A1 than those formed in step 2 (Fig. 3); 5) calcite and zeolite form, nearly closing the remaining pores of the rock. Following these alteration processes, the clay minerals, mainly smectites, may form at least 10 to 20 percent of the entire rock (Delmont, 1985; Humphris et al., 1979). A mixture of celadonite with stage 4 saponite, sometimes of green colour, may result from a rapid sequence of infilling of veins and amygdules. Intermediate chemical compositions of this material have sometimes been obtained using micro-analytical techniques. This need not imply the existence of minerals intermediate between celadonite and saponite. Following detailed study, our opinion is that the two minerals are different, and that they do not have an "interlayered' intermediate structure. Even though the iron content may appear similar in the two minerals, Mossbauer spectra show that the sites where iron occurs in the lattice are basically different in the two minerals (see p. 362365 below).

Geological setting of the host rock It is obvious that most of the mid-oceanic ridge basalts where celadonite has been observed were extruded in a marine environment. Some of them were emplaced in deep-sea conditions but some others were emplaced in much shallower (site 554) and subaerial (site 553) environments like on the Rockall Plateau and Faeroe Islands in Palaeogene time. Celadonite-bearing volcanic rocks which now occur on continents also formed in a marine environment. In the Teruel area, volcanics are bracketed with shallow water, ammonite-rich Toarcian limestones. In the Monte Baldo area, the limy sediments are neritic (very shallow facies) with large foraminifers both above and below the volcanic layer. In some outcrops of this area, and essentially at the top of the volcanics,

345 marine molluscs and large foraminifers may be found mixed with the upper 10 m of tuffs. In Cyprus, celadonite is found in volcanics where pillow lava structures are abundant and indicate subaqueous formation of the volcanics; their marine character is confmed by the presence of interlayered sedimentary lenses with radiolarians. In Sicily, the celadonite-bearing basalts occur with tuffs which contain marine fossils (Scherillo, 1938). From the above considerations, and the apparent lack of celadonite in strictly continental volcanics, it is suggested that the origin of celadonite in volcanic rocks is linked with their occurrence in a marine environment. Celadonite seems to crystallize significantly after extrusion of the volcanic formation. Consequently, it is suggested that celadonite genesis is probably not directly linked with submarine extrusion itself but instead requires the presence of the sea above the volcanic rocks at the time when celadonite forms. The presence of sea-water at the time when magma is extruded may physically or chemically induce the porosity within the rock (pillow lavas, breccia, cracks, fissures, and vesicles) which later will provide an adequate environment for celadonite formation. MINERALOGICAL PROPERTIES OF CELADONITE

Studies by Hendricks and Ross (1941), complemented later by Foster (1969) provide most of the available mineralogical information on celadonites. These references are especially useful for distinguishing between glauconitic minerals and celadonite.

Physical properties of celadonite Optical properties Knowledge of the optical properties of celadonites is potentially useful since in many outcrops which now occur on continents, and in oceanic basalts, celadonite is only available in small quantities. A thin section is thus an easy way to study the green clay. However, at present we do not have definitive optical criteria which always enable celadonite to be distinguished from evolved glauconitic minerals or, more important, from chlorite or green saponite. Therefore, a thin section study by itself cannot give sufficient information for the identification of a green clay. When celadonite is blue-green, the colour is quite diagnostic and a direct comparison with a reference section of an already identified celadonite may be helpful. When celadonite is green, chlorite may have a similar appearance. The crystallites of celadonite are sometimes oriented giving a fibrous-like structure. When oriented and blue-green, celadonite manifests a beautiful pleochroism: blue-green to green-yellow. Smear slides may help this observation. Fibroradiated or acicular structures are sometimes observed in vesicles (Kempe, 1974). The index of refraction is generally not diagnostic. The difference in facies is certainly the easiest way to distinguish sedimentary glauconitic minerals from volcanic celadonite. If no definitive method of

346

Figure 4. Smear slides of celadonite and glauconite (microscopic view, natural light). The black bars are 25 pm long. The celadonite to the left comes from a dark-green vesicle from San Valentino (Monte Baldo, Italy). The glauconite sample to the right is the reference material glauconite G L - 0 (Odin and coll., 1982) which has been gently crushed. (Photo. by Odin and Zelvelder)

investigation is possible, it is best to use the term green-cluy for description of the component when it is seen in a thin section. To give the name celadonite to a chlorite, or vice-versa, would introduce an error of interpretation of the rock due to difference of geological significance of these two minerals. It is possible to distinguish celadonite from chlorite with a microscope using a micro-sample extracted with a sharp pin. A smear slide of celadonite obtained after gentle crushing will typically show, at the highest magnification in oil, narrow rods or needles about 1 to 10 pm in length (Fig. 4); sometimes the needles reach 25 pm in length, much longer than observed by Beutelspacher and Van der Mare1 (1968) or by Sudo et al. (1981) using the scanning electron microscope. Chlorite will show larger and flatter blades; glauconitic minerals will show small blades or agglomerates which are much more irregular than chlorite. The recognition of celadonite using optical techniques is therefore possible. The acicular facies is not specific for celadonite; Mevel (1979), along with other authors, notes that the brown clay obtained from Legs 51 to 53 (saponite) commonly shows fibrous aggregates (spherulites).

347

Figure 5. Transmission electron microscopy of celadonite. On the left are two pictures of celadonite C2 from Cyprus; the black bars are 1 pm in length; the laths are much smaller and narrower than the ones from the Rockall Plateau shown to the right. The top sample from the Rockall Plateau was collected from a 5 mm wide vein from the lower part of the vesicular zone of a flow (Leg 81, site 553, core 47-1). The bottom sample is from site 553, core 40. (Pictures by A. Desprairies)

Electron microscopy It is easy to characterize celadonite using electron microscopy. Transmission electron microscopy (Fig. 5) will show more useful images than scanning electron microscopy (Fig. 6 ) due to the fact that celadonite readily shows individual laths after gentle crushing. Beutelspacher and Van der Mare1 (1968)

348

Figure 6. Celadonite viewed by scanning electron microscope. White bars are 10 pm in length; all samples are from the Monte Baldo area. Upper left comer: small vesicle, dark-green coloured, San Valentino; others: veins from near Brentonico (pictured by M. Zelvelder at the Mulum National d Histoire Naturelle). Note the exceptional thickness (about 0.3 pm) of some laths in the bottom right photomicrograph.

and Sudo et al. (1981) have obtained pictures of laths usually smaller than 1 p m in length. The laths are 2 to 5, sometimes 2 to 10 times longer than they are wide (Fig. 5). As noted above, in the Monte Baldo area we have observed laths reaching 25 pm in length; the thickness of these crystallites is only about 200 8, (Fig. 6). In the Faeroes, Delmont (1985) has observed very narrow laths from 0.03 pm x 0.6 pm to 0.2 pm x 1.5 pm. The fact that the microcrystallites are nicely shaped is probably related to very good conditions for crystal growth. Factors favouring the reaction of crystallization may be either a temperature higher than is usual for the sedimentary clays, or a relatively high ionic concentration. These must be kept in mind when we discuss conditions and environment of formation of this green clay.

349

I

A I

GLAUCONITIC M I C A

K2O*88

I . . 36. . . .32 . . . . . . . . . . 2 0. . . . . . . . . . 28

24

16

12

8

4'

Figure 7. Comparison of X-ray diffraction patterns for potassium-rich celadonite and glauconitic mica. The celadonite is from a vein at Monte Baldo; the glauconitic mica is from the Late Cretaceous of southeasteam Fmce (Conard et al., 1982). When similar conditions of diffraction and recording are used, the difference between the two minerals may be observed by looking at 1) the width of the 001 peak; 2) the width of the 003 peak; 3) the width of the 112 and 112- peaks; 4) the shape of the 111- and 021 peaks; 5) the relative height of peaks 023 and 130. (After Odin and Delamare, 1986)

We also note that the microcrystallites are small compared to many other clays, and especially to the most evolved glauconitic minerals. We suggest that this might be related to the occurrence of very favourable conditions of growth only during a very short period of time compared to the time needed for crystallization of glauconitic mica.

X-ray diffractometry (XRD) As for glauconies, clays extracted from volcanic rocks, and celadonite especially, are best characterized using randomly oriented mounts (a slurry of clay and acetone is deposited on a glass slide and evaporated in a few seconds). Celadonites collected from different outcrops in the Monte Baldo area give good X-ray diffraction patterns. Very high and narrow peaks are obtained. Comparison with a glauconitic mica allows us to emphasize a number of specific characteristics (Fig. 7).

350 To make this comparison, both minerals must have been studied using similar conditions of X-ray emission and data recording . The positions of the peaks are similar for the two minerals, but the shapes of the peaks are really very different. The ratio of height-to-width (at mid-height) is 2 to 5 times greater for the 001 peak of celadonite compared to glauconitic mica. The peak for celadonite appears, therefore, much sharper. A similar situation exists for the 003 peak as well as for the two peaks on each side of 003: 112 and 112-. The two small peaks following 020 (111- and 021) are more clearly separated from background for celadonite than for glauconitic mica. The diffractogram of glauconitic mica shows that the 111- peak begins to be drawn before the 020 peak is completed, while the 021 reflection in turn begins before the 111- peak is completed (at a chart speed of 1"8per minute). All the above criteria, as well as slightly higher 002 and 023 peaks for celadonite, indicate that the celadonite has better defined reflections than those shown on a glauconitic mica diagram, and, therefore, a much more regular (ordered) crystallographical structure. This observation confirms that the conditions of crystallization are generally more favourable for celadonite than for glauconitic mica. Table 2. Comparison of celadonite and glauconitic mica. Our values are based on the data from Figure 7. Diffraction This work Wise & Eugster (1964) This work Warshaw (1957) Glauconite Glauconite Celadoni te Celadoni te hkl 100

47

100

100

002

10

-

5

0

020

40

85

40

80

iii

15

42

10

20

021

15

37

10

10

115

40

80

40

40

60

70

50

60

112

40

80

40

40

115

5

10

5

5

023

30

75

15

10

130...

35

100

40

100

30

75

30

60

001

003

...

132 ~

Finally, there is another criterion which allows a distinction between 10 A green clays. The ratio of heights of reflections 023 versus 130..., is noticeably higher, and sometimes reaches one in diagrams for celadonite compared to diagrams for glauconitic mica where the ratio is constantly lower than 0.2.

35 1 This summary shows that there are definite ways to distinguish between natural glauconitic mica and celadonite using the randomly oriented powder diffractogram technique (Odin and Delamare, 1986). However, after superficial weathering, the X-ray diagram of celadonite tends to become similar to the one obtained from glauconitic mica.

Discussion of physical properties of celadonite Based on the above observations, it should be easy to distinguish natural celadonite from natural glauconitic minerals. Toward this end, electron microscopy and X-ray diffractometry are the best techniques. Other physical properties are of little use. The density of celadonite of the Monte Baldo has been measured by de Brignoli (1820) and by Delesse (1848); they obtained ~, Celadonite from the Tertiary of values of 2.83 and 2.907 g . ~ m -respectively. Washington was measured at 2.95 to 3.05 g . ~ m by ' ~ Wise and Eugster (1964). These values are similar to those of evolved glauconitic grains. Paramagnetism of celadonite is equal to that of glauconitic grains. X-ray diffraction patterns are not always as easy to interpret as is the pattern from the Monte Baldo sample or other potassium-rich and pure celadonites (Fig. 7). Although this pure, well-crystallized and dark-green celadonite is common in nature, as opposed to the relatively rare glauconitic mica, there are examples where the celadonite shows less well-developed peaks for diverse reasons. Some celadonites are intimately mixed with other minerals which relatively diminishes the height of the peaks of celadonite. The main change is that the 001 peak may decrease in height without changing width. In other cases, the 0 0 1 peak, with a top still at 10 A, becomes enlarged at the base towards the smaller angles; this indicates layers thicker than 10 A. The other peaks also become smaller than for well-crystallized celadonite but their width never reaches that of a glauconitic mineral (Fig. 7). On the contrary, the peaks remain clearly recognizable above the background. We have also observed that the ratio of peaks 023 versus 130 becomes much smaller than one following alteration. The ratio usually remains higher than that obtained for all glauconitic minerals, however. We emphasize that celadonite samples have never exhibited the diffractometric series shown by glauconitic minerals which range from a smectite type to a mica type. In the great majority of natural glauconies, selection of different fractions allows us to observe to a certain degree this series of minerals (see Part C of this volume, Fig. 7, p. 310; Fig. 9, p. 313). For celadonites from a single outcrop, variation in X-ray diagrams is almost never observed. This is considered meaningful. It seems that the mineral we observe today formed in a single episode, probably short in time. This may also be interpreted as reflecting particularly favourable thermodynamic conditions which, during a single short episode, directly crystallized the well-ordered potassium-rich mineral. This remark reinforces the similar suggestion above regarding the form of well-shaped crystallites. X- ray diffraction patterns intermediate between mica and smectite have not been found, but it is not rare to obtain an X-ray diffraction pattern of smectite in

352

K 32

28

24

20

16

12

8

Cu 4

Figure 8. X-ray diffraction patterns of various celadonites. 1) Celadonite vein, Cyprus (dark-green portion of sample C2); 2) Same sample as 1) above (light-green fraction); 3) Celadonite vein, Cyprus (with admixture of zeolite); 4 ) Celadonite film, geode from Teigahorn, Iceland (with admixture of zeolites); 5 ) Celadonite replacement of augite (replacement also includes quartz (Q) and chlorite (C).The association celadonite-chlorite is not rare (e.g., Teruel area, Val di Fassa, Iceland). Compare this figure to Figure 7.

the volcanic green clays. However, in that case, it appears that this represents another phase, distinct from the celadonitic one. Similarly, the apparently homogeneous green pigment collected with a pin from a basalt may contain a mixture of celadonite with chlorite (aluminous or ferric), or with various zeolites, as well as with calcite.

353 Physico-chemical properties of celadonite We have gathered under this heading physically measured properties which

are partly due to the chemical composition of the mineral. Infra-redabsorption spectra of celadonite and related minerals Infra-red spectroscopy mainly provides information on the composition of the octahedral sheet of the layer silicates. The sensitivity of the IR response to cation changes in the hydroxyl-cation bonds makes clay mineral spectra useful for identification of similar minerals. Table 3. Assignment of the absorption bands for celadonite and glauconitic minerals in the region of high and medium kquencies. Assignment w A13+.Mg2+-OH

Celadonite 3603*

w A 13+. Fe2+-OH

3577-3575*

w Fe3+. Fe3+-OH

absent

w Fe3+. Mg2+-OH

3558

w Fe3+. Fe2+-OH

3535

w H20

3400**

Tetrahedral sheet ( w Si-0)

6 A13+.Fe2+-OH 6 A 13+.PAg2+-OH 6 Fe3+. Fe3+-OH 6 Fe3+. Mg2+-OH

1110

G lauconitic minerals 3604*

absent 3560 3534-3544** 3440**

absent

1070-1080

1070-1 080

955-970

990-1025

absent

880*

840*

835*

absent

815-818

800

absent

680

679

or/and 6 Fe3+.Fe2+-OH

*sometimes lacking; **broad; V: stretching vibration modes; 6: bending vibration modes

Infia-red spectra of celadonites fi-om the Monte Baldo (Italy), Nevada, and Oregon (U.S.A.) are given by Van der Mare1 and Beutelspacher (1976). A more detailed study has been undertaken on the celadonites from the Faeroe Islands by Delmont (1985). -Procedure A few milligrammes of fiie-grained powder (grain-size < 2 pm) are requi-

354 red for analysis of materials like minerals or rocks. Great care should be taken in the particle size control after grinding, in order to perform each analysis under exactly identical conditions (Frohlich, 1981; Pichard and Frohlich, 1986). The samples are mechanically ground in a volatile liquid (e.g., acetone), next mixed and diluted with KBr (0.25%), and then converted into a disc (300 mg) by pressing in a die under vacuum. The IR spectra shown here were obtained using a Pye Unicam SP 2000 spectrophotometer (resolution better than 2 cm-' at 3000 cm-'). -1nfra-red spectra of celadonite Within the range 3700-3500 cm-', the IR spectra of celadonites exhibit very sharp and intense OH-stretching absorption bands (Fig. 9). Such bands, with a half-width (width, in cm-', of a band at half absorbance) of less than 10 cm-', are quite exceptional in the dioctahedral clay series where this half-width is usually about 90 cm-' for the OH-stretching bands. Natural celadonites exhibit s h T OH bands (Fig. 9) at 3558 cm-l and 3535 cm-l, and sometimes at 3603 cm . In addition, there is a wide OH-stretching band between 3400 cm-' and 3500 cm-l due to adsorbed H 0 (Table 3). Farmer (1974) mentioned that l absorption at 3577 cm-'. This natural celadonites have a fourt?OH-stretching band, which is sometimes seen at 3580 cm-' (Slonimskaya et al., 1986), or at 3578 cm-' (Buckley et al., 1978) or near 3585 cm-' (Delmont, 1985), does not BM 1921

\\ m

0

m

(D

C446

i' 6m

n 0) n m

cm -1

36m

3400

Figure 9. Infra-red spectra of four celadonite minerals in the region of the OH-stretching frequencies; 0.75 mg celadonite analysed.

355

appear on the spectra collection of van der Mare1 and Beutelspacher (1976). Absorption modes of the tetrahedral and the octahedral sheets as well as OH-bending vibrations are responsible for the absorptions in the 1200 cm-' to 400 cm-' range (Farmer, 1974). Compared with the other dioctahedral minerals, spectra of celadonites have absorption bands which are more numerous and sharper (Fig. 10). Most bands are considered to be constant in absorbance and in frequency. The band at 1070 cm-' is sometimes poorly resolved from the strong band at 1110 cm-'. In these conditions, the strong band at 955 cm-'

Figure 10. Infra-red spectra of celadonite minerals (same as for Fig. 9) in the low frequency region; 0.75 mg analysed.

356 shifts towards higher frequencies (965 cm-' to 970 cm-'). The absence of the 840 cm-' absorption correlates with better resolution of the s ectra. As this 840 cm-l absorption decreases in intensity, the band at 492 cm- shifts towards 488 cm-'. Finally, the two absorptions at 800 cm-l and 840 cm" seem to be inversely related (Fig. 10); if one is strong, the other is weak, and vice versa. -1nfra-red spectra of glauconitic minerals The IR spectra of glauconitic minerals exhibit a few absorption bands. These bands are broader and less intense than for celadonite. In the OHstretching region, only two strong bands are seen, near 3540 cm-l and at 3560 cm-l, with a third subsidiary, weaker band at 3604 cm". The f i s t two bands are poorly resolved or even unresolved Fig. 11). The intensity of the H20 band (OH-stretching mode) near 3430 cm-(I is extremely variable. In the region of the spectrum from 3400 cm-' to 3700 cm-', glauconitic minerals can be unambiguously distinguished from celadonites.

P

Table 4. Low frequency absorption bands for celadonite and glauconitic minerals. Celadonite

Glauconitic minerals

(748)

(763)*

absent

660*

492-488

495-489

46 0

452

436

434

387

(389)*

(355)

absent

345

33 5-345

320

absent

290

(288)

* sometimes lacking; (748) weak absorption. The difference between celadonite and glauconitic minerals is also obvious in the low frequency part of the spectrum (Fig. 12). The strong absorption near 960 cm-' for celadonite is seen at 990 cm-' for glauconitic minerals. The rather sharp band at 1 1 10 cm-' for celadonite is absent on spectra of glauconitic minerals. The 1080 cm-' band occurs for both celadonite and glauconitic minerals, and is very variable in intensity. This band is sometimes not distinct from the 990 cm-' maximum on the spectra of glauconitic minerals. In this case, as seen by Manghnani and Hower (1964), the maximum of absorption shifts from 990 cm-' towards 1010 cm-', and even 1025 cm-' (Table 3). The weak band at 748 cm-l for celadonite is usually lacking on spectra for glauconi-

357 tic minerals. However, another band is seen near 763 cm-' for the latter. The strong and sharp band at 679 cm'' for celadonites is less intense and is broadened on the spectra of glauconitic minerals. Below 500 cm-', celadonite and glauconitic mineral spectra are similar although all absorptions are weaker for glauconitic minerals (Table 4). In summary, compared to celadonite, glauconitic minerals exhibit a general broadenin of the absorption bands, some shifting in frequency, absence of the 3577 cm- , 1110 cm-', 800 cm-' and 320 cm-' bands, and a stronger H 2 0 absorption near 3420 cm-'.

F

Figure 11. OH-stretchingabsorptions of glauconitic minerals; 0.75 mg analysed.

-Band assignment IR spectra of natural celadonites have been widely described and discussed in an outstanding book by Farmer (1974). The sharpest absorption bands obtained from celadonite, as for the other layer silicates, are in the 3400-3700 cm-' region. These are due to stretching vibrations of the hydroxyl groups. According to Farmer (1974), these frequencies strongly depend on the nature of the octahedral cations. The occupancy of the octahedral sites by various trivalent cations usually gives rise to a single, broad and undifferentiated absorption band. In the case of celadonites, the two to four sharp bands are related to a coupling between trivalent and divalent cations with the hydroxyl groups. Each coupling of cations should produce stretching and bending vibration modes at specific frequencies. According to Farmer (1974 the vibration A13+.Mg2+-OHis at 3603 cm-'; A13+.Fe2+-OHat 3577 cm-'; Fe +.Mg2+-OHat 3558 cm-' and Fe3+.Fe2+-OHat 3535 cm-'. Recently, Slonimskaya et al. (1986) proposed new assignments based on deconvolution of the spectra. This model implies coupling between

b

358

Figure 12. Infra-red spectra of glauconitic minerals in the low frenquency region; 0.75 mg analYSed.

trivalent and divalent cations, but also between divalent cations (Mg2+.Mg2+; Fe2+.Fe2+)and between trivalent cations (A13+.A13+;Fe3+.Fe3+).These new computed bands do not appear on the expanded spectra (Fig. 13) and the occurrence of R3+.R3+-OHand R2+.R2+-OHcoupling is not consistent with the details of absorption in the region of the OH-bending frequencies on celadonite spectra and with the OH-stretching and OH-bending absorptions of other dioctahedral minerals where these bands are well known. The strongest bands of the celadonite spectra between 950 cm-' and 1110 cm-I are related to vibrations within the tetrahedral sheet. The absorption at 1110 cm-' should correspond to a Si-0 perpendicular vibration, and the two bands near 1080 cm-' and 960 cm-l are assigned to in-plane Si-0-Si stretchin modes (Farmer, 1974). The Si-0 bending vibration mode is seen at 460 cm- . The other absorptions seen between 500 cm" and 200 cm-' on the spectra of celadonite are not well understood. Between 540 cm-' and 400 cm", Si-0 bending vibrations contribute to the absorption, but the specific frequencies are

B

359 greatly influenced by the composition of the octahedral sheet (Farmer, 1974). The OH-bending modes are obvious on the celadonite spectra in the 950 to 600 cm-' region (Fig. 10). The absorption bands due to the octahedral sheet are well known and give an accurate picture of the octahedral site occupancy. Farmer (1974) assigns the 840 cm-l band to an A13+.M 2+-OHbending mode, and the 800 cm-' band, to both Fe3+.Mg2+-OHand Feg.Fe2+-OH. The strong absorption at 680 cm-', characteristic of Fe-Mg dioctahedral micas, is assigned by Farmer (1974) to an out-of-plane bending mode of the hydroxyl group.

Figure 13. Expanded spectra of glauconitic mica (G563A) and celadonite (BM 1921) in the region of the OH-stretching frequencies. Transmission scale is xl; wave-number scale is x10; compare with Figures 9 and 11.

Celadonite and glauconitic minerals are both Fe-Mg dioctahedral micas. The absorptions of the Si-0 and Si-0-Si vibration modes are nearly identical; however, most OH frequencies differ. Farmer (1974) has not identified these bands for glauconitic minerals. Slonimskaya et al. (1986) give similar interpretations for both celadonite and glauconitic OH-stretching bands, but their computed bands do not appear on the expanded spectra of glauconitic minerals (Fig. 13). The latter exhibit an A13+.Fe3+-OHbending mode at 880 cm-' (Fi . 12) as do other dioctahedral minerals (Farmer, 1974). The weak 835 cm- absorption shown by glauconitic minerals (Fig. 12) is probably related to the 3604 cm-' band, and as for celadonite, is to be assigned to an A13+.Mg2+-OHbending vibration. Absorption bands at 815-835 cm-l and 3560 cm-' are present on the spectra of glauconitic minerals and this is a common property with nontronite spectra (Fig. 11; van der Mare1 and Beutelspacher, 1976; Buckley et al., 1978).

P

360 According to Buckley et al. (1978) the 818 cm-' band could be assigned to an Fe3+.Fe3+-OHbendin mode. The 3560 cm" band seems to be related to both Fe3+.Fe3+-OHand Feii'.Mg2'-OH stretching modes. The coalescence or poor resolution of the 3560 cm-' and 3540 cm-' bands on the spectra make interpretation difficult. The 3540 cm-' band, sometimes shifting towards 3534 cm-l, is likely due to Fe3+.Fe2+-OH.When the glauconitic spectrum exhibits a single band in the OH-stretching region with a maximum of absorption centered at 3540 cm-', the 815 cm-' and 680 cm-' bands are still well defined. However, a new weak absorption occurs at 660 cm-', apparently because Fe3+, Fe2+and Mg are likely to be less ordered in the octahedral sheet for glauconite than for celadonite. This disorder induces the broadening and coalescence of the absorption bands, and thus the OH-stretching absorptions cannot be distinguished. As for celadonite, the 680 cm-' band for glauconitic minerals should be assigned to the out-of-plane bending vibration of the hydroxyl groups which are in coordination with Fe3+and Mg cations. -Infra-red spectra and the crystal chemistry The shifting of the Si-0 stretching band from 955 cm-' for celadonite towards higher frequencies, and up to 1025 cm-' for some glauconitic minerals, is likely caused by Al-for-Si substitutions in the tetrahedral sheet (Manghnani and Hower, 1964). Wise and Eugster (1964) also report that for the dioctahedral mica series, the Si-0 stretching bands (at 1110 cm-', 1075 cm-' and near 960 cm-l) are directly influenced by the substitution of trivalent cations for silicon atoms in the tetrahedral sheet. Tetrasilicic celadonites (four Si atoms for four tetrahedral positions) are considered one end-member of the Fe-Mg dioctahedral mica series, with the best resolved Si-0 stretching bands (Fig. 10) and the lowest Si-0-Si stretching frequencies. These frequencies range from 955 cm-' (well-resolved Si-0 stretching bands) to 970 cm-' (poorly-resolved bands) which must denote additional A1 in the tetrahedral positions. At the other end of the mica series, with a single Si-0 stretching band at frequencies ranging from 980 cm-' to 1025 cm-', glauconitic minerals are close to trisilicic dioctahedral micas. Therefore, as indicated by Si-0 stretching bands, celadonite and glauconitic minerals are separate species. The vibrations of K+ interlayer cations, give rise to absorption bands in the far infra-red, near 100 cm-' (Farmer, 1974), but usual spectrophotometers do not allow observations at such frequencies. Although A1 substitutions in the tetrahedral sheet of glauconitic minerals give rise to a shifting of the Si-0 frequencies, there is no significant relationship between the potassium content and the tetrahedral Si-0 stretching frequencies in these minerals (Fig. 14). In the octahedral sheet of celadonite, one trivalent cation and one divalent cation occupy two of the octahedral sites and the third is vacant. This R3+.R2+ coupling gives rise to very sharp absorption bands. Absorbances have been calculated for OH-stretchin bands (3500 cm-' - 3600 cm-') and OH-bending bands (600 cm-' - 900 cm- ) for the four celadonite spectra represented in Figures 9 and 10. Figure 15 displays good linear relationship between bending and stretching bands. The data points obtained for each cation pair (A13+.Mg2+;

5

361 1025

low

h

L

3

.-. ¶SO

950

r

01

h

*

OH

362 Mossbauer spectra of celaabnites and related minerals The samples studied are secondary minerals which fill vesicles and veins in altered basalts recovered from holes drilled during D.S.D.P. Leg 81.XRD data indicate that the samples contain a smectite (trioctahedral)-mica (dioctahedral) assemblage. Both minerals are well-crystallized. Based on micro-analyses (Table 5 ) the smectites are Fe-rich saponites and the mica is celadonite. Pure saponite samples can be obtained by hand-picking. In these rocks, the celadonite samples always contain, in more or less small proportions, a smectite phase which could be either saponite or beidellite (Desprairies et al., 1984). Four samples of saponite-celadonite assemblages from four independent vesicles and cracks were studied by Mossbauer spectroscopy. Mossbauer spectra were recorded at room temperature and then fitted using a least-squares computer program assuming Lorentzian line shapes. The Mossbauer spectra of the two minerals are different. Those of saponite contain two broad doublets

A

V 4

-2

0

2

4

I

I

I

I

I

Velocity (mm/s)

Figure 16. MCissbauer spectra of celadonite (A) and saponite (B).

363

A

B

54.8

51.5

A1203

1.2

Fe203

Si02

B

A

Si

3.87

3.44

5.3

A1 IV

0.1

0.42

14.8

14.7

A1 V1

3.97

3.86

FeO

6.8

2.2

Fe3+

0.78

0.74

MgO

7.8

16.2

Fez+

0.4

0.12

CaO

0.2

1.6

Mg

0.82

1.61

Na2O

2.9

7.8

Ca

0.02

0.06

0.7

Na

0.39

1.01

K

1.03

0.11

11.5

K20

100

Tota 1

100

Table 6. Miissbauer parameters of coexisting saponites and celadonites. I.S. = Isomer shift (relative to iron metal); Q.S. = quadrupole splitting; r = peak width (all values in mm/s); s = relative intensity. Samples from zoned vesicles in D.S.D.P., Leg 81 basalts, Site 553 (1) 47-1,33-35 cm (2) 41-3,3-6 m (3 and 4) 40-1, 15-18 cm (Desprairies et al., 1984). Fez*

Fe3+ Sample Mineral Sap.

Site

I.S.

Q.S.

r

s

M1

0.394

1.286

0.576

0.53

M2

0.346

0.653

0.60

0.33

0.369

0.385

0.368

0.66

I.S.

Q.S

r

1.12

2.615

0.36

0.14

1.087

1.814

0.38

0.23

1.12

2.528

0.49

0.11

1.096

2.548

0.353

0.16

1.055

2.04

0.328

0.33

1.133

2.624

0.521

0.14

1.132

2.652

0.321

0.14

1.263

1.894

0.507

0.15

1.132

2.652

0.327

0.05

2.63

0.35

0.14

0.143

1 Cel.

Sap.

M1 M2 M1

0.388

1.261

0.541

0.49

M2

0.343

0.58

0.612

0.35

M2

0.368

0.356

0.333

0.53

Ml

0.393

1.297

0.571

0.60

M2

0.33

0.771

0.47

0.26

0.354

0.361

0.357

0.80

0.34

0.16

2

Cel.

Sap.

M1

0.47

0.14

3 Cel.

Sap.

M1 M2

M1

0.38

1.18

0.60

0.53

M2

0.38

0.48

0.40

0.33

1.02 1.067

1.991

0.63

0.43

0.341

0.337

0.373

0.43

1.131

2.632

0.34

0.14

0.20

0.14

4 Cel.

MI M2

0.57

364

A

.O

0 OH A Fe3+

Figure 17. Schematic representation of iron-site occupancy in octahedral layers of (A) saponite, (B) celadonite. Other octahedral atoms are not shown and the actual atomic ratio in M1 (trans) and M 2 (cis) sites is not correctly drawn.

characteristic of octahedral Fe3+,and one doublet corresponding to octahedrally coordinated Fe2+(Table 6, Fi 16). The celadonite spectra show only one narrow doublet attributed to Fe5+in the octahedral site, and two doublets for octahedral Fe2+.No detectable amount of iron in tetrahedral coordination was found for all the studied samples (Bancroft, 1973; Coey, 1980). The atomic ratio Fe2+/Fe2++Fe3+is higher in celadonite than in saponite. Saponites are trioctahedral minerals. For these samples, the octahedral sheet contains about 2.5 cations per formula unit (Table 5), and both M1 and M2 sites are occupied (Fig. 17). According to previous studies (Coey, 1980; Heller-Kallai and Rozenson, 1981) the Fe3+doublet with the smaller quadrupole splitting (Q.S.=0.6 m d s ) in saponite spectra is attributed to Fe3+ions in the M2 site, while the other Fe3+doublet corresponds to Fe3+in the M1 site; Fe2+ions are located in the M2 sites. Celadonites are mainly dioctahedral (about 2 octahedral cations per formula unit) and have few substitutions in the tetrahedral sheet (Table 5). In Mossbauer spectra for glauconitic minerals and celadonite, assignments of doublets are conflicting (Kotlichi et al., 1981). All authors attribute the Fe3+doublet (Table 6) to ferric iron in M2 sites (Dyar and Bums, 1986). The outer Fe2+ doublet with larger quadrupole s litting is usually assigned to ferrous iron in M2 sites, the inner doublet to FeE in M1 sites; however, other authors (Hogg and Meads, 1970; Mc Conchie et al., 1979) propose the opposite assignment. Despite this unclear situation, the celadonites analysed are characterized by Fe3+ ions located in M2 site only and Fe2+ions distributed in both M1 and M2 sites.

365 Therefore, iron site occupancy is quite different in saponites and celadonites collected from Leg 81 basalts. A conversion or transformation of saponite into celadonite by progressive interlayenng has been suggested by Andrews (1980) and Andrews et al. (1983). This hypothesis would seem to mimic the expenmental formation of micas from vermiculite by fixation of potassium or cesium (Sawhney, 1967). However, in the case of the celadonite-saponite assemblages under study, such a transformation involves com lex, non-realistic, changes in iron site occupancy, such as transition of Fegions from the M1 to the M 2 site, and partial in situ reduction of these iron atoms. Consequently, if a conversion of saponite to celadonite occurs, it implies a dissolution-reprecipitation mechanism and there are no intermediate minerals.

Chemical properties of celadonites

Major element analyses on large samples In evaluating chemical data available for celadonite, it is necessary to distinguish between data obtained from relatively large samples (several grams) and those obtained by micro-chemical techniques. Data obtained on large samples are generally more precise and accurate by a factor of 10 compared to micro-chemical analyses. Macro-samples also allow us to determine the degree of purity of the sample before analysis, and therefore be more confident in the results obtained. Furthermore, macro-samples may be more representative and thus may give better information on the conditions of formation and significance of the facies. However, published analyses of macro-samples sometimes represent marginally useful data on unknown mixtures of minerals because no parallel physical study was published. With these problems in mind, we have tried to select useful and representative data from the literature (Table 7). The first useful summary of analytical data for celadonites was by Hendricks and Ross (194 1) who compiled 10 analyses. They reported three of the four very similar analyses published by Heddle (1879); these were obtained from one Irish and three Scottish celadonites. There was no physical mineralogical control available at that time. Saponite may be present in the Scottish samples as Heddle describes apple-green saponite in the same formation. Moreover, the same Scottish celadonites all show a high weight loss after heating (more than 10.4°/~).Two of them have very abnormal low Fe3+/Fe2+ ratios and high Mg contents (8.3 and 8.5%). Finally, the three Scottish celadonites have an apple-green or lighter colour and their specific gravity is low, between 2.574 and 2.598. Thus we will not consider the Scottish celadonites any further. The Irish celadonite from the basalt south of the Giant's Causeway appears to have higher specific gravity (2.63). normal water content, and brilliant dark apple-green colour. We will accept its analysis as representative of a celadonite (Table 7). The celadonite from Madagascar (Lacroix, 1916) included in the report by Hendricks and Ross is also mineralogically very dubious. Maegdefrau and Hofmann (1937) reported data for celadonite from Vesuvius. They included thermal analyses and X-ray diffraction data. There are, however, problems with this "celadonite"; 1) the X-ray pattern shows an aspect similar to that of a glauconitic mineral: the

366 023 peak is very small in spite of high K 2 0 (9.38%); 2) there is no other report of macroscopic samples of celadonite in the Vesuvius area to our knowledge; 3) the lavas from Vesuvius are continental and were never beneath the sea and 4) the A1203/Fe203ratio is very unusual (10.9/6.95, greater than 1). The fourth problem may have an easy explanation: in the same table of results, Maegdefrau and Hofmann (1937) give an analysis of a normal glaucony from New Jersey which also has an unusual A1203/Fe203ratio: 19.30n.52. We suggest that these A1203/Fe203values have been reversed in their table. We have good confidence in the data obtained for celadonites from Monte Baldo (Brentonico) by several authors because all of our mineralogical analyses have shown that pure celadonite is easy to obtain from that area. Our own new data confirm the earlier results. According to Hendricks and Ross (1941), the two analyses proposed by Wells (1937) were also apparently obtained from pure well-crystallizedceladonites. Eighteen previously published analyses are given by Wise and Eugster (1964), but they were obtained from older literature which did not include a discussion of the corresponding mineralogy. We will, however, tabulate the original analysis of celadonite from Washington given by these authors because the material was carefully examined prior to the chemical analysis. The X-ray diagram shows a typically high 023 peak. Foster (1969) reported 13 analyses, but mineralogical data were lacking especially for the results compiled from Russian authors. We include the analysis from Scherillo (1938) for celadonite from the basalt of Giuliana (Sicily). This sample was incorrectly reported to be from Contessa Entellina by Foster (1969). Hendricks and Ross (1941) reported analytical results from the same reference which do not correspond to the analysis published by Scherillo. Although no precise mineralogical details are available, the description by Scherillo is reasonably good. He cleaned the sample with dilute acetic acid in order to remove calcite without modifying the celadonite. This Sicilian celadonite probably provides the least reliable data in our table since the basalts are said to be very weathered, and this may have modified the celadonite. Celadonite from Turkey (Bayramgil et al., 1952) has an X-ray diagram with a 001 peak at 10.4 8, and a very small 023 peak;a relatively high 113 peak is also present. This mineral is probably weathered because weathering provokes an enlargement of the 001 peak as we have observed in several samples from the Monte Baldo area. Weathering also causes celadonite to become lighter green and finally reddish-brown for the most altered samples. We give the analyticaldata for the Turkish celadonite material. We have considered the two analyses from Cyprus by Bear (1963) and Courtois and Desprairies (1978) because pure celadonite is readily available from the volcanic area of this island. The values by the latter authors must be viewed with more confidence since a detailed X-ray diffraction study was done before major element analysis.

-+-+-+

Massif- Kambia, large pockets; XRD control; pure celadonite. 12. Buckley et al. (1978). Museum sample without reference (possibly Monte Baldo); XRD control; pure celadonite. 13.Buckley et al. (1978). Bohemia, Most massif; X-ray control; celadonite plus nontronite. 14a. Kempe (1974). Deep-sea off Madagascar, Leg 26 site 250; Coniacian XRD control. Celadonite plus smectite. 14b. Kempe (1974). Micro-analysis of the same sample for comparison. 15. Bayramgil et al. (1952). Turkey-Zonguldak; XRD control; slightly altered celadonite.

367 Table 7. Major element chemical analyses of macro-samplesof natural celadonites. Sample

SiO2

A1203

1

51.3

7.25

Fez03

.

2

53.2

2.13

20.5

4.14

3a

54.3

5.08

14.8

3b

50.5

1.40

.

4.82

3c

51.3

2.81

19.1

2.44

3d

53.1

1.87

18.9

3.41

(20.7)

(20.3)

FeO

MgO

CaO

Na2O

K20

.

5.98

1.9

6.2

. (6.67) .

5.67

-

-

7.95

.

6.05

0.8

(3.8)

(4.85)

5.45

0.1

0.0

8.95

5.28

0.6

0.5

7.33

5.83

0.4

0.4

.

7.80 8.43

HzO'

H2°-+C02

(6.18)

.

. (5.62) . (n.d.)

. (10.4) . . (7.6) .

8.94 4

56.1

2.14

14.1

5.1

5.9

0.6

-

8.83

5.43

1.36

5

55.6

0.79

17.2

4.02

7.26

0.21

0.2

10.03

4.40

0.48

6

50.6

4.2

14.1

3.3

6.4

3.1

0.4

8.7

6.9

1.4

7

55.3

(10.9)

(6.95)

3.54

6.56

0.47

0.0

9.38

5.21

1.3

8

51.4

1.7

23.7

1.4

6.34

0.5

1.3

6.62

5.87

1.11

9

54.4

5.41

14.2

3.56

6.40

0.42

0.0

9.23

4.80

1.16

10

52.3

4.02

.

.

5.79

1.2

6.71

5.17

5.05

11

54.0

4.10

18.3

0.75

6.17

0.3

0.1

7.95

12

53.0

2.90

16.4

3.80

6.20

0.33

0.38

9.38

4.59

0.75

13

49.7

1.90

21.9

2.30

3.70

1.95

0.0

6.65

5.78

4.0

14a

55.0

5.02

12.9

3.56

6.77

0.15

0.7

7.43

5.64

3.23

14b

55.8

5.76

12.3

3.42

6.78

0.36

0.8

9.28

nd

nd

15

56.5

9.1

12.4

2.19

5.98

1.13

0.9

6.49

5.32

nd

(19.2)

'

. (6.66) .

1.Delesse (1848). Italy-Monte Baldo. Brentonico; Late Eocene - veins. 2. Levi (1914). Brentonico; Late Eocene - veins. 3a. Lacroix (1916). Brentonico; Late Eocene - veins. N%O and K 2 0 incorrect. 3b. Buckley et al. (1978). Brentonico; Late Eocene - veins. Micro-analysis for comparison. XRD control. 3c, 3d. Unpublished new data. Brentonico; two different veins - X-ray fluorescence by M. Lenoble (Paris); potassium re-measured by M. Zelvelder in Brest (atomic absorption). The lighter green vein (3c) contains less q0than the darker green vein (3d). 4. Heddle (1879). Ireland - Giants'Causeway; filling druses. 5. Wells (1937). USA-Nevada-Reno; vesicular basalt. Large error bars due to small quantity analysed. 6. Wells (1937). USA-New Mexico-Sandoval; vesicular basalt. 7. Maegdefrau and Hofmann (1937). Italy-Vesuvius(?); A1203and F 9 0 3 probably inverted; XRD control. 8. Scherillo (1938). Sicily - Giuliana; a single void in Jurassic basalt. Mixed calcite removed. 9. Wise and Eugster (1964). USA-Washington - Budweiser Creek springs; Early Tertiary cavity fillings in basalt flow. 10. Bear (1963). Cyprus - T r d o s Massif - terre verte of Kambia, large pockets; calcite plus celadonite. 11. Courtois and Desprairies (1978). Cyprus - T r d o s... (Caption follows bottom p. 366)

368 Finally, analyses obtained at the British Museum on macro-samples following mineralogical studies (Kempe, 1974; Buckley et al., 1978) have been included; in addition, for comparison there are some micro-analyses on material similar to the analysed macro-samples.

We believe that we have compiled most of the analytical data presently available on celadonite macro-samples for which reasonably good mineralogical and geological control are available. The samples come from basalts of various ages although none are older than approximately 160 Ma. We consider that our collection, although very restricted, is representative of naturally occurring celadonite.

Discussion of chemical analyses on macro-samples The aluminium content is remarkably low (Table 7). Along with Hendricks and Ross (1941), we may conclude that 'aluminium is not a necessary constituent of celadonite'. This is the most obvious characteristic of celadonites and leads to several other considerations. The calculation of formulae shows a very high silica content in the tetrahedra: 3.9 to 4.0 ions for four sites (Hendricks and Ross, 1941; Foster, 1969). Hendricks and Ross remark that in celadonites 'the minimum value... for Si, is ... greater than the maximum value for glauconites'. This, in turn, determines the main mineralogical difference between celadonites and glauconitic minerals: the layer charge is mainly due to octahedral substitution in celadonites and to tetrahedral substitution in glauconitic minerals. Another property pointed out by Hendricks and Ross (1941) is that 'The magnesium content is surprisingly constant in celadonites ... in none of the samples is the magnesium content as low as the highest value found for gluuconites '. This is also true in our selected data if we justifiably disregard the celadonite-nontronite mixture from Bohemia. We fully agree with Hendricks and Ross who claim that 'the near constancy in the magnesia content of glauconites of 0.32 (for four silicon ions) reflects the essentially unchanging nature of the environment'. For celadonites the calculation of formulae gives a Mg content (for four silicon ions) between 0.6 and 0.7. This also reflects a very restricted variety of environments of formation, different from that where glauconitic minerals form. There is no marked difference in iron in celadonites compared to glauconitic minerals. Total iron is possibly slightly lower for celadonites (18-25% total Fe,O,) than for glauconitic minerals (20-27%), but the percentage of trivalent iron compared to total iron is very similar: 80-96% in celadonites and 8595% in glauconitic minerals. This is a confirmation of the interpretation of the Mossbauer spectra. This similarity of iron compositions indicates comparable Eh conditions for the formation of both celadonite and glauconitic minerals. For glauconitic minerals, it is usually accepted that the general environment (sea-water) is slightly oxidizing while the local microenvironment (the grain interior) is more reducing at the beginning of the glauconitization process. This emphasizes the role of the local environment of growth compared to the general properties of the circulating fluids. We suggest that the fluids responsible for celadonite formation were generally oxidizing when in large cracks or fissures,

369 while the more confined environment in the rock itself (local veinlets or vesicles) was nearly neutral, if not slightly reducing. This allows the iron to enter the crystal structures. To summarize, when the fluids circulate quickly, the environment is too oxidizing for celadonite crystal growth; but when the iron-bearing fluids circulate more slowly or are trapped, then the environment is dominated by the reducing property of the basalt. This suggests that the fluids come from the sea. The potassium content is exceptionally and constantly high in these analyses. Nearly all of the values below 8% in Table 7 may be due to the presence of impurities. We have analysed five more celadonites from the Monte Baldo area for potassium. The light-green vesicle fillings (possibly slightly weathered) from San Valentino contained 8.24% (K,O) and the dark-green vesicle fillings contained 9.33%; a sample from Castellano contained 9.00°/~(K20); two more samples from Brentonico, one obtained from the British Museum (B.M. 1921 = sample 3b in Table 7, courtesy B. Buckley), and one from the Ecole des Mines (E.M. 22436, courtesy M. Poullen) collections gave 9.68% and 9.52% K,O, respectively (unpublished data by M. Lenoble). Compared to values obtained for glauconitic minerals, these data are significantly higher and have a narrow range of values. Amongst approximately one hundred potassium values measured for clean glauconies (Odin and Matter, 1981; unpublished personal data) we have found only one value as high as 9.0% (K 0);moreover, the majority of the values reported as between 8.7 and 8.9% ( 0) by Odin and Matter (samples from northwestern Spain) are known to y to be slightly (about 5%) lower. The relatively limited range of potassium contents in celadonites plus the high values, is again a reflection of very favourable and homogeneous environmental factors everywhere celadonite formed. In summary, the mineral celadonite appears chemically very homogeneous from one outcrop to another. When pure, its potassium content is very high. This component does not exhibit a series similar to that shown by the glauconitic minerals. Celadonite is a mica which is characterized by a very small replacement of aluminium for silicon: celadonite is specifically a tetrmilicic clay mineral. This can be ascertained from chemical analyses where silica contents of 52 to 56% are higher than for other mica-like clays. A characteristic chemical formula follows:

53

In addition to the above formula, we propose mean oxide contents for typical celadonite: SiO = 54% ; A1 0, = 3.5% ; Fe,O, = 17% ; FeO = 4% ; MgO = 6% ; K 2 0 = 9 2 ; H 2 0 = 6.8%.

Micro-chemical analysesfor major elements In the case of celadonite, and other clay minerals from similar outcrops, micro-analyses are especially helpful because colour and other properties shown in thin sections are of little help in distinguishing celadonite from green

370 saponite or nontronite or chlorite. These micro-analytical techniques are necessarily less precise and do not allow the analyst to verify the crystallographic composition and purity of the material analysed. Precision is usually better than f 10% (20) for major elements but decreases rapidly for elements which make up less than 3% of the total rock. Large errors may also occur because clays, when present in basalts, cannot be prepared with well-polished surfaces, and these are important for obtaining reliable micro-chemical analyses. In these circumstances, the results from Buckley et al. (1978) obtained both by microchemical and macro-chemical techniques on six samples of material similar to celadonite are very significant. The differences between the results obtained using their micro-che&cal technique and analyses of large samples are shown in Table 8. Table 8. Comparison of data obtained by micro-analysis and macro-analysis* (calculated from results by Buckley et al., 1978, as a percentage of results of large-sample analysis). A positive value indicates a higher result by the micro-chemical technique and a lower result is shown as negative.

Element

Sample 14D

18b

34D

35L

D

BM

Si02

+ 6

+ 4.5

+ 2.3

+ 6. 1

+ 0.8

+ 5.2

A 203

+ 0.9

- 4.8

- 1. 4

- 48

- 13

- 11.6

Fe2°3

+ 17.5

+ 21

- 3.5

+ 31

+ 3.9

- 1.3

MgO

- 2.3

- 7.3

+ 8.6

- 0.8

+ 4.0

- 9.7

K20

+ 2.3

+ 4.8

- 5.0

+ 42

+ 9.6

+ 1.0

CaO

small quantity greatly underestimated by microchemistry

* Note that a minimum of six points were analysed by two different analysts for each microchemical datum. The data are sometimes systematically biased as is seen for SiO values. One sample (35L) yields significantly different values for most e ements. Calcium content is underestimated by micro-chemistry by as much as 90%. Iron results are very inconsistent in spite of the minerals having high contents (18.5 to 25.5%, Fe,O,). These comparisons will have to be considered seriously when using the micro-chemical analyses for mineralogical interpretation. From our point of view, it appears very imprudent to systematically calculate chemical formulae from such data.

f

Table 9 summarizes analytical results compiled from eight different sources which discuss deep-sea basalt alteration products. Kasmer and Gieskes (1976) analysed a sample from Leg 35 site 323 in the South Pacific Ocean. The basalt, of early Tertiary age, shows a 10 A clay mineral as veins and replacement of pyroxene crystals. Buckley et al. (1978) reported four ana-

37 1 lyses from deep-sea basalts of north-central Pacific and the northwestern part of the South Atlantic, and six analyses from continental outcrops. The results are very homogeneous with most magnesium contents between 5% and 7.4% and all but one potassium value between 7.7% and 10%. Some of the samples were X-rayed but the published diagrams are not very convincing. Seyfried et al. (1978) analysed a green micaceous mineral from a vein (filling of open crack) in a basalt of Leg 34, site 321 in the southeastern Pacific using a 100 pm beam. This analysis shows a large amount of iron as do the two analyses published by Andrews (1980). The latter author analysed 27 light- to dark-green minerals but gives details only on four samples. Of these, he assumes that two are mixtures of celadonite and saponite. However, Andrews suggests the existence of a complete mixed-layer saponite-celadonite series to interpret his data; in our opinion, this is not correct. Most likely, he has analysed mixtures of different minerals. Andrews explains the high iron content in his analyses of "celadonite" as due to contamination with extraneous iron oxide. We suggest, however, that this is unlikely since iron oxide would have produced a recognizable brownish colour, and the silica content would have been reduced by such contamination. We believe either that iron is analytically overestimated (see Table 8) or that deep-sea celadonites are particularly rich in iron. Laverne and Vivier (1983) do not refer to celadonite but to 'dark-green smectite' in their work on secondary minerals from basalts collected near the Galapagos Ridge (Leg 70, site 510). However, from later discussion with the authors (Laverne, 1986, personal communication to the editor) and evaluation of the published analyses, it is obvious that what they called 'dark-green mineral' is more or less pure celadonite. In the absence of X-ray study, a mixture with saponite must be considered. Finally, we point out that these celadonites, in basalts 2.8 Ma old, are amongst the youngest known today. Alt and Honnorez (1984) reported one analysis of vein celadonite from a basalt collected from the northwestern Atlantic (site 417). Two to six analyses of 5 to 25 areas were used to calculate the mean values for this dioctahedral mica. According to X-ray studies, this sample is the richest in celadonite amongst three celadonite-nontronite mixtures. The K 2 0 content (7.2%) shows that the nontronite content is relatively low and the high iron and low aluminium contents are diagnostic of celadonite. The Fe3+/Fe2' ratio has been chemically analysed at 0.89, a typical value for celadonite. Desprairies et al. (1984) reported six results for celadonites from the Palaeogene Rockall Plateau in the North Atlantic (Leg 81-site 553). These celadonites were separated from vesicle fillings where saponite was also present; the two samples with the lowest potassium values ( K 2 0 = 5.7%) concurrently showed relatively high Mg contents (12.5% and 9.7%). The other samples show more usual K 2 0 contents for celadonites, between 9.6% and 8.1%, and Mg contents between 7.4% and 8.4%. One Fe3+/Fe2+ratio was measured and is 0.72. Note that values from Desprairies et al. were calculated without considering water content. Realistic errors were assumed to be between 5% and about 20%. Micro-chemical values published by Delmont (1985) for celadonites collected from the Faeroes in the North Atlantic are also given in Table 9. These results are heterogeneous and several extreme values are not representativeof the general composition of the seven samples that were analysed. Five samples contain more than 7.6% K20; the two others must contain some smectite. Five samples have A1203 values of 2.1 to 4.7%; six samples show 5.3 to 6.7% MgO. The formulae calculated by Delmont suggest that five samples have almost no aluminium in the tetrahedral layers.

372 Table 9. Range of values (in percent) from micro-chemical analyses of presumed celadonites; single values that deviate are given in parentheses. F e z 0 3 or F e O

Si02

*l2O3

1 (1 sample)

52.2

5.5

15.8

MgO

CaO

6.9

K20 9.3

2 (10 samples)

52-55

1.7-5.8

12.5-24

(3.3)-7.4

0-1

(6.7)-10

3 (1 sample)

50.4

1.8

29.7

4.7

0.2

6.3

4 (2 samples)

53.5

0.2-0.3

27

5.8-6.1

0.1-0.2

7.1-7.3

5 ( 5 samples)

51-53

0.3-1.5

26-29

4.0-5.5

0.3-0.6

(5.0)-7.8

6 (1 sample)

51.6

0.3

25.5

6.0

0.7

7.2

7 (6 samples)

55-59

0.7-7.5

16-23

7.4-(12.5)

0.2-1.2

(5.7)-9.6

8 ( 7 samples)

50-58

2.1-(7.2)

12.5-23

(3.1)-6.7

0.1-1.0

(5.4)-8.9

(45.5)-51

1.9-(9.2)

22.5-25

32-6.1

10 ( 7 s a m p l e s )

9 ( 9 samples)

49.5-52

1.7-3.9

20-25.5

4.0-8.2

(6.3)-7.9 ( 6 .O ) - 7 . 5

11 ( 1 3 s a m p l e s )

(45.5)-53

2.5-(11)

(19.5)-( 28)

12 ( 5 samples)

45.7-51

3.1-7 .O

22.6-26.5

3.3-7.2 4.6-5.3

5.1-7.8 4.8-7 .O

-

1 Kastner and Gieskes (1976), microprobe analysis without perfect polishing; semi-quanti-

tative. Buckley et al. (1978; Table 4 ), X-ray control. Six or more spots per datum. 3 Seyfried et al. (1978 ;Table 1: sample 6), X-ray control: green mica; mean of 50 spots. 4 Andrews (1980, Table 4: samples 3,4), no mineralogy; iron as FeO. 5 Laverne and Vivier (1983; Table 6: no 6,8,9, 11, 13), dark-green clays; iron as FeO. 6 Alt and Honnorez (1984, Table 1: sample 2), X-ray control: celadonite f little nontronite; iron as FeO. 7 Desprairies et al. (1984, Table 1: sample 2), X-ray control: mainly celadonite; data calculated assuming samples to be water free. 8 Delmont (1985), X-ray control: celadonite + illite-smectite; mean of 6 to 12 areas, error bar f 5 to 10%. 9 Humphris et al. (1979, Table 2: samples 1 to 6, 10, 11, 13), X-ray control: diffuse 10 A peak; iron as FeO. 1 0 Humphris et al. (1979, Table 4), green rims; probable mixtures. 1 1 Mevel(l979, Leg 417 D-418 A, Table 4). bright-green clays: iron as FeO. 1 2 Mevel(l979, Leg 417 A, Table 15), green clays; iron as FeO. 2

Humphris et al. (1979) published two series of microprobe analyses using a defocused beam. The first was a study of green material for which the X-ray analyses showed a 10 A mineral but the 001 peak was diffuse; these data probably represent either mixtures of minerals, or poorly crystallized or altered celadonite. The second series of analyses involved the green rims of zoned vesicles, including celadonite mixed with small proportions of other clays. We note that the SiO, contents are generally low as are most of the values for MgO. In the same D.S.D.P. volume, Scheidegger and Stakes (1979) give good X-ray evidence that practically all green veins and cavity fillings are mixtures of minerals including dominant celadonite. From the same sites, Mevel(l979) has published two series of micro-chemical analyses

373 using a similar defocused beam. The results from Humphris et al. (1979) and Mevel (1979) compare favourably. The single difference is the generally lower K 2 0 contents given by Mevel. We have only used the analyses of "bright-greenclays" by this author because those that are 'light-green' show less than 4% K 2 0 and more than 14% MgO which indicates a mixture of celadonite and saponite.

The values in Table 9 represent a large variety of geographic occurrences of celadonite mainly from oceanic basalts. Compared to the data compiled in Table 8 for analyses of large samples, there is generally good agreement: high potassium and iron, low aluminium, and constant MgO are the main chemical characteristics of celadonite. Potassium oxide contents lower than 7%0,and magnesium oxide lower than 4% and higher than 7%, are not representative of pure celadonites; the high MgO values are probably due to a mixture (not an interlayered structure) of celadonite with saponite. The scattered results obtained for iron contents are probably linked with problems of polishing. In spite of this difficulty however, the very high values of more than 25% found by several authors possibly represent an iron-rich celadonite variety particularly characteristic of some deep-sea basalts. If so, iron ranging between about 16% and about 29% (Fe20,) would be the most variable component in celadonite. Aluminium contents show smaller variations which partly counterbalance the range in iron contents. Table 10. Probable range in composition of pure celadonite: major elements; comparison with glauconitic minerals. Note that there is no overlap for SO2 and MgO contents; the two minerals are chemically distinct. C eladonite* Si02

52 - 56

G l a u c o n i t i c minerals** 47.5 - 5 0 . 0

- 11.0

A 2O3

0.5 - 6.0

Fe203

16 - 28

19 - 27

MgO

5 - 7

2 . 6 - 4.6

K20

7.0 - 9.5

3.0 - 8 . 5

3.5

* according to Table 7 and 9. ** according to Odin (1975a); 54 to 75 analyses. In summary, we suggest that the chemical composition of pure natural celadonite is within the range shown in Table 10 based on the analyses we have compiled from the literature. When values are outside the proposed range, impurities can be suspected. Celadonite has higher Si and Mg contents than natural glauconitic minerals. The potassium content is generally higher in celadonite which has a narrower range of KzO contents than glauconitic minerals. Because of their different compositions, it seems clear that the environment and

374 processes of genesis of celadonite and glauconitic minerals must be different.

Trace and rare earth elements There is very little data published concerning the trace elements present in celadonite. Courtois and Desprairies (1978) have measured some trace elements in celadonite from Cyprus (no11 in Table 7). These authors note that there is a much lower trace element content, including rare earth elements (R.E.E.), in the celadonite compared to the average basalt within which it forms. In pure celadonite, rubidium often is above 200 ppm; Hart and Staudigel(l979) quote a value of 226 ppm. Strontium content is low, less than 15 pprn when the green clay is purified. This allows us to consider using this mineral for determination of Rb-Sr ages (see section on radiometric dating: p. 382-392). Table 11. Trace and rare earth elements in celadonites Rb 1 basalt Cyprus 2 celadonite 3 K-rich celadonite 4 K-poor celadonite 5 K-rich celadonite 6 K-poor celadonite

Sr

V

Cr

Co N i

Cu Zn Total R E E

-

- 271

7

34

78

63

210

2-4 100

3

19

10

195

72

14

19

1

-

20

-

17

-

-

29

68

90

-

-

-

-

-

-

-

-

-

3

119

-

-

-

-

12

67.5 50.4

-

-

117

0

51

30

72

32

Courtois and Desprairies (1978), mean values for the basalt in Cyprus. Courtois and Desprairies (1978), celadonite of Cyprus; Rb and Sr values from the K-rich celadonite of Staudigel et al. (1986). 3 to 6 Desprairies et al. (1984), celadonites from D.S.D.P. Leg 81, Rockall Plateau, extracted from vesicles. 1 2

Data from Fleet et al. (1980) show that celadonite contains 2 to 10 times less of each R.E.E. compared to glauconitic minerals. This could provide a way to distinguish between the two facies. Fleet et al. also noted that the glauconitic grains have R.E.E. contents similar to continentally derived sediments. This shows that R.E.E. are more easily incorporated from detrital sediments into glaucony than from basalts into celadonite formed in voids within the basalt. This possibly suggests that celadonite crystallization is a rapid phenomenon compared to the glauconitization process: the celadonite crystals, quickly crystallized, have no time to incorporate trace elements from the surrounding rocks but include only those from the fluids present in the immediately surrounding cavity.

375 ENVIRONMENT OF FORMATION OF CELADONITE

Environment of formation of celadonite based on petrography The following occurrences of celadonite-bearing basalts were shown to have a marine history of emplacement: Monte Baldo area, Val di Fassa, Teruel area, Cyprus Island, Turkey, Faeroe Islands, and most D.S.D.P. basement rocks. In some cases, however, it appears that basalts were erupted in a sub-aerial to continental environment (e.g., Rockall Plateau). In that case, the basalts were first subjected to aerial alteration. Petrographical study of these altered basalts shows that celadonite was produced at a later time, after the basalts had been inundated by sea-water. This happened to samples from site 553 (Leg 81) although it had previously been interpreted that weathering occurred after celadonite genesis (Desprairies, work in progress). The depth at which celadonite formed probably is very different from one occurrence to another: a few tens of metres at Monte Baldo to several thousand metres for oceanic ridge basalts; celadonite is not a depth indicator. Celadonite preferentially crystallizes in voids in basalts and replaces diverse phenocrysts which are subjected to alteration. Vesicles or altered phenocrysts are not fully closed chemical systems because addition of ions is necessary to fill an empty vesicle, or exchange of cations is necessary to replace plagioclase by Fe-rich clays. The D.S.D.P. studies indicate that this circulation of fluids sometimes occurred to a depth of 600 m in the basalt (Andrews, 1980). However, celadonite is usually concentrated in the most physically disturbed portions of the rock; the more vesicular top and bottom of a flow and pillow lavas allow fluids to intensively circulate in the deposited formation. The largest voids of a basalt (and especially the large fissures) are, however, less likely to be filled with celadonite than are the smallest openings. We suggest that, in addition to permeability, it is also important that the environment where celadonite forms remains somewhat confined in order to allow ions time to combine: celadonite does not simply precipitate (as the carbonate probably does) but needs a certain confinement in order to crystallize in large volumes. We suggest that this is a new example of semi-confinement as described for the formation of glauconitic grains (Chapter C4). The origin of the circulating fluids has been considered using the petrologic data. Seyfried et al. (1976) noted the presence of Late Eocene coccoliths in veins of smectite formed in basalts of Leg 34. They thus suggest that cold seawater had circulated through the basalts. Andrews (1980) also noted that some fissures are filled with nannofossil-bearing sediments and said that this implies introduction of sea-waterfrom the top to the bottom, into the volcanic layers. Introduction of sea-water into the basalt may also be postulated because the stage of alteration of the basalts is an oxidizing one. The margins of the fissures appear to be oxidized before deposition occurs. The oxidizing processes presumably are related to massive sea-water introduction. The initial circulation of oxidizing waters is only one amongst several oxidizing episodes occumng during the long process of alteration of the deep-sea basalts. Celadonite itself

376 may be oxidized after formation (it becomes yellow or brown and is enriched in Fe203) or the deposition of brown clays mixed with Fe203may occur in the centre of vesicles in which celadonite rims first formed. In spite of evidence supporting penetration of sea-water into basalt, as discussed above, Delmont (1985) recognized what he considered to be hydrothermal channels which implied the circulation of fluidsfrom the bottom to the top. Hydrothermal fluids may be assumed to be reducing since in hydrothermally active submarine areas, deposits are mainly sulfides (see mid-Atlantic ridge, Leg 109 in: Episodes, May 1986). Formation of pyrite which is frequently observed at different stages of alteration of deep-sea basalts, may indicate reducing conditions caused by hydrothermal (i.e., juvenile) fluids coming out of deep basalt layers. Thus, there is evidence for fluids moving in both ways. It is difficult to ascertain whether or not a preferential direction of water movement dominates at different stages of the alteration process. However, celadonite forms at a given time in the sequence of alteration. The first stage involves an oxidizing process with formation of black halos and iron oxides. Silica could be deposited in a very preliminary stage of the formation of the secondary minerals. Smectite, especially the trioctahedral saponite, is frequently the first clay formed; the initial black saponite is followed by late light-green saponite (Desprairies et al., 1984). Celadonite forms later, and is followed by a second stage of saponite formation. Zeolite and calcite are products of the later alteration of the volcanics. This sequence of mineral formation might be explained in terms of redox properties with oxygen-rich fluids present at the beginning of the sequence and more reducing conditions present later. Pyrite seems to form at a late stage and is sometimes found mixed with celadonite as well as with all other secondary minerals. This sequence may also be considered in terms of confinement of voids and concentrations of ions. The oxidizing stage and formation of silica deposits would represent an open environment with significant circulation of fluids from the sea to the volcanics. Secondary clay minerals, would form under conditions of less influx of sea-water, higher ion contents in the water, and greater confinement. Finally, zeolite and calcite would precipitate from ion-rich fluids.

Environment of formation of celadonite based on mineralogy Similarities between celadonite and glauconitic minerals The mineralogical data reported herein can be used to compare the genesis of celadonite with that of glauconitic minerals (Chapter C4).The mineralogical similarity of celadonite and glauconitic minerals is obvious; both are ferric 2:1 phyllosilicates. This implies that the two facies formed in a relatively confined environment where fluids were rich in cations. However, the environment cannot be very reducing and extremely confined because the iron in the mineral is mainly ferric. The circulating fluids must usually be oxidizing, as indicated by the formation of black halos before celadonite formation, and by the fact that

377 celadonite is sometimes oxidized after formation. However, at the precise place and time where and when celadonite forms, the fluids have lost most of their oxidation potential. Sea-water, at relatively shallow depth, appears to be an especially good candidate for such a circulating fluid both because it has a relatively high ion content, and because it contains oxygen, the proportion of which can be lowered during the initial oxidation of the basalts. Deep-sea waters with their low temperature and oxidation potential also could produce celadonite. Continental meteoric fluids would generally be too oxidizing and have too low a cation content to favour the genesis of celadonite. Table 12. Comparison of mineralogical observations for celadonite and glauconitic minerals. Factor

Celadonite

Glauconitic minerals

Mineralogy

l O A sheet silicate

lOA sheet silicate

Main octahedral cation

Iron, mainly ferric

Iron, mainly ferric

Inter layer cations

High content

from low to high

K 2 0 maximum

9.5-10.0%

8.5-9.0%

X-ray diagrams

Sharp peaks

Broad peaks

1.R. spectra

Sharp bands some are specific

Broad bands some are specific

Crystal form

We1 1-shaped laths

G lobu 1es to

irregu 1ar b 1ades Crystal size

Max. 15 to 30 p m (length)

Max. 5 to 10 p m (diameter)

Tetrahedral A1

L o w , CO.214 sites

0 . 2 5 to 0.5014 sites

Octahedral Mg

0.6-0.812 sites

0.3-0.512

sites

DifSerences between celadonites and glauconitic minerals The maximum potassium content reached by celadonite is clearly higher (9.5 to 10.0% K 2 0 ) than the one reached by the glauconitic mica (8.5 to 9.0% K,O). In our opinion, this observation shows that celadonite crystallization is due to much more favourable thermodynamic factors than are present in the "grain/sea-water" system of glauconitic minerals discussed in Odin (1975a) Odin and Matter (1981) and chapter C4. The potassium-rich celadonite formed in equilibrium with the circulating fluid in a single reaction; the glauconitic mica formed in successive reactions, the first of which is crystal growth of Fe-rich smectite followed by several recrystallization processes which trapped more and more potassium.

378 We suggest that there are environmental factors which account for the differences between celadonite and glauconitic minerals. One factor is temperature. Glauconitic minerals form at temperatures below 15" or even 10°C at the sea-bottom. An increase of only 10" to 20°C would represent much more favourable conditions for the formation of celadonite. A second suggestion, is that the circulating fluid feeding the growing celadonite crystals is probably richer in ions than the pore fluids in a green grain at the sea-bottom (this seems likely if one remembers that glauconitic minerals form near the top of the sediments, possibly a few centimetres or decimetres from the sea-water and sediment interface). On the other hand, sea-water circulating in a very thick volcanic pile may become nearly saturated with ions which are added to the growing celadonite. A long history of circulation of such modified seu-wuter through basalts appears to be a possible favourable factor for celadonite formation. A third suggestion able to explain the morphology and good crystallinity of celadonite, has to do with the amount of space where the reaction occurs. Instead of very small micropores, like those where glauconitic minerals grow, the celadonite crystals usually develop in voids where plenty of room is available for growth. Therefore, euhedral crystals may grow in these micro-geodes. In a very few cases, micro-geodes exist in grains which are undergoing glauconitization; under these circumstances, morphologically well-shaped tiny sand rosettes of glauconitic minerals are formed (Odin, 1974). On the other hand, celadonite that replaces phenocrysts seems to differ from that described above. The few diffractometric and microchemical analyses available show less well-crystallized minerals (often mixed with other secondary products) and a generally lower potassium content. Celadonite crystals are frequently smaller than the microcrystals observed in evolved to highly-evolved glauconitic minerals. We have suggested above that this could be explained by the short time usually available for celadonite growth; however, given the generally long time available for alteration of a basalt at the sea-bottom, this needs further discussion. To say that there is only a short time during which celadonite can grow is to suggest that at least one of the favourable factors is short lived. We suggest that favourable conditions occur when the fluids undergo changes in either chemical composition or rate of circulation. Another possible explanation would be the presence of supersaturated solutions which undergo a decrease in temperature. In summary, celadonite probably needs specific conditions to grow: 1) voids; 2) circulation of ion-enriched water; 3) comparatively elevated temperature. These conditions are only met during one or possibly several short periods during the alteration of a basalt at the sea-bottom. Further remarks on the environment of genesis of celadonite may be based on the chemical data. Compared to glauconitic minerals, celadonite has higher silica and magnesium contents, and lower alumina content. This composition suggests that the fluid causing the growth of celadonite crystals is not just sea-water. A significant enrichment in silica is very probable. The availability of silica is indicated by the occurrence of siliceous deposits such as those around the celadonite veins in Cyprus Island or the Monte Baldo area.

379 l80 Isotopic

study on celadonite

We have suggested that the temperature factor could have played a significant role in the genesis of celadonite. The data provided by oxygen isotopes can partly be interpreted in terms of the palaeotemperature of the fluids from which celadonite is crystallizing. This interpretation is based on a certain number of assumptions which are difficult to verify. They are: 1) the fractionation factor for the analysed mineral is known; 2) isotopic equilibrium between the fluid and the mineral is reached; and 3) the isotopic composition of the fluid is known. If all these assumptions are valid, then the temperature may be determined. The fractionation factor for celadonite has never been measured; but study has been made of similar minerals like illite and smectite. The illite fractionation factor would be the best one to use if we assume that fractionation factors are similar for nearly identical minerals; the use of the fractionation factor calculated from glauconitic minerals would be even better. Savin and Epstein (1970) have measured oxygen compositions of various glauconies from the present-day sea-floor. However, in absence of mineralogic data, it is unknown whether or not inherited light isotopes were still present in their samples (Chapter C1, p. 238). According to what we know about glaucony from off the California coast (Odin and Stephan, 1981), the glaucony from the Santa Monica Bay analysed by Savin and Epstein (1970) is probably Recent and evolved (6l80= 23.8%0). We also evaluated additional evolved glauconies which possibly reached isotopic equilibrium with sea-water and thus lost most of their li ht oxygen which was inherited from the initial substrate. The measured 6 0 values are quite similar: between 22.5%0 and 24.0%0 (Table 13). These values are for a component which formed in equilibrium with sea-water (6l80= Oo/o0 although this is not a certain value during Quatern time), at a temperature probably close to 10°C. This means that if we obtain l 8 0 values near 23%0on Recent celadonite which formed in waters with an isotopic composition similar to that of the sea-water, then that celadonite formed at about 10°C. A lower

6

7

Table 13. Oxygen isotopes in evolved glauconies. (Data from Keppens et al., in press)

Samples

K20

6 l 8 0 SMOW ~~

6941

<3 M a off California

high

23.8 +

G . 585

0 M a o l d off California

7.5

22.95

ML. 279

6 Ma o l d off Spain

8.0

22.47 o

ML. 253

6 M a o l d off Spain

8.9

23.98

o

ML. 253

6 M a o l d off Spain

8.9

22.78

o

ML. 212

6 Ma o l d off Spain

8.8

23.23

o

o

380

6l80 will imply higher temperatures for Recent celadonites; a correction is necessary for older samples because the 6l80 of sea-water varies with time. Even though there are difficulties in interpreting the data, we may profitably consider the 6l80values for celadonites and related minerals which have been published and interpreted in the literature (Table 14). Kastner and Gieskes (1976) reported a 6l80 value of 21% for celadonite collected from the Indian Ocean (Leg 35-site 323) from a Palaemene or older basalt. These authors calculated a maximum temperature of 26 k 5'C based on an isotopic composition of 0% for the oxygen of the fluid from which the celadonitecrystallized. Seyfried et al. (1978) reported a similar value of 21.9"& for celadonite from basalt collected off Peru. They also pointed out, following the suggestion of Kastner and Gieskes (1976) and Lawrence et al. (1975), that the water circulating in the basalts was most probably depleted in l80and thus the temperature of formation was almost certainly less than the calculated 22°C. Finally, calcite which precipitated as a late phase in veins near the celadonite has a 6l80 value which suggests a maximum temperature of formation of 7'C. From basalt at the same location, Stakes and ONeil (1982) obtained nontronite and celadonite from vesicles and measured 6l80values of 22.9% and 19.5%, respectively. They suggest temperatures of formation of about 35'C for the two minerals. Celadonites from Cyprus yielded formation temperatures of 15-20"C(6l80 = 23.8% to 23.2%) and 25'C (6l80=34.4%) for silica from borders around the celadonite veins. BCIhlke et al. (1984) summarize results for several secondary minerals extracted from Aptian basalts from the Bermuda Rise in the North Atlantic Ocean. The authors calculate formation temperatures of 53'C for saponite; 35.4'C for chert with celadonite and calcite; 25.6'C for celadonite-nontronitemixture (?) from a vein; 13.8'C for celadonite with minor nontronite, and 28.6'C to 7.4'C for five different calcites. The temperatures are calculated assuming that the fluids from which the minerals crystallized are similar to the Cretaceous sea-water. The illite fractionation factor is used for the celadonite mixtures; the use of a smectite fractionationfactor would increase the apparent temperatures. Andrews (1980) measured 6l80values of 24 to 25% for two smectites associated with celadonite and suggested a maximum temperature of formation of 50'C. However, Desprairies et al. (1984) have published significantly smaller values of 17.1% for celadonite and 18.5"A for saponite which were extracted fiom the vesicular upper and central portions, respectively, of Palaeocene basaltic flows in the North Atlantic. New data indicate a regular increase in the 6l80values both for celadonite (13.4% to 17.1%) and saponite (18.5% to 23.6%) from the central part to the top of flows. Consequently,estimated temperatures, assuming equilibrium with water at 0% and using,the magnesium smectite/water fractionation curve of Escande (1983) plus the fractionation factor for glauconite of Savin and Epstein (1970), range from 90'C to 6O'C for celadonite and from 110'C to 65'C for saponite (Table 14). Interpretations of data are hindered by the assumption that the aqueous phase in e uilibrium with clay minerals was ordinary sea-water. A negative or positive shift in the 6?8 0 of the water, would result in lower or higher temperatures than those suggested.

Even if the calculated temperatures are biased, the relative values are correct. It is therefore interesting to note that, in a typical basalt alteration sequence, the first formed secondary minerals, chert and saponite, have apparent temperatures of formation higher than celadonite (more or less mixed with smectites: saponite or nontronite) and calcite. This is the temperature sequence shown for

38 1 Table 14. Oxygen isotopic data for celadonite and associated minerals; corresponding approximate temperatures are quoted. 6 180

Minerals

Calculated T ("C)

ShlOW,% 6 1 8 0 ~= 0%

Comments

6 180Wa = -l%o

1 Cel adoni t e

21.4

26b

Vein, with g o e t h i t e and c a l c i t e

2 C el adoni t e

21.9

22b

Vein, with f e r r i c oxides

2 Calcite

33.2

7c

3 C el adoni t e

19.5

35d

Zoned v e s i c l e f i l l i n g s of c e l a d o n i t e

3 Nontronite

22.9

35e

and nontronite

In last phase

4 C el adoni t e

23.2

9 b

C e l a d o n i t e l e n s e s in s h e e t f l o w s , with

4 C el adoni t e

23.8

g b

minor nontronite

4 Q uar t z

34.4

20h

Massive c h e r t from borders

5 C el adoni t e

23.7

1qd

In vein with minor nontronite

5 Green c l a y

21.4

26d

Vein, mixture celadonite-nontronite

5 Calcite

30.6

9 f

Fibrous c a l c i t e in reopened celadonite-nontronite vein

5 Calcite

26.4

2gf

In i n t e r p i l l o w area with celadonite-nontronite and c h e r t

5 Q uar t z

32.1

35g

Massive c h e r t in i n t e r p i l l o w area with c e l a d o n i t e , nontronite and c a l c i t e

-

-

53e

5 Saponite

-

Scoriaceous t o p of b a s a l t i c f l o w : arnygdules f i l l e d with c e l a d o n i t e ; c h e r t with c e l a d o n i t e and c a l c i t e

-

Upper p a r t of Lesicular zone : v e s i c l e s f i l l e d with intergrowth of c e l a d o n i t e a n d saponite

6 Cel adoni t e

17.1

52b

6 Q uar t z

28.5

52h

6 C el adoni t e

16.2

59b

6 Saponite

23.6

65i

6 C el adoni t e

14. 6

72b

Veins and v e s i c l e

6 Saponite

21.3

84i

l o w e r p a r t of t h e v e s i c u l a r zone

6 C el adoni t e

13. 4

84b

Vesicular zone

6 Saponite

20.9

8Ei

Vesicular zone

f i l l i n g s in t h e

1)Leg 35, D.S.D.P., Bellingshausen Abyssal Plain (Kastner and Gieskes, 1976); 2) Leg 34, D.S.D.P., Nazca Plate (Seyfried et al. 1978); 3) Peru-Chile Trench (Stakes and ONeil, 1982); 4) Troodos ophiolite, Cyprus (Desprairies, unpublished); 5) Legs 51-53, Bermuda Rise (Btihlke et al., 1984); 6) Leg 81, D.S.D.P., Rockall Plateau (Desprairies et al., 1984 and unpublished). a) Assuming formation in equilibrium with Cretaceous sea-water (Shackleton and Kennett, 1975). Fractionation factors for b) glauconite (Savin and Epstein, 1970); c) calcite (Craig, 1957); d) - e) illite and smectite (Yeh and Savin, 1977); f) calcite (ONeil et al. 1%9); g) quartz (Knauth and Epstein, 1976); h) quartz (Clayton et al., 1972); i) Mg-smectite (Escande, 1983).

382 all published series of data; therefore, the temperature of the fluids is decreasing when celadonite forms. This may well create the super-saturated solutions favourable for celadonite crystal growth. The temperature of formation may have been different in different areas. Values around 25°C were common but higher apparent temperatures, between 50" and lOO"C, have been reported here. These comparatively high apparent temperatures would need to be confirmed elsewhere. A similar variability in the apparent temperature of formation is supported by some l80data published for calcites in veins of sites 417 (Bermuda Rise). Lawrence (1979) measured 6l80 values in these calcite veins collected from Lower Cretaceous basalts. He calculated temperatures of 14.6" to 33.5"C which is only slightly more than the values published later by Bohlke et al. (1984); he, however, notes that these values were higher than the values calculated from calcite from Legs 37 or 45 (apparent temperature less than 15°C).

Time of formation of celadonite Radiometric dating applied to celadonite Celadonite and basalts collected from the Bermuda Rise have been investigated geochronologically. The basalts are known to have been emplaced at the time of magnetic anomaly M, which is considered by palaeomagneticians to be of Early Aptian age. Moreover, sediments of Early Aptian age were observed above the basalts at site 417 D and 418 A (Donelly et al., 1979). Seventeen fission-track ages were measured on glasses separated from the basalt. The glass samples were fresh and common in pillow margins (Mevel, 1979); they are thermally rejuvenated but a correction factor has been applied successfully by Stoner and Selo (1979). The results scatter between 100 Ma and 118 Ma, but a mean age of 108.3 Ma was suggested. This was recalculated later (change in decay constants) to 109.2 f 3.6 Ma (Selo, 1983, personal communication to the editor). This numerical age also corresponds to the age estimate by Channel (1982) for that anomaly. Hart and Staudigel(l979) obtained one celadonite from 60 m below the top of the basaltic pile at site 417A, and nine other clay minerals, reported to be smectites, from site 418A. However, Hart and Staudigel(l986) consider that at least two clay separates ("smectites") are mixtures of celadonite and smectite. On the isochron plot, the isotopic data for these two clays give two points between the celadonite point, very far from the origin, and a group of somewhat scattered points near the origin for the other smectites (Fig. 18). These four "points" define a line corresponding to an age of 111 k 7 Ma (Hart and Staudigel, 1979). This line can be considered a mixing line between pure celadonite and pure smectites, with two intermediate mixtures. In these circumstances, it may be better to calculate a two point isochron using the most reliable (K-rich) geochronometer (celadonite) and an assumed isotopic composition for the fluid from which the mineral crystallized. If we use the Aptian sea-water strontium isotopic ratio of 0.7072 (Faure, 1982) the age calculated is 109 f 4 Ma, a value that is analytically the same as that obtained above. The use of a different initial

.80

383

I

Site 417 celadonite

/

, /

5 (D

+//

$a3

/

109 Ma

/ ,

I

5

10

ik Rb/Sr

20

*

Figure 18. Rb-Sr radiometric data for smectites and celadonite from the Bermuda Rise. The detail of the area near the y axis is shown along with age reference lines corresponding to 109 Ma and 11 1 Ma. The scattered data points obtained for smectites are very near the y axis, and the choice of an isotopic initial ratio for strontium greatly influences model ages for these samples. (Modified from Hart and Staudigel, 1979)

isotopic ratio would not significantly modify the calculated age of the celadonite due to the very high Rb/Sr ratio of this clay. In conclusion, there is no measurable difference between the age of emplacement of the Bermuda Rise basalts and the apparent Rb-Sr age of the celadonite obtained from these basalts. The smectite isotopic data from Hart and Staudigel (1979) imply an initial strontium ratio of about 0.7062 which is significantly less than the ratio of about 0.7072 for Aptian sea-water (see enlarged corner of Fig. 18). However, the isotopic data obtained by the same authors on seven calcites show values around 0.7072, essentially identical to the Aptian sea-water. Therefore, calcites which formed after celadonite were crystallized in an environment similar to sea-water as far as the 87/86 strontium isotopic ratios are concerned. The analytical results obtained from smectites are of uncertain value and Hart and Staudigel mention the possibility of the clays being contaminated by altered basalt. It seems quite possible that the celadonite crystallized in a fluid similar to sea-water at least with respect to strontium isotopic composition. The second example of geochronological study concerns samples from the Monte Baldo area (Odin et al., 1986, and Fig. 19). The celadonite-bearing volcanics, about 100 m thick, are located between bioclastic marine limestones which contain numerous large foraminifers; these allow biostratigraphers to assign a Biarritziarw age to the extrusive event. This stratigraphic position pro-

384 I

44

Figure 19. Sample locations for celadonite-bearing volcanics from the Monte Baldo area. Volcanics are dotted. The area where celadonite is abundant is dashed. It is called Terre Verdi on old maps. Samples 65 and 71 were collected more than 40 years ago from now disused mines.

bably corresponds to the 38-40 Ma time period (according to the time scale by Curry and Odin, 1982) as the unit is entirely pre-Priabonian and is supposedly late- to post-Lutetian. The basalt itself is recognizably altered but a reasonably well-preserved whole-rock sample has been analysed, resulting in a K-Ar age of 36.5 f 2.0 Ma. Well-preserved plagioclase was separated from a contemporaneous second basalt and a K-Ar age of 40.4 f 0.8 Ma was calculated. Celadonite is present locally in the topmost layers of the tuff and lava pile as vesicle fillings; deeper in the volcanics, veins are present (de Brignoli, 1820). Seven celadonite samples were obtained: three direct from vesicular basalts at San Valentino (Monte Baldo) and Castellano (18 km to the north of San Valentino; Fig. 19), two come from underground workers who obtained them

385 from near the locality of Brentonico, and one sample came from each of the British Museum (London) and Ecole des Mines (Paris) mineralogical collections (because they were more pure than the material from miners and were larger than the samples from vesicles). The preliminary sample provided by the Ecole des Mines (of uncertain origin, however) was dated by P.Y. Gillot in Gif at 22.5 Ma; the other results are in Table 15. One may note the generally high potassium content of the celadonites, all of which are well-crystallized. All of the celadonite ages are 8 to 12 Ma and up to 16 Ma younger than the probable age of the Biarritziano. Table 15. K-Ar ages of celadonite and volcanics stratigraphically interlayered in the Biarritziano from Monte Baldo (Southern Alps). Geochronometer

K %

% rad.

Apparent age 2 2 d 36.5 2 2 . 0

Whole-rock b a s a l t San Valentino

609.7

0.66

63.7

S a m e s a m p l e as above

609.7

0.66

49.3

*

35.5

1.8

*

(2)

Plagioclase Besagno

609.1

1.26

96.0

*

40.4 i 0.8

*

(1)

Vesicular dark C e l a d o n i t e San Valentino

C64a

7.74

84.8

29.7 i 0.8

(1)

Vesicular l i g h t C e ladonite San Valentino

C64b

6.72

57.1

27.0 2 1 . 4

(2)

C e l a d o n i t e vein ( l i g h t ) Brentonico

C71b

6.47

70.9

30.1

1.5

(2)

C e l a d o n i t e vein ( d a r k ) Brentonico

C71a

7.41

75.7

31.6

1.6

C e l a d o n i t e vein "Brentonico" ( B M 1 9 2 1 )

C65

8.04

86.1

Vesicu 1 ar C e ladonite Castellano

C44a

7.35

71.0

*

*

(1)

*

30.1 2 0.8

*

28.5 i 1 . 4

(2) (2)

*

(2)

1) Odin et al. (1986); 2) unpublished argon analyses by M. Zelvelder in Brest.

* mean of 2 measurements.

To help determine if these ages could be related to the time of genesis of the celadonite, or to a regional rejuvenation caused by heating of the geochronometer, some of the samples have been submitted to an independent Rb-Sr study. The analytical results are shown in Table 16. They show a very high rubidium content for the celadonite, especially the two samples from near San Valentino (Monte Baldo sensu stricto) which give Rb-Sr apparent ages of 29.1 Ma and 28.3 Ma which is quite similar to the K-Ar ages of 29.7 Ma and 30.1 Ma. The sample from Castellano shows a Rb-Sr age of 24.8 Ma which is lower than the K-Ar age of 28.5 Ma. The general agreement between K-Ar and Rb-Sr ages in the Monte Baldo area allows us to suggest that these ages are significant and represent the time of crystallization of the celadonite. It is worth

386 noting that the 87/86 strontium isotopic ratios of the carbonates from the sediment and from the volcanics are very similar, but both are slightly lower than is usually accepted for Eocene sea-water (0.708 according to G. Faure, 1982). TablelC. Rb-Sr ages of Eocene celadonites from the Monte Baldo area. Rb ( p p m )

Sr ( p p m )

87Rb/86Sr

87Sr/86Sr

Apparent a g e i 2 U

~~

Vesicular cel adoni t e C a s t e l l a n o , C44a

199.8

4. 3s

133.4

0.75452

24.8

i 0.6

(1)

Vesicular dark cel ado nite San Valentino, C64a

398.7

5.05

230.8

0.80264

29.1 2 0.6

(2)

41.9

0.72416

28.3 2 1.0

(2)

C e la doni t e vein Brentonico, (British Museum s a m p l e ) , C65

412.5

28.6

(1) using an initial 87/86 Sr ratio of 0.70750 measured on calcite glomerules of the basalt (2) using an initial 87/86 Sr ratio of 0.70736 measured on calcite from overlying sediment *(Data by P.D. Fullagar in Chapel Hill)

From this second example, it may be concluded: 1) that apparent radiometric ages about 10 Ma lower than the time of extrusion are obtained on celadonite while the radiometric age of the basalt apparently gives the time of the extrusion; 2) the apparent age of celadonite may significantly vary from one outcrop to another in a single formation; and 3) the general agreement between K-Ar and Rb-Sr ages suggests that these are ages of crystallization. The third example of geochronological study involves the K-Ar dating of celadonites from the Rockall Plateau (Leg 81), undertaken by Desprairies et al. (1983) and complemented here (Table 17). The basalts are now known to have been extruded in a shallow marine to subaerial environment prior and close to the Palaeocene-Eocene boundary which corresponds to an age of 50 to 53 Ma (Curry and Odin, 1982). Probably equivalent basalts were dated in eastern Greenland and an age of about 50 Ma was suggested (Odin and Mitchell, 1983). The oceanic basalts on the Rockall Plateau itself were submitted to K-Ar investigations by MacIntyre and Hamilton (1984). As is usually the case with basalts from the oceanic floor, the analytical results have to be very carefully interpreted due to large variations in the apparent ages. A good example of this variation of K-Ar apparent ages in deep-sea basalts is given by the results proposed independently by several authors for the Leg 34 basalts (in another basin); ages of 8 Ma to 25 Ma were found for samples from site 319, and ages ranging from 23 Ma to 4 4 Ma for site 521 (Hogan and Dymond, 1976; Seidemann, 1976; Lanphere and Dalrymple, 1976; Reynolds, 1976). Some of these ages are younger than the accepted time of extrusion of about 16 Ma for site 319 and about 40 Ma for site 321. These young ages are usually considered to result from long-term alteration including recrystallization of potassium-bearing glass. In addition, apparent ages slightly to considerably older than the time of extrusion were also measured, especially in the case of Leg 8 1. These old apparent ages, obtained for deep-sea basalts, are usually considered to be due to the presence of significantand heterogeneous excess argon.

387 In spite of this difficulty, MacIntyre and Hamilton suggested an age of about 53 Ma for the basalts collected at site 555 (Leg 81) which is in agreement with the time scale of Odin and Cuny (1982) Celadonite was formed after the basalt was inundated by sea-water following subsidence of the sea-bottom. The green clay was scraped from different layers in the lava pile, and K-Ar ages measured by H. Bellon in Brest. The results are compiled in Table 17. It may be noted that four of the five "celadonites" have a comparatively low potassium content and might be mixtures of celadonite and saponite; the finer fraction (0.5 to 2.0 pm) of the celadonite from the uppermost flow was also separated and analysed. It shows a lower potassium content, probably linked with a higher saponite content; the lower apparent age compared to the bulk celadonite could be explained by poor argon retention of the minerals. Regardless, there is a pattern to the apparent ages; the oldest sample is from the centre of the volcanic pile and the youngest comes from the top of the sequence. Table 17 indicates that there is a good correlation between apparent ages and palaeotemperature with youngest age corresponding to lowest temperature. From this example, it may be concluded that the K-Ar apparent ages of celadonite may, but need not, correspond to the time of extrusion of the lava. There is a tendency for younger ages to be related to a sample's proximity to sea-water. H. Bellon suggests that celadonites from the exterior may have crystallized (or recrystallized) a long time after the celadonite from the interior of the flow. This interpretation assumes that the apparent ages of the celadonite samples really correspond to their time of genesis. If this is the case, some celadonite samples from the Rockall Plateau formed at essentially the same time as the basalts, and others crystallized from fluids penetrating the rock as much as 20 Ma after the basalt emplacement. Table 17. K-Ar radiometric ages of Palaeogene celadonites from the Rockall Plateau and 6l80 values obtained from corresponding authigenic clays. K (95) % r a d .

age

2cr 2.7

Centre o f f l o w

4.81

73.8

53.8

Middle part of f l o w

5.39

27.6

47.9 i 2.4

As above ( f i n e f r a c t i o n ) 4 . 9 8

64.3

42.1 2 2.1

External Part of f l o w (top)

7.82

52.7

42.5 i 2.1

Scoriaceous t o p of f l o w (most external part)

5.64

56.6

3 4 . 5 +-

6 l8O

%O

celadonite

6 l80 %o saponite

13.4 (84°C)2

20.9 (88°C)2

1 4 . 6 (72'C)

2 1 . 3 (84OC)

16.2 (59°C)

23.6 (65'C)

The fourth example of radiometric investigation is on celadonite from Cyprus which was investigated independently by several research groups. The Troodos Massif is described by Desprairies and Lapierre (1973) as a plutonic massif covered by pillow lavas and cut by dykes. Celadonite is present

388 in the upper pillow lava formation which is overlain by the Peradephi Formation, including umbers, chert and mudstone. The age of the umbers has been much debated. Based on field evidence, they are certainly older than the foraminifer-bearing Kanniviou Formation of well-accepted Maastrichtian age (Blome and Irwin, 1985). They were long considered to be of Campanian age. However, the umberiferous strata contain radiolarian faunas considered to be middle to late Turonian by Blome and Irwin (1985). This age assignment seems somewhat too precise considering the relatively poor knowledge of the time distribution of the faunas (P. de Wever, personal communication). According to De Wever, a pre-Coniacian to post-Albian age can be suggested for the pillow lava formation. This corresponds to a numerical age older than 87 Ma and younger than 96 Ma (Kennedy and Odin, 1982). The volcanics from the Troodos Massif have been repeatedly dated (Table 18). The good review by Blome and Irwin (1985) gives a number of wholerock basalt ages, mainly obtained in Geneva by M. Delaloye. Most of these ages are in the interval of 75 to 79 Ma, with a maximum value at 83 k 3 Ma. This is in disagreement with the biostratigraphic conclusions above but may be (and was) explained by the high degree of alteration of most of the analysed basalts. Staudigel et al. (1986) published a similar age of 84.7 ?r: 2.3 Ma, for a relatively fresh andesite from the Pediaeos river banks. In summary, the wholerock geochemical systems seem to have been rejuvenated after crystallization of the volcanics. Staudigel et al. (1986) report that two celadonite veins were sampled. One of these veins is formed of clay with 4.0 to 4.5% K20, and 53 to 140 ppm of strontium, and cannot be a pure celadonite. Unfortunately, no X-ray study was reported and thus the significance of this vein is doubtful. From the isotopic data obtained, Staudigel et al. (1986) draw a single isochron line for the two celadonite veins. This probably is not valid because it cannot be demonstrated that the two veins are cogenetic. Moreover, the K-poor celadonite data are so near the diagram origin (see enlarged corner Fig. 20) that the age for this sample is very dependent on the initial isotopic ratio assumed for the strontium. After examining their isochron diagram, we note that their analytically older celadonite samples (8A, 123) are below their reference line and their analytically younger celadonite (8N108) is above (the reverse should be correct). There is something wrong either in their figure or their data table. We have reinterpreted their data by calculating model ages using the sea-water strontium composition accepted for the Middle Cretaceous (0.7075; Faure, 1982). The three fractions of the comparatively K-rich celadonite vein give model ages of 86.6 Ma ( K 2 0 = 7.34%), 88.9 Ma (K2.0 = 6.49%), and 90.8 Ma (K,O = 5.96 wt%). Using an initial strontium isotopic ratio of 0.703, a value appropriate for the surrounding basalts, would slightly increase these ages to 88.5 - 92.0 Ma. Figure 20 redrawn using the data in Staudigel et al. (1986), shows that the three data points for the K-rich samples plot far from the y axis. Rb-Sr model ages also were calculated for the two most radiogenic samples of the K-poor mixtures (sample 9, Fig. 20). Values of 74 Ma and 77 Ma were obtained using an initial 87/86 strontium isotopic ratio of 0.7075, but ages were

389 SAMPLE 8

-

K-RICH CELADONITE

FRACTIONS

8N 12:

9 = K-POOR I00

0 90

Figure 20. Rb-Sr radiometric data for celadonites and zeolites from the Troodos Massif (Cyprus). Crosses are for celadonites, stars are for zeolites. The very small area along the y axis is enlarged to the right, and shows that the age of these celadonites, which have very low contents of K (and Rb) is highly dependent on their initial 87/86 Sr isotopic ratio. (After data from Staudigel et al., 1986)

much older (more than 85 Ma) using a slightly lower initial ratio of 0.7065. It is therefore impossible to know whether or not there is a significant age difference between the two veins in absence of precise knowledge of the initial strontium isotopic composition. Staudigel et al. also published K-Ar ages on the same material. The K-rich celadonite gives two K-Ar ages of 87.1 0.8 Ma (analyses by I. Mc Dougall, Canberra). Two dates were also obtained at the Oregon State University on K-poor celadonite. However, these apparent ages, 79.8 Ma and 76.5 Ma, were obtained after baking the clays for line cleaning to 180°C, a temperature inappropriately considered to be low by Staudigel et al. (1986): K-poor clays may well lose a portion of their argon at such a temperature (Odin and Bonhomme, 1982; Zimmermann and Odin, 1982); thus, according to us, the actual analytical ages are probably slightly underestimated by an unknown amount. The K-Ar and Rb-Sr ages for K-rich celadonites obtained by Staudigel et al. (1986) are very similar (Table 18); consequently, they most probably represent the time of crystallization of the analysed green clay. The calculated ages are about the same as or perhaps a few million years younger than the generally accepted estimate for the age of extrusion of the pillow-lavas. It is interesting to

*

390 note that this K-rich celadonite may have behaved more like a closed system than did the whole-rock basalts. We have analysed new samples of celadonite from large pockets in the Troodos Massif in order to determine if the apparent ages of the material are reproducible from one location to another. Also, these pockets contain a celadonitic filling which has higher potassium contents than the previously dated materials. Table 18 gives these new K-Ar and Rb-Sr results. The calculated K-Ar and Rb-Sr apparent ages are nearly similar for a given material. This suggests that these celadonites probably crystallized about 70 Ma ago, which was more than 15 Ma after the basalt eruption. These ages are much younger than the apparent ages of the K-rich celadonite from veins dated by Staudigel et al. (1986). This difference suggests that celadonite, in the Troodos Massif, might have formed at different times during a long period following the basalt eruption. Table 18. Radiometric data on pre-Coniacian/post-Albian celadonites and volcanic rocks from the Troodos Massif (Cyprus). Sample

K - A r age

K %

W . K . basalts

(altered)

75 t o 79 Ma

W . R . andesite

(fresh)

84.7

K-richest celadonite vein

6 . 1 to 4 . 9

87.1

Ce ladonites*

Rb-Sr age

(1)

2.3

t-

(2) about 90 Ma

0.8

(2)

K-Ar age i 2 0

rad. A r ( % )

( K 96)

Reference

C2a

6.95

78.6

70.0

3.5

(3)

C2a

6.95

78.0

67.9

3.4

(3)

C2b

7.37

89.0

66.0

3.3

(4)

c3

6.65

83.6

68.9 2 3.5

(3)

Rb (pprn)

Sr (pprn)

87Rb/86Sr

87Sr/86Sr

Kb-Sr age 2 2 0

C2b

142

4.61

89.6

0.79881

71.3 2 1.4

(5)

c3

114

3. 20

104.0

0.80704

67.0 2 1.4

(5)

* Sample C2 is a celadonite chemically analysed by Desprairies and Lapierre (1973). 1) review by Blome and Irvin (1985); 2) Staudigel et al. (1986); 3) H. Bellon in Brest (unpublished, 1983); 4) mean of two analyses by M. Zelvelder (1986) in Brest; 5) P.D. Fullagar (unpublished, 1987) initial 87/86 strontium isotopic ratio assumed to be 0.708.

A fifth example of radiometric dating of celadonite is on samples from the Toarcian volcanics from the Teruel area (Spain). A basaltic formation is precisely bracketed in a very fossiliferous series allowing good biostratigraphical calibration (Gautier and Odin,1985). These Toarcian volcanics should have formed at about 185 Ma (scale of Kennedy and Odin, 1982). Preliminary

39 1 K-Ar analyses showed that the basalts were altered, perhaps in Late Jurassic time or at about 135 Ma (Table 19). K-Ar and Rb-Sr ages measured on samples from one K-rich celadonite veinlet are consistent around 150 Ma. They are about 35 Ma younger than the presumed age of the Toarcian. The celadonite age is older than the apparent age of the altered whole-rock basalts, suggesting that this clay mineral is a geochemically better closed system than is a wholerock basalt system under certain conditions of alteration. The strontium isotopic ratios measured on carbonates from fossils immediately overlying the volcanic formation of the Teruel area (0.7074 and 0.7072) are similar to the ratio accepted for the sea-water at that time (0.7073 according to Faure, 1982). A carbonate vein in the tuff near the celadonite vein, has a lower ratio (0.70656) which indicates the (moderate) influence of the low strontium isotopic ratio of the basalt on the composition of the fluids circulating in its fissures. Table 19. Radiometric ages of Toarcian volcanics and celadonite of the Teruel area (Spain). Sample

K (%)

rad. Ar ( % )

K-Ar age

+- 2

0

Reference

Whole rock basalts

1.51

82

1 3 2 t o 138 Ma

(1)

Celadonite : C59

7.16

94

153 2 6

(2)

Rb (ppm)

Sr (ppm)

87Rb/86Sr

242.7

7.68

93.2

Sample Celadonite : C59

87Sr/86Sr

Rb-Sr age

0.90307

147.7

3.0

2

(3)

1) personal communication, B. Donville in Gautier and Odin (1985; three ages); 2) unpublished results, two analyses by M. Zelvelder in Brest; 3) data by P.D. Fullagar in Chapel Hill; initial 87/86 strontium isotopic ratio assumed to be 0.7074. Three carbonate fractions were measured for 87/86 strontium isotopic ratio; results were: 0.70740 (ammonite); 0.70723 (belemnite) and 0.70656 (vein in tuff).

In conclusion, the radiometric studies of celadonite have been very informative. The main observation is that there is not a single interpretation which can be applied to all outcrops. As a general rule, it seems to be easier to interpret results obtained from pure material with high potassium content. In contrast to glauconies, which sometimes yield ages older than the time of sedimentation (Odin and Dodson, 1982), we know of no example of celadonite giving an age older than the time of extrusion of the host basalt. Some of the Rb-Sr and K-Ar calculated ages do not indicate a significant interval of time between the basalt extrusion and celadonite crystallization. However, most of the calculated ages show a significant difference between the time of basalt extrusion and the apparent time of crystallization of celadonite. This difference may well be real with respect to the time of crystallization of celadonite in the basalt. On the other hand, the possibility cannot be ruled out that a tectonic or a post-extrusive hydrothermal event has reset the age of the green mica. For example, late Oligocene extrusive events are known in the southern Alps around the Monte Baldo area, several tens of kilometres from the celadonite-bearing basalts.

392 However, we presently are convinced that celadonite may crystallize a very long time after basalt emplacement, perhaps as much as 20 Ma. This conclusion agrees with the new data gathered by Hart and Staudigel (1986) and obtained from deep-sea basalts although these authors generally overestimate the interval of time concerned based on overestimated numerical age assignment for lava emplacement.

Formation of celadonite in young oceanic basalts A second approach which may provide information on the interval of time between basalt emplacement and the growth of celadonite is based on observation of present-day, progressively older basalts, extruded at oceanic ridges, and in way of alteration in deep-sea basins sampled during DSDP cruises. This approach has been considered by Laverne and Vivier (1983), and by Alt and Honnorez (1984). Laverne and Vivier (1983) studied the basement basalts from the Galapagos spreading-centre (D.S.D.P. sites 507 to 510) and gave a description of the alteration of crack walls. The fissures in 0.54 Ma old basalt show a thin film (1 mm) of alteration: 'iddingsite' and 'yellowish-green smectite'. Fissures in 0.69 Ma old basalt have similar but thicker alteration (5- 10 mm). The 0.85 Ma old basalt contains cracks which exhibit twice as thick bands of the same type of alteration as in the 0.69 Ma old basalt, plus a small amount of 'dark-green smectite'. In the 2.73 Ma old basalts, the altered vein is not any thicker but it includes a larger proportion of 'dark-green smectite' followed by 'pale-brown smectite'. The names indicated by quotation marks were used by Laverne and Vivier. In absence of mineralogical data, these were tentative. However, according to micro-chemical analyses the 'dark-green smectite' (with 7.8% K,O) probably is celadonite. 0.54Ma

. lhicknesb 0

. 2

4

wall alteration and crack filling

6 1

8 0

2

10

B

3-

14mm

12

4 m

5

B

Figure 21. Sequence of alteration products in basalt fissures from the Galapagos spreading centre. (Modified from Laverne and Vivier, 1983) 1) brown material (black halo with iddingsite; mixture of smectite and iron oxide; 2) yellowish green smectite (green saponite); 3) celadonite; 4) pale-brown smectite (brown saponite); 5) calcite. The thickness of the alteration products increases with the estimated age of the basalts based on their distance from the spreading centre.

393 The sequence of alteration is well described (Fig. 21). Laverne and Vivier also observed that carbonates are common in the 2.73 Ma old basalts, very rare in the 0.85 Ma old basalts and absent in the younger basalts (the numerical ages of the basalts are extrapolated values from palaeomagnetic data, assuming constant rate of extrusion). From these observations, it appears that celadonite crystallized 1 to 2.5 Ma after the basalt extrusion in the Galapagos area. Alt and Honnorez (1984) compared the data by Laverne and Vivier with their own observations and other published data on basalts collected near spreading-centres. They concluded that the formation of black halos requires about 1 Ma, and that zeolite-calcite formation stages occurred in basalts 3.5 to 10 Ma old; these older ages are therefore maximum ages for the time of crystallization of celadonite after basalt extrusion. However, the authors have noted that basalt alteration may be continued for longer than 10 Ma in other examples. NATURE AND GEOLOGICAL SIGNIFICANCE OF CELADONITE

General characteristics of celadonite The celadonitefacies Celadonite-bearing volcanics are widespread. The mineral was first identified from a limited number of continental outcrops; however, it is now believed to be linked with the submarine alteration of volcanic material. The presence of celadonite is therefore a good indication of submarine alteration. Celadonite is not related to rocks of special age, or from a particular depth: it is present in shallow as well as in deep-sea volcanics. We have not yet found evidence that celadonite is more abundant in a particular kind of basalt. The localization of celadonite in basalt seems to be the consequence of physical characteristics (cracks, fissures, vesicles, brecciated habit). Celadonite occurs in five ways: 1) infilling of vesicles; 2) replacement of phenocrysts; 3) diffuse; 4) film around voids; 5) veins. Each one of these occurrences indicates a particular type of primary void and alteration in the massive rock. The kind of initial void determines the celadonite habit . Mineralogicalproperties Celadonite is frequently present in microscopic quantities, and it is often either mixed with another clay mineral or altered during a secondary process of alteration of the host basalt. In spite of this, it has been shown above that celadonite is remarkably homogeneous. Crystallographically, the celadonite structure is well-ordered. Chemically, this micaceous clay is characterized by its tetrasilicic nature. The octahedral layers are specific, too, with their very low aluminium content. The iron content is high with a dominance of the trivalent form. Another aspect of homogeneity is the relatively constant magnesium content. These characteristics suggest that celadonite samples must have formed in a chemically homogeneous environment since structures and compositions are alike in spite of celadonite formation at different ages and different depths.

394 This in turn reinforces the idea that sea-water is the common factor for celadonite formation. There could be variation in the Al-Fe content of the octahedral site. Some celadonites appear to be nearly aluminium free, while in other cases the aluminium content seems to reach a maximum of about 2/5 of the iron content. However, this variation is based on micro-chemical analyses for which large uncertainties remain, both as to the analytical results (results on iron contents are poorly reproducible), and on the precise nature of what has been analysed (frequent admixtures). Some investigators have suggested that the potassium content of celadonite is variable and that this variability suggests the existence of a mineralogical series of celadonitic minerals. Our own observations as well as data from the literature indicate that all pure celadonites are potassium-rich. In summary, celadonite is a homogeneous mineral, rich in potassium, and silicon and poor in aluminium.

The genesis of celadonite Celadonite formation is a widespread phenomenon both in time and space. This facies is the result of the interaction of sea-water on volcanic rocks. Celadonite is usually a subsidiary component of most altered basalts compared to other clay minerals such as saponite. A necessary factor for the growth of celadonite is the presence of voids which allows the crystals to be dominantly euhedral. Celadonite is thus a mineral which essentially crystallizes in usually small geode-like systems. The frequently observed acicular aspect of celadonite is similar to that of many zeolites or even to that of saponite found in geodes from basalts. The role of redox potential appears to be important. There are indications of the presence of oxidizing fluids: before celadonite formation, black halos form at the beginning of alteration of the basalt; during celadonite formation, iron is incorporated in the ferric state; after celadonite formation brown clays form mixed with iron hydroxides, and sometimes there is alteration and oxidation of the celadonite itself. This oxidation can be related to the presence of sea-water as the main circulating fluid. Another interesting aspect of celadonite formation is when it occurs after the extrusion of the volcanics. The following remarks are relevant: 1) there is usually a geologically significant interval of time between basalt extrusion and crystallization of celadonite; 2) this interval of time may vary within a single formation; 3) its duration may be radiometrically undetectable (less than 1 Ma after emplacement of the basalt) to very easily measurable (10 Ma or more). Thus far, no example of apparent age older than the time of emplacement of the basalt has been published for celadonite. This observation can be explained by the fact that celadonite forms in a rock which, at f i t sight, does not contain inherited components as do the sediments where glauconitic minerals form (the inherited initial substrate of glauconitic grains); but inherited argon may be present in deep-sea basalts (see p. 386), consequently, the absence of apparent ages older than the time of eruption in celadonite is most probably related to the fact that celadonite actually forms in voids; by this process, there is no possible

395 inheritance. Therefore, in the cases where pure celadonite may be collected in enough quantity, it may help in obtaining an age estimate for the time of alteration of the basalt, and therefore, a minimum age for the basalt extrusion. Finally, the question of the temperature of formation of celadonite is important. The isotopic composition of oxygen in celadonite indicates that celadonite, usually, forms at temperatures colder than 5 0 T , down to temperatures similar to that of sea-water. On occasion, celadonite may form at temperatures between 50°C and 90°C. Since the term hydrothermal is considered to refer to 'watery fluids of diverse origin ranging in temperature from 50" to 700°C' (Bates and Jackson, 1980), it seems inappropriate to attribute celadonite formation to hydrothermal alteration processes sensu smcto as is usually done. Furthermore, hydrothermal juvenile fluids favour the deposition of material typical of a very reducing environment. For example, during ODP Leg 109 (Episodes, 9, 1986, p. 37), a hydrothermal area was investigated near active vents at the axial rift of the Mid-Atlantic Ocean. Lenses of massive sulfides were repeatedly cored as well as layered deposits of Fe, Cu, and Zn sulfides in a clay mamx. This typical hydrothermal environment is very different from that where celadonite is usually described. The fluids from which celadonite forms might better be described as interstitial with a dominantly marine character. In this view, celadonite can be considered to be a synsedimentary to early diagenetic mineral resulting from re-equilibration, in modified sea-water, of the elements from basalts which were altered by circulation of initially marine fluids. In this water-rich, low temperature environment, elements of the fluids combine in a sequence of new minerals. Silica dominates initially, and carbonates form at the end of the sequence. This suggests a long term evolution of the chemical composition of the interstitial fluids. If celadonite forms from fluids of marine origin, as we beleive it, this might suggest that circulation occurs from the top to the bottom of the celadonitebearing lava pile. However, there is also evidence which implies different circulation. Fluids involved appear to be significantly modified sea-water and this requires a long period of circulation; this could be the case if a thick volcanic pile had been crossed; however, because celadonite is usually more abundant at the more porous top and bottom of flows, the modified sea-water most probably has undergone a long horizontal circulation. In several cases, the temperature of the fluids has been suspected to be higher than that of the sea-water. However, because celadonite usually forms a very long time after basalt emplacement, the basalt itself is cold; therefore, a high temperature for the fluids seems possible only if the fluids circulate deep in the volcanic pile fist, and then toward the top of the pile where celadonite forms.

Comparison between celadonite and glauconitic minerals Similarities In terms of general characteristics, celadonite and glauconitic minerals are fundamentally linked with the marine environment. The two minerals occur in

396 facies that are the result of interaction of sea-water with an existing formation at the sea-bottom: a granular substrate or a porous volcanic flow. The two parageneses are located at the boundary between synsedimentary and early diagenetic processes. The two facies are characterized by neoformed green 2:1 (mica-like) phyllosilicates in which iron is abundant. Moreover, this iron is essentially trivalent, and the clays are dioctahedral. The environment indicated by the two facies suggests a redox potential that is slightly oxidizing allowing crystallization of ferric clays; however, it is clear that reactions mostly occur because we are at the boundary between oxidizing sea-water and more reducing interstitial fluids. The growth of the two green clays involves the circulation of dominantly marine fluids which are enriched by ions removed from the substrate within which the minerals form. In addition to circulation of fluid, a certain amount of closure of the spaces where the clay minerals crystallize is necessary. The two facies are therefore the result of the growth of clays in semi-confned spaces which allows exchanges of ions with dominantly marine waters but at such a rate that the interstitial waters with a higher ion content are protected against the diluting influence of the marine reservoir. The general porosity of the substrate within which the clays form is a common favourable factor for the growth of these two clays.

DifSerences Although both are the result of primarily marine geochemical reactions, the celadonite-bearing facies is significantly different than the glaucony facies. The former is unequivocally linked with long term alteration of submarine basalts. Therefore, celadonite (known only from volcanic outcrops) has a chemical composition which depends on that of sea-water as well as that of basalts which supply ions to the interstitial waters. Also, celadonite often crystallizes in volumes larger than the ones usual for glauconitic minerals. While glaucony is commonly in the form of green grains (due to the fact that grains are the most common and favourable substrate for the growing of glauconitic minerals), celadonite never occurs as grains. It often forms vesicles or vein fillings, and glaucony does not. Another important difference between the two facies is how the marine ions are transported to the growing minerals. For glaucony, most of the ions are carried by fluids mainly formed by sea-water coming from a few decimetres or metres above the glauconitized grains. The fluids have, therefore, a composition which is not very different than that of sea-water. For celadonite, the sea-water does not travel just short distances through the deposit, but circulates very deeply along fractures and has time to be modified by interaction with the rock before reaching the voids where crystal growth occurs. The resulting fluids will be influenced both by the composition of the sea-water and by that of the volcanics. This contrasts with glaucony where the nature of the substrate does not noticeably influence the composition of the glauconitic minerals. Finally, it may be noticed that the rocks and minerals which occur with celadonite and glaucony are different. Celadonite is part of a sequence of in situ formed minerals such as saponite or zeolites, which are not found in glauconitic

397 sediments. Thus, the glaucony facies and the celadonite-bearing facies are different geologically and must be considered separately. From crystallographical and chemical points of view, celadonite and the glauconitic minerals are significantly different. The crystallinity (ordering of the crystal structure) seen on X-ray diffraction patterns plus infra-red spectra, as well as the crystal shapes shown by microscopy, allow us to easily distinguish most celadonites from most glauconitic minerals. The chemical composition is known to be distinctly and consistently different: celadonite is constantly richer in silica, poorer in aluminium and contains a little more magnesium. Thus, the tetrasilicic celadonite differs from the glauconitic minerals. This is a reflection of different environments. Differences in the processes of formation is also indicated by the potassium content of the two materials. Glauconitization involves several stages with growth of smectite followed by both recrystallization and crystal growth (see Chapter C4). Celadonite formation seems to be a one-stage process with the direct crystal growth of a potassium-rich micaceous clay. Moreover, this potassium-rich clay frequently shows a potassium content which is never reached in glauconitic minerals. This reflects a more favourable environment of crystallization for celadonite compared to that available for the glauconitization process. In this sense, the geode system dispersed in the basaltic rock is more suitable for the growth of well-formed minerals than are the micropores in the grains near the sea-water/ sediment interface. The time when crystal growth occurs is another useful comparison. The glauconitic minerals begin to form almost as soon as the granular substrate is available at a favourable depth on the sea-bottom. After a few thousand years the grain is already green and shows the presence of the typical glauconitic smectite. On the contrary, it seems that a million years or more may lapse before the first celadonite appears in the altered volcanics. We suggest that this long time is due to the need of a certain maturation of the sea-water introduced into the cracks of the basalts before that fluid becomes favourable for the crystal growth of celadonite. That "maturation" is mainly an enrichment of ions coming from the volcanics by way of alteration. Alternatively, one may suggest that during the first phase of alteration of the submarine basalts, water circulation may be too rapid to allow celadonite to form in this "too open" environment. In conclusion, celadonite is the ferric clay associated with basalts altered by sea-water just as the glauconitic minerals constitute the ferric clay series of sediments altered at the bottom of the open sea. Although mineralogically very similar, the two components have compositional and crystallographical properties which are clearly distinct and represent different conditions of formation. All considerations indicate that it is proper to use two distinct names for these geologically distinct natural products.

398 ACKNOWLEDGEMENTS

The editor thanks the members of the Geochronology Unit, University of Berne, and especially J.C. Hunziker and A.J. Hurford for their help while measuring Ar isotopes in their laboratory. Dr Buckley (Curator, British Museum, London) and Dr Poullen (Curator, Ecole des Mines, Paris) are thanked for providing samples from the mineralogical collections; Dr Sorbini also provided us with samples from the museum of Verona. In the field, the help of D. Andreolli and E. Pachera, old miners who provided celadonite from Brentonico, and of Don Elio Nizzero, welcoming priest of Priabona (Lessini Mounts, Italy), is appreciated. We thank Miss Madeleine Lenoble for her careful chemical analyses, and Olivier Fay for expert photographic processing. The grant J.E. 4577 "ArchCologie" devoted to "the study of ancient pictural pigments" by the CNRS France partly financed this research.

399

CONCLUSION TO THE STUDY OF GREEN MARINE CLAYS by G.S. Odin Amongst the diverse aspects of the geology of clay touched on in the previous chapters, several merit to be recalled in this postcript to the study of green marine clays.

Variety of green marine clays Table 1 summarizes the variety of clay minerals which have been identified in the four facies described in this volume. Table 1 indicates that there are similarities between the clay minerals characterizing the two shallow marine facies on the one hand, and between the two open-sea facies on the other hand. Clay minerals from the ironstone fucies are specifically 1) very rich in ferrous iron, 2) comparatively poor in silicon and magnesium, and 3) comparatively rich in aluminium. These minerals do not fully represent the sedimentary geochemical environment of deposition because their excellent crystallinity probably results from diagenetic reactions. Clay minerals from the verdinefucies are specifically 1) dominantly ferric and magnesian, 2) comparatively rich in silicon (nearly tetrasilicic) and magnesium, and 3) comparatively poor in aluminium. These minerals are equally di- and trioctahedral. The 7 8, phase has been obtained pure enough from New Caledonia and Guinea for detailed analyses to be undertaken. This phase is a new mineral species of which the pure 7 8, spacing has been verified by HRTEM. Furthermore, this technique has shown that 14 8, microcrystals exist (pure material has not been obtained up to now); to the editor's opinion, this clay is also a new mineral species. The mixed-layer (7 A-14 8,) is rare but has been distinctly observed using HRTEM. The hybrid clay quoted in Table 1 has a behaviour intermediate between a smectite and a swelling-chlorite and gives poor figures using HRTEM. This phase is predominant in a small number of samples and appears to represent a particular variety of the verdine facies. The significance of the 10 A phase is not entirely clear. The 10 8, spacing is very evident using HRTEM; preliminary micro-analyses indicate that the corresponding mineral is aluminium-rich (in contrast to all other phases of the facies), poor in iron, and poor in potassium which is surprising for a mica. The three phases: 14 8, phyllite V, (7 8,-14 A) mixed-layer and 10 %, clay mineral, probably result from a synsedimentary recrystallization of the previously formed 7 8, phyllite V (in contact with sea-water) in relict sediments. Clay minerals from the gluuconyfucies represent a continuous series from a potassium-poor smectite (glauconitic smectite) phase to a potassium-rich mica (glauconitic mica) phase. All members of the family are rich in ferric iron. Post-sedimentary recrystallizations may well provoke chemical differentiation

400 in this series (aluminium enrichment for example) but, at the origin, it seems that the series is very homogeneous and clearly distinct from other ferric iron-rich 2: 1 clay minerals. Celadonite-bearing rocks have essentially been considered for their micaceous clay mineral, which is characterized by a much higher silicon content (nearly tetrasilicic tetrahedra) and slightly higher magnesium content compared to glauconitic minerals. The common point is the high ferric iron content. Celadonite is usually potassium-rich but alteration occasionally modifies this situation. In addition to the dioctahedral celadonite, the trioctahedral saponite (magnesian smectite) is more common in the celadonite-bearing facies; its colour may be green or chocolate brown; nontronite is sometimes quoted as well. Mixing of these clays frequently gives the impression that a continuous series smectite-mica exists; but this has not been proved. In addition to the four facies quoted above, the sea-bottom also harbours a number of other parageneses where authigenic clay minerals are formed (they are not discussed in this volume) e.g., bentonites (smectitic) horizon or fibrous clays in volcanic areas, and red clays in deep-sea basins. Table 1. Variety of clay minerals in green marine clays

Structure 1:l (7

A)

Ironstone berthierine (trioctahedral )

2 : l (mica) dioctahedral

Verdine 7

Celadonite-bearing rocks

A

phyllite V (di-tri)

present

2 : l (smectite) dioctahedral 2 : l (smectite) trioctahedra 1

g 1 auconitic mica

c e ladonite

glauconitic smectite

nontronite

saponite

A

2 : l : l (chlorite)

chamosite (trioctahedral )

14

2 : 1 :1/2 :1 hybrid-clay

swe 1 1 ing-ch 1 orite

phyllite C

1 :1/2 :1 :1 mixed-layer

Glaucony

phyllite V (di-tri)

present

Concerning the variety of marine ferric clay minerals, the significance of the diverse clays can be assessed in terms of evolutionary stages in the case of glaucony and verdine facies. The initial stage of the glauconitization process appears to be a comparatively rapid crystallization of smectite; this reaction gathers most necessary cations. The system then evolves towards greater stability with little mineralogical change (potassium addition) following a recrystallization process; but recrystallizations occur together with continuous crystal growth. Similarly, it appears that the 7 A phyllite V is the first stage of the verdinization process gathering most of the necessary cations while the second stage seems to be a slower reaction resulting in crystallographic re-organization

401

and finally leading to a composite assemblage including diverse mineral phases as listed above (Table 1).

The relationship between green marine clays and their local environment There are three types of sea-water/rock interactions resulting in the growth of clay minerals. The first is characterized by a large body of fluid phase wholly favourable for crystal growth; this leads to oolitic structure and diffuse clay formed by precipitation in this fluid. This relation is well illustrated by the ironstone facies. The second type of relation is characterized by clay minerals which only grow in a localized microenvironment, usually the micro-voids in previously deposited sediment particles. As above, there is no direct chemical relation between the nature of the deposited substrate and that of the authigenic minerals in the sense that neither cations from the substrate, nor a fortiori its mineralogical composition influence the nature of the neoformed minerals. The third type of relation is characterized by clay minerals which grow from a fluid whose composition is partly influenced by the immediately surrounding rock. This case is illustrated by the celadonite-bearing facies or by clay minerals formed from subaerially erupted volcanic ashes (bentonitic clays). Obviously, there are mixed relations between these three types. In this volume, the clay minerals grown in localized microenvironment have been illustrated in detail with the glaucony and verdine facies. This relation between a general environment which is not favourable and a restricted porous physical substrate constitutes a very efficient circumstance for genesis of a variety of clay minerals. The notion of semi-confinement in a localized microenvironment is a predominant factor in this connection. Other marine minerals such as phosphate or pyrite as well as minerals formed later in the domain of diagenesis also depend more on the nature of the local semi-confined microenvironment than on the general environment. This microenvironment at the boundary between fluids characterized by different pH and Eh conditions, allows combination of ions mostly originating from the large reservoir constituted by seawater (alternatively interstitial waters) and combining into sheltered restricted volumes. These micro-voids are favourable for crystal growth but need to be fed with external cations; in this sense, the semi-confined microenvironment acts as a trap slowly extracting cations from the externai fluids by immobilizing them in stable crystallized structures.

Activity of iron in the sea The study of marine green clays is one of the bearings of geochemical activity of iron in the marine environment, more precisely at the interface between sea-water and rocks. A large proportion of iron is present in all green clay minerals; therefore, the availability of that element is one of the determining factor when dealing with significant quantities of authigenic clays.

402

There are two sources for iron in the sea: fluvial waters and volcanic rocks. The study of the verdine facies has clearly indicated the general relationship between fluvial iron influx to the sea and extent of the paragenesis on the one hand, and the local relation with weathered volcanics on the other hand. The continental origin of iron for most oolitic ironstones has long been accepted; the influence of local volcanic outcrops has also been reported. Concerning glaucony, the link with fluvial iron is less obvious because the facies only develops at depth usually located far from a river mouth on passive continental margins. The link with volcanic iron may be observed in active margins such as along the western coast of the Americas or volcanic highs in the Indian Ocean. Finally, celadonite utilizes volcanic iron alone. In order to represent these two origins, a scheme has been suggested and improved during the last ten years (Odin, 1975b; Odin and Matter, 1981; Fig. 1). Figure 1 is an improved version of the model which shows the two sources (noted Fe) and resulting two paths of iron in the sea. The nature of the two main sources is different; one is punctual (i.e., river mouth) and the resulting zonation is reasonably clear; the other is more diffuse and comprises (amongst others) all mid-oceanic ridges and active margins, so that the resulting zonation is less clear. The model proposes five main zones often characterized by specific green clays but also other compounds. For the fluvial source the zones are as follows. Zone I is the area of deposition of detrital iron near the continent. In the submarine delta, iron is partly reduced where Eh is low; it becomes soluble, available for local recombination (pyrite, siderite), or migrates to feed other zones. Zone 2 is the zone of positive Eh where oxi-hydroxides of iron form through biochemicalkhemical precipitation or in situ alteration of other iron-rich minerals. This zone mostly concerns the intertropical coast line where very oxygenated waters are present between 0 m and 10 m depth. The difference between the iron of zone 1 and zone 2 is that the former remains within the deposited sediment while in the latter, iron has migrated out of the deposited body. Zone 3 favours the verdinization process at depth mostly between -20 m and -60 m also in a tropical climate. The oolitic ironstone facies could be located at the boundary between zone 2 and zone 3 but persisting problematic factors other than depth and distance to the river mouths (notably the precise Eh conditions) remain to be identified. Zone 4 is characterized by the glauconitization process occurring deeper than 60 m and well removed from the fluvial input. Zone 5 is the area where iron is incorporated in deep-sea clays (often iron smectites) or other iron-rich compounds e.g., ferro-manganese nodules or amorphous silico-ferriferous complexes such as those identified by Frohlich (1980, 1982) in the Indian Ocean. Concerning the juvenile source of iron the depth factor mainly influences development of zone 3; therefore, a volcanic island has been shown in Figure 1. The equivalent of zone 1 is shown as 1b in Figure 1, and corresponds

403

Figure 1. Geochemical paths of iron in sea-water. Two sources of iron (fluvial input to the sea and volcanic rocks altered in contact with sea-water) determine two paths which both end in deep-sea basins. Along those paths iron is geochemically active and contribute to the formation of a variety of clay minerals.

to the extruded volcanic rocks at the direct contact with sea-water; the rest of the ridge may well be covered with sediments. Sea-water circulates in these rocks which are altered; minerals equivalent to pyrite in deltaic sediments are formed in the newly deposited body; celadonite is in this zone with other clay minerals and oxides of the same paragenesis. The equivalent of zone 2 i.e., minerals crystallized out of the deposited body from ions originating from it, could be identified in sulfides or clays (smectites) deposited near deep oceanic hydrothermal sources; this is shown as 2b at the contact with zone lb. The equivalent of zone 3, shown as 3b in Figure 1, is easily identified in present seas around volcanic oceanic islands and other volcanic outcrops on passive margins (New Caledonia and Cap Vert, Senegal, respectively). Zone 4b is illustrated by glaucony formed on rising parts of the sea-floor (Chatham Rise, N.Z.), on active margins (West Americas), or around volcanic islands (Kerguelen). Zone 5 can be considered to be fed both with juvenile and continental iron and is common to the two paths.

Mechanism of genesis of clay minerals The detailed observation of marine green clays leads us to reject the hypothesis of a transformation process (i.e., the modification of inherited clays by cation exchange in a permanent crystal lattice) as a possible mechanism of origin. This is particularly significant since glauconitization has long been considered a good example supporting that mechanism. Particular arguments resulting from the comparison of the diverse facies have been gathered in this volume. The comparison between verdine and glaucony is highly demonstrative since the same sorts of granular substrate are generally involved as the guests for two different mineral geneses. For example, kaolinite-rich faecal pellets and mica flakes i.e., a 1:1 and a 2:1 layer silicate structure respectively, were observed in the two facies. Kaolinitic substrates harbour the genesis of 2:1 glauconitic smectite while micaceous substrates harbour the genesis of 1:1 or 2: 1:1 verdine minerals. Therefore, the mineralogical nature of the substrate is not a key factor and the transformation process hypothesis is not valid here. In contrast, the crystal growth process is able to explain all our observations. We foresee therefore a progressive abandoning of the transformation hypothesis which was based on too superficial studies. This revision of the ideas on the main process of clay genesis in the sea was already supported twelve years ago by the present editor (Odin, 1975a). Thanks to observations initially made on glaucony, more and more specialists justifiably consider this new approach on clay genesis, i.e. process of neoformation better than transformation, in other domains of the geochemistry of the earth surface.

Relationship between facies and minerals The present study of green marine clays has proposed a means of providing proper designation of the green material observed in sedimentary formations. A distinction has been made between the facies i.e., the morphological, mineralogical, and sedimentological characters observed or deduced from a rock, and the minerals involved in these facies. One kind of mineral corresponds to a single facies but this facies may favour the formation of several minerals (Table 1). Specific names have been suggested (see glossary), others could be preferred but the important thing is to avoid the use of a single name for the two geological and mineralogical notions. We may briefly recall a case where the lack of nomenclatural distinction has led to confusion and misunderstanding. This concerns "glauconite". Most English writing geochronologists use the word "glauconite" to designate both the green grains and their mineralogical components. As a result, the green grains are said to be more or less potassium-rich as a whole, the duality of significance between inherited components of the substrate and the authigenic green clays within the green grains is not conceived and the radiometric ages calculated from this material are highly erratic. The chronometer is thus considered whimsical and therefore, practically unreliable for dating purposes. The fact is that the lack of nomenclatural distinction firstly between green grains

405 and their dual mineralogical (and geochronological) significance and secondly, between the various mineralogical compositions of the authigenic phase (the more or less closed and evolved green clay minerals) did not allow English speaking authors 1) to easily understand the different reasons why "glauconites" do not fit the common geochronological model used to calculate an age and 2) to select those "glauconites" which are the most able to fit this model. The selection of the geochronometer necessary for any radiometric study is, therefore, usually not properly done. This has been discussed elsewhere (Odin, 1982d), and this discussion was only possible after a complete understanding of the geological significance of the diverse materials previously similarly called "glauconite" i.e., green grains, substrate, authigenic components of the green grains, and more or less evolved (K-rich) authigenic phase. Another aspect of the "facies-mineral"relationship in green marine clays, is the specificity of the mineralogical components as a function of the environment of genesis i.e., of the rock facies. Each facies has its own mineral species and the same species is not common to different facies. The false problem of the celadonite/glauconitedistinction repeatedly approached in the literature (Hendricks and Ross, 1949; Foster, 1969 and many others) always leads to the same conclusion when the study is seriously done: the green clay(s) formed in granular substrates at the sea-bottom is mineralogically different from the green clay formed in voids of submarine basalts. In this sense, the hypothesis made by specialists of nomenclature, that if a green clay with the composition of celadonite was formed in marine green grains, that clay should be termed celadonite (and not glauconite), is not relevant. This is because glauconitic minerals are the geochemical answers to the conditions present in granular substrates, at the sea-bottom, and these conditions are different from those where celadonite forms. As a result, there is no reason for a celadonite-like mineral to form where glauconitic minerals usually form. This claim only assumes that the variety of environmental conditions is so wide on earth that there is little probability for the realization of precisely similar local conditions in distinctly different general environments.

Acknowledgements The present volume results from the collaboration of a number of colleagues and collaborators with the editor. Their help is deeply acknowledged. Some of these colleagues have co-authored the preceding chapters; the editor would also like to mention that: - he appreciates the honour to begin this volume with a Foreword written by Professor Georges Millot whose teaching of the Geology of Clays allowed me to discover a fascinating subject about twenty years ago; - the contribution of sample collectors in many areas, all over the world, was a prerequisite contribution to the present work because the editor has never obtained himself the necessary financial help for sampling personally in any of the numerous tropical countries studied in this volume; these contributions have been quoted in the relevant chapters and are greatly appreciated;

406 - detailed studies in regional sedimentology, mineralogy, and isotope geology were made in coordination with specialists whose expertise was essential; the technical assistance of Madeleine Lenoble, Eliane Lebrun and Gilbert Richebois allowed proper chemical definition of the minerals studied; - in the absence of specific financial support for preparation of this volume, English improvement of manuscripts was a problem with which Professors S.W. Bailey and P.D. Fullagar especially were kind enough to help the editor; the introduction and conclusion to the volume, the glossary, and Chapter C1 were improved by John F. C. Thompson who translated the French Priface; - the practical realization was made possible thanks to Solange da Matha who kindly typed all preliminary manuscripts, and to Chantal Odin who perfectly prepared the tables; the quality of the photographic illustration is due to Olivier Fay's expertise; Jaqueline Carbone nicely contributed to the reproduction of the successive manuscripts and was responsible for mailing; - the editor is alone responsible for the final presentation; for preparation of the latter, he is indebted to RenC Utolle and Andri Mariotti for permission to use their versatile Macintosh facility during the twenty four months during which this volume was prepared.

It is a pleasure to the editor to deeply thank all these friends, colleagues, and collaborators.

407

GLOSSARY berthierine Berthierine and chamosite are two clay minerals of similar chemical composition. Berthierine is the 1:1-type mineral and its composition was reviewed by Brindley (1982; see this volume Table 2, p. 180). Brindley describes berthierine as 'a ferrous, aluminian analogue of lizardite' i.e., a trioctahedral clay mineral. The tetrahedral sheet is made of 1.1 to 1.5 Si cations for two sites; the rest is aluminium. The octahedral sheet is mostly made of Fe2+ (1.5 to 1.8 cations) and A1 (0.6 to 1.0 cations); Fe3+(0.0 to 0.3) and Mg (0.1 to 0.6) fill the three sites of the sheet. The mineral usually shows well shaped 001 peaks on X-ray diffraction patterns (see Fig. 3, p. 41); berthierine is often present in oolitic ironstone formations of Mesozoic age. burial Burial is the process by which a deposited sediment is covered by a new sediment. Practically, a buried marine sediment is a sediment which is no longer in contact with sea-water. Before burial, cationic exchanges exist between the deposited sediment particles and overlying fluids. Burial separates the synsedimentary geochemical reactions (those mostly concerned in this volume) from the diagenetic ones. celadonite The green clay-size fraction rich in ferric iron, poor in aluminium, rich in potassium, and found in volcanic rocks is a tetrasilicic clay mineral called celadonite. Celadonite is a dioctahedral mica which differs from other clay minerals in the same paragenesis such as green saponite: a trioctahedral smectite (p. 343; 362-365). Celadonite also differs slightly but constantly from glauconitic minerals which are not tetrasilicic, slightly poorer in magnesium, and more or less potassium-rich (see Table 10, p. 373). Ideal celadonite composition is:

K (Fe3+,Mg) Si, O,, (OH), less than 0.1 atom of tetrahedral A1 and less than 0.3 atom of octahedral Al per formula unit can be added to this formula (see also mean formula proposed in p. 369). Infra-red spectra (sharp peaks) are specific compared to glauconitic minerals (see Bailey, 1980; Fig. 9 and 10: p. 354 and 355 in this volume). X-ray diffraction patterns are also specific with sharp peaks and d-060 smaller than 1.51 8, (see X-ray diffraction p. 349, and a general comparison between celadonite and glauconitic minerals in Table 12, p. 377). Celadonite often presents idiomorphic microcrystals in the form of thin, more or less elongated laths up to 25 pm in length (see Fig. 5, p. 347). The mineral is usually found filling vesicles, veinlets, and other small voids of the

408 porous portion of lava flows which have been altered in submarine conditions. Sometimes, the mineral appears to form very late after the emplacement of the volcanic rocks. The mineral was used as a painting pigment in ancient times; it is known more today as a common product of alteration of deep-sea basalts. The term celadonite should be restricted to those clay minerals with a known composition as defined above because in natural samples, it corresponds to geologically and thermodynamically restricted conditions of genesis. The term celadonitic minerals is recommended for tetrasilicic clay minerals with slightly different compositions or synthetic minerals for which the composition is rarely known with precision. In case of doubt between celadonitic or glauconitic minerals or femc illite the term green clay is the only acceptable.

c hamosi t e Chamosite is a 2:l:l chlorite of chemical composition similar to that of berthierine (see X-ray diffraction patterns in Fig. 3, p. 41 and Fig. 4, p. 43; chemistry in Tables 2 and 3, p. 19 and 20). Therefore, chamosite is a ferroan chlorite; it has been defined from the iron ore near Chamoson (Chapter Al). Mineralogically, chamosite is a trioctahedral chlorite for which Fe2+cations are dominant. The end member may be written: (Fe2+5,Al) (Si, Al) O,, (OH), (Bailey, 1980). Formulae have been calculated in this volume for a chamosite collected from Chamoson and a bavalite from Bas Vallon (see p. 20). The relationships between the chemical compositions of chlorites, chamosite, berthierine, and green clays from the verdine facies are given in Chapter B5 (Fig. 13, 19,20, p. 182, 201, and 202; see also p. 19). The term chamosite is generally used to designate chloritic green clays from oolitic ironstone formations.

clay mineral structures Phyllosilicates related to clay minerals are of four types i.e., 1:1,2:1,2: 1:1, and fibrous. The layer of a 1: 1-type clay mineral is made of two sheets: one tetrahedral sheet and one octahedral sheet. When trivalent cations occupy the octahedral sites, the six charges available are compensated by twocations and the mineral is said to be &octahedral (kaolinite); when bivalent cations occupy the octahedra three cations are needed and the mineral is said to be octahedral (serpentine). The 2: 1-type clay minerals are made of three sheets: two tetrahedral sheets surrounding one octahedral sheet. In this group, illite contains cations (potassium) in the interlayer and has a 001 distance of 10 A like micas; smectites contain water in the interlayer and have a wider 001 distance, usually near 14 A. The 2:l: 1-type clay minerals are made of four sheets; three of them form a 2:l layer, the fourth is an hydroxide sheet (with an octahedral-like arrangement) in the interlayer; the d-001 is about 14 A; this group is formed of chorites; most of them are wholly trioctahedral (p. 41), but a few mixed di-tri structures are known (see p. 194). Fibrous clays are formed of ribbons instead of sheets and are not discussed in this volume. Intermediate clay minerals exist

409 (see swelling-chlorite, below). Most of these structures were observed and described from authigenic marine minerals from present-day sea-bottom (see Table 1, p. 400). From a purely mineralogical point of view, an interesting feature of the verdine facies is that it mainly comprises clay minerals which are intermediate structures e.g., di-tri serpentine, di-tri chlorites (see phyllite V below) and intermediate smectite/swelling-chlorite(see phyllite C below).

crack The green grains found in marine sediments often show deep cracks at their surface. These cracks correspond to micro-crevices which narrow towards the interior of the green grain. They develop as a result of a differential crystal growth process more efficient in the centre of the grain (more confined than at the surface) rather in the way that cracks develop at the surface of a rising loaf (e.g., Fig. 4 and Fig. 6, p. 253 and 257). Cracked green grains represent an evolved stage of glauconitization and result from a volume increase of the initial grains (p. 232, 246, 320). Cracks may be filled by a late genesis of green clays; in such a case, the green grains become highly-evolved. crystal growth The term crystal growth is used here to describe genesis of minerals by crystallization from ions extracted from a fluid. This process was found to be at the origin of the verdinization (see Chapter B6) and of the glauconitization (see p. 244, 293, or 319). The nature and extent of crystal growth is strictly influenced by the immediately surrounding conditions i.e., the physico-chemical conditions such as temperature, pH, Eh, and the availability of cations (presence, relative proportions) and geochemical activity of cations. The cations used for crystal growth become depleted in the microenvironment; therefore, the possibility of renewing these cations by exchange between the microenvironment of growth and the general environment is an important factor influencing crystal growth.

diagenesis We have employed the term diagenesis in its restricted sense usual in French literature (Dunoyer de Segonzac, 1969). It designates the physical and chemical reactions occurring in a sediment after burial and before heating, tectonic pressure, or weathering. Diagenetic reactions occur when predominant exchanges with fluids overlying the sediments (meteoric water or sea-water) have ceased. This volume mostly considers clay minerals formed before diagenesis. There is an interesting aspect of the relationship between diagenesis and marine green clays. Diagenesis (and metamorphism) are usually considered to be generally homogenizing processes. In contrast, green marine clays tend to become more heterogeneous during diagenesis. Green marine clays are initially in equilibrium with sea-water (a very homogeneous environment) therefore, their chemical composition is initially very homogeneous. During diagenesis, the presence of diverse detrital components surrounding the authigenic phase will diversely influence the chemical nature of the fluids immediately surroun-

410 ding the green clays, whose composition will tend to be re-equilibrated with this new environment. For example, the previously homogeneous glauconitic minerals will tend to become heterogeneous (essentially for the Fe/Al ratio) during diagenesis. It is of importance to distinguish this diagenetically acquired heterogeneity from the proper characteristics of the green marine clays.

evolved The green grains of the glaucony facies result from the evolution of granular substrates within which neoformed glauconitic minerals are formed. Depending on the stage of evolution, the grains are said to be nascent, slightly-evolved, evolved, and highly-evolved (see Fig. 9, p. 242; Fig. 12, p. 320). Initially dominant (nascent stage) the substrate progressively disappears and the authigenic minerals become predominant and exclusive (highly-evolved stage). The composition and geological significance of the green grains differ, (duration of evolution and concomitant lack of deposition less to more important) when slightly-evolved to highly-evolved grains are concerned. This point is fundamental when radiometric dating is applied to glaucony with the aim of estimating the numerical age of a sediment (see Fig. 13, p. 322). faecal pellet Excrement of animals frequently observed in sediments where mud eaters live takes the form of faecal pellets. In this volume, the term designates ellipsoidal 0.1 to 2 mm large pellets probably generated by worms. They constitute a favourable substrate for verdissement in the glaucony (Fig. 4, p. 230; Fig. 15, p.328) and verdine (Fig. 7, p. 92; Fig. 7, p. 145; Fig. 6, p. 213) facies. An equivalent word is coprolite, but this term is more general and may include larger elements originating from vertebrates and implies a mineralization transforming the original excrement into a lithic object. glauconite The term glauconite has been used to designate various geological notions including 1) the green grains in the marine sediments (the glaucony facies); 2) the minerals composing these green grains (the glauconitic minerals); 3) the micaceous end-member of the glauconitic minerals (the glauconitic mica). The term glauconite should be used with caution and it is recommended to adopt the terms suggested in parentheses and discussed below, to avoid confusion (see Table 1, p. 295). glauconitic mineral The clay minerals characterizing the glaucony facies are called glauconitic minerals. These are specifically 2: 1 dominantly dioctahedral clay minerals, rich in ferric iron, with more than 0.2 Al cation per formula unit (4 Si cations) in the tetrahedral sheets; in contrast, there is less than 0.1 A1 cation for celadonite, (see above), and with potassium in the interlayer (a mean formula is given p. 314). Chemical results are reported Table 1, p. 236 and Table 1, p. 268. Glauconitic mica shows specific infra-red spectra with broad peaks (see p. 353

41 1 to 359, Fig. 11, 12, and 13) and X-ray diffraction patterns with generally broad peaks (see Fig. 7 and 9, p. 310 and 313), d-060 larger than 1.51 A, and 023 diffraction peak much smaller than 130 peak (see Fig. 7, p. 349, and Table 2, p. 350). Glauconitic minerals constitute a family whose members differ chemically mostly in their potassium content with a K - p r glauconitic smectite and a K-rich glauconitic mica as end-members. The K-poor members are said to be open and the K-rich, closed; d-001 varies between 14 A and 10 A respectively. The internal structure of the glauconitic minerals may be disordered (generally for K-poor minerals and unburied K-rich minerals) or ordered (for K-rich minerals probably affected diagenetically). Glauconitic minerals are only known from marine sedimentary formations and are specific to the glaucony facies at the time of genesis. After early burial, glauconitic minerals are modified following deep burial, tectonic or heat processes. The resulting modified minerals show a compositional range wider than the original synsedimentary glauconitic minerals (especially A1 increase concomitant with Fe decrease). These diagenetic minerals should be distinguished from glauconitic minerals sensu smcto and we suggest for them the terms "diagenetic green clays" or diagenetically affected glauconitic minerals.

glauconitization The process by which a substrate deposited at the sea-bottom is progressively modified to a mass of glauconitic minerals, is called glauconitization. Glauconitization frequently occurs from granular substrates but also at the surface of boulders or of hardgrounds deposited at the sea-bottom. Glauconitization consits of both morphological (Fig. 9,p. 242; p. 27 1) and mineralogical evolutions (Fig. 6, p. 234; p. 243; Fig. 11, p. 266). During glauconitization, there is both a progressive alteration of the initial substrate and a genesis of green marine clays of which the proportion increases modifying the initial colour of the substrate. The synsedimentary marine green clays formed are the glauconitic minerals; they form in cold (less than 10°Cto 15'C), comparatively deep waters (more than 60 m depth) and well removed from continental detrital input to the sea (see p. 318 to 330). Glauconitization is a comparatively long process which may need several lo5 years to be entirely completed (p. 241242; Fig. 4, p. 304; Fig. 12, p. 320; p. 327-330). glaucon y (Word root: gluucos (Greek) green to blue-green colour). The marine facies characterized by a green pigment made of 2:1-type clay minerals (glauconitic minerals) is called the glaucony facies. The corresponding adjective should be glauconious (or glauconiferous; "glauconieux" or "glauconifkre" in French) but the term is not yet in English usage, and glauconitic (which should be restricted to the minerals) is used instead. Glaucony facies develops by interaction between sea-water and favourable deposited substrates; this is usually called glauconitization; to be coherent with the term glaucony, this process should preferably be called: glauconization. Depending on the form of the initial substrate, glaucony may be in the form of

412 green grains or of thin films at the surface of boulders or hardgrounds (see examples of habits in Chapter C2, p. 252-264, and Chapter C3, p. 283-289; classification of habits in Table 2, p. 297).

glossary Dictionary explaining specialized words. In this volume, the glossary serves as some sort of summarized review of the problems discussed with our proposed answers. Various words have been used in this volume with a signification which is possibly different from the usual one, other terms could be preferred by others. In general, definitions here correspond to specific geological notions which are useful to explain and understand the green marine clays but which do not claim to represent a definitive terminology. green clay The green pigment observed in sediments or rocks submitted to marine alteration and made of clay minerals is called here green clay. This general term should be used as long as a mineralogical study has not solved the problem of the precise nature of that clay (see p. 9; 43; 346). Various green clays were submitted for study to the present editor during recent years (material usually suspected to be glaucony) these included: glauconitic minerals, verdine minerals, chloritized biotite, inherited chlorite, and celadonitized pyroxenes. green grain Green grain is a term preferred in this volume when sand-sized constituents of a sediment have not been mineralogically analysed. This allows us to avoid confusion between authigenic glauconitic grains, authigenic green grains of the verdine facies, inherited green grains. X-ray diffraction study easily solves the question of the precise nature and geological significance of the green grains; green grains constitute a particular habit of the green pigment. green pigment In the first stage of the study of a green sediment or rock, green pigment is the term which should be used before more mineralogical precision is made available. For example, the green colour of a sand can be due to ferrous oxides or to clay minerals; it is important to use a sufficiently wide term for description in order to avoid confusion on the geological significance of the green material. The green pigment can be subdivided into two habits: granular or diffuse. infilling In this volume, infilling designates an internal mould of microtest. This test is a favourable semi-confined microenvironment where authigenic green clays may form whatever verdine (Fig. 7, p. 92) or glaucony (Fig. 6, p. 257; Fig. 5 and 6, p. 286-287) facies is concerned. The test is filled with this neoformed material by crystallization in voids; but the green clay could also replace a fine grained constituent previously accumulated within the test. After dissolution of the test (carbonate or silica), the green clay may form free green grains.

413

iron ore An ironstone from which iron may be profitably extracted is an iron ore; today, a content of 15% of iron is usually accepted as the limit between iron ore and ironstone.

ironstone A rock containing a substantial proportion of iron is called ironstone. One of these rocks contains crystallized iron compounds under the form of ooids and is sedimentary in origin; this is the oolitic ironstone, the facies discussed in part A of this volume. The oolitic ironstone facies is known to be formed at the boundary between continent and Ocean in marine water (Fig. 5, p. 45); ooids as well as the groundmass may be formed of trioctahedral green clays of three types: berthierine, chamosite and swelling-chlorites (see these words). Some oolitic ironstones are iron-rich enough to be used as iron ores; a well known example is the "minette" of Lorraine (France).

neoforma t ion The de novo formation of minerals (i.e., from ions) is called neoformation. The neoformation is a process generally opposed to the transformation (see below). Neoformed minerals form by crystal growth. This is the case for the glauconitic smectite (see p. 244) or the phyllite V at the sea-bottom. These crystal growths trap ions indirectly coming mostly from the sea-water reservoir through the filter made by the porous substrate where minerals are formed. The recrystallization process invoked as the second stage of glauconitization (see p. 273) is also considered as a neoformation. A similar process (recrystallization) is acce ted for the evolution of the 7 A phyllite V into other clay minerals such as 14 phyllite V (p. 218) , or the postulated modification of the initial minerals of the ironstone facies into berthierine and chamosite. During recrystallization, the new minerals appear to form from ions but these ions are mostly taken from the pre-existing minerals. For example, the early glauconitic smectite is locally dissolved and immediately recrystallized in a more stable (K-richer) form.

K

ooid

(definition by S . Guerrak) The term ooid (also spelt as oolith, or oolite which is more correctly used to describe an oolitic sedimentary rock) is originally a French term (oolithe) derived from two greek words: oon = a roe of a fish, and lithos = a stone. Ooid is applied to spherical and ellipsoidal grains composed of a nucleus surrounded by an envelope formed of concentric laminae. Therefore, ooid definition implies an external morphology and an internal structure. These coated grains show different shapes and have various mineralogies. Ooids with a radially, tangentially, or randomly arranged envelope are mainly found in carbonates (Richter, 1983). Ooids with only tangentially developed cortices are common in oolitic ironstones of Proterozoic and Phanerozoic age pig. 1, p. 32), and uncommon in Recent sediments. They may be composed of chamosite, berthierine, hematite, magnetite, goethite (in Recent

414 sediments) or maghemite, siderite or secondary calcite. Ooids may be phosphatic or manganiferous; oolitic phosphorites and oolitic manganese deposits are common; sulfidic (pyritic) or cherty (siliceous) ooids are rare and are commonly considered to have originated from replacements. Nuclei of ooids consist of detrital or bioclastic grains, or oolitic intraclasts made up of broken ooids or entire pre-existent ooids. The size of the ooids generally ranges from about 0.1 to 2.0 mm. An ooid with a diameter between 2 and 10 mm is called a pisoid (or pisolith). Partly following Carozzi (1960), several morphological types of ooids can be distinguished. 1) Superficial ooids consist of a nucleus surrounded by a single (or double) concentric layer; this layer can be discontinuous as a result of incomplete accretion or abrasion phenomena. These ooids are different from pseudo-ooids defined as ellipsoidal or spheroidal grains, associated with common ooids, but with no coating of concentric layers. 2) Common ooids can be concentric or eccentric as defined by Knox, (1970), they show successive layers forming an envelope distributed concentrically or eccentrically around a single nucleus or a pre-existing m i d (superficial, common, or broken). 3) Composite ooids (or complex, or multiple) consist of several ooids embedded together in a new envelope (see Fig. 1, p. 33 and 2, p. 35). 4) Polynucleate ooids have several distinct nuclei of the same or different nature. 5) Spastoliths are ooids showing a flattened, distorted or sinuous form. They most probably were generated by the compaction of sediments when they were still in a plastic state. They might be the result of diffusion phenomena during diagenesis.

oolitization (definition by S. Guerrak) Oolitization is the process by which any nucleus is coated to obtain an mid. This involves several successive phases of coating and suspension, during which a superficial ooid is built, deposited, reworked and coated, deposited again and so on, to a final mid. Each period of oolitization records the physicochemical conditions of the immediately contemporaneous environment. Various hypotheses for oolitization are summarized in this volume (p. 30-39). perigenic A new term proposed by Lewis (1964). A sedimentary constituent formed at the same time as the rock of which it constitutes a part but not at the specific location in which it is now formed in that rock. Green grains formed at the top of the slope but transported deeper soon after, are perigenic (see p. 229, 283, and 326). phyllite C Phylitte C is a marine green clay mainly characterized by a huge 14 8, first order 001 peak on X-ray diffraction patterns (Fig. 10, p. 95) which indicates a great deficiency of scattering power in the interlayer hence mainly a vermiculite

415

or smectite. The expansion on solvation with ethylene-glycol indicates mainly smectite, but the eak does not go all the way to 17 A nor does it collapse completely to 10 on heating. The mineral is therefore different from phyllite V, chlorite, or glauconitic minerals which never show that behaviour (Fig. 8, p. 170). The chemical composition indicates characters common to phyllite V with high Fe3+and Mg contents, low aluminium, and very low Fe2+contents (see Table 3, p. 183). Chemical1 , phyllite C appears intermediate between a glauconitic smectite and the 14 phyllite V. Its precise nature remains to be discovered; no similar characters have been found for any natural clay described in the literature. Today, this particular green clay mineral has only been found nearly pure in restricted areas: Recent sediments off Dakar (Baie de Rufisque) and in the Casamance Estuary.

1

1

phyllite V Phyllite V is a term suggested to designate a green marine clay (a synsedimentary product) found in the verdine facies. Both the tetrahedral and octahedral sheets have cationic compositions different in phyllite V compared to berthierine (see Table 2, p. 180), chamosite or any known chlorite (see Table 2, p. 19; Fig. 7, p. 21; Table 1, p. 74; Fig. 13, p. 182; Fig. 20, p. 202). In phyllite V, the tetrahedral sheet contains less than 0.2 A1 cation for two sites; the octahedral sheet is equally trioctahedral and dioctahedral (about 2.35 to 2.5 cations for three possible sites). The predominant octahedral cations are Fe3+(0.75to 1.0 cation for three sites) and Mg (0.75 to 1.0 cation for three sites). Calculated cationic compositions are given in Table 9: p. 198). In contrast to that, Fe3+ and Mg are the two less abundant cations in berthierine and chamosite (see these words above). Although chemically homogeneous in all outcrops from which it was extracted the green clay called phyllite V has been shown to be mineralogically heterogeneous (see Table 11, p. 203). Some monomineralic green grains have been found to be made of more than 95% of a 1:l layer-type clay, which is now called 7 A phyllite V (see HRTEM pictures in Fig. 9, p. 173; X-ray film powder pattern interpretation in Table 8, p. 195). Other green grains have shown the presence of a 2: 1:1 layer-type clay which is now called I 4 A phyllite V (see HRTEM pictures in Fig. 10, p. 174) and has apparently the same chemical composition as the 7A phyllite V . A character common to the 7 A and 14 A phyllite V is their fragility on heating or with acid treatments (see Fig. 9, 11, and 13, p. 70, 72, and 76 respectively; Fig. 4 and 5, p. 165-166, and interpretation p. 196-197). recrystallization Recrystallization may occur as a synsedimentary or a diagenetic process (see neoformation above, and p. 42, 50, 218,243, 245, 312,316, 321-323). relict A relict sediment has been deposited and modified in situ; however, the initial environmental conditions have changed and are unrelated to the present conditions even though the sediment remains unburied. Pleistocene or older

416 relict sediments are frequent on present-day continental shelves (p. 272,328) or on the top of oceanic highs (p. 282). Relict glaucony is usually well preserved in Quaternary sediments (see p. 227, 229, 292, 316, 321) although altered at depth smaller than 10 m; but verdine is often oxidized (see p. 112, 124, 134, 137, 146, 150, 152, 153) partly because it is usually lying at depths smaller than 20 m.

sedimen togenesis The period of sedimentary evolution following deposition and preceding diagenesis is called here sedimentogenesis. Glaucony and verdine are two facies developed during sedimentogenesis. Sedimentogenesis implies dominant influence of cationic exchanges with the sea-water reservoir. semi-confined microenvironment The microenvironment where only restricted exchanges with the general environment are possible is said to be semi-confined. In this semi-confined microenvironment the local physico-chemical factors are preserved but slow exchanges with the exterior occur. For example, a microtest filled with organic matter and completely closed will allow crystallization of only a very minute quantity of mineral (e.g., pyrite) in a reducing environment: the microenvironment is confined. If this microtest is perforated with minute holes, exchanges with the exterior (oxidizing environment with abundant cations present) are possible; the microenvironment becomes less reducing, cations may penetrate in order to compensate the lack due to trapping by crystal growth: the microenvironment is semi-confined and verdine minerals (see p. 78-79; 102; 214) or glauconitic minerals (see p. 273-274, 323-324) will be able to grow in appreciable quantity. If the holes are too wide, then the exchange with the exterior will be intense, the microenvironment becomes unconfined (e.g., oxidizing) and the material within the void will possibly be oxidized (see the term relict, above), no more crystal growth is possible for green marine clays. substrate The material deposited at the sea-bottom which serves to shelter the formation of authigenic minerals is called here the substrate. Its size and origin varies (verdine, p. 212 and Fig. 6, p. 213; glaucony, Fig. 3, p. 302), its role is physical in the sense that being microporous, it determines the prerequisite semi-confined microenvironment (verdine, p. 212-214); but its chemical components are not used by the growing minerals (glaucony, p. 301-302). When evolution proceeds, the material of the substrate is usually (but not necessarily) destroyed and replaced by the authigenic minerals. The nature of the substrate determines the habit of the future green grain (glaucony, Table 2, p. 297). swelling-chlorite A swelling-chlorite is a clay mineral resembling chlorite but which expands on solvation: ethylene glycol treatment swells the 001 distance to 16 A or 17 A. Strong heating treatment (above 500°C) usually reduces slightly the 001 distan-

417 ce to less than 14 A. This behaviour indicates that the hydroxide sheet in the interlayer is not complete. The mineral is intermediate between a smectite and a chlorite. The variety of behaviour of swelling-chlorite is large however. Our phyllite C appears to have a behaviour intermediate between that of a swellingchlorite and that of a smectite (p. 169-170; 200).

synsedimentary The geochemical processes occurring during sedimentogenesis are said to be synsedimentary (e.g. green clays of the ironstones: p. 29, 39). At the seabottom, synsedimentary processes are predominantly influenced by sea-water. For this reason, the results of a synsedimentary process in the sea are generally homogeneous in time and space. At the sea-bottom, the contact between generally oxidizing sea-water and generally reducing interstitial water privileges the activity of iron which is involved in many synsedimentary geochemical reactions. Geochemical changes occurring during the synsedimentary period are sometimes gathered under the term of syndiagenesis. transformation The process by which a clay-mineral is progressively modified to another clay mineral by slow cation exchange in a permanent general structure is called transformation. This process implies the prerequisite presence of a phyllosilicate to form a new clay. This hypothesis was generally accepted in the sixties and seventies. For example, glauconitic minerals (2: 1-type clay minerals) were thought to be derived from other 2: 1 phyllosilicates like illite or biotite by A1 for Fe or Fe2+for Fe3+substitution or from smectite by potassium absorption in the interlayer and concomitant layer change re-equilibration (see p. 3 18). Similarly, the green marine clay from the verdine facies (7 8, phyllite V) was thought once to derive from kaolinite by this process. This hypothesis was mainly based on apparently progressive general changes in X-ray diffraction patterns during the evolution of the analysed material. However, the transformation hypothesis is not supported by the recent detailed studies, especially on glauconitic (p. 244; 257-260; 316; 318-319) and verdine (p. 212) minerals. verdine (Word root: vert = green in French).The term verdine (Odin, 1985a) has been created to designate a sedimentological facies different from both the ironstone and the glaucony facies. The facies may be defined 1) by its habit: green grains or green pigment in microtest chambers, faecal pellets, or various pores of mineral debris; this factor widely resembles the glaucony facies and differs from the ironstone facies; 2) by its mineralogical composition: various specific marine green clays, first distinguished during the present study, such as phyllite V or phyllite C (see mineralogy of the verdine facies, Table 11, p. 203); and 3) by its geological significance: the result of the interaction between sea-water and a deposited sediment near an abundant sourcg of iron at shallow depth (20 to 60 m) in tropical areas; this widely resembles the ironstone facies

418 and differs from the glaucony facies. Although still little known today, the facies is widespread all over the world (Fig. 15, p. 157).

verdinization The process by which a substrate deposited at the sea-bottom is progressively modified to a green object related to the verdine facies is called verdinization. Verdinization occurs for various substrates including carbonate bioclasts (New Caledonia), mineral debris (French Guiana) or faecal pellets (Senegal, Gulf of Guinea). During verdinization, the substrate is partly altered and marine green clays are formed. These clay minerals are dominantly ferric and magnesian and differ from those formed in the glaucony or oolitic ironstone facies. Verdinization occurs in a subtropical, shallow, marine, warm water environment fed with an abundant source of iron (Chapter B6, Fig. 3,4, 5, p. 209, 210, 211). verdissement This French substantive means: becoming green. This could possibly be translated by "greening" but this word is not commonly used in English whereas verdissernent is usual in French. In this volume verdissement specifically designates the change in colour of a substrate following the formation of a green pigment within or around it. The progressive change in colour has been observed to result sometimes from the genesis of various green clays: in shallow marine waters, the process results from crystal growth of phyllite V or phyllite C, in deeper marine waters, glauconitic minerals are concerned. The term has been required in order to designate the general colour change observed for sedimentary particles when no mineralogical study has shown the origin of the colour change which may result from glauconitization, verdinization, or chloritization amongst others.

419

GENERAL LIST OF REFERENCES Ahn, J.H. and Peacor, D.R., 1985. TEM study of diagenetic chlorite in Gulf Coast argillaceous sediments. Clays Clay Min., 33: 228-236. Ah, J.C. and Honnorez, J., 1984. Alteration of the upper oceanic crust, DSDP site 417: mineralogy and chemistry. Contr. Miner. Petrol., 87: 149-169. Allen, G.P., Laurier, D. and Thouvenin, J., 1979. Etude sedimentologique du delta de la Mahakam. Notes et Mkmoires, 15, Cle franq. Pktroles Publ., Paris, 156pp. Amouric, M., Mercuriot, G. and Baronnet, A., 1981. On computed and observed HRTEM images of perfect mica polytypes. Bull. Mineral., 104: 298-313. Andrews, A.J., 1980. Saponite and Celadonite in Layer 2 basalts DSDP Leg 37. Contrib. Mineral. Petrol., 73: 323-340. Andrews, A.J., Dollase, W.A. and Fleet, M.E., 1983. A Mossbauer study of saponite in layer 2 basalt, Deep Sea Drilling Project Leg 69. In: J.R. Cann et al. (Edit.). Init. Rep. DSDP, 69, U.S. Govt. Print. Off., Washington, pp. 585-588. Aubry, M.P. and Odin, G.S., 1973. Sur la nature minkralogique du verdissement des craies: formation dune phyllite apparentke aux glauconies en milieu semi-confink poreux. Bull. SOC.gkol. Normandie, 61: 11-22. Badaut, D., Besson, G., Decarreau, A. and Rautureau, R. 1985. Occurrence of a ferrous, trioctahedral smectite in Recent sediments of Atlantis I1 Deep Red Sea. Clay Min., 20: 389-404. Bailey, S.W., 1980. Summary of recommendations of AIPEA Nomenclature committee. Clays Clay Min., 28: 73-78. Bailey, R.J. and Atherthon, M.P., 1969. The petrology of a glauconite sandy chalk. Sedim. Petrol., 39: 1420-1431. Bailey, S.W. and Brown, B.E., 1962. Chlorite polytypism: Regular and semi-random one-layer structures. Am. Mineral., 47: 819-850. Baltzer, F., 1969. Les formations vkgktales associkes au delta de la Dumbka (Nouvelle Calkdonie) et leurs indications kcologiques, gkomorphologiques et skdimentologiques. Cah. O.R.S.T.O.M., 1: 59-84. Baltzer, F. and Trescases, J.J., 1971. Premikre estimation du bilan de l'altkration, de l'krosion et de la skdimentation sur pkridotites sous le climat tropical de la Nouvelle Calkdonie. C. R. Acad. Sci. Paris, 273: 2034-2037. Bancroft, G.M., 1973. Mossbauer spectroscopy; an introduction for inorganic chemists and geochemists. McGraw Hill Publ., Maiden Head, England, 252 pp. Barusseau, J.P., Giresse, P., Faure, H., Lezine, A.M. and Masse, J.P., (in press). Marine sedimentary environments on tropical and equatorial atlantic margins of Africa during the Late Quaternary. Continent. Shelf Res., Pergamon Publ. Bass, M.N., Moberly, R., Rhodes, J.M., Chi-Yu Shih and Church, S.E., 1973. Volcanic rocks cored in the Central Pacific, Leg 17. In: E.L. Winterer

420 et al. (Edit.). Init. Rep. DSDP, 17, U.S. Govt. Print. Off., Washington, pp. 429. Bates, R.L. and Jackson, J.A., 1980. Glossary of Geology. _ _ Am. Geol. Inst. Publ., Falls Church, 749 pp. Bayramgil, 0,Hiigi, T. and Novacki, W., 1952. Uber ein Seladonite vorkommen in Gebiete von Zonguldak (Turkei). Schweizer. Mineral. Petrol. Mitt., 3: 242-250. Bear, L.M., 1963. The mineral resources and Mining industry of Cyprus, Terre verte. Bull. Geol. Sum. Cyprus, 1: 166-168. Bell, D.L. and Goodell, H.G., 1967. A comparative study of glauconite and the associated clay fraction in modem marine sediments. Sedimentology, 9: 169-202. Bellon, H., Fabre, A., Sichler, B. and Bonhomme, M.G., 1986. Contribution to the numerical calibration of the Bajocian-Bathonian boundary: 4OW40Ar and palaeomagnetic data from Les Vignes basaltic complex (France). In: G.S. Odin (Edit.). Phanerozoic time scale calibration. Chem. Geol. (Isot. Geosc. Sect.), 59: 155-162. Bentor, Y. and Kastner, M., 1965. Notes on the mineralogy and origin of glauconite. J. Sedim. Petrol., 35: 155-166. Berg-Madsen, V., 1983. High alumina glaucony from the middle Cambrian of Oland and Bornholm, southem baltoscandia. J. Sedim. Petr., 53: 875-893. Berthier, M., 1820. Sur la nature du minerai de fer magnCtique de Chamoison (Valais). Ann. Mines, 5: 393-396. Beudant, F.S., 1832. Trait6 de MinCralogie, 2nded, pp. 128-129. Beutelspacher, E.H. and Van der Marel, H.W., 1968. Atlas of electron microscopy of clay minerals and their admixtures. Publ. Amsterdam. Elsevier, 333 pp. Bezrukov, P.L. and Senin, K.M., 1970. Sedimentation on the West African shelf. In: The Geology of the East Atlantic continental margin. Rep. Inst. Geol. Sci. 70/16, pp. 1-8. Bhattacharyya, D.P., 1983. Origin of berthierine in ironstones. Clays Clay Min., 31: 173-182. Bichelonne, J. and Angot, P., 1939. Le Bassin Ferrifhre Lorrain. BergerLevrault, Paris, 464 pp. Bignot, G., 1974. L'observation des tissus minCralis6s au microscope Clectronique ?i balayage; les coquilles de Lamellibranches actuels. Trav. Labor. Micropal., Univ. Paris, 3: 87-131. Bignot, G., 1976. L'observation des tissus organiques minCralis6s au microscope Clectronique ?i balayage; le squelette des Echinodermes vivants et fossiles. Trav. Labor. Micropal., Univ. Paris, 5: 207-245. Birch, G.F., Willis, J.P. and Rickard, R.S., 1976. An electron microprobe study of glauconites from the continental margin off the west coast of S. Africa. Mar. Geol., 22: 271-384. Bjerkli, J. and Ostmo-Saeter, J.S., 1973. Formation of glauconie in foraminiferal shells on the continental shelf off Norway. Mar. Geol., 14: 169-178. Blancaneaux, P., 1981. Essai sur le milieu nature1 de la Guyane franqaise.

421 Trav. Doc. ORSTOM, 137,126 pp. Blome, C.D. and Irwin, N.P., 1985. Equivalent radiolarian ages from ophiolitic terranes of Cyprus and Oman. Geology, 13: 401-404. Bohlke, J.K., Alt, J.C. and Muehlenbachs, K., 1984. Oxygen isotope-water relations in altered deep-sea basalts: low temperature mineralogical controls. Can. J. Earth Sci., 21: 67-77 Boichard, R., Burollet, P.F., Lambert, B. and Villain, J.M., 1985. La plate-forme carbonatCe du Pater Noster (Indonhie). MCm. 20, Total Cle franc. PCtroles Publ., 103 pp. Bongo Passi, G., 1984. Contribution, h 1'Ctude lithostratigraphique du delta profond du fleuve Congo. T h h e 3e cycle, Univ. Perpignan, offset, 215 pp. Bornhold, B.D. and Giresse, P., 1985. Glauconitic sediments on the continental shelf off Vancouver Island (Canada). J. Sedim. Petrol., 55: 653-664. Bouysse, Ph., Kudrass, H.R. and Le Lann, F., 1977. Reconnaissance stdimentologique du plateau continental de la Guyane francaise (mission Guyamer, 1975). Bull. BRGM, 4: 141-179. Bradshaw, M.J., James, S.L. and Turner, P., 1980. Discussion of Kimberley, M.M.: Origin of oolitic ironstone formation. J. Sedim. Petrol. 50: 295-302. Brindley, G.W., 1961a. Chlorite minerals. In: G. Brown (Edit.). The X-ray identification and crystal structures of clay minerals. Mineralogical Society Publ., London, pp. 242-296. Brindley, G.W., 1961b. Kaolin, serpentine and hindered minerals. In: G. Brown (Edit.). The X-ray identification and crystal structures of clay minerals. Mineralogical Society Publ., London, pp. 5 1 - 13 1. Brindley, G.W., 1982. Chemical composition of berthierines, a review. Clays Clay Min., 30: 153-155. Brindley, G.W. and Ali, S.Z., 1950. An X-ray study of thermal transformations in some magnesian chlorite minerals. Acta. Cryst., 3 : 25-30. Brindley, G.W., Bailey, S.W., Faust, G.T., Forman, S.A. and Rich, C.I., 1968. Report of the Nomenclature committee (66-67) of the Clay Mineral Society. Clays Clay Min., 16: 322-324. Brindley, G.W. and Youell, R.F., 1953. Ferrous chamosite and ferric chamosite. Mineral. Mag., 30: 57-70. Bubenicek, L., 1961. Recherches sur la constitution et la rkpartition du minerai de fer dans 1'AalCnien de Lorraine. Sciences Terre, Nancy, 8: 5-204. Bubenicek, L., 1964. Etude sCdimentologique du minerai de fer oolithique de Lorraine. In: G.C. Amstutz (Edit.). Sedimentology and ore genesis. Developments in sedimentology, Elsevier Publ., Amsterdam, 2. Buckley, A., Bevan, J.C., Brown, K.M. and Johnson, L.R., 1978. Glauconite and Celadonite: two separate mineral species. Mineral Mag., 42: 373382. Burnett, W.C., 1980. Apatite glauconite association off Peru and Chile; palaeooceanographic implications. J. geol. Soc.London, 137: 757-764. Burst, J.F., 1958a. Mineral heterogeneity in glauconite pellets. Am. Mineral, 43: 481-497. Burst, J.F., 1958b. "Glauconite" pellets: their mineral nature and applications

422 to stratigraphic interpretations. Bull. Am. Ass. Petrol. Geol., 42: 310-337. Caillkre, S. and Giresse, P., 1966. Etude mineralogique de diverses "glauconies" actuelles; nouvelle contribution h la gen&se des minerais de fer skdimentaires. C. R. Acad. Sci., Paris, 263: 1803-1807. Caillkre, S. and HCnin, S., 1960. Relation entre la constitution cristallochimique des phyllites et leur tempcramre de deshydratation; application au cas des chlorites. Bull. Soc. franG. CCram, 48: 63-67. Caillkre, S., HCnin, S. and Rautureau, M., 1982. MinCralogie des argiles. Masson Publ., Paris, 2 vol., 373 pp. Caillkre, S. and Kraut, F., 1954. Les gisements de fer du bassin lorrain. MCm. Museum. nation. fist. natur., Paris, 4, pp. 1-176. Caillkre, S. and Lamboy, M., 1970a. Presence de berthiirine sur le plateau continental au nord de la pCninsule ibCrique. C. R. somm. Soc. gCol. France, 6: 218-220. Caillkre, S. and Lamboy, M. 1970b. Etude minkralogique de la glauconite du plateau continental au Nord-Ouest de 1'Espagne. C. R. Acad. Sci., Paris, 270: 2057-2960. Calvet, F. and Julih, R., 1983. Pisoids in the caliche profiles of Tarragona (NE Spain). In: P.M. Peryt (Edit.). Coated grains. Springer Verlag, Berlin, pp. 456-473. Caratini, C., Bellet, J. and Tissot, C., 1975. Etude microscopique de la matikre organique: palynologie et palynofaciks. Orgon 11, CNRS Publ. Paris, pp. 1 57 -203. Carozzi, A.V., 1960. Microscopic sedimentary petrography. John Wiley Publ., Chichester, 485 pp. Carroll, D., 1966. Clay Minerals: a guide to their X-ray identification. Geol. Soc. Am., Spec. Paper 126,60 pp. Caspari, W.A., 1910. Contributions to the chemistry of submarine glauconite. Roc. Roy. Soc. Edinburgh, 30: 364-373. Cayeux, L., 1897. Contribution h 1'Ctude micrographique des terrains skdimentaires, chap.4, Le Bigot Publ., Lille, pp. 163-184. Cayeux, L., 1909. Evolution min6ralogique des minerais de fer oolithiques primaires de France. C. R. Acad. Sci., Paris, 149: 1388-1390. Cayeux, L., 1916. Introduction h 1'Ctude pktrographique des roches sCdimentaires. (Glauconie, pp. 241-252). Imprimerie Nationale, Paris, 524 pp. Cayeux, L., 1932. Interpretation des phosphates de chaux draguCs sur 1'Agulhas Bank, au Sud du Cap de Bonne EspCrance. C. R. Acad. Sci., Paris, 194: 926-929. Chagnaud, M., 1984. Etude des sCdiments carottis du plateau continental de la Guyane franGaise. Recherches sur la nature et l'origine des phyllites authigknes. Apports h la connaissance du Quaternaire terminal, MCm. DEA, Bordeaux, offset, 115 pp. Champetier, Y., Darlet, P., Hamdadou, E. and Hamdadou, M., 1987. Les ClCments biogknes, supports de minCralisation diagCnCtiques. 4th Europ. Union Geosci. Congr., Abstracts, Terra Cognita, 7: 207. Champetier, Y., Hamdadou, E. and Hamdadou, M., 1987. Examples of bioge-

423 nic support of mineralization in two oolitic iron ores Lorraine (France) and Gara Djebilet (Algeria). Sedim. Petrol., 51: 249-255. Champetier, Y. and Joussemet, R., 1979. DCcouverte de nubiculaires et doncolithes en tant qu'ClCments phosphates. C. R. Acad. Sci. Paris, 288: 673-675. Channel, J.E.T., 1982. Palaeomagnetic stratigraphy as a correlation technique. In: G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 81-103. Chauvel, J.J. and Guerrak, S., 1986. Oolitization of iron formation: examples of North African deposits. 7* Region. Meet. Sedimentology, Krakow, pp. 44-45. Chauvel, J.J. and Massa, C., 1981. PalCozoYque de Lybie occidentale; signification des niveaux ferrugineux oolithiques. Notes MCm. CompIe fr. PCtroles, Paris, 16 : 25-66. Chevillon, C., 1986. Les sCdiments de la corne SE du lagon nCo-calkdonien: recueil de donnCes. Rapp. Scient. Techn. ORSTOM NoumCa, 40,13 pp. Cimbalnikova, A., 1970. Index of refraction and density of glauconites. Cas. Mineral. Geol. Ceskosl., 15: 335-345. Cimbalnikova, A., 197 la. Chemical variability and structural heterogeneity of glauconites. Am. Mineral, 56: 1385-1392. Cimbalnikova, A., 197 1 b. Cation exchange capacity of glauconites. Cas. Mineral. Geol. Ceskosl., 16: 15-21. Clayton, R.N., O'Neil, J.R. and Mayeda, T.K., 1972. Oxygen isotope exchange between quartz and water. J. Geophys. Res., 77: 3057-3067. Coey, J.M.D., 1980. Clay minerals and their transformations studied with nuclear techniques: the contribution of Mossbauer spectroscopy. Atomic Energy Rev., 18: 73-124. Collet, L.W, 1908. Les dCp8ts marins. (Chap. 2, La glauconie), Doin Publ., Paris, 325 pp. Collet, L.W. and Lee, G.W., 1906. Recherches sur la glauconie. Roc. Roy. Soc. Edinburgh, 26: 238-278. Conard, M., Kreuzer, H. and Odin, G.S., 1982. Potassium argon dating of tectonized glauconies. In: G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley Publ., Chichester, pp. 321-332. Correns, G.W., 1939. Pelagic sediments of the North Atlantic ocean. In: P.D. Trask (Edit.). Recent Marine Sediments, pp. 373-395. Coudray, J., 1976. Recherches sur le NCoghne et le Quaternaire marin de Nouvelle Calddonie. In: Exp6dition franpise sur les RCcifs coralliens de la Nouvelle CalCdonie, Fondat. Singer-Polignac Publ., 8: 5-275. Courtois, C. and Desprairies, A.,1978. Les terres rues et quelques ClCments de transition dans les minCraux argileux issus de deux processus daltbration de roches basiques. C. R. somm. Soc. gCol. Fr., 5: 242-245. Craig, H. 1957. Isotopic standards for carbon and oxygen and correction factors for mass-spectrometric analysis of carbon dioxide. Geochim. Cosmochim. Acta, 12: 133-149. Cullen, D.J., 1967. The age of glauconite from the Chatham Rise, East of New

424 Zealand. N. Z. mar. Fresh-wat. Res., 1: 399-406. Curtis, C.D., 1985. Clay mineral precipitation and transformation during burial diagenesis. Phil. Trans. Roy. Soc. Lond., 315: 91-105. Curry, D. and Odin, G.S. 1982. Dating of the Palaeogene. In: G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley Publ., Chichester. pp. 607-630. Dahanayake, K. and Krumbein, W.K., 1986. Microbial structures in oolitic iron formations. Mineral. Depos., 21: 85-94. Dangeard, L., 1928. Observations de gCologie sous-marine et docCanographie relatives B la Manche. Ann. Inst. OcCanographique, 6, 1: 199-211. Daynyak, L.G., Drits, V.A. Kudryavtsev, D.I., Simannovich, I.M. and Slonimskaya, M.V., 1983. Crystallochemical specificity of trioctahedral smectite containing Fe3+, the secondary alteration products of oceanic and continental basalts. Doklady SSSR, Earth Sciences, 259: 176-179 (English transl., orig. 1981: 1458-1462). Debenay, J.P., 1985. Recherches sur la skdimentation actuelle et les thanatocoenoses de grands Foraminiferes dans le lagon S.O.et sur la marge insulaire sud de Nouvelle CalCdonie. Th2se, Univ. Aix-Marseille, offset, 250 pp. Debenay, J.P. and KonatC, S . , (in press). Les Foraminifkres actuels des iles de Los (Guide). Premier inventaire et comparaison avec les microfaunes des rCgions voisines. De Brignoli de Brunnhoff, G., 1820. Sur la chlorite ou tern verte de VCrone. J. Phys. Chim. Hist. natur. Arts, 90: 355-360. and 423-442. Delesse, M., 1848. Composition chimique de la Terre verte de VCrone. Ann. Mines, 14 : 74-78. Delaloye, M.F., 1966. Contribution B 1'Ctude des silicates de fer sCdimentaires. Le gisement de Chamoson (Valais). MatCriaux GCologie Suisse, GCotechnique, 9, 71 pp. Delmont, P., 1985. Smectites et produits daltkration des basaltes tertiaires des iles Faeroe (Atlantique Nord Est). Genhe, Cvolution et contribution & la skdimentation ocCanique. Thkse, UniversitC de Bordeaux I, offset, 490 pp. Denis-Clocchiatti, M. 1982. Sedimentation carbonatke et palCoenvironnement dans l'OcCan Indien au CCnozo'ique. MCm. Soc. gCol. Fr., 60,92 pp. De Paolo D.J. and Ingram B.L., 1985. High resolution stratigraphy with strontium isotopes. Sciences, 277: 939-941. Desprairies, A., Bonnot, C., Jehanno, C., Vernhet, S . and Joron, J. L., 1984. Mineralogy and geochemistry of alteration products in Leg 81 basalts. In: D.G. Roberts et al. (Edit.). Init. Rep. DSDP, 81, U.S. Govt Print. Off, Wahington, pp. 733-742. Desprairies, A., Decarreau, A., Bonnot, C., Trauth, D., Vernhet, S . , Joron, J.L., Bellon, H., Labeyrie, L., Escande, M. and Jehanno, C., 1983. Les paragen2ses 'smectite-cCladonite' dans les basaltes du Leg 81, site 553: donnCes minCralogiques, gdochimiques, radiomktriques et isotopiques. C . R. RCun. A.T.P.-G.G.O. Soc. gtol. Fr., Brest, DCcembre 1983, offset. Desprairies, A. and Lapierre, H., 1973. Les argiles likes au volcanisme du Massif de Troodos (Chypre) et leur remaniement dans sa couverture. Rev.

425 GCog. phys. GCol. dynam., 15: 499-510. Desprairies, A., Tremblay, P. and Laloy, C., 1988. Secondary mineral assemblages in a volcanic sequence drilled by ODP Leg 104 in the Norwegian Sea. In: 0. Eldholm, J. Thiede, E. Taylor et al. (Edit.). Proceed. Init. Rep., Part B, O.D.P. 104, (in press). DCverin, L., 1944. Origine des oolites chamositiques dans les minerais de fer oolithiques de la Suisse. Migration de la chamosite. Bull. SOC.Miner. Petrogr., 18: 672-673. DCverin, L., 1945. Etude pktrographique des minerais de fer oolithiques du Dogger des Alpes Suisses. MatCriaux GCologiques Suisse, Giotechnique, 2, 80 pp. Donnelly, T., Francheteau, J. Bryan, W. et al., 1979. Introduction. Init. Rep. D.S.D.P., 51-52-53, U.S. Govt. Print. Off., Washington, pp. 1-10. Dreesen, R.J.M., 1987. Oolitic ironstones as event-stratigraphical marker beds within the Upper Devonian of the Belgian Ardenne shelf. Phanerozoic ironstones, Symposium, Abstracts and Progr., pp. 18. Dugas, F., 1974. La skdimentation en baie de St Vincent (C6te ouest de la Nouvelle CalCdonie). Cah.ORSTOM, 6: 41-62. Dugas, F. and Debenay, J.P., 1978. Carte sCdimentologique et carte annexe du lagon de Nouvelle CalCdonie 1/50.000e, feuille Mont Dore. ORSTOM Publ., Paris, Notice explicative n076,20 pp. Dugas, F. and Debenay, J.P., 1980. Carte sCdimentologique et carte annexe du lagon de Nouvelle CalCdonie, 1/50.000e, feuille Prony. ORSTOM Publ., Paris, Notice explicative n091, 35 pp. Dunoyer de Segonzac, G., 1969. Les minCraux argileux dans la diagenkse. Passage au mCtamorphisme. MCm. Serv. Carte gCol. Als. Lorr., Strasbourg, 29, 320 pp. Durand, A., 1974. Stratigraphie des terrains d'bge palCogkne supCrieur et nCogkne du plateau continental basque et asturien d'aprks 1'Ctude des Foraminiferes planctoniques. Thkse 3kme cycle, Rennes, offset, 118 pp. Dyar, M.D. and Bums, R.G., 1986. Mossbauer spectral study of fenuginous one layer trioctahedral micas. Am. Miner., 71: 955-965. Ehlmann, A., Hulings, N. and Glover, E., 1963. Stages of glauconite formation in modern foraminifera1 sediments. J. Sedim. Petrol., 33: 87-96. Eisma, D. and Van der Marel, H.W., 1971. Marine muds along the Guyana coast and their origin from the Amazon basin. Contr. Mineral. Petrol., 31: 321-334. Eldholm, O., Thiede, J., Taylor, E. et al., 1986. RCsultats prdliminaires de la campagne 104 Joides Resolution (O.D.P.) sur le Plateau de Voring (Est de la Mer de Norvkge). C. R. Acad. Sci. Paris, 303: 1467-1472. Emelyanov, E.M., 1970. The composition of the glauconitic and hydrogoethite-chamosite-glauconite sediments of the West-African shelf. In: The Geology of the East Atlantic Continental Margin. Rep. Inst. Geol. Sci. 70/16, 97-103. Erlank, A. J. and Reid, D. L., 1974. Geochemistry mineralogy and petrology of basalts Leg 25. In: R. Schlich et al. (Edit.). Init. Rep. DSDP, 25, U.S.

426 Govt. Print. Off., Washington, pp. 543-552. Escande, M., 1983. EchangeabilitC et fractionnement isotopique de l'oxygkne des smectites magnksiennes de synthkse. Etablissement dun gCothermomktre. Thkse, UniversitC de Paris Sud, 150 pp. Farmer, V.C., 1974. The layer silicates. In: V.C. Farmer (Edit.). The infrared spectra of minerals. Miner. SOC.,monogr. n04, London, pp. 331-365. Faure, G., 1982. The marine strontium geochronometer. In: G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley Publ., Chichester, pp. 73-79. Faure, H. and Elouard, P., 1967. SchCma des variations du niveau de l'OcCan Atlantique sur la c6te de l'Ouest de 1'Afrique depuis 40.000 ans. C. R. Acad. Sci. Paris, 265: 784-786. Figukres, G., Martin, J.M. and Meybeck, M., 1978. Iron behaviour in the ZaiYe estuary. Netherl. J. Sea Res., 12: 329-337. Fleet, A.J., Buckley, H. A. and Johnson, L.R., 1980. The rare earth element geochemistry of glauconites and celadonites. J. geol. SOC.London, 137: 683-688. Foster, M.D., 1962. Interpretation of the composition and a classification of the chlorites. U.S. Geol. Surv. Prof. Paper, 414-A, 33 pp. Foster, M.D., 1964. Water content of micas and chlorites. U.S. Geol. Surv. Prof. Paper, 474-F, 15pp. Foster, M.D., 1969. Studies of celadonite and glauconite. U.S. Geol. Surv. Prof. Paper, 614-F, 17 pp. Freeman, T., 1963. Quiet water ooliths from Laguna Madre, Texas. J. sedim. Petrol., 32, 475-483. Frohlich, F., 1980. NCoformation de silicates femfkres amorphes dans la skdimentation pClagique rCcente. Bull. Mineral., 103: 596-599. Frohlich, F., 1981. Les silicates dans l'environnement pClagique de l'ocCan Indien au CCnozoTque. MCm. MusCum nation. Hist. natur., 46, MusCum Paris Publ., 206 pp. Frohlich, F., 1982. Evolution minkralogique dans les dCp6ts azo'iques rouges de l'Oc6an Indien. Bull. Soc. gCol. Fr., 24: 563-571. Frohlich, F., 1983. Rapp. Mission: MD35/D.R.A.K.A.R., bord du Marion Dufresne, Mars 1983. Publ. Miss. Rech. T.A.A.F. 83-02. Paris, 93pp. Frohlich, F., 1986. PrCsence de dCp6ts phosphatCs sur le plateau de Kerguelen-Heard (OcCan Indien). C. R. Acad. Sci., 303: 167-170. Frohlich, F., Caulet, J.P. et al. 1983. Mise en Cvidence d'une sCrie pClagique du PalCogkne et du CrCtacC supCrieur sur le plateau de Kerguelen-Heard: resultats prdliminaires de la campagne MD35DRAKAR (mars 1983) du N.O. Marion Dufresne. C. R. Acad. Sci., Paris, 297: 153-156. Gale, M.J., 1980. Mineralogy and petrology of very low metamorphic grade Archean banded iron formations, Weld Range, Western Australia. Am. Mineral., 68: 8-25. Galliher, E.W., 1935. Geology of glauconite. Bull. Am. Assoc. Petrol. Geol., 19: 1569-1601. Gautier, F. and Odin, G.S., 1985. Volcanisme jurassique du Sud de 1'Aragon

427 (Espagne). Bull. Liais. Inf., I.G.C.P. Project 196, offset Paris, 5: 34-38. Gehring, A.V., 1985. A microchemical study of iron ooids. Eclogae Geol. Helv., 78: 451-457. Gehring, A.V., 1986. Mikroorganismen in kondensierten Schichten der DoggerMalm-Wende in Jurader Nordostschweiz. Eclogae Geol. Helv., 79: 12-18. Gibbs, R.J., 1967. The geochemistry of the Amazon River system. Part. I. The factors that control the salinity and the composition and concentration of the suspended solids. Geol. SOC.Amer. Bull., 78: 1203-1232. Gibbs, R.J., 1973. The bottom sediments of the Amazon shelf and tropical Atlantic Ocean. Mar. Geol., 14 : M39-M45. Gibbs, R.J., 1976. Amazon River sediment transport in the Atlantic Ocean. Geology, 4: 45-48. Giresse, P., 1965a. Observation sur la presence de "glauconie" actuelle dans les sCdiments ferrugineux peu profonds du bassin gabonais. C. R. Acad. Sci., Paris, 260: 5597-5600. Giresse, P., 1965b. Oolithes fermgineuses en voie de formation au large du Cap Lopez (Gabon). C. R. Acad. Sci. Paris, 260: 2550-2552. Giresse, P., 1969. Etude des differents grains fermgineux authigknes des sCdiments sous-marins au large du delta de l'OgoouC (Gabon). Sciences Terre, Nancy, 14: 27-62. Giresse, P., 1976a. Etude perspective des glauconies sous-marines du Golfe de GuinCe en tant que fertilisant potentiel des sols tropicaux. Ann. Univ. Brazzaville, 13: 29-39. Giresse, P., 1976b. MatCriaux pour la connaissance de la glauconitisation sur la plate-forme du Congo. Ann. Univ. Brazzaville, 13: 4 1-62. Giresse, P. and Kouyoumontzakis, G., 1973. Cartographie sCdimentologique des plateaux continentaux du Sud du Gabon, du Congo, du Cabinda et du Zaire. Cah. ORSTOM, GCol., 2: 235-257. Giresse, P., Lamboy, M. and Odin, G.S., 1980. Evolution gComCmque des supports de glauconitisation; application B la reconstitution du palbnvironnement. Oceanol. Acta, 3: 251-260. Giresse, P. and Odin, G.S., 1973. Nature minkralogique et origine des glauconies du plateau continental du Gabon et du Congo. Sedimentology, 20: 457-488. Giret, A., 1983. Le plutonisme ocCanique intraplaque. Exemple de l'archipel Kerguelen, Terres Australes et Antarctiques Franpises. Thkse, Univ. Paris, n"83-33. 135 pp. Goslin, J., 1981. Etude geophysique des reliefs asismiques de 1'ocCan Indien occidental. Thkse, Univ. Louis Pasteur, Strasbourg. 267 pp. Gosselet, J., 1901. Observations gCologiques faites dans les exploitations de phosphate de chaux. Ann. SOC.gCol. Nord, 49: 208. Gruner, J.W., 1922. The origin of sedimentary iron formation. Econ. Geol., 51: 565-695. Guerrak, S., 1987. Time and space distribution of Palaeozoic oolitic ironstones in Tindouf Basin (Algerian Sahara). In: Phanerozoic ironstones,

428 Symposium, Abstract and Progr., pp. 26-27. Guerrak, S. and Chauvel, J.J., 1985. Les mineralisations ferrifkres du Sahara algCrien: le gisement de fer oolithique de Meh'ri Adbdelaziz. Mineral. Depos., 20: 249-259. Gygi, R.A., 1969. Zur Stratigraphie der Oxford-Stufe (oberes Jura-System) der Nordschweiz und des suddeutschen Grentzgebeites. Beitr. geol. Karte Schweiz (N.F.), 136. Gygi, R.A., 1981. Oolitic iron formations: marine or not marine. Eclogae Geol. Helv., 74: 233-254. Gygi, R.A., 1986. Eustatic sea-level changes of the Oxfordian and their effect documented in sediment and fossil assemblages of an epicontinental sea. Eclogae Geol. Helv., 79: 455-491. Gygi, R.A. and Mac Dowell, F.W., 1970. Potassium argon ages of glauconites from a biochronologically dated upper Jurassic sequence of northern Switzerland. Eclog. Geol. Helv., 63: 111-118. 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. Hallimond, A.F., 1925. Iron ores: bedded ores of England and Wales; Petrography and chemistry. Geol. Surv. Memoirs: Spec. Rep. Mineral Resources Great Britain. H. M. S. O., London. Harder, H., 1964. On the diagenetic origin of berthierine (chamositic) iron ores. In: 22* Intern. Geol. Congr. New-Dehli, 5: 193-198. Harder, H., 1978. Synthesis of iron layer silicate minerals under natural conditions. Clays Clay Min., 26: 65-72. Harder, H., 1987. Mineral formation in hydroxide gels for understanding sedimentary iron ore genesis. In: Phanerozoic ironstones, Symposium, Abstract and Progr., pp. 28. Hardjosoesastro, R., 1971. Note on chamosite in sediments of the Surinam shelf. Geol. Mijnb., 50: 29-33. Harrison, R.K., Knox, R.W.O'B. and Morton, A.C., 1979. Petrography and mineralogy of volcanogenic sediments from DSDP Leg 48, Southwest Rockall Plateau, sites 403 & 404. In: L. Montadert, et al. (Edit.). Init. Rep. DSDP 48. U.S. Govt. Print. Off. Washington, pp. 771-785. Hart, S.R. and Staudigel, H., 1979. Ocean crust-sea water interaction sites 417 and 418. In: T. Donelly et al. (Edit.). Init. Rep. DSDP, 51-52-53, U.S. Govt. Print. Off., Washington, pp. 1169-1176. Hart, S.R. and Staudigel, H., 1986. Ocean crust vein mineral deposition : Rb/Sr ages, U-Th-Pb geochemistry and duration of circulation at DSDP sites 261,462 and 516. Geochim. Cosmochim. Acta, 50: 2751-2761. Hayes, J.B., 1970. Polytypism of chlorite in sedimentary rocks. Clays Clay Min., 18: 285-306. Heddle, M. F., 1879. The minerals from Scotland: Celadonite. Trans. Roy. SOC.Edinburgh, 29: 101-104. Hein, J.R., Allwardt, A.O. and Griggs, G.B., 1974. The occurrence of glauconite in Monterey Bay, California. Diversity, origins, and sedimentary

429 environmental significance. J. Sedim. Petrol., 44: 562-57 1. Heller-Kallai, L. and Rozenson, I., 1981. The use of Mossbauer spectroscopy of iron in clay mineralogy. Phys. Chem. Minerals, 7: 223-238. Hendricks, S.B. and Ross, C.S., 1941. Chemical composition and genesis of glauconite and celadonite. Am. Miner., 26: 683-708. Hogan, L. and Dymond, J., 1976. K-Ar and 39Arp0Ar dating of site 319 and 321 basalts. In: R.S. Yeats et al. (Edit.). Init. Rep. DSDP, 34, U.S. Govt. Print. Off., Washington, pp. 439-442. Hogg, C.S. and Meads, R.E., 1970. The Mossbauer spectra of several micas and related minerals. Miner. Mag., 37: 606-614. Houbolt, J.J.H.C., 1957. Surface sediments of the Persian gulf near the Qatar Peninsula. Mouton & CO Publ., La Haye, 113 pp. Houtz, R.E., Hayes, D.E. and Markl, R.G., 1977. Kerguelen Plateau bathymetry, sediment distribution and crustal structure. Mar. Geol., 25: 95-130. Hower, J., 1961. Some factors concerning the nature and the origin of glauconite. Am. Miner., 46: 313-334. Hower, R.J. and Mowatt, T.C., 1966. The mineralogy of illite and mixedlayer illite montmorillonite. Am. Miner., 5 1: 825-854. Humphris, S.E., Thompson, R.N. and Marriner, G.F., 1979. The mineralogy and geochemistry of basalt weathering, holes 417A and 418A. In: T. Donelly et al. (Edit.). Init. Rep. DSDP., 51-52-53, U.S. Govt. Print. Off., Washington, pp. 1201-1217. Iijuma, A. and Matsumoto, R., 1982. Berthierine and chamosite in coal measures of Japan. Clays Clay Min., 30: 264-274. Ireland, B.J., Curtis, C.D. and Whiteman, J.A., 1983. Compositional variation within some glauconites and illites and implications for their stability and origins. Sedimentology, 30: 769-786. James, H.L., 1966. Chemistry of the iron-rich sedimentary rocks. U.S. Geol. Surv. Prof. Pap., 440-W, 61 pp. Jarrige, F., Radok, R., Krause, G. and Rual, P., 1975. Courants dans le lagon de NoumCa (Nouvelle CalCdonie). Rapp. ORSTOM NoumCa, 6pp. Jeantet, D., 1982. Processus sedimentaires et Cvolution du Plateau Guyanais (Guyane franqaise) au cours du Quaternaire terminal. Thhse 32me cycle, Bordeaux, 336 pp. Joseph, P. and Beaudouin, B., 1983. MicrosCquences intra-oolithiques dans le mineral de fer ordovicien Normand (Llamvirn). Nouvelle hypothhse de genkse des oolithes ferrugineuses. C.R. Acad. Sci., Paris, 296: 1533-1537. Jouanneau, J.M. and Pujos, M., 1986. Variations annuelles des concentrations en suspension et estimations des dCbits solides des fleuves Maroni et Mahury (Guyane franipise). 1 l 2 Conf. GCol. CaraYbe, La Barbade, Abstr. Juignet, P., 1974. La transgression crCtacde sur la bordure orientale du Massif Armoricain (Aptien, Albien, CCnomanien de Normandie et du Maine; le stratotype du CCnomanien). Thhse, Univ. Caen, offset, 806 pp. Karig, D.E. and Ingle, J.C. Jr. et al. 1975. Init. Repts. D.S.D.P., U.S. Govt. Printing Office, Washington, 31,927 pp. Kastner, M. and Gieskes, J.M., 1976. Interstitial water profiles and sites of

430 diagenetic reactions, Leg 35, DSDP, Bellingshausen abyssal Plain. Earth Planet. Sci. Lett., 33: 11-20. Kearsley, A., 1987. Mineralogy microfabric and the evolution of iron-rich ooids. In: Phanerozoic ironstones, Symposium, Abstract , pp. 38. Keller, G.M. and Richards, A.F., 1967. Sediments of the Malacca Strait, SE Asia. J. Sedim. Petrol., 37: 102-127. Kempe, D.R., 1974. Petrology of basalts Leg 26. In: T.A. Davis et al. (Edit.) Init. Rep. DSDP, 26, U.S. Govt. Print. Off., Washington, pp. 465-504. Kennedy, W.J. and Odin, G.S., 1982. The Jurassic and Cretaceous time scale in 1981. In: G.S. Odin, (Edit.). Numerical Dating in Stratigraphy. John Wiley Publ., Chichester, pp. 557-592. Keppens, E., Obradovich, J.D., Odin, G.S., O'Neil, J.R. and Pasteels, P., (in press). Radiogenic and stable isotopes in glauconies: early diagenetic evolution. Chem. Geol. (Isot. Geosc. Sect.). Keppens, E. and O'Neil, J.R., 1984. Oxygen isotope variations in glauconies. Terra Cognita, Sp. Iss., Abstracts ECOG VIII, Braunlage, pp. 42. Keppens, E. and Pasteels, P., 1982. A comparison of rubidium-strontium and potassium-argon ages on glauconies. In: G.S. Odin (Edit.). Numerical dating in Stratigraphy. John Wiley Publ., Chichester, pp. 225-239. Kerforne, F., 1908. Sur l'iige des minerais de fer de la foret de Lorges (C8tes du Nord). C. R. Acad. Sci. Paris, 147: 1007-1008. Kimberley, M.M., 1974. Origin of iron ore by diagenetic replacement of calcareous oolite. Nature, 250: 319-320. Kimberley, M.M., 1978. Paleoenvironmental classification of iron formations. Econ. Geol., 78: 215-229. Kimberley, M.M., 1979. Origin of oolitic iron formations. Sedim. Petrol., 49: 111-132. Knauth, L.P. and Epstein, S., 1976. Hydrogen and oxygen isotope ratios in nodular and bedded cherts. Geochim. Cosmochim. Acta, 40: 1095-1108. Knox, R.W. O'B., 1970. Chamosite ooliths from the Winter Gill Ironstone (Jurassic) of Yorkshire, England. J. sedim. Petrol., 40: 1216-1225. Knox, R.W.O'B and Fletcher, T.P., 1987. Genesis of a mixed silicate-oxide oolitic ironstone: The Frodingham ironstone (Sinemurian) of Eastern England. In: Phanerozoic ironstones, Symposium, Abstract, pp. 40. Kossovskaya, A.G. and Drits, V.R., 1970. Micaceous minerals in sedimentary rocks. Sedimentology, 15 : 83-101. Kotlichi, A., Sztzyrda, J. and Wiewiora, A., 1981. Mossbauer study of glauconites from Poland. Clay Min. 16: 221-230. Lacroix, A., 1895. MinCralogie de la France, I. Baudry & Cie Publ., Paris, Chamosite, pp. 397-403. Lacroix, A., 1916. Sur le minCral colorant le plasma de Madagascar et sur la ckladonite. Ann. Mines: 90-95. Lagaaij, R., 1973. Shallow water Bryozoa from deep sea sands of the Principe Channel Gulf of Guinea. In: G.P. Larwood (Edit.). Living and Fossil Bryozoa. London Acad. Press, London, 250 pp. Lamboy, M., 1967. RCpartition de la "glauconie" sur le plateau continental de

43 1 la Galice et des Asturies (Espagne). C. R. Acad. Sci., Paris, 265: 855-857. Lamboy, M., 1968. Sur un processus de formation de la glauconie en grains i partir des dCbris coquilliers. Rdle des organismes perforants. C. R. Acad. Sci., Paris, 266: 1937-1940. Lamboy, M., 1975. La glauconie du plateau continental au Nord-Ouest de 1'Espagne dCrive danciens dCbris coquilliers. C. R. Acad. Sci., Paris, 280: 157- 160. Lamboy, M., 1976. GCologie marine du plateau continental au N.O. de l'Espagne, Th&se,Univ. Rouen, offset, 283 pp. Lamboy, M. and Odin, G.S., 1974. RCvision et aspects nouveaux concernant les glauconies rCcentes au Nord et i 1'Ouest de la pCninsule ibCrique. C. R. Acad. Sci., Paris, 279: 2007-2010. Lamboy, M. and Odin, G.S., 1975. Nouveaux aspects concernant les glauconies du plateau continental Nord-Ouest espagnol. Rev. GCogr. phys. GCol. dynam. 17: 99-120. Lanphere, M. A. and Dalrymple, G. B., 1976. Potassium-argon age of a basalt from hole 319a, Leg 34. In: R.S. Yeats et al. (Edit.). Init. Rep. DSDP, 34, U.S. Govt. Print. Off., Washington, pp. 443-444. Launay, J., 1972. La skdimentation en baie de DumbCa (C8te ouest, Nouvelle CalCdonie). Cah. O.R.S.T.O.M., 4: 25-5 1. Laverne, C. and Vivier, G., 1983. Petrographical and chemical study of basement basalts from the Galapagos spreading center, Leg 70. In: J. Honnorez et al. (Edit.), Init. Rep. DSDP., 70, U.S. Govt. Print. Off., Washington, pp. 375-389. Lawrence, J.R., 1979. Temperatures of formation of calcite veins in the basalts from deep sea drilling project hole 417A-417D. In: T. Donnelly et al. (Edit.). Init. Rep. D.S.D.P., 51-53, U.S. Govt. Print. Off., Washington, pp. 1183-1184. Lawrence, J.R., Gieskes, J.M. and Broecker, W.S., 1975. Oxygen isotope and cation composition of D.S.D.P. pore waters and the alteration of Layer I1 basalts. Earth Planet. Sci. Lett.,27: 1-10. Leneuf, N., 1962. Les pseudo-oolites ferrugineuses des plages de C8te d Ivoire. C. R. somm. Soc. gCol. Fr., 4: 145-146. Levi, G. 1914. Sulle celadoniti di alcune localiti venete. Riv. Min. Cristall. ital., Padova, 43: 72-75. Lewis, D.W., 1964. Perigenic, a new term. J. Sedim. Petrol., 34: 875. Logvinenko, N.V., 1982. Origin of glauconite in the recent bottom sediments of the ocean. Sedim. Geol., 31: 43-48. Loreau, J.P. and Purser, B.H., 1973. Distribution and ultrastructure of Holocene ooids in the Persian Gulf. In: B.H. Purser (Edit.). The Persian Gulf. Springer Verlag, Berlin, pp. 279-328. Lorimer, G.W., 1987. Quantitative X-ray microanalysis of thin specimens in the transmission electron microscope; a review. Min. Mag., 5 1: 49-60. MacIntyre, R. M. and Hamilton, P. J., 1984. Isotopic geochemistry of lavas from sites 553 and 555. In: D. G. Roberts et al. (Edit.). Init. Rep., DSDP, 81, U.S. Govt. Print. Off, Washington, pp. 775-781.

432

Mackinnon, I.D.R. and Kaser, S.A., 1987. Microanalysis of clays at low temperature, Microbeam Analysis, San Francisco Press, (in press). Maegdefrau, E. and Hofmann, U., 1937. Glimmerartige Mineralien als Tonsubstanzen. Zeits. Krist., A98: 31-59. Manghnani, M.H. and Hower, J., 1964. Glauconite cation exchange capacities and infrared spectra. Part 11: infrared absorption characteristics of glauconites. Am. Miner., 49: 1631-1642. Martin, L., 1970. The continental margin from Cape Palmas to Lagos: bottom sediments and submarine morphology. In: The Geology of the East African Continental Margin, a Symposium, 70/16, Cambridge, pp. 79-95. Martin, L., 1973. Morphologie, sCdimentologie et palCogCographie au Quaternaire rCcent du plateau continental ivoirien. Thkse, UniversitC Paris, offset, 339 pp. Masse, J.P., 1967. Sur la prksence d'une thanatocoenose B AmphistCgines diige quaternaire B la base du plateau continental de la region de la presqu' ile du Cap Vert. Act. VIh Congr. Panafi. PrCh. Dakar, pp. 359-370. Masse, J.P., 1968. Contribution B 1'Ctude des sediments actuels du plateau continental de la region de Dakar, Rap. Labo. GCol. Fac. Sci. Univ. Dakar, 23, 81 pp., 38 Plates. Masse, J.P., 1970. Contribution B 1'Ctude de la cartographie skdimentaire du plateau continental sCnCgalais. Cons. Intern. Explor. Mer, Rapp. 159, pp. 12-14. Mathieu, R., 1968. Les sCdiments du plateau continental atlantique entre Dar-Bou-Azza et Mohammedia. Bull. Inst. Peches marit. Maroc, 16: 65-76. Maynard, J.B., 1986. Biochemistry of oolitic iron ores, an electron microprobe study. Econ. Geol., 81: 1473-1483. McConchie, D.H., Ward, J.B., McCann, V.H. and Lewis, D.W., 1979. A Mossbauer investigation of glauconite and its geological significance. Clays Clay Min., 27: 339-348. McMaster, R.L., Lachance, T.P., Ashraf, A., and Boer, J. de, 1970. Geomorphology, structure and sediments of the continental shelf and upper slope off Portuguese Guinea, Guinea and Sierra Leone. In: The Geology of the East Atlantic continental margin, a Symposium., Part IV, Cambridge, pp. 107-119. McRae, S.G., 1972. Glauconite. Earth Sci. Rev., Netherl., 8: 397-440. Mevel, C., 1979. Mineralogy and chemistry of secondary minerals in altered basalts from legs 51-52-53. In: T. Donnelly et al. (Edit.). Init. Rep. DSDP legs 51-52-53, U.S. Govt. Print. Off., Washington, pp. 1299-1312. Milliman, J.D. and Barretto, H.T., 1975. Relict magnesian calcite oolite and subsidence of the Amazon shelf, Sedimentology, 22: 137-145. Millot, G., 1964. GCologie des argiles. Masson Publ., Paris, 499 pp. Millot, G., 1970. Geology of Clays, Springer Verlag., Heidelberg, 419 pp. Moguedet, G., 1973. Contribution B 1'Ctude des sCdiments superficiels du plateau continental de la Guyane franpise, Thkse 3emecycle, Univ. Nantes, offset, 143 pp. Moniod, F., 1966. Carte de pdcipitations annuelles, Nouvelle CalCdonie. Car-

433 te 1/400.000 et Notice explicative, O.R.S.T.O.M. Publ., Paris, 11 pp. Moreau, C., Robineau, B., Bah, M.S. and Camara, I.S., 1986. Caractkres pktrographiques et gComktriques dune structure annulaire de sytnites nCphCliniques: l'archipel des fles de Los (Guide). C. R. Acad. Sci., Paris, 303: 71-74. Morelock, J., 1972. Guiana-Orinoco continental shelf sediments. Bol. Inst. Oceanogr. Univ. Oriente, 11: 57-61. Moore, M.B., 1939. Faecal pellets in relation to marine deposits. In: P.D. Trask (Edit.). Recent Marine Sediments. Am. Assoc. Petrol. Geol., Tulsa, pp. 516-524. Morton, A.C., Merriman, R.J. and Mitchell, J.G., 1984. Genesis and significance of glauconitic sediments of the SW Rockall Plateau. In: D.G. Roberts, et al., (Edit.). Init. Rep. D.S.D.P. 81, U.S. Govt. Print. Off., Washington, pp. 645-652. Murray, J. and Chumley, J., 1924. The Deep-sea deposits of the Atlantic Ocean. Trans. Roy. SOC.Edinb., 54: 252. 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 1872-1876, Londres, 525 pp. (Chapter VI, pp. 378-391). Nahon, D., Carozzi, A.V. and Parron, C., 1980. Lateritic weathering as a mechanism for the generation of ferruginous ooids. Sedim. Petrol., 50: 1287-1298. Nonis, R.M., 1964. Sediments of Chatham Rise. N. Z. Dept. scient. indust. Res. Bull., 159, 40 pp. Nota, D.J.G., 1958. Sediments of the western Guiana shelf, Publ. Min. Geol. Inst. RijksUniv. Utrecht, lOlbis, 98 pp. Nota, D.J.G., 1969. Geomorphology and sediments of western Surinam shelf. Geol. Mijnb, 48: 185-188. Odin, G.S ., 1967. Etude minkralogique et gkochronologique des formations glauconieuses dans le CrCtacC et le Tertiaire du Bassin de Paris. Dipl. Ctud. sup., UniversitC de Paris, offset, 106 pp. Odin, G.S., 1969. MCthode de sCparation des grains de glauconie, intCr6t de leur Ctude morphologique et structurale. Rev. GCogr. phys. Gtol. dynam., 11: 171-174. Odin, G.S., 1971. Recherche sur la glauconie ? l'aide i du microscope Clectronique h balayage; relations entre propriCtCs cristallographiques et structure fine. Rev. GCogr. phys. GCol. dynam., 13: 379-382. Odin, G.S., 1972a. Observations nouvelles sur la structure de la glauconie en accordCon; description du processus de genkse par nCoformation. Sedimentology, 19: 285-294. Odin, G.S., 1972b. Modalitis du passage continu du sCdiment argileux au mineral glauconite dans les formations Cocknes du Rodeberg (Flandres occidentales). C. R. Acad. Sci., Paris, 274: 660-663. Odin, G.S., 1973. RCpartition, nature minkralogique et genkse des granules verts recueillis dans les sCdiments marins actuels. Sciences Terre, Nancy, 18: 79-94.

434 Odin, G.S., 1974. Application de la microscopie Clectronique par rkflexion h 1'Ctude des minCraux argileux: exemple des minCraux des glauconies. Trav. Lab. Micropal. UniversitC Pierre et Marie Curie, Paris, 3: 297-313. Odin, G.S., 1975a. De glauconiarum, constitutione, origine, aetateque. Thkse, UniversitC Pierre et Marie Curie, Paris, offset, 280 pp. Odin, G.S., 1975b. Migrations du fer des eaux continentales jusqu'aux eaux ocCaniques profondes. C. R. Acad. Sci., Paris, 281: 1665-1668. Odin, G.S., 1982a. Ar behaviour in clays and glauconies during preheating experiments. In: G.S. Odin (Edit.). Numerical Dating in Stratigrapy. John Wiley & Sons Publ., Chichester, pp. 333-343. Odin, G.S., 1982b. How to measure glaucony ages? In: G.S. Odin (Edit.). Numerical dating in Stratigraphy. john Wilef& Sons Publ., Chichester, DD. 387-403. Odli, G.S., 1982c. Effect of pressure and temperature on clay minerals K-Ar ages. In: G.S. Odin (Edit.). Numerical dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 307-320. Odin, G.S., (Editor), 1982d. Numerical Dating in Stratigraphy. John Wiley & Sons Publ., Chichester, 2 vol., 1094 pp. Odin, G.S., 1985a. La verdine, faciks granulaire vert, marin et cetier, distinct de la glauconie: distribution actuelle et composition. C. R. Acad. Sci. Paris, 301: 105-108. Odin, G.S., 1985b. Significance of green particles (glaucony, berthierine, chlorite) in arenites. In: G.G. Zuffa (Edit.). Provenance of Arenites. NATO AS1 Ser.C, 148, D. Riedel Publ., Dordrecht, pp. 279-307. Odin, G.S., 1986. The origin of clays on the Earth. In: A.G. Cairns-Smith and H. Hartman (Edit.). Clay Minerals and the origin of life. Cambridge University Press, Cambridge, pp. 81-89. Odin, G. S. et al. (35 collaborators), 1982a. Interlaboratory standards for dating purposes. In: G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 123-150. Odin, G.S. and Bonhomme, M.G., 1982. Argon behaviour in clays and glauconies during preheating experiments. In: G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 333-343. Odin, G.S., Debenay, J.P., Froget, C. and Rigolot, P., 1987. Le faciks verdine: nCoformation dune phyllite en milieu subrkcifal. GCodynamique, (in press). Odin, G.S. and Delamare F., 1986. Nature et origine des phyllites vertes utilisCes comme pigment dans les peintures murales romaines en Gaule: ckladonite et glauconie. C . R. Acad. Sci. Paris, 302: 745-750. Odin G.S. and Dodson, M.H., 1982. Zero isotopic age of glauconies. I n : G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 277-305. Odin, G.S. and Giresse, P., 1972. Formation de minCraux phylliteux (berthitrine, smectites ferrifkres, glauconite ouverte) dans les sCdiments du Golfe de GuinCe. C. R. Acad. Sci. Paris, 275: 177-180. Odin, G.S. and Giresse, P., 1976. Essai de chronomitrie de la glauconitisation

435 dans le Golfe de GuinCe, ComplCments et remarques. C. R. somm. SOC. gCol. Fr., 3: 108-111. Odin, G.S., Hernandez, J. and Hunziker, J. C., 1986. Le volcanisme du "Biarritziano" de VCnCtie (Italie): Bges K-Ar sur basalte, plagioclase et cCladonites. In: G.S. Odin (Guest Edit.). Phanerozoic Time Scale Calibration. Chem. Geol. (Isot. Geosc. Sect.), 59: 171-180. Odin, G.S. and Hunziker, J.C., 1982. Radiometric dating of the AlbianCenomanian boundary. In: G.S. Odin (Edit.) Numerical dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 537-556. Odin, G.S. and Lamboy, M., 1975. Sur la glauconitisation d'un support carbonate dorigine organique: les dCbris dEchinodermes du plateau continental nord-espagnol. Bull. SOC.gCol. Fr., 17: 108-115. Odin, G.S. and LCtolle, R., 1980. Glauconitization and phosphatization environments: A tentative comparison. In: Y.K. Bentor (Edit.). Marine Phosphorites. SOC.Econ. Pal. Min. Spec. Publ., 29: 227-237. Odin, G.S. and Matter, A., 1981. De glauconiarum origine. Sedimentology, 28: 611-641. Odin, G.S. and Mitchell, J.G., 1983. Dating of the Palaeocene-Eocene Blosseville Group basalts, Scoresby Sund, East Greenland: a review. Newsl. Stratigr., 12: 112-121. Odin, G.S. and Morton, A., (1988, in prep.). Authigenic green particles from marine environment. In: G.V. Chilingarian and K.H. Wolf (Edit.). Diagenesis 11, Developments in Sedimentology Series, Chapt. 4, 1988. Odin, G.S., Renard, M. and Vergnaud-Grazzini, C., 1982b. Geochemical events as a mean of correlation. In: G.S. Odin (Edit.). Numerical dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 37-71. Odin, G.S. and Rex, D.C., 1982. K-Ar dating of washed, leached, weathered and reworked glauconies. In: G.S. Odin (Edit.). Numerical dating in Stratigraphy. Jonn Wiley & Sons Publ., Chichester, pp. 363-385. Odin, G.S. and Stephan, J.F. 1981. The occurrence of deep water glaucony from the Eastern Pacific: the result of in situ genesis or subsidence? In: J.S. Watkins, J.C. Moore et al. (Edit.). Init. Rep. D.S.D.P., 66, U.S. Govt. Print. Off., Washington, pp. 419-428. Odin, G.S., Velde, B. and Bonhomme, M.G., 1977. Radiogenic argon retention in glauconites as a function of mineral recrystallization. Earth Planet. Sci. Letters, 37: 154-158. Ojakangas, R.W. and Keller, W.D., 1964. Glauconitization of rhyolite sand grains. J. Sedim. Petrol. 34: 84-90. O'Neil, J.R., Clayton R.N. and Mayeda, T.K., 1969. Oxygen isotope fractionation in divalent metal carbonates. J. Chem. Phys., 5 1: 5547-5558. Orcel, J., 1927. Recherches sur la composition chimique des chlorites. Bull. SOC.franG. MinCral., 50: 75-456. Orcel, J., HCnin, S. and Caillkre, S., 1949. Sur les silicates phylliteux des minerais de fer oolithiques. C. R. Acad. Sci. Paris, 229: 134-135. Pa&, J., Debenay, J.P. and Lebrusq, J.Y., 1987. L' environnement estuarien de la Casamance. Rev. Hydrobiol. tropicale, in press.

436 Parry, W.T. and Reeves, C.C., 1966. Lacustrine glauconitic mica from pluvial Lake Mound, Lym and Terry counties, Texas. Am. Miner., 51: 229-235. Pichard, C. and Frohlich, F., 1986. Analyses IR quantitatives des sCdiments. Exemple du dosage du quartz et de la calcite. Rev. I.F.P., 41: 809-819. Pinson, J., 1980. Les environnements skdimentaires actuels et quaternaires du plateau continental sCnCgalais. Th&se,Univ. Bordeaux I, offset, 125 pp. Polevaya, N.I., Murina, G.A., and Kazakov, G.A., 1961. Glauconite in absolute dating. Ann. N.Y. Acad. Sci., 91: 298-310. Porrenga, D.H., 1965. Chamosite in Recent sediments of the Niger and Orinoco deltas. Geol .Mijnb., 44: 400-403. Porrenga, D.H., 1966. Clay Minerals in recent sediments of the Niger Delta. Proc. of 14th Nat. Conf. Clays Clay Min., Pergamon Press, Oxford, pp. 221-233. Porrenga, D.H., 1967a. Clay mineralogy and geochemistry of recent marine sediments in tropical areas. Publ. Fysish-Geograf. Labor. Univ. Amsterdam, 9, 145 pp. Porrenga, D.H., 1967b. Glauconite and chamosite as depth indicators in the marine environment. Mar. Geol., 5: 495-501. Pratt, W.L., 1963. Glauconite from the sea floor off Southern California. Essays in marine geology in honor of K.O. Emery, Los Angeles Univ. Southern Calif. Press, 97: 119. Pryor, W.A., 1975. Biogenic sedimentation and alteration of argillaceous sediments in shallow-marine environments. Bull. geol. SOC.Am. 86: 1244-1254. Pujos, M. and Odin, G.S., 1986. La sddimentation au Quaternaire terminal sur la plateforme continentale de Guyane frantpise. Oceanol. Acta, 9: 363-382. ReniC, O., 1983. Sedimentation ditritique et bioghe, authigenkses femfires du plateau continental de la Guynane franpise: apports i~ la palCogCographie du Quaternaire rCcent MCm. DEA, Bordeaux, offset, 88 pp. Reynolds, P. M., 1976. 39Ar/40Ardating of Leg 34 basalts. In : R.S. Yeats et al. (Edit.). Init. Rep. DSDP, 34, U.S. Govt. Print. Off., Washington, pp. 449-450. Richter, D.K., 1983. Calcareous ooids: a synopsis. In: T.M. Peryt, (Edit.). Coated grains. Springer Verlag, Berlin, pp. 7 1-99. Robert, C. and Odin, G.S., 1975. Niveaux glauconieux dans les sCdiments rCcents du seuil Nord-EgCen. Bull. Gr. franG. Argiles, 27: 1-1 1. Roche, M.A., 1977. L'estuaire du Kourou en Guyane fraqaise; possibilitCs dalimentation en eau dune usine de phe de bois et risques de pollution de rejets industriels dans l'estuaire. Rapport O.R.S.T.O.M., Centre de Cayenne, 80 pp. Roeleveld, W., 1969. Pollen analysis of two sections in the young coastal plain of Surinam. Geol. Mijnb, 48: 215-224. Rohrlich, V., Price, N.B. and Calvert, S.E., 1969. Chamosite in the recent sediments of Loch Etive (Scotland). J. Sedim. Petrol., 39: 624-631. Rougerie, F., 1986. Le lagon sud ouest de Nouvelle CalCdonie: spCcificitC hydrologique dynamique et productivitC. Trav. Doc. Sci. Techn., O.R.S.T.

437 O.M., Paris, (sous-presse). Rozenson, I. and Heller-Kallai, L., 1977. Mossbauer spectra of dioctahedral smectites. Clays Clay Min., 25: 94-101. Ruffman, A., Meagher, L.J. and Mc Steward, G., 1977. BathymCtrie du talus, et du plateau continental du SCnCgal et de la Gambie, Afrique de 1'Ouest. Geomarine Assoc. ltd., New Scotia, Canada, 130 pp. Santos, M.E., 1972. Paleogeografia do quaternario superior na plataforma continental norte brasileira, Anais do 26 Congress0 brasileiro de geologia, Belem, pp. 267-288. Savin, S.M. and Epstein, S., 1970. The oxygen and hydrogen isotope geochemistry of clay minerals. Geochim. Cosmochim. Acta, 34: 25-42. Sawhney, B.L., 1967. Interstratification in vermiculite. Clays Clay Min., 27: 75-84. Scheidegger, K.F. and Stakes D.S., 1979. X-Ray diffraction and chemical study of secondary minerals from holes 417A-417D. In: T. Donelly, et al. (Edit.). Init. Rep. DSDP, 51-52-53, U.S. Govt. Print. Off., Washington, pp. 1253-1263. Schellmann, W., 1969. Die Bildungsbedingungen sedimentarer Chamosit und Hamatit-Eisenerze am Beispiel der Lagerstatte Echte. N. Jb. Mineral, 111: 1-31. Scherillo, A. 1938. Sulle rocce eruttive preterziarie della Sicilia Occidentale. Period. Miner. Roma, 9: 61-84. Seed, D.P., 1965. The formation of vermicular pellets in New-Zealand glauconites. Am. Mineralogist, 50: 1097-1 106. Seed, D.P., 1968. The analysis of the clay content of some glauconite oceanic sediments. J. Sedim. Petrol., 38: 229-231. Seidemann, D. E., 1976. K-Ar dates for basaltic rocks from site 319 and 321, Leg 34. In: R.S. Yeats et al. (Edit.). Init. Rep. DSDP, 34, U.S. Govt. Print. Off., Washington, pp. 445-447. Seyfried, W.E., Shanks, W.C. and Bischoff, J.L., 1976. Alteration and vein formation in site 321 basalts. In: R.S. Yeats et al., (Edit.). Init. Rep. DSDP, 34, U.S. Govt. Print. Off., Washington, pp. 384-392. Seyfried, W. E., Shanks, W. C. and Dibble, W. E. 1978. Clay mineral formation in DSDP Leg 34 basalt. Earth Planet. Sci. Lett. 41,265-275. Shackleton, N.J., and Kennett, J.P., 1975. Paleotemperature history of the Cenozoic and the initiation of Antarctic glaciation: oxygen and carbon isotope analyses in DSDP sites 277, 279, and 291. In: J.P. Kennett, R.E. Houtz et al. (Edit.).Init. Rep. D.S.D.P., 29, U.S. Govt. Print. Office, Washington, pp. 743-755. Shirozu, H., 1958. X-ray powder patterns and cell dimensions of some chlorites in Japan. Mineral. J. (Japan), 2: 209-223. Sholkovitz, E.R., Van Grieken, R. and Eisma, D.,1978. The major element composition of suspended matter in the Zaire River and estuary. Netherl. J. Sea Res., 12 : 407-413. Shutov, V.D., Katz, M.Y., Drits, V.A., Sokolova, A.L. and Kasakov, G.A., 1970. Crystallo-chemical heterogeneity of glauconite as depending of the

438 conditions of its formation and postsedimentary change. Intern. Clay Conf., Madrid, I: 327-339. Siehl, A. and Thein, J., 1978. Geochemische trends in der Minette (Jura, Luxembourg, Lothringen). Geol. Rundsch., 67: 1052-1077. Siehl, A. and Thein, J., 1987. Origin of minette-type ironstones. In: Phanerozoic ironstones, Symposium, Abstracts and Progr. pp., 44-45. Simpson, E.S.W., 1970. The geology of the south west African continental margin: a review. In: The geology of the East Atlantic Continental margin. Rep. Inst. Geol. Sci. 70/16, pp. 157-170. Slatt, R.M., Hodge, R.A. and Uzuakpunwa, A., 1972. Glauconite in surficial sediments as an indicator of underlying CretaceousiTemary bedrock on the northeast Newfounland continental shelf. Can. J. Earth Sci. 9: 1441-1446. Slonimskaya, M.V., Besson, G., Dainyak, L.G., Tchoubar, C. and Drits, V.A., 1986. Interpretation of the IR spectra of celadonites and glauconites in the region of OH-stretching frequencies. Clay Min., 21: 377-388. Smulikowski, K.,1954. The problem of glauconite. Polska Akad. Nauk, Komitet Geol., Arc. Min. Warsaw, 18: 21-120. Sorby, H.C., 1856. On the origin of the Cleveland Hill ironstone. Geol. Polytechnic Soc.West Riding Yorkshire Roc., 3: 457-461. Staudigel, M. Gillis, K. and Duncan, R. 1986. K-Ar and Rb-Sr ages of celadonites from the Troodos ophiolite, Cyprus. Geology, 14: 72-75. Stakes, D.S. and O'Neil, J.R., 1982. Mineralogy and stable isotope geochem i s t r y of hydrothermally altered oceanic rocks. Earth Planet. Sci. Lett., 57: 285 -304. Storzer, D. and Selo, M., 1979 Fission track age of magnetic anomaly M-zero and some aspects of sea water weathering. In: T. Donelly et al. (Edit.). Init. Rep. DSDP, 51-52-53, U.S. Govt. Print. Off., Washington, pp. 11291133 Strasser, A., 1986. Ooids in Purbeck limestones of the Swiss and French Jura. Sedimentology, 33: 71 1-727. Sudo, T., Shimoda, S., Yotsumoto, M. and Aita, S., 1981. Electron micrographs of Clay Minerals, Developments in Sedimentology, 31 : 141-142. Takahashi, J.I. and Yagi, T., 1929. The peculiar mud grains in the recent littoral and estuarine deposits, with special reference to the origin of glauconite. Econ. Geol., 24: 838-852. Talbot, M.R., 1973. Major sedimentary cycles in the Corallian beds. Palaeogeogr. Palaeoclimat. Palaeoecol., 14: 293-317. Talbot, M.R., 1974. Ironstones in the Upper Oxfordian of southern England. Sedimentology, 21: 433-450. Taylor, J.H., 1949. Petrology of the Northampton sand ironstone formation. Geol. Surv. Mem., G.B., Dept. Scient. Ind. Res., 1 1 1 pp. Tomita, K. and Takahashi, H., 1986. Quantification curve for the X-ray powder diffraction analysis of mixed-layer kaolinite/smectite. Clays Clay Min., 34: 323-329. Tooms, J.S., Summerhayes, C.P. and Mc Master, R.L., 1970. Marine geological studies on the North West African margin: Rabat-Dakar. In: The

439 Geology of the East Atlantic Continental margin. Rep Inst. Geol. Sci., 70/16, pp. 9-25. Trauth, N., Sommer, F. and Lucas, J., 1969. Evolution gCochimique dune sCrie sCdimentaire palCogkne dans le Bassin de Paris. Bull. Serv. Carte gCol Als. Lon., 22: 279-310. Triplehorn, D.M., 1966. Morphology, internal structure and origin of glauconite pellets. Sedimentology, 6: 247-266. Triplehorn, D.M., 1967. Morphology, internal structure and origin of glauconite pellets, A Reply. Sedimentology, 8: 169-17 1. Ushakov, P.V., 1970. Observations sur la rbpartition de la faune bentique du littoral guinCen. Cah. Biol. Mar., 11 : 435-457. Vali, H. and Koster, H.M., 1986. Expanding behaviour, structural disorder, regular and random irregular interstratification of 2: 1 layer-silicates studied by high resolution images of transmission electron microscopy, Clay Min., 21: 827-859. Van Houten, F.B. and Purucker, M.E., 1984. Glauconitic peloids and chamositic oids; Favourable factors, constraints and problems. Earth Sci. Rev., 20: 211-243. Van der Marel, H.W. and Beutelspacher, H., 1976. Atlas of Infra-Red Spectroscopy of clay minerals and their admixtures. Elsevier Publ., Amsterdam, 396 pp. Velde, B. and Odin, G.S., 1975. Further information related to the origin of glauconite. Clays Clay Min., 23: 376-381. Von Gaertner, H. R. and Schellmann, W., 1965. Rezente sedimente in Kustenbereich der Halbinsel Kaloun, Guinea. Tscherm. miner. Petr. Mitt., 10: 349-367. Warshaw, C.M., 1957. The mineralogy of glauconite. Ph. D. Thesis, Penns. State Univ., University Park, 155 pp. Webb, A.W., McDougall, I. and Cooper, J.A., 1963. Retention of radiogenic argon in glauconites from Proterozoic sediments, Northern Territory Australia. Nature, 199: 270-27 1. Wells, R.C., 1937. Analyses of rocks and minerals 1914-1936. Bull. Geol. Surv., 878: 102. Wermund, E.G., 1961. Glauconite in early tertiary sediments of gulf coastal province. Bull. Am. Assoc. Petr. Geol., 45: 1667-1696. Wicquart, E., 1983. Modkle lithostratigraphique du plateau de KerguelenHeard (OcCan Indien). Thkse 3eme cycle, Univ. P. & M. Curie, Paris, no 83-37. 135 pp. Wicquart, E. and Frohlich, F., 1986. La sidimentation sur le plateau de Kerguelen-Heard. Relations avec 1'Cvolution de l'OcCan Indien au CCnozoi'que. Bull. Soc. gCol. Fr., 8: 569-574. Wise, W.S. and Eugster, M.P., 1964. Celadonite: synthesis, thermal stability and occurrence. Am. Miner., 49: 1031-1083. Yeh, H.W. and Savin, S.M., 1977. The mechanism of burial metamorphism of argillaceous sediments, 3. Oxygen isotopic evidence. Geol. SOC.Amer. Bulletin, 88: 1321-1330.

440 Yerskova, Z.P., Nikitina, A.P., Perfilev, Y.D. and Babeshkin, A.M., 1976. Study of chamosites by Mossbauer spectroscopy. In: S.W. Bailey (Edit.). Proceedings Intern. Clay Conf. 1975, pp. 21 1-219. Zimmerman, J. L. and Odin, G. S., 1982. Kinetics of the release of Ar and fluids from glauconies. In: G.S. Odin (Edit.). Numerical Dating in Stratigraphy. John Wiley & Sons Publ., Chichester, pp. 345-354. Zumpe, H., 1971. Microstructures in Cenomanian glauconite from the Isle of Wight (England). Mineral. Mag., G.B., 38: 215-224.

441

INDEX of COLLABORATORS This work has been realized with the help of the following collaborators quoted in the text for either providing samples, or obtaining original results, or discussion and improvement of the manuscript, or contribution to the practical realization of the volume. : 219 Lebrun E. : 123, 130, 179,204,236,405 Aharon P. Andreolli D. : 398 Ledoux R. : 28 Arbey F. : 21,28, 52 LenobleM. : 74, 104, 123, 130, 140, 179, Bailey S.W. : 81,405 183, 204,236,247,275, 366, Blondeau A. : 275 369, 398, 405, Bonhomme M.G.: 247 LCtolle R. : 406 Brindley G.W. : 204 Lowe D.R. : 219 Mahabaleswar B.: 52 Buckley H.A. : 369,398 ChagnaudM. : 111, 130 Mariotti A. :406 Martin L. : 149, 157, 158,247 Champetier Y. : 38, 52 Chevillon C. : 59, 63, 81 Masse J.P. : 60, 81, 225 Curator British Museum: 340,385 Mathieu R. : 225 Matter A. : 259,263,298 Curator Museum Paris: 20,40, 52, 340, 385 Da Matha S . : 406 Millot G. : 204 Debenay J.P. : 59, 60, 81, 103 MoguedetG. : 105, 157, 158 Monteillet Y. : 85 Delaloye M.F. : 5 , 52, 388 De Wever P. : 388 MoreauC. : 138, 158 Nizzero Don Elio: 398 Dupeuble P.A. : 275 : 294 Odin Ch. : 406 Ehrlich A. Faugtxes J.C. : 225,294, 328 0' Neil J.R. : 238 Faure H. : 101, 241 Pachera E. : 398 : 28, 52, 104, 204, 219, 247 Person A. : 52 Fay 0. : 81 PetzoldM. 294, 398,406 : 85 Pinson J. Fripiat J.J. : 204 Poullen J.F. : 369, 398 Froget C. : 60, 81, 157 : 204 Prost R. Frohlich F. : 147 Pujos M. : 157 Fullagar P.D. : 239, 333,405 : 104, 144, 158,294 RCcy J. : 81 Gayet J. : 111, 130 RCniC 0. : 52 GehringU. : 247 Rex D.C. GeyhH.A. : 109, 130 Richebois G. : 236,247,275,406 Gillot P.Y. : 385 Giresse P. : 157, 158,247 Richer de Forges B.: 59,63,64,81 : 59,63,64,81 Rigolot P. : 59, 63, 64,81 Gout B. RobineauB. : 138, 158 : 28 GuerrakS. Hunziker J.C. : 247, 275, 397 Rouget G. : 219 : 382 Selo M. Hurf0rdA.J. : 397 : 103 Skelton P. Jakob J. : 20 Keppens E. : 238, 247, 307 Sorbini L. : 398 Thomassin B. : 59,63,64,81 KreuzerH. : 247 Thompson J. : 247,406 Kmmmenacher G.: 19,20 : 28 Waibel A. LaverneC. : 371

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443

GEOGRAPHICAL INDEX Aegean Sea

Africa Algeria Amazon R. Ar(ienne hk0M

Atlantic Ocean Australia Bahamas Baie de Rufiique Bas Vallon Bellingshausen Bermuda Rise Bohemia Borneo (shelf off) Brazil California Cap Vert Casamance R. Casamance (shelf) Chamoson Chatham Rise China Comoe R. Congo R. Congo (shelf off) Columbia Cyprus England Faeroe French Guiana Gabon Galapagos Ridge Great Britain Greenland Guinea Gulf Coast Gulf of Guinea Guyana

: 321 : 101,212,221,225,249, 277 : 34, 35, 38,46,47 : 53, 105 to 107, 110, 125 to 130, 135, 157,226 : 46 : 340 : 83, 157, 158, 207, 221, 225, 245, 246, 277, 326, 338 to 342, 371,

376,380, 395 : 57, 158,277,296 : 48 : 84 to 86,90,92,94,99, 153, 161, 169, 171, 182,205,210,211,415 : 20,21,408 : 381 : 340,380 to 383 : 338, 339,366,368 : 45, 142 to 144, 156 : 109, 127, 135,340 : 258,312,318, 325,327,340,379 : 83 to 92, 155,205,211,225,403 : 153 to 157, 169, 171,205 to 207,415 : 225,327, 328 : 2, 5 to 28, 33, 39 to 43, 202,408 : 270,321, 325,326,328,403 : 296 : 149 to 151, 157,206,225,226 : 145 to 149, 157, 179, 185, 186,207,208,225,226,228 : 158,227 to 247,306 to 308,312,321,329 : 125 : 338 to 345, 347, 352, 366, 367, 374, 375, 378, 380, 381, 387 to 390 : 36,39,41 : 338 to 344,348,353, 371, 375, : 2, 53, 105 to 130, 132, 135, 140, 148, 161 to 170, 177, 179, 185, 187, 190, 197, 199, 200,205 to 209,212,214,272,418 : 45, 131, 158,214,226,227,229, 291 : 371, 392,393 : 27,28 : 386 : 138 to 142, 148, 158, 159, 163, 171 to 173, 179, 193, 195, 196, 199, 205 to 207,225 : 196 : 146, 147, 150, 183, 225 to 247, 267, 270, 271, 273, 307, 318, 325, 418 : 109, 129, 134

444 Iceland : 338,339,341,352 Indianocean : 152, 158,277,278,293,326,339,380,402 Indonesia : 277,325 bland : 338,339,365,367 : 45, 183, 184,214,225 to 228,235,240,241,243 Ivory coast Japan (margin) : 326 Java : 277 Kerguelen (Plateau): 277 to 294, 299, 324 to 326,403 : 138, 139, 141, 153, 156, 157, 179,206,208, 209,226 Konkoure R. :48 LagunaMadre Limburg : 301 : 9,27,28,36, 38 to41,44,46, 160,413 Lorraine Maciagascar : 152, 153, 157,365,366 : 96,97, 144, 158,277,294,297,324 to 326 Makassar S n i t Malac€a : 144 Malvinas (islands) : 277 : 106 to108, 112, 126, 127 Maroni R. Mauritania : 225 Mayotte (island) : 152, 153, 157, 206, 207, 210 : 337 to 349,351,353,366,367,369,375,383 to 386,391 Monte Baldo Morocco 225,297 Nevada : 340, 342,353,367 : 2, 53, 57 to 81, 92, 99, 101, 103, 112, 114, 124, 125, 128, 129, 159 NewCaledonia to 173, 179, 185, 186, 190, 193, 196, 197, 199, 205. 207, 210, 214, 403,418 New Jersey : 340,366 New Mexico : 342,367 New Zealand : 340 : 134 to 139, 150, 152, 156, 157,225,228,233 Niger R. (delta) Nigeria (shelf off) : 229,240,243 North Africa : 28 Norway (shelf off) : 297, 325, 326 : 131 to 134, 139, 145 to 149, 156, 157, 183,205,207,225,226,228 Ogooue R. Oregon : 353 : 105, 106, 129, 130, 134, 135, 156, 157 Orinoco R. : 106, 107, 1.12, 126, 127 Oyapock R. : 57, 157, 158,277,293, 339, 340, 370, 371,402,403 Pacific Ocean Panama : 327 Paris Basin : 260,262,263,298,329 Persian Gulf : 48 Peru, Chile Trench : 380, 381 : 299, 326. 338, 342 to 344, 347, 362, 363, 365, 371, 374, 375, 381, RockallPlateau 386,387 Sarawak (shelf off) : 142 to 144,207, 318 Scotia Ridge : 277,293,325,326 Scotland : 5,39, 338,339, 365

445 : 83,87, 157 : 43, 53, 83 to 87, 112, 114, 124, 129, 138 to 140, 148, 155, 161 to 171, 174 to 179, 183, 185, 187 to192, 196, 199,206, 207. 211, 213, 241, 327,403,418 Sicily : 338 to 340,345,366, 367 South Africa : 270,277,291,312,321,325, 328 South America : 105,212 Spain northwestern : 138,221,249 to 275,299,312,318,328,369 South China Sea : 156 Sumatra : 144, 277 Surinam : 105, 109, 128, 129 Switzerland : 31, 32, 36, 37, 39,45,46 Teruel : 338,339, 341,343, 344,352,375,390, 391 Tirnor : 277 Turkey : 337, 339,366,375 Tyrol : 339 United States (East): 297, 325 U.S.S.R. : 296, 340, 366 : 338,340, 342,352,375 Val di Fassa Vancouver Island : 321, 325 : 109, 129 Venezuela Vesuvius : 365 to 367 Voring Plateau : 338, 342, 344 Washington : 340, 351,366,367 World (distribution): 157, 325 : 32, 39 Yorkshire

Senegal R. Senegal shelf

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