Green Marine Clays (Developments in Sedimentology)

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.

This Page Intentionally Left Blank

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

This Page Intentionally Left Blank

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.

This Page Intentionally Left Blank

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

This Page Intentionally Left Blank

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