Clay Mineral Cements in Sandstones - media control · 2013-07-23 · for clay mineral cements in sandstones P.J. Hamilton Chlorite case study 291 Chlorite authigenesis and porosity

SPECIAL PUBLICATION NUMBER 34 OF THE INTERNATIONALASSOCIATION OF SEDIMENTOLOGISTS

Clay Mineral Cements in Sandstones

EDITED BY

Richard H. Worden and Sadoon Morad

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CLAY MINERAL CEMENTS IN SANDSTONES

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SPECIAL PUBLICATION NUMBER 34 OF THE INTERNATIONALASSOCIATION OF SEDIMENTOLOGISTS

Clay Mineral Cements in Sandstones

EDITED BY

Richard H. Worden and Sadoon Morad

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© 2003 International Association of Sedimentologistsand published for them byBlackwell Science Ltda Blackwell Publishing company

350 Main Street, Malden, MA 02148-5018, USA108 Cowley Road, Oxford OX4 1JF, UK550 Swanston Street, Carlton, Victoria 3053, AustraliaKurfürstendamm 57, 10707 Berlin, Germany

The rights of Richard Worden and Sadoon Morad to be identified as theAuthors of the Editorial Material in this Work has been asserted inaccordance with the UK Copyright, Designs, and Patents Act 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, except aspermitted by the UK Copyright, Designs, and Patents Act 1988, withoutthe prior permission of the publisher.

First published 2003

Library of Congress Cataloging-in-Publication Data

Clay cements in sandstones/edited by Richard H. Worden and SadoonMorad.

p. cm. — (Special publication number 34 of theInternational Association of Sedimentologists)

Includes bibliographical references and index.ISBN 1-40510-587-9 (pbk.: alk. paper)1. Clay minerals. 2. Sandstone. I. Worden, Richard H.

II. Morad, Sadoon. III. Special publication . . . of the International Association of Sedimentologists; no. 34.

QE389.625.C517 2002549′.6—dc21

2002070932

A catalogue record for this title is available from the British Library.

Set in 9.5/12pt Meliorby Graphicraft Limited, Hong KongPrinted and bound in the United Kingdomby MPG Books Ltd., Bodmin, Cornwall

For further information onBlackwell Publishing, visit our website:http://www.blackwellpublishing.com

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Contents

vii Introductionviii Acknowledgements

Review papers

3 Clay minerals in sandstones: controls onformation, distribution and evolutionR.H. Worden and S. Morad

43 Predictive diagenetic clay-mineral dis-tribution in siliciclastic rocks within asequence stratigraphic frameworkJ.M. Ketzer, S. Morad and A. Amorosi

63 Oxygen and hydrogen isotopic composi-tion of diagenetic clay minerals in sand-stones: a review of the data and controlsS.H. Morad, R.H. Worden and J.M. Ketzer

93 Palaeoclimate controls on spectralgamma-ray radiation from sandstonesA.H. Ruffell, R.H. Worden and R. Evans

109 Smectite in sandstones: a review of thecontrols on occurrence and behaviourduring diagenesisJ.M. McKinley, R.H. Worden and A.H.Ruffell

129 Patterns of clay mineral diagenesis ininterbedded mudrocks and sandstones:an example from the Palaeocene of theNorth SeaH.F. Shaw and D.M. Conybeare

147 Cross-formational flux of aluminium andpotassium in Gulf Coast (USA) sedimentsM. Wilkinson, R.S. Haszeldine and K.L.Milliken

161 Silicate–carbonate reactions in sediment-ary systems: fluid composition control andpotential for generation of overpressureI. Hutcheon and S. Desrocher

177 Experimental studies of clay mineraloccurrenceD.A.C. Manning

191 Effect of clay content upon some physicalproperties of sandstone reservoirsP.F. Worthington

213 Quantitative analysis of clay and otherminerals in sandstones by X-ray powderdiffraction (XRPD)S. Hillier

253 A review of radiometric dating techniquesfor clay mineral cements in sandstonesP.J. Hamilton

Chlorite case study

291 Chlorite authigenesis and porositypreservation in the Upper Cretaceousmarine sandstones of the Santos Basin,offshore eastern BrazilS.M.C. Anjos, L.F. De Ros and C.M.A.Silva

Kaolinite case studies

319 Origin and diagenetic evolution of kaolinin reservoir sandstones and associatedshales of the Jurassic and Cretaceous,Salam Field, Western Desert (Egypt)R. Marfil, A. Delgado, C. Rossi, A. LaIglesia and K. Ramseyer

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vi Contents

343 Microscale distribution of kaolinite inBreathitt Formation sandstones (middlePennsylvanian): implications for massbalanceK.L. Milliken

361 The role of the Cimmerian Unconformity(Early Cretaceous) in the kaolinitizationand related reservoir-quality evolution inTriassic sandstones of the Snorre Field,North SeaJ.M. Ketzer, S. Morad, J.P. Nystuen andL.F. De Ros

383 The formation and stability of kaolinite inBrent sandstone reservoirs: a modellingapproachÉ. Brosse, T. Margueron, C. Cassou, B.Sanjuan, A. Canham, J.-P. Girard, J.-C.Lacharpagne and F. Sommer

Illite case studies

411 Illite fluorescence microscopy: a newtechnique in the study of illite in theMerrimelia Formation, Cooper Basin,AustraliaN.M. Lemon and C.J. Cubitt

425 Geochemical modelling of diageneticillite and quartz cement formation inBrent sandstone reservoirs: example ofthe Hild Field, Norwegian North SeaB. Sanjuan, J.-P. Girard, S. Lanini, A.Bourguignon and É. Brosse

453 The effect of oil emplacement on diage-netic clay mineralogy: the Upper JurassicMagnus Sandstone Member, North SeaR.H. Worden and S.A. Barclay

Glauconite case study

473 Application of glauconite morphology in geosteering and for on-site reservoirquality assessment in very fine-grainedsandstones: Carnarvon Basin, AustraliaJ.P. Schulz-Rojahn, D.A. Seeburger andG.J. Beacher

489 Index

Colour plates facing p. 296, p. 392 and p. 424.

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Clays are one of the most important groups of minerals that destroy permeability in sand-stones, but they also react with drilling and com-pletion fluids and induce fine-particle migrationduring hydrocarbon production. They are a verycomplex family of minerals that commonly aremutually intergrown and contain a wide rangeof solid solutions and form by a wide range ofprocesses. They form under a wide diversity ofpressure and temperature conditions, as well asrock and fluid compositional conditions.

In this volume, clay minerals in sandstonesare reviewed in terms of their mineralogy andgeneral occurrence, their stable and radiogenicisotope geochemistry and their relationship tosequence stratigraphy and palaeoclimate. Theirrelationship to the petrophysical properties ofsandstones and their use in drilling technolo-gies are also covered. The controls on the vari-ous clay minerals are addressed and a variety ofgeochemical issues, including the importanceof mass flux, links to carbonate mineral diagen-esis and linked clay diagenesis in interbeddedmudstone–sandstone are explored. The thornyissue of the quantification of clay mineralsusing XRD (X-ray diffraction) is tackled from a variety of technical angles. A number of casestudies are included for kaolin, illite and chlorite clay cements, and the occurrence of

smectite in sandstone also is reviewed. Clayminerals grow at rates that are only just becom-ing known; experimental data on clay cementsin sandstones are thus reviewed and there aretwo model-based case studies that address therates of growth of kaolinite and illite.

This volume follows a volume on CarbonateCementation in Sandstones (IAS SpecialPublication 26, edited by Sadoon Morad) and avolume on Quartz Cementation in Sandstones(IAS Special Publication 29, edited by RichardWorden and Sadoon Morad). With the publica-tion of this Special Publication, the majority ofmineral cements in sandstones will have beenaddressed in books sponsored by the IAS.

Amongst its readership this volume willattract a wide range of scientists and techno-logists including: (i) sedimentologists and petrographers who deal with the occurrence, spatial and temporal distribution patterns andimportance of clay cements in sandstones, (ii)geochemists who are involved in unravellingthe factors that control clay cement formationin sandstones and (iii) petroleum geoscientistswho need to predict areas of a play that maycontain formation-damaging clay minerals. The book also will be of interest to geologistsinvolved in palaeoclimate studies, basin ana-lysis and basin modelling.

Introduction

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Acknowledgements

Crucial to all refereed publications are the referees themselves. For this service, they receive noreal recognition and yet their efforts commonly lead to much improved final products. Wewould therefore like to acknowledge the following reviewers for their hard work, dedicationand care in reviewing the 21 papers published in this volume.

Alastair RuffellAlessandro AmorosiAndrew HoggAndrew HurstAnne GrauAtilla JuhaszBruce YardleyCathal DillonChris RochelleChristoph SpötlCraig SmalleyDave DoffDave ManningDavid AwwillerDewey Moore

Earle McBrideEtienne BrosseGhazi KraishanGiles BergerJean-Francois DeconinckJeff GrigsbyJim BolesKevin TaylorKitty Lou MillikenKnut BjørlykkeLou MacchiLyndsay KayeMark WilkinsonMichael Wilson

Mike MayallNorbert ClauerNorman OxtobyPatrick CorbettPer AagaardReinhard GauppRichard EvansShelagh BainesShirley DuttonSteve FranksStuart HaszeldineSusanne SchmidTony FallickWarren Dickenson

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Review papers

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INTRODUCTION

The amount, distribution pattern and mor-phology of clay minerals have significant effectson sandstone properties in terms of porosity,permeability, density, natural radioactivity,electrical conductivity, the water content of

petroleum fields and reactivity to variousenhanced oil recovery practices. Prior to theroutine use of the scanning electron micro-scope (SEM) in petrographic examination ofsandstones, clay minerals were often wronglyassumed to be detrital in origin, being co-deposited with the primary host sand.

Clay minerals in sandstones: controls on formation,distribution and evolution

R.H. WORDEN1 and S. MORAD2

1 Department of Earth Sciences, University of Liverpool, Brownlow Street, Liverpool L69 3GP, UK, e-mail: [email protected]

2 Department of Earth Sciences, Uppsala University, Villa vägen 16, S-752 36 Uppsala, Sweden, e-mail: [email protected]

ABSTRACT

This paper addresses the origin, distribution pattern and burial diageneticevolution of clay minerals in sandstone: kaolin, smectite, illite, chlorite,berthierine, glauconite and mixed-layer illite–smectite and chlorite–smectite. Clay minerals may be co-deposited with sand grains as sand-sizedargillaceous intra- and extra-clasts and as flocculated clays. These sand-sized argillaceous clasts are deformed by mechanical compaction into claypseudomatrix. Detrital clay minerals may be incorporated into sandydeposits by bioturbation and infiltration of muddy waters. Diagenetic clayminerals form by alteration of unstable detrital silicates and by transforma-tion of detrital and precursor diagenetic clay minerals. The most commoneogenetic clay minerals are kaolinite, dioctahedral and trioctahedral smec-tite, berthierine, glauconite and, less commonly, Mg-rich clay mineralssuch as palygorskite. The distribution of eogenetic clay minerals is stronglyrelated to depositional facies and sequence stratigraphic surfaces. Illite andchlorite dominate the mesogenetic clay minerals and usually grow at theexpense of eogenetic clay minerals and detrital feldspars and lithic grains.Mesogenetic illite and chlorite can result from widely different reactantsand processes. Clay minerals usually are assumed to be detrimental tosandstone reservoir quality because they can plug pore throats and someclay minerals promote chemical compaction. However, coats of chlorite on sand grains can preserve reservoir quality because they prevent quartzcementation. Adding oil to a sandstone stops clay diagenesis if the sand-stone is oil-wet but probably only slows clay reactions if the sandstone is water-wet. Sandstones tend to be more oil-wet as the Fe-bearing clay content of the sand increases and as oil becomes more enriched in polarcompounds.

Int. Assoc. Sedimentol. Spec. Publ. (2003) 34, 3–41

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4 R.H. Worden and S. Morad

However, laws of hydrodynamics tend to causeseparation of the clay- and sand-sized fractions,implying that post-depositional processes mustbe responsible for the incorporation of mostclay minerals into sandstones. The intention ofthis paper is to review the following:1 composition and mineralogy of clay mineralsin sandstones;2 how clay minerals are incorporated into sandsprior to diagenesis;3 early diagenetic (eogenetic) origin of clayminerals in sandstones;4 burial diagenetic (mesogenetic) origin of clayminerals in sandstones;5 uplift-related diagenetic (telogenetic) originof clay minerals in sandstones;6 effects of clay minerals on sandstone permeability;7 effect of petroleum emplacement on clay diagenesis in sandstones.

Definitions

The word ‘clay mineral’ refers to diverse groupsof minerals that are members of the hydrousaluminous phyllosilicates, whereas the word‘clay’ is strictly a grain-size term, classically forparticle diameters less than 3.9 µm (Wentworth,1922). Unfortunately, in sedimentary petrology,the term ‘clay’ is frequently used synonymouslywith ‘clay mineral’.

Eodiagenesis includes all processes thatoccur at or near the sediment surface, where the geochemistry of the interstitial waters iscontrolled mainly by the depositional environ-ment. Eodiagenesis also can be defined in termsof temperature, and depth, where the uppertemperature limit is < 70°C, typically equival-ent to about 2 km burial (Morad et al., 2000).

Mesodiagenesis occurs during burial andincludes all diagenetic processes followingeodiagenesis and through to the earliest stagesof low-grade metamorphism (as defined byChoquette & Pray, 1970). In many cases, thisincludes sediments buried to depths withequivalent temperatures of about 200 to 250°C.The main factors that influence mesogeneticchanges include the time–temperature history,

the primary mineralogy and fabric, local eo-genetic modifications, extent of material loss and gain to neighbouring lithologies, geochem-istry of the pore water and the presence ofpetroleum-related fluids.

Telodiagenesis occurs in inverted basins thathave experienced an influx of surface (usuallymeteoric) waters. Such water has the capa-city to cause significant geochemical changes,including feldspar dissolution and alteration tokaolinite.

There are eight main ways that clay mineralsare incorporated into sandstones:1 clay-rich rock fragments formed in the hinterland (extraclastic, allochthonous);2 clay-rich clasts formed within the sediment-ary basin (intraclastic, autochthonous);3 flocculated mud particles and faecal pellets;4 inherited clay rims on sands grains;5 post-depositional incorporation of detritalmud into the sandstone by bioturbation andclay infiltration;6 eogenetic reaction products in sandstone;7 mesogenetic reactions in sandstones;8 telogenetic reactions in sandstones.

CLAY MINERALS IN SANDSTONES:SUMMARY OF CHEMISTRY ANDSTRUCTURE

General aspects of clay mineral structure

Clay minerals are hydrous aluminosilicatesthat belong to the phyllosilicate group of miner-als (Deer et al., 1998). In addition to aluminiumand silicon they also may contain other cations,including alkali, alkaline earth and transitionmetals. Clay minerals have a sheet-like struc-ture in which the building blocks are eithertetrahedra or octahedra linked to each otherinto planar layers by sharing oxygen ionsbetween Si or Al ions of the adjacent tetrahedraor octahedra (Bailey, 1980; Fig. 1). The tetrahe-dra result from the close packing of four O ions,with the space between them occupied by a Si4+ ion or, to a lesser extent, an Al3+ ion. Theoctahedra result from the close packing of six

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Clay minerals in sandstones 5

anions that are dominantly oxygen but also caninclude some hydroxyl (OH) ions. The Si andAl ions mainly occupy the space between theoxygen octahedra and tetrahedra but othercations, such as iron, calcium, magnesium andpotassium, are required in the clay structure toensure charge balance. Tetrahedral and octahe-dral sheets are bound to each other in layersthat extend for tens to thousands of nanometres(nm) in the a and b crystallographic directions.The layers are stacked on top of each other in the c-axis direction. Figure 2 is a schematicphase diagram that incorporates most of theimportant clay minerals. This figure is subdi-vided between K–Al-rich, Al- (and ferric iron)rich and Mg–Fe2+-rich clay minerals.

Clay minerals can be classified based on thetypes of ions occupying the octahedral sites. Ifthe ions are trivalent (Al, Fe3+), the clay miner-als are said to be dioctahedral because only twoions are needed to provide six positive charges.

If the ions are divalent (Mg, Fe2+), they are said to be trioctahedral because three ions areneeded to provide six positive charges. Limitedsubstitution of trivalent ions in dioctahedralclay minerals and divalent ions in trioctahedralclay minerals is possible. Thus typically Mg-and Fe2+-rich clay minerals are trioctahedralwhereas Al- and Fe3+-rich clay minerals aredioctahedral. Interlayer cations are dominatedby potassium. Ammonia (NH4

+) can be presentin small quantities in the interlayer site in illite(Williams et al., 1992).

There are five dominant groups of clay minerals in sandstones: kaolin, illite, chlorite,smectite and mixed-layer varieties. A less com-mon clay mineral in sandstones is palygorskite.Polytypes, or polymorphs, of clay mineral have the same composition but different crystalstructures. The crystallographic differencestypically occur in response to different tem-perature conditions.

Fig. 1 Schematic diagram showingthe structures of the common clay minerals (a) kaolin, (b) illite (c) chlorite and (d) dioctahedralsmectite. The triangular motifrepresents tetrahedral layers. Thesolid grey bars represent octahedrallayers.

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Kaolin–serpentine series clay minerals,including berthierine

Kaolin–serpentine series clay minerals arecomprised of one tetrahedral layer linked toone octahedral layer with no interlayer cationsand are termed 1 : 1 layer structures connectedby O–H–O bonds. The chemical formula of kaolin is Al2Si2O5(OH)4, whereas the Mg end member serpentine has the formulaMg3Si2O5(OH)4. Serpentine can have Fe2+

substitution for Mg. Serpentine has not beenreported in sandstones whereas berthierine[Fe2

2+Al(Si,Al)2O5(OH)4] is a common clay mineral of the solid solution series between Fe-rich serpentinite and kaolin.

Kaolinite is the low temperature form,whereas dickite and nacrite are thought to bethe high temperature forms of kaolin. Kaolinitehas a unit cell of one octahedral–tetrahedral–octahedral package (unit cell of about 0.7 nm),whereas dickite has a unit cell made up of twoof these packages (thus with a unit cell of about1.4 nm) and nacrite has a unit cell made of six of these packages (unit cell of 4.3 nm).

Kaolinite tends to form pseudohexagonal platesthat commonly are stacked, in a book- or worm-like vermicular habit, whereas dickite tends toform small rhombic crystals (Fig. 3).

X-ray diffraction (XRD) analysis can be usedto discriminate between dickite and kaolinite.However, more accurate distinction betweenthe kaolin polymorphs can be made by deter-mining the position and relative intensity of theOH-stretching bands in the 3600–3700 cm−1

region of infra-red spectra (Ruiz-Cruz, 1996;Hassouta et al., 1999). Differential thermal ana-lysis also can be used to discriminate betweenkaolin polymorphs, because they have con-siderably different dehydration temperatures(e.g. Beaufort et al., 1998). Xia (1985) claimedthat a variation in the relative solubilities of the kaolin polymorphs towards hydrofluoricacid allows a quantitative analysis of theirabundances.

Illite and glauconite

Illite and glauconite are K-rich dioctahedralclay minerals comprised of one octahedral

Fig. 2 Phase diagram showing the main clay minerals in terms of divalent, trivalent and alkali element (monovalent)ion proportions. The predominant alkali earth is potassium although smectites can contain measurable sodium. The main divalent ions are magnesium and iron so that a very important range of solid solutions (Fe–Mg) cannot berepresented on this diagram. (a) Main minerals represented with areas of solid solution indicated. (b) Schematicrepresentation of typical compositions of a variety of different sandstones from clean arenites through to highly lithicsandstones. Note that thermodynamic equilibrium cannot be assumed in low temperature sediments and it ispossible to have many clay minerals together in one sample.

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layer sandwiched between two tetrahedral layers and so are termed 2 : 1 structures. O–K–Obonds connect two opposing tetrahedral layers.The interlayer K+ is required for charge balanceaccompanying the partial substitution of Al3+

for Si4+ in the tetrahedra and the substitution of divalent cations for Al3+ in the octahedra(Bailey, 1984). The O–K–O bonding is strongand prevents swelling behaviour in illite and

glauconite mica. Illite has octahedral sites dominated by Al, whereas glauconite has octa-hedral sites with abundant Fe3+.

The general chemical formula for illite isKyAl4(Si8−y,Aly)O20(OH)4 (Velde, 1985), wherey is typically significantly less than 2. Illite canoccur as flakes, filaments or hair-like crystals(Fig. 4). Illite occurs as polytypes that reflectdifferent ways in which layers are stacked.

Fig. 3 Scanning electron microscope (SEM) micrographs illustrating the progressive burial diagenetic transformationof kaolinite into dickite. Note that dickite and kaolinite may abut the quartz overgrowths, which often lead to amisinterpretation of the paragenetic relationship between these two minerals. (a) Disordered vermicular kaolinite. (b) Well-ordered kaolinite in discrete booklets. (c) Euhedral dickite crystals: D, dickite; Q, quartz cement. (d) Illitizedkaolinite (indicated by white arrow) with euhedral dickite preserved intact.

(a)

(c)

(b)

(d)

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Fig. 4 Scanning electron microscope (SEM)micrographs illustrating the progressive burialdiagenetic transformation of grain-coating, infiltrateddioctahedral smectite into illite, via mixed-layerillite–smectite (I/S). (a) Initial stages of I/S, curling claymineral flakes tangential to the detrital grain with 90%dioctahedral smectite, 10% illite. (b) Similar to (a) butwith 80% smectite: (c) 60% smectite, (d) 20% smectite,(e) < 10% smectite. Note the change from tangentialflakes to perpendicular fibres.

(c)

(a) (b)

(d)

(e)

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1M and 1Md polytypes are prevalent for lowtemperature (i.e. diagenetic) illites and also forglauconites. The one common at lower diagene-tic temperatures is the 1Md polytype, in whichthe crystallography repeats for every singlepackage of octahedral–tetrahedral–octahedralsheets (plus interlayer cations) and is dis-ordered. The 1M polytype tends to form athigher diagenetic temperatures, and is moreordered than the 1Md polytype. At high gradediagenetic to low grade metamorphic tempera-tures (> 200–250°C) illite is typically the 2Mpolytype, with a unit cell of about 2.0 nm comprised of a pair of tetrahedral–octahedral–tetrahedral layers.

Glauconite is a term that Odin & Matter(1981) have suggested be restricted to occur-rences of dark green, Fe-rich, mica-type clayminerals of marine origin and with K2O > 6%.Glauconite has the formula (K,Na,Ca)1.2–2.0(Al,Mg,Fe)4(Si7–7.6Al1–0.4O20)(OH)4 · nH2O. Theterm glaucony is thus recommended as a faciesterm that typically includes Fe-rich marine clay minerals that range in composition fromglauconitic smectite to glauconitic mica.

Smectite

Smectite is a group of 2 : 1 clay minerals withone octahedral layer sandwiched between twotetrahedral layers. Smectite has the general for-mula (0.5Ca,Na)0.7(Al,Mg,Fe)4(Si,Al)8O20(OH)4· nH2O. Trioctahedral smectite has octahedralsites dominated by divalent metals (Fe2+, Mg,Ca), whereas dioctahedral smectite has octahe-dral sites dominated by trivalent metals (Fe3+,Al). There is less binding of opposing tetra-hedral layers by K+ than in illite, with interlayerwater bound by weak van der Waal’s forces.Cations present between layers are exchange-able and reflect the chemistry of the aqueousmedium with which the smectite was last incontact. Interlayer cations are variably hydrated,resulting in the swelling characteristic of smectitic clay minerals. Smectites are definedby their tendency to swell when exposed to organic solvents, which can be absorbedbetween interlayers. Smectite usually occurs

as flakes curling up from an attachment zone on the detrital sand grain surface (Fig. 4).

Chlorite

Chlorite has a 2 : 1 : 1 structure comprised of anegatively charged 2 : 1 tetrahedral–octahedral–tetrahedral layered structure interlayered withan additional octahedral layer that is positivelycharged and comprised of cations and hydroxylions (e.g. brucite layers; Mg3(OH)6). A generalformula for chlorite is (Mg,Al,Fe)12[(Si,Al)8O20](OH)16. Solid solution is possible on all sites,leading to a very complex mineral group.Chlorite can exist as different polytype includ-ing the 1b polytype and the 2b polytypes with1.4 and 2.8 nm basal spacings, respectively. Fe-rich diagenetic chlorite (e.g. chamosite) istypically the 1b polytype, whereas the moreMg-rich diagenetic varieties (e.g. clinochlore)are typically the 2b polytype. It has been pro-posed that the type 1 polytype may be prevalentat lower diagenetic temperatures, with the type2 polytype forming as diagenetic temperaturesapproach low grade metamorphic conditions(Bailey & Brown, 1962). However, recent workhas concluded that there is no well-defined linkbetween temperature and chlorite polytypes(e.g. De Caritat et al., 1993; Walker, 1993).Chlorite occurs in a variety of morphologiesalthough classic chlorite occurs as a grain coat-ing boxwork, with the chlorite crystals attachedperpendicular to the grain surface (Fig. 5).

Mixed-layer clay minerals

Mixed-layer clay minerals result from the inter-stratification of different mineral layers in a single structure (Srodón, 1999). Most mixed-layer clay minerals contain smectite as a swell-ing component, and include illite–smectiteand chlorite–smectite (abbreviated to I/S andC/S respectively). During progressive burialdiagenesis I/S becomes more illite-rich and C/Sbecomes more chlorite-rich.

The stacking of layers in I/S usually is disordered (randomly interstratified) at the time of deposition and during eodiagenesis.

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Randomly interstratified mixed-layer clay minerals are labelled according to the types oflayers involved, with the most abundant layertype listed first (Reynolds, 1980). The term‘Reichweite’, denoted by R, is used to describeordering types. R = 0 describes totally randominterstratification of smectite and other clayminerals. R = 1, R = 2 and R = 3 describe progressively more ordered intercalations(Reynolds, 1980; Wilson, 1999). The degree ofdisorder in I/S decreases and the proportion ofillite increases in a semi-predictable mannerduring heating and burial diagenesis. Twotypes of ordered (R = 1, 1 : 1) mixed-layer clay

minerals have been identified and given dis-crete names: corrensite (chlorite–smectite) andallevardite (illite–smectite).

Palygorskite

Palygorskite is comprised of laterally con-tinuous two-dimensional trioctahedral sheets(dominated by Mg with OH ions) but does not have continuous SiO4 tetrahedral sheetsbetween the octahedral sheets. The tetrahedralpart of the structure occurs as ribbons, two silica tetrahedra wide, with infinite length,between which there are systematic tunnel-like

Fig. 5 Scanning electron microscope (SEM) micrographs illustrating the progressive burial diagenetic transformationof grain-coating, infiltrated trioctahedral smectite into chlorite via mixed-layer chlorite–smectite. (a) and (b) representa predominance of trioctahedral smectite present as a fine perpendicular honeycomb fabric. (c) Mixed chlorite–smectite showing a mixture of flakes and crystals. (d) Perfect Fe-chlorite occurring as perpendicular rosettes of well-formed crystals on detrital grain surfaces.

(a)

(c) (d)

(b)

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gaps. The tunnels have a width equivalent to twotetrahedra. Interlayer cations, required to keepcharge balance, are exchangeable with organicligands, resulting in an expandable structure.

DETRITAL CLAY MINERALS IN SANDSTONES

Detrital mud matrix

Laws of hydrodynamics control sedimenttransportation and deposition. The main vari-ables are the velocity of flow and the graindiameter, shape and density (Allen, 1997). Thefine grain size (< 0.004 mm) dictates that clayminerals are transported more easily than sandgrade material and that during waning flow,primary deposition of sand is followed by claydeposition only at very low flow rates (Blatt etal., 1980). Laws of hydrodynamics suggest thatdiscrete clay grains should not be deposited atthe same time as sand grade material, althoughincorporation of mud into turbidite, glacial anddebris-flow sediments may occur during deposi-tion if flow velocity decreases very rapidly.

Mud intraclasts

Intraclasts are autochthonous grains composedof aggregated clay minerals derived from theerosion of floodplain mud during fluvial chan-nel migration, and typically are deposited as alag in channel bases. In marine deposits theyform at ravinement surfaces marking trans-gression and the initiation of sea-level rise.Intraclasts are thus typically associated withfluvial and deltaic depositional environments.Shelf incision by valleys at sequence bound-aries also can result in mud intraclasts fromsediments that originally were deposited asmarine sediments.

Flocculated mud

Most suspended clay particles carried by riverspass out of estuaries and into the open marineenvironment. Estuaries are areally restricted

and tend not to accumulate suspended clayminerals, especially during times of sea-levelfall. However, coalescence and deposition ofsuspended clay minerals into sand-sized part-icles can occur by two mechanisms: (i) floc-culation of riverborne, suspended clay-sizedparticles upon mixing with seawater (vanOlphen, 1977), and (ii) biogenic agglomeration.As the clay aggregates are weak, they tend to bebroken up in high velocity currents (Krone,1978). Flocculated particles settle out in lowenergy environments or at times in the tidalcycle when there is little particle motion.

Clay minerals carried in colloidal suspensionin rivers have a negative surface charge that isneutralized by a layer of positively chargedcations, known as the Gouy layer. This resultsin a positively charged cation concentrationaround each clay particle relative to the riverwater. The repulsive power created betweenclay particles is greater than attractive van derWaal’s forces. Thus, colliding clay particlesrepel one another and tend not to coalesce in freshwater. However, when riverborne clayparticles encounter seawater, the concentrationof cations in the water increases. This leads to acollapse of the Gouy layer, reducing the repul-sion between clay grains. Thus, in brackishwater cohesive van der Waal’s forces becomestronger than the cation-layer-induced repul-sion, and cohesion, or flocculation, of clay particles can occur. Flocculation of clay isenhanced by very high velocity gradientsencountered in estuaries, which promotes agreat degree of collision between the sus-pended clay particles (Krone, 1978).

Flocculation and coalescence of riverborneclay particles is controlled not only by theincrease in aqueous salinity owing to mixingwith seawater but also by the type of clay mineral involved. The sequence in which clay minerals have a decreasing tendency to floc-culate is: kaolinite, illite and, finally, smectite(Edzwald et al., 1974; Krone, 1978). Thus theorder of deposition of clay aggregates will bekaolinite closest to the shore, followed by illiteand then smectite furthest from the shore. Dif-ferential flocculation becomes less important

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when the clay particles have organic (particu-larly humic acid) or iron oxyhydroxide coat-ings that modify the surface properties of theclay minerals (Gibbs, 1977).

In salt-wedge estuaries, where river flowdominates over tidal forces, little mixing occursbetween the seaward flowing river water in theupper layer, and the inward-flowing salt wedgebelow. This results in transfer of suspended,riverborne clay out to the sea to be deposited in deltas or offshore, in submarine canyons(Meade, 1972; Eisma et al., 1978).

Biogenic agglomeration occurs when suspen-sion feeders, such as bivalves, form faecal pellets.Deposited faecal pellets may be incorporatedsubsequently as grains in the sediment or dis-aggregated and resuspended (Pryor, 1975).

Clay-rich rock fragments (extraclasts)

Allochthonous clay-rich sand grains includesedimentary and low-grade metamorphic rock fragments. Additionally, high temperatureminerals and rocks can be converted into lowtemperature clay minerals (i.e. argillized) with-out complete physical disaggregation into com-ponent clay crystals. In these cases, erosion andtransport of such material will lead to lithicgrains, dominated by clay minerals, being co-deposited with other sand grains. In somecircumstances, detrital lithic fragments becomeclay-mineral rich only during subsequent dia-genesis (e.g. Worden et al., 1997, 2000) mak-ing it difficult to determine whether the claywas depositional (as clay-rich lithic fragments)or diagenetic (Whetton & Hawkins, 1970) in origin. Sediment source terrains dominated bymetapelite and metabasites will be especiallylikely to lead to clay-rich rock fragments.

Inherited clay coats

Clay coats on sand grains can be detrital in ori-gin, and occur most commonly within embayedsurfaces of sand particles (Pittman et al., 1992).In arid environments, clay coats form whenwind-blown clay material adheres to moistsand grains (Krinsley, 1998). Inherited clay

coats can be discriminated from diagenetic claycoats by: presence at point contacts betweendetrital sand grains, widely varying rim thick-ness, absence on diagenetic mineral surfacesand preferential occurrence in sediments re-sulting from lower energy sedimentary envir-onments (Wilson, 1992).

POST-DEPOSITIONALINCORPORATION OF DETRITALCLAY MINERALS IN SANDSTONES

Bioturbation

Bioturbation operates as a post-depositionalmechanism that mixes clay minerals from mud-rich layers with sand. Organisms can burrowinto sand (seeking food or shelter) enabling thephysical mixing of under- or overlying mud-rich layers with sand-rich layers. Faunal biotur-bation is prolific in shallow-marine sandstones,where the supply of nutrients is high and thereis plenty of light.

Infiltration of clay minerals

If water rich in suspended clay percolatesthrough the vadose zone of a sandy aquifer,then clay will be filtered out of the water ascoatings on sand grains (e.g. Moraes & De Ros,1990). It is likely to occur in unconfinedaquifers with substantial vadose zones that typify semi-arid climates and/or very thicksand sequences. Mud-rich surface waters areprevalent in floodplain and deltaic environ-ments (Dunn, 1992). Multiple episodes of clayinfiltration into coarse-grained, braided riversand deposits result in thick infiltrated graincoatings. Infiltrated clay coats are common insandy meandering river deposits, but tend to bethinner than in braided river deposits.

EOGENETIC CLAY MINERALS

On deposition, the primary sand comprises amixture of minerals that were formed under a

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wide range of conditions (e.g. temperature,pressure, oxidation state, water composition).Much freshly deposited sand contains unstablegrains that survived weathering, erosion andtransportation. Therefore, the detrital mineralassemblage may be inherently unstable andthus will tend to react with the ambient waterduring eodiagenesis.

The main eogenetic clay minerals are kao-linite, glauconite, berthierine, verdine, di- and trioctahedral smectite, I/S, C/S and Mg-clayminerals (palygorskite), which are formed by:(i) precipitation from pore waters, (ii) replace-ment of framework sand grains and (iii)replacement of precursor detrital or diageneticclay minerals. Illite and chlorite do not form in eogenetic environments and they are deposi-tional, rather than diagenetic, in origin wherefound in soils and sediments that have not suffered deep-burial diagenesis (Wilson, 1999).The formation of diagenetic clay minerals insands at near-surface conditions and duringshallow burial is controlled strongly by deposi-tional facies, detrital composition of the sand-stones and climatic conditions.

Smectite, mixed-layer clay minerals and palygorskite

Smectite and I/S form as grain-hugging flakes in sands under semi-arid climate. Mg-smectites(e.g. saponite, sepiolite) and palygorskite formas fibres and fibre bundles during near-surfaceeodiagenesis of lacustrine, fluvial and aeoliansediments, and, less commonly, coastal sabkha(Hover et al., 1999; Pozo & Casas, 1999), whichis subjected to strong evaporation conditions.Evaporitic conditions lead to the formation ofhypersaline pore waters enriched with Mg2+

and dissolved SiO2, but depleted in HCO3−,

SO42− and Cl−, through the precipitation of cal-

cite, aragonite, gypsum, anhydrite and halite.Trioctahedral smectite and palygorskite formeither by precipitation from hypersaline porewaters or through the transformation of a claymineral precursor or Mg–Si-rich gels (Mayayoet al., 1998). They can be associated closelywith pedogenic deposits developed on dolomitic

and basaltic bedrocks, which act as a local sourceof Mg2+ (Karakas & Kadir, 1998). Dioctahedralsmectite and I/S form in less evaporitic environ-ments compared with trioctahedral smectites.

The most typical Fe-rich smectite is nontron-ite, which commonly forms on the abyssalplain of deep ocean basins, and in the vicinityof mid-oceanic ridges. An elevated Si content,owing to the presence of biogenic silica, and Fe content, owing to the presence of Fe-oxyhydroxides, in conjunction with loworganic matter content in these sedimentsaccount for this Al-poor and Fe3+-rich smectite.Kohler et al. (1994) suggested that Fe oxidationin nontronite from submarine hydrothermalchimneys of the Galapagos Rift and MarianaTrough is bacterially mediated.

Sepiolite, palygorskite and atapulgite areclosely associated with Fe-oxides/oxyhydrox-ides, being typically formed by the replace-ment of the detrital ferroan silicates (biotite,pyroxene and amphibole) and volcanic rockfragments (Walker et al., 1978; Surdam & Boles,1979). The Mg-rich clay minerals are rare inancient sediments, as these are sensitive tochloritization during subsequent mesodiagene-sis (Stein et al., 1990).

Green clay minerals: berthierines glauconiteand verdine

Berthierine is an aluminous Fe2+-rich 1 : 1 claybelonging to the kaolinite–serpentine series ofminerals. Verdine represents a group of green-ish, metastable aluminosilicates with a widevariety of crystal structures and chemical com-positions. The most common minerals in thisgroup include: (i) phyllite C, which is a 1.4–1.5nm, Fe3+- and Mg-rich clay mineral (Odin,1990), and (ii) phyllite V (odinite), which is a0.72 nm, Fe3+- and Mg-rich clay mineral, with a 1 : 1 serpentine type structure intermediatebetween di- and trioctahedral structures(Bailey, 1988). Glauconite is an Fe3+-rich diocta-hedral clay, which upon primary formation hassmectite-like swelling behaviour but whichadopts mica-like (non-swelling) characteristicsupon ageing (Odin & Matter, 1981).

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Berthierine and verdine occur as small (< 5 µm), lath-shaped grain coatings and crys-tals, as coats (fringes or tangentially arranged),pellets, ooids and void-fillings, or form by thereplacement of detrital grains (e.g. silicates and carbonate bioclasts). They form during dia-genesis below the sediment–water interface in deltaic–estuarine deposits (Hornibrook &Longstaffe, 1996; De Hon et al., 1999), primarilyin tropical to subtropical seas (Odin, 1990;Kronen & Glenn, 2000; Thamban & Pur-nachandra, 2000). Berthierine, verdine andglauconite are often closely associated withauthigenic apatite (Morad & Al-Aasm, 1994;Purnachandra et al., 1995), suggesting growthin nutrient-rich coastal waters.

Berthierine authigenesis is favoured in vol-canogenic sediments deposited in estuarine–coastal-plain environments (Jeans et al., 2000).Verdine formation is favoured on shelves offriver mouths, at water depths of < 200 m, underconditions characterized by low sedimentationrates (Kronen & Glenn, 2000; Vaz, 2000). Owingto the domination of Fe3+ over Fe2+ in verdine,its formation is anticipated to be at depths ofcentimetres or decimetres below the seafloor,being favoured by suboxic conditions. Theseare mildly reducing conditions that correspondto depths below the seafloor where nitrate ormanganese reduction, or initial iron reductionoccur. Conversely, the predominance of ferrousiron in berthierine suggests that it formed understrongly reducing conditions (i.e. pore watersrelatively depleted in dissolved oxygen) com-pared with verdine, such as during iron-reducing suboxic conditions. Authigenesis ofboth verdine and berthierine occurs prior to the burial depths in sediments where bacterialsulphate reduction dominates, and where Fe2+ becomes incorporated preferentially insulphide minerals.

Upon ageing and shallow burial to a few hundred metres, phyllite C is transformed intoodinite (Purnachandra et al., 1993), which inturn is transformed into berthierine (Odin,1988). Odinite thus occurs only in Recent sediments, whereas berthierine occurs in anci-ent sedimentary rocks.

Berthierine and verdine typically form inancient and Recent marine-shelf settings.However, a few fresh- and brackish wateroccurrences of berthierine have been reported,such as in Wealden (Early Cretaceous) sedi-ments of southeast England (Taylor, 1990).However, even aeolian and fluvial deposits that become flooded by marine water during a transgression may be subjected to diageneticalterations that result in growth of berthierineand glauconite (Ketzer et al., this volume, pp. 43–61).

Berthierine occurs in estuarine and coastal-plain sediments whereas glauconite formsexclusively in open marine sediments (Odin &Matter, 1981), decimetres or metres below theseafloor. Berthierine forms in shallower watersand under more strongly reducing conditions,i.e. from pore waters richer in Fe2+, relative toglauconite.

Kaolin

Typically, eogenetic kaolinite has a vermicularand book-like habit (Figs 3 & 6; Ketzer et al.,this volume, pp. 361–382). Kaolinite formsunder humid climatic conditions in continen-tal sediments by the action of low-pH ground-waters on detrital aluminosilicate mineralssuch as feldspars, mica, rock fragments, mudintraclasts and heavy minerals (Emery et al.,1990). During forced regression and lowstand,large areas of marine sediment are subaeriallyexposed on the shelf, leading to an enlargementof the area of meteoric recharge. The basinwardmigration of the meteoric zone promotes theflushing of shallow-marine sediments and evendeep-water turbidites in some cases.

The amount and distribution pattern of kaolinite is influenced by the amount of un-stable detrital silicates, annual precipitation,hydraulic conductivity and rate of fluid flow in the sand body. Eogenetic grain dissolution is most prevalent in permeable sediments, such as channel sand deposits. Humid conditionsresult in the availability of greater amounts ofmeteoric waters, and hence promote eogenetickaolinite.

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Some of the earliest reactions to occur are thereplacement of albite and calcic plagioclase bykaolinite.

2NaAlSi3O8 + 2H2O + 2H+

albite

⇒ Al2Si2O5(OH)4 + 4SiO2 + 2Na+ (R1)kaolinite

CaAl2Si2O8 + H2O + 2H+ ⇒ Al2Si2O5(OH)4 + Ca2+

anorthite kaolinite(R2)

Calcic plagioclase and albite tend to be moresusceptible to kaolinitization than K-feldspar.It is possible that the protons are supplied bypartial dissociation of carbonic acid so thatreactions may be accompanied by carbonateprecipitation.

ROLE OF CLIMATE ANDDEPOSITIONAL ENVIRONMENTON EOGENETIC CLAY MINERALS

Depositional environment is a master controlon eodiagenesis because it controls the typeand amount of water present in sediment, waterinflux versus evaporation rate, temperature,exposure to atmospheric oxygen, plant-derivedCO2 and organic matter content. The role ofdepositional environment on clay mineral patterns in sandstones will be divided herebetween subaerial and marine systems. Con-tinental environments will be split betweenwarm and wet versus arid.

Warm and wet continental environments

Pore waters in humid, warm (subtropical totemperate) environments are dilute (less than afew hundred parts per million) and dominatedby Ca2+ and HCO3

− and slightly acidic (Fig. 7).Warm, wet, typically verdant, eogenetic envir-onments also have an abundance of organicmatter that undergoes bacterially mediateddecay (Berner, 1980). Fe-bearing minerals inthe sediment are readily reduced to aqueousFe2+ by redox processes, which typically is available for siderite growth because very lowconcentrations of SO4

2− leads to the absence ofFe-sulphides. The clay minerals in this envir-onment are typified by kaolinite because its formation requires low ionic concentrations inpore waters.

Arid continental environments

Dry, typically but not always hot, continentalenvironments (such as braid plains, playa margins and deserts) often have a low organic

Fig. 6 Schematic representation of the development of eogenetic kaolinite and the illitization of detritalsmectite. On deposition, the model arkosic sedimenthas a smectite grain coating. During eodiagenesis,feldspars are converted to kaolinite. Mesodiagenesis at temperatures > 110°C sees dickite growth fromkaolinite, quartz overgrowths on detrital sand grains,conversion of grain-coating smectite into illite andconcomitant growth of Fe dolomite.

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matter content, deep water tables and fully oxidized sediments (Fig. 8). In sediment belowtopographic depressions, brines often evolveunder strongly evaporitic conditions, such as inrift lakes. The pore waters are concentrated andtend to be dominated by Na+, Ca+, Mg2+ andHCO3

− and, to a lesser extent, SO42− (Eugster &

Hardie, 1978). Iron tends to be fully oxidized(ferric) and often coats minerals as a hydroxideor sesquioxide. Evaporation often exceedsmeteoric influx, leading to an upward flux ofgroundwater, evaporation and the consequentdevelopment of various smectite (e.g. montmo-rillonite, saponite) and Mg-rich clay minerals.

Marine environments

Marine environments are characterized byslightly alkaline waters (seawater pH is 8.3) andNa+–Cl− dominated water (with subordinateSO4

2−, HCO3−, Ca2+ and Mg2+), with a salinity of

about 35 000 ppm (Fig. 9).Bacterially catalysed processes characterize

marine eodiagenesis. The interaction of organicmatter and oxidizing inorganic solutes (e.g.SO4

2−) and minerals (e.g. Fe3+ minerals) causesrapid eogenetic alteration of shallow buriedsediments. The bacteria and a potent mix ofoxidizing and reducing material leads to an

Fig. 7 Clay eodiagenesis in warm wet eogenetic environments. Kaolinite grows in areas actively flushed by riverwater as the ion activities are never given the chance to increase, thus leaving the water in the kaolinite stability field.Areas more remote from the main channel axis are able to stagnate and thus increase in ion concentration, takinggroundwaters into the stability fields of smectite minerals. The relative absence of sulphur species minimizes growthof pyrite and so may allow available iron to enter clay minerals (mainly smectites). However, local availability oforganic matter leads to increased alkalinity (following decay and oxidation) and thus Fe-rich carbonate growth(Curtis et al., 1986). Thick soil horizons in abandoned channels will allow organic (humic) acids to build up,encouraging extreme feldspar weathering and leading to bleaching and growth of kaolinite deep in the soil profile.

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authigenic suite of minerals, including Fe sul-phides, various carbonate cements and a rangeof typically green sheet-silicates (‘glaucony’)including glauconite, berthierine and smectite(Stonecipher, 2000). Elevated concentrations ofiron, responsible for berthierine or glauconiteformation, are thought to be the result of enrich-ment of the detrital sediments in metabolizableorganic matter and/or brackish pore waters that have lower concentrations of dissolvedsulphate than marine pore waters. The loworganic-matter content in abyssal, deep-seasediments prevents the reduction of Fe-oxyhy-droxides derived primarily from hydrothermal

vents about mid-oceanic ridges, and hencefavours the formation of Fe3+ smectite (typic-ally nontronite).

SEQUENCE STRATIGRAPHY ANDCLAY MINERALS IN SANDSTONES

Relative sea-level changes and the rate of sedi-ment progradation influence strongly the posi-tion of the strand line, and hence the degree andpatterns of mixing between marine and contin-ental waters, and eodiagenesis of coastal andshallow-marine deposits (Morad et al., 2000).

Fig. 8 Clay eodiagenesis in arid eogenetic environments. Pedogenesis will lead to smectite growth in soils on sands inthe upper and middle reaches of hydrological basins. In the unsaturated zone of aquifers, clay infiltration will occur if percolating waters contain suspended clay minerals (Moraes & de Ros, 1990). Clay infiltration will occur whengroundwater evaporation exceeds the rate of percolation, leaving the suspended clay minerals coated on sand grains.Kaolinite may form from detrital feldspars in the aquifer if the groundwater is flowing and being recharged. In thelower reaches of the hydrological basin, where the Mg/Ca ratio has increased following calcrete formation, Mg-richclay minerals, as well as dolocrete, may form in the sediment.

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A schematic representation of the link betweenrelative sea-level change and clay eodiagenesisis depicted in Fig. 10.

Relative sea-level changes are controlled by a combination of eustatic changes, subsidence/uplift of basin floor and/or substantial varia-tions in the rate of sediment supply and resultin three major stratigraphic surfaces that sub-divide the depositional sequence into geneticpackages known as systems tracts (VanWagoner et al., 1988). These surfaces include:1 the transgressive surface (TS), which is thefirst significant marine-flooding surface acrossthe shelfathis surface forms the boundary

between the lowstand systems tract and over-lying transgressive systems tracts;2 the maximum flooding surface (MFS), whichcorresponds to the highest relative sea-levelreachedathis surface separates the transgres-sive systems tract from the overlying highstandsystems tract;3 the sequence boundary (SB), which forms as a response to relative sea-level fall. Asequence boundary is characterized by sub-aerial exposure and erosion associated withstream rejuvenation and truncation of theunderlying strata as well as a basinward shift infacies. The SB, marked by unconformities and

Fig. 9 Clay eodiagenesis in marine eogenetic environments. Marine sediments with a significant freshwater influxwill have limited sulphate, and thus sulphide, and limited growth of pyrite. Berthierine forms close to the sedimentsurface but requires isolation from oxidized water to prevent re-oxidation (Odin & Matter, 1981). As primary Fe3+ isreduced, the subsequent Fe2+ is free to form the clay berthierine in the methanogenic zone and then siderite in thedeeper decarboxylation zone (Pye et al., 1990). Fully marine sediments have abundant pyrite and thus minimalberthierine in the methanogenic and sulphate reduction zones. Glauconite forms on, or near to, the sediment surfacein condensed sequences that have time enough for biological alteration of detrital clay minerals (Pryor, 1975). In thedeeper decarboxylation zone Fe-rich carbonates can form, as in the case of brackish pore waters. Flocculation leads to distinct zonation of geochemically aggregated clay minerals, with kaolinite floccules forming closest to shore,smectite floccules in the ocean basin and with illite intermediate (Edzwald et al., 1974).

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early LST

late LST

TST

HST

high

low

R.S.L.

high

low

R.S.L.

high

low

R.S.L.

high

low

R.S.L.

basin-floor fan:clay mineral matrix+ diagenetic smectite+ diagenetic kaolinite

proximal lowstand wedge:berthierine

condensed section:glaucony

parasequence boundaries:glaucony

parasequence boundaries:odinite/berthierine

transgressive surface:greensand

HST delta front:berthierine

distal HST fluvial:thin infiltrated clay coats+ diagenetic kaolinite

proximal HST fluvial:thick infiltrated clay coat+ diagenetic smectite

early HST parasequence boundaries:glaucony/odinite

coastal plain/shoreface deposits

fluvial deposits offshore deposits

Fig. 10 Illustration of the location of clay minerals in sandstones in terms of the sequence stratigraphic organizationof sands. The systems tracts developed during a cycle of sea-level change is adapted from Posamentier & Vail (1988).HST, highstand systems tract; LST, lowstand systems tract; TST, transgressive systems tract.

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their correlative surfaces, bounds the deposi-tional sequence.

Each cycle of relative sea-level changesresults in lowstand, transgressive and thenhighstand systems tracts as defined by VanWagoner et al. (1988). It should be noted thatthe original sequence model of Van Wagoner et al. (1988) and Posamentier & Vail (1988) hasundergone substantial re-evaluation since itspublication, although discussion of its evolu-tion is beyond the scope of this paper. The dis-tribution of clay minerals in sandstones can beconstrained within these three types of systemstracts (for details see Ketzer et al., this volume,pp. 43–61), as follows (Fig. 10).

Lowstand systems tract (LST)

Deposition of a LST occurs as a response to afall and slow rise in the relative sea-level. TheLST can be divided into two: early (lowstandfan) and late (lowstand wedge and valley infill)deposits (Van Wagoner et al., 1988). The low-stand basin-floor fan is comprised of sedimentsthat bypassed the shelf through incised valleysand is deposited on the slope and in the basinin the form of submarine fans. The clay mineralsare predominantly detrital interstitial mineralsdeposited by turbidity currents and as sand-sized, mud- and glauconite intraclasts erodedby valley incision of highstand systems tractsediments exposed on the shelf. Burial andmechanical compaction of these ductile intra-clasts result in the formation of pseudomatrix.

As valley incision is superseded by valleyfilling, the quantity of coarse-grained sedimentsdelivered to the shelf decreases, resulting in the deposition of a lowstand wedge on theslope by a levee-channel complex with rhyth-mic turbidites, and later by delta progradation.These deposits are characterized by detritalclay minerals that are similar to those in thelowstand fan.

Clay minerals formed in sediments exposedon the shelf (i.e. valley-filling and exposedolder strata) are related to the prevailing climatic conditions. Kaolinite forms undersemi-humid to humid conditions. Clay infiltra-

tion occurs in semi-arid conditions. Stronglyarid, evaporitic climates result in the formationof Mg-rich clay minerals (e.g. palygorskite,saponite). The interfluve sediments will be sub-ject to pedogenesis. Kaolinite formation owingto meteoric water incursion commonly extendsto include the sandy facies of the precedinghighstand systems tract.

Transgressive systems tract (TST)

Deposition of a TST represents an abrupt land-ward shift of facies (i.e. transgression), whichoccurs as a response to a rapid rise in relativesea-level. A transgressive surface (TS) is devel-oped on top of the incised valleys and theirinterfluves, and also at the upper boundary of the proximal lowstand-wedge. The valleysfilled initially with fluvial deposits during lateLST, begin to be filled with estuarine and, sub-sequently, shallow-marine sediments. Hence,berthierine may form in the upper parts of theincised valley sediment fill.

The TS is characterized by elevated concen-trations of glauconite and verdine intraclaststhat locally may result in the formation ofgreensand deposits on the coastal plain envir-onments. A rise in relative sea-level is accom-panied by a decrease in sedimentation rate onthe shelf, because most of the coarse-grainedsediments are entrapped landward. Shelf sediments are thus dominantly fine-grained andshow a progressive upward increase in theamounts of authigenic glauconite and verdine,with maximum concentrations occurring alongthe MFS. The MFS is characterized by sedimentstarvation (i.e. condensed section) in the outerand middle shelf. In some cases, the formationof glauconite and verdine is favoured withinbioturbation structures. Within the inner shelf,berthierine is formed in estuarine and shallow-marine environments.

Highstand systems tract (HST)

Deposition of the HST occurs in response togradual slowing of relative sea-level rise andlater during the initial stages of sea-level fall.

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