Deformation of sediments and sedimentary rocks

Deformation of Sediments and Sedimentary Rocks Geological Society Special Publications Series Editor K. COE GEOLOG...

Author: Mervyn E. Jones | R. M. F. Preston | Geological Society of London

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Deformation of Sediments and Sedimentary Rocks

Geological Society Special Publications Series Editor

K. COE

GEOLOGICAL

SOCIETY

SPECIAL

PUBLICATION

Deformation of Sediments and Sedimentary Rocks EDITED

BY

M. E. J O N E S & R. M. F. P R E S T O N Department of Geological Sciences, University College London, London

1987

Published for The Geological Society by Blackwell Scientific Publications OXFORD LONDON EDINBURGH BOSTON PALO ALTO M E L B O U R N E

N O 29

Published by Blackwell Scientific Publications Osney Mead, Oxford OX2 0EL 8 John Street, London WC1N 2ES 23 Ainslie Place, Edinburgh EH3 6AJ 52 Beacon Street, Boston, Massachusetts 02108, USA 667 Lytton Avenue, Palo Alto, California 94301, USA 107 Barry Street, Carlton, Victoria 3053, Australia First published 1987 © 1987 The Geological Society. Authorization to photocopy items for internal or personal use, or the internal or personal use of specific clients, is granted by The Geological Society for libraries and other users registered with the Copyright Clearance Center (CCC) Transactional Reporting Service, provided that a base fee of $02.00 per copy is paid directly to CCC, 27 Congress Street, Salem, MA 01970, USA. 0305-8719/87 $02.00

DISTRIBUTORS USA and Canada Blackwell Scientific Publications Inc. PO Box 50009, Palo Alto California 94303 Australia Blackwell Scientific Publications (Australia) Pty Ltd 107 Barry Street Carlton, Victoria 3053 British Library Cataloguing in Publication Data Deformation of sediments and sedimentary rocks.--(Geological Society special publication; ISSN 0305-8719;no. 29) 1. Sedimentation and deposition 2. Rocks, Sedimentary 3. Rock deformation I. Jones, M.E. II. Preston, R. M. F. III. Series 551.3'04 QE571 ISBN 0-632-01733-3 Library of Congress Cataloguing-in-Publication Data Deformation of sediments and sedimentary rocks.-(Geological Society special publication; no. 29) Bibliography: p. Includes index. 1. Rock deformation. 2. Sediments (Geology) 3. Rocks, Sedimentary. I. Jones, M. E. (Mervyn E.) II. Preston, R. M. F. III. Geological Society of London. IV. Series. QE604.D44 1987 551.8 86-24496

Typeset, printed and bound in Great Britain by William Clowes Limited, Beccles and London

ISBN 0-632-01733-3

Contents Acknowledgments

JONES,M. E. & PRESTON,R. M. F. Introduction

vii 1

Part I: Theory and experimental OWEN, G. Deformation processes in unconsolidated sands

11

GRATIER, J. P. Pressure solution-deposition creep and associated tectonic differentiation in sedimentary rocks

25

MANDL, G. & HARKNESS,R. M. Hydrocarbon migration by hydraulic fracturing

39

CLAYTON, C. R. I. & MATTHEWS,M. C. Deformation, diagenesis and the mechanical behaviour of chalk

55

ALLISON,R. J. Non-destructive determination of Young's modulus and its relationship with compressive strength, porosity and density

63

MALTMAN, A. A laboratory technique for investigating the deformation microstructures of water-rich sediments

71

MALTMAN,A. Shear zones in argillaceous sediments--an experimental study

77

Part II: Processes UNDERHILL, J. R. & WOODCOCK,N. H. Faulting mechanisms in high-porosity sandstones;

91

New Red Sandstone, Arran, Scotland PETIT, J.-P. & LAVILLE,E. Morphology and microstructures of hydroplastic slickensides in sandstone

107

GUIRAUD, M. & SI~GURET,M. Soft-sediment microfaulting related to compaction within the fluvio-deltaic infill of the Soria strike-slip basin, northern Spain

123

LEEDER, M. Sediment deformation structures and the palaeotectonic analysis of sedimentary basins, with a case-study from the Carboniferous of northern England

137

LABAUME, P. Syn-diagenetic deformation of a turbiditic series related to submarine gravity nappe emplacement, Autapie Nappe, French Alps

147

SCHACK PEDERSEN, S. A. Comparative studies of gravity tectonics in Quaternary sediments and sedimentary rocks related to fold belts

165

FARRELL, S. G. & EATON, S. Slump strain in the Tertiary of Cyprus and the Spanish Pyrenees. Definition of palaeoslopes and models of soft-sediment deformation

181

CLIFFORD,P. M., RICE, M. C., PRYER,L. L. & FUETEN,F. Mass transfer in unmetamorphosed carbonates and during low-grade metamorphism of arenites

197

Part III: Descriptive PICKERING, K. T. Wet-sediment deformation in the Upper Ordovician Point Leamington Formation: an active thrust-imbricate system during sedimentation, Notre Dame Bay, northcentral Newfoundland

213

vi

Contents

BRODZIKOWSKI, K., GOTOWALA, L., KASZA, L. & VAN LOON, A. J. The Kleszcz6w Graben (central Poland): reconstruction of the deformational history and inventory of the resulting soft-sediment deformational structures

241

BRODZIKOWSKI,K., GOTOWALA,R., HALUSZCZAK,A., KRZYSZKOWSKI,D. & VAN LOON, A. J. Soft-sediment deformations from glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben (central Poland)

255

BRODZIKOWSKI,K., KRZYSZKOWSKI,D. & VAN LOON, A. J. Endogenic processes as a cause of penecontemporaneous soft-sediment deformations in the fluviolacustrine Czy~6w Series (Kleszcs6w Graben, central Poland)

269

BRODZIKOWSKI, K. & HALUSZCZAK,A. Flame structures and associated deformations in Quaternary glaciolacustrine and glaciodeltaic deposits: examples from central Poland

279

BRODZIKOWSKI, K., HALUSZCZAK,A., KRZYSZKOWSKI,D. & VAN LOON, A. J. Genesis and diagnostic value of large-scale gravity-induced penecontemporaneous deformation horizons in Quaternary sediments of the Kleszcz6w Graben (central Poland)

287

DAVENPORT, C. A. & RINGROSE, P. S. Deformation of Scottish Quaternary sediment sequences by strong earthquake motions

299

ALEXANDER,J. Syn-sedimentary and burial related deformation in the middle Jurassic nonmarine formations of the Yorkshire Basin

315

FITCHES, B. Aspects of veining in the Welsh Lower Palaeozoic Basin

325

INDEX

343

Acknowledgments It is impossible to thank, by name, all who contributed to the success of the conference and the production of this volume. Without financial backing from the Geological Society and the generous active support of Professor Mike Audley-Charles the whole project would have been still-born. Before and during the conference, the bulk of the day-to-day administration was handled by Dr Judith Rowbotham and a dedicated band of helpers. We know that everybody at the conference greatly appreciated their efforts. We wish to acknowledge those speakers who are not represented in this volume but who made valuable contributions on the day, and finally we thank the reviewers for their rapid and efficient scrutiny of the manuscripts.

vii

Contents Acknowledgments

JONES,M. E. & PRESTON,R. M. F. Introduction

vii 1

Part I: Theory and experimental OWEN, G. Deformation processes in unconsolidated sands

11

GRATIER, J. P. Pressure solution-deposition creep and associated tectonic differentiation in sedimentary rocks

25

MANDL, G. & HARKNESS,R. M. Hydrocarbon migration by hydraulic fracturing

39

CLAYTON, C. R. I. & MATTHEWS,M. C. Deformation, diagenesis and the mechanical behaviour of chalk

55

ALLISON,R. J. Non-destructive determination of Young's modulus and its relationship with compressive strength, porosity and density

63

MALTMAN, A. A laboratory technique for investigating the deformation microstructures of water-rich sediments

71

MALTMAN,A. Shear zones in argillaceous sediments--an experimental study

77

Part II: Processes UNDERHILL, J. R. & WOODCOCK,N. H. Faulting mechanisms in high-porosity sandstones;

91

New Red Sandstone, Arran, Scotland PETIT, J.-P. & LAVILLE,E. Morphology and microstructures of hydroplastic slickensides in sandstone

107

GUIRAUD, M. & SI~GURET,M. Soft-sediment microfaulting related to compaction within the fluvio-deltaic infill of the Soria strike-slip basin, northern Spain

123

LEEDER, M. Sediment deformation structures and the palaeotectonic analysis of sedimentary basins, with a case-study from the Carboniferous of northern England

137

LABAUME, P. Syn-diagenetic deformation of a turbiditic series related to submarine gravity nappe emplacement, Autapie Nappe, French Alps

147

SCHACK PEDERSEN, S. A. Comparative studies of gravity tectonics in Quaternary sediments and sedimentary rocks related to fold belts

165

FARRELL, S. G. & EATON, S. Slump strain in the Tertiary of Cyprus and the Spanish Pyrenees. Definition of palaeoslopes and models of soft-sediment deformation

181

CLIFFORD,P. M., RICE, M. C., PRYER,L. L. & FUETEN,F. Mass transfer in unmetamorphosed carbonates and during low-grade metamorphism of arenites

197

Part III: Descriptive PICKERING, K. T. Wet-sediment deformation in the Upper Ordovician Point Leamington Formation: an active thrust-imbricate system during sedimentation, Notre Dame Bay, northcentral Newfoundland

213

vi

Contents

BRODZIKOWSKI, K., GOTOWALA, L., KASZA, L. & VAN LOON, A. J. The Kleszcz6w Graben (central Poland): reconstruction of the deformational history and inventory of the resulting soft-sediment deformational structures

241

BRODZIKOWSKI,K., GOTOWALA,R., HALUSZCZAK,A., KRZYSZKOWSKI,D. & VAN LOON, A. J. Soft-sediment deformations from glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben (central Poland)

255

BRODZIKOWSKI,K., KRZYSZKOWSKI,D. & VAN LOON, A. J. Endogenic processes as a cause of penecontemporaneous soft-sediment deformations in the fluviolacustrine Czy~6w Series (Kleszcs6w Graben, central Poland)

269

BRODZIKOWSKI, K. & HALUSZCZAK,A. Flame structures and associated deformations in Quaternary glaciolacustrine and glaciodeltaic deposits: examples from central Poland

279

BRODZIKOWSKI, K., HALUSZCZAK,A., KRZYSZKOWSKI,D. & VAN LOON, A. J. Genesis and diagnostic value of large-scale gravity-induced penecontemporaneous deformation horizons in Quaternary sediments of the Kleszcz6w Graben (central Poland)

287

DAVENPORT, C. A. & RINGROSE, P. S. Deformation of Scottish Quaternary sediment sequences by strong earthquake motions

299

ALEXANDER,J. Syn-sedimentary and burial related deformation in the middle Jurassic nonmarine formations of the Yorkshire Basin

315

FITCHES, B. Aspects of veining in the Welsh Lower Palaeozoic Basin

325

INDEX

343

Introduction Mervyn E. Jones & R. M. F. Preston In recognition of the increasing interest in the subject of deformation of sediments and sedimentary rocks shown by Earth scientists in recent years, a major international conference with this theme was held at University College London in April 1985. This volume contains the texts of those contributions to the Conference that were submitted for publication. The collection of papers presented is not a complete record of the proceedings, as some contributors chose not to submit a manuscript. However, most important subject areas are represented, and the papers provide both a review of the present state of the art and pointers for future investigation. The articles have been grouped into three main divisions: experimental and theoretical, process orientated, and descriptive of particular areas or localities. Within those groupings there is no particular significance in the order of printing except that articles with aspects in common have been placed near to one another. Studies of naturally deformed sedimentary rocks repeatedly indicate that much of the observed deformation resulted from processes active before the rock was lithified. An understanding of the origins of these structures cannot be established using the principles of rock mechanics and crystal physics commonly employed in 'hard-rock' structural geology (Rutter 1976; White 1976), the principles of the engineering discipline of soil mechanics being more appropriate. Alternatively, the post-lithification deformation of sediments is a typical rock mechanics problem and may involve an understanding of elasticity (Jaeger & Cook 1969), fracturing (Price 1966; Barton 1976), crystal plasticity (Turcotte & Schubert 1982) and diffusion based deformation mechanisms (Rutter 1976,1983). The student of sediment deformation must therefore be conversant with all aspects of rock and soil deformation. The Conference, with its contributions by structural geologists, sedimentologists, geotechnical engineers and those conversant with specific aspects of rock and soil mechanics, provided a coverage of this very wide subject area. This is reflected in the contents of the volume. The bringing together of scientists and engineers representative of a number of different subject areas leads to problems of terminology

which result from the different usage of similar terms in different disciplines. For example, the term consolidation has a very specific meaning in engineering soil mechanics whereas in geology it is often used as an approximate synonym for compaction or lithification of a sediment. Conversely, the stress dependent reduction in volume termed compaction by geologists is generally referred to as compression by soils engineers. Compression has a very different meaning in structural geology. Because of these problems of terminology the definitions of a number of expressions that now appear in the geological literature, and which have conflicting or uncertain usage in engineering and geology, are summarised below. Additionally, a number of terms recently introduced to sedimentary and structural geology, from the engineering literature, are defined. The editors acknowledge the correspondence in which A. Maltman (pers. comm.) both proposed the need for a standard terminology and provided several of the definitions adopted below.

Cam-Clay theory. A unified theory for the behaviour of uncemented, weakly lithified sediments and soils which has been developed by the Cambridge soil mechanics group, and which is described in Atkinson & Bransby (1978). CamClay theory relates the volume of an uncemented clastic sediment to the effective stresses acting upon it, and to its previous stress history. The theory can be modified to describe the post-yield behaviour of porous sedimentary rocks. Cementation. The filling of the pore space in a sediment with mineral matter deposited from solution in the pore fluid. Compaction. The reduction in volume of a sediment as a result of increasing effective stress (compression of Atkinson & Bransby (1978) and other soil mechanics literature), (Fig. 1). Compression. The deformation resulting from the application of confining or constrictional normal stresses on a sediment. Such stresses are normally also referred to as 'compressive stresses' and are generally taken to be positive in the geological literature (note that this convention is not always followed in the engineering literature). The soil mechanics use of the term compression (Atkinson

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformationof Sediments and SedimentaryRocks, Geological Society Special Publication No. 29, pp. 1-8.

2

M . E . Jones & R. M. F. Preston

~

intersection of a stress path with the critical state line for a material is often taken to represent failure of the material in soil mechanics (Fig. 3). The critical state line and the related Roscoe and Hvorslev surfaces (Fig. 3) define a figure in effective stress/volume space that describes the possible volume/stress states for a particular unlithified clastic sediment. In practical terms, Fig. 3 is typically represented by projections of the critical state figure on to the mean stress/ volume and mean stress/deviator stress planes.

Subsequent burial urial

Uplift

FIG. 1. Schematic representation of the change in volume of a clastic sediment in response to changes in effective stress. Both normally consolidated and overconsolidated conditions are represented in the figure. & Bransby 1978) which is the same as compaction in geology, should not be introduced into the geological literature. Consolidation. The time dependent dissipation of an excess pore fluid pressure and the related decrease in volume of the sediment. The rate of consolidation is determined by the permeability of the sediment and the length of the drainage path (Terzaghi 1943; Taylor 1948; Lambe & Whitman 1979; Jones & Addis 1986a), (Fig. 2). Critical state line. A line in effective stress/volume space (Atkinson & Bransby 1978; Jones & Addis 1985, 1986c) at which constant volume shear deformation due to frictional sliding of grains takes over from compaction and consolidation due to de-watering under increasing load. The

Eq

Diffusive-mass-transfer. This term is used to describe the re-distribution of mineral matter within a rock as a result of diffusion. The diffusion pathways may be along grain boundaries (Coble Creep or Pressure Solution), or through the grains (Nabarro Herring Creep). In deforming rocks, ions diffuse from regions of high compressive normal stress and are precipitated in regions of low normal stress (Fig. 4). Diffusive-masstransfer has been widely studied in structural geology (Rutter 1976, 1983; McClay 1977; De Boer et al. 1977) but is a slow strain-rate process not normally recognized in geotechnical engineering. Effective stress. The law of effective stress was

q

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Roscoe surface e~ Criticalstate line \~ Impossible , ~ states

Normal \ consolidation line Impossible~'

r

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states

,, /

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v = I + e (Specific volume) Time

FIG. 2. Schematic representation of the change in volume of a clastic sediment caused by its consolidation following a change in total stress. The figure portrays the effect of two increments of loading, with the sediment achieving its equilibrium volume in each case.

FIG. 3. Schematic state boundary diagram in effective stress/volume space that forms the basis of the CamClay theory. The deformation of all normally consolidated clastic sediments occurs on the Roscoe surface. Overconsolidated sediments follow paths within the volume of the figure, or on the Hvorslev surface.

Introduction

~bl;rCrseerp ~

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FIG. 4. Diffusion pathways between two grains in a sediment. Generally, in the presence of an aqueous pore fluid, the pressure solution path is by far the most important. first stated by Terzaghi (1936) and can be expressed as follows: The stresses in any point of a section through a mass of soil can be computed from the total principal stresses al, 0"2, and 0-3 which act at this point. If the voids of the soil are filled with water under a stress u the total principal stresses consist of two parts. One part u acts in the water and in the solid in every direction with equal intensity. It is called the neutral stress (or the pore pressure). The balance 0 - t l = 0-1 - - U , 0 ' ' 2 = 0 - 2 - - U and 0 " 3 = 0"3 - - U represents an excess over the neutral stress u and it has its seat exclusively in the solid phase of the soil. This fraction of the total principal stress will be called the effective principal stress. All measurable effects of a change of stress, such as compression, distortion and a change of shearing resistance, are exclusively due to changes in the effective stresses (Atkinson & Bransby 1978). The law of effective stress is fundamentally important, changes in effective stress being responsible for changes in volume (both elastic and compactional) in porous sediments. Price (in Fyfe et al. 1978) has discussed refinements to the basic law, but these effects are of minor importance in the context of sediment deformation.

Excess pore fluid pressure. Any pore fluid pressure that exceeds the normal hydrostatic fluid pressure for the depth of burial of the sediment. Excess pore fluid pressures can be caused by increases in total stress (due, for example, to the accumulation of more sediment) and incomplete consolidation,

3

or as a result of dehydration reactions in the sediment during diagenesis, or as a result of the invasion of the pore space in the sediment with a fluid at higher pressure from another environment (e.g. hydrocarbon migration). Excess pore fluid pressure is synonymous with the term 'overpressure' which is widely used in petroleum geology and petroleum reservoir engineering (Fig. 5). In particulate materials which are not cemented, the changes in effective stress which are consequent upon changes in excess pore fluid pressure will be reflected by changes in the volume of the sediment. In such materials the excess pore fluid pressure will be supporting an increased pore volume (porosity or void ratio). Even in weakly cemented materials this effect is important, as evidenced by the substantial compaction of the Chalk hydrocarbon reservoirs of the Ekofisk Field in the Norwegian North Sea. This compaction has led to a significant settlement of the sea floor (Barton 1985; Jones 1986).

Lithification. General term to describe the conversion of a loose particulate sediment into a strong rock. The processes involved can include compaction, cementation, diffusive-mass-transfer, di-

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ure

FIG. 5. Increase in total stress and pore fluid pressure as a function of depth. The difference between the hydrostatic line and the pore pressure gives the magnitude of the excess pore pressure, and the difference between the pore pressure and total stress the magnitude of the effective stress at any depth. The shape of the pore pressure curve varies with the environment, the stratigraphy and the geological history of the formation.

4

M . E . Jones & R. M. F. Preston

agenetic phase change and recrystallization (Addis & Jones 1985). Overconsolidation. The state achieved by a particulate material which has been normally consolidated to one effective stress and then allowed to re-equilibrate to a lower effective stress (Fig. 2). Overpressure. Synonym of excess pore fluid pressure which is more commonly used in petroleum geology. Porefluid pressure. The hydrostatic stress exerted by the fluid in the pore space of a sediment against the mineral skeleton. Pore space. General term for the voids in a sediment or sedimentary rock which contain the pore fluid. Pore throat. The aperture between the voids which may be much smaller than the voids and which will often control the permeability of the sediment. Rock mechanics. The study of the mechanical behaviour of rocks. Rock mechanics concentrates on the investigation of the behaviour of rock materials that exhibit a measurable elastic response. Much of theoretical rock mechanics concentrates on the application of elasticity theory to rock engineering problems (Jaeger & Cook 1969). In 'hard-rock' structural geology high temperature and high pressure rock deformation experiments have been used to study the post-yield behaviour of lithified rocks and individual minerals (Turcotte & Schubert 1982). Shear deformation. The deformation resulting from the application of shear stresses. In lithified rocks, shear deformation is often localized to preexisting planes of weakness or discontinuities, whereas in particulate materials it may be either a pervasive or localized deformation depending on the water content of the sediment. In materials exhibiting Cam-Clay behaviour, the onset of shear deformation with no change in volume (either localized as at the base of a mudslide, landslip or in a fault plane, or penetrative as in a mudflow, slump or shale diapir), is taken to indicate that failure of the material has occurred. Soil mechanics. The study of the mechanical behaviour of unlithified particulate materials including clastic sediments. Soil mechanics has until recently been entirely an engineering discipline, and has developed sophisticated material testing and analytical procedures. The subject now has an enormous experimental data base (a result of the comparative simplicity of materials testing of soils compared with the difficulties of testing rock) and has established a strong theoret-

ical base (Terzaghi 1943; Taylor 1948; Atkinson & Bransby 1978; Lambe & Whitman 1979; Bishop & Henkel 1982). Recently, soil mechanics theory and testing procedures have been adopted by geomorphologists studying surface processes (Hutchinson 1970; Brunsden 1979; Allison pers. comm.) and by geologists and reservoir engineers interested in sediment behaviour (Jones 1978; Jones & Addis 1984, 1985, 1986a,b; Barton 1985). Stress path. The path followed by an element of a material in deviator stress/mean stress space during a deformation (Fig. 6). Stress paths can be determined for natural deformations, and modelled in laboratory experiments using special testing apparatus, or investigated numerically using appropriate finite element codes. Stress paths may be described in terms of total or effective stresses and are an important practical benefit of the Cam-Clay theory although they are by no means unique to that theory (Atkinson & Bransby 1978). Total stress. The total load carried by an element of sediment and its pore fluid in any given direction. Deformation of the element of sediment is generally controlled by changes in the way the total stress is shared between the mineral skeleton of the sediment and its pore fluid (see effective stress and excess pore fluid pressure) (Atkinson & Bransby 1978; Lambe & Whitman 1979). Void ratio. The ratio of the volume of voids to the volume of solids in a sediment. This is equivalent to porosity (the ratio of the volume of voids to total volume of the sediment), but is more commonly used in engineering. Uniaxial stress path. The stress path in which a cylindrical sample of sediment experiences an axial shortening whilst maintaining a constant diameter. This stress path is important in nature,

706050~'40j407

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FIG. 6. Stress paths recorded during three deformation experiments on samples of chalk.

Introduction being that followed by an element of sediment undergoing burial and vertical compaction without lateral deformation. This stress path is referred to as the Ko stress path in soil mechanics where the deformation is entirely compaction, but a similar deformation can be recorded for partially cemented porous rocks, in which the stress path to maintain no radial strain has three distinct sections with different slopes (elastic, elasto-plastic also referred to as yield or pore collapse, and plastic also referred to as compaction) (Fig. 7). Uniaxial stress ratio. The ratio of horizontal to vertical effective stress necessary to maintain the uniaxial stress path. For normally consolidated particulate materials this ratio is a constant (typically 0.6-0.8 for clays and 0.3 for sands). In overconsolidated particulate materials the ratio may exceed 1 and for some highly overconsolidated clays may reach 4 or 5 (Brooker & Ireland 1965). For partially cemented materials, the ratio is variable. In the Chalk, for example, the ratio is 0.35 below the yield point, increases to 1.1 at the yield point and then decreases to 0.64 once the yield stress is exceeded (Jones 1986). The uniaxial stress ratio provides a method of estimating the magnitude of the horizontal effective stress in bodies of sediment undergoing compaction during burial with no lateral deformation (Jones & Addis 1984), (Figs 8 & 9). Yield surface. The projection of the Mohr envelope (Jaeger & Cook 1969) for a material into three-dimensional stress space, or into stress path

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astic

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1~3 20 3'0 413 5'0 610 7fO 810 9'0 l(JO p'(aPa) FIG. 7. Stress path recorded in a carefully controlled uniaxial deformation experiment on a porous Chalk sample• This curve shows the different paths followed during elastic deformation, yield and compaction• During elastic deformation the path is controlled by Poisson's ratio and during compaction by the grain boundary friction. When the sample is yielding, the stress path is controlled by a combination of Poisson's ratio, frictional and fracturing processes•

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FIG. 8• Effective stress ratio plots for Ko compaction experiments on two shale samples. The break in the slopes is a function of overconsolidation during the samples' natural stress history. Only data recorded at higher stresses than this can be taken as being representative of their compaction and consolidation characteristics• space. For partially cemented porous materials, certain stress paths should intersect both the yield surface (elastic limit) and the critical state line (failure) (Jones & Addis 1986b) (Fig. 10). This list of definitions is, however, not complete. From the preceding discussion and from the contributions in this volume, it is immediately apparent that sediments exhibit two rather different types of deformational behaviour: that which predates lithification and that which postdates it. The pre-lithification deformations have often been referred to as soft-sediment deformations and dismissed by structural geologists. Maltman (1984) has already fully addressed the problem of terminology associated with this weak behaviour. We will not pursue this discussion

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6

M . E . Jones & R. M. F. Preston Decreasing initial porosity

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FI6.10. A schematic representation of the critical state diagram showing the effect of decreasing sample porosity on the position of the elastic yield surface in critical state space. further here, except to point out that the behaviour of unlithified sediments is controlled largely by effective stresses. Low effective stresses can occur in any geological environment where pore fluid pressures are large. Weak, substantially undercompacted sediments can thus occur at great depths. Examples are the Chalk Oilfields of the Ekofisk Group in the North Sea, and the mobilization of shale at substantial depths in deltas to form shale diapirs and related structures. Time since sedimentation, and depth of burial must therefore be excluded from any definition of 'soft-sediment' deformation. Perhaps in this same context, the definitions of sediment and sedimentary rock which are currently taken as synonyms by many geologists should be reviewed. It might be sensible, at least when referring to clastic deposits, to restrict the use of the term sediment for description of the unlithified material and sedimentary rock for its lithified equivalent. Soft-sediment deformation then becomes the deformation of sediments regardless of age or environment. Obviously the different behaviour of an unli-

thified sediment and its fully lithified equivalent represent two ends of a spectrum. Both are covered by the present theories of soil mechanics and rock mechanics, but the partially lithified (normally partially cemented) intermediate states that the material passes through during lithification exhibit intermediate mechanical properties. This behaviour is shown schematically in Fig. 10, and mechanical data for the Chalk which exhibits this type of behaviour are shown in Figs 11, 12 and 13. Understanding this intermediate behaviour is essential if deformation processes in sedimentary basins, and the engineering problems associated with production of hydrocarbons from high porosity/highly overpressured reservoirs, are to be fully understood. The contributions which follow this introduction consider, in different ways, the problem of deformation and the generation of structures in sediments. The authors of these papers have used laboratory, theoretical and field observation techniques to approach their individual problems, but repeatedly, discussion returns to the themes of compaction, lithification, pore fluid pressures and effective stress. The theory that links these phenomena, once the territory of the civil engineer, is now an essential part of sedimentology and structural geology, and finds applications in both academic and industrial projects. Recently, another group of earth scientists, the petroleum reservoir engineers, have found a need to integrate rock and soil mechanics theory in their investigations of reservoir produc-

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18

FIG. 11. Stress/strain curves for various Chalk samples showing the effect of sample porosity at the outset of the experiment.

Introduction

400 380 360 340320 300 280 260 240 -

• ko compression test chalk J • 30 MPa .

.

.

7

• • • o

70qAt1 60

.

P. limestone Chalk 1 Chalk 2 Chalk 3

I °|'o

50- •A~o e

~. 220-

~v200-~180b 160140 120100 80 60 40 20 0 0

~ZE40- ! ; • ~

•

20

• \ •

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I

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I

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'1

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FIG. 12. The relationship between the yield strength of Chalk and its porosity.

.10-

\

\

\

%0 • ••°¢t Oo

•

0%© 0

0

I

10

I

20

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0% (porosity) tivity. If this trend continues, geologists and geotechnical engineers will be asked to provide increasingly accurate information on the behaviour of sediments in the little-investigated medium pressure/medium temperature environment where hydrocarbon reservoirs are located. Both geologists and engineers need to develop an

FIG. 13. The relationship between Young's modulus and porosity for Chalks from a number of localities and for Portland Limestone. understanding of the ways in which sediments and sedimentary rocks behave in this environment if this need is to be met.

References ADDIS, M. A. & JONES, M. E. 1985. Volume changes during diagenesis. Mar. Petrol. Geol. 2, 241-46. ATKINSON,J. H. & BRANSBY,P. L. 1978. The Mechanics of Soils: an Introduction to Critical State Soil Mechanics. McGraw Hill, London.

BARTON, N. 1976. Shear strength of rock and rock joints. Int. J. Rock Mech. Min. Sci. & Geom. Abs, 13, 255-79. ,1985. Ekofisk reservoir subsidence: an investigation using continuum and discontinuum computer modelling. Unpub. Rep. Norwegian Petroleum Directorate. 241 pp. BISHOP,A. W. & HENKEL,D. J. 1982. The Measurement of Soil Properties in the Triaxial Test. Arnold, London.

BROOKER, E. W. & IRELAND, H. O. 1965. Earth pressures at rest related to stress history. Can. GeotechnicalJ. 11, 1-15. BRUNSDEN,D. 1979. In: EMBLETON,C. E. & THORNES, J. (eds) Processes in Geomorphology. DE BOER,R. B., NAGTEGAAL,P. J. C. & nuYvlS, E. M. 1977. Pressure solution experiments on quartz sand. Geochim. Cosmochim. Acta. 41,257-64. FYFE, W. S., PRICE, N. J. & THOMPSON,A. B. 1978. Fluids in the Earth's Crust. Elsevier, Amsterdam. HUTCHINSON, J. N. 1970. A coastal mudflow on the London Clay cliffs at Beltinge, North Kent. Geotech. 20, 412-38. JAEGER, J. C. & COOK, N. G. W. 1969. Fundamentals of R o c k Mechanics. Methuen, London.

8

M . E . Jones & R. M. F. Preston

JONES, M. E. 1978. Progress report on the deformation of deep oil well argillaceous sediments, submitted to Koninklijke Shell Exploratie en Produktie Laboratorium, Shell Research B.V., The Netherlands. ,1985. Deformation mechanisms in Chalk. Proc. Chalk Res. Symp. Stavanger, May, 1985. Paper 3, book 2 16 pp. ,1986. Analysis of the Ekofisk oilfield subsidence. Unpub. Rep. Norwegian Petroleum Directorate. 177 PP. & ADDIS, M. A. 1984. Volume change during sediment diagenesis and the development of growth faults. Mar. Petrol. Geol. 1, 118-22. & ADDIS, M. A. 1985. On changes in porosity and volume during burial of argillaceous sediments. Mar. Petrol. Geol. 2, 247-53. & ADDIS, M. A. 1986a. The mechanical behaviour of the Ekofisk overburden: additional experimental data. Unpub. Rep. Norwegian Petroleum Directorate. 40 pp. & ADDIS, M. A. 1986b. Mechanical behaviour of chalks and shales. Proc. Norwegian Soc. Chart. Eng., Norwegian Soc. Rock Mech. Application of rock mechanics to reservoir problems seminar, Stavanger, June 1986. & ADDIS, M. A. 1986c. The application of stress paths and critical state analysis to sediment deformation. J. struct. Geol. 8, 575-80. LAMBE,T. W. & WHITMAN, R. V. 1979. Soil Mechanics, SI version (2nd ed.). Wiley, New York. -

-

-

-

-

-

-

-

-

-

MCCLAY, K. R. 1977. Pressure solution and coble creep in rocks; a review. J. geol. Soc. London, 134, 5770. MALTMAN, A. 1984. On the term 'soft-sediment deformation'. J. struct. Geol. 6, 589-92. NICOLAS, A. & POIRIER, J. P. 1976. Crystal Plasticity and Solid State Flow in Metamorphic Rocks. Wiley, London. PRICE, N. J. 1966. Fault and Joint Development in Brittle and Semibrittle Rocks. Pergamon Press, New York. RUTTER, E. H. 1976. The kinetics of rock deformation by pressure solution. Philos. Trans. R. Soc. London, A283, 203-19. , 1983. Pressure solution in nature, theory and experiment. J. geol. Soc. London, 140, 725-40. TAYLOR, D. W. 1948. Fundamentals of Soil Mechanics. Wiley, New York. TERZAGHI, K. 1936. The shearing resistance of saturated soil and the angle between the planes of shear. Proc. 1st Int. Conf Soil Mech. & Foundn. Engng. 1, 54-6. Harvard, Mass. --, 1943. Theoretical Soil Mechanics. Wiley, New York. TURCOTTE, D. L. & SCHUBERT, G. 1982. Geodynamics Applications of Continuum Physics to Geological Problems. Wiley, New York. WHITE, S. 1976. The effects of strain on the microstructures, fabrics and deformation mechanisms in quartzites. Philos. Trans. R. Soc. London, A283, 69-86.

M. E. JONES • R. M. F. PRESTON, Department of Geological Sciences, University College London, Gower Street, London WC 1E 6BT.

Deformation processes in unconsolidated sands G. Owen S U M M A R Y: Deformation in unconsolidated sands requires the action of a deformation mechanism to reduce sediment strength and a driving force to induce deformation. Deformation mechanisms include liquefaction and fluidization and are reflected in the style of deformation and grain orientation fabrics. They are initiated by a trigger, including groundwater movements, wave action and seismic shaking. Driving forces include gravitational body force, unevenly distributed loads, unstable density gradients and shear forces, and are reflected in the geometry of deformation. These components are combined to produce a genetic classification of soft-sediment deformation processes and structures.

Sediments may become preserved in the geological record in a way which shows that they had been disturbed while still unconsolidated and close to, or at, the sediment surface (Fig. 1). Primary sedimentary structures are deformed into soft-sediment deformation structures, ranging from essentially in situ deformation achieved through relative grain displacements (e.g. overturned cross-bedding, load structures) to deformation involving bulk transport of an entire sediment mass (e.g. slumps, clastic intrusions). Soft-sediment deformation structures are significant in providing a fossilized record of processes which acted on unconsolidated sediments prior to, or soon after, burial. Many interpretations have been made of the origins of specific soft-sediment deformation structures, but little consensus exists about details

of deformation processes. Compare, for example, discussions of recumbently overturned crossbedding by McKee et al. (1962), Allen & Banks (1972) and Hendry & Stauffer (1975), or interpretations of convolute lamination by Kuenen (1953), Sanders (1960), Williams (1960), Nagtegaal (1963), Lowe (1975), Allen (1977) and Visher & Cunningham (1981). Furthermore, no comprehensive classification exists for soft-sediment deformation, either in terms of the morphology of structures, or the processes responsible for their formation. This paper provides a brief discussion of the physical conditions under which soft-sediment deformation can occur, leading to the erection of a genetic classification for soft-sediment deformation. The emphasis is on essentially in situ deformation in unconsolidated sands, as typified

FIG. 1. Intraformational recumbent fold in cross-bedded sandstone, Carboniferous Fell Sandstone, Bowden Doors, Northumberland (G R NZ070326). Foresets deformed into a parabolic recumbent fold, overturned in the down-current direction (right), can be traced laterally into undisturbed foresets. Bedding planes and sediment above and below are undisturbed. Deformation can be inferred to have taken place while the affected set was at the sediment surface in an unconsolidated state. From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 11-24.

II

G. Owen

I2

by overturned and otherwise deformed crossbedding (Fig. 1), convolute lamination and some load structures (see Owen 1985).

TABLE 1. Summary of deformation mechanisms A. Exceed yield strength B.

Deformation mechanisms The structures mentioned above involve the deformation of primary sedimentary structures by ductile processes, occurring through flow as either a plastic substance stressed beyond its yield strength or as a viscous fluid (Fig. 2; e.g. Gillott 1968). For flow to occur in a material which under normal circumstances behaves as a solid (e.g. sand), some process must operate which enables a normally solid-like material to exhibit essentially liquid-like behaviour. This process can be termed a deformation mechanism--that is, a mechanism whose operation enables deformation to take place.

stress

A

\

r

strain rate

stress /

f

strain rate FIG. 2. Conditions for ductile deformation. (A) Plastic substance stressed beyond its yield strength. (B) Viscous substance (fluid)--no yield strength.

C.

Reduce yield strength: 1. long time 2. high temperature 3. increased moisture content 4. increased pore fluid pressure Liquidize: 1. thixotropy 2. high sensitivity 3. liquefaction 4. fluidization

1| I plastic

t viscous

Several deformation mechanisms are relevant to geological materials (Table 1). A first group permits deformation in plastic materials by exceeding their yield strength. A stress may increase in magnitude so that it exceeds the normal yield strength of the material, causing brittle failure of elastic materials and ductile deformation of plastic substances. Alternatively, the yield strength of a material may be reduced to a level below the magnitude of stresses acting on it. Such a reduction in yield strength can be achieved, for example, through an increase in temperature, application of a stress over a long period of time, causing deformation by creep (e.g. McClay 1977), an increase in the moisture content of clay-rich sediments (Gillott 1968) or an increase in the pore fluid pressure, causing a reduction in the effective strength of the material (Hubbert & Rubey 1959, Terzaghi & Peck 1967). A second group of deformation mechanisms enables a normally solid-like substance of significant yield strength to behave temporarily in a liquid-like manner and deform as a viscous fluid, or as a plastic substance of negligible yield strength. This group is especially important in particulate materials such as unconsolidated sediments. A general term for such changes of state is liquidization (Allen 1977, 1982). Liquidized sediment is capable of deforming in response to stresses which would be too weak to induce deformation of the sediment in its normal, solid-like state. Four types of liquidization will be considered--thixotropy, sensitivity, liquefaction and fluidization. Thixotropic behaviour is a property of certain cohesive materials, including some clays, by which the material temporarily loses its strength when disturbed or shaken. This is due to a change in the packing of platy particles from a loose, flocculated structure to one in which they are aligned and dispersed in the pore fluid (Boswell 1961). Strength is gradually restored as the flocculated structure re-forms. Quick clays exhibit a high sensitivity, or ratio

Deformation processes & unconsolidated sands of strengths before and after disturbance. When disturbed, a loose, metastable packing of particles breaks down, the system becomes over-saturated and shows a dramatic loss of strength (Gillott 1968; Torrance 1983). In contrast to thixotropic systems, strength is only restored if moisture is lost, allowing a stable packing to form in equilibrium with new prevailing conditions. Cohesionless materials, such as dry or fully saturated sand, possess strength as a result of particle interlocking and friction at grain contacts. They are susceptible to liquidization through liquefaction and fluidization. Liquefaction is a loss of strength related to an increase in pore fluid pressure to a level which equals the overburden pressure (Fig. 3; Seed & Lee 1966; Lowe 1975, 1976b; Seed 1979; Allen 1982; Owen 1985). The overburden pressure is due to the weight of overlying grains, so on liquefaction this weight is transferred from particle contacts to the pore fluid. Particles become temporarily dispersed in the pore fluid, the strength of the system is reduced virtually to zero, and it behaves as a viscous fluid. Three types of liquefaction can be distinguished (Allen 1982; Owen 1985). Static liquefaction is induced by an increase in pore fluid pressure related to groundwater movements in a permeable layer beneath an impermeable or lowpermeability cap. Impulsive liquefaction is induced by a single large impulse which destroys a loose, metastable grain packing. Cyclic liquefaction is induced by a gradual build-up of pore fluid pressure to a value which equals the overburden pressure, in response to the repeated application of a stress whose magnitude is less than the shear strength of the sand (Seed & Lee 1966). Once a sediment becomes liquefied particles begin to settle through the pore fluid under their

own weight, restoring sediment strength. During this process a sedimentation front passes upwards from the base of the affected layer, separating solid-like sand with restored grain contacts beneath from a liquid-like liquefied dispersion above. The duration of the liquefied state at any level within the layer is controlled by the velocity of this sedimentation front, and depends on the thickness of the layer, the fall velocity of the grains and the fractional volume concentration of grains in the sedimenting dispersion and in the deposit (Lowe 1976b; Allen 1982). For a 1 m thick bed of quartz sand in water, sedimentation times for the whole layer are of the order of some tens of seconds to a few minutes. Such times are comparable with the duration of deformation in experiments involving the deformation of sands by shaking (Selley & Shearman 1962; Owen 1985). Liquefaction is a time-dependent loss of strength occurring within a closed system, with no net loss or gain of pore fluid. The fractional volume concentration of grains in the liquefied dispersion equals that in the metastable sediment immediately prior to liquefaction, and corresponds to the critical void ratio (Terzaghi & Peck 1967). Hence grain separations during liquefaction are small in comparison with grain diameters. Fluidization occurs in an open system of cohesionless granular material, requiring a continuous flow of fluid through the sediment (Fig. 4). It describes a condition in which the upward component of fluid drag equals or exceeds the downward-acting particle weight (Lowe 1975, 1976b; Allen 1982). The superficial fluid velocity is defined as the discharge through unit area of the bed and the minimum fluidization velocity is reached when the upward component of fluid drag exactly balances the particle weight. The excess f uid

a.

LOOSE DEPOSITIONAL PACKING SHEAR STRENGTH IMPARTED BY GRAIN-GRAIN CONTACTS.

b.

13

LIQUEFACTION

SHEAR STRENGTH (i.e. RESISTANCE TO DEFORMATION) OF SEDIMENT GREATLY REDUCED.

FIG. 3. Cartoon showing fabric changes during liquefaction.

c.

pore~

~.~a-mO~

RESEDIMENTATION TO T I G H T E R P A C K I N G .

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ooourl.

FIG. 4. Simple cartoon illustrating the process of fluidization in a cohesionless granular material. U is the superficial fluid velocity, Umfthe minimum fluidization velocity and U~, the particle fall velocity. The superficial fluid velocity increases from (A) to (D). effective overburden pressure then becomes zero and the system behaves as a viscous fluid. Lower values of superficial fluid velocity represent the condition of seepage, in which the overburden pressure, and hence the sediment strength, are partially compensated by fluid drag. If the superficial fluid velocity increases above the minimum fluidization velocity, grain separations increase and the apparent viscosity of the system decreases. Hence fluidization involves bed expansion. If the superficial fluid velocity exceeds the particle settling velocity, particles are elutriated, or transported by the fluid (Lowe 1975, 1976b; Allen 1982).

influence pore fluid pressures and superficial velocities include artesian groundwater movements and the escape of excess pore water from underlying sediment which is compacting or has liquefied. Triggers which provide an impulsive stress include rapid sediment deposition, slope failure in response to undercutting or oversteepening, breaking waves and flood surges. Triggers providing a cyclic stress include earthquakes, pressure fluctuations associated with waves, alternate freezing and thawing in permafrost layers and turbulent pressure fluctuations associated with flow separation.

Triggers

Deformation mechanisms in sands and their effects

Deformation mechanisms involving changes of state in particulate materials need to be initiated by some external agent, which can conveniently be termed a trigger. Triggers which directly

Stresses in the sedimentary environment may increase to magnitudes which exceed the normal shear strength of sand. These produce deformation by essentially solid-state intergranular shear

D e f o r m a t i o n p r o c e s s e s in u n c o n s o l i d a t e d s a n d s

along discrete shear planes, and a rotation of the long axes of elongate grains to become imbricated at roughly 45 ° to the shear direction (Hamilton et al. 1968; Oda & Konishi 1974). Examples of stresses which may induce deformation by such a mechanism include drag by glaciers or debris transported by strong currents (McKee et al. 1962; Brodzikowski & van Loon 1980; Chakrabarti 1981), in situ crystal growth (Wardlaw 1972) or organic activity such as animal burrowing or rootlet penetration. Under certain circumstances sand will behave as a cohesive substance which may deform by brittle failure, creep or plastic flow associated with a reduced yield strength due to increased moisture content or the compression of trapped air (de Boer 1979). These circumstances include sand which is damp, but not saturated, sand which has dried out after being wet, partially cemented sand, frozen sand, muddy sand, or sand grains with tacky surfaces (e.g. algal coatings). Under the action of a suitable trigger, watersaturated sand and coarse silt grade sediment will liquefy. Seismic shaking represents probably the most readily available trigger for regionally extensive liquefaction. Liquefaction effects are well documented following recent earthquakes (e.g. Seed 1968; Ambraseys & Sarma 1969; Youd & Hoose 1977; Youd 1978; Seed et al. 1981). Liquefaction is also recorded in response to groundwater seepage (e.g. Jeyapalan et al. 1983), ocean waves (e.g. Henken 1970; Seed & Rahman 1978; Nataraja & Gill 1983), breaking waves (Dalrymple 1979, 1980) and in experiments involving shaking (e.g. Kuenen 1958; Selley & Shearman 1962; Selley 1969; Finn et aI. 1970; Owen 1985). Several effects can be attributed to the liquefaction of sands. The sediment shear strength is greatly reduced, permitting ductile deformation in response to otherwise ineffective stresses. In consequence, irregularities on the surface of a liquefied layer will tend to become levelled. Sedimentation from the liquefied dispersion may result in a closer grain packing than in the initial state, and any such settling to a closer packing will be associated with the expulsion of excess pore water (water escape). This may cause fluidization in overlying sediment, including upper parts of the liquefied bed. The times available for deformation are of the order of tens of seconds or a few minutes. Lamination is likely to be preserved, although perhaps slightly blurred, on account of the relatively small grain separations involved (Fig. 5; Doe & Dott 1980). Finally, since grains may temporarily lose contact with one another, depositional grain fabrics may

15

o_eoo._o QQQ@O Q

®do3oSoSo FIG. 5. Deformation of a lamina produced by each grain being displaced upwards by two grain diameters relative to its neighbour on the left (after Doe & Dott 1980). be destroyed and new liquefaction-related fabrics formed, for example with particle long axes lying horizontally in the field of gravity (Owen 1985). Two circumstances in which sand exhibits plastic behaviour may be associated with liquefaction. Firstly, since any increase in pore fluid pressure causes a corresponding reduction in sediment strength, systems which are initially under stress (e.g. slopes) may deform before full liquefaction, when the pore fluid pressure equals the overburden pressure, is reached. Such failure under partial liquefaction can be considered as a type of cyclic mobility (Seed 1979). Secondly, as grain contacts are progressively re-established at the rising sedimentation front, sediment at a given level undergoes a phase of gradually increasing strength, behaving temporarily as a plastic substance of reduced yield strength compared with normal sand (Fig. 6; Owen 1985). Residual stresses remaining after deformation may temporarily exceed this yield strength, causing brittle failure or plastic flow, and producing minor structures such as flow folds and normal faults (Fig. 7; see Woodcock 1976 and Lindholm 1982 for structures which may have formed by such a mechanism). A grain fabric characteristic of liquefaction may become overprinted with one related to intergranular shear at this stage. Water-saturated sands and coarse silts can also be fluidized, provided that a flow of fluid through the sediment is available. Geological circumstances in which this might occur include groundwater movements, the escape of trapped gas, and the release of excess pore water from underlying sediment which has liquefied.

I6

G. Owen LIQUEFACTION STRENGTH sedimentation

ndisturbed

r e s t o r a t i o n of g r a i n / ~ ontacts;dilatationf if u n d e r g o i n g

I I "-~

-~TIM E

I

DEFORMATION MECHANISM

solid

viscous

fluid

[plastic

solid

I DEFORMAT!ON STYLE

none

flow

I f l o w or ]fracture

none

FIG. 6. Changes in strength and deformation mechanism associated with sedimentation from a liquefied state (after Owen 1985). the system and, if significant grain separations are involved, lamination may be totally destroyed. Elutriation may occur, producing structures such as disrupted laminae (type B pillars of Lowe 1975), clastic intrusions and sand volcanoes. In contrast to liquefaction, which affects sediment pervasively, fluidization is likely to occupy discrete zones or channels, and its effects are likely to be irregularly distributed within a layer. A discrete grain orientation fabric may be developed, perhaps with long axes at a high angle to bedding. In summary, several mechanisms may enable deformation to proceed in unconsolidated sands, but changes of state (liquidization) by liquefaction or fluidization are the most likely. Geologically feasible circumstances under which such liquidization will occur can be envisaged, and distinct effects are associated with each mechanism in addition to a drastic reduction in sediment strength.

Driving force systems FIG. 7. Normal faults formed during liquefactioninduced deformation in a shaking table experiment. The faults cut cross-bedding deformed by gravitational collapse, which is picked out by carborundum laminae within fine/medium-grained quartz sand. The faults are inferred to have formed during a late state of deformation, with the sediment behaving as a brittle solid (see Owen 1985). In addition to the loss of sediment strength, the effects of fluidization include several which contrast with those of liquefaction. The fluidized state can persist for as long as fluid is supplied to

Deformation will proceed if particles are elutriated during fluidization or if an applied stress exceeds the normal yield strength of a material. In the latter case, stresses are necessarily large for the deformation of unconsolidated sand and produce deformation of rigid blocks along discrete shear zones (McKee et al. 1962; Oda & Konishi 1974) or may initiate erosion or grain flow processes (Lowe 1976a). Under most circumstances, however, the action of a deformation mechanism alone will cause little or no deformation. Liquefaction of a homogeneous, isotropic sand layer with a level free surface and composed of spherical particles,

D e f o r m a t i o n p r o c e s s e s in u n c o n s o l i d a t e d s a n d s for example, may result in a small volume reduction if the particles settle to a closer packing, but will cause no other distortion or deformation of any layer or point markers. In general, a further condition is necessary for deformation to occur, namely a deviatoric stress must act while the deformation mechanism operates (cf. Lowe 1975). The stress may act continuously, but only while the deformation mechanism operates is it effective in causing deformation. For sands and sandstones, therefore, most soft-sediment deformation in the geological record is likely to have occurred in response to relatively weak stresses while the sediment was liquidized. The style of a deformation structure will depend on the initial sediment geometry and certain characteristics of the deformation mechanism, but is mainly determined by the orientations of the deforming stresses. The physical system responsible for these stresses can be called a driving force system. A relatively small number of driving force systems need be considered in the context of sedimentary environments. Each may relate to one or more geological agents or situations. These can account, either singly or in combinations, for most soft-sediment deformation structures occurring in sands and sandstones (Owen 1985). Firstly, a gravitational body force acts vertically downwards on all sloping surfaces. The strength of a solid enables slopes to exist, but a liquid, with no strength, attempts to attain a horizontal surface to restore gravitational equilibrium. Hence the liquidization of sediment resting on, or in possession of, a sloping surface, may induce deformation, with components of stress and displacement both parallel and perpendicular to the slope. A zone of basal decollement may confine deformation to a relatively thin, near-surface layer, producing various types of sediment gravity flows (Middleton & Hampton 1976; Lowe 1979, 1982). Alternatively, entire sediment masses with sloping surfaces, such as sand dunes or bars, may subside and spread under gravity upon becoming liquidized (Owen 1985). Such an origin can be proposed for some complex, large-scale types of deformed cross-bedding (e.g. Allen 1982). Secondly, an unevenly distributed confining load exists where a sediment surface possesses relief or where one sediment layer of variable thickness rests on another. Upon liquidization the lower layer loses its bearing capacity in much the same way as sand foundations fail when liquefied in earthquakes (e.g. Seed & Lee 1966; Youd 1978). The surface load sinks vertically into the substrate, but if the load itself becomes liquidized, it will also collapse and spread under

17

a gravitational body force, inducing lateral displacements as well. Soft-sediment deformation structures formed in this way include load-casted ripples (e.g. Dzulyfiski & Slaczka 1965; Anketell et al. 1970; Allen 1982), complex large-scale dune collapse structures (Horowitz 1982) and folds and faults related to laterally spreading sand bodies (Rettger 1935; Jones 1972). Thirdly, a gravitationally unstable density gradient exists where relatively dense sediment overlies relatively less dense sediment. Upon liquidization such a system is unstable, representing a Rayleigh-Taylor instability (Ramberg 1981 ; Allen 1982) and even an initially planar interface between the layers of contrasting density deforms into a series of anticlines and synclines, representing various types of load structure (e.g. Kelling & Walton 1957; Anketell et al. 1970; Allen 1982). The characteristics of the deformed structure are controlled by several variables, including viscosity contrast between the layers, the number of layers involved, whether or not the denser layers become pierced, and whether both layers, or just the lower, become liquidized (Anketell et al. 1970; Allen 1982; Owen 1985). Sources of density variation in unconsolidated sands include variations in material density (e.g. quartz and pumice sands), packing (e.g. sand on mud), grain size, degree of saturation and nature of the pore-filling medium (e.g. water and gas). An unstable density gradient may also be continuous within a single layer, due to inverse grading, or set up temporarily during sedimentation after liquefaction of a normally-graded layer. The latter may be a mechanism responsible for the formation of some convolute lamination (Allen 1977). Fourthly, shear stresses acting within, or at the surface of, the sediment may induce deformation. Tangential shear in the sedimentary environment is readily available from water currents, but other sources include the drag of debris flows, glaciers or debris carried in a current. Displacements related to tangential shear are parallel to bedding. Current shear across the surface of a liquefied sand bed can produce overturned cross-bedding comparable with that observed in the geological record (Fig. 8A; Owen 1985). Both components are essential to such deformation--liquefaction and tangential shear. Tangential shear across the surface of a non-liquefied bed will cause either erosion of grains, or deformation of a thin, nearsurface layer by intergranular shear, to produce a recumbent fold characterized by a sharp hinge zone, straight limbs, and subordinate shear dislocations (Fig. 8B; McKee et al. 1962; Hendry & Stauffer 1975).

I8 A

G. O w e n

.........~,~. . ~ 0 ~ ~. ......i.'~.... ~ ~...... ~ 'L~¸~ % '~~~ ' . . . . ~

FIG. 8. Intraformational recumbent folds in cross-bedding produced experimentally. (A) Fold produced by drag of water current over surface of sand liquefied on a shaking table (Owen 1985). This fold resembles that in Fig. 1. (B) Folds produced by drag of mass of sand carried across sediment surface by strong current (McKee et al. 1962). These folds are confined to the upper parts of sets, and are characterized by sharp hinge zones, straight limbs and subordinate shears, suggesting that the deformation mechanism was intergranular plastic shear. Grid in (A) is 5 cm square; scale in (B) in inches. Vertical shear is most likely to be associated with fluid drag during fluidization. Resulting vertical displacements may form pipe- or cusplike water-escape structures, often localized in extent, or irregularly distributed within a layer (Lowe 1975). Finally, other agents of physical, chemical or biological origins may drive soft-sediment deformation. These include the contraction of clays on desiccation, raindrop impacts, burrowing and crawling animals, plant roots, and stresses associated with the growth of crystals or concretions. Such driving forces will produce variablyorientated displacements, but evidence for their operation will typically be preserved in the sedimentary record.

Classification of soft-sediment deformation Attempts have been made to classify softsediment deformation using both morphological schemes relevant to the recognition of preserved structures (e.g. McKee et al. 1962; Pettijohn & Potter 1964; Allen 1977), and genetic schemes designed to clarify processes (e.g. Elliott 1965; Nagtegaal 1965; Allen 1982). Others have combined both types of approach (e.g. Lowe 1975; Brenchley & Newall 1977). Most classifications are of limited value, however. Genetic schemes are based on inferred, often poorly understood or contentious parameters, and typically provide no definite criteria by which the genetic parameters are to be related to characteristics of actual deformation structures. Hence they are difficult to apply to soft-sediment deformation structures preserved in the geological record. Morphological schemes are typically incomplete. They are applicable only to a limited range of structure

types or to a single stratigraphical occurrence and have little or no genetic or interpretive significance (see Dott 1963 and Brenchley & Newall 1977 for discussion). The erection of a workable classification for soft-sediment deformation is a desirable aim for several reasons. A genetic classification which can be applied to preserved structures would improve the understanding of their origins and enhance their diagnostic value in terms of sedimentary processes. It would identify structures which differ in morphology but not in process, and conversely identify those structures which differ in their process of formation but are presently grouped under a single name. In so doing it would provide unambiguous terminology and nomenclature, clearing up much confusion which exists, for example, where structures which clearly differ in morphology, and probably also in origin, have been given the same name (e.g. convolute lamination of Allen 1977 and Visher & Cunningham 1981), or where one type of structure has been described under a variety of different names (e.g. various names for ball-and-pillow structure summarized by Potter & Pettijohn 1977). Finally, it would elucidate genetic relationships between structures and identify families of structures with common origins. The discussion of the preceding sections has shown that two components--a deformation mechanism and a driving force system--are necessary for soft-sediment deformation to occur. These can be used as the parameters for a genetic classification of soft-sediment deformation, shown in Fig. 9, in which deformation mechanisms and driving force systems have been expanded in the light of the earlier discussion. This classification is comparable with those of Elliott (1965) and Allen (1982), but is more comprehensive and rigidly defined in its distinction and definition of mechanisms and forces.

pi.... d

o~

om

current drag

biological

chemical

physical

other vertical

tangential

/ /

//

/

/

/~/

/ / /

/ crystal growth

i

~

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/

//// // /

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/o

~

~//

,/~

///

///

/ /

~

increased pore-fluid pressure

YIELD

STRENGTH

/

slumps

ball-and-pillow

pseud0nodule s

fluidised

simple RF

dishstructure: cusps

ball-and-pillow

contorted HMB

pseudonodules

load casts

dish stroctllreY

slumps complexRF loadedripples clastic dykes convolute lamination?

slumps

liquefied

LIQUIDISE thixotropic sensitive

load casts

= heavy mineral bands.

~////

increased enhanced long time moisture temperature (creep) content

REDUCE

FIG. 9. E x p a n d e d classification for soft-sediment deformation. R F = recumbent fold, H M B

0

:

<~

m

l~yers

not pi....d

not pi....d

D~ ~

layers

within singlelayer a

coati.....

~DZ

!~

plastic

loaded / ripples

~

brittle

UNEQUALCONFININGLOAD

~

// ////slides /// slumps ~ ~

~

EXCEED YIELD STRENGTH

GRAVITATIONALBODYFORCE

DI~IVING FORCE SYSTEM

DEFORMATION MECHANISM

20

G. O w e n

The main problem with this scheme, in common with other genetic classifications and as evident from Fig. 9, is a difficulty in accurately and unambiguously locating real structures. This can be attributed to a lack of understanding of the origins of many structures and to structures which are presently identified under a single name having several possible origins. These problems highlight the need for controlled experimental investigations of driving forces and deformation mechanisms (Owen 1985). Some attempt can be made, however, to determine deformation mechanisms and driving forces responsible for deformation structures. Driving force systems can often be inferred with some confidence by determining the orientations of displacements and the sedimentological setting of deformed structures. Greater uncertainty lies in identifying deformation mechanisms--there is no accepted way to distinguish deformation in cohesive sediments caused by enhanced plasticity, thixotropy, sensitivity or an increase in stress, for example, and the processes of liquefaction and fluidization are commonly confused in the literature. However, the preceding section on deformation mechanisms outlined some of their effects, which can help in their distinction. In particular, deformation by liquefaction and fluidization can be confidently distinguished in many cases. One promising line of investigation into distinctions between deformation mechanisms lies in the study of microfabrics, combined with experimental deformation under controlled driving force conditions (Owen 1985; see also Hamilton et al. 1968; Yagishita & Morris 1979; Yagishita et al. 1981 ; Boulter 1983). Until deformation mechanisms can be more confidently identified, the classification in Fig. 9 may be substituted by a simplified form, shown in Fig. 10. In this scheme, brittle and ductile behaviour are distinguished and ductile behaviour is sub-divided according to whether it occurs in cohesive (plastic) materials--muds, and sands under certain circumstances---or in cohesionless materials--sands and gravels. Three types of deformation mechanism are distinguished for cohesionless materials--liquefaction, fluidization (both viscous) and intergranular shear. This division of deformation mechanisms, each associated with certain effects, replaces previous schemes in which plastic and viscous flow are distinguished, the criterion for viscous flow being that stratification is destroyed (Ferm & Huddle 1955; Dott 1963). This criterion is not valid, however, since simple experiments show that deformation of liquefied sediment, essentially as a viscous fluid, need not destroy lamination, on account of the small grain separations involved

(Figs 5 and 8). Moreover, the distinction between plastic and viscous behaviour--the existence of a finite yield strength--cannot be determined, given only that sediment has become deformed, as it need not affect the style of deformation. Hence the problem of whether liquefied sediment behaves rheologically as a viscous fluid or as a plastic substance of very low yield strength is not relevant to the classification of soft-sediment deformation. The division of deformation mechanisms in Figs 9 and 10 is also an improvement on Elliott's (1965) distinction of solid, quasi-solid, hydroplastic and quasi-liquid sediment behaviour, since it is more closely related to actual physical processes within sediments. The establishment of criteria by which deformation mechanisms and driving force systems can be determined from the characteristics of deformed sediments, points the way towards the erection of a complementary morphological classification of soft-sediment deformation structures, based on morphological criteria with known genetic significance (Owen 1985). A major advantage of the proposed genetic classification over existing schemes, then, is that it is based on parameters which can be recognized from the morphological characteristics of deformed structures, and therefore it has potential for relating deformation structures to deformation processes. Hence it provides a sound physical framework from which to study soft-sediment deformation, particularly when describing structures in the field. Furthermore, it is a versatile and flexible scheme which can be adapted to suit specific needs. The scheme proposed in Figs 9 and 10 is biased towards soft-sediment deformation in sand-dominated sequences, but the choice of deformation mechanisms could be restricted or expanded, a single driving force system could be isolated and split into a number of classes, or driving force systems could be replaced by stress systems or by specific geological agents. Finally, this classification is by no means intended as a complete statement of the genetic significance of soft-sediment deformation. One notable absence is that it does not specify geological agents which might have triggered deformation. A justification for this omission is that, at least in the case of liquefaction, once the deformation mechanism has been initiated--by whatever trigger--the characteristics and processes of deformation are controlled by the driving force system and the deformation mechanism, not by the trigger. Hence the identity of triggering agents should be sought not in deformation structures themselves, but more in their sedimentological contexts. In identifying defor-

R

I

F

g

~

~< ~

~

m

SM

pi....d

pierced

other drag

current drag

pierced

pierced

not

biological

chemical

physical

tangential

layer

multi-

laver

2-

not

within a single layer

UNEQUAL CONFINING LOAD

GRAVITATIONAL BODY FORCE

~

DEFORMATION

faults

sediment

soft-

BRITTLE FAILURE

-

bJoturbation

concretions crystal growth

desiccation cracks rain prints

EL (2c)

(2a)

clay diapirs

cryoturbation

CL

I

(3f]

i

simple RE

dish structure(3a) cusps and pillars

(i)

(i)

ball-and-pillow

heavy mineral contortions

FLUIDIZATION

sand volcanoes clastic dykes

(i)

pseudonodules

deformation .......

load casts

dish structure (3b)

CL

(I)

complex RF CL (2e)

slumping

LIQUEFACTION

- sands

disrupted lamination

cohesionless materials

coherent lamination

loaded ripples

(2b)

CL (2d)

CL

COHESIVE FLOW

cohesive materials muds

(McKee)

.

burrow collapse

RF

I

INTERGRANULAR SHEAR

shear dislocation

FIG. 10. Working classification for soft-sediment deformation, particularly suited to deformation in sands.

v

H

D

FORCE SYSTEM

~

discontinuities faulting in muds and sands

3.

1977.

E. Williams, Allen,

1956.

&

1982a.

Lowe & Lopiccolo, Allen,

3a 3b

1960.

1974.

1959; de Boer, 1979.

ten Heal,

1963. 1953; Sullwold,

Kuenen,

Nagtegaal,

Dish structure:-

2f

2e

2d

2c

2b

Sanders, 1960; Dzulynski Smith, 1963°

:-

Convolute lamination 2a

2.

horizontal displacements vertical displacements convolute lamination r e c u ~ e n t folds

Systems involving both mud and sand may deform by a combination of mechanisms.

= = = =

i.

H v CL RF

KEY:

r~

r~

$

G. Owen

22

mation mechanisms, however, the classification does contribute to solving the problem by clarifying whether a triggering agent would have been necessary at all in any particular case: if so, then its identity can be sought further. Several works exist in which soft-sediment deformation structures have been attributed to specific triggers. These include earthquakes in both Recent and ancient sediments (Rascoe 1975; Sims 1975; Weaver 1976; Okusa & Anma 1980; Hempton & Dewey 1983), water waves in Recent sediments (Dalrymple 1979; Okusa & Yoshimura 1981)and other, sometimes unspecified, aseismic triggers in Pleistocene and Recent deposits (van Loon & Wiggers 1976; Douglas 1982; Allen 1985).

Conclusions A deformation mechanism is a process which enables a normally solid-like substance to become deformed in a liquid-like (ductile) manner. Several characteristics of deformed sediments can be used to identify deformation mechanisms, including the lithology, whether deformation is pervasive or localized, preservation of lamination and style of deformation. Liquefaction and fluidization are potentially the most likely deformation mechanisms in unconsolidated sands, although under certain circumstances sand may behave as a brittle solid or as a plastic substance. Particle long-axis orientation fabrics may provide a useful tool for distinguishing deformation mechanisms in sands. Several deformation mechanisms are initiated by the action of a trigger or geological agent causing disturbance, although this is unlikely to

influence the deformation further. Seismic shaking represents perhaps the most readily available and regionally extensive trigger, but others include breaking waves, groundwater movements, rapid sediment deposition and pressure fluctuations associated with storm waves, flood surges or flow separation. Fluidization and an increase in stress to a level at which it exceeds sediment strength are deformation mechanisms which provide a force in themselves to drive deformation. In general, however, a driving force system is also necessary for deformation to occur. A relatively few driving force systems are likely to operate in the sedimentary environment. These are a gravitational body force, unevenly distributed confining load, a gravitationally unstable density gradient and vertical and tangential shear stress. The driving force system responsible for a particular structure can generally be determined by considering the orientations of displacements and the sedimentary setting. A genetic classification has been based on combinations of deformation mechanisms and driving force systems. It is presently of limited applicability to preserved soft-sediment deformation structures, due to a poor understanding of their origins. It is hoped that the classification will improve this understanding by providing a sound, physically-based framework for a rational approach to soft-sediment deformation structures in the geological record. ACKNOWLEDGEMENTS:This work represents part of the author's PhD study, carried out at the University of Reading under NERC studentship no. GT4/80/GS/ 122. The helpful supervision of Professor J. R. L. Allen is gratefully acknowledged.

References ALLEN, J. R. L. 1977. The possible mechanics of convolute lamination in graded sand beds. J. geol. Soc. London, 134, 19-31. - - , 1982. Sedimentary Structures ."their Character and Physical Basis. Elsevier, Amsterdam. ,1985. Wrinkle marks: an intertidal sedimentary structure due to aseismic soft-sediment loading. Sediment. Geol. 41, 75-95. -& BANKS, N. L. 1972. An interpretation and analysis of recumbent-folded deformed crossbedding. Sedimentology, 19, 257-83. AMBRASEYS,N. & SARMA,S. 1969. Liquefaction of soils induced by earthquakes. Bull. seismol. Soc. Am. 59, 651-64. ANKETELL, J. M., CEGLA, J. t~¢DZULYNSKI, S. 1970. On

the deformational structures in systems with

reversed density gradients. Rocz. Pol. Tow. geol. Krakowie, 40, 3-30.

BOSWELL, P. G. H. 1961. Muddy Sediments. Heifer, Cambridge. BOULTER,C. A. 1983. Compaction-sensitive accretionary lapilli: a means for recognizing soft-sedimentary deformation. J. geol. Soc. London, 140, 78994. BRENCHLEY,P. J. & NEWALL,G. 1977. The significance of contorted bedding in upper Ordovician sediments of the Oslo region, Norway. J. sediment. Petrol. 47, 819-33. BRODZIKOWSKI,K. & VANLOON,A. J. 1980. Sedimentary deformations in Saalian glaciolimnic deposits near Wtostbw (Zary area, western Poland). Geol. Mijnbouw, 59, 251-72.

Deformation processes in unconsolidated sands CHAKRABARTI, A. 1981. Kink-like structures and penecontemporaneous thrusting in the foreset laminae of mega-ripples. Indian J. Earth Sci. 8, 2934. DALRYMPLE, R. W. 1979. Wave-induced liquefaction: a modern example from the Bay of Fundy. Sedimentology, 26, 835-44. ,1980. Wave-induced liquefaction: an addendum. Sedimentology, 27, 461. DE BOER, P. L. 1979. Convolute lamination in modern sands of the estuary of the Oosterschelde, the Netherlands, formed as the result of entrapped air. Sedimentology, 26, 283-94. DOE, T. W. & DOTT, R. H. 1980. Genetic significance of deformed cross-bedding--with examples from the Navajo and Weber sandstones of Utah. J. sediment. Petrol. 50, 793-812. DOTT, R. H. 1963. Dynamics of subaqueous gravity depositional processes. Bull. Am. Assoc. Petrol. Geol. 47, 104-28. DOUGLAS, T. D. 1982. Periglacial involutions and the evidence for coversands in the English Midlands. Proc. Yorkshire geol. Soc. 44, 131-43. DZULYlqSKI, S. & SLACZKA, A. 1965. On ripple-load convolution. Bull. Acad. Pol. Sci., Ser. Sci. Geol. Geogr. 13, 135-9. - & SMITH, A. J. 1963. Convolute lamination, its origins, preservation, and directional significance. J. sediment. Petrol. 33, 616-27. ELLIOTT, R. E. 1965. A classification of subaqueous sedimentary structures based on rheological and kinematical parameters. Sedimentology, 5, 193209. FERM, J. C. & HUDDLE, J. W. 1955. Slumps and mud flows in rocks of Pennsylvanian age in the Appalachian Plateau. Bull. geol. Soc. Am. 66,1557. FINN, W. D. L., EMERY, J. J. & GUPTA, Y. P. 1970. A shaking table study of the liquefaction of saturated sands during earthquakes. Proc. 3rd European Symposium on Earthquake Engineering, Sofia, September 14-17, 1970, 253-62. GILLOTT, J. E. 1968. Clays in Engineering Geology. Elsevier, Amsterdam. HAMILTON, N., OWENS, W. H. & REES, A. I. 1968. Laboratory experiments on the production of grain orientation in shearing sand. J. Geol. Chicago, 76, 465-72. HEMPTON, M. R. & DEWEY, J. F. 1983. Earthquakeinduced deformational structures in young lacustrine sediments, East Anatolian Fault, southeast Turkey. Tectonophysics, 98, T7-T14. HENDRY, H. E. & STAUFFER,M. R. 1975. Penecontemporaneous recumbent folds in trough cross-bedding of Pleistocene sands in Saskatchewan, Canada. J. sediment. Petrol. 45, 932-43. HENKEL, n . L. 1970. The role of waves in causing submarine landslides. Geotechnique, 20, 75-80. HOROWITZ, n . H. 1982. Geometry and origin of largescale deformation structures in some ancient windblown sand deposits. Sedimentology, 29, 155-80. HUBBERT, M. K. & RUBEY, W. W. 1959. Role of fluid pressure in mechanics of overthrust faulting; 1. Mechanics of fluid filled porous solids and its

23

application to overthrust faulting. Bull. geol. Soc. Am. 70, 115-66. JEYAPALAN, J. K., DUNCAN, J. M. & SEED, H. B. 1983. Investigation of flow failures of tailings dams. Proc. Am. Soc. Civil Eng., J. geotech. Eng. 109, 17289. JONES, B. G. 1972. Deformation structures in siltstone resulting from the migration of an Upper Devonian aeolian dune. J. sediment. Petrol. 42, 935-40. KELLING, G. & WALTON, E. K. 1957. Load-cast structures: their relationship to upper-surface structures and their mode of formation. Geol. Mag. 94, 4 8 1 - 9 0 .

KUENEN, P. H. 1953. Graded bedding with observations on the Lower Palaeozoic rocks of Britain. Verh. K. ned. Akad. Wet. (I), 20, 1-47. , 1958. Experiments in geology. Trans. geol. Soc. Glasgow, 23, 1-28. LINDHOLM, R. C. 1982. Flat stratification: two ancient examples. J. sediment. Petrol. 52, 227-31. LOWE, D. R. 1975. Water escape structures in coarsegrained sediments. Sedimentology, 22, 157-204. , 1976a. Grainflow and grainflow deposits. J. sediment. Petrol. 46, 188-99. --, 1976b. Subaqueous liquefied and fluidized sediment flows and their deposits. Sedimentology, 23, 285-308. --, 1979. Sediment gravity flows: their classification and some problems of application to natural flows and deposits. Spec. Publ. Soc. econ. Paleontol. Mineral. Tulsa, 27, 75-82. --, 1982. Sediment gravity flows: II. Depositional models with special reference to the deposits of high-density turbidity currents. J. sediment. Petrol. 52, 279-98. - - & LOPICCOLO,R. D. 1974. The characteristics and origins of dish and pillar structures. J. sediment. Petrol. 44, 484-501. MCCLAY, K. R. 1977. Pressure solution and Coble creep in rocks and minerals: a review. J. geol. Soc. London, 134, 57-70. MCKEE, E. D., REYNOLDS,M. A. & BAKER, C. H. 1962. Experiments on intraformational recumbent folds in crossbedded sand. In: Short papers in geology, hydrology and topography. Prof. Pap. U.S. geol. Surv. 450-D, D155-60. MIDDLETON, G. V. & HAMPTON, M. A. 1976. Subaqueous sediment transport and deposition by sediment gravity flows. In: STANLEY, D. J. & SWIFT, D. J. P. (eds). Marine Sediment Transport and Environmental Management, pp. 197-218. Wiley, New York. NAGTEGAAL, P. J. C. 1963. Convolute lamination, metadepositional ruptures and slumping in an exposure near Pobla de Segur (Spain). Geol. Mijnbouw, 42, 363-74. ,1965. An approximation to the genetic classification of non-organic sedimentary structures. Geol. Mijnbouw, 44, 347-52. NATARAJA, M. S. & GILL, H. S. 1983. Ocean waveinduced liquefaction analysis. Proc. Am. Soc. Civil Eng., J. Geotech. Eng. Div. 109, 573-90. ODA, M. & KONISHI, J. 1974. Microscopic deformation

24

G. Owen

mechanism of granular material in simple shear. Soils and Foundations, 14, 25-38. OKUSA, S. & ANMA, S. 1980. Slope failures and tailings dam damage in the 1978 Izu-Ohshima-Kinkai earthquake. Eng. Geol. 16, 195-224. & YOSHIMURA,M. 1981. Possibility of submarine slope failure due to waves. (In Japanese). J. Ocean Studies Dept., Tokai Univ. 14, 227-34. OWEN, H. G. 1985. Mechanisms and controls of deformation in unconsolidated sands. an experimental approach. Unpublished PhD thesis, University of Reading. PETTIJOHN, F. J. & POTTER, P. E. 1964. Atlas and Glossary o f Primary Sedimentary Structures. Springer-Verlag, New York. POTTER, P. E. & PETTIJOHN, F. J. 1977. Paleocurrents and Basin Analysis (2nd ed). Springer-Verlag, New York. RAMBERG,H. 1981. Gravity, Deformation and the Earth's Crust (2nd ed). Academic Press, London. RASCOE, B. 1975. Tectonic origin of preconsolidation deformation in upper Pennsylvanian rocks near Bartlesville, Oklahoma. Bull. Am. Assoc. Petrol. Geol. 59, 1626-38. RETTGER, R. E. 1935. Experiments on soft-rock deformation. Bull. Am. Assoc. Petrol. Geol. 19, 271-92. SANDERS,J. E. 1960. Origin of convoluted laminae. Geol. Mag. 97, 409-21. SEED, H. B. 1968. Landslides during earthquakes due to soil liquefaction. Proc. Am. Soc. Civil Eng., J. Soil Mech. Fdns. Div. 94, 1053-122. , 1979. Soil liquefaction and cyclic mobility evaluation for level ground during earthquakes. Proc. Am. Soc. Civil Eng., J. Geotech. Eng. Div. 105, 201-55. , ARANGO, I., CHAN, C. K., GOMEZ-MASSO, A. & ASCOLI, R. G. 1981. Earthquake-induced liquefaction near Lake Amatitlan, Guatemala. Proc. Am. Soc. Civil Eng., J. Geotech. Eng. Div. 107, 501-18. & LEE, K. L. 1966. Liquefaction of saturated sands during cyclic loading. Proc. Am. Soc. Civil Eng., J. Soil Mech. Fdns. Div. 92 (SM6), 105-34. -& RAHMAN, M. S. 1978. Wave-induced pore pressure in relation to ocean floor stability of cohesionless soils. Marine Geotechnology, 3, 12350. SELLEY, R. C. 1969. Torridonian alluvium and quicksands. Scott. J. Geol. 5, 328-46.

-

-

-

-

& SHEARMAN, D. J. 1962. The experimental production of sedimentary structures in quicksands. Proc. geol. Soc. London, 1599, 101-2. SIMS, J. D. 1975. Determining earthquake recurrence intervals from deformational structures in young lacustrine sediments. Tectonophysics, 29, 141-52. SULLWOLD, H. H. 1959. Nomenclature of load deformation in turbidites. Bull. geol. Soc. Am. 70, 12478. TEN HAAF, E. 1956. The significance of convolute lamination. Geol. Mijnbouw, 18, 188-94. TERZAGHI, K. & PECK, R. B. 1967. Soil Mechanics in Engineering Practice (2nd ed). Wiley, New York. TORRANCE, J. K. 1983. Towards a general model of quick clay development. Sedimentology, 30, 54755. VAN LOON, A. J. & WIGGERS, A. J. 1976. Primary and secondary synsedimentary structures in the lagoonal Almere Member (Groningen Formation, Holocene, The Netherlands). Sediment. Geol. 16, 89-97. VISHER, G. S. & CUNNINGHAM, R. n . 1981. Convolute laminations--a theoretical analysis: example of a Pennsylvanian sandstone. Sediment. Geol. 28, 17588. WARDLAW, N. C. 1972. Sedimentary folds and associated structures in Cretaceous salt deposits of Sergipe, Brazil. J. sediment. Petrol. 42, 572-7. WEAVER, J. n . 1976. Seismically-induced load structures in the basal Coal Measures, South Wales. Geol. Mag. 113, 535-43. WILLIAMS, E. 1960. Intra-stratal flow and convolute folding. Geol. Mag. 97, 208-14. WOODCOCK, N. H. 1976. Structural style in slump sheets: Ludlow Series, Powys, Wales. J. geol. Soc. London, 132, 399-415. YAGISHITA, K. & MORRIS, R. C. 1979. Microfabrics of a recumbent fold in cross-bedded sandstones. Geol. Mag. ll6, 105-16. , WESTGATE, J. A. & PEARCE, G. W. 1981. Remanent magnetization in penecontemporaneous structures of the Pleistocene Scarborough Formation, Ontario, Canada. J. geol. Soc. London, 138, 549-57. YOUD, T. L. 1978. Major cause of earthquake damage is ground failure. Civil Engineering, April, 47-51. & HOOSE, S. N. 1977. Liquefaction susceptibility and geologic setting. Proc. 6th World Conf. Earthquake Engrg., New DelhL India, 6, 37~42.

-

-

-

-

G. OWEN, Department of Geology, University College Swansea, Singleton Park, Swansea SA2 8PP.

Pressure solution-deposition creep and associated tectonic differentiation in sedimentary rocks J. P. Gratier S U M M A R Y. Several models of pressure solution-deposition have been established by using various hypotheses: on the rate-limiting process of the deformation (kinetics of the solid/fluid reaction or rate of mass transfer); on the driving force for mass transfer (difference in normal stress or elastic, plastic or surface energies) and on the mechanism of mass transfer (diffusion or infiltration). These creep relations have been tested experimentally and in nature. In the last case the observations show that the pressure solution-deposition process is always associated with a chemical differentiation of the rocks when these rocks are initially composed of several minerals with various mobilities under stress. The aim of this paper is to discuss the development of this chemical (and mechanical) differentiation. The first part deals with the initiation and the development of the zone of dissolution (relation between stress and dissolution, effect of the initial heterogeneities, etc.) and of the zone of crystallization (effects of the nature of the rocks, processes of deposition, etc.). The second part deals with the development of chemical differentiation during the progressive deformation. The geometry and the equilibrium composition of the differentiated layer are fixed by various factors such as: the size of the dominant heterogeneities (initial or tectonically induced), the possibilities of mass transfer, the state of stress, and the nature of the solid and that of its solution. The wavelength of some stylolites could also be imposed by these various factors. The deformation of rocks by pressure solutiondeposition is one of the most common mechanisms of deformation in the upper crust from diagenetic to low-grade metamorphic environments (20-400°C) in the presence of an intergranular fluid phase which is a solvent of the solid. After Sorby (1863) and Gibbs (1877) (and considering a compressive stress as positive), we know that some minerals such as quartz, or calcite, are frequently dissolved in zones of maximum stress (with the development of pitted pebbles, stylolites, solution-cleavage seams), these minerals being redeposited in zones of minimum stress (in veins or pores of the rocks). This mass transfer leads to a change of shape (or a change of density) of the rocks. This deformation may be analysed as a creep mechanism (or a densification mechanism) by establishing the relation between the strain-rate (or the densification-rate) and the various parameters of the deformation. Several models of pressure solutiondeposition have been established by using the following assumptions: The deformation rate is dependent on the kinetics of the successive processes: the kinetics of the solid/fluid reaction (dissolution or crystallization) and the rate of mass transfer (by diffusion or by infiltration). If one of these processes is much slower than the others the rate of the deformation is imposed by the slowest process. This leads to the establishment of various models

of solution-deposition. On the other hand, the driving force for the mass transfer is usually linked to the difference in normal stress between the solid/fluid interface with dissolution and the solid/fluid interface with crystallization, but may be imposed by other forces such as the difference in elastic or plastic strain energy (Paterson 1973), or the difference in surface energy (linked to the difference in the curvature of the surface, Kingery et aL 1976), as in the model R, Fig. 1A. Considering only the difference in normal stress as the driving force (for simplification), the creep relation for the change of shape of a cube of a solid, by pressure solution-deposition in a closed system, is always of the following form: e" = f ( k , c, Aan, w/d", T),

where e'= strain rate, k = coefficient of transfer, c = solubility of solid in solution, Aan = difference in normal stress, w/d ~= geometric factors depending on the path of mass transfer, T = temperature. More precisely: k is the coefficient of transfer, depending on the limiting process (see above) and on the path of the transfer (see Fig. 1). After Gratier (1984, 1986) k is the kinetics of the solid/ fluid reaction (model R), the permeability coefficient (model I with mobile fluid), the diffusion coefficient (model D or D' with fixed fluids). In this last case the diffusion can occur either around a continuous solid, along grain boundaries, with trapped fluid, able to support shearing stress

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 25-38.

25

26

J. P. Gratier

initial

V•'3

deposition ~,~

• P[ 1

I deformed state

--~ (A)

@ (B)

(C)

(D)

FIG. 1. Various types of mass transfer in pressure solution-depositionprocesses (in closed system). (A) Diffusion through free fluid in a small cavity. (B) Diffusion around a continuous solid (along grain boundaries). (C) Diffusion through a solid aggregate (along high speed diffusion paths). (D) Infiltration through porous aggregates. Depending on the limiting process (kinetics of the reaction or rate of mass transfer) various creep relations have been established and tested experimentally and in nature for these different cases (models R, D, D', I). (model D), Rutter (1976); or through an aggregate, along high speed paths of mass transfer (model D'). These four types of models were tested experimentally and in nature : model R for the experimental change of shape of small cavities with diffusion through a free fluid phase, Fig. I(A) (Gratier & Jenatton 1984); models D and D' for the experimental and natural deformation of compacted grains, with diffusion along grain boundaries saturated with a trapped fluid phase (Fig. 1B, C), Grafter (1984, 1986); model ] for natural deformation with long distance of transfer (Fig. 1D). The mean distance of mass transfer is higher with infiltration (10-102 m) than with diffusion (10-4-10 - 1 m). These various coefficients of transfer k' (reaction), D (diffusion), K (permeability) are also dependent on the temperature (T), but in different ways. For example the activation energy varies from 90 to 10 kJ mol-1 K-1 between k' and D (through free fluid) (Barns et al. 1976; Robinson & Stokes 1959). c is the solubility of the solid in its solution. The effect of this parameter was shown in natural deformations by comparing the relative mobility of quartz and calcite with change in temperature, since these minerals have respectively a normal and retrograde solubility. With fluid inclusion studies for the determination of pressure and temperature, (see below), we found that at about 275°C, 145 MPa quartz is more mobile than calcite, but it is the reverse at about 145°C, 80

MPa (Gratier 1984). This effect also appeared in experimental deformations when using various fluids in the deformation of quartz or calcite by pressure solution-deposition. In this case, significant values of the change of shape of small cavities, or grains, cannot be obtained without using very powerful solvents around the solid (NaOH or NH4CI solutions, respectively, for quartz or calcite) (Grafter & Jenatton 1984; Gratier 1986). w/d" are geometric factors: w is the width of the path of transfer (for example the thickness of the grain boundaries), d is the mean distance of transfer (for example the length of a cube with mass transfer in a closed system). Depending on the model, w can be equal to d (model D') and the value of the power of a~ can vary from 1 (models R, I) to 2 (model D') or 3 (model D), Gratier (1984, 1986). Aa, is the difference in normal stress between the zone of dissolution and the zone of crystallization. Depending on the model, this can be the difference in normal stress along the same solid/ fluid interfaces (for example with mass transfer around a continuous solid, Fig. 1B), or the difference in fluid pressure (or in mean normal stress on the solid) between two zones of a porous aggregate (models D' or I). For all these models the solid is supposed to be composed of only one soluble species. But of course this is not generally the case in natural deformations. In such cases the problem is more

Pressure solution-deposition in sedimentary rocks complicated. W e shall discuss successively the initiation a n d d e v e l o p m e n t of dissolution and deposition zones in natural rocks, and a process almost always associated with the d e v e l o p m e n t of such zones w h i c h is the tectonic differentiation of the rocks (Soula & D e b a t 1976).

Initiation and development of the zones of dissolution W h e n a rock filled with a fluid phase, and composed of several minerals of various solubili-

z7

ties in this fluid, is d e f o r m e d by pressure solutiondeposition, a segregation of the minerals always appears. T h e relatively insoluble species are passively c o n c e n t r a t e d in the zone of depletion of the soluble species. C o m p a r a t i v e c h e m i c a l analyses b e t w e e n the zone of dissolution and the initial rock allow us to calculate the value of the volume decrease in the solution zone (Gratier 1983). For example, the distribution of various elements n e a r such a dissolution zone, (obtained by successive profiles with a microprobe), are given in Fig. 2(A). A solution cleavage seam is m a r k e d by the zone of high content of elements

V-3-T-

12

Ca

13 14

AI

8

(c)

(B)

stratification 0

I mm

(A)

I, l

li!

(iil

llii)

(D) FIG. 2. (A) Chemical distribution (obtained by successive profiles with microprobe) of various elements in a slaty matrix around a rigid indentor submitted to a maximum compressive stress value (al) horizontal and parallel to the initial layering. This natural deformation leads to a tectonic differentiation of the matrix: A1 (in illite or chlorite) is passively concentrated in the zones of dissolution. Ca (in calcite) is removed from these zones (and not redeposited in the pressure shadow), Si (in quartz and layered silicates) has an intermediate behaviour, but the deposition of quartz in the pressure shadow sector clearly appears. (B) Theoretical distribution of mean stress values in a viscous matrix around a rigid object (after Stromgard 1973). (C) Theoretical distribution of mean stress values in a viscous matrix around a soft inclusion (after Cosgrove 1976). In the two cases the zones of maximum compressive stress are black-coloured. (D) Schematical development of a zone of dissolution near a rigid object. (i) Initiation in the zone of maximum compressive values. (ii) Development of a soft zone by the passive concentration of insoluble minerals (see B): layered silicates (black). (iii) Propagation of zone of dissolution in its own plane by preferential removal of matter in sectors with maximum compressive stress around the soft zone (see C).

28

J. P. Gratier

such as A1, Fe, K (belonging to insoluble minerals such as illite, chlorite, etc.) associated with a depletion in content of other elements such as Ca, Si (belonging to soluble minerals such as quartz or calcite). The rock concerned here was naturally deformed with a maximum compressive stress (al) parallel to an initially horizontal layering. The zone of the maximum dissolution is localized near the top of a rigid indentor composed exclusively of Ca and embedded in the slaty matrix. With such a state of stress it is clear that the dissolution zone appeared in the sector of the matrix submitted to the maximum compressive stress, compared with the theoretical model of stress distribution around a rigid object (Fig. 2B, Stromgard 1973). But the shape of the curves of equal content in A1 (which corresponds to the curve of equal values of volume decrease) has not the same shape as the theoretical curves of, for example, the mean stress values. The zone of dissolution is a very narrow seam, shaped like parenthesis, largely extended beyond the rigid indentor (Fig. 2A), in contrast with the theoretical stress distribution curves which are more rounded and with maximum values localized near the top of the indentor (Fig. 2B). Such a difference has already been observed, and explained, by an internal deformation of the rocks (Casey 1976), but it is not the case here. The rock between the solution cleavage seams is almost undeformed (spaced cleavage). The difference between volume change and stress curves arises because in such a composite fine-grained rock as soon as one or two grains are dissolved in the zone of the maximum compressive stress a localized heterogeneity is induced in the rocks. The removal of soluble grains (quartz-calcite), which constitute the skeleton of the rock, disorganizes this rock and leads to a passive concentration of the insoluble species (here layered silicates with adsorbed or free associated fluids). This induces a relative increase in fluid content and a relative softening of the rock, with both chemical and mechanical effects. Following several authors (Heald 1956; Weyl 1959) the layered silicates catalyze the process of dissolution, either by increasing the kinetics of the reaction (model R) or by constituting some high speed paths for mass transfer (model D), since the diffusivity through these silicates is relatively high (depending on the fluid content), Calvet (1973). On the other hand, with the passive concentration of the layered silicates the dissolution zone becomes relatively less competent than the initial rock (an aggregate with both soluble and insoluble minerals, e.g. slate, is always less competent than an aggregate with only soluble minerals, e.g. quartzite or limestone). By comparison with the theoretical

distribution of the stress values in a viscous matrix around an elliptical soft inclusion (Cosgrove 1976), two zones of maximum compressive stress appear in this matrix near the tip of the inclusion. Consequently, as pointed out by Cosgrove, the dissolution is concentrated in these two zones with a slow but continuous propagation of the zone of dissolution in its own plane. Fletcher & Pollard (1981) named the development of such a solution zone perpendicular to the o1 direction 'anticrack propagation'. The driving force for mass transfer must, however, be discussed. Following the thermodynamic analysis (see, e.g. Paterson 1973), the major driving force is the difference, in the components of the stress normal to the solid/fluid interfaces, between the solution and deposition zones. Around a void filled with a free fluid phase and surrounded by a continuous solid, the normal stress value on the solid/fluid interface is constant (whatever the stress state in the solid) and cannot be the driving force for such a propagation. In this case the driving force must be a difference in elastic, plastic, or surface energy in the solid. A difference in normal stress (on the solid/fluid interface) can only be the driving force for mass transfer around a soft zone in a rock if this rock is an aggregate filled with its solution. In such a case, it is the increase in the 'trapped fluid' pressure (equivalent to the increase in the mean normal stress on the solid/fluid limit of the grain) which increases the dissolution of some grains in the rock near the tips of the parenthesis shaped solution seams (Figs 2 and 5). Another problem is the development of zones of dissolution on pre-existing surfaces (such as bedding, cleavage, faults etc.) oblique to the trl direction. This is attested by the observation of stylolite peaks oblique to the surface of dissolution (Arthaud & Mattauer 1969), and by the observation of the relations between two successive solution cleavages. In such a case it is generally observed that when two cleavages cross each other at various angles (in heterogeneous deformations) there is a minimum value for the angle of crossing of these planes (for example 30 ° in the Bourg d'Oisans basin, Gratier & Vialon 1980). This means that, during the development of the second (spaced) cleavage, the dissolution had locally continued on the first cleavage plane, without slipping on this plane. This observation first confirms the assumption made by Rutter (1983) of the possibility, for a dissolution surface, to support shearing stress. Secondly, this allows us to calculate an approximate value for the deviatoric stress during the pressure solutiondeposition process. By comparison with mechanical experiments on the studied rock (anisotropic

Pressure solution-deposition in sedimentary rocks slates) we have found that a low deviatoric stress value (about 10-25 MPa) is needed to avoid slipping on the cleavage when the angle between the normal to the cleavage plane and the al direction varies from 0 to 30°. All the examples studied on various rocks have shown that it is the initial heterogeneity of the rocks which imposes the position of the zone of dissolution for both chemical and mechanical reasons. The development of such a zone (schematically shown in Fig. 5) emphasizes this heterogeneity. A tectonic differentiation is thus a normal evolution for rocks deformed by pressure solution-deposition when these rocks are composed of minerals with various solubilities in the fluid phase.

Initiation and development of the zones of deposition The initiation and the development of the zones of deposition are very different from those of the zone of dissolution. Considering various examples of chemical differentiation around rigid objects, and observing the distribution of several minerals (by microprobe analysis), we have found that, if the dissolution zones have almost always the Same shape (of parenthesis) the deposition zones have a more varied aspect, depending on the nature of the rigid object. For example, in deformed slates the cubes of pyrite have large pressure shadow overgrowths, clearly observable even with the naked eye (Fig. 3A) whereas fossils, of the same shape and size, are accompanied either by scattered microscopic overgrowths or by sealed tectonic fractures within the fossil (revealed by cathodoluminescence studies, Gratier 1984). This can be explained by comparison with experimental deformation of a viscous matrix around a rigid object. When such a rigid object is not stuck to the matrix a large void appears in the zone of minimum stress (possibly tensional stress). But if the rigid object is well stuck to the matrix either a stretching of the matrix, with associated microcavities in the zone of minimum stress, or the fracture of the rigid object can appear. The same difference can appear in nature. Usually, the deposition thus occurs in open cavities (of various size) probably filled with free fluid (see below). Another difference between deposition zones and dissolution zones is that the former rarely show a large propagation in their own plane, as do the latter. The ratio 'thickness over width' of the zone of deposition is always much higher than that of the zone of dissolution. This perhaps occurs because

29

the development of the zones of deposition is often limited by the development of the zone of dissolution which constitutes some soft barriers perpendicular to the veins. But the sealing of a tectonic vein with the most competent mineral of the rock (quartz-calcite) also induces a hardening of the rock which facilitates the development of new veins 'en 6chelon' in the 'protected' zone of the first vein. We have thus commonly observed the schematical development given in Fig. 5. Another difference between dissolution and deposition is their evolution with time. The aspect of a crystal within a sealed cavity gives some indication of the mechanism of opening of the cavity. Euhedral crystals indicate growth within a free fluid phase in permanent open cavities (Fig. 3), whilst fibrous crystals parallel to the direction of maximum elongation are developed by repeated increments of microcrack opening, followed by sealing of the microcracks by deposition of the mineral from the dissolution zone (Ramsay 1980). Each sequence of growth is expressed by a network of fluid inclusions. The study of these inclusions gives some indications about the nature of the fluid phase, its pressure and temperature. We have found that the deformation of rocks by pressure solution-deposition generally occurs between 0.1-250 MPa and 20-400°C, water (with some content of NaC1) being the fluid phase. The comparison of the P,T values between successive veins associated with successive cleavages also gives an indication of the behaviour of quartz and calcite with various P, T conditions (see in the introduction the effect of c). Jenatton (1981) has shown that even with a single vein the P, T conditions may have evolved during the progressive deposition. But all these studies give the general evolution of P,T values with time. Another interesting result, obtained by Mullis (1975), has shown that even during the growth of an euhedral crystal in a permanent open cavity the fluid pressure evolves in a relatively discontinuous manner. As the crackseal process is also clearly discontinuous with time, it seems that in nature, and in opposition to the dissolution process, the deposition process, either with euhedral crystals, or with fibrous crystals, is rather discontinuous both in space (growth on solid/free fluid interface) and in time (periodic change of pressure values). This discontinuous process may be explained by successive changes in the volume of the cavities. A sudden opening of a vein (linked to an irregular fracture propagation) may induce a sudden decrease in the fluid pressure, which consequently induces the precipitation of the solute species and this during all the time of the re-equilibration of the pressure by the fluid flow

(A)

I*, o,

b

Iio

b

(F)

Zlta

b

+~JT--,~, iJ.-~

(C)

IV*

b

~_.

GEN

;l':

:

.: i

.........

......

.

.

.

.

~-~ :-: ~ ~

.

.

.

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

.

(G)

[ ...................

......

. . . . . . .

(D)

4

th

fluid

fluid"

(E)

step

step

.,,Hng

1

I

fracture

step

]~"

TT

I

FIG. 3. Various aspects of tectonic overgrowth: (A) Schematic view. (B) Photograph (from Pyrenees). Around pyrite = large pressure shadow, around fossil = scattered sealed microvoids and/or sealed brittle fracture in the fossil. Various aspects of crystals in tectonic cavities : (C) Growth of euhedral crystals in permanent open cavities. (D) Growth of fibrous minerals by successive microcracks immediately sealed. (E) Successive growth processes. Development with time of the deposition processes: (F) Periodic changes in fluid pressure for euhedral crystals (after Mullis 1976). (G) Self-healing of fracture during the crack seal process (fibrous crystals).

0

Soo

iooo

tSO0

P bar

i

0

Pressure solution-deposition in sedimentary rocks into the cavity. As for dissolution, a difference in fluid pressure or, more generally, a difference in the stress components normal to the solid/fluid limit, can be the driving force for the deposition (Fig. 1B, C, D). But experimental work on the healing of fractures (Gordina & Neverov 1967; Wiederhorn & Bolz 1970) has shown that this process might be developed, even without such a difference in normal stress, by the simple effect of the surface energy reequilibration along the fracture. Such a slow process of self-healing of a fracture was reproduced on quartz at 400°C, 200 MPa in water. After several weeks a cluster of small, fluid inclusions simply marked the position of the initial fracture (Smith & Evans 1984; Wilkins et al. 1985). In all the studied cases the tectonic veins developed during the pressure solution-deposition process have the aspect of a brittle fracture. Taking into account the relatively high confining pressure (up to 250 MPa) this means that the fluid pressure must remain not very different from the lithostatic pressure (Brace 1972): either a little above (hydraulic fracture) or a little below (assisted fracture). Such a brittle fracture is developed perpendicular to the tr~ direction in initially isotropic materials (Paterson 1978). However, it is well known that such tectonic veins are sometimes developed parallel to preexisting planes such as cleavage planes. Let us consider anisotropic rocks, with quite different tensile strength values perpendicular and parallel to the cleavage (e.g. respectively 1 and 30 MPa), submitted to low deviatoric stresses ((r~-o'3 < 30 MPa, tr~ and 0-3 being respectively perpendicular and parallel to the cleavage). With a slow increase in fluid pressure, the pressure value needed to open a fracture parallel to the cleavage (P > o"1 --~l) is inferior to that necessary to open a fracture perpendicular to this cleavage (P > o"3 +

3[

30). Fractures perpendicular to trl can thus appear more easily than fractures parallel to al in such anisotropic rocks (Fig. 4). As for dissolution, the situation of the zone of deposition can thus be imposed by the pre-existing structures of the rocks. A comparison of the evolution of the zone of dissolution and deposition has shown significant differences in the paths of propagation, in their continuity with time and in their driving forces. On the other hand, the initial heterogeneity (structure) of the rocks plays a role in the two processes. In all cases the dissolution appears in the relatively less competent sector (with the highest content in layered silicates) whereas the deposition zone appears in the more competent sector (with the highest content in soluble species). Both processes emphasized the initial difference, they are self-amplified during a progressive deformation.

Processes of tectonic differentiation Development of the processes Considering the development of the zones of dissolution and of the zones of deposition schematically represented in Fig. 5, the process of pressure solution-deposition naturally leads to the development of a tectonic layering. Several authors have discussed the development of such differentiated layers in metamorphic environments (see Robin 1979, for a pertinent review). Several explanations have been given, based either on chemical or on mechanical effects. Following this author, the initial (or tectonically induced) heterogeneities of the rocks lead to a heterogeneous distribution of stresses. In rocks with alternatively competent and incompetent

~3 ' ~.~] :zi I'L")"

I:r ii'L':",! iJi:'"~:,~"t :I ®

I~.31~ !;~ ~!~!~:l'~!1:iI I.iiii~ .:.:iiiliS,ili/ i-iii ®

Pf>O'I+T 1 if (o'3+X(~> (0"1 +TI)

il:~ ~").3 . . . . . l.,lll Jl~lj .lj.'~I.I

(9

'

®

FIG. 4. In anisotropic rocks schematic development of veins perpendicular to 0-1.After the development of a cleavage (1 ~ 2), an hydraulic fracturing can appear (3) with particular condition for the deviatoric stress value (0"1 -- 0"3) and for the tensional strength values of the rock (T1 and T3 respectively perpendicular and parallel to the cleavage). At (4) the fracture is sealed (with a decrease of the fracture pressure Pf), and after (5) the vein is progressively boudinaged always with the same state of stress (0"zperpendicular to the vein).

32

J. P. Gratier 4,

4,

4,~1

deposition

in veins,

or in voids

dissoluti~

4 . . . . . . . ~. .:. . . :: . 1~ .,:i ~,~:~ , : ~:, , • : ::: ~ !,,

~:~/:'~:"

~(.: : ::,~: .F~P.I :L:.. I

iii ii!iiiii l soft

:

i!ii! ilii!!il iiiiii!i!iii!!iiii!!!!iiiiii!i!! i¸i¸iiii/~i:!~ii:i~i! !!i¸ii! iiii¸i¸iUiii!;i~ii/~i!~i~!Si~iiii~ii~iiiiii!i!i~iliiiiii!~il !!! i iii!!i¸i!i ii iii'tiiiii!iiiiiiiiiiii!!!i 1 ~i~ii!i!~!iiii:~,

..............,,~,,~................

object

FIG. 5. Schematic comparison between the development of zones of dissolution and of zones of deposition. The propagation of the zones of dissolution can be almost infinite in its own plane, whereas the propagation of the zones of deposition is always limited in space, with induced en bchelon veins in relatively competent sectors 'protected from dissolution and possibly hardened by scattered depositions'. This simultaneous development leads to a tectonic layering of natural rocks.

layers (or flattened domains), submitted to a compressive maximum value of s t r e s s (o'1) perpendicular to the banding, the minimum values of stress (0"3) parallel to the banding vary from incompetent to competent layers. Consequently a mass transfer occurs from the surfaces perpendicular to 0"t to the surfaces perpendicular to o"3 (see also Stromgard 1973). The driving force for the mass transfer is proportional to O'l-0.3 (Fig. 6A). Theoretical models for differentiation have been established with such a mechanical approach by Fletcher (1982) and Van der Molen (1985). These models deal with simple instantaneous cases. O f course when mass transfer proceeds for a long time, the large change of physical properties associated with the chemical differentiation must be taken into account. The problem is more complicated when the a~

direction is parallel (or oblique) to an initial banding since folding of the layer is often associated with the chemical differentiation. The mass transfer occurs from limb to hinge. If the zones of removal of matter are the axial plane solution cleavage of the fold (Fig. 6B1) the problem is not very different from the preceding. But the limit of the layer (or some interstratified incompetent layers) can also be the zone of dissolution (Fig. 6B2). In this case various explanations are given based on the fact that the value of the stress component normal to the layer increases with the rotation of this layer (Fig. 6C) (Gray & Durney 1979), or on the fact that layered silicates oblique or perpendicular to a~ (in the limb) catalyse the pressure induced transfer of the more mobile species (Robin 1979). To test these hypotheses the mass transfer balance

~0~ t'~

-----~.,~

.,.~. , - h i . . . . . .

"1\

(BJ)

~3~ = ~

=-"~-I~

I

7:-

I

\

\\ 0

45 9O Ar~le 8

(A2) (B2) (C) FIG. 6. (A) Distribution of principal stress values (0"1,a3) in a heterogeneous material with alternately competent (c) and incompetent (i) layers (A1), or with competent domains embedded within incompetent matrix (A2), after Robin (1978). (B) Two types of dissolution seams in fold limb: axial plane cleavage (B1), or initial bedding joint (B2). (C) Evolution of the normal to bed stress values with the increase of the 0 angle (after Gray & Durney 1979).

Pressure solution-deposition in sedimentary rocks through the folds can be made with chemical comparative analyses (Gratier 1983). But we have observed that the volume of the closed system sometimes changed during the folding process (with the decrease of the hinge/limb angle ~). The closed volume decreases from the size of the limb/ hinge pair (low ~ values) to the size of the microlayers within the limb (high ~ values) (Gratier 1984). Depending on several factors such as the state of stress, the relative orientation of favourable surfaces and the possibilities of mass transfer, the mean distance of mass transfer can change during a progressive deformation. In all cases however, a rule is respected: the zone of dissolution occurs in incompetent sectors, the zone of deposition in competent sectors, as evidenced by .the many examples of solution-deposition associated with boudinage, ptygmatic folding, mass transfer around rigid objects etc. This rule could seem surprising since the competent sector usually has a higher content of soluble species than the incompetent sector. An explanation of this paradoxical behaviour is that the development of the dissolution needs the presence of both a soluble species and its solution. In competent sectors, the high content of soluble species is not usually associated with an equivalent high content of solution around the soluble grains. The fluid phase is rather localized in voids or in brittle fractures frequently perpendicular to 0-3 and then badly oriented to enhance the dissolution (but well oriented for the growth of the reprecipitated minerals). In incompetent sectors, with lower contents of mobile species, the soluble grains are well surrounded by layered

silicates (micas, clays), associated with free, or adsorbed, fluid and acting as privileged paths for the mass transfer. The chemical differentiation leads to changes both in the mechanical properties and in the possibilities of mass transfer. These two variations lead to the same result: an increase of insoluble content (clays or layered silicates) softens the rocks and connects privileged paths for mass transfer normal to the al direction, then promotes the dissolution. An increase of soluble content (quartz, calcite) hardens the rocks and facilitates the development of open veins (or microcavities) normal to the 0-3 direction then promotes deposition (Fig. 5). The chemical differentiation is then self-induced. A similar feedback mechanism was proposed by Merino et al. (1983) between porosity and solution-rate. However we did not find a large change of porosity during the dissolution of slates (Gratier 1983).

Compositional limits of the differentiated layers The question now is to know if the differentiation may lead to a complete partition between soluble and insoluble species or if there are compositional limits for two differentiated layers. In the example of a heterogeneously deformed matrix around a ptygmatic fold (Fig. 7) we have been able to estimate both the volume change values (A) and the internal deformation values (X-Z) (Gratier 1983). When plotting the A and X Z values along six layers of the slaty matrix on ratio

E,-E,f

33

s°luble/insolubla min.

y),, ~. /

f

/

.....

competent,

f

6 1

% incompetent

layer.

D-

1

2

E 2 -E 3

initial state

(A)

(B)

eq'uilib rium

It)

FIG. 7. (A) Evolution of values of volume decrease along several layers (of the same competence) in a slaty matrix. (B) (shaded area ---open system) around a ptygmatic fold (hatched area). The layers less folded (upper layers) show a linear relation between the volume change and the internal deformation values. The layers most deformed (lower layers) do not show such a linear relation since the volume decrease value is limited (-0.5) before the complete depletion of soluble species. (C) Schematic chemical evolution of two contiguous domains, by mass transfer in closed system, from an initial state (with slight difference) to an equilibrium composition of both competent and incompetent differentiated domains.

34

J. P. Gratier

a Flinn diagram (see also Ramsay & Wood 1974) it is possible to follow the evolution of the deformation (the path of deformation) by assuming that the spatial evolution along each layer is equivalent to a temporal evolution. We may note that a linear relation appears between volume change and internal deformation for the less deformed layers. But for the most folded, just in the vicinity of the competent strata, it seems that a maximum value of the volume change cannot be passed (A ~<0.5). This limit value is not linked to the complete removal of the soluble species, which remains about 20 and 18% respectively for quartz and calcite. It thus seems that all the soluble minerals cannot be 'extracted' from the zone submitted to the maximum compressive stress. It is observed, for example, that deformed Liassic slates of the French Alps always have the same limiting composition in the zone of dissolution. This perhaps occurs since the soluble minerals are too scattered in the matrix. This means that there is probably an equilibrium composition in the zone of dissolution which depends on the various factors influencing the rate of mass transfer (see the introduction and the following section). For the zone of deposition it is different as such zones can be completely composed of soluble species. However, if the temperature and the fluid pressure remain constant during the progressive deformation, the ratio of the quartz/calcite content is dependent on these P,T conditions since the solubility of these minerals are respectively normal and inverse with the temperature (see the introduction and Gratier 1984).

Spacing of the tectonic layers The preceding discussion can be extended to the problem of the spacing of the tectonic layers. Two cases must be distinguished: (i) when the volume of the closed system is largely superior to the spacing of the zones of dissolution, the system is said to be open, all the mobile species can move out of the studied sector; (ii) when the volume of the closed system is of the same order as the size of the tectonic layer, the system is closed, all the soluble species removed from the zones of dissolution are reprecipitated in the neighbouring deposition zone.

Opensystem In all the natural examples studied, we have observed the same rule for the position of the zone of dissolution. These zones are always associated with a heterogeneity in the rocks (see

the schematic relation in Figs 2-5). In slaty, finegrained rocks, the spacing of the solution seams is clearly always linked to the size of the heterogeneity of the rocks (Fig. 8A). Usually these heterogeneities are selectively more rigid than the matrix (Fig. 2A), but they can also be relatively less rigid than the matrix (Fig. 2C-5). They can be associated with the sedimentary process (fossils, grains or polycrystalline domains with contrast of competence with the matrix), but can also be induced by some process of deformation such as the folding (with dissolution in the limb of the fold, see Fig. 6). In limestone or quartzite the problem is more complicated because dissolution markers (often stylolites) appear with the same size as the grain size (see the section on stylolites).

Closedsystem If the soluble species are reprecipitated in the immediate vicinity of the zone of dissolution (with a closed system between neighbouring solution and deposition zones) the spacing of the differentiated layers is fixed by two successive effects. At the initiation of the process of differentiation the spacing of the zone of dissolution is always linked to the size of the heterogeneities of the rock (see the preceding section). However, as soon as the mobile species are reprecipitated in the zones of deposition these zones become harder than the initial rock. They then act as rigid indentors and induce, at their limits, new zones of dissolution. Depending on the mean distance of mass transfer, there are two possibilities. If the mean distance of transfer is inferior to the size of the initial heterogeneity (Fig. 8B top) this heterogeneity is progressively destroyed. If the mean distance of transfer is superior to the size of the initial heterogeneity (Fig. 8B bottom) this heterogeneity is progressively integrated within the zone of deposition. With such a closed system a regular spacing of the differentiated layers is acquired. The spacing value depends on the various factors of the rate of deformation (see the introduction). These factors are the driving force Act, (often proportional to the deviatoric stress value) and the nature of the solid and fluid (effect on c). Depending on the limiting process of the deformation rate (see the introduction) two other factors are probably more important: the effects of the factors of transfer (kinetics of reaction, diffusion coefficient, permeability), and the geometry of the path of transfer (or of the surface reaction). For example, the natural rate of deformation is often limited by the rate of diffusion (Gratier 1984). At, for example, 200°C,

Pressure solution-deposition in sedimentary rocks

I iil )

35

":-.'?'l

~:',"." o e

:cS:/:-:i:i.... :.: .i~i!::}l

(A)

(B)

Fio. 8. Relation between spacing of zone of dissolution and size of heterogeneities in rocks. (A) In an open system with removal of the soluble species out of the studied sector: the spacing between the solution seams is fixed by the size of initial (or tectonically induced) heterogeneities (fossil, relatively rigid (or soft) domains, folds, etc.). (B) In a closed system the redeposition of the soluble species in the vicinity of the zone of dissolution hardens the sectors of deposition. This sector, becoming relatively more competent than the matrix, induces dissolution zones at their limits. Depending on the mean distance of mass transfer the spacing between the induced solution plane may be superior (bottom), or inferior (top) to the spacing between the initial solution plane. the values of this coefficient range from 10 -9 to 10-16 m 2 s-1 depending on the fluid content in the rocks (Fisher & Elliot 1974), whereas the deviatoric stress values probably ranged from 1 to 102 MPa (the maximum value being of the order of the plastic yield point for most rocks). This means that usually the possibilities for mass transfer through rocks are the crucial factors which impose the rate of deformation and the size of the regular spacing of the differentiated layers. A schematic comparison clearly illustrates this fact. The mean thickness of the differentiated layers in metamorphic environments ranges from 10 -4 to 10-2m, with high temperature and pressure but relatively low fluid content, whereas the mean thickness of the differentiated layers in tectonic environments is larger (ranging from 10 -3 to 10-1m) with low temperature and pressure but relatively high fluid content. In each domain, the scattering of the values also depends on geometric factors: the ratio between the volume of the path of transfer and the total volume which is schematically linked to the grain size value and to the content of layered silicates (with high diffusion values). When considering the values of the mean

thickness observed in naturally deformed rocks (at relatively low temperature and pressure: 100400°C, 50-250 MPa) it is also clearly apparent that some of these values (10 -1 m) are of the same order of magnitude as the mean thickness of many stratigraphic bandings. We may wonder if the very regular spacing sometimes described in geological formations could not be, at least partially, linked to a mass transfer by pressure solution-deposition from competent to incompetent layers, with a composition limit and equilibrium thickness for each layer dependent on the conditions of diagenesis. An accurate study of the evolution of the composition and of thickness ofstratigraphic banding, with depth, for example, could perhaps answer this question.

Development of stylolites Two characteristics of the appearance of stylolites must be discussed namely the amplitude and the wavelength of their peaks (Bathurst 1975). The amplitude of the peaks is often used to estimate the relative displacement of the blocks limited by the dissolution surface (Stockdale 1922). Yet

J. P. Gratier

36

many natural observations have shown that when the process of dissolution becomes important, the peaks tend to be destroyed (Gratier 1976). The stylolite zone becomes successively an irregular surface and then possibly a fine incompetent layer. In this case the only method of estimating the volume removed from the solution zone is to make a comparative chemical analysis of the insoluble species (between solution zone and initial rock). The wavelength of tectonic stylolites is often of the same order of magnitude as that of the grain size of limestones or quartzite. In this case we may think that the relative difference in competence of neighbouring grains induces localized dissolution, the mean wavelength being fixed by the mean size of the heterogeneities (the grains) (Fig. 9). On the contrary, diagenetic stylolites often have a wavelength value largely superior to those of the mean grain size. This can be explained if the pressure solution-deposition occurs in rocks with an initial porosity and on pre-existing surfaces of stratigraphic origin. If zones of dissolution appear along such a surface, and if the matter removed from these zones is reprecipitated in the immediate vicinity of the initial surface, the sealing of the voids of the rocks will tend to consolidate these zones with deposition (Fig. 9). These zones, becoming

progressively harder than the initial rock, will act as rigid indentors and induce localized zones of dissolution at their limits, i.e. on the initial surface (Fig. 9). With such an evolution the mean wavelength of the stylolite is of the order of magnitude of the size of the consolidated zones of deposition. It is a process of differentiation not very different from the process which leads to a tectonic layering, but it is rather a tectonic differentiated cementation. In this case, the wavelength of the stylolites gives an indication of the possibilities for mass transfer at the moment of their development.

Conclusion The deformation of rocks by pressure solutiondeposition is always associated with the development of a tectonic layering when the rocks are initially composed of minerals with various mobilities. A feedback mechanism is proposed between clay- (or layered silicate-) content (insoluble species) and the solution-deposition rate. The mechanical properties and the possibilities for mass transfer both evolve with the change of insoluble content, the major effect being given by the second factor in most natural large deformations.

::~:;i::~i!:::' ~

...i'.'i:.i,--. -'.~';" .-:-!!;':'~':'ii~':.

:ii:ii!ii::i::~:;i~' :: :"::: ~

............. :=::' ~:-:~::~

• ;....--.:..... ": . .'.-:-:-.--.-=-~ "::: ~~ ::/;;!~

~ .

styIolite

~ ~ :,~:

,. :~ ,~:~]~.~,,-~::,.~. -.~. ,~ '~":" ~,v'-'°!:z"-i '-. ~. ~ ,~'~iP~*~.-";~',~.~ ,~.~:::,:,a~,~:~.:,-,,~:-;~..

• .~I,,..,.~~,..~.'.~.,

1

• .::'~<;',.,",..'?,

.-

.... ~;;..,M-'.. '~'L.'-;.~ ;'..~

.....

.

.

.

.

.

2

.

.

ep°siti°n ~ ~!' ....... ,~, ,.-. ~...~,,,~ .... ~...~. ~'-',--P ",°".'.;"-;"; ..,"~;'.~'""~'41.'~.'."~'~.~, ;r~,'. ~;, ,'~i "'~ .... ~.-",~.;",<' .... ~':,"/':..;;~:',',','::;,~ ,'.: ¢;'.¢~,:,~::. 't:.~r,:~':.:.,~,{,~7.,'.~"~;{~5,2~1-~,*.~::~-- '

• ~N." ," '~ ,, .A,C.~,,,,,.-.~

.~'.~',:~,.~;..',-'f~,',~;:

!:@i~:; :21:1 ~:: : :!~-~ ...i:'

~,~7~ ,.:~

.i

:l~..';".

-."1,~_~' "" ".,. :. ".. . . . . . . .

.~~." -.

""

.....

indentor

3

FIG. 9. Development of peaks of stylolites and acquisition of their characteristic wavelength. (A) Effect of the shape of the limit of the grain for tectonic stylolites. (B) Effect of the size of the heterogeneities along an initial surface for diagenetic stylolite. These heterogeneities could be linked to the local hardening of the initial rocks by localized deposition in the vicinity of the initial dissolution. The mean size of the hardened sector depends on the possibilities of mass transfer through the rocks, and imposes the mean wavelength of the stylolites.

Pressure solution-deposition in sedimentary rocks

37

(A)

oI

c8)

/

(c) FIG. 10. Mechanical and chemical differentiation of rocks under stress by pressure solution-deposition. The paths of the transformation depends on the initial arrangement of the rocks. From an unstable initial state (initial layers perpendicular to trl, scattered heterogeneities or initial layers parallel to trl) we may obtain, under a deviatoric stress, an equilibrium state (both in the composition and the size of the differentiated layers) for given conditions of deformation (possibilities of mass transfer, nature of the solid/fluid phases, state of stress, temperature and pressure, etc.).

D e p e n d i n g on the g e o m e t r y of the initial heterogeneities of the rocks and of the orientation of favourable surfaces to the o1 direction, various progressive d e v e l o p m e n t s of the differentiated layers are possible such as t r a n s f o r m a t i o n of an initial layering (Fig. 10A), d e v e l o p m e n t of successive layers a r o u n d dispersed heterogeneities (Fig. 10B), differentiated layering associated with folding (Fig. 10C). I n all cases the differentiated layers tend to an equilibrium composition a n d thickness d e p e n d i n g on the possibilities of mass transfer, the driving force of this mass transfer,

the nature of the soluble and solvent m a t t e r a n d the t e m p e r a t u r e a n d pressure conditions. As the same type of differentiation could a p p e a r along an initial well-oriented surface leading to differentiated c e m e n t a t i o n of the rocks, this process could control the size of diagenetic stylolites. ACKNOWLEDGEMENTS: I would like to thank G. Buffet, F. Thouvenot, P. Vialon, G. Vivier for their help, and M. Jones for his careful revision. This work received financial support from the Petrologic ATP no. 8158 and the transfert ATP no. 1512. (CNRS).

References ARTHAUD,F. & MATTAUER,M. 1969. Les d6formations

CALVET,R. 1973. Hydratation de la montmorillonite et

naturelles. Essai d'6valuation des conditions pression-temp6rature de diff6rents types de d6formation. Rev. Ind. Min. no Spbcial, juillet 1969. BARNS, R. L., KOLB, E. O., LAUDISE, R.A., SIMPSON, E. E. & KROUPA, K. M. 1976. Production and perfection of Z face quartz. J. Cryst. Growth, 34, 189-97. BATHURST,R. C. G. 1975. Carbonate Sediment and their Diagenesis: Amsterdam, Elsevier, pp. 658. BRACE, W. F. 1972. Pore pressure in geophysics. In: HEARD, H. C., BORG, I. Y., CARTER, N. L. & RALEIGH,C. I . (eds) Flow and Fracture of Rocks, Geophys. Monogr. 16, Amer. Geophys. Union.

diffusion des cations compensateurs. Ann. AAron. 24, 77-217. CASEY,M. P. 1976. Application offinite element analysis to some problems of structural geology. Unpublished PhD thesis, Univ. London. 303 pp. COSGROVE, J. W. 1976. The formation of crenulation cleavage. J. geol. Soc. Lond. 132, 155-78. FISHER, G. W. & ELLIOT, D. 1974. Criteria for quasisteady diffusion and local equilibrium in metamorphism. In: HOFMANN, A. W., GILETTI, B. J., YODER, H. S. & YUND, R. A. (eds) Geochemical Transport and Kinetics, Carnegie Inst., Washington.

38

J.P. Gratier

FLETCHER, R. C. & POLLARD, D. D. 1981. Anticrack model for pressure solution surfaces. Geology, 9, 419-24. GIBBS, J. W. 1877. On the equilibrium of heterogeneous substances. Transactions of the Connecticut Academy. 3, 108-248, 343-524. Also published in: The scientific papers o f Gibbs, J. IV., Vol 1. Thermodynamics, pp. 55-353. Longman, Green and Cy., 1906. GORDINA, Y. V. & NEVEROV, V. V. 1967. Healing of cracks in rock salt crystals. Soy. Phys. Cryst. 12, 421-4. GRATIER, J. P. 1976. D6formation et changement de volume dans un marbre ~ stylolites de la r6gion de Rabat (Maroc). Bull. Soc. Gbol. Fr. 7, XVIII, 14619. & VIALON, P. 1980. Deformation pattern in a heterogeneous material: folded and cleaved sedimentary cover immediately overlying a crystalline basement (Oisans, French Alps). Tectonophysics, 65, 151-80. --, 1983. Estimation of volume changes by comparative chemical analyses in heterogeneously deformed rocks (folds with mass transfer). J. struct. Geol. 5, 329-39. & JENATTON, L. 1984. Deformation by solutiondeposition and re-equilibration of fluid inclusions in crystals depending on temperature, internal pressure and stress. J. struct. Geol. 6, 189-200. --, 1984. La dbformation des roches par dissolutioncristallisation : aspects naturels et exp~rimentaux de ce fluage avec transfert de matikre dans la cro~te sup~rieure. Th6se d'Etat, Universit6 de Grenoble. --, 1986. Experimental pressure solution-deposition on quartz grains: the crucial effect of the nature of the fluid. J. struct. Geol. in press. GRAY, D. R. & DURNEY, D. W. 1979. Investigation of the mechanical significance of crenulation cleavage. Tectonophysics, 58, 35-79. HEALD, M. T. 1956. Cementation of Simpson and St. Peter sandstone in part of Oklahoma, Arkansas and Missouri. J. Geol. 64, 16-30. JENATTON, L. 1981. Microthermombtrie des inclusions fluides des cristaux associbs ~ l'ouverture de fentes alpines. Th6se du 36me cycle, Univ. Grenoble. KINGERY, W. D., BOWEN, H. K. & UHLMANN, D. R. 1976. Introduction to Ceramics. (2nd ed). (Wiley.). MERINO, E., ORTOLEVA, P. & STRICKHOLM, P. 1983. Generation of evenly-spaced pressure-solution seams during (late) diagenesis: a kinetic theory. Contrib. Mineral. Petrol. 82, 360-70.

-

-

-

-

MULLIS, J. 1975. Growth conditions of quartz crystals from Val d'Illiez (Valais, Switzerland). Schweiz. Mineralp petrogr. Mitt. 55, 419-29. PATERSON, M. S. 1973. Non-hydrostatic thermodynamics and its geological applications. Rev. geophys. Space Phy. 11, 355-90. --, 1978. Experimental Rock Deformation. The Brittle Field. Springer-Verlag. RAMSAY, J. G. 1980. The crack-seal mechanism of rock deformation. Nature, 284, 135-9. -& WOOD, D. S. 1974. The geometric effects of volume change during deformation processes. Tectonophysics, 16, 263-77. ROBIN, P. I. 1979. Theory of metamorphic segregation and related processes. Geochim. Cosmochimica Acta, 43, 1587-600. ROBINSON, R. A. & STOKES, R. M. 1959. Electrolyte Solution ( 2nd edn). Butterworth, London. RUTTER, E. H. 1976. The kinetic of rock deformation by pressure-solution. Philos. Trans. Roy. Soc. Lon. A283, 43-54. , 1983. Pressure solution in nature, theory and experiment. J. geol. Soc. Lond. 140, 725-40. SMITH, D. L. & EVANS, B. 1984. Diffusional crack healing in quartz. J. geoph. Res. 89, 4125-36. SORBY, H. C. 1863. On the direct correlation of mechanical and chemical forces. Proc. Roy. Soc. Lond. 12, 538-50. SOULA, J. C. & DEBAT, P. 1976. Developpement et caract6res des litages tectoniques. Bull. Soc. geol. Ft. XVIII, 1365-744. STOCKDALE, P. B. 1922. Stylolites. Their nature and origin. Thesis, Indiana Univ. Studies, 9, 97. STROMGARD, K. E. 1973. Stress distribution during formation of boudinage and pressure shadows. Tectonophysics, 16, 215-43. VAN DER MOLEN, 1985. Interlayer material transport during layer-normal shortening, part 1: model. Tectonophysics, 115, 275-97. WEYL, P. K. 1959. Pressure solution and force of crystallization, a phenomenological theory. J. geophys. Res. 64, 2001-25. WIEDERHORN, S. M. & BOLZ, L. H. 1970. Stress corrosion and static fatigue of glass. J. Am. Ceram. Soc. 53, 543-8. WILKINS, R. W. I., GRATIER, J. P. & JENATTON, L. 1985. Experimental observation of the healing of cracks and the formation of secondary inclusions in halite and quartz. Eur. Curr. Research on Fluid Inclusions, Gottingen.

J. P. GRATIER, IRIGM, University of Grenoble 1 B.P. 68, 38402 Saint Martin d'Hbres, France.

Hydrocarbon migration by hydraulic fracturing G. Mandl & R. M. Harkness The fracturing of cap rocks of very low permeability--in particular shales--by the intrusion of hydrocarbons under various in situ stresses has been considered. The conditions are discussed for the domination of the migration mode by the formation of either hydrocarbon dykes or sills, or a combination of both. Application of concepts of pro-elasticity and fracture mechanics to the process of dyke formation indicates the key roles of capillary pressure, premigration water pressure and in situ rock stresses, and provides conditions for maintaining open dykelets, for the change in width and spacing of growing fractures, and their closure, and explains differences in the behaviour of oil- and gas-filled dykelets. In discussing the situations where the state of in situ stresses favours the development of bedding-parallel intrusion fractures, various mechanisms are suggested which may allow interconnecting dykelets to form and feed hydrocarbon sills at stratigraphically higher levels.

The history of hydrocarbon movements after expulsion from source beds is part of any appraisal of the hydrocarbon potential of a basin or subbasin. The route taken by hydrocarbons during this 'secondary' migration phase in a given geological situation will obviously depend on the modes of transport available. Frequently, the stratigraphic and structural setting at the time of source maturation does not allow a direct connection between source and potential traps via permeable carrier beds, especially when the source beds are directly overlain by continuous rock layers--particularly shales---of very low or even zero permeability. Under what conditions could hydrocarbons migrate through such a barrier? Or, addressing a related problem, when would hydrocarbons leak through the cap rock seal of a hydrocarbon trap ? We shall assume in this paper that prior to the maturation of the source rock, or the charging of a trap, the overlying barrier rock is intact, i.e. not transected by faults or open fractures which would provide extra permeability. We shall discuss the possibility of hydrocarbons intruding and penetrating the cap rock by generating natural hydraulic fractures. Obviously, this mechanism would provide the most efficient mode of hydrocarbon migration through rocks of very low or zero permeability. It has long been recognized that fluids in the crust can produce fissures through which fluid can escape. Secor (1965), N. J. Price (1966, 1977, 1978), Phillips (1972) and others have discussed the mechanism of hydraulic fracturing by water. Snarsky (1962) suggested hydraulic fracturing as a mechanism for hydrocarbon expulsion from source rocks, while Dallmus (1955) suggested hydraulic fracturing along bedding planes as a

possible mechanism for lateral oil migration. More recently, du Rouchet (1981) has dealt with the vertical transfer of hydrocarbons from source beds through cap rocks by the opening of vertical fractures, emphasizing the role of the capillary pressure difference between the intruding hydrocarbons and the pore water. Early evidence of oil migration through networks of fractures was reported by Schnaebele (1948) from the oil mines (now closed) in Pechelbronn, France, where fractures crossing marl beds were found to be lined by oil films or, in many cases, by a greenish halo. Natural hydraulic fractures are extension fractures. This means that the rock is separated by the fluid pressure, which therefore has to overcome the tensile strength T (positive) and to push aside the rock walls against the action of the total rock stress. The fractures may be concordant or discordant with the bedding and may be referred to as fluid 'sills' or 'dykelets' in accordance with the nomenclature for igneous instrusions, for which Anderson (1942) presented a first mechanical theory. With reference to the mechanical origin of these fluid-filled extension fractures we distinguish between two types: 'internal' fractures and 'intrusion' fractures. The formation of internal fractures requires that an effective normal stress (total stress minus pore-fluid pressure) becomes tensile, i.e. the pore-fluid pressure exceeds the total stress component, as shown for a horizontal bed in Fig. 1(A). In contrast, intrusion fractures may be formed in a bed, under effective compression, by the injection of fluids whose pressure exceeds a component of total normal stress plus the tensile strength. This condition is illustrated in Fig. I(B). The tensile strength perpendicular

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks,

Geological Society Special Publication No. 29, pp. 39-53.

39

40

G. Mandl& R. M. Harkness INTERNAL EXTENSION FRACTURES

~rH' :O-H-p ~ O

the breakage of cohesive bonds along a plane perpendicular to the bedding. If the beddingparallel (total) normal stress which acts just prior to the formation of dykelets is denoted by 0-11, then the well-known fracture condition is: ph~ >~a ll + TII.

OVERBURDEN

Phc )O'H+ TH(Verlical}

(1)

Since it was assumed that the cap rock is under effective compression before the onset of intrusion fracturing, the pore-water pressure p at the moment dykelets start to form will have to satisfy

Phc)- O'v + T.k(HOrizontal ) (B)

FIG. 1. Hydraulic fracturing. (A) Cap rock under extension, (B) Cap rock under effective compression. to the bedding (T±) is usually much smaller than that parallel to the bedding (Ttl) and may be neglected with respect to the total overburden stress av. Whereas internal fractures have to be filled with fluid from the surrounding lowpermeability pore space, intrusion fractures are fed from outside. This affects the growth, width and spacing of the fractures differently. In particular, intrusion fractures may continue across beds of somewhat different lithology, while discordant internal fractures usually terminate at bed boundaries. In hydrocarbon intrusion the pressure of the non-wetting fracture fluid will exceed the porewater pressure by a capillary pressure difference. Since the hydrocarbon/water interfaces in the pores of the fracture walls are self-adjusting within certain limits, the capillary pressure may vary between zero and a maximum value p~. In this discussion we are concerned with the formation of hydrocarbon intrusion fractures in cap rocks of very low (or zero) permeability which are under effective compression and therefore not transected by open internal fractures. In general, this situation will exist in sediments that have never been subjected to extensional strain. First, we deal with the process of discordant or dykelet-type of intrusion fracturing.

Hydrocarbon dykelets The fracture condition

For a set of parallel hydrocarbon dykelets to form, the hydrocarbon pressure Ph~ has to overcome the smallest total normal stress that acts parallel to the bedding prior to hydrocarbon intrusion, plus the bedding-parallel tensile strength Tll, i.e. the resistance of the rock against

p
(2)

Although condition (1) is widely used, some comments should be made before using it in the present context. Firstly, it should be emphasized that condition (1) is little more than a fairly crude approximation of reality. The fluid pressure inside an intrusion fracture exerts a leverage action upon the fracture walls which increases with fracture length. Consequently the fracture may even penetrate into a region where all >Ph~ (Lachenbruch 1962). Moreover, the fracture-fluid pressure required to break the cohesive bonds in the tip region will also decrease as a monotonic function of increasing fracture length, as increasing leverage allows the same tensile stresses at the fracture tip to be produced with less fluid pressure. The so-called 'tensile strength' T is therefore not a real material constant. However, in spite of this, fracture condition (1) may suffice to define the in situ stress states that will most likely promote intrusion fracturing perpendicular to the direction of the least-compressive stress. When dealing with growth, width, spacing and interaction of fractures, however, one has to resort to the more sophisticated theory of elastic fracture mechanics. Secondly, when condition (1) is satisfied, dykelets will form, but after a short transition period they may close and stop propagating. This may occur when the cap rock has some permeability which allows the fracture pressure and the pore pressure in the wall rock to communicate. Then the pore pressure in the wall rock will be raised by the higher pressure of the hydrocarbon fluid in the dykelets. This rise in pore-water pressure will oppose the increase in effective rock stress that is necessary to compress the wall rock in order to allow opening of the fractures. It is a transient phenomenon which depends on the drainage conditions of the wall rock (permeability, bed thickness, fracture spacing and conditions at the bed surface) and on the capacity of the rock to take up, by compression of the rock skeleton, the skeleton material and the porewater, fluid that has penetrated the fracture walls.

Hydrocarbon migration by hydraulicfracturing In shaly cap rocks of very low permeability this pressure rise will take place during the period when the pore water cannot drain off the rock between parallel hydrocarbon dykes. Therefore, the pressure difference between the non-wetting hydrocarbon fluid and the preferentially wetting water phase will tend towards a constant value that equals the capillary pressure established at the fracture wails. Hence, our first concern is whether the hydrocarbon pressure can maintain separation of the dykelet walls. This requires specifying the water pressure in the cap rock. Here we have to distinguish between two situations. The cap rock may directly overlie a maturing source bed, allowing the water pressure in the cap rock to communicate with the fluid pressures in the source bed. Or the cap rock may be isolated from the source bed by a thin impermeable barrier at its base or by such layers somewhere inside. In the latter case, the pore pressure in the cap rock will not be affected by the gradual build-up of fluid pressure in the source bed as it approaches maturity. In the former case, however, the pressure of the cap-rock water is bound to rise-at least in the basal part of the cap r o c k - - i n line with the rise in fluid pressure of the maturing source rock. Thus in this case, the hydrocarbon dykelets will have to form in a zone with highly overpressured pore water. This pressure may be considerably higher than that in the cap rock prior to the onset of hydrocarbon generation. The pre-fracturing rise in pore-water pressure is most likely to be caused by water expelled from the maturing source rock before the generation and migration of hydrocarbons inside the source rock have reached the state where oil or gas is 'ponding' beneath the cap-rock base. These hydrocarbons will then cause the final rise in pore-water pressure of the cap rock. The process of pore-pressure rise in the cap rock is sketched in Fig. 2, where the original water pressure in the cap rock, i.e. the pressure not yet affected by the onset of maturation of the source rock, is denoted by pO, while pi is the pore pressure just before hydrocarbon 'ponding'. In Fig. 2 it is assumed that the pressurephc in the ponding hydrocarbons equals the total overburden stress av. Actually, it might not quite reach that value when dykelets form. Unfortunately, our knowledge of the expulsion mechanisms that operate in various types of source rocks is insufficient to ascertain the value ofp i. We therefore do not know whether the difference between Phc and pi allows the maximal capillary pressure difference Pc to be established between the water and the hydrocarbon fluid inside a preferentially water-wet cap rock. If this uncertainty is taken into account by

4I

PRESSURE

HIGH PERMEABILITY ROCK

WA

OVERBURDEN STRESS

,,~cP~

Z

~ P

\

C A P ROCK

MATURE SOURCE ROCK

Phc=~v

FIG. 2. Cap rock with hard overpressure in pressure contact with maturing source rock. introducing an undetermined factor 0
(0 =<~ __<1).

(3)

It is important to note that the effect of this rise in pore pressure on the stresses in a laterally constrained cap rock may be considered to be of an essentially elastic nature. Stresses, pore pressure and strains that existed in the rock prior to this rise in pore pressure were of an inelastic nature since they were caused by inelastic geological processes such as compaction and diagenesis. Accordingly, we separate the total stresses and the pore pressure acting in the cap rock just prior to dykelet formation into parts tx° and pO, respectively, and into a part oa caused by the poro-elastic response to the rise Ap in pore pressure: 0"il=0"°11-[-0"i

and

p=p°+Ap.

(4)

In discussing the conditions required for the hydrocarbon pressure to maintain separation of the fracture walls of dykelets, we shall, for convenience, consider horizontal or nearly horizontal cap rock and source beds, and indicate the direction of the greatest and smallest horizontal compressive stresses by the subscript H and h, respectively. Compression of the rock between parallel vertical fractures that are opening implies a positive strain increment (shortening is counted as positive): AeH > 0.

(5)

The fractured bed as a whole is not allowed to extend in any horizontal direction. In the absence of tectonic extension, friction between individual beds will generally be sufficient to provide this

42

G. Mandl& R. M. Harkness

lateral constraint. Therefore, there will be no change in horizontal strain parallel to the fracture walls: aeh=0.

(6)

We assume that the rock responds as an isotropic, linearly elastic material to the straining (5) and (6), and that the compressibility of the skeleton material is negligibly small in comparison with the compressibility of the skeleton as a whole ('bulk compressibility'). The relationships between strains and effective stresses (o-'= O--p) then become formally identical with the linear elastic stress-strain relations for a compact material. If the compressibility of the skeleton material cannot be neglected, the effective stresses a' have to be replaced by the stresses O-x= a' + tip where fl is the compressibility of the skeleton material divided by the bulk compressibility. If a cap rock is in fluid/fluid contact with the maturing source rock, the compression (5) induces an elastic stress in addition to that already imposed on the non-extending layer by the prefracture rise (3) in pore pressure. Because of (5) and (6), the poro-elastic stress/strain relationships then yield for the cumulative elastic changes in effective horizontal and vertical stresses the inequality: • + Ao'i'H > 1 -v v [O-i, v + Ao- i, v] al'"

(7)

where v = Poisson's ratio. Since the overburden is not constrained in the vertical direction and since the total overburden stress O-v therefore remains unchanged when the pore pressure is raised, we have O'i'v+ Ao-i'v= -]~-p

(8)

where ~-p is now the sum of the pre-fracture rise (3) and the additional rise in pore pressure that is eventually induced by the shortening (5) of the cap rock between opening fractures. The hydrocarbon pressure acting upon solid and fluid parts of the fracture walls will raise the total horizontal stress to Phc. Consequently, the total elastic change in effective horizontal stress becomes: O-i'H+ AO-i'H=Phc -- O'°H-- ~ "

(9)

The inequality (7) may now be written as: 1-2v Phc - O-°H> - f : ~ -v " #"

(10)

The total pore pressure rise ~-p cannot be less than Phc--Pc--P° nor more than Phc--P°. It may therefore again be expressed by (3) with

generally attaining a somewhat different value. We insert this expression into (10) and also introduce the following ratios for the original cap rock stresses, which existed at the onset of source rock maturation: K o = O-°'n/o-°'v and

2 =p°/o-°v, (o-°v= O-v) (11)

from which follows: a°n = [2 + Ko. (1 - 2)]- O-v.

(12)

We allow for a reduction in Ph0 which may be caused by the low rate of hydrocarbon generation and the relatively low flow resistance of open dykelets, by introducing: p h c = # . O-v

(13)

where 2~/~

1

(14)

since generally Ph~ does not exceed the weight of the overburden column of unit cross-section and since it cannot be smaller than pO, the pore-water pressure in the cap rock before maturation of the source rock*. With (3), (11), (12), (13) the inequality (10) may be written in the form: av

l_--~v(1--2)

K o - ] - ~ _v ~

.

(15)

It should be noted that Poisson's ratio v in these formulae refers to the bulk behaviour of the rock under completely drained conditions (i.e. pore pressure is kept constant). Its value will therefore hardly ever be outside the range 0.1 < v < 0 . 3 . In contrast, the parameter Ko for sediment that has been deposited in a tectonically quiescent environment and has not subsequently been subjected to extensional tectonics will rarely have a value smaller than, say, 0.45. From soil mechanical experience and in situ stress measurements at * The generally held view that pore fluid pressures cannot exceed the weight of the overburden column of unit cross-section is not strictly correct. It implies that pore-fluid pressures and overburden stress are averaged over a large area. Locally, fluid pressures in excess of the overburden column may very well be balanced, provided that in other parts fluid pressures are correspondingly lower, so that the overall balance is not violated. A maturing source rock of small lateral extent, for instance, may develop fluid pressures exceeding try, since sub-vertical frictional stresses will be mobilized inside the rock above the margins of the overpressured source rock and will assist the overburden weight in balancing the local excess pressure. Hydrocarbon pressures in excess of the overburden stress may also be expected in certain cases where the maturing source rock is contained in dipping beds under horizontal tectonic compression. This will be discussed in connection with Fig. 11.

Hydrocarbon migration by hydraulicfracturing depth, we may infer that typical Ko values for tectonically undisturbed beds will lie between 0.45 and 0.75. It should be emphasized that/20 is n o t related to Poisson's ratio, as is often erroneously assumed, since it is mainly inelastic geological processes of compaction and diagenesis that control the development of horizontal rock stresses in response to the effective overburden load. If the total overburden stress is expressed in terms of an average specific weight of the rock of, say, 2.5 g cm-3, and the overburden thickness is expressed in kilometres, condition (15) may be rewritten as:

4- X O

_j__,,.....

//

! .....

43

0.05

£,/¢ =

i

O - - V -

--

O O-

z (km) p~ (Mpa)

cS-

1-2v < 0.04 oc(1 -- 2) [Ko (1 - v) - v(,u - 2)/(1 - 2)1 (16) Thus, this condition for tectonically undisturbed cap rocks of very low permeability imposes an upper limit on the admissible overburden thickness. If the overburden were thicker, not even the highest possible hydrocarbon pressure could maintain separation of vertical fracture walls or re-open the fractures, near the cap rock base. Although derived for cap rocks in fluid pressure communication with the maturing source rock, condition (15) or (16) also applies to cap rocks that were isolated from the maturing source rock by some impermeable barrier. In this case, the maximum capillary pressure will generally have been established along the fracture walls (a = 1). With a = l and # = 2 the right side of (16) becomes the greater. Therefore, systems of open dykelets will certainly not form if the overburden thickness exceeds: 1

-

2v

po

z (km) = 0.04 - 1 - 2 Ko ( l - v ) - v

(17)

A graphical representation of this upper bound, showing the influence of Ko, overpressuring (1 2<0.55), and Pc (MPa), is given in Fig. 3. Note also that high overpressures, i.e. values of 2 close to 1, in the cap-rock water prior to the onset of source rock maturation are required to allow an overburden thickness of practical interest. It should be emphasized at this point that an overburden thickness in excess of the critical value shown in equation (17) does not exclude the possibility that the formation of a dyke system is initiated. It merely excludes the possibility for such dykes to remain 'open' or to be re-opened. In particular, where a cap rock is separated from the maturing source rock by an impermeable interlayer, the difference between the pressure of

O-

%'.~ o'.8

t'.t

1.4

i

i

i

i

i

,

t.7

2.0

2.5

2.6

2.9

5.2

5.5

Pc (MPa) FIG. 3. Maximum overburden (km) for open hydrocarbon dykes in tectonically undisturbed cap rock. (Original pore pressure/overburden ratio 2=0.95.)

the water in the wall rock on first contact with the higher hydrocarbon pressure in the new fracture will, in general, be higher than po, and a certain transition period will elapse until the rise in water pressure of the wall rock is practically completed and has approached the value A-~ in equation (10). Hence, irrespective of overburden thickness, open dykes may form. They will, however, eventually begin to close again if the overburden thickness exceeds the critical value. This transient phenomenon of open dykelets growing under conditions violating formulae (16) and (17) is controlled by 'diffusion' of the pressure build-up from the fracture walls into the wall rock (the 'diffusion coefficient' is the product of the permeability and stiffness modules of the fluid-filled rock). A sufficiently small diffusion coefficient and a sufficiently large distance between neighbouring parallel fractures may even allow fractures to reach a cap rock top before their growth is stopped by closure. It may be of practical interest that in strong brittle rocks this fracture closure will not be perfect and misfits between the contracting walls may even allow such 'closed' fractures to act as migration paths. In contrast, fracture closure in 'soft' clays may be nearly perfect. Moreover, 'soft' clays will often be associated with a high 'primary' overpressure, prior to first contact with the hydrocarbons, which may prohibit an appreciable period of pressure build-up in the wall rock.

44

G. Mandl&

R. M. Harkness

Obviously the chances of open-dyke formation become much less when, as may commonly be the case, the cap rock was in fluid/fluid contact with the maturing source rock. In fact, the rise in water pressure (3) caused by the maturing source rock may easily induce, inside the laterally constrained cap rock, horizontal normal stresses in excess of the vertical one. This is illustrated by the Mohr diagram of Fig. 4. As long as these high pore-water pressures in the basal part of the cap rock prevail, no hydrocarbon dykes can form. Instead, the hydrocarbon fluid, accumulated at the cap rock base, will enter the pores of the rock at a rate limited by the upward drainage flow of cap-rock water, and the pressure drop Pc (i.e. ct= 1) will be established across the displacement front. The associated draw-down in water pressure will then be (Fig. 2): 6p = pi -- (Phc " Pc) = pi + pc -- trv

(18)

where we have assumed that Phc becomes equal to the total overburden stress, in accord with the fact that 'sill' formation would prevent the generation of higher hydrocarbon pressures. The net rise in water pressure caused by the maturing source rock will then amount to (Fig. 2): Ap = p i _ 6p _pO = tr, _pO -Pc.

(19)

This pressure rise will elastically change the effective horizontal stress to (Fig. 4): a'H = U°'H-- ~

V

Ap.

(20)

The total stress (7u = (7,n +pO +

Ap

(21)

may then be expressed in terms of the overburden stress by introducing (11):

I

v

trH=trv /(o(1--2)+1-- 1--v --

1--

Pc"

(22)

This total horizontal stress has to be smaller than t r , - TH in order to allow the preferential formation of hydrocarbon dykes. By means of (22) this condition may be written: 1-v Pc > - 1-2v

One will immediately notice that high 'primary'

PRIORTOHYDROCARBON - INDUCED RISE IN POREPRESSURE~.P: HORIZONTAL EFE STRESSo-~'=Ko.O-v°' MOHR CIRCLETANGENT FOR ELASTICU N L O A D I N G ~ AFTER RISEIN POREPRESSURE ~ AP= Phc- Pc " P°

\k I

l

~

)

t-v -AP

i I

FIG. 4. Pore pressure rise in laterally confined horizontal layer causing interchange of principal stresses. overpressures and high capillary pressures are required to satisfy this relationship. Inserting, for instance, the values v=0.25, Ko=2/3, 2=0.95, TH = 1 MPa and the specfic overburden weight 2.5 g cm -3, the overburden thickness z (km) would require a capillary pressure in excess of [1.5 + 0.625z (km)] Mpa. When comparing expression (23) with condition (15) for maintaining open dykelets, and recalling the assumption of a fully established capillary transition zone (ct= l) and the highest possible hydrocarbon pressure (/~= l) available for fracturing the cap rock, we notice that satisfying equation (23) also implies that equation (15) is satisfied. The hydrocarbon dykelets, once formed in a cap rock that was in pressure communication with the maturing source rock, can be kept open. This is, of course, not surprising since in generating dykelets not only do the rock walls have to be pushed aside, but the tensile rock strength also has to be overcome. Because of TH in expression (23) condition (15) may even remain satisfied when the hydrocarbon pressure in the dykelets drops a little (p < 1). In fact, one has to expect that in view of the very low rate of hydrocarbon generation and the relatively high permeability of a system of open dykelets, the pressure in the growing dykelets will drop below the value required to keep the fractures open. The dykelets will then temporarily close and stop growing until the hydrocarbon pressure has been built up again. In general, dykelets may be expected to grow in a spasmodic fashion. The conditions for the formation of open dykelets clearly demonstrate the key roles of the capillary displacement pressure and a high overpressure in the cap rock prior to source rock maturation. Both parameters have to attain high

Hydrocarbon migration by hydraulicfracturing values to make large-scale dyke formation possible in tectonically undisturbed cap rocks at depths of practical interest. In this connection it should be recalled that the capillary displacement pressure in a water-wet rock is inversely proportional to the square root of permeability, and that the interfacial tension of a gas/water system is much less reduced by an increase in temperature than is the interfacial tension of an oil/water system. Consequently, we conclude that largescale systems of hydrocarbon dykelets in undisturbed cap rocks will be restricted to rocks of very low permeability and, moreover, that gas dykelets will have a better chance to develop than oil dykelets. This preference for gas as intrusion fluid might be particularly pronounced in shaly cap rocks with a high content of organic matter relative to which oil would be the preferentially wetting phase (Goff 1983). So far we have considered the upward growth of dykelets. The conditions discussed above for the formation of open dykelets also apply to the downward growth of dykelets in a low-permeability rock that separates a mature source rock from an underlying reservoir rock (Fig. 5). Since, however, overburden stress and horizontal stresses tend to increase with depth, while the hydrocarbon fracture pressure remains limited by the overburden stress at the source rock base, the penetration depth of downward growing dykelets will be rather limited. For the same reason, hydrocarbon sills cannot develop at all inside the low-permeability substratum. Determination of the penetration depth of downward growing dykes is complicated by the history of the fluid pressure in the barrier bed. It might, for instance, be quite common that a shaly barrier rock, directly underlying the source rock, is still compacting by dewatering downward into highly permeable drainage beds when maturation starts in the basal part of the source rock. These are at any moment the horizontal total stress in the compacting shale will be related to the water pressure by:

an(Z) = g'n (z) +p(z)

As an illustration in Fig. 6 we consider such a shaly barrier bed which, at the time of source rock maturation, is still under a very high pore water pressure (2=0.95). The pore pressure profile (1) is conjectural and intended to illustrate the pressure distribution in the compacting shale at the moment a hydrocarbon 'blanket' of sufficient saturation has been formed at the base of the source rock to interrupt inflow of water from the source rock. We assume that the maximum capillary pressure pC (2 MPa) is sufficient to prevent bulk flow of the hydrocarbon phase into the subjacent shale, and to allow further draw-down of the water pressure until a static pressure distribution is approached (2). This requires that the pressure in the drainage bed is maintained at a high level. If this were not so, drainage flow would continue and lower the water pressure at Za below Phc--Po and initiate hydrocarbon bulk flow into and, possibly, through the shale. From the aH profile, determined by (24), one may immediately read-off the maximum depths of fracture penetration for gas and oil in a shale without tensile strength. Typically, the depth values amount to a few metres only and these would be less if a tensile strength (TH < 1 KoPo) had to be taken into account. After this discussion of the various restrictions of dykelet formation it should be emphasized that the low-permeability rocks were assumed to be in a tectonically undisturbed state. Large-scale formation of hydrocarbon dykes may be drastically advanced by a regional or local tectonic extension that reduces the horizontal stress a°i~ and, consequently, the value of the parameter Ko. We may expect that strong rocks, such as strongly lithified shales, will respond with a drastic drop in/Co when subjected to horizontal extension. In contrast, clays which have not undergone overcompaction and lithification will respond to horizontal stretching at a low rate with only a very minor reduction in/Co. Practically no such reduction may be expected when clay, which is still in the process of vertical compaction, is extended in a lateral direction at a rate comparable with the rate of vertical compaction. The dykelet-promoting effect of a reduction in

(24)

Xo)p(~).

= K o ~ v ( z ) + (1 -

OVERBURDEN --- ~

Z

I t

;

/

/

/

/

/

/

/

/

/

/

OVERBURDEN

/

/

~

---

TOP SEAL--- ~ Y / / ' ~ S

SOURCE ROCK

....

S-C-S:= SHALE - - - 7---3

45

SOURCE ROCK

__~

-----_--~'- 7 SHALE-_- : - ~

RAIN AGE

~ @ _ - _ <-~:

~ -- ~ i ~

~ :

:~2 ~

DRAINAGEBED i L >

FIG. 5. Downward migration into reservoir rock separated from source by very low permeability layers.

G. M a n d l & R. M. Harkness

46 P(MPa)

]

"

~:

Pc

2F-----'~ M;S AXP D 'E E NP ETR HA\G TO I N\

~c

~F R \~OILA ~-~-C T UG PR AE SS ESU \R ~EIN

P,z, ,STAGNA,,\

\P(z, '? \

\

!

v

20 . . . .

SOURCEROCK

DRAINAGEBED (HIGHPOREPRESSURE) Z(m)

\

\ \.

i

i

t

2

\

~v(z)

i

1

3

4

Ko=0.6,TH=O

~'oil--0.8g/cmB

AT Z2: ° ' v - P = S M P a Pc=2MPa

~'gas = 0 . 2 g l c m 3

5

MPo

( X = 0 . 9 5 AT Z I = 4 k m )

FIG. 6. Stresses in compacting highly overpressured shale bed which separates source rock from subjacent drainage beds. (Stagnation of downward flow by hydrocarbon accumulation at source rock base.) the value of Ko is also clearly expressed in conditions (15), (16) and (17) for maintaining open dykelets, and by condition (23) which formulated the impeding effect of the porepressure rise induced in the cap rock by maturation of the source material. A sufficient reduction of Ko may even render the right sides of expressions (15) and (23) zero or negative, indicating that the formation of open dykelets becomes independent of overburden thickness and capillary pressure.

The dyke-dominated process The fracture conditions dealt with so far are not concerned with very local, small-scale changes in the in situ stresses caused, for instance, by bedding plane irregularities or by local differences in compaction. Such perturbations may result in very local reductions in bedding-parallel compressive stresses which will locally facilitate the formation of short discordant intrusion fractures. The role of such minor dykelets will be discussed later. Here we are still concerned with the operation of systems of major subvertical intrusion fractures which, by traversing cap rocks, would play a dominant role in the migration process. Assuming that the tectonic stress regime, capillary displacement pressure and initial overpressuring of a cap rock satisfy the conditions for preferential dyke formation, we may apply a few elementary results of fracture mechanics to obtain

a picture of the actual process. Approximate expressions for the growth rate of dykes, their width and spacing are given in Fig. 7. The width of a fracture, which is not hindered by neighbouring fractures, is proportional to the excess ofphe over the horizontal total rock stress (ri~ prior to fracturing (22), and proportional to the fracture length and the reciprocal of Young's modulus. Typically, a few bars of excess pressure in a fracture 10 m long would produce an average fracture width of about 1 ram. The rate at which an intrusion fracture grows is mainly controlled by the rate of fluid supply to the fracture tip. If the injection pressure Phc (O) at the cap rock base (Fig. 7) is maintained for a sufficient period, an approximate expression for the growth rate of an intrusion fracture is obtained by combining the law for steady laminar flow of a viscous fluid between parallel plates with the expression for the average fracture width. The growth rate is then proportional to the fracture length and to the cube of the excess pressure in the fracture, and inversely proportional to the squared Young's modulus of the rock and to the viscosity #ho of the fracture fluid. Inserting into the formula for the growth rate (Fig. 7) some typical values, say, E = 5 x 104 Mpa, ~Uhc=3 cP andph~ -- o n = 1 MPa, the fracture would increase in length by one-tenth in about one second. Obviously, the hydrocarbon supply from a mature source rock will hardly keep pace with such rapid growth to allow fractures to propagate continuously over several metres. Therefore, as men-

Hydrocarbon migration by hydraulicfracturing d max

t

47

~;

d

--

p (o) hc

RATE OF GROWTH OF VERTICAL INTRUSION FRACTURES: 3 V ' ~ ( Phc- °'H)3 L / ('r/hcE2) AVERAGE WIDTH OF THE FRACTURES W'-'3 ( Phc- °'H) L / E ( GEERTSMA 8~ DE KLERK, t969 ) FRACTURE SPACING 3 L d max=i - ~2

FIG. 7. The growth of hydrocarbon dykes. tioned before, dyke growth will be rather spasmodic. The shortening AeH between two parallel fractures immediately after a first phase of fracture growth will be equal to fracture width W divided by the fracture distance d (Fig. 7). Since the fracture width is propo~'tional to the fracture length, but AeH is independent of it, the fracture spacing will be proportional to the fracture length. Therefore, the process of large-scale development of hydrocarbon dykelets may be envisaged as follows. The fractures start at the cap rock base, closely spaced and initiated at points of stress concentrations such as bedding plane irregularities. Naturally, the fractures will differ in length with the longer fractures growing faster than the shorter ones and the shorter fractures being stopped by compression of the rock ahead of them (Fig. 7). Hence the hydrocarbons will enter the cap rock in a diffuse way through many fractures. But, while proceeding through the cap rock, the fractures will decrease in number and increase in width. Eventually they may reach the top of a thick cap rock at widely spaced discrete spots. The downward branching of the Gilsonite dykes in the Altamont Bluebell field of the Uinta Basin, as described by Lucas and Drexler (1975), might serve as an example for this concept. Again, gas dykelets will have a better chance of growing than oil dykelets because of the lower viscosity of the gas and the higher average fracture pressure associated with the lower specific weight. At the same injection pressure, gas dykelets will therefore exceed oil dykelets both in width and rate of growth. One might

therefore easily envisage situations where, in a thick cap rock, gas dykelets stop the migration of oil. Another interesting aspect of the process is its termination when the hydrocarbon supply declines and the hydrocarbon pressure drops. Consider first the case when the hydrocarbon dykes have not yet traversed a cap rock whose horizontal stress all, before dyke formation, decreased in the upward direction (Fig. 8A). In general, the rate of this decrease will exceed the rate at which the pressure of the fracture fluid decreases in the upward direction. For convenience we assume that both gradients are constant. When the hydrocarbon pressure has dropped sufficiently at the cap rock base, the dykes will close at the base (Fig. 8B). Actually, this requires that Phc attains a value somewhat smaller than aH to overcome the leverage of the excess pressure in the fractures, which allows fractures to extend a certain distance into a region where aH >Pho. Although the hydrocarbon pressure inside the fractures will have decreased and will have caused a reduction in aperture, the maximum capillary entry pressure may still be sufficient to prevent complete closing and concomitant intake of the hydrocarbon fill by the wall rock. If the fractures are of sufficient length, the difference in hydrostatic pressure gradient of the fracture fluid and the gradient of an will induce tensile stresses in the upper tip regions. These will break the interparticle bonds of the rock and allow the fractures to extend in the upward direction. Concomitantly, the fluid pressure inside the fractures will decrease and cause further closure

48

G. Mandl& R. M. Harkness \

! ---

0

0

CLAY

\

(A) Source Rock

'" '~. . . . . . . . . . . . . . . . . . . . . . .

(B)

-~ pd-

o

Phe

,%

=

d~21:5

Ig, -,:m,, i

Frocture

(c)

FIG. 8. The upward migration of hydrocarbon-filled

fractures. of the fractures at the lower end (Fig. 8C). In a cap rock with uniform strength and constant gradient of o"H we may expect the upwards migrating fractures to maintain a certain length L¢ which is just sufficient to overcome the tensile strength of the rock ahead. In fracture mechanics the tensile strength of the material is characterized by the so-called 'fracture toughness' (or critical stress intensity Kc). In order to overcome the strength of the material and to rise, while at the same time closing at its end, the fracture has to have the critical length:

L~-~ '/3 L@/dz dz J as has been discussed by Secor & Pollard (1975) in studying the stability of open hydraulic fractures, following earlier work by Weertman (1970) on the migration of water-filled crevasses in glaciers. Secor & Pollard assumed Ko = 100 MPa mm 1/2 for granite. Unfortunately, Kc values for normally consolidated clays and stronger shales are not yet available, but guided by reported values for weak carbonate rocks it might seem reasonable to expect Kc values of a few MPa mm ~/2 at the most for normally consolidated clays with no, or only very weak, permanent bonds. For a gas dykelet in an overpressured clay a typical value for the difference in gradients would be about 2 x 10 -5 MPa

m m - 1 and Lo would then become 1.3 x ge 2/3 m where Kc is expressed in M P a m m 1/2. For Kc = 2 MPa mm 1/2 a gas dykelet with a length of slightly more than 2 m would migrate towards the cap rock top as shown schematically in Fig. 9. Having traversed a cap rock, an upward migrated dykelet will leave a trace of zero tensile strength (TH=0) behind (Fig. 8C). Therefore, any restoration of the hydrocarbon pressure to the level of 17u at the cap rock base will discharge hydrocarbons into the cap rock. These then travel upwards along the fracture trace as thin sheets of arbitrarily short length (K~ = 0 in equation 25). Similarly, TH vanishes along dykes that extend from the base of a cap rock to its top. Again, closing of such dykes caused by a drop in hydrocarbon pressure at the cap rock base will proceed from the base upwards and the dykes will empty into strata overlying the cap rock. Any hydrocarbon 'pockets' left behind would be driven upwards along fracture planes, irrespective of the pocket length. Likewise, the repetitive drop and build-up of the hydrocarbon pressure that is likely to occur in the dyke-dominated migration process will cause individual thin hydrocarbon pockets to travel upwards along the paths of zero tensile strength that have been established by the first dykes that have broken through to the cap rock top. Considering how dykelets close when the hydrocarbon supply has ceased, it is not surprising that evidence for migration through systems of hydrocarbon dykes is rare. It should also be recalled that the spacing of ascending dykes increases with dyke length. Therefore, in a cap rock of sufficient thickness the spacing may eventually approach the value dmax (Fig. 7) for parallel dykes which do not interfere with each other. Using the formula for the width W of independently growing fractures (Fig. 7) and Hooke's law to describe undrained compression of the rock between opening fractures, the distance dmax between such non-

T.>T~O

|/ :::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::Phc=Gv o :::::::::::::::::::::;::::~:~:~:::!: :::: :~:~: :i:~!~Zi~:~:T:;::~:!i!!!iiiii:::::: :i~~{:i~iT~:~:~::::.:.:.:8~~i~i~i~iii~i:i:~ ...... T:~:i:~:~:??'-'-:.

FIG. 9. The 'sill-dyke' mechanism (not to scale).

Hydrocarbon migration by hydraulicfracturing interfering fractures is found to be about three times the fracture length (Fig. 7). It would therefore seem possible that some reservoir rock embedded in the upper part of a thick overpressured shale sequence is by-passed by hydrocarbon dykes.

The sill-dyke mechanism In beds where the horizontal total stress o-H exceeds tTv-Tn large-scale dyke formation is impossible. Only sill-type fractures can be formed, provided the hydrocarbon pressure equals the total overburden stress. Such a situation may arise quite often as, for instance, when the low-permeability beds were in fluid/fluid contact with source beds from which they received substantial additional overpressuring during the maturation process. Then the capillary pressure Pc and the degree of 'primary' overpressuring (2 > 0.45) would not have satisfied condition (23). Another tectonically undisturbed environment where sill-type intrusion fracturing may predominate exists in erosionally unloaded areas where unloading has left high horizontal residual stresses locked in the upper few hundred metres of overburden. Bedding plane intrusion of hydrocarbons should also form in the toe regions and along slope-parallel glide planes on delta slopes (Fig. 10A), in the limbs of compressional folds and, of course, in other compressional tectonic settings where the overburden stress is the smallest compressive stress. As long as hydrocarbons are supplied from the source rock, lateral spreading of a bedding plane sill can only be stopped by lithological barriers, by an increase in the total rock pressure on the bedding plane, or by leakage from the sill into the cap rock. In this respect a hydrocarbon sill differs from an igneous sill whose lateral expansion is stopped by solidification of the fracture fluid.* Naturally, spreading will be preferentially in the direction of decreasing total stress on the bedding plane, which usually will coincide with the direction of decreasing overburden. Leakage from a hydrocarbon sill into the cap rock will take place when the capillary pressure is not balancing the difference between the overburden carrying hydrocarbon pressure and the water pressure in the roof of the sill, be it because the maximum capillary pressurepc still has to be built * It would therefore not seem appropriate to calculate the shape of the roof of a hydrocarbon (or water-) sill and the associated stresses by applying the model of a thick elastic plate, clamped to a horizontal rigid base and pressurized from underneath, as has been applied by Johnson (1970) to magmatic intrusions.

49

up (Fig. 2) or that the water pressure has to be raised by the invasion of hydrocarbons. The growth rate of an expanding sill will, of course, slow down since the rate of cumulative leakage losses increases. However, in view of the extremely adverse drainage conditions for a thick low-permeability cap rock and the low saturation at which a non-wetting fluid enters this rock, the rate of leakage will remain small enough to allow bedding plane migration over considerable distances. Obviously a quantitative estimate would require data on the rate of hydrocarbon supply from the source rock. The next question is: will hydrocarbon migration essentially remain confined to bedding planes, or will the fluid break through to stratigraphic higher levels and form stacks of sills similar to the stack of water sills postulated by Price (1978)? As illustrated in Fig. 9, such a step-up to higher bedding plane sills would require the formation of interconnecting dykelets and is made possible by local reduction of the horizontal (beddingparallel) compressive stress. A first mechanism by which this can be accomplished has been proposed by Price (1977) with respect to the rise of hydrothermal water. Horizontal bedding planes under a non-uniform overburden, or dipping beds, will in general experience some shear stress. When a fluid sill forms along such a bedding plane the effective compressive stress across the bedding plane will be reduced by the pressure of the injected fluid which, according to Coulomb's friction law, implies a proportional decrease in shear strength. Consequently, the beds may no longer be able to carry the full shear stress, and the bed-parallel normal stresses will have to change in order to maintain static equilibrium. This is shown schematically in Fig. 10 (A, B) for gently dipping beds on a long delta slope. On such slopes the sub-horizontal normal stress increases downslope and over a wide distance may exceed the overburden stress. This also holds along a glide base on a slope ( M a n d l & Crans 1981). A maturing source rock may, therefore, be expected to produce hydrocarbon sills in this environment. The local reduction in shear stress z by the high-pressure hydrocarbons will cause a drop in bed-parallel rock pressure o-ii at the upslope part of a sill, which is compensated for by an increase of this stress in the down-dip direction to maintain slope stability (as long as critical stress conditions are not reached). The reduction in all will allow dykes to form which crossbeds to feed new sills at stratigraphically higher levels. In this way a process of upslope silldyke migration may be envisaged (Fig. 10B).

5o

G. Mandl & R. M. Harkness EXTENSION COMPRESSION

"-~--~~l/~ ~

.. ,',','-'-- ','~ ~,,,,,,,,,,,,,,,,,,\,,\,,~'~/

(A)

- u MA X

~ j~____~

g

~

"C'z-----

~

Phc Pwoter(H) >7''H cos

( 7 " SUBMERGED SPECIFIC WEIGHT OF BULK MATERIAL)

(B)

", 'z-~z~/

,'//

- . -. . -. . -. . -. . -. .~ . .S . .- . .Z . .- . . E . -. . -. . -. . S. . -. .-. .- . .- . .- . .S . .- . . . . -. . -. . _. . -. . . . ~ ; _ ; ; ; ; ; ; ; _ ; ; ; ; _ ; _ ; ; ; _ ~ ; ; _ - ; ; ; ; ; ; ; ; _ - ; _ - ; ; ; _ - _ ;;_;;;;;;-_;; .

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.

.

.

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.

.

SHALE

~

SAND

: -_-_--..-_-..-_-_-_-_-_-_-_-_-_-_-_-_-_-_-_-_-_-_-_-_-_--_-_-_-~ . . . . . . . . . . . . . . . . . . . . . . . . . .

~:::~-,~: :,:::::::

: : : : : : : : : :~', ~ , ,

:

.

.

~

.

.

.

.

.

.

.

~ .

.

.

~ . .

.

~ . .

.

-

.

.

.

:÷~,,i,,!::,:,:,',',"~','~ _ ~ - ~ , ' - , ~ ~ ~ ~

.

.

.

-

-_-

.-_-

-_-_-

.

- ~

(C) S P = STAGNATION POINT L

d = 'a' FOR LONG STRIP (: WIDTH / 2 )

\ 1~

\1' \ t

FOR o,sc \

(o -- RADIUS)

'X//\d~ HYDRAULIC HEAD " X~\ O N L HYDRAULIC HEAD X ~ j OF U O,S U B O

\\

FLOW

L

(D)

HYDRAULIC HEAD CAUSING FLOW L

FIG. 10. Formation of hydrocarbon dykes in sill-promoting stress regime. (A) Extension and compression on slope. (B) Long slope with static pore water; updip migration by 'sill-dyke' mechanism. (C) Extension by differential compaction. (D) Disturbance of hydraulic potential by hydrocarbon sill.

Hydrocarbonmigration by hydraulicfracturing When the fracture system enters extensional environment higher up on the slope, dykes may become the dominant fractures, interconnected by short sills. Another mechanism which may locally reduce bed-parallel compressive stresses to allow dykelets to branch off from hydrocarbon sills is differential compaction. Layers may be locally stretched by this, as depicted in Fig. 10(C). The process should operate particularly well in compacting shales with interspersed sand lenses. Finally, we consider a horizontal hydrocarbon sill inside an overpressured compacting shale with steady drainage flow in a vertically upward direction (Fig. 10D). The sill has the shape of a strip with half-width a and the capillary displacement pressure p~ of the system is assumed to be of sufficient magnitude to prevent any water flow across the sill. The water flow will then be disturbed by the presence of this fiat barrier and will completely stagnate at its centre (S.P. in Fig. 10D). Associated with this stagnation is an increase in hydraulic potential on the lower side and a decrease on the upper side of the sill. In undisturbed Darcy-flow of the pore water along the vertical z-axis, the hydraulic gradient would be: 1 dp

Vho= 1 + - . - -

7f dz

(26)

where ho = z q-p/yf (yf= specific weight of pore water). From potential theory we take the result (Milne-Thomson 1949) that the jump [hi in hydraulic potential across the fiat barrier at the stagnation point (S.P.) is [h] = - 2Vho. a.

(27)

Because of (26) the associated jump in water pressure is: [p]= - 2 a • (]~f+ ~Pz).

(28)

In highly overpressured shales the pressure gradient of vertical drainage flow often approaches closely the gradient of overburden stress - ) , (~, is the average specific overburden weight). The jump in water pressure across the centre of the hydrocarbon sill then becomes [p]=2a • (~-~f)--~0.3a(m) • 103 kPa.

(29)

If the hydrocarbon sill has a disc shape with radius a, the parameter a in the formulae above has to be replaced by (2/n)a. According to equation (29), a strip-shape sill with half width a = 50 m would produce at its

5I

upper side a central pressure draw-down of 750kPa and therefore increase the effective overburden stress by the same amount. The associated compaction is accompanied by an increase in effective horizontal normal stress of 750 Ko kPa while the total horizontal stress decreases by (750-750 Ko) kPa. Hence large hydrocarbon sills in combination with high capillary pressure and high overpressuring of a compacting shale could sufficiently disturb the drainage flow so as to facilitate the formation of dykelets. Although the detailed dynamics of the stepping-up of hydrocarbon sills are not well known, the foregoing discussion may justify the conclusion that even under conditions that exclude large-scale dyke formation, sill-type intrusion fractures produced by the hydrocarbons expelled from a mature rich source rock may ascend to the cap rock top. One may, of course, envisage (exceptional) situations where a fairly uniform thick cap rock, as yet unfractured, is subject to such all-round tectonic compression that even the smallest horizontal stress is .so far in excess of the overburden stress that the formation of dykelets is practically suppressed and hydrocarbon flow will be confined to one or a few bedding planes near the cap rock base. It may be also reasonable to expect that the more the growth of secondary dykelets is hindered by high horizontal stresses, the larger will be the hydrocarbon volumes stored in the sill system of the cap rock before breakthrough. A further point worthy of note is the dip of bedding planes. Although our discussion has been restricted to horizontal, or only slightly dipping beds, it is fairly obvious that in a non-compressional regime a more pronounced dip of cap rock will facilitate the updip formation of bedding plane sills. This is simply because the total normal stress tr± across such planes is smaller than the overburden stress. Less obvious is the possibility of hydrocarbon intrusion into dipping beds in a compressional regime, as sketched in Fig. 11. Both cap rock and source rock beds dip in the same direction and are compressed perpendicularly to their strike by a stress trH > try. The tensile strength across bedding planes, both in cap rock and source rock, is again negligibly small with respect to the tensile strength across any other plane. Without entering the complex field of hydrocarbon generation and transport inside a mature source rock, we may assume that the anisotropy in tensile strength inside the source rock will, in the present case, allow the hydrocarbon pressure inside the source rock to rise above the overburden stress o'v. The Mohr circle construction in Fig. 11 then demonstrates that

G. Mandl & R. M. Harkness

52

( O-H - crv ) sin2~

% TENSILE STRENGTH ZERO ALONG BEDDINGPLANES; ON ANY OTHER PLANE T>O

%>~,, > % > %

EXCLUSIVE FORMATION OF BEDDING PLANE 'SILLS' REQUIRES:

o'j.<- Phc < O"v + T WHICH IS ONLY POSSIBLE IF: (O'H- o"V )

sin2~<

T

FIG. l 1. Hydrocarbon intrusion into dipping beds under horizontal compression. sills will form along the dipping cap rock beds as long as the dip angle does not exceed the value defined in the figure. Once a sill-dyke system has penetrated a cap rock it will experience a drastic reduction in fracture width and lateral extent of the sills. Since the pressure in the sills at any level is at least equal to overburden stress, the interconnecting dykelets have maximum opening. It may be safely assumed that the associated permeability of the fracture system would require, after breakthrough, unrealistically high rates of hydrocarbon supply. Both sills and dykelets will therefore adjust their widths--and, most likely, almost close--to reduce the permeability of the fracture system to bring the available hydraulic head and the rate of hydrocarbon supply in balance.

Conclusions In general, hydrocarbons can only have migrated through an intact, preferentially water-wet, thick barrier rock of very low (or zero) permeability by fracturing the rock. Although such intrusion fracturing by highly pressurized hydrocarbons is more complex than expected at first sight, and a quantitative description of the process is impeded by many uncertainties regarding the geological and rock-mechanical parameters, some general trends emerge from our discussion. The most efficient mode of penetrating a rock barrier of very low permeability would seem to be the formation of a system of hydrocarbon

dykelets, fed through many small 'root' dykelets at the source rock contact and joining into branches which grow in width and spacing with increasing distance from the source ('invertedtree' model). Almost certainly less hydrocarbon volume will be required in this process to traverse a cap rock than when sill-type intrusions are also part of the process. Assuming that the hydrocarbons, expelled from a rich mature source rock, are overburden bearing, one might expect major dykelet systems to develop in most tectonically undisturbed cap rocks. However, large-scale dyke formation is severely restricted by two processes: Fluid/fluid communication between cap rock and maturing source rock very likely will raise the water pressure significantly in the basal part of the cap rock. This may be associated with an increase of total horizontal stresses beyong the overburden stress and hence obstruct dyke formation. When dykes have been initiated, the hydrocarbon pressure in the dykelets will cause the water pressure inside the low-permeability wall rock to rise. This in turn will tend to close the fractures again. It has been shown that for given values of the capillary displacement pressure and the pore pressure/overburden ratio before onset of maturation, open dykelets will be restricted to cap rocks carrying overburdens of a limited thickness. Open dykes at depths of practical interest require high overpressuring of the cap rock and/or capillary displacement pressures of several MPa. While,

Hydrocarbon migration by hydraulicfracturing however, the closure of dykelets in 'soft' clayey rocks may be complete, 'closed' dykelets in brittle rocks may still serve as migration paths. Horizontal tectonic extension will promote dyke formation in strong rocks such as strongly lithified shales, while it may have little effect on compacting clays. For various reasons (higher capillary pressure, lower viscosity and specific weight) gas dykelets have a better chance of growing than oil dykelets and might, in certain situations, even stop oil dykelets from growing. When the hydrocarbon supply and pressure decline, the dykes will close from the base upwards, leaving little trace of their activity. Downward migration through a barrier rock by hydraulic fracturing cannot involve sills. The penetration depth of downward-growing dykelets depends on the fluid-pressure history of the barrier rock and underlying drainage beds. For highly overpressured barrier rocks it was argued that the penetration depth will hardly exceed a few metres.

53

Where large-scale dyke systems cannot form, overburden-bearing hydrocarbons will form bedding plane intrusions. These sills will not only allow updip migration along bedding planes, but may 'step up' to stratigraphically higher levels by the formation of small interconnecting dykelets. Various mechanisms for such step-up have been discussed. Qualitatively speaking, the hydrocarbon volumes required to establish final breakthrough of a sill-dominated fracture system through a thick cap rock may be much larger than in a dyke-dominated system. The cap rock thickness may therefore become a controlling factor in sill-type migration, while it would hardly seem of relevance in dyke-dominated migration. After having traversed a cap rock, any dyke or dyke-sill system will drastically reduce the fracture width to adjust the permeability of the system to the low rate of hydrocarbon supply. ACKNOWLEDGEMENTS: The authors gratefully acknowledge the stimulation received from the work of Neville Price, and the discussions with him.

References ANDERSON, E. M. 1942. The Dynamics of Faulting. Oliver & Boyd, London. DALLMUS, K. F. 1955. Mechanics of basin evolution and its relation to the habitat of oil in the basin. AAPG Symposium 'Habitat of Oil" (ed. G. WEEKS), 1958, pp. 883-93. DU ROUCHET, J. H. 1981. Stress fields: a key to oil migration. Bull. Am. Assoc. Petrol. Geol., 65, 7485. GEERTSMA, J. & DE KLERK, F. 1969. A rapid method of predicting width and extent of hydraulically induced fractures. J. Petr. Techn. Dec. 1969. GOFF, J. C. 1983. Hydrocarbon generation and migration from Jurassic source rocks in the E. Shetland Basin and Viking Graben of the northern North Sea. J. geol. Soc. London, 140, 445-74. JOHNSON, A. 1970. Physical Processes in Geology. San Francisco. LACHENBRUCH, A. H. 1962. Depth and spacing of tension cracks. J. geoph. Res. 66, 4273-92. LUCAS, P. T. & DREXLER, J. M. 1975. Altamont Bluebell: a major fractured and overpressured stratigraphic trap, Uinta Basin, Utah. In: North American Oil and Gas Fields. Mem. Am. Assoc. Petrol. Geol. 34, 121-35. MANDL, G. & CRANS, W. 1981. Gravitational gliding in deltas. In: MCCLAY,K. R. & PRICE, N. J. (eds.) Thrust and Nappe Tectonics, Spec. Publ. geol. Soc. London, 9, 41-54.

MILNE-THOMSON,L. M. 1949. TheoreticaI Hydrodynamics, pp. 167, 477. PHILLIPS,W. J. 1972. Hydraulic fracturing and mineralisation. J. geol. Soc. London, 128, 337-59. PRICE, N. J. 1966. Fault and Joint Development in Brittle and Semi-Brittle Rock. Pergamon Press, Oxford. • 1977. Aspects of gravity tectonics and the development of listric faults. J. geol. Soc. London, 133, 311-27. - - , 1978. Fluids in the Earth's Crust, by FYFE, W. S., PRICE, N. J. and THOMPSON, A. B., Elsevier, Amsterdam. SCHNAEBELE, R. 1948. Monographic G6ologique du Champ P6trolif6re de Pechelbronn, Strasbourg. M~m. Carte Gbol. Alsace Lorraine, No. 7. SECOR, D. T. 1965. Role of fluid pressure in jointing. Am. J. Sci. 263, 633-46. - - & POLLARD,D. O. 1975. On the stability of open hydraulic fractures in the earth's crust. Geoph. Res. Lett. 2, No. 11. SNARSKY,A. N. 1962. Die primi~re Migration des Erd6ls, Freiburger Forschungshefie C 123 (with references to earlier work by the same author). WEERTMAN,J. 1970. Can a water-filled crevasse reach the bottom surface of a glacier? In: Proc. Cambridge Syrup. on Hydrology of Glaciers, Cambridge, England, September 1969.

G. MANDL& R. M. HARKNESS,Koninklijka/Shell Exploratie en Produktie Laboratorium, Rijswijk, The Netherlands.

Deformation, diagenesis and the mechanical behaviour of chalk C. R. I. Clayton & M. C. Matthews S U M M A RY: The chalk of England is a highly variable material resulting from numerous deformation processes of which the most important are diagenetic or tectonic in nature. The engineering behaviour of the chalk is influenced greatly by these processes thus giving rise to the necessity of carrying out studies of physical properties to aid the prediction of engineering performance. The results of these studies are of value when examining deformation mechanisms associated with diagenesis and tectonism. In this paper use is made of results obtained from engineering laboratory index tests to study deformation mechanisms associated with early diagenesis and the influence of overburden and tectonism.

In terms of its porosity, the chalk of England is a highly variable material. This variability results from a wide variety of diagenetic mechanisms, together with structural deformations associated with tectonism, and may be related to stratigraphic horizon or geographic location. The engineering behaviour of the chalk is influenced greatly by these deformation processes and has thus given rise to studies of porosity, plasticity, rate and amount of consolidation, hardness, and strength. The major objective of such studies is to enable the prediction of the performance of the material in such primary areas of civil engineering interest as earthworks, foundation design and the stability of natural slopes and excavations. The results thus obtained can be of value, however, when examining the deformation mechanisms associated with diagenesis and tectonism.

Composition and early diagenesis of the chalk It is widely known that the chalk is a biomicrite, its principal calcite components being derived from Inoceramus prisms, foraminifera, shell fragments other than Inoceramus, Spheres, Cayeux's organisms and the remains of coccolithophoridae (Black 1980). The coccolithophoridae provide the finest fraction of the chalk. In most chalks this dominates the behaviour of the mass. The chalk is normally remarkably pure, with calcium carbonate accounting typically for 98.5% of its dry weight in the Senonian. In soft chalks almost all of this material is organically derived and consists of coccolith debris sized between 0.5 and 10 t~m. Clay minerals are rare, and the material is essentially granular and weakly cemented, although the bonds between individual

grains are very difficult to see on scanning electron micrographs of soft chalks. A considerable number of deformation and diagenetic mechanisms have been postulated for the chalk. These include consolidation (gravitational compaction), reworking by benthos, slumping of bioherms, chemical solution processes, alteration by meteoric water and high thermal gradients, and pressure solution and shear distortion induced by tectonism. At the time of its deposition the chalk was essentially a granular material, with the unusual feature of already having some grains bonded together as part of the structure of the coccoliths. Intact coccolith accumulations will form sediments of very high porosity, but few coccoliths survive intact the process of sedimentation and earlier burial, in which case the sediment achieves a much closer state of packing. The depositional environment is thus important in determining the initial porosity of a chalky sediment. Bioturbation will cause further breakdown of coccolith material (into its smallest fraction). The extent of this reworking would almost certainly have depended on the rate of deposition, such that where deposition was slow higher degrees of disaggregation and densification would be expected. The maximum porosity found (by the authors) in the English chalk is of the order of 50% (~d = 1.33 Mg m -3, msat= 38%--see Appendix) indicating that this process did not necessarily produce a very compact mass. This is especially so when it is considered that this level of porosity probably reflects the influence of relatively high effective stress levels imposed by the subsequent overlying sediment pile and (in some areas) by tectonism. Fig. 1 shows the results of a survey of the English chalk to determine its variability in terms of dry density, the results of which are also given in terms of porosity. It can be seen that, for

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 55-62.

55

C. R. I. Clayton & M. C. Matthews

56

•

0

D (M g/m 3)

Porosity (%)

2.5 Upper Bound

20 2.0

// I

/

&

~i~- - .-~.

~X/

I

II

........4....

1.5

IJ&X N \

Average

30

//

\\

j-"

.4"

Average 1 Std. Dev.

. + .........÷"

-

40

-

50

Lower Bound

../ '-...... :.I.,.Y

50 1.0

I

I

I

B. A. H. H. muc. ~lu. fesf. ca.

[

1

ii. of.

H. p.

I

I

I

T. Rh. H, Iota. c. sub.

I S. v.

ZoBe

FIG. 1. Variability of the chalk in terms of dry density (?d) and porosity (n).

convenience, the material has been divided into the zones suggested by Jukes-Browne & Hill (1903, 1904), and that the range of minimum dry density in the white chalk (i.e. the Turonian and Senonian) is smaller than the range of dry density within any one zone. This tends to suggest that consolidation in the sense applied by geotechnical engineers to cohesive materials was not significant in reducing the porosity of the majority of chalks. Consolidation involves the transient sharing of total stresses between the pore fluid and the mineral skeleton of a deposit. In essence, the excess pore pressures (i.e. pore pressures greater than hydrostatic) which are developed to provide some of the support for the applied total stresses cause a delay in the compression process (the 'hydrodynamic time lag') and result from the fact that a relatively fast rate of load application, a relatively low permeability material, or a relatively thick deposit will not allow the rapid escape of interstitial pore fluids (Terzaghi 1923). If no excess pore pressures develop, then drained compression, rather than consolidation, will occur. The compression of the deposit is then controlled by the rate of load application and the rate of compression of the material with respect to time at constant effective stress level, rather than the rate at which the excess pore pressure can dissipate. This process is sometimes termed

'secondary compression' by geotechnical engineers. If either consolidation or drained compression were the controlling processes in the chalk, as Carter & Mallard 1974 have suggested, then it would be expected that the minimum dry density of the deposit would increase systematically with increasing burial depth. Although such a trend has been observed in the lower zones of the chalk (Fig. 1) it is created by: (i) the occurrence of extensive hardgrounds in the Turonian chalks, which are known to be associated with penecontemporaneous hardening (Kennedy & Garrison 1975; Bromley 1968); (ii) the inclusion of significant amounts of clay in the Cenomanian chalks. The idea that consolidation (i.e. compression associated with overpressuring, or positive excess pore pressures in soil mechanics parlance), was not a significant contributor to the process of confined (Ko) deformation and densification of the chalk is supported by three facts. Firstly, it is widely understood in geotechnical engineering that once interparticle contacts are established in granular materials confined compression, such as would occur in laterally extensive deposits, leads to particle locking. This process is observed for sands (e.g. Lambe & Whitman 1969; Dusseault

Mechanical behaviour of chalk & Morgenstern 1979) where increases in vertical stress have been shown to have relatively little influence on porosity until stress levels rise to the point at which grain crushing occurs (e.g. Vesic & Clough 1968). Similar effects can be seen when one-dimensional (confined) compression tests are carried out on crushed chalk. Fig. 2 shows that for chalk slurry the porosity at any vertical effective stress level is a function of the initial porosity of the specimen, and that despite the very high initial porosities of the specimens, a unique porosity/effective stress relationship (the virgin compression curve) is not achieved even at pressures such as would be imposed by some 100 m of submerged sediment. Secondly, it is well known that some of the hardest and densest levels of the chalk, such as Chalk Rock, were formed whilst the material was exposed on the sea bottom. The intense bioturbation (Kennedy & Garrison 1975) and the adaptation of the benthos to seafloor lithification (Bromley 1968), which are seen in a hardground, probably indicate that similar processes would have had a significant influence even in locations where deposition was occurring at a much more rapid rate. The shear deformation and loss of internal restraint associated with burrows and burrowing would be expected to reduce locking and allow collapse into denser, more stable structures. It would also lead to progressive disaggregation of coccolith debris. This densification process was evidently also associated with the formation of authigenic calcite, which further acted to cement and density the deposit. The source of the calcite for initial cementing was probably from aragonite tests of macro and microfossils, but cementing may also have oc-

57

curred as a result of the remoulding process itself leading to the solution of small calcite particles and their redeposition as aragonite. This process has been observed in compacted chalk fill. Fig. 3 shows a scanning electron photomicrograph of a compacted Senonian chalk fill forming part of the embankment for the south-eastern sector of the M25 motorway. The presence of clearly defined aragonite crystals in the material (which showed no signs of aragonite before excavation) is a surprise, particularly since aragonite is the high temperature polymorph of calcium carbonate and would be expected to degrade to calcite fairly rapidly at the temperatures and pressures found in an embankment. This observation does, however, suggest that the activity of benthos could contribute to the availability of cement. Thirdly, the rates of primary consolidation observed in chalk are much higher than those observed in clays. Confined (Ko) compression tests in the oedometer, and similar tests on chalk slurries carried out on 250 mm diameter specimens, indicate coefficient of consolidation (Cv) values of the order of 40 m z yr- 1. On the basis of Terzaghi's consolidation theory (Terzaghi 1923) assuming a sudden penecontemporaneous increase in vertical stress applied to a 200 m thick deposit of soft chalk, 95% consolidation (i.e. an average of 95% dissipation of induced excess pore pressures throughout the layer) would take about 1000 years. This is a very small period in the geological time-scale and, moreover, it is small in relation to the rate of sedimentation for the chalk, which has been estimated as between 2 and 3 cm per 1000 years and 15 cm per 1000 years (Kennedy & Garrison 1975; Hakansson et al. 1974). The consolidation of chalk is a rapid

50 Void Ratio (e)

I.~iU/~~ ~ /.8

0.g

Initial moisture c o n t e n t

/

/

46

31%

9*/,

0.8 19%

9%

o

-

_

0.7 40

23%

0.6 10

I

100 Vertical Effective Stress (kN/m2)

FIG. 2. Compression of chalk slurry in the oedometer test.

38 1000

Porosity

(%)

58

C. R. I. Clayton & M. C. Matthews weak Senonian chalk from Needham Market, Suffolk (70 = 1.38 Mg m - 3, msat = 36% n = 49%). Two phases are apparent: (i) bedding and 'elastic' compression; (ii) yielding of interparticle bonds, and collapse to a more stable structure.

FIG. 3. Scanning electron micrograph of compacted chalk fill, from an embankment on the M25 motorway, showing original calcite, and authigenic aragonite. process, but a further factor, not seen in chalks, acts to slow down the consolidation of clays. The porosity and permeability of soft clays are both significantly decreased by an increase in effective stress, so that those parts of a thick deposit of clay near to the drainage boundary may act to 'seal' the deposit during the consolidation process thus trapping excess pore fluid and retarding the dissipation of excess pore water pressures in the centre of the deposit (see, e.g. Pane et al. 1983): As has been demonstrated (Fig. 2) this does not occur in chalk except where seafloor cementation occurs.

The vertical effective stress levelat which yielding and 'collapse' occurs should be greater than all past in situ vertical effective stress levels (if it is assumed that recent alteration and weathering have not significantly reduced the strength of bonding in the material). It can be seen from Fig. 4 that for the very soft chalk of Needham Market yielding and collapse occurred at about 1700 kN m-2. This is equivalent to the vertical effective stress imposed by a sediment pile of approximately 200 m thickness. The maximum past effective overburden pressure applied to this chalk can only be estimated by studying the likely stress history undergone by this material. This is complicated by the major periods of erosion that occurred in this area during late Cretaceous and early Tertiary times and during the Pliocene and Pleistocene periods. As far as can be reasonably deduced from the regional stratigraphy and palaeogeography, the stress history may be divided into the following stages: (i) Deposition of some 110 m of chalk above the zone ofActinocamax quadratus (Boswell 1927). (ii) Uplift and peneplanation of chalk in late Cretaceous and early Tertiary times, leading Vertical Effective Stress 100

Deformation under the influence of overburden, and tectonism Since the strength of soft white chalks derives from bonds at particle contacts (probably caused by calcite overgrowths of the original grains) the breakdown of the bonds might be expected to lead to significant densification of the deposit. This would occur at relatively high effective normal stress and shear stress levels compared to those acting on the material as it lay close to the seafloor. Since very soft chalks are still to be found, apparently having undergone little change, it is interesting to compare the maximum effective stresses applied by present and past levels of overburden with the 'collapse' pressures obtained from laboratory tests. Fig. 4 shows the result of a one-dimensional compression test on an intact specimen of very

Iv~vv~~Im~

(kN/m ~} 10000 I

'I \, 0 05

0.10

\

/

\ \ \

015

\, ,\ \ \

Vertical Strain

\

FIG. 4. One-dimensional consolidation on soft intact chalk from Needham Market, Suffolk.

Mechanical behaviour of chalk to the removal of about 100 m of chalk at this locality. (iii) Deposition of Thanet sands, Reading Beds, and London Clay amounting to about 240 m of sediment. There is no evidence to suggest that any material younger than the London Clay was deposited in this area during the Palaeogene period. The presence of such material would, of course, have a significant effect on the estimate of maximum past effective stress. (iv) Erosion of early Tertiary sediments and the deposition of some 30 m of Pliocene and early Pleistocene Crag Sediments. The Red Crag deposits in the Ipswich area are known to overlap on to the Chalk, suggesting the complete removal of the early Tertiary sediments. The relatively shallow water environment of deposition of the Crag deposits suggests that it is unlikely that the total thickness of these much exceeded the maximum preserved thickness of about 30 m. (v) Erosion of the Crag deposits and the deposition of glacial sediments during the Pleistocene. The problem here is in estimating the full thickness of ice in this area and whether the full thickness of ice and sediment load should be considered in view of the evidence of the existence of large subglacial streams in this region (Woodland 1970). The maximum thickness of boulder clay in this region is between 30 and 60m. The ice sheet together with its sediment load would have had to be equivalent to a sediment pile of between 170 and 200 m thick in order to cause an increase in vertical effective stress above that imposed by the early Tertiary sediments. The above stress history is highly simplified and takes no account of tectonism. It will be seen from this, however, that the most significant event giving rise to the maximum past overburden pressure was the deposition of the early Tertiary strata. The estimated 240 m of early Tertiary sediments would have led to a maximum vertical effective stress of about 2000 kN m -2 on the chalk under consideration. This estimate of maximum past effective overburden pressure is based on the average thickness of the preserved Tertiary and Quaternary sediments in this area and an ice thickness of about 150 m. Further, it assumes that overpressuring in the Actinocamax quadratus zone of the chalk was either insignificant or did not last for long enough to prevent compression. Had a greater thickness of Tertiary sediment and Pleistocene ice been present at this

59

locality, then the actual maximum past effective overburden would have been greater than the estimate given above, and hence in the absence of overpressuring this estimate (i.e. 2000 kN m -z) must represent a lower bound value. Had significant overpressuring occurred over a long period of geological time, this would have had the effect of reducing the actual maximum past effective overburden pressure. If this above estimate of maximum past effective overburden pressure is compared with the collapse pressure observed in Fig. 4, it will be seen that the values are of similar magnitude. This may, therefore, indicate that the present density of the chalk at Needham Market has been brought about by the maximum past effective overburden pressure applied to it and that this pressure caused a densification above its initial value. Furthermore, if the estimate of maximum past effective overburden pressure is correct, it suggests that if overpressuring occurred it did not last long enough to prevent full compression of the sediment under the load of the overlying strata. Although it would seem that drained compression under the weight of the superimposed sediment pile may have been the factor controlling porosity at Needham Market, it should be remembered that this is chalk which is far from typical. It has the highest porosity found for the chalk in England, and has been relatively unaffected by seafioor diagenesis and tectonism. Some further light may be thrown on these deformation problems by observing the influence of tectonism on the properties of the chalk. The influence of tectonism on the hardness of the chalk has been noted by Strahan (1898) and more recently by Mimran (1975) and Clayton (1983). Strahan noted that in areas of high dip in Dorset the hardness of the chalk increased, and Mimran carried out a survey in which he correlated measured bulk densities of intact chalk with angle of dip for the Dorset chalk. Steeply dipping chalks also occur on the Isle of Wight, and in the area of the Hog's Back in Surrey. Sampling locations in the West Surrey area are shown in Fig. 5. Samples from these locations were subjected to tests to determine dry density (V~) and hence saturated moisture content (msat) and porosity. The results of these tests are shown in Fig. 6 where the distribution of density in this area is compared with the distribution for the more widespread sampling given in Fig. 1, and Clayton (1983). When the results are plotted as a function of dip, similar trends to those observed by Mimran emerge. The minimum density of the material sampled in West Surrey does not, however, fall as low as

C. R. I. Clayton & M. C. Matthews

60

"~"Chert say

t

•'l'Weybridge

-I'Bagshot "~"Camberley

"~"Esher

"~'Croydon

~..-""~o 1,9 "Sutton

+Cobham

/

"4-Waking

/ / 41Epsom l T • 4118 .18

"I'~F~2 0[ey O21

o13

"t'Frimley ~

"Coulsdon

"Leatherhead

"l-Fornborough O17

"l-Aidershot •

~ ~ ~iG°ul[dford

o's

&l

Brick

119

..~ "1"Redhill Reigate

"1"Dorking

111

J ~

• Farnham

Limit of Chalkoutcrop s Sampling location

i

5km

j

FIG. 5. Sampling locations for West Surrey chalk. that for the wider population. This suggests that there was a more general hardening in the area, not associated with highly dipping chalk and that either:

(ii) the small shear strains likely to be experienced by the chalk, even in areas of low dip, were sufficient to cause some breaking of bonds and collapse into a more stable state.

(i) tectonism induced high lateral compressive stresses on the deposit, leading to pressure solution. This is not supported by work done by authors such as Plint (1982); and/or

It appears that the influence of shear strain on the modification of the deposit is much more significant than that of confined compression. It might be argued that overpressuring should be just as effective in preventing density increases

Dry density

(Mg/~l

M~tximmum e o s u , red dry density (see Fiql.} o

2.5

Z

•

•

• ~ T r ind observed • ~O by Mimran (1975)

2"

2.0 .

1.5 t

1.0

0

Minimum measured dry density (see Fig.1.)

I

I

I

30

60

90

Dip (degrees) • Chalk from W. Surrey(see Fig. 5) o Chalk from the Isle of Wight & Dorset

FIG. 6. Results of density tests on chalk from West Surrey, Isle of Wight and the Isle of Purbeck.

Mechanical behaviour of chalk in the case of shearing as under confined compression. But as has been observed in soil mechanics tests on other granular soils, shearing of loose deposits causes a particle rearrangement which makes densification inevitable, because interparticle contact must be re-established.

Conclusions 1 The rapid rate of consolidation observed in confined compression tests on slurry chalks indicates that overpressuring was unlikely during the deposition of the chalk in Cretaceous times. 2 Seafloor lithification and tectonism has often had a much more significant effect on dry density than confined compression resulting

61

from the increase in vertical effective stress due to the overlying sediment pile. It seems likely that the density of the softest chalks found may give a reflection of the maximum past effective stresses, but the geographical and stratigraphical extent of such material is limited. Very small shear deformations, which may be associated with the boundaries of tectonically active areas, appear to have produced significant changes in the chalk, particularly in terms of its engineering behaviour. In areas of high tectonic activity such as the Isle of Wight, the Hog's Back and the Isle of Purbeck, large shear strains associated with an increase in thickness of the sediment pile have created chalks which, in some cases, are as hard as hardgrounds such as the chalk rock.

References BLACK, M. 1980. On Chalk, Globigerina ooze and aragonite mud. In: JEANS,C. V. & RAWSON,P. F. (eds) Andros Island, Chalk and Oceanic Oozes. Occasional Publication, Yorkshire Geological Society, 5, 54-85. BOSWELL, P. G . H. 1927. The geology of the country around Ipswich. Memoirs of the Geological Survey of Great Britain. HMSO, London, 121 pp. BROMLEY, R. G. 1968. Burrows and boring in hardgrounds. Meddr dansk geol., Foren. 18, 247. CARTER, P. G. & MALLARD, D. J. 1974. A study of strength compressibility and density trends within the chalk of south-east England. Q.J. Engng Geol. 7, 43. CLAYTON,C. R. I. 1983. The influence of diagenesis on some index properties of chalk in England. Geotechnique, 33, no. 3,225-41. DUSSEAULT, M. B. & MORGENSTERN, N. R. 1979. Locked sands. Q.J. Engng Geol. 12, no. 2, 177-81. HAKANSSON,E., BROMLEY, R. G. & PERCH-NIELSON, K. 1974. Maastrichtian Chalk of North-West Europe--a pelagic shelf sediment. In: HsO, K. J. & JENKYNS,H. C. (eds) Pelagic Sediments: on Land and Under the Sea. Spec. Pub. Int. Ass. Sediment. 1,211. JUKES-BROWNE,A. J. & HILL, W. 1903. The Cretaceous rocks of Britain--Vol. 2. The Lower and Middle Chalk of England. Memoirs of the Geological Survey of the United Kingdom. HMSO, London. -& HILL, W. 1904. The Cretaceous rocks of Britain--Vol. 3. The Upper Chalk of England. Memoirs of the Geological Survey at the United Kingdom. HMSO, London.

KENNEDY,W. J. & GARRISON,R. E. 1975. Morphology and genesis of nodular chalks and hardgrounds in the Upper Cretaceous of Southern England. Sedimentology, 22, 311. LAMBE,T. W. & WHITMAN,R. V. 1969. Soil Mechanics. Wiley, New York. 553 pp. MIMRAN, Y. 1975. Fabric deformation induced in Cretaceous Chalks by tectonic stresses. Tectonophysics, 26, 309. PANE,V., CROCE,P., ZNIDARCIC,D., KO, H. Y., OLSEN, H. W. & SCHIFFMAN, R. L. 1983. Effects of consolidation on permeability measurements for soft clay. Geotechnique, 33, 67-72. PLINT, A. G. 1982. 'Eocene sedimentationand tectonics in the Hampshire Basin'. J. geol. Soc. London, 139, 249-54. STRAHAN,A. 1898. The geology of the Isle of Purbeck and Weymouth. Memoirs of the Geological Survey of the United Kingdom. HMSO, London. TERZAGHI, K. 1923. Die berechnung der Durchlassigkeitsziffer des Tones aus dem Verlauf der hydrodynamischen Spannungerscheinungen. Sitzungsberichte (Abt IIa) Akademit der Wissenschaften Vienna, Part 20 32 (3/4) 125-38. VESIC, A. S. & CLOUGH, G. W. 1968. Behaviour of granular materials under high stresses. Proc. Am. Soc. Civil Engrg. J. Soil Mech. Found. Div. 94, No. SMD, 661-88. WOODLAND,A. W. 1970. The buried tunnel-valleys of East Anglia. Proc. Yorks. geol. Soc. 37, Pt 4, No. 22, 521-78.

C. R. I. CLAYTON& M. C. MATTHEWS,Department of Civil Engineering, University of Surrey, Guildford GU2 5XH.

62

C. R. I. Clayton & M . C. M a t t h e w s

Appendix: Definitions of soil mechanics terminology Dry density (Td). The ratio of the mass of dry solids of a specimen to its total volume (including voids). Saturated moisture content (msat). The moisture content of a specimen when the voids are completely filled with water. Moisture content (m). The ratio of the mass of water in a specimen to the mass of its dry solids. Total (normal) stress (o-). The force per unit area transmitted in a normal direction across a plane, imagining the deposit to be a solid, single phase material. Pore water pressure (u). The pressure of the water filling the void space between the solid particles. Effective (normal) stress (tr'). The stress transmitted through the materials skeleton only (i.e.

0"-- U). Excess pore water pressure (ue). Pore water pressure in excess of its hydrostatic or steady state seepage value.

Coefficient of earth pressure at rest (Ko). The ratio of the effective horizontal normal stress to the effective vertical normal stress for a condition of no lateral strain (i.e. confined compression). Virgin compression curve. For a given sediment type under Ko consolidation, a unique plane in void ratio, effective normal vertical stress, time space for first-time loading. Coefficient of consolidation (Cv). A constant defining the rate at which consolidation occurs. In the Terzaghi equation for one-dimensional consolidation.

6q2Ue CqUe 6qO'v Cv-~z 2 = ~---~- ~t where:

cv = lc/(mv . ~,w) ue = excess pore pressure z = coordinate in the vertical direction av = total vertical stress t=time k = coefficient of permeability ~'w= unit weight of water mv = coefficient of volume compressibility, the slope of the porosity/vertical effective stress confined compression curve.

Non-destructive determination of Young's modulus and its relationship with compressive strength, porosity and density R. J. Allison S U M M A R Y : The Grindosonic apparatus, new equipment capable of indirectly determining rock compressive strength, is discussed. This utilizes the principle that elasticity theory can be applied to rock masses (Attewell & Farmer 1976; Selby 1982) and directly measures the fundamental vibration frequency of a rock sample of regular dimensions following shock excitation. Dynamic Young's modulus and a variety of other parameters can be established. Samples of Upper Cretaceous Chalk and Upper Jurassic Portland Limestone are used to demonstrate the apparatus and its application. Test specimens were prepared and analysis conducted on material extracted at a number of locations throughout the Isle of Purbeck in Dorset, UK. Samples suitable for deformation in triaxial compression were also prepared and correlations drawn between compressive strength, dynamic Young's modulus, porosity and density.

The Grindosonic apparatus The Grindosonic apparatus (Fig. 1) is new equipment capable of determining rock strength. This utilizes the principle that elasticity theory can be applied to rock masses and directly measures the fundamental vibration frequency of a rock sample of known dimensions following shock excitation. The measured elastic material properties can then be calibrated to determine rock hardness and strength. The device, an example of ultrasonic pulse velocity testing equipment (Neville 1981), analyses the transient vibration of a test specimen. The sample is struck to set up a mechanical vibration pattern rather than being subjected to continuous flexure. This pattern is converted to an electronic signal via

FIG. 1. Grindosonic apparatus.

either a piezo-electric detector held in contact with the test piece surface or a microphone placed directly beneath the sample. The signal is amplified by the apparatus before being fed to the instrument input. If it exceeds the predetermined minimum level required for analysis, the time of eight wave passes is measured. A short interval between striking the sample and measurement prevents the analysis of spurious initial wave patterns which have complex harmonics and which occur when the test piece is initially struck. The lapsed time appears as a result in the equipment display panel and is known as the r-value. The vibrating sample experiences damping relative to its elastic properties. The decay of a vibration pattern set up in a hard, rigid test piece will take much longer than the same flexure in a soft material of similar dimensions. The result given by Grindosonic therefore constitutes a direct measure of sample rigidity or hardness. However, the natural frequency of vibration is determined by specimen shape and several other physical constants. Consequently, by utilizing details of the dimensions and weight of a specimen together with its r-value determined by the Grindosonic, Young's modulus of elasticity, shear modulus, Poisson's ratio, bulk density, seismic velocity and a variety of other parameters can be calculated. Additionally, these properties can be measured as a function of variables such as temperature, degree of saturation, chemical alteration and weathering.

From:JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformationof Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 63-69.

63

64

R. J. Allison

Sample preparation, constraints and testing Grindosonic testing relies on accurate sample preparation. Test pieces can be cut to a variety of shapes including bars, cylinders and circular discs. Some variables used in the analysis do have specific nominal limits. For example, when preparing bar-shaped specimens the length to thickness ratio should be greater than three and the width of the bar should be less than one third the length (Fig. 2). Beyond these limits the calculations gradually lose accuracy. Careful sample preparation is of the utmost importance because the variations in cross-sectional area and non-square edges cause significant changes in the vibration pattern. Grindosonic test pieces were carefully machined to a regular rectangular shape by cutting, approximately in the first instance, with a diamond saw and then by grinding plane faces using an engineering vice and precision toolroom grinder. All test pieces were cut in the same direction relative to bedding, dried at 60°C to constant weight and left to cool in a desiccator. Tests were conducted with samples resting on a foam rubber mat. The top face of each bar was struck at its central point with a thin circular glass rod. The piezo-electric detector was held in contact with the sample at the centre of one of the side faces. Two samples were extracted from each field location and two test specimens prepared from each sample. This permitted the examination of within sample-site variability as well as between sample-site differences. Rock samples were collected from Cretaceous Chalk and Jurassic Portland Limestone throughout the Isle of Purbeck. The spatial distribution of sample sites was

chosen such that the results would be representative of these rock types throughout the region. Joint bounded blocks of Chalk were removed at intervals along the coast from three quarries where the Chalk outcrop runs inland between Worbarrow and Studland, and at different points from north to south across the outcrop, providing 26 test pieces in all were provided. Specimens of Upper Jurassic Portland Limestone were removed following the same procedure from all localities where the material outcrops at the coast and from one quarry where the unit swings inland between Gad Cliff and St Aldhelm's Head, providing 68 test pieces.

Study area The Isle of Purbeck in Dorset forms an individually identifiable geological area of Late Jurassic to Upper Cretaceous rocks (Arkell 1947), (Fig. 3). The rock units run approximately parallel to the coastline (Fig. 4) and topographic variations conform closely to changes in geology, reflecting differences in the engineering behaviour of the rocks (Jones et al. 1983). The geological structure at shallow depths is a monocline with near vertical dips in its northern

¢

2 DORSET ..... •

,,

beck

FIG. 2. Fundamental flexural vibration of a square shape bar.

?

FIG. 3. Location of the Isle of Purbeck.

~?

~0o

Determination of Young's modulus 'l'J'll

-

[~

Tertiaryto recent ~ Chalk ~ ~ : ~ Greensand& Gault ~ Wealdon Beds

~

65 4

Purbeck Beds

PortlandStone PortlandSand Kimmeridge Clay

o_~,ooo m

FIG. 4. Geological map of the Isle of Purbeck.

limb and a gentle plunge to the east (Melville & Freshney 1982). The general trend of the monocline is oblique to the coast with the dip of bedding at the coast increasing from east to west. Details of the evolution of the monocline and its relationships to major structures at depth are still disputed (Plint 1982; Stoneley 1982).

Laboratory procedure Tests were conducted to establish the accuracy of Grindosonic in determining dynamic Young's modulus as an indication of rock compressive strength and the relationships between dynamic Young's modulus, compressive strength, porosity and dry density. Data were obtained by applying the following techniques:

(i) The Grindosonic was used to determine dynamic Young's modulus. Twenty r-value tests were conducted on each sample using the piezo-electric detector. The spread of these results seldom varied by more than 0.5% The mean was taken and used together with details of sample geometry and mass to determine dynamic Young's modulus. (ii) Triaxial tests, using an E.L.E. Hoek Cell, provided rock strength data (Brown 1981). Test pieces with a diameter of 38 mm and height/diameter ratio of 2.0-3.0 were cored from blocks and their ends ground flat to within ___0.02mm. These specimens were prepared from each block and tested at confining pressures of 15, 30 and 60 MN m-2. The axial stress was increased at a constant rate, this and axial strain being recorded on a short regular time base until failure. (iii) The Liquid Saturation technique was em-

ployed to establish material porosity and dry density (Carter & Matthews 1977). The bulk volume, grain mass and saturated surface dry mass were determined. Saturation was conducted under vacuum over 24 hours. Porosity and dry density were calculated from these results.

Relationship between porosity and dry density Materials such as Chalk often have a wide range of porosities, (5-80~), which can be expected to influence other parameters including compressive strength and dry density. Plots of dry density (g cm-3) against porosity (%) for Chalk (Fig. 5) and Portland Limestone (Fig. 6) show an inversely proportional linear relationship, indicating that a decrease in porosity is associated with an increase in density. Variations between porosity and dry density are noticeably different between these two materials. This is reflected by the different gradients of the two regression lines, illustrating a greater change in the porosity of the Chalk than in Portland Limestone for an equivalent drop in dry density. Nevertheless, there is a close association between their two parameters and it can be suggested that, although dry density and porosity have seldom been regarded as important indices, they may well be useful material parameters.

Relationship between porosity and dynamic Young's modulus Separate plots were initially drawn of porosity (~) against dynamic Young's modulus

66

R. J. Allison

50-

40-

E Z 30-

o ~ 20._o E

°%o

,~

3b

2% Porosity

(%)

FIG. 5. Plot of dry density (g cm-3) against porosity (%) for chalk. (kN mm -2) for Chalk (Fig. 7) and Portland Limestone (Fig. 8). Chalk shows an inversely proportional curvilinear relationship between these parameters. Although high porosities demonstrate only small differences in dynamic Young's modulus, small variations in porosity in the range 5-15% result

3

~

0 o

,~

2'0

!

30

.S

i

~o

6'o

Porosity (%)

FIG. 7. Plot of dynamic Young's modulus (kN mm- z) against porosity (%) for chalk. in a large change in dynamic Young's modulus. Portland Limestone again shows an inversely proportional relationship, although this appears to be linear, with the material exhibiting lower porosity values than those for the Chalk and correspondingly high dynamic Young's moduli. An initial assumption would be that each data set demonstrates a different relationship due to the different material characteristics. However, a plot of both sets of results on the same axes (Fig. 9) indicates a more fundamental relationship exists. Both sets of data lie on the same regression line of the equation:

:E

g

10-

0

o 2.

qb=m-(l/b) lnEmod

(1)

where:

1

2

3

4

5 Porosity (%)

6

7

8

9

FIG. 6. Plot of dry density (g cm-3) against porosity (%) for Portland Limestone.

~b= m= b= Emo~=

porosity constant = 68.48 constant -= 0.0633 dynamic Young's modulus.

Thus not only does a relationship exist between

67

Determination of Young's modulus

Q

•

•o

~

-

~

.

~

•

i •

40

!

•

2 O

C~ 20

.

0

0

~ '

;

~r

4,

',

6,

',

;

9l

Porosity ~q,'0)

Fro. 8. Plot of dynamic Young's modulus (kN mm- 2) against porosity (~) for Portland Limestone.

porosity and dynamic Young's modulus for specific materials but also, on a much broader scale, there is a general correlation for these carbonate sediments. Assuming that dynamic Young's modulus provides a measure of material strength, porosity can thus be used as an indirect index of rock strength and can be used to compare different materials.

z

? 40•

Chalk

•

Portland

Limestone

O

=

20-

Relationship between compressive strength and dynamic Young's modulus Plots of compressive strength (MPa) at 60 M P a confining pressure against dynamic Young's modulus (kN mm -2) for Chalk (Fig. 10) and Portland Limestone (Fig. 11) establish the accuracy of both dynamic Young's modulus as an indication of rock compressive strength and Grindosonic as a technique for determining dynamic Young's modulus. Both materials demonstrate a linear decrease in material strength with reduced dynamic Young's modulus values. Overall moduli are higher for Portland Limestone than they are for Chalk, with a corresponding increase in strength values. Scatter about the regression line is due partly to inaccurate sample preparation and partly to undetected sample-specific variations such as microcracks invisible to the naked eye. It is important to note that both sets of data lie on the same regression line, suggesting that not only is there a relationship between these two parameters

0

o

i

i

,o

~o

~'o

;o

~'o

!

~o

P o r o s i t y (%)

FIG. 9. Plot of dynamic Young's modulus (kN mm-2) against porosity (%).

of equation: ff = C. Emod

(2)

where: a = compressive strength e = constant = 6.505 Emod = dynamic Young's modulus but that this is constant for different carbonate sedimentary rocks. Progressively stronger materials show an increase in dynamic Young's modulus of approximately 0.14 k N mm -2 M P a - 1.

R. J. Allison

68

//

(b) 300"

200

/

I

30O

2OO

lOO-

/ /

, 10

i 20

30

i 40

i 50

D y n a m i c Y o u n g ' s M o d u l u s (kN m m "2)

,~

~'o

~'o

~

s~

~

/o

~

9'o

FIG. 10. Linear plot of yield strength (MPa) against dynamic Young's modulus (kN mm- 2) at 60 MPa confining pressure for chalk. (a) Data from Jones & Bedford (unpublished). (b) Data from Jones & Gilligan (unpublished).

FIG. 11. Plot of material yield strength (MPa) against dynamic Young's modulus (kN mm - 2) at 60 MPa confining pressure for Portland Limestone.

Discussion

where:

Dynamic Young's Modulus (kN mm-:')

The conclusions drawn from these tests have important implications for the relationship between compressive strength, dynamic Young's modulus, porosity and density of sediments and the application of Grindosonic. A clear relationship exists between porosity and dry density. Little has been made of this in the past, but data suggest these parameters can provide useful additional information on the strength of rock materials. From the close relationship between porosity and dynamic Young's modulus it can again be inferred that porosity can be used to provide useful additional information on rock material strength. Since Chalk and Portland Limestone data both conform to the same regression equation (1) results for different materials can be accurately compared. Using plots of yield strength at 60 MPa versus dynamic Young's modulus it is possible to derive an expression for compressive strength in terms of porosity: cry = c. exp(a - b. ~b)

(3)

ay = compressive strength c = c o n s t a n t (confining pressure dependant) a = constant = 4.335 (for Chalk and Portland Limestone) b = constant = 0.663 (for Chalk and Portland Limestone) = porosity. This holds true for both Chalk and Portland Limestone and therefore not only confirms a general relationship between Young's modulus and compressive strength for carbonate sediments, but also establishes Grindosonic as a technique capable of accurately determining dynamic Young's modulus of materials, as well as other parameters, such as seismic velocity, not discussed here. The Grindosonic technique has a number of benefits including the comparative ease of sample preparation, test simplicity, result reproducibility and the advantage of being non-destructive and hence permitting sample retesting under different conditions. Therefore, not only do these tests establish important relationships between dynamic Young's modulus, compressive strength,

Determination of Young's modulus porosity and density, but they also suggest that the Grindosonic apparatus has wide application in the field of rock characterization, testing and monitoring. Data presented here have been restricted to two rock units. Work on other materials is in progress. ACKNOWLEDGEMENTS: The author expresses his thanks to Dr M. E. Jones and Professor D.

69

Brunsden for encouragement with the field work and valuable discussions during the course of this work. The research was commenced while in receipt of a N E R C research studentship and is being continued with funding from the Addison Wheeler Fellowship. The figures were prepared with dexterity by R o m a Beaumont and Gordon Reynell in the Geography Department at King's College, London.

References ARKELL,W. J. 1947. The Geology of the Country around Weymouth, Swanage, Corfe and Lulworth, p. 386. HMSO, London. ATTEWELL, P. B. & FARMER,I. W. 1976. Principles of Engineering Geology, p. 1045. Chapman & Hall, London. BROWN, E. T. 1981. Rock Characterisation, Testing and Monitoring, p. 211. Pergamon Press, Oxford. CARTER, M. E. ~¢ MATTI-IEWS, A. P. 1977. The measurement of porosity of irregular vulgar rock samples using a modified liquid saturation technique. Geotechnique, 27, 435-7. JONES, M. E., ALLISON,R. J. & GILLIGAN,J. 1983. On the relationships between geology and coastal

landforms in central southern England. Proc. Dorset nat. Hist. archaeol. Soc. 105, 107-18. MELVILLE, R. V. & FRESrtNEY, E. C. 1982. British Regional Geology, the Hampshire Basin and Adjoining Areas, p. 146. HMSO, London. NEVILLE, A. M. 1981. Properties of Concrete, p. 779. Pitman, London. PLINT,A. G. 1982. Eogene sedimentation and tectonics in the Hampshire Basin. J. geol. Soc. London, 139, 249-54. SELBY, M. J. 1982. Hillslope Form and Process, p. 264. Oxford University Press, New York. STONELEY,A. 1982. The structural development of the Wessex Basin. J. geol. Soc. London. 139, 543-54.

R. J. ALLISON,Department of Geography, King's College London, Strand, London WC2R 2LS. Present address: Department of Geography, University of Durham, South Road, Durham DH1 3LE.

A laboratory technique for investigating the deformation microstructures of water-rich sediments Alex Maltman S U M M A R Y : Experimental deformation is a particularly valuable means of investigating the mechanical behaviour of water-rich sediments. Techniques have been developed for preparing and deforming wet sediments in the laboratory and petrographically examining the resulting microstructures. Clays marketed for the ceramic industry are more useful than natural clays because of their consistent properties. They are mixed with sea-water, aged, and consolidated to the desired value either by self-weight or in a simple levered oedometer. Argillaceous sediments with water contents of about 50% and lower are routinely produced. Wetter materials are extremely delicate and vulnerable to non-testing artefacts. In an attempt to circumvent this problem a device has been designed which enables slurries to be sedimented directly on to the base of a triaxial compression cell. The prepared specimens are tested in a conventional 'direct' shear box, or in 'distributed' shear which subjects the whole block to shear displacement, or in triaxial compression. The latter tests can be undrained, or partially/intermittently drained, from the top and/or the bottom of the sample. After testing, slabs are cut from the specimen with a wire-saw and impregnated with Carbowax P.E.G. 8000 to allow thin-section preparation and examination.

There is growing recognition of the range of geological circumstances in which near-surface sediments can be subject to deformation by external forces. The responses of the sediments have, so far, been little investigated. Direct observation is difficult, and because any structures formed are vulnerable to modification before they are preserved in rocks, the standard approach of interpreting deformation behaviour from the present appearance of structures is fraught with difficulties. Experimental deformation, therefore, provides a particularly valuable means of investigating the mechanical behaviour of water-rich sediments. A notorious problem in structural geology is the distinction between early, pre-lithification movements and structures formed after lithification (see discussion in Maltman 1984). The essential difference between the two is in the deformation mechanisms. Water-rich sediment, in contrast to sedimentary rock, deforms predominantly by grain boundary sliding. Therefore, it is at the microfabric scale that criteria might be discovered which would enable the general distinction of structures formed in this way from those produced in rock. For these reasons techniques have been developed for deforming water-rich sediments in the laboratory and petrographically examining the microstructures generated in the experiments. Wet sediments are delicate and difficult to work with. In dealing with these seemingly capricious materials numer-

ous practical 'knacks' have been learned which cannot easily be communicated in writing. The following is an outline of the main practical methods which have been evolved for preparing, deforming, and examining the various sediments.

Materials Much of the experimental programme so far has been concerned with argillaceous sediments, and a range of these has been utilized. Muds from modern estuaries (e.g. Mawddach, Wales, Grid Reference SH696186; Leri, Wales, SN616934; and West Cleddau, Wales, SM974115) and glacial clay (Hengwm, West Wales SN778946) have been used, either disaggregated and resedimented, or cut out in blocks and trimmed to the form required for testing. These materials provide some interest for comparative purposes but they present severe problems of compositional variation between samples. In addition, the presence of organic matter not only introduces a further influence on the mechanical properties but can cause the material to smell during storage. It is generally more satisfactory to purchase bulk lots of clays, such as kaolinite and bentonite, from suppliers to the ceramic industry. A good petrographic simulation of rock is provided by ball clay and this has been used extensively in the experiments. Various coloured clays (named pot clay and coloured body clay in the ceramic

From: JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 71-76.

7I

72

A. Maltman

industry) have been found useful for introducing marker laminations into the specimens. For materials with greater clast content a disaggregated Silurian shale used for brickmaking has been used, but it is heavily coloured. Washed sand from builders' merchants has been used in various mixtures with clays. However, where the proportion of admixed clay is substantially reduced, the cohesion between the sand grains tends to be lowered to the extent that the specimen does not remain intact.

can influence the mechanical properties of the resulting sediment (e.g. Einsele et al. 1974) and its microfabric (Maltman 1981). Therefore, the preparation has been restricted to a few selected procedures in the interests of efficiently producing samples with reasonably consistent starting properties. The ceramic clays mentioned above are marketed as dry powders and so they have first to be mixed with water. An arbitrary procedure is followed. The powder is mixed with clean seawater in a ratio of about one to two by volume, mechanically stirred for 24 hours and then allowed to age in covered plastic buckets for a month or more. The effects of varying this practice have not been explored. The resulting slurry is formed into specimens for testing in a variety of ways. The most straightforward method is simply to

Preparation It is clearly impossible to simulate closely the processes by which partially-consolidated sediments are produced in nature. Moreover, the preparation methods adopted in the laboratory

®

o

i

/

~-:.-,~.-~.~.,..7,~-.:.-,r.,....7-. ~-..--...-.=...~ .

: 51

ii®! ®

-Q

il

@

.5 cm

5 cm

~

:~

(A)

!

(c)

~

I

f-

5 cm

(a)

@

FIG. 1. Schematic drawings of devices for sample preparation. (A) Settling tank. Slurry (a) settles under its own weight; supernascent water (b) can be syphoned off through pipes (c); water leaves base through porous bauxilite plate (d) and removed through pipe (e). (B) In situ settling apparatus. (1) is base of high-pressure triaxial cell (shown in more detail in Fig. 5); (2) is one of three screw-rod assemblies which secure perspex settling cylinder (3) and stainless steel 3-way split former (4) to cell base. Removal of (2), (3), and (4) leaves material (5) settled in situ in rubber sleeve (6), and available for triaxial testing without disturbance. (C) Load consolidation apparatus. Load is applied by adding weights (A, on right-hand side of figure) at a pre-determined distance along lever (B), which acts on pivoted piston (C). Specimen (D) is dewatered through porous polythene discs (E) and drilled aluminium caps (F), and water is abstracted through pipes (G). Nylon straps (H) secure the perspex settling cylinder to the apparatus. Latch (I) enables lever to be held clear of the settling cylinder during assembly.

Deformation microstructures of water-rich sediments pour the slurry into an oiled vessel and allow it to settle under its own load. If the base of the vessel is constructed with a porous plate (Fig. 1A), water can be drained from below as the material consolidates and supernascent water can be syphoned off from time to time. Should the water be completely removed from the surface, airdrying commences and the material acquires cohesion rapidly. It shrinks slightly and acquires a porosity greater than its water content. Material prepared in this simple way is not observably different from that produced by the more sophisticated methods subsequently described. When the settled sediment is judged to have acquired sufficient cohesion a further slurry can be introduced. This can be either of the same clay to give a 'bedding plane' in the sediment, or of a contrasting clay to produce a marker band. Addition of further slurries enables a layered sediment to be built up. The settling vessels are perspex boxes for producing blocks and perspex tubes for preparing cylindrical specimens. Samples can be produced ready for testing if the vessels are obtained or constructed at the appropriate size. Alternatively, the material can be cut and pared to the right shape and size for testing. The block produced in the settling tank shown in Fig. 1(A) is of sufficient size to enable cylinders to be carved from it with their long axes parallel to the horizontal direction of the block. In this way the sedimentary fabric and any laminations are parallel to the long axis of the cylinders and

73

therefore to the 'vertical', or principal, compression in triaxial testing (cf. Maltman, this volume, fig. 2). Clay specimens can be produced in the ways mentioned above with water contents approximately in the range 20-45~. Air-drying is needed to reduce the water content further; specimens at the upper end of the range or higher are so weak and close to their liquid limits that they are very sensitive to handling. In an effort to circumvent this latter problem, slurries have been settled in situ directly on the base of high-pressure triaxial cells (Figs 1C and 2A). The stainless steel jacket is in three vertically-divided segments which are held together by ring clamps. A latex rubber sleeve (which will be needed in the triaxial test to isolate the specimen from the confining fluid) can be stretched inside the jacket. A perspex cylinder fits closely on top, and the whole assembly is secured to the cell base by three compound screwrod devices, one of which is depicted in Fig. 1(B). The slurry is poured into the jacket and perspex cylinder and allowed to settle under its own load. When it has consolidated sufficiently, the screw rods and perspex cylinder are removed and the jacket unclamped. The jacket then falls away leaving the specimen, already in a rubber sleeve, ready for testing with minimal disturbance. So far, however, this method has not been entirely satisfactory. It is still difficult not to jiggle the specimen. The top cap assembly has to be put in position, and the cell-base moved onto the platen

FIG. 2. Photographs of laboratory devices. (A) Arrangement for settling sediments in situ on base of high pressure triaxial cell. (B) Typical appearance of specimen assembly before high effective pressure triaxial test. (C) Device for inducing 'distributed' shear. (D) Load consolidation apparatus for shear box specimens.

74

A . Maltman

of the triaxial apparatus. Such is the sensitivity of very wet sediments that even the lightly stretched rubber sleeve can introduce disturbance. For most of the tests that are conducted, specimens produced by any of the methods outlined above will be under-consolidated in the engineering sense. For experiments that are meant to simulate deformation at burial depths of tens of metres and more it is more realistic to consolidate the slurry under an applied load. The conventional consolidometer of soil mechanics could be used for this (e.g. Akroyd 1964). However, if reasonably slow consolidation is desired together with interruptions from time to time to introduce different slurries, a number of consolidometers would have to be operated simultaneously in order to produce a useful throughput of samples. Therefore, a simple levered apparatus has been utilized in the present work. It is depicted and explained in Fig. I(C). There are various angular and frictional forces which prohibit use of the device in any rigorously quantitative way but it is suitable for the present purpose. An analogous device for producing blocks for shear-box testing is shown in Fig. 2(B). The design and construction of these consolidation devices is sufficiently simple to justify several of them being built and in constant use. Specimens consolidated to water contents less than about 40~, under loads up to the equivalent of 250 m burial, are routinely produced.

Deformation Shear tests

Tests in which the prepared partially-consolidated sediment is subjected to bulk simple shear are of two kinds, termed here direct and distributed shear. Direct shear is the conventional shear box test of soil mechanics (e.g. Akroyd 1964). It is shown schematically in Fig. 3. A 6 x 6 × 2 cm sample is placed within a horizontally-split rigid stainless-steel frame. By applying a constant rate of displacement to the bottom half while holding the top half stationary, the upper part of the specimen is caused to slide past the lower half, creating within it a zone of shear. A vertical load, normal to the shear plane, can also be applied. The bulk shear force and displacement are recorded. Thus the force needed to cause shear failure of the material is readily determined and the test finds widespread engineering application. However, because the plane of displacement is largely pre-determined by the frame design rather

FIG. 3. Schematic drawing of 'direct' shear box. (A) = Specimen. (B) = Locating bolt. (C) = Point of application of force to cause displacement. (D) = Point of application of vertical (normal to displacement) load. (E) = Point of monitoring shear force and displacement sustained by specimen. (F) = Separating bolt. In assembling the experiment the top and bottom halves of the shear box are aligned by bolts (B) and slightly separated, to avoid sliding friction, by bolts (F); the bolts are removed during the test.

than any intrinsic property of the specimen, this system has been little used for the present purposes. It is considered more realistic, geologically, to subject the entire specimen to the shearing force, and for this a 'distributed shear' device has been constructed. In this design (Fig. 4) the specimen, around which a thin rubber sleeve is lightly stretched, fits closely within a stack of thin, lubricated aluminium plates (Fig. 2C). The sleeve is clamped to toothed aluminium pads which fit into the top and bottom of the specimen. The bottom pad sits snugly in the recessed base of an aluminium carrier to which, as in the direct shear test, a constant rate of displacement can be applied while holding stationary the upper pad and arm assembly. The bottom surface of the specimen is displaced with respect to the top surface, but the sleeve and aluminium plates act together to distribute the shear force throughout the sample. This can then respond as its properties dictate, either by some bulk change of shape, or by concentrated shear anywhere within the specimen. As in the direct shear mode, loads can be applied normal to the shear direction. The bulk shear force and displacement are monitored. A weakness of the design is that at low normal loads there is a tendency for the displacement to be taken up preferentially by the top toothed-pad gouging across the top surface of the specimen. Both shear systems are operated on conventional motorized stands and the loads applied and measured in the ways standard to soil engineering laboratories.

Deformation microstructures of water-rich sediments Triaxial compression tests These experiments are also, in principle, routine in soil mechanics (e.g. Bishop & Henkel 1962; Akroyd 1964), and the basic equipment readily obtainable. The modifications for the present purpose lie in the assembly techniques, the way the experiments are conducted and the parameters investigated. The high pressure cell (Fig. 5) allows oil to be introduced as a confining fluid. The confining pressure is provided and maintained by a separate constant pressure apparatus. The specimens are cylindrical and measure 107 mm in height and 46 mm in diameter. They are normally placed on the bottom pedestal of the cell but, as explained earlier, the base of the cell is designed so that sediments can be settled in situ. Irrespective of the preparation method, water-rich specimens are so delicate that assembly of the test without disturbance is only feasible after some practice. Fitting the rubber sleeve to the specimen (to isolate it from the confining fluid), the 'O-ring' seals, the top cap assembly, the top drainage pipe, etc., are all manoeuvres which can easily shake the specimen. For drained tests at higher confining pressures (above about 200 kPa) it is necessary to use thicker rubber sleeves to prevent the confining fluid exploiting weaknesses. The particular weak spots where the specimen ends meet the top and bottom caps have to be reinforced by winding round adhesive plastic tape (Fig. 2D). With drained tests on high watercontent samples, filter paper and a bauxilite disc with a smooth metal outer rim are fitted between the specimen ends and the end caps to allow drainage of water without any egress of clay material. Drainage is achieved through holes in the end caps which are connected by pipes to pressure transducers in order to allow monitoring of the fluid pressure. Klinger valves allow the specimen to be sealed, or drained from the top and/or the bottom of the specimen, partially or fully, and continuously or intermittently. In these ways various geological dewatering situations can be simulated.

Examination After dismantling the deformation experiment, the specimen can readily be examined macroscopically but steps have to be taken to allow thin-section preparation. Air-drying followed by vacuum impregnation by various epoxy resins has been employed for fairly dry material, such as that which has completed a dewatering experiment, but clearly if the microfabrics associated with redistribution of water within the

75

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FIG. 4. Schematic drawing of device for inducing 'distributed' shear. (A) = Specimen. (B) = Rubber sleeve. (C) = Sleeve clamps. (D) = Lubricated plates. (E) = Balancing weight. (F) = Point of monitoring shear force and bulk displacement sustained by specimen. (G) = Point of application of vertical (normal to displacement) load. (H) = Point of application of force to cause displacement.

material are of interest, then any preparation method which involves shrinkage is not satisfactory. The most convenient and acceptable methods have utilized impregnation by the high molecular weight polyethylene glycol marketed as 'Carbowax'. A routine procedure is as follows. A slice 1-1.5 cm thick is cut from the central part of the deformed specimen with a fine wire stretched in a fret-saw frame. The slab is wrapped in cling-film and stored in a humidity cabinet until several specimens are available for impregnation. Material adjacent to the slice is used for determination of bulk water content, porosity, etc., although a y-ray attenuation device is under construction at present for assessing microvariations of water content in traverses across the slice. Areas of the slab of which a thin-section is required are selected and cut into blocks. These 5 × 3 × 1.5 cm blocks are immersed, usually about six at a time, in a Carbowax bath at 63°C, and held, with occasional turning, for about 6 days. After cooling to room temperature, a smooth surface is cut through the block using a precision power saw with Buehler Isocut fluid to prevent blade clogging. The surface is polished to flatness with wet-and-dry carborundum paper and paraffin as a coolant. The surface is cleaned using Loctite cleaner and mounted on a pre-ground slide using Loctite 15-12. The mounted slice is cut on the precision saw to a thickness of around 150 txm, thinned on a Jones-Shipman surface grinder and finished by hand on 800 grade carborundum. The slide is trimmed, cleaned, and a cover-slip fixed using Histomount. The thinsections can then be examined petrographically

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List of UK suppliers of specialized materials

® 5¢rrl

I

•

FIG. 5. Schematic drawing of high pressure cell for triaxial testing. (A) = Specimen. (B) = Rubber sleeve. (C) = Confining fluid. (D) = Pore water egress. and stored in the normal way. Examples of photomicrographs of fabrics induced by the deformation experiments outlined here can be seen in Maltman (1982).

Ceramic clays: Wengers Ltd, Etruria, Stoke-onTrent, Staffs ST4 7BQ; Harry Frazer Ltd, Vauxhall Street, Longton, Stoke-on-Trent, Staffs ST34 H D ; Potclays Ltd, Brickkiln Lane, Etruria, Stoke-on-Trent, Staffs ST4 7BP; Harrison Mayer Ltd, Meir, Stoke-on-Trent, Staffs ST3 7PX. Brick clay." Castle Brick Company, Buttington, near Welshpool, Powys, Wales. Perspex tubes." Visijar Plastics Ltd, Portland Street, Aston, Birmingham B6. Deformation equipment: Wykeham-Farrance Engineering Ltd, Weston Road Trading Estate, Slough, Bucks; ELE International Ltd, Eastman Way, Hemel Hempstead, Herts HP2 7HB. Carbowax: UNALCO, Maybrook Industrial Estate, Maybrook Road, Brownhills, Walsall, West Midlands W58 77DG. (Carbowax 6000 was used until 1982 when the product was discontinued, since then Carbowax 8000 has been used.) Precision saw: Constructed from a design kindly made available by the Department of Earth Sciences, University of Leeds. Isocut blade lubricant: Banner Scientific Ltd, Binns Close, Torrington Avenue, Tile Hill, Coventry CV4 9TB. Loctite adhesive: K. L. Whiston, New Mills, Stockport SK 12 4PT. Histomount adhesive: National Diagnostics, 45 Long Plough, Aston Clinton, Bucks HP22 5HD.

ACKNOWLEDGEMENTS. Many of the innovations in the equipment design and construction described here have been made possible by the unfailing interest, ingenuity and skill of Julian Bird. Details of the thin-section preparation method have been evolved by Tommy Ridgeway. The help and cooperation of these and other technicians in the Department of Geology, UCW, is gratefully acknowledged.

References AKROYD, T. N. W. 1964. Laboratory Testing in Soil Engineering. Soil Mechanics, London. BISHOP,A. W. & HENKEL,D. J. 1962. The Measurement of Soil Properties in the Triaxial Test. Edward Arnold, London. EINSELE, G., OVERBECK, R., SCHWARZ, H. V. & UNSOLD,G. 1974. Mass physical properties, sliding and erodibility of experimentally deposited and

differentially consolidated clayey muds (Approach, equipment, and first results). Sedimentology, 21, 339-72. MALTMAN,A. J. 1982. Experimentally deformed clay. In: BORRADAILE,G. J., BAYLY,M. B. & POWELL, C. McA. (ed) A tlas of Deformational and Metamorphic Rock Fabric's, 426-33. Springer-Verlag, Berlin. - - - , 1981. Primary bedding-parallel fabrics in structural geology. J. geol. Soc. London, 138, 475-83.

ALEX J. MALTMAN,Department of Geology, University College of Wales, Aberystwyth, Wales SY23 3EE, Wales.

Shear zones in argillaceous sediments

an experimental study

Alex Maltman S U M M A RY: Experimental deformation of water-rich argillaceous sediments has shown that they deform not by pervasive homogeneous flow, as has sometimes been surmised in the past, but by intense slippage within very narrow, discrete zones of shear. These shear zones enable large total strains to be accomplished whilst leaving the bulk of the material undisturbed. Macroscopically the zones are shiny, finely-grooved planes which may be stepped where sub-structures intersect them. Under the microscope the zones are seen to result from pronounced particle reorientation into sub-parallelism with the zone margins, presumably by slippage at the grain scale. The zones commonly curve and anastomose, and various sub-fabrics may be discernible. Although the details vary, the shear zones are strikingly consistent throughout the range of experimental conditions. In specimens with 15~ water content, the shear zones lack cohesion and are analogous to shear fractures in brittle rocks. Between approximately 15-45~ water content the same overall geometry persists, but the shear zones maintain cohesion. At greater water contents, up to at least 60~, the clay sediments are extremely weak and although they may appear to be undergoing pervasive flow, microscopic examination reveals that even here arrays of short, narrow zones of concentrated displacement are being generated.

It has been surmised in the past that water-rich argillaceous sediments deform chiefly by pervasive homogeneous flow. Intuitively it seems reasonable that the water should so reduce the friction between the clay particles that virtually every one would contribute to the bulk movement. There have, moreover, been suggestions that this flow would produce a pervasive fabric, which, if recorded by lithification, would be seen in a rock as slaty cleavage. As part of an investigation of how water-rich sediments behave during geological deformation, argillaceous sediments have been deformed experimentally in the laboratory (see introductory discussion in Maltman, this volume). Results of some of this work are reported here and these strongly suggest that argillaceous sediments do not deform by this pervasive flow mechanism. Rather, the deformation takes place by intense movement concentrated in narrow, discrete zones of shear. The bulk of the material is left undisturbed. Even high water-content clays, which in hand specimen appear to be undergoing bulk 'plastic' flow, on closer inspection reveal arrays of narrow zones of slip. These discrete zones are referred to here as shear zones. The purpose of this paper is to describe the nature and appearance of the shear zones in various experimental configurations, and to emphasize their important role in the laboratory deformation of argillaceous sediments. The occurrence of the shear zones in natural sediments and rocks will be considered in a sequel article.

Previous work The stress-strain behaviour of clays has been closely studied in soil mechanics, although relatively little attention has been given to the deformation mechanisms (e.g. Lambe & Whitman 1969, p. 313). From a geological point of view, however, and for conditions which are more relevant geologically (e.g. high initial water contents; greater possible confining pressures) there has been little more than intuitive speculation. Publications of some geological relevance are three pioneering German papers by H. Cloos, Hvorslev and Riedel (references given in Tchalenko 1968) and a series of papers from Imperial College, London, in the late 1960s. Among the latter, Skempton (1964) showed how the strength characteristics of clays depend upon the development of shear-induced fabrics, and compared them to tectonic shear zones (Skempton 1966). Tchalenko (1970) pursued further the comparison between shear zones in clays and analogous structures of greater magnitude. Morgenstern & Tchalenko (1967a) examined in careful detail the shear zones produced in pure kaolinite during experimental simple shear, and also in natural clay landslips (Morgenstern & Tchalenko 1967b). Maltman (1977) summarized a variety of microstructures produced experimentally in wet argillaceous sediments, including the shear zones to be described here. Maltman & Fitches (1982) noted occurrences of similar zones in the geological record. The shear zones have transpired,

From:JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 77-87.

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during the present experimental programme, to be the chief mechanism by which argillaceous sediments deform, and to be important in a wide range of geologically relevant conditions. The laboratory work is now reported more fully.

specimens can then be examined petrographically in the normal way.

Appearance of the zones Macroscopic appearance

Experimental method The methods employed in the experiments are described in a companion paper (Maltman, this volume) and only the briefest outline is given here. Disaggregated clays have been obtained in bulk from ceramic suppliers in order to achieve consistency of starting material. Some natural clays have also been sampled and the specimens deformed intact in the laboratory. The ceramic clays are mixed to a slurry with sea-water, aged, then poured into settling vessels. Some materials have been allowed to settle under gravity, others have been consolidated under applied loads. Both procedures were commonly interrupted to allow the introduction of contrasting slurries to produce laminated specimens. None of the laboratory materials discussed here was consolidated to a load greater than that of the subsequent deformation test, i.e. all the laboratory-prepared sediments were under or normally consolidated. Sediments with water contents up to about 50% have been routinely produced in these ways. At higher water contents, the materials approach their liquid limits and are extremely delicate and difficult to work with. Presumably some air is entrained in the sediments, because where porosity has been measured, it was in every case somewhat higher than the water content. Variation in preparation procedure appears to have no effect on the main results described in this paper. The sediments have been deformed chiefly by 'direct' shear using the conventional shear box of soil mechanics, 'distributed' shear in which the entire block is allowed to undergo simple shear, and triaxial testing of cylindrical specimens. Unless specified otherwise, all the tests reported here were undrained. (Some partially drained and fully drained tests have been conducted, and the main conclusions to be presented here appear to apply to these also.) In the tests, the main variables have been lithology, water content, strain rate and confining pressure. The last is simulated in the shear experiments by imposing a load normal to the bulk displacement plane of a confined specimen. The deformed specimens are cut with a fine wire saw and the central slice impregnated with Carbowax to allow thin-section preparation. The

In most of the experimental conditions employed here, the clays are highly ductile, that is they accomplish bulk strains of at least 30% without signs of fracture. However, the majority of the stress-strain curves for the clays do show a more or less marked peak, typically in the stress range 50-200kPa (Fig. 1). As the deforming clay approaches this peak region, fine lines, or very narrow corrugations, become visible on the exterior of the specimen. (They are observed by interrupting or prematurely terminating the test.) If laminations are present in the material, these begin to be displaced along the 'lines', and the boundaries of the specimen may also be displaced (Fig. 2). It can sometimes be seen, especially after additional strain, that in three dimensions the displacement is taking place along planes within the clay (Fig. 2). These are the shear zones of the present paper. In triaxial tests, the shear planes are obliquely inclined to the long axis of the cylinder. The sense of displacement along the planes corresponds to that of normal faulting, although localized reverse movement has been observed rarely. The planes themselves are shiny and finely grooved. In drier specimens (about 15-25% water content) the zones are flat-planar and regularly oriented, at angles as low as about 30 ° from the cylinder axis. In wetter materials (25-50% water content) the shear planes are typically about 40 ° from the cylinder axis, but here they are typically of a more complex arrangement, with no single representative angle. Corrugations representing the conjugate set of shear planes can usually be discerned, but there is commonly insufficient movement along this set for laminations to be visibly displaced. Probably, deviations from orthogonal geometry of the starting cylinder influence the extent to which the conjugate set is developed. The geometry of the shear zones is therefore analogous to shear fractures in rocks, but, except in very dry materials (less than about 15% water content), the shear zones described here do not exhibit fracture. At this scale of observation (and normally under the optical microscope also), there is no loss of cohesion across the planes. In direct shear tests, the macroscopic results have much in common with those of Morgenstern & Tchalenko (1967a) for pure kaolinite. They reported the progressive development of Riedel

79

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FIG. 1. Representative shear stress-shear strain curves to show the behaviour of various argillaceous materials in differing experimental conditions. (A) The possible effect of strain rate. Laminated ball clay in distributed shear; normal stress = 600 kPa. Groups of curves are shown stippled. The small range in water content of the lower curves, 27.5-31%, spans that of the upper curves, 28-29%, suggesting that the difference between the two groups of curves is due to strain rate. (B) Effect of normal stress (load normal to bulk displacement), values as shown. Laminated ball clay-kaolinite blocks in distributed shear; strain rate = 3.4 x 10- s s- 1. (C)-(F). The behaviour of various clays, under different normal stress conditions. The marked peaks in (D) and (E) are probably due to the sediments being, respectively, overconsolidated and of low water content. (C) 'Brick clay' (crushed Silurian shales from Buttington brick works, near Welshpool, Wales, GR SJ263099) in direct shear, strain rate = 3.4 × 10-4 s- 1. (D) 'Estuary clay' (cut from recent deposits on banks of the Leri, Ynyslas, west-central Wales, GR SN616934) in direct shear, strain rate = 3.4 x 10 -4 s- 1. (E) Kaolinite in direct shear, strain rate = 3.4 x 10 - 4 s - x, water content = 12-16%; (F) ball clay in distributed shear, strain rate = 3.4 x 10 - 5 s - 1. shears, 'thrust' shears, and, finally, a principal d i s p l a c e m e n t shear, all w i t h i n a n a r r o w central zone that allowed the d i s p l a c e m e n t o f the b o t t o m h a l f o f the kaolinite block past the top half. T h e utilisation o f listric Riedel shears m e r g i n g into principal shears to a c c o m p l i s h the bulk shear is

well illustrated in the present work, w h e r e contrasting coloured l a m i n a t i o n s w e r e i n t r o d u c e d during the sample p r e p a r a t i o n (Fig. 3A). ' T h r u s t ' shears h a v e only very rarely b e e n observed. T h e Riedel shears f o r m preferentially at the b o u n d a ries o f the specimen. T h e y t e n d to h a v e a listric

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A. Maltman

FIG. 2. Macroscopic appearance of shear zones in triaxially-deformed cylinders. Principal compression along the cylinder axis, parallel to marker laminations. (A) Water content = 28%, confining pressure = 3000 kPa, strain rate = 8 x 10- 2 s- 1. (B) Water content = 30% confining pressure = 2000 kPa, strain rate = 1 x 10- s s- 1.

FIG. 3. Macroscopic appearance of sheared blocks. (A) Photograph of a thin-section. Vertical marker laminations are displaced along shear zones (not themselves apparent in ordinary light), some of which have been indicated by black lines to show their overall listric Riedel geometry. Bulk displacement as shown. (B) Sheared specimen of laminated ball clay/ brown pot clay broken open along shear zone to show polished, striated surface of zone.

shape, even where the bulk shear is perpendicular to the sedimentary fabric and laminations. The same geometry of shear zones arises in the distributed shear tests. The bulk shear is largely achieved by the generation within the specimen of Riedel and principal shears. However, presumably because of the difference in test configuration, the Riedel shears are not usually concentrated at the specimen boundaries and the influence of any laminations present is greater. The planes maintain an overall listric shape, but may refract across laminations. In low watercontent samples, only one or two shear planes are produced, and cohesion may be lost across them. They are then able to accomplish large additional displacements without further shear zones being produced. In more ductile material, numerous Riedel shears are generated; they are visible on the specimen surface as very fine crinkles. At water contents of greater than 50% or so, the shears may be indiscernible macroscopically, in the absence of marker laminations, and the material can appear to be flowing homogeneously. Microscopic examination, however, reveals that the shear zone mechanism is still operating. If a sheared specimen is manually pulled apart, the shear surfaces appear to be polished and very finely grooved in the direction of movement (Fig. 3B). The surface is commonly not exactly planar, but crossed by fine offsets. These can be related to sub-structures within the shear zone which are not discernible macroscopically.

Shear zones in argillaceous sediments Microscopic appearance The geometric arrangement of the shear zones in a thin-section reflects the orientations observed macroscopically, and depends upon the kind of deformation test that produced the zones. In triaxial tests, the shear zones tend to form parallel to the planes of principal shear, and Riedel shears are of little significance. Therefore, because in these tests the sedimentary fabric was arranged to be either parallel or perpendicular to the principal compression, the shear zones tend to form oblique (at 45 ° or somewhat less) to the fabric. Shear zones produced in shear box experiments tend chiefly to be listric Riedel shears, and, if the sedimentary fabric is parallel to the bulk displacement, under the microscope they can be seen curving into parallelism with the fabric. The shear zones are approximately planar. In brittle conditions (especially in materials of low water content) they are sharply defined planes of fracture. In the ductile conditions of most of the present experiments, the zones bifurcate and anastomose, and at high bulk strains can develop a complex appearance. Some examples are shown in Fig. 4. In the materials tested, the shear zones typically have a width in the range 20-50 ~tm. This is probably chiefly a function of grain size (see 'Lithology' section). In clay samples with a greater content of clastic grains, the zones are less well defined, and as ductility of the argillaceous material is increased (e.g. greater water content) the zone margins become less sharp; but even here the total width rarely exceeds 100 ~m. Higher strains are accomplished by further movement along existing planes, or by the generation of additional zones, rather than by increasing the width of a zone. For example, in one sheared ball clay specimen a measurable displacement of 750 ~tm was produced along a single shear zone, but even at this very large strain the zone remained less than 100 ~tm in width. The fine-grained nature of the material makes observation of the fabrics within the shear zones difficult. A primary sedimentary fabric is commonly evident in the bulk of the specimen (Maltman 1981) from its approximately uniform extinction when viewed through crossed-polars. The extinction patterns within the shear zones, however, show that here the clays have achieved contrasting orientations. The sedimentary fabric swings sharply into the shear zone, reflecting the normal sense of displacement, and it increases in intensity. The internal shear zone fabric is subparallel, or at a low angle, to the shear zone margins. Consequently, the shear zones are

81

conspicuous when viewed in crossed-polars as brightly illuminated, narrow domains crossing the dark host material (e.g. Fig. 4A, 4B). Cracks which sometimes arise during the thin-section impregnation procedure seem to follow the sedimentary fabric of the bulk clay and these, together with occasional coarse phyllosilicate grains, also indicate a sigmoidal swinging of fabric across the shear zones. This pattern has much in common with the ductile shear zone fabrics of high-grade metamorphic rocks (e.g. Ramsay 1980). In addition, sub-fabrics are discernible within the shear zones. Spaced domains of aligned clay lie at low obliquity to the shear zone margins, pointing in the direction of movement (Fig. 5A). They are geometrically very similar to the shear bands of highly-strained metamorphic rocks (e.g. White et al. 1980). Some examples of shear zones into which the sedimentary fabric swings with a reverse sense of movement have been observed. It is not clear whether these are local accommodation effects within the deforming clay or artefacts of the testing arrangement. They could, for example, result from the elastic force of the stretched rubber confining sleeve acting on the deformed and weakened specimen. The vast majority of the shear zones seen during the present work showed normal displacement.

Effect of physical conditions Lithology The most striking feature of the shear zones is their consistency of appearance in a range of deformed argillaceous materials, although there are differences in detail. Homogeneous specimens produce a simpler pattern of shear zones than those with laminations. Presence of the latter generates strain heterogeneities which influence the timing of initiation, distribution and, to a smaller extent, orientation of the zones. Homogeneous materials develop shear zones when strained to near their peak strength. Laminated specimens which have been sheared, but terminated well before peak strength has been reached, show sets of small shear zones developing in the vicinity of the lamination, presumably because the strains are concentrating there. Shear tests with bulk shear parallel to laminations, tend to form shear zones near to a lamination boundary, into which the shear zone commonly curves. Only in poorly ductile and brittle conditions do the shear zones

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FIG. 4. Microscopic appearance of shear zones. Note that in (A), (C) and (D) the conspicuous black lines are cracks which formed during the impregnation procedure. They follow and help indicate the primary sedimentary fabric. (A) Single shear zone viewed in crossed-polars showing as a narrow bright domain of reoriented fabric. Kaolinite after triaxial deformation, water content = 30% confining pressure = 700 kPa but no drainage, strain rate = 8 x 10- 5 s- 1. (B) as (A), ball clay. Water content = 34~; confining pressure = 700 kPa but no drainage, strain rate = 1.6 x 10 -4 s- 1. (C) Triaxially-deformed ball clay and kaolinite. Shear zones have bifurcating pattern, are narrower and more sharply defined in ball clay. Experimental conditions as (A), but water content = 36% (D) Anastomosing pattern of shear zones in ball clay. Triaxial deformation, water content = 13~, confining pressure = 25 MPa but no drainage, strain rate = 5 x 10-7 s- 1. (E) Widely splaying pattern of shear zones in modern estuary clay. Direct shear, normal stress = 28 kPa, strain rate = 3.4 x 10- 3 s - 1. (F) Shear zones at high angle to laminations. Material and deformation conditions as (E).

cross the l a m i n a t i o n s . T h e s e effects are less n o t i c e a b l e in triaxial d e f o r m a t i o n . A l t h o u g h cylinders h a v e b e e n p r e p a r e d w i t h t h e i r long axes parallel to the s e d i m e n t a r y fabric, in b o t h this a n d t h e n o r m a l configuration, triaxial testing i n d u c e s shear zones at angles o f a p p r o a c h i n g 45 ° to the l a m i n a t i o n s , a n d at this o b l i q u i t y the p r e s e n c e o f the latter seems to h a v e little effect.

It s e e m s that, a l t h o u g h t h e difference in m e c h a n ical p r o p e r t i e s b e t w e e n t h e s e d i m e n t in a lamin a t i o n a n d t h a t in a d j a c e n t m a t e r i a l is very small, it n e v e r t h e l e s s p r o d u c e s a m e c h a n i c a l a n i s o t r o p y w h i c h is sufficient to influence t h e location a n d o r i e n t a t i o n o f the shear zones, as long as t h e y are f o r m i n g at low obliquity, p e r h a p s 20 ° or less. I n an organic-rich r e c e n t clay, cut f r o m the Leri

Shear zones & argillaceous sediments

83

FIG. 5. Aspects of the microscopic appearance of shear zones. (A) Detail of shear zone in ball clay, showing subfabric resembling shear bands. Crossed-polars. (B) 'Brick clay' (as Fig. 1C) showing shear zones in relatively fine matrix. (C) Shear zone in sheared ball clay with marker lamination. Macroscopic cohesion was maintained in that specimen remained intact after test, but at this scale of observation shear zone is brittle. Distributed shear, water content before test = 28~, but intermittent drainage, normal stress = 600 kPa, strain rate = 3.4 x 10- 5 s- 1. (D) Array of shear zones, including conjugate set, in 'slumped' specimen. Water content = 57%

estuary (Ynys-las, west-central Wales), and deformed by direct shear, the Riedel shears are highly oblique to the bulk shear direction (see Fig. 4E, F, and Discussion). The markedly peaked stress-strain behaviour (Fig. 1D) is probably due to this exhumed clay being overconsolidated. The width of the shear zones depends, to some extent, on grain size. The kaolinite used has a dominant grain size of 2 I.tm, whereas the ball clay is on average slightly finer: the shear zones are narrower and more sharply defined in the ball clay (Fig. 4C). Where the proportion of coarser particles is greater, such as in some of the pot clays used for the laminations, the shear zones are slightly wider, are more diffusely bounded and are more difficult to recognize. In these coarser clays, sub-structures within the shear zones have not been observed. Sheared samples of a brick clay which incorporates a range of grain sizes (25-2000 ~m) only revealed shear zones in what appears to be finer matrix (Fig. 5B) and, of the materials reported here, showed the least tendency to display peaked stress-strain curves (Fig. 1C).

Water content

The water content of argillaceous sediments largely governs whether the deformation is brittle or ductile. It also influences the tendency to show a peak strength and is the dominant control on the magnitude of that strength. Yet it is clear from the experiments that over a wide range of water contents, the fundamental deformation mechanism remains the production of shear zones. The characteristics of the zones vary in detail with different water contents, but it is their consistency of appearance that is more striking than any variability. In a few low strain rate unconfined triaxial tests, clays of very low water content (4-7%) have failed by axial splitting. Almost always, however, clays of water content less than approximately 15~ fail on inclined fractures in a pattern exactly analogous to faulting in rocks. In shear tests on sediments of this water content, listric Riedel fractures are generated. The fractures in both configurations are almost always very few in number. Stress-strain curves show very marked peaks (Fig. 1E).

84

A. Maltman

In these low water content materials, the failed specimen can literally fall apart when the experiment is disassembled. At slightly higher water contents, the fractured surfaces hold together, presumably by effects such as capillary forces and surface tension, although under the microscope at least some of the displacement surfaces are separated by voids(Fig. 5C). The water content above which brittle behaviour is not observed depends, to a small extent, on factors such as pre-existing fabric, strain rate and confining pressure, but generally, in the circumstances described here, fracture is not observed in deformed specimens of more than about 2 0 ~ water content. Here the strain is taken up on shear zones which, at the scale observed in the optical microscope, are able to maintain cohesion. There is no further fundamental change of mechanism as water content is increased, to at least about 60% In triaxial tests there is a general tendency for the inclination of the shear zones from the long axis of the cylinder (principal compression) to become close to 45 ° as the water content increases. An analogous change of angle has not been consistently observed in the shear tests. In all testing modes, the shear zones become more numerous and more closely-spaced in wetter specimens, and thus the materials acquire marked ductility (Fig. 6). The shear zones may form in closely-spaced pairs or complex domains (Fig. 4E, F), but it must be emphasized that even here the shear zones affect only a restricted part of the specimen, and the bulk of the material is undeformed. In specimens with layers of different

150

water contents the shear zones are produced preferentially in the wetter, weaker parts. Above about 5 0 ~ water content, ball clay is so weak that cylinders for triaxial testing are unable to hold their shape. The materials are vulnerable to damage during handling. Nevertheless, although it has so far not been possible to carry out rigorous tests at these very wet conditions, there are strong indications that the ductility is acquired by the production of shear zones closely similar to those in less wet materials. For example, a cylinder was only partially consolidated and dewatered, so that the central part retained a water content of 57% On removal from the consolidating apparatus, this central part 'slumped', and a thin section showed that this change of shape took place on shear zones (Fig. 5D). They are sufficiently numerous to form an array which also includes good development of the conjugate set, but, even with the macroscopic 'slumped' appearance, the flow is produced along the narrow zones of shear, and the larger part of the sediment is unaffected. Bulk stress and strain

Representative stress-strain curves are given in Fig. 1. Most of them show a peak, subject to the circumstances mentioned in the preceding sections. This peak strength varies from about 50 to 200 kPa and is typically reached at bulk strains in the range 5-15~. It is in this peak strength region that the shear zones are generated. As discussed earlier, there are macroscopic indications of this, and a series of shear tests which

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85

Shear zones in argillaceous sediments were terminated at various stages of progress and examined microscopically, clearly showed the shear zones arising as peak strength is approached (Fig. 7). Very weak and ductile sediments, such as those with high water content, show shear zones beginning to form as the fiat, ductile region of the stress-strain curve is approached, even though there is no peak. At strains greater than peak strength, most of the argillaceous sediments investigated here exhibit a residual strength. Here, additional increments of strain are undertaken by the material while sustaining a fairly constant stress difference. This is achieved by continued use of shear zones--either by further slippage within existing zones or the generation of new ones. This fiat region of the stress-strain curves exists until at least 30% bulk strain, the maximum that can be produced by the present equipment. No indications have been observed of strain hardening. Strain rate and confining pressure

It has so far been impossible, in practice, to produce suites of specimens of exactly identical water content, and such is the dominance of this variable that the roles of other factors are difficult to identify. The indications are that lower strain rates and lower confining pressures encourage more numerous, smaller shear zones, that is greater macroscopic ductility. Increased strain rate may prompt less ductile behaviour (as suggested in Fig. 1A), but its role and that of confining pressure has not yet been clearly defined. Drained tests at 60000 kPa on ball clay produced ductile shear zones not noticeably

different from drained tests without confining pressure. In shear experiments, increasing loads applied normal to the bulk shear direction, which because of the constraint on the specimen boundaries produced a quasi-confining pressure, increase the apparent strength of the specimen (Figs 1B-F, 6A) and lower the strain needed for shear zones to appear.

Discussion In the experiments, argillaceous sediments subjected to a stress approximately similar to their strength attempt to change shape by forming shear zones. In triaxial tests, the zones follow the planes of principal shear strain and in shear tests they tend to be induced in a Riedel shear orientation. The contrast is thought simply to be due to the geometric constraints imposed by the differing experiments. Similarly, the location of the zones is largely governed in experiment by the shape of the test specimen. For example, in triaxial tests the zones tend to pass through the central part of the cylinder; if conjugate pairs are generated they tend to be symmetrical about the central axis of the cylinder. Apart from these geometric constraints, the location of the shear zones is influenced by laminations in the specimen and any perturbations in them or the sedimentary fabric. Presumably the precise location is governed by the position of suitably oriented clay domains, as discussed below. The orientation of the shear zones is also influenced by the presence of any planes of mechanical anisotropy, such as the boundaries of laminations, especially when viewed at the

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(A) (B) (C) FIG. 7. Shear stress-shear strain curves to show the production of shear zones as the deforming material reaches its peak strength. Tests were terminated at various strains and the materials examined in thin-section for the presence or absence of shear zones. Width of stippled curves reflects some irreproducibility in successive tests. (A) Kaolinite in distributed shear. Water content = 34-36%, normal stress = 70 kPa, strain rate = 3.4 x 10- 5 s- 1. (B) Laminated ball clay in distributed shear. Water content= 31-33%, normal stress = 70 kPa, strain rate = 3.4 x 10- 5 s- 1. (C) as (B) but water content = 27.28%, normal stress = 600 kPa.

86

A. Maltman

microscopic scale. The macroscopic orientation is perhaps most simply interpreted by the Coulomb criterion (e.g. Lambe & Whitman 1969). The angle between the shear zones and the planes of principal shear stress is generally lower in wetter argillaceous sediments, which might be expected to have lower internal friction. Thus in shear tests, the zones tend to be nearer to horizontal in wetter sediments, and in triaxial tests they are nearer to 45 ° from the principal compression. The sheared natural clay, with shear zones at an unusually high angle to the bulk displacement (Figs 4E, F), is rich in organic material, including interlacing rootlets and plant debris which will increase the amount of interlocking and the interparticle friction. Higher effective pressures might be expected to produce a comparable effect, but in the drained tests so far carried out, this has not been clearly observed. Notwithstanding these differences in geometric results between shear and triaxial tests and various experimental conditions, the shear zones themselves are remarkably consistent in appearance. A general scenario of their production can be envisaged. The first increments of bulk strain in a deforming argillaceous sediment are probably taken up by slip and rotation of particles, or domains of particles, in an effort to increase packing efficiency. There will be greater scope for such adjustments in wet materials, and so the onset of other mechanisms will be delayed (e.g. the differing position of the peak strengths in Fig. 6A). However, this increased packing can only be achieved by a reduction in void ratio and water expulsion, which itself rapidly reduces the permeability of the sediment and thus curbs further water loss. Hence, the efficiency of this mechanism will be short-lived and alternative mechanisms will have to be sought. In some cases, microstructures not considered here might be produced. For example, a kaolinite with good sedimentary fabric oriented perpendicular to the bulk shear, appears to have accomplished some of the initial strain by forming kink-bands, which are seen to be displaced by the subsequent shear zones (Maltman 1977, Fig. 5B). In the anisotropic stress conditions induced by the tests, some clay particles will find themselves oriented close to the planes of maximum shearing strain, so that when the applied stress becomes sufficient to overcome the inter-particle friction they will start to slip. Once the displacement is established, it will present the easiest mechanism for further deformation and the bulk strain will be focussed on the slipping domains, forcing propagation of the displacement and the generation of a shear zone.

Other clay particles may be suitably oriented for rotation into the planes of maximum shear strain and to also take up slippage. This ability to rotate will be greater in high water content specimens and hence more shear zones are produced in wetter material. In less ductile conditions such as lower water-content sediments and faster strain rates, there is less opportunity for suitably oriented domains to be produced and exploited. There are consequently fewer shear zones and the specimen has a less ductile overall appearance. This effect is seen at its extreme in brittle conditions such as low water content sediments where particle rotation will be difficult, and the first clay domain to find itself suitably oriented for shear will take up much, or even all, of the bulk strain, with loss of cohesion across the zone. As the shear zone develops, the sedimentary fabric is deflected into the zone, thus further facilitating slip. A form of geometric strain softening arises (Platt & Vissers 1980). Large strains can then be accomplished along the shear zone. With additional strain, the obliquity between the shear zone margins and the internal fabric is reduced and the fabric, judging by the strength of its aggregate polarization pattern when viewed in crossed-polars, becomes more intense. The question of volume loss and water movements within the shear zones is the subject of current work. The sub-zones inclined to the shear zone margins (Fig. 5A) can be interpreted, in view of their geometric arrangement and the large strains within the shear zones, as shear bands (White et al. 1980; Platt & Vissers 1980), although such features are usually associated with the deformation of metamorphic rocks. It is suggested that in the sediments being considered here, at the particle scale, cohesion is lost during grain slippage and the behaviour is essentially brittle with Riedel shears being produced within the shear zones. Clay flakes and aggregates align themselves within these Riedel shear zones and are manifest under the optical microscope as the inclined domains. In this view, the shear zones are hybrid in mechanical nature, appearing ductile at the optical microscopic and larger scale, but being due to brittle behaviour at the particle scale. The fine offsets observed on the surfaces of shear zones therefore reflect the tiny local displacements along the Riedel shears within the main zone. The shear zones do not appear to grow in width during progressive strain and may actually decrease as the internal fabric intensifies. Therefore, additional strain can only be achieved by propagation of the length of the zone. If for some

Shear zones in argillaceous sediments reason this ceases to be possible, then presumably the zone locks and another shear zone is utilized for further movement. It may be speculated that eventually new shear zones are generated, although their nuclei were oriented unfavourably for initial selection. This effect has not been verified in the present experiments, which are restricted by their design to bulk strains less than about 30%, but it is predicted that material which has been subjected to significantly higher bulk strains will show more numerous shear zones. The flatness of the residual strength region of the stress-strain curves presented here suggests that any locking of existing shear zones is nicely balanced by the initiation of new ones. To summarise, although it has been surmised that water-rich argillaceous sediments deform by

87

pervasive homogeneous flow, the present experimental work strongly suggests that this is not the case. Instead, the particle slippage is concentrated in narrow, discrete zones, termed here 'shear zones'. The zones are able to accomplish very large displacements, so that bulk strains of at least 30% are taken up while leaving the greater part of the sediment undisturbed. No pervasive fabrics are generated. The shear zones are strikingly consistent in appearance over a range of clay materials and experimental configurations. Even clays with water contents of nearly 60~, which are extremely weak and ductile, deform in this way. It is proposed that the chief mechanism by which water-rich argillaceous sediments deform is the production of discrete shear zones.

References LAMBE,T. W. & WHITMAN,R. V. 1969. Soil Mechanics. Wiley, New York. MALTMAN,A. J. 1977. Some microstructures of experimentally deformed argillaceous sediments. Tectonophysics, 39, 417-36. , 1981. Primary bedding-parallel fabrics in structural geology. J. geol. Soc. London, 138, 475-83. , 1985. A laboratory technique for investigating the microstructures of deformed near-surface sediments. This volume. - - & FITCHES, W. R. 1982. Pre-cleavage shear zones in rocks and experiments. Mitt. Geol. Inst. E T H Univ. Zurich, Neue Folge, 239, 179-81. MORGENSTERN, N. R. & TCHALENKO, J. S. 1967a. Microscopic structures in Kaolin subjected to direct shear. Geotechnique, 17, 309-28. & TCHALENKO, J. S. 1967b. Microstructural observations on shear zones from slips in natural clays. Geotechnical Conf. Osol Proc. 1, 147-52.

PLATT, J. P. & VISSERS, R. L. M. 1980. Extensional structures in anisotropic rocks. J. struct. Geol. 2, 397-410. RAMSAY,J. G. 1980. Shear zone geometry: a review. J. struct. Geol. 2, 83-99. SKEMPTON, A. W. 1964. Long-term stability of clay slopes. Geotechnique, 14, 77-101. - - , 1966. Some observations on tectonic shear zones. 1st. Cong. Internat. Soc. Rock Mechanics Proc., Lisbon, 1,329-35.

TCHALENKO, J. S. 1968. The evolution of kink-bands and the development of compression textures in sheared clays. Tectonophysics, 6, 159-74. - - , 1970. Similarities between shear zones of different magnitudes. Geol. Soc. Amer. Bull. 81, 1625-40. WHITE, S. M., BURROWS, S. E., CARRERAS, J., SHAW, N. D. & HUMPHREYS,F. J. 1980. On mylonites in ductile shear zones. J. struct. Geol. 2, 175-87.

ALEX. J. MALTMAN,Department of Geology, University College of Wales, Aberystwyth SY23 3DB, Wales.

Faulting mechanisms in high-porosity sandstones; New Red Sandstone, Arran, Scotland John R. Underhill & Nigel H. Woodcock S U M M A R Y : Faults in the 'New Red' aeolian sandstones of Arran are unusual, firstly for occurring as closely-spaced (less than 1 m) often conjugate sets affecting large volumes of rock, and secondly for forming upstanding fault zones with numerous anastomosing strands of granulated rock, each preserving a small increment of slip. Anisotropy, such as bedding and cross-bed sets, has no discernible effect on fault behaviour. In contrast to the underlying Carboniferous rocks, large displacements are rarely concentrated on a single fault plane within the high-porosity sandstones. The proposed cause is slip-hardening of each fault after a very small displacement (less than 10 mm) causing the next slip increment to be taken up through undeformed rock rather than on the original plane. The common factor in recent records of similar faults elsewhere is their occurrence in high-porosity sandstones. Because of the low grain-contact strength, these rocks are partly analogous to unconsolidated sediment. The high porosity promotes high grain-contact stresses which induce rapid cataclasis during initial slip. Grain fracture and spalling of iron oxide coatings and quartz overgrowths produce a seam with reduced grain size, poorer sorting, higher angularity and lower porosity than the unfaulted rock. These factors collectively strengthen the seam because the coefficient of friction is increased, even though cohesion is reduced. This results in a Mohr failure envelope that lies outside the envelope of the undeformed rock for most stress states. A transient pore pressure increase in the fault seam may be important during slip. Rocks deformed by this slip-hardened faulting preserve a record of each increment of strain. If the displacement on each individual fault seam is the same, the geometry of the total fault systems is directly related to the bulk strain. Quadrimodal systems observed by us in Arran, and by others elsewhere, are probably a response to triaxial strain and show that bimodal 'Andersonian' fault systems are only special plane strain cases. If the bulk strain is irrotational, both the orientation and relative magnitude of the principal strains might be estimated.

Distinctive faults in ' N e w Red' aeolian sandstones on the Isle of Arran match faults recently described from high-porosity sandstones in North America. In this paper we introduce the Arran examples and discuss how they contribute to understanding faulting mechanisms in such rocks.

Terminology and previous work Previously described examples of faulted highporosity sandstones come from the Jurassic Entrada and Navajo sandstones of Utah (Aydin & Johnson 1978, 1983; Aydin & Reches 1982; Aydin 1978), the Ordovician Simpson Group in Oklahoma (Pittman 1981) and the Upper Triassic Wingate Sandstone in Colorado (Jamison & Stearns 1982). Recently, Bevan (1985) has described examples from Tertiary sands in southern England. A further example is known to us from Zakinthos, western Greece, where faults occur in

Pliocene sandstones that have been intruded by diapirs of Triassic evaporite (Underhill, in press). All these areas contain numerous separate granulated fault strands, approximately 1 m m in width and tens of metres in length, on which displacements are small (less than 10 mm). Subparallel, closely-spaced strands may form zones, millimetres to several metres wide, across which offset is usually less than a metre. These zones may occur in sub-parallel sets spaced tens of centimetres to metres apart and several nonparallel sets may make up the total fault system (see example from Arran in Fig. 2F). Aydin & Johnson (1983) termed the separate strands 'deformation bands' and aggregates of two or more parallel bands 'zones of deformation bands'. Pittman (1981) refers to the separate strands as 'granulation seams' while Jamison & Steams (1982) have termed them 'microfaults' or 'gouge zones'. In addition, discrete through-going surfaces of displacement are associated with some zones of highly concentrated strands. Aydin & Johnson

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 91-105.

91

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J . R . Underhill & N. H. Woodcock

(t983) termed these 'slip surfaces'. They are commonly striated in the direction of net slip and displacements across them are in the order of several metres or tens of metres. In view of the confusing current terminology, partly genetic, partly descriptive, we refer to separate granulated strands simply as 'faults' and to two or more closely-spaced and parallel faults as 'zones of faults'. The term 'slip surface' is retained for through-going discontinuities associated with zones of faults. The spatial relations between these three components suggest that they develop sequentially and all authors deduce that this results from strain-hardening. This led Aydin (1978) to study the faults in thin section. He recognized that the mineralogical composition of the fault was the same as that of the surrounding parent rock, but that there had been a grain size reduction within it. Using the scanning electron microscope, he found a constant chemical composition across each fault, but a greater concentration of major elements per unit area in the fault zone as a result of the large porosity reduction with respect to the parent sandstone. In addition, Aydin proposed a three-fold division of a fault. Our own observations in Arran confirm this subdivision and will be discussed later. A further important feature of faults in highporosity sandstones is their tendency to occur in regular multimodal systems (e.g. Aydin & Reches 1982). Four sets are sometimes seen rather than the two conjugate sets classically developed in plane strain (Anderson 1951). This is probably a response to triaxial boundary conditions, and is only recorded in strain-hardening fault systems. We discuss later this important clue to the general geometry of fault systems.

Regional setting of the Arran faults The New Red Sandstone in western Scotland and N W England was deposited in a series of ensialic basins controlled mainly by N W - S E striking faults (Fig. 1; McLean 1978). It ranges in age from possible uppermost Carboniferous through Permian and Triassic. The onshore Arran New Red Sandstone is mostly contiguous with that of the West Arran Basin (separated from the East Arran Basin by the Brodick Bay Fault, which probably runs onshore in NE Arran). Control on Permo-Triassic sedimentation by active faults has been postulated by Astin & MacDonald (1983), Clemmensen & Abrahamsen (1983) and McLean (1978), particularly on N E - S W striking faults such as the Highland Boundary Fault and the Plateau Fault (Fig. 1).

The sedimentology of the mainly continental New Red Sandstone rocks has been described by Gregory (1915), Piper (1970), Lovell (1971) and Clemmensen & Abrahamsen (1983). The faulting described in this paper is almost totally restricted to well-sorted aeolian sandstones dominating the lower part of the sequence. Faults in poorlysorted or muddy facies, or in underlying Carboniferous and Devonian rocks, are dominantly more conventional, widely-spaced fault zones each with large displacement. These spaced faults can be easily mapped in pre-Permian rocks (Fig. 1), but displacements become rapidly diffused upwards in the aeolian sandstones. Our view (Woodcock & Underhill, in press), is that this faulting mostly occurred in Tertiary time, during the intrusion of the North Arran Granite (Fig. 1, inset). The ballooning diapir reactivated the NE and N W striking fault systems that had earlier controlled Permo-Triassic sedimentation. Localized displacement on these faults propagated up into the aeolian sandstones and diffused upwards as a normal-fault dominated system in the stretched carapace to the intrusion.

Character of the Arran faults Field description Faults or zones of faults occur as conspicuous upstanding ribs in multiple sets (Fig. 2A-G). A common dip and strike is shared by members of each set, and they are separated by undeformed cross-bedded sandstone (Fig. 2G). Faults, or zones of faults, are not deflected by anisotropy formed by bedding, cross-bed sets, or reactivation surfaces. The fault zones serve to compartmentalize areas of undeformed sandstone which show varying degrees of cementation. This suggests that the zones act as barriers to the migration of fluids and in effect segment what is otherwise a superb reservoir rock. Slip surfaces are rare, but occur along the margins of some thick (greater than 0.5 m) zones of faults (Fig. 2D). Slickenside Iineations on these surfaces indicate oblique-slip movement. The amount of displacement on these surfaces was in all cases greater than 2 m. Displacements on the zones of faults range up to 1 m or more.

Thin section description Thin sections cut across faults in Arran are shown in Fig. 3(A, B). A three-fold subdivision (Fig. 4)

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similar to that described by Aydin 0978) can be recognized from them. The host sandstone consists of rounded, loosely packed grains with iron oxide coatings and a grain-contact cement. Its porosity is high and porespace is interconnected, suggesting high permeability. The outer zone of the fault is generally 0.5-0.75 mm in width. It contains some grains showing fracture or undulose extinction, but the majority are undeformed. The grain-contact cement has been ruptured and grains are more tightly packed than in the host sandstone. Porosity is reduced, but the outer zone is still well sorted. The inner zone of faults is generally less than 5 mm in width and is characterized by poor sorting and low porosity with a few large rounded grains sitting in a fine-grained matrix of spalled grain fragments, overgrowths and cement. Most of the grains within the zone show grain fracture. Point-counts have been carried out on thin sections cut across faults to assess quantitatively

the changes in grain size, porosity and angularity that occur across them (Fig. 5). A drastic reduction in the percentage of grains greater than 0.1 mm in diameter is seen when moving from the undeformed parent sandstone towards the inner zone (Fig. 5A, E). This decrease corresponds to an increase in the percentage of grains less than 0.03 mm in diameter (Fig. 5C, E). However, there does not seem to be any significant change in the abundance of grains between 0.03 and 0.1 mm in diameter (Fig. 5B, E). Porosity has also been evaluated by point-counts and by counting epoxy-resin impregnated pore space. A reduction in the percentage of pore space in the outer and inner zones of faults is apparent (Fig. 5D, E). Finally, angularity changes across faults have also been assessed by qualitatively assigning grains into three classes and carrying out pointcount traverses (Fig. 5F). In this case, a sharp rise in angularity is recognized in the inner zones of faults.

94

J. R. Underhill & N. H. Woodcock

FIG. 2. Photographs of fault zones within the New Red Sandstone of Arran. (A) Well defined zone highlighting anastomosing strands, Cock of Arran (NR956522). Scale bar (10 cm).

FIG. 2. (B) Numerous anastomosing strands developed within a diffuse fault zone, Corrygills (NS042354). (C) Well defined zone striking NE, Corrygills (NS042354). (D) Slip surface developed on margin to fault zone, Corrie foreshore (NS026430). (E) Diffuse zone of dip-bimodal strands, all having normal offsets, Brodick Old Quay (NS018378). Scale bar (10 cm), rod (1 m).

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96

J. R. Underhill & N. H. Woodcock

FIG. 2. (F) Complex fault pattern development, Pirates Cove (NS030396). (G) Undeformed parent sandstone, Lower Permian, Corrie foreshore (NS026431). Rod (1 m).

Interpretation of the faulting mechanism Our observations support previous deductions that the unusual character of faults in highporosity sandstones results from strain-hardening of each fault after only a small increment of slip. Theoretical Mohr diagrams can be used to compare a rock that strain softens at failure (Fig. 6A) to one that strain hardens (Fig. 6B). The first Mohr diagram (Fig. 6A) shows a failure envelope for the host rock (A) and a stress circle at failure. In most lithologies the resulting fault zone would have an envelope such as (B) with a similar friction angle, and a lowered or

zero cohesive strength. When stresses build again, the stress circle will touch envelope (B) before (A) and failure will occur preferentially along the already faulted zone. In this case the fault has strain-softened. Exceptions occur only where the fault or the stress system have substantially rotated. A strain-hardened fault must be represented by a failure envelope that lies outside that of the host rock (A) in the stress range of interest. This can occur in two ways: (i) through an increase in the cohesive strength (C) during failure with or without a change in friction angle, as shown by envelope (C), (Fig. 6B); (ii) through an increase in friction angle, ~, during failure, with or without a change in cohesive strength, as shown

Faulting mechanisms in high-porosity sandstones

97

FIo. 3. Thin sections across individual fault strands highlighting grain size reduction and poorer sorting in faults. (A) Several thin individual strands occur (less than 0.5 m m wide). (B) Grain comminution occurs over a wider zone as a result of fault strand amalgamation.

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: Some grains

inner zone

: Most grains fractured. Occasional rounded grains survive. Cement and grains spalled to give finer grained matrix. Poor sorting, low porosity.

broken. Tighter packing due to grain boundary slip after destruction of g r a i n - c o n t a c t cement. Good sorting, m o d e r a t e porosity.

FTG. 4. General characteristics in an idealized fault zone, close to its propagating tip. Full explanation in text.

by envelope (D), (Fig. 6C). In both cases renewed failure would occur in virgin host rock rather than on the pre-existing fault. We now investigate which of these effects is the more likely for the highly porous sandstones in Arran.

Effect of faulting on cohesion and friction angle Assessment of the mechanical processes operating along a fault (Table 1) suggests that, immediately after slip, it has a lower cohesion but higher friction angle than the host rock. Within the fault, soil mechanics principles (e.g. Lambe & Whitman 1969), can be used. Cohesion is lost throughout the fault, mainly by progressive rupture of weak grain-contact cement. The friction angle is increased in the outer zone by tightening of grain packing and in the inner zone by increasing angularity of particles and poorer sorting giving more intimate interlocking. The change in friction angle may be minor (less than 10°) partly because the loose packing and low interlocking give a rather low friction angle in the host rock. With reference to the Mohr diagram analysis, it appears that increase in friction angle alone could explain the inferred strain-hardening of the faults. It is also possible, though unproved, that minor recementation of grain contacts along

the fault could occur between slip increments, causing some recovery of cohesive strength and further enhancing the hardening of the fault. Two other effects well known in granular aggregates are relevant here. Firstly, uniform coarse sands show a high rate of granulation during slip due to the low number of grain contacts per unit volume and resulting high graincontact stresses. As grain size and sorting decrease, there is a rapid fall off in grain contact stresses and granulation rate, perhaps explaining why a few large grains may survive within the inner zone of a fault once they are surrounded by fine material. Secondly, smaller grain sizes show narrower slip zones, explaining why slip becomes concentrated in a narrow zone in the centre of the fault.

Stress-strain path for one slip increment In Fig. 7, curve A is a typical stress-strain curve for a rock or a dense uncemented sand that strain softens, whereas curve B is for a loose-packed sand that strain hardens. For uncemented sands that are not granulating curves A and B converge at a packing density (the 'critical state') at which the sand can deform continuously without further porosity change (Lambe & Whitman 1969). Curves C, D and E represent the stress-strain

Faulting mechanisms in high-porosity sandstones

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FIG. 5. P o i n t - c o u n t d a t a across faults. (A) G r a i n sizes g r e a t e r t h a n 0.1 m m in d i a m e t e r . (B) G r a i n sizes 0 . 0 3 0.1 m m in d i a m e t e r . (C) G r a i n sizes less t h a n 0.03 m m in d i a m e t e r . (D) I m p r e g n a t e d p o r e space.

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(F) FIG. 5. (E) Gives compilation of point-counts (n = 1508). (F) Gives point-counts (n = 150) for angularity across faults. behaviour of progressively finer, more poorly sorted sediment. The postulated stress-strain path (over a fraction of a second) for a high porosity weakly-cemented sandstone initially follows curve A, with recoverable elastic behaviour. At strain a (Fig. 7), the grain contact cement starts to fracture and the grain framework to collapse. The newly formed fault zone initially strain-softens due to the progressive loss of cohesion and transient lower friction angle (factor

8, Table 1). A texture, like that in the outer zone, forms (Fig. 4). The fault would behave as though it were an uncemented sand and its deformation path would follow curve A, were it not that the burial depth was enough to cause grain fracture. In this case, when the grain framework can collapse no further (at strain b), grains start to fracture, initially at a high rate but decreasing as grain size decreases. Friction angle increases and the stress-strain path is driven up again by the

F a u l t i n g m e c h a n i s m s in high-porosity sandstones

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TABLE 1. Summary of physical factors influencing cohesion and friction angle in high-porosity sandstones. The probable

dominant factors are in upper case

COHESION (C)

is low in high-porosity sandstones because of ( a ) l o w ~' cement, ( b ) l o w ~o clays and ( c ) l o o s e packing giving low interlocking of grains. Outer zone (= low strain)

Factors that tend to increase C

Factors that tend to

1.

Denser packing gives greater grain interlocking and apparent cohesion.

2.

Any spelling of clay films to grams increases Z clay and resulting electrostatic cohesion,

3.

Possible recementation at grain contacts between slip increments.

4. RUPTURE OF GRAIN-CONTACT CEMENT DURING

decrease C

SLIP.

FRICTION

Factors that tend to increase

Inner zone (=high strain)

ANGLE

5.

11.

Grain size reduction increases ,~ clay and silt grades, and increased number of contacts per unit volume gives greater electrostatic cohesion,

12.

Possible recementation at grain contacts between slip increments.

13.

RUPTURE OF REMAINING GRAIN-CONTACT CEMENT.

14.

Grain size reduction gives less grain interlocking.

(0)

is lOw in high-porosity sandstones because of ( a ) l o o s e packing (above the critical void ratio) gives collapse on failure, (b) low interlocking of grains. .. Outer zone (=low strain) Inner zone (=high strain)

DENSER PACKING (WITHOUT GRAIN SIZE DECREASE) GIVES GREATER INTERLOCKING OF GRAINS.

15.

GRAIN FRACTURE GIVES MORE ANGULAR FRAGMENTS AND GREATER GRAIN-CONTACT FRICTION.

16.

DECREASE IN SORTING GIVES MORE INTIMATE GRAIN INTERLOCKING AND LOWER POROSITY.

17.

Possible increase in clay percentage.

m Factors that tend to decrease

6. Rupture of cement leaves grains free to roll. 7. Clay or iron oxide coatings may lower grain-contact friction. 8. Collapse of grain packing gives transient lower friction.

strain-hardening effect. It crosses curves C, D and E, at a strain appropriate to the grain size at the time. Finally, the fault zone locks at strain d as it strain-hardens to a point at which its frictional resistance exceeds the shear stress along it. An obvious problem with this stress-strain path is that the lock-up stress must exceed the initial failure stress, Ty, if net strain-hardening has occurred. At first sight, this stress overshoot seems unlikely. There are two possible explanations. The first is that stress overshoot does not occur, the fault locking up when the frictional resistance reaches the initial failure stress (at strain c), or at even lower stresses. The second is that apparent stress overshoot occurs because of transient pore pressure increase along the fault. In this case, it is possible for the shear stress on the fault to remain constant or even decrease but still promote slip because the effective normal stress is being lowered by the increasing pore pressure. On a Mohr diagram (Fig. 8), this

situation is represented by successive stress circles being driven to the left to track a series of failure envelopes with increasing friction angle. After failure, the excess pore pressure dissipates and the stress circle moves to the right, well away from the failure envelopes. Another effect of the increase in pore pressure is to promote extensional failure, or the development of conjugate hybrid shear fractures [in which dihedral angles are less than 60 ° (Hancock 1985)], in rocks where some cohesive strength remains. Whilst long-term abnormal pore pressures seem unlikely in these porous and permeable sandstones, transient high pore pressures along the faults seem very likely during slip. Pore fluid has to escape from both outer and inner zones of the fault as porosity decreases, and its escape is progressively slowed, especially in the inner zone, by decreasing permeability. Transient overpressuring is therefore a plausible cause of apparent stress overshoot.

J. R. Underhill & N. H. Woodcock

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(A)

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.<_ STRAIN ~ S T R A I N SOFTENING HARDENING

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FIG. 7. Stress-strain path during failure of a strainhardening fault compared with paths for (A) a rock of dense unconsolidated sand and (B, C, D, E) loose packed (progressively finer and more poorly sorted) sands that strain-harden.

__L__

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.

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or FIG. 6. Mohr stress diagrams for faults that (A) strain soften and (B, C) strain harden. Explanation in text.

Fault patterns and bulk strain Strain-hardened fault zones potentially record each increment of slip and each new fault will form in an 'ideal' orientation to adjust the bulk shape of the rock body. The fault system will preserve a full record of the bulk strain. In contrast, strain-softened faults may each accommodate many slip increments, and may slip when less than ideally oriented to the stress system. Our conclusions on faulting mechanisms in this paper are largely independent of the age or regional origin of the faults and are dependent only on a correct interpretation of the physical state of the host sandstones during deformation. The fault pattern in the Arran New Red Sandstone is rather complex and will be detailed elsewhere (Woodcock & Underhill, in press). We mention here only selected observations relevant to strain-hardened fault patterns in general. Two quantitative observations support previous views (Jamison & Stearns 1982) that the

frequency of faults should correlate with the magnitude of bulk strain. Firstly, the frequency of faults is greatest close to the northern granite (Fig. 1) and decreases radially away from the pluton to the SE and SW. This fits expectations of bulk strain decrease from finite-element models of diapirs (Dixon 1975). Secondly, there is a local increase in fault frequency in the New Red Sandstone above discrete faults mapped in the underlying pre-Permian basement (Figs 1 and 2). This effect is consistent with a strong control on local strain intensity and orientation by basement faults propagating up into the porous sandstones. Jamison & Stearns (1982) described a similar basement control on faults in the Wingate Sandstone, Colorado, and Bevan (1985) in Tertiary sandstones of the Isle of Wight. In Arran, as in those examples, the field evidence suggests

< 3

<:

friction angle 2 1

increasing

%ff

,

pore pressure

increasing

FIG. 8. Successive Mohr stress circles and failure envelopes (1, 2, 3) for a fault with rapid transient increase in pore pressure during failure.

Faulting mechanisms in high-porosity sandstones that the basement displacements diffuse upwards within a few tens of metres into much wider fault zones in the porous sandstones. The Arran faults typically occur in conjugate sets with a low dihedral angle. The average angle that can be resolved on the contoured stereoplots is 36 ° (one standard deviation interval of 25-47 °, Fig. 9A), but smaller angles are commonly observed directly in the field. These would equate with hybrid shear fractures of Hancock (1985) in which the dihedral angle is less than 60 ° . This confirms other observations on strain-hardened faults in experiments (22-50 °, Aydin & Reches 1982) and in nature (40 °, Aydin & Reches 1982; Jamison & Stearns 1982). These angles are lower than expected from conventional Coulomb analysis (Anderson 1951). Hancock (1985) proposes that such hybrid shear fractures form in the shearextension transition. It is likely (Jamison & Stearns 1982; Aydin & Reches 1982) that the fault patterns are a kinematic response to displacement boundary conditions and/or are controlled by poorly understood rheological laws (Reches 1978). The simplest fault systems in the New Red Sandstone of Arran give approximately bimodal patterns of poles to faults. These may be strikebimodal (Fig. 9B) or dip-bimodal (Fig. 9C) depending on the orientation of the bulk strain axes. However, at outcrop many approximately bimodal systems may be seen to be significantly

103

multimodal, with, for instance, each set of a dipbimodal system itself containing two subsets with slightly different strike. This gives a lozenge pattern on both horizontal and vertical surfaces (Fig. 2B, E). Sometimes the dihedral angles between sets are large enough to discriminate a multimodal pattern of poles on a stereoplot (Fig. 9D). Multimodal patterns seem to be a response to triaxial boundary conditions and bimodal patterns will only develop in plane strain conditions (Reches 1978; Aydin & Reches 1982; Hancock 1985). This can be explained by analogy with slip systems in crystals (Taylor 1938). Patterns with orthorhombic symmetry would probably develop in ideal pure shear, but with a simple shear component some modes may be suppressed (e.g. Jamison & Stearns 1982). We have integrated experimental data (Aydin & Reches 1982) and theoretical analysis (Reches 1978) with observations in Arran to produce a qualitative guide to the fault pattern for varying bulk strain (Fig. 10). This is shown for a bulk strain with the intermediate axis, Y, vertical (giving dominantly strike-slip on each fault) but the same geometric relationship between the fault pattern and the strain axes would hold for X vertical (dominantly reverse dip-slip) and Z vertical (dominantly normal dip-slip). The particular complexity of the Arran faults arises from their control by underlying basement faults that themselves form two sets striking at

°l Oo

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~

n=47

n =53

FIG. 9. (A) Histogram of dihedral angle between conjugate fault sets, measured from modal data on stereoplots. (B, C, D) Examples of fault patterns shown as poles to faults on lower hemisphere equal area projections. Contours are at 5 and 10 points per 100/n ~ area.

104

J. R. Underhill & N. H. Woodcock I

FIG. 10. Diagrammatic variation in contoured patterns of poles to faults with bulk strain. Bimodal patterns in plane strain (for zero volume change) become quadrimodal for other strains. high angles to each other. Different systems of basement faults are initiated by each of the b a s e m e n t orientations and the bulk strain accomm o d a t i o n is partitioned between the two systems. These complexities will be discussed elsewhere (Woodcock & Underhill, in press).

Conclusions The m a i n conclusions to be d r a w n from this study are:

1 Faults in the N e w R e d Sandstone of Arran m a t c h examples elsewhere in high-porosity sandstones in occurring as discrete strands of granulated rock each with small slip. 2 A textural change from unfaulted rock into the centre of each fault involves progressive rupture of grain contact cements, tightening of packing and reduction of grain size by fracture. This spatial change corresponds to a temporal change during fault propagation. 3 E a c h fault effectively strain-hardens because although cohesion is destroyed, the denser packing, decrease in sorting and more angular fragments increase the friction angle. 4 A transient pore pressure increase along the fault is probably important during propagation, but dissipates i m m e d i a t e l y after slip. 5 The geometry of the fault system is controlled by regional boundary conditions. Bimodal patterns reflect plane strain, but multimodal patterns are more c o m m o n and indicate a general triaxial strain. ACKNOWLEDGEMENTS; This project was partly selfsupporting but we are particularly grateful to British Petroleum Co. Ltd, for providing the majority of the financial support. Mr and Mrs E. J. W. Underhill and the Department of Earth Sciences, Cambridge also contributed. We thank Professor Mike Brooks, Dr Graham Williams, Dr Philip Allen, Dr Rod Gayer (all University College Cardiff), Dr Tim Bevan (Bristol) and Dr Geoff King (Cambridge) for their constructive criticism of earlier drafts of this paper. In addition, we would like to acknowledge Dr Alan Smith (Cambridge) for his comments and for helping to initiate this project, and Ms Rosemary Gigg for her logistic support. Mrs Margaret Millen is thanked for her help with the drafting.

References ANDERSON, E. M. 1951. The Dynamics of Faulting.

CLEMMENSEN, L. B. & ABRAHAMSEN, K. 1983. Aeolian

Oliver & Boyd, Edinburgh. ASTIN, T. R. & MACDONALD, D. I. M. 1983. Syndepositional faulting and valley-fill breccias in the Permo-Triassic of Arran. Scott. J. Geol. 19, 47-58. AYDIN, A. 1978. Small faults formed as deformation bands in sandstone. Pure appl. Geophys. 116, 91330. -& JOHNSON, A. M. 1978. Development of faults as zones of deformation bands and as slip surfaces in sandstone. Pure appl. Geophys. 116, 931-42. - & JOHNSON, A. M. 1983. Analysis of faulting in porous sandstones. J. struct. Geol. 5, 19-31. - & RECHES, Z. 1982. The number and orientation of fault sets in the field and in experiments. Geology, 10, 107-12. BEVAN, T. G. 1985. Tectonic evolution of the Isle of Wight: a Cenozoic stress history based on mesofractures. Proc. Geol. Ass. 96, 227-35.

stratification and facies association in desert sediments, Arran basin (Permian), Scotland. Sedimentology, 30, 311-39. DIXON, J. M. 1975. Finite strain and progressive deformation in models of diapiric structures. Tectonophysics, 28, 89-124. GREGORY, J. W. 1915. The Permian and Triassic rocks of Arran. Trans. geol. Soc. Glasgow, 15, 174-87. HANCOCK, P. L. 1985. Brittle microtectonics--principles and practice. J. struct. Geol. 7, 437-57. JAMISON, W. R. & STEARNS, D. W. 1982. Tectonic deformation of Wingate Sandstone, Colorado National Monument. Amer. Assoc. Pet. Geol. Bull. 66, 2584-608. LAMBE,T. W. & WHITMAN,R. V. 1969. Soil Mechanics. Wiley. LOVELL, J. P. B. 1971. Petrography and correlation of sandstones in the New Red Sandstone (PermoTriassic) of Arran. Scott. J. Geol. 7, 162-9.

Faulting mechanisms in high-porosity sandstones MCLEAN, A. C. 1978. Evolution of fault-controlled ensialic basins in northwestern Britain. In ."BOWES, D. R. & LEAKE, B. E. (eds) Crustal Evolution in Northwestern Britain and Adjacent Regions, 325-46. Seel House Press, Liverpool. - - & DEEGAN, C. E. (eds) 1978. The solid geology of the Clyde sheet (55 N/6 W). Rep. Inst. Geol. Sci. No. 78/9.

PIPER, D. J. W. 1970. Eolian sediments in the basal New Red Sandstone of Arran. Scott. J. Geol. 6, 295-308. PITTMAN,E. D. 1981. Effect of fault-related granulation on porosity and permeability of quartz sandstones,

I05

Simpson Group (Ordovician), Oklahoma. Amer. Assoc. Pet. Geol. Bull. 65, 2381-7. RECHES, Z. 1978. Analysis of faulting in three-dimensional strain field. Tectonophysics, 47, 109-29. TAYLOR, G. I. 1938. Plastic strain in metals. Jour. Inst. Metal. 62, 307-24. UNDERHILL,J. R. in press. Triassic evaporites and PlioQuaternary diapirism in western Greece. J. Geol. Soc.

WOODCOCK, N. H. & UNDERHILL, J. R. in press. Emplacement-related fault patterns around the Northern Granite, Arran, Scotland. Bull. geol. Soc. Am.

J. R. UNDERHILL, Department of Geology, University College, PO Box 78, Cardiff CF1 1XL, Wales. Present address: Shell Internationale Petroleum Maatschappij. B.V. c/o Shell (UK) Exploration and Production, Shell-Mex House (Annexe), Little Adelphi Building, 10 John Adam Street, Charing Cross, London WC2R 0DX. N. H. WOODCOCK, Department of Earth Sciences, University of Cambridge, Downing Street, Cambridge CB2 3EQ.

Morphology and microstructures of hydroplastic slickensides in sandstone Jean-Pierre Petit & Edgard Laville S U M M A R Y : The High Atlas Mountains (Morocco) provide local basins with good Triassic exposures. Basin formation was induced by normal and strike-slip synsedimentary faults, either sealed by Mesozoic and Cenozoic rocks, or reactivated during Alpine orogeny. Thus the Triassic outcrops in these fault zones show not only slickensides obviously due to rupture and friction mechanisms in brittle material (reverse and strike-slip faulting during Alpine orogeny), but also 'hydroplastic' slickensides due to faulting in incompletely lithified sediments (normal and strike-slip faulting during Mesozoic pull-apart basin formation). Microstructural and morphological comparison between these two types of slickensides enable us to describe various specific criteria for 'hydroplastic slickensides'.

Sedimentary basin formation induces deformations such as thickness variations, progressive unconformities, collapse structures, olistolites, change in lithofacies and so on. Most of these appear linked to synsedimentary fault movements. The role of normal faulting in grabens was first recognized, but more recently the 'pullapart' basin model (Crowell 1974) emphasized the interaction of dip-slip and strike-slip faulting, especially in intracontinental basins often associated with main strike-slip fault zones (Arthaud et al. 1977; Reading 1980). When this model is used for interpretation, one must differentiate between synsedimentary normal and strike-slip faulting. This is generally done on the basis of indirect data such as cartographic fault patterns, seismic refraction profiles, spatial organization of thickness variations, facies and currents (Steel & Gloppen 1980). Direct observations of slickensides and faults which occurred during basin formation have never been used as far as we know, although they provide one of the most important ways of checking whether subsidence is linked to dip-slip, strike-slip or intermediate faults. This involves finding out the orientation and directions of relative movement on proven synsedimentary fault planes. Moreover, systematic measurement of such planes may give the stress state determination during basin formation (Etchecopar et al. 1981). Despite its importance, this sort of slickenside has not often been described (Davies & Cave 1976), probably because in most cases the fault planes are either obliterated by more recent layers or disturbed by further movements. The aim of this paper is to show that they can be systemati-

cally recognized by visual and microscopic features in specific conditions. Our interest in various types of slickensides arose during faulting tectonic studies in the Tizi n'Test zone, High Atlas, Morocco (Petit 1976; Proust et al. 1977). This is a weak crustal zone with a dense fault network located between the Hercynian belt and the West African craton (Mattauer et al. 1972). The faults involved were active at least from Palaeozoic to Alpine times in various stress conditions. We tried to determine the successive reactivations of the fault pattern by systematic observation of slickensides on major and associated minor faults in successive stratigraphic levels. Corresponding regional principal strain axes were deduced by graphical (Arthaud 1969) or computer methods (Etchecopar et al. 1981) applied to the sets of slickensides. A great number of observations carried out in various rocks (except limestone) enabled us to draw up a set of criteria for the direction of movement on faults (Petit et al. 1983). Triassic rocks provide very good illustrations of these criteria, mainly on strike-slip and reverse faults linked to the Alpine shortening which obviously affected lithified sandstone. Similar criteria have been found on a few normal faults. Studying the Mesozoic distension (Trias and Lias) with the pull-apart basin model in mind, we looked for specific Triassic 'soft' fault plane marks in proven Triassic synsedimentary faults; thus we were able to distinguish between 'brittle' and 'soft' (or 'hydroplastic') rheological behaviour of material. We prefer the term 'hydroplastic' for reasons explained below. For the sake of convenience we will use it to describe the corresponding faults and slickensides.

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 107-121.

Io7

108

J.-P. Petit and E. Laville

Sedimentary and tectonic setting of the studied outcrops In the High Atlas, the Massif Ancien (mainly Precambrian and Palaeozoic rocks) and surrounding areas (Fig. 1) provide good Triassic exposures, easily observed because of semi-arid conditions. They are the remains of larger areas eroded as a result of the Alpine orogeny (block tectonics). Nevertheless, isolated lozenge-shaped local basins with grabens or asymmetrically infilled half-grabens are recognizable, showing massive red sandstones (Fig. 2). These form a central sandstone facies along an E - N E orientated trough linked to the Tizi n'Test fault zone, and are dated as Late Carnian (Cousminer & Manspeizer 1976); they were deposited in a continental to near-shore environment (Biron 1982). They are overlain by shallow marine siltstone capped by Lower Lias tholeiite flows. On the north and south margin, sandstone thickness drastically diminishes (Fig. 2). Northwards only evaporitic siltstone and tholeiites c a n be seen, forming a siltstone marginal facies. This central sandstone facies includes most of the faults we describe. Fault planes are also

TINE MAL [TyF] MAIN FAULT

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Horizontal bedding Planar cross bedding Trough cross bedding Channel Slump structure Angular pebble Salt mould Bird's eye Bioturbation Plants Concretion Calcrete Malachite

FIG. 2. Schematic logs and sedimentological features of sandstone formation with thickness changes in the Tine-Mal zone.

FIG. 1. Structural map of the studied Triassic outcrops. M. A. = Massif Ancien; C = Central (calcareous) High Atlas; dots = Triassic outcrops; horizontal lines = Palaeozoic and Precambrian; 3 and 5 refer to corresponding figures.

present in basal conglomerates (BC) as well as in the Lower (LS) and Upper Siltstone (US), but here it is difficult to see the difference between brittle and hydroplastic faults (see below). The main sedimentological features of the sandstone formation (Biron 1982; Beauchamp & Petit 1983) are shown in Fig. 2. These sandstones usually exhibit grain sizes in the range 100-700 ~tm. In thin section, rounded quartz grains (generally 80%) which show haematite coated grains and overgrowths are cemented by illite and chlorite. Moderately weathered K-feldspar and plagioclase, a few oxides and calcite are also present. The importance of dip-slip Triassic faulting in the formation of the sandstone trough was first

Morphology of 'hydroplastic slickensides' demonstrated (Proust 1962, Brown 1980, Biron 1982 and Tixeront 1973), mainly by comparisons between thicknesses and facies analysis of each Triassic block limited by N60-70 to N90 trend faults (as in Fig. 3B). Different stages of subsidence are apparent in the restored section on Fig. 3(A). One difficulty with this type of comparison is that most of these Triassic faults were more or less reactivated during the Alpine continental convergence in the West Mediterranean area. The recent Alpine tectonics (Mio-Pliocene) mainly consist of reverse faulting as in Fig. 3(B), with varying vertical throws (from 0 to 1000 m on scissor faults), so that no recent significant shifting of sandstone blocks is expected. However, previous post-Triassic left-lateral strike-slip faulting has been discovered in the nearby Central High Atlas (Fig. 1) during Jurassic pull-apart basin formation (Laville & Harmand 1982), and left lateral N60-70 strike-slip faulting in the Massif Ancien (throw less than 10 km) is thought to be a result of early Alpine shortening (Proust et al. 1977). Thickness variations can thus be a result of post-Triassic juxtaposition of originally distant sandstone deposits, so it is very important to find non-reactivated Triassic faults in the field. Non-reactivated Triassic dip-slip faults are characterized by block tilting and their throw diminishing upwards mainly in Upper Siltstone. A demonstrative site is the north border of the Talaat n'Yacoub local basin (Fig. 1) which shows resurrected fault scarps where hydroplastic slickensides can be observed. In the Tine-Mal outcrops (Figs 3 and 4) post-Triassic tilting left Triassic fault scarps nearly horizontal. In surrounding outcrops one can see that the most important throw is in the sandstone layers, but it dies out upwards in the Upper Siltstone; while siltstone facing the fault shows bed fragmentation and disharmonic folding at the base, the upper layers are unaffected and lie parallel to the general dip. Thus this throw appears localized during the deposition of the base of the upper siltstone. These dip-slip faults with throws of 10-100 m are conjugated with the much more important TineMal main fault (TMF, Fig. 3B), which controlled

S,,iMFFi ~ ~

FIG. 3. Cross-section in the Tine-Mal zone (A), at the end of Triassic deposits, and (B), after Alpine shortening; for letters see Fig. 2.

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.

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109

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lOOm FIG. 4. Block diagram of minor Tine-Mal Triassic fault scarps after post-Triassic southwards tilting; circles = basal conglomerates and lower siltstone; dots = sandstone. Right = cross-section when upper siltstone was not eroded. Arrow = zone where fault scarp has been completely eroded. Detailed stratigraphic section in Fig. 2, right. thickness variations in the massive sandstone. Similar structures showing the association of Triassic minor faults conjugated with major faults controlling thickness variations in sandstone are frequently found (Beauchamp & Petit 1983). Major faults are more often reactivated as they are more important zones of weakness. The importance of Triassic left-lateral strikeslip faulting in a pull-apart model was more recently shown on the basis of observation and measurement of dip-slip and strike-slip hydroplastic slickensides (Laville & Petit 1984). A good example of strike-slip syn-sedimentary slickensides which helped to establish the pull-apart model is provided by a main western branch of the Tizi n'Test fault, which dies out westwards in the Triassic layers of the Argana basin (Fig. 1), without cutting overlying Jurassic layers. A feather-shaped ending (Cloos 1955) associated with progressive unconformities and thickness variations can be seen (Fig. 5). Hydroplastic slickensides are present at the contact between the Palaeozoic rocks and the sandstone. We have observed such slickensides elsewhere along N60120 trends, so some of the N60-70 faults acted in both normal and strike-slip fashion. The superposition of these two types of hydroplastic slickensides has been observed on the same outcrops.

Morphological criteria for hydroplastic slickensides and sense of fault movements Most morphological types of slickensides we describe were observed in the Tine-Mal and

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FIG. 5. Structural map of the western end of the Tizi n'Test fault; aligned dots and dip symbols outline progressive unconformities and folding; dots = sandstone; broken lines = upper siltstone overlying sandstone; thick vertical lines = Palaeozoic basement. Argana sites mentioned above. They illustrate not only the hydroplastic state of the sediments but also definite criteria for relative fault movement. We were able to see the difference between hydroplastic slickensides formed in an incompletely lithified sediment and brittle slickensides

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of Alpine age formed in the corresponding solid rocks because we had already established criteria for the sense of relative movement for the latter (Petit et al. 1983). Morphological comparisons between both types of faults in the field help to point out rheological contrasts in the faulted material. The main points of comparison are summarized in Table 1, so we merely comment on them here. At first sight, the most striking general contrast is the patina of the fault plane: brittle faults typically show striated shiny smooth slickensides (Fig. 6A, C), or at least variously shaped white patches (Fig. 6B, D) which stand out in red sandstone; these are due to the crushing of quartz grains with a throw of less than 1 cm. In contrast, hydroplastic slickensides are the same colour and as rough or granular as the surrounding surfaces of strata. W h e n the throw is small (cm to dm) as is frequently observed on boundary faults of minor grabens (Fig. 7), the striation is poorly marked, if present at all. Moreover, the two compartments are always welded together (arrowed on Fig. 7) where the

FIG. I5

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FIG. 6. Main types of Alpine fault plane marks observed in solid sandstone and used as relative movement criteria (A, B). Numbers refer to the table and text. Arrows indicate the sense of movement of the missing block. (C) Right-slip fault of Alpine age in Triassic sandstone. The secondary faults dip at a slight angle in the direction of movement of the removed block as on (A) (opposite sense). The mean fault plane appears flat, striated and shiny. (D) Left-slip fault of Alpine age in Triassic sandstone. The white patches are facing the movement of the missing block as on (B).

Morphology of 'hydroplastic slickensides' TABLE 1.

III

ComparisonbetweenAlpine and Triassicfaults Faults in solid sandstone (brittle behaviour)

Hydroplastic faults (in the corresponding uncompletely lithified sediment)

Mean fault plane

On minor faults, compartments can usually be On minor faults, compartments cannot be split split apart apart Mean surface is usually planar Fig. 9: Mean fault planes are quite often distorted or dislocated, especially on strikeslip faults Fig. 14: Collapse (convex bending) of fault planes

Patina

Shiny or with white patches due to crushing of Most often the same as surrounding rock quartz grains surfaces; sometimes a muddy aspect

Striation

Well marked and straight, even for very small Absent or poorly marked for small throws (cm throws (mm to cm) to dm)

Minor structures * +__similar

*specific

Fig. 6(A): Frequent striated secondary R Fig. 9(A, B, C): Many secondary Riedel-type (Riedel-type) planar (1) or with concave (R) faults, often distorted and even overintersection towards the top (2), dipping at turned a slight angle in the direction of movement Fig. 9(E): A few grooves and trails left by clay of the removed block pellets (1) which have been torn out (2) and spread along (3, 4) Fig. 6(B): Very frequent striation with whiten- Fig. 9(D) : Occasional grooves (2) closing up ing due to crushing of quartz grains on the quickly behind a harder striating element side of asperities (1) facing the movement (1) of the removed block Fig. 9(F): Bumpy surfaces due to intersection of minor fracture patterns oblique to the slip direction Fig. 9(G): Pads (1) accumulated at the front of a now removed block (2)

Microstructures of shear zones Very thin for small throws or reduced to one More or less wide and spread out or a few sliding planes with microbreccia Fig. 16(A, B): Secondary (R) fault zones made of welded quartz fragments usually fairly wide, with grain size reduction and isolated intact quartz grains Fig. 18 : Tendency towards particle orientation with long axis of quartz grains __+perpendicular to local shortening Fig. 16(C): Some stylolite-type structures Microfracturing

Fig. 15" Under the surface, frequent tension Figs 17, 18: Angular crushed grains in shear micro-cracks dipping at 25°+ 10° towards zones, but no orientated fracturing on the the direction of the removed block walls

fault plane penetrates inside the rock; in some cases, therefore, hydroplastic planes are difficult to distinguish from joints cutting the strata. Large-scale hydroplastic slickensides (several metres square) have also been observed, for example on tilted dip-slip faults in the Tine-Mal area (with decametric throw) (Fig. 4), or on strike-slip faults in the A r g a n a area (with kilometric throw (Fig. 5). Some of the secondary structures w h i c h have developed are also found on brittle slickensides, for example Riedel-type

secondary faults (R dipping about 20 ° inside the wall in the direction of the m o v e m e n t of the missing block). On brittle faults (Fig. 6A, C) they are more or less planar and well-striated, with the m a i n fault surface practically fiat; on hydroplastic slickensides they are usually more m a r k e d : On the dip-slip faults we observed the m a i n fault surface is usually irregular because these secondary R shear m o v e m e n t s induce step-like structures (Fig. 8) or swollen c o m p a r t m e n t s b e t w e e n t h e m (Figs 9A and 10B). These shears are

I 12

J.-P. Petit and E. Laville

FIG. 7. Minor grabens associated with the Tine-Mal tilted faults; for curved arrow, see text.

sometimes overturned or dislocated (Figs 9B and 10A, B). On the strike-slip faults, dislocations are more pronounced, probably because of the more important throw. Dihedral repeated structures sometimes of metric scale (Figs 9C and 11 A, B) seem to be due to the intersection of decimetric R (synthetic shears dipping about 20 °) and R' (antithetic shears dipping about 70 °) Riedel shears, or due to the exaggerated step-like structures formed by R shears on pre-existing joints (Fig. 12). They are often contorted (Fig. l lB) and sometimes worn (Fig. l l C ) by the movement of the other compartment. This shows that the material with its pre-existing Riedel-type structures formed at an early stage of deformation could have been dragged along, dislocated, contorted or worn by the friction of the now removed block. This means that the material behaved in a 'soft' or 'plastic' way. Linear marks have occasionally been observed on dip-slip or strike-slip hydroplastic faults on various scales. While on brittle faults the grooves left by a striated tool are typically straight and elongated, here (Fig. 9D) the groove behind is short and irregular (2) because the affected material has been rapidly pushed back into it behind the striating tool (1). In Figs 9(E) and 13 striation has been caused by a clay pellet (1, Fig. 9E) which has been torn (2) forming a tapering

FIG. 8. Step- (Riedel-type) structures on a hydroplastic normal fault. groove, and spread along (3 and 4), showing that clay pellets were still plastic and capable of striating very soft material. Some structures are only found on hydroplastic slickensides, for example, when the surface shows conjugated sets of fractures oblique to the throw direction (Fig. 9F), which are reminiscent of Cloos' extensional experiments in clays. Stylolitetype structures and disorganized blocks are observed several centimetres beneath the surface. On strike-slip faults we observed (Fig. 9G) decimetric scale pads (1) are obviously due to the deformation of sandy clays before a now removed block (2). These two features (sets of fractures and pads), specifically found on hydroplastic slickensides, as well as the large dislocated Riedeltype structures described above, show that the hydroplastic state was not limited to a narrow shear zone. This is borne out by the collapse of slickensides several metres square seen on dipslip faults forming grabens of hectometric scale at the top of upper sandstone (UPS, Fig. 2). As recent erosion has removed overlying upper siltstone (US, Fig. 2), one can see convex fault planes with more or less cylindrical curvature perpendicular to the slip direction (Fig. 14). This collapse indicates a pervasive low viscosity of the

Morphology of 'hydroplastic slickensides"

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I 14

J.-P. Petit and E. Laville

FIG. 10. (A) Step- (Riedel-type) structures on a hydroplastic normal fault, but overturned (same fault as Fig. 8). (B) Hydroplastic slickenside on Triassic normal fault with Riedel-type structures deformed irregularly. material which formed the present sandstone block.

Microstructural observations. Comparison with brittle slickensides and experiments To establish microstructural differences between brittle and hydroplastic slickensides, we compared thin sections from an Alpine fault (Fig. 15 corresponding to detail of A & B fault type on Fig. 6) and a Triassic fault (Figs 16A, B, 17, 18 and 19 corresponding to details of a Riedel-type

secondary fault, example A in Fig. 9). Both samples were taken from the same sandstone in the Tine-Mal area. In the brittle fault section (Fig. 15), well orientated tension cracks dipping at 25°+10 ° towards the direction of movement of the removed block always appear under the sliding surface, even if a few crushed and recrystallized grains are present at the surface. These cracks seem homologous with 'microscopic feather fractures' in experimental faulting tests under triaxial stresses (Friedman & Logan 1979; Conrad II & Friedman 1976). Such orientated fracturing is rare in hydroplastic fault sections under the slickenside (compare

Morphology o f 'hydroplastic slickensides'

I 15

FIG. 11. Hydroplastic strike-slip faults in the Argana area. (A) Dihedral blocks formed by Riedel-type secondary faults; horizontal fractures are due to Alpine shortening. (B) Detail of 11A showing contorted secondary faults. (C) Wearing and distortion on the same faults. Arrows indicate the sense of the missing block.

with Fig. 17 located on Fig. 16A) or in any zone of the section. On the hydroplastic fault section (Fig. 16A, B) one can see Riedel-type shear zones of varying widths (rather than well defined friction planes) which appear darker internally than outside the zones. Non-fractured grains are present both inside and outside the shear zone, which is mainly composed of various-sized angular fragments (Figs 18 and 19). Imbricated structures are observed on the contact between

intact coarse elongated grains (Fig. 19). Realignment of the long axes of grains is observed within a wide shear zone, as well as outside it (Figs 16B and 18). The shear zone may show either a clear-cut boundary between crushed and intact grains (UTZ) or a transition zone (LTZ). Comparable features have been described in brittle fracture experiments on high porosity sandstone (Dunn et al. 1973). But applied confining pressure (25-100 MPa) and differential stress

I 16

J.-P. Petit and E. Laville

FIG. 12. Riedel-type secondary shear in the Tine-Mal area.

FIG. 13. Grooves resulting from clay pellets torn out (curved arrow) and spread along on a minor fault plane; upper compartment (bottom right) shifted to the right.

Morphology of 'hydroplastic slickensides'

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rounded shape could be the result of protection from stress concentrations by the smaller surrounding grains, as in Mandl's sugar pack experiments. But it is more likely to be the result of lubrication by clay at the contact between moving grains, especially as associated grains do not show Hertzian radial cracks at their contact. Also, imbricated structures of elongated grains are fluidal structures. The dark internal aspect of the shear zone could be due to grain size reduction and weathering rather than clay enrichment. In Fig. 18 the clear-cut limit (UTZ) of the shear zone could be a gliding plane. The lower limit with crushing gradually diminishing downwards (LTZ) could be a zone of fracturing arising from grinding because of more cohesive upper and lower zones in non-dilatant conditions. This grinding (cataclasis) could be due to high local

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stresses resulting from episodic local grain interlocking during shearing. It could be encouraged by hydraulic fracturing: water over-pressure can be expected because of the interbedded and overlying clays in the studied zone. In some cases (as in Fig. 16C, from sample F, Fig. 9) sections show dislocated fragments (df) and stylolites (s), indicating local shortening (z).

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The observations presented above demonstrate that the morphological features as well as microstructures imply a macroscopic ductility of the sheared material. It can only be explained by grain boundary sliding, mainly controlled by the clay phase in a water saturated material. For this reason, we have termed this behaviour 'hydroplastic', extending its use to hydroplastic faults or hydroplastic slickensides for convenience. It could include various rheological states during lithification, from the quasi-liquid to quasi-solid behaviour of Elliott's classification (1965) and is more specific than 'soft' in the case of argillaceous rocks. Most of the hydroplastic minor structures on slickensides presented above, form a set of consistent criteria for sense determination: several of them can be observed on the same fault and all of them indicate the same sense, as long as they result from the same throw. These criteria can help with tectonic analysis, but only if two precautions are taken. Firstly, one must carefully distinguish between tectonic hydroplastic faults (linked to major faults on the basement of basins) and faulting due to gravity sliding in unconsolidated layers (Fig. 20). These types are very similar for minor faults. Secondly, one cannot always presume early- (contemporaneous with sedimentation) or late-faulting from the observation of hydroplastic slickensides alone. Argillite interbedded with sandstone layers does not seem to show striking differences between Triassic and Alpine faults: in both cases slickensides are shiny with irregularly striated surfaces, and examples of hydroplastic slickensides due to recent tectonics have been observed in Triassic argillaceous Upper Siltstone. One must exclude observations in argillaceous rocks which cannot be truly lithified. Even if the material can be lithified, one must bear in mind that a more or less hydroplastic state may have persisted for several million years. Lithification may have been very slow indeed, especially if there was water over-pressure in sedimentary basins. One must have some observations of sealed faults and other structural data for more definite conclusions. The above exam-

FIG. i7. T h i n section showing the type o f grain f r a g m e n t a t i o n u n d e r a hydroplastic slickenside. S c a l e = 500 gm.

U,T.Z.

S.Z.

L.T.Z.

FIG. 18. T h i n section in a shear zone (located in Fig. 15A); U T Z = u p p e r transition zone; SZ = shear zone; L T Z = lower transition zone. Scale = 1 m m .

120

J.-P. Petit and E. Laville

FIG. 19. Thin section in a shear zone (located on Fig. 15). Scale = 500 p.m.

pies s h o w h o w r e a c t i v a t i o n c a n p a r a d o x i c a l l y help to reveal such n o n - r e a c t i v a t e d faults. T h e types o f slickensides we h a v e p r e s e n t e d do n o t f o r m an e x h a u s t i v e list. M o r e o v e r , descriptions a n d i n t e r p r e t a t i o n s h a v e to be c o m p l e t e d . H o w e v e r , we h o p e a n a l o g o u s s t r u c t u r e s will be f o u n d to serve as guides for structural geologists s t u d y i n g f o r m a t i o n m e c h a n i s m s o f clastic sedim e n t a r y basins. ACKNOWLEDGEMENTS. We would like to thank most sincerely the reviewers and also Drs W. D. Means and M. Seguret for all their encouragement and helpful suggestions. We are also most grateful to Mr B. Sanche for the thin sections. We are indebted to A. T. P. Plis et Failles for financial support of this work.

FIG. 20. Block diagram showing hydroplastic minor faults linked to tectonics (left) or to gravity sliding of the unconsolidated bed (right).

References ARTHAUD, F. 1969. M6thode de d&ermination graphique des directions de raccourcissement, d'allongement et interm6diaire d'une population de failles. Bull. Soc. gkol. France, 7, XI, 729-37. --, MEGARD, F. & SEGURET, M. 1977. Cadre tectonique de quelques bassins s6dimentaires. Bull. Centre Rech. Exploration Production, Elf-Aquitaine, Pau, 1, 147-88. BEAUCHAMP, J. & PETIT, J. P. 1983. S6dimentation et Taphrogen6se Triasique au Maroc: l'exemple du Haut Atlas de Marrakech. Bull. Centre Rech. Exploration Production, Elf-Aquitaine, 7, 389-97. BIRON, P. E. 1982. Le Permo-Trias de la Rbgion de

l'Ourika ( Haut Atlas de Marrakech, Maroc). Th~ge de Troisi~me Cycle, Grenoble, France. BROWN, R. H. 1980. Triassic rocks of Argana valley, southern Morocco, and their regional structural implications. Bull. Amer. Assoc. Petrol. Geologists, 64, 988-1003. CLOOS, E. 1955. Experimental analysis of fracture patterns. Bull. geol. Soc. Amer. 66, 241-6. CONRAD II, R. E. & FRIEDMAN, M. 1976. Microscopic feather fractures in the faulting process. Tectonophysics, 33, 187-98. COUSMINER, H. L. & MANSPEIZER, W. 1976. Triassic pollen date Moroccan High Atlas and the incipient

Morphology of 'hydroplastic slickensides' rifting of Pangea in Middle Carnian. Science, 191, 943-5. CROWELL, J. C. 1974. Origin of late Cenozoic basins in southern California. In: DICKINSON, W. R. (ed) Tectonics and Sedimentation: Soc. econ. Paleont. Mineral., Special Publication, 22, 190-204. DAVIES, W. & CAVE, R. 1976. Folding and cleavage determined during sedimentation. Sed. Geol, 15, 89-133. DUNN, D. E., LAFOUNTAIN, L. J. & JACKSON, R. E. 1973. Porosity dependence and mechanism of brittle fracture in sandstones. J. geophys. Res. 7 8 , 14, 2403-17. ELLIOTT, R. E. 1965. A classification of subaqueous sedimentary structures based on rheological and kinematical parameters. Sedimentology, 5, 193209. ETCHECOPAR,A., VASSEUR,D. & DAIGNERES, M. 1981. An inverse problem in microtectonics for determination of stress tensors from fault striation analysis. J. struct. Geol. 1, 51-65. FRIEDMAN, M. 1975. Fracture in rock. Rev. Geophys. Space Phys. 13, 3,352-8. & LOGAN, J. M. 1979. Microscopic feather fractures. Bull. geol. Soc. Amer. 81, 3417-20. LAVILLE, E. 1981. R61e des dbcrochements dans le m6canisme de formation des bassins d'effondrement du Haut Atlas marocain au cours des temps triasiques et liasiques. Bull. Soc. gbol. France, 7, 303-12. - & HARMAND, C. 1982. Evolution magmatique et tectonique du bassin intracontinental m6sozoique du Haut Atlas (Maroc): un mod61e de mise en place syns6dimentaire de massifs "anorogbniques" li6s fi des d6crochements. Bull. Soc. gbol. France, 7, 213-27. - & PETIT, J. P. 1984. Role of synsedimentary strikeslip faults in the formation of Moroccan Triassic basins. Geology, 12, 424-7. MALTMAN, A. 1984. On the term 'soft-sediment deformation'. J. struct. Geol. 6, 5, 589-92. MANDL, G., DE JONG, L. N. J. & MALTHA, m. 1977. Shear zones in granular material. Rock Mechanics 9, 95-144.

I2I

MATTAUER, M., TAPPONNIER, P. & PROUST, F. 1972. Major strike-slip faults of late Hercynian Age in Morocco. Nature, 237, 160-2. PETIT, J. P. 1976. La zone du d~crochement du Tizi n'Test ( Maroc ) et sonfonctionnement depuis le Carbonifbre. Th6se de Troisi6me Cycle, Montpellier, France. --, PROUST, F. & TAPPONNIER, P. 1983. Crit6res de sens de mouvement sur les miroirs de faille en roches non calcaires. Bull. Soc. gkol. France, 7, XXV, N ° 4, 589-608. PROUST, F. 1962. Etude stratigraphique, pktrographique et structurale du Bloc Oriental du Massif Ancien du Haut Atlas (Maroc). Th~se Science, Montpellier, France. --, PETIT, J. P. & TAPPONNIER, P. 1977. L'accident du Tizi n'Test et le r61e des d6crochements dans la tectonique du Haut Atlas occidental (Maroc). Bull. Soc. gbol. France, 7, XIX, 3, 541-51. READING, H. G. 1980. Characteristics and recognition of strike-slip fault systems. In: BALLANCE,P. F. ,ee READING, H. G. (eds) Sedimentation in ObliqueSlip Mobile Zones. Int. Assoc. Sediment. Spec. Publ. 4, 7-26. RIEDEL, W. 1929. Zur Mechanik geologischer Brucherscheinungen ein Beitrag zum Problem den Fiederpatten. Cent. Blatt Mineral. GeoL Paliiont., pp. 354-68. STEEL, R. & GLOPPEN, T. G. 1980. Late Caledonian (Devonian) basin formation, Western Norway: signs of strike-slip tectonics during infilling. In: BALLANCE,P. F. & READING, H. G. (eds) Sedimentation in Oblique-Slip Mobile Zones. Int. Assoc. Sediment. Spec. Publ. 4, 79-103. TCHALENKO, J. S. 1968. The evolution of kink-bands and the development of compression textures in sheared clays. Tectonophysics, 6, (2), 159-74. TIXERONT, M. 1973. Lithostratigraphique et min6ralisations cupif6res et uranif6res syng6n6tiques et famili6res des formations d&ritiques permo-triasiques du couloir d'Argana (Haut Atlas Occidental, Maroc). Notes du Service Gbologique du Maroc, 33, 147-77.

J. P. PETIT & E. LAVILLE,Laboratoire de G6ologie Structurale, U.S.T.L., Place E. Bataillon, 34000, Montpellier, France.

Soft-sediment microfaulting related to compaction within the fluviodeltaic infill of the Soria strike-slip basin (northern Spain) Michel Guiraud & Michel S~guret S U M M A R Y : As much as 8 km of fluvio-deltaic strata accumulated in the Sorio strikeslip basin during the Late Jurassic-Early Cretaceous (Wealdian). The strata were bent in a huge syn-sedimentary syncline related to the development of an extensional half-graben in the basement. The alluvial plain siltstones and clays are cut by a large number of microfaults of a special type in regard to their geometry and their morphoscopic characters. The fault surfaces concave upward dip at 10°-40 ° to bedding with a strong dispersion of fault plane strikes, providing a characteristic circular pattern. They are closely spaced and cut the material into small bi-pyramidal cone-shaped units with axes normal to bedding. The fault surfaces are glossy with a metallic glint and a fine striation without crystallization. From electron microscopic observations, this is due to phyllite re-orientation. Along the fault planes the displacement is always dip-slip and slight. Analysis of fault striation populations by a computer-aided method gives a unique stress tensor responsible for the striation with the maximum compressive stress al normal to the bedding plane along the whole section of the synsedimentary syncline and a R ratio (R = (a2 0-3)/(O"1 -- 0"3) ) close to 0. The orientation of the stress tensor, the R = 0, i.e. 0"2= 03 and the morphological aspect of the fault surfaces similar to shear planes affecting present cohesive soils lead us to conclude that this deformation is related to compaction in cohesive clays contemporaneous with water escape structures in sandstones. This deformation occurs in the evolution of the basin fill from soft sediments to metamorphosed rocks. In recent years, a new branch of geoscience has been concerned with the description of features formed by the deformation of non-lithified sediments containing an important percentage of water, and with differentiating them from geological structures affecting rocks. The clear distinction between the two types of deformation will provide several significant geological applications concerned with the structural evolution of sedimentary basins. On the one hand, deformations of unlithified sediments by water escape have been described and a relationship between these structures and seismic activity of faults has been proposed (Sims 1973, 1975; Montenat 1980). On the other hand, hydroplastic microfaults disrupting unconsolidated sediments have been defined in the fluvial Permo-Triassic deposits of the western Atlas range in Morocco (Petit 1976; Laville & Petit 1984; Petit & Laville, this volume), and described in the fluvial Saxonien deposits of the Permian basin of Lodeve in southern France (Santoui11980), in the fluvial deposits of the Soria Basin (Guiraud 1983) and in the Upper Cretaceous Helminthoid flysch of the Autapie Nappe in the French Alps (Labaume, this volume). In the Soria basin, the water escape structures in fluvial sandstone bodies are associated with hydroplastic microfaults in interlayered clayey siltstones.

It is our intention: (i) to describe water escape structures and their relationships with microfaults; (ii) to describe the geometry and the morphoscopic characters of the microfaults using macroscopic and electron microscopic observations; (iii) to characterize precisely the criteria for the recognition of hydroplastic microfaults in non-lithified silt and clay and for the discrimination of them from post lithification faults; (iv) to compare these faults with data provided by shearing experiments on present cohesive soils; (v) to determine the type of stress state related to hydroplastic microfaulting by statistical microtectonic analysis. It is concluded that this type of deformation is related to compaction and water escape. This specific deformation had the weight of the sediments as the driving force. An eventual triggering by earthquakes, consistent with the tectonic setting of the basin, is not documented.

Sedimentary and tectonic setting of the Soria Basin The Soria Basin is located in the north-western Iberic Range, now included in the Alpine Belt of Spain (Fig. 1). The basin developed during the

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 123-136.

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Soft-sediment microfaulting, northern Spain

extensional syn-depositional structures within the basin (Fig. 4) (Guiraud & S6guret 1985). A model of releasing overstep has been proposed with different stages in the development of the basin. The N 060°E margins acted as sinistral strike-slip faults and the N 135°E trends functioned as dip-slip faults locating Wealdian depocentres. Outside the Soria Basin, at the tips of each major strike-slip fault, the Demanda and Montcayo massifs (Fig. 1) were compressed and consequently uplifted, supplying detrital material of the basin (Guiraud & S~guret 1985).

series is reduced to 500 m maximum and even 0 m; and in many places, an important erosional phase occurred during Wealdian time as, for instance, in the Moncayo area. The Soria Basin displays a rhomboidal shape and is bordered by N 60°E and N 135°E trends. In the studied area, restricted to the northern part of the basin, strata of Wealdian cyclothems III and IV are deformed into a large asymmetric syncline whose axis, striking N 135°E, lies along the northern border. The sedimentary pile of the southern limb, 5000-6000 m thick is reduced to 700-1000 m alongside the northern limb (Fig. 3). These data, associated with progressive unconformities, demonstrate the syn-depositional development of the asymmetric syncline. According to pre-tilting extensional microstructures (stylolites, tension gashes, normal microfaults), we can identify an Upper Jurassic to Lower Cretaceous N 040°E extensional tectonic event inside the basin. Consequently, we assume that the basin developed as a synformal half graben basin related to half graben formation within the basement by normal faulting along the NE margin (see Jubera, Arnedillo Fault; Fig. 2) (Guiraud & S6guret 1985). On NE-SW sections A and B (Fig. 3), during cyclothem I sedimentation, the southwards displacement of depocentre suggests a progressive southwards back-faulting of the southern border. In contrast, the successive northwards migration of depocentre during cyclothems III and IV deposition involves a northwards back-faulting of the northern border. Depending on structural geological investigations, we characterize the temporal association of compressive synsedimentary folds located along the N 060°E borders of the basin and

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Water escape structures in sandstones The bedding surfaces of the coarse-grained sandstones in cyclothem III are deformed by upstanding ridges corresponding to sand volcanoes locally bounded by roughly circular troughs 40-80 cm in diameter. In vertical section, the lamination of the original trough cross-bedding and current ripple structures are upturned and overturned into decimetric to metric synforms. Sandstone synforms are loaded into underlying clay and siltstones which are upwardly injected into antiformal structures with their axial traces perpendicular to bedding (Fig. 6). These structures are very similar to water escape features as defined by Lowe (1975) and described in ancient sedimentary environments, especially sandy fluvial deposits (Selley 1969; Burne 1970; Hubert et al. 1972). According to Lowe (1975), expulsion of depositional pore fluid by seepage, liquefaction and fluidization is an efficient process for deforming unconsolidated layers.

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Hydroplastic microfaults in clays and siltstones Location Numerous curved microfault planes have been observed, cross-cutting the grey or red alluvial marly claystones of cyclothems III and IV. Geochemical data from an unpublished internal report (SNEAP) have demonstrated that the alluvial plain deposits are essentially composed as follows: quartz (5-35%), calcite (0-20%) and an argillaceous phase which is represented by dominant illite associated with ferriferous chlorite. A few fractures showing an analogous morphoscopic type of fault surface are also present

within some yellowish bioturbated dolomicrite layers of the uppermost cyclothem III.

Geometry The microfault surfaces, whose lengths range from a few centimetres to a few metres, are curved, generally concave upward and dip 50 °10° (Fig. 5). The dihedral angle between conjugated microfaults is always significant (110 °130°). Caused by noticeable striking distribution, these fractures define a characteristic wedgysheared geometry and form bi-pyramidal coneshaped slices with axes of the cones perpendicular to tilted bedding planes (Figs 5 and 7). Bi-pyramidal units fit together perfectly with the load casts and rounded troughs and necks of water escape structures which are located at the limits of the neighbouring sandstones layers (Fig. 8). The unequivocal geometrical relationship between these two types of deformation enables us to conclude that microfaults in clays and water escape load marks in coarse-grained sandstones are genetically related.

Morphoscopic characters of the microfault surfaces From macroscopic observations, fractures in claystones are characterized by bumpy, irregular

Soft-sediment microfaulting, northern Spain

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planes formed by curved narrow striated grooves, commonly a few millimetres in width and several decimetres in length (Figs 7 and 11). Lustrous microfault surfaces are glossy, displaying a characteristic metallic brightness. Brilliant striated irregular surfaces with pronounced divergent slickensides generally alternate with dark uniform areas defined by a diffuse striation. These microfault planes always show a patina similar to those of the sediments cropping out. Moreover, the fault surfaces are characterized by the lack of synkinematic fibrous crystal growth (quartz or calcite). From electron microscopic observations, the microfault surfaces provide a significant reorientation of ragged phyllites parallel to the slip plane. A well defined microscopic slickenside direction seems related to linear stepped re-

I z7

arrangement of clay particles and to the occurrence of straight grooves on the failure surfaces (Fig. 9). Well crystallized minerals are absent. Consequently, the typical shiny macroscopic aspect of these fractures is probably based on face to face disposition of phyllites along microfault planes. On sections perpendicular to the slip surface and parallel to macroscopic slickensides, the original disposition of the clay minerals, defined by a network of small aggregates connected by links of particles with associated pores, is preserved below the shear plane. In contrast, a generalized reorientation of flexured phyUites disposed parallel or nearly parallel to the microfault surface occurs within the shear plane (Fig. 10). All these macroscopic and microscopic characteristics allow us to discriminate these faults from usual brittle faults in rocks (Petit et al. 1983). Criteria of displacement sense on hydroplastic microfaults in claystones-siltstones

In order to establish the sense of movement on hydroplastic fractures, a set of three macroscopic fault plane marks and two microscopic tectoglyphs have been selected and used. They are: 1 The elongated prod mark as defined by Tjia (1968) and Laville & Petit (1984). Locally, some calcareous pedogenetic nodules have been displaced and elongated along the fault plane during shearing. Pluck marks due to ball detachment and push-up microridge related to ball stopping characterize dip-slip microfaulting (Fig. 11).

FIG. 6. Water escape structures on lower surface of a sand bed with load cast structures and upward injection of clay.

128

M. Guiraud & M. Sbguret

FIG. 7. Hand specimen of the upper half of a bi-pyramidal cone-shaped unit generated by hydroplastic microfaults in claystones. The axis of the cone is normal to the bedding. The cone is 60 cm width. Location: cyctothem III claystones, Yanguas (see Fig. 2).

2 The Riedel-like secondary microfault planes (Petit 1976; Petit et al. 1983 ; Petit & Laville, this volume). The principal microfault plane presents numerous secondary striated microfractures dipping at 15°-20 ° inside the main microfault plane. By analogy with'Riedel type' shear (Riedel 1929), the secondary microfractures dip in the direction of displacement of the absent block. 3 An original criterion analogous to the 'striated asperities' criterion in sandstones, as determined

by Petit et al. (1983). The surface microreliefs, located on the side facing the displacement are generally dark, uniform and associated with a diffuse striation (see Fig. 7) due to a significant crushing of phyllites. The asperities opposed to movement are irregular, lustrous and characterized by well defined curved slickensides linked to reorientation of clay particles by shearing. During the electron microscopic study of microfracture surfaces, we have defined and

FIG. 8. Cross-section of curved and decimetric spaced microfaults in claystones underlying a sandstone bed whose upper surface is deformed by sub-circular troughs associated with sand volcanoes. (Side of view is 3 m across.) Location: Cyclothem III, Yanguas.

Soft-sediment microfaulting, northern Spain

12 9

FIG. 9. Electron micrograph of hydroplastic failure surface showing a vertical striation. Claystones of cyclothem III, Yanguas. Scale is 60 ~tm long. detailed many microscopic fault marks providing other criteria of displacement sense. 4 Numerous prod micromarks related to clay particle displacement during microfaulting have been observed (Fig. 12). 5 Along sections perpendicular to the slip plane, several secondary microscopic faults dip inside the major striated plane in the direction of the movement of the removed compartment (a, Fig. 10). By comparison with macroscopic fault

marks, these secondary microfaults are defined as Riedel-type microscopic faults. All these criteria are mutually consistent and provide the same information. Displacements determined by microscopic and macroscopic tectoglyphs are always coherent and consistent. In the field, the three macroscopic criteria can be used either together or independently; however, criterion 1 is the most useful.

FIG. 10. Electron micrograph of a 3-D view of a section perpendicular to slip plane and parallel to the macroscopic striation. From macroscopic criteria, the missing upper fault compartment has been displaced to the left; a--secondary Riedel-type microfault. Claystones of cyclothem III, Yanguas. Scale is 10 ~tm long.

I3o

M. Guiraud & M. Sbguret

FIG. 11. Morphoscopic features of an hydroplastic slickenside. Striae are curved; there is no crystal growth; the surface has a metallic glint; the missing hangingwall moved towards the left; the cavity on the fault plane is dissymmetric with a pressure ridge on the left side.

Amount of displacement on hydroplastic microfaults The hydroplastic microfaults in the claystonesiltstone succession of the Soria Basin do not have a cartographic expression. At outcrop scale, the microfaults cutting through siltstones do not cut through underlying and overlying sandstones beds. The fault planes have a metric extension and a slight offset. Pedogenetic nodules have been lightly sheared on to the fault surfaces and

elongated along centimetric distances. In some places, root traces in palaeosoil surfaces provide other good displacement markers. Generally, the root traces are cut by hydroplastic microfaults with very low-angle dip (close to the perpendicular to the root trace, i.e. close to parallel to the bedding). The root traces are displaced with offset ranging from 0.1 to 1 cm. All these observations are consistent and show that the mean displacement along the microfaults is millimetric to centimetric.

FIG. 12. Electron micrograph of a hydroplastic slip plane in claystone with an elongated prod mark (centre of the figure). Related to the geometry of the pushed-microridge, the absent block was moving from bottom to top of the micrograph. Cyclothem III, Yanguas. Scale is 60 ~tm long.

Soft-sediment microfaulting, northern Spain Comparison and analogy with fault planes in shearing experiments on present cohesive soils Several studies in the field of soil mechanics have been concerned with shear testing of cohesive soils (Lambe 1959; Lambe & Whitman 1979; Institution of Civil Engineers 1975; Bolton 1979). A singular class of polished, slickensided slip surfaces has been observed in direct shear tests on cohesive soils after considerable strain (Lambe & Whitman 1979, p. 313), on present landslide surfaces affecting Pliocene cohesive soils (Skempton & Brown 1961; Skempton 1964) and on collapse structures in muddy palaeo-sediments (Wood 1935; Phillips 1938). These glossy failure planes exhibit the same geometry and specific restrictive morphoscopic characters as the hydroplastic microfaults in the Soria Basin: bumpy, shiny slip planes associated with curved divergent slickensides, a face-to-face reorientation of particles in the 10/~m thick failure zone. Shearing fractures in present argillaceous cohesive soils are commonly observed in the deposits of the superficial weathered zone (010 m) defined by a water content of about 30~o (Skempton 1964). By analogy with both deformations, the previous results can be used to characterize and restrict hydroplastic microfaulting in the Soria Basin as progressive failure dismembering of unlithified muddy sediments associated with a high water content. If we refer to the classification of subaqueous sedimentary structures related to the rheological behaviour of sediments (Elliot 1965), the Soria's

131

microfaulting can be classified in quasi-solid to hydroplastic behaviour. For this reason, we adopt the term of hydroplastic microfaults and hydroplastic slickensides to distinguish these structures quickly.

Microfaulting stress tensor Method

Recently, many authors (e.g. Carey 1979; Armijo & Cisternas 1978; Angelier & Goguel 1979; Angelier & Manousis 1980) have elaborated quantitative computer aided methods to interpret fault surface striations, considering small subareas of fracturing. Etchecopar et al. (1981) and Etchecopar (1984) proposed a computing technique separating dissimilar sets of slickensides related to different superimposed tectonic phases and characterizing each associated tectonic stress tensor. This method is presented as an inverse technique where sorting of data and computation of stress tensor are alternatively operated, this particular approach has been used to study microfaulting of Soria Basin claystones (and mudstones). Six microtectonic stations selected for presentation are located along a section of the synsedimentary syncline (Fig. 13). The measurement of slickensides for stations 1, 2, 4 and 6 was carried out on microfaults cross cutting alluvial plain siltstones-claystones of cyclothems III and IV. Stations 3 and 5 result from measurements of

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For each microtectonic station, we have measured a significant population (between 25 and 40 measurements) of striated planes and their associated slickensides. The high density of microfaults, their varied orientation and the restricted spatial surface of the microtectonic subareas (9-

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Results All the different results provided by statistical treatment of stations 1 to 6 are represented in Fig. 14. The stereographic projection on the lower hemisphere of a Schmidt stereonet of the microfault planes and their corresponding slick-

ensides (diagrams a, Fig. 14) mostly shows for each station a large spatial distribution of fractures. The associated inclined stratification plane So is also plotted. Note that the strike distribution of the microfaults would be better expressed by a rotation (if) corresponding to the dip. For each striated surface, a satisfactory stress tensor is calculated as near as possible to the final average stress tensor issued from the statistic computation of particular stress tensor populations.

I34

M. Guiraud & M. S~guret

The obtained principal stress tensor directions al (for maximum compressive stress), a2 and 0-3 are then plotted on Schmidt stereonets (diagrams b, Fig. 14). The hachured confidence patterns, at a confidence level of 95% for each of the major axes, are represented. All data are consistent and involve the same information: for each diagram b, the principal axis al and its restricted confidence domain is normal to the bedding and tilted; the main axes a2 and 0"3 and their associated confidence areas are generally elongated and distributed along the tilted bedding; using a rotation (~,) pulling the stratification plane (So) to its original position, the principal stress al is everywhere vertical and a2, 0"3 horizontal. All these results demonstrate that the stress tensor related to microfaulting in the Soria Basin represents a pre-tilting tensor with the maximum compressive stress al perpendicular to bedding (i.e. vertical). With histograms or residuals between observed and computed theoretical slickensides (diagrams c, Fig. 14) we can control the validity of stress tensor calculations. For each histogram, we can note that the angular deviations between theoretical and observed slickensides are very small. Moreover, all the histograms show an acute characteristic peak which represents a significant set of striations defined by little angular deviations with theoretical slickensides. The high chosen percentages (between 85 and 95%) of measurements accounted for stress tensor calculations, are optimum and involve stable and analogous solutions for different random computations. The calculated value of ratio R (R = (0-2 -0-3/ 0-1 -0-3) ) is given for each microtectonic station. It is possible to separate two classes of results. For stations 1, 4 (and 5), the ratio R is very close to 0 and 0.2 equals 0.3. Consequently, the confidence domains of 0-2 and 0-3 are located all along the titled bedding trace (see diagrams b, Fig. 14) and the stress tensor can be represented by an ellipsoid of revolution around 0-1. Considering microtectonic subareas 2, 3 and 6, the values of ratio R are between 0.21 and 0.35. The intermediate 0-2 and minimum stress 0-3 have slightly different values and the stress ellipsoid is similar to a revolution type. By representation on Mohr's circles (diagrams d, Fig. 14), the values of the principal stresses 0-1, 0"2, 0"3 and the precise location of the different microfaults are visualized. We can observe that generally 0"2 is very close to 0-3. The various fault planes are elongated along the edge of the Mohr's circle and may determine dihedral angles between

the directions of al and the rupture planes with values between 20 ° and 90 ° . If we plot all the microfaults on the same Mohr's circle (Fig. 15A) it results in an interesting distribution: most of the faults appear with the maximum resolved shear stress; they correspond to experimental clay behaviour. Some faults behave like experimental rocks or soils. But many faults are close to perpendicular to the maximum compressive stress and do not agree with Coulomb's law. Two types of explanations may be invoked: (i) the sedimentary fabric of the clay, resulting from the flat lying of clay particles in the bedding plane, induced a strong anisotropy and the faults tend to work parallel to this plane of anisotropy; (ii) the deformation resulting from the displacement along the 45 ° dipping microfaults is transformed into flattening in zones of low angle shear planes (Fig. 15B); this explanation is supported by the very low offset of the fault planes where they have a very low dip. In that area the slickensides decrease in intensity and even disappear.

Conclusion: discussion on relationships between microfaulting and compaction and evolution of the basin fiH From a comparison made between soil mechanics and microfaults of the Soria Basin, the latter are

0"3

0"1

(A)

.... daJ'-. /

(B)

/

Nf"

FIG. 15. (A) Position of the microfaults of the six stations on a Mohr's circle. (B) Transformation of fault displacement into flattening in a zone where there is a very low dip of the fault planes.

I

Soft-sediment microfaulting, northern Spain supposed to be related to the failure of unlithified moist deposits. Using statistical microtectonic analysis, the hydroplastic stress tensor is close to an ellipsoid type, with the maximum compressive stress perpendicular to the bedding. Consequently, this deformation pre-dates the tilting of the stratification (Fig. 16). According to the syndepositional character of the major asymmetric syncline (Figs 2 and 3), the tilting of the Wealdian strata occurred progressively during Late Jurassic-Early Cretaceous time and the structures described (water escapes and microfaults) develop during the early evolution of the basin. We propose that water escape and microfaulting represent a peculiar stage of compaction of the Wealdian deposits. The expulsion of pore fluid content in sandstones probably occurred by liquefaction of the sand beds rather than by fluidization, because of the poor mobilization of sand (lack of sand dykes, poor relief of the sand volcanoes). This water expulsion is a pre-lithification deformation. Geometric relationships between water escape structures in sandstones and microfaulting in claystones demonstrate a genetic relationship. Microfaulting in claystones evidences a normal to bedding shortening. Extension into the bedding plane (parallel to 0"3 and tr2) is unlikely because such an extension would imply a basin scale increase of strata surface in both E W and N-S directions which does not agree with limit conditions. Hence, we propose that microfaulting resulted from vertical shortening without horizontal extension, but with volumetric loss. This interpretation can also explain both the low offset and the close spacing of the faults: after a slight displacement, a conjugate set of faults could not work further because the limit of volumetric loss (for the applied strain) was reached in the zone of flattening and another set began to work. This deformation, induced by loading, is the first deformation observed in the basin fill. A later deformation, characterized by normal to bedding N 130 striking calcite tension gashes or quartz dykes, has been related to strain state providing basin development (Guiraud & S6guret 1985). This later deformation is also almost a prefolding deformation, but it occurred when silts, sandstones and carbonates were mostly lithified. Moreover, a metamorphism (T>4500C, 1 <

SW YANGUAS

~

I

135

ENCISO

JUBERA NE gm

]~

[~

TIT Upper Juressic to lower Cretaceous Cyclothems

m

Juros,sic Triossic

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~

Poleozol'c Basement

FIG. 16. Orientation of the hydroplastic stress tensors along the section of the syn-sedimentary syncline (for location see Fig. 13). P < 3 kB, gradient of 100-150°C km -1) developed in the basin fill during basin development (Guiraud & S6guret 1985), so the evolution of the basin fill from soft sediments to metamorphosed rocks can be described in three stages: 1 Water escaping from unlithified sands and microfaulting in silts with vertical shortening and no horizontal extension, but volumetric loss induced by a bi-axial type stress/strain ellipsoid related to loading. 2 Tension gashes and quartz dyke development in mostly lithified sandstones and limestones with vertical shortening and N 40°E striking extension related to strain conditions induced by the strike-slip faults overstep (Fig. 4; Guiraud & S6guret 1985). 3 Metamorphism development in static or synkinematic conditions. This evolution occurred during about 30 Ma, between the deposition time of cyclothem III (not accurately known) and the 100 Myr age of temperature climax determined by ¢°Ar/39Ar dating of syn-kinematic muscovites (H. Maluski, unpublished). ACKNOWLEDGEMENTS: This work was supported by Elf-Aquitaine. The basic contribution by A. Etchecopar in assisting with microtectonic analysis and critical advice in the writing of the paper is greatly acknowledged. J. M. Golberg provided a decisive contribution in metamorphism study and H. Maluski in radiometric dating. We are grateful to our colleagues F. Proust, J. P. Petit, P. Labaume and M. Seranne for assistance in many matters, M. F. Roch and J. Faure for typing and J. Garcia for the improvement of the figures.

References

-

ANGELIER, J. & GOGUEL, J. 1979. Sur une m6thode simple de d6termination des axes principaux des contraintes pour une population de failles. C.r. Acad. Sci. Paris, 288, 307-10. & MANOUSlS,S. 1980. Classification automatique -

et distinction des phases superpos~es en tectonique de faille. C.r. Acad. Sci. Paris, 290, 651~4. ARMIJO,R. & CISTERNAS,A. 1978. Un probl6me inverse en microtectonique cassante. C.r. Acad. Sci. Paris, 287, 595-8.

136

M. Guiraud & M. SOguret

BEUTLER, A. 1966. Geologische untersuchugen in Wealdian and Utrillas. Schichten im Westteil der Sierra de los Cameros (Nordwestlich Iberisch Ketten). Beih Geol. Ib., 44, 103-21. BOLTON,M. 1979. A Guide to Soil Mechanics. Macmillan Press. BURNE, R. U. 1970. The origin and significance of sand volcanoes in the Bude formation (Cornwall). Sedimentology, 15, 211-28. CAREY, E. 1979. Recherches de directions principales de contraintes associ~es au jeu d'une population de failles. Rev. Geol. Dyn. Geog. 21, 57-66. ELIOTT, R. E. 1965. A classification of subaqueous sedimentary structures based on rheological and kinematical parameters. Sedimentology, 5, 193209. ETCHECOPAR, A. 1984. Etude des btats de contraintes en tectonique cassante et simulations de d~formations plastiques (approche math~matique ). Th~se, Facult6 des Sciences, Montpellier. - - , VASSEUR,G. & DAIGNII~RES,M. 1981. An inverse problem in microtectonics for the determination of stress tensors from fault striation analysis. J. struct. Geol. 3, no. 1, 51-65. Geotechnique, Institution of Civil Engineers 1975. Milestones in Soil Mechanics. Thomas Telford Ltd for Institution of Civil Engineers. GUIRAUD, M. 1983. Evolution tectono-sedimentaire du bassin wealdien (crbtac~ infkrieur) en relais de dbcrochements de Logroho-Soria ( N W Espagne). Th6se 36me cycle, Facult6 Sciences, Montpellier. & Si~GURET,M. 1985. A releasing solitary overstep model for the Late Jurassic-Early Cretaceous (Wealdian) Soria strike-slip Basin (Northern Spain). In: CHRISTIE-BLICK,N. & BIDDLE, K. T. (eds.) Strike-slip Deformation, Basin Formation, and Sedimentation. Spec. Publ. SEPM, 37, 159-75. HELING, D. 1970. Micro-fabrics of shales and their rearrangement by compaction. Sedimentology, 15, 247-60. HUBERT, J. F., BUTERA, J. G. & RICE, R. F. 1972. Sedimentology of Upper Cretaceous, Cocly Parkman Delta. SW Powder River Basin Wyoming. Bull. geol. Soc. Amer. 83 (6), 1649-70. LAMBE,T. W. 1951. Soil Testing for Engineers. Wiley. -• WHITMAN, R. V. 1979. Soil Mechanics, S.I. version (2nd ed.). Wiley. LAVlLLE,E. & PETIT, J. P. 1984. Role ofsynsedimentary strike-slip faults in the formation of Moroccan Triassic basins. Geology, 12, 424-7. -

-

LOWE, D. R. 1975. Water escape structures in coarse grained sediments. Sedimentology, 22, 157-204. MONTENAT, C. 1980. Relations entre d6formations syns6dimentaires et paleosismicit6 dans le Messinien de San Miguel de Salinas (Cordill+res B6tiques, Espagne). Bull. Soc. gbol. France 7, t. XXII, no. 3, 501-9. O'BRIEN, N. R. 1970. Comparison of the fabric of a sensitive Pleistocene clay with laboratory flocculated clay using the scanning electron microscope. Mar. Sedim. 5, 58-61. PETIT, J. P. 1976. La zone de dbcrochements du Tizi N'Test (Maroc) et son fonctionnement depuis le Carboni~re. Th+se Facult~ des Sciences Montpellier. --, PROUST, F. & TAPPONNIER, P. 1983. Crit+res de sens de mouvement sur les miroirs de failles en roches non calcaires. Bull. Soc. geol. France, 7, t. XXV, no. 4, 589-608. PHILLIPS, D. W. 1938. Microscopical evidence of shearing in argillaceous rocks. Proc. Yorkshire geol. Soc. 24, 67-9. RIEO~L, W. 1929. Zur mechanik geologischer Brucherscheinungen ein Beitrag zum problem den Fiederspatten. Zbl. f Min. Geol. Pal. Abt. B, 354-68. SALOMON, J. 1983. Les formations continentales du Jurassique supbrieur-Crktack inf~rieur. Th6se, m~moire g6ologique de l'Universit6 de Dijon no. 6. SANTOUIL, G. 1980. Tectonique et microtectonique comparbe de la distension permienne et de t~volution posttriasique dans les bassins de Lodbve, St Afrique et Rodez (France SE). Th~se 3~me cycle, Facult6 Sciences Montpellier. SELLEY, R. C. 1969. Ancient Sedimentary Environments. Chapman & Hall. SIMS, J. D. 1973. Earthquake induced structures in sediments of Van Norman lake, San Fernando, California. Science, 182, 161-3. --, 1975. Determining earthquake recurrence intervals from deformation structures in young lacustrine sediments. Tectonophysics, 29, 141-2. SKEMPTON, A. W. 1964. Long-term stability of clay slopes. Geotechnique, 14, 77-101. & BROWN, J. D. 1961. A landslide in Boulder Clay at Selset, Yorkshire. Geotechnique, 11, 4, 280-93. TJIA, H. D. 1968. Fault-plane markings. XXlllinternational geological congress, 13, 279-84. WOOD, A. L. 1935. The origin of the structures known as guilielmites. Geol. Mag. 72, 241-5. -

-

M. GUIRAUD, G6ologie Structurale USTL, P1. Bataillon 34060 Montpellier, France. Present address: Department of Geology and Mining, University of Jos, P.M.B. 2084, Jos, Nigeria. M. SI~GURET,G6ologie Structurale USTL, P1. Bataillon 34060 Montpellier, France.

Sediment deformation structures and the palaeotectonic analysis of sedimentary basins, with a case-study from the Carboniferous of northern England Michael Leeder S U M M A R Y: There is a close link between syn-sedimentaryfaulting caused by earthquakes of magnitude > 5 and accompanying 'soft' sediment intrafolial deformation in active neotectonic settings. Certain of these intrafolial structures may also be produced by purely sedimentary processes and are termed 'autokinetic'. Those caused by earthquake-induced stresses are termed 'allokinetic'. Although certain structures may be caused by either mechanism, it appears that large amplitude dewatering pipes and recumbent-folded cross stratification are likely to be predominantly allokinetic in origin. These structures may be used to define the Allokinetic Deformation Number which can be systematically mapped out in particular lithofacies over outcrops in a sedimentary basin. Results of this mapping are then combined with data on lithofacies thickening and stacking (tectonically-controlled architecture), hangingwall rollover deformation etc., to establish an interdisciplinaryoutline of palaeotectonic evolution. This methodology of palaeotectonics is tested in the suspected extensional terrain of the Northumberland Basin with encouraging results.

It is clear from numerous recent studies that active faulting in neotectonic areas is accompanied by various kinds of near-surface 'soft' sediment deformation. This link between brittle failure along fault planes and soft sediment deformation also defines the science of earthquake engineering soil mechanics, pioneered by H. B. Seed in a series of influential papers. It is the purpose of the present paper to use such links to establish a methodology for investigating palaeotectonics in ancient sedimentary basins. It is probably no overstatement to say that, for most geologists, soft sediment deformation structures are interesting curios, since they have no particular use in establishing contemporary flow vectors or conventional estimates of palaeoenvironments because of their clear post-depositional origins. There is, though, a wider recognition of the value of 'slump' structures in palaeoslope analysis (e.g. Woodcock 1976; Gawthorpe & Clemmey 1985). A few authors, notably Allen & Banks (1972), have clearly recognized that liquefaction induced by earthquake ground motions is a most effective way of deforming sediments close to the contemporary depositional surface. Attempts to link soft sediment deformation structures to a particular fault were made by Weaver (1976), who mapped out 'load' structures around the Swansea Valley Fault in the Westphalian rocks of South Wales, and MayaU (1983), who undertook a similar exercise in the Rhaetic strata of the Somerset area. Montenat (1980) utilized dewatering pipes to recognize an active fault

zone in the Messinian of the Betic Cordillera, Spain. It is in the field of neotectonics that soft sediment structures have recently become a critical link in the chain of reasoning associated with earthquake prediction studies. Following the study of Sims (1975), who clearly demonstrated the use of foundered pseudo-nodules in Californian lake sediments in establishing earthquake recurrence intervals, Sieh (1978) and Russ (1982) have both used deformed sediment layers to establish a methodology for earthquake prediction. These studies took place close to the San Andreas and New Madrid active fault zones respectively (Wessen & Wallace 1985). It is clear from these recent studies that there is a juxtaposition of structures due to faulting and soft sediment deformation. The major problem in palaeotectonic analysis is that there may exist: (i) a variety of faults of various ages and (ii) soft sediment deformation structures not associated with earthquake-induced ground motions. These problems are most acute in ancient extensional basins where normal faults of the active extensional phase become seismically inactive during the thermal subsidence phase. The thinned crust is then particularly prone to later oblique or compressive deformation which may reactivate certain syn-depositional faults and produce new post-depositional structures. The site of the formerly active extensional basin will thus be deformed by structures showing a variety of trends. As noted above, it is the purpose of the

From : JONES, M. E. & PRESTON, R. M. F. (eds), 1987, Deformationof Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 137-146.

I37

138

M. Leeder

present paper to provide a methodology for the recognition of syn-depositional faults in such basins. Although restricted to extensional basins, some of the conclusions are applicable to basins formed in compressional regimes. Attributes of both brittle and soft structures will be used and the methodology will be tested in the formerly extensional (Lower Carboniferous) Northumberland Basin of northern England.

Intrafolial sediment deformation and neotectonics Intrafolial sediment deformations may be defined as those deformations affecting sediment layers which have occurred at or near the contemporary deposurface and which (eventually) are bounded by undisturbed overlying beds deposited soon after the deformation event. Seilacher (1984) has recently called those intrafolial structures (and redeposited beds), due directly to earthquake motions, 'seismites', and has recognized the considerable difficulties involved in unambiguously identifying ancient seismites. There are several problems here, one being that there are many natural gravity slopes which originate as depositional surfaces. Their failure under normal (non-seismic) stresses will produce structures and deposits identical to those induced by seismicity. Another problem occurs in areas of rapid drawdown, such as river sandbars, dunes and banks, where enhanced pore pressures may build up due to poor draining, thus encouraging both mass failure and local dewatering. Because of these problems, it may be useful to define intrafolial structures as either 'allokinetic' or 'autokinetic'. Allokinetic deformation is due to kinetic energy contributed from outside the deposystem by periodic fault plane motions. Autokinetic deformation arises from kinetic energy supplied from within the depositional system from intrinsically depositional mechanisms such as (i) sudden mass deposition causing oversteepening and failure; (ii) drawdown effects and poor drainage; (iii) motion down constructional deposurfaces; (iv) bottom stress variations due to the passage of overlying water waves. The origin of allokinetic structures is closely linked to the processes whereby sediment strength of saturated granular deposits is lost during earthquake-induced ground motions. These processes have been extensively investigated by H. B. Seed and co-workers (1966, 1967, 1968, 1971, 1976, 1979) and may be briefly summarized as follows. Loss of bulk sediment strength occurs because the effective stress, a', approaches zero

as the pore pressure, p, approaches the solid stress, a, since a ' = a - p . This condition of liquefaction thus arises when stress is transferred from solid contacts between grains to the pore fluid. Experimental data show that in situ liquefaction is caused by cyclic applied stresses and that in nature such stresses are induced by ground motions due to upward migration of shear waves. The effect of cyclic stress application is a compaction of the cohesionless deposit, in turn causing a stress transfer to the pore fluid. Loss of strength follows and the liquid-like slurry is then capable of flow or of being deformed internally by local gravitational or fluid stresses. Liquefaction is encouraged by: (i) loose deposit packing and hence high initial porosity/void ratio; (ii) low confining pressures; (iii) high values of applied cyclic shear stress; (iv) large numbers of applied stress cycles; (v) low initial shear stresses acting on the deposit; (vi) sand grain sizes (fine to medium) that prevent very rapid drainage and pore pressure dissipation. In a homogenous sand deposit, liquefaction will occur first at the surface and the liquefaction front will then migrate downwards. However, if the deposit is more loosely packed at depth then, although the confining pressure is higher, it is possible that the deep layer will liquefy first. Following the cessation of earthquake motions, model studies (Seed 1979) show that in fine to medium sands, pore water pressures at depth dissipate slowly. Nearer the surface, pore pressures continue to build up, with conditions of liquefaction existing for long periods (several to tens of minutes), with the occurrence of water boils and more diffuse upward flow to the surface. In coarse sands and gravels there is no development of high pore pressures because of efficient drainage caused by high permeability. Significant liquefaction potential occurs in sedimentary sequences containing looselypacked sand lenses within impermeable clays. The lateral margins of such layers are particularly prone to deformation. Interestingly enough, there is a reduced likelihood of liquefaction on slopes (Seed 1968). All other factors being equal, liquefaction depends on the magnitude and duration of the applied cyclic stresses. This is well illustrated by long-term records of failure around Niigata, Japan, where liquefaction has only occurred where the maximum ground acceleration was greater than 0.12 g (data of Kawasumi, in Seed & Idriss, 1971, fig. 1). Since ground acceleration decays away from an active hypocentre (see review by Seed et al. 1976) it is clearly possible to make empirical predictions concerning the areal distribution of liquefaction occurrences for earth-

Palaeotectonic analysis of sedimentary basins quakes of various magnitude (Fig. 1 ; Seed 1979; Youd & Perkins 1978). In fact, knowing the near surface distribution of sediment types, it is possible to produce maps of the so-called 'liquefaction-induced ground failure potential' (Youd & Perkins 1978). The process of liquefaction described above leads to the possibility of using certain of the resulting sedimentary structures as indices of aUokinetic deformation. Such structures must be chosen so as to avoid the possibility of confusion with autokinetic structures. Structures produced at the surface by allokinetic deformation include fault breaks and fissures, sand boils (volcanoes) fed by (sometimes) deep feeder dykes tapping sand lenses, sand flows and slope slumps and slides. Following liquefaction, the consolidation of the deposit into its new packing mode will cause settlement of magnitude s = - h ( C 2 - C 1 ) , where h is the initial thickness of liquefied sand; C1 is the initial (looser) packing and C2 is the final packing ratio. Many of the surface effects noted above have a high preservation potential, but, as noted previously, some may also be produced by autokinetic deformation. Fine illustrations of these various allokinetic structures are to be found in the extensive monographs dealing

'l

8

5

LIQUEFACTION ~e ,,..,

STABILITY

g ~s lkm

10

100

Distance to furthest significant liquefaction

FIG. 1. Distance from seismic source to the furthest significant liquefaction observed in surficial floodplain and deltaic sediments. The limiting envelope has a threshold magnitude of 5.0 (no liquefaction effects recorded in the literature below this threshold) and a cut-off at 150 km (after Youd & Perkins 1978). Although based on limited data, these authors consider the plot to be reliable. Data points: 1--1954, Fallon-Stillwater; 2--1964, Niigata; 3-1976, Guatamala; 4--1906, San Francisco; 5--1964, Alaska. N.B., Point 6 is a later point from the 1979 Imperial Valley 'quake (after Youd & Wieczorek 1982). Note the consistency of this point with the earlier discriminant line.

139

with the Alaskan 1964 earthquake (US Geological Survey Papers 543, 544) and the Imperial Valley 1979 earthquake (US Geological Survey Paper 1254).

Penetrative sediment deformation and neotectonics The process of active lithospheric extension is witnessed at the surface by the production of characteristic graben and half-graben/tilt blocks. These are bounded by major normal faults, often of listric geometry (see McKenzie 1978; Jackson & McKenzie 1983). Such features define active sediment dispersal and depositional zones of great complexity, since sedimentation is occurring at the same time as active extension (Leeder & Gawthorpe, in press). The result is a sedimentary basin bounded by major extensional fault zones, but also with a variety of intra-basinal synthetic and antithetic normal faults, some of which will be rotated during continued extension to form reverse faults (Morton & Black 1975; Jackson et al. 1982). These intra-basinal faults occur on a variety of scales. They are well developed in the Rio Grande graben (Kelley 1977), in the Corinth graben (Brooks & Ferentinos 1984) and in the recently inactive Locride and Megara graben fills of central Greece (Dufaure et al. 1979; Jackson et al. 1982; Leeder & Gawthorpe, unpublished). The majority of active faults cutting the synextensional sedimentary fill, show the characteristic listric geometry, often toeing-out within the sedimentary prism, or occurring as offshoots from major structures defining extensional duplexes (see Gibbs 1984). Intrabasinal faults may not be capable of causing ground accelerations of sufficient magnitude to cause widespread liquefaction, in contrast to the major graben-bounding structures. They will, however, influence sedimentary processes if they break through to the surface or cause tilting. The majority of such faults occur in the major graben hangingwalls. Fine examples occur in the Rio Grande rift of New Mexico bounding the Ladron horst and Velarde/San Felipe mini-graben systems, and also in the lower Rhine graben of western Germany. The identification of once-active faults formed during extension is easier in the larger structures. These may (i) preserve their listric geometry; (ii) show associated rollover anticlines and other reverse drag features; (iii) show marked positive influence upon sedimentary thickness as the hangingwall slope is descended; (iv) show marked and predictable facies changes associated

I40

M. Leeder

with the production of secondary tectonic gravity slopes; (v) show periodic production of adjacent earthquake-induced soft sediment deformation structures; (vi) encourage marked orientation down the hangingwall slopes of gravity structures such as slump folds, debris flows and sandflows. Furthermore, since lithospheric extension is a regional reaction to deviatoric extensional stresses at plate edges, the regional stress regime will cause a broad parallelism of both minor and major penetrative structures in any one basin. Some of the above points are illustrated in Fig. 2.

Sediment deformation and palaeotectonics--a methodology Once active extensional tectonics cease, faultcontrolled deformation gives way to regional thermal subsidence of 'sag' type. Subsequent geological evolution may involve shortening deformation, then further periods of extension, and so on, in response to events at near or distant plate boundaries. Analysis of sediment deformation in such ancient basins may thus become rather involved. Following the previous discussion of intrafolial and penetrative deformation, a successful methodology for palaeotectonic analysis must involve the following steps: 1 Identification of major thickness and facies changes using geological and geophysical data as available.

2 Definition of basin margins and adjacent footwall uplands from the data in 1. 3 Identification of basin margin faults and hingelines from 1 and 2 and the location of all marginparallel, intra-basin faults. 4 All faults outlined in 3 are analysed for their syn-sedimentary nature using the six criteria proposed in the previous section. Any data on the timing of fault activity are considered critical, so that a true chronology of extension can take shape. This aspect is often neglected by structural geologists. 5 Data on fault resurrection in subsequent extensional phases are used to assess the 'inheritance factor', i.e. the degree to which earlier crustal structures influence later ones. It must be recognised that soft sediment deformation will be critical in step 4. Some further discussion is thus needed before a case history is presented. From the discussion above, we must eliminate all gravity-controlled structures from the analysis, since these may be autokinetic, albeit down a secondary gravity slope produced by tectonic tilting. Such structures include all slides, slump folds, asymmetric load casts and contorted cross-strata indicating downslope motion relative to the depositional dip. In exceptional circumstances, some of these structures may be of great value, for example, if the structures occur at one stratigraphic level over a large area. None of the above disqualifying remarks lessens the vital role that gravity slope

FIG. 2. Diagram to illustrate the kinds of stratigraphic and sedimentological data needed to test whether a normal fault was active during deposition. Fault 1 was inactive until the interval a-b when facies and thickness changes and allokinetic soft sediment deformation structures indicate syndepositional ground motions. The fault was then inactive during deposition of units b-c, whose further fault motion occurred before deposition of the upper limestone. The central horst block shows facies changes (thick coals) in the interval equivalent to a-b deposition. Fault 2 was initiated after deposition of the clastic unit but during deposition of the limestone after position d. Relationships based on field data observable in NE England in the Carboniferous-Triassic successions. Note: the presence of a rollover is not evidence of syndepositional fault motion without other criteria being present.

Palaeotectonic analysis of sedimentary basins structures play in assessing the directional gradients of ancient basin floors. The remaining structures, listed below, indicate motions normal to the contemporary depositional surface and are considered to be valuable indicators of allokinetic deformation, once a number of provisos have been discussed.

• .

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Catastrophic dewatering pipes These are very common around modern active fault zones where the water table is close to the surface. They are illustrated in Figs 3 and 4. The structures are not uniquely allokinetic, however, since small water expulsion pipes (generally < 20 cm amplitude) and silt volcanoes have recently been observed to form as water level is lowered around modern river floodplain or bartop bedforms which are sometimes capped by impermeable muds or low permeability silts (C. Bristow, data from river Brahmaputra, pers. comm.). Major slump folds and slides may also be capped by water expulsion pipes and sand volcanoes (Gill & Kuenen 1958). This allowance of autokinetic examples may be lessened by taking a lower limit of 0.25 m for the vertical amplitude of pipe cores or flanking upturned laminae (see Fig. 4).

Recumbent-folded cross stratification The formation of these structures has never been

/

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FIG. 4. Field sketch of typical dewatering pipe. Note: (1) upturned marginal laminae; (2) structureless pipe fill (these features indicate very rapid fluid expulsion causing viscous shear of adjacent liquefied sands-unpublished work); (3) overlying erosion surface cut in the lee of an advancing dune bedform; (4) evidence for a period of subsequent dewatering; (5) liquefaction of a migrating small dune during stream flow, causing shear in the liquefied sands into the typical recumbentfolded ('omelette') structure. Locality: Burnt Tom's Crags, Bewcastle Fells (M. Border Group, Lower Carboniferous) in the Northumberland basin. directly observed, either in nature or in experiment. Allen & Banks model (1972) is much favoured by an explicitly allokinetic driving mechanism. Previous objections to the mechanism (e.g. Hendry & Stauffer 1975) seem less than convincing to the author, since no palaeotectonic analysis was carried out.

Sandstone ball-and-pillow

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.--""

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3. Diagram to show a possible sequence of effects following liquefaction of a loosely-packed and crossstratified fine to medium grained sand and the resulting expulsion of sand-laden pore water through discrete pipes and ridges. Zone A is unliquefied sand the top surface of which ascends with time. Zone B is a water-rich slurry the top surface of which is deformed into diapiric intrusions into the liquefied sands of ?higher density in Zone C. This diagram was inspired by the quantitative model of Housner (1958) and by unpublished field observations concerning the regularity of 'pipe' wavelength in vertical 2-D exposures and by their crudely polygonal form in plan view. FIG.

These foundered structures have been most used in previous palaeotectonic studies, yet they are probably the least reliable allokinetic indicators, since it has been clearly shown that they may form spontaneously as sediment is deposited upon overpressured or 'quick' clays (Ankatell et al. 1969). A further complicating factor is the likelihood that there will be a strong facies control upon the occurrence of allokinetic structures, thus making it imperative to compare similar facies when assessing regional variations in the 'degree' of soft sediment deformation.

A deformation index The final methodological proposal concerns the definition of a proposed new index of abundance of soft sediment deformation which may be systematically mapped out in an ancient sedimentary basin from basic logging data on sedimentary

I42

M. Leeder

structures and depositional environments. The index is termed the 'allokinetic deformation number' (ADN) and is defined as a thicknessweighted deformation length ADN =

Td •T Tu+l

where To is the total thickness of deformed lithologies, Tu is the total of undeformed lithologies and T is the total thickness. The A D N must be used in an internally consistent and objective manner, preferably applied to a given lithofacies containing large likely allokinetic structures of the kinds discussed above (dewatering pipes, overturned cross sets, ball and pillow). All occurrences of the defined lithofacies should be assigned an ADN, where no deformation exists then A D N = 0 at that locality. The A D N is most effectively applied to relatively thin lithofacies deposited over a small time interval which may be correlated within the basin. The reason for giving the A D N a thickness-weighting is to include some estimate of relative 'quake intensity, ~.~

l I}

since the degree of seismic shaking should decrease away from active faults causing thinner intervals of sediment to show deformation structures.

Sediment deformation studies in the Northumberland basin: A brief case history It has been proposed (Leeder 1982) that the Lower Carboniferous basins of the northern British Isles originated as extensional features, probably due to marginal basin kinematics associated with plate-edge forces around the developing European Hercynian orogenic belt (Leeder 1986). Within this remarkable extensional province, the Northumberland basin stands out as a fine example of an internally complex extensional structure draped by a thermal sagfill which subsequently suffered shortening deformation, with local thrusts and cleavage formation. Fig. 5 shows the distribution of

Mesozoic basin fills

20km

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FIG. 5. Faults and major folds affecting the Carboniferous basin fill of northern England. These are a mixture of syn- and post-depositional structures thought to result from Lower Carboniferous extension, late Upper Carboniferous shortening and Permo-Triassic extension.

Palaeotectonic analysis of sedimentary basins mapped structures as seen today in the basin fill. As part of a pilot study in palaeotectonic methodology, many of these structures have been assessed in terms of their likely extensional origins using the various criteria discussed previously. Field work has been undertaken, and published BGS maps and memoirs and all other relevant literature have been scoured in an effort to locate structures controlling or showing: (i) stratigraphic thickening; (ii) reverse drag/rollovers; (iii) facies changes; (iv) high ADNs. Concerning the use of the A D N , it was decided to restrict the quantitative study to channel sandstone facies of fine to medium grade with the A D N sample units restricted to large-scale ( > 10 cm) cross-stratification. Gravity effects arising from liquefaction-induced slumps were hopefully minimised by choosing dewatering pipes with vertical axes (relative to the depositional horizontal) and recumbent-folded crossI1t

143

sets as the relevant structures. The use of cross stratification in fine to medium sandstones ensures that the sampled thicknesses had a reasonably constant high original porosity in which rapid draining would have been minimized and liquefaction encouraged after cyclic shocking. The cross sets also serve to highlight any liquefaction-induced disturbances. A total of over 150 outcrops were examined in the course of the study and a further 50 or so taken from data in dependable modern literature. The sampling programme was not exhaustive and many exposures remain to be visited, particularly in the remote afforested Northumberland moorlands. Results of the study are summarized in Figs 68, where it can be seen that: 1 A number (at least nine) of possible extensional normal faults clearly existed within the basin. 2 These faults were particularly active within the

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%AESTON: BiOC" ' FIG. 6. Map to show : (i) the spatial and crudely temporal distribution of ADNs in the Carboniferous basin fill of the Northumberland basin and adjoining Alston Block; (ii) rollover structures; (iii) suspected extensional faults active during Carboniferous times; (iv) suspected extensional faults active during the post-Permian (possibly reactivated from buried examples of the previously mentioned structures). Numbered faults: 1 Bowden Doors; 2, 5 Antonstown; 3 Sweethope; 4 Sam's; 6 Beckhead/Binky Linn; 7 90-Fathom; 8 Stublick; 9 North Solway. This list likely to grow as analysis continues.

I44

M. Leeder

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FIG. 8. To show the small rectangular area outlined in the SE corner of Fig. 6, where there is a fine example of a rollover structure present in the hangingwall of the 90-Fathom Fault in the area to north of Newcastle-uponTyne. Strata contours and faults after Land (1974, figs 81 and 82). Arrowhead lines show the graphical directional derivatives drawn roughly normal to the strata contours down the local gradient. These highlight the large hangingwall rollover to the north and the backtilted footwall block to the south (my unpublished data with Harry Clemmey). The width of the fault zone shaded on the diagram gives an idea of throw, varying from a maximum of 400 m to a minimum of 150 m (data of Land 1974).

central basinal area during M - U Border and Liddesdale Group times (see the detailed map in Fig. 7) where they controlled development of channel sandbody stacking in hangingwall lows, stratigraphic thickness, rollover structures and extensive allokinetic deformation structures. 3 There is little evidence for extensive intraWestphalian fault movement from the A D N data, confirming the Coal Measures strata as a passive infill deposited during a thermal sag phase of subsidence. A notable exception is the spectacular soft sediment deformation some 10 km north of the 90-Fathom Fault near Stannington. 4 The southern basin margin is not as well constrained as some authors maintain from gravity studies (e.g. Johnson 1984) since the impressive and newly recognised rollover associated with the 90-Fathom Fault (Fig. 8) is clearly post-Carboniferous (involving Permian strata) and does not cause widespread intra-Carboniferous effects demanded by the allokinetic criteria outlined previously. These observations, together with the thick Lower Carboniferous succession cored by the Harton No. 1 wildcat (Ridd et al. 1970) on the footwall side of the structure, suggest extreme caution should be exercised when guessing the exact location of the draped and buried Lower Carboniferous footwall scarp along the SE basin margin with the Alston tilt block.

5 The status of the northern basin margin as a portion of the Southern Uplands tilted hangingwail block is confirmed by the analysis (see Leeder 1982). Most of the N W - S E structures in the Langholm area, for example, show little evidence for syndepositional motion. Future work will aim to amplify, extend and further test these preliminary conclusions.

Conclusions Penetrative and intrafolial sedimentary deformation structures in extensional basins are closely linked. Certain allokinetic intrafolial structures may be used to define an index of deformation which may be mapped out within the basin on a systematic basis. Together with data on facies architecture and thickening trends and fault hangingwall geometry, the deformation index may be used to assess the timing and location of those faults that were formerly active during basin extension. ACKNOWLEDGEMENTS: I gratefully acknowledge Texa-

co's support for this research. The manuscript has benefited from receiving the critical comments of Jan Alexander, Charlie Bristow, Richard Collier, Dave Ord and Geraint Owen.

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5km FIG. 7. Detailed survey of A D N s (relative scales as in Fig. 6) and thickening trends within the area of the rectangle indicated in Fig. 6 in the very complex intra-graben tilt blocks of the central Northumberland basin. Heavy lines: thickening trends into the Bewcastle and Bellingham intragraben lows defined by the E N E - W S W faults of the Antonstown-Sweethope-Sams-Beckhead/Binky Linn formerly extensional normal faults. Arrows : 1, 2 Middle Border Group; 3-5 Upper Border Group; 6 Liddesdale Group. Stipple: sandstone members, mostly of fluvial and fluvio-deltaic channel origins with palaeoflow mostly directed from N W to SE, i.e. axial drainage with respect to the basinal fault trends. Data from Day (1970), Frost & Halliday (1980) and extensive personal observations. Key as Fig. 6.

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Palaeotectonic analysis of sedimentary basins

I45

References ALLEN, J. R. L. & BANKS,N. L. 1972. An interpretation and analysis of recumbent folded deformed cross bedding. Sedimentology, 19, 257-83. ANKATELL, J. M., CEGLA, J. & DZULYIqSKI,S. T. 1969. Unconformable surfaces formed in the absence of current erosion. Geol. Comana, 8, 41-6. BROOKS, M. & FERENTINOS, G. 1984. Tectonics and sedimentation in the Gulf of Corinth and the Zakynthos and Kefallinia channels, Western Greece. Tectonophysics, 101, 25-54. DAY, J. B. W. 1970. Geology of the country around Bewcastle. Mem. Geol. Surv. G.B. 357 pp. DUFAURE, J. J., BOUSQUET,B. & PECHOUX, P. Y. 1979. Contributions de la g6omorphologie fi la connaissance du Quaternaire continental grec, en relation avec les &udes de n6otectonique. Rev. Geol. dyn. Geog. Phys. 21/1 29-40. FROST, D. V. & HOLLIDAY, P. W. 1980. Geology of the country around Bellingham. Mem. Geol. Surv. G.B. 112 pp. GAWTHORPE, R. & CLEMMEY, H. 1985. The geometry of sedimentary slides in the Bowland Basin (Dinantian) and their relationship to debris flows. J. geol. Soc. Lond. 142, 555-65. GIBBS, A. D. 1984. Structural evolution of extensional basin margins. J. geol. Soc. Lond, 141,609-20. GILL, W. D. & KUENEN, P. H. 1958. Sand volcanoes on slumps in the Carboniferous of County Clare, Ireland. Q. J. geol. Soc. Lond. 113, 441-60. HENDRY, n . E. & STAUFFER,M. R. 1975. Penecontemporaneous recumbent folds in trough cross-bedding of Pleistocene Sands in Saskatchewan, Canada. J. sed. Petrol. 45, 932-43. HOUSNER, G. W. 1958. The mechanism of sandblows. Bull. seismol. Soc. Amer. 48, 155-61. JACKSON, J. & MCKENZIE, D. P. 1983. The geometrical evolution of normal fault systems. J. struct. Geol. 5, 471-82. , KING, G. & VITA-FINZI, C. 1982. The neotectonics of the Aegean: an alternative view. Earth planet. Sci. Letts. 51,303-18. JOHNSON, G. A. L. 1984. Subsidence and sedimentation in the Northumberland Trough. Proc. Yorks. geol. Soc. 45, 71-84. KELLEY, V. C. 1977. Geology of the Albuquerque Basin. Mem. 33, New Mexico Bureau Mines, Mineral Res. LAND, D. H. 1974. Geology of the Tynemouth District. Mere. Geol. Surv. G.B. 176 pp. LEE, K. L. & SEED, H. B. 1967. Cyclic stress conditions causing liquefaction of sand. J. Soil Mech. Found. Div. Proc. A.S.C.E. 93, SMI, 47-70. LEEDER, M. R. 1982. Upper Palaeozoic basins of the British Isles--Caledonide inheritance versus Hercynian plate margin processes. J. geol. Soc. London, 139, 479-91. --, 1986. Tectonic and palaeogeographic models for Lower Carboniferous Europe. In: MILLER, J., ADAMS, A. E. & WRIGHT, V. P. (eds) European

Dinantian Environments. Geol. J. Spec. Issue, 12, 1 - 2 0 .

--

& GAWTHORPE, R. L. 1987. Sedimentary models for extensional tilt block/half graben basins. In: COWARD, M. P., DEWEY, J. F. & HANCOCK, R. (eds) Extensional Tectonics. Spec. Publ. geol. Soc. Lond. 139-52. MAYALL, M. J. 1983. An earthquake origin for synsedimentary deformation in a late Triassic (Rhaetian) lagoonal sequence, SW Britain. Geol. Mag. 120, 613-22. MCKENZIE, D. P. 1978. Some remarks on the development of sedimentary basins. Earth planet. Sci. Lett. 40, 25-32. MORTON, W. H. & BLACK, R. 1975. Crustal attenuation in Afar. In: PILGER, A. & ROSSLER,A. (eds) Afar Depression o f Ethiopia, 55-65, Deutche Forschung, Stuttgart. MONTENAT, C. 1980. Relation entre deformations synsedimentaires et paleoseismicite dans le messinient de san Miguel de Salinas (Cordilleres betiques orientales, Espagne). Bull. Soc. geol. France, 22, 501-9. RIDD, N. F., WALKER, D. B. & JONES, J. M. 1970. A deep borehole at Harton on the margin of the Northumbrian trough. Proc. Yorks. geol. Soc. 38, 75-103. Russ, D. P. 1982. Style and significance of surface deformation in the vicinity of New Madrid, Missouri. U.S. geol. Surv. Prof. Paper 1236H, 95114. SEED, H. B. 1968. Landslides during earthquakes due to soil liquefaction. J. Soil Mech. Found. Proc. A.S.C.E., 94, SM5, 1055-122. , 1979. Soil liquefaction and cyclic mobility evaluation for level ground during earthquakes. J. Geotech. Eng. Div. Proc. A.S.C.E. 105, GT2, 20155. & LEE, K. L. 1966. Liquefaction of saturated sands during cyclic loading. J. Soil Mech. Found. Div. Proc. A.S.C.E. 92, SM6, 105-34. & IDRISS, I. M. 1971. Simplified procedure for evaluating soil liquefaction potential. J. Soil Mech. Found. Div. Proc. A.S.C.E. 97, SM9, 1249-73. , MURARKA, R., LYSHER, J. & IDRISS, I. M. 1976. Relationships of maximum acceleration, maxim u m velocity, distance from source, and local site conditions for moderately strong earthquakes. Bull. seismol. Soc. Amer. 66, 1323-42. SEILACHER,A. 1984. Sedimentary structures tentatively attributed to seismic events. Mar. Geol. 55, 1-12. SIEH, K. E. 1978. Prehistoric large earthquakes produced by slope on the San Andreas fault at Pallett Creek, California. J. geophys. Res. 83B, 3907-39. SIMS, J. D. 1975. Determining earthquake recurrence intervals from deformational structures in young lacustrine sediments. Tectonophysics, 29, 141-52. WEAVER, J. O. 1976. Seismically-induced load structures in the basal Coal Measures, South Wales. Geol. Mag. 113, 535-43.

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-

-

-

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M. Leeder

WESSON, R. L. & WALLACE,R. E. 1985. Predicting the next great earthquake in California. Sci. Amer. 252, 23-31. WOOI~OCK, N. H. 1976. Ludlow Series slumps and turbidites and the form of the Montgomery Trough, Powys, Wales. Proc. Geol. Assoc. 87, 169-82.

YOUD, T. L. & WIECZOREK,G. F. 1982. Liquefaction and secondary ground failure. U.S. geol. Surv. Prof. Pap. 1254, 223-50. & PERKINS,n. M. 1978. Mapping liquefactioninduced ground failure potential. J. Geotech. Eng. Div. Proc. A.S.C.E. 104, GT4, 433-46.

-

-

M. R. LEEDER,Department of Earth Sciences, University of Leeds, Leeds LS2 9JT.

Syn-diagenetic deformation of a turbiditic succession related to submarine gravity nappe emplacement, Autapie Nappe, French Alps Pierre Labaume S U M M A R Y : The Autapie Nappe (French Alps) consists of a Late Cretaceous 'Helminthoid Flysch' turbiditic succession. It was emplaced in the external Alps during the Late Eocene under the form of a gravity submarine nappe sliding in a turbiditic foreland basin. The syn-diagenetic gravity spreading of the Helminthoid Flysch succession resulted in a pervasive extensional disruption of the sandstone and mudstone layers related to decollements in the marlstone and claystone layers. The deformation regimes and mechanisms

testify to: (i) the incomplete and inhomogeneous lithification of the sediments, and (ii) the occurrence of pore pressure gradients and, possibly, excess pore pressure during the deformation. In particular, the critical state of lithification of the sandstone allowed the occurrence of an unusual chronology of deformation mechanisms, the 'brecciation' resulting from a partial loss of cohesion of the sandstone in high speed escaping pore-water (ductile, 'soft-sediment' behaviour) post-dating calcite-filled veins (brittle, 'rock' behaviour). Similar deformation style might be expected to occur at modern convergent margins.

Soft-sediment deformation was first described for syn-sedimentary structures, but it is now known that it can occur under wider conditions (Fitches & Maltman 1978; Maltman 1984, this volume). Moreover, since soft-sediment deformation usually affects water-saturated sediments, intergranular water behaviour (migration and possible overpressuring) commonly influences the deformation style. These factors must be considered when the evolution of convergent margin accretionary prisms and submarine nappes is analysed since: (i) these structures usually involve recently deposited sediments; (ii) fluid behaviour is an important parameter of their evolution (Von Huene & Lee 1982; Von Huene 1984). Investigation of modern convergent margins provides direct evidence of active processes such as fluid circulation and overpressuring (Moore et al. 1982; Suess & Massoth 1984; Legett & Platt 1985; Boul~gue et al. 1985), but the study of the geometry and mechanisms of the deformation is difficult there because of the limits of seismic profile interpretation and core and dive observations. On the other hand, the active processes cannot be observed in similar inland ancient series, but their resulting structures can be analysed there. Therefore, the study of these ancient structures may furnish data useful in establishing deformation models for active margins. This paper is devoted to the Helminthoid Flysch succession of the Autapie Nappe, an ancient submarine nappe of the French Alps. We

describe an early stage of extensional deformation and discuss how the deformation features testify to the state of lithification of the sediments as well as the strain and pore-water pressure conditions prevailing during the deformation. The 'state of lithification' will be regarded as the state of achievement of all the processes (compaction and related water escape, pressure solution, recrystallization, cementation etc.) which progressively changed the sediments into rocks, reducing porosity and enhancing cohesion and viscosity.

Geological setting The Autapie Nappe belongs to a group of allochtonous units which now rest upon the Mesozoic and Eocene sedimentary cover of the external Alps (Dauphinois Zone) between the Argentera and Pelvoux external crystalline massifs (Fig. 1). These thrust-sheets form three principal 'nappes' (Kerckhove 1969): (i) the subBrian~onnais Units; (ii) and (iii) two Helminthoid Flysch nappes, the Autapie and Parpaillon Nappes. The Helminthoid Flysch is a Late Cretaceous (locally up to Palaeocene) turbiditic succession and was the last stratum deposited in the Tethyan Ocean (Kerckhove 1969; Debelmas 1975; Kerckhove et al. 1981 ; Caron et al. 1981 ; Homewood & Caron 1982). Its deposition occurred east of the contemporaneous obduction zone of the Tethyan

From: JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 147-163.

147

I48

P. Labaume

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FIG. 1. Geological setting of the Autapie Nappe. 1 : Parpaillon Nappe; 2: Autapie Nappe; 3 : SubBrian~onnais Units; 4: Other internal Alps units; 5: Meso-Cenozoic sediments of the Dauphinois zone (external Alps); 6: External crystalline massifs.

oceanic crust above the European continental margin (Piedmont zone) (Tricart 1984; Malavieille et al. 1984). The Autapie Nappe was emplaced during the Late Eocene (Priabonian) in the external Alps (sub-Brian~onnais and Dauphinois Zones) as a submarine gravity nappe sliding westwards (or north-westwards) with a frontal olistostrome in a turbiditic foreland basin (Kerckhove 1969; Debelmas & Kerckhove 1973; Kerckhove et al. 1978; Kerckhove et al. 1981 ; Caron et al. 1981). No metamorphism was associated with the nappe emplacement. During the Oligo-Miocene, the nappe and its substratum were involved in polyphased westward to south-westward vergent thrust and fold tectonics (including the emplacement of the sub-Brian~onnais units). The Parpaillon Nappe, also gravity-driven, was emplaced in the area during the Early Miocene (Kerckhove 1969; Merle & Brun 1984).

1 Light-grey massive and strongly lithified calcareous mudstone layers probably deposited by low-density turbidity currents (Te Bouma division). 2 Turbiditic beds displaying usually base-missing Bouma sequences. These fining upward sequences show a basal sandstone layer overlain by a marlstone layer. At the base, the sandstone may be locally coarse to very coarse but is more usually fine to very fine. It consists of siliciclastic grains (mainly quartz grains and minor amount of phyllites, feldspars etc.) and of carbonate bioclasts (mainly pelagic foraminifera). Today, the sandstone is strongly lithified and has a very low porosity, mainly due to carbonate diagenesis (pressure-solution, recrystallization and cementation). The marlstone becomes progressively richer in clay towards the top and it usually presents a layer-parallel fissility. 3 Dark hemipelagic claystone layers showing a layer-parallel fissility.

Bed thicknesses vary greatly from one bed to another (a few centimetres to a few metres). The relative proportion of the four basic lithotypes also varies widely. These facies associations, as well as the original geometry of the beds, indicate a basin-plain environment for the Helminthoid Flysch deposition. Calcium carbonate-free hemipelagic clay testifies to a deep-sea deposition, below the calcite compensation level [i.e. below 3500-5000m, according to Hesse & Butt (1976) and Sagri 1979)].

IlOcm to lm ctaystone rnartstone sandstone

2

Lithology The Autapie Helminthoid Flysch consists of an originally regular succession of plane and parallel turbiditic beds and interlayered hemipelagics (Kerckhove 1969; Sagri 1979; Caron et al. 1981). Four basic lithotypes form three types of beds (Fig. 2):

1

3 I I

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1

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FIG. 2. Bed types and lithotypes of the Helminthoid Flysch of the Autapie Nappe. 1: Low-density carbonate turbidites; 2: Turbidites; 3 : Hemipelagics.

Syn-diagenetic deformation, Autapie Nappe, French Alps I49 Structural style of the deformation claystone are intensely disrupted and show important offsets along the shear planes (in many stage studied The polyphased evolution of the nappe is recorded by a complex structure pattern. An early stage of deformation is characterized by a pervasive stratal disruption resulting from multidirectional extension parallel to the layering (Fig. 3): at the outcrop scale, extensional ramps (normal faults) cut the sandstone and mudstone layers and flatten upwards and downwards as decollements in the marlstone and claystone layers (Fig. 4). Beds, or groups of beds, are thus truncated in lens-shaped, fault-bounded blocks of centimetric to decametric scale. There is no oblique cleavage associated with these structures. Variations in the stratal disruption intensity led Kerckhove (1969) to define in the nappe three basic deformation facies corresponding to increasing disruption, the 'normal', 'dissociated' and 'ultra-dissociated' facies, respectively. However, many bed ruptures also occur within the normal facies itself. These variations are spatially very closely related to the marlstone and claystone ratio: the parts of the succession poor in marlstone and claystone are relatively weakly deformed; on the other hand, the parts rich in marlstone and

cases more important than the outcrop dimensions, so that the finite extension usually cannot be measured) locally resulting in a 'block-inmatrix' or 'melange' texture (dissociated and ultra-dissociated facies, Fig. 3). Because of the geometry of the deformation and of the monogenetic constitution of the succession, there is no doubt that this block-in-matrix geometry results from intense shearing and is not of sedimentary origin (olistostrome) (Hsii 1974).

Structures in the sandstone layers Three types of microstructures may be associated with the extensional ramps cutting the sandstone layers: normal microfaults, calcite-filled veins and 'breccia' (Figs 4 and 5). All of these correspond to layer-parallel extension and layernormal shortening. They can be formed up to some decimetres from the ramps and progressively disappear further away (Fig. 4). Hereafter, we describe and discuss these various microstructures and then examine their geometrical and genetical relationships.

Calcite-filled veins Most of these veins (b in Fig. 5) are set at high angles to the bedding. They display a fibrous calcite-fill, indicating layer-parallel extension. The veins are very thin, usually less than 1 mm wide, exceptionally up to 5 mm. The vein opening very rarely induced breakage of grains, but rather occurred by grain to grain separation. Some grains have been completely separated from their neighbours and are isolated in the calcite fill. For this reason, some veins appear locally more like zones of diffuse increase of intergranular calcite cement than like well defined veins. These features suggest that the sandstone lithification was incipient when the veins developed: grain contact strength (resulting from compaction and probably from incipient cementation) was strong enough for the sand to behave as a cohesive and brittle material, but weak enough for the vein opening to be easier by grain to grain separation than by grain breakage.

Microfaults

FIG. 3. Typical outcrop of the Autapie Nappe, showing pervasive extensional stratal disruption.

These microfaults (a in Fig. 5) usually show a dip of 45°-60 ° with respect to the layering and induced millimetric to centimetric offsets. Faults are either synthetic or antithetic to the extensional ramps. Faults rarely occur alone but usually

150

P. Labaume

FIG. 4. Sandstone layers cut by extensional ramps (a). Note the breccia-like deformation (b) associated with layer rupture along one of the ramps.

FIG. 5. Deformed sandstone layer showing association of genetically related different types of microstructures (large-size thin section). (a) Microfaults; (b) pre-breccia calcite-filled vein; (c) breccia; (d) sand dyke formed during the brecciation (frame: Fig. 1l). The sandstone becomes progressively finer and more phyllite-rich towards the top of the layer. constitute anastomosed networks with millimetric to centimetric spacing of individual faults. A fault across a sandstone layer consists of a deformation band 0.1-2 mm wide and does not display a discrete slip plane. On the other hand, where a fault separates the sandstone layer from marlstone or claystone, the sandstone surface displays a striated slickenside.

Deformation bands Fig. 6 shows a deformation band 0.5 mm wide

which induced an 8 mm offset. Outside the band, the long axis of grains lies parallel to the bedding. Inside the band, most of the elongated grains have been reoriented parallel to the band walls (a in Fig. 6), evidencing deformation by important grain to grain movements. Progressive reorientation of these grains (particularly in folded phyllites, b in Fig. 6) close to the band walls resulted in a shear zone geometry for the band and indicates the sense of movement. Inside the band, small grains are more abundant than in the

S y n - d i a g e n e t i c deformation, A utapie Nappe, F r e n c h A l p s

151

reorientation of grains or grain fragments. This indicates that the grain contact strength was weak and thus the faults formed in un- or poorlylithified sediments. Similar natural deformation bands have been described by Aydin (1978), Cowan (1982), Aydin & Johnson (1983), Petit & Laville (this volume) and Underhill & Woodcock (this volume), who reached analogous conclusions. These conclusions are well supported by the strong analogy existing between the natural structures and those obtained experimentally both in: (i) unlithified sand (Borg et al. 1960; Friedman et al. 1980) or other granular material (Mandl et al. 1977) and (ii) in partly cemented porous sandstone (Handin et al. 1963; Dunn et al. 1973). From these experiments, the grain cataclasis is thought to result from stress concentration at grain contacts. Therefore, its intensity would be representative of the intergranular strength opposed to the grain to grain movements and is principally controlled by:

FIG. 6. Deformation band (microfault) in sandstone (thin section), a: reoriented elongated grains; b : folded phyllite; s: layering. Dark colour of the band results from post-faulting alteration by clay minerals.

undeformed sand, and calcite grains are rare. This suggests that intergranular movements were accompanied by grain cataclasis, mainly undergone by the calcite grains. This assumption is supported by the observation of some cracked grains outside the band. The band is brown in colour, due to a later alteration with the development of clay minerals. Another deformation band, 0.5 mm wide and with a centimetric offset, is shown in Fig. 7. In this case, the grain size reduction by cataclasis inside the band was more intense than in the previous case and very few large grains remain. In particular, no calcite crystal can be found inside the band. Between the grains, a 'dirty', dark and optically irresolvable matrix is probably made of the smallest siliciclastic grain fragments and the completely crushed calcite grains. Outside the band, many cracked grains can be found, crack abundance increasing towards the band walls, These deformation bands can be considered as ductile faults, the sandstone ductility primarily corresponding to the relative displacements and

1 The effective confining pressure, which is the difference between the confining pressure carried through the rock (burial) and the pore fluid pressure, as established by Terzaghi (1943) and applied to rock fracturing by Hubbert & Willis (1957), Hubbert & Rubey (1959) and many others. Handin et al. (1963) have shown that effective confining relief (i.e. pore pressure increase under constant burial or burial relief under constant pore pressure) tends to reduce cataclasis and to limit it to the deformation band. 2 The porosity. Dunn et al. (1973) have shown that high porosity (i.e. weak lithification) also tends to reduce cataclasis and to increase the deformation band width. Therefore, the weaker cataclasis in the case of the microfault of Fig. 6 than that in the case of Fig. 7 may be due to a lower burial, a higher pore pressure or a weaker lithification. Nevertheless, it may also be due, all other things being equal, to the larger amount of: (i) calcite grains which preferentially underwent cataclasis because of their strength being inferior to that of quartz grains, thus 'protecting' the latter, and (ii) phyllites which acted as a 'lubricant', facilitating the intergranular movements along their borders.

Slickensides The slickensides exhibit characteristic features (Fig. 8): 1 They are usually undulated. 2 They are rough, with the same weathering aspect as the non-deformed sand. Nevertheless,

15 2

P. Labaume

FIG. 7. Deformation band (microfault) in sandstone (thin section), s: layering. Note the strong grain size reduction by cataclasis inside the band.

locally, they may display a glossy aspect resulting from the spreading of phyllites on the fault plane during faulting. 3 There are no synkinematic fibres of calcite or quartz. 4 Striae indicating the direction of movement consist of grooves, often curved, of millimetric width developed in the sand itself. 5 The sense of movement is given by secondary shear-planes (Riedel planes) and by asymmetric grooves associated with dragged grains. The hanging-wails of the R iedel planes usually present rounded and irregular, locally breccia-like (a in Fig. 8), geometries.

Fro.8. Slickenside in sandstone. (a) Riedel-like shear plane giving the sense of movement (the missing block was moved upwards; note the brecciated texture of the hanging-wall); (b) claystone-filled vein.

In thin section perpendicular to a slickenside, the fault in the sandstone appears as half of a deformation band, the slickenside being in the deformed sandstone. The Riedel structures also consist of deformation bands. These slickenside features are probably related to ductile deformation of the sandstone by grain to grain movements associated with slip along the slickensides. Therefore, these features would be specific to faults developed in un- or poorlylithified sediments. Laville & Petit (1984) and Petit & Laville (this volume) have described similar slickensides formed in unlithified fluviatile sandstone. Such faults can easily be discriminated from those developed in strongly lithified sandstone, which are characterized by plane slickensides exhibiting a polished aspect resulting

Syn-diagenetic deformation, Autapie Nappe, French Alps from grain abrasion, synkinematic fibrous crystallizations and sharp geometry of the Riedel planes (Petit et al. 1983). The deformation bands across sandstone do not display slickensides because, in this case, the movement along the fault was only accommodated by ductile deformation in the width of the band, without loss of cohesion along a discrete slip plane. Such rupture planes were only initiated where the faults separated the sandstone layers from adjacent marlstone or claystone layers. Aydin & Johnson (1983) have noted that slip surface initiation inside a deformation band across porous sandstone (i.e. 'change in the style of deformation from continuous and zonal to discontinuous and planar') occurred only when displacement along the fault was of large (metric) amplitude.

Breccia Breccia corresponds to a more complex and penetrative deformation associated with extensional ramps than the deformation induced by the microfaults. At the outcrop, the boundary of the deformed layer exhibits sandstone 'clasts' of millimetric to decimentric size, in many cases in relief, in a sandstone 'matrix' (c in Fig. 9). Usually, marlstone or claystone has flowed between the clasts and, at very deformed layer boundaries, some

153

clasts or groups of clasts have been separated from the layer itself. Such deformation is difficult to analyse at the outcrop, because of its chaotic geometry and, locally, because of the diffuse transition between sandstone and marlstone or claystone. In thin section, one can see that the clasts are formed by sand in which sedimentary laminations have been preserved (c in Fig. 5 and Figs 9 and 10). Although the clast shape can be irregular, many of them are subrectangular, with borders respectively subperpendicular and subparallel to the bedding. Clasts are usually rounded with either sharp or more diffuse borders. Internal layering of the clasts indicates, for some of them, important body rotation; some elongated clasts have been folded. Around the clasts, the matrix corresponds to sand deformed mainly by grain to grain movements: all sedimentary structures disappeared and shear movements occurred, as indicated by the local preferred orientation of elongated grains. Evidence of quartz grain cataclasis is very rare; on the other hand, cataclasis of calcite grains occurred, as calcite grain size is usually reduced in the deformed sand and foraminifera truncated at clast borders have been observed. Geometrical and compositional relationships between clasts and matrix show that the latter is originated from both the loss of cohesion of the immediately adjacent clasts and from the injec-

- .~

FIG. 9. Early calcite-filled veins and later breccia in sandstone (large-size thin section), a: pre-breccia calcite, filled vein opened and injected with sand during the brecciation (frame: Fig. 10); b: clast of undeformed sandstone; c: matrix of sandstone deformed by grain to grain movements (its dark colour is due to post-breccia pressure solution of calcite and alteration by clay minerals); d: microfault related to the opening of the vein a; e: post-breccia calcite-filled vein; s: layering.

154

P. Labaume

tion between the clasts of sand issued from more distant (up to some centimetres) disaggregated internal layers of the bed. Such injections gave 'sand-dyke' geometries, as shown in Fig. 11, where the injected sand originated from a thin phyllite-rich layer interbedded between thicker phyllite-poor layers. The injected sand shows a strong fabric of grain long-axis, parallel to the dyke borders. Close to the dyke borders, the enclosing sand locally shows disorganization and progressive disappearance of the layering, evidencing incipient brecciation by in situ grain mobilization (b in Fig. 11). The deformed sand of the breccia matrix is usually brown in colour, thus darker than the undeformed sand. This is due to strong post-breccia pressure-solution of

calcite and alteration with development of clay minerals. These features suggest that the lithification was weak enough to allow a loss of cohesion of part of the sandstone, but strong enough to allow the preservation of the clasts. The lithification was also inhomogeneous inside individual beds, since the sand of some internal layers was completely mobilized, while large clasts were preserved in others. As in the case of the microfaults, occurrence of phyllites played a major role in facilitating the grain mobility (i.e. sand dyke of Fig. 11). The grain size was also an influential parameter, as shown by the increased brecciation intensity linked with the grain size increase illustrated in the lower part of the layer

FIG. 10. Early calcite-filled veins and later breccia in sandstone (thin section, situation in Fig. 9). a: pre-breccia calcite-filled vein; b : layered sandstone clast; c: matrix; d : both sides of the same vein, opened and injected by the matrix; e : calcite clast torn away from the neighbouring vein-fill during the matrix flow.

Syn-diagenetic deformation, Autapie Nappe, French Alps in Fig. 5. This was probably due to a lower volume of intergranuIar pores and a higher number of grain contacts in the finest sandstone, facilitating, in the latter, a more rapidly efficient lithification. Sand injections are likely to result from fluidization [grain transport by high-speed escaping pore fluid (Lowe 1976)], implicating porewater pressure gradients and high strain rate. This suggests that both these parameters controlled the partial loss of cohesion of the sandstone characterizing brecciation. Microstructure relationships and deformational sequence The time and space relationships between the microstructures described above indicate the following chronological sequence: (i) normal microfaults and calcite-filled veins; (ii) normal microfaults and breccia. During the first stage, the calcite-filled veins never constituted enbchelon precursor sets to the microfaults but, in many cases, they corresponded to a layer-parallel opening at the tip of microfaults. During the second stage, many clast borders and sand injections were initiated by rupture along early calcite-filled veins. These relationships are well illustrated by Figs 9 and 10, where the opening of a vein (a in Fig. 9) was related to the slip along a fault (d in Fig. 9). The first stage of the vein opening was accompanied by calcite crystallization, and the second by sand injection during the brecciation. During the latter stage, some calcite crystals, or groups of crystals, tore away from the vein-fill and were incorporated into the flowing sand as clasts (e in Fig. 10); some of them underwent cataclasis during this process. Many other remains of early calcite-filled veins can be found in the breccia, either across clasts, at clast borders, or as isolated crystals, or groups of crystals, in the matrix (black in Fig. 10). Another example of microstructural association is given in Fig. 5, where a set of normal faults developed in the upper part of the sample while breccia occurred at the base, with increasing intensity towards the bottom. In the transitional medium area, early calcite-filled veins were reactivated as faults during the brecciation. These veins were formed subperpendicular to the bedding and later shear movement along them was accompanied by rigid body rotation of the microlithons. The sand dyke of Fig. 11 was formed during the brecciation along one of these early veins; calcite-fill cataclasis during the sand injection occurred, as in the case presented in Fig. 9. Some remains of an early calcite-filled

155

vein are also found in the intensely brecciated lower part. On the other hand, no early veins can be found in the upper part, where only faults developed. We can thus infer a two-stage deformation: layer-parallel extension was accomodated: (i) in the upper part by normal faulting during the whole deformational sequence, and (ii) in the lower part by calcite-filled vein opening during the first stage of deformation and by brecciation during the second one. This second stage induced more important finite extension than the first one. These vein/breccia relationships, as well as the common clast geometry (borders subparallel and subperpendicular to the bedding), suggest that brecciation was initiated as tensile layer-parallel extension in continuity with opening of the previous calcite-filled veins, but with a drastic change in the deformation mechanisms. This change consisted of a partial loss of cohesion of the sandstone between the two stages, probably resulting from an increase in strain rate inducing pore pressure gradients: with high strain rate, calcite crystallization could not follow the vein opening any longer; this would have induced a relief of the horizontal confining, as well as the pore pressure in the opening zones, thus draining intergranular water towards them. In the less lithified sandstone, this resulted in immediate vein wall disaggregation and collapse; where the cohesion was strongest (particularly where the opening was along early calcite-fill giving rigidity to the walls), actual openings allowed injections of grains from adjacent areas. As deformation continued, shear movements occurred in the matrix, inducing important body rotations of the clasts and cataclasis of the calcite grains. This in turn helped to 'erode' the clasts which were rounded and reduced in size while the matrix developed. This resulted in the complex brecciated texture with abundant matrix and largely occulted initial rupture geometry. This vein/breccia succession gave an unusual chronology of deformation mechanisms, since brecciation can be referred to as partially ductile, 'soft-sediment' behaviour while the early calcitefilled veins indicate brittle behaviour, classically characteristic of 'rock deformation'. The distribution of shear and tensile structures is usually not random: Fig. 5 shows that the microfaults preferentially developed in the phyllite-rich layers, and the tensile microstructures in the phyllite-poor layers. This may be due to a higher tensile strength and a lower shear strength of the former than of the latter: as shown in Fig. 12, the shape of the failure envelope of the phyllite-rich sandstone probably tends toward that of a shale (De Sitter 1966, fig. 14). In

P. Labaume

156

i~,11 IIIIII

I

F16.11. Upward injection structure (sand dyke) of phyllite-rich sand formed along early calcite-filled veins during the brecciation (thin section, situation in Fig. 5). a: phyllite-poor undeformed layered sandstone; b: sandstone affected by grain to grain movements in the enclosing sandstone of the sand dyke and veins; c : upward injected phyllite-rich sand; d: early calcite-filled vein broken during the injection. consequence, for a given state of stress, rupture could occur in the shear field for the phyllite-rich sandstone (A in Fig. 12) and in the tensile field for the phyllite-poor sandstone (B in Fig. 12). Since the influence of the phyllite grains consisted of facilitating grain movements along their borders, such control of the sandstone behaviour indicates a weak state of lithification when the deformation occurred. In conclusion, all microstructure features in the sandstone indicate a weak lithification state at the deformation time. They suggest that early cementation was probably already achieved, but was weak enough to be locally broken, its breakage being easier than that of the grains

themselves. This allowed granulometric and compositional variations to firmly control the sandstone behaviour (i.e. the fault-tensile structure distribution, the cataclasis intensity inside the deformation bands and the intensity of the grain mobility during the brecciation), even where these variations are weak (Fig. 5). Completely unlithified sand would probably have undergone more complete grain mobilization, while strongly lithified sandstone would show no grain mobilization and poor influence of the sand composition. This latter point is well illustrated by the later structures of the nappe (see below). The deformation mechanisms also indicate the occurrence of the role played by pore-water. We

Syn-diagenetic deformation, Autapie Nappe, French Alps

/

lb I \

I

0"3 eff,

0"~ eff.

(-T)

(3T)

or,..

FIG. 12. Possible envelopes of phyllite-rich (A) and phyllite-poor (B) weakly lithified sandstones. For a given state of stress (Mohr circle), rupture occurred in the tensile field for the latter (b) and in the shear field for the former (a). al effand a3 eft: greatest and least principal effective stresses, respectively; T: tensile strength of the phyllite-poor sandstone.

have discussed how changes of the effective confining pressure (i.e. of pore pressure or of burial) and of the strain rate may have influenced the deformation mechanisms in addition to the influence of the lithological variations. Nevertheless, since all these parameters were likely to vary both in time and space throughout the nappe, an accurate determination of their respective influence remains speculative.

157

Figure 14 illustrates a case in which a vein network in mudstone was injected by sand and clay grains originating from neighbouring siliciclastic layers, the injection having probably occurred during brecciation of the latter. This injection is thus similar to those presented in Figs 10 and 11, but, in this case, it did not remain internal to the original bed. The calcite fill of the veins was broken during the injection and many calcite crystals, or groups of crystals, have been integrated into the injected sediment; although some mudstone clasts remained angular in shape, they were also probably partially disaggregated (mud flow), since some of them are rounded, and the injected sediment is rich in carbonate mud. This example emphasizes: (i) the role of porewater pressure gradients which induced intergranular water movements fluidizing unlithified sediments, and (ii) the difference of lithification state between the mudstone and the sandstone during the deformation.

Marlstone and claystone layers The marlstone and claystone layers acted as decollement levels while the sandstone and

Structures in the mudstone layers In contrast to the sandstone layers, the mudstone layers present, almost everywhere, a strongly brittle behaviour. The normal faults display synkinematic fibrous calcite crystallization. The calcite-filled vein networks are strongly developed near to the extensional ramps, giving a brecciated texture to the bed (Fig. 13). Except for one observed case, the calcite-filled veins are never post-dated by mud flow and the clasts are always angular in shape. These features indicate that, unlike the sandstone, lithification of the mudstone was advanced when the deformation occurred. Although the vein network geometries can be complex, the vein opening basically corresponds to layer-parallel extension. As in the case of the sandstone layers, many veins teminate along shear-planes, the vein opening being related to slip along them. Some of these shear-planes are oblique with bedding normal faults but most of them are parallel to the bedding plane (b in Fig. 13). Layer-normal shortening was thus reduced and no related pressure-solution (stylolitization of the bedding planes) was observed.

FIG. 13. Calcite-filled vein network (a) giving a brecciated texture in a mudstone layer, b: layering, which acted as decoUement levels.

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P. Labaume

FIG. 14. Sand and clay grain injection structure formed in a calcite-filled vein network in a mudstone layer (thin section), a: mudstone; b: pre-injection calcite-filled vein; c : injected sediment. Arrow: sense of injection. This injection occurred during the brecciation of siliciclastic layers close to the mudstone layer. mudstone layers were truncated by ramps; this indicates a viscosity difference between these layers, the latter showing a greater competency than the former. Plastic behaviour of the marlstone and claystone is evidenced by some marlstone and claystone-filled veins in the sandstone and mudstone layers (b in Fig. 8 and a in Fig. 15). This indicates that marl and clay were still unlithified when the deformation occurred. These veins are equivalent to the calcite-filled veins described above, but they joined the bed surfaces (base or top), so that adjacent marl or clay could flow (upward and downward) into them, up to several decimetres from the bed surface. A layerparallel fissility in the marlstone and claystone layers probably results from both compaction and tectonic layer-normal shortening.

Relationships of the mierostructure development with strain and pore fluid pressure in the series The microstructures described above are associated with the extensional ramps across the sandstone and mudstone layers and therefore

probably result from internal deformation of these layers prior to the ramp initiation (Fig. 16). The development of a tensile rupture regime during this deformation implies that, locally: (i) 0-3 eff (the least principal effective stress) was layer-parallel and negative, and (ii) 0"1 e f t - 0"3 eft (the differential stress) was small [theoretical values are 0"3 eft= - T and al eft-o- 3 eft= 4T, T being the tensile strength of the rock, according to Secor (1965), a in Fig. 12]. This may result from the fact that the initiation zones of the ramps are analogous to releasing overstep of faults, where mathematical simulations have shown that tensile rupture conditions can be realized (Segall & Pollard 1980; Liu 1983). The strain rate increase which may explain the veinbreccia transition may have been immediately prior to the bed rupture itself, as observed in many rock rupture experiments ['tertiary creep' (Price 1966)]. High pore pressure conditions may have assisted the rupture by shifting the Mohr circle leftward, according to the effective stress principle (Hubbert & Rubey 1960), but this was not a necessary condition, considering the evolution of the stress conditions prevailing in the overstep areas.

Syn-diagenetic deformation, Autapie Nappe, French Alps

FIG. 15. Claystone-filled veins (a), partially emptied at outcrop, in sandstone. Note later calcite-filled veins (b), developed when the clay was no longer plastic enough to flow into the veins.

The bulk deformation of the succession basically consisted of layer-parallel decollements in the claystone and marlstone layers. This may result from high pore pressure conditions, reducing friction and thus facilitating movements along these low-angle shear planes (Hubbert & Rubey 1959). We can thus postulate the occurrence of high pore-pressure conditions which could not be easily constrained from the microstructure discussion. High pore-pressure could develop because in this incompletely lithified and dewatered succession, the marl and clay layers were barriers to the water escape. This assumption is consistent with the fact that the more deformed parts of the succession are the ones richer in marlstone and claystone. Therefore, the deformation was contemporaneous with compaction and water escape and it can also be considered as a local 'catastrophic' increment of these processes since it involved important mobilization of both the intergranular water and the grains. Although the structures that we describe show some soft-sediment features, they do not corre-

159

spond to slumps, i.e. intrabasinal, superficial synsedimentary gravity sliding (Gawthorpe & Clemmey 1985): (i) soft-sediment features are here restricted to very specific structures, the sediments being more lithified than in most classical slumps in basinal series; (ii) no evidence of synsedimentary structures can be found anywhere, i.e. no turbiditic deposits cover earlier structures. The structures are clearly post-sedimentary and must be regarded as classical tectonic features. Because of its relationship with lithification, this deformation can be referred to as syn-diagenetic. Because of the local importance of the finite extension, this deformation stage cannot easily be expected to have occurred in a laterally confined area. It is more probable that it resulted from gravity spreading of the succession, probably associated with the first stages of the gravitydriven emplacement of the nappe. Deformation occurred in very superficial conditions: as indicated above, the Helminthoid Flysch was the last stratum deposited in the Tethyan Ocean and the nappe was emplaced as a submarine nappe, without having undergone any metamorphism. Its present thickness does not exceed 750 m, and it is generally much less than that. Its original thickness was greater, considering the local importance of extension, but it is unlikely that the structures described above developed to a depth of more than 2 or 3 km.

Diagenesis and deformation mechanisms post-dating the early bed ruptures We have seen that where the grain cataclasis was weak, the deformed sandstone in the microfaults and breccia matrix was preferentially affected by later pressure-solution in calcite and by alteration by clay minerals. This probably resulted from easier fluid circulations (i.e. increased permeability) in the deformed sandstone than those in the undeformed one, due to a porosity increase resulting from: (i) destruction of early cements, and (ii) dilatancy during grain to grain movements. Therefore the immediate effect of the sandstone flow was to decompact the grains, but this in turn facilitated a stronger diagenesis in this deformed sandstone than in the undeformed one. In contrast, the sandstone in the highly cataclastic deformation band of Fig. 7 is less altered than the undeformed sandstone, probably because of a porosity (i.e. permeability) reduction resulting from the severe grain size reduction in this case.

I6o

P. Labaume

MUDSTONE" strongly lithified

1- brittle microfaults and calcite-filled veins

2- id.

3- bed

rupture

SANDSTONE" weakly lithified

1 ] - d u c t i l e microfaults and calcite-filled veins

2- d u c t i l e m i c r o f a u l t s and semi-ductile breccia

3- bed rupture

FIG. 16. Model of development of stratal disruption in the Helminthoid Flysch of the Autapie Nappe. The internal deformation of the mudstone and sandstone layers is related to the formation of extensional ramps joining decollements in the claystone or marlstone layers. The injection of siliciclastic grains inside a calcite-filled vein network in a mudstone layer presented in Fig. 15 occurred during stage 2, the sandstone and mudstone layers being adjacent in this case.

Later structures developed when all the sediments were strongly lithified and they emphasize the specificity of the structures described above. In the sandstone : (i) the slickensides exhibit grain abrasion and fibrous calcite and quartz crystallization, and (ii) calcite and quartz-filled vein opening induced grain breakage and no grain mobilization (no brecciation, e in Fig. 9); lithological changes of the sandstone did not control its behaviour any longer. In the marlstone and claystone, brittle behaviour is indicated by the presence of wide veins opened by boudinage, which exhibit calcite and quartz fill and which frequently remained partially empty; the marlstone and claystone were then no longer plastic enough to flow into the veins opened in the mudstone and sandstone layers (b in Fig. 15). Therefore, discrimination of the earlier and later stages of deformation by the determination of deformation mechanisms is a very useful tool for studying the structural evolution of the nappe.

Conclusion We have presented an example of deformation resulting from extensional strain associated with the gravity driven emplacement of a submarine nappe. The deformation regimes and mechanisms were closely controlled by the nature and the state of lithification of the sediments, the strain rate and the occurrence of pore-water pressure gradients. The latter developed during tensile ruptures which may have resulted from

horizontal confining relief in decollement overstep areas. Actual excess pore pressure cannot be proved, but was likely to occur in this watersaturated and marlstone and claystone-rich series. It could explain why the bulk deformation of the succession basically consisted of layer-parallel decollements. Detailed discussions have shown that the parameters controlling the deformation regimes and mechanisms can be appreciated, but that accurate determination of their respective influences remains speculative because they are likely to have varied, either together or separately, in both time and space through the nappe. Some aspects of the sandstone, marlstone and claystone behaviour places such deformation in the category of 'soft-sediment' deformation. But there is also evidence that the sandstone was partially lithified and the mudstone strongly lithified. This deformation can thus be referred to as 'syndiagenetic' and represents an intermediate stage between classical soft-sediment deformation on the one side and rock deformation on the other. We have shown that this was a critical state for the sandstone, thus enabling an unusual deformation mechanism chronology to occur; partially ductile soft-sediment structures (brecciation) post-dating brittle, rock structures (calcitefilled veins). Therefore, our study sheds light on the progressive lithification of the various components of a basinal succession and emphasizes that: (i) a wide range of structures and associated deformation mechanisms exist among the structures referred to as 'soft-sediment' deformation, and (ii) soft-sediment deformation is not neces-

Syn-diagenetic deformation, Autapie Nappe, French Alps sarily restricted to synsedimentary structures, but can also result from purely post-depositional, tectonic, deformation. Since the deformation involved intergranular water and grain mobilization and affected later fluid circulations in the poorly lithified sandstone, it can also be considered as a catastrophic increment of dewatering and compaction of the succession. Analogy can be made with convergent plate margins where tectonic deformation has been recognized to increase dewatering and compaction (Caron et al. 1982; Bray & Karig 1985; Fowler et al. 1985). Deformation processes described here may be expected to be encountered at modern convergent margins: many accretionary prisms are composed of very similar partially lithified turbiditic succession and the importance of fluid overpressuring in the prism evolution is now well established (Von Huene & Lee 1982; Von Huene 1984; Shi & Wang 1985), either from seismic data (BijuDuval et al. 1982; Westbrook & Smith 1983), core interpretation (Caron et al. 1982), in situ core measurement (Moore et al. 1982) and outcrop observations, both on land (Legget & Platt 1985) or during deep-sea dives (Suess & Massoth 1984; Boul/~gue et al. 1985). Specific deformation mechanisms resulting from these soft-sediment and high pore-pressure conditions are probably the cause of aseismic deformation in the prisms (Chen et al. 1982).

161

Analogy can also be made between the Autapie Nappe and these units, commonly referred to as 'melanges'. Cowan (1985) has recently pointed out that this term covers a wide range of geological units of sedimentary as well as tectonic origin, which developed under various conditions of pore-pressure and states of lithification of the sediments. He also postulated that melanges may develop in various tectonic settings, their common character being that they characterize convergent margins (see Cowan for specific bibliography). We have shown that, in our case: (i) accurate assumptions can be made concerning both the general tectonic setting and the deformation geometry and mechanisms, and (ii) discrimination of the various stages of deformation by determination of deformation mechanism changes is a useful tool in studying the structural evolution of this kind of structural unit.

ACKNOWLEDGEMENTS"This study was realized with the financial support of the Centre National de la Recherche Scientifique, France, as part of the programme 'A.T.P. Seismog~n~se, Plis, Failles'. A. Etchecopar, M. S6guret and J. M~gard-Galli are greatly acknowledged for stimulating discussions and critical reviews of the manuscript, B. Sanche for making the thin-sections, M. F. Roch and J. Faure for help with the typewriting and M. Mareschal for improvement of the English.

References AYDIN, A. 1978. Small faults formed as deformation bands in sandstone. Pure App. Geophys, 116, 91330. & JOHNSON,A. M. 1983. Analysis of faulting in porous sandstones. J. struct. Geol. 5--1, 19-31. BIJU-DUVAL,B., LE QUELLEC,P., MASCLE,A., RENARD, V. & VALERY, P. 1982. Multibeam bathymetric survey and high resolution seismic investigations on the Barbados Ridge Complex (Eastern Caribbean): a key to the knowledge and interpretation of an accretionary wedge. Tectonophysics, 86, 275304. BORG, J., FRIEDMAN,M., HANDIN, J. & HIGGS, D. V. 1960. Experimental deformation of St. Peter sand: a study of cataclastic flow. In: GRIGGS, D. & HANDIN, J. (eds) Rock Deformation. Geol. Soc. Amer. Mem. no. 79, 133-91. BOULI~GUE, J., LE PICHON, X. & IIYAMA, J. T. 1985. Pr6vision des tremblements de terre dans la r6gion du Tokai (Japon). Comptes Rendus des Sbances de l'Acadkmie des Sciences, Paris, 301 (S6rie II)-16, 1217-9. BRAY, C. J. & KARIG, D. E. 1985. Porosity of sediments in accretionary prisms and some implications for -

-

dewatering processes. J. geophys. Res. 90--B1,76878. CARON, C., HESSE, R., KERCKHOVE,C., HOMEWOOD, P., VAN STUIJVENBERG,J., TASSE,N. & WINKLER, W. 1981. Comparaison prGliminaire des flyschs ~t Helminthoides sur trois transversales des Alpes. Eel. Geolog. Helvet. 47-2, 369-78. CARSON, B. H., VON HUENE, R. & ARTHUR, M. 1982. Small scale deformation structures and physical properties related to convergence in Japan trench slope sediments. Tectonics, 1(3), 277-302. CHEN, A. T., FROHLICH, C. & LATHAM, G. 1982. Seismicity of the forearc marginal wedge (accretionary prism). J. geophys. Res. 87-B5, 3679-90. COWAN, D. S. 1982. Deformation of partly dewatered and consolidated Franciscan sediments near Piedras Blancas Point, California. In: LEGGETT, J. K. (ed) Trench-Forearc Geology: Sedimentation and Tectonics on Modern and Ancient Active Plate Margins. Spec. Publ. geol. Soc. London, 10, 439-

57. 1985. Structural styles in Mesozoic and Cenozoic melanges in the Western Cordillera of North America. Geol. Soe. America Bull. 96, 451-62. DEBELMAS, J. 1975. RGflexions et hypothGses sur la --,

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pal6og6ographie cr6tac6e des confins alpino-apenniniques. Bull. Soc. Gkol. de France. (7), XVII-6, 1002-12. - & KERCKHOVE, C. 1973. Large gravity nappes in the French-Italian and French-Swiss Alps. In: DE JONG, K. A. & SCHOLTEN, R. (eds) Gravity and Tectonics. 189-200. DE SITTER, L. U. 1966. Structural Geology. McGrawHill, New York. DUNN, D. E., LAFOUTAIN,L. J. St; JACKSON,R. E. 1973. Porosity dependence and mechanisms of brittle fracture in sandstones. J. geophys. Res. 78-14, 2403-17. FITCHES, W. R. & MALTMAN, A. J. 1978. Conference report: Deformation of soft sediments. J. geol. Soc. London, 135-2, 245-51. FOWLER, S. R., WHITE, R. S. & LOUDEN, K. E. 1985. Sediment dewatering in the Mackran accretionary complex. Earth planet. Sci. Lett. 75, 427-38. FRIEDMAN, M., HUGMAN, R. H. H. I!I & HANDIN, J. 1980. Experimental folding of rocks under confining pressure, Part VIII--Forced folding of unconsolidated sand and of lubricated layers of limestone and sandstone. Geol. Soc. America Bull. 91, 30712. GAWTHORPE, R. L. 8¢ CLEMMEY,H. 1985. Geometry of submarine slides in the Bowland Basin and their relation to debris flows. J. geol. Soc. London, 142, 555-66. HANDIN, J., HAGER, R . V . Jr, FRIEDMAN, i . & FEATHER, J. N. 1963. Experimental deformation of sedimentary rocks under confining pressure: pore pressure test. Amer. Assoc. Petrol. Geol. Bull. 47-5, 717-55. HESSE, R. & BUTT, A. 1976. Paleobathymetry of Cretaceous turbidite basins of the east Alps relative to the calcite compensation level. J. Geol. 84, 50533. HOMEWOOD, P. & CARON, C. 1982. Flysch of the western Alps. In ."Hsf3, K. J. (ed) Mountain Building Processes. 2-2, 157-68. Academic Press, London. Hsf2, K. J. 1974. Melanges and their distinction from olistostromes. In: DOTT, R. M. J. 8¢ SHAVER,R. H. (eds) Modern and Ancient Geosynclinal Sedimentation. Spec. Publ. Soc. Econ. Paleont. Mineral. 19, 321-33. HUBBERT, M. K. & RUBEY, W. W. 1959. Role of fluid pressure in mechanics of overthrust faulting, I: mechanics of fluid-filled porous solids and its application to overthrust faulting. Geol. Soc. Amer. Bull. 70-2, 115-66. - & RUBEY, W. W. 1960. Role of fluid pressure in mechanics ofoverthrust faulting--reply. Geol. Soc. Amer. Bull. 71-5, 611-28. -& WILLIS, D. G. 1957. Mechanics of hydraulic fracturing. Trans. Amer. Inst. Mech. Engineers, 210, 153-68. HUENE, R. VON, 1984. Structural diversity along modern convergent margins and the role of overpressured pore fluids in subduction zones. Bull. Soc. Gdol. France, 7, XXVI-2, 207-19. -& LEE, H. 1982. The possible significance of pore fluid pressures in subduction zones. In: WATKINS, J. S. & DRAKE, C. L. (eds) Studies in Continental

Margin Geology. Amer. Assoc. Petrol. Geol. Mem. no. 34, 781-91. KERCKHOVE, C. 1969. La 'zone du Flysch' darts les nappes de l'Embrunais-Ubaye (Alpes occidentales). Gdol. Alpine, 45, 5-204. - - , BEBELMAS,J. & COCHONAT,P. 1978. Tectonique du soubassement parautochtone des nappes de l'Embrunais-Ubaye sur leur bordure occidentale, du Drac au Verdon. Gdol. Alpine, 54, 67-82. --, CARON, C., CHAROLAIS,J. & PAIRIS, J.-L. 1981. Panorama des s6ries synorog~niques des Alpes occidentales. In: AUTRAN, A. & DERCOURT, J. (eds) Evolutions G~ologiques de la France. Bur. Rech. G6ol. MiD. M6m. no. 17, 234-55. LAVILLE,E. 8¢ PETIT, J. P. 1984. Role ofsynsedimentary strike-slip faults in the formation of Moroccan Triassic basins. Geology, 12, 424-7. LEGGETT, J. K. & PLATT, J. 1985. Soft-sediment deformation and mud diapirism, Makran accretionary prism, paper presented at the conference on 'Deformation Mechanisms in Sediments and Sedimentary Rocks', University College London, April 15-17. LIU, X. 1,983. Perturbations des contraintes libes aux structures cassantes dans les calcairesfins du Languedoc. Observations microtectoniques et simulations mathbmatiques. (Unpublished thesis) University of Montpellier, France. LOWE, D. R. 1976. Subaqueous liquified and fluidized sediment flows and their deposits. Sedimentology, 23, 285-308. MALAVIEILLE,J., LACASSIN,R. ~¢ MATTAUER,M. 1984. Signification tectonique des lin6ations d'allongement dans les Alpes occidentales. Bull. Soc. G~ol. France, 7, XXVI-5,895-906. MALTMAN, A. 1984. On the term 'soft-sediment deformation'. J. struct. Geol. 6-5, 589-92. MANDL, G., DE JONG, L. M. J. & MALTHA, A. 1977. Shear zones in granular material. Rock Mechanics, 9, 95-144. MERLE, O. & BRUN, J. P. 1984. The curved translation path of the Parpaillon Nappe (French Alps). J. struct. Geol. 6-6, 711-9. MOORE, C. J., BIJU-DUVAL, B. & DSDP Leg 78A Scientific party. 1982. Offscraping and underthrusting of sediments at the deformation front of the Barbados Ridge: Deep Sea Drilling Project Leg 78A. Geol. Soc. Amer. Bull. 93, 1065-77. PETIT, J. P. & LAVILLE,E. Morphology and microstructure of hydroplastic slickensides. This volume. - - , PROUST, F. & TAPPONNIER, P. 1983. Crit6res de sens de mouvement sur les miroirs de failles en roches non calcaires. Bull. Soc. gdol. France, 7, XXV-4, 589-608. PRICE, N. J. 1966. Fault and Joint Development in Brittle and Semi-Brittle Rock. Pergamon Press, Oxford. SEGALL, P. & Pollard, D. O. 1980. Mechanics of discontinuous faults. J. geophys. Res. 85-B8, 433750. SAGRI, M. 1979. Upper Cretaceous carbonate turbidites of the Alps and Apennines deposited below the calcite compensation level. J. sed. Petrol. 49-1, 238.

Syn-d&genetic deformation, Autapie Nappe, French A lps SECOR, D. T. 1965. Role of fluid pressure in jointing. Amer. J. Sci. 263, 633-46. SHI, Y. & WANG, C. Y. 1985. High pore pressure generation in sediments in front of the Barbados Ridge complex. Geophys. Res. Lett. 12-11,773-6. SUESS, E. & MASSOTH, G. J. 1984. Evidence for the venting of pore waters from subducted sediments in the Oregon continental margin. EOS Trans. Amer. Geophys. Union, 65-45, 1089. TERZAGHI, K. W. 1943. Theoretical Soil Mechanics. Wiley, New York.

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TRICART, P. 1984. From passive margin to continental collision: a tectonic scenario for the western Alps. Amer. J. Sci, 284, 97-120. UNDERHILL, J. R. & WOODCOCK, N. H. Faulting mechanisms in high-porosity sandstones; PermoTriassic, Arran, Scotland. This volume. WESTBROOK, G. K. & SMITH, i . J. 1983. Long decoUements and mud volcanoes: evidence from the Barbados Ridge Complex for the role of high pore-fluid pressure in the development of an accretionary complex. Geology, 11,279-83.

P. LABAUME, Laboratoire de Tectonique (u.a. CNRS 266), Universitb des Sciences et Techniques du Languedoc, Place E. Bataillon, 34060 Montpelier Cedex, France.

Comparative studies of gravity tectonics in Quaternary sediments and sedimentary rocks related to fold belts S. A. Schack Pedersen S U M M A RY: Two different modes of gravity tectonics have been active as deformation mechanisms in the Quaternary of Denmark, namely 'gravity gliding' and 'gravity spreading'. Examples of gravity gliding can be found in the fairly slowly moving coherent landslides which characterize many coastal cliffs with lithologies intercalated with Tertiary marine clay or Quaternary interglacial clay. The gravity spreading mechanism generated folds and dislocation structures characteristic of glacial tectonics, mainly during the Weichselian glacier advances of the Pleistocene glaciation. Imbricate dislocation structures formed by diapiric rise of water-saturated clay into sand occur. This structure type is proposed to be characteristic for the gravity spreading model. Analysis of these gravity deformed Quaternary deposits serve as a model for structures occurring in large-scale tectonic settings. Among the megascopic tectonic settings mentioned, two regions are discussed. Firstly, the transverse thin-skinned thrust fault belt in the southern part of the North Greenland Fold Belt is demonstrated to contain the characteristics of a gravity gliding system. Secondly, the gravity spreading structures of the glacio-dynamic deformations in Denmark are suggested to have a megascopic equivalent in the platetectonic convergence zone of the Barbados Ridge.

The concept of gravity gliding is familiar to scientists studying landslides and related deformations in near-surface sediments, but it was first fully accepted as a model for deformation in thrust fault belts and folded mountain ranges after Hubbert & Rubey (1959) published their classic paper on the importance of anomalous high pore-fluid pressure to overcome the frictional resistance along the sole of the thrust block. About ten years later the gravity gliding model was widely rejected in favour of the gravity spreading concept (Bucher 1956; Price & Mountjoy 1970). Another ten years passed and Ramberg (1980) wrote, in a paper dealing with gravity spreading: 'It is generally accepted, though not definitely proved, that the large-scale structure of orogens is related to continental drift and seafloor spreading, and hence that the chief forces behind orogenesis are the same deep-seated ones which supposedly keep crustal plates moving'. Bally (1981) and Price (1981) expressed similar plate-tectonic thoughts and described the formation of thrust fault belts as being due to subduction and compression. However, discussions of gravity tectonics have continued and have resulted in the description of numerous structures as being characteristic of either gravity gliding or gravity spreading (Cooper 1981; Graham 1981; Mandl & Crans 1981). In this paper, a set of criteria is listed which permits the distinction between the deformation due to gravity gliding and that due to gravity

spreading. Although lack of adequate exposure may render distinction between the two mechanisms difficult, and bearing in mind both may act simultaneously, this distinction is essential for the study of deformation in sedimentary rocks.

Gravity gliding In the gravity gliding model, the deformation is caused by a (sedimentary) body moving downslope due to its own weight: the 'body force' principle. The structures characteristic of gravity gliding deformation are (Fig. 1): 1 Listric normal faults at the trailing edge of the basal thrust surface and a related truncation of the internal stratigraphy of the individual thrust sheets--the roll-over anticline (Kehle 1970; Hose & Danes 1973; Cooper 1981; M a n d l & Crans 1981; Wernicke & Burchfiel 1982). 2 Flat-lying coherent thrust sheets occasionally displaying extensional features, mainly grabens (Hose & Danes 1973) and related synthetic faults. 3 The diverticulation phenomenon (Lemoine 1973; Cooper 1981), whereby the stratigraphical uppermost units are displaced further than the lower units. 4 The exposure of a 'peel off' regime in the rear part of a thrust fault region. A very important relationship which is a consequence of the first criterion is the 'younger-on-

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 165-180.

x65

I66

S. A. Schack Pedersen

GRAVITY GLIDING

SPREADING

1

2

3

4 J FIG. 1. Schematic illustration of the characteristic features related to gravity gliding and gravity spreading. The four criteria set up for gravity gliding are given in the column to the left and are: (1) listric normal faults and extensional fault imbrication at the trailing edge of the basal thrust surface; (2) fiat-lying coherent thrust sheets disturbed by extensional graben faulting; (3) diverticulation phenomenon, whereby the stratigraphical uppermost units are displaced further than the lower units; (4) exposure of a peel-off regime in the rear part of the thrust fault region. In the column to the right four criteria of gravity spreading are pictured: (1) an imbricate fan formed by listric splay thrust faults in front of the overlying spreading mass, (2) a duplex of imbricate thrust sheets developed in the deformed and overthrust sediments. Due to loading, boudins may be formed beneath the spreading unit, (3) in the frontal part of the gravity-spreading deforming-system, syn-tectonic basins are formed with progressively younger sediments away from the front of the spreading mass, (4) due to increasing fluiddynamic overpressure in the decollementunit water-escape structures and mud-diapirs occur in the frontal region of the gravity-spreading system.

older' stratigraphic units (Hose & Danes 1973) in relation to the fiat-lying thrust faults.

Gravity spreading Gravity spreading denotes deformation caused through loading by an overlying medium which, during spreading, produces a lateral stress component that presses the underlying body forward and upwards along listric reverse thrust fault

surfaces. The concept of gravity spreading is based on a wax model used by Bucher (1956) in which solid wax spread outwards over interlayered wax and grease. The gravity spreading model was applied by Bucher (1956); Price & Mountjoy (1970) and Price (1973) for the interpretation of the structures in the foreland thrust belt of the Canadian Rocky Mountains. According to Cooper (1981 ), the gravity spreading model requires upwelling and lateral spreading of a mobile infrastructure in the core of the orogen.

167

Studies of gravity tectonics However, the concept of gravity spreading requires only the spreading of a body over an underlying body, for example a glacier spreading over underlying sediments. This situation was considered by Gry (1940), who experimented with a wooden plate and snow in order to reproduce dislocation structures formed by glacial tectonics. He described the mechanism as the 'ice weight hypothesis', but since the concept of gravity spreading is so well known, I prefer to use this term. The characteristic structures produced by gravity spreading are as follows (Fig. 1): 1 Listric splay thrust faults. 2 A duplex of imbricate thrust sheets developed as the sediments are deformed and overthrusted with nappe boudins being formed due to gravity collapse (Ramberg 1980). 3 The development, away from the margin of the loading body, of syn-tectonic basins filled with progressively younger sediments. 4 The development of soft-sediment waterescape structures and clay diapirs in the syntectonic sedimentary basins. The diapirism and upward moving structures, due to thermal convection within orogens (Talbot 1977; Ramberg 1980), are beyond the scope of the structures considered here, and are regarded as belonging to a separate class of thermal-gravity structures.

Gravity gliding settings The most obvious gravity gliding system in nature is the landslide. In Denmark landslides occur along the coastal cliffs where Tertiary marine clay or glacial laminated clay and silt crop out. There is a gradual transition from landslides to mudslides. The morphology and processes of mudslides are described by Prior (1977), and are regarded as sedimentary processes, whereas landsliding of coherent sedimentary bodies is a tectonic process. In the N W part of Denmark (the Limfjord region), landslide processes are very common. Two different lithologic sequences are here subjected to landsliding: (i) a Palaeogene sequence of diatomite with ash-layers grading into a black clay-rich unit at the bottom. This black clay contains the decollement level for the sliding; (ii) a glacio-lacustrine clay and silt sequence grading up into glacio-fluvial sand, often overlain by a till horizon. In Fig. 2 an example of a landslide active in the years 198283 is illustrated. At the head of the slide normal listric faults are developed. In the central part, the allochthonous mass exhibits horst and graben structures due to the extensional forces. In the

early spring of 1983 the toe of the slide was beautifully exposed due to low tide and calm weather, which prevented erosion of the frontal part of the slide. Here, isoclinal recumbent folds were observed with a 'fold nappe' thickness of 30-60cm, with the amplitude of the folds reaching the order of 10 m. Syn-tectonically deposited beach sand and shells had been overthrust by the toe and formed white or light grey-blue cataclastic sheets, or slides, between the 'fold nappes' and even infolded layers 1-5 cm thick in the sole of the fold structures. It is questionable whether these isoclinal, recumbent toe-folds are formed by gravity gliding or gravity spreading. However, it is obvious that such features may exist within a gravity gliding system. The landslide described above belongs to the group of slides with rather steep original surfaces (10°-15°). Another group of slides is characterized by original surfaces with dips less than 5° . In eastern Peary Land, North Greenland, in an area covered by elevated Neogene marine clay and sand deposits, the c. 100 m thick Kap Kebenhavn Formation (Funder et al. 1984), the area is dominated by recent landslides and mudslides. The landslide (Fig. 3) was active during the Kap Kobenhavn expedition's visit to the area in 1983 (Funder et al. 1984). The slide was 20 m wide and 50 m long. The sliding sheet consisted of finegrained sand, silt and clay. The thickness was between 40 and 60 cm, and the decollement surface was situated on the top of the permafrost level. It is obvious that the melting of permafrost had a dominant influence on the sliding process. The displacement of the sheet was about 25 m, which is the size of the peel-off area at the head

/ /

/ . ~ ~ .~'~\' ~

c. 25 m

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. ~ ~ ~

~

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FIG. 2. Block diagram of a recent landslide in the Limfjord region, NW part of Denmark. The landslide slipped along listric curved normal faults. The main allochthonous mass contains graben-structures, the

formation of which is due to extension in the lateral displacement direction. In the distal part of the slide a toe was formed with curved tongues of isoclinal recumbent folded clay and silt with infolded beach ridges of shells, sand and gravel, deformed into cataclastic slide-sheets, possibly due to a gravity spreading effect of the allochthonous mass itself.

I68

S. A. Schack Pedersen

FIG. 2. (A-D) Photographs. (A) Landslide of coastal cliff in the Limfjord region NW Denmark. (B) Head of slide with groove mark lineation. (C) Recumbent isoclinal folded clay infolding and overthrusting the beach deposits at the toe of the slide. (D) Landslide overthrusting the beach ridge consisting of sand, gravel and shell debris deposited during the winter storms. Intervals on measuring pole: 20 cm.

Studies o f gravity tectonics

of the slide (C in Fig. 3). In the proximal part of the main allochthonous body, the sheet shows extensional 'chocolate tablet' structures. Towards the more distal part of the slide, small lakes characteristically appear, a feature which is common in landslides. These lakes are a direct indication of the abnormally high pore-water pressures in the sediments which are required for gliding to occur. Along the sides of the landslide (D in Fig. 3), two small wrench faults limit the central part of the slide. On both sides of the fault, synthetic en echelon folds developed, with an angle of 30 ° between the crests and the fault line. To examine larger gravity sliding systems we must turn to submarine slides. Large-scale gravity slides and slump masses are reported from continental slopes and rises, but they are generally inferred from seismic profiles (Lemoine 1973; Reineck & Singh 1975, p. 387). The geometry of gravity slides and slump sheets as described by Lewis (1971) and Moore et al. (1970) shows a proximal denuded source area, a central area with flat-lying sheets, and a distal area characterized by listric thrust faults and frontal anticlines. The largest reported gravity slide thrust sheets are described by Dingle (1977; 1980) from the passive ocean margins of South Africa. The average thickness of these allochthonous bodies is 250 m and longest fault about 180 km. A single thrust sheet c. 200 m thick and 80 × 20 km in size,

16 9

glided for a distance of about 40 km (Dingle 1980). In fold belt terrains, gravity gliding is described in the Maritime Alps by Graham (1981), and is mainly identified by the normal listric fault structures at the trailing edge of the basal thrust surface. The thrust fault region of the Naukluft Mountains in SW Africa was originally described as having been deformed due to gravity gliding, the initiation of this being caused by basement

2 FIG. 3. Landslide developed on permafrozen slope in the Kap Kobenhavn area in East Peary Land, North Greenland (for location see Fig. 5). In the proximal part of the slide the main escarpment is steeply dipping from the earth surface to the permafrost surface. In the peel-off regime the frozen sediment of the permafrost ground is exposed, but it will successively be covered by mudslides and mudflow deposits. In the trailing edge of the allochthonous mass an 'iceland' is left in the peel-offarea.

FIG. 3. (A-C) Photographs of Peary Land Landslide. (A) Compressional folds and small-scale thrusts at the toe of the slide. (B) Central part of the slide containing the main allochthonous mass. The little lake in the foreground is formed due to the release of the porepressure. At the trailing edge of the allochthonous mass the sliding sheet is broken up into a chocolate tablet pattern due to extensional forces. Note the groove mark lineation on the strike-slip surface at the side of the slide. (C) Tear fault bordering the slide on the northern side of the main allochthonous mass. Note the en echelon folds on both sides of the vertical fault line.

~v

O

Studies of gravity tectonics uplift to the NE in the Damara orogen, and the Naukluft Nappes being displaced down-slope towards the SW (Korn & Martin 1959). The geotectonic setting of this mountain range has later been discussed (Martin & Porada 1977; Behr et al. 1981; Weber & Ahrendt 1983), and the thin-skinned nappes have now been described as a compressional foreland thrust belt. However, the cross-sections of the structures drawn by Korn & Martin (1959), show imbricate structures in the rear part of the thrust fault complexes which are most easily interpreted as normal thrust fault duplexes (Fig. 4). In the southern part of the North Greenland Fold Belt, a thrust fault region which in many aspects is comparable to the Naukluft Nappes has recently been discovered and mapped by the author (Pedersen 1980; 1981a). This thin-skinned thrust fault belt is related to a late Caledonian event and is interpreted as having been formed by gravity gliding (Pedersen 1981 b). This thrust fault region serves as a model for gravity gliding systems, with a scale between the large submarine slides and thrust fault deformations related to major compressional fold belt tectonics. A brief description is given below.

171

1980; Hhkansson & Pedersen 1982; Pedersen & Holm 1982). The lowermost unit in the region is a 400 m thick purple and green Cambrian mudstone. The main basal decollement zone was situated in the mudstone unit, which, in accordance with the criteria given for a gravity gliding lithology, mentioned by Guth et al. (1982), has all the qualities of a suitable sliding unit. Above the mudstone unit, two Ordovician groups appear. The lowermost consists of a 600 m thick sequence of coarsening upwards sandy turbidites, and above is a 400 m thick unit of black and cherty shales, irregularly interbedded with thick, redeposited limestone-conglomerates. The uppermost Ordovician shaly unit is also a convenient lithology for thrust fault fiats to develop. The uppermost stratigraphical unit is a sequence of Silurian flysch up to 3 km thick. This unit is regarded as a loading unit which created an increase of the pore pressure in the buried watersaturated clay and mudstone units. The Volvedal thrust fault region is divided into five zones according to structural characteristics (Fig. 6). Zone L

Situated in the foreland to the westward directed thrust faulting. To the south (Fig. 6), zone I is unconventionally termed a sideland, along which the thrust sheets were parallelly displaced. The sideland consists of a Lower Palaeozoic carbonate platform covered by Silurian flysch (Dawes 1976; Christie & Peel 1977). The thrust faulting did not affect the carbonate platform lithologies, and the border between deformed and undeformed regions may be caused either by a change in the Cambrian mudstone lithology (Surlyk & Hurst 1984) or in the sedimentological anisotropy created along the fossil carbonate platform margin.

The Vslvedal thrust fault deformation The thrust fault structures in the southern part of the North Greenland Fold Belt in Peary Land (Fig. 5), form a transverse thrust fault belt in relation to the main E - W trending EllesmerianNorth Greenland Fold Belt. The deformation has affected a sequence of Lower Palaeozoic unmetamorphosed clastic sediments about 4.5 km thick. These were deposited in the deep-water clastic trough of North Greenland (Friderichsen et al. 1982; Surlyk & Hurst 1984). The deformation event, called the Volvedal Orogeny (HS.kansson & Pedersen 1982; Pedersen & Holm 1982), is dated as latest Silurian/earliest Devonian time, and the structures were deformed by the early Carboniferous Ellesmerian folding (Pedersen ~/

Z o n e 11.

In this marginal zone, the thrust faults cut up into the Silurian flysch (Fig. 7) where the stress is j

/

/

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FIG. 4. Detail of trailing edge structures in the distal, compressional part of a gravity gliding system. Modified after Korn & Martin (1959).

I72

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S 20 °

...,¢.£~qLAND F O L D ~ ~u--

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FIG. 5. Location map of North Greenland. The position of the Volvedal thrust fault region is shown by hatched area. KK marks the location of the Kap Kobenhavn landslide area. absorbed in numerous thin bedding parallel thrust faults and slickensided bands. The Silurian flysch generally forms coherent flat-lying sheets, cut by ramp faults in only a few places, No long distance transport over fiats at the top of the flysch has been recognized. Along the southern margin the frontal anticlines of the thrust sheets are bent into parallelism with the sideland trend.

flat-lying sheets. The thrust sheets consist of the 150 m thick competent limestone-conglomerate units deposited within the black shales of the uppermost Ordovician group. In the rear part of the thrust sheet complex of zone IIIb, the only well exposed normal listric faults appear and examples of younger stratigraphical units thrust over older are present (Fig. 8).

Zone IlL

Zone IV.

This is dominated by thrust sheets consisting of the Ordovician shale, turbidites and redeposited conglomerates. In the northern part of the zone, IIIa (Fig. 6), imbricate thrust sheets appear steeply dipping towards the east, and a continuous N-S strike can be followed for about 10 km. The structural complex is regarded as a frontal imbricate fan and no mega-scale duplex, due to overthrusted nappes, is present. To the west, the imbricated thrust sheets are stacked against a steeply east-dipping ramp to the flat-lying Silurian flysch of zone II.

The thrust fault structures of this zone are flatlying, except in the western part of the area, where they become steeply dipping to the east due to the listric nature of the thrusts. In relation to this structural pattern there is a marked lithological distribution: in the eastern part of the zone the Cambrian mudstone unit dominates, whereas in the western part the Ordovician units form the main part of the steeply dipping thrust sheets. This framework is interpreted as being an indication of the diverticulation phenomenon. Zone V.

Zone IIIb.

This is characterized by long (c. 10 km) coherent

This is the easternmost area of the Volvedal thrust fault region (Fig. 6). The dominant

Studies o f gravity tectonics 3;"w

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FIG. 6. Map of the zonal division and the main trend of the thrust faults in the Volvedal thrust fault region. sedimentary rock in this zone is the Cambrian mudstone, and the area is regarded as part of a peel-off regime in the most proximal part of the deformation system. The remaining part of the peel-off regime, as well as the hinterland to the thrust fault region, is unknown because a large regional transcurrent fault has displaced the northern part of the thrust fault belt and its hinterland region.

FIG. 7. Thrust faults affecting the Silurian flysch in the middle part of zone II b of the Volvedal thrust fault region. The black unit folded in a hanging-wall anticline is Ordovician shales with limestoneconglomerate. The location is situated on longitude 34°W on the south coast of the main E-W trending fjord, see Fig. 6.

The most important field relationship of zone V occurs in the south-eastern part, where a thrust sheet of Silurian fiysch is thrust over the Cambrian mudstone. The thrust sheet is bent into a rollover anticline at the trailing edge, and in the uppermost part of the fiat-lying mudstone unit a mud supported shale-clast conglomerate appears. The shale-clasts are sheared, due to the overthrusting, but they can easily be recognized as having been derived from the Ordovician shale units. The conglomerate is interpreted as an olistostromic conglomerate deposited in the peeloff regime. The relationship between the Cambrian mudstone and the Silurian flysch thrust sheet demonstrates the younger-on-older fault relationship, thus indicative of a gravity gliding system.

Discussion of gravity gliding in thrust fault regions Bally (1981) stated that if thrusting and nappe formation should be interpreted as being due to

FIo. 8. Flat-lying, coherent thrust sheets in the proximal part of zone IIIb of the VNvedal thrust fault region. Lithostratigraphic units are F: Cambrian mudstones, V: lower Ordovician turbidites, K: Middle Ordovician limestone-conglomerateunit, black is the middle to upper Ordovician shale unit. Note the listric normal faults in the trailing edge of the thrust sheets. The locality is situated between longitude 34°W and 35°W on the north coast of the main E-W trending fjord, see Fig. 6.

I74

S. A. Schack Pedersen

FIG. 9. Five depositional stages in the development of the 'banana' shaped basin. In the column to the right the relation between the 'banana' basin and the tectonic evolution is illustrated, I: initial stage of weak undulation of the mudstone surface, III: intermediate stage with the first two sand bodies folded in a syncline appearing between two thrust sheets, V: final stage with the sand bodies folded in the syncline, thus forming the 'bananas'. The asymmetric syncline is formed by a combination of the tilting of the thrust sheet to the right and a drag of the flank in the foot-wall position to the left. gravity gliding, it needed the presence of extensive listric normal faulting. In the analogous example of gravity gliding from the Quaternary of the Kap Kobenhavn area, very few normal listric faults are present, possibly no more than the one at the main glide plane scar at the head of the slide. In the transverse thrust fault belt of the North Greenland Fold Belt, as well as in the Naukluft Nappe complex, no extensive normal faulting is present. The peel-off regime seems, instead, to be the important criterion of the gravity gliding system. It might be convenient to distinguish between gravity gliding and extensional tectonic systems; a distinction not made by Bally (1981). The seismic sections shown by Bally (1981) illustrate that extensional tectonic systems are confined to the early and main phase of a syn-sedimentary basin development. In these systems a dominant vertical displacement component is present. In contrast, gravity gliding systems characteristically develop during a late, mainly erosional, phase in the basin evolution, and the system is characterized by long distance lateral displacement.

Gravity spreading settings Several coastal cliffs in Denmark display imbricate thrust sheet complexes of Tertiary and Quaternary sediments, previously referred to as the dislocated cliffs [translated from Dan-

ish,'dislocerede klinter', discussed by Jessen (1931) and Gry (1940, 1941)], who interpreted the imbricate thrust sheets as having been generated in front of a prograding ice-sheet. This glacial tectonic model was mainly based on the studies of recent glacial tectonics on Spitsbergen by Gripp (1929). The most exposed and best developed imbricate fan occurs in the c. 40 m high coastal cliff section of Lonstrup Klint on the west coast of North Jutland, Denmark. The structures and sediments were described by Jessen (1931) and were interpreted as thrust sheets deformed by an ice push from the north. Today the Lonstrup Klint profile may be described as a duplex complex, where the remnant of the overthrust nappe is a thin sandy tillite. The lowermost lithological unit is an early Weichselian marine laminated clay unit. The basal decollement zone is situated in this expedient sliding unit. The glaciomarine clay grades up into a glaciolacustrine sequence of silty mudstone, interlayered with fine-grained climbing ripple laminated sand. The uppermost unit is a medium to coarse-grained sand, which shows features interpreted as indicative of syn-tectonic deposition. These features are 'banana' shaped basins (Fig. 9) in between the thrust sheets, growth faults, syn-sedimentary thrusting, erosion of the thrust fault surfaces, and diapir structures. The structural framework of Lonstrup Klint constitutes a distal part with gently dipping thrust

FIG. 9. (A-D) Photographs of the Lonstrup Klint profile. (A) Frontal anticline in the distal part of the Lonstrup Klint profile. The folded beds are fine- to medium-grained sand, and the deformation is very irregular, a combination of thrust sliding and flexural slip folding. (B) A 'banana' shaped basin in climbing ripple-laminated and cross-bedded glacio-fluvial sand. The sand is gradually infilled between two thrust sheets of mudstone one of which is seen in the left hand side of the photograph. Size of pole: 1 m. (C) Medium-size mud-diapir in the frontal nose of a thrust wedge. Two parallel thrust faults are running diagonally across the photograph from the upper right to the lower left corner. Below the thrust is a syn-tectonically deposited 'banana' basin of coarse-grained sand. (D) Contact within thrust sheet between intrusive structureless mud and thin-bedded to laminated glacio-lacustrine clay, silt and fine-grained sand. Thrust sheet in proximal part of the Lonstrup Klint profile. Intervals on the measuring pole : 20 cm. The pole is standing just in front of the intrusive mobilized mud.

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FIG. 10. The listric imbricate thrust fault fan of the Lonstrup Klint profile. The three Wulff net stereograms, lower hemisphere, show the average orientation of the thrust faults in the distal, the central, and in the proximal part of the system.

FIG. 11. Development of mudlump diapiric structures due to increasing hydraulic mud-water pressure. In stage I incipient folding of the upper mudsurface and initiation of listric fractures in the laminated and thinbedded mud appear. During stage II the thrusting propagates along the listric faults and 'banana' shaped syntectonic sand bodies are deposited between the steepening thrust sheets. In stage III the hydraulic mud-water pressure is too high to be sealed and mobilized mud penetrates the mud/sand surface and forms an injected mud diapir. In the mud layer the clay, silt and fine-grained sand are mobilized and form multiple ball and pillow structures or the lithologies are completely mixed to form a structureless mass. Vertical ruling is the hydraulic mud. In stage IV the mud diapir is dewatered and solidified, and the structure is tilted due to continuous thrusting along the steepening listric fault. In the thrust sheet to the left the diapirism has not broken through the mud/sand surface, but an intrusive contact between mobilized mud and laminated clay, silt and fine-grained sand is formed (compare with Fig. 9D).

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Studies of gravity tectonics faults (Fig. 10), bordered by a frontal anticline at the margin of the undeformed foreland (Fig. 9A). In the central part of the thrust fault complex, the thrust faults dip moderately to the north, and in the proximal part, the thrust faults are steeply dipping or vertical. North of the steeply dipping imbricate structures, the hinterland of the imbricate fan constitutes a zone of sandy tillite containing sheared out bands of mudstone and shear-folded beds of the clay-mudstone units. Diapiric structures vary in size, from waterescape structures within single beds, through multiple water-escape structures related to a sequence of beds, to diapirs affecting the entire glaciolacustrine unit with mushroom-shaped tops rising into the overlying syn-tectonic sand. Diapir-mudlump structures, where the whole clay-mudstone unit forms one large diapir, also occur. The diapirs are easily observed on a scale of up to 5 m, whilst on a larger scale, the mudlump structures are characterized by a structureless mudstone lithology, whereas in the thrust sheets the bedding is well preserved. In some of the mudlump structures a combination of thrust and mud diapir deformation exists. Here, the uppermost part of the mud-sheet is preserved, whereas the lower part consists of a structureless mud diapir which intrudes the uppermost part (Fig. 9D). This relationship is interpreted as indicating that the mud-water pressure forced the sheet upward along steep listric thrust surfaces (Fig. 11). It is still an open question whether the syntectonic sand unit had an influence on the development of the mudlump structures. However, it is unquestionable that gravity spreading deformations due to prograding sand loading do occur. The best example of such deformations are the mudlump structures in the Mississippi delta (Morgan et al. 1968). Here postglacial shelf clay deposits and prodelta clay are pressed up in mudlump structures with a diapiric rise of up to 150 m by a gravity spreading bar sand unit about 100 m thick (Morgan et al. 1968). An indication of the distance from the gravity spreading front to the rising diapir is given by a comparison of Figs 10 and 11 in Morgan et al. (1968). Here the mudlumps are rising about 1 km in front of the top of the advancing foreset slope. Gravity spreading mud-diapir structures are not normally identified as being related to fold belt terrains. However, one should expect these to occur in front of convergence zones where an accretionary prism is advancing over an ocean plate with deep-sea mudstones, as described by Westbrook & Smith (1983). They described a mud-diapir with a vertical rise of 8 km and a diameter of 2 km from the recent plate-tectonic

177

convergence zone of Barbados Ridge. The diapir occurs a few kilometres in front of the accretionary wedge with an internal framework comparable to a compressional foreland thrust fault belt. The formation of the mud-diapir in the foreland is regarded as being due to the hydrodynamic lateral pressure produced by the advancing thrusting (Westbrook & Smith 1983). However it is remarkable that the diapir structures appear in the southern part of the Barbados Ridge complex. Here a large sedimentary body is deposited in front of the Orinoco River delta. An alternative suggestion for the formation of these diapirs is that it was caused by gravity spreading and is related to the deposition of an additional sand body in the Orinoco delta.

Discussion of the gravity spreading model The gravity spreading model has been shown to apply to glacial tectonic, to syn-sedimentary deformations in deltas and to plate-tectonic convergence zones. In the development of gravity spreading structures, two features seem to be important: (i) a metastable situation with heavier sediments overlying units with a smaller specific gravity (i.e. sand on clay); (ii) a continuous advance of a wedge shaped body resulting in an increasing pore pressure in the sliding unit. These two dynamic situations will create the structures typical for the gravity spreading model (Fig. 1).

Conclusion Structures due to gravity tectonic deformations are divided into gravity gliding and gravity spreading settings. Deformations due to these gravity mechanisms are identified in small to medium scale settings and may be traced into large-scale patterns. It is not argued that gravity gliding and gravity spreading are the driving mechanisms in formation of fold and thrust fault belts, but that some syn- or epi-sedimentary tectonic settings exist, and that these settings may be exposed due to later orogenic activity. ACKNOWLEDGEMENTS: Grants from the University of

Copenhagen and from the Carlsberg Foundation are gratefully acknowledged. Data from the North Greenland Fold Belt are published by permission of the Director of the Geological Survey of Greenland. Technical assistance was provided by Pernille Andersen, Ole Bang Berthelsen and Ren6 Madsen. I direct my best thanks to the above mentioned.

I78

S. A. Schack Pedersen

References BALLY,A. W. 1981. Thoughts on the tectonics of folded belts. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 13-32. BEHR, H. J., AHRENDT, H., SCHMIDT,A. & WEBER, K. 1981. Saline horizons acting as thrust planes along the southern margin of the Damara Orogen (Namibia/SW-Africa). In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 167-72. BUCHER, W. H. 1956. The role of gravity in orogenesis. Bull. geol. Soc. Am. 67, 1295-318. CHRISTIE, R, L. & PEEL, J. S. 1977. Cambrian-Silurian stratigraphy of B~rglum Elv, Peary Land, eastern North Greenland. Rapp. Gronlands geol. Unders. 82, 1-48. COOPER, M. A. 1981. The internal geometry of nappes: criteria for models of emplacement. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 22534. DAWES, P. R. 1976. Precambrian to Tertiary of northern Greenland. In: ESCrlER, A. & WATT, W. S. (eds) Geology of Greenland, 248-303, Copenhagen. DINGLE, R. W. 1977. The anatomy of a large submarine slump on a sheared continental margin (S.E. Africa). J. geol. Soc. London, 134, 293-310. - 1980. Large allochthonous sediment masses and their role in the construction of the continental slope and rise off southwestern Africa. Marine Geol. 37, 333-54. FRIDERICHSEN, J. D., HIGGINS, A. K., HURST, J. M., PEDERSEN, S. A . S., SOPER, N . J. & SURLYK, F. 1982. Lithostratigraphic framework of the upper Proterozoic and Lower Palaeozoic deep water clastic deposits of North Greenland. Rapp. Gronlands geol. Unders. 108, 1-20. FUNDER, S., BENNIKE, O., MOGENSEN, G. S., NOENYGAARD, B., PEDERSEN, S. m. S. & PETERSEN, K. S. 1984. The Kap Kobenhavn Formation, a late Cainozoic sedimentary sequence in North Greenland. Rapp. Gronlands geol. Unders. 120, 918. GRAHAM, R. H. 1981. Gravity sliding in the Maritime Alps. ln." MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 335-52. GRIPP, K. 1929. Glaciologische und geologische Ergebnisse der Hamburgischen Spitzbergen Expedition 1927. Naturwissenschaftlichen Verin, Hamburg, B. 22. GRY, H. 1940. De istektoniske Forhold i Moleromrhdet. Med bem~erkninger om vore dislocerede Klinters dannelse og om den negative Askeserie. Medd. dansk geol. Foren. 9, 586-627. - 1941. Diskussion om vore dislocerede Klinters dannelse. Medd. dansk geol. Foren. 10, 39-42. GUTH, P. L., HODGES, K. V. & WILLEMIN, J. H. 1982. Limitations on the role of pore pressure in gravity gliding. Bull. geol. Soc. Am. 93, 606-12. H~,KANSSON, E. & PEDERSEN, S. A. S. 1982. Late

Paleozoic to Tertiary tectonic evolution of the continental margin of North Greenland. In: EMBRY, A. F. & BALKWILL, H. R. (eds) Arctic Geology and Geophysics. Mem. Can. Soc. Petrol. Geol. 8, 331-48. HOSE, R. K. & DANES, Z. F. 1973. Development of the Late Mesozoic to Early Cenozoic Structures of the Eastern Great Basin. In: DEJONG, K. A. & SCHOLTEN, R. (eds) Gravity and Tectonics, 429-42. Wiley, New York. HUBBERT, M. K. & RUBEY, W. W. 1959. Role of fluid pressure in overthrust faulting. I. Mechanics of fluid-filled porous solids and its applications to overthrust faulting. Bull. geol. Soc. Am. 70, 11566. JESSEN, A. 1931. L~nstrup Klint. Danmarks Geol. Unders., Rk. II, K~benhavn. KEHLE, R. O. 1970. Analysis of gravity gliding and orogenic translation. Bull. geol. Soc. Am. 81, 164164. KORN, H. & MARTIN, H. 1959. Gravity Tectonics in the Naukluft Mountains of South West Africa. Bull. geol. Soc. Am. 70, 1047-78. LEMOINE, M. 1973. About gravity gliding tectonics in the Western Alps. In: DEJONG, K. m. & SCHOLTEN, R. (eds) Gravity and Tectonics, 201-16. Wiley, New York. LEWIS, K. B. 1971. Slumping on a continental slope included at 1°-4 °. Sedimentology, 16, 97-110. MANDL, C. & CRANS, W. 1981. Gravitational gliding in deltas. In. MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 41-54. MARTIN, H. & PORADA, H. 1977. The intercratonic branch of the Damara Orogen in South West Africa. I. Discussion of geodynamic models. Precambrian Res. 5, 311-38. MOORE, T. C., ANDEL, T. H. VAN, BLOW, W. H. & HEATH, G. R. 1970. Large submarine slide off northeastern continental margin of Brazil. Bull. Am. Assoc. Petrol. Geol. 54, 125-8. MORGAN, J. P., COLEMAN, J. M. & GAGLIANO, S. M. 1968. Mudlumps: Diapiric structures in Mississippi delta sediments. In: BRAUNSTEIN, J. & O'BRIEN, G. D. (eds) Diapirism and Diapirs. Mem. Am. Assoc. Petrol. Geol. 8, 145-61. PEDERSEN, S. A. S. 1980. Regional geology and thrust fault tectonics in the southern part of the North Greenland Fold Belt, North Peary Land. Rapp. Gronlands geol. Unders. 99, 79-87. 1981 a. The application of computer-assisted photogrammetric methods in the structural analysis of part of the North Greenland Fold Belt. J. struct. Geol. 3, 253-64. - 1981b. Thrust fault tectonics along the Palaeozoic continental margin of North Greenland: the westernmost effect of the Caledonian orogenesis. USC abstract 130 in Terra Cognita 1, 72. -& HOLM, P. M. 1982. The significance of a Middle Devonian K/Ar age of an intrusive rock in the southern part of the North Greenland Fold Belt. Bull. geol. Soc. Denmark, 31, 121-7.

Studies of gravity tectonics PRICE, R. A. 1973. Large-scale gravitational flow of supracrustal rocks, southern Canadian Rockies. In: DEJONG, K. A. & SCHOLTEN, R. (eds) Gravity and Tectonics, 491-502. Wiley, New York. -1981. The Cordilleran foreland thrust and fold belt in the southern Canadian Rocky Mountains. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust andNappe Tectonics. Spec. Publ. geol. Soc. London, 9, 427-48. -& MOUNTJOY, E. W. 1970. Geologic structure of the Canadian Rocky Mountains between Bow and Athabasca Rivers--a progress report. Spec. Pap. Geol. Assoc. Canada, 6, 7-25. PRIOR, D. B. 1977. Coastal Mudslides Morphology and Processes on Eocene Clays in Denmark. Geografisk Tidsskr. 76, 14-33. RAMBERG, H. 1980. Diapirism and gravity collapse in the Scandinavian Caledonides. J. geol. Soc. London, 137, 261-70.

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REINECK, H-E. & SINGH, I. B. 1975. Depositional Sedimentary Environments. Springer-Verlag. SURLYK, F. & HURST, J. M. 1984. The evolution of the early Paleozoic deep-water basin of North Greenland. Bull. geol. Soc. Am. 95, 131-54. TALBOT, C. J. 1977. Inclined and asymmetric upwardmoving gravity structures. Tectonophysics, 42, 15981. WEBER, K. & AHRENDT, H. 1983. Mechanisms ofnappe emplacement at the southern margin of the Damara Orogen (Namibia). Tectonophysics, 92, 253-74. WERNICKE, B. & BURCHFIEL, B. C. 1982. Modes of extensional tectonics. J. struct. Geol. 4, 105-16. WESTBROOK, G. K. & SMITH, M. J. 1983. Long decollements and mud volcanoes: Evidence from the Barbados Ridge Complex for the role of high pore-fluid pressure in the development of an accretionary complex. Geology, 11,279-83.

STIG ASB.I~RN SCHACK PEDERSEN, Geological Institute, University of Copenhagen, Oster Voldgade 10, DK-1350 K~benhavn K, Denmark. Present address: Geological Survey of Denmark, Thoravej 31, DK-2400 K~benhavn NV, Denmark.

Slump strain in the Tertiary of Cyprus and the Spanish Pyrenees. Definition of palaeoslopes and models of soft-sediment deformation S. G. Farrell* & S. Eaton S U M M A R Y: Slumps exposed in Tertiary sediments in Cyprus and the Spanish Pyrenees vary in inferred translation distance from metres to hundreds and possibly thousands of metres. Contractional strain initiates open to tight upright to inclined folds with hinges that are sub-parallel to the strike of the palaeoslope. This strain is produced mainly by pure shear layer parallel shortening as the slump detachment propagates, but may involve a component of simple shear strain. Downslope translation of the failed unit commonly imparts a simple shear strain which tends to tighten folds and rotate fold hinges and fold axial surfaces. Fold hinges rotate towards the downslope direction, and fold axial surfaces rotate towards parallelism with the upper and lower surfaces of the slump. Coaxially refolded folds and sheath folds may also develop. Steeply plunging folds may initiate in slumps, due to the development of steeply dipping shear zones with margins sub-parallel to the downslope direction which accommodate differential downslope movements of segments of the slump. Fold axial surface attitude and fold profile may be used as a rough estimate of the degree of fold rotation.

In studies of sedimentary basins, slump fold hinge orientations are routinely used to infer palaeoslopes (Jones 1937; Woodcock 1976a; N aylor 1981). Palaeoslope estimations are usually made with the assumption that the mean fold hinge orientation is parallel to the strike of the palaeoslope (Woodcock 1976b). However, fold hinge rotation during slumping may complicate the interpretation of palaeoslope orientation (Lajoie 1972; Woodcock 1979), and with small data sets may give rise to highly inaccurate inferred palaeoslope orientations. The aim of this paper is to describe a variety of slump strain geometries from Tertiary rocks in the Spanish Pyrenees and Cyprus and through a consideration of fold genesis to infer and interpret possible rotations of fold hinge orientations during slumping.

Definition of palaeoslopes Palaeoslopes in the basins investigated have been defined by assuming that slump folds with gently plunging fold hinges initiate with a mean azimuth plunge that is sub-parallel to the palaeoslope on which failure occurred (Jones 1937; Woodcock 1979; Farrell 1984a). The direction ofpalaeoslope dip has been made assuming that folds predominantly verge downslope (Woodcock 1976a; Naylor 1981; Farrell 1984a). Basin palaeogeography *Steve Farrell died on 30 January 1987 after a brave two-year fight against cancer.

defined by palaeocurrents and sediment dispersal and distribution patterns has been used to corroborate palaeoslope estimates made using slump folds.

Geological setting The Ainsa Basin, Spanish Pyrenees Deep marine Paleocene and Eocene rocks in the south Pyrenees were deposited in the South Pyrenean basin during the Pyrenean orogeny (S6guret 1970; S6guret et al. 1983; Nijman & Nio 1975). The Ainsa basin is a compartment of the Palaeogene South Pyrenean basin and contains numerous slump horizons (Van Lunsen 1970; Mutti 1977) which are up to 120 m thick (Farrell 1984a). Slumped units in the Spanish Pyrenees, described in this paper, crop out in the N E of the Ainsa basin in the Marina and San Vicente formations (Van Lunsen 1970) (Fig. 1). In the NE Ainsa basin, slump strain has been observed in a variety of turbidite and hemipelagic lithologies. These include massive sandstones, thinly bedded sandstones and marlstones, interlaminated marlstones and siltstones and laminated marlstones. Slump strain is most easily observed in thinly interbedded (10-100mm) sandstones and marlstones. Slump units preserve evidence of various states of slope failure. Some slumps are traceable to units of undeformed sediment showing that movement on the basal

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 181-196.

181

S. G. Farrell & S. Eaton

182

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marlstone. These folds have fold limb lengths of up to 3 m and are usually tight to isoclinal, with limbs that are sub-parallel to underlying undisturbed bedding. High slump strains in these units have partially obliterated bedding and both extensional and contractional strain overprinting features are commonly developed. Fold hinge orientations are variable but are often sub-parallel to the inferred down-palaeoslope direction.

The Khalassa and Maroni Basins, Cyprus

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FIO. 1. Geological setting of the Ainsa basin showing the general stratigraphy and the position of the study area in the NE of the basin. detachment was limited (Farrell 1984a) and that the basal detachment did not cut up to the seafloor to delineate a discrete failed unit. Estimates of translation on the basal faults of these slumps can be made by graphically restoring a cross-section through the slump and is usually found to be less than 50 m (shortening is generally less than 25%). Many slumps of this type occur in the Molinos marls member and have been used to define the palaeoslope (Farrell 1984a). This inferred palaeoslope towards the SW is supported by palaeogeographical reconstructions of the basin (Nijman & Nio 1975). Slumps in other members are traceable over larger areas (up to 8 km 2) and represent major slope failure events (Farrell 1984b). Where slumps can be shown to have travelled only short distances, coherent trains of folds tend to be open to tight, with axial surfaces that dip more than 30 ° with respect to the underlying undisturbed bedding. In the larger slumps, translation distances are inferred to be one or two orders of magnitude greater than slumps in the Molinos marls member. The larger slumps contain isolated intrafolial folds of sandstone, siltstone or marlstone in a matrix of

Slumped horizons from southern Cyprus, described in this paper, occur in the Miocene sedimentary cover to the Troodos massif in the area to the east of the town of Limassol (Fig. 2). Miocene sediments have been divided into the Pakhna and Kalavasos formations (Fig. 2), and consist of marine chalks, marls, calcarenites, polymict conglomerates and gypsum beds (Bagnall 1960; Pantazis 1967). Miocene sedimentation was strongly influenced by the uplift of the Troodos ophiolite culminating in late Miocene times with the southward overthrusting of the Limassol Forest igneous block on to the sedimentary cover of the ophiolite (Robertson 1977). Southward thrusting generated the Yerasa fold and thrust belt and was accompanied by the development of the Khalassa and Maroni basins to the south and east of the Limassol Forest block. Slumps were examined at Amathus in the Khalassa basin and at Petrounda Point, Ayios Theodhoros and Tokhni Quarry in the Maroni basin (Fig. 2). A variety of fold styles and intensities of deformation and bedding disaggregation have been observed in Cyprus. Fold styles are similar to those observed in slumps in the Ainsa basin. The slumped unit at Amathus forms part of the Koronia Limestone (Pakhna Formation) and consists of channelized debris flows and graded bioclastic calciturbidites. The slump unit investigated (Fig. 3A) is 0.8-1.0 m thick and consists of thinly interbedded medium to coarse carbonate sands and silts. Slump folds are generally asymmetrical with axial surfaces that have a mean dip of 32° with respect to the underlying bedding. Rare recumbent sheath folds are also present (Figs 7B and 8A). The Lefkara Formation which crops out at Petrounda Point is Miocene in age and consists of chalks, marls, foraminiferal sands and lignitic marls which have been incorporated into a slumped unit with an exposed thickness in excess of 12 m. Slumps with similar deformation styles and in a similar stratigraphic position (Bagnall 1960) occur at localities up to 35 km apart, suggesting extensive development of major slope

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] Formation

Pre Middle M i o c e n e s e d i m e n t s Igneous basement

CAPE GATA

FIG. 2. Geological setting and stratigraphy of the Khalassa and Maroni basins, Cyprus. failure units. A variety of fold geometries is developed (Table 1) in the Petrounda slumps, with early thin slumps and debris flow deposits folded by later open to tight, upright to inclined, large scale folds which have fold limb lengths of up to 8 m. The early thinner slumps contain

isolated recumbent folds with curvilinear hinges that may take the form of recumbent sheath folds (Figs 7A and 8D). Another slumped unit in the Pakhna Formation crops out 750 m south of Ayios Theodhoros. This 5 m thick unit is composed of medium

Dm~

q~

c~o

Slump strain in the Spanish Pyrenees

185

FIG. 3. (A) Asymmetrical slump folds at Amathus in the Pakhna basin (Cyprus). This slump has a basal failure and extensional strain of gently dipping fold limbs. (B) Isoclinal recumbent slump folds exposed at Ayios Theodhoros, Maroni basin, Cyprus. (C) Upright symmetrical slump folds developed in gypsum at Tokhni Quarry in the Maroni basin, Cyprus. (D) Upright symmetrical folds, Molinos marls, Ainsa basin, Spain.

S. G. Farrell & S. Eaton

186 TABLE 1.

Summary of slump datafrom the Pakhna, Ayios Theoderos and NE Ainsa basins Cyprus

Ainsa basin

Spanish Pyrenees

r

,f

Thickness of

Up to

Slump units

3m

Upright to Inclined Early Folds

Up to 25m

1.0m

over 12m

1,5m

5m

"~

Recumbent Isoclinal Folds

~/

Sheath Folds

/

Refolded Folds Late Upright Folds

,/

bedded foraminiferal chalks with thin marly interbeds. Asymmetrical folds have gently dipping axial surfaces and curvilinear hinges (Figs 3B, 8C and 9D). Slump faults are common and cut many folds so they occur as isolated recumbent intrafolial folds. Slumped gypsum deposits at Tokhni Quarry (Figs 3C and 8B) are part of the Kalavasos Formation which has been correlated with Messinian evaporites elsewhere in the Mediterranean (Gass & Cockbain 1961; Baroz & Bizon 197a; Baroz & Bizon 1977; Pantazis 1978). The 1.21.5 m buckle folded slump sheet exposed at this quarry can be correlated with exposures up to 5 km distant. Fold styles in Cyprus and the Ainsa basin Coherent trains of symmetrical and asymmetrical folds occur in slumps where slump strain and the degree of bedding disaggregation is low to moderate. Near symmetrical folds with steep axial planes have been observed in both Cyprus and the Spanish Pyrenees (Figs 3C, D). Fold hinges plunge up to 26 ° and axial surfaces are approximately planar. These folds have a predominantly parallel geometry (class 1B of Ramsay 1967) with some flowage of sediment into the cores of the folds. Asymmetrical folds are also common in both the Ainsa slumps and the Miocene slumps on Cyprus. The steep limbs of folds may be overturned and fault splays from the slump basal detachment commonly terminate as tip lines (in the sense of Elliott 1976) in the cores of these folds (Fig. 4A). Some asymmetrical folds have extensionally strained shallow limbs.

Fig 4(B) from the Ainsa basin shows boudinage of the gently inclined limb of an asymmetrical fold with boudin axes that are sub-parallel to the slump fold hinge. Some gently dipping fold limbs at Amathus are cut by normal micro-faults. The steep limbs of folds may be reverse faulted, either by micro-faults or by larger listric reverse faults which root into the basal detachment. Fold style is generally parallel (class 1B of Ramsay 1967) although some folds may be similar in style (class 2). Folds which approach a similar geometry are generally tight and recumbent and often have attenuated lower limbs (Figs 5A, 6A). There is a general correlation of fold interlimb angle and the attitude of fold axial surfaces for folds with gentle hinge plunge azimuths (less than 30°). Tighter folds tend towards a recumbent geometry which bears out observations made by N aylor (1981). Folds with steeply plunging hinge azimuths (30-90 °) show no good correlation of fold tightness and dip of axial surfaces. Refolded folds Recumbent folds and refolded folds generally occur in slumps with a moderate to high degree of bedding obliteration. This is commonly the case in thicker, more aerially extensive slumps (Table 1). Refolded folds observed in slumps in the Ainsa basin are commonly either: (i) recumbent, tight to isoclinal early folds with fiat lying axial surfaces folded by upright, generally open, later folds with sub-vertical axial surfaces (Fig. 5B), or (ii) plane, non-cylindrical, approximately coaxial refolded folds with flat lying axial surfaces (Figs 5C, 6C and 7D). Both of these refolded fold

Slump strain in the Spanish Pyrenees

~

I87

'3 (B)

F[o. 4. (A) Asymmetrical slump folds developed in interbedded sandstones and marlstones, NE Ainsa basin (Spanish Pyrenees). The basal fault to this slump terminates as tip lines in the cores of asymmetrical folds. (B) Asymmetrical slump fold, NE Ainsa basin (Spanish Pyrenees). The gently dipping limb of the fold is extensionally strained (boudinage). styles have been observed in the Ainsa basin slumps and at Petrounda Point in Cyprus (Fig. 7D). They correspond to type two and type three refolded folds, defined by Ramsay (1967), and are similar to refolded slump folds described in California by Tobisch (1984). Fold hinge curvature is commonly observed in tight to isoclinal recumbent folds and varies from gentle undulations of the fold hinge to isoclinally bent hinges (sheath folds) (Figs 6B and 7A-C). Folds with strongly curvilinear hinges have hinge azimuth orientations that are consistently subparallel to the inferred down-palaeoslope direction (Fig. 8D, F). Plots of axial surface dip against fold hinge azimuth orientation, indicate that rotation of fold hinge azimuths towards parallelism with the gross transport direction often results in a tendency for folds to become recumbent (Fig. 9A, C, D, E). However, some slumps in the Ainsa basin contain folds with a variety of fold hinge azimuth orientations which show no correlation between fold hinge azimuth orientation and axial surface dip (Fig. 9B). Slumps with steeply plunging fold hinges also show no correlation of fold orientation and dip of axial surfaces.

Interpretation Fold development The presence of boudinage in some slumps (e.g. Fig. 4B) indicates that competency contrasts

existed between layers during slumping. Where there is a competency contrast between layers undergoing deformation, strain is accommodated by buckling (Ramsay 1967) and suggests that slump folds developed actively by buckling rather than passively by shear or flow folding. The comparative rarity of class 2 (similar folds) in slumps also indicates that during slumping folds develop principally by buckling (Woodcock 1976b). However, the occurrence of some similar folds indicates that although buckling may initiate folding, homogeneous or heterogeneous simple shear may modify some fold geometries by flow or shear folding during fold growth and rotation (Woodcock 1976b). However, the occurrence of some similar folds indicates that although buckling may initiate folding, homogeneous or heterogeneous simple shear may modify some fold geometries by flow or shear folding during fold growth and rotation (Woodcock 1976b). Upright, near symmetrical folds probably developed as buckle folds reflecting predominantly pure shear deformation during longitudinal shortening of the failed unit. These upright folds indicate that the x-y plane of the finite strain ellipsoid for the slump deformation was perpendicular or nearly perpendicular to bedding, and involved layer parallel shortening. Although asymmetrical folds can develop during either a bulk pure shear or simple shear deformation, asymmetrical folds developed during slumping are often a manifestation of tip strain associated

188

S. G. Farrell & S. Eaton

FIG. 5. (A) Similar style recumbent isoclinal slump fold, NE Ainsa basin (Spanish Pyrenees). (B) Refolded fold. An early tight to isoclinal recumbent fold with a gently dipping axial surface refolded by a later open upright fold with a steeply dipping axial surface. Upper slumps, Ainsa Basin, Spain. (C) A plane non-cylindrical approximately coaxial refolded fold. Lower slumps, Ainsa basin, Spain.

S l u m p strain in the Spanish Pyrenees

~

~

~

I~__~--.-~--~-~~ ~ I~,~ ~ < - ~ ' ~

18 9

AJ

B

~ ~ \1 ~ - ~ 1

,o

l

~

Curvilinear hinge

o

o;,rn

Fold hinge

l Curvilinear hinge

~ 0i

~.1

late axial trace

0i2m

FIG. 6. (A) Line drawing of Fig. 5(A); (B) Line drawing of Fig. 7(C). (C) Line drawing of Fig. 5(C). (D) Line drawing of Fig. 7(D).

with reverse fault imbricate fans (Farrell 1984a) produced during pure shear layer parallel shortening. Asymmetrical tectonic folds with long limbs showing extensional strain, and short limbs showing contractional strain, have been described by Ramsay et al. (1983) (Fig. 10) and interpreted as having developed principally by buckling during simple shear deformation. During simple shear deformation, material lines, such as fold limbs, lying in the contractional field of the incremental strain ellipsoid, will shorten. Material lines initially lying in, or subsequently rotated into, the extensional field of the incremental strain ellipsoid will extend (Ramsay 1967). The angular relationship of the shallow limbs of asymmetrical folds with the margins of the shear zone in which they are developed is such that they commonly lie in the extensional field of the incremental strain ellipsoid (Ramsay et al. 1983). Although the steep limbs of asymmetrical folds commonly begin to develop in the contractional field of the incremental strain ellipsoid, during progressive simple shear deformation and active buckling they may become rotated into the extensional field of the finite strain ellipsoid (Ramsay et al. 1983). Slump folds which show extensionally and/or contractionally strained fold limbs that are not the product of a separate overprinting deformation can be interpreted as developing during simple shear deformation (Fig.

4B). If slump fold limbs are extensionally strained the limbs may extend by normal faulting, boudinage or ductile thinning, depending on the rheologies and competency contrasts of the sediments that are slumping. Ductile thinning and attenuation of the lower limbs of folds may produce folds with a similar style (Figs 5A and 6A). The development of sheath folds has been linked to the accentuation of curvature of initially curved fold hinges during simple shear deformation (Escher & Waterson 1973; Ramsay 1980; Ghosh & Sengupta 1984). During simple shear deformation, linear elements such as fold hinges are progressively rotated towards the x-direction of the finite strain ellipsoid (Escher & Waterson 1974; Williams 1978; Ramsay 1980)and planar surfaces such as axial surfaces are progressively rotated towards the x-y plane of the finite strain ellipsoid. Plane coaxial non-cylindrical refolded folds (see Figs 5C and 7D) (type three of Ramsay 1967) have also been interpreted as having developed during progressive simple shear deformation (Ghosh & Sengupta 1984). From these observations of slumps it is suggested that slump folds initiate as open to tight parallel style folds with axial surfaces that dip at 30-90 ° with respect to the slump basal detachment. Fold hinges initiate sub-parallel to the palaeoslope (Jones 1937), although some initial curvature to the fold hinges may be generated as

FIG. 7. (A) Sheath fold, Petrounda Point, Maroni basin, Cyprus. (B) Sheath fold, Amathus, Khalassa basin, Cyprus. (C) Isoclinal fold refolded by a sheath fold, Lower slumps, Ainsa basin, Spain. (D) An early tight fold refolded by a later tight recumbent fold. Petrounda Point, Maroni basin, Cyprus.

t..,¢.

o~

e~

o

Slump strain in the Spanish Pyrenees

191

(A)

(B) N

N AMATHUS

~

n =

/

Fola

TOKHNIQUARRY

37

~

n

=

33

~ ~ MeanFold hinge th

\x=_

\

/

,n,~r~d

MeanFold hinge Azimuth /

Inferred

~ ~l~osl~e

N

AYIOSTHEODHORUS CYPRUS ~

]

L

~

•Mean Fold hinge / Azimuth

--1--

s

l

o

.

p

I

tt-,~ ~/

PETROUNDA POINT CYPRUS

~

N

n = 19

(D)

Inf . . . . d

e

~

~

Inferred ~Palaeoslope

(r)

(E) N

MOLINOSMARLS AINS

Palaeoslope

n =

159

uPPER SLUMPS

N

n=50

AI

~

~, Mean Fold hinge Azimuth

~ - Inferred ~ . Palaeoslope ~

+. "~ . X MeanFold hinge .ip+lt".+~ ~ Azimuth

FIG. 8. Equal area stereoplots of fold hinge azimuths from slumps in the N.E. Ainsa basin (Spanish Pyrenees) and the Khalassa and Maroni basins (Cyprus).

I9Z

S. G. Farrell & S. Eaton UPPER SLUMPS

MEMBER

WEST

(B)

UPPER SLUMPS MEMBER EAST

(m) 9o

n

n=38

[]

D°

o

[]

[]

60-

[]

[] Cl

[]

[]

t

I 300

250

O

I 350

°

40-

~)

20-

[]

D

[]

%

[]

palaeoslope

!

200

[] []

[]

~9

i

[3

[]

o

o

[]

[]

[] on

[] D

[] []

~I

I 090

040

t

[]

palaeoslope

palaeoslope I 140

6 o

i 190

2o0

t

I

i

250

300

I 040

! 35O

(C

AMATHUS

(D)

AYIOS THEODHOROS

90

n=31

n=37

B0--

60-

o

2O-

[]

0

) ,¢

80-

::) (/)

60-

[] % % o•

p ......

o

[

40-

c [] [] o o OO

[]

[] [] [] D

,lope t 50

[] c3

ooo

o

m~

[] []

O

D..

•D

20-Dn

[] ~ ° 0 palaeoslope

palaeoslope

D:o

100

090

FOLD HINGE AZIMUTH

FOLD HINGE AZIMUTH

40-

18

80-

03 )

=

i

!

i

150

200

250

t

!

"0

0

300

[]

'

1o0

[]

o O

palaeoslope

+

I

|

i

I

150

200

250

300

350

FOLD HINGE AZIMUTH

FOLD HINGE AZIMUTH

PETROUNDA

(E

POINT

n

[]

8o-

=

19

6oo ,~ x

o

[]

D

40tl

o

u o

20-

o

[]

palaeoslope

i:

palaeoslope

Cl o

0 5

I 100

I 150

I 200

FOLD HINGE

I 250

I 300

AZIMUTH

FIG. 9. Plots of dip of slump fold axial surface and fold hinge azimuth. Plots (A) and (B) are from the Upper slumps, NE Ainsa basin (Spanish Pyrenees); plot (C) is from Amathus, Khallasa basin (Cyprus); plot (D) is from Ayios Theodhoros, Maroni basin (Cyprus) and plot (E) from Petrounda Point, Maroni basin (Cyprus). Plots (A), (C), (D) and (E) show a decrease in dip of fold axial surface as the fold hinge azimuth tends towards the downpalaeoslope direction. Plot (B) shows no decrease in dip of fold axial surface with change of fold hinge azimuth. a result of the radial propagation of the slump basal failure (Farrell 1984a). During internal deformation of a slump, folds are progressively rotated so that axial surfaces become sub-parallel to the slump basal detachment and folds are tightened so that they may become isoclinal. Concomitant with the tendency for folds to be rotated towards a recumbent geometry, fold hinges tend to be rotated towards the downslope direction (gross transport direction). Curvilinear hinges may be accentuated into sheath folds. Refolding of folds producing up-

right, open late folds suggests that final minor strain increments may be due to a pure shear deformation with the slump thickening and shortening. Sediment compaction imparts a flattening strain on the slump. This strain will tend to rotate fold axial surfaces further towards parallelism with undisturbed bedding and may modify parallel recumbent folds so that they have similar geometries (Fig. 11). This proposed scheme for the genesis of slump folds can be interpreted as fold initiation during layer parallel shortening, and pure shear deformation followed by strain which is predominantly

Slump strain in the Spanish Pyrenees

193

Contractional Strain ! NORMAL FAULTS

Extensiona! Strain

(A)

II

I

/

III /, FAULT TqP

r ,,,tl

,,

II

INITIAL FAILURE (FOOTWALL) FAULT TIP

(8) •

FAULTS / TIAL FAILURE (HANGING WALL)

\ FIG. 10. Strain in the shallow limb (normal faults) and

steep limb of an asymmetrical fold developed in a shear zone (after Ramsay et al. 1983). due to simple shear deformation. In slumps with basal detachments that did not cut up to the seafloor (see Fig. 4A), contractional slump strain is predominantly in the form of folds and faults, indicating that deformation principally involved pure shear layer parallel shortening of the failed unit. This strain develops as a consequence of the propagation of the slump basal detachment (Farrell 1984a). Slip on the basal fault is accommodated as internal strain in the hanging wall to the detachment (Fig. 12A) and will take the form of a predominantly pure shear strain with a subvertical x-y plane to the finite strain ellipsoid. However, simple shear strains may also be developed during propagation of the basal detachment and will be superimposed on the pure shear strains (Fig. 12B). It is possible that some slumps initiate as zones of slope parallel simple shear, but unless a basal detachment develops, translation of the slumped unit will be limited.

COMPACTION

BEDDING

/

• : :

:

:: : ( : 5 .

SIMILAR STYLE FOLD

t FIG. 1I. Change of fold profile from a parallel to similar style by the superposition of a compactional flattening strain on a recumbent slump fold.

FAULT TIP

\\

\

INITIAL FAILURE (FOOTWALL) FAULT TIP

FIG. 12. (A) Contractional pure shear strain developed downslope of the initial slope failure as a consequence of the propagation of the basal slump detachment. Movement on the fault is taken up as strain in the hanging wall to the detachment. (B) Propagation of a basal slump detachment with a superimposed component of simple shear strain.

The development of extensional slump strain upslope of an initial slope failure and contractional strain downslope of the failure in many slumps and Holocene to recent submarine slope failures (Farrell 1984a), indicates that development of a basal fault and associated pure shear strain is a common mode of slope failure. It is likely that much of the variation in observed slump strain is due to the development of a spectrum of failure styles dependent on the mechanical properties of the sediment and the relative importance of the components of pure and simple shear. During translation of the failed unit, thickening or thinning (pure shear strain with a sub-vertical x - y plane) of the slump will occur whenever there are sequential up or downslope propagating velocity changes. These sequential velocity changes will be expected to develop during translation over changes in slope and when friction increases on the basal detachment by localised dewatering of the slump (Fig. 13) (Farrell 1984a). Final strain increments may be in the form of open buckle folds if the toe of the slump comes to rest before the rest of the failed unit. Where slumps are inferred to be far travelled, the common occurrence of fold rotation and coaxial refolding shows that simple shear is often an important consequence of downslope translation. Simple shear strains are likely to occur in a slumped body as a result of dewatering of the slump and loss of fluid from the basal fault. The shear strength of a sediment body is a function of pore-fluid pressure; high pore-fluid pressures are often considered to lead to a loss of shear strength

I94

S. G. Farrell & S. Eaton decrease in axial surface dip may occur in such steeply inclined shear zones. The strain ellipsoid for steep shear zones will have a sub-vertical x - y plane and folds initiated in such steep shear zones will tend to be steeply plunging. Fold rotation in such steep shear zones will tend to steepen the dip of fold axial surfaces.

,,

i

r;:

ace

FIG. 13. Pure shear strains develop in slumps during sequential velocity changes. Deceleration on the front of a slump generates contractional strains and may lead to the development of late upright folds which refold earlier structures.

and subsequent failure of the unit (Mandl & Crans 1981). Deceleration of a slump by recovery of shear strength may be caused by loss of excess pore fluid pressure. Recovery of shear strength will eventually result in significant friction on the detachment, and momentum and gravitational potential of the failed body may be expended in internally deforming the slump. If general dewatering occurs by upwards movement of water through the slump, the shear strength of the base of the slump may be recovered while the upper layers of the slump are still moving. This differential movement will generate simple shear strains. The movement of salt (Talbot 1981) and ice (Nye 1953) glaciers over obstructions in the glacier floor generates simple shear strains in the glacier and may be analogous to the movement of slumps over an irregular seafloor topography. The first irregularity encountered on the sea floor by the slump is the frontal toe scar to the slope failure (Schwartz 1982). Strain imparted to the slump as it moves over this scar may be analogous to the simple shear strains imparted to thrust sheets as they move over frontal thrust ramps (Fischer & Coward 1982). Gravitational collapse of unlithified sediments during translation may generate significant strains in the failed unit. Ramberg (1981) has suggested that gravitational collapse of long thin bodies will generate simple shear strains as the upper parts of the collapsing unit moves further than the lower parts. If differential downslope movements occur in segments of the slump then strike-slip and oblique faults and steeply dipping shear zones may develop. Differential downslope velocities are likely to occur because of local porefluid pressure related variations in friction on the slump basal fault. The development of steeply plunging folds and fold rotation without a

Definition of palaeoslopes The initial development of slump folds is related to the shortening of the slump as the basal fault propagates. This produces a strain ellipsoid with an x - y plane that is sub-parallel to the strike of the palaeoslope, and therefore slump folds initiate with a mean fold hinge azimuth sub-parallel to the strike of the palaeoslope. Therefore, small data sets of upright symmetrical or inclined asymmetrical folds may be used with more confidence to define a palaeoslope strike than data sets composed of recumbent isoclinal folds that may have been rotated during translation of the slump. Fold rotation may result in a wide spread of hinge azimuths, and therefore a palaeoslope defined from a small data set of recumbent isoclinal folds may be highly inaccurate. Large data sets recorded from many slump units eliminate many of the problems of having relatively unclustered fold hinge azimuth data. If it is not possible to measure a large data set, a rule of thumb estimate of the possibility of a particular fold having been rotated may be made by considering the fold shape. Recumbent isoclinal folds are likely to have rotated significantly in comparison with upright symmetrical folds. Unfortunately, it is not possible to equate the observed strain in a slump directly with the amount of translation that the slump has undergone. Translation on the basal detachment of a slump may not generate slump strain and relatively unstrained slumps may be translated for considerable distances on the sea bed. However, slumps which can be demonstrated to have been translated only small distances commonly have much less internal deformation than slumps which are inferred, from their size and lateral continuity, to be far travelled.

Conclusions Slumps in the Khalassa and Maroni basins in Cyprus, and the Ainsa basin in the Spanish Pyrenees, contain folds with a wide variety of fold styles. Fold style can be related to the inferred translation distance of a slump. Translation distance has been inferred by consideration of slump size, the degree of bedding disaggregation

Slump strain in the Spanish Pyrenees and whether a slump is transitional into unslumped sediment. Upright and inclined folds with fold hinges sub-parallel to the palaeoslope are frequently observed in slumps where translation distances are small. Recumbent isoclinal folds, sheath folds and refolded folds are common in slumps where translation distances are large. Fold hinge rotation towards a down-palaeoslope trend is also common in far travelled slumps. Slump strain can develop during the initiation, translation and termination phases of slump development. Contractional slump strain, initiated during the propagation of the basal failure to a slump, is predominantly buckle folding related to the longitudinal shortening of the slump. Folds produced during this initiation phase of slump development are generally upright to inclined and open to tight. During translation of the slump, strain is principally in the form of simple shear which may rotate and tighten folds and lead to the development of sheath folds and

195

refolded folds. Deceleration of a slump may impart a final pure shear strain to the slump producing upright open to tight folds. Fold rotation is, therefore, a common consequence of slump translation and small sets of fold orientation data measured from far travelled slumps may indicate spurious palaeoslope orientations. Thus the internal deformation of a slump should be taken into account when deciding the degree of confidence with which palaeoslope orientation estimates are made. ACKNOWLEDGEMENTS:S. Farrell acknowledges British

Petroleum for permission to publish this paper. Diagrams were drafted by B.P. special reprographics section. S. Farrell was supported by a Shell Postgraduate Scholarship at University College Cardiff. S. Eaton was supported by a NERC studentship under the supervision of A. H. F. Robertson. He also thanks the Cyprus Geological Survey Department, particularly C. Xenophontos and A. Panayiotou for their help in Cyprus.

References BAGNALL,P. S. 1960. The geology and mineral resources of the Pano Lefkara-Larnaca area. Cyprus Geol. Survey Dept. Mem. 5, 116 pp. BAROZ,F. & BIZON, G. 1974. Le Neogene de la Chaine du Pentadaktylos et de la partie nord de la Mesaorie (Chypre). Etude stratigraphique et micropaleontologique. Inst. Francais Petrole, Rev. 29, 327-59. -& BIZON, G. 1977. La couverture Tertiaire du flanc nord du massif du Troodos et de la partie meridionale de la Mesaoria. Etude stratigraphique et micropaleontologique. Inst. Francais Petrole. Rev. 32, 719-59. ELLIOTT,D. 1976. The energy balance and deformation mechanisms of thrust sheets. Philos. trans. Roy. Soc. A283, 289-312. ESCHER, A. & WATERSON,J. 1974. Stretching fabrics, folds and crustal shortening. Tectonophysics, 22, 223-31. FARRELL, S. G. 1984a. A dislocation model applied to slump structures. Ainsa basin, South Central Pyrenees. J. struct. Geol. 6, 727-36. -1984b. Slope processes and tectonism in Eocene marine sediments of the Ainsa basin, Spanish Pyrenees. Unpubl. Ph.D. thesis, Univ of Wales. FISCHER, M. W. & COWARD, M. P. 1982. Strains and folds within thrust sheets. An analysis of the Hellam thrust sheet, north western Scotland. Tectonophysics, 88, 291-312. GASS, I. G. & COCKBAIN, A. E. 1961. Notes on the occurrence of gypsum in Cyprus. Overseas Geol. Min. resources, VIII. 269-87. GHOSH, S. K. & SENGUPTA, S. 1984. Successive development of plane noncylindrical folds in progressive deformation. J. struct. Geol. 6, 703-9. JONES, O. T. 1937. On sliding and slumping of submarine sediments in Denbighshire, North

Wales, during the Ludlow period. Q. J. geol. soc. Lond. 93, 241-83. LAJOIE, J. 1972. Slump fold axis orientations, and indication of palaeoslope. J. sed. Pet. 584-6. LUNSEN VAN, H. A. 1970. Geology of the Ara-Cinca region, Spanish Pyrenees, province of Huesca. Geol. Ultraectina, 1-119. MANDL, G. & CRANS, W. 1981. Gravitational gliding in deltas. In : MCCLAY,K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 41-54. MUTTI, E. 1977. Distinctive thin bedded turbidite facies and related depositional environments in the Eocene Hecho Group (South Central Pyrenees). Sedimentology, 24, 107-31. NAVLOR, M. A. 1981. Debris flow (olistostromes) and slumping on a distal continental margin; the Palombini limestone-shale sequence of the northern Apennines. Sedimentology, 28, 837-52. NIJMAN, W. J. • NIO, S. D. 1975. The Eocene Montanana delta (Tremp-Graus basin, province of Lerida and Huesca, South Pyrenees, N. Spain). In." ROSELL, J. & PUIGDEFABREGAS, C. (eds) Sedimentary Evolution of the Palaeogene South Pyrenean Basin, 9th congress Sedim, Part B, 1-9, Nice, 1975, Field trip 19. NVE, J. R. 1953. The motion of ice sheets and glaciers. J. Glaciol. 3, 493-507. PANTAZlS, T. M. 1967. The geology and mineral resources of the Pharmakas-Kalavasos area. Cyprus Geol. Survey Dept. Mere. 8, 190 pp. 1978. Cyprus evaporites. In: Ross, D. A., NEPROCHNOV, Y. P., et al, 1978. Initial reports of the Deep Sea Drilling Project, 42, part 2, 118594 pp, Washington ( U.S. Government Printing Offlee).

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S. G. Farrell & S. Eaton

RAMBERG, H. 1981. The role of gravity in orogenic belts. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 125-40. RAMSAY, J. G. 1967. Folding and Fracturing o f Rocks. McGraw-Hill, New York. - 1980. Shear zone geometry: a review. J. struct. Geol. 2, 83-99. - CASEY, M. & KLIGFIELD, R. 1983. Role of shear in development of the Helvetic fold, thrust belt of Switzerland. Geology, 11,439-42. ROBERTSON, A. H. F. 1977. Tertiary uplift history of the Troodos massif, Cyprus. Geol. Soc. Am. Bull. 88, 1763-72. SCHWARTZ, H. O. 1982. Subaqueous slope failures, experiments and modern occurrences. Contrib. Sediment. 11, 116 pp. SEGURET, i . 1970. Etude tectonique des nappes et series decollees de la partie centrale du versant sud des Pyrenees, Publ Ustella, Montpellier. Ser. Geol. Struct. 21, 1-155. --, LABAUME, P. & MADARIAGA, R. 1983. Eocene

seismicity in the Pyrenees from megaturbidites in the South Pyrenean basin (N. Spain). Mar. Geol. 55, 117-31. TALBOT, C. J. 1981. Sliding and other deformation mechanisms in a glacier of salt, S. Iran. In: MCCLAY, K. R. & PRICE, N. J. (eds) Thrust and Nappe Tectonics. Spec. Publ. geol. Soc. London, 9, 173-83. TOBISCH, O. T. 1984. The development of foliation and fold interference patterns produced by sedimentary processes. Geology, 12, 441-4. WILLIAMS, G. O. 1978. Rotation of contemporary folds into the X-direction during overthrusting, Processes in Lakesfjord, Finnmark. Tectonophysics, 48, 29-40. WOODCOCK, N. H. 1976a. Ludlow series slumps and turbidites of the Montgomery trough, Powys, Wales. Proc. Geol. Ass. 87, 169-82. -1976b. Structural style in slump sheets, Ludlow series, Powys, Wales. J. geol. Soc. Lond. 132, 399415. -1979. The use of slump structures as palaeoslope orientation estimators. Sedimentology, 26, 83-99.

S. G. FARRELL,BP Petroleum Development Ltd, Farburn Industrial Estate, Dyce, Aberdeen, Scotland. S. EATON, Grant Institute of Geology, The University of Edinburgh, Edinburgh EH9 3JW, Scotland.

Mass transfer in unmetamorphosed carbonates and during lowgrade metamorphism of arenites P. M. Clifford, M. C. Rice, L. L. Pryer & F. Fueten S U M M A R Y: Study of stylolites in Silurian carbonates of southern Ontario, and of spaced cleavage in Lower Palaeozoic meta-arenites of Nova Scotia, shows that diffusional or infiltrational mass transfer has been the dominant mechanism for dimensional change. The carbonates show a quite specific pattern of stylolite growth in, and normal to, the mean plane of the stylolite, and there is always a low-porosity halo about each stylolite. Implied shortening usually exceeds the amplitude of any teeth in stylolites and may reach 5% of the 'local' sequence. In the meta-arenites, increased inequancy, decreased long-axis fluctuation of quartz grains, lithon to cleavage, coupled with virtually constant grain dimension parallel to cleavage, suggest gross loss of volume. Modal and chemical analyses suggest that an average cleavage zone represents a dilation of - 50%, expressed as shortening normal to cleavage. Bulk loss is

about 10%, almost all SiO2. In both cases, mass removal begins at zero or low levels of overall deformation. For the carbonates, total deformation is near zero; the process shuts down by self-suffocation. In the meta-arenites the process was probably partly driven by the conditions associated with greenschist-level metamorphism. In both cases, collapse across the surfaces of dissolution is nicely adjusted to keep the gross shortening approximately homogeneous.

Mass transfer (see Durney 1976) has yielded spaced solution zones (SSZ) in Silurian dolostones of southern Ontario, and spaced cleavage zones (SCZ) in meta-arenites of the Lower Palaeozoic Meguma Series of Nova Scotia. In our studies of these and other examples of rocks affected by mass transfer, we have inferred certain sequences of development which are basically the same in all cases, even where, as far as our two specific cases are concerned, the conditions are very different. Below we briefly outline some interim results, and comment on a variety of implications and consequences concerned with reduction of the rock volume, and motion of material.

Silurian dolostones The carbonate units of the Niagara Escarpment of Southern Ontario, interpreted as having been deposited in a shallow, warm marine environment, dominate the Silurian portion of the stratigraphic sequence. They are totally dolomitized. The Silurian Lockport Formation at Dundas, Ontario, consists of some 6 m of somewhat irregularly layered dolostone. The current texture is finely crystalline rather than bioclastic, though locally there is abundant fossil material, mostly broken up, and isolated fossils, now beautifully preserved by dolomite or pyrite. Maximum depth

of past stratigraphic cover is difficult to estimate, but may have reached 1000 m; all of the cover has been removed at Dundas, but up to 600 m of it remains in south-western Ontario. Pervading this Lockport dolostone is a multitude of spaced solution zones (SSZ). There is a variety of local geometries, but we perceive these to be linked in an evolutionary sequence. A schematic diagram of a typical spaced solution zone (Fig. 1) shows that: 1 At the lateral tips of the SSZ (a, Fig. 1) very thin films of insoluble material occur at almost all grain boundaries, but especially at those boundaries which are roughly bedding-parallel. These films are usually only one grain long; thus there is no extended film in any direction, but there is a great deal of film pervading a considerable volume of rock. 2 Proceeding laterally towards the central portion of the SSZ, the single-grain films become visibly connected over a great many grains, and acquire an irregular profile, sometimes dentate, sometimes undulate (b, Fig. 1). The films are also thicker, and become closer together. Clearly grains between films are losing mass and volume and the rock is undergoing overall collapse normal to bedding. The dentate films are irregular in thickness, maximum thickness occurring at the crowns of the teeth, there being little or no material along the sides. Some of the sides are

From: JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformationof Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 197-209.

I97

P. M. Clifford et al.

19 8

.... I-

2h

,-

,--,~.,

b

.

c

~, ,,

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~

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.____....~/--~-,-

t, C

~-~

~

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FIG. 1. Schematic diagram of the main elements of SSZ. (a) 'tip' zone; (b) microstylolite formed by amalgamation; (c'): stylolite; (c"): seam; inset: overall form, with wavelength (W) and amplitude (h) shown. t = active thickness, dashed zone = tip zone. grooved, suggesting differential motion of teeth as material is dissolved at the corners. These have been termed microstylolites (see Wanless, 1979). 3 From the microstylolites, two forms of central or mature SSZ occur (c', c", Fig. 1). One type is the classic stylolite, large-toothed, with substantial amounts of insoluble material capping the teeth and locally smeared down the sides. Commonly, one stylolite dominates, cutting off, and incorporating elements of, other stylolites (Fig. 2), while persisting in its highly dentate form. There must have been an accompanying solution of intervening material. The other type is a smoothly undulating to nearly rectiplanar zone of insoluble material, produced by a continuation of the removal of material between SSZ until two (or more) such SSZ amalgamate to form a thicker accumulation of insoluble residues (Fig. 2). We see this as an evolutionary sequence. The 'tips' are where rock solution and accumulation of insoluble material are inaugurated. The zone of removal of rock advances laterally into unaffected rock, and involves substantial local thicknesses at scales of the order of a centimetre. Continued collapse by removal of material from between SSZ leads to amalgamation of SSZ, which is a 'mature' stage of solution, occurring at the centre of a zone while the 'juvenile' activity proceeds at the tips.

The end result is the production of SSZ whose overall geometry is discoidal. In this case, the XY plane of the discoidal SSZ is bedding parallel, and X / Y is approximately one. A small discoidal SSZ may have a diameter of 10 cm, a mature seam thickness of 1 mm and a tip up to 3 cm across. Clearly, amalgamation can occur over very short ranges laterally. Large SSZ are seen mainly in section, but some must have areas very

FIG. 2. (A) Truncating stylolite; (B) seam amalgamation. Bar = 4 mm.

Mass transfer during metamorphism of arenites much in excess of 103 m 2, and are virtually all 'mature' forms; the best developed seem to be thick undulate zones which locally look like semiinfinite egg cartons on bedding surfaces. The amplitude of the zones is rarely more than the amplitude of local undulations or teeth. These amplitudes seem to be reached early in the solution process, and are low to trivial compared to the extent of the SSZ parallel to bedding. It is also obvious that the final local wavelengthamplitude relationship seen is not a function of a high rate of advance parallel to bedding compared to a low rate of advance normal to bedding. Local amplitude and wavelength seem to be set very early in the growth sequence, especially for the dentate form, and simply develop laterally, tooth by tooth. The amalgamation of SSZ by elimination of intervening layers contributes nothing to amplitude and may indeed reduce it locally; wavelength may increase slightly. However, the overall form of the discoidal zone is one for which the length (comparable to a half-wave) does grow to much greater distances than does the width of rock measured normal to bedding through which the SSZ develops. Locally, it is possible that teeth increase in wavelength, due to irregular advance of a tooth cap relative to its neighbours. Where one cap catches up with its neighbours, the result is to broaden the column without increasing its amplitude. There appears to be no consistent wavelength/ amplitude ratio for the overall form of the SSZ. Small SSZ may have an overall amplitude as great as any extensive thick seam. This again suggests that the bulk of growth activity via solution is in the tip zone.

Meguma meta-arenites* The Meguma Group of Nova Scotia is now regarded as Cambro-Ordovician in age. It is essentially unfossiliferous, and lithologically rather monotonous. The Halifax Formation consists of argillaceous rocks with lesser amounts of arenites; it is the upper formation of the Meguma Group. The Goldenville Formation is dominated by arenites with intercalated argillites (the latter usually called slates) and is the lower formation of the Meguma Group. For several decades, gold was extracted from rocks of the Goldenville Formation. Mines were set mainly in 'saddle reef' localities, and Faribeault spent many years mapping and interpret* arenite is used throughout. The rocks are actually metagreywackes.

199

ing the structure and the location of gold deposits within it (Faribeault 1899). The question of the source of the gold was, and still is, a matter of debate. There is no obvious plutonic source for gold-bearing fluids, and the deposits themselves are totally within the sedimentary rock pile. The question arises: is there any portion of that pile which might be a low grade gold-bearing host, from which gold might have been transferred to sites of useful concentration ? In the course of a team investigation of this possibility, we dealt with aspects of mass transfer in rocks of the Goldenville Formation. The Goldenville Formation is interpreted as having been deposited by turbidity currents (Schenk 1970). Outcrops show that the arenite layers are made up of some or all of Bouma divisions A to D; the 'slate' interbeds can be regarded as Bouma E divisions. A divisions predominate in the thick arenite layers and undisturbed plane bed divisions are very hard to identify. Some arenite layers are actually amalgamated beds. Conspicuous features in many outcrops are water-escape structures. On many layer surfaces, sand volcanoes are abundant, with well-preserved radial discharge patterns and a central vent. They are now much distorted, and suggest a flattening ratio of about 2:1 (Henderson 1983) with the local axial surface as the plane of greatest flattening. In sections through the arenite layers, the elutriation pipes which fed the sand volcanoes are well-displayed. Clearly, the arenites were zones through which, and presumably from which, water passed in substantial quantities. The other conspicuous features in the arenite layers are spaced cleavage zones (SCZ). These appear as phyllosilicate-rich zones which anastomose or taper off within an arenite layer. (The 'slate' layers between arenite layers have a slatey cleavage much more closely packed than the spaced cleavage; it is also inclined at substantial angles to the spaced cleavage, though there is also in some locations gradational refraction associated with an equally gradational grain size change from arenite to 'slate'.) The spaced cleavage is everywhere parallel to, though not coincident with, the elutriation pipes, in both fold core and fold limbs. At sites in the limbs, cleavage and pipes have the same angular relationship to bedding. We interpret this to mean that the cleavage was initiated as zones statistically normal to bedding on the grounds that the pipes to which cleavage zones are parallel would have been statistically normal to bedding both before and in the very early stages of deformation. There is the further inference that cleavage formation began very early in the

P. M. Clifford et al.

200

deformation history of these rocks. The current angular relationship between SCZ and bedding is not explicable either by simple bending of the beds alone, or by flexural slip alone; nor even by an episode of homogeneous flattening superimposed on an early-formed fold geometry. The development of the SCZ was dominated by mass transfer, and that affected the quartz more than anything else. Modal analysis across several cleavage-plus-lithon sets shows that the SCZ have a high phyllosilicate/quartz ratio (P/ Q) and a high muscovite-biotite (M/B) ratio, while lithons have a low P/Q ratio and a low M/ B ratio (Fig. 3). All SCZ are narrower than lithons by an average factor of 5. The orientation distribution of phyllosilicates appears about the same in both kinds of domains, suggesting no or limited rotation or recrystallization of phyllosilicates. Shapes of quartz grains, however, are definitive. Moving from lithons to cleavage zones, quartz grains generally become progressively more inequant and straight-sided, and the fluctuation of their long axes decreases. Fig. 4 shows a sequence of axial ratio versus fluctuation angle (AT/40 diagrams, lithon to mid-cleavage for two sample sites, representing mediocre development and good development of cleavage zones respectively. Similar diagrams have been prepared for several such traverses, and all show the same sequence. The implied changes in strain ellipsoids, calculated from two sections per site, are listed in Table 2. We interpret these ellipsoids, essentially of the type 21 = 22 >> 23, as having been brought about by one-dimensional removal of quartz along 23 . We see scant evidences of ductile shape change of individual grains. The certain sign of solution is that the range of lengths of quartz grains within the cleavage zones, measured parallel to the cleavage traces, is no more than the range of lengths in lithons; at high levels of deformation, it may be less. This means low or no elongations in directions parallel to the

80 60

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FIG. 3. Plots of variation of quartz (q), muscovite (m)

and biotite (b) in (A) sample G26-415 AC having well-developed cleavage, and in (B) sample G20-753 AC having poorly-developed cleavage. Dark bars indicate cleavage zones. Sections normal to cleavage, parallel to lineation. cleavage planes, but substantial negative elongations normal to the cleavage planes. Treating these negative elongations as being strictly due to loss in 23, n o gain in 21, 22, the volume losses are about 70% for CZ proper. It should be noted that the intermediate zones, IM, are zones which show considerable signs of pressure solution compared to lithons, but which are not as affected as CZ. If they were included

TABLE 1. Modal analyses, two sections for each of four samples, of Goldenville metagreywackes. > 2000 points per section

G26 415 AC G26 415 BC G26 353 AC G26 353 BC G19 354 AC G19 354 BC G20 753 AC G20 753 BC Overall averages (%)

18 mrn

Quartz

White mica

Dark mica

Opaques

Carbonate

Feldspar

68.6 58.8 68.1 68.7 68.3 58.0 67.0 73.2

26.0 34.0 24.9 23.6 23.0 20.4 23.5 17.7

1.4 2.2 2.9 3.3 2.7 6.2 4.9 4.5

2.0 1.9 1.1 1.3 1.0 0.8 1.6 1.3

1.7 2.2 1.7 2.5 3.6 3.5 1.4 2.1

0.3 0.9 1.2 0.6 1.3 1.0 1.4 1.1

68

24

4

1

2

2

Mass transfer during metamorphism of arenites TABLE 2. Calculated strain ellipsoids, Meguma

meta-arenites and inferred volume loss, Av Sample

Site

Xt

Yt

Zt

Av (%)

G26-415

CZ IM ML

1.15 1.31 1.30

1.0 1.0 1.0

0.25 0.53 0.81

72

G26-353

CZ IM ML

1.08 1.03 1.05

1.0 1.0 1.0

0.27 0.56 0.70

71

G19-354

CZ IM ML

1.07 1.01 1.03

1.0 1.0 1.0

0.30 0.45 0.71

68

G20-753

CZ IM ML

1.27 1.15 1.03

1.0 1.0 1.0

0.33 0.56 0.80

68

CZ: cleavage zone; IM: intermediate zone; ML: mid-lithon.

as part of SCZ (almost certainly the case chemically; and approximately the case modally), the effective volume loss would decline to values very like those from the chemistry cited below. Chemical analyses of samples of lithons and cleavage zones support the notion of large-scale material transfer. Analyses of lithons and cleavages at five widely separated sites are given graphically in Fig. 5. Within each population, compositions are somewhat alike. Between lithon and cleavage, the following differences are obvious: (i) a loss of SiO2, CaO, Na20, from SCZ; (ii) a gain of A1203, Fe203, MgO, K 2 0 and TiO2 to SCZ. Of these oxides, only SIO2, and A1203 are abundant; the rest are minor in amount. Clearly, the adjustments involve mainly the removal of SiO2 from SCZ. The shift of K 2 0 is presumably associated with the abrupt rise in the muscovite/biotite ratio in CZ as compared with the lithon. TiO2 is contained in tourmaline, which is simply concentrated as an 'insoluble residual' mineral in CZ. Adopting the approach of Gresens (1967), and taking the densities of cleavage and lithon as the same (such difference as there is does not materially affect the argument) we have that: Xn = lO0[fvn(CnC-C.L)], where X. is the actual amount of component n lost or gained; f~. is volume factor for component n; C. c, 6'. L are the amounts of component n in cleavage and lithon respectively. We calculated values of f~ for all oxides in the analyses (Fig. 6). We now choose a component or components which are judged least affected by transfer processes. Gresens (1967) suggests that these can be identified according to the following criteria:

201

1 A clustering of f~ for several components, suggesting these had been gained or lost to the same degree, or more likely, that they had been relatively immobile components. 2 A consistency of ratios of two or more oxides. Where such ratios are about the same in lithon and cleavage, the components were gained or lost to the same degree, or maybe not moved at all. In Fig. 6 the obvious cluster to use is [TiO2 + AI20 3 + Fe203]; of these, f~ (A1203) is taken as the base (true) volume factor for the system; it is the central one of the three, and is the most abundant, least influenced by minor gains or losses. This base volume factor for the five sites ranges from 0.4 to 0.6, and is a direct measure of what is left behind after transfer of other components. A volume loss of 40~o (f~=0.6) to 60~o (fv = 0.4) is implied for SCZ, relative to its parental rock volume. These SCZ are not uniform along their trace. In both outcrop and thin section, individual CZ have diffuse 'tips' and quite concentrated central sections. SCZ amalgamate with other zones in their own plane (rare, because of the scale) and with adjacent SCZ to yield an anastomosing pattern of SCZ enclosing lensoid volumes of lithons. It seems that SCZ grow in much the same way as solution zones in dolostones.

Deposition of material The inverse of the solution described above is deposition. In the dolostones, dolomite has been formed in at least two stages. The initial stage was as replacement of initial calcite (or possibly aragonite in fossil material)--early dolomite. A second stage occurs as a generally uniformly thick overgrowth on the first-formed dolomite crystals, nicely revealed by cathodo-luminescence. There are also dolomite linings to pore space; in some pores these fillings are only partial and geopetal. Finally, there are discrete, rather large dolomite crystals which grow into voids independently. This kind of dolomite precipitation suggests a substantial original porosity in many parts of the rock. Some of that porosity has been occluded by dolomite presumed derived from the solution which was left behind the SSZ. Patterns of porosity in the vicinity of some small SSZ are illustrated in Fig. 7. They suggest relatively short range transport of dolomite. Because we are uncertain how much material has been lost due to the formation of SSZ, we cannot tell how much of the late dolomite is to be credited to the solution process; but given the apparent short range of transport, it may well be all of it.

P. M. Clifford et al.

202 16

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Mass transfer during metamorph&m of arenites

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AVERAGE CLEAVAGE COMPOSITIONS OF GOLDENVILLE SAMPLE SETS

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FIG. 5. Graph of chemical variations in five sample sets. Sample G26-415 has the best development of cleavage.

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Mass transfer during metamorphism of arenites

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VOLUME

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FOR G O L O E N V I L L E

2. SAMPLE

--,v- - SETS

FIG. 6. Volume factor plot, after the method of Gresens (1967); points for Ti02, A 1203, Fez03--filled symbols.

The Meguma rocks are rather different. Our chemical and physical measurements suggest that an individual SCZ represents at least twice its current width/volume of original rock (the range, derived via three different methods of assessment, is from less than 50% loss up to 80% loss of material from SCZ). Mass balance calculations suggest that we should see between 10 and 25% 'new' quartz in lithons and SCZ if the transfer is only local, from SCZ to adjacent lithons. In fact, we see less than 5%, in the form of straight-fibre beard overgrowths of quartz in lithons and, to a lesser extent, in SCZ. This suggests that perhaps 75% of the SiO2 removed from SCZ is actually lost from the local system (Fueten et al. 1986). Henderson (1983) thinks that the quartz veins which are relatively abundant at higher structural and stratigraphic levels may be precipitated SiO2 from deeper SCZ. They may constitute up to 10% of the local rock volume. One implication of these values for the amount of SiO2 moved and the distances it may have travelled is that a substantial quantity of fluid is probably involved. Even at the highest, most favourable solubilities of SIO2, volumes of water of the same order of magnitude as the solid rock volume are suggested. This, allied to the inferred early initiation of cleavage, leads us to the conclusion that part of the process was associated with the large-scale dewatering of the sedimentary pile.

Bulk dimensional changes Estimates of volume loss may be translated into shortening of rock column or along layering. For SSZ in dolostones, the actual loss is variable. 'Dirty' rocks have a greater proportion of insolu-

ble residue to accumulate in SSZ; in their case, a millimetre of seam may be equivalent to a centimetre of original rock. For 'cleaner' rocks, the ratio is clearly higher. The width of rock covered by the diffuse 'tip' area is a better guide to the likely loss, which appears to be between 50 and 90% locally. In one exceptional instance, an orthoconic nautiloid has been trapped in a seam. Its internal structure is well-preserved and undistorted, and its original diameter of 15 mm lies in the plane of the seam; but its thickness is less than 1 mm. This implies a volume loss of more than 90% accomplished solely by solution. The irregularity of spacing makes any estimate of bulk shortening somewhat hazardous. In the Meguma rocks, bulk shortening on axes normal to SCZ based upon strain analysis via the fluctuation diagrams, and on mass balance calculations applied to modal analyses of cleavage and lithon, is about 10%. The first method ignores the possibility of any constant-volume ductile strain within quartz grains. At one site this may account for 10-15% of the total shape change of the quartz grains, but for other sites its contribution is negligible. The second method takes the current modal composition of the lithon as original. There may have been some small solution loss in the lithons; there is also some small depositional gain in CZ. The effect of these does not materially alter our conclusions. In both cases, the zones are not of constant character along their length. This means that estimates of local volume loss or of local shortening are not likely to be the same everywhere along these lengths. For cleavage zones in the Meguma, the variation is not very large. But for SSZ in dolostones, there may be lateral variation in the amount of loss from the stratigraphic column due to any one SSZ. For some rather small SSZ the

206

P. M. Clifford et al.

~!!:~!i~ ~i !~ ~

!~!!~!i~ii!~!ii~i!i~ ii~~i~!~!¸ ....... :~

1/i FIG. 7. Porosity variations in dolostones in vicinity of SSZ. Width of each section about 2 cm. Dark areas are epoxy, filling pore space. Note low porosity zone flanking maturer portions of SSZ. Total length of seams approximately 5 cm.

Mass transfer during metamorphism of arenites loss may range from zero to a centimetre over less than 10 cm extent (zero to 90% localized collapse). For larger SSZ, the lateral gradient in shortening is lower. Two things are certain, however. First, the amplitude of the structure is not a very useful guide to amount of column collapse. Second, the loss of material is nicely adjusted through the affected rock volume, so that, taking one column with another, the collapse is about the same in both. This is accomplished by successive seams not being in phase with each other, and by having seams frozen at slightly different values of material loss. Such local irregularity of loss may well accentuate local irregularities in the quality of layering. There are no signs of any differential settling as a result of solution in the Lockport Formation dolostones.

Discussion The most interesting difference between the dolostones and the meta-arenites is the apparent transport distance. In the dolostones there is substantial occlusion of porosity, up to complete elimination of pores in the vicinity of the seams. Points further from a seam than half the wavelength for small SSZ suffer little or no apparent porosity loss compared to points close to the seam, though there is much variability. The loss of porosity is clearly due in large part to deposition of second-generation dolomite, which can be seen to line the pores more or less evenly. Porosity loss may have occurred due to removal of material followed by subsequent physical collapse of pore space, but there is no clear evidence for this process, and its real effect is unknown. A plausible interpretation is that the secondgeneration dolomite was produced by solution, but travelled very short distances, centimetres only. If this were the case, then the exact process of transfer must have been such as to encourage deposition after a brief residence time in solution. This may be due to: (i) the near-saturated state of the local fluid phase prior to solution, so that little new material could be dissolved before some of it had to be reprecipitated elsewhere if the solution process was to continue; (ii) lack of sufficient fluid phase at rest, or a sufficient flux of fluid phase, to dissolve and allow movement of large quantities of material substantial distances; (iii) restricted permeability. The process appears to be somewhat analogous to case 4 of Durney (1976), in that individual crystals of dolomite are badly dissolved, or even removed entirely, in the vicinity of SSZ, only for

2o7

dolomite to appear by precipitation in a nearby pore space. The system is porous and permeable to a limited extent at the beginning of the process, but becomes less so locally as precipitation occurs. It may be close to diffusional transfer. The situation in the Meguma is different. If our estimates for the amount of SiO2 moved from SCZ are correct, then little of that SiO2 was reprecipitated locally, judging by the low abundance of beard overgrowths in nearby lithons. The conclusion is that an early stage of the transfer process took place in a porous, open system (case 4 of Durney 1976), during which stage much of the SiO2 left the local system. A second stage of beard quartz precipitation corresponds to case 1 of Durney (1976), suggesting the total occlusion of porosity at the end of stage 1, thereby converting the local process to a constantvolume process. This accords with the view of Fyfe (1976, p. 205) that ' . . . the process will be most important when fluids are being generated and porosity and permeability permit a continuous transport path . . . ' . It fits exactly with Durney's supposition that in a porous rock, solution-transfer should be accomplished by uniaxial shortening increments, and that pressure-growth positive elongations will commence as porosity reaches zero (Durney 1976, pp. 2367). From this discussion, we may conclude that material transfer in and through the Meguma rocks was a longer range process than transfer in the Silurian dolostones; this we ascribe to the enhanced role of water. This water was obviously able to circulate much more freely in the Meguma rocks than in the dolostones. Presumably, permeability was maintained for a proportionately longer time in the Meguma rocks. Also, the pressure would have helped in sustaining fluid movement through the rock. This would result in more rapid transport of solute, than in a static fluid, at low temperature and pressure, such as perhaps occurred in the dolostones. This process for the Meguma rocks is, therefore, dominantly one o f transfer by infiltration rather than by diffusion. For the Meguma, both mudrocks and arenites will have contained water. However, this is clearly volumetrically inadequate to the task, and some additional source is required. The only likely candidate is meteoric water. If this is a reasonable notion, it further reinforces the conclusion about transport early in the deformation process in a rock with a well-maintained permeability. A second difference between the rocks is mineralogical. The dolostones are virtually monomineralic and unmetamorphosed; the metaarenites are mainly quartz and phyllosilicates,

P. M . Clifford et al.

208 SSFROM TLOS YSTE]M

RESIDUALSSZ

BEARDS

PRESSURESOLUTION/ SiO2 MOVES H2OMOVES

/

"-,. \ J. . . . . . . . . . . .

I

'TIMEr LONG-RANGEITY

DOLOMITE I 0

TIME

DOLOMITE II ~

-

~SHORT-RANGU PERMEABILITY

Schematic evolution of spaced cleavage zones (left) and spaced solution zones (right)

FIG. 8. Schematic outline of evolution of SCZ (left) and SSZ (right). and are metamorphosed. In both case, phyllosilicates are present, and are essential to the process; solution of any consequence takes place only when clays or micas are adjacent to dolomite or quartz grains. The Meguma rocks, however, were relatively 'dirty' rocks. The biotite and muscovite now present have been recrystallized from a presumed clay parent. The most likely transformation is from clay to white mica, e.g. 6 Ca0.5(Na, K)o.sAI3MgSi7.sAlo.5020(OH4) + 4(Na, K) ÷ + 14H ÷ 3.5(Na, K)2A14Si6A12020(OH)4 + 3Ca 2÷ + 6Mg 2÷ + 12H20 + 24SIO2 (Beach & King 1978, equation 2). This liberates SiO2, H20, Ca, Mg. In our rocks the amount of MgO gained by the cleavage is trivial; we here discount it, as also we discount the loss implied by the above equation. The released SiO2 and the water may have short- or long-distance transport. We prefer short; the above reaction is probably metamorphic rather than diagenetic in these rocks, and so this release of SiO2 would be late in the process, perhaps after closure of pore space. Perhaps this is part of the source for the quartz beard overgrowths. However, if this reaction did proceed during the later stages of deformation of these rocks, one would expect to see a rather well marked preferred orientation to the new phyllosilicates. This does not seem to be the case. A third difference lies in the periodicity of SSZ and SCZ. In the dolostones, seams are developed to varying degrees in any given section. Although an average spacing can be calculated, it is not very meaningful. 'Dirty' rocks can be thoroughly pervaded by SSZ, and have an average spacing less than 1 cm. 'Clean' rocks may have very few SSZ, even none, and the average spacing rises to 10cm. We conclude, with many others, that whatever the precise mechanism, the most likely

places for substantial solution of dolomite and construction of SSZ are in clay-rich zones. These are set by primary depositional processes; the mechanism takes advantage of such clay-rich zones, and the overall periodicity (or lack of it) is established accordingly. We concur with Merino et al. (1983) that planes which separate layers of initially different textures are sites particularly susceptible to SSZ formation. Whether the further generation of SSZ by a porosity instability occurs, we cannot say. Periodicity of SCZ in Meguma rocks is a much more regular affair. Cleavage spacing is quite consistent at any one site; but the spacing decreases, proceeding from sites in fold limbs to sites in the fold hinge zone, a distance of 0.5 km, measured normal to mean axial surface. The change is gradual over that distance. Also the SCZ form at large angles to the layering, so the inauguration and growth of SCZ cannot be a consequence of fluctuations in properties produced in the sedimentation episode. The distribution of phyllosilicates in lithons, taken as representative of rock relatively unaffected by mass transfer, is essentially uniform, and provides no obvious incentive for preferential location of solution sites. Elutriation pipes may provide porosity changes along the bed, and so set up a mechanism of the kind that Merino et al. (1983) suggest. But the SCZ also occur in rock where there is no sign of such pipes, so the overall mechanism must involve some other factor. The regularity of spacing of SCZ, as compared to SSZ, is a function of the angular relationship to bedding. The SSZ in our dolostones is at the mercy of fluctuations from layer to layer. They may not occur at consistent spacing because local layer mineralogy or porosity/permeability are unfavourable. But for SCZ, at high angles to So, each zone has to cross the same basic variation in layer character. Though there may be changes along the zones due to such character changes,

Mass transfer during metamorphism of arenites

209

there is no incentive for development of local highly erratic spacing intervals.

losses from the system, this means a 10~ dimensional loss in directions normal to cleavage in the arenites. Bulk shortening of the dolostones is really unknowable, but we estimate it at least

Conclusions

5%.

Mass transfer in unmetamorphosed dolostones and in modestly metamorphosed arenites has led to the production of spaced solution zones and spaced cleavage zones. These appear geometrically similar in their general form. Relatively 'insoluble' materials are passively concentrated in such zones, with the bulk of the activity at any instant concentrated at the ends of such zones. In both cases the bulk of the transfer occurred early in the rock history. The higher porosity and permeability of the arenites, and the greater availability of water, promoted considerably more solution and transported material much further than the process acting in the dolostones. Local losses associated solely with the production of an individual zone and measured by three different methods, are anywhere from 50% shortening to 80% shortening. Translated into bulk

We note finally that in the dolostones the residual seams are essentially impermeable. Their growth increases the path length for transport of anything and markedly increases the anisotropy of bulk permeability. Both effects are liabilities if one is interested in extracting fluids from the rock; they are also deplorable if one is concerned with waste disposal in the subsurface. ACKNOWLEDGEMENTS: Discussions with J. H. Crocket, J. R. Henderson and P-Y. F. Robin have been both stimulating and helpful in developing some of our ideas about the Meguma rocks. The essential financial support has been provided by a Research Contract Award from Energy, Mines and Resources, Canada, for work on the Meguma rocks, and by a scholarship to MCR from Texaco Canada Limited to study solution features in carbonates; also by continued invaluable help from the Department of Geology, McMaster University. M. J. Thompson carried out modal analyses.

References BEACH, A. & KING, M. 1978. Discussion on pressure solution. J. geol. Soc. Lond. 135, 649-57.

DURNEY,D. 1976. Pressure-solution and crystallization deformation. Philos. Trans. R. Soc. London, A283, 229-40.

FARIBEAULT,E. R. 1899. The gold measures of Nova Scotia and deep mining. J. Min. Soc. Canada, 2, 119-28. FUETEN,F., CLIFFORD,P. M., PRYER,L. L., THOMPSON, M. J. & CROCKET,J. H. 1986. Formation of spaced cleavage and concurrent mass removal of SiO2, Meguma Group greywackes, Goldenville, Nova Scotia. Mar. Sedim. Atlantic Geol. 22, 35-50. FYFE, W. S. 1976. Chemical aspects of rock deformation. Philos. Trans. R. Soc. London, A283, 221-8.

GRESENS, R. 1967. Composition-volume relationships of metasomatism. Chem. Geol. 2, 47-65.

HENDERSON,J. R. 1983. Analysis of structure as a factor controlling gold mineralization in Nova Scotia. Current Research,part B. Geol. Sum. Canada; Paper 83-1B, 13-21. MERINO, E., ORTOLEVA,P. & STRICKHOLM,P. 1983. Generation of evenly-spaced pressure-solution seams during (late) diagenesis: a kinetic theory. Contrib. Mineral. Petrol. 82, 360-70. SCHENK,P. 1970. Regional variation of the flysch-like Meguma Group (Lower Paleozoic) of Nova Scotia, compared to recent sedimentation off the Scotian Shelf. Geol. Assoc. Can., Special Paper 7, 127-53. WANLESS, H. R. 1979. Limestone response to stress: pressure solution and dolomitization. J. sediment. Petrol. 49, 437-62.

P. M. CLIFFORD,M. C. RICE & L. L. PRYER,Department of Geology, McMaster University, Hamilton, Ontario L8S 4M1, Canada. F. FUETEN,Department of Geology, Erindale College, University of Toronto, Mississauga, Ontario, Canada.

Wet-sediment deformation in the Upper Ordovician Point Leamington Formation: an active thrust-imbricate system during sedimentation, Notre Dame Bay, north-central Newfoundland Kevin T. Pickering S U M M A R Y: The Upper Ordovician (late Caradoc-Ashgill) Point Leamington Formation is essentially composed of fine-grained, thin-bedded turbidites that accumulated in a deep marine environment below wave-base. Immediately underlying, and apparently conformable with the formation, there are about 120 m of red, red-green and grey bioturbated cherts overlain by late Llandeilo-early Caradoc black shales--these argillaceous, 'pelagic', sediments conformably overlie about 800 m of mafic pillow lavas and flows of the Lawrence Head Volcanics (middle Exploits Group). The Point Leamington Formation, up to 2200 m thick, together with the mainly fine-grained, thin-bedded turbidites, contain locally abundant wet-sediment deformation in horizons up to tens of metres thick, although most of the thickest horizons appear to represent multiple events. Conglomerate-filled gullies, channels and canyons occur within the finer grained Point Leamington Formation facies-associations. The formation is overlain, gradationally, by the deep-water conglomeratic Goldson Formation, interpreted as submarine canyon deposits; with the Ordovician-Silurian boundary occurring towards the top of the Point Leamington Formation or towards the base of the Goldson Formation, based on the presence of lower Llandovery corals in limestone boulders within an olistostrome near the gradational boundary between these formations. Wet-sediment deformation occurs as: (i) coherent folded layers; (ii) semi-coherent folded layers; (iii) chaotic balled or brecciated sediments; (iv) boudinaged layers; (v) faulted layers, involving normal, reverse and thrust faults, and best seen on a micro- to meso-scopic scale. Clastic dykes, other liquefaction and fluidization structures, together with convolute lamination, also occur either isolated from, or in association with, the above listed wetsediment deformation styles. The wet-sediment deformation generally appears to be related to gravity-controlled slope failure, in surface to near-surface mass failure or at unspecified depth of burial, on a margin with a regional south to south-eastward downslope dip, although some folds may be interpreted as thrust-related (tectonic) deformation in wet sediments. The overall stratigraphy, sedimentology and structure suggests Upper Ordovician to Lower Silurian deep marine sedimentation in small fault-bounded and fault-controlled basins, some of which contained relatively coarse-grained, small-diameter, submarine fans. Regional considerations are consistent with the succession having accumulated in a thrust imbricate system, active during sedimentation, with new evidence to indicate AshgillLlandovery volcanism and magmatism that can be related to initial subduction followed by possible crustal extension. Sometime during the Ashgill, subduction appears to have ceased as all the intervening oceanic crust was consumed and volcanic arc activity shut down. During late Ashgill-Llandovery time, thin continental crust of the 'Gander' terrane to the east then underplated the 'Dunnage Zone', probably leading to foreland basin development, analogous to the plate tectonic processes in the Banda Arc today. Thus, there was a reversal of subduction polarity from the eastwards subduction of pre-Caradoc times, such that the remnant backarc basin became, for a while, a forearc as the wide backarc or marginal basin floored, at least in part by oceanic crust, telescoped. Contemporaneous sinistral shear appears to have been important in the late Ordovician-Silurian, during which time unspecified 'suspect' terranes may have slipped in and out of the Dunnage Zone. The contemporaneous Ashgill-Llandovery bimodal igneous activity can be explained by phases of crustal transtension as re-entrants in opposing plates slid past each other, followed by phases of transpression to re-activate the thrust imbricate system. G r a v i t y - i n d u c e d s e d i m e n t failure involving both superficial and sub-surface (to depths of a few h u n d r e d metres) s e d i m e n t sliding is widely reported from most m o d e m continental margins, for example in the books edited by Doyle & Pilkey (1979), Saxov & N i e u w e n h u i s (1982),

W a t k i n s & D r a k e (1982), Stanley & M o o r e (1983), volumes 1, 2 a n d 3 by Bally (1983), and Stow & P i p e r (1984). In m a n y cases, the deform a t i o n and sliding involves partially lithified (undrained) wet-sediment. Clearly, it is inappropriate to define c o n t e m p o r a n e o u s 'soft-sediment'

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks,

~eological Society Special Publication No. 29, pp. 213-239.

213

K.T. Pickering

2 I4

deformation within the superficial, uppermost metres of a sedimentary succession, since gravityinduced sediment mass movement can involve substantial sediment thicknesses. Inter-stratal slip within sediment slides, between individual layers/beds and packets of layers/beds, will generate shear zones and resultant micro-/meso-structures that may resemble 'hard-rock' features, for example the listric fault patterns and 'roll-over' structures due to gravitational gliding in deltas (Crans et al. 1980; Mandl & Crans 1981). In modern continental margins and other basin slopes, it may be possible to elucidate the driving mechanism for sediment failure, determine the rheology of the sediments and appreciate whether or not failure was associated with over-pressured zones at depth. However, in ancient sedimentary successions such distinctions may not prove possible. Furthermore, in many ancient successions, arguments about 'tectonic' versus 'sedimentary' sediment deformation may be facile, given the complex interaction of tectonics and sedimentation, and the common occurrence of overpressured sediments at depth. These are problems that are particularly germane when considering active continental margins. Given the syntactic and semantic arguments that are generated when trying to define 'tectonic'

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versus 'sedimentary' soft-sediment deformation, and in agreement with Maltman (1984), who stressed the need to describe fully the kind of deformation structures, whether they be 'early' or 'late', and the generating force(s), the term 'wet-sediment deformation' is adopted in this study to describe deformation structures in unlithified, undrained sediments. The term 'wetsediment' is preferred to 'soft-sediment' because it tends to suggest deformation of sediments that still contain significant trapped pore fluids and does not have a depth connotation. All ductile deformation, whether early syn-sedimentary or late tectonic, with lithified sediments, must involve soft sediment. An attempt is made to differentiate between early (generally superficial) and late (at unspecified depths beneath the palaeomud-line) wet-sediment, gravity-induced deformation. Also, such deformation is described with reference to the lithified sediment, hard-rock, deformation that occurred within the succession. The purpose of this paper is to describe examples of wet-sediment deformation that are believed to have been mainly gravity-induced, and which occur within a Caradoc-Llandovery deep-marine succession in north eastern central Newfoundland, associated with an active plate margin. Since the succession contains abundant deformation of unlithified, semi-lithified and

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FIG. 1. Map modified from Colmann-Sadd & Swinden (1984) of Newfoundland to show the major tectonostratigraphic zones (after Williams 1978), distribution of ophiolitic rocks (black) and allochthonous rocks emplaced on to the Humber Zone (diagonal line legend) during the Taconic Orogeny. Study area in Dunnage Zone shown by boxed area. BV-BL = Baie Verte Brompton Line; LC-CF = Lobster Cove-Chanceport Fault; LA-SHF = Lukes Arm-Sops Head Fault; RF = Reach Fault; CRF = Cape Ray Fault; HF = Hermitage Flexure; D-HBF = Dover-Hermitage Bay Fault.

Wet-sediment deformation in the Leamington Formation TREMADOC

Passive

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K. T. Pickering

216

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Wet-sediment deformation in the Leamington Formation

217

FIG. 2. Summary cartoon of Tremadoc to Llandovery geological evolution of the Dunnage Zone, Newfoundland. Diagrams drawn with information from Dean (1978), Nelson (1981), Watson (1981), Arnott (1983a, b), Arnott et al. (1985), Colman-Sadd & Swinden (1984), and Jacobi & Wasowski (1985). Location of N e w Bay (including Point Leamington Formation) stratigraphy beneath central part of Llandeilo-Llandovery reconstructions. H Z = H u m b e r Zone; C H = Cow Head Group; FdeL = Fleur de Lys rocks; M H G = Moretons Harbour Group (shown in forearc, but could be intra-arc); TG = Twillingate Granidiorite; C G & LB = Cutwell Group & Lushs Bight Group (incl. Pacquet Harbour Group, Snooks Arm Group); W B G = Wild Bight Group (may incl. Frozen Ocean Group); L E G = Lower Exploits Group, including Tea Arm Volcanics (TAV) & N e w Bay Formation (NBF); SG = Summerford Group; L H V = Lawrence Head Volcanics; D M = Dunnage Melange; C A L = Cobbs Arm Limestone; B V - B L = Baie Verte-Brompton Line; L C - C F = Lobster Cove-Chanceport Fault; SAF = Shoal Arm Formation; LHS = Lawrence Harbour Shale; D H F = Dark Hole Formation; RCS = Rodgers Cove Shale; RAV = Roberts Arm Volcanic Belt; L A - S H F = Lukes Arm-Sops Head Fault; BG = Burlington Granidiorite; G I F = Gull Island Formation; P L F = Point Leamington Formation; M A F = Milliners Arm Formation; N F = N e w Bay Fault; TF = Toogood Fault; C A F = Cobbs Arm Fault; BoI = Bay of Islands and associated ophiolites; G F = Goldson Formation. Melanges such as the Boones Point Complex are associated with the Lukes Arm-Sops Head Fault, as olistostromes tectonized by the syn-sedimentary thrust faulting. Subduction directions are shown, together with backarc spreading centres. Arc-continent collision began in the late Llanvirn; this collision also involved the SG, D M and CAL but the 2-D sketches do not adequately show this--throughout this time, over 800 m of tholeiitic basalts (LHV) of the L E G were extruded in a submarine environment, implying continued backarc extension. Gander Zone underplating may have occurred from late Ashgill onwards. Sea-level shown by lines marked sl. During Caradoc, cessation of arc volcanism may have led to thermal contraction of the lithosphere and associated subsidence (chert and black shale deposition), possibly with global rise in sea-level. See text for explanation and Table 1 for stratigraphy. Time-scale after McKerrow et al. (1985). AshgillLlandovery (and Wenlock) history probably involved phases of crustal transtension with associated bimodal igneous activity of the Springdale Group and youngest parts of the Topsails igneous complex (Whalen et al. in press); but history is essentially crustal transpression under sinistral shear couple.

TABLE 1. Summary stratigraphy of Lower Palaeozoic Dunnage Zone, Notre Dame Bay, Newfoundland (adapted and modified from Arnott et al. 1985). Faults separating successions are : L C - C F = Lobster Cove-Chanceport; N B F = N e w Bay; VVF = Village Virgin; BIF = Boyds Island; BCF = Byrne Cove, and R F = Reach Fault. Stratigraphic units are: BC = Bobby Cove Formation; VB = Venams Bight Basalt; BB = Balsam Bud Cove Formation; R H = Round Harbour Basalt; SI = Stag Island Formation; P H = Pigeon Head Formation; BH = Burnt Head Formation; PP -- Parson Point Formation; LT = Long Tickle Formation; W B G = Wild Bight Group (? time equivalent to Frozen Ocean Group); SAF = Shoal Arm Formation; G I F = Gull Island Formation; BPC = Boones Point Complex; T A V = Tea Arm Volcanics; N B F = New Bay Formation; L H V = Lawrence Head Volcanics; LHS = Lawrence Harbour Shale; P L F = Point Leamington Formation; G F = Goldson Formation; CCI = Carters Cove Olistostrome; IHI = Intricate Harbour Olistostrome; J W A = Joe Whites Arm Shale; M A F = Milliners Arm Formation; J C M = Joeys Cove Melange; B = basalts; C A L - - Cobbs Arm Limestone; RCS = Rodgers Cove Shale; B M C - - B i g Muddy Cove Formation; LF = Lawrenceton Formation; W F = Wigwam Formation. The Indian Island Group, in the Botwood Zone, probably was contemporaneous with the uppermost part of the Davidsville Group and the lower part of the Botwood Group (Dean 1978). Timescale after McKerrow et al. (1985), with graptolite zonation after Williams et al. (1972). For detailed lithostratigraphy, see Dean (1978). Note, pre-Caradoc mainly volcanoclastic/pyroclastic successions interlayered with arc volcanics are differentiated from post-Caradoc sandstones rich in arc-derived volcanic clasts. Question marks denote considerable uncertainty about boundaries and precise age ranges. New data (Whalen et al. in press) suggests Ashgill-Llandovery phase of arc-related igneous activity in western Newfoundland, followed by a late Ashgill-Llandovery-Wenlock crustal extension phase to produce the bimodal youngest part of the Topsails igneous complex and the Springdale Volcanic Group

FIG. 3. Geological map of the N e w Bay area and western Exploits Bay area, Notre Dame Bay, Newfoundland. Point Leamington and Goldson Formation outcrop based on author's mapping, with oldest thick conglomeratic unit in Point Leamington Formation based, in part, on Helwig (1967). Other formation and group boundaries based on published geological map (1 : 50,000 Government of Labrador & Newfoundland, Dept. of Mines & Energy). See Table 1 for stratigraphy. Typical bedding dip and direction of dip shown, together with major faults. Cottrells Cove Group age uncertain--may be pre-Caradoc.

218

K. T. Pickering

lithified sediments, it is a useful field laboratory in which to study sediment deformation from wet-sediment to dry-sediment failure.

Regional setting The study area in NE central Newfoundland, bordering Notre Dame Bay, occupies part of the Lower Palaeozoic Dunnage Zone (Fig. 1) of the Central Mobile Belt (Williams 1978, 1979; Williams & Hatcher 1982, 1983). The stratigraphy and structure are believed to represent the vestiges of the Palaeozoic Iapetus Ocean (Williams 1964a, 1984) connected to the mid-European Tornquists Sea (Cocks & Fortey 1982; Neuman 1984). The western boundary of the Dunnage Zone is the Baie Verte-Brompton Line, marked by ophiolitic rocks, emplaced in the Llandeilo (Williams & St Julien 1982), while the eastern margin is defined by the Gander River Ultramafic Belt (GRUB). Marshall Kay and his students (Helwig 1967; Horne 1968; Eastler 1969, 1971; Kay 1976), together with Williams (1962, 1963, 1964b), regarded the structure of eastern Notre Dame Bay in terms of fault-bounded tectonostratigraphic units with distinctive and differing structure and stratigraphy. Research by Nelson (1981), Watson (1981), Arnott (1983a, b) and Arnott et al. (1985), supports the earlier interpretation of fault-bounded and fault-controlled basins. New evidence from the Topsails igneous terrain and the Springdale Group (Whalen et al. in press), in the Dunnage Zone, suggests Ashgill arc-related igneous activity about 438 Ma (Rainy Lake complex). Also, Pickering (unpublished data) has identified some hitherto unknown dacite/rhyo-dacite flows in the late CaradocAshgill Point Leamington Formation that appear to be essentially contemporaneous with sedimentation and would, therefore, correlate well with the new data of Whalen et al. (in press). Late Ashgill-Llandovery igneous activity in the Springdale Belt (about 429-427 Ma) and the youngest parts of the Topsails igneous terrain (about 417-415 Ma) could have formed during crustal extension lasting between 1 and 9 Myr (Whalen et al. in press). The history of the Dunnage Zone, from Arenig to Llandovery times, may be considered in terms of Llanvirn to Llandeilo arc-continent collision (Humberian Orogeny), with south-eastward-directed subduction, followed by the post-Caradoc telescoping of the remnant backarc or marginal basin(s), possibly eventually leading to foreland basin development during the later stages of collision in the late Ashgill-Wenlock. Major

sinistral shear appears to have been important from the Ashgill onwards, based on regional structural considerations, and by extrapolation from the British Caledonides (Murphy & Hutton 1986; Soper & Hutton 1985) to those in Newfoundland (Fig. 2). The late Ashgill-Wenlock igneous activity (see above) suggests phases of crustal transtension during essentially transpressional plate collision; probably as major promontories and re-entrants slid past each other in opposing plates. The geology of central Newfoundland is summarized by Dean (1978) and Kean et al. (1981), who note the diachronous influx of sand-rich turbidite successions and associated conglomeratic deposits (Table 1). The wet-sediment deformation described in this paper occurs within the late Caradoc to Ashgill Point Leamington Formation (Table 1 and Fig. 3), interpreted as basin-slope and terrace sedimentation. The Ashgill represents about 7 Myr (McKerrow et al. 1985), but there is considerable uncertainty about the Caradoc, with values ranging from about 26 Myr (Ross & Naeser 1984), through 12 Myr (McKerrow et al. 1985), to 10 Myr (Harland et al. 1982). Thus, the time period involved for deposition of the oldest parts of the Point Leamington Formation (late Caradoc), together with the immediately underlying Caradoc cherts and black shales, has considerable uncertainty.

Study area The late Caradoc to Ashgill (? Llandovery) Point Leamington Formation, initially proposed as the Point Leamington Greywacke in the Upper Exploits Group by Helwig (1967, 1969) and H o r n e & Helwig (1969), is an approximately 2200 m thick succession of mainly thin-bedded, fine-grained siliciclastic, deep-marine turbidites and other sediment gravity flow deposits (Fig. 4). Locally abundant wet-sediment deformation horizons within the formation are the focus of this paper. The Point Leamington Formation occurs in the New Bay area, and along the west coast of Exploits Bay, in Notre Dame Bay, north-central Newfoundland (Fig. 3). The formation is part of the Dunnage Zone (or Central Mobile Belt). Immediately underlying, and apparently conformable with, the Point Leamington Formation there are about 120 m of red, red-green and grey bioturbated cherts overlain by Caradoc, graptolitic, black shales (Lawrence Harbour Shale and 'Unnamed' Argillite of Helwig 1967, 1969)these 'pelagic' sediments conformably overlie about 800 m of mafic pillow lavas and flows of the Lawrence Head Volcanics in the Middle

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Wet-sediment deformation in the Leamington Formation

219

FIG. 4. Representative outcrop of Point Leamington Formation, showing mainly thin.bc'dde~ fine-grained siliciclastic turbidites. Beds young to left. East coast of Southwest Arm.

Exploits Group (Table 1). The Lower Exploits Group (Helwig 1967, 1969) is interpreted as backarc volcanics and volcanoclastic sedimentation associated with backarc extensional tectonics. There are some thick (up to tens of metres) lens-shaped conglomeratic units within the essentially fine-grained, thin-bedded Point Learnington Formation (Figs 3 and 5), especially towards the Ashgill/Llandovery boundary where there are Llandovery corals in limestone blocks within an olistostrome (Helwig 1967, 1969). The top of the formation is marked by an upward change into about 800 m of mainly pebble conglomerates assigned to the Goldson Formation (Table 1), the top of which is faulted against older rocks.

Wet-sediment deformation In the Point Leamington Formation, now generally metamorphosed to prehnite-pumpellyite facies, wet-sediment deformation is characterized by the following features occurring without associated hard-rock (lithified), tectonic micro-/ meso-structures such as vein arrays, slickensided surfaces and kink bands: 1 Sediment folds without a related cleavage. 2 Curvilinear fold hinges. 3 Continuous layers/beds laterally becoming totally disrupted and chaotic over distances of centimetres to decimetres. 4 Layers of chaotic, brecciated sediments. 5 Detached fold hinges with liquefied sediment

flow structures, both within hinge zones and attenuated fold limbs. 6 Clastic dykes, typically on a centimetre to decimetre scale, associated with deformed horizons, and showing no preferred orientation. 7 Brittle to ductile (including semi-brittle) deformation closely associated in the same horizons, for example faults, boudinaged layers (especially the sand/silt-rich layers), semi-coherent layers, folded coherent layers, and liquefied/fluidized 'homogenized' layers (particularly in the finest grained layers). 8 Successive decollement surfaces separating horizons of either similar or varying deformation style. Contacts appear welded. 9 The main, and first recognized, cleavage commonly is at an acute to larger angle to the fold axial surfaces and the cleavage is refracted through such fold hinge zones. 10 Variable fold style within apparently identical lithologic units, i.e. where grain size, layer thickness, sand/shale ratio and total deformed horizon thickness appear similar. 11 Draped sediments over deformed horizons with irregular bed thickness appearing to infill residual depressions to level any uneven topography. 12 Debris flows (pebbly mudstones) showing erosional surface into the upper parts of some deformed horizons with 'exotic' (olistostrome) material, for example extraformational limestone clasts. 13 In thin-section, there is an absence of cataclastic textures associated with deformation. 14 In thin-section, mud flowage fabric (inter-

K. T. Pickering

220 POINT LEAMINGTON - BRIMSTONE POINT SECTION

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FIG. 5. Representative sedimentary log of Point Leamington Formation facies and facies-associations between Passage Rocks and Brimstone Point on the western shore of Southwest Arm. This section occurs towards the uppermost part of the formation. See text for explanation.

preted from elongate silt grain long-axis alignments) along fault surfaces, and concentration of mud in zones less than 1 mm thick along fault surfaces, suggesting clay-mineral elutriation by escaping pore-waters along the faults during failure. Wet-sediment deformed horizons vary considerably in structural style and thickness (Table 2), but invariably involve the finest grained, thinbedded turbidites within the Point Leamington Formation. The following distinct styles of wetsediment deformation, within a continuum of deformation, are recognized (Table 2): (i) coherent folded beds, sometimes showing isoclinal folding (Fig. 6A) and in 3-D exposure typically reveal that the main (first) tectonic cleavage is at

an acute or greater angle to the fold axial surface (Fig. 6B); (ii) rotational slide folds (slumps sensu stricto) similar to those described by Laird (1968) (Fig. 6C); (iii) semi-coherent folded beds, typically with only up to a few metres of contiguous bedding (Fig. 6D), that break down laterally into; (iv) chaotic bedding (Fig. 6E) involving detached fold hinges (Fig. 6F) and sliver-like lenses of sediment; (v) boudinaged silty and sandy beds (Fi~. 6G) that may be associated with brittle deformation as early, wet-sediment faults (Fig. 6H), and (vi) early, wet-sediment faults, best seen as small-scale faults that may be closely-spaced normal faults (Fig. 6I), conjugate normal fault pairs and/or reverse faults. Within any of the above wet-sediment deformation horizons, or unrelated to such horizons in otherwise undisturbed silty and sandy beds, local liquefaction and fluidization structures may occur, such as small-scale elastic dykes and convolute laminae. Also, wet-sediment coherent folded beds may show local chaotic zones where the deformed sediments occur as well-rounded to rounded clasts (Fig. 6J). Cowan (1985) describes similar structures from Mesozoic/Cenozoic melanges of North America. The wet-sediment deformation horizons may be considered in two discrete populations--those whose genesis is demonstrably related to superficial sediment sliding (and slumping) because of their relationship with immediately overlying and underlying deposits (e.g. erosional surfaces ~.nd sediment draping), and those horizons that are sandwiched between apparently normallybedded deposits, only showing internal deformation, thereby giving no indication of the depth of burial at the time of sediment failure. Neither group appears to be associated with a specific style of wet-sediment deformation, as described in the above six categories. Sediment slides

Using a string grid of metre squares overlying the wet-sediment deformed horizons, it is possible to demonstrate that some of the horizons must have developed as superficial, sometimes multiple, gravity-induced slide events, for example immediately west of Passage Rocks (Fig. 7). The slides shown in Fig. 7 show a wide range of wetsediment deformation styles (enlarged area shown in Fig. 8), without a preferred orientation for the slide fold axes. This example is partially eroded by a prominent pebbly mudstone containing angular to rounded fossiliferous limestone and black, pyrite-rich, cherty mudstone clasts set in a black mudstone matrix (Fig. 9), interpreted as a debris flow. The characteristics of the debris

W e t - s e d i m e n t deformation in the L e a m i n g t o n Formation

22I

TABLE 2. Characteristics of wet-sediment deformation horizons in the Point Leamington Formation TYPICAL

INTERNAL STRUCTURE

N A T U R E OF

THICKNESS

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Essentially 1-10

continuous

layers,

n<~n-cylinderical,

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irregular,

folds

m

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hinges & limbs: ghost

LAYERS

irregular

-

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- smooth

smooth

non-erosive

stratigraphy

commonly

Rounded to angular

defineable

clasts

&

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BRECCIATED

- smooth

E r o s i v e - planar 1-10

BALLED

Irregular nen-erostva

detached

layers,

SEMI-COHERENT

CHAOTIC

UPPER S U R F A C E

- planar

m

LAYERS

FOLDED

Erosive

SURFACE

m

of a t t e n u a t e d

& deformed

non-erosive swirly

texture

Attenuated, BOUDINAGED LAYERS

- planar

layers,

LAYERS

0.1-1

m

to deformed

stretched

layers

layers,

some with well-developed

Pinch & swell

Pinch & swell

rod structure

Brittle f a i l u r e , scale FAULTED LAYERS

s e e n as s m a l l (some conjugate): Draped

outcrop

scale

of l a r g e Note: all styles

best

normal ~ults

-

of w e t - s e d i m e n t

deformation

may

occur

scale

prevents early

in s a m e

recognition

faults

horizon

flow are unique to the lower parts of the Point Leamington Formation since it contains extraformational clasts or olistoliths. The intimate association of the slides and debris flow favours a common triggering mechanism for the sediment failure. Overlying sediment slide horizons, bedding tends to be irregular and includes sediment draping, suggesting that post-failure sedimentation infilled residual depressions and smoothed uneven topography created by the slides (Fig. 10). Given the complex tectonic deformation history of the Lower Palaeozoic rocks in the Notre Dame Bay area, it has not been possible to demonstrate that the common subtle variations in bedding dip and strike are due to sediment draping across low-relief, relatively small-scale uneven topography: therefore, apart from obvious sediment accommodation immediately above slides, it is impossible to fully appreciate just how uneven the topography may have been. Wet-sediment deformation formed at indeterminate burial depth

Within the Point Leamington Formation, there are many horizons of wet-sediment deformation

& pass laterally

into each other

abruptly

or g r a d a t i o n a l l y

that show no clear indication of the depth of burial of the sediments during failure. Such zones may have formed at or near the palaeo-mud-line, or they may have developed at considerable depth within over-pressured sediments. Typically, such horizons display many of the 14 characteristics listed above, apart from those that are diagnostic of superficial sediment failure (i.e. 11, 12). Within the formation, there are thin (typically, less than 1 m thick) horizons of coherently folded silty sandstones in which the fold style, amplitude and wavelength vary considerably over decimetres (Fig. 11). Such pre-cleavage deformation could be explained by wet-sediment, intrafolial slip at indeterminate depth of burial, possibly by creep deformation of slope sediments as described by Hill et al. (1982) from the Canadian Beaufort Sea continental slope. Other pre-S2 cleavage structures include small-scale thrusts and associated imbricates (Fig. 12). Fig. 13, from a promontory on the east of Southwest Arm, opposite Passage Rocks, is an example of an horizon of wet-sediment deformation of uncertain genesis. The wet-sediment deformation occurs in a zone about 10 m thick, is sandwiched between apparently normally-bedded, undisturbed thin-bedded, fine-grained tur-

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K. T. Pickering

FIG. 6. Types of wet-sediment deformation: (A) Coherent folded strata. Note isoclinal fold immediately left of hammer (shaft is 35 cm long), and irregular, laterally discontinuous, strata to right of fold. Beds young to left. West coast of Southwest Arm near Passage Rocks. (B) 3-D exposure of fold hinge zone showing non-axial planar cleavage. Fold axial surface projects obliquely from plate towards left, whereas $2 penetrative slaty cleavage projects almost perpendicularly from plane of plate. 5 cm scale. West coast of Southwest Arm: This folded bedding locally becomes disrupted within a few metres laterally. (C) Rotational slide fold (slump) in thin-bedded silty turbidites, Passage Rocks, Southwest Arm. Hammer for scale in centre background. Beds young to left. 'Normal' bedding, almost vertical, in background. Incipient roll-over structure in coherently-folded bedding towards centre of plate. (D) Semi-coherent folded beds. Lens cap for scale. Note discontinuous bedding. West coast of Southwest Arm near Passage Rocks. (E) Semi-coherent folded beds (background), showing irregular (non-cylindrical) folds that abruptly break down into chaotic bedding (foreground), with detached fold hinges and fold limbs. Hammer for scale. East coast of Southwest Arm. (F) Close-up of chaotic bedding, showing detached fold hinge. Lens cap for scale. (G) Complexly deformed and essentially coherent bedding, showing boudinage of sand-rich beds (especially to right of hammer shaft). Note decollement surfaces cutting down through beds that (overall) young to left. West coast of Exploits Bay, immediately north of Lawrence Harbour. (H) Close-up of semi-coherent folded bedding to show geometric association of folds, boudinage and normal faults. Note fluid-like disruption of laminae within the sand-rich beds and mud injection into fold limb. 5 cm scale. West coast of Southwest Arm near Passage Rocks. (I) Early small-scale faults, typically showing conjugate pairs. Note fluid-like disruption of laminae. West coast of Southwest Arm near Passage Rocks. (J) Coherent folded bedding showing local chaotic (typically well-rounded) sediments (immediately left of hammer shaft). West coast of Southwest Arm.

Wet-sediment deformation in the Leamington Formation coherent

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FIG. 7. Scale-drawing of lower part of wet-sediment deformation horizon on west coast of Southwest Arm, near Passage Rocks. Note range of deformation styles, divergent fold axial directions with plunge given, erosional surfaces cutting down to right, and debris flow deposit at top. Close-up of small-scale folds and faults drawn on left of this diagram shown in Figs 6(H) and 8.

bidites, upper and lower bounding surfaces appear welded and show no obvious truncation of the deformed sediments, and displays a variety of deformation styles. Non-cylindrical folds with curvilinear hinge zones/axes, as coherently folded beds, abruptly disintegrate laterally into chaotic, brecciated, bedding (Fig. 13). Also, synformal and antiformal structures show terminations and bifurcations (Fig. 13). In some examples of wet-sediment deformation

in this category, the deformed sediments show asymmetric to monoclinal folds, with axial surfaces sub-parallel, or at somewhat larger acute angles, to the approximately parallel-sided bounding surfaces of the deformation zone. The geometry of such folds could be explained by wetsediment folding in a shear zone generated by inter-stratal slip, at unspecified depths of burial, possibly in over-pressured sediments. Invariably, such horizons show coherently folded sediment

FIG. 8. Example of range of wet-sediment deformation structures occurring in a small area: chaotic bedding at top; semi-coherent folds, with extremely attenuated limbs; boudinage and normal faulting at bottom of plate. This plate taken in wet-sediment deformation horizon shown in Fig. 7.

zz6

K. T. Pickering

FIG. 9. Pebbly mudstone (interpreted as a debris flow deposit) occurring at the top of the wet-sediment deformation horizon shown in Fig. 7. Clasts are mainly black, pyrite-rich, cherty mudstones or light-coloured carbonates. that changes laterally into semi-coherent and/or chaotic beds, i.e. deformational style varies over metres. Wet-sediment deformation in thin-section A detached fold hinge zone was collected from the sediment slide zone shown in Fig. 7 since it showed: (i) non-axial planar cleavage (at about 20 ° to the fold axial plane); (ii) apparent wetsediment flow structures as 'swirly' lamination around the fold; (iii) early, pre-cleavage, curved small-scale faults, radially disposed about the axial plane that die out vertically; (iv) very small-

e

scale mud injections along the pre-cleavage fault surfaces (Fig. 14A, B). A series of thin-sections, perpendicular to the axial plane, were sampled. In thin-section, the 'swirly' lamination is not associated with any cataclastic texture; the only 'tectonic' fabric being re-oriented clay minerals and minor slip and/or, pressure solution along the penetrative $2 cleavage. In thin-sections from the limbs of the fold, the penetrative $2 cleavage is at 20-30 ° to the fault surfaces, across which it cuts (Fig. 15A), thereby demonstrating that the faults pre-date the cleavage. Furthermore, most of the fault surfaces are associated with very thin (less than 0.5 mm thick) clay-rich layers that

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FIG. 10. Lenticular/irregular bedding immediately overlying a wet-sediment deformation horizon (right side of plate with hammer for scale about 45 cm from top of horizon). Note 'normal' bedding younging to left of plate, above irregular bedding.

Wet-sediment deformation in the Leamington Formation

227

FI6. 11. Small-scale example of intrafolial wet-sediment deformation that developed at unspecified depths of burial. Beds young to left. 15 cm scale. Note variable wavelength and amplitude folds between two prominent light-coloured sandy beds. This deformation locally becomes chaotic and passes laterally into apparently undisturbed bedding.

appear to have roots in mudstone laminae (Fig. 15B), i.e. mudstone clastic injections. Elongate silt-grade clasts appear parallel with bedding, but within the mudstone injections the long axes have been rotated through about 90 ° to parallel the fault surfaces. Where the faults juxtapose similar lithologies of very fine/fine-grained sandstone, the clay-rich injection is generally barely discernible or appears absent so that the fault appears annealed or welded. The thin-sections support the interpretation of wet-sediment ductile and brittle deformation structures that pre-date the over-printing $2 penetrative cleavage. The faults cut the swirly lamination and therefore post-date the wetsediment folding. The mudstone injections along the fault surfaces suggest that the faults formed in wet sediments. The transition from ductile to brittle deformation in the folds could be explained by a range of processes, including a change in the rheology of the sediments as pore waters were expelled during folding and/or, an increase in the strain rate. Whatever the origin of the faults, their radial disposition about the fold axial surfaces suggest that they are genetically related to the folding, albeit at a slightly later time. Wet-sediment fold: structural considerations

Bedding, the main penetrative $2 cleavage (see below) and fold data were plotted on equal area, lower hemisphere, projections (Figs 16, 17 and

20). An $1 cleavage is only recognized in the Lower Exploits GrOup (New Bay Formation) by Helwig (1967). The $2 cleavage is the first defined cleavage in the Point Leamington Formation. Wet-sediment fold axes and fold axial planes were rotated (Fig. 17) to: (i) remove the effects of the F2 tectonic folding, using the Beaver Brook Synclinorium fold axial plunge of 20/030, with an approximately vertical axial plane; (ii) restore bedding to horizontal. Wet-sediment fold data were rotated individually with respect to local 'normal' bedding dip, not from an average bedding dip value on the major fold limbs. Since the effects of any tectonic pre-F2, and post-F2 folding, appear relatively subtle, or are of uncertain extent, no account has been taken of such structures for data rotation (see section below for chronology of deformation events and their characteristics). Within the Point Leamington Formation, the $2 penetrative cleavage typically is vertical to sub-vertical (Fig. 16B) and commonly almost coincident, or at an acute angle, with bedding. Since many of the wet-sediment folds are isoclinal to tightly folded, with axial surfaces parallel/subparallel to bedding, the bedding-cleavage relationship in essentially 2-D outcrops appears to show $2 as an axial planar fabric to the wetsediment folds. However, 3-D exposures commonly show that $2 is non-axial planar to the wet-sediment folds (Fig. 6B). Figure 17 shows rotated wet-sediment fold

2z8

K. T. Pickering

FIG. 12. Wet-sediment (note pre-cleavage) thrust cutting through undisturbed bedding (foreground) from top right to bottom left of plate. Lens cap for scale (centre). Beds above thrust surface show incipient roll-over and small-scale imbricate wedge (immediately above lens cap). West of Southwest Arm.

data grouped into data sets from the west coast (Fig. 17A, B) and east coast (Fig. 17C, D) of Southwest Arm, West Arm (Fig. 17E, F), and Exploits Bay (Fig. 17G, H). The data have been plotted to show the down-plunge sense of fold axial surface rotation, however, without a rigorous palaeoslope determination using either the mean axis or separation arc methods, as outlined by Woodcock (1979). The orientation of the fold axes and axial surfaces, from the equal area

projections, as previously noted by Helwig (1967, 1970), suggest palaeoslopes dipping towards the southern sector, although the data from the eastern side of Southwest Arm are equivocal (Fig. 17C, D). Data from Exploits Bay (Fig. 17G, H) suggest a basin-slope dipping towards the ESE. Palaeocurrent data from Helwig (1967, 1970) and this study (Fig. 18) indicate a regional basinslope generally dipping southwards, in agreement with the interpreted palaeoslope deduced from

229

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wet-sediment fold data. The equivocal fold data from the eastern side of Southwest Arm are matched by anomalous palaeocurrents (Fig. 18) that suggest a local eastward-dipping slope (also noted by Helwig 1967, 1970). It should be noted, however, that the orientation of the wet-sediment fold axial planes is inconsistent with palaeocurrent data, since a local eastward-dipping slope, theoretically, should be associated with poles to axial planes concentrating in the NW, not SE, quadrant (Fig. 17D). This problem is considered in the discussion section (below).

Chronology of deformation events Table 3 summarizes the main deformation events found within the Point Leamington Formation and related rocks, and compares this scheme with those erected by previous researchers in this succession (Helwig 1967, 1970) and other approximately contemporaneous successions in Notre Dame Bay (Horne 1968; Eastler 1971; Nelson 1981; Karlstrom et al. 1982, 1983; Arnott 1983a, b). The studies by Eastler (1971) and Karlstrom et al. (1982), in contrast to the others listed above, were in the Lower Palaeozoic to the east of the Reach Fault, interpreted by McKerrow & Cocks (1977, 1986), and by Arnott et al. (1985), as the suture across which the Iapetus Ocean finally closed. Therefore, while apparently similar deformation events are identified both to the east and

west of the Reach Fault (Karlstrom et al. 1982), such events may prove diachronous or un-related. However, this study accepts the arguments of Karlstrom et al. (1982) that the Reach Fault, as exposed today, is a late stage strike-slip fault across which there is structural and, probably, stratigraphic continuity; i.e. any major postLlandovery suture must be located farther east, such as the Dover-Hermitage Bay Fault zone. In agreement with Helwig (1967, 1970), the first recognized phase of wet-sediment deformation is given an 'F' designation because the folds occur throughout the Point Leamington Formation, and other broadly contemporaneous successions from Notre Dame Bay, and they are volumetrically important within the formation. The wet-sediment superficial slide and slump folds are designated FO (also by Helwig 1970) and possess the attributes described above. The orientation of the FO fold axes and axial surfaces, in relation to the southerly-dipping regional basin slope, interpreted from palaeocurrents, favours gravity-controlled sediment failure. The F1 phase of folding (attributed to D1) was recognized by Helwig (1970), previously designated F2 by Helwig (1967), with an associated weak axial planar cleavage in mud-/silt-grade rocks. Definite F 1 mesoscopic structures are only recognized in the structural 'block' between Exploits Bay in the east, and east of Southwest Arm in the west (the Paradise Block--Helwig 1967). An F1 fold phase, producing mesoscopic

230

K. T. Pickering

10

FIG. 14. Polished slabs of wet-sediment deformation fold hinge west of Southwest Arm near Passage Rocks. (A) Fluid-like disruption of primary sedimentary structures in silty beds/laminae, radial fault pattern, small-scale mud injections along fault planes, and non-axial planar slaty cleavage (vertical). (B) Limb of fold shown in (A) (reverse side of slab) showing small-scale faults and fluid-like disruption of laminae. Cleavage at oblique angle to faults. See Fig. 15.

Wet-sediment deformation in the Leamington Formation

231

FIG. 15. Thin-sections from wet-sediment fold hinge shown in Fig. 14. (A) Early normal fault cut by penetrative slaty cleavage (parallel to bar); bedding is parallel with lower edge of plate. (B) Early normal fault with mud (clay-rich) injection that intruded vertically up fault surface from underlying mud. Bedding is parallel with lower edge of plate.

and macroscopic open folds and monoclinal flexures, could explain the development of the basin, with a downslope basement high to the ESE, in which most of the Point Leamington Formation accumulated. Palaeocurrents and facies analysis (unpublished data) suggest that the present-day Beaver Brook synclinorium was developing during the deposition of the Point Leamington Formation and controlled the locus of deep marine sedimentation in this basin (see also Helwig 1970). Furthermore, in a recent study of the Ordovician-Silurian of the New Bay area and of New

World Island, Nelson (1981) and Arnott (1983a, b), respectively, identified N N W dipping D1 thrusts that were developing during sedimentation and controlling the location of small slope basins. Such D1 thrusts, of post-Caradoc (? Caradoc) to Llandovery age may have been active farther to the west in the Point Leamington Formation basin(s) such as the early New Bay Fault, producing the topographic highs to the west and east of the present-day Beaver Brook synclinorium, now defined by the anticlinoria of Wild Bight Group and Lower Exploits Group, respectively.

232

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234

F2 mesoscopic and macroscopic folds are common within the Point Leamington Formation (Fig. 19), and refold the D1, F1, deformation. In many cases, the coincidence of F1 (non-rotated) and F2 fold axes (Fig. 20A, B) and axial planes, makes recognition difficult, especially in the absence of convincing differentiating features (as listed above). The post-Acadian, Silurian-Devonian or later, D3 deformation generated open folds, refolded $2, and produced either spaced or crenulation cleavage, and kink bands (Table 3). Since D3 involved hard-rock deformation and post-dates the Acadian orogeny, it is not considered further in this paper.

Some of the wet-sediment fold deformation that occurred at indeterminate depths of burial, interpreted as gravity-controlled sliding from the fold data plotted on stereographic projections (Fig. 17), may have occurred in response to developing thrusts at depth. Such early thrusting could have led to wet-sediment failure in overpressured zones at depth, followed by gravitycontrolled intrafolial slide deformation. Although this scenario must be considered as speculative, it could explain some of the arguments recently raised by Karlstrom et al. (1982) who dispute the early, superficial, wet-sediment slide interpretation of Helwig (1967, 1970), and this study, instead proposing an F1 phase of thrusting (see discussion below). The D2 deformation, producing the regional penetrative $2 cleavage, was the main 'tectonic' deformation to affect the Lower Palaeozoic in the Dunnage Zone. The main attributes of F2 folds are given in Table 3. D2, with associated F2 and $2, was the 'Acadian' orogeny, occurring during the Silurian. D2 deformation involved lithified sediments, and F2 folds commonly show welldeveloped fanned $2 cleavage in the hinge zones.

Discussion Helwig (1967, 1970) provided the first detailed insight into the wet-sediment fold deformation in the Point Leamington Formation, interpreting it as superficial slope failure 'slump folds'. The criteria used by Helwig (1970, p. 174) for recognizing 'slump folds' and associated struc-

TABLE 3. Summary of deformation history for the Lower Palaeozoic succession in Notre Dame Bay. Deformation terms most commonly used by other researchers in Notre Dame Bay are shown for comparison. See text Jbr explanation. Siluro-Devonian deformation may extend into the Carboniferous ~=

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235

tures include many of the attributes used in this paper for characterizing wet-sediment deformation. Both Helwig (1967, 1970) and this study favour gravity-controlled wet-sediment deformation for FO and probably F 1 folds (sliding and slumping), although, in contradistinction, this study suggests that some of the deformation may have occurred under considerable but indeterminate depths of burial in over-pressured zones, either gravity-controlled or thrust-related. Evidence in favour of either interpretation is sometimes equivocal. Recently, prior to this study, Karlstrom et al. (1982) disputed Helwig's 'slump fold' interpretation. In a study of Ordovician-Silurian sedimentary rocks east of the Reach Fault, Karlstrom et al. (1982, p. 2338) state that 'with the exception of convolute bedding in turbidites, we know of no proven soft-sediment deformation,' noting that 'ductile flow of hard rocks can produce the same structures' (p. 2329), as observed in superficial sediment sliding/slumping. Based on their observations and interpretations east of the Reach Fault, Karlstrom et al. (1982) reinterpret Helwig's (1967, 1970) slump folds as complex penetrative orogenic, generally isoclinal, commonly intrafolial folds formed during regional thrusting and folding events. Their F1 folds are overprinted by F2 (equivalent to F2 in this study) to produce coaxial and mushroom interference patterns. This feature and the sub-horizontal enveloping surfaces of F2 were used to deduce that F1 folds were recumbent prior to F 2 - - a geometrical argument that may be applied to the Point Leamington Formation. Furthermore, in addi-

FIG. 19. Tectonic (lithified sediment) anticline with well-developedaxial planar slaty cleavage, from west coast of Southwest Arm. F2 fold.

236

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FIG. 20. Lower hemispherical equal area projection of non-rotated wet-sediment and tectonic fold axis data from New Bay area. Note that wet-sediment fold axes appear to have been rotated towards plane of $2 slaty cleavage.

tion to other considerations, Karlstrom et al. (1982) recognized a zone of F1 folds separating Silurian turbidites from subaerial volcanics on the Port Albert Peninsula, and interpreted the F1 structures as related to macroscopic thrusting. Following open scientific debate about the proposed early thrusting (and by inference thrustrelated F1 folds) between Currie et al. (1983) and Karlstrom et al. (1983), the origin of the F1 structures has remained unresolved. The above arguments are dealt with in some depth because Karlstrom et al. (1982) demonstrate a similar deformation history across eastern Notre Dame Bay, with a general continuity of macroscopic F2 folds throughout the area. Thus, the nature and origin of any apparently similar F 1 structures across the area should be considered collectively, although various, geologically compatible, interpretations may result (as in this study). This study interprets pre-F2 (D1) structures as wet-sediment deformation, grouping the demonstrably superficial sediment slide folds as FO, and the slide folds that appear to have developed as intrafolial deformation at unspecified depths as F1. F0 folds and associated structures are always interpreted as gravity-controlled wetsediment deformation. F1 folds and related structures generally appear to be gravity-controlled (intra-slope and downslope) wet-sediment deformation, based on the relationship between palaeocurrents and the geometrical considera-

tions from the stereographic projections. However, such a simplistic interpretation appears untenable for the fold data from the eastern side of Southwest Arm (Figs 17C, D). The plunge of fold axes, with the associated down-p!unge sense of rotation of the axial planes, is ambiguous without clearly-defined populations, and the axial planes (generally dipping towards the NNE), if the folds were gravity-controlled, favours a southeastward-dipping basin slope. Palaeocurrents from this section, containing mainly very thin-bedded silty turbidites, suggest a local westward-dipping basin-slope (Fig. 18). In the Point Leamington Formation, on the east of Southwest Arm, early, syn-sedimentary tectonic reverse faulting or thrusting, along westward-dipping thrusts could explain: (i) the formation of a topographic (? submarine) 'high' on the regional southward/southeastward-dipping basin-slope, from which very thin-bedded, silty turbidites were shed; and (ii) many wetsediment FI fold axial planes showing a northwestward-dip, associated with fold axes that appear to form two populations, suggesting two 'slopes' dipping 180 ° apart. The two fold populations could be resolved as either gravity-controlled slide folds (superficial or at depth) or thrust-related folds: in both cases, deformation is believed to have affected wet-sediments, but in the latter interpretation the thrusting would have involved sediments in over-pressured zones, at depth. Today, the lower Caradoc black shales

Wet-sediment deformation in the Leamington Formation and overlying Point Leamington Formation, east of Southwest Arm, are upfaulted against the lower Exploits Group (Tea Arm Volcanics, New Bay Formation etc.--see Table 1) along the New Bay Fault (Fig. 3). It is possible that a precursor of this fault, and other parallel/sub-parallel (unnamed) faults, active during deposition of the Point Leamington Formation, may have been northwestward to westward-dipping thrusts. The projected extreme westward flexure (through the northern part of Southwest Arm), then northward flexure (through Little Northwest Arm) of the New Bay Fault (Fig. 3) is consistent with this fault having been a thrust as postulated here. Indeed, while Helwig (1967, p. 127) favours an interpretation of an early New Bay Fault as a normal fault during sedimentation of the Point Leamington Formation, he states (Helwig, pers. comm.) that the fault could have been a thrust during late Caradoc-Ashgill times. Certainly, the New Bay Fault is an early structure since it is offset by later D3 transcurrent faults (Helwig 1967). Furthermore, this hypothesis is consistent with recent research in essentially contemporaneous, nearby successions that have suggested early westward to north-westward-dipping thrusts (Nelson 1981; Karlstrom et al. 1982; Arnott 1983a, b) that may have been active during sedimentation (Nelson 1981 ; Arnott 1983a, b).

Conclusions The Point Leamington Formation is interpreted as a deep marine succession that accumulated in a series of small fault-bounded and fault-controlled basins at an active margin with considerable sinistral oblique-slip tectonics. The abundance of wet-sediment deformation, either as superficial slides and slumps, or sub-surface failure, together with the essentially fine-grained and thin-bedded siliciclastics, suggests sedimentation on a basin slope. Palaeocurrents suggest a regional slope towards the south to SE, with topographic highs somewhat oblique to the strike of the slope, acting as dams behind which sediments accumulated and, perhaps, acting as deflectors to sediment flows moving down the

237

regional slope. The conglomeratic, lens-shaped, packets of deep marine sediment gravity flows show erosional surfaces, sometimes through tens of metres of underlying beds, into the typical Point Leamington Formation lithologies and are interpreted as infilled slope gullies and canyons (unpublished data). Thus, the Point Leamington Formation is interpreted as a complex basin slope and terraces, probably thrust fault-controlled, and cut by submarine gullies and canyons. The wet-sediment deformation within the Point Leamington Formation shows the complex interaction of tectonics and sedimentation, emphasizing the problems in unravelling a realistic chronology of sediment deformation events. Depositional slopes frequently appear to have become unstable, mainly as a result of tectonic processes. However, as noted by Hill et al. (1982), creep, defined as strain with time under constant loading conditions, can produce small strain rates that may lead to significant cumulative displacements over geologically short time intervals. Thus, slope creep processes could have generated some of the wet-sediment folds and associated deformation (both on a micro- and macroscopic scale) observed in the Point Leamington Formation and interpreted to have formed at unspecified depths of burial. Many of the tectono-sedimentary features of the Caradoc-Llandovery rocks of Notre Dame Bay can be found in the offshore basins of the Californian Borderland (Nardin 1983; Aydin & Page 1984), or in arc-related accretionary prisms (Aoki et al. 1983). Clearly, the Lower Palaeozoic plate tectonic history of this area was very complex with phases of crustal transtension superimposed on essentially plate collisions, probably similar to the marginal basins of the Western Pacific or in the Banda Arc today. ACKNOWLEDGEMENTS: Grateful appreciation for the time spent in reviewing various versions of this manuscript must go to Drs Harold Williams, Scott Swinden, Stuart McKerrow and James Helwig. However, the ideas expressed in this paper are the author's own and not necessarilythose shared by these reviewers. The Royal Society is thanked for its generous financial support to undertake this research, with a University Research Fellowship.

References AOKI, Y., TAMANO, T. & KATO, S. 1983. Detailed structure of the Nankai Trough from migrated seismic sections. In: WATKINS,J. S. & DRAKE,C. L. (eds), Studies in Continental Margin Geology. Am. Assoc. Petrol. Geol. Mere. 34, 309-22. ARNOTr, R. J. 1983a. Sedimentologyof Upper Ordovi-

--

cian-Silurian sequences on New World Island, Newfoundland: separate fault controlled basins? Can. J. Earth Sci. 20, 345-54. 1983b. Sedimentology, structure and stratigraphy of northeast New World Island, Newfoundland. D.

Phil. thesis, Oxford Univ.

238 --,

K. T. Pickering

MCKERROW, W. S. & COCKS, L. R. M. 1985. The tectonics and depositional history of the Ordovician and Silurian rocks o f Notre Dame Bay, Newfoundland. Can. J. Earth Sci. 22, 607-18. AYDIN, A. & PAGE, B. M. 1984. Diverse PlioceneQuaternary tectonics in a transform environment, San Francisco Bay region, California. Geol. Soc. Am. Bull. 95, 1303-17. BALLY, A. W. (ed.) 1983. Seismic expression of structural styles. Am. Assoc. Petrol. Geol., Studies in Geology Ser. 15, 1, 2 & 3. COCKS, L. R. M. & FORTEY, R. A. 1982. Faunal evidence for oceanic separations in the Palaeozoic of Britain. J. geol. Soc. London, 139, 465-78. COLMAN-SADD,S. P. & SWINDEN, H. S. 1984. A tectonic window in central Newfoundland? Geological evidence that the Appalachian Dunnage Zone may be allochthonous. Can. J. Earth Sci. 21, 134967. COWAN, D. S. 1985. Structural styles in Mesozoic and Cenozoic melanges in the western Cordillera of North America. Bull. geol. Soc. Am. 96, 451-62. CRANS, W., MANDL, G. & HAREMBOURE, J. 1980. On the theory of growth faulting: a geomechanical delta model based on gravity sliding. J. petrol. Geol. 2, 265-307. CURRIE, K. L., PICKERILL, R. K. & PAJARI, G. E. Jr. 1983. Structural interpretation of the eastern Notre Dame Bay area, Newfoundland: regional postMiddle Silurian thrusting and asymmetrical folding: Discussion. Can. J. Earth Sci. 20, 1351-2. DEAN, P. L. 1978. The volcanic stratigraphy and metallogeny of Notre Dame Bay, Newfoundland. Memorial University of Newfoundland, Geology Report, 7, 204 pp. DOYLE, L. J. & PILKEY, O. H. (eds) 1979. Geology of Continental Margins. Soc. econ. Paleont. Min., Tulsa, 27. EASTLER,T. E. 1969. Silurian geology of Change Islands and eastern Notre Dame Bay, Newfoundland. In ." KAY, M. (ed.), North Atlantic-Geology and Continental Drift. Am. Assoc. Petrol. Geol. Mem. 12, 425-32. -1971. Geology o f Silurian rocks, Change lslands and easternmost Notre Dame Bay, Newfoundland. Ph.D. thesis, Columbia Univ., New York. HARLAND, W. B., COX, A. V., LLEWELLYN, P. G., PICKTON, C. A. G., SMITH, A. G. & WALTERS, R. 1982. Subdivisions of Phanerozoic Time. Card publ. by British Petroleum Co. p.l.c. & Cambridge Univ. Press. HELWIG, J. 1967. Stratigraphy and structural history of the New Bay area, north-central Newfoundland. Ph.D. thesis, Columbia Univ., New York. 1969. Redefinition of the Exploits Group, Lower Palaeozoic, northeast Newfoundland. In: KAY, M. (ed.), North Atlantic--Geology and Continental Drift. Am. Assoc. Petrol. Geol. Mem. 12, 407-21. 1970. Slump folds and early structures, northeastern Newfoundland Appalachians. J. Geol. 78, 17287. HILL, P. R., MORAN, K. M. & BLASCO, S. M. 1982. Creep deformation of slope sediments in the

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Canadian Beaufort Sea. Geo-Marine Letters, 2, 163-70. HORNE, G. S. 1968. Stratigraphy and structural geology of southwestern New Worm Island area, Newfoundland. Ph.D. thesis, Columbia Univ., New York. -& HELWm, J. 1969. Ordovician stratigraphy of Notre Dame Bay, Newfoundland. In- KAY, M. (ed.), North Atlantic--Geology and Continental Drift. Am. Assoc. Petrol. Geol. Mem. 12, 388-407. JACOBI, R. D. & WASOWSKI, J. J. 1985. Geochemistry and plate-tectonic significance of the volcanic rocks of the Summerford Group, north-central Newfoundland. Geology, 13, 126-30. KARLSTROM,K. E., VANDER PLUIJM, B. A. & WILLIAMS, P. F. 1982. Structural interpretation of the eastern Notre Dame Bay area, Newfoundland: regional post-Middle Silurian thrusting and asymmetrical folding. Can. J. Earth Sci. 19, 2325-41. & --, 1983. Structural interpretation of 'the eastern Notre Dame Bay area, Newfoundland: regional post-Middle Silurian thrusting and asymmetrical folding: Reply. Can. J. Earth Sci. 20, 1353-4. KAY, M. 1976. Dunnage melange and subduction of the protacadic ocean, northeast Newfoundland. Geol. Soc. Am. Spec. Paper 175. KLAN, B. F., DEAN, P. L. & STRONG, D. F. 1981. Regional geology of the Central Volcanic Belt of Newfoundland. In: SWANSON,E. A. & STRONG,D. F. (eds), The Buchans Orebodies: Fifty years of Geology and Mining. Geol. Assoc. Canada, Spec. Paper, 22. LAIRD, M. G. 1968. Rotational slump scars in Silurian rocks, western Ireland. Sedimentology, 10, 111-20. MCKERROW, W. S. & COCKS, L. R. M. 1977. The location of the Iapetus suture in Newfoundland. Can. J. Earth Sci. 14, 488-95. & -1986. Oceans, island arcs and olistostromes : the use of fossils in distinguishing sutures, terranes and environments around the Iapetus Ocean. J. geol. Soc. Lond. 143, 185-91. --, LAMBERT, R. ST-J. & COCKS, L. R. M. 1985. The Ordovician, Silurian and Devonian Periods. In: SNELLING, N. J. (ed.), The Chronology of the Geological Record. Geol. Soc. London Mem. 10, 73-80. MALTMAN, A. 1984. On the term 'soft-sediment deformation'. J. struct. Geol. 6, 589-92. MANDL, G. & CRANS, W. 1981. Gravitational gliding in deltas. In: MCCLAY, K. & PRICE, N. J. (eds), Thrust and Nappe Tectonics. Spec. Publ. Geol. Soc. London, 9, 41-54. MURPHY, F. C. & HUTTON, D. H. W. 1986. Is the Southern Uplands of Scotland really an accretionary prism? Geology, 14, 354-57. NARDIN, T. R. 1983. Late Quaternary depositional systems and sea level change--Santa Monica and San Pedro Basins, California Continental Borderland. Am. Assoc. Petrol. Geol., Bull. 67, 1104-24. NELSON, K. D. 1981. Melange development in the Boones Point Complex, north-central Newfoundland. Can. J. Earth Sci. 18, 433-42. NEUMAN, R. E. 1984. Geology and paleobiology of islands in the Ordovician Iapetus Ocean: Review -

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Wet-sediment deformation in the Leamington Formation and implications. Bull. geol. Soc. Am. 95, 1188201. Ross, R. J. & NAESER, C. W. 1984. The Ordovician time scale--new refinements. In: BRUTON, D. L. (ed.), Aspects of the Ordovician System. Universitetsforlaget, Oslo, Norway, 5-10. SAXOV, S. & NIEUWENHUIS, J. K. (eds) 1982. Marine Slides and Other Mass Movements. Nato Conf. Ser. IV: Marine Sciences, Plenum Press, New York. SOPER, N. J. & HUTTON, D. H. W. 1985. Late Caledonian sinistral displacements in Britain: implications for a three-plate collision model. Tectonics, 3, 781-94. STANLEY, D. J. & MOORE, G. T. (eds) 1983. The Shelfbreak: Critical Interface on Continental Margins. Soc. econ. Paleont. Min. Tulsa, 33. STOW, D. A. V. & PIPER, D. J. W. (eds) 1984. FineGrained Sediments: Deep-Water Processes and Facies. Spec. Publ. Geol. Soc. London, 15. WATKINS, J. S. & DRAKE, C. L. (eds) 1982. Studies in Continental Margin Geology. Am. Assoc. Petrol. Geol. Mem. 34. WATSON, M. P. 1981. Submarine fan deposits of the Upper Ordovician-Lower Silurian Milliners Arm Formation, New World Island, Newfoundland. D.Phil. thesis, Oxford Univ. WHALEN, J. B., CURRIE, K. L. & VAN BREEMEN, O. 1987. Episodic Ordovician-Silurian plutonism in the Topsails igneous terrace, Western Newfoundland. Trans. Roy. Soc. Edinburgh, in press. WILLIAMS, H. 1962. Botwood (west half) map-area, Newfoundland. Geol. Surv. Canada, Paper 62-9.

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, 1963. Twillingate map-area, Newfoundland. Geol. Surv. Canada, Preliminary Ser. Map 60-63. --, 1964a. The Appalachians in northeastern Newfoundland--a two-sided symmetrical system. Am. J. Sci. 262, 1137-58. 1964b. Botwood map-area, Newfoundland. Geol. Surv. Canada, Preliminary Set. Map 60-1963. (compiler), 1978. Tectonic-lithofacies map of the Appalachian orogen. Memorial University of Newfoundland, St John's, Nfld., Map No. 1. -1979. Appalachian orogen in Canada. Can. J. Earth Sci. 16, 792-807. 1984. Miogeoclines and suspect terranes of the Caledonian-Appalachian Orogen: tectonic patterns in the North Atlantic region. Can. J. Earth Sci. 21, 887-901. & HATCHER, R. D. Jr. 1982. Suspect terranes and accretionary history of the Appalachian orogen. Geology, 10, 530-36. & -1983. Appalachian suspect terranes. In: HATCHER, R. D., WILLIAMS, H. & ZEITZ, I. (eds), Contributions to the tectonics and geophysics of mountain chains. Geol. Soc. Am. Mere. 158, 33-53. -& ST. JULIEN, P. 1982. The Baie Verte-Brompton Line: Early Palaeozoic continent-ocean interface in the Canadian Appalachians. In." ST JULIEN, P. & BELAND, J. (eds), Major structural zones and faults of the northern Appalachians. Geol. Assoc. Canada, Spec. Paper 24, 177-207. WOODCOCK, N. H. 1979. The use of slump structures as palaeoslope orientation estimators. Sedimentology, 26, 83-99. -

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K. T. PICKERING, Department of Geology, University of Leicester, University Road, Leicester LE1 7RH.

The Kleszcz6w Graben (central Poland): reconstruction of the deformational history and inventory of the resulting soft-sediment deformational structures Krzysztof Brodzikowski, Roman Gotowala, Ludwik Kasza & Antonius J. Van Loon S U M M A R Y : There is a deep tectonic graben near the village of Kleszcz6w, some 50 km south of L6d~. A Mesozoic substratum has subsided in the graben. The infilling consists of unconsolidated Tertiary and Quaternary deposits that reach thicknesses of up to 300 m; locally they may be even more than 400 m thick. The equivalent deposits outside the graben area are 5-7 (or more) times thinner. The lowermost part of the graben contains the oldest infilling, probably indicating the very first moment of subsidence, and consists of late Oligocene deposits. The overlying sediments form an almost complete succession up to the Holocene and represent the best developed and most intensively studied Cenozoic section in Poland. The sections show a unique degree of exposure in continuous horizontal and vertical sections, well developed deformational structures and an exceptional thickness of the Cenozoic succession. There has been much interest in the geological history of the intensely deformed infilling of the graben. The reconstruction of the sedimentological and deformational history is the main topic of this contribution.

Geological setting T h e Kleszcz6w G r a b e n (Fig. 1) forms part of the Szczecin-L6d~-Miech6w synclinory. It separates the L6d~ depression in the north from the R a d o m s k o high in the south. T h e graben itself SIERADZ

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has an e a s t - w e s t direction and its length is s o m e w h a t m o r e t h a n 4 0 k m (Biernat 1968; Po2aryski 1971 ; Kossowski 1947a; Po2aryski & Brochwicz-Lewifiski 1978; Ciuk 1980). T h e regional geology is r a t h e r complex. T h e oldest strata, f o u n d only in boreholes over 4000 m

L

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FIG. 1. Principal tectonic elements of the Radomsko high and position of the Kleszcz6w Graben. 1 = Kleszcz6w Graben. 2 = Jurassic covered by Cenozoic. 3 = Cretaceous covered by Cenozoic. 4 = faults. 5 = depth of Zechstein base (simplified after Po~aryski 1971).

From:JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformationof Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 241-254.

24I

242

K. B r o d z i k o w s k i et al.

deep, consist of strongly disturbed Carboniferous sandstones, deformed during the Variscan orogeny. They are overlaid by salt-bearing Zechstein, and by sandstones, limestones, marls and shales of Triassic, Jurassic and Cretaceous age. That succession, almost 3000 m thick, was folded and faulted during the Laramide phase (Early Paleogene) of the Alpine orogeny. During that deformational period many roughly parallel synclines and anticlines formed with N W - S E trends, together constituting the Szczecin-L6d~-Miech6w synclinory. The Kleszcz6w Graben must have formed slightly later, probably during a final Laramide deformation phase. It was a result of faulting in the hinge zone between the L6d~ depression and the Radomsko high. The direction of the graben was almost identical to that of a structural trend in the deep substratum, but it cannot be excluded that it was a more direct result of an en echelon pattern (Fig. 2) that is connected with the Laramide folding (Baraniecka 1971; Kossowski 1974a). In the last 25 years much geological research has been carried out in the graben because of the presence of large browncoal deposits. The browncoal is now being exploited in an open-cast mine with a surface area of some 8 km 2 and a present depth of about 200 m. Detailed and systematic investigations have been made at the site and have provided a detailed three-dimensional picture of the deposits, especially since mining proceeds gradually parallel to one of the walls of the open-cast mine.

Causes of deformation The sedimentary succession within the graben (Fig. 3) has been strongly deformed (e.g. Van

KLESZCZOW

Loon et al. 1984) due to the endogenic activity that started over 25 million years ago and still continues (Baraniecka 1971, 1975; Brodzikowski & Gotowata 1980; Ciuk & Piwocki 1980; Brodzikowski & Kasza 1982; Gotowata 1982). Differential horizontal and vertical movements took place due to irregular displacements of the Mesozoic substratum which had been fractured into more or less independent blocks. These blocks were affected by the uplift of the L#kifisko anticline, resulting in the opening of hinge fissures. Halokinesis of the Permian salts locally led to considerable vertical displacements. This deformation has generally been attributed to Alpine folding in the Polish Trough (Btaszkiewicz et al. 1968; Derkacz 1968; Kossowski 1974a; Po2aryski & Brochwicz-Lewifiski 1978). Although most salt domes do not occur in the graben area in a strict sense, the diapir near D~bina (Fig. 2) lies clearly within the graben, more or less separating it into two different parts. There are indications that karst phenomena also contributed to the deformation of the graben sediments. Well developed caverns occur in Jurassic limestones all over the graben area. These are filled with typical karst sediments, suggesting that overlying materials have slid down. Other displacements due to gravityinduced mass movements must have occurred when rocks collapsed due to karst; indications of such displacements are the rather frequent local depressions, filled with younger material, in the limestone substratum. A probably much more important factor was the compaction of the Miocene browncoals, which even now are up to 200 m thick. There are indications that during the 12 million years that passed since the end of the organogenic deposition the average degree of compaction has been

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The Kleszcz6w Graben, central Poland

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some 50%. The influence of a thickness reduction of some 200 m upon the deformational history has been emphasized previously by various authors (Ciuk 1975; Ciuk & Piwocki 1980; Brodzikowski & Kasza 1982). Another factor that strongly affected the sediments was the Pleistocene glacial history. Several Scandinavian glaciations reached the Kleszcz6w area (Baraniecka & Sarnacka 1971; R6£,ycki 1978; Lindner 1982, 1984; Lindner & Grzybowski 1982). Seven or even eight different glacial cycles have so far been distinguished. They are separated by either erosional horizons or fluvial and lacustrine sediments (Brodzikowski et al. 1980; Brodzikowski 1982a; Hatuszczak 1982). Both sedimentological and structural observations indicate that the weight of the ice sheet influenced the periodic reactivity of some older dislocation zones. It also probably induced a local increase of the subsidence intensity and may have activated halokinesis at the ice margins (Kossowski 1974a; Brodzikowski 1982a, 1985; Gotowala 1982). Apart from the rather large-scale processes just mentioned, a number of less important exogenic processes have contributed to the soft-sediment deformation within the graben. Although such processes may have resulted locally in severe deformation, the scale of such deformation is usually small if compared to the structures that

formed due to the previously mentioned processes. It is possible to distinguish between metadepositional (penecontemporaneous), cryogenic and glacitectonic (Brodzikowski & Van Loon 1985a) deformations. Although the glacitectonic deformations have frequently been mentioned in the literature (Brodzikowski 1982b, 1985; Brodzikowski & Kasza 1982; Gotowata & Hatuszczak 1982), they are the least important type in the sediments considered here. A schematic diagram of the interrelations between the various deformational agents discussed above is presented in Figure 4.

Deformational history and resulting soft-sediment deformational structures As can be deduced from the oldest infillings in the deepest parts of the graben, the history of the latter probably began during the late Oligocene when marginal faults, determining the boundaries of the graben, were formed in the north and in the south. The faults developed due to tension in the hinge zone of the Lskifisko anticline. Differential displacement along the faults largely determined the sedimentary pattern and the depositional character.

z44

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Early Miocene The oldest Miocene sediments suggest that, after a first stage with rapid subsidence, the tectonic activity in the entire area became less intense, but the subsidence rate again increased from the middle part of the Early Miocene onward. This resulted in the deposition of mainly fluvial sands, but weathering took place on locally outcropping Mesozoic rocks and a slope cover formed strongly suggesting karst phenomena occurred simultaneously. Subsidence along Laramide dislocation zones continued slowly and repeatedly during the Early Miocene. Syntectonic sedimentation resulted in deposits with abrupt thickness changes at the borders between differentially subsiding blocks.

At the end of the Early Miocene the subsidence became more intense and various blocks subsided rapidly over 100-200 m. A new fault system developed, causing the top of the substratum in general to become 100-150 m lower than it was during the earliest Miocene (Fig. 5). The Lower Miocene cannot be observed in the field. Since cores of borings do not provide much reliable data with respect to deformational structures, it can only be assumed that the most common deformations are flexures related to both faults and flexures. Intraformational flowage structures may also be present within the uppermost fluviolacustrine and lacustrine deposits since similar phenomena have been observed elsewhere at tectonically active places with comparable sediments.

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The Kleszcz6w Graben, central Poland Middle Miocene

Paludal and lacustrine conditions prevailed in the Kleszcz6w Graben during the Middle Miocene. Consequently, thick layers of peat were deposited with a local fresh-water limestone facies (Ciuk 1975; Ciuk & Piwocki 1980). On average the thickness of the Middle Miocene amounts to 40-60 m, but may reach up to 150 m locally. The relatively high deposition rate was due to more or less continuous subsidence, originally slow but gradually becoming more rapid. Periods with more intense flexuring took place simultaneously in the bends of the L~kifisko anticline. This resulted in the reactivation of Laramide and Early Miocene fault systems which in turn led to the formation of local depressions. These were bordered by faults where lakes formed in which limestone was deposited, possibly because of the supply of large amounts of calcium carbonate in solution deriving from the surrounding walls. The thickness of the fresh-water limestones indicates that the maximum vertical displacements during the Middle Miocene amounted to some 100 m (Fig. 6). The tectonic subsidence was accompanied by erosion of the suddenly exposed fault walls, so that slope deposits were formed frequently, especially near the southern marginal-fault system. The more or less continuous subsidence and sedimentation in the graben, with gradually changing patterns because of locally differentiated tectonics, now makes it hard to define the many fault zones exactly, although lithologic facies changes may give clear indications. Horizontal facies transitions from fresh-water limestones into slope covers or paludal sediments occur within 100-300 m. The tectonic activity SLOPE

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decreased at the end of the Middle Miocene. As a result, the conditions became less favourable for peat accumulation and the Middle Miocene browncoal now has a sharp upper boundary. Late Miocene

During the Late Miocene the Kleszcz6w Graben became increasingly filled with alternating paludal and lacustrine sediments. The latter show both clayey and carbonate facies. The sedimentary pattern is quite regular because there were few differential displacements. It seems that the subsidence rate was rather low, since sediments built up more rapidly than subsidence took place, resulting in a gradual shallowing of the sedimentary basin. That tendency is clearly expressed by the fluvial sediments that represent the final part of the Late Miocene. Pliocene

The Early Pliocene is characterized by renewed and intense tectonic activity. In fact, the structural pattern of the Tertiary infilling was determined during this period. There were not only important vertical displacements of over 100 m, but horizontal movements obviously affected the mutual position of the individual substratum blocks that consist mainly of Cretaceous and Jurassic rocks. It even seems that the older, prevailing downward displacements changed into predominantly strike-slip or even wrench faulting. Overall subsidence was nevertheless still considerable and, together with the horizontal displacements, resulted in distinct horizontal compression (Fig. 7). Strike-slip movements caused en echelon fracturing and the development of conical folds. THICK LIGNITE /- UNIT / / LOWER MIOCENE

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K. B r o d z i k o w s k i et al.

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Some authors associate the renewed tectonic activity with reactivated halokinesis (Btaszkiewicz et al. 1968; Derkacz 1968; Kossowski 1974a, b; Ciuk 1975). That assumption fits well with the observed strike-slip movements along a number of older fault zones. The increased tectonic activity is also represented by a thick cover of coarse clastic deposits, overlying the Late Miocene sediments. The Pliocene cover consists of large fluvial deposits with a deltaic shape and of debris-flow deposits that pass gradually into typical fluvial deposits towards the axis of the graben. The structural and sedimentological development during the Pliocene gave rise to deformations with large genetic and geometrical variations (Fig. 8). The most common of these are faults and isolated dislocations, but largescale folds occur as well, often showing a diapirlike form. On an intermediate and small scale there are frequent subhorizontal shear planes, overthrusts and brecciation horizons.

Pre-glacial Quaternary stage It seems that the development of the Kleszcz6w Graben suddenly changed at the Tertiary/Quaternary boundary, since a period with intense degradation then started. Most probably that altered situation was related to the uplift of the Lgkiflsko anticline (Po~aryski 1971 ; Kossowski 1974a, b). Consequently, some 2 million years are not represented and the Quaternary starts with Cromerian sediments (Baraniecka 1971, 1975; Baraniecka & Sarnacka 1971 ; Baraniecka et al. 1980). There was, however, renewed activity, induced in the authors' opinion by isostatic movements

related to the deglaciation of the first large Pleistocene ice cap in Europe. That Podlazian glaciation (R6~ycki 1978) corresponds to the Menapian, Baventian, G/inz and Odessan glaciations elsewhere in Europe. It did not reach the Kleszcz6w graben proper (Brodzikowski 1982a), but old faults became reactivated and the axial zone subsided some 100-200 m. The secondary graben of Wola Grzymalina was thus formed (Baraniecka 1971). Its formation was compensated by vertical movements along many other displacement zones (Fig. 9). The depression became filled with fluvial sediments of the Przasznyski interglacial (R6~.ycki 1978), equivalent to Cromerian 2B3, Bestonian+ Cromerian, G - M interglacial and Morozovan. The pre-glacial Quaternary deformations seem to have greatly affected the graben sediments but it is difficult to distinguish them from disturbances directly related to subsequent glaciations. Isolated faults can be recognized, however, as well as brecciation layers in clayey sediments (Fig. 10). Horizontal stresses perpendicular to the graben axis must have played an important role. Whereas compression prevailed during the Middle Miocene, tension seems to have occurred more frequently afterwards. Fracturing of the anticlinal hinge zones resulted and may have occurred during this stage.

Glacial Quaternary Continental ice sheets have covered the Kleszcz6w zone at least seven or eight times. Glacial sediments are intercalated with interglacial and interstadial deposits; in some cases they are separated by erosional surfaces. The repeated glaciations and deglaciations

FIG. 8. Deformed Tertiary deposits in the graben. (A) Part of a large anticline in the Upper Miocene. (B) Reversed fault in clayey Upper Miocene. (C) Miocene coal intruding the Quaternary. (D) Faulted anticline in clayey Upper Miocene and fresh-water calcareous deposits.

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K. B r o d z i k o w s k i et al. PRE- GLACIAL QUATERNARY FLUVIAL FACIES

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FIG. 9. Graben development during the pre-glacial Quaternary. induced (glacio)isostatic movements and led to both renewed fault activity and halokinesis (Baraniecka 1971; Kossowski 1974a; Brodzikowski 1982a; Gotowata 1982). Although renewed displacements occurred in the graben, a new deformational pattern was formed, mainly because glaciogenic deposits were transported by mass movements towards the deepest parts of the depressions. Multilayered masses became folded, overthrusted and sheared; the amplitude of the folds exceeds 20 m in various cases. That value is more or less similar to that of the contemporaneous vertical displacement of the substratum (Fig. 11). The large variation in deformational structures (e. g. Brodzikowski et al. 1987 a ,b,c), their different scales and the unconformities present are strong indicators of multi-stage endogenic activity. Most of the activity took place during the Odranian glaciation (equivalent to Drenthian, Drzieprovian, Altriss, and the maximum extent of the Saalian), when both the secondary Wola Grzymalina graben and the D~bina salt diapir received what is more or less their present configuration. Field evidence shows that the previous morphology was connected with the continued subsidence of the secondary graben. The surrounding deposits dip towards the graben axis, mainly because of gravity-induced movements. The resulting deformations show amplitudes of up to some 50 m. Those deformations all show the same basic structural pattern (Fig. 12) and make other patterns barely discernible. In consequence, glaciotectonic structures in a strict sense may be difficult to distinguish (Brodzikowski 1982b, 1985; Brodzikowski & Kasza 1982; Gotowata & Hatuszczak 1982).

Large-scale folds and flexures are predominant whereas faults are rare. Tertiary sediments are found locally with normal faults, overthrusts and brecciation layers that may have been formed during this stage. Late Pleistocene and Holocene

Following the Odranian glaciation (some 280,000 years BP) tectonic activity clearly slowed down. The structural pattern of the younger deposits is therefore quite different: they are thinner and show greater horizontal continuity. Only local subsidence took place, resulting in rather abrupt facies changes, thicker deposits and an increased number of meso- and micro-scale gravity-induced deformations (Figs 13 and 14). Most deformations are found in glaciolacustrine deposits formed during the Wartanian (equivalent to Saalian II, Moskovian, Typeriss) and in fluviolacustrine deposits from Vistulian (equivalent to Weichselian, Wtirm) age (Brodzikowski I982a,b; Gotowata & Ha|uszczak 1982).

Conclusion The variety of deformational structures in the Kleszcz6w Graben is the result of a complex history. Several phases of intense activity occurred determining both the strain rates and the sedimentary conditions. Younger deformations both mask older deformations and have reshaped them into complex polyphase structures. Deformations took place syn-, meta- and postdepositionally, but during that entire time

FIG. 10. Second-order and small-scale deformations within the Upper Miocene and Pliocene graben infilling. (A) Faulted Pliocene sands. (B) Brecciated coaly clay horizon. (C) Infilling of a fault fissure. (D) Brecciated infilling of a tension fissure in an anticlinal hinge zone.

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FIG. 11. Graben development during the Quaternary glaciations. the sediments remained unconsolidated and the moisture content constituted a prime parameter. There are several structural floors (in the sense of Brodzikowski & Van Loon 1983, 1985b) separating units with different structural characteristics. Those structural floors sometimes coincide with local unconformities. The structural development of the graben is not yet finished. Local subsidence still takes place and some faults develop with an average displacement of 2-5 mm per month. Holocene changes

with respect to the depositional centres are additional evidence of ongoing tectonic activity. Since the large open-cast mine is still under exploitation more data will become available in forthcoming years and no final conclusions can be drawn, as yet, with respect to the details of the deformational history. The model established with the data so far obtained has, however, proved sufficiently accurate to explain the genesis of structures formed by both endogenic and exogenic agents. UPPER GLACIAL QUATERNARY AND POSTGLACIAL DRAINAGE AXIS

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FIG. 12. Deformational structures in Odranian (Drenthian) deposits. (A) Large-scale fold. (B) Flow folds in glaciodeltaic deposits. (C) Disharmonic structure in glaciolacustrine silts and clays. (D) Cleavage-like fault system in glaciofluvial sands.

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The Kleszcz6w Graben, central Poland

253

References BARANIECKA, M. D. 1971. Dorzecze Widawki na tie obszaru marginalnego stadiatu Mazowiecko-Podlaskiego (Warty) w Polsce (with English summary: The Widawka drainage basin as a part of the marginal area of the Masovian-Podlasie (Warta) stadial in Poland). Bull. Inst. Geol. 254, 11-36. 1975. Fazy tektoniczne w czwartorz~dzie w ~rodkowej cz~ci Ni2u Polskiego (with English summary: Quaternary tectonic phases in central part of the Polish Lowlands). In: Proc. 1st Symp. "Recent and neotectonic crustal movements in Poland', 185-194. - & S A R N A C K A , Z. 1971. Stratygrafia czwartorz~du i paleogeografia dorzecka Widawki (with English summary: The stratigraphy of the Quaternary and paleogeography of the drainage basin of the Widawka). Bull. Inst. Geol. 254, 157-259. , CIESLIiq'SKI, S., CIUK, E., DffBROWSKI, A., D#BROWSKA, Z., PIWOCKI, M. & WERNER, Z. 1980. Bodowa geologiczna rejonu belchatowskiego (with English summary). Przeglgd Geol. 7, 381-91. BIERNAT, S. 1968. Problemy tektoniki i morfologii stropu mezozoiku mi~dzy Betchatowem a Dziaioszynem (with English summary). Kwartalnik Geol. 12, 296-306. BLASZKIEWICZ, A., CIESLIIqSKI, S., D~BROWSKA, Z., KARCZEWSKI, L., KOPIK, J. & MALINOWSKI, L. 1968. Zarys stratygrafii i tektoniki potudniowej cz,~ci niecki L6dzkiej (rejon Betchatowa) (with English summary). Kwartalnik Geol. 12, 279-94. BP,ODZmOWSKI, K. 1982a. Problemy wyksztatcenia modelu sedymentacyjnego okres6w glacjalnych w rowie tektonicznym. In: Proe. 1st Symp. "Quaternary of Betehat6w Region', 66-103. -1982b. Przejawy subglacitektonizmu okresu warciafiskiego w przekroju 'Chojny'. In. Proc. 1st Symp. 'Quaternary of Betehatbw Region', 212-23. 1985. Glacial deformation environment in subsiding zone with the Kleszcz6w tectonic graben as an example. Quat. Studies in Poland (in press). -& GOTOWALA, g. 1980. Struktury deformacyjne osad6w czwartorzCdowych. In: Proe. 52nd Meeting Polish Geol. Sot., 309-14. , GOTOWALA, R. & HALUSZCZAK,A. 1980. Kompleksy osadowe odstoni~tej cz~ci nadktadu czwartorz~dowego. In: Proc. 52nd Meeting Polish Geol. Sot., 305-08. , GOTOWALA, R., HALUSZCZAK, A., KRZYSZKOWSKI, D. & VAN LOON, A. J. 1987a. Softsediment deformations from glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben (central Poland). This volume. , HALUSZCZAK, A., KRZYSZKOWSKI, D. & VAN LOON, A. J. 1987b. Genesis and diagnostic value of large-scale gravity-induced penecontemporaneous deformation horizons in Quaternary sediments of the Kleszcz6w Graben (central Poland). This volume. & KASZA, L. 1982. Geneza struktur deformacyjnych osad6w czwartorz~dowych w rowie Klesz-

-

-

-

czowa i ich znaczenie w prognozie geologicznog0rniczej. In: Proc. 1st Symp. "Quaternary of Betchat6w Region', 204-11. , KRZYSZKOWSKI, D. & VAN LOON, A. J. 1987c. Endogenic processes as a cause of penecontemporaneous soft-sediment deformations in the Kleszcz6w Graben (central Poland). This volume. -& VAN LOON, A. J. 1983. Sedimentology and deformational history of unconsolidated Quaternary sediments in the Jarosz6w Zone (Sudetic Foreland). Geol. Sudetica, 18, 121-96. & VAN LOON, A. J. 1985a. Penecontemporaneous non-tectonic brecciation of unconsolidated silts and muds. In: HESSE, R. (ed.): Sedimentology of siltstone and mudstone. Sediment. Geol. 41, 26982. -& VAN LOON, A. J. 1985b. Inventory of deformational structures as a tool for unravelling the Quaternary geology of glaciated areas. Boreas, 14, 175-88. CIUK, E. 1975. Geologiczne podstawy realizacji inwestycji belchatowskiej. In." Proc. Symp. 'Betchatowskie Zagt~bie W~glowe'. & PIWOCKI, M. 1980. Geologia trzeciorz~du w rowie Kleszczowa i jego otoczeniu. In: Proc. 52nd Meeting Polish Geol. Soc., 56-70. DERKACZ, J. t968. Trzeciorz~d strefy zapadliskowej Rz~t~nia-Kleszcz6w-Kamiefisk (with English summary). Przeglgd Geol. 11. GOTOWALA, R. 1982. Tektonika i wyksztatcenie strukturalne czwartorz~du w rejonach Piaski i BuczynaChojny. In." Proc. 1st Symp. 'Quaternary of BeL chat6w Region', 41-65. & HALUSZCZAK, A. 1982. Struktury glacitektonicne rejonu Piaski. In: Pro¢. 1st Symp. 'Quaternary of Betehat6w Region', 224-28. HALUSZCZAK, A. 1982. Zarys budowy geologicznej czwartorz~du w rejonach Piaski oraz BuczynaChojny. In: Proe. 1st Symp. "Quaternary of Betchat6w Region', 14-35. KOSSOWSKI,L. 1974a. Budowa geologiczna zto2a wegla brunatnego Belchat6w ze szczeg61nym uwzgl~dnieniem tektoniki podtoza (with English summary: The geological structure of Betchat6w brown coal deposit considering the tectonic of the underlayer). G6rnictwo Odkrywkowe, 10/ll, 336-44. -1974b. Powierzchenia erozyjna w utworach nadw~glowych zto2a 'Betchat6w'--geneza i charakterystyka (with English summary: Erosive surface in above-coal formations of 'Betchat6w' deposit-genesis and characteristics. G6rnictwo Odkrywkowe, 5, 119-22. LINDNER, L. 1982. South-Polish glaciations (Nidanian, Sanian) in southern central Poland. Acta Geol. Poloniea, 32, 163-77. 1984. An outline of Pleistocene chronostratigraphy in Poland. Acta Geol. Poloniea, 34, 27-49. & GRZYBOWSKI, K. 1982. Middle-Polish glaciations (Odranian, Wartanian) in southern central Poland. Acta Geol. Polonica, 32, 191-206. -

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POT.ARYSKI, W. 1971. Tektonika elewacji radomskowskiej (with English summary: Tectonics of Radomsko elevation). Ann. Geol. Soc. Pol. 41. --

& BROCHWICZ-LEwIIqSKI,W. 1978. On the Polish Trough. In: VAN LOON, A. J. (ed.): Key-notes of the MEGS-H (Amsterdam 1978). Geologie en Mujnbouw, 57, 545-58.

R67.YCKI, S. Z. 1978. From Mochty to a synthesis of

the Polish Pleistocene. Ann. Geol. Soc. Pol. 48, 445-78. VAN LOON, A. J., BRODZIKOWSKI,K. & GOTOWALA,R. 1984. Structural analysis of kink bands in unconsolidated sands. Tectonophysics, 104, 351-74. & -1985. Kink structures in unconsoli~tated fine-grained sediments. In." HESSE, R. (ed.): Sedimentology of siltstone and mudstone. Sediment. Geol. 41,283-300. -

-

K. BRODZIKOWSKI, R. GOTOWALA& L. KASZA, Department of Applied Geology, Institute of Geological Sciences, University of Wroctaw, Uniwersytecka 19/20, 50-145, Wroclaw, Poland. A. J. VANLOON, Julianaweg 5, 6862 ZN Oosterbeek, The Netherlands.

Soft-sediment deformations from glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben (central Poland) Krzysztof Brodzikowski, Roman Gotowala, Andrzej Haluszczak, Dariusz Krzyszkowski & Antonius J. Van Loon S U M M A R Y : The Kleszcz6w Graben, situated some 50 km south of L6d~, has a Tertiary and Quaternary infilling of up to 300-400 m. The graben is some 40 km long and a large exposure is formed by an open-cast browncoal mine. The well discernible structural pattern is due to frequent deformations, many of which are related to subsidence of the graben. The sediments within the graben are all unconsolidated and show a large variety of deformational structures with respect to both genesis and geometry. Various stages of more intense deformation can be discerned. The glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in particular show frequent deformations. They are interpreted as being induced by both endogenic and exogenic agents, often during several phases that affected the same sedimentary unit.

Geological outline The Kleszcz6w Graben in central Poland (Fig. 1) offers a well exposed succession of the uppermost 200 m of Tertiary and Quaternary sediments (Fig. 2). The main part of the exposed Tertiary consists of browncoal. Within the Quaternary, some 40% (locally up to 60%) consists of glaciodeltaic, glaciolacustrine and fluviolacustrine sediments. The deposits can be traced horizontally over hundreds of metres or even kilometres. Their thickness is clearly influenced by the movements of the substratum (Brodzikowski et al. 1987a). Some rather abrupt changes in facies thickness occur indicating that tectonic subsidence played a major role with respect to the sedimentary development of the basin. Since the hard-rock Mesozoic substratum was fractured and split into more or less isolated blocks, a large number of smaller vertical faults developed. This influenced the sedimentary pattern, especially since the position of the deepest point shifted repeatedly within each sub-basin. Consequently, relatively steep slopes could form suddenly, giving rise to unstable conditions that led to mass movements. Both the tectonic activity and the forthcoming mass movements induced other deformations (Brodzikowski et al. 1987c), but many other processes also played a deformational role (Brodzikowski et al. 1987b).

Sedimentology All sedimentary units considered here show great lithological variation. There are medium- and

fine-grained sands, silts, clays, peats and organic detritus. Another characteristic they possess is the good stratification, usually formed by laminae of alternating grain sizes. The succession may show cycles, sometimes with subcycles. Well developed cyclic sedimentation is found in varves, sometimes with interseasonal lamination (Fig. 3). More or less comparable characteristics are shown by grading-upward layers. These are most probably turbidites. Well developed cross-bedding within sandy deposits is rare but does occur, especially in the glaciodeltaic and fluviolacustrine facies (Fig. 4). More details regarding the sedimentology of the graben have been reported by Brodzikowski et al. (1980), Hatuszczak (1980, 1982), Brodzikowski (1982a, b,c), Brodzikowski & Baraniecka (1982), Brodzikowski & Hatuszczak (1982) and Hatuszczak & Brodzikowski (1982).

Deformational processes The frequently occurring rhythmic layering has greatly influenced the genesis of deformational structures. Rhythmic bedding is very sensitive to mechanical pulses, especially because of the large vertical anisotropy. Even weak pulses may result in complicated disturbances (cf. Dott & Howard 1962; D±uiyfiski & Smith 1963; Butrym et al. 1964; D2uiyfiski & Walton 1965; D±ulyfiski & Radomski 1966; Anketell et al. 1969; Danilov 1973). Consequently, a large number of both large- and small-scale deformations can be found in the Kleszcz6w Graben (Brodzikowski & Gotowata 1980; Brodzikowski et al. 1987a).

From:JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks,

Geological Society Special Publication No. 29, pp. 255-267.

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FIG. 1. Geological setting of the Kleszcz6w Graben (modified after Poiaryski 1971 and Kossowski 1974). (A) Main tectonic elements. 1 = Location of the graben. 2 = Jurassic under Cenozoic cover. 3 = Cretaceous under Cenozoic cover. 4 = Main faults. 5 -- Depth of the Zechstein. (B) Schematic Quaternary infilling of the graben. A-C -- Location of sections in figure 1C. Q = Quaternary. T = Tertiary. M = Mesozoic. (C) Tectonic sketch map of the Mesozoic in the Kleszcz6w area. 1 = Stratigraphic boundaries. 2 = Faults. 3 = Graben boundaries. 4 = Sections (see B). M a n y processes have been involved in producing d e f o r m a t i o n s including tectonic subsidence and horizontal displacements, e a r t h q u a k e s without a p p a r e n t displacement, collapses due to karst p h e n o m e n a , halokinesis, bioturbation, cryogenesis, c o m p a c t i o n and dewatering. T h e various types of resulting deformations will be described a n d analysed below.

Interpretation of deformational structures Most attention has been paid to the d e f o r m a t i o n s present in the Q u a t e r n a r y g r a b e n infilling. Alt h o u g h various types of structures can be discerned, it appears extremely difficult to group t h e m into well defined categories. This is espe-

cially true since various processes m a y have affected the same s e d i m e n t simultaneously, and since d e f o r m e d materials m a y have been reshaped several times by later deformational processes. E a c h classification must therefore be so detailed as to be unwieldy, or must be insufficiently unambiguous. A classification system was arrived at in spite of the problems encountered. The m a i n (practical) criteria were the geometry of the final structure and its vertical extent. The six groups that were thus distinguished can each be subdivided on the basis o f additional criteria.

Group I: large-scale folds Large-scale folds with amplitudes of 25-50 m can be found in the Pleistocene deposits formed during and b e t w e e n the first glaciations. It is

Soft-sediment deformations in the Kleszcz6w Graben interesting to note that the amplitude corresponds closely to the vertical displacement in the same time interval (Brodzikowski et al. 1987a). It therefore seems reasonable to assume a relation between the complex and differential tectonic movements. The validity of this assumption is confirmed by the frequent changes in the direction of the fold axes. Moreover, the vergence of the folds is usually in agreement with the inclination of the corresponding (now buried) morphological surface. Overturned and recumbent folds prevail. Wherever upright folds occur, they have a diapiric character (Fig. 5). Together these data indicate slow, large-scale gravity flowage of multilayered units. Other folds apparently formed more by chance when subsidence resulted in a topographic depression. Such mass movements were much less uniform and continuous, as shown by frequent and abrupt changes in the direction of the fold axes. Subhorizontal decollement planes are not uncommon in those types of gravityinduced large-scale folds. They indicate that the folds came to rest then slid down again, sometimes repeatedly. In such a case the fold axis of a second movement may differ from that of the first movement. Consequently, this type of fold often shows a rather chaotic interior, as is easily seen from the grain-size variations in the layers and laminae involved.

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The sediments under consideration react to strain in several different ways because of the rapid vertical grain-size alternations. Consequently, more or less uniform parts within large-scale folds behave as separate units reacting, not to the general conditions affecting the sediment, but to the specific conditions present within the largescale fold structure. A number of meso- and micro-scale deformational structures can therefore be found within the large folds. In many cases it is quite clear that the secondary deformations in turn created stress conditions that gave rise to still smaller structures. It is obvious that specific sedimentary units behaved in quite different ways depending on the physical conditions. In the glaciodeltaic sediments, for example, structures indicative of fluidization prevail, but systems of cracks are also common. Other structures that occur are cleavage in the axial zones of folds, and boudinage in the flanks. The secondary structures usually can be considered to be a direct consequence of diverging physical parameters, the main one being the pore-water content so strongly dependent on

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FIG. 2. Composite stratigraphic column of the Quaternary infilling of the Kleszcz6w Graben. A = Nidanian (lower Elsterian). B = Malopolanian. C = Sanian (upper Elsterian). D = Ferdynandovian. E = pre-Odranian (lower Saalian). F = Odranian (Drenthian). G = Pilician. H = Wartanian (upper Saalian, Warthian). I = Eemian. J = Vistulian (Weichselian). K = Holocene. 1 = Glacial tills. 2 = Glaciofluvial facies. 3 = Glaciodeltaic and glaciolacustrine facies. 4 = Fluvial facies. 5 = Fluviolacustrine facies. 6 = Organogenic sediments.

FIG. 3. Odranian glaciolacustrine sediments. (A) Lower level. (B) Middle level. (C) Upper level.

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Soft-sediment deformations in the Kleszcz6w Graben grain-size distribution. When the pore-water pressure increased, liquefaction may have taken place and elastic dykes formed. The process of liquefaction quite frequently led to sands and silts filling micro-scale tension cracks that occur where the folds bent most intensely. Group II: flowage and related structures The amplitude of the flowage and related folds varies between a few centimetres and about 10 m, so they may be classified as meso-scale deformations. A vast majority are recumbent and disharmonic, probably having developed due to gravitation-induced flowage over the (now buried) sedimentary surface. An interesting phenomenon is the (infrequent) occurrence of cleavage in the folds (Fig. 6). It appears that most such structures are found in the fluviolacustrine and glaciodeltaic deposits, so they seem to have formed metadepositionally since sedimentation seems to have continued without major interruptions in their environment. A subaerial origin of such structures might be considered in a few cases, but has never been proved on the basis of unambiguous data. The genesis of the structures is considered to be related to endogenic activity (earthquakes may have been the trigger mechanism), but exogenic processes certainly played a role. Exceeding the critical stability value of a slope (either by continuous sedimentation at that slope or by ice melting in the ice-contact zone) in particular may have induced such movements (cf. Dott 1963; Elliot 1965; Brodzikowski & Gotowata 1980; Brodzikowski 1982a; Brodzikowski & Kasza 1982; Brodzikowski & Van Loon 1983). Group III: large-scale load casts and related deformations The sediments considered here show specific mechanical properties (e.g. differential viscosity and density because of the vertical grain-size alternations), usually in the form of silty laminae that contain a large proportion of either sand or clay. Frequent load casts, drop structures, pseudonodules and related structures could form due to these conditions. Their sizes vary considerably and seem dependent on the pore-water pressure (and its gradients) and thickness of the layers or laminae involved. Long term conditions seem to have favoured the gradual growth of the structures, but the time interval during which the structures developed can only be estimated indirectly.

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This type of structure can be found almost everywhere, but a few horizons (Fig. 7) show exceptionally well developed specimens (Brodzikowski 1982b). In a few cases the vertical extent of the load casts is equal to the thickness of the underlying, less dense layer and in such cases the parent layer has been entirely reshaped into load casts. Such an extreme development might be due to the contemporaneous occurrence of tectonic activity (Brodzikowski & Hatuszczak 1982). Group IV: discontinuous structures and breccias The vertical anisotropy of the successions resulted not only in folds in response to exerted stress, but also in discontinuous structures, e.g. faults, complex fault zones, shear zones, breccias and fissures. Varves especially were affected by such types of deformation. The size of the disturbances varies from less than one centimetre to several hundreds of metres, the latter especially being faults formed by a subhorizontal decollement plane at their lower end, a single fault over the greater part gradually passing into a number of smaller faults at the upper part. Other faults show less marked characteristics, being only a few metres long and passing ultimately into translation flexures or comparable structures. Because of local tension, tension cracks have formed in several places. Those fissures are now filled with severely disturbed material, probably because infilling took place through gravityinduced collapse of overlying material when the fissures opened. In some cases the fissure filling has been brecciated (Fig. 8), but that is considered to be a still later type of deformation. Other types of brecciation must be due to the sliding of varved sediments over sedimentary interfaces. Such intraformational sliding could develop when the inclination of the sediments increased, due to faulting or folding activity within the graben. A third type of breccia is related to the lower parts (subhorizontal decollement planes) of the larger faults, and thus might be considered as tectonic brecciation of unconsolidated sediments.

Group V: small-scale convolution horizons and related deformations There exist abundant, rather small, penecontemporaneous deformations such as convolutions, deformed ripple marks, flames and other deformations resulting from unstable density gradients, etc. Since they occur in each lithostratigraphic unit it is clear that the condi-

FIG. 6. Flowage folds and related deformations in lower Saalian (Odranian and pre-Odranian) glaciodeltaic and glaciolacustrine sediments.

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tions in the sedimentary environments have always remained favourable to such disturbances. The water-saturated nature of the sediments, the relatively high silt content and the pore-water pressure gradients due to vertical anisotropy are especially important parameters. Despite the general conditions favouring smallscale deformations, endogenic activity must have provided additional possibilities, probably mainly in the form of a repeatedly active trigger mechanism. Within the 5 m thick, laminated glaciolacustrine unit, for example, there occur more than forty horizons which show even more abundant microdeformations than do the other parts of the unit. Since the frequency of deformation in those horizons remains exceptionally high even when gradual grain-size changes occur, an external influence such as an earthquake is assumed to have played a role.

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Group VI: polyphase deformations Various complex structures must have formed during several phases of development. The most interesting examples are the tension cracks that occur in the hinge zone of large anticlines consisting of Tertiary (Miocene) and Quaternary deposits (Fig. 9). It seems that the anticlines developed in several phases between the Pliocene and the Odranian ( = Saalian I), each phase being characterized by renewed activity of the graben substratum (Fig. 10). Early-diagenetic compaction of the sediments seems to have been involved in the deformational process as well, indicating that the present vertical dimensions of the cracks (up to 50 m) may have been even larger. Such cracks are connected with faults in the substratum or with decollement planes in the flanks of the anticlines. Filled-in tension cracks may resemble fossilized ice wedges, especially when seen in twodimensional cross sections (cf. Jahn 1975) but their three-dimensional structure shows clear differences.

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z65

Conclusions The lithological succession of the glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben shows layers and laminae with a rapidly alternating grain size. Consequently there is great vertical anisotropy. The contacts between the various layers and laminae in particular offered favourable places for microstructural deformations, but thicker successions as a whole were also affected by deformational processes. The resulting structures are both abundant and diverse. This hampers any strict classification that might have helped to unravel their genesis. A first, rough classification could only be based on geometry and size of the structures because so many are due to various deformation processes and/or shaped during various deformational stages. A more detailed analysis will be published when more data become available during the continuous exploitation of the browncoal overburden in the graben. One of the early conclusions that can be drawn is that 'classical' concepts for the explanation of hard-rock deformations cannot always be applied to the unconsolidated sediments dealt with here. The original sedimentary pattern might strongly affect the overall mechanical behaviour of a sedimentary unit.

FIG. 10. Details of collapse structures due to opened tension fissures in the hinge zone of a Tertiary anticline. The collapsed material belongs to the Czy2bw Series. (A) Uppermost part of the collapse zone. (B) Middle part of the collapse zone. (C) Various meso-scale collapse structures.

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Soft-sediment deformations & the Kleszcz6w Graben

267

References ANKETELL, J. M., CEGLA, J. & DZULYIQSKI,S. 1969. Unconformable surfaces formed in the absence of current erosion. Geol. Romana, 8, 41-6. BRODZIKOWSKI,K. 1982a. Deformacje osad6w nieskonsolidowanych w obszarach ni~owych zlodowacefi plejstocefiskich na przykladzie Polski potudniowozachodniej (with English summary : Deformations in unconsolidated sediments in areas glaciated during the Pleistocene with south-west Poland as an example). Acta Univ. Wratislaviensis 574, Studia Geogr. 34, 87 pp. -1982b. Problemy wyksztatcenia modelu sedymentacyjnego okres6w glacjalnych w rowie tektonicznym. In: Proc. 1st Syrup. 'Quaternary of Betchat6w Region', 66-103. 1982c. Wstqpne wyniki badafi sedymentologicznych i strukturalnych osad6w fluwiolimnicznych mi~dzymorenowej serii 'Czy~6w' (strefa Buczyna-Chojny). In. Proc. 1st Syrup. 'Quaternary of Betchat6w Region', 180-98. --& BARANIECKA,M. D. 1982. Strefa przekrojowa 'Chojny', interpretacja sedymentologiczna stanowisk osad6w organicznych. In. Proc. 1st Syrup. 'Quaternary o f Betchat6w Region', 124-40. -8~; GOTOWALA, R. 1980. Struktury deformacyjne osad6w czwartorz~dowych. In: Proc. 52nd Meeting Polish Geol. Soc. 309-14. ---, & HALUSZCZAK, A. 1980. Kompleksy osadowe odstoni~tej cz~gci nadktadu czwartorz~dowego. In: Proc. 52nd Meeting Polish Geol. Soc. 305-08. , KASZA, L. & VAN LOON, A. J. 1987a. The Kleszcz6w Graben (central Poland)" reconstruction of the deformational history and inventory of the resulting soft-sediment deformational structures. This volume. - - - & HALUSZCZAK,A. 1982. Deformacje metasedymentacyjne oraz wczesnodiagenetyczne w osadach rzecznych i zbiornikowych grodowisk depozycyjnych. In: Proc. 1st Symp. 'Quaternary of Betchat6w Region', 264-70. , KRZYSZKOWSKI,D. & VAN LOON, A. J. 1987b. Genesis and diagnostic value of large-scale gravity-induced penecontemporaneous deformation horizons in Quaternary sediments in the Kleszcz6w Graben (central Poland). This volume. & KASZA, L. 1982. Geneza struktur deformacyjnych osad6w czwartorz,dowych w rowie Kleszczowa i ich znaczenie w prognozie geologicznog6rniczej. In: Proc. Ist Symp. "Quaternary of Betchat6w Region', 204-11.

--,

KRZYSZKOWSKI,D. & VAN LOON, A. J. 1987c. Endogenic processes as a cause of penecontemporaneous soft-sediment deformations in the fluviolacustrine Czy~6w Series (Kleszcz6w Graben, central Poland). This volume. • VAN LOON, A. J. 1983. Sedimentology and deformational history of unconsolidated Quaternary sediments of the Jarosz6w Zone (Sudetic Foreland). Geol. Sudetica, 18, 121-96. BUTRYM, J., CEGLA, J., DZULYI;,ISKI,S. & NAKONIECZNY, S. 1964. New interpretation of periglacial structures. Folia Quaternaria, 17, 1-34. DANILOV, I. D. 1973. Subaqueous pseudomorphoses in Pleistocene deposits. Biul. Pervglacjalny, 22, 34045. DOTT Jr., R. H. 1963. Dynamics ofsubaqueous gravity depositional processes. Amer. Assoc. Petrol. Geol. Bull. 47, 104-28. ~; HOWARD, J. K. 1962. Convolute lamination in non-graded sequences. J. Geol. 70, 114-20. DP.ULYlqSKI,S. & RADOMSKI,A. 1966. Experiments on bedding disturbances produced by the impact of heavy suspensions upon horizontal sedimentary layers. Bull. Acad. Pol. Sci. 14, 227-30. & SMITH, A. J. 1963. Convolute lamination, its origin, preservation and directional significance. J. sediment. Petrol. 33, 616-27. -~; WALTON, E. K. 1965. Sedimentary features of flysch and greywackes. 274 pp. Elsevier. ELLIOT, R. E. 1965. A classification of subaqeous sedimentary structures based on rheological and kinematical parameters. Sedimentology, 5, 193209. HALUSZCZAK, A. 1980. Wst~pna charakterystyka facjalna poziom6w zastoisko wych. In: Proc. 52nd Meeting Polish Geol. Soc. 314-18. & BRODZIKOWSKI,K. 1982. Flamaj strukturoj kaj akompanaj metasedimentaj asambleoj de deformoj en glacilagaj sedimentoj. Geol. Internacia, 4, 11726. JAHN, A. 1975. Problems of the periglacial zone. PWN, 233 pp. KOSSOWSKI,L. 1974. Budowa geologiczna zto±a w~gla brunatnego Betchat6w ze szczeg61nym uwzgl~dnienem tektoniki podto~a (with English summary: The geological structure of Betchat6w brown coal deposit considering the tectonic of the underlayer). G6rnictwo Odkrywkowe, 10[11, 336-44. PO~,ARYSKI,W. 1971. Tektonika elewacji Redomskowskiej (with English summary: Tectonics of Radom-

-

-

-

-

-

K. BRODZIKOWSKI,R. GOTOWALA, A. HALUSZCZAK& D. KRZYSZKOWSKI,Department of Applied Geology, Institute of Geological Sciences, University of Wroctaw, Uniwersytecka 19/20, 50-145 Wroclaw, Poland. A. J. VAN LOON, Julianaweg 5, 6862 ZN Oosterbeek, The Netherlands.

Endogenic processes as a cause of penecontemporaneous softsediment deformations in the fluviolacustrine Czyi6w Series (Kleszcz6w Graben, central Poland) Krzysztof Brodzikowski, Dariusz Krzyszkowski & Antonius J. Van Loon S U M M A RY: Thick fluvial and lacustrine deposits of Quaternary age are rare in Europe. One such succession is found in the Kleszcz6w Graben. The high content of organogenic material is a specific characteristic. One of the major fluviolacustrine units in the graben is the Czy26w Series, formed during a Pleistocene interstadial. The series is characterized by abundant penecontemporaneous deformation structures whose formation was favoured by both the grain-size distribution (with a relatively high silt content) and repeated tectonic displacements that served as a trigger mechanism. The analysis of these structures forms the main topic of the present contribution. Endogenic activity within the graben emerged as the main cause not only of 'normal' tectonic deformations, but also of a large number of deformations that are usually considered to have a sedimentary or non-tectonic diagenetic origin.

Geological setting and sedimentology T h e Kleszcz6w G r a b e n in c e n t r a l P o l a n d , s o m e 50 k m south o f L6d~, has b e e n f o r m i n g since the Oligocene ( B a r a n i e c k a & S a r n a c k a 1971; Got o w a t a 1982; B r o d z i k o w s k i et al. 1987a) a n d the u n c o n s o l i d a t e d T e r t i a r y a n d Q u a t e r n a r y infilling

c o n t a i n s m a n y units t h a t are affected by deform a t i o n a l processes ( B r o d z i k o w s k i et al. 1987b,c). T h e infilling r e a c h e s a local t h i c k n e s s of u p to 400 m b u t the Pleistocene is restricted to s o m e 100-250 m ; t h a t is n e v e r t h e l e s s a b o u t three to five t i m e s t h e t h i c k n e s s o f the e q u i v a l e n t sedim e n t s o u t s i d e t h e g r a b e n (Fig. 1). S

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F]o. 1. Geological overview of the Kleszcz6w Graben (modified after Baraniecka 1980; Baraniecka & Sarnacka 1971). (A) Main structural elements. 1 = Permian Debina diapir. 2 = Sudetic monocline. 3 = Silesia-Cracow monocline. 4 = Depressions filled with Cretaceous sediments, covered by Cenozoic. 5 = Kleszcz6w Graben. 6 = Stratigraphic boundaries within the Cretaceous (Crc = Cenomanian and Turonian; Crs = Santonian and Coniacian; Crk = Campanian; Crm = Maestrichtian). (B) Simplified cross section through the graben near Wola Grzymalina (see IC). 1 = Mesozoic. 2 = Tertiary. 3 = Quaternary. (C) Palaeogeographic reconstruction of interglacial (now buried) valleys in the Kleszcz6w zone. 1 = ?Cromerian Wola Grzymalina valley. 2 = Mazovian (Holsteinian) Ruszczyn valley. 3 = Pilician (middle Saalian) valley. I = Location of section in figure 1B.

From" JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 269-278.

269

270

K. Brodzikowski, D. Krzyszkowski & A. J. Van Loon

The succession within the graben has been studied in detail because of the almost complete succession from the Oligocene up to the Recent, and has many features of special interest to stratigraphers. One of the features is the Czy~6w Series, which according to palynological analysis must have formed during one of the interstadials of the South-Polish glaciation (Baraniecka 1982; Hatuszczak 1982; Krzyszkowski 1984; Lindner 1984). The same unit is of interest to sedimentologists because of both the well traceable lateral and vertical facies transitions between lacustrine and fluvial deposits, and the abundantly present penecontemporaneous deformations.

Characteristics of the Czy~bw Series The basin where the Czy~6w Series was deposited (Fig. 2) was situated in the eastern part of the Kleszcz6w Graben. The basin was up to 3.5 km long and up to 1.3 km wide. The sediments reach a thickness of up to 40 m in the central part of the basin and comprise subunits with diverging characteristics. There occur laminated silts and fine-grained sands, silty-clayey rhythmites, medium- and coarse-grained sands with current ripples and coarse units which are mainly built up from cross-bedded sand layers. There is distinct cyclicity although the number of cycles depends on the exact position within the basin. More than ten cycles have been observed in some sections, each starting with coarse sands eroding the underlying fine-grained material (Fig. 3). The results of detailed sedimentological research have been only partially published so far (Brodzikowski 1982; Brodzikowski & Baraniecka 1982), but more data will soon be available (Krzyszkowski, in press). From the well exposed lithology (the graben is --,

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being exploited for Miocene browncoal) it appears that the Czy26w Series consists of fluvial and lacustrine sediments formed simultaneously in an ever changing sedimentary pattern. Consequently there occur many vertical and lateral facies transitions between the two main types of sediment. The lacustrine deposits within the fluviolacustrine series tend to be fine-grained and to contain organogenic layers (peats and carbonaceous silts), whereas the fluvial sediments generally are coarser.

Deformation structures Some deformation structures have clearly formed in a post-depositional stage, which means that the sediments had already been covered by younger layers before deformation started. In such cases the deformational process was not directly related to the process of sedimentation. The majority of structures, however, apparently formed during the depositional process and as a result of the continuing deposition, should be considered as penecontemporaneous or even metadepositional (formed before being covered). In many cases tectonic activity in the graben has influenced both the sedimentary pattern and the deformational process. The deformations found most frequently are load casts, ball-and-pillow structures, flowage folds, buckle and drag folds, various types of faults, tension fissures (sometimes in regular patterns) and large-scale flexures.

Load casts, ball-and-pillow structures and related deformations Although this group of (frequently complex) structures is quite common all over the Czy~6w Series, there exist specific horizons where they are particularly abundant. Such horizons usually belong to the lower- or uppermost parts of the lacustrine deposits and the structures can generally be found over their entire lateral extent. The vertical dimensions of the structures vary between a few cm and some 3 m (Fig. 4). 'Classical' load casts with well preserved, rather quiet internal lamination are most common. There are, however, large drop structures with a complex interior, formed during several deformational stages. The structures show frequent faults and material derived from the overlying sediments may be found inside them. These large structures seem restricted to specific horizons, fourteen of which have been found. Flame structures occur in a more scattered way and it cannot always be established whether they result from pressure exerted by load casts (or

Endogenic processes in the Czy~.6w Series

27I

FIG. 3. The middle Elsterian fluviolacustrine Czy~6w Series. (A-C) Overview of exposed sections in the opencast mine of Betchat6w. (D) Detail of a fluvial succession. (E-G) Details of the lacustrine successions. comparable structures) or whether they should be considered as a result of upward pressure in a zone of weakness (diapirs or clastic dykes in statu nascendi). Complete destruction of original structures due to liquefaction is much less common. The frequent occurrence of these structures, at

least in specific horizons with a large lateral extent (up to 1 km), suggests a specific origin. It seems that a relationship exists between the genesis of deformational levels and facies transitions. The structures dealt with in this section, however, throw no light on the kind of relation-

272

K. Brodzikowski, D. Krzyszkowski & A. J. Van Loon

FIG. 4. Horizons with large-scale load casts in the Czy26w Series. ship there must be between those two factors, but changes in the regional sedimentary pattern certainly played a role.

Buckle folds, drag folds and flowage structures This group of structures usually has an intraformational character.

Drag folds were found in the limbs of synclines formed in the Czy26w Series. The amplitude of the folds is up to some 50 cm and the vergence of the drag folds on one side of a larger-scale syncline is always opposite to the vergence of the drag folds at the other side of the larger structure (Fig. 5). This is in agreement with folding models for multilayered systems (Ramsay 1967).

Endogenic processes in the Czy26w Series

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found (Fig. 6) that have an amplitude similar to that of the drag folds. The occurrence of drag and buckle folds points to endogenic influences during the deformation of the Czy26w Series. Various other phenomena

FIG. 6. Drag folds in the multi-layered system of the Czy~6w Series (compare with figure 5).

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K. Brodzikowski, D. Krzyszkowski & A. J. Van Loon

such as rapid facies changes and widespread deformational horizons might be due to the same endogenic activity. Faults, fissures and fissure systems

The validity of the conclusion just drawn is confirmed by analysing the faults present in the Czy~6w Series. Most large normal faults occur in the limbs of the synclinal parts of the series. The position of those synclines is clearly dependent on zones in the Mesozoic and Tertiary substratum where tectonic displacements took place (Fig. 7). The relationship is expressed not only by similar directions of the faults in both the substratum and the Czy~6w sediments, but also by the position of the synclines just on top of the subsided parts of the substratum. There are, of course, completely independent faults and fissure systems as well. Most of them are related to pressures exerted during the formation of load casts and these two types of structures can often be proved to have formed contemporaneously. The throw of the faults mentioned here generally does not exceed a metre (Fig. 8). Nearby faults may show quite different throws, another indication of a syndepositional development.

have shown that these fissures only occur in the hinge zones of anticlines. They reach down to the Tertiary where they form part of larger deformational systems. The fissures may reach lengths of up to 50 m, cutting through large parts of the Quaternary. It is remarkable that they all end in the Czy~6w Series. It is well known that the series was formed during an interstadial, but the upper ends of the large fissures seem initially to indicate giant epigenetic fossil ice wedges with faulted margins (cf. Jahn 1975). The faults tend to mask the throw (of up to 4 m) that may be present between both sides of the fissure (Fig. 9). The wedges show various kinds of infilling. The uppermost part, which has the shape of a funnel, clearly shows how overlying sediments have slid down into the opening. The middle part resembles more a real wedge with infilling that does indeed resemble that of ice-wedge casts. The lower parts of the fissures have the character of a complex fault zone with homogeneous or brecciated material inside, without any original layering left in it. There seems no doubt that the origin of the fissures is related to tectonic movements of the substratum (Fig. 10). Moreover, the vertical length practically excludes the possibility of a non-tectonic origin.

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Large wedge-shaped tension fissures occur in the southern and eastern parts of the Kleszcz6w Graben. These are filled with sediments derived from the overlying layer(s). Detailed observations Q/T BOUNDARY

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Conclusions The deformational history of the Czy~6w Series had already begun at a metadepositional stage (penecontemporaneously s.s.) and continued post-depositionally, both during the deposition of younger layers of the same unit (contemporaneously s.1.) and later (even up to the present day). The depositional basin in the Kleszcz6w Graben was repeatedly affected by tectonic displacements, alternating with quieter periods. Local differences in tectonic activity also played a role in the resulting changes of the sedimentary pattern. The most obvious consequences were rapid lateral and facies transitions between fluvial and lacustrine sediments. Since the facies transitions were accompanied by differences in the average grain-size and density of the layers, unstable density gradients developed favouring deformational processes such as loading. Earthquakes may have been relatively common and possibly served as a trigger for such deformations that then became widespread over specific horizons (the sedimentary surface of that specific moment). The endogenic activity also resulted in large

FIG. 8. Meso-scale and micro-scale deformations in the Czy/~6w Series. (A) Normal faults. (B) Load casts and gravifossum (cf. Van Loon & Wiggers 1976). (C) Faulted load-cast horizon. (D) Flowage folds.

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Endogenic processes in the Czy26w Series

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synclines, due to subsidence of the cover on top of local subsidence in the graben substratum, and consequently in anticlines between the synclines. Drag and flowage folds developed in the limbs of the larger structures during the formation of the synclines and anticlines, whereas buckle folds were formed in the axial parts. All these smallscale structures are to be expected from the folding model of multilayered systems. Large faults developed during the deposition of the Czy~6w Series. These are found in flexure zones and in the limbs of the large synclines and anticlines. Other faults represent typical meta-

depositional stress systems resulting from various types of unstable-density deformations. The hinge zones of the anticlines contain the top part of large tension wedges that developed as a reaction to the folding caused by tectonic activity in the graben hard-rock substratum. The wedges are filled by overlying sediments. Endogenic activity in general must be considered to be responsible for most of the deformational structures within the Czy26w Series. In some cases the activity was a primary cause, but in others the deformations were a secondary consequence of graben activity.

References BARANIECKA, M. D. 1980. Geologia czwartorz~du dorzecza Widawki. In: Proc. 52nd Meeting Polish Geol. Soc., 71-84. - 1982. Dotychczasowe i bie2~tcebadania czwartorz~du rejonu Betchatowa. In: Proc. 1st Symp. 'Quaternary of Betchat6w Region', 1-13. - • SARNACKA,Z. 197l. Stratygrafia czwartorz~du i paleogeografia dorzecza Widawki (with English

summary: The stratigraphy of the Quaternary and paleogeography of the drainage basin of the Widawka). Bull. Inst. Geol. 254, 157-259. BRODZIKOWSKh K. 1982. Wst~pne wyniki badafi sedymentologicznychi strukturalnych osad6w fluwiolimnicznych mi~dzymorenowej serii Czy26w (strefa Buczyna-Chojny). In: Proc. 1st Syrup. "Quaternary of Betchat6w Region', 180-98.

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& BARANIECKA,M. D. 1982. Strefa przekrojowa 'Chojny', interpretacja sedymentologiczna stanowisk osad6w organicznych. In." Proe. 1st Symp. 'Quaternary of Betehat6w Region', 124-40. --, GOTOWALA, R., KASZA, L. & VAN LOON, A. J. 1987a. The Kleszcz6w Graben (central Poland): Reconstruction of the deformational history and inventory of the resulting soft-sediment deformational structures. This volume. - - , HALUSZCZAK, A., KRZYSZKOWSKI,n. & VAN LOON, A. J. 1987b, Soft-sediment deformations from glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben (central Poland). This volume. --, HALUSZCZAK, A., KRZYSZKOWSKI,D. & VAN LOON, A. J. 1987c. Genesis and diagnostic value of large-scale gravity-induced penecontemporaneous deformation horizons in Quaternary sediments of the Kleszcz6w Graben (central Poland). This volume. GOTOWA~A, R. 1982. Tektonika i wyksztatcenie strukturalne czwartorz~du w rejonach Piaski i BuczynaChojny. In: Proc. 1st Syrup. 'Quaternary of Betchat6w Region', 41-65. HALUSZCZAK, A. 1982. Zarys budowy geologicznej czwartorz~du w rejonach Piaski ora BuczynaChojny. In: Proc. 1st Symp. 'Quaternary of Betehat6w Region', 14-35.

-

-

JAHN, A. 1975. Problems of the periglacial zone. PWN, 223 pp. KRZYSZKOWSKI,D. 1984. Interpretacjapaleogeograficzna piozumu gleby kopalnej w serii Czy2dw na tle budowy geologicznej nadktadu czwartorzfdowego Kopalni Wfgla Brunatnego 'Betehatdw'. Ph.D. thesis Wroctaw Univ., 145 pp. - - , (in press). Rozw6j strukturalny osad6w czwartorz~dowej serii Czy~.6w w strefie L~kiflsko (R6w Kleszcz6wa) (with English summary: The structural development of Quaternary Czy26w Series near L~kifisko (Kleszcz6w graben area)). Kwartalnik Geol. --, (in press). Interpretacja paleogeograficzna poziomu gleby kopalnej wr osadach czwartorz~dowych kopalni 'Betchat6w' (with English summary: Paleogeographical interpretation of a Quaternary buried soil within the open-cast mine 'Betchat6w'). Roczn. Gleb. LINDNER,L. 1984. An outline of Pleistocene chronostratigraphy in Poland. Acta Geol. Polonica, 34, 28-49. RAMSAY, J. G. 1967. Folding and Fracturing of Rocks. McGraw Hill, 568 pp. VAN LOON,m. J. & WIGGERS,A. J. 1976. Metasedimentary 'graben' and associated structures in the lagoonal Almere Member (Groningen Formation, The Netherlands). Sediment. Geol. 16, 237-54.

K. BRODZIKOWSKI, D. KRZYSZKOWSKI, Department of Applied Geology, Institute of Geological Sciences, University of Wroctaw, Uniwersytecka 19/20, 50-145 Wroctaw, Poland. A. J. VAN LOON,Julianaweg 5, 6862 ZN Oosterbeek, The Netherlands.

Flame structures and associated deformations in Quaternary glaciolacustrine and glaciodeltaic deposits: examples from central Poland Krzysztof Brodzikowski & Andrzej Haluszczak S U M M A R Y : Glaciolacustrine and glaciodeltaic sediments in the Quaternary of the Kleszcz6w Graben show many types ofpenecontemporaneousdeformations. Their formation was due to the existence of important vertical anisotropies of the mechanical properties. Flame structures and associated types of deformations occur frequently. These structures are described in this contribution and their origins analysed.

The Klescz6w Graben in central Poland, some 50 km south of L6d~. (Fig. 1), has a thick Tertiary and Quaternary infilling (Brodzikowski et al. 1985a). The Quaternary is locally up to 300 m thick and shows seven glacial cycles. Abundant soft-sediment deformations occur, especially in glaciodeltaic, glaciolacustrine and fluviolacustrine deposits (Brodzikowski et al. 1987b), p a r t l y due to endogenic activity affecting both the substratum and the infilling of the graben since the Oligocene (Brodzikowski et al. 1987c). Various types of penecontemporaneous deformation structures resembling load casts are not due to unstable density gradients but to liquefaction of the underlying (often sandy) material (Brodzikowski et al. 1987d). Vertical anisotropy of the mechanical properties of the sediment played a major role, but structural instability was

GRABEN AREA I!

also caused by rapid lateral changes of the density gradients and the hydroplastic behaviour of the sediments. It is known that even small tangential current activity of overflowing water may disturb the equilibrium of sediments with sand-silt-clay lamination and cause deformations (Allen 1982). Other disturbances in (glacio)lacustrine and (glacio)deltaic sediments may be caused by turbidity currents (Kuenen 1951) and by shocks resulting from sedimentary or early-diagenetic processes, mass movements or earthquakes (D~tyfiski & Walton 1965; Mills 1983; Allen 1985). The deformation structures formed during such processes commonly include flame structures (Kelling 1958). These are frequently found in the glaciolacustrine and glaciodeltaic sediments in the Kleszcz6w Graben. T h e flames show a large

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Geological Society Special Publication No. 29, pp. 279-286.

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K. Brodzikowski & A. Hatuszczak

variety of shapes, and their sizes range from only a few centimetres (Brodzikowski et al. 1987b) to several metres (Hatuszczak & Brodzikowski 1982). Theories on flame structures

Kelling & Walton (1957) and Kelling (1958) mentioned flame structures which they interpreted as having been formed as a result of water flowing over freshly accumulated deposits in a hydroplastic state. The flames were considered to be deformed primary depositional structures developed at the contact between silt laminae and overlying sand laminae. Such flames look like small irregular diapirs, muddy clastic dykes or convolutions. Generally their vertical extent (amplitude) ranges from a few millimetres to several centimetres and is much larger than their width. It has been noticed that flame structures may show a distinct vergence. This gives a good indication of the palaeocurrent direction (Potter & Pettijohn 1963; Jerzykiewicz et al. 1976): the hinges of the usually inclined 'diapirs' are thought to be parallel to the palaeocurrent direction (Kelling & Walton 1957). Experiments with systems showing reversed density gradients (Anketell et al. 1970) have shown, however, that identical flame structures can develop when such systems are loaded differentially by means of a rolling cylinder. The different densities and kinematic viscosity of the media then cause longitudinal ridges that, in cross-section, show symmetrical flame structures (D~utyflski & Walton 1965). Other comparable structures have been proved to form as a result of load casting, the flames being constituted by the material that is pressed upwards (because of space problems) when relatively heavy sand 'balls' sink down in lighter material (especially silt and clay). If sand rests upon a water-saturated silt, a shock (e.g. an earthquake) may be sufficient to start the loading process. The longer the loading continues, the more the material in between is pressed upward, forming flames (D~utyfiski & Walton 1965). In such cases the hinges of the flame structures are discontinuous and irregular. Allen (1985) described structures from intertidal environments which he named wrinkle marks. He mentioned that their cross-sections parallel to the current direction show deformations that resemble flame structures, and he supposed that their origin is the loading of sandy current ripples in muddy sediments underneath. For that process the position upon a slope, reducing the stability of the sediment, might be important.

Recent investigations show an increasing number of structures that in specific cross-sections resemble typical flames, but have other origins. The present research is aimed at these particular deformations in two specific stratigraphic units within the Quaternary succession of the Kleszcz6w Graben, viz. the Saalian (Odranian, Dnieper, Altriss, Wolstonian) deposits and, at a lower level, the Elsterian (Sanian, Oka, Mindel, Anglian) deposits.

Geometry and origin of the deformation structures The deformations considered in this contribution are restricted to structures that in cross-section show a flame-like geometry. They all show a uniform vergence that can be related to the palaeocurrent direction, and all are developed at the vertical transition between silts and overlying sands. Small flame structures in fine-grained sediments

Characteristic flame structures, not complicated by additional deformations, have been found at the contact between silt and fine-grained sand laminae within Saalian glaciodeltaic deposits. The lamination in those sediments represents interseasonal rhythms within the summer siltysand layers (cf. Gustavson 1975; Ashley 1975). The flames are small tongues of different shapes and sizes penetrating into the overlying sandy material. Two types of flames can be distinguished. The first type consists of notch-like structures only 24 r a m high (Fig. 2A, B), resembling chevron marks due to tool skimming just above the sedimentary surface (cf. D~utyfiski & Walton 1965). Their vergence is not very clear but roughly perpendicular to, or slightly oblique with respect to, the palaeocurrent direction as determined from current ripples in the sets just on top and underneath the considered level. Most structures are found where the silty part is relatively thick and the overlying sandy part relatively thin. The second type also resembles notch-like muddy dykes. They are larger, however, (about 2 cm high) and show a rather regular 'wave length', also of some 2 cm (Fig. 2C). Their vergence is perpendicular to the palaeocurrent direction or oblique, sometimes even more than in the former case. The origin of both types is still in dispute. Both their regular appearance and their vergence seem to indicate that they should be considered as induced by currents, either directly or indirectly.

Flame structures in central Poland

281

FIG. 2. Micro-scaleand small-scaleregular flame structures from glaciodeltaic sediments.

On the other hand, it seems that the sandy top part, which apparently is indispensable, belongs to the next sedimentary unit. The small differences in vergence might be explained by processes related to differently directed post-depositional loading of the material surrounding the flames. It has not yet been established whether or not endogenic activity played a role in the genesis of these structures. Flame structures in sands

Flame structures with amplitudes of some 5 cm have been observed in Elsterian glaciodeltaic deposits at the point of contact between sandysilty laminae and the overlying (much more common) sandy laminae (Fig. 3A). The flames are strongly inclined and show vergences perpendicular to the palaeocurrent direction. They might be considered as little overturned and conical folds with lengths of up to 20 cm. The deformation of the silty laminae may have

started with reshaping of the ripples formed by relatively strong currents. This indicates a development which began during a syndepositional stage. Apparently horizontal tension took place subsequently in the hinge zones of the structures and resulted in either decollement or thrusting of the laminae. This is an indication for distinct meta- or post-depositional displacement and the present location of the flames, therefore, may be quite different from where the process started syndepositionally. Flame structures within fluidized systems

In the same Elsterian glaciodeltaic succession deformation structures have been encountered with flame-like geometry but with much larger sizes than flames show elsewhere in the Quaternary of the Kleszcz6w Graben (Fig. 3B). Two types can be distinguished. The first type occurs in poorly sorted sands with a few silt laminae. The flames form irregular

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FIG. 3. Flame structures in sands (A) and muddy flames in fluidized sand (B) within the Elsterian glaciodeltaic succession. muddy dykes with heights of some 20 cm. A horizontal section shows them as dendritic strips. Uniform vergence is present in their upper, but not in their lower, parts. This probably results from horizontal movements. These structures resulted from systems that lost their internal stability due to liquefaction of the sands just on top of the silty laminae. The second type occurs in distinctly laminated deposits. Their height can reach up to 15 cm and the structures seem to be the result of tangential forces during the deformational process. Both the properties of the silt laminae and the pattern of the flames suggest partial liquefaction and contemporaneous loss of stability. Since the upper parts have been eroded by high-density currents (cf. D:~utyfiski & Walton 1965; Allen 1982), the passage of that material may have been the trigger mechanism for the deformations.

Helictic flame structures

In the Saalian glaciodeltaic sediments, again at the boundaries between silty laminae and overlying sandy laminae, irregular flames occur that form diapirs or muddy dykes some 3 cm high and 1-2 cm apart (Fig. 4). In horizontal section they form discontinuous strips perpendicular or oblique to the palaeocurrent direction that can be reconstructed from current ripples. Contact planes indicate that the structures formed in a system with unstable density gradients. The configuration of silty laminae within the sands suggests a metadepositional origin. Water running over the sedimentary surface cannot be considered to have played a role in the genesis of these structures. It is more likely that shock waves, initiated by earthquakes or mass movements, were the trigger mechanism for a

FIG. 4. Helictic flame structures within glaciodeltaic sediments deposited during the maximum extent of the Saalian.

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Flame structures in central Poland process that started as a loss of internal sedimentary stability. Subsequently, the deformed structure became re-shaped when a turbidity current affected the sediment. Due to that current action the structures became inclined. In various cases it is clear that subsequent early-diagenetic processes made additional changes in the pattern of the structure, but that is not essential. The microscale irregularities in the diapiric flames are generally considered to result from differential mechanical properties of the overlying sands (Hatuszczak & Brodzikowski 1982). Isolated flames within other deformation horizons

In many deformation levels structures are found that have the same geometry as 'classical' flames (Figs 5 and 6). They occur in isolation or in groups. It is possible to distinguish a total of six types. The first type consists of structures that must be due to the formation of sandy load casts when strong turbidity currents locally deposited relatively thick sands, probably in the form of ripple trains. Silty diapirs formed in between the loaded sands and were orientated towards the same direction as the turbidity current that deposited the sand.

The second type also results from load casting. Here, sandy material loaded into the underlying mud laminae and later changed due to liquefaction destroying the deformation level but leaving the upper parts of the earlier formed silty diapirs intact. The vergence of these structures is variable and in practice only their geometry compares to flames. The third type consists of isolated flames developed in thin muddy laminae within glaciodeltaic rhythmites. They formed due to specific characteristics of palaeocurrents, in combination with particular mechanical properties of the muddy laminae. Variation in the geometry of the structures is considered to be a consequence of varying mechanical properties of the deformed laminae. The fourth type comprises of muddy laminae in the glaciodeltaic sediments that have been deformed by first loading and, subsequently, by currents. The distribution of the muddy material indicates the palaeocurrent direction. Flames of the fifth type have been formed by disturbance of a mud layer due to deposition of sand ripples on top of it. Consequently the deformations are typically current-induced. In the first stage the laminae became undulated in a way comparable to convolute lamination (Allen

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286

K. Brodzikowski & A. Hatuszczak

1982). Due to the permanent tangential stress, laminae broke down and decollement planes developed (cf. D~utyfiski & Walton 1965) as a result of 'plastic sliding' during which the beds became folded, and sometimes thrusted, one over another. In the last stage several decollement planes were re-shaped into flame-like structures. The last type consists of material between small sandy load casts caused by shocks. The structures were re-shaped by currents and afterwards by compaction. Only in some cases does the geometry resemble that of flames.

Conclusions The analysis of the flame structures in the Quaternary deformation levels of the Kleszcz6w Graben allows us to make three general remarks.

Firstly, the term 'flame structure' was introduced to indicate specific types of deformations. Since the term was restricted to diapirs or muddy dykes, there was a genetic implication. It is incorrect to use the name 'flame structure' only for the types described by Kelling (1958) and D~utyfiski & Walton (1965). Secondly, the variations in geometry and lithology allow a sub-division, with new descriptive terms, dealing also with the previous nature of what has been re-shaped into a flame structure. On this basis, it is possible to establish flamestructure systematics (Fig. 7). Finally, more detailed analyses of deformation levels and isolated deformation structures are required in order to get a better understanding of flame structures. Too many different interpretations exist at present, descriptions are too vague and systematics are, as yet, incomplete.

References ALLEN, J. R. L. 1982. Sedimentary Structures, their Character and Physical Basis, volume II, 663 pp.

ments of the Kleszcz6w Graben (central Poland).

Elsevier. ,1985. Wrinkle marks: an intertidal sedimentary structure due to aseismic soft-sediment loading. Sed. Geol. 41, 75-95. ANKETELL,J. M., CEGLA,J. & DZULYIqSKI,S. 1970. On the deformationstructures in systems with reversed density gradients. Ann. Geol. Soc. Pol. 40, 3-30. ASHLEY, G. M. 1975. Rhythmic sedimentation in glacial lake Hitchcock. Massachusetts-Connecticutt. In ." Glaciofluvial and glaciolacustrine sedimentation. Soc. econ. Paleont. Mineral., Spec. Publ. 23, 304-20.

D£ULY~rSKI, S. & WALTON, E. K. 1965. Sedimentary Features of Flysch and Greywackes, 274 pp. Elsevier. GUSTAVSON, T. C. 1975. Sedimentation and physical limnology in proglacial Malaspina Lake, Southern Alaska. In : Glaciofluvial and glaciolacustrine sedimentation. Soc. econ. Paleont. Mineral., Spec. Publ. 23, 249-63. HALUSZCZAK, A. & BRODZIKOWSKI,K. 1982. Flamaj strukturoj kaj akompanaj metasedimentaj asambleoj de deformoj en glacilagaj sedimentoj. Geologio Internacia, 4, 117-26. JERZYKIEWICZ,T., KIJEWSKI,P., MROCZKOWSKI,J. & TEISSEYRE, A. K. 1976. Geneza osad6w bialego sp~tgowcamonoklinyprzedsudeckiej (with English summary: Origin of the Weissliegendes deposits in the Fore-Sudetic monocline). Geol. Sudetica 11, 57-97. KEELING, G. 1958. Ripple-marks in the Rhinns of Galloway. Trans. Edinburgh geol. Soc. 17, 117-32.

BRODZIKOWSKI,K., GOTOWALA,R., KASZA, L. & VAN

LOON,A. J. 1987a. The Kleszcz6w Graben (central Poland): reconstruction of the deformational history and inventory of the resulting soft-sediment deformational structures. This volume. --,

, HALUSZCZAK, A., KRZYSZKOWSKI,D. &

VAN LOON, A. J. 1987b. Soft-sediment deformations from glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben (central Poland). This volume. --,

KRZYSZKOWSKI, D. & VAN LOON, A. J. 1987c.

Endogenic processes as a cause of penecontemporaneous soft-sediment deformations in the fluviolacustrine C~yz6w Series (Kleszcz6w Graben. central Poland). This volume. --,

HALUSZCZAK, A., KRZYSZKOWSKI, D. & VAN

LOON, A. J. 1987d. Genesis and diagnostic value of large-scale gravity-induced penecontemporaneous deformation horizons in Quaternary sedi-

This volume.

--

•

WALTON, E. K. 1957. Load-cast structures:

their relationship to upper surface structures and their mode of formation. Geol. Mag. 94, 481-90. KUENEN, P. H. 1951. Mechanics of varve formation and the action of turbidity currents. Geol. F6ren. Stockholm F6rhandl. 73, 69-84. MILLS, P. C. 1983. Genesis and diagenetic value of softsediment deformation structures--a review. Sediment. Geol. 35, 83-104. POTTER, P. E. & PETTIJOHN, F. J. 1963. Paleocurrents and Basin Analysis. 296 pp. Springer-Verlag.

K. BRODZIKOWSKI & A. HALUSZCZAK, Department of Applied Geology, Institute of Geological Sciences, University of Wroctaw, Uniwersytecka 19/20, 50-145 Wroctaw, Poland.

Genesis and diagnostic value of large-scale gravity-induced penecontemporaneous deformation horizons in Quaternary sediments of the Kleszcz6w Graben (central Poland) Krzysztof Brodzikowski, Andrzej Haluszczak, Dariusz Krzyszkowski & Antonius J. Van Loon S U M M A RY: Gravity-induced soft-sediment deformations are quite common throughout the Tertiary and Quaternary infilling of the Kleszczbw Graben. Only a limited number of horizons, however, are characterized by the abundant occurrence of large-scale load casts, ball-and-pillow structures, convolute lamination and related structures. It appears that the structures started to develop at a metadepositional stage and that the process was favoured by unstable density gradients. The further development was rather complicated, but endogenic activity in the graben must have played a role, though mainly as a trigger mechanism.

Penecontemporaneous deformation structures are not uncommon in sediments that contain a substantial amount of silt. It is well known that reconstruction of their genesis may help to unravel both the depositional history of the deposit (Nagtegaal 1965; D~utyflski & Walton 1965; Allen 1982; Mills 1983) and the rate and character of strain (Van Loon et al. 1984, 1985). A major part of the deformation is induced by gravity (sensu lato). In the well exposed Quaternary infilling of the Kleszcz6w Graben some horizons occur with relatively large-scale structures: load casts, ball-and-pillow structures, convolute lamination and comparable deformations, showing vertical dimensions of over 0.5 m and up to 3 m. Although more or less similar structures have been described elsewhere (Macar 1958; D~utyflski & Walton 1965; Anketell et al. 1970; Dionne 1971; Reineck & Singh 1973; Van Loon & Wiggers 1976; Alien 1982; Mills 1983), it seems exceptional that such a variety of structures of the large size mentioned occur in a few specific horizons within less deformed sediments.

Geological setting The structures discussed here were observed in the Kleszcz6w Graben, some 50 km south of L6d~ (Fig. 1). The graben was probably formed during the final stage of the Laramide folding phase (although the oldest infilling dates are from the Oligocene) and subsidence still continues (Baraniecka 1971, 1982; Baraniecka & Sarnacka 1971 ; Brodzikowski et al. 1987b). Up to 400 m of Cenozoic sediments have accumulated, with a local maximum of 300 m of Quaternary. This thickness allows a good reconstruction of the

depositional history for the last million years, but older Quaternary deposits are lacking (Baraniecka & Sarnacka 1971; Brodzikowski 1982b; Haluszczak 1982). At least seven glacial cycles have been recognized within the Quaternary. They are separated by either 'warmer' sediments (from interglacials or interstadials) or well developed erosional surfaces (Fig. 2). Many sedimentary deformations occur (Brodzikowski et al. 1987a); horizons with frequent large-scale penecontemporaneous deformations occur in both the 'cold' and the 'warm' sediments (Fig. 2).

Deformational structures Nine major deformational horizons have so far been detected. They occur in fluvial, fluviolacustrine, glaciolacustrine and glaciodeltaic deposits, mainly in silty, silty-sandy and organogenic beds, but also in sandy material. The sediments belong to the Piaski Series (Vistulian), the Chojny Series (Saalian), deposits from the Lower Saalian (Odranian) and Upper Elsterian (Sanian) Complexes, and the Czy~6w Series (middle Elsterian). The Kleszcz6w Graben is now being exploited and only parts of the deformation levels have been exposed to date. The proven surface area of the various levels is in no case less than about 2 km 2, and a surface area of some 10 km 2 has even been found in one case.

Vistulian Piaski Series The Vistulian (Weichselian, Wiirm, Valdai, Devensian) Piaski Series comprises a complex system with fluvial, fluviolacustrine and aeolian sediments. The series is up to 50 m thick and

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 287-298.

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F]o. 1. Geological setting of the Kleszcz6w Graben (modified after Baraniecka & Sarnacka 1971 ; Po2aryski 1971 ; Kossowski 1974). (A) Main regional structural elements. 1 = Kleszcz6w Graben. 2= Jurassic covered by Cenozoic. 3 = Cretaceous covered by Cenozoic. 4 = Main faults. 5 = Base of Zechstein. (B) Simplified crosssection through the graben near the villages of Czy26w and Piaski (see C). (C) Tectonic sketch map of the Mesozoic substratum. 1 = Stratigraphic boundaries. 2 = Faults. 3 = Graben boundary. I -- Location of crosssection in (B). must have formed only within some 34 000 years: from 44 000-10 000 years BP (Baraniecka 1980, 1982; Butrym et al. 1982). An upper, a middle and a lower deformation horizon can be identified (Fig. 3).

Upper deformation horizon The highest level in the succession showing abundant large-scale penecontemporaneous deformations has developed in silty, locally organogenic deposits with a fluvial overbank facies, in combination with the underlying sands that have been deposited in fluvial channels. The upper boundary of the deformed unit is undulating and is covered locally by late Vistulian sands or by Holocene channel deposits of the Widawka River. This deformed unit, which is 40-70 cm thick, shows the largest surface area (10 km 2) found to date.

The shapes of the deformations are irregular; silty and silty-sandy pseudonodules and drop structures (Fig. 4) being the most common. The process of loading must have differed considerably from place to place. In some locations rapid deformation of silty layers apparently resulted from sudden liquefaction of the sandy substratum; elsewhere loading apparently took place slowly and gradually, resulting in silty pseudonodules locally depressing the underlying sand surface.

Middle deformation horizon The middle deformation horizon of the Piaski Series is found in the upper part of a fluviolacustrine unit with silts, fine-grained sands and organogenic layers and is characterized by continuous silty laminae in the upper part. The maximum amplitude of the deformations is some

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Lower deformation horizon The lower deformation horizon strongly resembles the middle one, in its geological setting and in its deformations. The sediments consist of silty-organogenic layers and fine-grained sands, whereas the underlying material is built up of undeformed sandy silts and sands. The lateral extent of the horizon is about 2.5 km 2, identical to the extent of the depositional unit. The deformations present are mainly ball-andpillow structures, but isolated pseudonodules are also often found. Both the internal structure of silty load casts and their direct surroundings suggest that liquefaction occurred quite frequently in the deformation horizon, but was usually restricted in area. More cohesive parts of the horizon, especially the silty intervals, show

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~ FIG. 2. Schematic Quaternary stratigraphy in the graben. 1 = Fluvial or lacustrine deposits. 2 = Glaciogenic deposits. Black dots indicate horizons with intense deformations.

2.5 m and the horizon has a surface area of at least 3 k m 2, equal to the extent of the fluviolacustrine unit that possibly continues outside the area exposed by open-cast mining. The deformation structures tend to be rather complex (Fig. 5) and mainly comprise irregular load casts, ball-and-pillow structures and various types of structures due to flowage of cohesive material or liquefaction, especially of sandy layers. The lower parts of the deformations are usually the most regular, indicating that after a more or less chaotic deformation process the final stage was slower and more continuous.

~

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FIG. 3. Composite section of the Vistulian Piaski Series (partly based on Baraniecka 1980; Go~dzik 1980; Butrym et al. 1982). 1 = Eemian sands. 2 = Lower fluvial cycle. 3 = Lower organic layer (with C-14 dating). 4 = Middle lacustrine horizon. 5 = Middle organic layer. 6 = Lacustrine sandy silts. 7 = Horizon with large-scale deformations underneath the upper organic layer. 8 = Upper lacustrine silts (with C-14 dating). 9 = Upper fluvial cycle. 10 = Upper deformation horizon. 11 = Late Vistulian aeolian sands. Black dots (12) indicate other deformation horizons. Roman numbers indicate sedimentary cycles.

FtG. 4. (Upper) Details of the upper Vistulian deformation horizon (see Fig. 3). FIG. 5. (Lower) Details of the middle Vistulian deformation layer (see Fig. 3).

o

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Deformation horizons, Kleszcz6w Graben

P57000y BP [83 0 0 O y

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level (Fig. 7). The deformed horizons show irregular convolutions as well as load casts and less frequent ball-and-pillow structures and pseudonodules. The amplitude of the deformations amounts to O.7-2 m.

)RIZON IAMMUTHUS HROGONTERI POHLIG)

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: .-...~i

5 O00y BP

FIG. 6. Composite stratigraphic column of the Pilician (middle Saalian) Chojny Series. 1= Odranian (Drenthian) glacial till. 2 = Lower organic horizon (peat layer). 3 = Main organic horizon (site Chojny II). 4 = Laminated silts underneath the middle organic horizon (site Chojny III). 5 = Fluvial sands. 6 = Wartanian (late Saalian) fluvioglacial sands. 7 = Varved proglacial sands (dated by thermoluminescence). Black dots (8) indicate deformation horizons. Roman numbers indicate sedimentary cycles. TL indicates level with thermoluminescence dating.

almost no liquefaction, but flowage structures are common. Middle Saalian Chojny Series The Chojny Series (Fig. 6) was deposited 265 000200 000 yr BP (Butrym et al. 1982) and consists of fluvial (mainly channel) deposits 20-25 m thick. The lateral extent is some 3 km z, corresponding to the size of the depositional area. Fine-grained material and organogenic sediments are rare but occur in thin, very continuous intervals. These rather rare lithological units in particular show deformations. The most distinct disturbances occur in the uppermost organogenic

Lower Saalian and upper Elsterian glaciodeitaic sequences Various deformation horizons occur within the lower Saalian (Odranian, Drenthian, Dniepr, lower Wolstonian) and upper Elsterian (Sanian, Anglian, Mindel, Oka) successions (Fig. 8). Unfortunately these sediments are strongly deformed by endogenic activity (graben subsidence). Nevertheless, three main deformation horizons could be distinguished (Fig. 2), each with an area of some 2 km 2. The three horizons are very similar. Most deformations (Fig. 9) seem to have taken place at a metadepositional stage. They result mainly from unstable density gradients and thus could be considered as load structures sensu lato. Most structures are very complex and in some places their original lamination is completely destroyed due to liquefaction. Small diapirs and elastic dykes are another indication of fluidisation. Elsewhere the original lamination was preserved and only flowage took place. Some load casts and ball-and-pillow structures indicate a relatively short deformational period. The deformation horizons are occasionally intersected by faults. At such intersections there tends to be a discontinuity in the general deformational appearance. Consequently it seems that the faulting took place contemporaneously with the metadepositional deformations. Middle Elsterian Czy~bw Series The Czy~6w Series lies embedded between two Elsterian tills, but was deposited during an interstadial as a fluvial and lacustrine complex (Brodzikowski 1982a; Brodzikowski et al. 1987c; Krzyszkowski 1984). Various levels with intense deformations are present (Fig. 10), but only the two most important ones, those with metadepositional disturbances, are dealt with here. Upper deformation horizon

The deformations occur in a level of silts and peats, and continue into the underlying sands. Consequently the structures may reach amplitudes of 2.0-3.5 m. The horizon extends over some 2 km 2, equal to the depositional area (cf. Brodzikowski et al. 1987b). Deformations are so abundant that the horizon

FIG. 7. Details of the main Pilician deformation horizon (see Fig. 6).

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Deformation horizons, Kleszcz6w Graben iF

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SAALIAN CYCLES

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FIG. 8. Schematic Elsterian and lower Saalian glacial cycles (after Brodzikowski 1982b). 1 = Glacial till. 2 = Fluvioglacial sediments. 3 = Glaciodeltaic and glaciolacustrine sediments. 4 = Relative scale for intensity of endogenic activity. 5 = Endogenic activity. 6 -- Horizons with (metadepositional) deformations. 7 = Distinct changes in sedimentary conditions. 8 = Relative scale for abruptness of facies transitions. 9 = Facies transitions. 10 = Sedimentary cycle. 11 = Sedimentary subcycle. 12 = Horizon with large-scale deformations. has almost been split up into isolated structures, mainly narrow and vertically elongated load casts, separated by sandy diapirs (Fig. 11A). The diapirs apparently formed when loading took place on such a scale that the liquid limit of the underlying water-saturated sands was exceeded. Reconstruction shows that these penecontemporaneous disturbances formed simultaneously with a number of faults in the underlying unconsolidated sediments. The most convincing proof is the difference in frequency of occurrence of drop structures on either side of the faults. In their turn, the large faults in the underlying sediments are related to graben tectonics.

The horizon essentially contains deformations due to unstable density gradients (Fig. l lB), mainly pseudonodules developed in silts and peats. Lateral transitions into undeformed parts of the horizon indicate that the disturbances took place intraformationally, i.e. when the affected layer had already been covered by younger (nonaffected) sediments. There is a relationship between the deformations and faults in the underlying sediments similar to that in the upper deformational horizon of the Czy~6w Series.

Lower deformation horizon

It appears that many, quite different factors must be considered to be responsible for the formation of typical deformation horizons within the Kleszcz6w Graben. The vertical lithological alternations and the degree of pore-water pressure

The lower horizon (Fig. 10) is found in a layer about 1 m thick extending continuously over the 2 km 2 that has been investigated thus far.

Discussion

z94

K. Brodzikowski et al.

FIG. 9. Metadepositional deformations in Saalian glaciodeltaic (A, B) and glaciolacustrine (C) successions.

Deformation horizons, Kleszcz6w Graben

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metadepositional loading processes (s.1.) can easily develop (cf. Macar 1958; D2utyfiski & Walton 1965; Kuenen 1965; Allen 1982; Mills 1983). Under such conditions deformations can easily be triggered by a slightly different thickness of freshly accumulated material. Whereas increased accumulation on top of an unstable substratum results in loading, a relative lack of accumulation may result in diapirism. Reversed density gradients and kinematic viscosity are not essential for those processes (Anketell et al. 1970); the sedimentary micro-patterns may give sufficient pressure differences (Brodzikowski & Van Loon 1979, 1983; Brodzikowski 1982c).

oo • TU-,558000y BP

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,( FIG. 10. Composite stratigraphic column of the ?Malopolonian (middle Elsterian) Czy26w Series (based on Brodzikowski 1982c; Krzyszkowski 1984). 1= ?Nidanian (early Elsterian) glacial till. 2 = Lowermost organic horizon. 3 = Lower varvite succession. 4 = Lower lacustrine unit. 5 = Fluvial sands. 6 = Middle lacustrine rhythmites. 7 = Middle organic layer (peat and buried soils). 8 = Upper overbank sediments. 9 = Sanian (late Elsterian) glaciofluvial sands. 10 = Sanian glacial till. Black dots (11) indicate other deformation horizons. Roman numbers indicate sedimentary cycles. TL indicates level with thermoluminescence dating. (depending on the degree of water saturation) are two of the most important parameters. The former determines vertical anisotropy, whereas an increase in the latter results in a proportional increase of the sediment's ductility. Unstable density gradients The structural stability of water-saturated unconsolidated sediments is generally very low and

Graben tectonics as a trigger Many of the deformation horizons described in this contribution have different origins. Deformations were not caused by a difference in pressure exerted by overlying sediments (and especially 'heavy' material sinking down into 'lighter' sediments). This can be deduced from the frequent occurrence of opposite situations: 'light' silts and peats that formed load casts and comparable structures in 'heavy' sand. In such cases there are sufficient indications that the underlying 'heavy' material (usually sands or silty sands) lost its strength due to sudden liquefaction, whereas the overlying, finer-grained material remained cohesive. This phenomenon is explained by the sealing of pores at the upper sand boundary by the overlying silts and/or clays. The lithostatic pressure (pore-water pressure in the sands) increased due to the continuing accumulation of material, but the seal prevented water from escaping. As long as depositional uniformity was maintained, the pressure increased, but no deformation took place. The high viscosity of the sand-silt interface must be considered one of the main factors responsible for the relatively long-lasting equilibrium. Only locally weak zones in the fine-grained material might become intruded by the pressurized water. Small-scale diapirs, clastic dykes and waterescape structures only were formed in such cases. Situations must have occurred, however, during which the layers concerned broke up simultaneously and almost completely, everywhere. The large number of deformations indicates that intrusion of liquefied sand into a single weakness zone cannot be held responsible. In such a case the water pressure would have decreased by escape of water (and sand) at that specific point only. It may be that such processes happened simultaneously but not at hundreds or thousands of places and not repeatedly within a relatively short time interval. There must, therefore, have

296

K. B r o d z i k o w s k i et al.

FIG. 11. Metadepositional disturbances in the Czy26w Series. (A) General view. (B) Detail of (A). been a trigger mechanism repeatedly and simultaneously destroying the existing equilibrium over the entire surface. Shocks are the most logical explanation, especially since it is known that graben activity continued during the entire depositional period of the sediments. If earthquakes resulting from graben activity are accepted as a main trigger mechanism, many small- and large-scale deformations within the

Kleszcz6w Graben could be explained. Inhomogeneities within the sediment must be held responsible for the numerous locations where deformation (e.g. diapirism or loading) took place. During the liquefaction of sand, silty layers 'floated' on top of the liquefied sand and were folded according to the resistance provided by the sand. Repeated shocks resulted in more complex structures, and silty layers might have

Deformation horizons, Kleszcz6w Graben mQX

E

D I___

@

VII Vl t _ _ _

IV

B

.

III

A

i ¸ ~

/.

297

b e c o m e b r o k e n up and d e f o r m e d plastically with fold axes directed completely differently. If the liquefaction of the sand had lasted long enough, and if the overlying fine-grained material had been sufficiently deformed, structures would have formed that looked like load casts (but that should be considered as material squeezed up b e t w e e n d e f o r m e d r e m n a n t s of a more cohesive layer). T h e rather u n c o m m o n d e v e l o p m e n t of these 'load casts' was facilitated by a favourable succession of sediments (Brodzikowski 1982a, b) a n d the still active g r a b e n f o r m a t i o n resulting in rather frequent earthquakes. It is well k n o w n that the endogenic activity also resulted in other p e n e c o n t e m p o r a n e o u s d e f o r m a t i o n s (Brodzikowski et al. 1987a, b). F r o m the position of d e f o r m a t i o n horizons in the lithostratigraphic c o l u m n (Fig. 12) it can be d e d u c e d that there were alternating periods with more and less tectonic activity.

FIG. 12. Schematic composite stratigraphic column of the Quaternary. 1 = Glacial till. 2 = Fluvioglacial sediments. 3 = Glaciodeltaic and glaciolacustrine sediments. 4 = Fluviolacustrine sediments. 5 = Fluvial sediments. 6 = Organic sediments. 7 = Horizons with metadepositional deformations. 8 = Occurrence of earthquakes. 9 = Occurrence of faulting. 10 = Distinct subsidence of the graben substratum. 11 -- Duration of tectonic activity. I-II = ?Nidanian (early Elsterian) glacial cycles. A = ?Matopolonian (middle Elsterian) interglacial. III = Sanian (late Elsterian) glacial cycle. B = Ferdynandovian interglacial. IV-V = PreOdranian and Odranian (early Saalian) glacial cycles. C = Pilician (middle Saalian) interglacial. VIVII = Wartanian (late Saalian) glacial cycle. D = Eemian interglacial and Vistulian glacial cycle. E = Holocene.

.. ....'

i @

II [ .....

I C....

References ALLEN, J. R. L. 1982. Sedimentary Structures, their Character and Physical Basis, vol. II. pp 663. Elsevier. ANKETELL,J. M., CEGLA, J. & D~ULYNSKI,S. 1970. On the deformational structures in systems with reversed density gradients. Ann. geol. Soc. Pol. 40, 3-30. BARANIECKA, M. D. 1971. Dorzecze Widawki na tie obszaru marginalnego stadiatu Mazowiecko-Podlaskiego (Warty) w Polsce (with English summary: The Widawka drainage basin as a part of the marginal area of the Masovian-Podlasie (Warta) stadial in Poland). Bull. Inst. Geol. 254, 11-36. --, 1980. Osady zimnego zbiornika jeziornego zlodowacenia Vistulian. ln." Proc. 52nd Meeting Polish Geol. Soc. 318-22.

--,

1982. Dotychczasowe i bie~ce badania czwartorz~du rejonu betchatowa. In: Proc. 1st Syrup. 'Quaternary of Betchat6w Region', 1-13. & SARNACNA,Z. 1971. Stratygrafia czwartorz~du i paleogeografia dorzecza Widawki (with English summary: The stratigraphy of the Quaternary and paleogeography of the drainage basin of the Widawka). Bull. Inst. Geol. 254, 157-259. BRODZIKOWSKI,K. 1982a. Deformacje osad6w nieskonsolidowanych w obszarach ni~owych zlodowacefi plejstocefiskich na przyktadzie Polski potudniowozachodniej (with English summary: Deformations in unconsolidated sediments in areas glaciated during the Pleistocene with south-west Poland as an example). Acta Univ. Wratislaviensis 574, Studia Geogr. 34, pp 87.

298 --,

K. Brodzikowski et al.

1982b. Problemy wyksztalcenia modelu sedymentacyjnego okres6w glacjalnych w rowie tektonicznym, ln: Proc. 1st Syrup. 'Quaternary of Betchat6w Region', 46-103. --, 1982c. Wst~pne wyniki badafi sedymentologicznych i strukturalnych osad6w fluwiolimnicznych mi,dzymorenowej serii 'Cz%6w' (strefa BuczynaChojny). In: Proc. 1st Symp. 'Quaternary of Betchatdw Region', 180-98. --, GOTOWALA, R., HALUSZCZAK, A., KRZYSZKOWSKI, D. t~ VAN LOON, m. J. 1987a. Softsediment deformations from glaciodeltaic, glaciolacustrine and fluviolacustrine sediments in the Kleszcz6w Graben (central Poland). This volume. - - , KASZA, L. & VAN LOON, A. J. 1987b. The Kleszcz6w Graben (central Poland): reconstruction of the deformational history and inventory of the resulting soft-sediment deformational structures. This volume. --, KRZYSZKOWSKI, D. & VAN LOON, A. J. 1987c. Endogenic processes as a cause of penecontemporaneous soft-sediment deformations in the fluviolacustrine Cz%6w Series (Kleszcz6w Graben, central Poland). This volume. VAN LOON, A. J. 1979. Comparison of metasedimentary structures and their genesis in some Holocene lagoonal sediments in The Netherlands and Pleistocene (Mindel) glaciofluvial sediments of Poland. Bull. Acad. Pol. Sci. 27, 95105. --, 1983. Sedimentology and deformational history of the Jarosz6w Zone (Sudetic Foreland). Geol. Sudetica 18, 121-96. BUTRYM, J., BARANIECKA,M. D., KASZA, L., BRODZIKOWSKI, K., HALUSZCZAK,A., GOTOWALA, R. & JANCZYK-KOPIKOWA, Z. 1982. Datowanie bezwzgl~dne osad6w czwartorz~dowych g6rnego pi~tra strukturalnego w strefach Piaski-Buczyna-Chojny odkrywki betchatowskiej. Proc. 1st Syrup. 'Quaternary of Betchat6w Region', 150-57. DIONNE, J. C. 1971. Contorted structures in unconsolidated Quaternary deposits, Lake Saint-Jean and Saquenay regions, Quebec. Rev. Geogr. Montreal 25, 5-34. D:~ULYIQSKI, S. & WALTON, E. K. 1965. Sedimentary Features of Flysch and Greywackes. pp 274. Elsevier.

GOZDZIK, J. 1980. Osady i struktury peryglacjalne z plejstocenu okolic Belchatowa. In: Proc. 52nd Meeting Polish Geol. Soc. 322-25. HAZ.USZCZAK, A. 1982. Zarys budowy geologicznej czwartorz~du w rejonach Piaski i Buczyna-Chojny. In: Proc. 1st Symp. 'Quaternary of Betchat6w Region', 14-35. KOSSOWSKI, L. 1974. Budowa geologiczna zto~a w~gla brunatnego Betchatbw ze szczeg61nym uwzgl~dnieniem tektoniki podtoza (with English summary: The geological structure of Bdchat6w brown coal deposit considering the tectonic of the underlayer). G6rnictwo Odkrywkowe 10/11,336-34. KRZYSZKOWSKI, D. 1984. Interpretacja paleogeograficzna poziomu gleby kopalnej w serii Czy~dw na tie budowy geologicznej nadktadu czwartorzfdowego kopalni w¢gla brunatnego 'Betchat6w'. PhD. thesis Wroclaw Univ., 145 pp. KUENEN, P. H. 1965. Value of experiments in geology. Geologie Mijnb. 44, 22-36. MACAR, P. 1958. Les d6formations p6n6contemporaines de la s6dimentation. Rev. Quest. Sci. 19, 533. MILLS, P. C. 1983. Genesis and diagnostic value of softsediment deformation structures--a review. Sediment. Geol. 35, 83-104. NAGTEGAAL, P. J. C. 1965. An approximation to the genetic classification of non-organic sedimentary structures. Geologie Mijnb. 44, 347-52. PO~ARYSKI, W. 1971. Tektonika elewacji radomskowskiej (with English summary: Tectonics of Radomsko elevation). Ann. geol. Soc. Pol. 41. REINECK, H. E. & SINGH, I. B. 1973. Depositional Sedimentary Environments. pp 439. Springer-Verlag. VAN LOON, A. J., BRODZIKOWSKI,K. & GOTOWALA,R. 1984. Structural analysis of kink bands in unconsolidated sands. Tectonophysics, 104, 351-74. --, 1985. Kink structures in unconsolifine-grained sediments. In." HESSE, R. (ed) Sedimentology of Siltstone and Mudstone. Sediment. Geol. 41,269-82. - - . & WIGGERS, A. J. 1976. Metasedimentary'graben' and associated structures in the lagoonal Almere Member (Groningen Formation, The Netherlands). Sediment. Geol. 16, 237-54.

hated

K. BRODZIKOWSKI,A. HALUSZCZAK& D. KRZYSZKOWSKI,Department of Applied Geology, Institute of Geological Sciences, University of Wroctaw, Uniwersytecka 19/20, 50-145 Wroctaw, Poland. A. J. VAN LOON, Julianaweg 5, 6862 ZN Oosterbeek, The Netherlands.

Deformation of Scottish Quaternary sediment sequences by strong earthquake motions Colin A. Davenport & Philip S. Ringrose S U M M A R Y: Deformed Quaternary sand, silt, clay and peat sequences from four localities in Scotland are interpreted as seismites by the deformation history of liquefied, fluidized, slumped, and faulted units, and by comparison with seismites documented elsewhere. Two of the sequences consist of outwash sands and silts, deposited during the recession of the Devension ice cap in the region of Perth. Other deposits consist of glacial sands and interglacial peats in the Western Highlands, and glacio-lacustrine silts in the Glen Roy area; both probably deposited during the late stages of the Loch Lomond Readvance. Careful evaluation of internal geometries and deformation sequences reveals an intimate association of ball-and-pillow, fluidization, dish, load, flame and 'fault-grading' structures which are most satisfactorily explained by prolonged ground shaking produced by shallow earthquakes. In some cases, potential fault sources for the inferred earthquakes have been identified.

Engineering geological studies related to earthquake hazard assessments in Scotland include the mapping and analysis of soft-sediment deformation structures occurring in favourably preserved sequences of Late and post-glacial Quaternary lake sediments and kindred deposits (Davenport & Ringrose 1985). To date, four areas have been investigated; the lowlands of Perthshire and Angus, and the Central and Western Highlands (Fig. 1). In contrast with deposits which have been deformed by slope instabilities (monotonic*), simple sediment loading (static*), and periglacial forces (dynamic*), these deposits exhibit internal geometries, relationships and deformation sequences best explained by prolonged strong vibratory ground motion associated with shallow earthquakes (cyclic*), i.e. the deposits are palaeoseismites. (Asterisked terms are defined in Table 1.) Moreover, certain of the deformation structures compare closely with those known to have been produced in situ by earthquake activity or experimentally by induced shock, viz: (a) Ball-and-pillow structures observed in silty sediments of the lower Van Norman reservoir, correlated with known earthquakes in the San Fernando area, California (Sims 1975). (b) Load structures experimentally produced by shock inducement (Kuenen 1958). (c) A 'fault-grading' stratigraphy observed in the Miocene Monterey Shales, Santa Barbara, California, interpreted as a seismite (Seilacher 1969, 1984). A number of mechanisms of soft-sediment defor-

mation can be recognized in the Scottish deposits, involving liquefaction*, fluidization*, slope movements and structural changes*. Deformation styles and mechanisms are illustrated by detailed consideration of selected examples in the sections below. These sequences are believed to be the first reported Quaternary seismites in the UK.

fORm

FIG. 1. Map of Scotland showing the four localities of soft sediment deformation: K--Kinloch Hourn, G-Glen Roy, M--Meikleour, and A--Arrats Mill.

From: JONES,M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 299-314.

299

C. A. Davenport & P. S. Ringrose

300 TABLE 1. Explanation of terms

Overall phenomenon--LIQUID1ZATION--formation of liquidized sediment. (i.e. loss of shear strength in a two phase material: cohesionless particulate solid and fluid.) Mechanisms of liquidization change in state of the solid phase FLUIDIZATION:Change in mobility of fluid phase such leading to suspension of particles in fluid phase that the solid phase is supported by fluid drag. (when effective stress reaches zero).

LIQUEFACTION :

Means of achieving liquefaction STATIC: Solely by an increase in pore fluid pressure-DYNAMIC:Monotonic." change in solid phase structure solid phase de-stabilizing as effective stress reaches by sufficiently large single impulse load. Cyclic." zero. change in solid phase by increasing cyclic strain under vibratory loads Participating phenomena (in liquefaction) Pore water pressure, finite structural change (produced by gravitational instability), cyclic mobility, various regimes of fluidization, dewatering

Meikleour Site description A sequence of deformed structures has been exposed in an excavation through a late Devensian (Paterson 1974) outwash terrace near Meikleour, 6 km south of Blairgowrie, Perthshire. This locality was brought to our attention by Dr I. B. Paterson of the British Geological Survey, who also observed similar structures in a former gas pipeline excavation cut through the same deposit 1.5 km to the north. The terrace consists mostly of very fine sand p h i = 2 - 5 ) with occasional coarser sand units, with grains up to 2 mm in diameter (phi = 0), and horizons with pebbles up to 10 mm in diameter. Clayey silt material is observed in lower portions of the deposit. A general coarsening upwards sequence of well sorted, subangular grains, with cross bedding in coarser horizons is consistent with deposition in a prograding fluvio-glacial outwash plain. The terrace dips gently southwards with an average gradient of 2.5 m k m - 1 (Paterson 1974). At this exposure, undeformed bedding is very close to horizontal in the plane of the section (approximately E-W). The nearest kettleholes are over 3 km away from the site, to the north, and therefore local deformations associated with kettlehole formation would not be expected. Upon locating ball-and-pillow structures in a natural exposure at the site, a large face was excavated mechanically, in order to display the extent and relationship of the structures. No limit to the deformation features was exposed, despite extending the cuttings 23 m laterally, and 9 m vertically. A small cutting 150 m to the south in the same terrace did not reveal severe deforma-

tion, but contained minor faulting. To date, the stratigraphic relationship between this and the main exposure has not been established.

Deformation phenomena Portions of the main exposure are shown in Figs 2 and 3. Discussion below focuses on internal deformation relationships and micro-structures, within the large cut face, as a means of identifying deformation mechanisms. Structures are better preserved and more abundant in the upper portion of the exposure, as can be seen in Fig. 2. Detail of this upper portion is illustrated in Fig. 3. Key horizons are indicated in both figures, namely a thin cross-bedded sand layer (L) and a well preserved truncation surface (M). Vertical and horizontal grid lines shown on both figures are used for location, e.g. H3.0/V5.0 (H = horizontal; V = vertical; numbers in metres). In the sub-sections below, a number of examples of subaqueous deformation features and processes are given.

Loading Loading of the horizon L layer sands into silts beneath is observed throughout its length, for example at H1.9/V4.9 (Fig. 3). A detached sinking layer seen between H0.5/V4.9 and H1.5/ V4.9 (Fig. 3) shows extensional microfaulting (left) developing into complete detachment (right).

Loss of shear strength Destruction of fabric, loss of lamination, evidence of flow and relationship to preserved structures are collectively used to interpret portions of the

Deformation of Scott&h Quaternary sediment sequences

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deposit which substantially lost their shear strength during deformation, either by liquefaction or fluidization mechanisms. The structureless zone of silts immediately below the horizon L sand layer exemplify this phenomenon.

Fluid flow Sub-vertical channels are frequently observed, for example at: (a) H3.3/V5.0 (Fig. 3), where an undeformed layer of medium grained sand has been intruded by underlying fine sands and silts, causing upturning of the sand layering (material was channelled horizontally above horizon L before reaching this channel). and at, (b) H1.5/V4.5; H1.9/V4.5 (Fig. 3), where fine sand and silt has been channelled upwards around a large pillow, breaching the horizon M truncation surface which displays dark red oxide staining and more abundant clay (shown in heavy tone) in the vicinity of these channels.

Mass flow Three phenomena indicate mass flow in the L to M interval: (a) Overall thickening of this interval in the centre of its exposure with respect to the peripheral regions.

(b) Local thrusting due to horizontal movement (of not more than 20 cm) as indicated by the imbrication of horizon M at H3.1/V4.4 (Fig.

3). (c) The distribution of ball-and-pillow structures suggesting that a large volume of liquefied material has migrated upwards to form an upper structureless zone, through which sand layers have disaggregated and deformed as pillows; the pillows now being concentrated in the lower portion of the layer. It may be significant that the two expulsion channels above horizon L occur over the zone of maximum thickening, and also that local thrusting of horizon M occurs between centres of disruption through that horizon (at H1.9/V4.5 and H3.7/V4.4--Fig. 3). These phenomena tend to suggest injection of material and fluid into the L-M interval, loss of shear strength in finer layers, small lateral movements of portions of the layer and eventual expulsion of material and fluid into overlying sediment. It should also be noted that large scale lateral movements, as would be expected in soft-sediment slump processes, are not present in the exposure; moreover, pillows show little sign of rotation and thrusts are minor features.

Truncation surfaces Several surfaces truncating deformation structures are indicated in Fig. 2. These surfaces are only preserved in the upper layers, the best preserved being horizon M. The pillows concen-

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trated above horizon M generally have bases which remain intact, whereas pillows beneath are sharply truncated (e.g. H4.3/V4.3-Figs 3 and 4A). The only place where truncation of bases of pillows overlying horizon M is observed is where the truncation surface has developed imbricate thrusting (H3.2/V4.4-Fig. 3). These truncation surfaces are believed to be the result of erosion by fluid and liquefied sediment as they flow past and away from descending pillows, and also by later 'thrusting' action along the layering. The former mechanism results in truncation of the underlying material only, whereas the latter also results in truncation of pillow bases. Truncation horizons below M are less well preserved. This appears to correlate with increasing deformation with depth. This phenomenon may be explained by either: (i) simultaneous liquefaction of susceptible layers with the duration increasing with depth, or (ii) liquefaction having been initiated at depth, persisting, and subsequently developing at gradually shallower depths. In both cases, duration of liquefaction and the resulting deformation by gravitational instability and fluid flow would increase with depth.

Balls, pillows and tulips A great variety of form is observed in the balling of coarser material into a liquefied matrix. Detail of four pillow-like structures from different depths

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are illustrated in Fig. 4 (A, B, C, D), and discussed below. 1 Fig. 4(A)--H4.3/V4.3 : A suite of pillows made of medium sand shows original layering bent round parallel to the basal surface, and truncated at the top surface (horizon M). Homogeneous, slightly finer sandy layers below the pillows have probably been winnowed during liquefaction to provide the clay which coats the base of the pillows. Deformation and flow structures are seen between the pillows, indicating mass and fluid flow which produced and accommodated the pillows during formation. Continued channelling to the right of the central pillow has penetrated the horizon M truncation surface. The channel is lined with clay, and a trail of clay blebs below indicates injection from the structureless layer beneath. Immediately above the truncation surface, a thin layer of deformed silt occurs, thought to be the remnant of the liquefied material which flowed along the truncation surface as overlying pillows descended. 2 Fig. 4(B)-H4.3/V3.5: Two main pillow units, with deformed silts and secondarily deposited clays in between, have winnowed clay deposited at their bases. However, there are two additional clay layers, suggesting that two pillows have been destroyed and removed by the expulsion of material laterally. The remnants of one of the destroyed pillows are apparent in the thin sand layer above the clay layer second from the right. Later fluid channel development has locally disrupted the pillows causing them to begin to

C. A. Davenport & P. S. Ringrose

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FIG. 4. Detail of four pillows in the Meikleour section. Pillow locations given by co-ordinates. (Sand: light stipple, silt: unornamented, clay: black.) break up into separate balls. There is an increasing intensity of channelling from X to Y. 3 Fig. 4(C)--H3.4/V3.0: A detached tulipshaped pillow of medium sand shows some bedding features, and lies within a matrix of liquefied silt. The geometry and structure is more complex than the pillows A and B. The 'basal' clay deposit here mantles most of the pillow, and has been removed, broken or folded at the base. A truncation surface is evident at the top but has been broken by dewatering and injection of pillow material at X, and by channelling of liquefied matrix at Y. 4 Fig. 4(D)--H3.0/V0.9: A tulip-shaped pillow with a tightly-folded internal structure, appears to exhibit remnants of a truncation surface now curved round to form the top and right sides of the pillow. Vertical channel development, involving substantial mass transport of liquefied matrix, encapsulates the pillow. A disrupted clay raft, derived from depth, is seen to the left and displays original layering, injection by silt (for example at X) and erosion by channel flow (at Y). The tulip shape of the pillows in (C) and (D) results from the development of a 'pinched stem'

feature in the lower side. Its origin is uncertain, but possibly results from rapid movement of liquefied matrix, laterally and upwards, creating a lower pressure zone at the base. That the tulip shape is a later stage in pillow development is apparent from the fact that the winnowed clay deposit surrounding the pillow is drawn out with the stem, as seen in (D). Interpretation of deformation structures

The above deformation features have been described in detail in order to demonstrate a number of characteristics: (a) There is a large number of displacements involved which all occurred whilst the sediment was very soft and, at times, liquidized. (b) Changes in deformation style vertically are far greater than those occurring laterally. (c) Deformation appears to have been initiated in situ and involves very limited horizontal displacement. (d) Structures are rarely directional. (e) Periods of liquidization are likely to have been of the order of minutes to hours (as

Deformation of Scottish Quaternary sediment sequences suggested by permeabilities appropriate to sediment grain sizes). Accordingly, bulk downslope slumping, passive loading, and near-field fault displacements cannot be considered to be reasonable causative mechanisms. The above listed characteristics are, however, consistent with earthquake-induced ground shaking which triggers deformation by liquefying susceptible layers (i.e. fine sands and silts), and results in fluidization, dewatering, massflow and loading effects. To explain the overall deformation fabric and in particular the increasing deformation with depth, two scenarios are considered, and are outlined in Table 2: (a) Synchronous liquefaction initiation (Fig. 5A). (b) Diachronous liquefaction initiation (Fig. 5B). These require different interpretations for the complete sequence of deformation structures. However the two interpretations may well be complementary in terms of a more complex model. It is envisaged that the deformations are the result of ground motions produced by a single earthquake shock, typically lasting for a duration of tens of seconds. However, multiple events cannot be ruled out, either as aftershocks occurring within hours of a main event (whilst the sediment is still liquefied), or as much later main events.

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of water presumed to be very shallow. Each cycle of sediment contains strongly bedded sand and silt layers close to their tops. However only the layer above horizon L (Fig. 3) is capable of retaining its shear strength during liquefaction. This kind of stratigraphy is apparent in the less deformed zone between V4.8 and V5.2 (Fig. 3). Liquefaction is assumed to initiate simultaneously in all susceptible (fine sand and silt) layers, and the increase of deformation with depth is taken to suggest that post-ground-shaking movements continued longer and/or were more intense at depth due to decreasing pore water escape rates and increasing normal stress with depth.

Diachronous liquefaction initiation A similar stratigraphy to that outlined in the previous section is assumed. However, here liquefaction is considered to have been initiated at some depth where a susceptible layer occurs within a zone of maximum acceleration. [This zone of maximum response-acceleration is inferred from experimental work and is thought to occur at depths of between 3 and 10 m in thick cohesionless soils (Dikmen & Ghaboussi 1984).] Liquefaction is later achieved at other shallower susceptible layers as cyclic loading progresses, and as dewatering of lower earlier liquefied layers results in increased pore water pressures.

Arrats Mill

In this scenario it is envisaged that a sequence of coarsening upward cycles, varying in thickness from less than 0.5 to over 1.0 m, stand in a body A)

305

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FIG. 5. Alternative liquefaction models for the interpretation of the Meikleour deposit. [I : Liquefaction initiation curve, T: Termination of liquefaction curve, C: Increasing cohesion (assumed cohesionless above).] (A) Simultaneous liquefaction initiation. Termination curve primarily a function of efficiency of fluid escape. (B) Diachronous liquefaction initiation. Initiation curve resulting from response of soil profile to ground shaking; termination assumed simultaneous for simplicity.

Site description

An outcrop of outwash sands, presumed to be late Devensian (Paterson, pers. comm.), has been exposed at a landfill site, 4 km east of Brechin, County Angus. The site was reported to contain ball-and-pillow structures (Paterson, pers. comm.), and major excavations using on-site earth moving equipment, carried out during the summer of 1984, revealed a large volume of material deposited within a depression having exposed dimensions of 7 0 m wide and 100 m long. Cuttings are typically between 2 and 4 m in depth. The deposit infilling the depression contains material very similar to Meikleour, with very fine sands ( p h i = 2 - 5 ) accounting for most of the volume, and several layers of medium sand with grains up to 0.5 mm (phi = 1). Occasional pebbly layers are present, and thin clay bands are observed in lower layers. This deposit, however, displays many more abundant laminations than Meikleour, and cross-bedding has not been observed. This suggests the deposit may have

Events

Evidence Homogenous zones (loss of shear strength) Pillow morphology; clay mantles; detachment below horizon L; truncation surfaces

Overall geometry; greater fluid streaming at depth (where pressure difference is greatest) Channelling better preserved in upper layers; thrusting on horizon M

Advanced pillow deformation (tulip shape) and broken truncation surfaces at depth

Events

1. Simultaneous liquefaction of fine sands and silts,

2. Settling of upper layers of each sediment cycle as pillows; liquefied matrix flows upwards depositing winnowed clay on pillow bases; truncation surfaces produced by erosion

3. Inversion in each cycle throughout the sequence; upper cycles stabilising first where dewatering is most easily achieved

4. Later dewateringandexpulsionofliquidized sediment from deeper layers disrupts upper layers, which have regained some strength, and causes 'thrusting' and lateral displacement

5. Restricted dewatering of deeper layers resulting in greater strains at depth. Final dewatering and sediment expulsion may have occurred through large sand dykes (not revealed)

5. Regaining of shear strength with escape of fluids to surface or overlying high permeability layers

4. Lateral movement of upper layers resulting from expulsion of fluid and liquidized sediment from below

3. Upward escape of pore fluid from susceptible layer hastening and aiding liquefaction of sediment cycles above

2. Settling of upper layer of susceptible sediment cycle as pillows; upward flow of liquefied matrix depositing winnowed clay on pillow bases; truncation at base of susceptible layer

1. Liquefaction of initial susceptible layer at depth.

(B) Diachronous liquefaction initiation

(A) Synchronous liquefaction initiation

TABLE 2. Sequence o f events during liquefaction at Meikleour as envisaged in the two models

Fluid injection channels above horizon L

Thrusting on horizon M located between fluidization channels

Overall geometry; fluidization channels related to pillow morphology in upper layers

Pillow morphology and distribution with depth; relict pillows and truncation surfaces at depth

Assumed to be located within lower highly deformed zone

Evidence

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Deformation o f Scottish Quaternary sediment sequences been a small glacio-fluvial lake rather than a river channel. The deformation structures are observed in the laminated lacustrine material, and occur to depths of 3 m in the central region, and to progressively shallower depths towards the edges. The top of the deformed sediment is overlain by flat-lying undeformed laminated sand. In contrast to the Meikleour deposit, much more can be said here about the relationship of deformed material to surrounding undeformed deposits, as portions of both the upper, lower and lateral limits of deformation are exposed.

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Limits of deformation The base of deformed material through much of the section is immediately above a thin zone of clay layers interbedded with sand (at 0.73 m on Fig. 6). Degree of deformation increases rapidly up from this base, from slight up-turning and breakage (at 0.80m, Fig. 6), through more strongly upturned layers (at 0.85 m, Fig. 6) to strongly deformed layers (at 1.3-1.5 m, Fig. 6). It appears that the clay zone has acted as a fluid seal, restricting escape of pore water from the overlying sands. The top of deformed material can be seen at 2.8 m in Fig. 7 and in the detail of Fig. 8, and comprises a horizontal surface, containing broken clay clasts, overlain by undeformed, laminated, pebbly sands.

Pillow and dish structures A similar process of pillow formation to that described for the Meikleour deposit is envisaged. Pillows at Arrats Mill are less complex, and rarely develop into complete balls. A portion of a large pillow can be seen between 0.9 and 1.2 m in Fig. 6. Rotation of pillows and lateral movement are not apparent, and structures generally give the impression of being in the early stages of formation. At higher levels, the pillows become less curved and generally smaller, and eventually begin to resemble the dish structures described by Lowe & LoPiccolo (1974) among others. They can be seen between 2.4 and 2.6 m, Fig. 7 and in Fig. 9. They typically consist of a faint dirty layer up to 5 mm thick, with oxide staining and local blebs of clay and organic matter (e.g. bottom left

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corner of Fig. 8). Their relationship to original lamination is difficult to appreciate as they are found mostly in massive, possibly homogenized, sands. However, the fact that they form a continuum with the pillows beneath, which do have lamination parallel to bases, does suggest they also follow primary laminations. Furthermore, close to the top of the deposit, two semicontinuous laminae, which are almost certainly primary, appear to be incipient on dish formation (see centre of Fig. 8). Lowe & LoPiccolo (1974) attribute dish structures to gradual dewatering of underconsolidated beds. We would tend to attrib-

C. A. Davenport & P. S. Ringrose

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Flotation of organics Small blebs of clay-rich, organic matter are indicated by heavy dots in Figs 6-8. They tend to occur in sinuous vertical trails. Between 0.9 and t.5 m in Fig. 6 the trails can be seen to originate from layers within the pillow, and in Fig. 8 they appear to approach the top surface, with a few blebs at the surface itself; there is also a suggestion that they emanate from the top of the fault; hence, these trails give a strong suggestion of vertical dewatering throughout the deformed body.

Influence of slope The section shown in Fig. 9 displays flaming at the base of deformation and between incipient pillows. The flames get larger from left to right, and begin to face into the local trough seen in Fig. 9(B). Dish and pillow structures also tend to tilt towards this trough. There are reasons to suppose that this deformation was not induced by slope, but it is apparent that, once liquefaction had taken place, slopes did influence deformation style, such that structures have a tendency to face downslope.

Interpretation

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structures

The Arrats Mill section displays much simpler deformation structures than seen at Meikleour. A single deformation event is sufficient to explain the structure. Slope or sediment loads cannot be reasonably called upon as causes of liquefaction

Deformation of Scottish Quaternary sediment sequences

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ARRAT'S MILL - V e r t i c a l section through exposure at edge of deposit

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Glen Roy Site description The famous parallel 'roads' of Glen Roy, in the Western Highlands, are shorelines of a series of ice-dammed lakes which existed in Loch Lomond Readvance times (Sissons 1979). At their maximum extent the lakes covered an area of about 90 km 2 in the Glens of Roy, Gloy and Spean. Laminated silts and clays (varves) were deposited in the lake beds and typically achieve thicknesses of 1-3 m. Coarser sands and gravels form adjacent beds above and below the varves. Following the discovery of deformation structures in July 1-984 by one of us (P.R.), a sampling programme was initiated, which has currently identified 50 localities revealing deformation phenomena. The structures appear to have an areal distribution showing decreasing degrees of deformation outwards from a central zone into peripherally undisturbed sediment. The occurrence of a surface displacement in the parallel roads (Sissons & Cornish 1982), along with the presence of a number of contemporaneous landslips and a fault zone showing several signs of recent activity all within the central zone, suggest the possibililty of one or more fault-related earthquakes having caused all these phenomena. More detailed work on the nature and zonation of sediment deformation is currently in progress. The lake sediments consist of stiff laminated

silts and clays with occasional layers of sand and gravel. They occur on most of the steep-sided lake beds, but are best exposed on small perched delta fans, frequently capping broad outwash gravel fans. Only one locality (from the central zone) is considered here for the discussion of deformation mechanisms. It comprises a section a little over 2 m deep, and is illustrated in Figs 10-12. Sequences above and below the lacustrine sediments were not exposed.

Deformation phenomena

Slumping At the top of the exposure, a slump unit 0.4 m thick is seen (0.1-0.5 m--Fig. 10). The slumping can be seen to involve highly plastic flow, with preserved laminae showing tight folding. The base of the slump does not rest on a decollement surface, but rather involves gradual detachment above well preserved varves at the base. Excessive lateral movement is not apparent, and the amount of flow seems to decrease rapidly from the top of the slump downwards.

Faulting Small faults are seen through most of the exposure (Figs 10-12). They are mostly normal faults with displacements of up to 20 mm, and often varying rapidly along their length. However, a few reverse faults are present, for example at 0.7 m on Fig. 10, indicating that simple settlement or slump mechanisms are not sufficient to explain their presence.

Deformation of Scottish Quaternary sediment sequences

3II

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312

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Davenport& P. S. Ringrose

Flotation of organics This phenomenon is very similar to that described for the Arrats Mill section, but is particularly well displayed here. Organic varves in the lower portions of the section are clearly the sources for the disaggregated blebs higher up, as can be seen by debris trails leaving layers at 1.0 m (Fig. 10) and between 1.9 and 2.05 m (Fig. 11). Organicdebris trails are not seen in many portions of the section, but are seen immediately below the slump unit, where they appear to be floating up to collect on the 'pre-slump surface' (0.5 m--Fig. 10).

Fault grading stratigraphy This phenomenon was first described by Seilacher (1969, 1984) in the Miocene, Monterey Shales of Elwood Beech, Santa Barbara, California. He described a sequence showing undisturbed sediment, with occasional faults, grading upwards in a zone of dense microfaults, with throws increasing upwards, then breaking up into a rubble zone, and finally grading up into a homogeneous 'soupy' zone. He interprets the sequence as 'strongly suggesting a seismic origin'. The central portion of this Glen Roy section shows a stratigraphy remarkably similar to Seilacher's. This has been detailed in Fig. 12. All four elements of the fault grading stratigraphy are present: (i) faintly layered silts with occasional faults (1.7 m and below); (ii) a denser set of microfaults, with throws increasing upwards (1.75-1.65 m); (iii) a rubble zone, containing a broken organic layer and suspended clasts of clay (1.65-1.5 m); and (iv) a homogenized zone showing no structures apart from organic-filled fissures (1.5 m and above). A further interesting feature is the presence of small incipient sand and clay pillows with truncated tops between the fault segments at 1.63 m (Fig. 12).

Interpretation of deformation structures The presence of fault grading, flotation of organics to a free surface and faulting of both reverse and normal senses collectively suggest deformation by cyclic loading involving liquefaction, upward pore water migration and mass settlement. The slump unit at the top of the section may also result from earthquake activity. The styles of deformation in this area are different from those seen at Meikleour and Arrats Mill, primarily because the Glen Roy sediments are significantly more cohesive; hence deformation exhibited by fault-grading rather than by balland-pillows. Elsewhere in the Glen Roy area,

sand deposits do contain ball-and-pillow structures.

Kinloch Hourn Two other sites showing soft-sediment deformation have been investigated. They both occur in glacial sands and interglacial peats in close proximity to the Kinloch Hourn fault, Western Highlands. The deformation horizon is similar in both sites (7.5 km apart). It is 0.5-1.0 m thick and extends throughout exposures (i.e. for 50 and 100 m). Ball, pillow and flame structures similar to those described above are present. The top of deformation occurs within postglacial peat, which indicates that the deformation event occurred after the Loch Lomond Readvance. Their proximity to a presently active fault and similarity to other deposits again suggests a seismic origin to the deformation (Davenport & Ringrose 1985). Study of these deposits is still in the early stages, and will be discussed later elsewhere.

Concluding remarks Deformed sediment sequences within the late Quaternary of Scotland have been described, emphasising those features which are believed to be characteristic of seismites or, when taken together, suggest an initial liquefaction event. Each locality contributes to the picture. At Meikleour a number of sediment layers are involved in a complex deformation sequence which exhibits the sequential effects of loss of shear strength, loading, fluid and mass flow, and internal erosion and deposition. There is a clear increase of deformation with depth, which is considered to indicate that post-liquefaction movements continued for longer (and/or were more intense) at depth. In contrast, at Arrats Mill, the sequence represents a simpler sequence of a single sediment unit deforming and dewatering below a free (lake) surface. Here, dewatering of liquefied sediment has been achieved in a more uniform fashion, resulting in dish-structures at the top of the sequence. The thin deposits of the Glen Roy area provide a picture of lateral variations in deformation, characterized by faultgrading structures in a central zone (indicating stronger ground motions) and localized slumping, faulting and dewatering effects in a larger region. The styles of deformation differ from those of Meikleour and Arrats Mill as a result of somewhat greater quantities of clay and organic matter in the varves of the Glen Roy area, and because of their deposition on steep rock slopes.

Deformation of Scottish Quaternary sediment sequences

313

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Both Glen Roy and the less well documented localities at Kinloch Hourn occur close to surface fault zones which show evidence of being active earthquake sources, in terms of movement since glacial times and instrumental microearthquake activity (Davenport & Ringrose 1985). The deposits at Arrats Mill and Meikleour are large single exposures, 56 km apart, but are considered to have been influenced by seismic events, and possibly a common seismic source. Whereas in Glen Roy the largest deformation thicknesses involve slump material, the other

sequences reveal insignificant lateral movement. The presence of undisturbed stratigraphy above and between deformation zones, the preservation of stratigraphic order, the small amount of rotation and sliding, and the lack of slope failures, all suggest in situ liquefaction by cyclic loading. Further studies are required to demonstrate that the subject deformation structures can be formed uniquely by appropriate levels of earthquakeinduced vibratory ground motion. Our approach involves extensive excavation and logging, mathematical modelling, and natural-scale simulation

314

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of deformation tables.

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ACKNOWLEDGEMENTS: We gratefully acknowledge N E R C for the funding of this project (research

studentship--P.R.). Special thanks are offered to our colleagues in various departments of the University of Strathclyde. The co-operation of members of the British Geological Survey, Edinburgh, particularly Dr I. Paterson, along with Lord Lansdowne and North Angus County Council, has been most valuable.

References

DAVENPORT, C. & RINGROSE, P. 1985. Fault activity and palaeoseismicity during Quaternary time in Scotland--preliminary studies. In: Earthquake Engineering in Britain. Thomas Telford Ltd, London, 81-93. DIKMEN, S. & GHABOUSSI, J. 1984. Effective stress analysis of seismic response and liquefaction. J. geotech. Eng. 110, No. 5,628-58, ASCE. KUENEN, P. 1958. Experiments in geology. Geol. Mag. 23, 1-28. LOWE, D. & LoPICCOLO, R. 1974. The characteristics and origins of dish and pillar structures. J. sedim. Petrol. 44, 484-501• PATERSON, I. 1974. The supposed Perth Readvance in the Perth district. Scott. J. Geol. 10, 53-66.

SEILACHER, A. 1969. Fault-graded beds interpreted as seismites. Sedimentology, 13, 155-9. - - - , 1984. Sedimentary structures tentatively attributed to seismic events. Mar. Geol. 55, 1-12. SIMS, J. 1975. Determining earthquake recurrence intervals from deformational structures in young lacustrine sediments. Tectonophysics, 29, 141-52. SISSONS, J. 1979. Catastrophic lake drainage in Glen Spean and the Great Glen, Scotland. J. geol. Soc. Lond. 136, 215-24. -& CORNISH, R. 1982. Rapid localized glacioisostatic uplift at Glen Roy, Scotland. Nature, 286, 142-3.

C. A. DAVENPORT (~Z P. S. RINGROSE, Department of Applied Geology, University of Strathclyde, Glasgow G 1 lXJ.

Syn-sedimentary and burial related deformation in the Middle Jurassic non-marine formations of the Yorkshire Basin J. Alexander S U M M A R Y: The non-marine formations of the middle Jurassic Ravenscar Group consist of channel sandstone bodies set in interbedded mudstones, silts and thin sheet sandstones. These rocks contain syn-sedimentary and burial related deformation structures. Frequent discharge fluctuations are deduced to have enhanced inherent river bank instability and with occasional tectonic movements produced a wide range of bank collapse structures. These had a low preservation potential and are observed at various stages of erosion. Several faults associated with channel sandstones can be related neither to tectonic activity nor to bank collapse. They may be explained as compensation structures formed during differential compaction of the sandstones and adjacent mudstones. Differential compaction structures form on a smaller scale around plant material, early diagenetic nodules and sandstone dykes. Continued sedimentation on the alluvial plain, together with differential compaction, produces a positive topography over channel sandstones. This effect was reduced by variations in deposition rate and facies. Laterally continuous beds deposited on this topography form gentle folds that are exaggerated by differential compaction during burial. Although most of the deformation seen in the Ravenscar Group is related to alluvial processes or sediment burial, some is connected to syn-sedimentary tectonic activity. The non-marine formations of the Middle Jurassic Ravenscar Group (Fig. 1) were deposited on an extensive coastal plain that was periodically inundated by marine water depositing the volumetrically less important marine formations. The coastal plain had a low relief with large areas of saturated soil and shallow water table lakes. Locally levees, alluvial ridges and tectonic highs produced relief that was high enough to allow the development of well drained soils. R Scalby F o r m a t i o n A Scarborough Formation(m) iE Gristhorpe Mb Cloughton ~ o n A Formation ~ ~ G

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FIG. 1. Nomenclature of the Ravenscar Group (from Hemingway & Knox 1973) and location map. Evidence for a wide range of stream forms is preserved in these rocks; from extensive, complex sandbodies deposited by braided rivers, meanderbelts, or by the superimposition of channels, to isolated low and high sinuosity channel sandstone bodies. In all of these, as well as in the related

overbank deposits, deformation is abundant and mostly caused by syn-sedimentary and burial related mechanisms. Syn-sedimentary deformation is ubiquitous and highly variable in form and origin. Burial related deformation is less easily distinguished, but is of importance to sedimentological and architectural interpretations of these sediments. Post-depositional tectonic activity is of minor importance, expressed by a few faults, gentle warps and joints in the more competent rocks. This paper describes some of the styles of syn-sedimentary and burial related deformation observed in the Ravenscar Group and comments on the mechanism of their formation and occurrence.

Syn-sedimentary deformation Syn-sedimentary deformation is ubiquitous in the Ravenscar Group and includes a wide range of structures formed by: (i) fluid escape; (ii) density instability; (iii) slope instability and (iv) animal activity. It would be nice to be able to distinguish between forms related to inherent processes and those related to tectonic activity. This is not easy, however, as seismic shock generally leads to an intensification of inherent failure processes. It is possible to infer tectonic activity from the intensity and distribution of deformation. The small amounts of tectonic warping and faulting

From: JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks, Geological Society Special Publication No. 29, pp. 315-324.

315

316

J. Alexander

that occurred at the time when the Ravenscar Consolidation of fine sediments may be Group were deposited are shown by the thickness brought about by the deposition of sand; this and facies variations, palaeocurrent d i r e c t i o n s forces fluid out of the subjacent strata, the and erosional patterns seen in these rocks. Syn- escaping fluid producing liquefaction and fluidisedimentary soft sediment deformation results zation when forced into this newly deposited from four major causes: (i) fluid escape; (ii) sand (Lowe 1975). Observations from the 1984 density instability; (iii) slope instability and (iv) flood deposits of the Brahmaputra (Bristow, in animal activity. These are described below. preparation) demonstrated that this process may be responsible for the formation of clusters of small mud or sand volcanoes and collapse hollows Fluid escape (Fig. 2). Large sand volcanoes are also recorded Sediments are generally deposited in a saturated in relation to several recent earthquakes (Witkind state and lose water in a variety of ways. When 1960; Morse 1941; Steinbrugge & Moran 1986) the escape of water is rapid, laminae may be bent where seismic loading was the driving force. or fractured and particles may be transported These volcanoes may be random or linearly within the fluid. For forced liquid expulsion, an arranged and are related to facies distribution external force is required to pressurize the pore and water table configuration. In the Ravenscar fluid. This may be in the form of continued Group, injected sandstone dykes are common, sedimentary loading or dynamic loading, such as though sand cones are only rarely observed. The by seismic shock. When the pore water pressure distribution of these dykes is related to collapse cannot be dissipated at a sufficiently fast rate, structures and to linear features, implying that liquefaction may occur (liquefaction being de- inherent and seismic processes were active in fined by Yould et al. 1978, as 'the transformation their formation (Fig. 3). This is also implied by of a cohesionless material from a solid to a liquid the size variation observed. state as a consequence of increase in pore pressure and decrease of effective stress'). Ground failure Density instability due to liquefaction is a common occurrence during earthquakes and its extent is controlled by Density instability occurs where material with a local topography, ground water configuration, higher density is deposited on top of material sediment type and thickness. It is interesting to with a lower one. If the bearing capacity of the subjacent material is low the denser material will note also that shaking intensity measured on 'load' into the underlying material in an attempt alluvium saturated with water appears to be about one intensity unit greater than on dry to overturn the system. In rare cases, balls of sand will separate and sink through the lower strata. alluvium (Ziony & Tinsley 1983).

FIG. 2. Sand volcanoes and associated collapse hollow in flood deposits of the Brahmaputra. Scale bar 10 cm. (Photographer C. Bristow.)

Deformation in Middle Jurassic formations, Yorkshire Basin

317

FIG. 3. Sandstone dykes in the 'meander belt' of the Scalby Formation south of Scarborough. Scale bar 1 m.

This process and the resulting structures are the best illustrated form of soft sediment deformation structure (Weaver 1979; Hempton et al. 1983). Load structures are commonly found in the Ravenscar Group, but their distribution is problematical. Examples may be found where the density contrast and thickness ratios estimated for the time of deposition are the same, but loading has occurred in some instances only. It is clear that some other control exists; this may be hydraulic or in the form of an external trigger (Sims 1975; Mayall 1983; Weaver 1979). Another structure that is seen in the Ravenscar Group is best described as a 'squelch'. This is where an area of laminated sand and mud or sandstone has sunk into a hollow in the lower stratum. The edges of the dish are often fractured and may be associated with water escape structures (Fig. 4). These are usually small (50 to 200 cm in diameter) but may have a similar mechanism of formation to the earthquake craters described by Morse (1941) or the collapse hollows of Bristow (in preparation).

and the difficulty of recognizing partly eroded structures. In the Ravenscar Group, bank collapse structures are common and can be identified in all the channel forms. Small slumps and slides can be recognized in the finer channel deposits, while large rotational and complex slides are more easily recognized in the larger coarse-grained channel deposits. Laury considered that the type of bank failure most likely to be preserved in the stratigraphic record is that produced by largescale shear failure, in particular rotational sliding. In the rapidly migrating channels that deposited the fine sandstones of the meander belt (at location 1 on Fig. 1) of the Scalby Formation (Nami 1976) small collapse structures are common (Fig. 5). The large-scale bank failure structures that can be observed in the Ravenscar Group are often

Slope instability Slope instability is the major cause of soft sediment deformation in an alluvial setting due to the active processes of bank erosion. This has been well documented from modern river studies, but, as stated by Laury (1971), 'evidence of stream bank failure in the geological record is surprisingly meagre'. The main reasons for this are the low preservation potential of the failure structures

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FIG. 4. A sketch of a 'squelch' structure observed in finely interbedded sandstone and mudstone of the Scalby Formation, above the 'meander belt'.

318

J. Alexander

FIG. 5. Small bank collapse structure preserved in a lateral accretion unit in the Scalby Formation north of Scalby. The notebook is 30 cm long. related to down cutting in the associated channel, and the failure surface may in part be confined to weak material in the bank. One of the best exposed examples is in the Saltwick Formation south of High Hawsker (Fig. 6 at location 2 on Fig. 1). Here a large sand body deposited during the lateral migration of a channel cuts down southward. To the north of the thickest part of the sand body a plane of failure goes through the edge of the sandstone to a coaly layer below and passes into contorted material next to the deepest part of the channel. The sandstone above this major failure surface suffered minor faulting; these small faults are now the most obvious structures, but are difficult to interpret without first identifying the major failure surface. Earlier

bank failures are recorded within the major slide block by slump folds and small faults in laminated sand and mud. Continued migration of the channel allowed the preservation of these structures. Further to the south in the same sand body younger failure surfaces can be observed. Here the slide planes are confined to the sandstone and these smaller slides were partially eroded before being preserved. Discharge fluctuations or seismic activity may have triggered the small-scale bank failures seen in the channel deposits of the Ravenscar Group. Discharge changes appear to have occurred often, judging from the alterations in sedimentation of the sandstone bodies. In some places (for example near High Hawsker, in a channel deposit north

Deformation in Middle Jurassic formations, Yorkshire Basin

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of that described above) in situ slumped, laminated sandstone contains sub vertical root systems (Fig. 7). The colonization by plants occurred after the abandonment of that part of the channel. The good preservation of the structures suggests that they were formed around the time of abandonment and may have been triggered by the channel abandonment or by the event that caused the change in channel plan. Well preserved bank failure structures are quite common in the exhumed meander belt of the Scalby Formation near Burniston, and here, also, the best preserved examples appear to have occurred around the time of meander abandonment or river avulsion. A study of the occurrence of bank collapse structures on the meander belt shows that within a single meander bend the amount of collapse often increases down current. This may be a function of the increase of fine material within any one unit around the meander or the changing cross-sections. The pattern of sediment slumping on these point bars gives a component down the sedimentary dip and often also a deflection in the down-current direction, possibly related to fluid shear. There is also a random element that may result from small rheological changes, animal activity or a preferred tectonic stress direction. At Cromer Point (location 3 on Fig. 1) a series of cross sections may be examined through a sand body that was deposited by a relatively low sinuosity migrating stream. The relationships between slides, contorted strata and intraformational conglomerates are clearly shown in the complex exposures and the structures observed at different stages of erosion (Fig. 8). In the best preserved examples, contorted lamination overlies a discontinuity that is interpreted as a failure surface on which sediment slipped down towards the channel. The contorted rocks grade down the sedimentary dip into intraformational conglom-

319

erates that have a greater channel-parallel extent than the contorted material, with a major extension in the down current direction. Erosion produces isolated remnants of the component structures and leads to resedimentation of the intraformational conglomerates. In the larger sandstone bodies of the Ravenscar Group, cross-bedding is common but overturned cross sets are rare. The mechanism responsible for the formation of these structures has been discussed in detail (McKee et al. 1962; Allen & Banks 1972; Jones & Rust 1983), but with no clear consensus of opinion. The deformed crosssets in these rocks appear to be confined to areas near faults considered to have been active at the time of sandstone deposition. The best preserved examples can be seen in the Scalby Formation south of Scarborough and in the Saltwick Formation west of Whitby. Conclusions on soft sediment deformation

Syn-sedimentary, soft sediment deformation complicates the sedimentary picture and introduces new structures to the rocks. Failure surfaces produced by syn-sedimentary deformation may be reactivated or altered during burial or subsequent structural deformation. The control of facies on structural development has often been described and it seems likely that existing structures and weaknesses may have a similar influence. The occurrence and distribution of syn-sedimentary deformation seen in the Ravenscar Group is summarized in Fig. 9. This is a simplified picture of deformation on an alluvial plane that shows neither the complex interaction of the controlling factors, nor the detailed morphology of the structures produced. Burial related deformation

In general, during burial of homogeneous sediments the overburden stress will be evenly distributed, resulting in a laterally uniform rate of consolidation. Alluvial deposits are far from homogeneous and the resulting complex strain behaviour with increasing overburden stress is poorly understood. Due to the inherent lateral facies and depositional rate variations in alluvial settings, the overburden stress will be laterally variable. These effects, although only significant in the top few metres of the sediment and small in magnitude, have a significant effect on the floodplain topography and consequently on the resulting alluvial architecture. If there is a persistent lateral overburden stress variation, caused by stationary facies relationships controlled for example by

320

J. Alexander

FIG. 7. Slumped laminated sandstone with sub-vertical roots, from cliffs near High Hawsker. localized tectonic subsidence, the effects of differential compaction will be considerable. Consolidation of an inhomogeneous material produces structures resulting from variations in the amount of compaction. Small-scale structures may be seen around early diagenetic concretions, sandstone dykes, and plant material. Larger structures are produced by the greater volume reduction of over-bank fines than channel sand deposits during burial. Several authors (e.g. Raiswell 1971; Oertel & Curtis 1972) have observed that laminae in the host material bend conformably around the edges of siderite concretions and that within concretions the laminae curve towards the major axis of the nodule with increasing distance away from the centre. The thickness variation between

laminae may be used to measure the vertical shortening, assuming that the laminae were originally parallel. In addition, an estimate of compaction factor may also be obtained from the ratio between the volume percentage of cement within and outside the nodule. This assumes that the cement is non-displacive and pore filling (Raiswell 1971; Gautier 1982). One of the problems in both these methods is the uncertainty of the depth at which the nodule grew. Previous studies have used differential compaction structures to demonstrate an early diagenetic origin for concretions, but it is not so easy to judge where in the top few metres of sediment concretion growth started. These structures are useful however, to give a minimum measure of compaction of the host rock.

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FIG. 8. Fence diagram of Cromer Point to show the relationship between slides, contorted strata and intraformational conglomerate. Lateral variations in compressibility also exist where desiccation cracks or fissures of tectonic origin (Witkind 1964) occur in fine sediment and become infilled by coarser material. Downward tapering sandstone and silt dykes are common in the Ravenscar Group and are observed to have

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been buckled during sediment compaction. Shelton (1962), assuming the sandstone in such dykes to be incompressible and the dykes originally vertical, calculated the amount of compaction from the shortening of the dykes. It is clear from structural considerations that the best results from this method will be obtained from narrow dykes and that the results could be of great use, as they record the compression since the fissures were formed at the depositional surface. Estimates using this method on dykes in the Scalby Formation produce results that are much smaller (ranging between a decompaction number of 1.18 and 2) than those from theoretical predictions (about 2.4 based on considerations of the depth of burial), or from the observations of siderite nodules. This discrepancy may be explained by the fine nature of the dykes allowing internal deformation and by the small-scale bulging observed on the sides of the dykes. Dykes observed in more competent material, such as laminated silt and mudstone, give higher values of vertical shortening despite the lower compressibility of the laminated sediments. Where plant material (logs or branches) is preserved in the Ravenscar Group, it is often accompanied by disturbance of the surrounding material. In some instances this disturbance may be related to sedimentary processes, for example around log jams in channel sand bodies. More ~

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322

J. Alexander

often it is caused by differential compaction. The plant material has a much higher compressibility than the surrounding sediments (this effect has commonly been observed with regard to coal) and often results in small subsidence structures during burial (Fig. 10). Some of these may be demonstrated to have occurred at shallow burial depths and to have effected the sedimentation at the palaeosurface. Occasionally logs with a sediment infill may be observed in the Ravenscar Group and these may be used, as suggested by Elliott (in press) to calculate the amount of compaction that occurred up to the time of infill (by the variation from circular, assuming that an originally cylindrical log is observed in perpendicular cross-section) and the amount of subsequent compaction (from the distortion around the log). This is difficult to assess, as the log will initially compact more rapidly than the surrounding sediments, but later the rates may be reversed creating complex structures. Features not so well described are those that are caused by differential compaction around sandstone bodies. In the Ravenscar Group, the grain to grain contacts within the sandstones are rarely sutured and the inferred quantity of dissolution is negligible. The physical deformation of the sandstone is therefore insignificant. This would be expected for rocks that have not been deeply buried. Perrier & Quiblier (1974) observed that the porosity of the average sand does not decrease greatly down to 500-600 m of burial. The Ravenscar Group have not undergone much greater burial. The overbank fines have suffered volume reduction during burial demonstrated by the structures described above. As a result of this volume change large structures are formed, and the sandstone body geometry is altered during the burial history. The effects of differential compaction are most extreme around major sandstone bodies, particularly those of large vertical extent. This is usually difficult to observe in outcrop and it is often better to consider it theoretically. From consideration of alluvial architectural models (Bridge & Leeder 1979; Crane 1982), it is clear that sand body concentration will be produced when aggradation rates are low, when channel activity is effected by localized subsidence or under less predictable hydraulic situations. A low aggradation rate or a braided river system may lead to the deposition of a sheet sandstone and the surface topography will effect the deposition and compaction of the overlying sediments and isolate the compactional structures in the strata above and below. The effects of sandstone sheet surface topography may be observed above the 'meander belt' in the

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FIG. 10. Deformation around a buried log in the Scalby Formation south of Scarborough. Scalby Formation (Nami 1976). Hollows in the surface of the sand sheet were often the site of organic soil formation; this material has a high compressibility and may have led to the maintenance of the hollow and created an unusually thick deposit of this facies. Features similar to this are observed in the Nottinghamshire Coalfield where unusual thicknesses of coal occur in abandoned channels on the surface of larger sand bodies (Elliot 1965). The high compressibility of the coal or organic soil, in this environment, will lead to a large amount of differential compaction over the sand body sufficient to produce faulting at the edges of the depression. This effect was observed during the New Madrid 1811 earthquake (Morse 1941) when the seismic shaking caused consolidation of the abandoned channel fills and resulted in localized subsidence. Sandstone ridges produced topographic highs that affected the subsequent facies distribution and consolidation has led to an exaggeration of the topography in continuous strata over this surface in some places. This effect is often reduced by facies and deposition rate variations in the overlying deposit. Surface faulting due to differential compaction has also been observed in connection with water extraction. When water is artificially removed from an aquifer a volume reduction occurs. In places where there is a rapid thickness change of the aquifer surface subsidence rates vary sharply and this can create enough stress to cause faulting or fissuring of the land surface (Holzer 1984). Vertical stacking of sand bodies occurs adjacent to syn-sedimentary faults and in subsidence hollows. In both of these cases a difference in depositional thickness is likely. This will be exaggerated by the greater compactibility of the fine material deposited on the topographic highs (the reverse of the case for small topographic features produced by inherent floodplain processes, described above). If the upthrow side of a fault is dominated by fine sediments, and the hanging wall by channel sandstones, differential

Deformation in Middle Jurassic formations, Yorkshire Basin compaction will exaggerate the original thickness difference. This is further complicated by erosion, pedogenesis and warping. In the compaction of horizontal strata by uniform overburden pressures, the strains are unlikely to reach values sufficient to produce failure. If the loading is uneven or the compressibility is variable then the situation becomes more complicated. Where the thickness of a sandstone bed changes rapidly, or at the edge of a sand body, consolidation of the surrounding fine sediments will result in localized shear stress that may be enough to cause faulting. This is shown in Fig. 9 in a simplified way. Geometrical considerations of small faults near a sandstone body (location 4 on Fig. 1) in the Scalby Formation suggest that these may be compensation structures produced during burial (Fig. 11). It is probable that similar structures are produced around larger sand bodies, but these are difficult to recognize in the field. It is clear that failure will occur preferentially along existing planes of weakness and that existing structures will be complicated in the process. The geometry of basal river bank failures will be considerably altered and compensational

323

movement may occur along small growth faults. The angular relationships of structures will be altered during burial and the dip of faults will be reduced in compressed sediments. Although many different factors influence the compactional response of clay sediments to increasing overburden, it is clear that a general relationship between overburden and compaction exists (Curtis 1977). From this assumption it is possible to process measured sections and remove the expected amount of compaction. This may be further enhanced using a computer to calculate the effects of slight variations in compaction to reproduce the original geometry of structures. This is obviously a subjective process at present, but it makes it easier to assess the importance of variations in compressibility and test the origin of suspected differential compactional structures.

ACKNOWLEDGEMENTS: I would like to thank Dr M.

Leeder for supervision, C. Bristow for stimulating discussion and photographs from Bangladesh, Rob Gawthorpe and Sophie Clarke for helpful comments, and British Petroleum for financing my studentship.

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FIG. 11. Compensational structures formed during burial caused by variations in compaction rate around a small channel sandstone body in the Scalby Formation at Burniston Wyke. Rapid thickness variations seen in a 'decompacted section' (produced in this case by the use of a constant compaction factor for the mudstones) not formed by sedimentary processes, can be related to changes in the amount of compaction. In this example small faults have developed at the points of change in amount of compaction.

324

J. Alexander

References ALLEN, J. R. L. & BANKS, N. L. 1972. An interpretation and analysis of recumbent folded deformed crossbedding. Sedimentology, 19, 257-83. BRIDGE, J. S. & LEEDER, M. R. 1979. A simulation model of alluvial stratigraphy. Sedimentology, 26, 617-44. CRANE, R. C. 1982. A computer modelfor the architecture of the avulsion-controlled alluvial suites. Unpublished PhD. thesis. Reading University. CURTIS, C. D. 1977. Sedimentary geochemistry: environments and processes dominated by involvement of an aqueous phase. Philos. Trans. R. Soc. Lond. A. 286, 353-27. ELLIOTT, R. E. 1965. Swilleys in the coal measures of Nottinghamshire interpreted as palaeo-river courses. Mercian Geologist 1(2), 133-41. --, 1985. Quantification of peat to coal compaction stages, based especially on phenomena in the East Pennine Coalfield, England. Proc. Yorks. geol. Soc. 45(3), 163-72. GAUTIER, D. L. 1982. Siderite concretions: indicators of early diagenesis in the Gammon Shales (Cretaceous). J. sed. Pet. 52(3), 859-71. HEMINGWAY, J. E. & KNOX, R. W. O'B. 1973. Lithostratigraphical nomenclature of the Middle Jurassic strata of the Yorkshire Basin of Northern England. Proc. Yorks. geol. Soc. 39, 527-35. HEMPTON, M. R., DUNNE, L. A. & DEWEY, J. F. 1983. Sedimentation in an active strike-slip basin, Southeast Turkey. J. Geol. 91(4), 401-12. HOLZER, T. L. 1984. Ground failure induced by groundwater withdrawal from unconsolidated sediments. Geol. Soc. Am. Rev. Eng. Geol. 6, 67-101. JONES, B. G. & RusT, B. R. 1983. Massive sandstone facies in the Hawkesbury Sandstone, a Triassic fluvial deposit near Sydney, Australia. J, sed. Petrol. 53(4), 1249-59. LAURY, R. L. 1971. Stream bank failure and rotational slumping: preservation and significance in the geologic record. Bull. geol Soc. Amer. 82, 1251-66. LOWE, D. R. 1975. Water escape structures in coarsegrained sediments. Sedimentology, 22, 157-204. MAYALL,M. J. 1983. Earthquake induced syn-sedimen-

tary deformation in a Late Triassic (Rhaetian) lagoonal sequence, South West Britain. Geol. Mag. 120(6), 613-22. MCKEE, E. D., REJNOLD, M. A. & BAKER, C. H. 1962. Experiments on intraformational recumbent folds in crossbedded sands. U.S.G.S. Prof. Paper 450-D, 155-60. MORSE, W. C. 1941. New Madrid earthquake craters. Bull. seism. Soc. Amer. 31, 309-19. NAMI, M. 1976. An exhumed Jurassic meander belt from Yorkshire, England. Geol. Mag. 113, 47-52. OERTEL, G. & CURTIS, C. D. 1972. Clay-ironstone concretions preserving fabrics due to progressive compaction. Geol. Soc. Amer. Bull, 83, 2597-606. PERRIER, R. & QUIBLIER, J. 1974. Thickness changes in sedimentary layers during compaction history; methods for quantitative evaluation. Am. Asoc. Petrol. Geol. Bull. 58(3), 507-20. RAISWELL, R. 1971. The growth of Cambrian and Liassic concretions. Sedimentology, 17, 147-71. SHELTON, J. W. 1962. Shale compaction in a section of Cretaceous Dakota Sandstone, Northwest North Dakota. J. sed. Petrol. 32, 873-77. SIMS, J. D. 1975. Determining earthquake recurrence intervals from deformation structures in young lacustrine sediments. Tectonophysics, 29, 141-52. STEINBRUGGE, K. V. & MORAN, n . F. 1956. Damage caused by the Earthquakes of July 6 and August 23, 1954. Bull. seism. Soc. Amer. 56, 15-33. WEAVER, J. n. 1979. Seismically induced load structures in the basal coal measures, South Wales. Geol. Mag. 113(6), 535-43. WITKIND, I. J. 1960. The Hebgen lake, Montana, Earthquake of August 17, 1959. Billings Geol. Soc. Guidbook, l l th Ann. FieM Conj. 31-44. --, 1964. Reactivated Jaults north of Hebgen Lake. U.S. Geol. Survey Pro./'. Paper 435G, 37-50. YOULD, T. L., ASCE, M. & PERKINS, D. M. 1978. Mapping liquefaction induced ground failure potential. J. geotech. Enging. Div. Proc. Am. Soc. Cir. Eng., 104, 433-46. ZIONY, J. I. & TINSLEY, J. C. 1983. Mapping the earthquake hazards of the Los Angeles Region. Earthqu. Inform. Bull. 15, 134.

J. ALEXANDER,Department of Earth Sciences, The University of Leeds, Leeds LS2 9JT.

Aspects of veining in the Welsh Lower Palaeozoic Basin Bill Fitches S U M M A R Y: In the Welsh Lower Palaeozoic marginal basin, veins of diverse origin are widely distributed in space and time. The veins are categorized as pre-, syn- and post-tectonic with respect to folding and cleavage development; these sub-divisions denoting different types of veining processes. Among the pre-tectonic veins are those described as regional veins which were emplaced diachronously by hydraulic jacking and related processes due to overpressured pore fluids affecting newly lithified rocks while sedimentation continued above. Other pre-tectonic veins are those around Snowdon which were due to processes closely associated with the development of an Ordovician resurgent caldera. Syn-tectonic veins are uncommon. They include the auriferous lodes of Ogofau in south-central Wales. Post-tectonic veins are the most widespread type and are exemplified mainly by those of the Harlech and Plynlimon Domes. These late veins are attributed largely to regional extension and dewatering of the Lower Palaeozoic pile caused by the early Variscan extensional regime.

The Welsh Basin (Fig. 1) evolved throughout Lower Palaeozoic times as a marginal basin on the SE side of the Iapetus suture (Phillips et al. 1976). The sedimentation history of the E and SE flank of the basin has been reviewed by Woodcock (1984a, b) and the offshore environments in W Wales are discussed by Cave & Hains (in press). The magmatic activity which accompanied Ordovician stages of basin development is reviewed by Kokelaar et al. (1984). Most rocks were altered by anchimetamorphism during burial (Bevins & Rowbotham 1983) and in parts of N Wales

FIG. 1. Outcrop of the Welsh Lower Palaeozoic Basin, showing areas discussed in the text.

middle greenschist facies conditions were reached (Roberts 1981; Roberts & Merriman 1985). Almost all Welsh Lower Palaeozoic rocks have been at least slightly deformed. It is now recognized that deformation was a semi-continuous process throughout the evolution of the basin. Faulting and broad warping, for instance, have been shown to have accompanied sedimentation and volcanism in several areas (Kokelaar 1979; Campbell 1984; Reedman et al. 1984; Fitches & Campbell 1987). Woodcock (1984a, b) has proposed that strike-slip tectonics played an important role during Ordovician to Silurian times, causing the progressive development of folds and cleavage in deeper parts of the sedimentary pile while deposition continued above. Deformation which produced compressional structures culminated in late Silurian to early Devonian times, then gave way to an extensional tectonic regime. Veins are commonplace in the Welsh Basin. For the most part, however, they have received scant attention and usually only those containing minerals of economic interest have been studied in detail. An exception is the suite of nonmetalliferous veins documented from many parts of the basin by Fitches et al. (1986). Most studies have concentrated on particular vein provinces, have been concerned primarily with metallogenesis and have not considered the veins in the context of the tectonic development of the whole Welsh basin. This paper summarizes published and unpublished data on the better known examples of metalliferous and barren veining that have accrued especially during the last 15 years. The main objectives are to illustrate

From: JONES, M. E. & PRESTON,R. M. F. (eds), 1987, Deformation of Sediments and Sedimentary Rocks,

Geological Society Special Publication No. 29, pp. 325-342.

325

326

B. Fitches

from these examples the variety of different styles of veining, vein-forming processes and relationships between veining and deformation that are found in this marginal basin and also to examine aspects of the various roles and histories of the fluids involved in vein development. In the following sections, the veins are divided according to their age relationships with respect to regional tectonics, specifically fold and cleavage production. This categorization does not necessarily have absolute time connotations because regional deformation was perhaps diachronous. It does, however, have implications regarding the prevailing regional stress conditions, the sources and roles of fluids, and the mechanical properties of the host rocks during the production of veins.

Pre-tectonic veins The two suites of veins described in this category are very different in many respects. One suite is widely distributed in space and time, comprises non-metalliferous veins and developed diachronously during the progressive burial of clastic sediments; the veins of this suite are here termed Regional Veins. The other suite is confined to an Ordovician volcanic centre in Snowdonia in North Wales, contains metalliferous as well as barren veins and, to a large extent, owes its origin to magmatic processes.

Regional veins Regionally distributed veins with a common, pretectonic origin are described in detail by Fitches et al. (1986). They are recorded from late Ordovician rocks in W Wales, are particularly common in Llandovery strata in West and Central Wales and are also known in Wenlockian rocks of NE and East Wales. Most host rocks were deposited in or between submarine fans and are chiefly siliciclastic sediments. The veins have two main forms: by far the more common are bedding-parallel, between 0.5 mm and 50 cm in thickness and some exceed 25 m in length. In places they are abundant and spaced at intervals of a few centimetres over several metres of strata. In West and Central Wales, these veins comprise chiefly ferroan dolomite and quartz with variable amounts of chlorite, rare pyrite and exceptionally traces of galena. In contrast, the bedding-parallel veins of NE and East Wales are dominated by calcite, usually with some quartz and chlorite. Less common are bedding-normal veins, of similar dimensions to the other type and spatially

closely related to them. Quartz dominates the steep veins of West and Central Wales, although ferroan dolomite is abundant in some examples; in NE and East Wales, calcite predominates. Both types of vein are considered to be pretectonic because they were folded, cleaved and locally boudinaged with their host rocks. Fitches et al. (1986) inferred that the host strata had undergone partial compaction, dewatering and local slumping before veining. Some stages of diagenetic alteration also preceded the veins, notably the growth of various types of concretions and the mimetic enhancement of clay-mineral compaction fabrics by chlorite-white mica stacks. The bedding-parallel veins are regarded as products of hydraulic jacking by overpressured pore fluids (cf. Price, in Fyfe et al. 1978; Guth et al. 1982; Stoneley 1983). Upward migration of fluids through the sediment pile was impeded by the many layers of argillite, whose permeability had previously been greatly reduced by compaction and diagenesis. Fluid pressures reached high values in the argillites and, exploiting the low tensile strength of the bedding plane compaction and mimetic fabric, produced fractures parallel to bedding. The fluids deposited minerals on the roofs and floors of the cavities opened along the fractures. Most bedding-normal veins are attributed to fracturing of the cavity roofs as explained by Price (in Fyfe et al. 1978) or, where strata were slightly inclined, to down-slope extension (Price 1977). In both cases, high fluid pressure and local nearly horizontal tensile stress combined to create the fractures. By these processes, a stack of bedding-parallel veins, in part inter-connected by steep veins, propagated up through the pile during progressive burial. It is considered likely that the depth at which fracturing took place remained more or less constant so that successively younger strata were affected as they were depressed to appropriate levels. The depth of vein formation is not known. Cone-in-cone concretions, which had grown before the host sediments had reached depths required for fracturing, may give approximate minimum depths. Marshall (1982) inferred from oxygen isotope data that similar concretions in Mesozoic strata grew at depths of a few tens to a few hundreds of metres. A major component of the fluid which created the fractures and deposited the vein minerals is likely to have been sea-water which had been trapped in, and progressively buried with, sediment. Other components may have come from water released by dehydration of smectite and other clay minerals during diagenesis (Bruce

A s p e c t s o f veining in the W e l s h B a s i n 1984) at or below the level of fracturing, and from metamorphic dehydration processes operating deep in the sedimentary pile. Modification of fluid composition would have resulted from chemical interaction with the sediments. A conspicuous feature of the regional veins is the close correlation between their mineralogy and that of their host rocks; ferroan dolomite is common to the veins and host rocks of West and Central Wales, calcite dominates the veins and is abundant in the sediments of NE and East Wales. This correlation perhaps implies that the fluids responsible for veining had not circulated, or at least had not significantly leached sediments, more than a few hundred metres distant from the sites of vein deposition. Snowdonia veins Reedman et al. (1985) proposed that the metalliferous veins located in an area ofc.25 km 2 around Snowdon are related to a submerged resurgent caldera of late Soudleyan-Longvillian age (Fig. 2). Most veins are located in the top parts of the Lower Rhyolite Tuff Formation, a unit largely ponded within the caldera wall and composed chiefly of volcanic material erupted during initiation of the caldera, and also in basal parts of the Bedded Pyroclastic Formation erupted during waning stages of, and after, caldera resurgence. Veins also occur in Soudleyan sedimentary rocks below the volcanic units. The veins are steep, discordant to bedding and contain mainly quartz with smaller amounts of pyrite, chalcopyrite and sphalerite with local concentrations ofpyrrhotite, galena, magnetite and other metalliferous minerals. Almost all the veins lie within the caldera and are concentrated at the northern and southern margins and in the NE trending apical halfgraben produced during resurgence. Reedman et al. (1985) recognize two sets of veins. The veins of one set are mainly restricted to, and are aligned parallel with, the half-graben. Single veins are small, usually less than 30 cm in width and less than 30 m in length, but locally they form zones up to 200 m in width. The veins of the other set are more widespread, trend 280-320 ° and reach lengths of 300 m. In addition to the metalliferous veins described by Reedman et al. (1985), barren veins composed of quartz, usually accompanied by traces of chlorite, are abundant in places. These veins have been examined by the author in the Cwm LlanCwm Tregelan area to the S of Snowdon summit. The area lies near the apical graben fault system and the veins cut lower parts of the Lower Rhyolitic Tuff Formation and underlying coarse

327

clastic sedimentary rocks. The veins are small, mostly less than 2 cm in width and a few metres in length, but in places they form dense arrays occupying some 25% of the rock mass. They are steep, inclined at high angles to bedding and trend in two main directions, 050-080 ° and 340350 ° (Fig. 3). In most localities, one trend or the other predominates, but almost everywhere there is a wide range of trends. The veins cut and locally offset each other, but no systematic pattern has been recognized whereby veins with one trend consistently offset those with another trend. It seems, therefore, that veins with various orientations were emplaced more or less contemporaneously. The metalliferous and barren veins are pretectonic with respect to folds, which have steep NE to ENE axial planes and are mostly large scale, and to the strong axial planar cleavage which affects this part of Snowdonia. Reedman et al. (1985) quote some of the author's evidence for relationships between veining and deformation and further amplification is given here. Veins initially oriented at high angles to cleavage have been folded (Fig. 4A) and shortened by up to 40.5% (calculated by I. Wilkinson, pers. comm.), whilst those originally aligned close to cleavage have been boudinaged. Sulphides in the veins are also clearly pre-tectonic; some prismatic grains have been folded with their host vein (Fig. 4B), others have incipient (Fig. 4C), or well-developed, quartz-fibre pressure fringes (Fig. 4D) aligned in cleavage. Wall rock fragments in vein breccias contain a common, uniformly aligned cleavage. Reedman et al. (1985) considered the deformation to be a late Silurian-early Devonian event and thereby inferred that vein emplacement is constrained to a c. 50 Ma period after accumulation of the host rocks. Their preferred model of vein formation invokes a modified Kuroko-type process. Sea-water which permeated the calderafill was circulated by convection driven by heat from the same high level magma chamber, sited below the caldera, that led to resurgence and late stage rhyolite dome emplacement. Most metals in the veins are considered to have been leached from the volcanic rocks and underlying sedimentary rocks, although fluorine enrichment in some vein wall rocks (R. Fuge, pers. comm.) and the presence of traces of tungsten in some magnetite bearing veins may imply a magmatic component to the hydrothermal fluid. Of fundamental importance to vein formation is the caldera-setting. Reedman et al. (1985) have recorded the control on the locations and trends of the metalliferous veins by the caldera margin and the apical half-graben, which suggests that

328

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fracture development was governed by stress systems related to caldera evolution. The association of mineralized fractures with caldera development has been widely documented in other volcanic terrains (e.g. Steven & Ratte 1965; Smith & Bailey 1968; Lipman et al. 1976). Smith & Bailey (1968), in particular, drew attention to the radial and concentric fracture systems that accompany caldera subsidence, especially the resurgence stages when apical

graben may develop to give an additional linear fracture pattern. The fracture patterns are explained by principal tensile and intermediate stress trajectories assuming a radial and annular disposition with respect to the caldera margin during caldera collapse and again during the resurgence of the underlying sub-cylindrical magma chamber. The vein orientations and positions in the Snowdon caldera are broadly consistent with

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FIG. 3. Geological map of the Cwm Tregalan-Cwm Llan area south of Snowdon, simplified from unpublished 1: 10,000 maps prepared by A. J. Reedman of BGS. Rose diagrams show the orientations of pre-tectonic quartz veins measured at six sites. Synoptic rose diagram for all measured veins shown left centre. those of other calderas. Because most of the host rocks accumulated during or after caldera collapse, the majority of fractures are likely to be products of resurgent activity rather than initial subsidence. While the fractures appear to be readily explained in terms of stress systems associated with caldera evolution, the role of fluids and the processes whereby the fractures were mineralized is less obvious. To explain this type of mineralization, it is usually accepted that thermal convection was set up by magmatic heat beneath, or in, the volcanic pile. Sea-water percolated down through the pile, scavenging silica, sulphides and other mineral components e n r o u t e . Some fluid may have penetrated the top of the magma chamber by thermal cracking (Lister 1983; Sleep 1983) thereby collecting mineral components directly from the pluton. Downward passage of fluid may have been facilitated by fractures produced during caldera subsidence and resurgence, but unless the volcanic pile had been

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extensively lithified, most fractures would have readily rehealed. In the absence of numerous open fractures and the tendency for the fluid to leach rather than precipitate minerals as it was heated during descent, few veins are likely to mark the downward passage of fluid. Fluid descent would have been accomplished instead, initially at least, largely by migration through pores. During prolonged percolation, however, clogging of pore spaces, notably by silica, may well have taken place in higher temperature parts of the caldera-fill (Sleep 1983; Henneberger 1984). Then, any fractures that had remained open would have become increasingly important passage ways and may have become sites of mineral deposition. Convective rise would return the fluid, now mineralized and mixed with fluid of magmatic origin, toward the surface. Again, bulk migration probably took place through pore spaces, but with prolonged circulation and cooling of fluid during ascent, clogging would place increasingly strong constraints on migration. In these circumstances the free flow of fluids was impeded and fluid pressures could rise to values in excess of the normal fluid hydrostatic head. The effect of high fluid pressures in rocks by now extensively lithified, especially by silicification, and set in the tensile stress regime of the caldera, was to create hydraulic fractures and vein breccias marking the upward propagation of normal faults and tensile fractures (Phillips (1972); also discussed below, p. 335). Accordingly, most of the veins in the Snowdon area are considered to have been generated by the pressure of the fluid which mineralized them. The widespread vein breccias composed of dispersed and disjunct wall rock clasts cemented by quartz, sulphides and other minerals are the visible expressions of these processes. It is unlikely that the veins originated as open fractures formed soon after volcanic accumulation which were then used passively by later circulating fluids.

Syn-tectonic veins Almost all rocks in the Welsh Basin are at least gently folded and mildly cleaved. It might be expected that fluids would have been particularly mobile during deformation, producing prolific veins. Yet syn-tectonic veins appear to be much less common than those preceding or following deformation. However, the possibility remains that they are common, but have been largely ignored because they contain no minerals of economic interest. Moreover, demonstrating that

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Fro. 4. Pre-tectonic veins and sulphides from Cwm Tregalan-Cwm Llan area, south of Snowdon (see Fig. 3). (A) Folded veinlets in volcanogenic sandstone from Site ! (Fig. 3). (B) Photomicrograph of folded quartz veinlet and iron sulphides in cleaved meta-basalt, 500 m SE of Site 1 (Fig. 3). Bar scale 0.5 mm. Plane polarized light. (C) Photomicrograph of incipient pressure fringes (arrowed) on iron sulphides in meta-basalt, from same locality as (B). Bar scale 0.25 mm. Cleavage direction marked by black line. Crossed polars. (D) Photomicrograph of sphalerite grain (Sp) with quartz fibres (q) in pressure fringe in hydraulic fracture breccia, from same locality as 4B. Wr = wallrock fragment. Bar scale 0.5 mm. Crossed polars.

vein formation was synchronous with deformation can be difficult. In m a n y parts of the Welsh basin there are localized examples of small syn-tectonic veins; tension cracks fanning around outer arcs of smallscale folds, saddle reefs in fold hinge zones (e.g. Fitches t972, Fig. 2) and tension gash arrays produced by flexural slip in fold limbs. These veins comprise quartz, ferroan dolomite and chlorite in various proportions and, exceptionally, small amounts of pyrite and chalcopyrite. Other syn-tectonic veins are located in dilation zones along the faults which developed as accommodation structures in some fold cores. Syn-tectonic veins which are well-documented, mainly because of their economic history and potential, are the auriferous quartz veins of

Ogofau in South Wales (Steed et al. 1976) (Fig. 5).

Ogofau veins The host rocks of these veins are predominantly black shales, spanning the Ordovician-Silurian boundary, with minor intercalations of sandstone, siltstone and pyritic carbonaceous shale. Sulphides and gold in the pyritic carbonaceous shales are considered by Tater (1975) to be syndepositional or diagenetic. The rocks are in the north west limb of the Cothi Anticline, a subsidiary fold of the Towy Anticline, and are deformed by tight to isoclinal, mesoscopic and small folds with axial planes dipping steeply N W ; axial planar cleavage is strongly developed

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in the argillites. Thrusts and shear zones occur along, or close to, the axial surfaces of several folds. Annels & Steed [in Shrestha (1975)] inferred that later transverse folds, on ENE axial planes, were imposed on the limbs of larger folds, perhaps accompanying the development of thrusts. Still later normal faults strike NW-SE. The veins, composed exclusively of quartz with small amounts of pyrite, arsenopyrite and minor gold, are of four main types [Annels & Steed in Shrestha (1975)]. 1 Saddle reefs, concordant with bedding, in fold hinges. The biggest example is the Roman Lode, an S-shaped vein up to 3.5 m thick, located in the hinges and common limb of an anticline-syncline pair. Smaller veins with similar form are recorded by Tater (1975). 2 Irregular stringers discordant to bedding as exemplified in the footwall of the Roman Lode. 3 Tension gash arrays located mainly in fold limbs. 4 Podiform veins in shear zones axial planar to folds. Annels & Steed concluded that the veins were emplaced during or immediately after deformation, in fractures and along bedding planes dilated during folding and thrusting. The sulphides in the veins are probably derived from immediately adjacent host rocks because they are concentrated at intersections of veins with the pyritic shales. Gold, from the same source, migrated preferentially to anticlinal crests. More recently, Annels (1981) recorded veins from an area NE of Ogofau Mine which have

331

different characteristics from those described above and provide additional information on their origins. The host rocks and tectonic setting of the veins are similar to those of Ogofau. However, two generations of veins are distinguishable. Early veins comprise pyrite, arsenopyrite and carbonates derived, according to interpretation of stable isotope data, from diagenetic material in the host rocks. Some early veins and vein minerals were strongly deformed during the development of cleavage. Late veins contain quartz, sulphides of lead, copper and zinc, and carbonates, which Annels (1981) suggested were derived from Ordovician volcanic rocks or deeper and older rocks. The late veins are undeformed and were emplaced in fractures produced during or after cleavage formation. Filling temperatures in the range 318-372°C were obtained from fluid inclusion studies on vein quartz, after pressure correction for 5 km depth of burial. An alternative interpretation of veining, not considered in literature available to the author, is that the saddle reefs, such as Roman lode, and some steep veins, including the earlier set described by Annels (1981), are pre-tectonic and were generated by the processes advocated to explain the regional veins (p. 326). These processes, involving very local derivation of sulphides and gold as in other explanations, would have been particularly effective in the impervious argillite-dominated strata of Ogofau. They could have produced almost all documented features except the tension gash arrays and podiform veins which remain clearly syn-tectonic. Moreover, this model would explain the steep veins, located beneath the Roman Lode for example, for which no mechanism has been suggested. It would also alleviate the problem of explaining why dilation and veining should take place in fold hinges in ductile argillites and why the Roman Lode, for example, does not die out in the common limb of the anticline-syncline pair. The spatial relationship of concordant veins in fold hinges appears anomalous in this model but can be accounted for by folds nucleating on the earlier veins. To test this alternative model, detailed studies are required on, for example, vein quartz fabrics and on detailed relationships between cleavage development and vein formation. Also of interest in the Ogofau area are the veins along late N W - S E normal faults (Steed et al. 1976). In contrast to the earlier veins, these contain galena and sphalerite, perhaps implying that the fluids depositing the veins had scavenged more widely than those responsible for the Roman Lode and associated veins. In most respects these late veins have much in common with the posttectonic veins discussed below.

B. Fitches

332 Post-tectonic

(Allen & Jackson 1985), comprising sandstones, siltstones and mudstones. These rocks are cut by concordant and discordant sheets of microdiorite and microtonalite ('greenstones') probably related to the pre-Arenig Rhobell Fawr magmatism (Kokelaar 1979). The structure of the area appears to be the product of several stages of deformation (Allen & Jackson 1985): (i) localized N N E to north trending folds related to faulting (Kokelaar 1979) which preceded the pre-Arenig Rhobell Volcanic Group; (ii) late Tremadocian north-south folds, also probably fault-related; (iii) N W - S E folds found mainly in Cambrian rocks but locally also

veins

Veins which were emplaced after folding and cleavage development are the commonest type of vein documented in detail in the Welsh Basin. Those of the SE part of the Harlech Dome and of the Plynlimon Dome (Fig. 1) are discussed here as the best known examples and brief reference is made to those of the Migneint area of North Wales and of West Shropshire. South east Harlech dome

The host rocks of these veins are the Cambrian Harlech Grits and Mawddach Groups (Fig. 6) -.

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Aspects of veining in the Welsh Basin in Ordovician strata; (iv) N N E - N E folds most widely developed in Ordovician strata and attributed to end-Silurian deformation. The Dolwen pericline and Caerdeon syncline, the major folds of the region, were probably initiated in late Tremadocian times and accentuated at later stages. Cleavage is sporadically developed, trending approximately north-south or locally N W SE; its character, spatial and temporal relationships to folds have not been investigated in detail. The veins, up to 3 km in length, lie along some of the radially disposed normal faults and mostly trend 030-070 ° or less commonly 290-315 ° (Allen & Jackson 1985). The veined faults have a staircase geometry; they are steep, brittle structures in the relatively rigid greenstones and sandstones but flatten off in argillaceous layers (Gilbey 1968) where they are represented by ductile shear zones (Ashton 1981). These faults are cut by later north-south normal faults such as the major Trawsfynydd fault. Quartz is the dominant vein mineral with small amounts of sulphides (including pyrite, pyrrhotite, arsenopyrite, sphalerite and galena), carbonates, free gold and minor gold tellurides (Gilbey 1968). Ashton (1981) has considered the relationships between faulting and vein development at Gwynfynydd, 10 km north of Dolgellau (Fig. 6). He drew the important conclusion that, after fault initiation, mineralizing fluids played an active role in promoting fault dilation and further displacement, rather than passively filling preexisting faults as had been previously supposed. Ashton's model for Gwynfynydd, which is probably widely applicable to veining in other parts of the Harlech Dome, proposes initiation of normal faulting in a regional tensile stress regime without the aid of fluid pressure. The faults were then permeated by fluids which began to deposit vein minerals and to exert sufficient pressure to dilate not only the steep sections of the faults but also the low angle sections in the argillites, there exploiting the earlier shear zones. Ashton also gives evidence for repetitive minor shear displacement and dilation probably related to episodic changes in fluid pressure, differential stress, or both. The source of fluids responsible for fault dilation and mineral deposition is not known. Gilbey (1968) proposed that they were derived from juvenile fluids mixed with others rising from a hypothetical subjacent batholith, which he considered was partly responsible for creating the Harlech Dome. Shepherd & Allen (in Allen & Jackson 1985) suggested they were produced by dewatering of lowermost Cambrian and Precambrian rocks. The temperatures of the mineralizing fluids

333

reached 535°C according to Gilbey's (1968) estimates based on pyrite-pyrrhotite assemblages in some veins. He estimated a maximum pressure of 2.4kbar on the basis that by end-Silurian times, when veining was supposed to have taken place, the host rocks had been buried to a depth of about 9 kin. These temperature and pressure estimates are probably too high. Shepherd & Allen (in Allen & Jackson 1985) inferred from fluid inclusion data that temperatures were no more than 435°C and were probably less than 350°C (130-310°C according to Shepherd & Allen 1979). The pressure estimate is based largely on the assumption that this area was buried beneath the complete Ordovician-Silurian lithostratigraphic column, whose thickness was calculated from data outside the Harlech Dome. There is no way of evaluating this assumption, nor is there any proof that Silurian rocks were deposited in the area. The Harlech Dome veins are described here as post-tectonic, in accordance with the general opinion that they were emplaced after endSilurian regional deformation. Whilst this view may prove to be valid and is supported by regional considerations discussed in a later section, it is nevertheless important to record that the age of the veins and their relationships to deformation have not been closely defined. Ineson & Mitchell (1975) reported K/Ar isotopic ages in the range 397_ 5 to 368 ___5 Ma yielded by phyllosilicate (mainly illite) concentrates obtained from wall rock material sampled from mines in the main mineralized part of the Harlech Dome. Allen & Jackson (1985) provide additional K/Ar ages, obtained from similar material, in the range 410+13 to 390___12 Ma. If these isotopic ages were to date the wall rock alteration, then vein emplacement took place in late Silurian to midDevonian times. However, these K/Ar ages require cautious interpretation because of the possibilities, noted by Ineson & Mitchell (1975), of isotopic contamination by older material or argon-loss, for example. In effect, it is not clear how to interpret these ages. Among a variety of alternative interpretations which merit consideration, is that some or all of the veins and wall rock alterations are older than these ages: perhaps veins were emplaced in early Lower Palaeozoic times, were held at levels where temperatures remained sufficiently high and fluid flushing continually homogenized the isotopic systems until uplift or other events in late SilurianDevonian times led to conditions suitable for complete argon-retention. Another unresolved problem is the timing of vein emplacement with respect to deformation. That the veins cut across the trendsof axial traces

334

B. Fitches

and cleavages does not, on its own, demonstrate post-tectonic veining, nor are the strains observed in vein quartz (Ashton 1981; Shepherd & Allen 1979) necessarily due exclusively to late movement on the host faults. In Snowdonia, where tectonic strains are substantially higher than those in the Harlech Dome, only the thin and suitably oriented pre-tectonic veins are obviously folded, whilst veins the size of those in the Harlech Dome appear to be undeformed. Moreover, in the Snowdonia veins, demonstrably pretectonic for various other reasons, the microscopic signatures of strain, which can be unambiguously attributed to regional deformation rather than minor local movement, are not commonplace. Another problem concerns the age of the last stage of deformation, believed to have preceded veining and assumed to be end-Silurian. This assumption is based largely on the parallelism between the latest folds and cleavage and those in Silurian rocks south and east of the Harlech Dome. In the light of evidence for repeated deformation in the Harlech Dome, and elsewhere in the Welsh basin, structural congruence does not necessarily imply contemporaneous deformation.

Plynlimon dome Veins in mid-Wales were emplaced in Ashgillian to Upper Llandoverian strata (Jones 1922; BGS Central Wales Mining Field 1:100,000 Map, 1973; BGS Aberystwyth 1:50,000 Sheet 163, 1984) which were deposited in various submarine fan and basin plain facies (Cave & Hains, in press). These strata are deformed by several major open periclinal folds, of which the Plynlimon Dome (Fig. 7) is the largest, with steep N N E striking axial planes. Parasitic folds occur on all scales. Cleavage, approximately axial planar to folds, is widely developed in argillites, but is poor or absent in coarser rocks. Kink bands and crenulation cleavages are locally imposed on the other structures (Fitches 1972). The veins are grouped mainly in the SW and N W flanks of the Plynlimon Dome and also cross the Van Dome to the east. The veins are developed along ENE striking normal faults which dip north or south, reach a maximum length of 20 km and maximum displacement of 200 m. Toward their lateral terminations, the inclination of the faults increases, the throw decreases and branching takes place (Jones 1922; Phillips 1972). Quartz is the predominant vein mineral, with various amounts of pyrite, ankerite, chalcopyrite, sphalerite and galena; barytes and witherite are found in the veins of the north west flank of the

Plynlimon Dome and cobaltite, arsenopyrite and marcasite appear locally in some veins (Raybould 1974). The faults and veins are clearly post-tectonic with respect to folds and cleavage; vein breccias contain disoriented cleaved wall rock fragments whilst galena veins locally cut late folds (Fitches 1972, Fig. 2). Moreover, some veins deviate from their usual ENE trend into ESE 'a-c' joints (Raybould 1976), probably indicating vein emplacement during or after strain relaxation following the regional deformation. The intimate spatial relationship between faulting and vein mineralization has been elegantly explained by Phillips (1972). According to his model, the combination of pressure from fluids permeating the tip zones of faults and tensile stress related to regional crustal extension led to progressive upward normal fault propagation. The abrupt pressure drop immediately following tip propagation caused hydraulic brecciation and deposition of carbonate in the fault plane, closely followed by the precipitation of quartz and sulphides. Brecciated mineralized normal faults with 60 ° dips were produced at deeper levels in the sedimentary pile, but at successively higher levels, as confining pressure declined, the faults steepened to become vertical tensile fractures without dip slip displacement. The veins are estimated to have been emplaced at depths of 3-8 kin, where lithostatic pressures of 0.75-2 kbar would have been reached in the host rocks (Phillips 1972). These estimates are based largely on the assumption that during vein emplacement the full lithostratigraphic thickness of Silurian strata measured in this and adjacent parts of the basin had accumulated in the Plynlimon Dome area and had not been removed by erosion. With a geothermal gradient of 30°C k m - 1, host rock temperatures would have been in the range 90°-240°C which, according to Ashton (1981), is consistent with the low-medium temperature Pb-Zn vein mineralization. The age of vein emplacement is not known accurately because, as in the case of the Harlech Dome veins, the age of regional deformation is difficult to define and the significance of K/At isotopic ages is unclear. The faults are younger than late Llandoverian because they cut rocks of that age. The folds and cleavage cut by the veins are generally assumed to be end-Silurian to early Devonian structures, because late Silurian strata elsewhere in central Wales are deformed by structures with similar styles and trends to those of the Plynlimon Dome. Isotopic data seem to indicate their emplacement during late Lower Palaeozoic to early Upper Palaeozoic times. Ineson & Mitchell (1975)

Aspects o f veining in the Welsh Basin

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FIG. 7. Simplified geological map of the Plynlimon Dome showing the position of the major veins (based on fig. 3 of Phillips 1972; B. G.S. Central Wales Mining Field 1 • 100,000 Map, 1973 ; B. G.S. Aberystwyth 1 : 50,000 Sheet 163, 1984). obtained K/Ar ages in the range 375 + 6 to 334+ 6 Ma from 18 clay mineral samples taken from wall rocks. They considered the older and younger ages in the range may be due to contamination of samples by older material and argon-loss respectively and do not necessarily imply a prolonged period of veining. The mean age of 356 +_7 Ma obtained from the 10 ages agreeing within experimental error is in their view perhaps close to the true age of emplacement. The area invaded by metalliferous veins covers at least 1000 km 2 of central Wales. Feasibly, barren veins belonging to the same suite may be recognized over even wider areas, but have yet to attract attention. The fluids responsible for this extensive vein system are usually described as hydrothermal pore fluids and the metal content is attributed to leaching from the sedimentary pile. There are no igneous rocks exposed at the surface and no buried plutons are predicted from surface geology or geophysical data, so that magmatic fluid contributions are unlikely. However, the presence of early Ordovician or older volcanic assemblages deeper in the pile, which

might provide a source of metals for scavenging fluids, cannot be totally ruled out. The most likely sources of fluid are sea-water, trapped in and buried with the sediments, and water released by clay mineral alteration during burial (cf. Bruce 1984). Migneint

area

The Migneint area, lying about 25 km N N E of Dolgellau (Fig. 1), contains large numbers of quartz veins and mineral lodes which have received little attention in the literature. Lynas (1970, 1973) has described the Cambrian and Ordovician sedimentary, volcanic and minor intrusive rocks of the area, the deformation and has briefly discussed the veins. According to Lynas (1973), most veins trend east-west or less commonly N W - S E and comprise chiefly quartz with sphalerite, galena, copper sulphides, pyrite, haematite and calcite. He describes the veining as late in the deformation sequence and suggests a Hercynian age. The quartz veins around Manod Mawr, in the

336

B. Fitches

extreme N W part of the Migneint area, have been studied in some detail by the author. The host rocks are flat-lying rhyolitic welded and unwelded tufts, volcanogenic sandstones and slates of Lynas' (1973) mid-Ordovician Rhiw Bach Formation. Most rocks, especially the argillites, are strongly deformed by a gently dipping cleavage, lying closely to, but rarely parallel with, bedding, which contains a strong, nearly N-S grain elongation (Fig. 8). The cleavage is axial planar to exceptionally rare small folds of bedding which are tight to isoclinal and verge southward. These folds are best exposed in slate caverns. These structures and fabrics probably mark the NE extension of the northward-dipping shear zone which ramps over the Tan y Grisiau Granite in the Tremadoc-Blaenau Ffestiniog district to the SE (Campbell et al. 1985). This shear zone is perhaps the ductile bead in which the Tremadoc Thrust zone is situated and is considered by

Campbell et al. (1985) to be part of the main (endSilurian) Caledonian structure of N Wales. The veins are steep and most have an 075 ° strike, are rarely more than 1 m wide and 10 m long and are sporadically developed, mainly in the tufts and sandstones. Most are composed entirely of quartz, but some also contain carbonate, chlorite and, exceptionally, small amounts of pyrite and galena. The veins are post-tectonic, showing no signs of folding in the strong flattening fabric, or any effect on cleavage strength and attitude. Moreover, they are clearly developed along one of three sets of steep joints (Fig. 8) which are later than the penetrative deformation. The steep joints have a close geometrical relationship with the regional strain ellipsoid inferred from the cleavage and linear fabrics. The joints striking nearly north-south lie in the X Z plane of the ellipsoid whilst the 075 ° and 135 ° sets have a conjugate relationship to each other and intersect on the Z-axis of the ellipsoid. These N

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Aspects of veining in the Welsh Basin

337

West Shropshire

relationships appear to be best explained by strain relaxation after the end-Silurian deformation, the N-S joints forming as tensile fractures along the original extensional strain axis whilst penecontemporaneously the other sets developed as conjugate shear fractures. The emplacement of quartz veins only in the 075 ° joints and certain internal features of the veins are not readily explained in terms of strain relaxation. In most veins the quartz is massive, without discernible megascopic growth textures. Some veins, however, contain fibrous quartz. The fibres are either curved, and plunge east or west, or straight and elongated approximately normal to vein walls. The disposition of curved fibres suggests dextral oblique slip during fracture dilation and quartz growth, whilst the straight fibres indicate simple N N W - S S E dilation. Neither fibre pattern is consistent with the northsouth compression inferred from the joints, but imply extension in this direction. One possible explanation of these anomalies is that the veins developed later than the joints in response to a younger regional stress regime. Extension at this later time was approximately N N W thus favouring the dilation and or dextral movement of the 075 ° joints, but not the other sets. This point is taken up in later discussion.

Dines (1958) has given the most comprehensive account of the veins in West Shropshire, and only aspects relevant to this paper are summarized here. The veins were emplaced chiefly in Ordovician strata, but in places penetrated laterally into late Precambrian and Silurian strata (Fig. 9). They carry quartz, carbonates, sulphides of lead, zinc and copper, barium minerals and, locally, a little fluorite and bituminous material. The veins are steep and were emplaced along normal faults trending mainly ENE and NW. The major Habberly Fault, a N N E pre-Carboniferous fault, cuts some veins and their host faults whose ages therefore appear to be constrained to the period between the late Silurian and late Devonian. However, Woodcock (1984b) has considered that the Habberly (Pontesford-Linley) Fault and other faults in this district acted at different times as both dextral strike-slip and dip-slip faults throughout the period from at least mid-Ordovician to Triassic times. It is not known at what stage or stages in this long history of faulting the veins were emplaced, nor have the stress conditions prevailing during veining been established. Ineson & Mitchell (1975) obtained from this

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district six K/Ar isotopic ages in the range 375 _+ 6 to 326_ 6 Ma from wall rock clay fractions. It is not clear how these ages are to be interpreted.

Discussion Veins are widespread in the Welsh Basin and were emplaced at various times in response to a variety of processes (Fig. 10). The veins are categorized according to their temporal relationships to regional deformation, different relationships denoting different modes of emplacement. It is noted, however, that relationships to deformation and the timing of deformation are not invariably clearly defined. Regionally distributed pre-tectonic veins are considered to be products of deposition in fractures induced by overpressured fluids trapped beneath impermeable lithological units. The veins were produced at burial depths probably in excess of a few hundred metres, are diachronous and, in effect, propagated up through the sedimentary pile during progressive burial of the sediments. The fluids responsible for fracturing and mineral deposition were probably mainly trapped sea-water, whose composition had been extensively modified by interaction with buried sediment, and by addition of water produced by smectite dehydration. A very different type of pre-tectonic veining is confined to the Ordovician resurgent caldera of Snowdon. Here, veining was controlled mainly by magmatic processes: heat from the sub-caldera pluton circulated sea-water through the calderafill; the pluton contributed a magmatic component to vein depositing fluids; radial distension of the caldera during resurgent doming, coupled with fluid pressure, produced hydraulic fracturing above the pluton. In comparison to fluid activity associated with the regional veins, fluid circulation in the Snowdon caldera was probably much more vigorous and leaching of hot volcanic rocks, and, perhaps, also the subjacent pluton, was possible. These fluids therefore gave rise to veins enriched in sulphides and other minerals. Syn-tectonic veins are apparently rare, are mostly small-scale and formed in local dilational zones of folds and fractures. Fluids which deposited the vein minerals are probably very locally derived and for the most part produced barren veins. The Ogofau veins are perhaps particularly well-developed syn-tectonic veins deposited in fold crests and shear zones and owe their metal content partly to very local derivation from syngenetic sulphide layers and partly to influx of fluids which had leached deeper rocks.

An alternative interpretation, that some of the Ogofau veins are products of pre-tectonic hydraulic-jacking and other fluid overpressuring processes, is worthy of evaluation by detailed studies of vein-tectonic fabric relationships. If syn-tectonic veins are relatively uncommon, the reasons might include: the closing of many microfractures, especially steep ones, under compression thus reducing porosity and permeability; slow strain rates and a tendency for ductile styles of deformation rather than widespread brittle deformation and the production ofdilatant structures. Fluids have undoubtedly been active on a small scale during deformation, promoting pressure solution and growth of pressure fringes on minerals during cleavage development. Post-tectonic veins are the commonest type in the Welsh Basin and have received most attention in the literature because of their economic interest. Apart from following the development of folds and cleavage, these veins have many features in common: (i) they may use fractures initiated during strain relaxation after regional deformation; (ii) they are commonly aligned ENE; (iii) many are metalliferous; (iv) K/Ar isotopic ages obtained from wall rocks range from late Silurian to early Carboniferous. It is suggested that the post-tectonic veins are expressions of Variscan extensional processes known to have affected other parts of Wales and nearby regions, but unrecognized over most of the basin. In South Wales, for example, there is strong evidence that Caledonian compressional tectonics had given way by latest Silurian-early Devonian times to an extensional regime marking the onset of Variscan sedimentary basin development. The forerunners of the ESE-trending Retic, Benton and other faults were active then as normal faults and producing graben which strongly influenced Lower Devonian sedimentation and continued as growth faults well into Carboniferous times (Hancock 1973; Allen et al. 1978; Allen et al. 1981; Dunne 1983). Similarly, in S and Central Ireland the Devonian and Carboniferous sedimentation patterns were strongly influenced by syn-depositional faults (e.g. Gardiner & Sheridan 1981). In England too, notably in the Pennine districts, faulting was active during deposition of Carboniferous sediments (Anderton et al. 1979). Evidently, large areas of the crust surrounding the Welsh Basin were under extension or transtension in early Upper Palaeozoic times. In the absence of an Upper Palaeozoic cover to all but the periphery of the Welsh Basin, there is no direct evidence that the interior parts of the basin were similarly affected. It is most improbable, however, that the Welsh Basin would have escaped these Upper

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Palaeozoic events. Accordingly, any interpretation of post-tectonic veining in the basin must consider the likely role of early Variscan extensional processes. Until end-Silurian times, the Lower Palaeozoic rocks in Wales were under a dominantly compressive tectonic regime and fluids in the pile, although under pressure, were largely immobilized. Subsequently, this regime gave way to one of regional extension, caused initially perhaps by uplift related to isotatic recovery from Caledonian deformation, but increasingly, or even exclusively, to Variscan crustal extension. In this new regime, tensile stress, commonly augmented by fluid pressure, provided the means for widespread dewatering of the Lower Palaeozoic pile. In places, exemplified by West Shropshire and probably the Harlech Dome, deviatoric stresses were sufficiently high to initiate normal faulting without the aid of fluid pressure. Fluids, now migrating more freely upward and into lower pressure sites, accumulated in the earlier faults, repeatedly stimulating further displacement and causing several stages of vein deposition. In other areas, notably the Plynlimon Dome, tensile stress and fluid pressure acted together to produce veins synchronously with fault propagation. To some extent these processes were able to exploit the legacy of Caledonian deformation. Joints, geometrically tied to the earlier strain axes, were an almost inevitable consequence of relaxation during uplift, Variscan extension or both. In the Migneint area, joints probably preceded the veins which used them whilst in the Plynlimon Dome the 'a-c' joints used by veins were perhaps already in existence or feasibly developed and opened during veining in response to early Variscan tension and fluid pressure. In many parts of the Welsh Basin, however, there is no obvious geometrical relationship between the veins and the Caledonian structures. Whereas the orientations of the Caledonian folds and cleavage vary widely over the basin many, probably the majority, of the post-tectonic veins maintain a persistent E N E trend. To a large extent, therefore, it seems that, to emplace many of the veins, the regional stresses and fluid

pressures were required to generate new fracture systems and disregard earlier structures. The common E N E alignment of veins in the Harlech Dome, Plynlimon Dome, Migneint area, West Shropshire and elsewhere, and paucity of veined faults and joints in other directions suggest that the Welsh Basin was subjected to a stress regime which had a strong component of N N W SSE extension. It remains for future investigations to determine whether this regime was one of simple N N W - S S E extension or involved transtensional elements as advocated by Woodcock (1984b). When the post-tectonic veins are viewed in the context of early Variscan extensional processes, the isotopic data obtained from vein wall rocks from several parts of the basin, ranging from 410+ 13 to 326+6 Ma, begin to take on a new significance. It remains unreasonable to assume that the isotopic age of a wall rock is an accurate guide to the age of a vein. However, there is no particular reason to regard the upper and lower ages as anomolous, nor to seek meaningful average ages. This range of ages is to be expected if the veins of the Welsh Basin are an expression of prolonged dewatering during early Variscan extension. Finally, it might be noted that the post-tectonic veins are not necessarily exclusively early Variscan. At the periphery of the Welsh Basin, and more especially in central Ireland, veins are located in Carboniferous rocks. The Irish veins are considered to be related to strike-slip tectonic processes associated with the main, late Carboniferous Variscan deformation (Critchley et al. 1983). Consequently, although the great majority of veins in the Welsh Basin are probably related to early Variscan dewatering of the Lower Palaeozoic pile, the possibility remains that any main Variscan activity, not yet recognized but likely to have had some effect on the basin, would have mobilized residual fluids capable of depositing vein minerals. ACKNOWLEDGEMENTS" Mike Andrews, my colleague, is gratefully acknowledged for his advice on several aspects of this paper.

References ALLEN, J. R., ELLIOT,T. & WILLIAMS,P. B. J. 1978. The sequence of the earlier Lower Old Red Sandstone (Siluro-Devonian), north of Milford Haven, southwest Dyfed (Wales). Geol. J. 13, 11336. & -1981. Old Red Sandstone and Carboniferous fluvial sediments in South Wales in

Field Guides to modern and ancient fluvial systems in Britain and Spain (Ed. Elliot, T.). Univ. Keele.

ALLEN, P. & JACKSON,A. 1985. Geology of the country around Harlech. Mem. Br. Geol. Surv., Sheet 135. ANDERTON, R., BRIDGES, P. H., LEEDER, M. R. & SELLWOOD,B. W. 1979. A Dynamic Stratigraphy o f the British Isles. Allen & Unwin, London.

Aspects of veining in the Welsh Basin ANNELS, A. E. 1981. Gold mineralization at Ogofau, Dyfed, Central Wales. Abstracts for Mineral Deposits Studies Group Metting, Cardiff, 1-2. ASHTON, J. H. 1981. Wallrock geochemistry and ore geology of certain mineralised veins in Wales. Unpubl. Ph.D. thesis. Univ. Wales (Aberystwyth). BEVINS, R. E. & ROWBOTHAM, G. 1983. Low-grade metamorphism within the Welsh sector of the paratectonic Caledonides. Geol. J., 18, 141-67. BRUCE, C. H. 1984. Smectite dehydration--its relation to structural development and hydrocarbon accumulation in northern Gulf of Mexico Basin. Amer. Assoc. Petrol. Geol. Bull. 68, 673-83. CAMPBELL, S. D. G. 1984. Aspects of the dynamic stratigraphy (Caradoc-Ashgill) in the northern part of the Welsh marginal basin. Proc. Geol. Ass., 95, 390-91. , REEDMAN, A. J. & HOWELLS, M. F. 1985. Regional variations in cleavage and fold development in N. Wales. Geol. J. 20, 43-52. CAVE, R. & HAINS, B. A. in press. Geology of the country between Aberystwyth and Machynlleth (Sheet 163). Mem. geol. Surv. G.B. CRITCHLEY, M. F., PHILLIPS, W. E. A. & COLLER, D. W. 1983. Correlation of geological, geochemical and geophysical data with satellite imagery in central Ireland. Abstract for Mineral Deposits Studies Group Meeting, Manchester. B5. DINES, H. G. 1958. The West Shropshire mining region. Geol. surv. Gt. Brit. Bull. 14, 1-43. DUNNE, W. M. 1983. Tectonic evolution of S.W. Wales during the Upper Palaeozoic. J. geol. Soc. London, 140, 257-65. FITCHES,W. R. 1972. Polyphase deformation structures in the Welsh Caledonides near Aberystwyth. Geol. Mag. 109, 149-55. & CAMPBELL,S. D. G. 1987. Tectonic evolution of the Bala lineament in the Welsh Basin. In ."FITCHES, W. R. & WOODCOCK, N. H. (eds) Sedimentation and Tectonics in the Welsh Basin. Spec. Issue Geol. J. (in press). --, CAVE, R., CRAIG, J. & MALTMAN, A. J. 1986. Early veins as evidence of detachment in the Lower Palaeozoic rocks of the Welsh Basin. J. struct. Geol. 8, 607-20. FYFE, W. S., PRICE, N. J. & THOMPSON, A. B. 1978. Fluids in the Earth's Crust. Elsevier, Amsterdam. GARDINER, P. R. R. & SHERIDAN, D. J. R. 1981. Tectonic framework of the Celtic Sea and adjacent areas with special reference to the location of the Variscan front. J. struct. Geol. 3, 317-31. GILBEY, J. W. G. 1968. The mineralogy, paragenesisand structure o f the ores o f the Dolgellau Gold Belt, Merionethshire, and associated wall rock alteration. Unpubl. Ph.D. thesis. Univ. London. GUTH, P. L., HODGES, K. V. & WILLEMIN, J. H. 1982. Limitations on the role of pore pressure in gravity gliding. Geol. Soc. Amer. Bull. 93, 606 612. HANCOCK, P. L. 1973. Structural zones in Variscan Pembrokeshire. Proc. Ussher Soc., 2, 509-20. HENNEBERGER, R. C. 1984. Ohakuri fossil epithermal system, New Zealand in Epithermal Gold (compiled by HEDENQUIST, J. & REID, F.). Earth Res. Foundation, Univ. Sydney. 168-79. -

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INESON, P. R. & MITCHELL, J. G. 1975. K-Ar isotopic age determinations from some Welsh mineral localities. Inst. Min. Metall. Trans. (Sect. B: Appl. earth sci), 84, 7-16. JONES, O. T. 1922. Lead and zinc. The mining district of North Cardiganshire and West Montgomeryshire. Mere. geol. Surv. spec. Rep. Miner. Resour. Gt. Br. 20. KOKELAAR,B. P. 1979. Tremadoc to Llanvirn volcanism on the southwest side of the Harlech Dome (Rhobell Fawr) N. Wales. In. HARRIS, A. L., HOLLAND, C. H. & LEAKE, B. E. (eds) Caledonides of the British Isles: reviewed. Spec. Publ. geol. Soc. London, 16, 591-6. , HOWELLS, M. F., BEVINS, R. E., ROACH, R. A. & DUNKLEY, P. N. 1984. The Ordovician marginal basin of Wales. In : KOKELAAR,B. P. & HOWELLS, M. F. (eds) Marginal Basin Geology ." Volcanic and Associated Sedimentary and Tectonic Process in Modern and Ancient Marginal Basins. Spec. Publ. geol. Soc. London, 16, 245-69. LIPMAN, P. W., FISHER,F. S., MEHNERT, H. H., NAESER, C. W., LUEDKE, R. G. & STEVEN, T. A. 1976. Multiple ages of mid-Tertiary mineralisation and alteration in the western San Juan mountains, Colorado. Econ Geol. 71,571-88. LISTER, C. R. B. 1983. The basic physics of water penetration into hot rocks. In: RONA, P. A., BOSTROM, K., LOUBER, L. & SMITH, K. L. (eds) Hydrothermal processes at seafloor spreading centres. NATO Sci. Affairs Div., Plenum Press, New York, 141-68. LYNAS, B. D. T. 1970. Clarification of the polyphase deformations of North Wales Palaeozoic rocks. Geol. Mag. 107, 505-10. --, 1973. The Cambrian and Ordovician rocks of the Migneint area, North Wales. J. geol. Soc. London, 129, 481-503. MARSHALL, J. D. 1982. Isotopic composition of displacive fibrous calcite veins: reversals in pore-water composition trends during burial diagenesis. J. Sed. Petrol. 52, 615-30. PHILLIPS, W. E. A., STILLMAN, C. J. & MURPHY, T. 1976. A Caledonian plate tectonic model. J. geol. Soc. London, 132, 579-610. PHILLIPS, W. J. 1972. Hydraulic fracturing and mineralisation. J. geol. Soc. London, 128, 337-59. PRICE, N. J. 1977. Aspects of gravity tectonics and the development of listric faults. J. geol. Soc. London, 133, 311-27. RAYBOULD, J. G. 1973. Studies o f the variations in paragenetic sequence and zoning in mineral veins of Cardiganshire and Montgomeryshire. Unpubl. Ph.D. thesis. Univ. Wales (Aberystwyth). --, 1974. Ore textures, paragenesis and zoning in the lead-zinc veins of mid-Wales. Inst. Min. Metall. Trans. (Sect. B. Appl. earth sci.) 83, 112-19. REEDMAN, A. J., LEVERIDGE, B. E. & EVANS, R. B. 1984. The Arfon Group ('Arvonian') of North Wales. Proc. Geol. Ass. 95, 313-21. - - , COLMAN,T. B., CAMPBELL,S. D. G. & HOWELLS, M. F. 1985. Volcanogenic mineralisation related to the Snowdon Volcanic Group (Ordovician),

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Gwynedd, North Wales. J. geol. Soc. London. 142, 875-88. ROBERTS, B. 1981. Low grade and very low grade regional metabasic Ordovician rocks of Llyn and Snowdonia, Gwynedd, North Wales. Geol. MaR. 1 1 8 , 189-200. - & MERRIMAN,R. J. 1985. The distinction between Caledonian burial and regional metamorphism in metapelites from North Wales: an analysis of isocryst patterns. J. geol. Soc. London, 142, 61524. SHEPHERD, T. J. & ALLEN, P. M. 1979. Fluid inclusions as a guide to mineral exploration in the Harlech Dome. In: ALLEN, P. M., COOPER, D. C. & SMITH, I. F. (compilers) Mineral exploration in the Harlech Dome. Inst. Geol. Sci. Rept. 29, 34-41. SHRESTHA, P. L. 1975. A comparative study of the biogeochemical and soil geochemical exploration techniques in the vicinity of the Ogofau gold mine. Unpubl. M.Sc. thesis. Univ. Wales (Cardiff). SLEEP, N. H. 1983. Hydrothermal convection at ridge axes. In . RONA, P. A., BOSTROM,K., LAUBER,L. & SMITH, K. L. (eds) Hydrothermal processes at seafloor spreading centres. NATO Sci. Affairs. Div. Plenum Press, New York, 71-82.

SMITH, R. L. & BAILEY,R. A. 1968. Resurgent Calderas. In: COATS, R. R., HAY, R. L. & ANDERSON, C. A. (eds) Studies in Volcanology. Geol. Soc. Amer. Mem. 116 (Howel Williams Vol.). 613-62. STEED, G. M., ANNELS, A. E., SHRESTHA, P. L. & TATER, P. S. 1976. Geochemical and biogeochemical prospecting in the area of the Ogofau gold mines, Dyfed, Wales. Inst. Min. Metall. Trans. (Sect. B: Appl. earth sci). 85, 109-117. STEVEN, T. A. & RATTE, J. C. 1965. Geology and structural control of ore deposition in the Creede District, San Juan Mountains, Colorado. U.S.G.S. Prof Pap. 487, 90 pp. STONELEY,R. 1983. Fibrous calcite veins, overpressures and primary oil migration. Am. Ass. Petrol. Geol. Bull. 67, 1427-8. TATER, P. S. 1975. The Ogofau gold bearing lodes." their structural control and primary geochemistry. Unpubl. M.Sc. thesis. Univ. Wales. (Cardiff). WOODCOCK, N. n . 1984a. Early Palaeozoic sedimentation and tectonics in Wales. Proc. Geol. Ass., 95, 232-335. --, 1984b. The Pontesford Lineament, Welsh Borderland. J. geol. Soc. London, 141, 1001-14.

W. R. FITCHES, Department of Geology, Llandinam Building, The University College of Wales, Aberystwyth, Dyfed SY23 3DB.

Index Pages on which figures appear are printed in italic, and those with tables in bold

ADN (allokinetic deformation number) 142-3, 144 age of vein emplacements 333, 334-5,340 Ainsa basin 181-2, 186, 186, 187, 188, 191, 192, 194 allokinetic deformation 138, 139, 141-3, 144 amplitude of stylolite peaks and relative displacement 35-6 angularity changes across faults 93, 100 arlticrack propagation 28 appearance of shear zones 78-81,82, 83 aragonite in chalk 57 arenites see meta-arenites argillaceous sediments see clays Arran faults 91-104 Arrats Mill, outwash sands 305, 307-10, 312-3 asymmetrical slump folds 186, 187-9, 187, 193 Autapie Nappe 147-61 autokinetic deformation 138

ball and pillow structures as allokinetic indicators 141 earthquakes as triggers 299, 300, 302, 303-5 graben tectonics and 270-2 ball clay, experimental deformation and 71, 79, 82, 83, 84, 85 banana shaped basins 174, 174, 176 bank collapse structures 317-8 basal slump detachments 192 5 basins extensional and palaeotectonics 137-44 formation and slickensides 107-20 microfaulting and compactiqn 123-35 slump strain in 181-95 veins 325-40 see also banana shaped basins bedding angular relationship of spaced cleavage zones to 199-200, 208-9 regional veins and 326 wet-sediment folds and 227-9 benthos, activity in chalk 55, 57 bimodal patterns, of poles to faults 103 breccias 149, 152-5,261 brick clay, experimental deformation and 72, 79, 83 brittle behaviour 12, 20 fractures in pressure solution-deposition 30, 31 progressive lithification and 155, 160 slickensides 110-1,112, 114-8 water content in clays 83-4, 86 browncoalsin Kleszcz6w Graben 242 buckle folds 187-9, 272-4, 277 burial depth, indeterminate, and wet-sediment deformation 221-5,225, 226, 227, 228, 231-4, 237 burial related deformation 319-23

calcite-filled veins, Autapie Nappe 149, 155, 157, 158, 160 caldera development, Snowdonia veins and 327-9, 338 Cam-Clay theory 1, 2 capillary pressure, open dykelets and 39, 41, 44-5 carbonates, unmetamorphosed see dolostones Carboniferous, Lower, Northumberland Basin 142-4 cataclasis of grains, deformation bands and 151 causes of deformation see processes of deformation cavities, growth of crystals within 29-31 cementation 1 Central Mobile Belt see Dunnage Zone chalk experimental deformation of 4, 5, 65-8 mechanical behaviour of 6, 7, 55-61 ChojnySeries 287, 291,291 classification of deformational structures in the Kleszcz6w Graben 256-63 of soft-sediment deformation 17-21 clays consolidation of 58 experimental deformation of 71-2, 77-87 hydroplastic microfaultsin 126-35 syn-diagenetic deformation in the Autapie Nappe 158 cleavage relationship with bedding in wet-sediment folds 227-9 see also spaced cleavage zones closed systems and spacing of tectonic layers 34-5 closure of dykes in hydrocarbon intrusion 47-8, 53 CobleCreep 2, 3 coccolithsin chalk 55 coefficients of transfer, creep relation and 25-6 cohesion, effect of faulting on 98 cohesionless materials 13, 14, 20 cohesive materials 12, 20 sandas 15 shearing experiments on soils 131 collapse pressures in chalk 57-8, 60 compaction and microfaulting in the Soria Basin 134-5 definitions 1-2 differential see differential compaction compositional limits of differentiated layers 33-4 compression definitions 1-2 of chalk 56, 57-8, 59-61 compressive strength of chalk, and Young's Modulus 65, 67-9 computer-aided technique for determining stress tensors 133-4

343

344

Index

concretions, structures formed around in burial related deformation 320 confined compression in chalk 56, 57-8, 59-61 confining pressure effective and grain cataclasis 151 strain rate, shear zones and 85 consolidation definition 1, 2 in chalk 56-7 of samples for experimental deformation 74 convection, thermal in Snowdon caldera 329, 337-8 convergent margins, syn-diagenetic deformation at 161 convolute lamination 1I, 17 convoluted flame structures 280, 286 convolution horizons, small-scale in the Kleszc6w Graben 261-3 creep by pressure solution-deposition 25-37 indeterminate burial depth in wet-sediment deformation 237 creep relation 25 critical length of fractures 48 criticalstateline 2, 6 critical stress intensity 47-8 cross-bedding, deformed 11-12, 17, 319 crystals in cavities, growth of 29-31 currents causing flame structures 279, 280-6 see also palaeocurrents curvature of fold hinges 187, 189, 192 cycles, glacial, in the Kleszcz6w Graben 246, 287, 293 cyclic liquefaction 12 Czy~6w Series 270-7, 291,293,295, 296

decollements, layer-parallel, Autapie Nappe 158, 159, 160 definitions of terms 1-6 deformation bands 91 inAutapie Nappe sandstone 150-1 see also faults deformation mechanisms see mechanisms deformation processes see processes deformation structures see structures densification and creep 25 of chalk 56-7, 59-60 see also dry density density gradients, unstable 17, 22, 293, 295, 317 deposition extension during 139-40 Silurian dolostones 201 zones of 29-31, 32, 34, 36 dewatering pipes as allokinetic indicators 141 see also water escape diachronousliquefaction initiation 305, 306 diagenesis and deformation mechanisms 159-60 chalk 55-7 diapirs flame structures like 280, 282, 285-6

mudlump caused by gravity spreading 166-7, 176, 177 differential compaction burial related deformation 319-23 sills and dykelets 49, 50 differentiation, tectonic and pressure solutiondeposition 31-7 diffusion, open dykelets and 43 diffusive-mass-transfer 2, 3 dihedral angles, conjugate sets in Arran faults 99-100 dimensional changes see volume changes dip and density of chalk 59, 60 hydrocarbon intrusion and 51 dip-slip faulting, hydroplastic slickensides and 107, 108-9, 111-2, 113, 117 direct shear 74, 78-80 see also shear testing discharge fluctuations and bank failures 318-9 dish structures 307-8 displacement of microfaults see movement dissolution and stylolites 35-6 zones of 27-9, 32, 33, 34, 35, 36 distributed shear 73, 74, 75, 78 see also shear testing dolostones, Silurian 197-9, 201,205-9 downward migration of hydrocarbons 45, 53 drag folds, endogenic processes and 272-3, 277 drained compression in chalk 56, 59 driving force systems 16-17, 18, 20 and pressure solution-deposition 25, 28 in the Kleszcz6w Graben 293-5 dry density of chalk 55-6, 59, 60, 62, 65, 66 ductile behaviour Autapie Nappe sandstone 150-1,152-3,155,160 deformation mechanisms for 11-14, 20 water content in clays and 83, 84, 86 Dunnage Zone 214, 217, 218 dykes hydrocarbon 39-53 flame structures like muddy 280, 282, 286 quartz in the Soria Basin 135 sand see sand earthquake motions allokinetic deformation and 137, 138-9 and liquefaction of sand 14-15 bank failures triggered by 315-6, 318-9 graben activity and 295-7 in Scottish Quaternary 299-305, 310, 312, 313-4 see also endogenic activity effective confining pressure and grain cataclasis 151 effective stress 2-3 Elsterian Upper, glaciodeltaic sequence 291,293 see also Czy~6w Series endogenic processes 263,269-77 see also earthquake motions; subsidence estuary clay, recent 79, 82, 86 euhedral crystals, growth in open cavities 29, 30 examination of specimens, experimental deformation 75

345

Index excess pore fluid pressure see overpressure exogenic processes 261 experimental deformation of water-rich sediments 71-87 extension basins under 139-40, 142-4 gravity-driven emplacement of a submarine nappe and 147, 149, 160 gravity gliding and 174 hydrocarbon intrusion and 39, 45-6, 52 post-tectonic veins and 338-40 fault-grading stratigraphy, earthquakes and 299, 312, 314

fault planes brittle and hydroplastic 109-14, 115, 116 hydroplastic microfaults and shearing experiments on cohesive soils 123, 126-31 faulting during extension 139-40, 143 endogenic processes and 261,274, 276-7 in high-porosity sandstones 91-104 laterally spreading sand bodies 17 veins and 333,334, 337 see also fault planes; microfaulting feedback mechanism in tectonic differentiation 33, 36 fibrous crystals, crack-seal process and 29-31 fibrous quartz in Welsh Basin veins 337 fissures 261,274, 277 see also tension cracks flame structures 270-1,279-86 flotation of organics 308, 312 flowage structures 261,272-3,275, 277 fluid escape see water escape fluid flow, Meikleour deposits 302 fluid phase, crystals in cavities and 29, 31 fluidization 12, 13, 14, 300 flame structures and 281-2, 285 unconsolidated sands 15, 16, 17, 20, 22 fluids causing veins 326-7, 329, 333,335,338 fluvio-deltaic strata, microfaulting in 123-35 fluviolacustrine deposits 255-6, 261,265, 269-77 folds classification of, in the Kleszcz6w Graben 256-61 gravity tectonics and 165-77 laterally spreading sand bodies 17 slumps 181,186-95 wet-sediment 225-6, 227-37 fracture condition, hydrocarbon dykelets 40-6 fracture toughness 48 fractures brittle in pressure solution-deposition 30, 31 hydrocarbon migration and hydraulic 39-53 pre-tectonic veins and 326, 328-9 friction angle and faulting, high-porosity sandstone 96-8

glaciolacustrinedeposits 255-65, 279-86 Glen Roy 310-4 gold-bearing veins, Welsh Basin 330-1 Goldenville Formation 199 gouge zones 91 see also faults graben tectonics as trigger for deformation 295-7 grain mobilization and deformation bands 150-1, 156-7 grain size changes in New Red Sandstone 93, 95, 99-100

granulation seams 91 see also faults gravitational body force as driving force system 17, 22 gravitationally unstable density gradient see density gradient gravity gliding 155-74, 177 gravity-induced deformation 140-1 inthe Kleszcz6w Graben 287-97 similarity to hydroplastic faults 118,120 wet-sediment sliding 213-8, 220-1,229,234-5, 237 see also gravity gliding; gravity spreading gravity spreading 166-7, 174-7 Greenland, North, Fold Belt see Peary Land Grindosonic apparatus 63-4, 65, 68-9 growth rate of intrusion fractures 46-7 halokinesis, glacial cycles and 242, 243,246, 248 hardness of chalk 59-61, 63 HarlechDome 332-4, 340 helictic flame structures 282-5,285 Helminthoid Flyschturbiditic series 147-61 heterogeneity, initial and pressure solution-deposition 29, 31, 34-5, 37 High Atlas, hydroplastic slickensides in 107-20 high-porosity sandstones, faulting mechanisms in 91-104 histories of deformation Autapie Nappe 155-7 Kleszcz6w Graben 245-8 Point Leamington formation 229-34 Soria Basin 134-5 see also process of deformation Hog's Back, Surrey, chalk samples 59-61 Holocene deposits in the Kleszcz6w Graben 248-50 hydraulic fracturing, hydrocarbon migration by 39-53 hydraulic jacking, veinsand 326 hydrocarbons migration by hydraulic fracturing 39-53 reservoirs of 3, 6-7 hydroplasticmicrofaulting 123, 126-35 hydroplastic slickensides see slickensides ice weight hypothesis see gravity spreading imbricate structures, gravity spreading and

167, 174,

176

glaciation in the Kleszcz6w Graben 243,246 see also glaciodeltaic deposits; glaciolacustrine deposits glaciodeltaicdeposits 255-65,279-86

impulsive liquefaction 13 index of deformation (ADN) 142, 144 inherent failure processes and tectonic activity 315 injection structures 153-5, 156, 157, 158, 316

346

Index

interconnection of sills and dykelets 49, 50-3 internal extension fractures 39-40 intrafolial sediment deformation and neotectonics 138-9 intrusion extension fractures 39-53 irregular stringers (veins) 331 isolated flames 285-6 jacking, hydraulic 326 Jurassic, Middle see Ravenscar Group

kaolinite, experimental deformation of 71, 79, 82, 83, 85

karst phenomena in the Kleszcz6w Graben 236 Khalassa Basin 182-6, 189, 190, 191, 192, 194 KinlochHourn fault 312, 313 Kleszcz6w Graben 241-97 Kuroko-type process, modified for vein formation 327

laminations and shear zones 78, 79-80, 81-3, 85 landslides 167-9, 168, 169, 170, 171 layering sample preparation for experimental deformation 73 tectonic differentiation and 31-5, 36-7 limits of deformation, Scottish Quaternary 307 liquefaction 13,300 and deformation structures in the Kleszcz6w Graben 257-61,296-7 and fluid escape in syn-sedimentary deformation 316 earthquake-initiated 303,305, 306, 313 in sand 14-15, 20 intrafolial sediment deformation 138-9 liquidization 11-12, 14-17,300 see also fluidization; liquefaction lithification 3-4 and hydroplastic faults 118 mechanical properties during 6 of chalk on the seafloor 56-61 syn-diagenetic deformation and 154-5, 157-8, 160 lithology of shear zones 81-3 lithons and cleavage zones in meta-arenites 200-1, 204, 205 load, unevenly distributed confining 16-17, 21, 22 load casting in the Kleszcz6w Graben 261,270-2, 275, 296-7 flamed structures 280, 285,285 load structures density instability and 316 shock-induced 299 unconsolidated sands and 11-12, 17 loading in the Meikleour deposits 300 Lockport Formation see dolostones magmatic processes and veining 329, 337-8 marlstone layers in the Autapie Nappe 158-9 MaroniBasin 182-3, 190, 194 mass flow in the Meikleour deposits 302

mass transfer as mechanism for dimensional change 197-209 driving force for in pressure solution-deposition 25-6, 28, 32-3, 34-5, 37 see also oxides mechanicalbehaviour of chalk 57-61 mechanisms of deformation gravity tectonics 165-77 in chalk 57-61 inthe Autapie Nappe 147-61 in unconsolidated sands 11-13, 14-16, 18, 20 shear zones in water-rich clays 85-7 see also processes of deformation Meguma Group, meta-arenites 199-201,205, 207-9 Meikleour deposits 300-5, 313 melanges 149, 161 meta-arenites 199-201,205,207-9 metagreywackes see meta-arenites metalliferous veins, Snowdonia 327-9 metamorphism meta-arenites and mass transfer during 197-209 Soria Basin development 135 microfaulting Autapie Nappe 149, 150-3 hydroplastic in the Soria Basin 123, 126-35 microstructures experimental deformation and water-rich sediments 71-87 inthe Autapie Nappe 149-59 on hydroplastic slickensides 111, 114-8 microstylolites 197-8 Migneint area 335-7, 340 migration of hydrocarbons by hydraulic fracturing 39-53 mineralization of veins and faulting 334 Miocene sediments in the Kleszcz6w Graben 245 modal analysis of Goldenville meta-arenites 200 movement, criteria for sense of 109-14, 127-30 mudlumpdiapiric structures 176, 177 mudslides 167 mudstone injections in wet-sediment deformation 227 layersin the Autapie Nappe 157-8 multimodal patterns of poles to faults 92, 103 Nabarro Herring Creep 2, 3 nappes, formation by gravity gliding 167, 171,174 see also Autapie Nappe Needham Market, chalk samples from 57-9 neotectonics 137, 138-40 New Red Sandstone, Isle of Arran 91-104 Northumberland Basin 142-4 Notre Dame Bay, wet-sediment deformation in 218-37 Ogofau veins 330-1 open dykelets 43-5, 46, 52 open systems, spacing of tectonic layers 34 Ordovician, Upper see Point Leamington Formation organics, flotation of 308,312 orientation of shear zones 78, 81-3, 85-6 overburden changing and burial related deformation 319-20, 323

Index overburden (cont.) deformation of chalk 57-61 hydrocarbon intrusions and 41-5, 49, 52 overconsolidation 4 overgrowths, tectonic 29-31 overpressure definition 3 hydrocarbon intrusions and 41, 43, 44-6 transient and stress overshoot 101 oxides, gain and loss in mass transfer during metamorphism of arenites 201,205 palaeocurrents data from in Point Leamington Formation 228-9, 235, 237 flame structures and 280, 285 palaeoslopes 181,194, 195 palaeotectonics 140-4 Palaeozoic, Lower in the Welsh Basin 325-40 patterns of faults and bulk strain 102-4 peak strengths of argillaceous sediments and water content 83-5 Peary Land, gravity gliding in 171-4 penecontemporaneous deformation in the Kleszcz6w Graben 269-77, 287-97 penetration depths of downward growing dykes 45 penetrative sediment deformation and neotectonics 139-40 permeability and hydrocarbon intrusion 45, 52 pervasive homogenous flow 77, 87 pillow structures 302-4, 307-8 see also ball and pillow structures plant material and burial related deformation 321-2 plastic behaviour 12, 15, 20 Pleistocene, late, sediments in the Kleszcz6w Graben 248 Pliocene sediments in the Kleszcz6w Graben 246 Plynlimon Dome 334-5,340 podiform veins 331 point-counts in New Red Sandstone 93, 99-100 PointLeamington Formation 218-37 Poisson's ratio and hydrocarbon intrusions 42-3 poles to faults, patterns in New Red Sandstone 103 polyphase deformations in the Kleszcz6w Graben 263 'ponding' of hydrocarbons 41 pore fluid pressure 4 and deformation in the Kleszcz6w Graben 257-61, 293,295 and hydrocarbon intrusion 40-6, 47-8, 52 and microstructures in the Autapie Nappe 159, 160 transient increases in faults 101,104 pore space 4 pore throat 4 porosity 4 and behaviour of chalk 6-7, 55-6, 65-7, 68 and grain cataclasis 151 sandstone with high 91-104 variations in dolostones 201,206, 207 Portland Limestone 7, 65-8 post-tectonic veins 332-7, 338-40 pot clay, experimental deformation of 71-2, 83 pre-lithification deformation 5-6

347

see also soft-sediment deformation preparation of samples for experimental deformation 72-4, 78 for Grindosonic testing 64 pressure capillary see capillary pressure effect of confining and strain rate on shear zones 85 pore fluid see pore fluid pressure see also overpressure pressure solution diffusion 2, 3 pressure solution-deposition creep 25-37 pre-tectonic veins 326-9, 330, 338 processes of deformation faulting in high-porosity sandstones 96-104 hydraulic fracturing and hydrocarbon migration 39-53 hydroplastic faults 118-20 in the Autapie Nappe 158-61 in the Kleszcz6w Graben 242-3,248-50, 255-6, 274-7, 285-6, 293-7 in the Northumberland Basin 142-4 in the Point Leamington Formation 234-7 in the Scottish Quaternary 305,306, 310, 312 in the Yorkshire Basin 315-23 in unconsolidated sands 11-21 in water-rich clays 77-87 mass transfer and volume loss 207-9 microfaulting in the Soria Basin 134-5 pressure solution-deposition creep 25-37 slump folds 187-95 veins in the Welsh Basin 338-40 prod marks, microfaults and 127-9 Purbeck, Isle of 61, 64-5

quartz dykes in the Soria Basin 135 in metamorphism of arenites 200-1,202-3, 205, 207, 208 veins in the Welsh Basin 330-1 Quaternary sediments earthquake motions in 299-314 flame structures in 279-86 gravity tectonics and 165-77 penecontemporaneous deformation horizons in 287-97

rate of deformation and solution-deposition 25, 34-5 ratio of strengths 12-3 Ravenscar Group 315-23 recumbent-folded cross stratification as allokinetic indicator 141 recumbent folds in slumps 186-7, 192, 193, 195 tangentialshear causing 17 refolded folds in slumps 186-7, 192, 194-5 regionalpre-tectonic veins 326-7, 337 relative fault movements 109-14, 127-30 residual strength of argillaceous sediments 85 rigidity see hardness rock mechanics 4 rolling cylinder experiments and flame structures 280

348 rotation of folds 192, 194-5 of particles, and shear zones 86 Saalian Chojny Series 291,292 saddle reef veins 331 salt diapirs 242, 248 sample preparation see preparation sand deformation processes in unconsolidated 11-21 -dyke injections 288, 152-5,156 flame structures in 281 intrafolial sediment deformation in 138 volcanoes 316 see also sandstone sandstone bodies and burial related deformation 321-3 faulting mechanisms in high-porosity 91-104 hydroplasticslickensidesin 107-18 syn-diagenetic deformation of 149-57, 158-61 water escape structures in 125-6 saturated moisture content 59, 62 SCZ see spaced cleavage zones seafloor lithification of chalk 56-8, 61 secondary compression of chalk 56, 59 secondary microfaults on slickensides 111-2, 114, 115, 116, 128-9 sediment, restriction of term 6 sedimentary sediment deformation (as opposed to tectonic sediment deformation) 214 sedimentation from the liquefied state 15 sedimentation fronts 13 seismic processes see earthquake motions seismites 138 self-healing of fractures 30, 31 self-induction of chemical differentiation 33 sense of fault movements see movements sensitivity 12 settlement of sediments 72-3, 139 shear deformation 4, 14 and folds in slumping 187-9, 194, 195 shear testing clays 73, 74, 75, 78-80, 81, 86 cohesive soils 131 shear zones inthe Kleszcz6w Graben 261 in water-rich clays 77-87 on slickensides 111, 114-8, 118, 120 sheath folds 187, 189, 195 shock waves causing flame structures 279, 282-5 see also earthquake motions shortening bulk and volume loss 205, 207, 209 of slumps 187-9, 193, 195 Shropshire, West, post-tectonic veins in 337, 339 SiO2, removal from spaced cleavage zones 201,205, 207, 208 sills 39, 45, 48-53 siltstones, hydroplastic microfaults in 126-35 simulation of sediments in the laboratory 71-4 slickensides hydroplastic 107-20 hydroplastic microfaults and 127, 130, 131-4

Index microfaulting in the Autapie Nappe 150, 151-3 sliding, superficial sediment 213-4, 220-1,229, 234-5, 237 see also gravity-induced deformation slip in the New Red Sandstone 91-2, 98-9 of particles in shear zones 86, 87 slope influence in deformation 308 instability and bank erosion 317-9 slump folds and models of soft-sediment deformation 181-95 and wet-sediment deformation 234-5 slumping at Glen Roy 310 Snowdoniapre-tectonic veins 327-9, 337-8 soft-sediment deformation definition 5-6 wet-sediment deformation and 213-4 soil mechanics 4 solubility of a solid in its solution 25, 26, 33-4, 37 Soria Basin 123-35 spaced cleavage zones (SCZ) 197, 199-200, 205,207, 208-9 spaced solution zones (SSZ) 197-9, 200, 205-7,208 spacing of hydrocarbon dykes 46, 47, 48 of tectonic layers 34-5 specialized materials for experimental deformation 76 spreading, lateral, of bedding plane sills 49 'squelch' structures 317 SSZ see spaced solution zones staticliquefaction 13 strain and microstructures in the Autapie Nappe 158-9 and shear zones 84-5, 86-7 see also stress; stress-strain curves strain ellipsoids in meta-arenites 200, 201 strain-hardening faults 92, 96, 98-101,102-4 stratal disruption in the Autapie Nappe 149, 158-60 strength and water content in clays 83-4 measurement of compressive 65, 67-8 strengths, ratio of 12-3 stress deviatoric and pressure solution-deposition 25, 26, 28-9, 31-2 effective 2-3 history of Needham Market chalk 58-9 needed for deformation 16 see also strain; stress-strain curves stress paths for chalk 4-5 stress ratio 5 stress-strain curves for chalk 6, 58 for clays 79, 83-5 for strain-hardening faults 98-102 stress tensor, microfaulting 131-4 strike-slip faulting 107, 109, 112, 113, 115 structures, deformation and classification of soft-sediment deformation 18, 20 and palaeotectonics 140-4 characteristic of wet-sediment deformation 219-20

349

Index structures, deformation (cont.) earthquake-induced in the Scottish Quaternary 300-5, 307-12 inthe Autapie Nappe 149-58 in the Kleszczdw Graben 243-50, 256-63,270-7, 280-6, 287-97 in the Yorkshire Basin 315-23 see also microstructures stylolites 35-6, 37 see also microstylolites submarine gravity nappe emplacement see Autapie Nappe subsidence, tectonic in the Kleszcz6w Graben 242, 243-5,246, 248,255 see also endogenic processes superficial fluid velocity 14, 14 superficial sediment sliding see sliding symmetrical slump folds 186, 187, 194 synchronous liquefaction initiation 305,306 synclines in the Czy~6w Series 274, 274, 277 syndiagenetic deformation 159, 160-1 syn-sedimentary deformation 315-9 syn-tectonicveins 329-31,338

tangential shear stress, effects of 17, 18 tectonic activity and hardness of chalk 59, 60-1 in the Yorkshire Basin 315-6 see also earthquake motions; endogenic processes; subsidence tectonic sediment deformation (as opposed to sedimentary sediment deformation) 214 temperature and pressure solution-deposition creep 25, 29 tension cracks in brittle slickensides 111, 114 in the Kleszcz6w Graben 261,263,266, 274, 277 inthe Soria Basin 135 tension gash arrays of veins 331 termination of hydrocarbon intrusion 47-8 terminology, need for standardization 1 Tertiary sediments see Khalassa Basin; Maroni Basin; Miocene; Pliocene thickness-weighted deformation length see ADN thin sections Arran faults 92, 97 brittle and hydroplastic slickensides 114-8,117, 118, 119, 120

preparation after experimental deformation 75, 78 wet-sediment deformation 226, 227 thixotropy 12 thrust fault regions and gravity gliding 165, 169, 171-4 thrusting, and folding in wet-sediment deformation 236 time, evolution with, of deposition and dissolution 29-31 Tine-Mal Zone 108, 109, 111, 114 Tizi n'Test fault zone 107, 108, 109, 110 total stress 4 translation of slumps 193-4, 195 transport distance in dolostones and meta-arenites 207, 208,209

Triassic outcrops, slickensides in 107-20 triaxialcompression tests 74-5, 76, 78 and shear zones in clays 80, 81, 82, 84, 85-6 triggers 14, 15, 22 graben tectonics as 295-7 see also mechanisms of deformation; processes of deformation truncation surfaces in the Meikleour deposits 303-4 tulip structures 304 turbiditic series, syn-diagenetic deformation of 147-61 turbidity currents see currents unconsolidated sands 11-21 unevenly distributed confining load uniaxial stress path 4-5 uniaxial stress ratio 5

16-17, 21, 22

Variscan extension in the Welsh Basin 338-40 veins and pressure solution-deposition 31 in the Autapie Nappe 149, 155, 157, 158, 160 in the Welsh Basin 325-40 vertical shear stress, effects of 18, 19 vibration, Grindosonic apparatus and 63 viscous behaviour 12, 20 Vistulian Piaski Series 287, 288-9, 291 void ratio 4 volcanoes, sand see sand volume and stress 2 changes and internal deformation 33-4 losses in metamorphism of arenites 200-1,205, 207,209 of cavities and deposition 29 V~lvedal thrust fault region 171-3

water enhanced role in Meguma rocks 207, 209 flame structures caused by see currents water content of clays, deformation and 78, 79, 83-4 water escape structures as allokinetic indicators 141 in sandstones 123, 125-6, 135 in unconsolidated sands 18 syn-sedimentary deformation and 315 water-rich sediments 71-87 wavelength of tectonic stylolites 36 wax model for gravity spreading 166 wedge-shaped tension fissures 274-277 Welsh Basin 325-40 wet-sediment deformation and soft-sediment deformation 213-4 in the Point Leamington Formation 219-37 Wight, Isle of, chalk samples from 59, 61 wrinkle marks (flame structures) 280

yield strength deformation mechanisms and

11-12

Index

35o yield strength ( c o n t . ) of chalk 7,58 yield surface 5 Yorkshire Basin 315-23 younger-on-older stratigraphic units

Young's modulus for chalk 7, 63, 65-8

165-6,173

zones of deposition s e e deposition zones of dissolution s e e dissolution