Deformation of sediments and sedimentary rocks

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

o-P1--(T 3

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

Normal \ consolidation line Impossible~'

r

=

states

,, /

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/

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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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~

~f'fu~°nn

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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-

00

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i

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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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Hydrostatic 0

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1'0 2'0 30 40

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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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FIG. 9. Effective stress ratio plots for various uniaxial compaction experiment on Chalks. Note the changes in slope due to elastic yielding and the onset of normal compaction behaviour.

6

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

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Porosity @aW== Grains transported with fluid (elutrlstion).

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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

~

/ bi°turbation

/rain /prints //e.g.

/

//// // /

/ /

/ t

~

~

pierced

/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

:

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 ~~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

pPh~ (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.

(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 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.'~

5"20 ,

I

J iii!i ii!!i E

Fro. 1. Map of Arran and surrounding areas (after McLean & Deegan 1978). Strain-hardened faults occur in New Red Sandstone sandstone facies. Inset is cartoon at same scale during intrusion of the Northern Granite.

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).

5"

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.

98

J . R . Underhill & N. H. Woodcock *-500mm

t4

~.1

,~ ~50mm -~-~'~ :o °'~,.',~e. -.~¢:°_. inner ~ ~ ~:~;:~;-~

I

~

outer

zone undeformed

( ~ ~

parent sst.

C h a r a c t e r i s t i c s are

parent

sst.

~ ' ~

r I

1.

m

.6mm

±

:-

: Rounded grains

with limonite coatings and g r a i n - c o n t a c t cement, loose packing, high porosity.

outer zone

: 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

99

(%) 7O

I I I

60

50

40

30

> 0.1mm

20

10

o

0.125

I 0.50

/ O. 5

I 1.0

(mm)

(A)

(%)

I

I

I

I

I

,

O.03-0.1mm

I

2O 10 I I

0

0.25

0.50

(B)

O. 5

1 0

(ram)

(%) 70

I

50 40 30

I

o

2 O. 5

I 0.50

L 1 0.75

1.0

(mm)

(c)

(% 30

1

I

I I I

I I I

pore space

I ;I"

I I I 0.25

I

I I 0.150

(D)

I I 0.75

I0 1.

(mm)

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.

Ioo

J. R. Underhill & N. H. Woodcock (%) 70 !

I I I

0.55mm T

0.15mm T

1.05mm -1grain s i z e > O . l m m

60

\,,

/

',

20

grain size O.03-O. l m m 10

grain

I 0.1

0.2

0.3

0.4

0.6

0.7

OUTER ZONE

=

I N N E R ZONE

0.5

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1.0

1.1

size';.--~.;:_y,---,

/

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IC) FIG. 16. Drawings from thin sections (cut perpendicular to the fault planes and parallel to the movement) in hydroplastic faults; dots = shear zones; broken lines = grain alignment; A and B from sample A, Fig. 9; C from sample F. Scale = 1 cm, letters explained in text.

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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later Jurassic-early Cretaceous and was filled by fluvial dominated deltaic series. To the north, the basin fill is thrust over the Oligo-Miocene Molasses of the Ebro Basin, but the Alpine deformation remains light and the original internal geometry of the basin is preserved. To the south, the Tertiary conglomerates of the Iberic Meseta unconformably overlap the Wealdian deposits and the Palaeozoic rocks of the Moncayo Massif. Inside the basin, the total stratigraphic thickness is up to 8 km (Beutler 1966; Salomon 1983; Guiraud 1983; Guiraud & S~guret 1985). The Upper Jurassic-Lower Cretaceous infill can be subdivided into four major cyclothems (see Fig. 1, cyclothems I to IV), each one about 1500-3500 m thick. Each cyclothem is composed of a basal formation of fluvial meandering channel sandstone bodies sandwiched in silt and clay deposits of alluvial plain origin. This basal formation grades-up into an upper formation composed of lacustrine to restricted shallow marine carbonates (Fig. 2). The well defined basal discontinuity of each cyclothem is related to a rapid progradation (relative sea-level fall) of fluvio-deltaic sediments on to marine deposits. Outside the basin, the thickness of the Wealdian

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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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M. Guiraud & M. Sbguret

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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

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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
-:Ts,~:~ ~ - - ~

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flow

bedding

POINT LEAMINGTON FORMATION

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

F ,

~

.

z."

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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

Wet-sediment deformation in the Leamington Formation

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FIG. 13. Scale drawing of part of wet-sediment deformation horizon on east coast of Southwest Arm (opposite Passage Rocks). Note curvilinear fold axes (with local plunge shown), and bifurcation of synforms and antiforms. Also, note the abrupt transition from coherently folded to chaotic bedding towards lower part of diagram.

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

K. T. Pickering N

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FIG. 17. Lower hemispherical equal area projection of wet-sediment fold data. (A) & (B) from west of Southwest Arm; (C) & (D) from east of Southwest- Arm; (E) & (F) from West Arm, and (G) & (H) from Exploits Bay. Data are generally consistent with a southerly-dipping basin-slope, although data from east of Southwest Arm is ambiguous (see text for discussion). All data rotated to approximate pre-D2 orientation as outlined in text. Open circles show clockwise, and solid circles show anticlockwise, sense of rotation of axial plane looking down plunge.

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K. T. Pickering

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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Wet-sediment deformation in the Leamington Formation GROOVE MARKS POINT LEAMINGTON FORMATION

S O U T H W E S T ARM (WEST), WEST ARM,NORTHERN ARM & LITTLE NW. ARM N=39

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FIG. 18. Summary palaeocurrent data from Point Leamington, suggesting a generally southerly-dipping basin-slope. Compare with data in Fig. 17. White arrows show sector towards which currents flowed based on limited flute marks and current-ripples.

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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TECTONIC FOLD AXES NEW BAY

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-

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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

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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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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.

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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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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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