Deformation Micro Structures and Mechanisms in Minerals and Rocks

DEFORMATION MICROSTRUCTURES AND MECHANISMS IN MINERALS AND ROCKS

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Deformation Microstructures and Mechanisms in Minerals and Rocks by Tom Blenkinsop Department of Geology, University of Zimbabwe, Harane Zimbabwe

KLUWER ACADEMIC PUBLISHERS NEW YORK, BOSTON, DORDRECHT, LONDON, MOSCOW

eBook ISBN: Print ISBN:

0-306-47543-X 0-412-73480-X

©2002 Kluwer Academic Publishers New York, Boston, Dordrecht, London, Moscow Print ©2000 Kluwer Academic Publishers Dordrecht All rights reserved No part of this eBook may be reproduced or transmitted in any form or by any means, electronic, mechanical, recording, or otherwise, without written consent from the Publisher Created in the United States of America Visit Kluwer Online at: and Kluwer's eBookstore at:

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Contents Acknowledgements

ix

Symbols, Abbreviations and Units

xi

1 Introduction and Terminology 1.1 Introduction 1.2 Classifications of deformation microstructures and mechanisms 1.3 Deformation microstructures and mechanisms in the earth: Brittle-semibrittle-plastic transitions 1.4 The description of deformation: Scale, continuity, distribution, mechanism and mode 1.5 Ductility and the “brittle-ductile transition” 1.6 Character and classification of deformation zone rocks 1.7 Format and use of this book

1 1 1 3 4 4 5 5

2 Cataclasis 2.1 Introduction 2.2 Fundamental cataclastic deformation mechanisms 2.2.1 Microcracking 2.2.2 Frictional sliding 2.3 Microcracks 2.3.1 Classification, characteristics and observation 2.3.2 Microstructures and mechanisms 2.3.3 Impingement microcracks 2.3.4 Flaw-induced microcracks 2.3.5 Microfracturing of pre-existing flaws 2.3.6 Cleavage microcracks 2.3.7 Elastic mismatch microcracks 2.3.8 Plastic mismatch microcracks 2.3.9 Microfault-induced microcracks: Microscopic feather fractures (mffs) 2.3.10 Thermally-induced microcracks 2.3.11 Phase transformation-induced microcracks 2.4 Microfaults 2.4.1 Characteristics 2.4.2 Mechanisms 2.5 Deformation bands 2.5.1 Characteristics and classification 2.5.2 Mechanisms 2.6 Distributed cataclasis and cataclastic flow 2.7 Gouge zone microstructures 2.8 Microfracture surface features 2.9 Crystallographic fabrics 2.10 Pre-lithification deformation microstructures and mechanisms 2.11 Pseudotachylites and frictional melting 2.11.1 Characteristics 2.11.2 Origin 2.11.3 Misidentification v

7 7 7 7 10 10 10 12 12 13 13 13 13 14 14 15 16 17 17 17 18 18 18 18 19 19 22 22 22 22 22 23

vi

CONTENTS

3 Diffusive Mass Transfer by Solution 3.1 Introduction 3.2 Fundamental deformation mechanisms of diffusive mass transfer by solution 3.3 Grain surface solution textures 3.4 Indenting, truncating and interpenetrating grain contacts 3.5 Strain caps 3.6 Microstylolites 3.6.1 Characteristics 3.6.2 Formation and propagation 3.7 Diffusive mass transfer and cleavage 3.7.1 Classification 3.7.2 Spaced cleavages 3.7.3 Continuous cleavage 3.8 Grain surface deposition textures 3.9 Overgrowths, porosity reduction, pressure shadows and fringes, and mica beards 3.9.1 Characteristics 3.9.2 Mechanisms 3.10 Grain shape fabrics 3.11 Fluid inclusion planes 3.12 Microveins

24 24 24 25 25 27 27 27 27 28 28 29 30 30 32 32 33 33 33 35

4 Intracrystalline Plasticity 4.1 Introduction 4.2 Fundamental mechanisms of intracrystalline plasticity 4.3 Deformation twins 4.4 Undulatory extinction 4.5 Intracrystalline deformation bands, kink bands and subgrains: Recovery 4.6 Deformation lamellae 4.7 Grain shape fabrics and ribbon grains 4.8 New grains, core and mantle structure: Dynamic recrystallization 4.9 Crystallographic fabrics

39 39 39 39 41

5 Diffusive Mass Transfer and Phase Transformations in the Solid State 5.1 Introduction 5.2 Fundamental deformation mechanisms of solid state diffusive mass transfer and phase transformations 5.3 Grain shape fabrics and ribbon grains 5.4 Foam texture, static and secondary recrystallization 5.5 Decussate texture 5.6 Porphyroblasts and inclusion trails 5.6.1 Characteristics 5.6.2 Growth mechanisms 5.6.3 Relationship to deformation 5.7 Reaction rims, relict minerals, coronas and symplectites 5.8 Chemical zoning 5.9 Solid state phase transformation microstructures 5.10 Superplasticity

52 52 52 52 54 54 54 54 55 55 57 57 57 57

6 Magmatic and Sub-magmatic Deformation 6.1 Introduction 6.2 Fundamental deformation mechanisms and microstructures in rocks containing melt 6.2.1 Magmatic flow 6.2.2 Sub-magmatic flow 6.2.3 Magmatic and sub-magmatic flow and rheology 6.3 Mesoscopic evidence for magmatic and sub-magmatic flow 6.4 Magmatic microstructures 6.4.1 Grain shape fabrics 6.4.2 Crystallographic fabrics 6.5 Sub-magmatic microstructures 6.5.1 Grain shape fabrics

59 59 59 59 60 60 60 62 62 62 62 62

41 47 47 47 50

CONTENTS

vii

6.5.2 Intracrystalline plasticity 6.5.3 Diffusive mass transfer 6.5.4 Cataclasis 6.6 Other microstructures 6.7 Non-magmatic deformation

62 63 63 63 63

Microstructural Shear Sense Criteria 7.1 Introduction 7.2 Curved foliation 7.3 Oblique foliations and shape preferred orientations 7.4 Porphyroclast systems 7.4.1 Characteristics and classification 7.4.2 Mechanisms of formation and tails 7.4.3 Stair-step direction: 7.4.4 Faces of a tail tails 7.4.5 Deflection and embayments of 7.5 S-, C- and 7.5.1 Characteristics and classification 7.5.2 Formation and evolution 7.5.3 Curvature of S-foliation 7.5.4 Shear on C- or 7.6 Pressure shadows and fringes 7.6.1 Kinematics of pressure shadows and fringes in shear zones 7.6.2 Geometry of the last increment of growth 7.6.3 Shape 7.7 Mica fish 7.8 Porphyroblast internal foliations 7.9 Crystallographic fabrics 7.10 Asymmetric microboudins 7.11 Asymmetric microfolds and rolling structures 7.12 Shear sense criteria in rocks containing melt 7.12.1 Magmatic shear zones 7.12.2 Oblique grain shape fabrics 7.12.3 Tiling and imbrication 7.12.4 S-C fabrics 7.12.5 Sub-magmatic microfractures 7.13 Shear sense criteria for faults 7.13.1 Shear sense observations on faults 7.13.2 Displaced grain fragments 7.13.3 Risers and slickenfibres 7.13.4 Gouges 7.13.5 Jogs and bends

65 65 66 66 67 67 68 69 69 69 70 70 72 73 73 73 73 73 73 74 75 75 76 77 77 77 78 78 78 78 79 79 79 79 79 79

8 Shock-induced microstructures and shock metamorphism 8.1 Introduction 8.2 Shock mechanisms 8.3 Microfractures 8.4 Planar Deformation Features (PDFs) 8.5 Mosaicism 8.6 Diaplectic glass 8.7 High pressure polymorphs of quartz - Coesite and stishovite 8.8 Lechatelierite 8.9 Tectites, microtectites and spherules 8.10 Shock barometry and thermometry 8.11 Calibration of shock pressures from microstructures 8.11.1 Calibration of shock pressures from optical properties of quartz 8.11.2 Problems of shock barometry 8.12 Diagnostic impact microstructures

80 80 80 80 81 82 83 83 83 84 85 85 87 87 88

7

viii

CONTENTS

9 From Microstructures to Mountains: Deformation Microstructures, Mechanisms and Tetonics 9.1 Introduction 9.2 Failure criteria 9.2.1 Coulomb and Mohr failure criteria 9.2.2 Griffith failure criteria 9.3 Pore fluid pressure and faulting 9.4 Fracture mechanics and failure criteria 9.5 Frictional sliding laws 9.5.1 Byerlee’s law 9.5.2 Rate and state dependent frictional sliding 9.6 Flow laws 9.6.1 Diffusive mass transfer: Grain size sensitive creep 9.6.2 Intracrystalline plasticity 9.6.3 Empirical flow laws from experimental data 9.7 Polymineralic deformation 9.8 Deformation mechanism maps 9.9 Lithospheric strength envelopes 9.10 Palaeopiezometry 9.10.1 Methods and calibration 9.10.2 Recrystallized grain size 9.10.3 Subgrain size 9.10.4 Dislocation density 9.10.5 Twinning - differential stress 9.10.6 Deformation lamellae 9.10.7 Principal stress orientations from deformation lamellae 9.10.8 Principal stress orientations and strains from twins 9.10.9 General problems with palaeopiezometers 9.11 Geothermobarometry 9.11.1 Methods and calibration 9.11.2 Calcite twin morphology 9.11.3 Sutured quartz grain boundaries 9.11.4 Subgrain boundary orientation in quartz

90 90 90 90 90 91 91 91 91 92 92 92 93 94 94 97 97 98 98 98 100 100 100 101 103 103 104 104 104 104 105 105

References

107

Index

127

Color Plate Section

133

Acknowledgements Most of the photomicrographs were developed and printed by Cuthbert Banda, whose assistance, with that of other members of the staff of the Geology Department, University of Zimbabwe, was invaluable. James Preen guided the preparation of the TEX version of the text. Faith Samkange and Maxwell Matongo were able research assistants, supported by the University of Zimbabwe Research Board. The following are thanked for contributing photomicrographs or thin sections: P. Dirks (Plates 11, 30, 31, 32, 33, 35, 44), R. Fernandes (Fig. 2.9), H. Frimmel (Plate 18), S. Kamo (Figs. 8.4 - 8.6), H. Leroux (Figs. 8.2, 8.3), J.E.J. Martini (Figs. 8.7, 8.8), U. Reimold (Plate 46), J. Stowe (Plate 22), D. Van Der Wal and M. Drury (Fig. 4.2). Plates 1 - 4 and Figs. 2.5, 2.18, 2.19, 3.18 of core material from the Cajon Pass drillhole were made at the Institute for Crustal Studies, University of California, Santa Barbara, as part of research on deformation mechanisms with R. Sibson, supported by the National Science Foundation, U.S.A., under grant DAR-84-10924. The assistance of the technical staff at U.C.S.B. is gratefully acknowledged. Some research for this book was supported by the IUGS Commission on Tectonics, COMTEC. Detailed reviews of chapters from the following are greatly appreciated: P. Dirks, R. Hanson, H. Jelsma, W. Means, A. Ord, M. Paterson, U. Reimold, E.H. Rutter, A. Schmid Mumm. Fig. 8.2 was reproduced from Leroux et al. (1994), and Fig. 8.7, 8.8 from Martini (1991), all with kind permission of Elsevier Science - NL Sara Burgerhartstraat 25, 1055 KV Amsterdam, The Netherlands. Fig. 8.4 was reproduced from Krogh et al. (1996) with kind permission from the American Geophysical Union. Fig. 9.5 was reproduced from Burkhard (1993), and Fig. 7.20 from Goldstein (1988), both with kind permission from Elsevier Science Ltd. The Boulevard, Langford Lane, Kidlington OX5 1GB, U.K. Figs. 9.6, 9.7 were reproduced from Kruhl and Nega (1996), with kind permission from Springer Verlag. Full details of these publications are given in the references.

ix


Symbols, Abbreviations and Units Numbers in brackets give the chapters and sections where the symbols are used. Units are given if they are referred to in the text.

a A b B c

CL CMF d

D

DMT

E f

GBM

G ISA

J k

kC l L

LPO LSE m

n P PDF

Microcrack long axis (2.2.1), Rate and state variable friction law constant (9.3.2) Flow law constant, (9.4.1), or (9.4.3) S to C-surface angle (7.5), Flow law constant (9.4.3), Effective stress coefficient (9.2.3) Microcrack short axis (2.2.1), Flow law constant (9.4.3) Burger’s vector (4.2, 9.8) Particle velocity/unit potential gradient (3.2), Flow law constant (9.4.3) Microcrack or flaw length (2.2, 2.3, 2.4.2), concentration of particles (3.2) Cohesion, MPa (9.2, 9.3), Reference state solubility, mole fraction (9.4.1) Angles defining external asymmetry of an LPO (7.9) Cathodoluminescence (2, 3, 8) Critical melt fraction (6.2) Particle size (2.2.2), Flaw spacing (2.4.2), Grain size, m (9.4), mm (9.8) Subgrain size, mm (9.8.3) Fractal dimension (2, 9.9.3) Reference state diffusion coefficient for Pressure solution creep, (9.4.1) Reference state diffusion coefficient for Nabarro-Herring, Coble creep, (9.4.1) S to C- or angle (7.5), Grain boundary width, m (9.4) Diffusive mass transfer (1, 3, 5, 9) Twinning density, (9.8.5) Young’s modulus (2.2) Strain rate, strain rate at 0 K in the Dorn Law, (9.4) Volume fraction of phase in polymineralic flow law (9.5) Crystal fraction, fraction at critical packing (6.2.1) Microcrack extension force, Critical value (2.2) Grain boundary migration (4.8, 9.4, 9.8) Surface tension force (2.2), Angle between flaw and (2.3.4) Fracture toughness (2.2) Viscosity of suspension, Viscosity of pure melt (6.2.1) Instantaneous stretching axes (7.1, 7.7) Twinning incidence (9.8.5) Diffusive flux (3.2) Angle between flaw array and (2.4.2) Dislocation density palaeopiezometer constant (9.8.4) Stress intensity factors for mode I, II, and III opening (2.2) Threshold, Critical stress intensity factors (2.2) Velocity of deforming crystal face, (9.4.1) Model microcrack length (2.3, 2.4.2), Dislocation density palaeopiezometer constant (9.8.4) Critical slip distance (9.3.2), Length of grain boundary (9.9) Lattice preferred orientation (4.9, 5.10) Lithospheric strength envelope (9.7) Grain size exponent in flow law (9.4, 9.5), Grain size palaeopiezometer constant (9.8.2) Coefficient of friction (2, 9.2), Chemical potential (3.2), Shear modulus, GPa (9.8.3, 9.8.4) Coefficient of internal friction (9.2) Power law exponent for flow laws (9.4, 9.5) Mean stress, Pa (3.2, 9.4) Planar deformation feature (7) xi

xii

PPL

r R RF S

S

SEM ST

TEM u U V

XP () {} <>

SYMBOLS, ABBREVIATIONS AND UNITS

Pore fluid pressure (3.2, 9.2.3) Plane polarized light Flaw to microcrack angle (2.3.4, 2.4.2) Activation enthalpy, pressure solution and grain boundary diffusion, (9.4) Activation enthalpy for volume diffusion, and diffusion creep, (9.4) Length of stride in divider method (9.9.3) Gas constant, (9.4) Reflected light Density, (9.4.1), Dislocation density, (9.8.4) Deformation lamellae spacing, mm (9.8.6) Molar entropy (3.2) Cohesion, MPa (9.2) Oblique foliation, Bands parallel to shear plane (7.3) External, Internal foliations (5.6, 7.8) Scanning electron microscope Sensitive tint plate Differential stress, Pa or MPa (9.4) Maximum, intermediate, minimum principal stresses, Compression positive (2,3,5) Remote stress for microcrack closure (9.2.2), Remote applied stress (2.2.1) Normal stress (2,3, 9.2), Stress at 0 K in Dorn Law (9.4) Temperature, 0C and K, Melting temperature Uniaxial tensile strength, MPa (9.2.2) Transmission electron microscope Shear stress, MPa (2.2.2, 9.2, 9.3) Dislocation density exponent in dislocation density palaeopiezometer (9.8.4) Molar internal energy (3.2) Subgrain size exponent in subgrain size palaeopiezometer (9.8.3) Microcrack velocity (2.2.1), Sliding velocity (9.3.2), Activation volume, (9.4.3) Molar volume, (3.2, 9.4) Maximum twin volume, % (9.8.5) Thickness of a grain boundary fluid (9.4.1), Grain size palaeopiezometer exponent (9.8.2) Angles defining internal asymmetry of an LPO (7.9) Crossed polarized light State variable in dynamic frictional sliding law (9.3.2) Miller-Bravais indices of a crystal face Miller-Bravais indices of a representative face of a form Miller-Bravais indices of a crystal direction

Chapter 1

Introduction and Terminology 1.1 Introduction

changes that remains after stress is removed, as opposed to elastic (recoverable) deformation which is not seen directly Deformation microstructures in rocks and minerals are mi- in the geological record. Deformation microstructures can be croscopic features created by deformation. A deformation divided into three major categories: mechanism is a process on one scale that accommodates an 1. Microfractures, displacement, and/or rotation of rigid imposed deformation on a larger scale. This book describes particles with no permanent lattice distortion. The typmicrostructures and mechanisms at the scale of a thin section, ical microstructures seen in thin section are microfracthe scale that most geologists use for detailed petrography, tures and fragments surrounded by a fine-grained matbased on the premise that many of the fundamental mechanrix. isms can be inferred from microstructures at this scale. The book aims to be a guide to the recognition and interpretation 2. Microstructures showing material removal, transport of deformation microstructures and mechanisms, and should and deposition without fracturing, permanent lattice disbe used in conjunction with a petrographic microscope. tortion or melting. Examples of microstructures at sites Why is the study of deformation microstructures and of material removal are distinctive types of grain conmechanisms useful ? Deformation mechanisms are determtacts and microstylolites. Microstructures indicating ined by temperature, stress (both hydrostatic and deviatoric material deposition include microveins, overgrowths, components), strain rate, pore fluids, mineralogy, and the texpressure shadows and porphyroblasts. Many metature of the deforming rock (especially grain size and porosmorphic textures associated with deformation fall into ity). Recognition of deformation mechanisms from microthis category. structures allows limits to be placed on these variables. For example, microstructures called subgrains are formed by the 3. Permanent lattice distortion without fracturing. Exdeformation mechanism of intracrystalline plasticity, which amples of typical microstructures are undulatory exindicates deformation at temperatures above 250°C in quartz. tinction, subgrains, recrystallized grains, and crystalloThe size of the subgrains can be used to gauge the deviatgraphic fabrics. oric stress during deformation. These are essential pieces of information for tectonic analysis and for understanding the This simple classification scheme can be applied on the basis behaviour of the lithosphere by mathematical modelling (e.g. of optical microscope observations. Table 1.1 summarizes Kusznir and Park 1987, Molnar 1992, Beaumont et al. 1996). the scheme, gives examples of specific microstructures, and Microstructures and deformation mechanisms are a grow- refers to the relevant chapters and sections of the book. ing field of interest in the earth sciences. The application of Table 1.1 also shows the relation between this scheme and materials science to minerals and rocks has provided much of a classification of deformation mechanisms, which has three the new impetus. However, the literature is scattered through similar categories (e.g. Knipe 1989): journals in a large number of disciplines, and most structural 1. Cataclasis - deformation by microfracturing, sliding geology textbooks have limited coverage of the field. At least and/or rotation of rigid particles (Chapter 2). Brittle desome of this diverse literature is reviewed in this book, which formation is often used as a synonym for cataclasis, but contains a comprehensive reference list. Hopefully the book is more accurately defined as deformation by fracture or is written in sufficient depth to be useful at advanced undermicrofracture. graduate level and above. Familiarity with elementary concepts and terms in structural geology is assumed. 2. Diffusive mass transfer (DMT) - deformation by diffusion, the movement of lattice defects, ions, atoms or molecules in response to gradients of chemical potential 1.2 Classifications of deformation mi(Chapters 3 and 5).

crostructures and mechanisms

3. Intracrystalline plasticity - deformation by the movements of extra half-planes of atoms (dislocations) in a crystal lattice (Chapter 4).

Deformation microstructures in minerals or rocks are the record of permanent deformation, i.e. shape and/or volume 1

2

CHAPTER 1. INTRODUCTION AND TERMINOLOGY


A general term for both categories 2 and 3 is useful because they are often closely associated with each other: crystal plasticity or simply plasticity means deformation by either or both DMT and intracrystalline plasticity. DMT can be split into two major sub-divisions: diffusion via a solution (pressure solution, Chapter 3), and diffusion in the solid state (Chapter 5). Solid state phase transformations may occur during deformation, some involving DMT: these are included in Chapter 5. Two or more mechanisms may act simultaneously together within a single mineral. An example is the combination of sliding on grain boundaries, and mass transfer of material by diffusion to fill space created by sliding. This is is one type of superplasticity, which is a composite deformation mechanism; one in which two or more mechanisms are coupled together in the same mineral under the same conditions. An important composite deformation mechanism involves the interaction of microfractures and intracrystalline plasticity: this is called semibrittle deformation. The different components of a composite mechanism may have variable strain rates: the mechanism with the slowest rate determines the composite strain rate, and is said to be rate-limiting. It is fortunate that the non-genetic classification of the microstructures matches the genetic classification of the mechanisms so well, and this means that both classifications can be referred to by similar names. The chapter titles of this book use the names of the mechanisms for simplicity. The order of the chapters follows the general change in deformation mechanisms from low to high grade metamorphic conditions during deformation (see next section). Specific microstructures that are diagnostic for each mechanism are shown in bold in Table 1.1. Three chapters deal with relatively new developments in structural geology, which can involve all three categories of deformation microstructures and mechanisms. Magmatic and sub-magmatic deformation (Chapter 6) describes the deformation of rocks that contain melt. Important deformation mechanisms are flow of melt and crystals, with crystal deformation (sub-magmatic flow) or without crystal deformation (magmatic flow). This topic has great relevance to current debates about pluton emplacement mechanisms. Microstructural shear sense criteria (Chapter 7) provide clues to the displacement of rock masses during deformation on all scales: this is one of the most important types of tectonic information. Shock-induced microstructures and shock metamorphism are produced by meteorite impacts (Chapter 8), and are the focus of much current interest, because microstructural studies have a central role to play in the debates about mass extinction and other possible effects of large meteorite impacts on the earth’s evolution.

3

events, each of which may be associated with different mechanisms. One of the applications of this book should be to allow the unravelling of multiple deformations from their associated microstructures. Secondly, even within a single deformation, mechanisms and microstructures vary from mineral to mineral within a polymineralic rock: a common example is the intracrystalline plasticity of quartz in a shear zone at greenschist facies, that contrasts with cataclastic deformation of feldspar in the same conditions. Thirdly, one mechanism may be incapable of accommodating the imposed stress, strain or strain rate, so that other mechanisms are substituted or added during the same deformation: for example, faulting (cataclasis) may relieve high stress levels in a shear zone otherwise deforming by intracrystalline plasticity. The record of both the cataclasis and plasticity may be preserved in the microstructures. Mechanisms that can only accommodate deformation in a single direction, for example slip on a single fault set, are especially restrictive and invariably require supplementary mechanisms. The variation in conditions in the earth, particularly of temperature and pressure, causes a corresponding variation in deformation mechanisms. Two systematic changes occur with increasing depth: temperature increases by the geothermal gradient, and pressure increases due to the effect of gravity. Cataclasis is relatively insensitive to the variation in temperature, but is suppressed by pressure. The plastic deformation mechanisms (intracrystalline plasticity and solid-state DMT) behave in the opposite way: they are strongly promoted by temperature but relatively insensitive to pressure. As a result, cataclasis is generally restricted to the upper crust (where pressures are low), and crystal plasticity occurs in the lower crust and the rest of the lithosphere (where temperatures are high). The transition from cataclasis to plasticity is sometimes called the brittle-plastic transition. Experiements have identified an important intermediate regime of semibrittle behaviour where microfractures interact with intracrystalline plasticity (e.g. Carter and Kirby 1978, Kirby and Kronenberg 1984), leading to the concept of two transitions in deformation mechanism with increasing depth: firstly the brittle - semibrittle transition in the upper crust, and secondly the semibrittle - plastic transition in mechanisms at mid-lower crustal levels. These generalizations need to be heavily qualified because of the other variables that affect deformation mechanisms. Stress, strain rate and pore fluids are some of the most important additional variables to be considered. For example, high stresses or strain rates may cause cataclasis at greater depths or higher temperatures, which would otherwise be associated with plasticity, as in the above example. Pore fluids are essential for solution-assisted DMT, and may promote cataclasis by mechanical and chemical effects. Furthermore, 1.3 Deformation microstructures and the various minerals within a polymineralic rock have differtransitions from brittle to semibrittle to plastic behaviour, mechanisms in the earth: Brittle- ent and interactions between the minerals themselves also need to be considered as an independent variable (Chapter 9). These semibrittle-plastic transitions qualifications, particularly the variation in properties of difSeveral different deformation microstructures commonly oc- ferent minerals, mean that no unique depth or temperature cur together in rocks for three important reasons. Firstly, can be given for the brittle-semibrittle and semibrittle-plastic deformation microstructures may record several deformation transitions. However, the transitions are important concepts

4


for understanding the behaviour of the earth, and the conditions for the transitions can be predicted for specific models of the earth, as described in Chapter 9.

1.4

The description of deformation: Scale, continuity, distribution, mechanism and mode

Deformation of minerals or rocks should be described in terms of three fundamental attributes. Continuity is the connectivity of material points through the deforming body; deformation can be characterized as continuous if points remain connected, or discontinuous if not. The distribution of deformation can be described as localized (e.g. a shear zone) or distributed. The deformation mechanism can be described in one of the three categories given in Section 1.2. All three attributes depend on the scale of observation. Scales are loosely defined in this book as macro (greater than outcrop, i.e. > 10 m), meso (outcrop, i.e. 1 cm - 10 m) or micro (microscopic, i.e. < 1 cm). Permanent deformation of minerals or rocks involves breaking atomic bonds and is therefore is discontinuous at the atomic scale. As the scale of observation is increased, these atomic discontinuities can not be discerned, and deformation appears to be continuous. The top surface of Fig. 1.1 shows schematically how deformation continuity is a function of the scale of observation. An example of the discontinuous to continuous transition with increasing scale is the accommodation of a fold by cataclastic flow (Chapter 2.6). At micro to meso scales, the deformation is by discontinuous fracture, but at a

macroscopic scale, these discontinuities are not seen and the fold appears to be continuous. Continuity should be described at the time of deformation, but can easily be altered by subsequent events. Another problem with specifying deformation continuity is posed by features such as overgrowths, pressure shadows, and pressure fringes (Chapter 3.9), which may be continuous with the surface from which they are grow, and discontinuous with the surface towards which they grow. Boundaries between grains are perhaps best regarded as continuous on a microscopic scale during recrystallization, but after recrystallization they appear as discontinuities. These examples show that a certain amount of judgement may be necessary to describe continuity, which is a necessary shortcoming of any description that has to take into account the past history of deformation in minerals and rocks. Deformation distribution is also highly scale-dependent. As scale of observation is increased, an localized deformation may appear pervasive: for example, microfractures are highly localized deformation on a microscopic scale, but a network of microfractures can have the effect of a pervasive strain on the scale of an outcrop or a regional map. The combination of deformation distribution (localized/pervasive) and mechanism (cataclastic or plastic) was described as a “mode of failure” by Rutter (1986). This concept can be extended through the incorporation of deformation continuity and scale. The combination of continuity, mechanism and scale can be called a deformation mode, and represented on a diagram such as Fig. 1.1, where a mode is specified by deformation mechanism (z-axis), continuity (which can be qualified by distribution, x-axis) and the scale length (y-axis). The front of the diagram shows the field of possible deformation modes described in this book (i.e. at the microscopic scale), with examples of some microstructures. Figure 1.1 makes some important links between microstructures and mechanisms. Cataclastic deformation microstructures are discontinuous, and intracrystalline plasticity microstructures are continuous, at the microscopic scale. These observations point towards one of the major themes of this book: deformation microstructures can be used, albeit with care and a certain amount of ambiguity, to identify deformation mechanisms. This link is possible by induction and deduction from the microstructures and theoretical understanding of the mechanisms, and by comparison with experiments.

1.5 Ductility and the “brittle-ductile transition” Great confusion has been caused by the use of the terms ductile and brittle-ductile transition. Much of this confusion can be traced to the dichotomy between laboratory (and materials science)-based and field-based approaches to deformation (the subject is well reviewed in Evans et al. 1990 and Williams et al. 1994). For example, rock mechanics experiments use a maximum permanent strain before failure of more than 5% to define ductile behaviour (e.g. Paterson 1978). Similarly, Griggs and Handin (1960) defined

5


relative ductility as: “the amount of permanent deformation achievable prior to rupture”. These definitions of ductility based on stress-strain relationships are satisfactory and quantifiable, but the widespread application of the term outside the rock mechanics laboratory necessitates alternative definitions, since stress-strain relationships are never known for rocks in the field (e.g. Griggs and Handin 1960). The use of deformation continuity or distribution in the definition of ductile is implicit in the application of the terms by most geologists in the field. Rutter (1986) suggested that ductility is “the capacity for substantial, non-localized strain”, thus strictly excluding shear zones, and “is a concept which must be defined on a macroscopic scale”. However, “ductile shear zones” (e.g. Ramsay and Huber 1987) is a commonly used phrase. It is proposed here that continuity rather than distribution should be used to define ductility, and that a more satisfactory definition of ductility is “macroscopically continuous deformation”. This definition has the merit of encompassing all features that the field geologist usually refers to as ductile, including shear zones and macroscopic folds, and is preferred because continuity can be identified more precisely than distribution at any scale. Brittle and ductile are defined above by different criteria. Therefore it is possible for a rock to be both brittle and ductile: for example, a type of deformation band (Chapter 2.5) forms by fracture (it is brittle) but has macroscopic strain continuity (it is ductile). The concept of the term “brittle-ductile transition” is rendered questionable by these definitions. Many terminological problems can be avoided by using the concept of a deformation mode and the categories of deformation continuity, distribution and mechanism suggested in Section 1.4, and avoiding the use of the term brittle-ductile transition, which has no meaning using the above definitions.

1.6

Character and classification of deformation zone rocks

Rocks within deformation zones are commonly highly strained and may exhibit some of the best examples of deformation microstructures. Deformation reduces the grain size of some of the protolith to produce a matrix of finer grains surrounding remnant larger grains or grain aggregates, known as porphyroclasts or clasts, which is the typical microstructure of deformation zone rocks. Deformation mechanisms and microstructures may differ between the matrix and porphyroclasts. Some definitions and classifications of deformation zone rocks have advocated the use of deformation mechanisms as a classificatory tool. For example, Wise et al. (1984) proposed to include crystal plasticity as an essential element of the definition of mylonites. As for any other classification of observational data, the use of mechanisms should be avoided because subjective interpretations are required, that may change in the light of new knowledge. Nomenclature and classification of deformation zone rocks is extensively discussed in Snoke et al. (1998). The non-genetic classification scheme of Sibson (1977) is

the most widely used today. The scheme is based on the distinction between random fabric and foliated deformation zone rocks, as well as the cohesion of the rock (the degree to which it behaves as a continuous body during deformation) and proportion of matrix to porphyroclasts. The original scheme is slightly modified to allow for a range of foliation intensities in fault rocks in Table 1.2, by including the additional category of foliated cataclasites (e.g. Chester and Logan 1987), and by extending the gouge, breccia and pseudotachylite categories to the foliated category. Random fabric has been changed to Unfoliated to allow for deformation zone rocks that have some order to their fabric, for example fragments with a jigsaw texture, and Foliated has been changed to Strongly foliated in order to contrast with Unfoliated. Many deformation zone rocks have two or even three distinct foliations. All foliations should be considered when assessing foliation intensity, and the existence of different foliations can be used to further classify deformation zone rocks (e.g. S-C mylonites, Lister and Snoke 1984). The crush breccia series of the original classification has been omitted for simplicity, and because these terms have not found widespread use. The classification scheme in Table 1.2 is based on description at the hand specimen scale, and recognizes that larger fragments in a finer grained matrix is a fundamental character of most deformation zone rocks. However, as noted by Sibson (1977), any pidgeon-hole classification such as Table 1.2 suffers from the problem that the classification criteria may show a continuous range of variation. This is especially problematic in the assessment of foliation development, which may be difficult to quantify, or to judge objectively on a qualitative basis. Another problem with the classification is the definition of cohesion, which is specified in the original classification as primary cohesion, i.e. cohesion during deformation. Post-tectonic processes may decrease cohesion, for example by weathering, or increase cohesion by cementation. These processes may not be recognized easily or allowed for when assessing primary cohesion.

1.7

Format and use of this book

Chapters 2 to 5 deal with the major categories of deformation microstructures and mechanisms. Each chapter begins with a brief introduction to the fundamental mechanisms, and proceeds to descriptions of the characteristic microstructures, identifying which are diagnostic, and illustrating them with diagrams and black-and white photomicrographs embedded in the text. Colour photomicrographs are collected separately and referred to as plates in the text. Key words are emphasized where they occur for the first time in the chapter. Some mechanisms and microstructures, especially cataclasis, are dealt with in more detail than others because comprehensive descriptions are lacking in the literature. The mechanisms involved in the formation of each microstructure are interpreted from experimental and theoretical backgrounds. The final chapter shows how deformation microstructures and mechanisms can be used to make quantitative inferences about deformation conditions for tectonic analysis. Throughout the text, relatively new developments in the subject, and those that have not been described in previous textbooks, have been

6

more heavily referenced than other topics. The references are presented in two forms. The main list gives all references in full. This is followed by a list of abbreviated references collected by chapter, which shows important general sources for the chapter topics in italics. A general list of symbols, abbreviations and units prefaces this chapter. Abbreviations and conventions for all photomicrographs are as follows: PPL - Plane polarized light XP - Crossed polarized light RF - Reflected light


ST - Sensitive tint (gypsum) plate inserted SEM - Scanning electron microscope optical image TEM - Transmission electron microscope image CL - Cathodoluminescence image from the SEM Figures below the captions give the horizontal dimension of the image in mm. Shear sense (see Chapter 7) is given as sinistral or dextral (assuming that the shear or fault plane is vertical); all illustrations with shear senses given are perpendicular to the shear plane and parallel to the shear direction, and the shear plane is approximately parallel to the horizontal edge of the image.

Chapter 2

Cataclasis 2.1 Introduction

The second model treats the microcrack as flat with a sharp tip (e.g. Lawn and Wilshaw 1975a), and gives a versatile Microfractures, displacements and rotations of rigid particles solution for the stress field around the microcrack, in the polar with no permanent lattice distortion are microstructures coordinates of Figure 2.1b: formed by cataclasis. There are two fundamental cataclastic deformation mechanisms: microcracking (Section 2.2.1) and frictional sliding (Section 2.2.2). The single most distinctive cataclastic microstructure is the microfracture, defined The stress is thus specified by K, the Stress intensity factor, as a tabular or planar microscopic discontinuity. The term describing the intensity of the field around the microcrack, thus includes microfaults, microscopic deformation bands, and the stress distribution, described by radial factor microjoints, microcracks, microveins, and microscopic slip and a function of which depends on the propagation mode. surfaces. Microfractures can be sub-divided into microfaults The three microcrack propagation modes shown in Figure 2.2 (Section 2.4), which contain a fragmental matrix, and mi- are tensile or opening mode, (mode I), in which displacecrocracks (Section 2.3), which do not. Pseudotachylites are ments are perpendicular to the fracture plane, and two shear briefly discussed in Section 2.11 because they are associated modes, in which displacements are parallel to the fracture plane, sometimes called sliding and tearing modes (modes II with cataclastic mechanisms. and III, Fig. 2.2). The stress intensity factor depends on the propagation mode, the microcrack length and the applied 2.2 Fundamental cataclastic deforma- stress

tion mechanisms 2.2.1

Microcracking Given these descriptions of the stress around the flat microcrack, it is possible to deduce the failure criteria for brittle solids from Griffith’s postulate that a microcrack will extend when the total energy change with the propagation of the microcrack is negative or constant. The energy terms involved in microcrack propagation are release of mechanical energy, which must be equal to the energy required for creation of surface area.In terms of force, the microcrack extension force, G, must be greater than or equal to twice the surface tension force, (the doubling factor accounts for the two sides of the crack). This leads to the classical Griffith result for failure under a tensile remote stress in a solid with Young’s modulus E:

Microcracking involves microcrack nucleation followed by propagation. Nucleation is irrelevant in a geological context because of abundant heterogeneities in natural rocks and minerals. Dynamic microcrack propagation

Microcrack propagation from an initial flaw can be considered by two different models. The first model assumes an elliptical microcrack (Inglis 1913, Griffith 1924). For a uniaxial remote applied stress parallel to the microcrack axis, the tangential stress, on the microcrack surface varies from a negative (tensile stress) equal to the value of at the long axis of the microcrack, to a compressive stress at the short axis (Fig. 2.la). In a biaxial stress field is equal to at the long axis of the microcrack (Fig. 2.1a; Jaeger and Cook 1979). The stress state around an elliptical microcrack illustrates the essential concepts that large tensile stresses can develop at the tip of a microcrack surface in compression, and that the maximum tensile stress will develop in the direction of the maximum applied stress. These two concepts explain why extension microcracking perpendicular to the least principal stress is a widespread and fundamental cataclastic mechanism.

The well-known Griffith failure criterion and its derivatives can be determined from this relation (Section 9.2). The Griffith failure criterion thus implicitly assumes the existence of microcracks of length Griffith observed that the theoretical strength of solids is much greater than their actual strength, and postulated that real solids contain microcracks which concentrate stress leading to failure, thus explaining the discrepancy between theoretical and measured strengths. The existence of such microcracks or “Griffith flaws” is still accepted as the basis of most failure criteria. 7

8

CHAPTER 2. CATACLASIS

9


The model so far assumes elastic behaviour, but an additional energy term must be incorporated to account for breaking of atomic bonds at the crack tip, which can be done by postulating the existence of a small zone ahead of the microcrack tip in which non-linear forces are expended in breaking bonds. The additional energy involved is incorporated in the energy balance in the form of a new parameter the fracture toughness, to replace the surface energy term The energy balance condition is now:

This equation gives the condition for microcrack propagation in terms of a value for G usually known as the critical fracture toughness, which can be related to the critical stress intensity factor, for a given geometry. This analysis is valid provided that the size of the nonlinear zone is much smaller than the length of the microcrack i.e. it does not affect the elastic stress system as a whole: this assumption is called the small scale yielding (Rice 1968) or the small scale zone approximation (Lawn and Wilshaw 1975a). More detailed analyses can relate to the displacement on the microcrack if some function of the cohesive forces ahead of the tip with distance is postulated: these are the cohesive force models (e.g. Rudnicki 1980). The existence of such a non-linear zone ahead of the microcrack tip is known experimentally from ceramics and metals, where it is referred to as a process zone, and it has been detected from acoustic emission and microcracking in geological materials (e.g. Swanson 1981, Peck 1983, Labuz et al. 1987). Sub-critical microcrack propagation

The above analysis is restricted to microcracks that propagate at speeds that are typically significant fractions of the velocity of elastic waves in solids, or dynamic microcracking. However, microcracks may propagate under stress conditions well below the critical stress intensity factor, at rates that depend on temperature and chemical environment as well as stress intensity. This phenomenon is known as sub-critical microcrack growth and is potentially of enormous geological importance (e.g. Anderson and Grew 1977, Das and Scholz 1981, Atkinson 1982, Meredith 1983). A diagram showing microcrack velocity (V) as a function of mode I stress intensity factor illustrates the main features of sub-critical microcrack growth (Fig. 2.3), which have been demonstrated for a range of geological materials (e.g. quartz, granite, andesite, basalt, calcite, oil shale, sapphire and glass). The relationship falls into three parts: Region 1. Velocity is highly sensitive to water concentration and temperature. There may be a threshold stress intensity factor for microcrack growth to occur at all (Meredith 1983). Region 2. Velocity is dependent on water concentration and temperature, but not Region 3. Velocity increases extremely rapidly with microcrack growth becomes dynamic at

until

Five mechanisms of sub-critical microcrack growth have been proposed, but stress corrosion, a general term for environmentally influenced, stress driven, thermally activated chemical reaction allowing breaking of bonds is considered to be the dominant mechanism for geological materials in crustal conditions (Atkinson 1982). Hydrolysis of the Si-O-Si bond is likely to be responsible for the weakening. The microcrack velocity in Region 1 is controlled by the reaction rate of the hydrolysis, and transport-rate control occurs in Region 2. Other types of chemical reaction occur with sub-critical microcracking. An example is the reaction of plagioclase to increasingly sodic compositions and ultimately to laumontite, which creates a 60% volume increase (Blenkinsop and Sibson 1991). This alteration occurs along cleavage microcracks, resulting in a texture of expanded, matching fragments, derived from a single parent crystal (Plate 1). This texture suggests chemical alteration and sub-critical microcracking occurred in a linked process, which can be called alterationenhanced microcracking. Sub-critical microcrack growth has been incorporated into some models of crustal processes, for example to explain the difference between creep (stress intensity factor between and and seismic faulting (stress intensity factor equal to Rudnicki 1980), the barrier theory for earthquake rupture of Das and Scholz (1981), and features of magmatic intrusion and hydrofracture propagation (e.g. Anderson and Grew

10


Adhesive strength, asperity interlocking and increasing contact area have been combined in an elegant and simple model by Wang and Scholz (1994, 1995), which accounts for experimental results very well. Frictional sliding leads to the production of gouge by asperity failure and ploughing. Once formed, fragmentation within a gouge layer continues by mutual impingement of particles, leading to particle size distributions (PSDs) that are characteristically fractal (e.g. Blenkinsop 1991). A fractal distribution of particle sizes can be described by the relation:

1977, Atkinson 1982).

2.2.2

Frictional sliding

Amonton’s law that the steady state shear stress is proportional to the normal stress on the sliding surface via the coefficient of friction

is predicted by a simple adhesion theory of friction. The theory assumes that the rough surfaces contact at asperities (protruding irregularities), which will yield under normal stress until sufficient area of contact is established to support the normal load (Fig. 2.4). The contact area is considered to have an adhesive strength which must be exceeded over the entire area for sliding to occur. This model successfully predicts the normal-stress dependence of friction, but it underestimates the value of because there are other contributions to the shear strength in addition to the adhesive strength of the contacts (Scholz 1990). These include: 1. Interlocking of asperities. Oblique surfaces of contact are created between two asperities that come into contact after sliding (Fig. 2.4). Interlocking may be relieved by shearing through asperities (adhesive wear), or by sliding on the oblique contacts, which causes dilatancy. 2. Increasing area of contact due to asperity shearing. As sliding continues, asperities fail and the contact area increases.

where N (d) is the number of particles greater than size d, and D is the fractal dimension. D for particles in cataclastic rocks ranges from 1.88 to 3.08 (e.g. Sammis et al. 1987, Sammis and Biegel 1989, Olgaard and Brace 1983, Wang 1987), with many results for natural and experimental gouges around the value of 2.6 (Marone and Scholz 1989, Biegel et al. 1989). Sammis et al. (1987) showed that a distribution of spheres of unequal sizes reduces impingement stresses on individual fragments, and proposed that microcracking in a gouge will proceed in order to minimize the probability of neighbouring grains having equal sizes. This constrained comminution model predicts D = 2.58, which is very close to observed values (e.g. Sammis and Biegel 1989). Gouge formation (and cataclasis generally) may occur by at least two other processes of particle size reduction. Alteration, for example the laumontization of feldspar described above, may lead to lower values of D (Blenkinsop and Sibson 1991). At advanced stages of gouge formation, selective microfracture of larger particles takes place, creating PSDs with fractal dimensions greater than 2.58 (Blenkinsop 1991). Plates 1 to 4 show a sequence of cataclastic textures arranged in order of increasing fractal dimension of PSDs, which is the evolutionary sequence of textures with strain. A simple theory of wear can predict that the volume of gouge created by sliding, and hence the thickness of the gouge, will increase linearly with displacement (Scholz 1990), and some experimental results confirm this relationship (e.g. Teufel 1981). This relationship has also been claimed for faults (e.g. Robertson 1983, Hull 1988). However, there are at least two reasons the experimental data should not be applied directly to faults. Firstly, the roughness of laboratory prepared surfaces is not comparable to natural fault surfaces (Brown and Scholz 1985), and secondly, once a gouge layer has accumulated sufficient thickness to prevent interaction between the sliding surfaces, wear will no longer occur according to the simple model of surfaces in contact with each other. The reported data on fault gouge thicknessdisplacement relationships does not substantiate a proportionality (Blenkinsop 1989).

2.3

Microcracks

2.3.1

Classification, characteristics and obser-

3. Asperity ploughing. An asperity with a greater strength vation than the opposite surface will cut into the weaker material, generating a groove and wear fragments. This is Microcracks can be classified as intragranular (within single grains), transgranular (across two or more grains) and cirknown as abrasive wear.


11

12


cumgranular or grain boundary. The occurrence of these different types of microcrack depends on the microcrack mechanism (see below) and on the microstructure of the rock. Intragranular microcracks are characteristic of poorly cemented, highly porous rocks, whereas well-cemented, low porosity rocks have transgranular microcracks. This classification is useful in discriminating various microcrack mechanisms described in Sections 2.3.3 to 2.3.11. Tectonic trans- or intragranular microcracks commonly link contact points between adjacent grains, and are kinked or curved. Despite grain-scale irregularities in fracture geometry, microcracks generally have a strong preferred orientation. Several sets of microcracks may exist within one rock. Extension (mode I) microcracks are usually filled with material in optical continuity with the host grains, so that their importance may be underestimated or even completely overlooked. The only manifestation of a fracture left in the rock may be a line of fluid inclusions which are the trace of a fluid inclusion plane (Fig. 2.5, Section 3.11), and careful observations of well-prepared thin sections at high magnifications under the optical microscope are necessary to detect such microcracks if they contain small inclusions. Microcracks can have a wide range of aspect ratios and densities. Extension microcracks may have regionally systematic orientations because they form perpendicular to (Section 2.2.1.1), and have been used very effectively to deduce regional stress systems (e.g. Lespinasse and Pêcher 1986). Cryptic microcracks may be spectacularly revealed by cathodoluminescence (CL) (e.g. Smith and Stenstrom 1965, Sprunt et al. 1978, Sprunt and Nur 1979, Blenkinsop and Rutter 1986). Luminescence depends on silica polymorph, and chemical, thermal and mechanical histories (e.g. Seyedolali et al. 1997). Microfracture fillings which form under different conditions from the host grain, and experience only part of their tectonic history, therefore contrast in luminescence with the rest of the grain and usually have very low size, sorting, and grain shape. Impingement microfracturing luminescence (Fig. 2.6). can be understood from an analysis of the stress field created on loading a plane surface by a pointed (Boussinesq 2.3.2 Microstructures and mechanisms configuration) or spherical (Hertzian configuration) object Nine microcrack “mechanisms” can be distinguished, mainly (Fig. 2.8). The point load should produce radial microcracks from experiments, where extension microcracks, usually sub-perpendicular to the indenter, and such microcracks have known as axial microcracks, form from about half the peak been observed in indenter experiments (e.g. Lindquist et al. strength through to post-failure (e.g. Tapponier and Brace 1984). The spherical indenter (possibly more geologically 1976). These are secondary mechanisms compared to the realistic) will contact the plane surface along a spherical surfundamental physics of tensile microcracking described in face, inducing a region of compressive stresses immediately Section 2.2.1, and they mainly describe specific geomet- below the indenter, surrounded by tensile stresses near the ries that create microcrack tip tensile stresses (Krantz 1983, edge of the contact surface. An extension microcrack in the Atkinson 1982). Table 2.1 summarizes the characteristic fea- form of a cone (cone microcrack) forms beneath the indenter tures of the mechanisms, including the types of microcrack (Fig. 2.8) at a critical load. The critical indenter radius to gen(intra-, trans- or circumgranular) that can form by each mech- erate a tensile microcrack depends on the applied pressure via anism. a square root (Lawn and Wilshaw 1975b): very modest pressures for even quite large indenters can create microcracks. A still more geologically realistic configuration is the case 2.3.3 Impingement microcracks of two spheres loading each other: in this case, an extension Impingement microcracks link points of contact between ad- microcrack initiates at the edge of the contact plane between jacent grains, and are usually intragranular. They may link the two spheres at a critical load that depends on porosity, several pairs of contact points around grains. Four basic grain size, elastic moduli and fracture toughness, providing patterns are shown in Fig. 2.7 (Gallagher et al. 1974), that there are pre-existing flaws present in the loaded grains which depend on the boundary loads, packing arrangement, (Fig. 2.8, Zhang et al. 1990). The critical load measured in

13


several rocks (Wong 1990) follows the theoretical relationship and shows that the required pre-existing flaws have submicron dimensions, comparable to the flaws invoked in Griffith’s failure analysis (Section 2.2.1). Photoelastic experiments on convexo-concave contact surfaces which model indenting grain contacts due to pressure solution (Section 3.4) show that tensile microcracks should form normal to the indenter contact, and shear failure is predicted along curved trajectories that are approximately normal to the contact in its immediate vicinity, but deviate progressively with distance (McEwen 1981). The diagnostic feature for recognizing the impingement mechanism is linking of contact points by intragranular microcracks (Fig. 2.9, Table 2.1), although the contact points may not be visible in the plane of the section. Impingement microcracking is suppressed in well-cemented or low porosity crystalline rocks because impingement contacts are lacking and tensile stress concentrations are dramatically reduced by the cement (e.g. Wong and Wu 1995, Menéndez et al. 1996).

2.3.4

flaw (Fig. 2.6, Table 2.1). Such microcracks can be intra-, trans- or circumgranular.

2.3.5

Microfracturing of pre-existing flaws

Microfracturing of pre-existing flaws is known as an important and even dominant microcracking mechanism from experiments (e.g Biegel et al. 1992, Menéndez et al. 1996). It is considered to be at least as important as cleavage in controlling the overall microcracking process, and Atkinson (1982) asserts that it dominates the upper 20 km of the crust. The opening of grain boundaries (a type of pre-existing flaw) is well known to contribute to experimental cataclasis (e.g. Dunn et al. 1973, Hadizadeh 1980, Tapponnier and Brace 1976). The weakness of some natural grain boundaries, even when these are overgrown by optically continuous quartz cements, is apparent from the observation that overgrowths are a major component in the matrix of faults (Pittman 1981). Microfracturing of pre-existing flaws can be identified by microfracture of a cement, and may involve all three types of microcrack (Fig. 2.11, Table 2.1).

Flaw-induced microcracks

Flaw-induced microcracks are joined to flaws such as other microcracks, pores and grain boundaries. They form because of the tensile stresses that develop on the flaw surface when remote stresses are imposed. They are recognized in experiments on analogues and on rocks (e.g. Brace and Bombolakis 1963, Tapponnier and Brace 1976). Analytical solutions show that the microcracks will grow along curved trajectories from both ends of a flaw to produce the well-known “wing crack” geometry (Fig. 2.10, Horii and Nemat-Nasser, 1985, 1986, Kemeny and Cook 1987, Baud et al. 1996). For isolated microcrack growth, a flaw length 2c, orientated at to is assumed to have both cohesive and frictional resistance to shear, and tensile microcracks grow from both ends with length l (Fig. 2.10), making an angle with the flaw. The results show that microcracks will grow by tensile failure from the edges of the flaw along paths which fit experimental observations very well (Horii and Nemat-Nasser 1985). After a critical length (approximately 1.0) in even the slightest tensile value of microcrack growth becomes unstable, but remains stable in compression. Flaw-induced microcracking can be recognized by the connection of the microcrack to a

2.3.6

Cleavage microcracks

Microcracks in biotite are controlled by the basal (001) cleavage (Plate 42, Wong and Biegel, 1985). In feldspars, the major (001) cleavage and also (010), and (110) planes exert a strong influence on microcracks (Willaime et al. 1979, Brown and Macaudière 1984, Tullis and Yund 1992). Cleavage microcracking of feldspars is important during deformation of granitic rocks in the upper crust (Plate 1; Evans 1988). The fracture toughness of quartz is least along the rhombohedral planes, followed by the basal plane: these are preferentially exploited during fracture (e.g. Borg et al. 1960, Vollbrecht et al. 1991). Cleavage microcracks can be recognized because they occur in crystallographically controlled sets parallel to known cleavages within single grains (Table 2.1).

2.3.7

Elastic mismatch microcracks

Microcracks have been noted in quartz and feldspar grains at contacts with micas in experimental studies (Tapponier and Brace 1976, Wong and Biegel 1985), and in a naturally deformed quartzite (Hippert 1994). These microcracks are

14


thought to develop because of the difference in elastic strain across the quartz-mica or feldspar-mica boundaries due to the different elastic moduli of the two minerals in contact along a coherent interface (Plate 5, Fig. 2.11). They can be recognized by the localization of intragranular microcracks around contacts between grains of different mineralogy (Table 2.1), but may be difficult to distinguish from thermally-induced microcracks (see Section 2.3.10).

2.3.8

Plastic mismatch microcracks

Where intracrystalline plasticity is localized in one area (e.g. in twins, deformation lamellae and kinks, Chapter 4), microcracks may be initiated due to the strain incompatibility between the area of plastic deformation and the adjacent undeformed area (e.g. Olson and Peng 1976, Wong and Biegel 1985; Fig. 2.12). Microcracks have been observed along kink bands in naturally deformed enstatite, and normal to the kink bands in experimentally deformed quartz (Carter and Kirby 1978). Plastic mismatchs may account for the common microcracking of feldspar porphyroclasts surrounded by deformed quartz grains in quartzofeldspathic mylonites (e.g. Evans and White 1984; Fig. 2.13). Plastic mismatches may also occur within single phases or grains due to stress concentrations created by intracrystalline plasticity (e.g. Lawn and Wilshaw 1975a). Plastic mismatch-induced

microcracking is largely responsible for semibrittle behaviour (e.g. Carter and Kirby 1978). It can be recognized by the close association between intragranular microcracks and areas or individual microstructures of intracrystalline plasticity, such as subgrains, kink bands, deformation lamellae or twins (Table 2.1).

2.3.9

Microfault-induced microcracks: Microscopic feather fractures (mffs)

Mffs are intragranular microcracks found only adjacent to faults. They are characteristically wedge-shaped, opening towards the fault plane (Fig. 2.12c). They were identified in experimentally generated faults by Friedman and Logan (1970), who found them exclusively within 5-10 grain diameters of shear faults, and parallel to They did not occur adjacent to an incipient shear, and therefore formed in response to shearing. Conrad and Friedman (1976) defined mffs as microcracks occurring only within grains adjacent to a fault, dying out rapidly away from the fault and statistically close to or parallel to the applied direction of Microcrack density and length increase with displacement and normal stress (confining pressure) (Conrad and Friedman 1976, Teufel 1981). Tectonic analogues of mffs have been observed associated with shear surfaces between pebbles in contact with each other (McEwen 1981). T fractures as described by Petit (1987)


have the characteristics of mffs. Mffs are created by tensile fracture at contact points along the sliding surface (Teufel 1981). The relations between microcrack density, length, displacement and normal stress observed in the experiments are all consistent with the formation of mffs due to contact stresses on the sliding surface. Mffs are intra- or transgranular microcracks that can be recognized by their localization adjacent to the fault plane, their inclinations of 20-50° to the fault plane, and their wedge-shape opening towards the fault plane (Table 2.1). Mffs can be distinguished from Riedel microfractures, which may also be associated with fault planes (e.g. Petit 1987), by the shear offset along the latter.

2.3.10 Thermally-induced microcracks Microcracks can relieve stresses caused by differential thermal expansion or contraction between adjacent minerals. Such microcracks may form in grains of one mineral surrounded by another during heating or cooling. If heating or cooling are accompanied by pressure changes, elastic mismatch microcracks may also form. Thermally-induced microcracking can only be distinguished from elastic mismatch-induced microcracking if the P-T path is known. The case of cooling granite has been considered in some detail by Bruner

15

16


(1984) and Vollbrecht et al. (1991). Granite can be treated as a composite of quartz surrounded by feldspar. Two extreme cases can be considered during cooling and uplift. In isothermal decompression, greater elastic expansion of quartz may cause microcracking of feldspar, while isobaric cooling may lead to microcracking of quartz due to its greater thermal contraction. The critical geothermal gradient for equilibrium between thermal and elastic strains in quartz and feldspar is 10°C/km. Crystal anisotropy may be important on the grain scale; quartz has a maximum coefficient of thermal expansion perpendicular to the c-axis, favouring microcracks at low angles to the c-axis. Regional stresses can also be important: microcracking will be favoured in those grains with appropriate orientations relative to the regional stress.(e.g. Vollbrecht et al. 1994). The general interaction between thermal and elastic stress around inclusions has been modelled by D’Arco and Wendt (1994), and specifically for garnet by Whitney (1996). Thermally-induced microcracking in granites can be recognized by intragranular microcracks concentrated in quartz surrounded by feldspar. Thermal microcracks in quartz may have a preferred orientation parallel to the c-axis.

2.3.11

Phase transformation-induced microcracks

The strain associated with solid state phase transformation can produce distinctive microcracks. For example, the

transformation involves a volume increase of 11%. Quartz inclusions in garnet or omphacite are surrounded by radial extension microcracks (e.g. Chopin 1984, Smith 1984, Wang et al. 1989), and indeed this texture has been used as evidence for the former presence of coesite (e.g. Wang and Liou 1991). Although it is clear that the transition has occurred in these rocks because relict coesite can be found in some inclusions, the question arises whether such microcracks could be due to elastic mismatches between the silica phase and the host. The key evidence for phase transition microcracking is the observation that there is no microcracking around other types of inclusion, including rutile. The extensional nature of the microcracks and their origin from tips of inclusions are consistent with the mechanism. Fracture surface energy measurements in the conditions of the quartz phase transition imply that microfractures could also be associated with this transformation, and other minerals undergoing similar phase transformations (Kirby and Stern 1993). Radial microcracks have also been observed around calcite inclusions which have replaced aragonite, a transformation that involves a 8.5% volume increase (Wang and Liou 1991). The distinctive features of phase transformation microcracking are the association of intragranular microcracks with evidence for phase transformation. In the case of the transition, radial microcracks around inclusions of quartz after coesite are distinctive.


17

2.4 Microfaults 2.4.1 Characteristics Microfaults are shear microfractures that contain grain fragments formed by cataclasis (Plate 3). Displacement parallel to a microfault surface can be identified from displaced grain boundaries or fragments. However, caution is needed in positively identifying such shear displacements, because a purely extensional microcrack can appear to have a shear displacement when viewed in a section which is oblique to the opening direction of the microcrack. A useful indication of true shear displacement is the consistent sense of offset of a number of grain boundaries at a variety of angles to the microcrack (Plate 45). Two components can comprise microfault matrix: fragments derived from the wall rock, and a precipitated cement. The fragments are usually angular and very poorly sorted, ranging in size from micrometres up to the width of the microfault. It is sometimes possible to identify the parent grain from which the fragments have been derived (the sensitive tint plate is useful for this purpose; Boldt 1995), and fragment displacement or rotation may also be detected (Plates 2, 4, 45). The proportion of cement may vary from relatively little, in which the fragments may form a jigsaw texture, to a majority, in which the fragments will be matrix-supported and isolated from each other. Precipitated cements can be identified from euhedral crystal faces overgrowing primary grains (e.g. Pittman 1981) or from CL studies (e.g. Stel 1981) because the precipitated cement has a different luminescence from the grains in the matrix (Section 3.9). A cyclic history of cementation and shear can sometimes be deduced from the presence of resheared matrix. The edges of microfaults may be planar and truncate adjacent grains cleanly, or they may be irregular, perhaps comprising a large number of circumgranular microfractures.

2.4.2

Mechanisms

Experiments identify the shear failure mechanism as the linking of extension microcracks once a critical microcrack density is achieved (Krantz and Scholz 1977, Costin 1983). Several mechanisms of microcrack linkage have been suggested. Peng and Johnson (1972) proposed that axial microcracks became linked when individual beams, bounded by

axial microcracks, buckled at a critical fibre strain, allowing the first through-going fault plane to form (Fig. 2.14). This linking mechanism can be contrasted with the behaviour of other siliclastic rocks, in which grain boundary microcracks form first, to be linked by axial microcracks (Hadizadeh 1980, Menéndez et al 1996). The detailed sequence of microcracking and linkage probably depends on the initial microstructure of the rock (Hadizadeh 1980, Blenkinsop and Rutter 1986). A third type of linking mechanism is the direct interaction of microcrack stress fields. Such interactions were classified into en-echelon and en passant types by Krantz (1979) (Fig. 2.15). The former occur between two straight, sub-parallel microcracks, which are linked by a third straight microcrack (Fig. 2.15 a, b). En passant interactions involve curvature of microcrack paths because the microcrack tip stress fields influence each other (Fig. 2.15 b, c). The definitive theoretical model of microcrack interactions is based on solutions for the behaviour of isolated microcracks in compression (Horii and Nemat-Nasser 1985). To model shear failure, the microcracks interactions are considered in an array of parallel flaws with individual angles and overall angle to and spacing d (Fig. 2.16). The results show that axial compression at first increases with for all values of i.e. microcrack growth is stable. However, at lower values of microcracks interact unstably at a certain stress. In general the model predicts that and are different, suggesting that a failure plane will consist of oblique microcracks linked by axial microcracks. The model correctly predicts the effect of confining pressure on ultimate strength and is validated by experiments on resin blocks. Natural microfaults seem to evolve by microcrack linking at a critical microcrack density

18


guished between deformation bands with no cataclasis, deformation bands with cataclasis, and deformation bands with clay smearing. Deformation bands tend to cluster together to form zones of bands, often in an anastomosing pattern (e.g. Engelder 1974, Blenkinsop and Rutter 1986). The thickness of these larger-order features depends on the number of individual bands which comprise them, and their total displacement is given by the sum of the individual components. Deformation zones may contain discrete surfaces with large discontinuous displacements that cut through all other features, and may have striations on their surfaces. These were described as slip surfaces by Aydin and Johnson (1983), and are the latest features to form in a sequence from deformation bands to zones to slip surfaces. The distinguishing characteristics of deformation bands are shear displacements which preserve material continuity on a mesoscopic scale, and their occurrence in porous, granular rocks.

2.5.2

in similar ways to experiments and models (e.g. Blenkinsop and Rutter 1986).

2.5 Deformation bands 2.5.1

Characteristics and classification

A distinctive type of localized deformation occurs in porous granular materials. Deformation bands are roughly planar features from millimetres to a hundred metres long, from fractions to a few millimetres wide, and with net slips from fractions to tens of millimetres. Displacement across deformation bands is continuous at a mesoscopic scale, which distinguishes deformation bands from faults. Porosity in deformation bands may be greater or less than the wall rocks; in the more common case of porosity reduction, the bands contain a layer of fine-grained matrix partly comprised of grain fragments. Intragranular extension microcracks are found within this type of deformation band. There is a continuum between faults and deformation bands, with many descriptions of faults having some characteristics of deformation bands (e.g. Engelder 1974, House and Gray 1982, Jamison and Stearns 1982, Underhill and Woodcock 1985, Narahara and Wiltschko 1986, Blenkinsop and Rutter 1986, Cruikshank et al. 1991, Zhao and Johnson 1991). Deformation bands were explicitly described by Aydin (1978), Aydin and Johnson (1978, 1983), and most comprehensively by Antonellini et al. (1994), who distin-

Mechanisms

Dilatancy caused by grain sliding, and impingement microfracturing are important mechanisms in deformation band formation. Experiments, theory and observations suggest that the differences between the three categories of deformation band are related to confining pressure and initial microstructure (Antonelli et al. 1994). Lack of grain microcracking in the first category may indicate low confining pressures during deformation compared to the bands in the second category. Initial porosity determines whether a deformation bands dilates or compacts: low initial porosity leads to dilation, and vica versa. These relationships are predicted by critical state theory for granular materials (e.g. Schofield and Wroth 1968). The third category forms as a result of the higher clay content in these rocks. The localization theory of cataclasis accounts for the formation of microfaults, deformation bands, zones and slip surfaces (e.g. Rudnicki and Rice 1975, Aydin and Johnson 1983). Localization is defined as a difference in strain rate between a band and the matrix, and is a highly appropriate way to describe deformation bands since strain appears continuous across the bands. The theory predicts the formation of deformation bands at or before peak stress, and their subsequent spread due to strain hardening. The development of slip surfaces can be understood from the same theory as a process in which stress within the deformation band becomes too great to be accommodated with the result that a discontinuity develops.

2.6 Distributed cataclasis and cataclastic flow A number of experiments have produced distinctive cataclastic microstructures in samples that have maintained strength without localization of deformation. The microstructures are characterized by distributed fracture and displacement of fragments, but the deformation is macroscopic-

19


ally continuous: this is called cataclastic flow, which can be defined as deformation by cataclastic mechanisms leading to continuous flow at a given scale. Cataclastic flow occurs at higher confining pressures than faulting in equivalent rocks, and is favoured by lower differential stresses. Extensively damaged grain boundaries, perhaps produced by fracture under sliding indenters (microfault-induced microcracks, Section 2.3.9), and wear products accumulated in the sliding zone and in pores are characteristic microstructures of cataclastic flow in quartzites. (Rutter and Hadizadeh 1991). Compaction and filling of pores by crushed material seems to be typical of cataclastic flow of siliclastic rocks in general (e.g. Menéndez et al. 1996). Axially orientated microcracks are dominant. Microstructures produced during experimental cataclastic flow of feldspar aggregates consist of intragranular shear microcracks mainly along feldspar cleavages producing blocky fragments between the cleavage microcracks at 90° to each other (Tullis and Yund 1987, 1992). The spacing of the microcracks is as low as Deformed grains have a puckered appearance consisting of areas of patchy extinction, caused by slight mismatches between adjacent blocks, which resemble subgrains or even recrystallized grains formed at higher temperatures by intracrystalline plasticity. However, TEM observations established that no dislocation processes contributed to the block rotation. Microcrush zones < wide have sharp boundaries and strong grain rotation. A crystallographic preferred orientation and a strong grain shape fabric developed by slip on the closely-spaced grain-scale faults that are geometrically similar to crystal plastic slip systems, and mechanical twins were formed. Similar features have been observed in cataclastic flow of anorthosite, and lamellar features interpreted as bundles of cleavage microcracks were also observed (Hadizadeh and Tullis 1992). Twinning is evidence for the operation of semibrittle deformation mechanisms. It is likely that quartz does not undergo the same cataclastic flow regime observed in feldspars, but requires a thermallyactivated mechanism to stabilize flow (Hirth and Tullis 1991). The difference between feldspar and quartz behaviour is due to the excellent cleavages in feldspar. Pyroxenes and amphiboles may behave in a similar fashion to feldspar (Tullis and Yund 1992). Cataclastic flow has also been reported from experiments on basalts (Mogi 1965; Shimada 1986), dunite and gabbro (Byerlee 1968). A maximum of 30% shortening has been achieved in experimental cataclastic flow, after which faulting generally occurs, and therefore it is unclear whether cataclastic flow can be stable to higher strains. The problem is complicated because faulting is artificially initiated by the experimental configuration. If porosity reduction is necessary for cataclastic flow in porous rocks (e.g. Rutter and Hadizadeh 1991), this must limit the amount of strain that can accumulate by cataclastic flow. At the opposite extreme of scale, cataclastic flow has been suggested for continents (Gallagher 1981), and at intermediate scales such as kilometric scale folds (Stearns 1968, Hadizadeh and Rutter 1983, Blenkinsop and Rutter 1986). In all these cases, the deformation is discontinuous at a smaller scale than the scale of observations.

2.7 Gouge zone microstructures Faults formed at low grade conditions commonly contain gouge with a distinctive set of microstructures, which have been closely replicated by experiments (e.g. Logan et al. 1981, Rutter et al. 1986). Most of the features of gouge zones can be understood in the context of a zone of simple shear with boundaries parallel to the shear plane (Fig. 2.17). P-foliation is formed by grains (most commonly phyllosilicates) aligned at an oblique angle to the gouge zone boundary (Fig. 2.17, Plate 6). It may be inosculating and rather variably developed. It is one of the first microstructures to form in experimental gouges. Shears parallel to the P-foliation with the same sense of shear as the gouge zone are known as P-shears. Riedel shears (R) form at an oblique angle in the opposite direction to the P-foliation (Fig. 2.17, Plate 6). They are discrete shears with the same sense of shear as the gouge zone and have geometrical similarities to extensional crenulation cleavages and (Section 7.5, cf. Platt and Vissiers, 1980). Conjugate Riedel shears form at a large angle to the gouge zone boundary with the opposite sense of shear (Fig. 2.17). They are generally much less prominent in gouge zones than Riedel shears. T fractures or extension fractures are inclined to the gouge zone boundary in the opposite direction to the P-foliation (Fig. 2.17). They may be localized in more competent rocks in the gouge zone. Competent rock may form boudins, which are commonly asymmetric (Section 7.10) and separated along Riedel shears (Fig. 2.17, Plate 6). Ductile stringers (Fig. 2.17) consist of relatively large, rigid, asymmetric clasts elongated in the direction of the P-foliation, with stepped tails that may become detached from the main clast (Logan et al. 1981). Asymmetric folds are also found in gouge zones: they may fold earlier boudins. The vergence of the folds may be the same as the shear sense of the gouge zone (Fig. 2.17), although opposite vergence can occur (Section 7.11). The orientation of fold axes within gouge zones may define a girdle distribution parallel to the shear plane. Y-shears are discrete, relatively long shears that cut through the gouge zone with the same sense of shear as the gouge zone (Fig. 2.17, Plate 7). There is usually only one Y-shear in the central part of the gouge zone. In experiments, Y-shears are the last feature to form. The orientation of all features except Y-shears shown in Figure 2.17 is variable since their initial orientation is likely to change with finite strain, because material lines rotate during simple shear. Features rotate clockwise in a dextral gouge zone such as illustrated in Figure 2.17. Components of pure shear may complicate this general rule. The foliation in a foliated cataclasite may originate as a fabric along any of the above shears, or as a P-foliation in a gouge zone (Plate 8).

2.8 Microfracture surface features Microfracture surfaces may show a variety of microstructures that convey useful information about cataclastic deformation. Intact specimens of microfracture surfaces can be easily observed under a binocular microscope, and at greater magnification in the SEM. It is also useful to examine the cross-

20

sections by cutting thin sections perpendicular to the microfracture surface, preferably parallel and perpendicular to the slip direction. There is a rich field of microscopic studies on extension microfracture surfaces in the material science literature (e.g. Bansal 1977, Quakenbush and Frechette 1978, Michalske and Frechette 1980) which suggests that fracture surface features could be used to reveal rates of fracture propagation and even fracture stresses (e.g. Scholz 1972, Martin and Durham 1975, Swanson 1981, Norton and Atkinson 1981, Meredith 1983, Cox and Atkinson 1983). Surface microfracture features such as mirror, mist, velocity hackle and forking, and Wallner lines, which are known to be diagnostic of dynamic microcrack propagation in glass, have not been found on fracture surfaces in rocks. This has been taken to indicate that the studied rock fractures were never dynamic (Kulander and Dean 1995). This conclusion is supported by observations such as irregular rib marks that imply stable propagation at low velocities, again by analogy with glass. A shear microfracture with a smooth or shiny surface is a slickenside, which is commonly marked by parallel lines known as slickenlines or slickenside striations. Many slickensides are the surfaces of microfaults formed during tectonic deformation. The lines are parallel or tapering ridges or grooves parallel to the slip direction. There may be more than one set of slickenlines on a microfault surface in different directions. It is sometimes possible to distinguish overprinting relationships between different generations of slickenlines. Probably the most common mechanism of slickenline formation is asperity ploughing (Section 2.2.2), which can be identified by the preservation of the asperity on one side of the microfracture and a matching groove on the opposite surface. The type of slickenline produced is called a wear groove (Fig. 2.18), tool track or mark, or prod mark (Tija 1967, Hancock 1985). Ridges of gouge may be formed by wearing down asperities or by accumulating gouge around a hard particle: this mechanism is debris streaking. Erosional sheltering occurs where a solid ridge of microfracture surface is preserved in the down-slip direction of a particle. A


category of slickenlines consists of equally developed ridges and grooves with perfectly matching (or nested) profiles on both sides of a slickensided surface (Means 1987). The ridge and groove structure can be formed by localized shear associated with a grain shape fabric, and the length of individual ridges and grooves may be greater than the displacement on the shear zone (Wil and Wilson 1989). Ridge and groove structures can also be formed by inosculation of shear surfaces (C-surfaces) in an S-C mylonite (Section 7.5, Lin and Williams 1992). Another type of slickenside which may also be formed by continuous deformation consists of very fine grained quartz (0.01 to ) with a strong crystallographic preferred orientation (Power and Tullis 1989). The preferred orientation may have resulted from orientated growth by diffusive mass transfer through a solution during the interseismic, low strain rate part of the seismic cycle. Microfracture surfaces may be offset along discontinuities approximately perpendicular to slickenlines, known as risers or steps (Fig. 2.19). Risers are generally less dense and coarser features than slickenlines, and are commonly more irregular. They may be approximately linear or crescent shaped. Risers are described as incongruous if the offset opposes the direction of movement of the opposite block, and congruous if the opposite is true. Crystal fibres on fault surfaces are known as slickenfibres, and risers created by the crystal terminations are accretion steps (Norris and Barron 1968). They are identified by the presence of fibres that are commonly monocrystalline. Risers may form either by initial irregularities in the fault surface, or by the intersection of secondary fractures with the fault surface. Possible geometries of secondary fractures can be separated into P fractures, R, T and fractures using the same terminology as introduced above for gouge zone features (Petit 1987). Slickenfibres form by crystal growth in the dilatant gap adjacent to congruous risers.


21

22

2.9


Crystallographic fabrics

Experimental studies have shown that crystallographic fabrics can be produced by cataclasis (e.g. Borg and Maxwell 1956, Borg et al. 1960). Crystallographic fabrics in feldspars have been produced by slip and rotation on multiple grain-scale faults (Tullis and Yund 1987, 1992). The faults are geometrically analogous to slip systems in intracrystalline plastic deformation (e.g. Allison and La Tour 1977, Section 4.9). Cataclastically-formed crystallographic fabrics can be distinguished from intracrystalline plastic fabrics by the presence of multiple, grain-scale faults such as those illustrated by Tullis and Yund (1992).

2.10

Pre-lithification deformation microstructures and mechanisms

The microstructures of pre-lithification deformation, and even the deformation itself, may be difficult to recognize. A grain shape fabric can be produced by frictional grain sliding and rotation, known as independent particulate flow (IPF) because the deformation is independent of fracture (Borradaile 1981). This deformation mechanism is sometime regarded as separate from cataclasis, because there is no fracturing; here, however, cataclasis is used in a broader sense to encompass IPF. High pore fluid pressures may be essential in pre-lithification IPF (Knipe 1986, 1989, Maltman 1994). The diagnostic feature of a grain shape fabric produced by IPF is the absence of microcracks and lack of internal deformation of grains, but intracrystalline plasticity can also be involved during pre-lithification deformation (e.g. kinking in biotite; Morrit et al. 1982). The fabric of scaly clays (e.g. Agar et al. 1988) may also be characteristic of pre-lithification deformation.

2.11

Pseudotachylites and frictional melting

2.11.1 Characteristics Pseudotachylite is a glassy or very fine-grained rock occurring in veins and associated with deformation zones (Plates 911). It generally consists of a dark matrix which encloses angular to rounded fragments of the wall rock (Plates 9, 11). The fragments are very poorly sorted and may have a high clast/matrix ratio. Clasts in the range to 10 mm in some pseudotachylites have a fractal PSD, with a fractal dimension of 2.5 (Shimamoto and Nagahama 1992). Quartz and feldspar fragments may be more common in the pseudotachylite than in the wall rock, while hydrous minerals such as biotite may be less common. Flow banding defined by colour variations may be seen in the matrix (Plates 10, 11), which may contain spherulites, dendritic crystals, crystallites, and microphenocrysts, and may vary from opaque to colourless and isotropic in thin section. Other microscopic features include optically isotropic rims around clasts, embayments of clasts, sub-microscopic crystals, and amygdales.

Fine grained margins may occur at the contact with the wall rock. Pseudotachylite often occurs in planar veins a few mm to 1 m thick from which veinlets may branch and penetrate the wall rock. Planar-convex lenses of pseudotachylite may occur on either side of the vein. The pseudotachylite may occur in a volume between two parallel planar zones, and surround large fragments of wall rock (e.g. Grocott 1981, Magloughlin 1989). The wall rock fragments can form a jigsaw-textured breccia with pseudotachylite matrix. Joints may form perpendicular to the vein walls. The pseudotachylite veins cross-cut any earlier fabrics in the wall rock, and may contain fragments of previously-formed pseudotachylite, indicating cyclical generation (e.g. Sibson 1980, Passchier et al. 1990). All these features may occur on outcrop to thin section scale.

2.11.2 Origin The origin of pseudotachylite has usually been discussed in terms of the two mutually exclusive hypotheses of crushing (e.g. Wenk 1978) or frictional melting (e.g. Sibson 1977). There is conclusive evidence that many pseudotachylites involved frictional melting during faulting. The intrusive textures of pseudotachylite veins together with their flow banding demonstrate that they formed in a fluid state. The spherulites and dendritic crystals and crystallites formed by cooling from a melt or from devitrification of a glass formed by quenching, by analogy with structures formed in rapidly cooled igneous rocks. The fine-grained margins of the veins can be interpreted as chilled margins, and the joints perpendicular to the veins walls can be interpreted as cooling fractures (e.g. Camacho et al. 1995). Some pseudotachylites have the same bulk chemistry as their hosts, and the pseudotachylite composition may change with that of the local wall rocks through which they pass (e.g. Maddock 1986, Maddock et al. 1987), demonstrating that the melt was locally derived and not transported by an intrusive dyke. In other pseudotachylites, there is a consistent difference in the bulk chemistry of the pseudotachylite and wall rock which can be interpreted by fusion of the lowest melting point fraction of the rock. Preferential melting of the hydrous mafic components of the host rock can generate a pseudotachylite matrix which is more mafic than the host rock, and explains why hydrous and mafic minerals are less common in the matrix than the host rock (e.g. Magloughlin 1989, Maddock 1992, Camacho et al. 1995). The localization of pseudotachylite on and near planar surfaces suggests that it is generated by sliding (hence these surfaces are called generation planes). The rounding of the fragments may be due to thermal spalling. A link between slickensides, slickenlines and pseudotachylite has been suggested by Spray (1989) from observing mechanical excavation of sandstone, which created thin layers of melt with shiny and striated surfaces. Pseudotachylite has also been successfully created in frictional sliding experiments (e.g. Spray 1987, 1995). Filaments of frozen melt drawn out across fractures “like warm mozarella” are a delicate feature of experimentally-produced pseudotachylite which has not so far been observed in tectonic pseudotachylite, but could be a useful diagnostic fea-

23


ture. The experiments resolve the controversy about the origin of pseudotachylites: both crushing and frictional melting are likely to be involved, depending on the strain rate (Spray 1995). Crushing occurs in experiments at strain rates of and melting at strain rates from to The passage of seismic or shock waves, with a rapid increase in strain rate, may involve three sequential processes: initial fracture, comminution, and melting when heat can not be dissipated fast enough. Together with results from thermal modelling, experiments show that pseudotachylites should only form in dry rocks at relatively high temperatures, because the presence of water decreases the effective stress (Section 9.2.3) on the fault plane so that insufficient heat is generated to melt the rock (e.g. Spray 1987, 1988). On the other hand, Magloughlin (1989) has argued that pseudotachylites were generated from a hydrous cataclasite with a low melting temperature. The permeability of the wall rocks is an important factor in determining whether excess fluid pressures can build up during faulting and suppress melting (Mase and Smith 1985). Experiments confirm the widely held view that pseudotachylites can form during seismic slip at strain rates greater than Some pseudotachylites are associated with features such as craters, shatter cones and shock metamorphism which show that they have been produced during meteorite impacts, for example the Vredefort Dome in South Africa, from where Shand (1916) first coined the term pseudotachylite, and

the Sudbury structure (Spray and Thompson 1994). Other pseudotachylites have been described from the bases of large landslides (e.g. Masch et al. 1985).

2.11.3 Misidentification Veins of tourmaline, chlorite, ultracataclasite and fine-grained igneous dykes can easily be mistaken for pseudotachylite. Passchier and Trouw (1996) suggest some microstructures that allow the distinction of pseudotachylite from other types of veins. The diagnostic microstructures that identify pseudotachylite as a melt in distinction from hydrothermal veins are the quench and devitrification textures. The distinctive chemistry, reflecting the wall rocks, is also diagnostic and distinguishes pseudotachylite from ultracataclasites. Distinction from igneous dykes relies on evidence for faulting associated with the generation of pseudotachylite, and the different chemistry of the dykes. Pseudotachylites should be distinguishable from ultramylonites by their lack of evidence for extensive intracrystalline plasticity. There are nevertheless many cases in which detailed evidence must be sought to prove the origin of a pseudotachylite by frictional melting. It is a relatively uncommon rock type in the field. The requirement of formation in hot anhydrous conditions effectively limits the host rocks to igneous or metamorphic rocks.

Chapter 3

Diffusive Mass Transfer by Solution 3.1 Introduction Microstructures that show evidence for material removal, transport and deposition without fracturing or lattice distortion form by diffusive mass transfer (DMT). It is often difficult to identify positively the nature of the diffusive flux, and the evidence for the role of a solution may be indirect. Nevertheless, it is a safe generalization that up to amphibolite facies, by far the majority of DMT occurs via solution because solid-state diffusion occurs very slowly at these temperatures (e.g. Rutter 1983).

3.2 Fundamental deformation mechanisms of diffusive mass transfer by solution Diffusion of material occurs in response to gradients in chemical potential, as summarized by Fick’s law which states that the flux of the material (J) is proportional to the chemical potential gradient

B is the particle velocity per unit potential gradient and c is the concentration. The chemical potential of a component is given by:

where U is the molar internal energy, T the temperature, S the molar entropy, is the normal stress, is the pore fluid pressure, and is the molar volume (e.g. Green 1980). Therefore variations in normal stress and pore fluid pressure can establish chemical potential gradients necessary for DMT to occur; internal strain energy may also play a role (Wintsch and Dunning 1985, Bell and Cluff 1989). The diffusion will be from sites of high to low normal stress; a gradient in normal stress is essential to cause material flux (Raj 1982). Hence deformation accommodated by DMT through an intergranular solute is often called pressure solution, despite the fact that solubility itself is not significant to the thermodynamics of DMT. However, recent experiments that demonstrate a crystallographic dependence of solubility may pose some problems for current pressure solution models based on this thermodynamic approach (Becker 1995, Den Brock 1996).

The exact geometry of the diffusion path is still unclear. Rutter (1976, 1983) has suggested that a water film on the order of nm thick must be present along grain boundaries to allow DMT to occur. This concept is supported by observations of arrays of fluid inclusions along grain boundaries in naturally deformed halite, interpreted as evidence of a syntectonic continuous film of fluid along grain boundaries that subsequently healed into discrete fluid inclusions (Urai et al. 1986), and by observations of grain boundary pores which probably formed in equilibrium with a fluid in quartz (Hippert 1994). Alternatively, islands of solid material may support normal stresses and allow diffusion to occur in a solution within intervening channels (Raj and Chyung 1981, Raj 1982, Spiers and Schutjens 1990). It may be necessary for different phases to be adjacent to each other for a continuous fluid film to be present (Hickman and Evans 1991). The undercutting mechanism proposed for pressure solution is completely different and does not require diffusion in a solute film between surfaces under high normal stress gradients (e.g. Bathurst 1958, Tada and Siever 1986). In this mechanism, grain contact deformation occurs by intracrystalline plasticity or cataclasis, while dissolution of free surfaces around the contacts maintains small contact areas and therefore the high stresses necessary for the solid-state grain deformation. The relative importance of solute film diffusion versus undercutting depends on temperature, grain size and free grain surface area. Calculations by Tada et al. (1987) suggest that undercutting will provide faster strain rates under most circumstances, but that solute film diffusion will become important for very small free grain surface areas, as in the final stages of compaction, and in dense aggregates. The distribution of fluid in a rock, and therefore the effectiveness of pressure solution, is dependent on surface energy (measured by the dihedral angle, ) and deformation. Fluids with greater than 60° will occur as isolated pores at equilibrium, as interconnected channels at triple grain junctions for and along all grain surfaces for (e.g. Watson and Brennan 1987). The first two cases clearly constrain the potential for pressure solution because the diffusion pathways are limited, and probably apply to fluids in most static geological situations. However, experiments and theory show that fluids can wet grain boundaries completely under deformation, leading to enhancement of diffusion creep by orders of magnitude (e.g. Urai 1983, Copper et al. 1989, Heidug 1991, Tullis et al. 1996). Moreover, dihedral angles may be a function of pressure and temperature, as established

24

CHAPTER 3. DIFFUSIVE MASS TRANSFER BY SOLUTION

25

by experiments on brines in halite, which show that the dihedral angle reduces below the 60° threshold at elevated pressures and temperatures (Lewis and Holness 1996). Unfortunately experiments also suggest that evidence for fluids along grain boundaries may be rapidly removed after deformation ceases. DMT is often inferred from microstructures that show a systematic relation to surfaces which may have been under higher normal stress (e.g. surfaces parallel to a shape fabric). However, it is the principal strains which are apparent from most microstructures, and thus there is always an ambiguity about interpreting DMT from the such microstructures (cf. Groshong 1988). The above inference relies on an assumption that the principal stress and strain axes are parallel, which will be true for coaxial deformation. In this chapter, finite shortening and extension directions will be emphasized because they can be inferred from the microstructures. Any links made between DMT microstructures and principal stress orientations or gradients in normal stress must involve some assumptions about the nature of the deformation, such as coaxiality.

3.3 Grain surface solution textures Solution creates distinctive grain surface textures which are clearly observed in the scanning electron microscope. The typical texture consists of pits on the grain surface, giving a corroded appearance. Irregular pores with channel and cave morphologies 1 to long and deep have been described in a naturally deformed micaceous quartzite (Hippert 1994); similar channel morphologies have been produced during experimental deformation of quartz by DMT (Den Brock and Spiers 1991). Atomic-scale details of the solution process can be seen using atomic force microscopy (e g. Gratz et al. 1991).This technique shows that calcite dissolution and precipitation occurs in layers, and that solution generates crystallographically-controlled etch pinholes less than 5 nm deep, and rhombic etch cores more than 90 nm deep (Hillner et al. 1992). Quartz solution occurs along nm-sized pits and ledges parallel to the and directions (Gratz et al. 1991).

confidence in sediments with rounded grain shapes that allow the pre-solution grain shape to be accurately delineated. Truncated contacts may appear superficially like flat contacts formed by mutual concordance between overgrowths (Section 3.9), but can be distinguished by the evidence for material removal.

3.4 Indenting, truncating and interpenetrating grain contacts

The formation of indenting or truncating grain contacts can be understood from theoretical and photoelastic treatments of the stress distribution in contacting grains. The values of the principal stresses, and the mean and differential stresses, increase with the radius of curvature, suggesting that DMT will preferentially occur in grains with larger radii of curvature, creating indenting grain contacts (e.g. McEwen 1981). Truncating grain contacts are expected from approximately equal rates of DMT between two grains of equal curvature.

Grain contacts affected by DMT are distinctive. Where two grains of similar composition and orientation but with different shapes are in contact, the grain with the smaller radius of curvature typically penetrates into the grain with the larger radius of curvature (Figs. 3. la, 3.2), creating an indenting grain contact. A flat or slightly curved contact is observed between similar grains with approximately equal radii of curvature. Such truncating contacts are common on longer edges of grains which are parallel to any shape fabric (Fig. 3.3). Material removal is demonstrated by truncation of the original grain shape (Plate 12). The amount of material removed can be estimated by reconstructing the missing grain boundaries (Fig. 3.1b; Onasch 1994). This can be done with reasonable

Interpenetrating grain contacts are mutually interlocking protrusions of grains into each other, which have a similar morphology to microstylolites (Section 3.6), and indicate material removal along the contacts (Figs. 3.1c, 3.4). Sutured grain contacts may look superficially like interpenetrating grain contacts, but these are formed by intracrystalline plasticity (Section 4.8). Sutured contacts can usually be distinguished from interpenetrating contacts by the presence of subgrains, which may form promontories along a sutured contact that are slightly misoriented with respect to the lattice of the rest of the grain (Fig. 3.1d). By contrast, interpenetrating contacts formed by DMT usually occur between grains with uniform lattice orientations.

26

CHAPTER3. DIFFUSIVE MASS TRANSFER BY SOLUTION

27


3.5 Strain caps Strain Caps are strongly foliated domains enriched in micas or less soluble minerals around opposite surfaces of relatively rigid objects (Passchier and Trouw 1996). The foliation trace in the strain cap is concordant with the margins of the object, and grades into continuity with the bulk rock foliation with increasing distance from the object, so that it appears to be deflected around the object (Plate 13). The concentration of less- soluble minerals such as micas in strain caps suggests that they form by removal of the more soluble matrix around the rigid object, which can be predicted from the thermodynamic considerations above, since the object will concentrate normal stress on some surfaces due to its greater rigidity.

3.6 Microstylolites 3.6.1

Characteristics

Microstylolites are microscale discontinuities with offset or wave-like shapes that may truncate grains or other markers such as bedding planes or fossils (Fig. 3.5, Plate 14). Microstylolites are readily observed under the microscope, and have most features in common with mesoscale stylolites. Individual protrusions on a microstylolite are called teeth or columns, which may have flat crowns at the top and steep walls along their sides (Fig. 3.5). The offset or wavelike geometry is observed in all sections, showing that the three dimensional shape is a forest of columns and matching pits approximately perpendicular to the plane of the mi-

crostylolite. Microslickolites are distinguished from microstylolites by having teeth that are oblique to the plane. The oblique teeth create a lineation on the microslickolite surface. Microstylolites may follow grain boundaries, especially of insoluble grains. Microstylolites can be conspicuous because of coatings or fillings of iron oxides, hydroxides, phyllosilicates (which may have a strong grain shape fabric) or organic matter, which distinguish them from the wall rock. The thickness of the filling is commonly highly variable. The morphological classification of cross-sections through stylolites by Guzetta (1984) into sharp peak, rectangular, wave, smooth and composite types is useful for microstylolites (Fig. 3.6a, Plate 15). Parameters for describing microstylolites quantitavely are shown in Fig. 3.6b (from Andrews and Railsback 1997).

3.6.2 Formation and propagation The geometrical similarities between stylolites and microstylolites suggests that they have the same genesis. The discussion in this section refers to sylolites, because most previous work has focussed on the mesoscale features. Truncation of markers is direct evidence that material has been removed along a stylolitic surface. The minimum thickness of material removed along a stylolite is approximately twice the amplitude, or the height of the teeth (Tada and Siever 1989). The walls of the teeth are parallel to the direction of movement between the opposite sides of the stylolite. Railsback and Andrews (1995) have shown that the orientation of teeth is more constant than the orientation of the plane

28

of the stylolite, and is useful for kinematic analysis because it shows the direction of maximum finite shortening. The obliquity of the teeth of slickolites shows that there has been a component of shortening parallel to the slickolite surface (Fig. 3.6c). The coatings or fillings on stylolite surfaces are usually interpreted as insoluble residues that did not diffuse and were concentrated at the sites of solution. The relative insolubility of these coatings is good evidence that diffusion was via solution. This is further supported by a correlation between the thickness of the coating and the concentration of this material in the adjacent layers of rock (e.g. Borradaile et al. 1982), and by the influence of heterogeneities on the stylolite trace, which show that solution was concentrated in areas of higher inferred normal stress, and was impeded by insoluble material. The concentration of insoluble material on a stylolite surface compared to its concentration in the host rock can be used to give an estimate of the amount of shortening (Railsback and Andrews 1995). However, in some cases the stylolite filling material is not found in the host rock, and is therefore the product of metamorphic reactions (Beach 1979). Reaction products can also mark grain boundaries to indicate where fluid transport has occurred (McCaig 1987). A stylolite can be regarded as an anticrack, or an ellipsoidal volume removed from the rock (Fletcher and Pollard 1981). If the resulting hole is closed up by elastic deformation, stresses are induced around the anticrack tip, which cause the stylolite to propagate in its own plane. Therefore a stress concentration under the appropriate conditions is suf-

CHAPTER. 3 DIFFUSIVE MASS TRANSFER BY SOLUTION

ficient to cause a stylolite to develop and propagate under its own stresses. Anticrack tip stresses may be dissipated in a zone at the stylolite tip which can be referred to as a process zone, analogous to the non-linear zone ahead of a microcrack tip (Section 2.2.1.1). Carrio-Schaffhauser et al. (1992) found evidence for a process zone in the form of enhanced porosity at stylolite tips in limestones, and suggested that material removed from the process zone was deposited in a zone of reduced porosity adjacent to the stylolite. Stylolites may form on previous fracture surfaces (e.g. Petit and Matthauer 1995, Railsback and Andrews 1995). Stylolites and pressure solution features may be localized along authigenic clay layers and around pre-existing phyllosilicates. The proposal that contacts between different phases are necessary for continuous fluid films (Hickman and Evans 1991) suggests a good reason for enhanced DMT around phyllosilicates, and may also apply to the formation of strain caps.

3.7 Diffusive mass transfer and cleavage 3.7.1

Classification

Diffusive mass transfer via solution is an important and even dominant deformation mechanism in cleavage formation. Unfortunately there has been a strong tendency to invoke genetic nomenclature for cleavage formation (e.g. solution

29


the rock, forming a zonal cleavage. Another important aspect of description is the transition from cleavage domains to microlithons, which may be sharp or gradational. Many cleavages referred to as solution cleavages belong in this category. Disjunctive cleavage domains are often compositionally distinct from microlithons, having a higher proportion of opaque minerals and phyllosilicates. Zonal cleavages with marked compositional contrasts between cleavage domains and microlithons constitute a new compositional layering, which has been called solution striping (e.g. Williams 1972). Truncation of markers in cleavage domains suggests that material loss by DMT has occurred (Fig. 3.8), although shear displacements on fractures may create a similar effect. Material removal can be distinguished from the latter effect in some cases by inconsistent offsets of a folded layer on several cleavage surfaces, and large variations in thickness of a marker in adjacent microlithons (Fig. 3.8). The amount of material removal can be estimated from truncation of markers and other effects such as imbrication of chert pebbles and offset of bedding. The volume loss may vary from a few percent for a weak cleavage to over 35% for very strong cleavage in cherts (e.g. Alvarez et al. 1976), and 30-40% for disjunctive cleavage in coarse siltstones and sandstones (Murphy 1990). Material transfer by diffusion can also be inferred from the compositional contrasts between the cleavage domains and the microlithons. A solution can be inferred as the diffusing phase because of the relatively low grades of cleavage formation. Crenulation cleavage

cleavage), which hinders discussion about deformation mechanisms, but Powell (1979) and Borradaile et al. (1982) give an excellent alternative, non-genetic classification, which is followed here. The main subdivision is made between spaced cleavage, consisting of cleavage surfaces that are separated by tabular bodies of rock called microlithons, and continuous cleavage, in which cleavage is penetrative throughout the whole body of the rock. Since most cleavages are concentrated into domains at some scale, this distinction is evidently scale-dependent. Powell (1979) suggests that any cleavage with domains spaced less than 0.01 mm (i.e. at the limit of optical microscopic resolution) should be called continuous at a microscopic scale.

3.7.2

Spaced cleavages

Disjunctive cleavage

Disjunctive cleavage is a spaced cleavage in which the cleavage domains cut across the previous fabric of the rock (Fig. 3.7). The cleavage surfaces may be smooth, rough (discontinuous) or wave-like, and in relation to each other they may be subparallel, anastomosing, or trapezoidal/conjugate. Their spacing may be controlled by the rock type. The width of the cleavage domains may be a substantial proportion of

This is a spaced cleavage in which microlithons contain an earlier fabric that is systematically related to the fabric of the cleavage domains. The cleavage domains are usually axial planar to folds (crenulations) in the earlier fabric (Plates 16, 17). The crenulations may be symmetric or asymmetric, and rounded or angular, and the transition from cleavage to microlithon may be sharp or gradational. The crenulation wavelength may be determined by the thickness of folded layers as predicted by buckling theory (cf. Price and Cosgrove 1990), or the grain size may dictate a minimum wavelength. The cleavage domains are commonly localized along the limbs of the folds, especially those associated with asymmetrical folds, in which the cleavage domains are usually along only one of the two unequal limbs. Most crenulation cleavages are zonal because the cleavage domains are compositionally quite distinct from the microlithons. Phyllosilicates and opaques are concentrated in cleavage domains and quartz, calcite and feldspar are concentrated in hinges. While the essential role of buckling in crenulation cleavage formation is evident, the importance of DMT is clear from the compositional zoning. Cleavage domains in typical crenulation cleavages are depleted in Si, Ca, Na, P, Mg, Fe, and F, and enriched in Al, K, Rb, Y, Ce, Sc, and Ba, and there is a net volume loss from the cleavage domains that is approximately balanced by volume gain in the microlithons (Manktelow 1994). The zoning can be explained in a simplistic way by the normal-stress dependence of diffusion: crenulation hinges, where the primary foliation lies at a high angle to the maximum principal stress, may be sites of low

30


normal stress. It is also possible that anisotropy in diffusion pathways may play an important role in the differentiation during crenulation (Marlow and Etheridge 1977). Other lines of evidence that can be used to show the importance of DMT are the narrow widths of quartz and feldspar grains in cleavage domains relative to microlithons, truncation of grains, the association of crenulation cleavage with overgrowths and mica beards, and the lack of other deformation mechanisms (Gray 1977, Borradaile et al. 1992).

3.7.3

Continuous cleavage

The most characteristic continuous cleavage is slaty cleavage, defined as a preferred orientation of phyllosilicates too fine to be seen with the unaided eye (Fig. 3.9). Slaty cleavage commonly has cleavage domains with spacings of 0.1 mm or less, which are easily visible under the optical microscope or the SEM. Mechanical rotation of grains (cataclasis), kinking (intracrystalline plasticity) as well as DMT are involved in the formation of slaty cleavage (e.g. Wood 1974). Electron microscope studies show that phyllosilicates and quartz are mobilized by DMT in the development of cleavage domains (Knipe and White 1977,1979, White and Knipe 1978, Knipe 1979). Structural studies using a variety of strain markers including reduction spots, trace fossils and grain boundaries suggest that slaty cleavage formation is accompanied by volume loss (e.g. Ramsay and Wood 1973: 20%, Wright and Platt 1982: 50%, Wright and Henderson 1992: 40-60%),

which must have occurred by DMT via solution at these metamorphic grades. However, bulk volume loss is not permitted by several geochemical studies on the same rocks, as discussed below (e.g. Erslev and Ward 1994). Slaty cleavage domains are dominated by phyllosilicates, resulting in increased K, Al, Ti, Ba, and U, Na, Ca, Mn, Mg, Fe, and Si are often depleted due to the loss of albite, carbonate, quartz and feldspar (Borradaile et al. 1982). The depleted elements may be redistributed by DMT into microlithons, suggesting that there is little bulk volume change in the rock (e.g. Groshong 1976). Large-scale volume fluxes are also unlikely at low grades because average shale and slate compositions are very similar (Erslev and Ward 1994). The reasons for the contradiction between this conclusion and the large volume losses suggested by the structural studies is presently unclear. Whether or not large scale volume changes accompany cleavage formation, the pronounced local compositional differentiation is clear evidence that DMT processes are essential to slaty cleavage development.

3.8 Grain surface deposition textures Crystal growth from solution occurs by two different mechanisms with different microstructures (Bennema and van der Eerden 1987). Atomically flat crystal faces first nucleate steps to which atoms can attach, and growth then occurs by accretion along the steps in crystallographically controlled planes. Dislocations (Section 4.2) may provide important


31

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nucleation sites for both growth and solution on crystal surfaces (e.g. Casey 1995). Above a critical temperature known as the roughening transition, growth rates are isotropic, and sub-spherical crystals may be produced. Deposition by the first mechanism can be recognized from the euhedral nature of grain surfaces seen under the microscope; the second may be responsible for botryoidal textures.

3.9 Overgrowths, porosity reduction, pressure shadows and fringes, and mica beards 3.9.1

Characteristics

Overgrowths are mineral deposits surrounding a grain in a rim. The rim and grain are usually the same mineral. The rim is commonly in crystallographic continuity with the grain, and may therefore be difficult to detect. The overgrowth may, however, be separated from the original grain by a zone of inclusions (Fig. 3.10), and CL can distinguish subtle features of overgrowths (Plate 18). A distinctive microstructure formed by overgrowths in well cemented rocks are relatively straight grain boundaries and 120° triple junctions between three grains (Fig. 3.11a). This texture arises from mutual impingement of overgrowths growing at equal rates from adjacent grains. It can easily be confused with granoblastic polygonal textures (Section 5.4) unless the overgrowth can be distinguished from the grain.

Porosity reduction may occur by mechanical compaction, solution and precipitation: the latter two are DMT processes, and they can be recognized by the loss of material at strain markers and the precipitation of cements in pores in the form of overgrowths (e.g. Carrio-Schaffhauser and Gaviglio 1990). Crystal growth in pore spaces typically results in a convoluted pore-grain interface which is a fractal curve. Models suggest that the fractal dimension of the curve increases with the ratio between precipitation and dissolution rates, and also increases from about 2.5 to about 2.75 in natural sandstones as porosity reduces during diagenesis (Aharonov et al. 1997). Pressure shadows and fringes are domains of secondary mineral growth adjacent to grains (Plate 19). Pressure shadows lack distinctive internal crystal forms, while pressure fringes have fibrous mineral fillings. Mica beards are a distinctive type of pressure fringe, consisting of fibrous mica defining a good shape fabric. Passchier and Trouw (1996) prefer to use strain instead of “pressure” on the basis that it is nongenetic; however, they point out that the term strain shadow is also potentially misleading as it incorrectly implies that strain is low in the shadow zone. Analysis of steady-state pure shear around a rigid sphere shows that there are two volumes of low pressure at the ends of the object perpendicular to the maximum applied stress, but that differential stress and strain are high in these areas (Masuda and Mizuno 1995). With this insight, and since pressure shadow is so firmly entrenched in the literature, it is probably the better choice of terminology, despite its genetic connotations. The mineralogy in a pressure shadow or fringe may be the same as the grain, the same as the matrix, or different from

33


low for fibrous microveins (e.g. Durney and Ramsay 1973, Elliot 1973, Ramsay and Huber 1983, Passchier and Trouw 1996). The use of fibres as shear sense indicators is discussed in Section 7.7.

3.10 Grain shape fabrics Grain shape fabrics are a common DMT microstructure. They may form by material removal in the shortening direction (e.g. indenting, truncating or interpenetrating grain boundaries), and by material addition (e.g. overgrowths) in the extension direction (Fig. 3.11b). Grain shape fabrics have been produced experimentally by DMT (e.g. Schutjens 1991). Lack of microfractures distinguish DMT grain shape fabrics from cataclastic grain shape fabrics, and the lack of undulatory extinction, sub-grains or recrystallized grains (Chapter 4) distinguishes them from intracrystalline plastic grain shape fabrics.

3.11 Fluid inclusion planes

both. Crystallographic continuity may be maintained with the grain, host, or neither. Fibres may be perpendicular to the grain boundary or oblique. They may be found on more than two sides of polyhedral grains. Fibres may be straight, curved, deformed or undeformed. All of these basic aspects of pressure shadow and fringe description are important in interpreting the mechanism of formation.

3.9.2

Mechanisms

Pressure shadows and fringes are considered to form by DMT because they are a new mineral growth under low grade conditions. Diffusion towards sites of low mean stress is predicted by the thermodynamic approach of Section 3.2, which explains why pressure shadows and fringes form at the ends of grains perpendicular to the inferred maximum principal stress. However, the geometry of the pressure shadows and fringes is controlled by a number of other factors including strain, grain shape, whether growth occurs at the grain interface (antitaxial) or the matrix interface (syntaxial), whether the orientation of fibres is controlled by the grain surface (face controlled) or the incremental extension direction (displacement controlled), and whether the shadows or fringes are deformed or not. Some of these factors are illustrated for a coaxial deformation in Fig. 3.12. Given due consideration of these factors, fibres can be used to deduce the extensional strain, the deformation path, and the type of flow, as discussed be-

Fluid inclusions consist of nm to mm sized cavities filled by fluids, which may also contain solid material. They occur as isolated inclusions, in clusters, and in planes. They are common in thin section, where they may appear as dark inclusions which reveal their fluid contents on examination at higher magnifications. The most common fluids are aqueous, saline or with possible admixtures of sulphur compounds, or more complex hydrocarbons. An important distinction is made between primary fluid inclusions that form during growth of the original minerals of a rock, and secondary inclusions that form subsequently (e.g. Roedder 1984). Primary inclusions can be recognized because they are disposed on euhedral crystal forms, whereas secondary inclusions cut across crystal growth features. Fluid inclusion planes (FIPs) are planes of fluid inclusions that often have a strong preferred orientation on a microscopic scale, and may also have regionally consistent orientations (Fig. 3.13, Fig. 2.5). The fluids in any set of FIPs are generally compositionally homogeneous, and may be distinct from fluids in other sets of FIPs that occur in the same rock. FIPs form by trapping of a fluid during precipitation of microcrack fillings. The microcracks are usually extension microcracks formed by the mechanisms described in Section 2.3. Healing of microcracks in quartz can occur as rapidly as micrometres/day (Smith and Evans 1984). The healing rate depends on temperature, concentration of the fluid, and microcrack dimensions (e.g. Brantley 1992). Healing leaves a plane of cylindrical or spherical fluid inclusions along the former microcrack, firstly by forming cylindrical tubes of fluid parallel to the microcrack tip (necking down), and pinching off or ovulation to form approximately spherical or negative crystal shapes (Fig. 3.14). Isolation of spheres occurs because grain boundary migration rates depend on the thickness of the fluid phase (the microcrack can be considered as a type of grain boundary): slower migration rates occur in thicker fluid films. A local thickening of the microcrack will slow

34



35

the grain boundary migration rate at that point, isolating a fluid inclusion behind the rest of the more rapidly migrating boundary (e.g. Urai 1983). FIPs are primary evidence for the importance of fluids and DMT during deformation. They form perpendicular to and can be very useful in kinematic analysis, because they can constrain both the orientation of and the fluid pressure (e.g. Lespinasse and Cathelineau 1995). They can also give evidence for fluctuating fluid pressures associated with stress cycling and therefore probably with earthquake faulting (e.g. Robert et al. 1995).

3.12

Microveins

Microveins are microscopic tabular zones of secondary mineral growth. The rich variety of microvein textures observable under the microscope can allow very detailed interpretations of their formation. CL is valuable for analysis of microveins, because variations in fluid chemistry and temperature during microvein filling can be reflected in the luminescence, delineating delicate growth features (e.g. Dietrich and Grant 1986, Urai et al. 1991). One of the most important features of a microvein is the orientation of the opening vector, the vector which connects points that were originally joined before microvein opening. The opening vector may be perpendicular to the wall (an extension microvein) or have components of displacement both perpendicular and parallel to the walls (a shear microvein). The opening vector can be determined from displaced markers, or from some features of microvein fillings, as described below. The opening vector can be considered for the net opening history of the microvein (the cumulative opening vector) or for individual increments of opening (the incremental opening vector). Common microvein fillings are carbonates, quartz, chlorite and epidote. The texture of the filling may be massive, equant (blocky; Plate 20), fibrous, laminated, euhedral (idiomorphic), or botryoidal. Filling textures are a function of nucleation and growth kinetics, the rate of microvein opening, the geometry of the opening vector and the geometry of the walls. Euhedral crystal terminations and botryoidal textures are diagnostic of growth into open space, and are useful in ore paragenetic studies for discriminating epithermal environments. Euhedral crystal faces can be recognized by growth zoning in the crystal, which may be defined by fluid or solid inclusions. Wilson (1984) suggests that considerable variation in crystal

orientation and grain size is characteristic of such “free-face” growth. In addition to the major filling phases, many microveins also contain inclusions of the same mineralogy as the wall rock. Lines of inclusions parallel to the microvein margins are known as inclusion bands, while those at higher angles to the margins are inclusion trails (Figs. 3.15, 3.16). Inclusion bands and trails may form by two mechanisms: overgrowth of wallrock fragments, and fracturing of the wall rock followed by incorporation of fragments into the filling. The latter mechanism may occur where irregularities on microvein walls obstruct opening, and must be broken off for opening to occur (e.g. Urai et al. 1991). The presence of numerous inclusion bands is evidence for a cyclic process of microcrack opening followed by filling, a process known as crack-seal (Ramsay 1980). Each inclusion band represents one cycle. The width between inclusion bands is for many rock types, and there may be up to thousands of bands in a microvein (Ramsay and Huber 1987). Inclusion trails are markers of the position of particular points on the microvein margin at successive opening positions: they are therefore parallel to the opening vector (Fig. 3.15b). Paradoxically, microstylolites sub-parallel to microvein margins have also been described, particularly along inclusion bands (e.g. Cox 1987). The shortening demonstrated by such microstylolites could be part of the crack-seal cycle if fluid pressures decreased to less than lithostatic in part of the cycle, causing the microvein to close and experience compressional stress. These microstylolites may also be due to a later deformation, unrelated to the microvein formation. Laminated microveins have planar bands, often composed of phyllosilicates, subparallel to the margins. The bands

36



may form as inclusion trails where the opening vector has a large component parallel to the margin (Cox 1987). In this situation there is no distinction between inclusion trails and bands. Fibrous microveins are particularly rewarding for kinematic interpretation. Fibrous fillings are common and form by progressive microvein opening at a rate that can be matched by crystallization of the filling. Fibres can grow by at least five different filling sequences: these need to be carefully established before kinematic interpretations can be made. Syntaxial growth means growth from the wall rock towards the microvein centre on both sides of the fracture (Fig. 3.17a, Plate 21). The diagnostic features of syntaxial growth are two separate bands of fibres on either side of a central suture. The fibres are not continuous across the suture. The microvein fill is similar to the wall rock and may be in crystallographic continuity with it. Fibre widths generally increase in the direction of growth due to competition between fibres, which ensures that the faster growing and therefore larger fibres overgrow and eventually isolate the slower and smaller fibres from the precipitating solution (e.g. Smith 1964). Therefore in syntaxial growth, fibre widths may increase towards the suture. By contrast, antitaxial growth occurs from the microvein towards the wallrock, and may occur symmetrically on both sides of the microvein (Fig. 3.17b) or asymmetrically on only one side (Fig. 3.17c, 3.18). The diagnostic feature of antitaxial growth is fibre continuity across the microvein. Antitaxial growth is characterized by fillings of different material from the wall rock, and no crystallographic relationship between the filling and wall rock. Sym-

37

metric antitaxial growth may result in a median line of inclusions along the centre of the microvein, unlike asymmetric growth. Fibre widths can also be used to distinguish symmetric growth (fibre widths increase symmetrically in two opposite directions towards the wall rock) from asymmetric growth, in which the widths increase unidirectionally across the microvein. Composite growth means both syntaxial and antitaxial protions in the same microvein (Fig. 3.17d). Nonsystematic growth (“ataxial” - Passchier and Trouw 1996) histories may involve fracture at any point in the crystal fibres, which are termed stretched crystal fibres by Durney and Ramsay (1973). They exhibit none of the systematic features described for the other four categories of growth above (Fig. 3.17e), and the sides of the fibres have distinctive interlocking teeth, dividing then fibres into tablets (Plate 22). The lack of directional growth indicators (e.g. fibre widening direction) in these microveins is diagnostic. In many fibrous microveins, the fibres grow parallel to the incremental opening vector: Such tracking or displacementcontrolled fibres can be used to deduce incremental strain histories with great effect. They can be recognized because fibres connect markers across the microvein, because they are parallel to inclusion trails (Urai et al. 1991), or because they have a constant orientation between microvein walls of variable shape (Plate 23). The strain history deduced from the fibres can be plotted on a diagram showing rotation of the incremental strain on the horizontal axis against strain on the vertical axis (a cumulative incremental strain history or cish diagram, e.g. Fisher and Anastasio 1994, Hedlund et al. 1994).

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Non-tracking fibres do not track the incremental opening vector (e.g. Durney and Ramsay 1973). Non-tracking fibres can be recognized because they do not connect markers across the microvein (Fig. 3.19). However, markers may not be connected even by tracking fibres if a microvein has an early history of shear displacement before filling


occurred. Inclusion trails may be used to distinguish tracking and non-tracking fibres in this case. Non-tracking growth occurs because fibre growth directions are determined by the orientation of the growth surface, like face-controlled pressure fringes. Fibre boundaries are either perpendicular to the growth surface (Fig. 3.19), or along the bisector of two adjacent growth surfaces. The fibre boundaries have no fixed relationship to the opening vector. The tracking efficiency, or degree of match between the opening vector and the fibre, is determined by the angle between the opening vector and the growth surface and the shape of the growth surface (Fig. 3.19, Urai et al. 1991). These fibres can only be parallel to the opening vector when when the tracking efficiency is 1. Fibres in microveins with any of the above growth histories are often curved. The curvature can be primary (formed during crystal growth), or secondary (due to subsequent deformation). Primary curvatures can form in both tracking and non-tracking fibres due to rotation of the incremental opening direction with respect to the previously formed part of the fibre, and can be recognized because the curved fibres show no evidence of strain or recovery. All low-temperature microvein fillings form by precipitation from solution, and are therefore excellent evidence for DMT via solution. Microveins are a conspicuous feature of greenschist-facies and lower grade deformation because both cataclasis (fracture) and DMT via solution are required for their formation. At higher grades, other deformation mechanisms operate, and microvein textures have low preservation potential because of recrystallization.

Chapter 4

Intracrystalline Plasticity 4.1 Introduction Permanent distortion of a crystal lattice without fracture occurs by intracrystalline plasticity. The definitive feature of all intracrystalline plastic deformation mechanisms is the involvement of dislocation motion (Section 4.2), but solid state diffusion (Chapter 5) is an integral part of some mechanisms considered in this chapter. Intracrystalline plasticity has a long history of research and a vast literature, particularly in the materials science field, which is necessarily rather condensed in this chapter. Useful references that provide more explanation and detail are Hull (1975) and Nicolas and Poirier (1976).

4.2 Fundamental mechanisms of intracrystalline plasticity A dislocation is one type of imperfection or defect in a crystal lattice. A useful way to understand dislocations is to imagine cutting into the lattice, stretching the lattice apart along the cut, and inserting an extra plane of atoms into the cut. The dislocation is the line along the edge of the extra plane of atoms (Fig. 4.1). The stretching of the lattice to accommodate the extra atoms causes a distortion, the orientation and size of which is measured by the Burgers vector which can be perpendicular to the dislocation line (an edge dislocation), parallel to the line (a screw dislocation) or oblique (a mixed dislocation) (Fig. 4.1). If has a magnitude of one lattice unit, the dislocation is described as perfect, but partial dislocations have with magnitudes of fractions of lattice units. The lattice distortion causes a stress field around the dislocation, and is one way of storing strain energy in the crystal. A direct image of dislocations can be obtained in the transmission electron microscope (TEM) where they usually appear as dark lines (Fig. 4.2). They can also be revealed by the technique of etching, in which a flat crystal surface is exposed to acid, creating a small depression (etch pit) at the intersection of a dislocation with the surface of the crystal. The etching occurs due to dissolution which is enhanced by the lattice distortion and strain energy of the dislocation. The density of dislocations can be used to measure past stress fields (Section 9.8.4). Movement of a dislocation is accomplished by breaking bonds in the intact lattice ahead of the dislocation and re-forming the bonds behind the dislocation, causing it to advance through the lattice (Fig. 4.3).

Dislocations move along planes in the crystal called slip planes, in a direction within the plane known as the slip direction; this process is called dislocation glide. The combination of the slip plane and slip direction is known as the slip system, which is specified by the crystallographic orientations of the slip planes and directions. For example, slip along the prism planes of quartz in the < a > direction is annotated by Breaking bonds during glide requires energy, which can be provided by heat, so the stress necessary to cause glide decreases with increasing temperature. Slip systems in some directions are more active than in others because of anisotropy in crystal properties. The relative ease of glide between different slip systems changes with temperature. Glide is impeded by impurities in the crystal lattice, and by the stress fields associated with other dislocations. These obstacles may be overcome by dislocations changing their slip plane. Screw dislocations can accomplish this by a process called cross-slip (Fig. 4.4). Edge dislocations can change their slip plane if a lattice vacancy replaces the last atom in the half-plane of the dislocation (Fig. 4.4). This process of dislocation climb therefore involves diffusion, although it is an intracrystalline plastic deformation mechanism because dislocation motion occurs. Deformation by a combination of dislocation glide and climb is called dislocation creep, and occurs at higher temperatures, and lower strain rates than dislocation glide (Nicolas and Poirier 1976).

4.3 Deformation twins A twin is a region of a crystal that is rotated or reflected with respect to the rest of the crystal (the host). Twins may form during crystal growth or deformation: the latter are known as deformation or mechanical twins. Growth twins are generally straight and of constant thickness, but deformation twins have variable thickness (thinning and branching towards the edge of a crystal), and are commonly bent (Plate 24, Fig. 4.8). Deformation twins are common in carbonates (where their morphology is temperature-dependent; Section 9.9.2), and feldspars. Deformation twins form by shear of the crystal lattice with respect to the host lattice along the twin plane, together with minor rearrangements of the twinned lattice points. The twin plane is a mirror plane comprising an array of partial or twinning dislocations. The strain due to twinning can be measured, and deformation twins can also be used to measure stress and temperature during their formation

39

40

CHAPTER 4. INTRACRYSTALLINE PLASTICITY


(Sections 9.8.5, 9.8.9, 9.9.2).

4.4

Undulatory extinction

Undulatory extinction is visible in cross-polarized light when a single crystal has variable extinction positions (Fig. 4.5). The extinction position may change consistently from one end of a crystal to the other, producing a continuous sweep of extinction as the stage is rotated. Irregular patches of distinct extinction positions can occur, and sectors of different extinction positions may radiate from the centre of the crystal. Undulatory extinction is caused by distortion of the crystal lattice by dislocations of a consistent orientation (Fig. 4.6a). It forms at low strains during intracrystalline plastic deformation, and as such it is a sensitive indicator of intracrystalline plastic deformation.

4.5

Intracrystalline deformation bands, kink bands and subgrains: Recovery

Intracrystalline deformation bands are tabular low-strain domains within crystals, separated from other parts of the crystal along approximately planar boundaries across which there is a slight change in lattice orientation. “Intracrystalline” is used to distinguish these smaller scale features from cataclastic deformation bands (Section 2.5). Kink bands are sim-

41

ilar to intracrystalline deformation bands, but have sharper, more planar boundaries. The change in lattice orientation across an intracrystalline deformation or kink band may be visible from changes in extinction position (Fig. 4.7) or changes in the orientation of twin or exsolution lamellae in feldspars or calcite (Fig. 4.8). Kinks are different from twins because the change in orientation across a kink plane is not a fixed amount, and the kink plane is not a mirror plane. Subgrains are seen under the microscope as areas within grains of slightly different crystallographic orientation, separated by boundaries subparallel to crystal planes. The variation in crystallographic orientation is visible in quartz as slight changes in extinction position which affect discrete areas, in contrast to the progressive change in extinction position characteristic of undulatory extinction. Subgrains in quartz are usually elongate parallel to the prism planes, which form the subgrain boundaries (Fig. 4.9). Tabular subgrains thus have a similar geometry to kink bands. An alternative pattern in quartz consists of approximately square subgrains with boundaries parallel to both prism and basal planes, known as a chessboard pattern (Fig. 4.10). Basal subgrain boundaries are only visible in grains with their c-axes subparallel to the plane of the section, in contrast to prismatic subgrain boundaries. Lattices with high dislocation densities possess a large internal strain energy (Section 4.2), which can be lowered by dislocation movement into surfaces surrounding relatively dislocation-free volumes (Fig. 4.6). This process of recovery results in a microstructure of low-energy volumes (subgrains) surrounded by walls across which there is a slightly different lattice orientation: hence the walls are sometimes called lowangle boundaries. Undulose extinction, deformation and kink bands, and subgrains are a sequence of microstructures that form with progressive strain, by generation and movement of dislocations into low-angle boundaries (Fig. 4.6; White 1976). The change in lattice orientation across the boundaries, and the dislocation density in the boundaries, increases with strain. Subgrain boundaries are orientated approximately perpendicular to the glide direction of the moving dislocations and therefore the orientations of the subgrain boundaries

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45

46



47 walls made up of well-ordered arrays of two to three sets of dislocations (Blenkinsop and Drury 1988, McLaren 1991). The change in orientation across the subgrain walls is less than 2°. The lamellae may have high fluid inclusion densities, and variable dislocation densities which suggest a highly recovered structure (Blenkinsop and Drury 1988). These deformation lamellae probably formed by recovery of dislocation slip bands, which leads to a variable decrease in dislocation density and precipitation of water in fluid inclusions (Drury 1993). Deformation lamellae can be used as a crude paleopiezometer (Section 9.8.6). Planar deformation features (PDFs) are different type of planar feature caused by shock metamorphism (Section 8.4, 8.11).

4.7 Grain shape fabrics and ribbon grains

may change with temperature. Most natural quartz subgrain boundaries formed in the stability field are prismatic, while in the stability field, both prismatic and basal subgrain boundaries (chessboard pattern) form (Kruhl 1996). These observations do not fully agree with experimental results, or with the suggestion that basal subgrains are indicative of water-present deformation (Mainprice et al. 1986), but the observations are robust empirical evidence, allowing the occurrence of chessboard patterns to be used as a geothermobarometer (Section 9.9.4). The size of subgrains can be used as a palaeopiezometer (Section 9.8.3).

One of the most characteristic microstructures formed by intracrystalline plasticity are flattened or elongated grains with a preferred orientation, resulting in a grain shape fabric, or shape preferred orientation (Fig. 4.12). The effect of movement of dislocations through a crystal is to change its shape towards that of the strain ellipsoid. Apart from the distortion of individual grains, intracrystalline plastic grain shape fabrics can also form by coalescence of grains of similar orientation during dynamic recrystallization (Section 4.8; Means and Dhong 1982). Extreme strain can result in monocrystalline grains with very large aspect ratios known as ribbon grains (Fig. 4.13); they may have intracrystalline plastic deformation features such as undulatory extinction or subgrains. Grain shape fabrics can also be produced by both cataclasis and DMT. Intracrystalline plastic grain shape fabrics can be distinguished by their association with other intracrystalline plastic microstructures.

4.6 Deformation lamellae

4.8 New grains, core and mantle structure: Dynamic recrystallization

Deformation lamellae are crystallographically orientated planar features wide (Drury 1993). Deformation lamellae can be seen under the optical microscope using a high magnification and a narrow diaphragm to enhance relief contrasts. In quartz, they have a slight extinction or refractive index contrast with the adjacent host grain (Fig. 4.11), and they commonly have a sub-basal orientation, sub-perpendicular to prismatic deformation bands (e.g. Spang and Van der Lee 1975, Drury 1993). Deformation lamellae are commonly slightly curved, and have a single, consistent orientation within a grain. They are known from both experimentally and naturally deformed rocks, especially in quartz but also in olivine and plagioclase (e.g. Den Brock and Spiers 1991). The nature of deformation lamellae is enigmatic: observations include slip bands, walls of tangled dislocations, twin boundaries and planes of glass (e.g. McLaren et al. 1967, Christie and Ardell 1974, Twiss 1974). However, TEM studies have shown that typical deformation lamellae in quartz consist of elongate subgrains wide with a sub-basal orientation bounded by curved dislocation

New grains, usually equant and strain free, are common around and within larger original grains in moderately or highly strained rocks. The new grains may have a strong crystallographic fabric which is often systematically related to adjacent older grains. The new grains may be concentrated in a mantle that partly or completely surrounds the older grains, described as core-and-mantle structure (Fig. 4.14), and they may have similar sizes and orientations to adjacent subgrains. These features are characteristic of dynamic recrystallization, or recrystallization that occurs syntectonically. The lack of strain in the new grains shows that they are at an advanced stage of recovery. The new grains form by two main mechanisms: subgrain rotation (SGR), and grain boundary migration (GBM). SGR and GBM are structural transformations which do not per se accommodate deformation, and therefore they are not strictly speaking deformation mechanisms (Urai et al. 1986), but merely structural rearrangements. Movement of dislocations into subgrain walls during recovery causes progressive rotation of the subgrains (see

48

previous section) leading to the formation of a new grain (Fig. 4.15). The collection of dislocations into the subgrain walls may occur by subgrain walls sweeping through the crystal, or by movement of the dislocations towards the walls, and some GBM probably accompanies SGR (Urai et al. 1986). There is a continuum between subgrains and new grains formed by this process, so that the definition of a new grain as distinct from a subgrain is somewhat arbitrary. A generally used criterion is that the new grain lattice orientation differs from its host by more than 10°, but other values have been used (e.g. Urai et al. 1986). SGR is readily inferred from the coexistence of new, smaller grains with subgrains of similar size within the older grains (Fig. 4.16). Other diagnostic microstructures for SGR include clusters of new grains with similar orientations, which they inherit from a single parent grain (Urai et al. 1986), and the coincidence of one or more lattice directions in adjacent subgrains. GBM occurs by the movement of a grain boundary from one grain into another, and finally by the closure of the boundary to isolate a new grain in a different lattice orientation from the host (Fig. 4.17). The grain boundary migrates in response to an internal strain energy gradient from a less-deformed grain with a lower dislocation density to a more highly strained one. The final closure of the boundary may be achieved by an intermediate stage consisting of a bridging subgrain boundary that accommodates progressively more misorientation (Means 1981). GBM can be recognized by the presence of



49

50

new grains along grain boundaries, together with the serrated and lobate shapes of grain boundaries (Fig. 4.18, Plate 25). The characteristic shape of these grain boundaries, also called sutured grain boundaries, is caused by the fact that the boundaries migrate away from their centre of curvature (e.g. Hirth and Tullis 1992). A bimodal distribution of grain sizes may develop between the old and new grains (e.g. Lloyd and Freeman 1994). New grains often develop where there is large lattice distortion and changes in lattice orientation, such as deformation bands and kink planes (Fig. 4.5). This occurs because grain boundaries migrate more rapidly at a misorientation of about 5°. Drury et al. (1985) emphasize the importance of these special sites by proposing a third category of dynamic recrystallization: sub-boundary migration, which is defined as subgrain growth in areas with large gradients of strain and orientation. New grains also form in high stress sites around porphyroclasts. The size of recrystallized grains can be used as a paleopiezometer (Section 9.8.2). Dynamic recrystallization in naturally-deformed quartz occurs by either a mixture of 50% SGR and GBM (Fig. 4.19), or by SGR alone (Drury et al. 1985). Both mechanisms may be observed because they occur sequentially and cyclically (Lloyd and Freeman 199la, b, 1994). Three regimes of dynamic recrystallization in quartz, depending on deformation conditions, have been indentified from experiments (Hirth and Tullis 1992). The Hirth and Tullis classification of dy-


namic recrystallization in quartz is given in Table 4.1, and is commonly applied to naturally deformed rocks.

4.9 Crystallographic fabrics A crystallographic fabric, or lattice or crystal preferred orientation (L or CPO) is a concentration of lattice orientations in one or a limited number of directions. Crystallographic fabrics can be recognized quickly under a optical microscope by observing that many grains have the same extinction position or interference colour. A useful way to detect the presence of a crystallographic fabric in more detail is to insert the sensitive tint plate or a quarter-wavelength plate, which distinguish the different optical axes and therefore provides a more precise picture of the lattice orientations than extinction position or interference colour. This is particularly useful for quartz because of its low order interference colours. Detailed measurements of crystallographic orientation are made with the universal stage under the optical microscope, by electron diffraction, channelling, and backscattering techniques under the electron microscope, and by X-ray and neutron diffraction. Crystallographic fabrics can be produced in rocks by at least four distinct mechanisms (e.g. Hobbs et al. 1976, Mainprice and Nicolas 1989). The first two mechanisms, anisotropic crystal growth (Chapter 5.3) and rigid-body rotation


of grains in a flowing matrix, can be recognised because individual grains are not deformed. A simple way in which intracrystalline plasticity can produce a crystallographic fabric is shown in Fig. 4.20, in which a grain is deformed by the operation of a single slip system. As the grain rotates, the slip plane becomes orientated normal to the finite shortening direction and the slip direction rotates towards a bulk shear plane (cf. Allison and LaTour 1977). Although this is a simplified case, it illustrates that lattice reorientation can occur by dislocation glide. Fourthly, crystallographic fabrics may be created by recrystallization. The exact mechanism is unclear, but experiments demonstrate that the mechanism of dynamic recrystallization determines the type of LPO (Gleason et al. 1993). A great deal of work has gone into theoretical attempts to predict crystallographic fabrics, and to compare them with experimental results and natural LPOs (see reviews by Law 1990, Wenk and Christie 1991). The general development of a crystallographic fabric by intracrystalline plasticity can

51

be understood through relatively complex models that allow for multiple slip systems and make additional assumptions, such as strain compatibility between grains, uniform stress, or viscoplastic self consistent theory, which minimizes stress and strain differences from an average value (e.g. Wenk and Christie 1991). The work of Jessell (1988a, b, Jessell and Lister 1990) for quartz incorporates effects of both the intracrystalline plastic mechanisms referred to above, lattice rotation and recrystallization. These models show that the dominant slip plane does not necessarily align with the bulk shear plane, and the dominant slip direction is not necessarily parallel to the shear direction, in contrast to the single slip system model above. The simulations also show that the LPO is affected by the type of flow during deformation, the finite strain magnitude and type, and the temperature, which determines what slip systems are active. Crystallographic fabrics carry large amounts of structural information, and have been used to analyze shear sense (Section 7.9), finite strain, and grade of deformation.

Chapter 5

Diffusive Mass Transfer and Phase Transformations in the Solid State 5.1

Introduction

Material removal, transport and deposition without fracturing, lattice distortion or melting at metamorphic grades at and above amphibolite facies suggest diffusive mass transfer (DMT) in the solid state. The major evidence for solid state as opposed to fluid assisted DMT is provided by phenomena that can be explained in terms of known solid state diffusion coefficients, including many metamorphic textures. Recent experimental evidence suggests that a variety of transformations, some involving DMT, may also be important deformation mechanisms in the solid state (Section 5.9). There has been considerable recent discussion about the potential tectonic importance of superplasticity, which involves a composite solid-state deformation mechanism: these microstructures and mechanisms are described in the last section, 5.10.

5.2

ture without bond breaking and with minor shape change. An example is the phase transition in quartz.

2. Martensitic-like transformations are coherent transformations involving dominantly shear strain. geological examples include:

Important and

3. Coherent exsolution may involve both dislocation movement and DMT to effect crystallographic and chemical rearrangements. Exsolution of clinopyroxene from orthoenstatite is one of the best studied examples.

4. Order-disorder transformations occur by disordering of cation site occupancies, for example in Mg spinels.

Fundamental deformation mech- 5.3 Grain shape fabrics and ribbon anisms of solid state diffusive mass grains transfer and phase transformaA grain shape fabric of unstrained grains is an important mitions crostructure in metamorphic rocks at higher grades (Plate 26).

The grain shape fabric can be produced by anisotropic grain boundary migration recrystallization, although the details of the mechanism are not clear (e.g. Jessell 1987), and by anisotropic crystal growth (e.g. Shelley 1989a, b). Both of these processes are important DMT mechanisms. Quart-mica rocks at higher metamorphic grades commonly have a distinctive microstructure of flat grain boundaries between quartz and mica parallel to the mica basal plane, and quartz-quartz grain boundaries approximately perpendicular to the mica flakes (Plate 27). A related effect is pinning of quartz-quartz grain boundaries at the end of mica flakes. These microstructures attest to impeded grain boundary movement, and can be attributed to a greater surface energy between quartz and mica than between quartz and quartz grains. Where one phase has a strong preferred orientation, the impeding effect on grain boundary growth may create a grain shape fabric in the other phase, which grows parallel to the shape fabric (Plates 26, 27). Monocrystalline quartz ribbon grains (Fig. 5.1) in high grade gneisses may form by the above mechanism, and 1. Displacive transformations are changes in crystal struc- are characteristically free of internal structures (by contrast

DMT in the solid state may occur though the crystal lattice by the movement of lattice defects. This is known as volume diffusion, and the resulting deformation is NabarroHerring creep (e.g. Nicolas and Poirier 1976). Coble creep is a different type of deformation resulting from solid state DMT through grain boundaries, which have different diffusion characteristics from the grain interiors (e.g. White and White 1981). Deformation by either or both types of creep is known by the general term diffusion creep. Diffusion occurs in response to gradients in chemical potential (Fick’s law) which may be created by variations in normal stress, internal strain energy, and grain boundary configuration (cf. Section 3.2). Solid state transformations can be subdivided into four types (Green 1985, Kirby and Stern 1993), which may be important in deformation because they are sensitive to deviatoric stress. The transformed polymorph is related to the original by crystallographic rules in the first three types, which are known as coherent transformations.

52

CHAPTER 5. DIFFUSIVE MASS TRANSFER AND PHASE TRANSFORMATIONS IN THE SOLID STATE

53

54


to lower grade ribbons; Section 4.7, Fig. 4.13). Gower and Simpson (1992) proposed that the geometry of quartzfeldspar grain boundaries in ribbon grains is largely controlled by a combination of dislocation creep and diffusion that results in a microstructure of straight quartz-feldspar boundaries perpendicular to the shortening, and cusps pointing in the extension direction. However, recently MacKinnon et al. (1997) have proposed that the textures of some ribbon grains may form by filling of microfractures, based partly on the evidence that grains adjacent to the ribbons appear to be truncated by the ribbon boundaries, and that the ribbons contain wall rock particles that have very similar geometries to inclusion bands (Section 3.12). Some of these features are illustrated in Fig. 5.1. Both models account for many of the observed features of high-temperature ribbon grains. A possible approach to distinguishing the two mechansims may lie in careful examination of the ribbon tips, which could yield evidence for incipient microcracking or diffusion processes.

than pinning the crystal boundaries as described in Section 5.3, which are then free to grow at a faster rate. This has been distinguished as a separate type of grain boundary migration mechanism called fast or free grain boundary migration by Urai et al. (1986).

5.5 Decussate texture Decussate texture consists of randomly orientated, interlocking elongate crystals (Plate 28). It is especially common in amphiboles and micas, and arises because of unequal growth rates in different crystallographic directions. The growth anisotropy modifies the ideal foam texture to favour grain boundaries parallel to the faster growth directions, and may lead to randomly orientated acicular crystals for particularly high anisotropies.

5.6 Porphyroblasts trails 5.4 Foam texture, static and secondary recrystallization 5.6.1 Characteristics Monomineralic metamorphic rocks, particularly at higher metamorphic grades, may have a distinctive microstructure consisting of approximately hexagonal grain sections with straight or curved grain boundaries and strain-free grain interiors, known as foam or granoblastic polygonal texture (Fig. 5.2). This microstructure is readily interpreted in terms of grain boundary migration and DMT in the solid state driven by surface energy. The process of achieving a minimum energy configuration under hydrostatic stress is static recrystallization. The lowest surface energy will occur for the smallest surface area to volume ratio. The minimum surface energy configuration for equal-sized, space-filling polyhedra are rhomb dodecahedra (truncated octahedra). However, this figure does not have an isotropic distribution of surface energies, because angles between grain edges around four grain junctions are unequal. Grain edges and boundaries may curve to allow all four grain edges to intersect in the same angle of 109.5°, resulting in a structure of polyhedra with curved faces, similar to liquid films and as observed in annealed alloys and chromitites in ultramafic intrusions (Smith 1964). In two dimensions, 120° triple junctions are common between aggregates of the same phase. The surface area to volume ratio can also be reduced by an increase in grain size, which often accompanies static recrystallization, and is known as Ostwald ripening or exaggerated grain growth. This microstructure can be useful in revealing a post-tectonic period of relatively high temperatures. Grain shape fabrics can be produced even under static recrystallization by at least two processes. An older fabric can control grain boundary mobility by surface energy effects (Section 5.3) or a new mineral growth may overgrow a preexisting fabric: this is known as mimetic crystallization. Relatively large, irregular crystals which commonly contain inclusions of secondary phases are formed by the process of secondary recrystallization. This occurs when second phases become incorporated into the growing crystal, rather

and

inclusion

Porphyroblasts are single crystals grown during metamorphism with a larger size than the adjacent grains in the matrix. They are only widespread in rocks that have been at upper greenschist facies or higher, and are most common in metapelites or metabasites. Chlorite, chloritoid, biotite, garnet, cordierite, sillimanite, kyanite, andalusite, and staurolite are common porphyroblastic minerals. Porphyroblasts commonly contain inclusions, the most common of which are opaque minerals, aluminium-rich phases such as sillimanite or spinel, quartz, zircon, apatite, rutile, and sphene. A porphyroblast with a very high density of inclusions is a poikiloblast. Inclusions may have a variety of textures that contain important microstructural information, and need to be described carefully. The shape of individual inclusions may be euhedral, platy, linear, or rounded. Inclusion trails are aligned inclusions that define a fabric which is given the symbol where the refers to an internal fabric (compared to for fabric external to the porphyroblast). usually has a two-fold rotational symmetry about an axis parallel to the length of the porphyroblast. may be straight or curved. It may be continuous and curve smoothly from the core of the porphyroblast to the rim; it may be sharply deflected or cut off along deflection and truncation surfaces respectively, which are sub-parallel to more external parts of the inclusion trail. Special types of inclusion trails include snowball textures (Plate 29), which are spiral-shaped inclusion trails that curve through more than 180°, with the total curvature decreasing from the centre towards either end of the rotation axis (e.g. Powell and Treagus 1970, Busa and Gray 1992). Millipede texture consists of a straight, parallel which is deflected into at the porphyroblast margin. Helicitic texture consists of folds defined by Inclusions may be crystallographically controlled by the porphyroblast: this can give rise to a number of distinctive textural zoning patterns such as sector zoning and re-entrant zones. The relation between and is very


55

5.6.3 Relationship to deformation Porphyroblast textures can yield detailed information about the relation between porphyroblast growth (and thus P-T conditions) and deformation. This is best analyzed by looking at the relationship between inclusions and which can be classified into four non-interpretive categories (Fig. 5.3): 1. Inclusions are randomly orientated, is commonly also curved around the porphyroblast (Fig. 5.3a).

important: can be continuous with or truncated by at the porphyroblast boundary. This forms the basis of the interpretation of inclusion trail patterns (Section 5.6.3).

5.6.2 Growth mechanisms Porphyroblasts are a mineral growth and therefore form by DMT. In many cases there is no evidence for a fluid phase and it is assumed that diffusion was in a solid state. This can sometimes be demonstrated by the observation that the most mobile components are those with the highest solid state diffusion coefficients. Porphyroblasts grow in response to thermodynamic considerations including composition, temperature, internal strain, and surface energy. Unusually large porphyroblasts, especially those which contain high densities of inclusions, may have grown by secondary recrystallization (Section 5.4). The size and distribution of porphyroblasts reflects a balance between the energy required for nucleation and that for growth. High ratios of nucleation to growth energies will favour the formation of fewer and larger porphyroblasts.

2.

is discontinuous with is often also curved around the porphyroblast (Fig. 5.3b).

3.

is continuous with but and shapes or orientations (Fig. 5.3c).

have different

4.

is continuous with and and shapes and orientations (Fig. 5.3d).

have similar

These categories are objective descriptions, but also allow straightforward interpretations based on the original ideas of Zwart (1960, 1962), and the recent comprehensive treatment of Passchier and Trouw (1996). The first category suggests growth of a porphyroblast before deformation i.e. pretectonic (e.g. Borradaile et al 1982, p. 438-441), but exceptions are noted by Passchier and Trouw (1996). The curvature of is due to deformation of the matrix around the more rigid porphyroblast (Plate 30). The second category is called intertectonic by Passchier and Trouw (1996), and indicates porphyroblast growth following the deformation event that generated but preceding another event that formed (Fig. 5.4). The two deformations that created and may be separated by other deformations that are not recorded by or (e.g. Johnson and Vernon 1995). An intertectonic interpretation is reliably demonstrated when is folded and is straight (Plate 31, e.g. Borradaile et al. 1982, p. 441, 453). The third category indicates syntectonic porphyroblast growth (Plate 32). A common example of different geometries assumed by and is a coarser grain size in than which can be interpreted to show that the matrix was overgrown by the porphyroblast relatively early and subsequently underwent grain size reduction. Coarsening of towards the margin of a porphyroblast indicates growth during prograde conditions. A particularly complex pattern may be produced by overgrowth of one type of in a pressure shadow and another type of in a strain cap, both Si being formed in the same event (Shoneveld 1977). The fourth category indicates post-tectonic growth (Plate 33). However, even in the case of concordant and with no obvious differences in geometry, may be curved around the porphyroblast, suggesting either that the last increment of deformation outlasted porphyroblast growth, or that low strain deformation occurred subsequently. Detailed analyses shows that some porphyroblasts can not be accommodated into the above simplified scheme. Truncation of within the porphyroblast may indicate two phases of porphyroblast growth and foliation development. Truncation surfaces can also be formed by post-tectonic overgrowth of an early foliation which has been truncated by a strain cap

56



during a later deformation, or a later stage of a progressive deformation (Passchier et al. 1992). Bell et al. (1992) have argued that truncation surfaces form in the strain shadow at the end of a porphyroblast during progressive deformation. The same interpretations may apply to deflection surfaces. Problems of interpreting curved inclusion trails are discussed in Section 7.8.

5.7 Reaction rims, relict minerals, coronas and symplectites Reaction rims are rims of altered mineralogy around grains. This common metamorphic texture testifies to DMT since metamorphic mineral growth requires mobility of components. Relic minerals are mineral grains that have been mostly replaced by reaction rims. Reaction rims that completely surround a grain are coronas (Plate 34), which are more common in high grade rocks. Symplectites are lamellar or vermicular intergrowths that are common in reaction rims (Plate 35). The intimate mixture of the phases involved, the fine grain size and the disequilibrium grain shape shows that diffusion was not possible over long distances.

5.8 Chemical zoning

57

ural examples are so far virtually undescribed. The quartz phase transition can be viewed as Dauphiné twinning on the scale of the unit cell, so that relict Dauphiné twins as well as microfractures (Section 2.3.11) could constitute microstructural evidence of this deformation mechanism (Kirby and Stern 1993). Martensitic-like transformations can occur during cooling under hydrostatic stress, but in the case of the transformation, deformation-induced transformation can be distinguished by lack of twinning in the clinoenstatite (e.g. Kirby and Stern 1993). The best known microstructures of Martensitic-like transformation are those of the olivine-spinel system. “Fingers” or lobes of spinel grow into olivine parallel to in experiments on (Vaughan et al. 1984), and a similar phase transformation microstructure of aragonite crystals has been observed growing into calcite parallel to (Hacker and Kirby 1993). These microstructures are consistent with the theoretical analysis of Green (1985), which suggest that ellipsoids of the stable phase should grow with long axes parallel to possibly coalescing into ellipsoids separated by cusps. However, later experiments on olivine showed lens shaped spinel inclusions perpendicular to These can be accounted for by stress redistribution at a microscopic scale between stronger olivine and weaker spinel (Green and Burnley 1989). Exsolution of Ca-clinopyroxene from orthoenstatite is similar to the transformation described above, with the addition of Mg, Fe and Ca diffusion. Cooling exsolution may be distinguished from deformation exsolution by the orientation of the exsolved phase (Champness and Lorimer 1974).

Metamorphic minerals, especially garnets, amphiboles, pyroxenes and feldspars, commonly show systematic compositional variations from core to rim, or a relatively constant core composition and a quite different rim composition. These variations can be seen in the colours of some minerals (e.g. blue-green colour variations in amphiboles), but are most accurately revealed by electron microprobe studies. 5.10 Superplasticity Zoning is clear evidence for DMT, and may form in two difSuperplasticity was first used to describe the behaviour of ferent ways. metals deformed in extension to large strains without failure Growth zoning. Preferential partitioning of an element into (e.g. Langdon 1982). The dominant deformation mechana mineral during growth can cause growth zoning by deism in these experiments is grain boundary sliding, with inpletion of that element within an equilibrium volume. compatibilities between grains mainly relieved by solid state This describes how growth zoning may occur at conDMT around grain boundaries. Unfortunately the term has stant temperature and pressure: however, the P-T conbeen used in at least three different ways in a geological conditions may change during mineral growth, which will text (Gilotti and Hull 1990): also change the distribution coefficients for elements 1. As a deformation mechanism consisting of grain boundbetween mineral and matrix. This is probably the most ary sliding accommodated by DMT and/or intracrystalcommon way in which growth zoning is produced. line plasticity. Reaction zoning. After a crystal has formed, DMT may oc2. As a set of deformation conditions and mechanical recur with neighbouring minerals in response to a change sponses in rock. In particular, dependence of strain rate in P-T conditions from the formation of the original metamorphic rock. Commonly the rims of minerals have on an inverse power of grain size, high temperatures, and proportion between strain rate and a low power of stress compositions that reflect lower metamorphic grades than the cores: the rims have equilibriated during the cooling have been regarded as characteristic (Section 9.4.3). of a rock. Such zoning is known as retrograde zoning. 3. As a description of strain e.g. “continuous, homogeneous deformation to very large strain”. This “phenomenological” definition proposed by Gilotti and Hull (1990) 5.9 Solid state phase transformation is not in widespread use, and the more common first geomicrostructures logical definition will be followed here. Microstructures characteristic of rocks considered to have Experiments suggest that phase transformations under deviatoric stress could have distinctive microstructures, but nat- been deformed by Superplasticity have been summarized by

58


Fliervoet and White (1995). They include: 1. A very fine grain size (less than in quartzites, more in carbonates) (e.g. Behrmann 1985). 2. Equiaxed grains with diamond or blocky shapes (e.g. Drury and Humphreys 1988). 3. Alignment of grain boundaries over several grains (e.g White 1977). 4. An inverse correlation between finite strain and grain size (e.g. Evans et al. 1980). 5. A weak crystallographic fabric (Rutter et al. 1994). 6. An inverse correlation between dislocation density and grain size (e.g. Behrmann 1985). 7. High dislocation densities and voids at grain triple junctions (e.g. White 1977). 8. Large mismatches between the crystallographic orientations of adjacent grains (e.g. Fliervoet and White 1995). 9. Rotations between grains that can not be explained by rotation axes perpendicular to known Burgers vectors (e.g. Fliervoet and White 1995).

Several of these microstructures are ambivalent in their interpretation. For example, Fliervoet and White (1996) describe an exceedingly fine-grained quartz mylonite which deformed exclusively by dislocation creep with no grain boundary sliding. A fine grain size is apparently a necessary but not sufficient condition for superplasticity. Crystallographic fabrics may be developed during superplastic flow of finegrained calcite rocks, with the implication that they could in principle become strong after sufficient strain (Schmid et al. 1987, Rutter et al. 1994). Fig. 5.5 shows an ultramylonite that shows some of the characteristic features of superplasticity. The clear definition and recognition of superplasticity in rocks is still a difficult matter. An important example of superplasticity related to phase transformation may occur when olivine transforms to spinel. This is known as the anticrack theory of phase transformation faulting, which suggests that olivine transforms into spinel in anticracks which localize into faults within which grain boundary sliding occurs on fine grained spinel (Green and Burnley 1989, Burnley et al. 1991). The theory accounts well for earthquakes that occur at depths too great for frictional behaviour, and is well supported by experimental evidence (e.g. Tingle et al. 1993).

Chapter 6

Magmatic and Sub-magmatic Deformation 6.1 Introduction Identification of deformation microstructures and mechanisms in rocks containing melt has several important tectonic implications, notably for the problem of melt extraction and in the interpretation of pluton ascent and emplacement mechanisms. However, microstructural criteria to distinguish magmatic, sub-magmatic and non-magmatic deformation microstructures and mechanisms are not well established, and the distinction is best made using a combination of mesoscopic and microscopic evidence. The mesoscopic evidence is briefly described in Section 6.4.

6.2 Fundamental deformation mechanisms and microstructures in rocks containing melt 6.2.1 Magmatic flow Flow of magma (i.e. melt and crystal phases) by transport of rigid crystals is often regarded as the typical deformation mechanism in melt-bearing rocks. Magmatic flow has been defined as flow by displacement of melt and rigid-body rotation of crystals without sufficient interaction to cause crystal plastic deformation (Paterson et al. 1989). As shown below, crystal interaction in melts may lead to deformation by cataclasis and diffusive mass transfer (DMT) as well as intracrystalline plasticity. A more useful, general definition for magmatic flow is flow of melt and crystals without crystal deformation; this definition allows for the possibility that crystals may be deformed by processes other than crystal plasticity, and also describes the flow of a suspension. It leads naturally to a definition of magmatic microstructures as those which indicate melt-present deformation without crystal deformation. Some experimental studies suggest that viscosities of meltladen systems reduce abruptly by orders of magnitude when the proportion of melt increases beyond a value known as the critical melt fraction, CMF (Arzi 1978, Van der Molen and Paterson 1979). The CMF is commonly taken to be 30%, but may be as much as 50% (Vernon et al. 1988) or as little as 10-20% for gabbroic rocks (Nicolas et al. 1988). Experiments on two silicate melts reported an increase in viscosity of three orders of magnitude, and a change from Newtonian to non-Newtonian behaviour, as melt fraction increased from

40 to 60% (Lejeune and Richet 1995). The importance of the CMF is further suggested by the observation that the maximum proportion of phenocrysts in volcanic rocks is 55-65%: volcanic rocks with larger proportions of phenocrysts may not be able to erupt because their viscosities are too high (Marsh 1981, Wickham l987). A minimum melt proportion for magmatic flow can also be deduced from the critical packing density of crystals to bring them into a coherent mass. The critical packing density depends on crystal shape, size, size distribution, packing arrangement, and amount of compaction. Table 6.1 summarizes porosity at critical packing density, which is equal to the minimum melt proportion, for some combinations of these factors in geometrical models, and estimates for magmas. Surface energies of the crystal and liquid phases may be important in flow because they can affect melt distribution (e.g. Jarewicz and Watson 1984, 1985), and indeed an explicit relationship between the CMF and surface energy can be formulated (Riley 1990). A more fundamental parameter than melt fraction for determining the rheology of melt-laden systems may be the contiguity (the fraction of grain surface area in contact with other grains), because contiguity affects surface energy and the resistance to shearing at the average surface. A load bearing framework of crystals breaks down at contiguities of less than 0.15-0.2 (Miller et al. 1988), which correspond to variable equilibrium melt fractions depending on surface energies and grain size and shape distributions (German 1985). These factors may explain the variation in estimates of the CMF. Viscosities of suspensions depend on fluid (melt) composition, pressure, temperature, and proportion, size and shape distribution of solids. The viscosity of melts containing spherical crystals is often approximated by the EinsteinRoscoe equation (Roscoe 1952):

where is the viscosity of the pure melt, and is the crystal fraction for coherent packing. Values of 0.6 for seem to represent silicate systems well. n has a theoretical value of 2.5 for variably-sized spheres, which is also a good fit to data for silicate systems, but the viscosity-melt fraction relationship has a slight temperature dependence which can be allowed for by letting n vary between 2.0 and 2.5 with temperature (Lejeune and Richet 1995). Another relationship between viscosity and crystal fraction includes the effect of grain size (Sherman 1968). At lower melt fractions, the rhe-

59

60

CHAPTER 6. MAGMATIC AND SUB-MAGMATIC DEFORMATION

ology of the magma may be at least partly controlled by the 6.2.3 Magmatic and sub-magmatic flow and mechanical properties of the solid phases or by the factors afrheology fecting diffusive mass transfer (DMT) through the melt phase It appears that two fundamental changes in deformation (see below). mechanisms occur as melt fraction is increased from the While some experiments and observations suggest a signisolid to liquid states, and these correspond to reductions in ficant mechanical change at the CMF, other experiments cast strength or viscosity. An order of magnitude decrease in doubt on the existence of a large change in viscosity over a strength occurs as melt fraction rises from zero to only a narrow range of melt fraction (e.g. Rushmer 1995), and it few percent melt, and solid-state deformation mechanisms has been argued that the abrupt change in viscosity observed change to melt-enhanced diffusion creep, with other deformin some previous experiments was due to a combination of ation mechanisms (cataclasis, grain boundary sliding, and intemperature effects and increase in water contents of melts tracrystalline plasticity) also possibly accommodating crys(Rutter and Neumann 1995). There may be a continuous detal deformation. As melt fraction increases still further, suscrease in magma viscosity as melt fraction increases, and no pension flow can occur, reducing strength by three orders CMF, if the change in water content is allowed for. Yet other of magnitude, and crystal deformation ceases. Figure 6.1 experiments suggest that the change in strength at the CMF shows these two transitions in mechanism schematically, but is due to dilatancy hardening (cf. Brace and Martin 1968) at it should be emphasized that the values and range of melt fraclow melt fractions (Renner et al. 1999). tions over which the transitions occur, and the magnitudes of the changes in viscosity, vary with magma composition, and are presently the subject of some debate.

6.2.2

Sub-magmatic flow

Sub-magmatic flow can be defined as deformation involving flow of melt and crystals with crystal deformation. Submagmatic microstructures are accordingly those indicating melt-present deformation with crystal deformation, and correspond to the pre-full crystallization fabric of Hutton (1988). Crystal deformation by DMT through the melt phase is probably the dominant deformation mechanism at low melt fractions and lower strain rates. The thermodynamic considerations for DMT through melt are the similar to those described for DMT through solution in Section 3.2. Experiments show that strength decreases by an order of magnitude, and intracrystalline plasticity changes to melt-enhanced diffusion creep, when melt proportion increases from zero to only 3-5% (Cooper and Kohlstedt 1984, Dell’Angelo and Tullis 1988, Dell’Angelo et al. 1987). The weakening occurs due to enhanced grain boundary diffusion though the melt. Experiments on rock analogues by Park and Means (1996) also demonstrate the importance of this process in some systems at low melt fractions. Surface energy considerations may be important in sub-magmatic deformation. Most diffusion creep models assume isotropic surface energy, but measurements suggest an order of magnitude anisotropy in olivine, for example (Cooper and Kohlstedt 1982). This may affect strength by allowing a continuous film of melt along two-grain boundaries instead of limiting melt to three or more grain junctions, as usually assumed (Hirth and Kohlstedt 1995).

6.3

Mesoscopic evidence for magmatic and sub-magmatic flow

It is important to use mesoscopic evidence in conjunction with microscopic observations to distinguish magmatic, submagmatic and non-magmatic flow. Fabrics in magmatic rocks can be defined by phenocryst alignment, matrix mineral shape fabrics, compositional banding, enclave or schlieren shape fabrics, or magnetic anisotropy. Phenocrysts may be tiled or imbricated (e.g. Blumenfeld 1983, Blumenfeld and Bouchez 1988). The fabrics can have any shape from prolate to oblate. The mere existence of any of these fabrics is not diagnostic of melt-present flow, but detailed examination on the mesoscopic scale may provide some strong indicators to distinguish magmatic/sub-magmatic from non-magmatic flow. The following criteria are proposed as diagnostic of magmatic or sub-magmatic flow because they suggest extremely large, non-systematic strain gradients in an homogeneous rock on a meter scale, which are mechanically unlikely in the nonmagmatic state. Microscopic studies should be used to back up these lines of evidence. 1. Abrupt and non-systematic changes in fabric orientation. 2. Abrupt and non-systematic changes in fabric intensity. 3. Irregular, polyclinal, disharmonic or rootless folds in


61

massive rocks that have no sign of mechanical aniso- compositions. These foliations may have formed in a short tropy or shear surfaces (e.g. McLellan 1984). interval of cooling after the melt has been emplaced but before final solidification. Phenocryst density and size distributions may be affected The relationship between igneous layering and fabrics can by magmatic flow. Interaction between phenocrysts in conalso be used to suggest magmatic/sub-magmatic flow. For centrations greater than 8% in a moving fluid creates a “grain example, Pons et al. (1995) describe centimetre-scale cycles dispersive pressure” which is proportional to the velocity with sharp bases consisting of a layer of fine-grained amgradient in the magma and is therefore greatest at the margins phiboles to a medium grained amphibole-plagioclase layer to of an intrusion (Bagnold 1954, Komar 1972a, b). Phenocrysts a slightly porphyritic plagioclase-K-feldspar-quartz layer in are concentrated into the centre of the intrusion, and may also alkali granites. The layers are parallel to the preferred orientcoarsen in this direction. The velocity distribution and phenoation in any of these minerals, and the fabric never cross-cuts cryst concentration should be plug-shaped with a central rethe layering. The layering and fabric can be explained by gion of high density and an abrupt decrease in density towards trapping of a mafic cumulate layer from a slowly convecting the sides of an intrusion, due to the effect of phenocryst conmagma, leaving an increasingly felsic rich magma to cryscentration on viscosity, and this is indeed observed in many tallize as the upper layers. By contrast, Paterson and Vernon (1995) discuss several examples of an apparently mag- natural examples (e.g. Ryan 1995). Discordance between fabrics in a xenolith and a surroundmatic foliation that cross-cuts both gradational compositional changes and contacts between phases of different magmatic ing igneous rock is often interpreted as evidence for mag-

62


matic flow, but this criteria is not diagnostic: fabrics may be preserved in xenoliths during non-magmatic deformation due to competence contrasts between the xenolith and subsolidus matrix. Sub- or non-magmatic fabrics on this scale are demonstrated by deformation of individual crystals such as feldspar megacrysts. However, concordance between fabrics within enclaves and the host rock can be evidence for magmatic/sub-magmatic flow if the enclaves can be interpreted as partially molten during deformation (e.g. Vernon and Paterson 1993). Shear zones may develop during intrusion by magmatic, sub-magmatic and non-magmatic mechanisms (e.g. Guineberteau et al. 1987, Pons et al. 1995). Magmatic shear zones can be identified by fabrics defined by unstrained igneous minerals (e.g. Miller and Paterson 1994) or reorientation of a magmatic fabric in the shear zones (e.g. Pons et al. 1995). Non-magmatic shear zones in the latter study had pressure shadows filled by epidote around megacrysts, and ribbon grains formed by intracrystalline plasticity in quartz. Magmatic shear zones were identified in the experiments of Park and Means (1996) by analyzing movements of solid inclusions, but there was very little direct microstructural evidence to distinguish the shear zones from the unsheared walls. Criteria which are sometimes misused to demonstrate magmatic flow include imbrication, which may occur in mylonites (Section 7.11.3), magnetic anisotropy, which is well known in metamorphic rocks, and S-C fabrics, which may form in either sub-magmatic or non-magmatic deformation (Blumenfeld and Bouchez 1988).

formed inclusions in annealed minerals, and by a granoblastic texture in mineral aggregates (especially quartz) which themselves define a shape fabric. Primary igneous fabrics can also be recognized by random spatial relationships between different phases. By contrast, solid state deformation produces higher frequencies of contacts between like grains because new grains preferentially nucleate on the same phase (e.g. Ashworth and McLellan 1985).

6.4.2

Magmatic crystallographic fabrics are due to orientation of euhedral, inequant phenocrysts during flow. In mafic rocks, (010) crystal faces are parallel to the magmatic foliation in olivines, pyroxenes and feldspars (Benn and Allard 1989). The [001] direction of olivine and clinopyroxene is parallel to the magmatic lineation, and [100] has been observed parallel to the magmatic lineation in feldspars within gabbros and tonalites (Benn and Allard 1989). Such crystallographic fabrics can be distinguished from fabrics due to intracrystalline plasticity by complete lack of evidence for intracrystalline plastic microstructures (Chapter 4), and in the case of olivine, by the fact that high-temperature non-magmatic deformation leads to [100] parallel to the lineation rather than [001], as observed for magmatic fabrics (Benn and Allard 1989).

6.5 Sub-magmatic microstructures 6.5.1

6.4 Magmatic microstructures 6.4.1

Grain shape fabrics

The typical microstructure of magmatic flow consists of a preferred orientation of euhedral phenocrysts, in an isotropic matrix that shows igneous textures (Plate 36). Grain shape fabrics are commonly defined by feldspars and micas in felsic rocks, and may be defined by feldspar, olivine and pyroxene in mafic rocks (e.g. Benn and Allard 1989). This is diagnostic of magmatic flow if there is no sign of any other deformation mechanism. Since intracrystalline plastic deformation microstructures are readily visible in quartz after only small strains, the lack of deformation in quartz (no undulose extinction, subgrains, or kink bands) is a reliable criterion for magmatic flow. Lack of sub-solidus deformation can also be checked from inclusions (e.g. rutile) in quartz: the inclusions should be undeformed themselves and randomly orientated (e.g. Mitra 1976, Stel 1991). Other primary, undeformed igneous features such as strongly euhedral crystals, igneous (e.g. idiomorphic) zoning, growth twins, and ophitic texture can be used to demonstrate the absence of intracrystalline deformation. A fabric in igneous rocks can only be confirmed as a magmatic microstructure by lack of any evidence for other deformation microstructures or mechanisms, including any of the microstructures described in previous Chapters. It may be difficult to distinguish magmatic deformation from non-magmatic deformation followed by static recrystallization. Post-deformational annealing can be revealed by de-

Crystallographic fabrics

Grain shape fabrics

Grain shape fabrics may be expected to form in submagmatic flow by both rigid body rotation and crystal deformation by DMT or intracrystalline plasticity. The crystal deformation combined with evidence of melt such as matrix with igneous textures is diagnostic of sub-magmatic flow (e.g. Quick et al. 1992). Localization of melt in structures that are coeval with the foliation constitutes good evidence for submagmatic deformation: this sort of evidence is more readily seen at outcrop scale. As for magmatic flow, static recrystallization may give the matrix the false appearance of a primary igneous texture.

6.5.2

Intracrystalline plasticity

Intracrystalline plasticity in quartz may be expected in the presence of melt under appropriate differential stresses (>1 MPa) and temperatures (700 to 800°C) on the basis of granite and quartz flow laws (Rutter and Neumann 1995). The very common appearance of undulose extinction in quartz within granites suggests that intracrystalline plasticity occurs during sub-magmatic deformation (Plate 37). Subgrain formation or recrystallization of quartz has been taken as an indicator of sub-magmatic deformation where an overall igneous texture is preserved and no other non-magmatic deformation event is known (e.g. Bouchez and Gleizes 1995). Deformation twins and bent twins in feldspar may also hint at sub-magmatic intracrystalline plasticity (Plate 38). However, these microstructures on their own do not demonstrate that melt was present during deformation.

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6.5.3 Diffusive mass transfer Despite the experimental evidence for the importance of meltaided diffusion creep, microstructural evidence has remained elusive, probably because melt films have almost no potential for preservation in the geological environment. However, experiments suggest some features that could be used: truncated, embayed, scalloped or overgrown grain boundaries in igneous rocks may be analogous to some of the features discussed in Chapter 3 that indicate DMT through a fluid phase (e.g. Dell’Angelo et al. 1987, Park and Means 1996). The potential importance of surface energy anisotropy in melt distribution needs to be investigated and may have some microscopic expression in the form of differences between crystal faces.

6.5.4

Cataclasis

Paradoxically, some of the clearest sub-magmatic microstructures involve cataclasis of the crystals, such as microfractures which are healed by melt (e.g. Hibbard 1987, Bouchez et al 1992, Karlstrom et al. 1993). Microfractures in plagioclase can be demonstrated to have been filled by melt from the following criteria: 1. The microfractures are intragranular. This allows that the plagioclase crystals were in contact with melt. 2. The microfracture filling is compositionally and crystallographically continuous with the same phase in the igneous groundmass of the rock. 3. The composition of the microfracture filling is compatible with the later stages of the igneous petrographic history of the rock. Plagioclase microfracture fillings may have lower anorthite contents than their host crystals, consistent with progressive evolution towards a minimum melt composition (Bouchez et al. 1992). The relation between quartz and plagioclase in the microfracture fillings also suggests a residual melt: feldspars are on the walls or tips of the microfractures. 4. Early crystals (e.g. biotite, sphene) are trapped within the microfracture fillings.

Cataclastic microstructures observed in experiments confirm the potential importance of cataclasis in sub-magmatic deformation in the experiments of Rutter and Neumann (1995). Up to 10% melt, axially-orientated cracks formed and filled with melt, and the sample was faulted. Between 10 and 45% melt, cataclastic flow occurred with pore collapse. Axially orientated microcracks formed by high melt pressures have been observed in granitic aggregates containing 2-15% melt (Dell’Angelo and Tullis 1988). Connolly et al. (1997) have demonstrated that microcracking caused by volume increase during melting is a viable way to create permeable fracture and melt-pool networks in a muscovitebearing quartzite. The syn-kinematic experiments by Park and Means (1996) also recorded fracture, localized along a kink band boundary. On a larger scale, several cataclastic features are commonly associated with melt in migmatites. Metatextites often consist

of a competent body sub-divided by fractures that are filled by melt. Melt may form in pressure shadows at the ends of boudins, and fill faults (e.g. Quick et al. 1992). Quartz-feldspar neosomes have been described accumulated under impermeable refractory layers such as amphibolite sheets which are boudinaged, allowing the neosomes to rise into boudin necks, and forming a geopetal structure (Burg 1991). Segregations in shear zones and along axial planes of folds in migmatites are common. Localization of the melt in these structures indicates that the melt was syntectonic. This sort of evidence has great relevance to the problem of extracting melt to form plutonic bodies (e.g. Wickham 1987).

6.6

Other microstructures

Park and Means (1996) introduced the term “contact melting” to describe melting at contact points between grains observed in their experiments, and suggested that indented boundaries and truncation of growth zoning might be indicative of the process in nature. Another suggestion from their experiments is that filter pressing and expulsion of melt might be recognizable from layers characterized by intracrystalline plasticity adjacent to layers of less-deformed or undeformed rock formed by crystallization of the expressed melt. Overgrowths of feldspar along low-stress boundaries have been considered as indicators of magmatic or sub-magmatic deformation (Hibbard 1987), but they may equally be be formed by subsolidus DMT processes (Paterson et al. 1989).

6.7

Non-magmatic deformation

Non-magmatic microstructures reflect deformation without any melt present, which has also previously been known as sub-solidus or solid state deformation. Non-magmatic is preferred to “solid state” because it allows for the presence of fluids other than melt. Two sort of evidence can be used to suggest that non-magmatic deformation has continued as part of the same deformation event shown by magmatic or submagmatic features. The first sort is evidence for high temperature deformation, e.g. prism slip in quartz (recognized by quartz c-axes with an orientation close to the lineation e.g. Law et al. 1992, Lagarde et al. 1994), basal subgrain boundaries (Section 4.6), or albite exsolution lamellae indicating exsolution above the alkali feldspar solidus. A distinctive grain boundary shape in quartz consisting of reticular grain boundaries, resulting in a mosaic pattern, indicates crystallographic control of grain boundaries, extreme grain boundary mobility and therefore high temperatures (Gapais and Barbarin 1986). However, a difficulty with this sort of evidence is that the high temperature deformation may be a later event which is entirely unrelated to the magmatic deformation (Paterson et al. 1989). The second line of evidence is kinematic; if magmatic/submagmatic microstructures are kinematically compatible with non-magmatic microstructures, they can plausibly be related to the same event. This type of evidence is probably more reliable. One example is non-magmatic S-C fabrics recording the same shear direction and sense as magmatic/sub-

64

magmatic flow (Blumenfeld and Bouchez 1988, Miller and Paterson 1994). Another case is the formation of quartz rods, folds and boudins in a non-magmatic state along an intrusive interface, while the bulk of the intrusion was still magmatic, as suggested by the fact that all these structures die out away from the interface (Stel 1991). The non-magmatic simple shear zones described by Ramsay (1989) have a tangential extension direction around an intrusion: these can be related to ballooning strains during successive phases of expansion of the pluton.


Myrmekite is sometimes taken as evidence for nonmagmatic deformation, but this is unreliable as myrmekite can form by direct crystallization (Paterson et al. 1989). Hibbard (1987) has suggested that magmatic myrmekite can be distinguished by growth in dilational sites (“pressure shadows”) around phenocrysts, in contrast with tectonically induced myrmekite which forms in volumes perpendicular to the shortening direction (e.g. Simpson 1985).

Chapter 7

Microstructural Shear Sense Criteria 7.1 Introduction Shear sense can be defined formally as the rotation sense (sign) of the average angular velocity of material lines with respect to the directions of maximum and minimum stretching rate, the instantaneous stretching axes, ISA (e.g. Hanmer and Passchier 1991). This rotation is also known as the shearinduced vorticity or internal vorticity (Means et al. 1980, Williams et al. 1994). Features such as pressure shadows, porphyroclast inclusion trails and fibrous vein fillings may preserve some information about the orientation of the ISA, and can be used to determine internal vorticity. However, shear sense is commonly used in a much looser way to indicate the displacement sense of a fault or shear zone, usually specified by one or a combination of the usual terms for fault or shear zone movement i.e. normal, reverse, sinistral, dextral. In this wider usage, shear sense is defined with respect to a shear plane, which is a useful external reference frame. Rotation relative to this frame of reference is external vorticity, consisting of the sum of internal vorticity and any rotation of the ISA with respect to the external frame of reference, or spin, which should be accounted for in a rigorous shear sense evaluation. Shear sense should only be analyzed in a section that is perpendicular to the shear plane and parallel to the shear direction, which is known as the shear sense observation plane or the vorticity profile plane (Robin and Cruden 1994). The shear direction is often assumed to be parallel to mineral stretching lineations or slickenlines, and the shear plane to foliations or fault planes; hence the shear sense observation plane would be parallel to the lineation and perpendicular to the foliation (Fig. 7.1). Other sections can give incorrect shear senses when using marker displacements as shear sense indicators. However, the assumption that mineral foliation and lineation are parallel to the shear plane and shear direction breaks down in two circumstances: when the total shear strain is low, and in complex types of three-dimensional strain. Transpression with a high ratio of pure shear to simple shear is one example of the latter. In this case, the maximum finite stretching axis (tracked by the mineral lineation) may come to lie perpendicular to the shear direction (Fig. 7.2; Tikoff and Fossen 1993). This may explain some puzzling cases where shear sense indicators are not seen in the plane parallel to the lineation and perpendicular to the foliation, but in planes perpendicular to the lineation (e g. the Larder LakeCadillac deformation zone, Canada; Robert 1989). The vorticity vector in these cases must be near to the maximum fi-

nite stretching axis, and the vorticity profile plane must be perpendicular to the lineation. The vorticity vector in general may lie at any angle to the maximum ISA, depending on the amount of transpression, and this relationship can change across a single shear zone (Robin and Cruden 1994, Tikoff and Greene 1997). In these circumstances, the shear direction can only be established reliably by using shear sense indicators other than displaced markers, observed in a variety of sections. The shear direction is within a plane across which the shear sense appears to reverse. These complications do not apply to any of the examples used in this chapter, in which all illustrations are in the plane perpendicular to the foliation and parallel to the lineation, which is the correct shear sense observation plane in these cases. A convention used in many studies is to take the shear plane as perpendicular to the plane of the illustration, and the shear direction as horizontal within the plane of the illustration, so that the shear sense can be specified by dextral (clockwise) or sinistral (anticlockwise). Unfortunately this convention is not always explicitly recognized, sometimes creating the misleading impression of horizontal movement on dipslip faults or shear zones. The convention is used throughout this book, so that dextral or sinistral can be used to describe shear sense. Microscopic observations such as shown in the plates in this chapter are often invaluable aids to shear sense determination, especially in fine-grained rocks. This is because of the finer detail visible in thin sections, and because the sections can be cut accurately in the shear sense observation plane, which may not be visible in outcrop. Contradictory shear sense indicators are common in outcrop or thin section. There are good theoretical reasons for this, including the fact that material lines rotate in opposite directions under general shear (e.g. Passchier and Trouw 1996). Therefore it is often necessary to evaluate a large number of shear sense indicators for reliability. This may mean examining many thin sections, but a useful technique can be to cut a number of hand specimens on a rock saw in the shear sense observation plane. The shear sense may become visible due to the smooth surface of the saw cut and because it is the correct observation plane, and the procedure allows rapid examination of a large number of specimens. Specimens must be orientated correctly in the field as discussed by Prior et al. (1987) or Passchier and Trouw (1996). Orientation errors can occur in the thin section laboratory, but errors can be checked if all rock fragments created during sectioning are carefully preserved, so that they can be reconstructed to the pre-sawing configuration.

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CHAPTER 7. MICROSTRUCTURAL SHEAR SENSE CRITERIA

ing, as suggested by the observation that the rock has no fabric outside the shear zone. A pre-shear foliation may not exhibit this characteristic curvature, and is a less reliable shear sense indicator.

7.3 Oblique foliations and shape preferred orientations

7.2 Curved foliation Many shear zones have a foliation that curves smoothly across the shear zone as shown in Fig. 7.1. The angle between the foliation trace and the shear plane decreases from about 45° or less to approach zero at the centre of the shear zone (Plate 39). The rotation of the foliation from the outside to the centre of the shear zone (clockwise in Fig. 7.1) is the same as the shear sense of the shear zone (dextral). This criterion is one of the most common and reliable for deducing shear sense, and can be applied to shear zones on any scale. This sort of foliation follows the orientation of the XY plane of the finite strain ellipsoid, and the curvature is due to increased intensity of finite strain towards the centre of the shear zone (Ramsay and Graham 1970). Such foliations are referred to as strain-sensitive by Hanmer and Passchier (1991) and constitute one of the most common and reliable shear sense indicators. An important caveat to the use of such foliations is that the foliation should be formed during shear-

Oblique foliations are grain shape fabrics that maintain an approximately constant, acute angle to the shear plane across a shear zone. The shear plane in this case is often represented by a compositional banding, so that the fabric consists of compositional bands within which there is an oblique foliation. The shear sense is given by the acute rotation from the oblique foliation to the compositional banding (clockwise or dextral in Fig. 7.3). Oblique foliations and the bands parallel to the shear plane have been called and respectively (Law et al. 1984). The geometry of oblique fabrics has strong similarities with some S-C fabrics (see Section 7.5). The oblique foliations described above are examples of shape preferred orientations, which can be used as shear sense indicators by a different method (Shelley 1995). The aspect ratio of the grain long axes is measured and plotted against their angle with respect to an arbitrary reference orientation. Frequency curves can be plotted for the highest and the mean aspect ratios, in class intervals of 5-10°. The angular distribution of aspect ratios is typically skewed, and the shear sense is given by two criteria: the rotation sense from the longer part of the curve towards the highest aspect ratio orientation, and from the mean orientation of all grains towards the orientation of the highest aspect ratio orientation. Shape preferred orientations can result from passive behaviour of grains that behave as strain markers (i.e. they flatten and rotate towards the shear plane as described in Section 7.2), from anisotropic mineral growth (Chapter 5.3), or from rotation of grains behaving as rigid objects in a fluid (e.g. Shelley 1989b). Development of shape preferred orientations may be counteracted by dynamic recrystalliza-


tion of new equant grains; oblique foliations probably reflect a balance between shaping and homogenizing processes which can lead to an equilibrium state or steady-state foliation (Means 1981). These foliations are distinguished as straininsensitive because they do not track the finite strain ellipsoid closely (Hanmer and Passchier 1991).

7.4

Porphyroclast systems

7.4.1 Characteristics and classification A porphyroclast system comprises a porphyroclast with a mantle and tails (or wings) of grains attached and derived from the porphyroclast (Passchier and Simpson 1986). The mineralogy of the mantle and tails may be the same as the porphyroclast (a “mantled porphyroclast”; Passchier and Trouw 1996), or may be related to the porphyroclast by a metamorphic reaction, such as a feldspar porphyroclast breaking down to tails of quartz and mica. Two tails are usually developed around a porphyroclast, symmetrically related to each other by a 180° rotation about an axis parallel to the shear plane and perpendicular to the shear direction. Fig. 7.4 shows four categories of porphyroclast system geometry, three of which are named from their similarity to the Greek letters: and (Passchier and Simpson 1986, Passchier 1994). and have monoclinic symmetry (Figs. 7.5-7.7), while have orthorhombic symmetry. Complex types (Fig. 7.8) consist of more than one tail on each side of the porphyroclast (Passchier and Simpson 1986). The geometry of porphyro-

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68


clast systems can be described with respect to a reference plane that passes through the centre of the porphyroclast parallel to the planar part of the tail far from the porphyroclast, and that contains a symmetry axis (Fig. 7.4; Passchier and Simpson 1986). An important distinction is made between in-plane tails that both lie on the reference plane, and stairstepping tails that lie on opposite sides of the reference plane (Fig. 7.3). types have been subdivided into which are found around isolated porphyroclasts in an homogeneous matrix, and which are are found in S-C or mylonites, with short tails that curve into C- or (Passchier and Simpson 1986). tails were originally defined by the criterion that an individual tail crossed the reference plane near the porphyroclast, and therefore the two tails stair-stepped (Passchier and Simpson 1986). However,the characteristic feature of a is commonly taken to be a deflection near the porphyroclast which produces a feature known as an embayment (Fig. 7.9), so that systems do not necessarily stair step. There is a great deal of variation in the detailed morphology of porphyroclast systems, and all gradations exist between and (Fig. 7.10).

7.4.2

Mechanisms of formation

The formation of porphyroclast systems is now understood to some extent from experiments and numerical simulations (e.g. Passchier and Simpson 1986, Passchier et al. 1993, Passchier 1994, Passchier and Sokoutis 1993, Bjørnerud and Zhang 1995, Ten Brink and Passchier 1995). Tails are con-

sidered to develop by dynamic recrystallization of the porphyroclast to form a mantle of finer grains around the clast, which subsequently become entrained in the flow of the matrix. The geometry of tails depends on the rate of recrystallization relative to the rate of rotation of the porphyroclast, the shape of the porphyroclast, the rheology of the tails and matrix, the shear strain, and the degree to which the tails and matrix adhere to the porphyroclast (the coupling). Many of these factors can be simplified by analyzing the relation between the recrystallizing mantle and the “separatrix”, which is a surface around the porphyroclast that separates closed from open flow lines. The separatrix is either eye-shaped or bowtie shaped in simple shear, depending on a variety of factors including the rheology of the matrix. The shape of the tails is governed by the shape of the separatrix and the location of the separatrix relative to the mantle. A few generalizations can be made from the experimental work of Ten Brink and Passchier (1995): A mantle entirely within the separatrix will not develop tails. A mantle that intersects the separatrix will first form tails which then develop into tails, thus explaining why there are transitional types between the two endmembers. A mantle that encloses the separatrix will first form tails, which develop into type tails for an eyeshaped separatrix, but remain as tails for a bowtie shaped separatrix.

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tails will stair step in a bow-tie shaped separatrix, but not in an eye-shaped one.

7.4.4 Faces of a

tail

The evolution of tails is incompletely understood at present, but these approaches offer the possibility that analysis of tails will yield considerable information on rheology and deformation conditions. There are three simple guidelines for using porphyroclast systems as shear sense indicators. It should be emphasized that shear sense can only be determined from tails with monoclinic symmetry, i.e. and tails.

The asymmetry of an individual tail can be used to define shear sense (Simpson and Schmid 1983). An individual tail has two faces in contact with the matrix around the porphyroclast, one of which is approximately parallel to the shear plane, and the other which is curved and oblique (Fig. 7.9a). Provided the face parallel to the shear plane can be identified, the curvature and obliquity of the other face can be used to give the shear sense: it is convex in the direction of movement of the adjacent matrix, and the acute rotation of the curved face to the shear plane is in the same sense as an oblique foliation. It is usually possible to check this criterion by stair-step direction, but it is sometimes useful if only one tail is well-developed (Fig. 7.5).

7.4.3 Stair-step direction:

7.4.5 Deflection and embayments of tails

tails only develop with high coupling.

and

tails

The stair-step direction can be defined by taking an imaginary walk along one tail towards the porphyroclast in the shear sense observation plane: the stair-step direction is the direction of the step that needs to be taken over the porphyroclast to reach the opposite tail. Thus the tails on the left of Fig. 7.4 step to the left. The stair-step direction defines the shear sense unambiguously: the shear sense is the opposite to the stair-step direction i.e. for a left stair-step direction, the shear sense is right-handed (dextral) and vice versa (Figs. 7.5 to 7.10). This is probably the most reliable way in which porphyroclast systems can be used for shear sense determination.

The deflection of a tail from the distal end towards the porphyroclast can be used to give the rotation sense of the porphyroclast (Passchier and Simpson 1986). The deflection of the tail is in the opposite sense to the rotation of the porphyroclast. For example, the deflection of the tails in Fig. 7.4 is to the left as the porphyroclast is approached, giving a clockwise or dextral rotation and sense of shear. An embayment is the term for the approximately triangular region of matrix between a tail and the adjacent porphyroclast (Fig. 7.9b; Passchier and Simpson 1986). The interface between the porphyroclast system and the matrix in the em-

70

bayment has the shape of a fold: the direction of fold closure is the same direction as the rotation of the porphyroclast (Figs. 7.7, 7.9b). Ideal conditions for the use of porphyroclast systems as shear sense indicators are that the grain size of the matrix is small with respect to the porphyroclast, the fabric of the matrix is homogeneous, the porphyroclast systems were formed in one phase of deformation, the tails are long enough to define a reference plane, and observations are made in the shear sense observation plane (Passchier and Simpson 1986). Porphyroclasts should be spaced widely enough not to interfere with one another. Structures similar to tails can develop in a passive matrix around a rotating porphyroclast, but these have formed in a different manner from the porphyroclast systems described above and should not be used in the same way for shear sense determination (Section 7.10).

7.5 S-, C- and 7.5.1 Characteristics and classification


nis and Secor 1987, 1990). Figure 7.11 shows the essential geometry of these fabrics and the angular relationships between them, which are useful for description and classification. S-surfaces intersect C- or in a line perpendicular to, or at a high angle to, the shear direction, as defined by the lineation on C- or surfaces. Angle is defined as that between the shear plane and the local orientation of S-surfaces between C or and angle is that between the shear plane and which are inclined to the shear plane in the opposite direction to the S-surfaces. C-surfaces are parallel to the shear plane. In addition an angle can be defined, which is the angle between S- and C- or is equal to between S- and C-surfaces, and between S- and S-, C- and have a wide variation in morphology (Fig. 7.12, Plates 40, 41). Two major subdivisions of S-C mylonites were proposed by Lister and Snoke (1984), but problems with this classification have been pointed out (Passchier and Trouw 1996). An alternative approach to description and classification (Blenkinsop and Treloar 1985) uses: 1. The angles

and

2. The structures that define the S-surfaces, Many shear zones contain two or three planar fabrics that can be divided into a sigmoidal penetrative foliation (an S3. The spacing of the C- or and foliation), and discrete planar shears that displace the fo4. The relative strengths of S- and C- or liation, which are known as C- or (Berthé et al. 1979), extensional crenulation cleavage or ecc (Platt This classification is descriptive, yet uses features that have and Vissers 1980), shear bands or shear planes (White et possible genetic significance. Three end members (porphyroal. 1980, Simpson 1986), or normal-slip crenulations (Den- clastic, megacrystic and banded types) identified using the


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localized shear analogous to Y-shears and Riedel shears respectively, as observed in simple shear experiments and fault gouges (e.g. Rutter et al. 1986, Evans and Dresden 1991) and in strike-slip fault systems (e.g. Sylvester 1988). Some may form in the orientation of a Coulomb failure surface. It is likely that anisotropy controls the formation and the orientation of i.e. the value of (Platt and Vissers 1980, White et al. 1980, Platt 1984, Passchier 1984).

scheme are shown in Fig. 7.13 and the relations between and for the different fabric types are summarized in Fig. 7.14. This scheme seems to have wide applicability, based on other published descriptions.

7.5.2 Formation and evolution Variations in the values of and are controlled by strain and microstructure, and can be used to distinguish between models for the formation of S-, C- and A common feature of several models is that S-surfaces are parallel to the XY plane of the local finite strain ellipsoid (Ramsay 1967, 1980, Berthé et al. 1979, Lister and Snoke 1984, Blenkinsop and Treloar 1995). C- and are zones of

If S-surfaces are parallel to the axes of the local finite strain ellipsoid, then the value of should decrease with progressive strain in these domains. However, C- and probably do not not rotate with increasing strain. This is suggested by the planar nature of C- and if significant rotations occurred in a material with discontinuous C- and strain incompatibilities would distort initially planar surfaces (Shimamoto 1989). If decreases and remains constant with progressive strain, then should also decrease, which has been reported in many studies (e.g. Zee et al. 1985, Burg 1987, Ghisetti 1987, Bal and Brun 1989, Scheuber and Andriessen 1990, Rykkelid and Fossen 1992), but is not compatible with steady-state models (e.g. Dennis and Secor 1987, 1990). Volume changes and “stretching” shear zones that extend parallel to their length may be important in the formation of some S-, C-, and (Behrmann 1984, Passchier 1991). Complications in the use of S-, C-, and as kinematic indicators have been pointed out by Behrmann (1987), but the general rules given below (Sections 7.5.3 and 7.5.4) can be used for all fabrics. These relationships may be seen particularly clearly in thin section.

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7.5.3

Curvature of S-foliation

geometry that can be used in the same way as porphyroclasts (Section 7.4.3). It may be difficult to distinguish pressure Many S-foliations curve into like the curved foliation de- shadows from porphyroclasts (Plate 44). Rotation of the porscribed in Section 7.2, and the rotation sense from the S- phyroclast relative to the ISA and deformation of the pressure foliation to the C- or gives the sense of shear re- shadow or fringe complicate this simple geometry. Despite liably (Fig. 7.12, Plates 39, 40). the large variety of intricate structures that may be produced (e.g. Etchecopar and Malavieille 1987, Aerden 1996) there are two reliable methods of determining shear sense from 7.5.4 Shear on C- or more complex pressure shadows and fringes. The sense of shear on C- or is always the same as the overall sense of shear. It can sometimes be determined 7.6.2 Geometry of the last increment of growth independently from the curvature of S-foliations by identifying features such as veins that are offset by shear on C- or The last increment of growth may be virtually undeformed and can therefore indicate the direction of the last maximum ISA. In the case of antitaxial growth, the last increment will be closest to the inclusion, while the opposite will be true for 7.6 Pressure shadows and fringes syntaxial growth. The sense of rotation from the last maximum ISA to the shear plane give the shear sense. Fibre ori7.6.1 Kinematics of pressure shadows and entation in pressure fringes can be used to indicate the ISA only if the fibres are displacement-controlled. fringes in shear zones The formation and geometry of pressure shadows and fringes has been discussed in Chapter 3.9, and here it needs to be re-emphasized that pressure shadows and fringes grow in the volume of low mean stress around the inclusion, which approximately coincides with the direction of the ISA. Therefore in the simplest case of undeformed pressure shadows, the pressure shadow or fringe will be oblique to the shear plane, and the shear sense can be deduced from the acute rotation from the pressure shadow/fringe to the shear plane (Fig. 7.15). Some pressure shadows/fringes have a stair-step

7.6.3

Shape

Fortunately a simple empirical rule for shear sense determination is clear from computer simulations and natural examples, even without detailed considerations of the mechanism of pressure shadow or fringe formation. The shape of the whole pressure shadow or fringe, defined either by its median line or by its enveloping surfaces, forms a doubly-inflected curve which is either an S or a Z shape. S shapes imply dextral rotation of the core objects and therefore dextral shear

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sense, and vice versa (Figs. 7.16, 7.17). One complication that may occur in the use of pressure shadows or fringes as shear sense indicators is that the shadow/fringe may be rotated with the inclusion (White and Wilson 1978). Pressure shadows formed in the extension field may be rotated into the shortening field and therefore give the incorrect shear sense. Longer pressure shadows can come to resemble porphyroclast tails by this process (Hanmer and Passchier 1991).

7.7 Mica fish Mica fish are a distinctive type of porphyroclast consisting of elongated single crystals of mica, often associated with tails (e.g. Lister and Snoke 1984, Hanmer 1986, Passchier and Trouw 1996). Stair-stepping of the tails can be used as a shear sense indicator following Section 7.4, and the sense of the acute rotation from the long axis of the fish to the shear plane also gives the shear sense directly (Plate 42). The formation of mica fish is not well understood, but may involve development of tails by dynamic recrystallization as well as cataclasis, and rotation of the fish. Other minerals (e.g. horn-


75


blende, Plate 43) may show the same geometrical features as understand the porphyroblast growth mechanism, especially whether the porphyroblast is in a domain of coaxial or nonmica fish. coaxial flow.

7.8 Porphyroblast internal foliations Curved foliations within porphyroblasts (internal foliations, usually defined by inclusion trails) have been generally considered to form by rotation and growth of porphyroblasts with respect to the matrix foliation and the flow axes during non-coaxial flow (Fig. 7.18a). However, Bell and coworkers (e.g. Bell 1985, Bell and Johnson 1989, Bell et al. 1992, Aerden 1995) have argued that many form by porphyroblasts overgrowing a new foliation that formed in a different orientation relative to an external frame of reference, and that the porphyroblasts themselves do not have an external vorticity. The latter interpretation leads to the opposite sense of shear to the former (Fig. 7.18). A third possibility for interpreting curved is that a pre-existing foliation rotates during coaxial flow (Fig. 7.18b; Ramsay 1962). A fourth possibility for generating curved inclusions with continuous and is by post-tectonic overgrowth of folds (helicitic structure). The last case can be distinguished by the presence of folds with a similar geometry to the inclusion trail in the matrix away from the porphyroblast. The above ambiguities are sufficient to preclude the use of porphyroblast inclusion trails as a shear sense indicator unless the mechanism by which they formed is well known. It is important to know the regional kinematic framework to

7.9 Crystallographic fabrics Crystallographic fabrics are usually represented by lower hemisphere, equal area stereographic projections of Crystallographic directions of individual grains or sub-grains. It is standard practice to use the shear sense observation plane as the projection plane, with the foliation as a vertical eastwest plane, and the lineation as a horizontal east-west line. This section focuses on the use of quartz c-axis Crystallographic fabrics, which are readily obtained from the universal stage following the procedures in Turner and Weiss (1963) or Passchier and Trouw (1996). Quartz c-axes in non-coaxial shear are known to have an asymmetrical distribution relative to foliation and lineation from experiments, numerical models and natural examples. The asymmetry has been demonstrated as a reliable shear sense criterion (e.g. Behrmann and Platt 1982, Bouchez et al. 1983, Simpson and Schmid 1983, Law 1986, 1987, 1990, Law et al. 1994) but important exceptions are known (e.g. Passchier 1983). Crystallographic fabrics should only be used for shear sense determination when deformation is the result of intracrystalline plasticity, the foliation and lineation are clearly defined and related to the finite strain, the deformation is homogeneous, and the mineral used for the determination is

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Quartz a-axes can also be used for shear sense determination, but they require X-ray measurement. Calcite c-axes develop a girdle in non-coaxial shear that is inclined to the foliation in the opposite direction to the quartz c-axis girdle (e.g. Schmid et al. 1987, Wenk et al. 1987). Asymmetric girdles in calcite determined by X-ray diffraction have proven to be reliable shear sense indicators (Erskine et al. 1993). Plagioclase, olivine and orthopyroxene asymmetric dominant in volume (Bouchez et al. 1983). Crystallographic fabrics have also been used to determine shear sense (Mercier fabrics should always be used in conjunction with other shear 1985, Gapais and Brun 1981, Passchier and Trouw 1996). sense indicators. The crystallographic fabric asymmetry is best defined by the fabric skeleton, which are the lines that connect the max- 7.10 Asymmetric microboudins imum concentrations of c-axes (Behrmann and Platt 1982, Law 1987). Two types of asymmetry are distinguished: ex- Microboudins can be observed in thin section and share the ternal asymmetry is defined by the obliquity of the central same characteristics as the mesoscale examples described in girdle to the foliation (angle in Fig. 7.19; angle nomen- outcrop or experiments. Boudins in non-coaxial deformation clature after Law 1987), and by the difference between the are asymmetric, due either to modification of initially symangles and (Fig. 7.19). Internal asymmetry is defined metrical boudins (types I and II in Fig. 7.20; Hanmer 1986), by the angles and between the arms of the crossed or to development of asymmetric boudins ab initio (type III girdle and the central girdle (Fig. 7.19). The most important in Fig. 7.19, Goldstein 1988). The three types of asymmetric shear sense criterion is the obliquity of the central girdle to boudin can be used as shear sense criteria according to the the foliation (Fig. 7.19). The shear sense is given by the dir- following rules: Type 1 asymmetric boudins have two faces ection of the acute rotation from the girdle to the shear plane. parallel to the shear plane, and two inclined to the shear plane The difference between the values of and or between such that an acute rotation from the inclined faces to the shear and can also be used (Fig. 7.19, cf. Law 1987, 1990, plane gives the shear sense. Types II and III boudins are sepLaw et al. 1994). At high temperatures where prism slip arated on shears that are inclined to the shear plane, and disoperates, a completely different type of fabric forms, consist- place the boudins in the same sense as the overall shear zone ing of a point maximum of c-axes near the lineation. The (analogous to and Riedel shears, Section 7.5, Plate 6). acute rotation from the point maximum of the c-axes to the However, the geometry of boudins in non-coaxial shear delineation is in the opposite direction to the shear sense (Main- pends on the initial shape and orientation of the boudinaged price et al. 1986). layer (Goldstein 1988). Boudins can be separated by shears

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antithetic to the dominant sense of shear, even when layers lie in the extensional field (Fig. 7.21). “Bookshelf sliding” also involves antithetic shearing. Therefore microboudins can only be reliably used for shear sense determination if their initial shape and orientation are known: this may be possible from observations outside shear zones.

7.11

Asymmetric microfolds rolling structures

and

The vergence of asymmetric folds in shear zones is often the same as the dominant shear sense (e.g. Fossen and Hoist 1995), but fold vergence opposed to the shear sense is also well documented (e.g. Krabbendam and Leslie 1996). Oppositely verging folds can be generated by shear of buckle folds in a layer inclined in the shortening direction of noncoaxial flow (e.g. Ramsay et al. 1983, Little et al. 1994) or by heterogeneous simple shear (Fig. 7.22; Krabbendam and Leslie 1996). Folds in simple shear zones may be highly noncylindrical, and this also complicates the use of fold vergence for shear sense determination. Asymmetric microfolds, like asymmetric microboudins, should not be used for shear sense determination unless the way in which they formed is known. Reliability can be enhanced when the vergence of several layers in different orientations can be observed, or when the initial orientation of the folded layer can be determined, for example by observations outside the shear zone. A special type of asymmetric fold is developed by the interaction between rigid inclusions and layering. This type of fold is localized around the inclusion, and has the same vergence as the sense of shear. Such structures developed at high shear strains by interaction between porphyroclast and matrix are referred to as rolling structures (Van Den Driessche and Brun 1987).

7.12

Shear sense criteria in rocks containing melt

7.12.1 Magmatic shear zones Shear planes and shear directions that existed during deformation of melt-bearing rocks may be difficult to identify because they leave no microstructural imprint (e.g. Park and Means 1996). Compositional layering may function as a shear plane because of its rheological anisotropy (Benn and Allard 1989). It has been suggested that magmatic flow planes and directions may correspond to the average orientation of planar and linear shape fabrics respectively (e.g. Guineberteau et al. 1987). However, magmatic foliation defined by megacrysts is not generally parallel to the shear plane (e.g. Paterson and Vernon 1995, Yoshinobu and Paterson 1996) because megacrysts rotate at variable rates depending on their aspect ratios and the rheology of the matrix (e.g. Tikoff and Teyssier 1995). Indeed the first shear sense criterion suggested below is based on the obliquity between grain shape fabrics and the shear plane. Possible dangers of the interpretation of magmatic or sub-magmatic foliations as

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shear planes are pointed out by Tikoff and Teyssier (1995).

3. Clasts are passive markers, but trains do not rotate (“March-fixed train" model).

7.12.2

Oblique grain shape fabrics

Grain shape fabrics of olivine, pyroxene and feldspar in gabbros have been used to determine sense of shear (Benn and Allard 1989) on the basis that they form a fabric at an acute angle to the shear plane. The shear sense is given by the acute rotation from the fabric to the shear plane. However, since elongate grains may rotate past the shear plane, this criterion needs to be used with caution, and a large number of observations should be collected.

7.12.3 Tiling and imbrication Tiled and imbricated clasts have been proposed as shear sense indicators in rocks containing melt (e.g. Blumenfeld 1983, Paterson 1989), and as diagnostic of melt-present, noncoaxial deformation. Imbrication has been comprehensively modelled by Tikoff and Teyssier (1994), who suggested three possibilities corresponding to magmatic, sub-magmatic and solid state deformation respectively:

In all cases, the imbrication of clast trains gives the shear sense according to the rule that the acute rotation from the train long axes to the shear plane is the same as the shear sense. Individual clast long axes can not be used reliably as shear sense indicators because they may rotate past the shear plane in the first two models, even in simple shear. The results of numerical modelling show that clast interaction depends strongly on clast density, and that the March models (2 and 3), possibly corresponding to sub-magmatic and solid state deformation, are more effective at imbricating clasts. Therefore clast imbrication is not diagnostic of melt-present deformation.

7.12.4

S-C fabrics

S-C fabrics have been described in rocks that have no evidence of solid state flow (Blumenfeld and Bouchez 1988, Benn and Allard 1989), and these fabrics can be used as a shear sense indicator as described in Section 7.5. Euhedral, little deformed plagioclase grains in S- and C-orientations have 1. Rigid clasts in a fluid, which form imbricated trains of been described by Miller and Paterson (1994), suggesting that clasts that rotate as a unit when they have coalesced the S-C fabric formed in a sub-magmatic state. (“Jeffrey rotating train” model, after Jeffrey’s (1922) analysis of rigid objects in a fluid). 7.12.5 Sub-magmatic microfractures

2. Clasts behave as passive markers in a fluid; trains rotate as a unit (“March-rotating train” model, after March’s (1932) description of the rotation of passive markers in a fluid). March-like behaviour of rigid clasts is suggested by the experiments of Ildefonse and Manktelow (1993).

Sub-magmatic microfractures can be used as a shear sense indicator. They are inclined to the shear plane in the same orientation as T fractures (Fig. 2.17); the acute rotation from the fractures to the shear plane opposes the shear sense (Bouchez et al. 1992).


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7.13 Shear sense criteria for faults Determining the shear sense (sensu lato) of faults requires observations in the shear sense observation plane which can only be identified if the net slip vector of the fault is known. In practice this almost always requires the presence of slickenlines to identify the shear direction within the fault plane. Incohesive fault rocks are particularly difficult to sample and section. It may be possible to collect intact and orientated specimens by removing them in a metal casing, and to prepare them for thin sectioning by impregnation with adhesive.

been suggested as shear sense indicators by observations on faults with known shear senses in specific conditions (e.g. Petit 1987, Doblas et al. 1997). It is not clear how widespread these features are, and many of them could have ambiguous interpretations on faults with unknown shear sense. Until further work is reported, the only generally reliable fault surface kinematic indicator at present are risers created by slickenfibre terminations (or accretion steps), which are always congruous (Fig. 7.23). When using this criterion, it is useful to examine fault surfaces in more detail under a binocular microscope such as those in common use for micropalaeontology, or under the SEM. This technique is particularly good for dealing with incohesive specimens.

7.13.2 Displaced grain fragments

7.13.4

Small displacements on faults can be measured by matching fragments of distinctive grains in the cataclastic fault matrix to their parents in the walls of the fault, or by matching individual grains on either side of the fault (Fig. 7.23, Plate 45). The sensitive tint plate is useful for this task. However, the sense of shear determined from fracturing of porphyroclasts or boudins is ambiguous, because both synthetic and antithetic fractures can form (Section 7.10).

The characteristic structures of gouges shown in Fig. 2.17 can be used to identify shear sense on a microscopic scale, including P-foliation, Riedel and conjugate Riedel shears, Tfractures and ductile stringers.

7.13.1

7.13.3

Shear sense observations on faults

Risers and slickenfibres

Risers (Section 2.9) on fault surfaces have been used historically to determine shear sense, making the assumption that they are congruous, but the occurrence of incongruous steps is well known. A variety of other fault surface features have

7.13.5

Gouges

Jogs and bends

Compressional or dilational volumes are created at bends or jogs in faults. On a microscopic scale, compression can often be recognized by pressure solution, and dilation by creation of open space or mineral deposition (e.g. Gamond 1987). The shear sense can be determined from the sense of step and the type of deformation: Right-handed bends or jogs are compressional for sinistral faults and dilational for dextral faults. The opposite rules apply for left-handed bends or jogs.

Chapter 8

Shock-induced microstructures and shock metamorphism 8.1

Introduction

A number of distinctive natural microstructures have recently been accepted as characteristic of shock metamorphism on the basis of field observations, detailed microstructural studies, and comparisons with impact experiments. The products and processes of shock deformation are distinct from other geological microstructures and mechanisms, which justifies a separate Chapter for their description. Shock effects have well-established links with meteorite impact structures on earth, which form under pressures ranging from <5 to > 100 GPa, temperatures up to 10 000 °C, and typical strain rates of to (Fig. 8.1; Stöffler and Langenhorst 1994, Reimold 1995). These pressures and strain rates are far greater than values for tectonic crustal deformation (using the term tectonic to imply endogenous processes). There has been vigorous debate about which microstructures are diagnostic of meteorite impact: this is discussed in Section 8.11.

8.2

Shock mechanisms

Four fundamental shock mechanisms have been identified from experimental work that simulates shock conditions: cataclasis, intracrystalline plasticity, solid state phase transformation, and melting (e.g. Huffman et al. 1993, Stöffler and Langenhorst 1994). Cataclasis results in irregular, randomly orientated microfractures, and sub-parallel sets of planar microfractures (Section 8.3). Intracrystalline plasticity is shown by twinning and, possibly, shock mosaicism (Section 8.4), but the formation of the latter is not well-understood. The most distinctive shock-induced microstructures and mechanisms are associated with phase transformation at high compressive stress and with pressure-release melting. Quartz undergoes a phase transformation to an amorphous state in which Si coordination with O changes from four-fold to sixfold under pressures of approximately 5-35 GPa (Stöffler and Langenhorst 1994). The phase transformation is heterogeneous at relatively low temperatures and high strain rates, and is localized on crystallographic planes leading to the formation of one type of planar deformation feature (PDF, Section 8.4). At higher temperatures and strain rates, the amorphization is homogeneous. Quenching following melting forms

a distinctive type of glass called diaplectic glass at moderate pressures (25-50 GPa; 8.6), and a fused glass called lechatelierite at higher pressures (Section 8.8). Coesite and stishovite, the high pressure polymorphs of silica, may be formed during shock metamorphism by solid state transformation or direct crystallization from melt (Section 8.7). These shock microstructures can be classified into high and low pressure types (Table 8.1).

8.3 Microfractures Non-planar, unorientated microfractures are common in shocked materials. Planar, open microfractures (termed planar fractures) parallel to low index crystallographic planes in quartz are a distinctive shock microstructure. They occur parallel to (0001) and planes, with a spacing of (Stöffler and Langenhorst 1994). In experimentally shocked dunite, intragranular microfractures formed in sets of variable orientation with separations of less than and more planar fractures with a spacing of formed in subregions of single crystals (Reimold and Stöffler 1978). Their orientation is crystallographically controlled and a function of the direction of shock wave propagation, and their density increases linearly with shock pressure in a

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CHAPTER 8. SHOCK-INDUCED MICROSTRUCTURES AND SHOCK METAMORPHISM

81

given rock type. PDFs (next Section) do not cross the microfractures, which has been interpreted to indicate that the and PDFs have spacings from 1 to and thicknesses fractures formed earlier than the PDFs during the shock event of under the TEM. Four different types of PDFs can (Stöffler and Langenhorst 1994). be distinguished in the TEM (Goltrant et al. 1991): 1. Brazil twins in the basal plane.

8.4 Planar (PDFs)

Deformation

Features

Planar Deformation Features are intracrystalline zones of optical and/or crystallographic contrast with the host crystal, typically on the order of micrometres wide and with spacings of 2 to as seen under the optical microscope (Fig. 8.2, Plate 46). Quartz is the mineral in which shock characteristic deformation is best developed and has been most extensively studied. PDFs in quartz occur in sets parallel to the crystallographic directions given in Table 8.2 Other orientations include and As many as 18 different sets may form in one grain. The orientations of PDFs in quartz can be related to the shock pressure as discussed in Section 8.10. Planar features with the dimensions of PDFs in shocked quartz are spectacularly revealed in SEMCL images (Seyedolali et al. 1997), though it is not known yet certain that they correspond to PDFs. A fundamental distinction can be made under the optical microscope between decorated PDFs, which are marked by fluid inclusions or pores, and non-decorated PDFs. It is necessary to turn to the TEM to study PDFs in more detail. About ten times more PDFs are visible in the TEM than under the optical microscope (Goltrant et al. 1991),

2. Bands of high dislocation density, with concentrations of voids or bubbles in rhombohedral planes. 3. Bands of variable proportions of glass and crystalline

quartz in the form of microcrystallites ~ 10 nm in size in rhombohedral planes. 4. Glass lamellae in rhombohedral planes.

The basal orientation of the Brazil twins (type 1, Fig. 8.3) distinguishes them from Brazil growth twins which are on rhombohedral planes, and their association with dislocations shows that they are mechanical twins (Leroux et al. 1994). Types 2 and 3 are thought to be secondary features due to post-shock annealing of type 4 PDFs, which are recognized, together with basal Brazil twins, as characteristic of fresh shock microstructures, and are the only type of lamellae found in experimentally shocked material. The glass lamellae can be subdivided into narrow transformation lamellae (< 10 nm wide) which are only visible in the TEM, and wide transformation lamellae (50-500 nm wide) which are the optically visible lamellae, and which may contain a finer scale sub-lamellar structure of pillars at high angles to the lamellae boundaries, which is clearly revealed by etching (Gratz et al. 1996).

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A number of models have been put forward for PDF formation in quartz, which have largely been consolidated in the work of Huffman et al. (1993), Langenhorst (1994), and Huffman and Reimold (1996). Huffman et al. (1993) proposed that PDFs formed by heterogeneous solid-state transformation to amorphous silica along planes of progressively lower compressibility, from basal to m, z, and a orientations, with increasing pressure. Optically visible PDFs are coalesced TEM-scale PDFs in domains between microfaults. Langenhorst (1994) considered the temperature profiles associated with shock fronts to show that, at relatively low pressures (<25 GPa), melts were unlikely to form and PDFs should consist of short-range order (quasi-diaplectic) silica glass formed by solid-state transformation. At intermediate pressures (25-35 GPa), thin bands of melt may form in addition to the solid state transformation, and at 35-50 GPa, large-scale melting during compression may form diaplectic glass. At >50 GPa, melting continues after shock compression to form lechatelierite by quenching under low pressure. This model successfully accounts for the presence of both amorphous silica in PDFs at temperatures too low for melting, and diaplectic glass quenched from a melt. Fresh, impactinduced PDFs are either Brazil twins or amorphous silica lamellae formed by solid state transformation or quenching of melt. Post-shock annealing may have several effects on PDFs. The bands of high dislocation density in quartz (type 2, decorated PDFs) are likely to be due to post-shock annealing (Fig. 8.2). Annealing may be recognized by a distinct type of fluid inclusion, which is typically much smaller (tens of nm) than inclusions trapped during crystal growth (Goltrant et al. 1991), although annealed PDFs may also contain inclusions up to diameter (Leroux et al. 1995, Leroux and Doukham 1996). Fluid inclusions may form by diffusion along dislocations during recrystallization of PDFs, because water is much more soluble in amorphous silica than in crystalline quartz. The tiny crystallites in the interior regions of PDFs (type 3 of Goltrant et al. 1991) are a characteristic annealing texture (e.g. Langenhorst and Clymer 1996). Zircon, like quartz, is a favorable mineral to investigate shock effects because of its transparency, lack of important twinning or cleavage, and uniaxial optical character. Furthermore, it is a very refractory mineral which is resistant to alteration and thermal overprint. Shock features in zircon can be seen after etching in the SEM (Bohor et al. 1993). Planar features continuous across the whole grain occur with a spacing of or less (Fig. 8.4, Kamo et al. 1996). It

is uncertain whether these features are analogous to PDFs in quartz: recent TEM analyses of experimentally shocked zircon single crystals suggests that these features are microcleavage, not PDF-type lamellar defects (Leroux et al. 1998). Zircon grains may also have a granular, polycrystalline texture of zircon crystals (“strawberry texture”), which has been observed in zircon from laboratory experiments and at several confirmed impact sites, including the Vredefort dome, but never in other circumstances: this texture may therefore be diagnostic of impact (Figs. 8.5, 8.6).

8.5 Mosaicism Mosaicism is a pattern of domainal lattice misorientation seen as a mottled extinction pattern, which is distinctly different from undulatory extinction, due to the sub-microscopic size of the domains (Plate 48, Stöffler and Langenhorst 1994). Mosaicism can be described from spot or line broadening observed in X-ray diffraction (XRD) patterns from single crystal grains of quartz. As experimental shock pressures rise, the single crystal XRD pattern becomes increasingly similar to a powder diffraction pattern. The technique can be used to determine that domain size decreases from >3000 nm to <200 nm with pressure. Single zircons with planar features can also show asterism (Bohor et al. 1993). The deforma-


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density (Table 8.3). Langenhorst’s (1994) model for the formation of diaplectic glass suggests that it forms by quenching of a melt before shock pressure is completely released. Its distinctive properties are due to some long range order which is inherited from the solid state. Diaplectic glass in natural impact rocks recrystallizes during annealing to characteristic textures which depend on the shock pressure reached. For shock pressures less than 35 GPa, the glass becomes brown. Microstructures of spherical single crystals, all in the same orientation, called ballen, form at higher pressures. At still higher pressures, the ballen have different crystallographic orientations (Grieve et al. 1996, Plates 47, 48).

8.7 High pressure polymorphs of quartz - Coesite and stishovite Coesite in some shocked samples appears as colourlessbrown, sized aggregates of individual grains less than in size, typically twinned parallel to (010), often enclosed by diaplectic glass or isotropic quartz with abundant remnants of PDFs (Grieve et al. 1996). The aggregates may be aligned on crystallographic planes. Coesite associated with pseudotachylite veins from the Vredefort Dome is colourless, with high relief, and very low birefringence (Fig. 8.7; Martini 1991). It is either massive or forms radiating needles up to long. Coesite has been observed at impact craters and artificial explosion craters, but never in shock experiments: this is attributed to the much shorter duration of the pressure pulse in experiments compared to natural events (Stöffler and Langenhorst 1994). Stishovite, the higher pressure polymorph of silica, can be distinguished from coesite by a higher birefringence and a brownish colour (Martini 1991). It occurs in the same two habits described above for coesite (Fig. 8.8). It appears in TEM as extremely fine-grained aggregates parallel to PDFs in experimentally shocked quartz (Ashworth and Schneider 1985). The difficulty of producing stishovite in shock experiments suggests that relatively long shock durations may be required, and that it may crystallize from a high-pressure melt (Stöffler and Langenhorst 1994). The preservation of stishovite appears to require rapid quenching because it is transformed to amorphous quartz at relatively low temperatures.

tion mechanism responsible for shock mosaicism is unclear. Although mosaicism appears similar to subgrain formation, dislocations may not be mobile on the time scales of shock events: therefore a direct comparison with subgrain formation by recovery may not be valid (Stöffler and Langenhorst 1994). Remnants of planar features associated with mosaicism in quartz suggest that the mosaic domains may have formed by the intersection of two sets of PDFs (Leroux and Doukham 1996). Mosaicism has also been described from 8.8 Lechatelierite other minerals, for example feldspars, olivine and pyroxene (Hörz and Quaide 1973). A patchy optical extinction pattern Lechatelierite is a silica glass which may contain flow strucsimilar to mosaicism in feldspar can be formed by growth tures and vesicles, and may occur as veins within quartz with zoning (Sharpton and Schuraytz 1989). PDFs. The optical properties of lechatelierite are summarized in Table 8.3. The vesicles and flow structures demonstrate that lechatelierite forms as a quench product from a melt under low pressure, in distinction from diaplectic glass. Melting 8.6 Diaplectic glass to form lechatelierite requires pressures in excess of those for Diaplectic glass is a distinctive type of silica glass that oc- whole rock melting, and therefore it is only preserved in veins curs in homogeneous patches and associated with some types where pressures were locally increased by shock reverberaof PDF. It can be distinguished from either synthetic quartz tion, or as inclusions in whole rock melts or tectites (Stöffler glass or lechatelierite (Section 8.8) by optical properties and and Langenhorst 1994).

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8.9


Tectites, microtectites and spherules

MgO and ratios, and higher Cr, Ni, and Co (Koeberl 1990). These characteristics are similar to sediments.

Tectites occur over large areas of the earth’s surface called strewn fields, demonstrating that they were deposited from the atmosphere. Terrestrial and lunar volcanism, and meteorite impact, have been proposed as possible origins for tectites, but a number of arguments favour the impact hypothesis (Glass 1990). Flow structures demonstrate that tectites 1. Splash forms: spheres, disks, dumbell and teardrop were molten: the flanges of ablated forms were formed durshapes. ing a second period of melting. Lechatelierite inclusions suggest temperatures of 2000°C, far greater than volcanic tem2. Ablated forms, which are similar to splash forms but peratures. The chemical composition of tectites shows that with an additional flange. they were derived by melting of sediments and not from terrestrial or lunar igneous rocks. Furthermore, the tectites in 3. Muong Nong types, which are larger, layered chunks of two strewn fields can be linked to proven impact craters: for tectite glass. example, tectites in Czechoslovakia came from the Ries crater Tectite glass has flow structures and abundant inclusions of in Germany, while the Ivory Coast strewn field is derived from the Bosumtwi crater in Ghana. lechatelierite. Baddeleyite, coesite, metallic spherules, and corundum may occur as inclusions, as well as quartz, zirMicroscopic tectites (microtectites) have been described con, rutile, chromite and monazite in the Muong Nong types from several sites at the Cretaceous-Tertiary boundary, and (Glass 1990). The chemistry of tectites is similar to felsic vol- associated with impact at the Chicxulub site in Mexico (e.g. canic glasses, with contents of >65%, but with higher Smit et al. 1992, Olsson et al. 1997). Many of these miTectites are rounded silicate glass bodies usually less than a few centimetres in diameter (Glass 1990). They are usually black, but may be translucent, brown or green. Three major types are recognized:


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crotectites are spherical, with diameters of 0.2 to 5 mm, and to mosaicism and PDFs, followed by diaplectic glass, coescomposed of clays or calcite. Some of these spherules contain ite and stishovite, and lechatelierite at the highest pressure. a core of glass, suggesting that the clays formed by devitrific- However, the accurate pressure calibration of this general seation. Flow structures in the glass, and associated shocked quence is subject to the effects of a number of other variables grains, demonstrate the impact origin of these spherule beds. (8. 10.3) which are not yet quantified with the exception of the initial temperature of the target rocks and porosity. Figure 8.9 shows onset pressures for shock microstructures 8.10 Shock barometry and thermo- in quartz and feldspar based on experiments on single crystals of quartz or non-porous crystalline rocks. The lines showing metry decrease in onset pressure with temperature were determined by experiments on granite and quartzite (Reimold and Hörz Establishing shock pressures has fundamental implications 1986, Huffman et al. 1993, Huffman and Reimold 1996). for establishing the origin of shock microstructures and is Microfracturing occurs at the lowest pressures: planar microuseful to estimate the rate of shock pressure attenuation for fractures form at pressures greater than 5 to 10 GPa (Huffman modelling impact processes. Shock pressures have been et al. 1993). Mosaicism in quartz is first seen from pressures calibrated experimentally using microstructures, the density of 6 to 10 GPa. The onset pressures for PDFs in these experiof quartz, and the optical, X-ray diffraction, spectroscopic ments fall between 15 and 18 GPa, but this is well above figand thermoluminescent characteristics of quartz (Stöffler and ures given by other workers, who suggest formation of nonLangenhorst 1994, Grieve et al. 1996). Shock temperatures basal PDFs (i.e. not Brazil twins) at pressures as low as 10 can be calculated in non-porous rocks from shock pressures GPa. These lower pressures are shown on Fig. 8.2 by labels (e.g. Langenhorst 1994). In the following Sections, two giving the orientations of the dominant PDFs (after Stöffler methods of shock wave barometry using optical microscopy and Langenhorst 1994, Grieve et al. 1996). The nature and (microstructures and optical properties of quartz) are briefly orientations of PDFs in quartz changes as a function of presdescribed. sure. Brazil twins form at the lowest pressures, followed by PDFs. PDFs are generally considered to form at pressures of ~20 GPa, and become the dominant orientation at ~25 8.11 Calibration of shock pressures GPa (e.g. Langenhorst and Deutsch 1994). Widths of experimentally created PDFs at the TEM scale change in thickness from microstructures from <100 nm at <25GPa to ~200 nm at >25 GPa. IniMicrostructures form with increasing shock pressure in a tial (i.e. target) temperature has an effect on PDF developsequence that has been clearly established by experimental ment. Approximately equal numbers of grains show single work. In quartz, from the lowest pressure, fracturing evolves and multiple sets of PDFs in experiments at 18 GPa and 250C

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CHAPTER 8. SHOCK-INDUCED MICROSTRUCTURES AND SHOCK METAMORPH1SM


(Huffman and Reimold 1996). The number of grains with PDFs decreases as temperature increases, and multiple sets are preferentially eliminated; at 750°C, there are very few grains with single sets of PDFs in samples shocked at this pressure. The shape of the PDFs also changes from narrow with sharp boundaries at low temperatures to broad with wavy boundaries at higher temperatures. There is some evidence for a change from dominantly to orientations as temperature is increased from 25°C to 440°C (Huffman et al. 1993). Quartz phase may affect the character of PDFs: samples shocked at 20-25 GPa and 540°C showed dominantly PDFs, compared to a broad range of PDF orientation at 630°C, corresponding to the change from to stability in the target rocks (Langenhorst et al. 1992, Langenhorst and Deutsch 1994). Experimentally shocked and are also distinguished by the observation that positive and negative planes of the same form never coexist in samples shocked at less than 573°C, the transition temperature, as expected from the trigonal symmetry of This is in distinction from PDFs parallel to the hexagonal pyramids of samples shocked in the stability field, and has been used as a geothermometer for the pre-shock temperature of the target rocks: for example the temperature of the Ries crater and Eocene-Oligocene impact targets can be constrained to less than 573°C (Engelhardt and Bertsch 1969, Langenhorst and Clymer 1996). A pressure calibration for porous rocks (Table 8.4) has been established from work on the Barringer Crater in Arizona (Kieffer 1975, 1981, Kieffer et al. 1976). In general, the onset pressure for any microstructure decreases with increasing porosity, because stress concentrations can occur on the grain scale. PDFs seem to be less abundant in porous targets than crystalline ones, and there is some evidence that different orientations occur in porous rocks: and may be important (e.g. French et al. 1974). A sequence of microstructures similar to those in quartz and feldspar with increasing shock pressures has been established in olivine and pyroxene. Planar fractures and mosaicism develop in olivine at 10 to 15 GPa, intergranular brecciation at 30 GPa, and glass, annealing and melting occur above 50 GPa. In pyroxenes, twinning occurs at 5-10 GPa, mosaicism and PDFs at 20-25 GPa, and glass, melting and recrystallization above 75 GPa (Huffman and Reimold 1996). Details of the techniques for measuring PDF orientations are described in Engelhardt and Bertsch (1969), Stöffler and Langenhorst (1994) and Grieve et al. (1996). There are various methods of assigning an average shock pressure to orientation measurements from a sample, which should consist of as many grains as possible. According to the method of Robertson (1975), individual grains are assigned the value of the highest pressure PDF observed within the grain according to some experimental calibration. The shock pressure is taken as the average of individual grain pressures. Variations on this method which account, respectively, for grains with no PDFs and grains converted to diaplectic glass, have been proposed (Fel’dman 1994). Mosaicism, the first appearance of diaplectic glass, 100% diaplectic glass, and lechatelierite, as well as the existence of coesite and stishovite can all be used as quartz barometers

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from Fig. 8.7. The effect of temperature on the high pressure phases is not well known, however. Microstructures in feldspar show a similar pattern with the notable difference that mosaicism occurs at considerably lower pressures. The best pressure estimates are given by the coexistence of two or more microstructures.

8.11.1

Calibration of shock pressures from optical properties of quartz

The refractive indices and birefringence of quartz at room temperature decrease continuously from samples shocked at 25 GPa to those shocked at 35 GPa, which corresponds to the transition to diaplectic glass (Fig. 8.10). The refractive index of diaplectic glass decreases slightly as pressure increases further to 50 GPa. The effect of increasing pre-shock temperature is to restrict the pressure interval over which the decreases occur: at 630°C, the change occurs between 25 and 26 GPa (Stöffler and Langenhorst (1994).

8.11.2

Problems of shock barometry

Experiments have also shown that strain rate, pulse duration, and rate of pressure increase may also be important variables that should be considered in an accurate calibration of shock pressure (e.g. Huffman et al. 1993, Huffman and Reimold 1996). Different calibrations between experiments with different pulse lengths suggest that this is a factor which has not been adequately explored, and raises the problem of extrapolation from experiment to nature (Huffman and Reimold 1996). Grain size, grain size distribution, porosity and modal mineralogy may also have effects (e.g. Robertson and Grieve 1977, Reimold and Stöffler 1978, Grieve et al. 1996). Basal PDFs appear to be more resistant to post-shock annealing than other types of PDF (Leroux et al. 1994), rendering shock pressure calibrations from PDFs unreliable where there has been post-shock annealing. Other barometers may be questionable in this situation: for example, diaplectic glass is a metastable phase and may not be preserved. Stishovite

88


inverts in only 3 days at 250°C in the presence of water, and in impact rocks, so that the third criterion is not general. Multhus has limited preservation potential (Gigl and Dachville tiple, intersecting sets of tectonic lamellae are also known 1968). (e.g. Lyons et al. 1993), but these lamellae are distinctly different to PDFs, mainly because they are non-planar (Reimold 1994). Several studies of PDF orientations in shocked targets 8.12 Diagnostic impact microstruc- show that criterion 4 is false: in particular, shock-induced basal PDFs (Brazil twins) are common at low shock prestures sures. Straightness (planarity), and sharpness or definition remain important characteristics for identifying PDFs, but have The identification of microstructures that are diagnostic of not been quantitavely defined: this lack of objectivity has impact has profound implications for the understanding of the evolution of the earth. The subject has been very controver- caused considerable controversy in the past. Optical studies alone appear to be incapable of fully distinguishing impact sial, with much debate on the extent to which shock-induced from tectonic lamellae. microstructures could be produced by volcanic eruptions or SEM and TEM-scale features do allow a clear distinction other endogenous processes as alternatives to meteorite imbetween fresh PDFs and tectonic deformation lamellae. Etchpact. PDFs have been associated with shock metamorphism on ing with HF liquid or vapour followed by SEM examination the basis of experiments and observations of their distribu- is a more widely available technique than the TEM (Gratz et tion in suspected impact target rocks (e.g. McIntyre 1962, al. 1996). Relatively wide, glass-filled PDFs from impacts Carter 1965, Hörz 1968, Robertson and Grieve 1977). How- and shock experiments are revealed as sharp, planar features wide after etching, often containing sublamelever, PDFs can be superficially similar to deformation lamel- less than lar structure (see Section 8.4). By contrast, etched tectonic lae formed by tectonic processes as described in Section 4.6, lamellae are much wider than and often sinuous, similar and this raises the controversial questions of whether PDFs can be diagnostic of shock metamorphism, and whether shock to their appearance under the optical microscope (e.g. Joreau et al. 1997; Fig. 4.11). Table 8.5 summarizes the differences features can be produced in volcanic eruptions. Commonly used criteria for identifying impact-induced between shock-induced PDFs and deformation lamellae in quartz. Post-shock annealing reduces PDFs to a dislocation PDFs include the following (e.g. Alexopoulos et al. 1988): structure similar to deformation lamellae, making the distinc1. Well-defined, sharp features. tion between the two difficult (cf. Goltrant et al. 1992). Basal Brazil twins have only ever been observed in quartz from 2. Straight. impact structures and in experiments with deviatoric stresses over ~4 GPa at room temperature (Leroux et al. 1994). These 3. Often occur in multiple sets per grain. observations were used by Leroux et al. (1994) to argue that basal Brazil twins are strong evidence for shock. 4. 90% are in the and directions. Stöffler (1972) suggested that mosaicism is the most comHowever, single sets of PDFs are developed in shock experi- mon manifestation of shock. However, mosaicism is not ments under higher temperature conditions and are observed diagnostic of shock effects, and may also be caused by tec-


tonic deformation or even chemical effects, especially in feldspars (e.g. Sharpton and Schuraytz 1989, Officer and Carter 1991). The occurrence of coesite is known from inclusions in high-pressure metamorphic rocks and kimberlites, but the very fine-grained, polycrystalline nature of coesite in impact related rocks is diagnostic (Grieve et al. 1996). Diaplectic glass can be formed by static compression at 25-35 GPa, but this is much higher than any pressures recorded by rocks at the surface of the earth (Fig. 8.1): diaplectic glass remains a diagnostic indicator for shock pressures. Explosive volcanism and deep-seated explosive activity are possible alternatives to meteorite impact as a source of shock waves. Controversy exists over what pressures may be obtained in volcanic explosions; estimates range from values

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within the field of crustal metamorphism (0.5 GPa) to values as much as 10 GPa. Planar microstructures, mosaicism and Brazil twins in volcanic rocks have been claimed as shock microstructures (e.g. Carter et al. 1986, 1990, Huffman and Reimold 1993), but much of this evidence has been strongly refuted (Sharpton and Schuraytz 1989). The case for shock features related to volcanism or deep-seated explosive activity is not strong. In summary, Brazil twins, glass-filled PDFs, diaplectic glass, the distinctive very fine grained, polycrystalline occurrence of coesite, and shock features in zircon are diagnostic of impact shock. As stressed by Reimold (1994), a thorough evaluation of the case for shock metamorphism should examine a variety of minerals and evidence on macroscopic to microscopic scales.

Chapter 9

From Microstructures to Mountains: Deformation Microstructures, Mechanisms and Tectonics 9.1

Introduction

9.2

Failure criteria

9.2.1

Coulomb and Mohr failure criteria

Deformation microstructures and mechanisms can be used to deduce conditions of deformation, because they imply particular relations between stress, strain rate (and possibly strain), temperature, fluids, microstructure, and mineralogy. These relations are derived from laboratory experiments and theories of material deformation. There are three types of relationship: failure criteria (Section 9.2), which give stress conditions for faulting, frictional sliding laws (Section 9.3), which give stress or stress-strain rate conditions for frictional sliding, and flow laws (Section 9.4), which give stress-strain rate relationships for flow accommodated by intracrystalline plasticity, diffusive mass transfer (DMT), or composite deformation mechanisms. Relationships between flow laws are usefully illustrated on a deformation mechanism map (Section 9.6) which shows fields of stress, temperature and grain size in which particular mechanisms give the highest strain rate. The rheology of the lithosphere is often shown by a plot of stress against depth, or a lithospheric strength envelope (Section 9.7).

The simplest failure criterion, known as the Coulomb criterion, is that failure occurs when the shear stress on a plane equals a material property called the cohesion and the normal stress multiplied by the coefficient of internal friction,

Although failure criteria and flow laws are essential tools in tectonic interpretation, they are based on laboratory experiments which have serious limitations when extrapolated to nature (e.g. Carter and Tsenn 1987, Paterson 1987). Two of the most important problems are the limited scale of triaxial tests, which means that mesoscopic heterogeneities in rocks are not accounted for, and the fast strain rates of experiments (typically to ) compared to most tectonic deformation ( to ), which mean that the deformation mechanisms may differ between nature and experiment. Microstructures offer a potential solution to these problems: if deformation microstructures and mechanisms can be identified in nature that correspond to experiments, then laboratory results can be extrapolated with more confidence. This Chapter provides some basic failure criteria, frictional sliding laws, and flow laws to be used in conjunction with microstructural observations to make tectonic interpretations. The primary intention of the Chapter is to provide quantitative data based on the concepts in the previous Chapters.

3. High porosity sedimentary and some low grade metamorphic rocks.

The criterion in terms of the principal stresses is:

where is the uniaxial compressive strength. The value of is around 0.58 for many rocks, while varies widely for different rock types. A useful generalization of rock types in terms of their uniaxial compressive strengths is given by Paterson (1978): 1. Igneous and high-grade metamorphic rocks.

2. Low-porosity sedimentary and low-medium grade metamorphic rocks.

4. Low porosity dolomites and quartzites. However, the relationship between shear and normal stress of real rocks is not linear as suggested by the Coulomb criterion. The Mohr criterion describes the relationship by an empirical fit to data from rock mechanics experiments, and so is limited to materials that have been extensively tested. The Griffith criterion (next Section) is a better approximation to rock behaviour in most situations.

9.2.2 Griffith failure criteria Griffith’s thermodynamic approach to failure outlined in Section 2.2.1 leads to a failure criterion of the form:

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CHAPTER 9. FROM MICROSTRUCTURES TO MOUNTAINS

is the uniaxial tensile strength. This simple criterion is widely applicable, but limitations are the difficulty of measuring (which is dependent on the type of test used) and the observation that in most rocks, the uniaxial compressive strength is 10 to 12 times the value of (Jaeger and Cook 1979). The Griffith criterion predicts that Values of for use in the Griffith criterion can be calculated from the generalized values of given in Section 9.2.1. The simple Griffith criterion does not predict any effect of on failure: this was accommodated in the Extended Griffith criterion by Murrell (1963), giving the failure criterion:

and may also depend on the elastic moduli of the solid and pore fluid (Bernabé 1987, Warpinski and Teuffel 1993). generally lies between the value of porosity and 1; typical values for reservoir rocks are 0.5-0.8 (e.g. Detournay et al. 1989, Kumpel 1991). is approximately 1 for many crystalline rocks. Departures from this value occur when the pore fluid has a chemical effect, when permeability is low enough to prevent uniform pore fluid pressures during experiments, and when porosity is not interconnected (Paterson 1978).

9.4 This criterion implies that in better agreement with observed values. A more rigorous treatment of triaxial stress acting on a body containing randomly orientated ellipsoidal cavities leads to a failure criterion which is independent of and implies lower values of than the Griffith failure criterion (Murrell and Digby 1970). The effect of frictional sliding on closed Griffith flaws with a coefficient of friction and stress required for closure was dealt with by McClintock and Walsh (1962) in the Modified Griffith criterion:

is a function of Poisson’s ratio and the axial ratios of the flaws: for penny-shaped flaws and for biaxial stress. The value of is (Murrell 1965); a value of 100 MPa was used for in lithospheric failure modelling by Kusznir and Park (1984, 1987). If is taken to be negligible, this criterion becomes equivalent to the Coulomb criterion and implies that which generally has little relationship to the observed values of (Jaeger and Cook 1976). However, this criterion does provide a better fit to stresses measured at compressive failure than either the original or extended Griffith criteria.

9.3

Pore fluid pressure and faulting

The principle of effective stress (Terzaghi 1943) describes the important mechanical effect of pore fluid pressure on failure in many experiments (e.g. Paterson 1978). The principle states that normal stresses are reduced by the value of the pore fluid pressure, so that effective stresses can be substituted in any of the failure criterion above. The effective stresses (“conventional effective stresses” of Paterson 1978) are calculated from:

Fracture mechanics and failure criteria

Shear failure occurs by linking of axial tensile microcracks (Section 2.4). This fundamental observation has been incorporated into failure criterion (e.g. Peng and Johnson 1972) through the application of fracture mechanics (e.g. Lawn and Wilshaw 1975a, Nemat-Nasser and Horii 1982, Horii and Nemat Nasser 1985, 1986, Kemeny and Cook 1987, Hallam and Ashby 1990, Baud et al. 1996). Unfortunately the resultant expressions for failure are complex and depend on the geometry and size of the initial flaws, which are not readily measured microstructural properties.

9.5

Frictional sliding laws

9.5.1 Byerlee’s law Amonton’s Law (Section 2.2.2) describes frictional behaviour of rocks well at low normal stresses, but at higher normal stress, a threshold shear stress must be applied before any movement occurs: this is called the cohesion, so that frictional sliding in rocks is described by:

The form of this equation is the same as used for the Coulomb criterion; however, the cohesion and coefficient of friction of intact rock (the coefficient of internal friction) are different from those of frictional sliding. Byerlee (1978) discovered that values of in most rocks are in the range of 0.5-1, with an average of approximately 0.75, and that frictional sliding in most rocks could be described by two simple equations, depending on the normal stress.

These two equations have become known as Byerlee’s law, and are approximately independent of rock type, roughness More detailed experiments and poroelasticity theory indicate of the sliding surface and temperature. Byerlee’s Law is ofthat the normal stress is not reduced by the exact value of the ten assumed to describe the state of stress in the upper crust pore fluid pressure. The effective stresses are given by (e.g. Burov and Diament 1996) which implies that failure in the upper crust is governed by frictional sliding of previously formed faults in ideal orientations for slip. Marone et al. (1992) have proposed that distributed dewhere is known as the effective stress coefficient and has values less than or equal to 1 for elastic properties. deformation in a gouge zone has a separate frictional sliding creases with permeability, porosity and confining pressure, law, because stress is re-orientated within the sheared layer

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so that the maximum possible shear stress is supported. This localization on shear surfaces, which always has a negative results in a sliding law of the form: velocity dependence (Marone et al. 1992). Seismological observations support the above picture, but there is a lack of corresponding knowledge about natural microstructures produced by seismic slip. Pseudotachylites where is the coefficient of sliding on a discrete (Section 2.11) are the only direct evidence for seismic slip surface. (e.g. Sibson 1977), but seismic slip can also be tentatively suggested by textures such as crack-seal vein fillings which indicate repetitive fracture followed by fracture sealing (Sec9.5.2 Rate and state dependent frictional slidtion 3.12). Particle size distributions of gouge may give an ining dication of its mechanical behaviour: velocity strengthening behaviour changed to velocity weakening when the fractal diDynamic effects during frictional sliding, such as stick slip mension of the particle size distribution of the gouge attained behaviour, are not described by Amonton’s and Byerlee’s a value of 2.6 in experiments (Biegel et al. 1989). Another Laws. Stick-slip behaviour involves a steady rise in stress potential microstructural indicator of seismic slip is the evidwith no movement followed by a sudden rapid episode of ence of continuous deformation alternating with discontinustress drop and sliding, and is analogous to earthquake faultous deformation in slickenside surface material (Power and ing. Stick slip behaviour is favoured over stable sliding in Tullis 1989). rocks with higher quartz contents, by sliding along rougher surfaces, by the lack of gouge, by high normal stresses, and low temperatures (e.g. Paterson 1978). Dynamic frictional sliding can be modelled by making the coefficient of friction in Amonton’s Law depend on the rate of frictional sliding V:

where is the coefficient of friction at a reference velocity V*, a and b are empirically determined constants, and is a state variable (e.g. Rice and Gu 1983). L is the critical slip distance for sliding to return to a stable rate after a change in sliding velocity. It is possible to expand the above expression with more than one state variable. The value of a expresses the dependence of on a change in velocity, and b expresses a change in that decays with time. This behaviour is found to describe laboratory experiments very well. The stability of a sliding process depends on the value of if this is negative, increasing velocity will lead to a reduction in and the fault is said to show velocity weakening, and hence be capable of nucleating earthquakes. On the other hand, if is positive, velocity changes will lead to increased friction and the sliding will be stable. The rate dependent friction law can be applied to the nucleation of earthquakes in the crust (Tse and Rice 1986, Scholz 1990). is expected to be positive in the upper few km of well-developed faults due to the presence of a layer of gouge (which is characterized by positive values of ). Therefore earthquakes are not expected to nucleate in the top part of the crust except in areas lacking well-developed faults. Below the level of gouge, there is an upper stability transition to negative values of which are the conditions for earthquake nucleation (Scholz 1990). becomes positive with increasing temperature, so that there is also a lower stability limit below which earthquakes will not nucleate: this is probably approximately at the 300°C isotherm in continental crust, which explains the observation that earthquakes generally nucleate in the top 15 km of the crust (e.g. Sibson 1982, Meisner and Strehlau 1982). However, this picture is complicated by the relation between velocity dependence and displacement: positive velocity dependence at small displacements changes to negative dependence with increasing displacement, due to

9.6

Flow laws

9.6.1

Diffusive mass transfer: Grain size sensitive creep

The fundamental flow laws for deformation accommodated by diffusive mass transfer can be derived from theoretical principles for three cases: diffusion through a grain boundary fluid (pressure solution, Section 3.2), diffusion through the grains (Nabarro-Herring creep, Section 5.2), and diffusion through grain boundaries (Coble creep, Section 5.2). All flow laws for diffusion-accommodated deformation have one distinguishing characteristic: strain rate is a function of grain size. Pressure solution creep Diffusive mass transfer through a solution requires three steps: solution, diffusion and precipitation. The strain rate may be controlled by the rate of diffusion in the fluid, the rate of solution or precipitation at the source or sink, or the rate of diffusion at the source or sink (Paterson 1995), and the flow law depends on which process is rate-limiting. The following flow law can be derived theoretically for diffusion control (e.g. Spiers et al. 1990, Lehner 1990):

where is the strain rate A is a dimensionless parameter varying from 40 to 140 that depends on a geometrical model for the arrangement of the grains, is the molar volume is the thickness of the grain boundary fluid (m), is the reference state solubility of the solid in the fluid (mole fraction), is the reference state diffusion coefficient in the grain boundary fluid is the activation enthalpy for pressure solution R is the gas constant T is the temperature (K), and is the differential stress (Pa). The reference state for and is It is conventional to quote flow laws for the uniaxial

93


shortening strain rate under axisymmetric stress, and the differential stress relations between other components of the stress and strain rate tensors can be obtained from appropriate geometrical modifications. The terms are sometimes gathered together as a phenomenological coefficient representing effective grain boundary diffusivity (e.g. Spiers et al. 1990). Diffusion control is relevant for salt in which solution rates are very fast (e.g. Spiers et al. 1990), and has been reported for experiments on quartz (e.g. Gratier and Guiget 1986). This flow law is only valid for dense polycrystals: porosity is a very important factor in pressure solution creep, which requires a specific geometrical model (e.g. Lemée and Guegen 1996). Paterson (1995) presents a model for pressure solution creep in porous materials based on an analogy with particulate flow in soil mechanics, leading to equations for each of the three possible rate limiting processes. Diffusion-limited pressure solution creep in quartz has been described by the following flow law (modified by Angevine and Turcotte 1983 after Rutter 1976):

Nabarro-Herring creep The flow law for Nabarro-Herring creep is similar to the previous law, with the following important differences: the grain size dependence is to the square of d there is no grain boundary thickness to take into account, and the diffusion coefficient and activation enthalpies are different:

where is the reference state volume diffusion coefficient, and is the activation enthalpy for volume diffusion. Parameters for the flow law are given in Table 9.1. A flow law for diffusion creep that combines both Coble and Nabarro-Herring creep (Ashby 1970, Raj and Ashby 1971) is:

The value of A may differ in all the above expressions since it depends on the preferred geometrical model of the grains. where A is in is the effective normal stress in Pa, is the solubility of quartz at the reference state in and is the density in A flow law for the solution/deposition-controlled case is given by :

where kC is the velocity of the deforming crystal face e.g. Raj 1982). Solution/deposition controlled pressure solution has been described in quartz (Gratier and Jenatton 1984), but seems to be less common than diffusion control. Paterson’s (1995) model suggests that for quartz and rocksalt aggregates, source/sink diffusion control is likely to be rate limiting under geological conditions, but in high-temperature laboratory conditions, precipitation control is effective for quartz. Experimentally determined values for the relevant parameters are given in Table 9.1. Coble creep The flow law for Coble creep has a very similar form to that for pressure solution creep since both processes involve diffusion around grain boundaries. The major differences are the lack of a term for solubility, and differences in the values of the diffusion coefficient and activation enthalpy:

is the reference state diffusion coefficient for grain boundary diffusion, is the grain boundary width, and is the activation enthalpy for grain boundary diffusion. The specification of is difficult because it depends on the diffusing species. The strain rate will however be limited by the diffusion rate of the slowest species, which is normally quoted. In olivine, the values of the diffusion coefficients increase in the order Si < O < Fe; the value for Si is rate limiting. Values for Coble creep flow laws are given in Table 9.1.

9.6.2

Intracrystalline plasticity

Two different flow laws can be formulated from theory to describe flow controlled by dislocation glide (abbreviated to dislocation glide; Section 4.2) and flow controlled by dislocation climb, known as dislocation creep (Section 4.2). Dislocation glide Theoretical flow laws for dislocation glide are characterized by an exponential relationship between strain rate and stress, known as the Dorn Law, of the form:

where and are the strain rate and stress at 0 K, and F is a constant (e.g. Dorn 1954). Dislocation glide is not considered to be important in most minerals under geological conditions because of the relatively high stresses required for the process. Dislocation creep At lower stresses, higher temperatures and slower strain rates, a power law relationship describes the strain rate-stress relationship: where A is a pre-exponential constant, is the activation enthalpy for dislocation creep, and n is the stress exponent (Weertman 1968). A depends on the volume diffusion coefficient and the Burgers vector of the active dislocation system. Unfortunately there are severe problems in using this simple flow law. One of the most important variables that is not explicitly taken into account is the effect of water. For example, “Dry” quartz, with H contents of less than Si, is an order of magnitude stronger than “wet” quartz (the

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considerable amount of literature on this subject is summarized in Paterson and Luan 1990). The effect is probably due in part to replacement of Si-O-Si bonds by weaker Si-OH-OHSi (hydroxyl) bonds, known as hydrolytic weakening (Griggs and Blacic 1965, Griggs 1967), which is similar to the mechanism that controls sub-critical microcrack propagation (Section 2.2.1.1). Hydrolytic weakening enables easier dislocation glide and creep, although other mechanisms may also play a role in the weakening effect of water. This suggests that the flow law should incorporate a function of the water activity (e.g. Kronenberg and Tullis 1984, Ord and Hobbs 1985). Experiments show that the power law exponent depends on water content, at least for low water contents (e.g. Blacic and Christie 1984). In any event, it is difficult to determine water contents at the time of deformation, although Kronenberg et al. (1990) have shown how quartz water contents measured across a natural shear zone vary with shear strain, possibly reflecting a variation during deformation. Mean stress is another variable that should appear in the flow law, which may have important rheological effects by increasing water diffusion rates, increasing water solubility, and increasing activation enthalpy. Quartz phase appears to affect both the activation enthalpy ( is higher for than and the power law exponent n which is lower for (e.g. Linker and Kirby 1981). The quartz phase affects diffusion coefficients in a remarkably similar way to activation enthalpies of power law creep, showing that the creep enthalpies are essentially the same as the enthalpies of diffusion (Gilletti and Yund 1984, Dennis 1984). Trace amounts of impurities in quartz may also cause considerable weakening, probably through their control on glide velocity (e.g. Jaoul et al. 1984).

9.6.3 Empirical flow laws from experimental data


ation energy P is the mean stress (Pa), V is an activation volume often ignored due to its relatively small contribution to the exponential term), and m and n are general dimensionless grain size and stress exponents. The values of m and n suggest interpretations of the deformation mechanism from experimental data by analogy with the theoretical relationships: and indicates dislocation creep, and indicates pressure solution or Coble creep, and and indicates Nabarro-Herring creep. Superplasticity is characterized by and (e.g. Boullier and Guegen 1975, Schmid et al. 1977). Tables 9.29.4 are a compilation of some results for minerals and rocks, divided into grain-size sensitive Table 9.2), and grain-size insensitive flow laws ( Table 9.3 for minerals and Table 9.4 for rocks). When using these tables or any such rock mechanics data, it is advisable to consult the original reference to find any important specific conditions of the experiments which may limit the applicability of the flow laws. At high stress levels, the power law relationship does not fit the experimental data (“power law breakdown”, Carter and Tsenn 1987). Two alternative empirical relationships may describe the data well:

is an exponential law, with B and as empirical constants, suggesting that flow controlled by dislocation glide is important (Tsenn and Carter 1987). An expression that fits the intermediate and high stress behaviour is :

e.g. Heard (1963), with C, constants.

9.7

and n as the empirically-derived

Polymineralic deformation

Experimental data can be fitted to empirical laws of the folWhile flow laws such as those in Section 9.4.3 can describe lowing general form for low stress levels: the bulk rheology of rocks, microstructural evidence shows that strain is distributed in a complex manner between minerals with different rheologies, which should be accounted for where A is a pre-exponential constant or in a detailed understanding of deformation. An empirical apdepending on the units of ), Q is an activ- proach was taken by Tullis et al. (1991), who derived a flow


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96 law for a bimineralic rock that weighted the flow law parameters of each mineral by the volume fractions of the weak and strong minerals as follows:

where n, Q, and A are the parameters in the dislocation creep flow law (Section 9.4.2.2), and the subscripts and s are for composite, weak, and strong minerals respectively. This gives a reasonable fit to experimental data. A more fundamental way to treat polymineralic rocks is to model the relationship between the behaviour of different minerals. Handy (1990, 1994) has suggested that for two minerals of contrasting strength, the microstructure may be described as a load bearing framework (LBF) if the stronger mineral supports stress around pockets of a weaker mineral, or as an interconnected weak layer (IWL) if the weaker mineral forms a matrix around microboudins of the stronger mineral. The behaviour of the composite depends on the contrast on rheological properties of the minerals, and on their volume fraction. Quantitative observations can be used to distinguish LBF and IWL microstructures. In granite deformed under amphibolite facies, for example, the LBF case may be characterized by similar grain shapes and sizes in all minerals, and strain magnitudes, as estimated approximately by grain shapes, are low (Schulmann et al. 1996). Quartz has large recrystallized subgrains, and K-feldspars are fractured, while plagioclase recrystallizes. There is a transition to an IWL with increasing strain, which is manifest from higher strain in quartz than feldspar, and the appearance of S-C fabrics. The ultimate stage of the process is a transition to another type of IWL microstructure in which there is little contrast between the quartz and feldspar grains: all minerals have recrystallized completely in a mylonite. Two flow laws are necessary to describe the LBF and IWL cases separately, which can be derived from considering that the total rate of viscous strain energy dissipation is the sum of the rates in the weak and strong minerals. In the LBF case, both minerals deform at the same rate and the shear strength of the composite is given by the sum of the shear strengths of each mineral multiplied by their volume fractions: where and are the shear strengths and volume fractions of the weak and strong minerals respectively. For the IWL case,

where a function x is introduced that describes the sensitivity of the strain rate partitioning in the rock to the strength contrast of the minerals. The nature of this function can be suggested by considering limiting values. One possibility suggested by Handy (1994) is:


The distinction between LBF and IWL behaviour occurs at the conditions where both give the same bulk strength. On either side of this equality, the more stable microstructure is the one that dissipates less strain energy. For a quartzfeldspar composite with dislocation creep flow laws, the LBF microstructure is only stable for very low volume fractions of quartz. This model shows that strength of the composite always lies between the limiting strengths of the strong and weak phases, but the relative strengths of the composite compared to the end members are a complex function of their strength contrasts and their volume fractions. In a granite, 10-20% volume fraction of quartz reduces the composite strength by one half compared to the strength of pure feldspar, assuming IWL behaviour, and that the deformation occurs in the conditions appropriate for dislocation creep flow laws. Quartz effectively governs the strength of the composite at volume fractions over 10%. Feldspars in gabbros significantly weaken the composite strength from a pure pyroxene rock up to 900°C: at higher temperatures, pyroxene is weaker than feldspar, and the composite has an intermediate strength. Both the Tullis and Handy approaches outlined above are limited to two phase rocks. A generalization to any number of phases is given by Ji and Zhao (1993) for two extreme assumptions. If all phases are loaded equally (isostress criterion), the flow law is given by taking the volume average of the strain rates of individual phases:

where n, A, and Q are the flow law parameters, the subscripts r and i refer to the isostressed aggregate and the individual phases (total number m and volume fractions f) respectively. An alternative assumption is that all phases have the same strain and strain rate. This isostrain criterion leads to the following aggregate properties:

An average flow law is considered to lie between the physically unrealisitic extremes of the isostress and isostrain cases. The average is determined by minimising the difference between two composite flow laws, that each allow for variable, complementary proportions of the isostress and isostrain flow law parameters. The results achieve reasonable fits to some experimental data and the method is advantageous when the effects of more than two phases need to be considered. In another approach to the deformation of polymineralic rocks, Ji and Zhao (1994) have used a fibre loading model to derive an expression for the elastic solution to the problem of rigid fibres in a weaker matrix, and generalized from the elastic solution to a viscoplastic case. The equation they derive includes an expression for the aspect ratio of the fibres. Their model describes some experimental results better than the Tullis model, but they point out that it assumes uniformly


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distributed fibres with a constant orientation, which may limit haviour of the lithosphere and have important applications to its geological application. Flow laws for polymineralic rocks large-scale tectonic problems (e.g. Burov and Diament 1995, require more theoretical and experimental research. 1996). The upper part of the LSE is usually constructed using Byerlee’s Law for frictional sliding to represent cataclastic 9.8 Deformation mechanism maps behaviour. Coulomb or Griffith failure criteria can also be used to model the behaviour of intact lithosphere (cf. SibThe rheology predicted by the flow laws in Section 9.4 can son 1983, Kusznir and Park 1987). A power law for dislocabe compared on a deformation mechanism map (Fig. 9.1), tion creep is usually assumed for the lower, plastic part of the which shows the relationships between deformation mech- LSE, justified by observations that suggest that lithospheric anism, stress, strain rate, temperature, and grain size (e.g. stress/temperature/grain size conditions are appropriate for Stoker and Ashby 1973, Ashby and Verrall 1978). These this deformation mechanism. Dislocation glide can also be maps may be shown in plots of stress normalized by the included (e.g. Burov and Diament 1995). shearmodulus against temperature normalized by meltThe LSE is based on a model of the compositional structure ing temperature (homologous temperature, ). The shear of the earth. Gross simplifications are necessary due to the modulus normalization allows for the systematic dependence limited availability of rock mechanics data, and the difficulty of the flow law constant A on and for the temperature deof applying them to the earth. The upper crust is commonly pendence of the modulus (cf. Tsenn and Carter 1987). The modelled by wet quartzite, the mid to lower crust by feldspar melting temperature normalization allows for better comparor diorite, and the mantle as olivine. A temperature-depth ison between different materials because transitions between profile is chosen for the crustal model. Appropriate steady different rheological regimes depend on These norstate geotherms can be calculated from surface heat flux (e.g. malizations are relevant for materials-science applications, Kusznir and Park 1984, 1987), but Burov and Diament (1995) but the direct values of stress and temperature are more useful have modelled an evolving geotherm due to cooling after orofor tectonic studies. geny. The fundamental assumptions for constructing these maps The LSE is delineated by the lesser of the cataclastic or are that deformation will occur by the mechanism that plastic strengths. The stresses can either be calculated for provides the fastest strain rate, and that steady state conditions a given strain rate, or by applying a constant stress at the are obtained. Deformation mechanism fields in stress and boundaries of a model that analyses the stress and strain with temperature space are contoured for strain rate using the flow time. The latter approach shows that stress applied uniformly law of the deformation mechanism with the fastest strain rate; at the boundaries of the lithosphere is redistributed into the boundaries between fields correspond to loci of equal strain more competent layers which deform elastically, a process rate. Different maps are constructed for different grain sizes; referred to as stress amplification (Kusznir 1982). Eventualternatively grain size can be added as a third dimension. ally the elastic strength of all layers is exceeded, leading to The possibility of contributions to strain rate by more than whole lithosphere failure, and deformation continues by fricone mechanism can be allowed for (e.g. Stoker and Ashby tional sliding and intracrystalline plasticity (Kusznir and Bott 1973). For example, both diffusional creep mechanisms can 1977, Kusznir and Park 1982, 1984, 1987). Hence steady be combined as in the last equation in Section 9.4.1.3, and difstate LSEs have no part that is controlled by elastic strength. fusion and dislocation processes can also combine. However, The shape of the simplest LSE (for a lithosphere of homoflow by dislocation glide and dislocation creep are mutually geneous composition) consists of a rapid linear increase in exclusive. strength with depth from the surface of the earth governed by Deformation mechanism maps for geological materials have been produced for only quartz, olivine and calcite (e.g. the high normal-stress dependence of Byerlee’s Law or a failRutter 1976, White 1976, Carter and Tsenn 1987), because a ure criterion (e.g. the top part of Fig. 9.2). The cataclastic range of flow laws for different mechanisms are not available part of the LSE depends on the orientation of the failure surfor other minerals or rocks. Microstructures are a guiding face: differential stresses for faulting or frictional sliding at principle to the construction and use of deformation mechan- any level of the crust increase in the order normal - strikeism maps: similarity in experimental and natural microstruc- slip - thrust fault. The strength of the lithosphere reaches tures and mechanisms allows for the correct choice of flow a maximum where the cataclastic strength is equal to the flow stress for intracrystalline plasticity. Below this point the law parameters (e.g. Busch and Van der Pluijm 1995). flow stress is less than the cataclastic strength, and continues to decrease with depth as temperature increases. LSEs for layered models of the earth generally show the above pattern 9.9 Lithospheric strength envelopes repeated in each layer (Fig. 9.2). There are dramatic strength A Lithospheric Strength Envelope (LSE) shows the relation- contrasts across layers, especially between the crust and the ship of lithospheric strength with depth derived from elasti- mantle. This suggests that detachment could occur at these city, failure criteria, frictional sliding laws and flow laws layer boundaries, which are sometimes misleadingly called (Fig. 9.2). The term yield strength envelope is commonly “brittle-ductile transitions”, but are more accurately referred used (e.g. Goetze and Evans 1979), but since the upper part to as brittle-plastic transitions (Section 1.3, 1.5). However, saw-tooth LSEs such as those in Fig. 9. 2 can of the envelope is governed by cataclastic failure, LSE is a better term. LSEs are a powerful way of describing the be- be criticized because the validity of Byerlee’s law has not

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been established at mid to lower crustal pressures, and the high stresses predicted by extrapolation of the law to these conditions are unrealistic (Ord and Hobbs 1989). Moreover, the semibrittle deformation regime is not recognised, probably because of the paucity of published flow laws. More realisitic possibilities for the shape of the LSE at mid crustal depths are either a vertical line of constant stress with a value equal to the stress at the base of the seismogenic zone (Hobbs et al. 1986), or non-linear, pressure-dependent functions derived from Mohr-Coulomb constitutive laws for permanent deformation (Ord 1991). The experimental basis for the latter laws is very limited, and they have the disadvantage of predicting strain-dependent values of stress.

9.10

Palaeopiezometry

9.10.1 Methods and calibration Palaeopiezometry is the determination of past stress fields, which can provide major constraints on tectonic models. Microstructural methods have concentrated mainly on measuring differential stress (Sections 9.10.2-9.10.6), and the ori-

entation of the stress tensor (Sections 9.10.7-9.10.8). Palaeopiezometers are calibrated by experiments and theory, but there are large variations in calibrations of the same palaeopiezometer. Furthermore, different palaeopiezometers yield considerably different values when applied to the same rocks. The latter variations can potentially lead to further tectonic insights (Section 9.10.9). Determination of the mean stress is considered in the separate Section on geothermobarometry (Section 9.11).

9.10.2

Recrystallized grain size

The size of grains recrystallized during deformation (d, mm) can be related to differential stress ( MPa) by an expression of the form: where m and are constants, which can be determined experimentally or derived theoretically (e.g. White 1979). Values of m and are given in Table 9.5, and the calibrations for quartz are plotted in Fig. 9.3. The differential stresses for quartz predicted by different calibrations differ by up to four orders of magnitude, for at least seven possible reasons,


99

100


where and are calibration constants, is the shear modulus, and is the most common Burgers vector. Values of 1 and 2 for have been proposed on theoretical grounds (Twiss 1977). Another possible relationship proposed by Twiss (1986) is:

which are worth examining in detail because they highlight the problems of recrystallized grain size palaeopiezometry. 1. Recrystallized grain size appears to depend on the recrystallization mechanism: subgrain rotation (regime 2) gives smaller grains than grain boundary migration (regime 3) (e.g. Drury et al. 1985, Drury and Urai 1990, Hirth and Tullis 1992). 2. The importance of the effect of water is obvious from comparing the dry and wet calibrations of Ord and Christie (1984). However, it does not appear to be important in olivine (Van der Wal 1993). 3. The possibility that the recrystallized grain size is sensitive to temperature is suggested by some data on metals and olivine (e.g. Mercier et al. 1977, Ross et al. 1980), and predicted by theory (Mercier 1980a).

4. The possibility that the recrystallized grain size is dependent on quartz phase is suggested by the flow law dependence on phase (9.4.2.2).

where C and are constants, implying that there is a stable subgrain size at zero stress. However, the second relationship does not fit data for olivine and quartz better than the first. This may be due to the lack of accurate, low stress measurements (Twiss 1986). Until these are available, empirically derived constants for the first relationship are probably the most satisfactory calibration of the subgrain size palaeopiezometer, as given in Table 9.6. Problems with the application of this palaeopiezometer include a possible temperature dependence, the sensitivity of subgrain size to water known for olivine (Twiss 1986, Van der Wal 1993) and the problem of subgrain size measurement. Subgrain sizes may be measured under the optical microscope in reflected light after etching (see Ord and Christie 1984 for the method) or by electron microscopy. Optically determined subgrain sizes are commonly an order of magnitude larger than subgrain sizes measured by electron microscopy (Ord and Christie 1984). Available data at present may not allow reliable extrapolation outside the range of subgrain sizes used to calibrate the palaeopiezometer, since alternative calibrations fit the experimental data almost equally well, yet give enormous differences when extrapolated (Twiss 1986).

9.10.4

Dislocation density

The dislocation density written as:

relationship can be

where k and u are constants, and and are defined as above. 5. Kinetic effects may severely hamper determination of A theoretical value of u is 0.5 (Weathers et al. 1979), but the calibration constants if equilibrium grain size is not measured values have a range from 0.45 to 3.33 (Twiss 1986). attained during experiments (Twiss 1977). The lack of An alternative form of the relationship sometimes used is: such equilibrium for grain sizes on the order of mm is strongly suggested by the experiments of Kronenberg and Tullis (1984); hence recent experiments have been on much finer grain sizes (e.g. Post and Tullis 1999). where D and are constants, implying a steady state free dislocation density. However, data for olivine and quartz 6. Equilibrium grain size may not be achieved if grain are not fitted any better by the second relationship. Table 9.7 boundaries are pinned by impurities (e.g. Evans and presents calibration constants for quartz and olivine using the White 1984). first expression. Problems with this palaeopiezometer are 7. Recrystallized grain size has been measured by various possible temperature dependence and the difficulty of meastechniques and may be specified by different parameters: uring dislocation density accurately. Invisibility of dislocaa standard practice is to measure the mean linear inter- tions under the TEM may cause an underestimate in dislocacept grain size and multiply by 1.5 to allow for truncation density of 30% (Ord and Christie 1984). tion and sampling effects (e.g. Christie and Ord 1980). However, this will not be correct for grains with a fabric.

9.10.5 Twinning - differential stress

9.10.3

Subgrain size

Subgrain size of the form:

can be related to stress by an expression

Three different approaches have been proposed to relate twinning in carbonates to differential stress. Two of these (Jamison and Spang 1976 and Laurent et al. 1981, 1990) take no account of the fact that grain size is known to have an effect on the twinning, and are therefore less suitable than

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the approach of Rowe and Rutter (1990), who have calib- a non-twinning deformation mechanism. In principle, grain rated three twinning palaeopeizometers from laboratory ex- size distribution could have an effect on this palaeopiezoperiments, which all appear to be independent of temperature, meter because stress distributions in grains are strongly instrain rate and strain. fluenced by the size of neighbouring grains (Newman 1994), but this has not yet been evaluated or demonstrated. This palaeopiezometer has been calibrated only for calcite (other Twinning incidence, carbonates twin at different stress levels), and since the calis the proportion of grains with optically visible twins in ibration experiments were carried out at 400°C and above, it any grain size class. It can be related to differential stress is probably not appropriate for low temperature deformation (MPa) and grain size (d, mm) by the following equation: (Burkhard 1993). The best results will be obtained in samples that have experienced a single, coaxial strain event (the conditions of the experimental calibration), otherwise the method In practice, stresses are calculated from for each grain size will overestimate stress (Rowe and Rutter 1990). There is some experimental evidence that twinning in class and plotted against grain size: they should lie on lines of constant slope for different values of The standard error pyroxene could be used as a palaeopiezometer: Tullis (1980) suggests that the minimum stress needed for twinning in in the experimental calibration was 31 MPa. pyroxenes is 100 MPa. Twin density, is the number of twins per mm. This is calculated by the slope of the relation between number of twins in a grain size interval and grain size. appears to be independent of grain size, and is related to with a standard error of 43 MPa, by:

Maximum twin volume, is the maximum volume proportion (%) of twinned material in a grain size class. In the experiments, the area fraction of twins was taken as equal to the volume fraction, and was measured from a two-dimensional thin section. is related to with a standard error of 41 MPa, by :

As with the previous palaeopeizometers, there are a number of potential problems. Clearly if twinning is the only deformation mechanism, the amount of twinning measured by any of the three parameters will be a function of strain. The method thus assumes that after twinning, equilibrium is reached with

9.10.6 Deformation lamellae The appearance of deformation lamellae (Section 4.7) in experiments is characteristic of flow laws with an exponential relationship between stress and strain rate, which occur at relatively high stresses. This observation leads to the suggestion that there may be a critical differential stress for the formation of deformation lamellae, dependent on the type of bonding and crystal structure of the material (Blenkinsop and Drury 1988). Values of 100-200 MPa have been suggested for quartz. However, it is possible that the critical stress may be inversely temperature-dependent, so that at present deformation lamellae in quartz should only be taken as qualitative indicators of relatively high stress levels (Drury 1993). Spacing of deformation lamellae (s, mm) may constitute a palaeopiezometer (Koch and Christie 1981, McLaren 1991). The relation derived by Koch and Christie is ( MPa):

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103

9.10.7 Principal stress orientations from de- (e.g. Turner 1953, Weiss 1954). Twinning in calcite occurs on the e planes in the direction of the plane containformation lamellae ing the pole to the twin and the c-axis (Fig. 9.4). and are therefore at 45° to the twin plane in the plane of the c-axis and the pole to the twin - a similar geometry to that used for the quartz deformation lamellae method. However, in contrast to that method, is in the opposite quadrant from the c-axis (Fig. 9.4). A similar construction can be applied to f twins in dolomite, but here is in the same quadrant as the c-axis (Fig. 9.4). A development of this technique is to use the right dihedra method to determine the and orientations that are compatible with the largest numbers of twins. The value of the maximum number of compatible twins (the MAX number) gives a measure of the degree to which the data can be fitted by a single stress orientation (Pfiffner and Burkhard 1987). This method has the advantage that it does not assume a fixed angle between and the twin plane. A further development is the determination of principal stress orientations and absolute magnitudes by assuming homogeneous stress distribution and a critical resolved shear stress for twinning (Laurent et al. 1981, 1990). It is possible to calculate a finite strain tensor by incorporating the shear strain necessary for twinning into the orientation analysis (e.g. Groshong 1972, 1974, 1984). A computer program to carry out the analyses is described by Evans and Groshong (1994). The measurements necessary for the determination include the c-axis orientation, twin set orientation, average thickness and number of twins, and grain width. 25 grains in two perpendicular thin sections should be measured. Twins which do not fit the calculated strain or stress fields are a problem in all these methods. Two basic philosophies to deal with this problem have been adopted. The first is to recognize that some twins will inevitably develop in incompatible orientations due to stress inhomogeneity. The proportion of such twins (negative expected values, NEV, Groshong 1972) can be used to test the homogeneity of the sample, and the incompatible grains can be removed from the data set. 9.10.8 Principal stress orientations and strains Alternatively, the incompatible twins are presumed to reflect from twins different superimposed stress fields, and they can be used to Twins have been used to deduce principal stress orientations separate out a number of different stress fields (e.g. Lacombe by assuming that the twin plane is a shear plane and the twin et al. 1990). This method has been criticized by Burkhard direction is parallel to the maximum resolved shear stress (1993) because it has yielded unrealistic results, and the asTechniques of using deformation lamellae in quartz to deduce principal stress orientations were developed by Carter and Friedman (1965) and Carter and Rayleigh (1969). Of the three methods proposed by Carter and coworkers, the “arrow” method is the easiest and yields consistent results. The orientations of the c-axis and deformation lamellae are first measured from individual quartz grains. The plane containing the pole to a lamella and the c-axis is constructed (Fig. 9.4). and are located within this plane at 45° to the lamella plane. is distinguished from by the fact that the c-axis lies closest to An arrow from the c-axis to the lamella pole will point from to (Fig. 9.4). Unique and orientations can be derived by eigenvector analysis from a number of such measurements (e.g. Spang and Van der Lee 1975). This construction implies that the deformation lamellae form in a plane of maximum resolved shear stress with a maximum shear stress direction parallel to the projection of the c-axis onto the lamellae plane. The evidence presented in Section 4.7 shows that recovery is an integral part of lamellae development, so that a simple analogy with a slip system for the formation of lamellae may not be entirely valid. However, the strength of the method is based on its empirical success (Pavlis and Bruhn 1988, Drury 1993), even if its theoretical basis remains unclear. The principal stress axes from deformation lamellae are in good agreement with those calculated from the carbonate twin method from the same samples (e.g. Spang and Van der Lee 1975). It is possible to consider other angular relationships between and the lamella plane, but some results suggest that the stress axes do not change appreciably from the 45° constraint (Spang and Van der Lee 1975). It is also possible to analyze for the principal stresses without using a fixed angular relationship between and the lamella plane, by using the right dihedra technique (cf. Angelier 1984).

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for subgrain sizes and perhaps 30% for recrystallized grain sizes; dislocation densities may therefore record later tectonic events with low strains such as stress during uplift (e.g. White 1979a). This problem can potentially be overcome by plotting stress levels derived from one palaeopiezometer against another. If the stresses are recorded by both palaeopiezometers from the same part of the stress-time path, they will lie on a line with a unit slope, while deviations from this line indicate records from different parts of the stress path.

9.11

Geothermobarometry

9.11.1 Methods and calibration The determination of past temperature and mean stress (geothermobarometry) is traditionally the domain of thermodynamics and mineral chemistry. However, there are a number of promising new developments in the application of microstructures to geothermobarometry, which are particularly robust because they are calibrated from naturally deformed samples. In addition to these new, quantitative approaches, experiments and thermodynamic results allow some generalizations about the minimum temperatures for plasticity. Plasticity in the form of twinning can occur even at room temperature in calcite, but appears to require a temperature of 300°C in dolomite. Plasticity (sensu lato) in quartz and mica is generally restricted to temperatures greater than 200°C. The onset of plasticity in feldspars is commonly taken as 450°C, but this value is not well constrained. The common mafic minerals (clinopyroxene, amphibole, orthopyroxene and olivine) may be plastic from 500°C or above. These numbers must however be treated with great caution, because there are many other variables involved in the mechanism transitions (notably stress/strain rate, see Section 1.3), and because of the problems of extrapolating from experiments to nature (Section 9.1).

9.11.2

Calcite twin morphology

Twin morphology in calcite appears to be strongly temperatsumption of stress homogeneity is clearly false in a polysized ure dependent, and has been proposed as a geothermometer grain aggregate. (Burkhard 1993). Four types of twins can be distinguished (Fig. 9.5), together with appropriate temperature ranges for deformation: 9.10.9 General problems with palaeopiezomet-

ers

Type I. Thin (1 mm), straight, rational twins form at temperatures of less than 200°C. A general problem with all palaeopiezometers is the interpretation of the results. The significance of the stress level recorded can only be evaluated in the context of an evolving stress Type II. Thicker, straight and rational twins (> 1 mm), which are slightly lens shaped, form at 150-300°C. field, analogous to the problem of interpreting results from geothermobarometry within a P-T path. Interpretations become especially problematic when considering results from Type III. Thick, curved, irrational twins, which may contain twins within twins, form at temperatures greater than the same sample that differ according to the method used. For 200°C. example, stresses measured by dislocation density, subgrain size and recrystallized grain size may differ because they have variable dependence on strain magnitude. Dislocation dens- Type IV. Thick, irrational and patchy twins, breaking up into trails of small grains, form at temperatures over 250°C. ities may be reset after less than 1% strain, compared to 5%


9.11.3

105

Sutured quartz grain boundaries

were limited to to and maximum strains of 31%. Ultimately the validity of a temperature - only dependThe geometry of sutured grain boundaries (Section 4.8) ence of D rests on the empirical evidence. formed by grain boundary migration depends on the temperature during deformation. With increasing temperature, the length of segments, serrations or lobes increases (Fig. 9.6, 9.11.4 Subgrain boundary orientation in Kruhl and Nega 1996). The geometry of a grain boundary can quartz be characterized by its fractal dimension, D, which is best The orientation of subgrain boundaries in quartz appears to measured by the “divider” method (Kruhl and Nega 1996). be controlled by the phase present during deformation (SecA grain boundary is divided into linear segments (“strides”) tion 4.6). The presence of chessboard patterns, indicating of individual length r. The length of the grain boundary L both basal and prism-parallel sub-grain boundaries, is restricis equal to the product of the number of segments and their ted to deformation in the field (Kruhl 1996). This length. The process is repeated over a range of values of r. means temperatures above 573°C at 0 MPa, and 825°C at For a fractal grain boundary, L is related to r by: 1000 MPa (Gross and Van Heege 1973). Chessboard patterns can only be seen in grains with c-axes subparallel to the plane of the section (high interference colours) and must be distinD is derived from the gradient of a log-log plot of L against guished from three types of pseudochessboard pattern (Kruhl r. Kruhl and Nega (1996) established the relation shown 1996). in Fig. 9.7 empirically from measurements of natural su1. Two sets of prismatic subgrain boundaries visible in tures with known temperatures of deformation. The relation grains with c-axes at high angles to the section (i.e. low between T (°C) and D is approximately: birefringence).

D values of sutures produced in experimental deformation appear to be affected by the experimental strain rate as well as temperature (Takahashi et al. 1997). However, the extent to which strain rate may be important in natural deformation is unknown, especially because the experimental strain rates

2. Rectangular kink band boundaries that are not crystallographically orientated. 3. Two sets of rhombohedral subgrain boundaries in grains with c-axes subparallel to the section, which can be distinguished by their inclined extinction positions.


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Abbreviated references collected by chapter Italics indicates important general sources for the chapter topic. Chapter 1. Beaumont et al. 1996, Byerlee 1968, Carter & Kirby 1978, Chester & Logan 1987, Evans et al. 1990, Griggs & Handin 1960, Kirby & Kronenberg 1984, Knipe 1989, Kusznir & Park 1987, Lawn & Wilshaw 1975a, Lister & Snoke 1984, Molnar 1992, Patterson 1978, Ramsay & Huber 1987, Rutter 1986, Sibson 1977, Snoke et al. 1998, Williams et al. 1994, Wise et al. 1984.

123

Krantz 1979, 1983, Krantz & Scholz 1977, Kulander & Dean 1995, Labuz et al. 1987, Lawn & Wilshaw 1975a, b, Lespinase & Pêcher 1986, Lin & Williams 1992, Lindquist et al. 1984, Logan et al. 1981, Maddock 1986, 1992, Maddock et al. 1987, Magloughlin 1989, Maltman 1994, Marone & Scholz 1989, Masch et al. 1985, Mase & Smith 1984, McEwen 1981, Means 1987, Menéndez et al. 1996, Meredith 1983, Michalske & Frechette 1980, Mogi 1965, Morrit et al. 1982, Narahara & Wiltschko 1986, Norris & Barron 1968, Norton & Atkinson 1981, Olgaard & Brace 1983, Olsson & Peng 1976, Passchier & Trouw 1996, Passchier et al. 1990, Peck 1983, Peng & Johnson 1972, Petit 1987, Pittman 1981, Platt & Vissiers 1980, Power & Tullis 1989, Quakenbush & Frechette 1978, Rice 1968, Robertson 1983, Rudnicki 1980, Rudnicki & Rice 1975, Rutter & Hadizadeh 1991, Rutter et al. 1986, Sammis & Biegel 1989, Sammis et al. 1987, Schofield & Worth 1968, Scholz 1972,1990, Seyedolali et al. 1997, Shand 1916, Shimada 1986, Shimamoto & Nagahama 1992, Sibson 1980, Sibson et al. 1975, Smith 1984, Smith & Stenstrom 1965, Spray & Thompson 1994, Spray 1987, 1988, 1989, 1995, Sprunt & Nur 1979, Sprunt et al. 1978, Stearns 1968, Stel 1981, Swanson 1980, 1981, Tapponier & Brace 1976, Teufel 1981, Tija 1967, Tullis & Yund 1987, 1992, Underhill & Woodcock 1985, Vollbrecht et al. 1991, 1994, Wang 1987, Wang & Liou 1991, Wang & Scholz 1994, 1995, Wang et al. 1989, Wenk 1978, Whitney 1996, Will & Wilson 1989, Willaime et al. 1979, Wong 1990, Wong & Biegel 1985, Wong & Wu 1995, Zhang et al. 1990, Zhao & Johnson 1991. Chapter 3. Aharonov et al. 1997, Alvarez et al. 1976, Andrews & Railsback 1997, Bathurst 1958, Beach 1979, Becker 1995, Bell & Cluff 1989, Bennema & Van Der Eerden 1987, Borradaile et al. 1982, Brantley 1992, CarrioSchaffhauser & Gaviglio 1990, Carrio-Schaffhauser et al. 1992, Casey 1995, Cox 1987, Den Brock 1996, Den Brock & Spiers 1991, Dietrich & Grant 1986, Durney & Ramsay 1973, Elliot 1973, Erslev & Ward 1994, Fisher & Anastasio 1994, Fletcher & Pollard 1981, Gratz et al. 1991, Gray 1977, Green 1980, Groshong 1976, 1988, Guzetta 1984, Hedlund et al. 1994, Heidug 1991, Hickman & Evans 1991, Hillner et al. 1992, Hippert 1994, Knipe 1979, Knipe & White 1977, 1979, Lespinasse & Cathelineau 1995, Lewis & Holness 1996, Manktelow 1994, Marlow & Etheridge 1977, Masuda & Mizuno 1995, McEwen 1981, McCaig 1987, Murphy 1990, Onasch 1994, Passchier & Trouw 1996, Petit & Matthauser 1995, Powell 1979, Price & Cosgrove 1990, Railsback & Andrews 1995, Raj 1982, Raj & Chyung 1981, Ramsay 1980, Ramsay & Wood 1973, Ramsay & Huber 1983, 1987, Robert et al. 1995, Roedder 1984, Rutter 1976, 1983, Schutjens 1991, Smith 1964, Smith & Evans 1984, Spiers & Schutjens 1990, Tada & Seiver 1986, 1989, Tada et al. 1987, Tullis et al. 1996, Urai 1983, Urai et al. 1986, 1991, Watson & Brennan 1987, White & Knipe 1978, Williams 1972, Wilson 1994, Wintsch & Dunning 1985, Wood 1974, Wright & Platt 1982, Wright & Henderson 1992.

Chapter 2. Agar et al. 1988, Allison & La Tour 1977, Ameen 1992, Anderson & Grew 1977, Antonellini et al. 1994, Atkinson 1982, Aydin & Johnson 1978, 1983, Aydin 1978, Bansal 1977, Baud et al. 1996, Biegel et al. 1989, 1992, Blenkinsop 1989, 1991, Blenkinsop & Rutter 1986, Blenkinsop & Sibson 1991, Boldt 1995, Borg & Maxwell 1956, Borg et al. 1960, Borradaile 1981, Brace & Bombolakis 1963, Brown & Macaudiere 1984, Brown & Scholz 1985, Bruner 1984, Byerlee 1968, Camacho et al. 1995, Carter & Kirby 1978, Chopin 1984, Conrad & Friedman 1976, Cooper et al. 1989, Costin 1983, Cox & Atkinson 1983, Cruickshank et al. 1991, D’Arco & Wendt 1994, Das & Scholz 1981, Dunn et al. 1973, Engelder 1974, Evans 1988, Evans & White 1984, Friedman & Logan 1970, Gallagher 1981, Gallagher et al. 1974, Griffith 1924, Grocott 1981, Hadizadeh 1980, Hadizadeh & Rutter 1983, Hadizadeh & Tullis 1992, Hancock 1985, Hippert 1994, Hirth & Tullis 1991, Horii & Nemat-Nassar 1985, 1986, House& Gray 1982, Hull 1988, Inglis 1913, Jaeger & Cook 1979, Jamison Chapter 4. Allison & LaTour 1977, Blenkinsop & Drury & Stearns 1982, Kemeny & Cook 1987, Knipe 1986, 1989, 1988, Christie & Ardell 1974, Den Brock & Spiers 1991,

124

Drury et al. 1985, Drury 1993, Gleason et al. 1993, Hirth & Tullis 1992, Hobbs et al. 1976, Hull 1975, Jessel 1988a, b, Jessel & Lister 1990, Kruhl 1996, Lloyd & Freeman 1991a, b, 1994, Mainprice & Nicolas 1989, Mainprice et al. 1986, McLaren 1991, McLaren et al. 1967, Means 1981, Means & Dhong 1982, Nicolas & Poirier 1976, Spang & Van der Lee 1975, Twiss 1974, Urai et al. 1986, Wenk & Christie 1991, White 1976. Chapter 5. Behrmann 1985, Bell et al. 1992, Borradaile et al. 1982, Burnley et al. 199l, Busa & Gray 1992, Champness & Lolimer 1974, Drury & Humphreys 1988, Evans et al. 1980, Fliervoet & White 1995, Gilotti & Hull 1990, Green 1986, Green & Burnley 1989, Gower & Simpson 1992, Hacker & Kirby 1993, Jessell 1987, Johnson & Vernon 1995, Kirby & Stern 1993, Langdon 1982, MacKinnon et al. 1977, Nicolas & Poirier 1976, Passchier & Trouw 1996, Passchier et al. 1992, Powell & Treagus 1970, Rutter et al. 1994, Schmid et al. 1987, Shelly 1989a, b, Shoneveld 1977, Smith 1964, Tingle et al. 1993, Urai et al. 1986, Vaughan et al. 1984, White 1977, White & White 1981, Zwart 1960, 1962. Chapter 6. Arzi 1978, Ashworth & McLellan 1985, Bagnold 1954, Brace & Martin 1968, Benn & Allard 1989, Blumenfield 1983, Blumenfield & Bouchez 1988, Bouchez & Gleizes 1995, Bouchez et al. 1992, Burg 1991, Conolly et al. 1997, Copper & Kohlstedt 1982, 1984, Dell’Angelo & Tullis 1988, Dell’Angelo et al. 1987, Finney 1970, Gapais & Barbarin 1986, German 1985, Gray 1968, Guineberteau et al. 1987, Hibbard 1987, Hirth & Kohlstedt 1995, Hutton 1988, Jarewicz & Watson 1984, 1985, Karlstrom et al. 1993, Komar 1972a, b, Lagarde et al. 1994, Law et al. 1992, Lejeune & Richet 1995, Marsh 1981, McBirney & Murase 1984, McLellan 1984, Miller & Paterson 1994, Miller et al. 1988, Mitra 1976, Nicolas et al. 1988, Park & Means 1996, Paterson et al. 1989, Paterson & Vernon 1995, Pons et al. 1995, Quick et al. 1992, Ramsay 1989, Renner et al. 1999, Riley 1990, Roscoe 1952, Rushmer 1995, Rutter & Neumann 1995, Ryan 1995, Sherman 1968, Simpson 1985, Stel 1991, Van der Molen & Paterson 1979, Vernon & Paterson 1993, Vernon et al. 1988, Wickham 1987. Chapter 7. Aerden 1995, 1996, Balé & Brun 1989, Behrmann 1984, 1987, Behrmann & Platt 1982, Bell 1985, Bell & Johnson 1989, Bell et al. 1992, Benn & Allard 1989, Berthé et al. 1979, Bjørnerud & Zhang 1995, Blenkinsop & Treloar 1995, Blumenfield 1983, Blumenfield & Bouchez 1988, Bouchez et al. 1983, 1992, Burg 1987, Dennis & Secor 1987, 1990, Doblas et al. 1997, Erskine et al. 1993, Etchecopar & Malavielle 1987, Evans & Dresden 1991, Fossen & Holst 1995, Gamond 1987, Gapais & Brun 1981, Ghisetti 1987, Goldstein 1988, Guineberteau et al. 1987, Hanmer 1986, Hanmer & Passchier 1991, Ildefonse & Mancktelow 1993, Jeffrey 1922, Krabbendam & Leslie 1996, Law 1986, 1987, 1990, Law et al. 1984, 1994, Lister & Snoke 1984, Little et al. 1994, Mainprice et al. 1986, March 1932, Means 1981, Means et al. 1980, Mercier 1985, Miller & Paterson 1994, Park & Means 1996, Passchier 1983, 1984, 1991, 1994, Passchier & Simpson 1986, Passchier & Sokoutis 1993,

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Passchier & Trouw 1996, Passchier et al. 1993, Paterson & Vernon 1995, Paterson et al. 1989, Petit 1987, Platt 1984, Platt & Vissers 1980, Prior et al. 1987, Ramsay 1962, 1967, 1980, Ramsay & Graham 1970, Ramsay et al. 1983, Robert 1989, Robin & Cruden 1994, Rutter et al. 1986, Rykkelid & Fossen 1992, Scheuber & Andriessen 1990, Schmid et al. 1987, Shelley 1989b, 1995, Shimamoto 1989, Simpson 1986, Simpson & Schmid 1983, Sylvester 1988, Ten Brink & Passchier 1995, Tikoff & Greene 1997, Tikoff & Fossen 1993, Tikoff & Teyssier 1995, Turner & Wiess 1963, Van den Driesche & Brun 1987, Wenk et al. 1987, White & Wilson 1978, White et al. 1980, Williams et al. 1994, Yoshinobu & Patterson 1996, Zee et al. 1985. Chapter 8. Alexopoulos et al. 1988, Ashworth & Schneider 1985, Bohor et al. 1993, Carter 1965, Carter et al. 1986, 1990, Engelhardt & Bertsch 1969, Fel’dman 1994, French et al. 1974, Gigl & Dachville 1968, Glass 1990, Goltrant et al. 1991, 1992, Gratz et al. 1996, Grieve et al. 1996, Hörz 1968, Hörz & Quaide 1973, Huffman & Reimold 1996, Huffman et al. 1993, Joreau et al. 1997, Kamo et al. 1996, Kieffer 1975, Kieffer et al. 1976, Koeberl 1990, Langenhorst 1994, Langenhorst & Clymer 1996, Langenhorst & Deutsch 1994, Langenhorst et al. 1992, Leroux & Doukham 1995, 1996, Leroux et al. 1994, Lyons et al. 1993, Martini 1991, McIntyre 1962, Officer & Carter 1991, Reimold 1994, 1995, Reimold & Hörz 1986, Reimold & Stöffler 1978, Reimold et al. 1998, Robertson 1975, Robertson & Grieve 1977, Seyedolali et al. 1997, Sharpton & Schuraytz 1989, Stöffler 1972, 1984, Stöffler & Langenhorst 1994. Chapter 9. Angelier 1984, Angevine & Turcotte 1983, Ashby & Verrall 1978, Avé Lallement 1978, Baud et al. 1996, Berckhemer et al. 1979, Bernabé 1987, Biegel et al. 1989, Blacic & Christie 1984, Blenkinsop & Drury 1988, Boland & Tullis 1986, Boullier & Guegen 1975, Burkhard 1993, Burov & Diament 1995, 1996, Busch & Van der pluijm 1995, Byerlee 1978, Caristan 1982, Carter & Friedman 1965, Carter & Rayleigh 1969, Carter & Tsenn 1987, Carter et al. 1993, Christie & Koch 1982, Christie & Ord 1980, De Bresser 1988, De Bresser & Spiers 1990, Dennis 1984, Detournay et al. 1989, Dorn 1954, Drury 1993, Drury & Urai 1990, Drury et al. 1985, Evans & Groshong 1994, Farver & Yund 199la, b, Fowler 1990, Gilletti & Yund 1984, Goetze & Evans 1979, Gratier & Guiget 1986, Gratier & Jenatton 1984, Griggs 1967, Griggs & Blacic 1965, Groshong 1972, 1974, 1984, Gross & Van Heege 1973, Hallam & Ashby 1990, Handy 1990, 1994, Hansen & Carter 1982, 1983, Heard 1963, Hirth & Kohlstedt 1995, Hobbs et al. 1986, Horii & Nemat-Nasser 1985, 1986, Jaeger & Cook 1979, Jamison & Spang 1976, Jaoul et al. 1984, Ji & Zhao 1993, 1994, Karato & Wu 1993, Kemeny & Cook 1987, Kirby & Kronenburg 1984, Koch & Christie 1981, Kollé & Blacic 1983, Kronenberg & Tullis 1984, Kronenburg et al. 1990, Kruhl 1996, Kruhl & Nega 1996, Kumpel 1991, Kusznir 1982, Kusznir & Bott 1977, Kusznir & Park 1982, 1984, 1987, Lacombe et al. 1990, Laurent et al. 1981, 1990, Lawn & Wilshaw 1975a, Lehner 1990, Lemée & Guegen 1996, Linker & Kirby 1981, Marone et al. 1992, McClintock & Walsh 1962, McLaren 1991, Meisner & Strehlau 1982,

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Mercier 1980a, Mercier et al. 1977, Muller & Briegel 1978, Murrell 1963, 1965, Murrell & Digby 1970, Nemat-Nasser & Horii 1982, Newman 1994, Ord 1991, Ord & Christie 1984, Ord & Hobbs 1985, Ord & Hobbs 1989, Paterson 1978, 1987, 1995, Paterson & Luan 1990, Pavlis & Bruhn 1988, Peng & Johnson 1972, Pfiffner & Burkhard 1987, Post & Tullis 1999, Power & Tullis 1989, Raj 1982, Raj & Ashby 1971, Raterron & Jaoul 1991, Rice & Gu 1983, Ross et al. 1980, Rowe & Rutter 1990, Rutter 1976, Schmid et al.

125

1977, 1980, Scholz 1990, Schulmann et al. 1996, Shelton et al. 1981, Sibson 1982, 1983, Spang & Van der Lee 1975, Spiers et al. 1990, Stoker & Ashby 1973, Takahashi et al. 1997, Terzaghi 1943, Tse & Rice 1986, Tsenn & Carter 1987, Tullis 1980, Tullis et al. 1991, Turner 1953, Twiss 1977, 1986, Van der Wal 1993, Walker et al. 1990, Warpinski & Teuffel 1993, Weertman 1968, 1978, Weiss 1954, White 1976, 1979a.


Index (Bold numbers refer to plates or figures)

abrasive wear 10 accretion steps 20 activation enthalpy 92-96 activation volume 94 adhesive strength, wear 10 alteration 9, 10, Plate 1 Amonton’s law 10, 91 amphibolite facies 24, 52, 96 amygdale 22 annealing 62, 82, 88 anticrack 58 antitaxial fibre growth 33, 37, 3.17, 3.18 arrow method 102, 103 asperities, asperity ploughing 10, 2.4, 20 asterism 82 ballen 83, Plates 47, 48 barrier theory 9 bend 79 bookshelf sliding 77 botryoidal texture 32, 35 boudin 19, 2.17, Plate 6 Boussinesq configuration 12 Brazil twin 81, 8.3, 85, 88, 89 breccia 5, 6 brittle 1,3 brittle-ductile transition 4, 5 brittle-plastic transition 3, 97 Burger’s vector, 39, 4.1, 58 Byerlee’s law 91, 97 calcite twin morphology 104 cataclasis 1-3, 5, 7-24, 30, 38, 47,60,63,74 cataclasite 6 foliated cataclasite 6, Plate 8 protocataclasite 6 ultracataclasite 6, 23 cataclastic flow 4, 18, 19, 63 cathodoluminescence 2.6, 12, 17, 32,35, 81 cement, cementation 17, 32 chemical potential 24, 52 chemical zoning 2, 57 chessboard pattern 41, 4.10, 47 Chixculub structure 84 Cish diagram 37 cleavage 2, 13, 28-30

classification 28, 29 continuous 29 crenulation 29, 30, Plates 16, 17 disjunctive 29, 3.7 domain 29, 30 slaty 30, 3.9 spaced 29, 30 zonal 29 coefficient of friction, coefficient of internal friction 90, 91 coesite 16, 80, 83, 8.7, 85-88 coherent exsolution 52 coherent transformation 52 cohesion 5, 90, 91 composite deformation mechanism 3, 52, 90 composite fibre growth 37, 3.17 constrained comminution 10 contact melting 63 contiguity 59 core-and-mantle structure 4.14, 42, 49 corona 2, 57, Plate 34 crack-seal 35, 92 crater 23 creep 92-94 Coble 52, 92-94 diffusion 52, 54, 60, 97 dislocation 93-97 grain size sensitive 92, 93 Nabarro-Herring 52, 92, 94 pressure solution 92, 93 critical melt fraction 59 critical packing, packing density 59, 60 critical slip distance 92 cross-slip 39, 4.4 crystal plasticity 4, 5, 19, 59, 104 crystallites 22 crystallographic fabric, preferred orientation 1, 2, 19, 22, 50, 51, 4.20, 58, 62, 75, 76, 7.19 curved foliation 66 Dauphine twin 57 debris streaking 20 decussate texture 2, 54, Plate 28 defect 52 deflection surface 54, 57 deformation bands 2, 5, 7, 18, 41, 4.6, 47, 50 deformation lamellae 2, 14, 4.11, 47, 49, 101-103, 9.4 127

128

deformation mechanism 1, 2, 3, 5, 90 classification 1, 2 map 90, 97, 98 deformation microstructure 1-3, 5, 90 classification 1, 2 deformation (mechanical) twin 2, 39-41, 4.8, Plate 24, 62, 81, 87, 102-104, 9.4, 9.5 deformation zone rocks 5 deformation continuity 4, 5, 18, 19,57 distribution 4, 5, 18 mechanism and mode 4, 5 scale 4 dendritic crystals 22 diaplectic glass 80-83, 85, 87, 89 diffusion 24, 29, 53, 57, 82, 94-94 coefficient 52, 92-94 creep - see creep grain boundary 93 volume 93 diffusive mass transfer 1-3, 47, 63, 90, 92 by solution 3, 24-38 in melts 60 solid state 52-58 dihedral angle 24 dilatancy hardening 60 dislocation 1, 30, 39, 4.1, 4.2, 4.3, 4.4, 47 climb 39, 93 creep 39, 5.1, 54, 93-97 density 39, 41, 47-49, 58, 81, 82 glide 39, 51,93-95 displaced grain fragments 79, 7.23, Plate 45 displacement control 33 displacive transformation 52 divider method 105 Dorn law 93 ductile stringer 19, 2.17 ductility 4, 5 dynamic recrystallization 47-50, 4.19 effective stress 91 Einstein-Roscoe equation 59 enclave 60 erosional sheltering 20 etching, etch pit 39, 88 exaggerated grain growth 54 exponential law 93-97 extensional crenulation cleavage 70 external asymmetry 76, 7.19 fabric skeleton 76 face control 33, 38, Plate 19 failure criteria 90, 91, 97 Coulomb and Mohr failure criteria, 90, 97, 98 and fracture mechanics 91 Griffith criteria 91,97 fault gouge - see gouge fault breccia - see breccia fibre 32, 33, 37, 65

INDEX

fibre strain 17 growth histories: antitaxial, ataxial, composite, nonsystematic, syntaxial 37, 3.17, 3.19 non-tracking 38, 3.19 tracking 37, 38, Plate 23 tracking efficiency, 38, 3.19 Fick’s law 24, 52 finite shortening, extension 25 flow laws 90, 92-97 fluid inclusions planes 2, 12, 24, 2.5, 33-35, 3.13, 3.14, 47 foam texture 2, 54 fold 4, 19, 2.17, 60, 63, 70 foliation 54-57, 65, 75 curved 66, Plate 39 external foliation 54, 55, 5.3, Plates 29-33 internal foliation, 54, 55, 5.3, Plates 29, 31-33 oblique foliation, 66, 67, 7.3 steady-state, strain insensitive 61 strain sensitive 66 fractal dimension 10, 22, 92, Plates 1-4 fracture 1 fracture toughness 9 fragmentation 10 frictional sliding 7, 10, 2.4, 90-92, 97 generation plane 22 geotherm 97 geothermobarometry 104, 105 gouge and gouge zone microstructures 5, 6, 10, 17, 2.17, Plate 6, 72 grain boundary migration 4.12, 47, 4.17, 48-50, 4.18, 4.19, 95, 99, 100 grain boundary width - fast or free 54 grain dispersive pressure 61 grain shape fabric 2 cataclastic 19, 33 diffusive mass transfer 33, 3.11, 52, Plates 26, 27, 54 intracrystalline plastic 33, 4.12, 47 magmatic 62, Plate 36 sub-magmatic 62 grain size sensitive creep 92, 93 grain surface deposition textures 2, 30 grain surface solution textures 2, 25 granoblastic polygonal texture 32, 54, 62 growth twins 39, 62, 81 growth zoning 57, 83, Plate 18 helicitic texture 54 Hertzian configuration 12 homologous temperature 97 hydrolytic weakening 9, 94 igneous zoning 62 imbrication 60, 62, 78 inclusion 35, 3.15, 3.16, 62, 84 band 35, 3.15, 3.16 trail 35, 3.15, 3.16, 37, 38, 54, 65 indenting grain contacts 2, 25, 3.1, 3.2, 33, 63 independent particulate flow 2, 22

129

INDEX

instantaneous stretching axis, axes 65, 73, 7.15, 7.17 interconnected weak layer 96 internal asymmetry 76, 7.19 internal strain energy 4, 48, 52, 55 interpenetrating grain contacts 2, 25, 3.1, 3.4, 33 intracrystalline plasticity 1, 3, 4, 14, 19, 22-24, 30, 39-51, 57, 60, 62, 63, 75, 90, 93, 97 intracrystalline deformation bands 2, 41 island and channels 24 isostrain criterion 96 isostress criterion 96 jog 79

kink, band 2, 14, 30, 41, 4.6, 4.7, 4.8, 50, 62 landslide 23 lattice preferred orientation 58, see crystallographic fabric lechatelierite 80-85 lineation 66, 7.1, 7.2, 75 lithospheric strength envelope 90, 97, 98, 9.2 load bearing framework 96 low angle boundaries 41 lower stability limit 92 magmatic flow 3, 59-63, Plate 36 magmatic microstructures 59, 62 magmatic shear zones 62, 77, 78 magnetic anisotropy 60, 62 Martensitic-like transformation 52, 57 mean stress 94 megacryst 62 metatextite 63 mica beard 32 mica fish 74, Plate 42 microboudins, asymmetric 76, 7.20, 7.21, Plate 6 microcrack 2, 7, 2.1, 2.5, 2.6, 10-16, 33 axial 12, 17, 2.14, 19, 63,91 characteristics and observation 10-12 circumgranular 10-12 classification 10-12 cleavage 13, 19 cone 12 dynamic propagation 7 elastic mismatch 13, 14, 2.11, Plate 5 en-echelon 17, 2.15 en passant 17, 2.15 extension 12, Plate 2 extension force 7 flaw-induced 2.6, 13, 2.10 grain boundary 13 impingement 2.6, 12, 13, 2.7, 2.8, 2.9 intragranular 10, 2.5, 12, 13, 2.9, 18, 19 microfault- induced 14, 19 microscopic feather fracture 14, 15, 2.12 modes I, II, III 7, 2.2, 12 phase transformation-induced 16 plastic mismatch 14, 2.12, 2.13 refracturing 13, 2.11

sub-critical propagation 9, 94 thermally-induced 15, 16 transgranular 10, 2.5, 12, 13, 16 wing 13, 2.10 microcrystallite 81, 82 microfault 2, 7, 17, 18, 2.14-2.16, 7.23, Plates 3, 45 microfold 77, 7.22 microfracture 1, 2, 4, 7, 80, 85, 86, Plate 4 mirror 20 mist 20 surface features 2, 19, 20 velocity hackle 20 Wallner line 20 microlithon 29, 30 microphenocryst 22 microslickolite, 27, 28, 3.6 microstylolite 1, 2, 27, 28, 3.5, Plate 14, 3.6 characteristics 27 filling 24-28 formation and propagation 27, 28 teeth, walls, crown 27, 3.5, 28 microvein 1, 2, 7, 35-38 free-face growth 35 fibrous 33, 35, 37, 38 laminated 35 opening vector 35 migmatite 63 millipede texture 54 mimetic crystallization 54 mode of failure 4 molar entropy 24 molar internal energy 24 molar volume 24, 92-94 mosaicism 81-83, Plate 48, 85-89 mylonite 5, 6, 62 protomylonite 6 ultramylonite 6 necking down 33, 35, 3.14 neosomes 63 new grains 47, 50 non-magmatic deformation 63, 64 non-systematic fibre growth 37, 3.17, Plate 22 normal slip crenulation 70 oblique grain shape fabric 78 ophitic texture 62 order-disorder transformation 52 Ostwald ripening 54 overgrowth 1, 2, 4, 13, 32, 33, 3.10, Plate 18, 3.11, 4.11 P-foliation 19, Plate 6, 79 P-shear, fracture 19, 20 paleopiezometry, palaeopeizometer, 98-101 deformation lamellae 101-103 dislocation density 100, 101 general problems 104 maximum twin volume, 101 principal stress and strains 103, 104, 9.4

130

INDEX

recrystallized grain size 98-100, 9.3 pseudotachylites 2, 5-7, 22, 23, Plates 9-11, 83, 92 subgrain size 100, 101 characteristics 22 misidentification 23 twinning - differential stress 100, 101 twinning density, 101 origin 22, 23 twinning incidence, 101 rate and state dependent frictional sliding 92 particle size distributions 10, 22, 92 reaction rims 2, 57 perlitic texture Plates 47, 48 reaction zoning 57 permeability 91 recovery 2, 41, 47, 82 phase transformation microstructures 2, 57, 58, 80-92 recrystallized grain size 49, 50 phenocryst 60-62 reentrant zoning 54 phenomenological coefficient 93 relict minerals 2, 57 phyllonite 6 pinching off 33, 3.14 rheology 90 planar deformation feature 47, 8.2, 80-89 rhomb dodecahedra 54 ribbon grain 2, 4.13, 52-54, 5.1 decorated 81,82 non-decorated 81, 82 ridges and grooves 20 sub-lamellar structure 81, 88 Riedel, conjugate Riedel shear 2.17, 19, Plates 6, 7, 20, 72, plastic, plasticity 3, 14 79 fractures 15 poikiloblast 54 Poisson’s ratio 97 risers 20, 2.19, 79 polymineralic deformation 94-97 congruous/incongruous 20, 2.19 rolling structures 77 pore fluid, pressure 3, 22-24, 91 porosity 12, 13, 18, 87, 88, 93 roughening transition 32 reduction 19, 32 S-, C- and 5, 20, 62, 63, 66, 70-73, Plates 40, 41 porphyroblasts 1, 2, 54-57 characteristic and classification 70-72, 7.11, 7.13, 7.14 characteristics 54, 55 curvature of S-foliation 73 growth mechanisms 55 formation and evolution 72 internal, external foliations 54, 75, 7.18 shear on C-or 19, 73 intertectonic 55-57, 5.3, 5.4, Plate 31 shear sense from 73 plate 33 scaly clays 22 posttectonic 55-57, 5.3 scanning electron microscope 2.6, 19, 2.18, 2.19, 25, 30, 50, pretectonic 55-57, 5.3, Plate 30 79, 82, 88 relationship to deformation 55-57, 5.3 schlieren 60 shear sense indications 75, 7.18 secondary recrystallization 54, 55 syntectonic 55-57, 5.3, Plate 32 sector zoning 54 porphyroclast, porphyroclast systems 5, 14, 67-70 semibrittle 3, 14, 19 characteristics and classification 67, 68, 7.4 complex type 67, 7.4, 7.8 sensitive tint plate 2.6, 12, 17, 50, Plates 1-4, 45, 79 separatrix 68, 69 67-70, 7.4, 7.7, 7.9, 7.10, 74 deflection, embayments 68-70, 7.7, 7.9 shape preferred orientation 47, 66 faces of a tail 69, 7.9 shatter cones 23 in-plane 67, 68 shear band 70 67, 7.4 shear modulus 97 mantle 67 shear sense 65-79 criteria 65-79 mechanisms of formation 68, 69 for faults 79, 7.23, Plate 45 67-69, 7.4, 7.5, 7.6, 7.9 stair-step 67, 68 in rocks containing melt 77-79 tail, wing 67 observation plane 65, 71 power law 93-97 shock-induced microstructures 3, 80-89, 8.1 breakdown 94 shock mechanisms and metamorphism 23, 80 pre-exponential constant 93-97 shock wave barometry and thermometry 85-88 pre-lithification deformation 2, 22 calibration 85-87 problems 87, 88 pressure shadows and fringes 1, 2, 4, 32, Plate 19, 33, 3.12, 63, 65, 73,74 slickenfibre 20, 2.19, 79 last increment of growth 73, slickenline 20 kinematics in shear zones 73, 7.15 slickenside 20, 92 shape 73, 7.16, 7.17 slip 39 process zone 9, 28 direction 39, 51 prod mark 20, 2.18 plane 39, 4.1, 51

131

INDEX

system snowball texture 54, Plate 29 solid state transformation 52-58, 80-82 solubility 24, 92-94 spherulites 22, 83, 84 spin 65 state variable in frictional sliding law 92 static recrystallization 54, 62 stick-slip behaviour 92 stishovite 80, 83, 8.8, 85-87 strain cap 2, 27, Plate 13 strain rate 90, 92-97 strawberry texture 82, 8.5, 8.6 stress 1 amplification 97 corosion 9 exponent 93-97 intensity factor 7, 9 stretched crystal fibre 37, 3.17 strewn field 84 sub-boundary migration 50 sub-magmatic flow 3 microstructures 3, 59-63, Plates 37, 38, 78, 79 subgrain 1, 2, 14, 19, 32, 41, 47, 4.9, 4.13, 47, 4.16, 50, 62 boundary orientation in quartz 41, 63, 105 rotation 47, 4.15, 4.16, 48, 49, 4.19, 50, 99, 100 size 1, 47 superplasticity 3, 52, 57, 58, Fig. 5.5, 94 surface tension force, energy 7, Plate 27, 54, 55 sutured grain boundary 25, 50, 3.1, 4.19, 104, 105, 9.6, 9.7 symplectite 2, 57, Plate 35 syntaxial fibre growth 33, 37, 3.17, Plate 21 T fracture 19, 2.17, 20, 79 tectite, microtectite 83, 84

tiling 60, 78 tool track or mark 20, 2.18 transmission electron microscope 19, 39, 4.2, 81-83, 88, 101 transpression 65, 7.2 triple grain junction 54, 5.2 truncating grain contact 2, 25, 3.1, 3.3, Plate 12, 30, 33 truncation surface 30, 3.7, 54, 55 twin, twinning 47, 57, see also deformation (mechanical) twin, growth twin, paleopiezometry undercutting mechanism 24 undulatory extinction 1, 2, 33, 41, 4.5, 4.6, 4.13, 49, 62, 82 uniaxial compressive strength 90, 91 uniaxial tensile strength 90, 91 upper stability transition 92 velocity weakening, strengthening 92 vesicle 84 viscosity 59-61 volume loss 29, 3.8, 30 vorticity 65 external 65 internal 65 profile plane 65, 7.1 shear-induced 65 vector 65, 7.1 wear groove 20, 2.18 whole lithosphere failure 97 Y-shear Plate 7, 19, 2.17, 72 Young’s modulus 7 zircon, shocked 82, 8.4-8.6, 89


Color Plate Section


COLOR PLATE SECTION

135

Plate 1. Alteration-enhanced microcracking. The plagioclase crystal in the light blue colour is fragmented along cleavage planes without significant rotation or displacement of the fragments. Alteration of the feldspar to more sodic compositions (and ultimately to laumontite) occurs along the microcracks, showing that chemical reaction and deformation were linked in alteration-enhanced microcracking. The fractal dimension of the particle size distribution (PSD) is ~2.0, which is characteristic of this type of cataclasis. (Chapters 2.2.1, 2.3.6). XPL, ST, 4.3 mm, Biotite granite, Cajon Pass drillhole, California, U.S.A. Plate 2. Extension microcrack. Grain boundaries between the yellow and pink grains in the wallrock can be matched across the microcrack, demonstrating extension in the plane of the section. The microcrack is filled by angular fragments of quartz, biotite and feldspar in random orientations which can not be matched with adjacent grains in the wall-rock. These textures suggest that the fragments were transported in a single, chaotic manner from some distance, probably in a fluid matrix. The fractal dimension of the PSD is approximately 2.8, slightly higher than typical for extension microcracks. (Chapter 2). XPL, ST, 2.2 mm, Granodiorite, Cajon Pass drillhole, California, U.S.A. Plate 3. Microfault. The great variety of colours of the fragments (mostly quartz) in the matrix under the sensitive tint plate shows that they have been rotated and probably derived from grains at least outside the field of view. Many fragments are equant and sub-angular, suggesting some wear during shear. The fractal dimension of the PSD is approximately 2.6, a common value for fragments in shear fractures, and also the value predicted by the constrained comminution model. (Chapters 2.2.2, 2.4). XPL, ST, 4.3 mm, Granite, Cajon Pass drillhole, California, U.S.A.

136

COLOR PLATE SECTION

Plate 4. Selective microfracture of larger fragments. Angular, randomly orientated fragments were produced by microfracture during cataclasis. A large feldspar grain at lower right is separated into two fragments by a microfault. The fractal dimension of the PSD is approximately 2.9. (Chapter 2.2.2). XPL, ST, 4.3 mm, Granite, Cajon Pass drillhole, California, U.S.A. Plate 5. Elastic mismatch strains. Intracrystalline plastic strains around the ends of an inclusion of biotite are dramatically shown by the areas of quartz in extinction, compared to the birefringent host grain elsewhere. The strained area has the same geometry as the area of high stress predicted from modelling the elastic stress field around a weak inclusion, suggesting that the elastic stresses resulted in lattice distortion. See Fig. 2.11. (Chapter 2.3.7). XPL, 0.5 mm, Chilimanzi granite, Great Zimbabwe, Zimbabwe. Plate 6. Riedel shears, asymmetric boudin and P-foliation in a sinistral gouge zone. Riedel shears trend from top right to bottom left, and cut across a Pfoliation which trends in the opposite direction. The two features surround an asymmetrical boudin. All three features give a sinistral shear sense. Compare with Figs. 2.17, 7.20. (Chapters 2.7, 7.13.4). PPL, 2 mm, Tarskavaig thrust zone, Tarskavaig, Isle of Skye, Scotland, U.K.

COLOR PLATE SECTION

137

Plate 7. Y-Shear in a gouge zone. The central dark line parallel to the top/bottom edges of the photomicrograph is a Y-shear. Above and below are Riedel shears similar to those in the previous plate, which give a sinistral sense of shear. Compare with Fig. 2.17. (Chapters 2.7, 7.13.4). PPL, 2 mm, Tarskavaig thrust zone, Tarskavaig, Isle of Skye, Scotland, U.K. Plate 8. Foliated cataclasite. The foliation (planar feature from upper right to lower left) is defined by bands of quartz and calcite. Calcite-filled microfractures at high angles to the foliation demonstrate the brittle component of the deformation. (Chapter 1.3). XP, 1 mm, Foliated cataclasite, Punchbowl fault zone, California, U.S.A. Plate 9. Pseudotachylite. Dark, fine grained matrix containing clasts fills a triangular-shaped veinlet branching from a narrow vein of the same material parallel to a fault surface. The veinlet cross-cuts an older foliation in the wall rock. Sub-angular fragments of wall rock in the veinlet have a large range of sizes. All these features are typical of pseudotachylite. See Plates 10, 11. (Chapter 2.11). PPL, 2 mm, Pseudotachylite in grey gneiss mylonite, Buhwa, Zimbabwe.

138

COLOR PLATE SECTION

Plate 10. Pseudotachylite. The same view as plate 9 in crossed polars. The opacity of the matrix suggests a glassy nature. Slight alteration to a very fine grained, moderately birefringent phyllosilicate picks out a weak foliation in bands which are concave towards the apex of the veinlet. These bands suggest a primary flow foliation. See Plates 9, 11. (Chapter 2.11). XL, 2 mm, Pseudotachylite in grey gneiss mylonite, Buhwa, Zimbabwe. Plate 11. Pseudotachylite. Branching veins contain an opaque, fine grained matrix suggesting glass. Poorly sorted, sub-rounded fragments define a weak foliation parallel to a feint colour banding in the matrix. The foliation and banding are folded about an axis parallel to the margin of one vein, suggesting a flow foliation. See Plates 9, 10. (Chapter 2.11). PPL, 2 mm, Pseudotachylite in tonalitic granulite, Bollingen Islands, Antarctica. Plate 12. Truncated Ooid. Concentric banding in the elliptical ooid is truncated at the upper and lower surfaces by quartz grains along approximately planar surfaces. The length of the original ooid can be estimated as approximately 0.6 mm: its present length is 0.3 mm. (Chapter 3.4). XPL, 1 mm, Bargate Stone, U.K.

COLOR PLATE SECTION

139

Plate 13. Strain cap. The porphyroclast of feldspar at centre has foliae of muscovite wrapped around the top and bottom surfaces. Flakes of muscovite distributed sparsely through the matrix around the porphyroclast define a weak foliation. The presence of muscovite in the matrix suggests that the muscovite was concentrated in the strain cap by diffusion of the other matrix components away from the strain cap. (Chapter 3.5). XPL, 1 mm, Mylonite, Moine thrust zone, Scotland, U.K. Plate 14. Microstylolite truncating crinoid fragment. The microstylolite has a filling of opaque iron oxides/hydroxides and truncates the crinoid fragment, showing material removal. See Fig. 3.8. (Chapter 3.6). PPL, 1 mm, Crinoidal limestone, Derbyshire, U.K. Plate 15. Microstylolites. The wider stylolite is approximately planar, and the narrow stylolite is sinusoidal. Both are zones of concentrated opaque minerals, which are found in low concentrations throughout the adjacent rock, showing that they have been concentrated by removal of the more soluble carbonates along the stylolites. Shell fragments are truncated by the stylolites. The amplitude of the sinusoidal stylolite suggests approximately 0.3 mm of shortening perpendicular to the stylolite trace. (Chapter 3.6). PPL, 2 mm, Carboniferous limestone, Somerset, U.K.

140

COLOR PLATE SECTION

Plate 16. Incipient crenulation cleavage. A early fabric defined by muscovite and chlorite is folded into open folds. A weak second fabric is seen on the fold limbs, defined by muscovite and chlorite alignment parallel to the axial surfaces of the folds, and by concentrations of opaque minerals which are found elsewhere throughout the rock, suggesting that diffusive mass transfer has contributed to the formation of this cleavage by removal of the more soluble phases from the fold limbs. (Chapter 3.7.2.2). PPL, 2 mm, Albite chlorite schist, Perthshire, U.K. Plate 17. Advanced crenulation cleavage and fabric transposition. Two distinct types of layers are visible in this rock. Layers consisting of strongly aligned biotite also contain a few small elongate grains of quartz parallel to the biotite fabric (S2, not labelled). Layers between the S2 domains contain larger, equant quartz grains and a higher proportion of quartz, which defines a compositional banding (S1, not labelled) of quartz layers alternating with biotite layers, which also have a strong fabric parallel to S1. S1 is folded into tight to isoclinal asymmetric folds; small scale fold hinges can be seen in some of the biotite grains and in the compositional layering. Depletion of quartz in the S2 layers shows that it has diffused away from the S2 domains. S1 has been almost completely transposed into S2. (Chapter 3.7.2). XP, 1 mm, Biotite schist, Cornwall, U.K. Plate 18. Precipitation and solution textures in hydrothermal quartz. This SEM-CL false colour image shows a zoned quartz crystal with euhedral concentric growth zones shown by different colours disrupted by solution. The complex zonation patterns show variations in fluid chemistry during growth. (Chapter 3.9). SEM-CL, 0.9 mm. Hydrothermal quartz vein cross-cutting the Ventersdorp Contact Reef, Klerksdorp goldfield, Witwatersrand, South Africa.

COLOR PLATE SECTION

141

Plate 19. Pressure fringe. Fibrous quartz overgrows euhedral faces of pyrite crystals. The quartz fibres are perpendicular to several faces, showing a typical face-controlled geometry. (Chapter 3.9). XPL, 1 mm, Quartz-sericite schist, Sabi mine, Zvishavane, Zimbabwe. Plate 20. Blocky microvein filling. Blocky grains of calcite filling this microvein suggest a single opening and filling event. (Chapter 3.12). XPL, 2 mm, Mylonite, Freda Rebecca mine, Bindura, Zimbabwe. Plate 21. Syntaxial vein filling. The photomicrograph shows one side of a syntaxial quartz and calcite vein filling. A narrow band of partly fibrous quartz lines the vein adjacent to the dark wall rock. Some fibres overgrow quartz grains in the wall rock, and the quartz grain size increases away from the wall, showing that the quartz grew syntaxially. The calcite consists of curved but unstrained calcite fibres, which also coarsen away from the wall rock. The calcite fibres grow towards a median suture which is parallel to the vein wall. A second group of fibres growing from the opposite side of the vein meets the first group at the median suture, but the two groups are not continuous. The continuity between the vein quartz and the wall rock, the increase in grain size away from the wall, and the joining of two separate groups of fibres at the median suture are all distinctive features of syntaxial growth. The curved but undeformed nature of the fibres shows that the curvature is a primary growth feature. Since the fibres grow towards a median suture, wall rock control of their orientation is unlikely, and the fibres probably tracked the incremental opening vector, which rotated with respect to the vein. See Fig. 3.17a. (Chapter 3.12). XPL, 4.3 mm, Quartz-calcite vein, locality unknown.

142

COLOR PLATE SECTION

Plate 22. Non-systematic (stretched) fibres. Individual quartz fibres stretch all the way across the microvein. There is no symmetry or suture, and the sides of the fibres have interlocking teeth. They are subdivided into tablets perpendicular to their length, each of which represents a growth increment. These are the typical features of non-systematic growth histories. See fig. 3.17e. (Chapter 3.12). XPL, 4 mm, Semi-psammite, Anglesey, U.K. Plate 23. Tracking fibres. Undeformed biotite fibres grow obliquely across the microvein, and remain parallel through a small jog in the margin (centre). These observations suggest that the fibres were tracking an incremental opening direction that was oblique to the margin, since the change in the margin orientation at the jog has no effect on the fibre orientation. (Chapter 3.12). XPL, 1 mm, Serpentinite, Sabi mine, Zvishavane, Zimbabwe. Plate 24. Deformation twins in calcite. Characteristic features of deformation twins are variable thickness, pinching out, and bending of individual twins. These twins are type 2 of Burkhard’s classification, and therefore formed between 150 and 300°C. Also see Fig. 4.11. (Chapters 4.3, 9.11.2). XPL, 0.25 mm, Limestone, Mushandike, Zimbabwe.

COLOR PLATE SECTION

143

Plate 25. Grain boundary migration. Advanced stages of grain boundary migration. As well as highly convoluted grain boundaries, blue grains can be seen entirely surrounded (in the plane of the section) by the yellow grain. Also see Figs. 4.17, 4.18. (Chapter 4.8). XP, ST, 0.5 mm, Deformed granitic rock, Unknown locality. Plate 26. A strong grain shape fabric defined by hornblende and biotite shows no evidence of internal strain features, and formed by solid state diffusive mass transfer. The fabric in the mafic minerals has controlled the shape of the quartz grains, imparting a similar fabric. Also see Plate 27. (Chapter 5.3). PPL, 4 mm, Hornblende schist, Makaha greenstone belt, Zimbabwe. Plate 27. Grain boundary control in a two-phase aggregate. Quartz grain boundaries are controlled by muscovite grains. Boundaries between two quartz grains are perpendicular to muscovite basal planes, and often pinned at the ends of the muscovite grains (centre). Both these effects are due to the higher surface energies between quartz and mica than between quartz grains. The resultant quartz grains are elongate parallel to the muscovite fabric. Also see Plate 26. (Chapter 5.3). XP, 2 mm, Mica-garnet schist, Miami district, Zimbabwe.

144

COLOR PLATE SECTION

Plate 28. Decussate texture. Randomly orientated, interlocking, elongate actinolite crystals show a good example of this texture, produced by diffusive mass transfer processes in crystals with anisotropic growth rates. (Chapter 5.4). XPL, 2 mm, Biotite-actinolite schist, Arcturus mine, Zimbabwe Plate 29. Snowball texture in garnet. Inclusions trails of quartz define an internal foliation that curves through more than 180°. The external foliation is defined by a strong grain shape fabric in adjacent muscovite. See Fig. 7.18. (Chapters 5.6, 7.8). XPL, 4 mm, Muscovite-garnet schist, Makaha greenstone belt, Zimbabwe. Plate 30. Pretectonic porphyroblast. The andalusite porphyroblast contains randomly orientated opaque inclusions. An external fabric defined by biotite wraps around the porphyroblast. This is the typical texture of pre-tectonic porphyroblasts (Fig. 5.3c). Asymmetric quartz pressure shadows (light areas) occur on either side of the porphyroblast. See Plate 44 for another example from this rock. (Chapter 5.6). PPL, 2 mm, Biotite-andalusite schist, Sharriva greenstone belt, Zimbabwe.

COLOR PLATE SECTION

145

Plate 31. Intertectonic porphyroblasts. The two biotite porphyroblasts contain an internal fabric of relatively large quartz grains. The external fabric is continuous with but has a smaller grain size The strong fabric localized on the limbs of folds in and the folds themselves, are not present in the biotite porphyroblasts. The porphyroblasts grew after was formed, but before reached its present grain size and was folded, showing the diagnostic features of intertectonic porphyroblast (Fig. 5.3b). (Chapter 5.6). PPL, 2 mm, Metapelite, Longman Mountains, Sichuan province, China. Plate 32. Syntectonic porphyroblast. The external foliation is continuous with the internal foliation, which is curved within the staurolite porphyroblast. This is the typical texture of a syntectonic porphyroblast (Fig. 5.3c). (Chapter 5.6). XPL, 4 mm, Staurolite schist, Dindi greenstone belt, Zimbabwe. Plate 33. Post-tectonic porphyroblast. The kyanite porphyroblast at centre grows across a strong biotite-muscovite fabric. The internal fabric external fabric are continuous, and is undeformed around the porphyroblast, indicating post-tectonic growth (See Fig. 5.3d). (Chapter 5.6). PPL, 2 mm, Kyanite staurolite schist, Reynolds range, Arunta block, Australia.

and the

146

COLOR PLATE SECTION

Plate 34. Corona. The central hornblende grain is surrounded by an intergrowth of orthopyroxene, plagioclase and magnetite, a typical texture produced by prograde reaction of hornblende. (Chapter 5.7). PPL, 1 mm, Mafic granulite, Northern Marginal Zone, Chief Bota, Zimbabwe. Plate 35. Symplectite. Intergrowth between orthopyroxene (clear) and plagioclase (clouded appearance). Vermicular intergrowth of the two phases occurs because diffusion distances during the reaction were too small to allow an equilibrium microstructure to form. (Chapter 5.7). PPL, 2 mm, Mafic granulite, Datong-Huai’an, China. Plate 36. Magmatic grain shape fabric. The three large euhedral plagioclase phenocrysts and several smaller ones define a grain shape fabric trending from upper right to lower left. The phenocrysts are completely unstrained and set in a matrix of pyroxene. Undeformed igneous zoning can be seen in the largest phenocryst. No evidence for deformation is seen in the hand specimen, which also shows the alignment of euhedral phenocrysts. The complete lack of microstructural evidence for strain demonstrates that the fabric is magmatic. (Chapter 6.4). XPL, 2 mm, Basalt, Bembezi river, Zimbabwe.

COLOR PLATE SECTION

147

Plate 37. Magmatic/submagmatic grain shape fabric. Euhedral plagioclase and biotite crystals define a grain shape fabric parallel to the top/bottom of the plate. The only other evidence for deformation is slight undulatory extinction in quartz. Primary, undeformed zoning is visible in the dark plagioclase crystal on the right. The biotite and feldspar crystals define a moderate grain shape fabric in the hand specimen. There is no evidence for recrystallization (e.g. a quartz aggregate shape fabric). The deformation was either magmatic, with a later non-magmatic deformation that created undulatory extinction in the quartz, or sub-magmatic. In either case, the fabric was created with melt present. (Chapter 6.5). XPL, 4 mm, Tonalite, Fort Rixon, Zimbabwe. Plate 38. Sub-magmatic fabric. The euhedral outline of the feldspar megacryst is parallel to a biotite and hornblende grain shape fabric. A euhedral feldspar megacryst alignment is a conspicuous feature of the outcrop from which this specimen was collected. The feldspar megacryst is recrystallized along its margins, and slight bending of the feldspar twins is visible. This granite is part of a suite that is syntectonic with a major deformation event. The shape fabric and the intracrystalline deformation features can therefore be interpreted as submagmatic. (Chapter 6.5). XPL, 2 mm, Razi Granite, Mavizhu, Zimbabwe. Plate 39. Curved foliation. Foliation defined by muscovite lies at angle of about 45° to the upper edge of the plate. It curves smoothly to become parallel to the lower edge of the plate, which is the orientation of the shear plane. The clockwise rotation shows a dextral sense of shear (Fig. 7.1). (Chapter 7.2). XPL, 1 mm, Granite, Mushandike, Zimbabwe.

148

COLOR PLATE SECTION

Plate 40. fabric or extensional crenulation cleavage. A penetrative S-foliation from upper left to lower right is defined by hornblende grains. S is deflected into a narrow inclined in the opposite direction. The sinistral shear sense is clear from the curvature of S into (See Figs. 7.12, 7.13). (Chapter 7.5). XP, 1 mm, Hornblende Schist, Unknown locality. Plate 41. S-, C- and Three fabrics are visible. The pervasive S-foliation inclined from upper right to lower left is defined by muscovite and quartz grain shapes and compositional banding. Discrete shears parallel to the top and bottom edges are C-surfaces. A few discrete shear surfaces are inclined to the shear plane from upper left to lower right; these are surfaces. Both S-C porphyroclastic fabrics and banded fabrics and are visible (see Fig. 7.11 for definitions of and The dextral shear sense is clearly given by the curvature of S towards C- and (See Figs. 7.12, 7.13). (Chapter 7.5). PPL, 1 mm, Mylonite in grey gneiss, Buhwa, Zimbabwe. Plate 42. Mica fish. The lozenge shape of this single crystal of biotite is typical of mica fish. Short tails of biotite extend from the end of the fish, and appear to have formed by microfracture along basal cleavage planes. The basal cleavage is parallel to the long axis of the fish. Recrystallization occurs along the left and right margins of the fish. The shear sense is clear from the left-stepping tails, and the clockwise rotation from the long axis of the fish to the shear plane (parallel to the top/bottom edges of the plate). Also see Plate 43. (Chapter 7.7). XPL, 0.5 mm, Mylonite, Shamva greenstone belt, Zimbabwe.

COLOR PLATE SECTION

149

Plate 43. Hornblende fish. The single hornblende crystal has a similar geometry to the mica fish in the previous plate. The dextral shear sense is clear from the stair stepping of the tails and the obliquity of the fish long axis to the shear plane. Also see Plate 42. (Chapter 7.7). XPL, 2 mm, Granite mylonite, Zivuku, Zimbabwe. Plate 44. Asymmetrical pressure shadow. This andalusite porphyroclast gives a dextral shear sense from the left-stepping tails. The clear areas are pressure shadows of quartz which also step up to the left. Plate 30 shows another example from this rock. (Chapter 7.6). PPL, 2 mm, Biotite schist, Shamva greenstone belt, Zimbabwe. Plate 45. Matching of grains across a microfault. The use of the sensitive tint plate allows grains of similar colour to be matched across the microfault, showing a sinistral separation of ~ 0.1 mm in this section. The microfault matrix contains fragments of different colours from the adjacent grains, suggesting rotation or transport from other grains, as well as a quartz cement which is continuous with the wall grains (Fig. 7.23). (Chapter 7.13). XPL, ST, 1 mm, Barrios quartzite, Cantabrian zone, Villamanin, Spain

150

COLOR PLATE SECTION

Plate 46. Multiple sets of PDFs in quartz. The central grain shows at least three sets of PDFs, visible from slight contrasts in optical properties. The typical planar, sharp and parallel (within individual sets) geometries of PDFs contrasts with the non-planar appearance of deformation lamellae, shown in Fig. 4.13. (Chapter 8.4). XPL, 0.45 mm, Aruonga Impact crater, Chad. Plate 47. Suevite. Two types of clast can be distinguished. The large clear clast on the left has prominent circular features which are ballen structure. The right hand clast is feldspar. The dark matrix anastomoses around the clasts, defining a weak fabric. (Chapter 8.6). PPL, 1 mm, Suevite, Bosumtwi Impact crater, Ghana. Plate 48. Suevite. The same view as the previous plate in cross polarized light. The clear clasts are non-birefringent. The boundaries between individual ballen are filled by a birefringent phase, probably a clay mineral, suggesting that the ballen structure formed in a similar way to perlitic texture by hydration of a glass. Traces of albite twins can be seen in the feldspar clast, which has the mottled extinction characteristic of mosaicism. However, in view of the evidence for alteration, the mosaicism can not be unambiguously attributed to shock. (Chapter 8.6). XPL, 1 mm, Suevite, Bosumtwi Impact crater, Ghana.

Deformation Micro Structures and Mechanisms in Minerals and Rocks

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