Deformation Micro Structures and Mechanisms in Minerals and Rocks

DEFORMATION MICROSTRUCTURES AND MECHANISMS IN MINERALS AND ROCKS

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Deformation Microstructures and Mechanismsin Minerals and Rocks

by

Tom BlenkinsopDepartment of Geology,University of Zimbabwe, Harane Zimbabwe

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

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Contents

Acknowledgements ix

Symbols, Abbreviations and Units xi

1 Introduction and Terminology 11.11.21.31.41.51.61.7

Introduction 1134455

7777

Classifications of deformation microstructures and mechanismsDeformation microstructures and mechanisms in the earth: Brittle-semibrittle-plastic transitionsThe description of deformation: Scale, continuity, distribution, mechanism and modeDuctility and the “brittle-ductile transition”Character and classification of deformation zone rocksFormat and use of this book

2 Cataclasis2.12.2

IntroductionFundamental cataclastic deformation mechanisms2.2.12.2.2

MicrocrackingFrictional sliding 10

101012121313131314141516171717181818181919222222222223

2.3 Microcracks2.3.12.3.22.3.32.3.42.3.52.3.62.3.72.3.82.3.9

Classification, characteristics and observationMicrostructures and mechanismsImpingement microcracksFlaw-induced microcracksMicrofracturing of pre-existing flawsCleavage microcracksElastic mismatch microcracksPlastic mismatch microcracksMicrofault-induced microcracks: Microscopic feather fractures (mffs)

2.3.102.3.11

Thermally-induced microcracksPhase transformation-induced microcracks

2.4 Microfaults2.4.12.4.2

CharacteristicsMechanisms

2.5 Deformation bands2.5.12.5.2

Characteristics and classificationMechanisms

2.62.72.82.9

Distributed cataclasis and cataclastic flowGouge zone microstructuresMicrofracture surface featuresCrystallographic fabrics

2.102.11

Pre-lithification deformation microstructures and mechanismsPseudotachylites and frictional melting2.11.12.11.22.11.3

CharacteristicsOriginMisidentification

v

vi CONTENTS

3 Diffusive Mass Transfer by Solution3.13.23.33.43.53.6

Introduction2424242525272727272828293030323233333335

3939393941

4147474750

5252525254545454555557575757

595959596060606262626262

Fundamental deformation mechanisms of diffusive mass transfer by solutionGrain surface solution texturesIndenting, truncating and interpenetrating grain contactsStrain capsMicrostylolites3.6.13.6.2

CharacteristicsFormation and propagation

3.7 Diffusive mass transfer and cleavage3.7.13.7.23.7.3

ClassificationSpaced cleavagesContinuous cleavage

3.83.9

Grain surface deposition texturesOvergrowths, porosity reduction, pressure shadows and fringes, and mica beards3.9.13.9.2

CharacteristicsMechanisms

3.103.113.12

Grain shape fabricsFluid inclusion planesMicroveins

4 Intracrystalline Plasticity4.14.24.34.44.5

4.64.74.84.9

IntroductionFundamental mechanisms of intracrystalline plasticityDeformation twinsUndulatory extinctionIntracrystalline deformation bands,kink bands and subgrains: RecoveryDeformation lamellaeGrain shape fabrics and ribbon grainsNew grains, core and mantle structure: Dynamic recrystallizationCrystallographic fabrics

5 Diffusive Mass Transfer and Phase Transformations in the Solid State5.15.25.35.45.55.6

IntroductionFundamental deformation mechanisms of solid state diffusive mass transfer and phase transformationsGrain shape fabrics and ribbon grainsFoam texture, static and secondary recrystallizationDecussate texturePorphyroblasts and inclusion trails5.6.15.6.25.6.3

CharacteristicsGrowth mechanismsRelationship to deformation

5.75.85.9

Reaction rims, relict minerals, coronas and symplectitesChemical zoningSolid state phase transformation microstructures

5.10 Superplasticity

6 Magmatic and Sub-magmatic Deformation6.16.2

IntroductionFundamental deformation mechanisms and microstructures in rocks containing melt6.2.16.2.26.2.3

Magmatic flowSub-magmatic flowMagmatic and sub-magmatic flow and rheology

6.36.4

Mesoscopic evidence for magmatic and sub-magmatic flowMagmatic microstructures6.4.16.4.2

Grain shape fabricsCrystallographic fabrics

6.5 Sub-magmatic microstructures6.5.1 Grain shape fabrics

CONTENTS vii

6.5.26.5.36.5.4

Intracrystalline plasticityDiffusive mass transferCataclasis

6.66.7

Other microstructuresNon-magmatic deformation

7 Microstructural Shear Sense Criteria7.17.27.37.4

IntroductionCurved foliationOblique foliations and shape preferred orientationsPorphyroclast systems7.4.17.4.27.4.37.4.47.4.5

Characteristics and classificationMechanisms of formationStair-step direction: and tailsFaces of a tailDeflection and embayments of tails

7.5 S-, C- and7.5.17.5.27.5.37.5.4

Characteristics and classificationFormation and evolutionCurvature of S-foliationShear on C- or

7.6 Pressure shadows and fringes7.6.17.6.27.6.3

Kinematics of pressure shadows and fringes in shear zonesGeometry of the last increment of growthShape

7.77.87.9

Mica fishPorphyroblast internal foliationsCrystallographic fabrics

7.107.117.12

Asymmetric microboudinsAsymmetric microfolds and rolling structuresShear sense criteria in rocks containing melt7.12.17.12.27.12.37.12.47.12.5

Magmatic shear zonesOblique grain shape fabricsTiling and imbricationS-C fabricsSub-magmatic microfractures

7.13 Shear sense criteria for faults7.13.17.13.27.13.37.13.47.13.5

Shear sense observations on faultsDisplaced grain fragmentsRisers and slickenfibresGougesJogs and bends

8 Shock-induced microstructures and shock metamorphism8.18.28.38.48.58.68.78.88.9

IntroductionShock mechanismsMicrofracturesPlanar Deformation Features (PDFs)MosaicismDiaplectic glassHigh pressure polymorphs of quartz - Coesite and stishoviteLechatelieriteTectites, microtectites and spherules

8.108.11

Shock barometry and thermometryCalibration of shock pressures from microstructures8.11.18.11.2

Calibration of shock pressures from optical properties of quartzProblems of shock barometry

8.12 Diagnostic impact microstructures

6263636363

656566666767686969697070727373737373737475757677777778787878797979797979

808080808182838383848585878788

viii CONTENTS

9 From Microstructures to Mountains: Deformation Microstructures, Mechanisms and Tetonics 9090909090919191919292929394949797989898

9.19.2

IntroductionFailure criteria9.2.19.2.2

Coulomb and Mohr failure criteriaGriffith failure criteria

9.39.49.5

Pore fluid pressure and faultingFracture mechanics and failure criteriaFrictional sliding laws9.5.19.5.2

Byerlee’s lawRate and state dependent frictional sliding

9.6 Flow laws9.6.19.6.29.6.3

Diffusive mass transfer: Grain size sensitive creepIntracrystalline plasticityEmpirical flow laws from experimental data

9.79.89.9

Polymineralic deformationDeformation mechanism mapsLithospheric strength envelopes

9.10 Palaeopiezometry9.10.19.10.29.10.39.10.49.10.59.10.69.10.79.10.89.10.9

Methods and calibrationRecrystallized grain sizeSubgrain size 100

100100101103103104104104104105105

107

127

133

Dislocation densityTwinning - differential stressDeformation lamellaePrincipal stress orientations from deformation lamellaePrincipal stress orientations and strains from twinsGeneral problems with palaeopiezometers

9.11 Geothermobarometry9.11.19.11.29.11.39.11.4

Methods and calibrationCalcite twin morphologySutured quartz grain boundariesSubgrain boundary orientation in quartz

References

Index

Color Plate Section

Acknowledgements

Most of the photomicrographs were developed and printed by Cuthbert Banda, whose assistance, with that of other membersof the staff of the Geology Department, University of Zimbabwe, was invaluable. James Preen guided the preparation of theTEX version of the text. Faith Samkange and Maxwell Matongo were able research assistants, supported by the University ofZimbabwe Research Board. The following are thanked for contributing photomicrographs or thin sections: P. Dirks (Plates11, 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 CrustalStudies, University of California, Santa Barbara, as part of research on deformation mechanisms with R. Sibson, supportedby the National Science Foundation, U.S.A., under grant DAR-84-10924. The assistance of the technical staff at U.C.S.B. isgratefully 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 ofElsevier Science - NL Sara Burgerhartstraat 25, 1055 KV Amsterdam, The Netherlands. Fig. 8.4 was reproduced from Kroghet 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 SpringerVerlag. 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 thetext.

aA

Bc

CLCMFd

D

DMT

E

GBM

G

ISA

J

kClLLPOLSEm

nPPDF

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

b

f

k

xii SYMBOLS, ABBREVIATIONS AND UNITS

PPL

rRRF

S

S

SEMST

TEM

u

U

V

(){ }< >

Pore fluid pressure (3.2, 9.2.3)Plane polarized lightFlaw 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)ReflectedlightDensity, (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 microscopeSensitive tint plateDifferential 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 temperatureUniaxial tensile strength, MPa (9.2.2)Transmission electron microscopeShear 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 lightState variable in dynamic frictional sliding law (9.3.2)Miller-Bravais indices of a crystal faceMiller-Bravais indices of a representative face of a formMiller-Bravais indices of a crystal direction

XP

Chapter 1

Introduction and Terminology

1.1 Introduction

Deformation microstructures in rocks and minerals are mi-croscopic features created by deformation. A deformationmechanism is a process on one scale that accommodates animposed deformation on a larger scale. This book describesmicrostructures and mechanisms at the scale of a thin section,the scale that most geologists use for detailed petrography,based on the premise that many of the fundamental mechan-isms can be inferred from microstructures at this scale. Thebook aims to be a guide to the recognition and interpretationof deformation microstructures and mechanisms, and shouldbe used in conjunction with a petrographic microscope.

Why is the study of deformation microstructures andmechanisms useful ? Deformation mechanisms are determ-ined by temperature, stress (both hydrostatic and deviatoriccomponents), strain rate, pore fluids, mineralogy, and the tex-ture of the deforming rock (especially grain size and poros-ity). Recognition of deformation mechanisms from micro-structures allows limits to be placed on these variables. Forexample, microstructures called subgrains are formed by thedeformation mechanism of intracrystalline plasticity, whichindicates deformation at temperatures above 250°C in quartz.The size of the subgrains can be used to gauge the deviat-oric stress during deformation. These are essential pieces ofinformation for tectonic analysis and for understanding thebehaviour of the lithosphere by mathematical modelling (e.g.Kusznir and Park 1987, Molnar 1992, Beaumont et al. 1996).

Microstructures and deformation mechanisms are a grow-ing field of interest in the earth sciences. The application ofmaterials science to minerals and rocks has provided much ofthe new impetus. However, the literature is scattered throughjournals in a large number of disciplines, and most structuralgeology textbooks have limited coverage of the field. At leastsome of this diverse literature is reviewed in this book, whichcontains a comprehensive reference list. Hopefully the bookis written in sufficient depth to be useful at advanced under-graduate level and above. Familiarity with elementary con-cepts and terms in structural geology is assumed.

1.2 Classifications of deformation mi-crostructures and mechanisms

Deformation microstructures in minerals or rocks are the re-cord of permanent deformation, i.e. shape and/or volume

changes that remains after stress is removed, as opposed toelastic (recoverable) deformation which is not seen directlyin the geological record. Deformation microstructures can bedivided into three major categories:

1.

2.

3.

Microfractures, displacement, and/or rotation of rigidparticles with no permanent lattice distortion. The typ-ical microstructures seen in thin section are microfrac-tures and fragments surrounded by a fine-grained mat-rix.

Microstructures showing material removal, transportand deposition without fracturing, permanent lattice dis-tortion or melting. Examples of microstructures at sitesof material removal are distinctive types of grain con-tacts and microstylolites. Microstructures indicatingmaterial deposition include microveins, overgrowths,pressure shadows and porphyroblasts. Many meta-morphic textures associated with deformation fall intothis category.

Permanent lattice distortion without fracturing. Ex-amples of typical microstructures are undulatory ex-tinction, subgrains, recrystallized grains, and crystallo-graphic fabrics.

This simple classification scheme can be applied on the basisof optical microscope observations. Table 1.1 summarizesthe scheme, gives examples of specific microstructures, andrefers to the relevant chapters and sections of the book.

Table 1.1 also shows the relation between this scheme anda classification of deformation mechanisms, which has threesimilar categories (e.g. Knipe 1989):

1.

2.

3.

Cataclasis - deformation by microfracturing, slidingand/or rotation of rigid particles (Chapter 2). Brittle de-formation is often used as a synonym for cataclasis, butis more accurately defined as deformation by fracture ormicrofracture.

Diffusive mass transfer (DMT) - deformation by diffu-sion, the movement of lattice defects, ions, atoms ormolecules in response to gradients of chemical potential(Chapters 3 and 5).

Intracrystalline plasticity - deformation by the move-ments of extra half-planes of atoms (dislocations) in acrystal lattice (Chapter 4).

1

2 CHAPTER 1. INTRODUCTION AND TERMINOLOGY

CHAPTER 1. INTRODUCTION AND TERMINOLOGY 3

A general term for both categories 2 and 3 is useful becausethey are often closely associated with each other: crystalplasticity or simply plasticity means deformation by eitheror both DMT and intracrystalline plasticity. DMT can besplit 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 occurduring deformation, some involving DMT: these are includedin Chapter 5.

Two or more mechanisms may act simultaneously togetherwithin a single mineral. An example is the combination ofsliding on grain boundaries, and mass transfer of material bydiffusion to fill space created by sliding. This is is one typeof superplasticity, which is a composite deformation mech-anism; one in which two or more mechanisms are coupledtogether in the same mineral under the same conditions. Animportant composite deformation mechanism involves the in-teraction of microfractures and intracrystalline plasticity: thisis called semibrittle deformation. The different componentsof a composite mechanism may have variable strain rates: themechanism with the slowest rate determines the compositestrain rate, and is said to be rate-limiting.

It is fortunate that the non-genetic classification of the mi-crostructures matches the genetic classification of the mech-anisms so well, and this means that both classifications can bereferred to by similar names. The chapter titles of this bookuse the names of the mechanisms for simplicity. The order ofthe chapters follows the general change in deformation mech-anisms from low to high grade metamorphic conditions dur-ing deformation (see next section). Specific microstructuresthat are diagnostic for each mechanism are shown in bold inTable 1.1.

Three chapters deal with relatively new developments instructural geology, which can involve all three categoriesof deformation microstructures and mechanisms. Magmaticand sub-magmatic deformation (Chapter 6) describes the de-formation of rocks that contain melt. Important deformationmechanisms are flow of melt and crystals, with crystal de-formation (sub-magmatic flow) or without crystal deforma-tion (magmatic flow). This topic has great relevance to cur-rent debates about pluton emplacement mechanisms. Micro-structural shear sense criteria (Chapter 7) provide clues tothe displacement of rock masses during deformation on allscales: this is one of the most important types of tectonicinformation. Shock-induced microstructures and shock meta-morphism are produced by meteorite impacts (Chapter 8), andare the focus of much current interest, because microstruc-tural studies have a central role to play in the debates aboutmass extinction and other possible effects of large meteoriteimpacts on the earth’s evolution.

1.3 Deformation microstructures andmechanisms in the earth: Brittle-semibrittle-plastic transitions

Several different deformation microstructures commonly oc-cur together in rocks for three important reasons. Firstly,deformation microstructures may record several deformation

events, each of which may be associated with different mech-anisms. One of the applications of this book should be toallow the unravelling of multiple deformations from their as-sociated microstructures. Secondly, even within a single de-formation, mechanisms and microstructures vary from min-eral to mineral within a polymineralic rock: a common ex-ample is the intracrystalline plasticity of quartz in a shearzone at greenschist facies, that contrasts with cataclastic de-formation of feldspar in the same conditions.

Thirdly, one mechanism may be incapable of accommod-ating the imposed stress, strain or strain rate, so that othermechanisms are substituted or added during the same de-formation: for example, faulting (cataclasis) may relieve highstress levels in a shear zone otherwise deforming by intracrys-talline plasticity. The record of both the cataclasis and plas-ticity may be preserved in the microstructures. Mechanismsthat can only accommodate deformation in a single direction,for example slip on a single fault set, are especially restrictiveand invariably require supplementary mechanisms.

The variation in conditions in the earth, particularly of tem-perature and pressure, causes a corresponding variation in de-formation mechanisms. Two systematic changes occur withincreasing depth: temperature increases by the geothermalgradient, and pressure increases due to the effect of gravity.Cataclasis is relatively insensitive to the variation in temper-ature, but is suppressed by pressure. The plastic deformationmechanisms (intracrystalline plasticity and solid-state DMT)behave in the opposite way: they are strongly promoted bytemperature but relatively insensitive to pressure. As a res-ult, cataclasis is generally restricted to the upper crust (wherepressures are low), and crystal plasticity occurs in the lowercrust and the rest of the lithosphere (where temperatures arehigh). The transition from cataclasis to plasticity is some-times called the brittle-plastic transition. Experiements haveidentified an important intermediate regime of semibrittle be-haviour where microfractures interact with intracrystallineplasticity (e.g. Carter and Kirby 1978, Kirby and Kronen-berg 1984), leading to the concept of two transitions in de-formation mechanism with increasing depth: firstly the brittle- semibrittle transition in the upper crust, and secondly thesemibrittle - plastic transition in mechanisms at mid-lowercrustal levels.

These generalizations need to be heavily qualified becauseof the other variables that affect deformation mechanisms.Stress, strain rate and pore fluids are some of the most im-portant additional variables to be considered. For example,high stresses or strain rates may cause cataclasis at greaterdepths or higher temperatures, which would otherwise be as-sociated with plasticity, as in the above example. Pore flu-ids are essential for solution-assisted DMT, and may promotecataclasis by mechanical and chemical effects. Furthermore,the various minerals within a polymineralic rock have differ-ent transitions from brittle to semibrittle to plastic behaviour,and interactions between the minerals themselves also need tobe considered as an independent variable (Chapter 9). Thesequalifications, particularly the variation in properties of dif-ferent minerals, mean that no unique depth or temperaturecan be given for the brittle-semibrittle and semibrittle-plastictransitions. However, the transitions are important concepts


for understanding the behaviour of the earth, and the condi-tions for the transitions can be predicted for specific modelsof 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 interms of three fundamental attributes. Continuity is the con-nectivity of material points through the deforming body; de-formation can be characterized as continuous if points remainconnected, or discontinuous if not. The distribution of de-formation can be described as localized (e.g. a shear zone) ordistributed. The deformation mechanism can be described inone of the three categories given in Section 1.2. All three at-tributes depend on the scale of observation. Scales are looselydefined in this book as macro (greater than outcrop, i.e. > 10m), meso (outcrop, i.e. 1 cm - 10 m) or micro (microscopic,i.e. < 1 cm).

Permanent deformation of minerals or rocks involvesbreaking atomic bonds and is therefore is discontinuous at theatomic scale. As the scale of observation is increased, theseatomic discontinuities can not be discerned, and deformationappears to be continuous. The top surface of Fig. 1.1 showsschematically how deformation continuity is a function of thescale of observation. An example of the discontinuous to con-tinuous transition with increasing scale is the accommodationof a fold by cataclastic flow (Chapter 2.6). At micro to mesoscales, the deformation is by discontinuous fracture, but at a

macroscopic scale, these discontinuities are not seen and thefold appears to be continuous.

Continuity should be described at the time of deformation,but can easily be altered by subsequent events. Another prob-lem with specifying deformation continuity is posed by fea-tures such as overgrowths, pressure shadows, and pressurefringes (Chapter 3.9), which may be continuous with the sur-face from which they are grow, and discontinuous with thesurface towards which they grow. Boundaries between grainsare perhaps best regarded as continuous on a microscopicscale during recrystallization, but after recrystallization theyappear as discontinuities. These examples show that a certainamount of judgement may be necessary to describe continu-ity, which is a necessary shortcoming of any description thathas to take into account the past history of deformation inminerals and rocks.

Deformation distribution is also highly scale-dependent.As scale of observation is increased, an localized deforma-tion may appear pervasive: for example, microfractures arehighly localized deformation on a microscopic scale, but anetwork of microfractures can have the effect of a pervasivestrain on the scale of an outcrop or a regional map.

The combination of deformation distribution (local-ized/pervasive) and mechanism (cataclastic or plastic) wasdescribed as a “mode of failure” by Rutter (1986). Thisconcept can be extended through the incorporation of deform-ation continuity and scale. The combination of continuity,mechanism and scale can be called a deformation mode, andrepresented on a diagram such as Fig. 1.1, where a modeis specified by deformation mechanism (z-axis), continuity(which can be qualified by distribution, x-axis) and the scalelength (y-axis). The front of the diagram shows the field ofpossible deformation modes described in this book (i.e. at themicroscopic scale), with examples of some microstructures.

Figure 1.1 makes some important links between micro-structures and mechanisms. Cataclastic deformation mi-crostructures are discontinuous, and intracrystalline plasti-city microstructures are continuous, at the microscopic scale.These observations point towards one of the major themesof this book: deformation microstructures can be used, al-beit with care and a certain amount of ambiguity, to identifydeformation mechanisms. This link is possible by inductionand deduction from the microstructures and theoretical un-derstanding of the mechanisms, and by comparison with ex-periments.

1.5 Ductility and the “brittle-ductiletransition”

Great confusion has been caused by the use of the termsductile and brittle-ductile transition. Much of this confu-sion can be traced to the dichotomy between laboratory (andmaterials science)-based and field-based approaches to de-formation (the subject is well reviewed in Evans et al. 1990and Williams et al. 1994). For example, rock mechanicsexperiments use a maximum permanent strain before fail-ure of more than 5% to define ductile behaviour (e.g. Pa-terson 1978). Similarly, Griggs and Handin (1960) defined

CHAPTER 1. INTRODUCTION AND TERMINOLOGY 5

relative ductility as: “the amount of permanent deformationachievable prior to rupture”. These definitions of ductilitybased on stress-strain relationships are satisfactory and quan-tifiable, but the widespread application of the term outsidethe rock mechanics laboratory necessitates alternative defin-itions, since stress-strain relationships are never known forrocks in the field (e.g. Griggs and Handin 1960).

The use of deformation continuity or distribution in thedefinition of ductile is implicit in the application of the termsby most geologists in the field. Rutter (1986) suggested thatductility is “the capacity for substantial, non-localized strain”,thus strictly excluding shear zones, and “is a concept whichmust be defined on a macroscopic scale”. However, “ductileshear zones” (e.g. Ramsay and Huber 1987) is a commonlyused phrase. It is proposed here that continuity rather thandistribution should be used to define ductility, and that a moresatisfactory definition of ductility is “macroscopically con-tinuous deformation”. This definition has the merit of encom-passing all features that the field geologist usually refers to asductile, including shear zones and macroscopic folds, and ispreferred because continuity can be identified more preciselythan distribution at any scale.

Brittle and ductile are defined above by different cri-teria. Therefore it is possible for a rock to be both brittleand ductile: for example, a type of deformation band(Chapter 2.5) forms by fracture (it is brittle) but has macro-scopic strain continuity (it is ductile). The concept of the term“brittle-ductile transition” is rendered questionable by thesedefinitions. Many terminological problems can be avoidedby using the concept of a deformation mode and the cat-egories of deformation continuity, distribution and mechan-ism suggested in Section 1.4, and avoiding the use of theterm brittle-ductile transition, which has no meaning usingthe above definitions.

1.6 Character and classification of de-formation zone rocks

Rocks within deformation zones are commonly highlystrained and may exhibit some of the best examples of de-formation microstructures. Deformation reduces the grainsize of some of the protolith to produce a matrix of finergrains surrounding remnant larger grains or grain aggregates,known as porphyroclasts or clasts, which is the typical micro-structure of deformation zone rocks. Deformation mechan-isms and microstructures may differ between the matrix andporphyroclasts.

Some definitions and classifications of deformation zonerocks have advocated the use of deformation mechanisms asa classificatory tool. For example, Wise et al. (1984) pro-posed to include crystal plasticity as an essential elementof the definition of mylonites. As for any other classifica-tion of observational data, the use of mechanisms should beavoided because subjective interpretations are required, thatmay change in the light of new knowledge. Nomenclatureand classification of deformation zone rocks is extensivelydiscussed 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 dis-tinction between random fabric and foliated deformation zonerocks, as well as the cohesion of the rock (the degree to whichit behaves as a continuous body during deformation) and pro-portion of matrix to porphyroclasts. The original scheme isslightly modified to allow for a range of foliation intensities infault rocks in Table 1.2, by including the additional categoryof foliated cataclasites (e.g. Chester and Logan 1987), andby extending the gouge, breccia and pseudotachylite categor-ies to the foliated category. Random fabric has been changedto Unfoliated to allow for deformation zone rocks that havesome order to their fabric, for example fragments with a jig-saw texture, and Foliated has been changed to Strongly fo-liated in order to contrast with Unfoliated. Many deforma-tion zone rocks have two or even three distinct foliations. Allfoliations should be considered when assessing foliation in-tensity, and the existence of different foliations can be usedto further classify deformation zone rocks (e.g. S-C mylon-ites, Lister and Snoke 1984). The crush breccia series of theoriginal classification has been omitted for simplicity, and be-cause these terms have not found widespread use.

The classification scheme in Table 1.2 is based on descrip-tion at the hand specimen scale, and recognizes that largerfragments in a finer grained matrix is a fundamental characterof most deformation zone rocks. However, as noted by Sibson(1977), any pidgeon-hole classification such as Table 1.2 suf-fers from the problem that the classification criteria may showa continuous range of variation. This is especially problem-atic in the assessment of foliation development, which may bedifficult to quantify, or to judge objectively on a qualitativebasis. Another problem with the classification is the defini-tion of cohesion, which is specified in the original classific-ation as primary cohesion, i.e. cohesion during deformation.Post-tectonic processes may decrease cohesion, for exampleby weathering, or increase cohesion by cementation. Theseprocesses may not be recognized easily or allowed for whenassessing primary cohesion.

1.7 Format and use of this book

Chapters 2 to 5 deal with the major categories of deform-ation microstructures and mechanisms. Each chapter beginswith a brief introduction to the fundamental mechanisms, andproceeds to descriptions of the characteristic microstructures,identifying which are diagnostic, and illustrating them withdiagrams and black-and white photomicrographs embeddedin the text. Colour photomicrographs are collected separatelyand referred to as plates in the text. Key words are emphas-ized where they occur for the first time in the chapter. Somemechanisms and microstructures, especially cataclasis, aredealt with in more detail than others because comprehensivedescriptions are lacking in the literature. The mechanisms in-volved in the formation of each microstructure are interpretedfrom experimental and theoretical backgrounds. The finalchapter shows how deformation microstructures and mech-anisms can be used to make quantitative inferences about de-formation conditions for tectonic analysis. Throughout thetext, relatively new developments in the subject, and thosethat have not been described in previous textbooks, have been


more heavily referenced than other topics. The references arepresented in two forms. The main list gives all references infull. This is followed by a list of abbreviated references col-lected by chapter, which shows important general sources forthe chapter topics in italics.

A general list of symbols, abbreviations and units prefacesthis chapter. Abbreviations and conventions for all photomic-rographs 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 ofthe image in mm. Shear sense (see Chapter 7) is given assinistral or dextral (assuming that the shear or fault plane isvertical); all illustrations with shear senses given are perpen-dicular to the shear plane and parallel to the shear direction,and the shear plane is approximately parallel to the horizontaledge of the image.

Chapter 2

Cataclasis

2.1 Introduction

Microfractures, displacements and rotations of rigid particleswith no permanent lattice distortion are microstructuresformed by cataclasis. There are two fundamental cataclasticdeformation mechanisms: microcracking (Section 2.2.1) andfrictional sliding (Section 2.2.2). The single most distinct-ive cataclastic microstructure is the microfracture, definedas a tabular or planar microscopic discontinuity. The termthus includes microfaults, microscopic deformation bands,microjoints, microcracks, microveins, and microscopic slipsurfaces. Microfractures can be sub-divided into microfaults(Section 2.4), which contain a fragmental matrix, and mi-crocracks (Section 2.3), which do not. Pseudotachylites arebriefly discussed in Section 2.11 because they are associatedwith cataclastic mechanisms.

2.2 Fundamental cataclastic deforma-tion mechanisms

2.2.1 Microcracking

Microcracking involves microcrack nucleation followed bypropagation. Nucleation is irrelevant in a geological contextbecause of abundant heterogeneities in natural rocks and min-erals.

Dynamic microcrack propagation

Microcrack propagation from an initial flaw can be con-sidered by two different models. The first model assumesan elliptical microcrack (Inglis 1913, Griffith 1924). For auniaxial remote applied stress parallel to the microcrackaxis, the tangential stress, on the microcrack surface var-ies from a negative (tensile stress) equal to the value of atthe long axis of the microcrack, to a compressive stress at theshort 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 ellipticalmicrocrack illustrates the essential concepts that large tensilestresses can develop at the tip of a microcrack surface in com-pression, and that the maximum tensile stress will developin the direction of the maximum applied stress. These twoconcepts explain why extension microcracking perpendicularto the least principal stress is a widespread and fundamentalcataclastic mechanism.

The second model treats the microcrack as flat with a sharptip (e.g. Lawn and Wilshaw 1975a), and gives a versatilesolution for the stress field around the microcrack, in the polarcoordinates of Figure 2.1b:

The stress is thus specified by K, the Stress intensity factor,describing the intensity of the field around the microcrack,and the stress distribution, described by radial factorand a function of which depends on the propagation mode.The three microcrack propagation modes shown in Figure 2.2are tensile or opening mode, (mode I), in which displace-ments are perpendicular to the fracture plane, and two shearmodes, in which displacements are parallel to the fractureplane, sometimes called sliding and tearing modes (modes IIand III, Fig. 2.2). The stress intensity factor depends on thepropagation mode, the microcrack length and the appliedstress

Given these descriptions of the stress around the flat mi-crocrack, it is possible to deduce the failure criteria for brittlesolids from Griffith’s postulate that a microcrack will extendwhen the total energy change with the propagation of the mi-crocrack is negative or constant. The energy terms involvedin microcrack propagation are release of mechanical energy,which must be equal to the energy required for creation ofsurface area.In terms of force, the microcrack extension force,G, must be greater than or equal to twice the surface tensionforce, (the doubling factor accounts for the two sides ofthe crack). This leads to the classical Griffith result for fail-ure under a tensile remote stress in a solid with Young’smodulus E:

The well-known Griffith failure criterion and its derivativescan be determined from this relation (Section 9.2). The Grif-fith failure criterion thus implicitly assumes the existence ofmicrocracks of length Griffith observed that the theoreticalstrength of solids is much greater than their actual strength,and postulated that real solids contain microcracks whichconcentrate stress leading to failure, thus explaining the dis-crepancy between theoretical and measured strengths. Theexistence of such microcracks or “Griffith flaws” is still ac-cepted as the basis of most failure criteria.

7

8 CHAPTER 2. CATACLASIS

CHAPTER 2. CATACLASIS 9

The model so far assumes elastic behaviour, but an addi-tional energy term must be incorporated to account for break-ing of atomic bonds at the crack tip, which can be done bypostulating the existence of a small zone ahead of the micro-crack tip in which non-linear forces are expended in breakingbonds. The additional energy involved is incorporated in theenergy balance in the form of a new parameter the fracturetoughness, to replace the surface energy term The energybalance condition is now:

This equation gives the condition for microcrack propagationin terms of a value for G usually known as the critical frac-ture toughness, which can be related to the critical stressintensity factor, for a given geometry.

This analysis is valid provided that the size of the non-linear zone is much smaller than the length of the microcracki.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 Wilshaw1975a). More detailed analyses can relate to the dis-placement on the microcrack if some function of the cohesiveforces ahead of the tip with distance is postulated: these arethe cohesive force models (e.g. Rudnicki 1980). The exist-ence of such a non-linear zone ahead of the microcrack tip isknown experimentally from ceramics and metals, where it isreferred to as a process zone, and it has been detected fromacoustic 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 propag-ate at speeds that are typically significant fractions of the ve-locity of elastic waves in solids, or dynamic microcracking.However, microcracks may propagate under stress conditionswell below the critical stress intensity factor, at rates thatdepend on temperature and chemical environment as well asstress intensity. This phenomenon is known as sub-criticalmicrocrack growth and is potentially of enormous geologicalimportance (e.g. Anderson and Grew 1977, Das and Scholz1981, Atkinson 1982, Meredith 1983). A diagram showingmicrocrack velocity (V) as a function of mode I stress intens-ity factor illustrates the main features of sub-critical mi-crocrack growth (Fig. 2.3), which have been demonstrated fora range of geological materials (e.g. quartz, granite, andesite,basalt, calcite, oil shale, sapphire and glass). Therelationship falls into three parts:

Region 1. Velocity is highly sensitive to water concen-tration and temperature. There may be a threshold stressintensity factor for microcrack growth to occur atall (Meredith 1983).

Region 2. Velocity is dependent on water concentration andtemperature, but not

Region 3. Velocity increases extremely rapidly with untilmicrocrack growth becomes dynamic at

Five mechanisms of sub-critical microcrack growth have beenproposed, but stress corrosion, a general term for environ-mentally influenced, stress driven, thermally activated chem-ical reaction allowing breaking of bonds is considered to bethe dominant mechanism for geological materials in crustalconditions (Atkinson 1982). Hydrolysis of the Si-O-Si bondis likely to be responsible for the weakening. The microcrackvelocity in Region 1 is controlled by the reaction rate of thehydrolysis, and transport-rate control occurs in Region 2.

Other types of chemical reaction occur with sub-criticalmicrocracking. An example is the reaction of plagioclase toincreasingly sodic compositions and ultimately to laumontite,which creates a 60% volume increase (Blenkinsop and Sibson1991). This alteration occurs along cleavage microcracks,resulting in a texture of expanded, matching fragments, de-rived from a single parent crystal (Plate 1). This texture sug-gests chemical alteration and sub-critical microcracking oc-curred in a linked process, which can be called alteration-enhanced microcracking.

Sub-critical microcrack growth has been incorporated intosome models of crustal processes, for example to explain thedifference between creep (stress intensity factor betweenand and seismic faulting (stress intensity factor equal to

Rudnicki 1980), the barrier theory for earthquake ruptureof Das and Scholz (1981), and features of magmatic intru-sion and hydrofracture propagation (e.g. Anderson and Grew


1977, Atkinson 1982).

2.2.2 Frictional sliding

Amonton’s law that the steady state shear stress is propor-tional to the normal stress on the sliding surface via thecoefficient of friction

is predicted by a simple adhesion theory of friction. The the-ory assumes that the rough surfaces contact at asperities (pro-truding irregularities), which will yield under normal stressuntil sufficient area of contact is established to support thenormal load (Fig. 2.4). The contact area is considered to havean adhesive strength which must be exceeded over the entirearea for sliding to occur. This model successfully predictsthe normal-stress dependence of friction, but it underestim-ates the value of because there are other contributions tothe shear strength in addition to the adhesive strength of thecontacts (Scholz 1990). These include:

1.

2.

3.

Interlocking of asperities. Oblique surfaces of contactare created between two asperities that come into contactafter sliding (Fig. 2.4). Interlocking may be relieved byshearing through asperities (adhesive wear), or by slid-ing on the oblique contacts, which causes dilatancy.

Increasing area of contact due to asperity shearing. Assliding continues, asperities fail and the contact area in-creases.

Asperity ploughing. An asperity with a greater strengththan the opposite surface will cut into the weaker ma-terial, generating a groove and wear fragments. This isknown as abrasive wear.

Adhesive strength, asperity interlocking and increasingcontact area have been combined in an elegant and simplemodel by Wang and Scholz (1994, 1995), which accounts forexperimental results very well.

Frictional sliding leads to the production of gouge by as-perity failure and ploughing. Once formed, fragmentationwithin a gouge layer continues by mutual impingement ofparticles, leading to particle size distributions (PSDs) that arecharacteristically fractal (e.g. Blenkinsop 1991). A fractaldistribution of particle sizes can be described by the rela-tion:

where N (d) is the number of particles greater than size d, andD is the fractal dimension. D for particles in cataclastic rocksranges from 1.88 to 3.08 (e.g. Sammis et al. 1987, Sammisand Biegel 1989, Olgaard and Brace 1983, Wang 1987), withmany results for natural and experimental gouges around thevalue of 2.6 (Marone and Scholz 1989, Biegel et al. 1989).Sammis et al. (1987) showed that a distribution of spheresof unequal sizes reduces impingement stresses on individualfragments, and proposed that microcracking in a gouge willproceed in order to minimize the probability of neighbour-ing grains having equal sizes. This constrained comminutionmodel predicts D = 2.58, which is very close to observed val-ues (e.g. Sammis and Biegel 1989).

Gouge formation (and cataclasis generally) may occur byat least two other processes of particle size reduction. Alter-ation, for example the laumontization of feldspar describedabove, may lead to lower values of D (Blenkinsop and Sib-son 1991). At advanced stages of gouge formation, selectivemicrofracture of larger particles takes place, creating PSDswith fractal dimensions greater than 2.58 (Blenkinsop 1991).Plates 1 to 4 show a sequence of cataclastic textures arrangedin order of increasing fractal dimension of PSDs, which is theevolutionary sequence of textures with strain.

A simple theory of wear can predict that the volumeof gouge created by sliding, and hence the thickness ofthe gouge, will increase linearly with displacement (Scholz1990), and some experimental results confirm this relation-ship (e.g. Teufel 1981). This relationship has also beenclaimed for faults (e.g. Robertson 1983, Hull 1988). How-ever, there are at least two reasons the experimental datashould not be applied directly to faults. Firstly, the roughnessof laboratory prepared surfaces is not comparable to naturalfault surfaces (Brown and Scholz 1985), and secondly, oncea gouge layer has accumulated sufficient thickness to preventinteraction between the sliding surfaces, wear will no longeroccur according to the simple model of surfaces in contactwith each other. The reported data on fault gouge thickness-displacement relationships does not substantiate a proportion-ality (Blenkinsop 1989).

2.3 Microcracks

2.3.1 Classification, characteristics and obser-vation

Microcracks can be classified as intragranular (within singlegrains), transgranular (across two or more grains) and cir-


cumgranular or grain boundary. The occurrence of these dif-ferent types of microcrack depends on the microcrack mech-anism (see below) and on the microstructure of the rock. In-tragranular microcracks are characteristic of poorly cemen-ted, highly porous rocks, whereas well-cemented, low poros-ity rocks have transgranular microcracks. This classificationis useful in discriminating various microcrack mechanismsdescribed in Sections 2.3.3 to 2.3.11.

Tectonic trans- or intragranular microcracks commonlylink contact points between adjacent grains, and are kinkedor curved. Despite grain-scale irregularities in fracture geo-metry, microcracks generally have a strong preferred orienta-tion. Several sets of microcracks may exist within one rock.Extension (mode I) microcracks are usually filled with ma-terial in optical continuity with the host grains, so that theirimportance may be underestimated or even completely over-looked. The only manifestation of a fracture left in the rockmay be a line of fluid inclusions which are the trace of a fluidinclusion plane (Fig. 2.5, Section 3.11), and careful obser-vations of well-prepared thin sections at high magnificationsunder the optical microscope are necessary to detect such mi-crocracks if they contain small inclusions. Microcracks canhave a wide range of aspect ratios and densities. Extensionmicrocracks may have regionally systematic orientations be-cause they form perpendicular to (Section 2.2.1.1), andhave been used very effectively to deduce regional stress sys-tems (e.g. Lespinasse and Pêcher 1986).

Cryptic microcracks may be spectacularly revealed bycathodoluminescence (CL) (e.g. Smith and Stenstrom 1965,Sprunt et al. 1978, Sprunt and Nur 1979, Blenkinsop andRutter 1986). Luminescence depends on silica polymorph,and chemical, thermal and mechanical histories (e.g. Seye-dolali et al. 1997). Microfracture fillings which form underdifferent conditions from the host grain, and experience onlypart of their tectonic history, therefore contrast in lumines-cence with the rest of the grain and usually have very lowluminescence (Fig. 2.6).

2.3.2 Microstructures and mechanisms

Nine microcrack “mechanisms” can be distinguished, mainlyfrom experiments, where extension microcracks, usuallyknown as axial microcracks, form from about half the peakstrength through to post-failure (e.g. Tapponier and Brace1976). These are secondary mechanisms compared to thefundamental physics of tensile microcracking described inSection 2.2.1, and they mainly describe specific geomet-ries that create microcrack tip tensile stresses (Krantz 1983,Atkinson 1982). Table 2.1 summarizes the characteristic fea-tures of the mechanisms, including the types of microcrack(intra-, trans- or circumgranular) that can form by each mech-anism.

2.3.3 Impingement microcracks

Impingement microcracks link points of contact between ad-jacent grains, and are usually intragranular. They may linkseveral pairs of contact points around grains. Four basicpatterns are shown in Fig. 2.7 (Gallagher et al. 1974),which depend on the boundary loads, packing arrangement,

size, sorting, and grain shape. Impingement microfracturingcan be understood from an analysis of the stress field cre-ated on loading a plane surface by a pointed (Boussinesqconfiguration) or spherical (Hertzian configuration) object(Fig. 2.8). The point load should produce radial microcrackssub-perpendicular to the indenter, and such microcracks havebeen observed in indenter experiments (e.g. Lindquist et al.1984). The spherical indenter (possibly more geologicallyrealistic) will contact the plane surface along a spherical sur-face, inducing a region of compressive stresses immediatelybelow the indenter, surrounded by tensile stresses near theedge of the contact surface. An extension microcrack in theform of a cone (cone microcrack) forms beneath the indenter(Fig. 2.8) at a critical load. The critical indenter radius to gen-erate a tensile microcrack depends on the applied pressure viaa square root (Lawn and Wilshaw 1975b): very modest pres-sures for even quite large indenters can create microcracks.

A still more geologically realistic configuration is the caseof two spheres loading each other: in this case, an extensionmicrocrack initiates at the edge of the contact plane betweenthe two spheres at a critical load that depends on porosity,grain size, elastic moduli and fracture toughness, providingthat there are pre-existing flaws present in the loaded grains(Fig. 2.8, Zhang et al. 1990). The critical load measured in


several rocks (Wong 1990) follows the theoretical relation-ship and shows that the required pre-existing flaws have sub-micron dimensions, comparable to the flaws invoked in Grif-fith’s failure analysis (Section 2.2.1).

Photoelastic experiments on convexo-concave contact sur-faces which model indenting grain contacts due to pressuresolution (Section 3.4) show that tensile microcracks shouldform normal to the indenter contact, and shear failure is pre-dicted along curved trajectories that are approximately nor-mal to the contact in its immediate vicinity, but deviate pro-gressively with distance (McEwen 1981).

The diagnostic feature for recognizing the impingementmechanism is linking of contact points by intragranular mi-crocracks (Fig. 2.9, Table 2.1), although the contact pointsmay not be visible in the plane of the section. Impingementmicrocracking is suppressed in well-cemented or low poros-ity crystalline rocks because impingement contacts are lack-ing and tensile stress concentrations are dramatically reducedby the cement (e.g. Wong and Wu 1995, Menéndez et al.1996).

2.3.4 Flaw-induced microcracks

Flaw-induced microcracks are joined to flaws such as othermicrocracks, pores and grain boundaries. They form becauseof the tensile stresses that develop on the flaw surface whenremote stresses are imposed. They are recognized in experi-ments on analogues and on rocks (e.g. Brace and Bombola-kis 1963, Tapponnier and Brace 1976). Analytical solutionsshow that the microcracks will grow along curved trajectoriesfrom both ends of a flaw to produce the well-known “wingcrack” geometry (Fig. 2.10, Horii and Nemat-Nasser, 1985,1986, Kemeny and Cook 1987, Baud et al. 1996). For isol-ated microcrack growth, a flaw length 2c, orientated at to

is assumed to have both cohesive and frictional resistanceto shear, and tensile microcracks grow from both ends with

a critical length (approximately 1.0) in even the slightesttensile value of microcrack growth becomes unstable, butremains stable in compression. Flaw-induced microcrackingcan be recognized by the connection of the microcrack to a

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 import-ant and even dominant microcracking mechanism from ex-periments (e.g Biegel et al. 1992, Menéndez et al. 1996).It is considered to be at least as important as cleavage incontrolling 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 Brace1976). The weakness of some natural grain boundaries, evenwhen these are overgrown by optically continuous quartz ce-ments, is apparent from the observation that overgrowths area major component in the matrix of faults (Pittman 1981).Microfracturing of pre-existing flaws can be identified by mi-crofracture of a cement, and may involve all three types ofmicrocrack (Fig. 2.11, Table 2.1).

2.3.6 Cleavage microcracks

Microcracks in biotite are controlled by the basal (001) cleav-age (Plate 42, Wong and Biegel, 1985). In feldspars, the ma-jor (001) cleavage and also (010), and (110) planes ex-ert a strong influence on microcracks (Willaime et al. 1979,Brown and Macaudière 1984, Tullis and Yund 1992). Cleav-age microcracking of feldspars is important during deforma-tion of granitic rocks in the upper crust (Plate 1; Evans 1988).The fracture toughness of quartz is least along the rhombo-hedral planes, followed by the basal plane: these are prefer-entially exploited during fracture (e.g. Borg et al. 1960, Voll-brecht et al. 1991). Cleavage microcracks can be recognizedbecause they occur in crystallographically controlled sets par-allel to known cleavages within single grains (Table 2.1).

2.3.7 Elastic mismatch microcracks

Microcracks have been noted in quartz and feldspar grainsat contacts with micas in experimental studies (Tapponierand Brace 1976, Wong and Biegel 1985), and in a naturallydeformed quartzite (Hippert 1994). These microcracks are

length l (Fig. 2.10), making an angle with the flaw. The res-ults show that microcracks will grow by tensile failure fromthe edges of the flaw along paths which fit experimental ob-servations very well (Horii and Nemat-Nasser 1985). After


thought to develop because of the difference in elastic strainacross the quartz-mica or feldspar-mica boundaries due to thedifferent elastic moduli of the two minerals in contact alonga coherent interface (Plate 5, Fig. 2.11). They can be recog-nized by the localization of intragranular microcracks aroundcontacts between grains of different mineralogy (Table 2.1),but may be difficult to distinguish from thermally-inducedmicrocracks (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), mi-crocracks may be initiated due to the strain incompatibilitybetween the area of plastic deformation and the adjacent un-deformed area (e.g. Olson and Peng 1976, Wong and Biegel1985; Fig. 2.12). Microcracks have been observed alongkink bands in naturally deformed enstatite, and normal tothe kink bands in experimentally deformed quartz (Carterand Kirby 1978). Plastic mismatchs may account for thecommon microcracking of feldspar porphyroclasts surroun-ded by deformed quartz grains in quartzofeldspathic mylon-ites (e.g. Evans and White 1984; Fig. 2.13). Plastic mis-matches may also occur within single phases or grains dueto 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 bythe close association between intragranular microcracks andareas or individual microstructures of intracrystalline plasti-city, such as subgrains, kink bands, deformation lamellae ortwins (Table 2.1).

2.3.9 Microfault-induced microcracks: Micro-scopic feather fractures (mffs)

Mffs are intragranular microcracks found only adjacent tofaults. They are characteristically wedge-shaped, opening to-wards the fault plane (Fig. 2.12c). They were identified in ex-perimentally generated faults by Friedman and Logan (1970),who found them exclusively within 5-10 grain diameters ofshear faults, and parallel to They did not occur adja-cent to an incipient shear, and therefore formed in responseto shearing. Conrad and Friedman (1976) defined mffs as mi-crocracks occurring only within grains adjacent to a fault, dy-ing out rapidly away from the fault and statistically close to orparallel to the applied direction of Microcrack density andlength increase with displacement and normal stress (confin-ing pressure) (Conrad and Friedman 1976, Teufel 1981). Tec-tonic analogues of mffs have been observed associated withshear 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 mi-crocrack density, length, displacement and normal stress ob-served in the experiments are all consistent with the formationof mffs due to contact stresses on the sliding surface. Mffs areintra- or transgranular microcracks that can be recognized bytheir localization adjacent to the fault plane, their inclinationsof 20-50° to the fault plane, and their wedge-shape openingtowards the fault plane (Table 2.1). Mffs can be distinguishedfrom Riedel microfractures, which may also be associatedwith fault planes (e.g. Petit 1987), by the shear offset alongthe latter.

2.3.10 Thermally-induced microcracks

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


(1984) and Vollbrecht et al. (1991). Granite can be treatedas a composite of quartz surrounded by feldspar. Two ex-treme cases can be considered during cooling and uplift. Inisothermal decompression, greater elastic expansion of quartzmay cause microcracking of feldspar, while isobaric coolingmay lead to microcracking of quartz due to its greater thermalcontraction. The critical geothermal gradient for equilibriumbetween thermal and elastic strains in quartz and feldspar is10°C/km. Crystal anisotropy may be important on the grainscale; quartz has a maximum coefficient of thermal expan-sion perpendicular to the c-axis, favouring microcracks at lowangles to the c-axis. Regional stresses can also be important:microcracking will be favoured in those grains with appropri-ate orientations relative to the regional stress.(e.g. Vollbrechtet al. 1994). The general interaction between thermal andelastic stress around inclusions has been modelled by D’Arcoand Wendt (1994), and specifically for garnet by Whitney(1996).

Thermally-induced microcracking in granites can be re-cognized by intragranular microcracks concentrated in quartzsurrounded by feldspar. Thermal microcracks in quartz mayhave a preferred orientation parallel to the c-axis.

2.3.11 Phase transformation-induced micro-cracks

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

transformation involves a volume increaseof 11%. Quartz inclusions in garnet or omphacite are sur-rounded by radial extension microcracks (e.g. Chopin 1984,Smith 1984, Wang et al. 1989), and indeed this texture hasbeen 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 be-cause relict coesite can be found in some inclusions, the ques-tion arises whether such microcracks could be due to elasticmismatches between the silica phase and the host. The keyevidence for phase transition microcracking is the observationthat there is no microcracking around other types of inclusion,including rutile. The extensional nature of the microcracksand their origin from tips of inclusions are consistent withthe mechanism. Fracture surface energy measurements in theconditions of the quartz phase transition imply thatmicrofractures could also be associated with this transforma-tion, and other minerals undergoing similar phase transform-ations (Kirby and Stern 1993). Radial microcracks have alsobeen observed around calcite inclusions which have replacedaragonite, a transformation that involves a 8.5% volume in-crease (Wang and Liou 1991). The distinctive features ofphase transformation microcracking are the association of in-tragranular microcracks with evidence for phase transform-ation. In the case of the transition, radialmicrocracks around inclusions of quartz after coesite are dis-tinctive.


2.4 Microfaults

2.4.1 Characteristics

Microfaults are shear microfractures that contain grain frag-ments formed by cataclasis (Plate 3). Displacement paral-lel to a microfault surface can be identified from displacedgrain boundaries or fragments. However, caution is neededin positively identifying such shear displacements, because apurely extensional microcrack can appear to have a shear dis-placement when viewed in a section which is oblique to theopening direction of the microcrack. A useful indication oftrue shear displacement is the consistent sense of offset of anumber of grain boundaries at a variety of angles to the mi-crocrack (Plate 45).

Two components can comprise microfault matrix: frag-ments 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 mi-crofault. It is sometimes possible to identify the parent grainfrom which the fragments have been derived (the sensitivetint plate is useful for this purpose; Boldt 1995), and frag-ment displacement or rotation may also be detected (Plates 2,4, 45). The proportion of cement may vary from relativelylittle, in which the fragments may form a jigsaw texture, to amajority, in which the fragments will be matrix-supported andisolated from each other. Precipitated cements can be identi-fied from euhedral crystal faces overgrowing primary grains(e.g. Pittman 1981) or from CL studies (e.g. Stel 1981) be-cause the precipitated cement has a different luminescencefrom the grains in the matrix (Section 3.9). A cyclic historyof cementation and shear can sometimes be deduced from thepresence of resheared matrix. The edges of microfaults maybe planar and truncate adjacent grains cleanly, or they may beirregular, perhaps comprising a large number of circumgran-ular microfractures.

2.4.2 Mechanisms

Experiments identify the shear failure mechanism as the link-ing of extension microcracks once a critical microcrack dens-ity is achieved (Krantz and Scholz 1977, Costin 1983). Sev-eral mechanisms of microcrack linkage have been sugges-ted. Peng and Johnson (1972) proposed that axial micro-cracks became linked when individual beams, bounded by

axial microcracks, buckled at a critical fibre strain, allow-ing the first through-going fault plane to form (Fig. 2.14).This linking mechanism can be contrasted with the beha-viour of other siliclastic rocks, in which grain boundary mi-crocracks form first, to be linked by axial microcracks (Had-izadeh 1980, Menéndez et al 1996). The detailed sequenceof microcracking and linkage probably depends on the ini-tial microstructure of the rock (Hadizadeh 1980, Blenkinsopand Rutter 1986). A third type of linking mechanism is thedirect interaction of microcrack stress fields. Such interac-tions were classified into en-echelon and en passant typesby Krantz (1979) (Fig. 2.15). The former occur betweentwo straight, sub-parallel microcracks, which are linked bya third straight microcrack (Fig. 2.15 a, b). En passant inter-actions involve curvature of microcrack paths because the mi-crocrack tip stress fields influence each other (Fig. 2.15 b, c).The definitive theoretical model of microcrack interactions isbased on solutions for the behaviour of isolated microcracksin compression (Horii and Nemat-Nasser 1985). To modelshear failure, the microcracks interactions are considered inan array of parallel flaws with individual angles and overall

axial compression at first increases with for all values ofi.e. microcrack growth is stable. However, at lower values

of microcracks interact unstably at a certain stress. In gen-eral the model predicts that and are different, suggestingthat a failure plane will consist of oblique microcracks linkedby axial microcracks. The model correctly predicts the effectof confining pressure on ultimate strength and is validatedby experiments on resin blocks. Natural microfaults seem toevolve by microcrack linking at a critical microcrack density

angle to and spacing d (Fig. 2.16). The results show that


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

2.5 Deformation bands

2.5.1 Characteristics and classification

A distinctive type of localized deformation occurs in porousgranular materials. Deformation bands are roughly planarfeatures from millimetres to a hundred metres long, fromfractions to a few millimetres wide, and with net slips fromfractions to tens of millimetres. Displacement across deform-ation bands is continuous at a mesoscopic scale, which distin-guishes deformation bands from faults. Porosity in deforma-tion bands may be greater or less than the wall rocks; in themore common case of porosity reduction, the bands containa layer of fine-grained matrix partly comprised of grain frag-ments. Intragranular extension microcracks are found withinthis type of deformation band.

There is a continuum between faults and deformationbands, with many descriptions of faults having some charac-teristics of deformation bands (e.g. Engelder 1974, Houseand Gray 1982, Jamison and Stearns 1982, Underhill andWoodcock 1985, Narahara and Wiltschko 1986, Blenkinsopand Rutter 1986, Cruikshank et al. 1991, Zhao and John-son 1991). Deformation bands were explicitly described byAydin (1978), Aydin and Johnson (1978, 1983), and mostcomprehensively by Antonellini et al. (1994), who distin-

guished between deformation bands with no cataclasis, de-formation bands with cataclasis, and deformation bands withclay smearing.

Deformation bands tend to cluster together to form zonesof bands, often in an anastomosing pattern (e.g. Engelder1974, Blenkinsop and Rutter 1986). The thickness of theselarger-order features depends on the number of individualbands which comprise them, and their total displacement isgiven by the sum of the individual components. Deforma-tion zones may contain discrete surfaces with large discon-tinuous displacements that cut through all other features, andmay have striations on their surfaces. These were describedas slip surfaces by Aydin and Johnson (1983), and are thelatest features to form in a sequence from deformation bandsto zones to slip surfaces.

The distinguishing characteristics of deformation bands areshear displacements which preserve material continuity on amesoscopic scale, and their occurrence in porous, granularrocks.

2.5.2 Mechanisms

Dilatancy caused by grain sliding, and impingement micro-fracturing are important mechanisms in deformation bandformation. Experiments, theory and observations suggestthat the differences between the three categories of deform-ation band are related to confining pressure and initial mi-crostructure (Antonelli et al. 1994). Lack of grain micro-cracking in the first category may indicate low confining pres-sures during deformation compared to the bands in the secondcategory. Initial porosity determines whether a deformationbands dilates or compacts: low initial porosity leads to dila-tion, and vica versa. These relationships are predicted by crit-ical state theory for granular materials (e.g. Schofield andWroth 1968). The third category forms as a result of thehigher clay content in these rocks.

The localization theory of cataclasis accounts for the form-ation of microfaults, deformation bands, zones and slip sur-faces (e.g. Rudnicki and Rice 1975, Aydin and Johnson1983). Localization is defined as a difference in strain ratebetween a band and the matrix, and is a highly appropriateway to describe deformation bands since strain appears con-tinuous across the bands. The theory predicts the formationof deformation bands at or before peak stress, and their sub-sequent spread due to strain hardening. The development ofslip surfaces can be understood from the same theory as a pro-cess in which stress within the deformation band becomes toogreat to be accommodated with the result that a discontinuitydevelops.

2.6 Distributed cataclasis and cata-clastic flow

A number of experiments have produced distinctive cata-clastic microstructures in samples that have maintainedstrength without localization of deformation. The micro-structures are characterized by distributed fracture and dis-placement of fragments, but the deformation is macroscopic-


ally continuous: this is called cataclastic flow, which can bedefined as deformation by cataclastic mechanisms leading tocontinuous flow at a given scale. Cataclastic flow occurs athigher confining pressures than faulting in equivalent rocks,and is favoured by lower differential stresses.

Extensively damaged grain boundaries, perhaps producedby fracture under sliding indenters (microfault-induced mi-crocracks, Section 2.3.9), and wear products accumulated inthe sliding zone and in pores are characteristic microstruc-tures of cataclastic flow in quartzites. (Rutter and Hadizadeh1991). Compaction and filling of pores by crushed materialseems to be typical of cataclastic flow of siliclastic rocks ingeneral (e.g. Menéndez et al. 1996). Axially orientated mi-crocracks are dominant.

Microstructures produced during experimental cataclasticflow of feldspar aggregates consist of intragranular shear mi-crocracks mainly along feldspar cleavages producing blockyfragments between the cleavage microcracks at 90° to eachother (Tullis and Yund 1987, 1992). The spacing of themicrocracks is as low as Deformed grains have apuckered appearance consisting of areas of patchy extinction,caused by slight mismatches between adjacent blocks, whichresemble subgrains or even recrystallized grains formed athigher temperatures by intracrystalline plasticity. However,TEM observations established that no dislocation processescontributed to the block rotation. Microcrush zones <

wide have sharp boundaries and strong grain rota-tion. A crystallographic preferred orientation and a stronggrain shape fabric developed by slip on the closely-spacedgrain-scale faults that are geometrically similar to crystalplastic slip systems, and mechanical twins were formed.Similar features have been observed in cataclastic flow ofanorthosite, and lamellar features interpreted as bundles ofcleavage microcracks were also observed (Hadizadeh andTullis 1992). Twinning is evidence for the operation ofsemibrittle deformation mechanisms.

It is likely that quartz does not undergo the same cataclasticflow regime observed in feldspars, but requires a thermally-activated mechanism to stabilize flow (Hirth and Tullis 1991).The difference between feldspar and quartz behaviour is dueto the excellent cleavages in feldspar. Pyroxenes and am-phiboles may behave in a similar fashion to feldspar (Tullisand Yund 1992). Cataclastic flow has also been reported fromexperiments on basalts (Mogi 1965; Shimada 1986), duniteand gabbro (Byerlee 1968).

A maximum of 30% shortening has been achieved in ex-perimental cataclastic flow, after which faulting generally oc-curs, and therefore it is unclear whether cataclastic flow canbe stable to higher strains. The problem is complicated be-cause faulting is artificially initiated by the experimental con-figuration. If porosity reduction is necessary for cataclasticflow in porous rocks (e.g. Rutter and Hadizadeh 1991), thismust limit the amount of strain that can accumulate by cata-clastic flow. At the opposite extreme of scale, cataclastic flowhas been suggested for continents (Gallagher 1981), and atintermediate scales such as kilometric scale folds (Stearns1968, Hadizadeh and Rutter 1983, Blenkinsop and Rutter1986). In all these cases, the deformation is discontinuousat a smaller scale than the scale of observations.

2.7 Gouge zone microstructures

Faults formed at low grade conditions commonly containgouge with a distinctive set of microstructures, which havebeen closely replicated by experiments (e.g. Logan et al.1981, Rutter et al. 1986). Most of the features of gougezones can be understood in the context of a zone of simpleshear with boundaries parallel to the shear plane (Fig. 2.17).P-foliation is formed by grains (most commonly phyllosilic-ates) aligned at an oblique angle to the gouge zone boundary(Fig. 2.17, Plate 6). It may be inosculating and rather variablydeveloped. It is one of the first microstructures to form in ex-perimental gouges. Shears parallel to the P-foliation with thesame sense of shear as the gouge zone are known as P-shears.Riedel shears (R) form at an oblique angle in the oppositedirection to the P-foliation (Fig. 2.17, Plate 6). They are dis-crete shears with the same sense of shear as the gouge zoneand have geometrical similarities to extensional crenulationcleavages and (Section 7.5, cf. Platt and Vissiers,1980). Conjugate Riedel shears form at a large angleto the gouge zone boundary with the opposite sense of shear(Fig. 2.17). They are generally much less prominent in gougezones than Riedel shears.

T fractures or extension fractures are inclined to the gougezone boundary in the opposite direction to the P-foliation(Fig. 2.17). They may be localized in more competent rocksin the gouge zone. Competent rock may form boudins,which are commonly asymmetric (Section 7.10) and sep-arated along Riedel shears (Fig. 2.17, Plate 6). Ductilestringers (Fig. 2.17) consist of relatively large, rigid, asym-metric clasts elongated in the direction of the P-foliation, withstepped tails that may become detached from the main clast(Logan et al. 1981). Asymmetric folds are also found ingouge zones: they may fold earlier boudins. The vergenceof the folds may be the same as the shear sense of the gougezone (Fig. 2.17), although opposite vergence can occur (Sec-tion 7.11). The orientation of fold axes within gouge zonesmay define a girdle distribution parallel to the shear plane.

Y-shears are discrete, relatively long shears that cut throughthe 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 thecentral part of the gouge zone. In experiments, Y-shears arethe last feature to form.

The orientation of all features except Y-shears shown inFigure 2.17 is variable since their initial orientation is likely tochange with finite strain, because material lines rotate duringsimple shear. Features rotate clockwise in a dextral gougezone such as illustrated in Figure 2.17. Components of pureshear may complicate this general rule. The foliation in afoliated cataclasite may originate as a fabric along any of theabove shears, or as a P-foliation in a gouge zone (Plate 8).

2.8 Microfracture surface features

Microfracture surfaces may show a variety of microstructuresthat convey useful information about cataclastic deformation.Intact specimens of microfracture surfaces can be easily ob-served under a binocular microscope, and at greater magni-fication in the SEM. It is also useful to examine the cross-


sections by cutting thin sections perpendicular to the micro-fracture surface, preferably parallel and perpendicular to theslip direction.

There is a rich field of microscopic studies on extensionmicrofracture surfaces in the material science literature (e.g.Bansal 1977, Quakenbush and Frechette 1978, Michalskeand Frechette 1980) which suggests that fracture surface fea-tures could be used to reveal rates of fracture propagation andeven fracture stresses (e.g. Scholz 1972, Martin and Durham1975, Swanson 1981, Norton and Atkinson 1981, Meredith1983, Cox and Atkinson 1983). Surface microfracture fea-tures such as mirror, mist, velocity hackle and forking, andWallner lines, which are known to be diagnostic of dynamicmicrocrack propagation in glass, have not been found on frac-ture surfaces in rocks. This has been taken to indicate thatthe studied rock fractures were never dynamic (Kulander andDean 1995). This conclusion is supported by observationssuch as irregular rib marks that imply stable propagation atlow velocities, again by analogy with glass.

A shear microfracture with a smooth or shiny surface isa slickenside, which is commonly marked by parallel linesknown as slickenlines or slickenside striations. Many slick-ensides are the surfaces of microfaults formed during tec-tonic deformation. The lines are parallel or tapering ridges orgrooves parallel to the slip direction. There may be more thanone set of slickenlines on a microfault surface in different dir-ections. It is sometimes possible to distinguish overprintingrelationships between different generations of slickenlines.

Probably the most common mechanism of slickenlineformation is asperity ploughing (Section 2.2.2), which canbe identified by the preservation of the asperity on one sideof the microfracture and a matching groove on the oppositesurface. The type of slickenline produced is called a weargroove (Fig. 2.18), tool track or mark, or prod mark (Tija1967, Hancock 1985). Ridges of gouge may be formed bywearing down asperities or by accumulating gouge around ahard particle: this mechanism is debris streaking. Erosionalsheltering occurs where a solid ridge of microfracture sur-face is preserved in the down-slip direction of a particle. A

category of slickenlines consists of equally developed ridgesand grooves with perfectly matching (or nested) profiles onboth sides of a slickensided surface (Means 1987). The ridgeand groove structure can be formed by localized shear asso-ciated with a grain shape fabric, and the length of individualridges and grooves may be greater than the displacement onthe shear zone (Wil and Wilson 1989). Ridge and groovestructures can also be formed by inosculation of shear sur-faces (C-surfaces) in an S-C mylonite (Section 7.5, Lin andWilliams 1992). Another type of slickenside which may alsobe formed by continuous deformation consists of very finegrained quartz (0.01 to ) with a strong crystallographicpreferred orientation (Power and Tullis 1989). The preferredorientation may have resulted from orientated growth by dif-fusive mass transfer through a solution during the interseis-mic, low strain rate part of the seismic cycle.

Microfracture surfaces may be offset along discontinuit-ies approximately perpendicular to slickenlines, known asrisers or steps (Fig. 2.19). Risers are generally less denseand coarser features than slickenlines, and are commonlymore irregular. They may be approximately linear or cres-cent shaped. Risers are described as incongruous if the off-set opposes the direction of movement of the opposite block,and congruous if the opposite is true. Crystal fibres on faultsurfaces are known as slickenfibres, and risers created by thecrystal terminations are accretion steps (Norris and Barron1968). They are identified by the presence of fibres that arecommonly monocrystalline. Risers may form either by ini-tial irregularities in the fault surface, or by the intersection ofsecondary fractures with the fault surface. Possible geomet-ries of secondary fractures can be separated into P fractures,R, T and fractures using the same terminology as intro-duced above for gouge zone features (Petit 1987). Slicken-fibres form by crystal growth in the dilatant gap adjacent tocongruous risers.


2.9 Crystallographic fabrics

Experimental studies have shown that crystallographic fab-rics can be produced by cataclasis (e.g. Borg and Maxwell1956, Borg et al. 1960). Crystallographic fabrics in feld-spars have been produced by slip and rotation on multiplegrain-scale faults (Tullis and Yund 1987, 1992). The faultsare geometrically analogous to slip systems in intracrystal-line plastic deformation (e.g. Allison and La Tour 1977, Sec-tion 4.9). Cataclastically-formed crystallographic fabrics canbe distinguished from intracrystalline plastic fabrics by thepresence of multiple, grain-scale faults such as those illus-trated by Tullis and Yund (1992).

2.10 Pre-lithification deformation mi-crostructures and mechanisms

The microstructures of pre-lithification deformation, andeven the deformation itself, may be difficult to recognize. Agrain shape fabric can be produced by frictional grain slidingand rotation, known as independent particulate flow (IPF) be-cause the deformation is independent of fracture (Borradaile1981). This deformation mechanism is sometime regardedas separate from cataclasis, because there is no fracturing;here, however, cataclasis is used in a broader sense to en-compass IPF. High pore fluid pressures may be essential inpre-lithification IPF (Knipe 1986, 1989, Maltman 1994). Thediagnostic feature of a grain shape fabric produced by IPF isthe absence of microcracks and lack of internal deformationof grains, but intracrystalline plasticity can also be involvedduring 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 deforma-tion.

2.11 Pseudotachylites and frictionalmelting


Pseudotachylite is a glassy or very fine-grained rock occur-ring in veins and associated with deformation zones (Plates 9-11). It generally consists of a dark matrix which encloses an-gular to rounded fragments of the wall rock (Plates 9, 11).The fragments are very poorly sorted and may have a highclast/matrix ratio. Clasts in the range to 10 mm insome pseudotachylites have a fractal PSD, with a fractal di-mension of 2.5 (Shimamoto and Nagahama 1992). Quartzand feldspar fragments may be more common in the pseudot-achylite than in the wall rock, while hydrous minerals suchas biotite may be less common. Flow banding defined bycolour variations may be seen in the matrix (Plates 10, 11),which may contain spherulites, dendritic crystals, crystal-lites, and microphenocrysts, and may vary from opaque tocolourless and isotropic in thin section. Other microscopicfeatures include optically isotropic rims around clasts, em-bayments of clasts, sub-microscopic crystals, and amygdales.

Fine grained margins may occur at the contact with the wallrock.

Pseudotachylite often occurs in planar veins a few mm to1 m thick from which veinlets may branch and penetrate thewall rock. Planar-convex lenses of pseudotachylite may oc-cur on either side of the vein. The pseudotachylite may occurin a volume between two parallel planar zones, and surroundlarge fragments of wall rock (e.g. Grocott 1981, Magloughlin1989). The wall rock fragments can form a jigsaw-texturedbreccia with pseudotachylite matrix. Joints may form perpen-dicular to the vein walls. The pseudotachylite veins cross-cutany earlier fabrics in the wall rock, and may contain frag-ments of previously-formed pseudotachylite, indicating cyc-lical generation (e.g. Sibson 1980, Passchier et al. 1990). Allthese features may occur on outcrop to thin section scale.

2.11.2 Origin

The origin of pseudotachylite has usually been discussed interms 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 in-volved frictional melting during faulting. The intrusive tex-tures of pseudotachylite veins together with their flow band-ing demonstrate that they formed in a fluid state. The spher-ulites and dendritic crystals and crystallites formed by cool-ing from a melt or from devitrification of a glass formedby quenching, by analogy with structures formed in rapidlycooled igneous rocks. The fine-grained margins of the veinscan be interpreted as chilled margins, and the joints perpen-dicular to the veins walls can be interpreted as cooling frac-tures (e.g. Camacho et al. 1995).

Some pseudotachylites have the same bulk chemistry astheir hosts, and the pseudotachylite composition may changewith that of the local wall rocks through which they pass (e.g.Maddock 1986, Maddock et al. 1987), demonstrating that themelt was locally derived and not transported by an intrusivedyke. In other pseudotachylites, there is a consistent differ-ence in the bulk chemistry of the pseudotachylite and wallrock which can be interpreted by fusion of the lowest meltingpoint fraction of the rock. Preferential melting of the hydrousmafic components of the host rock can generate a pseudot-achylite matrix which is more mafic than the host rock, andexplains why hydrous and mafic minerals are less common inthe matrix than the host rock (e.g. Magloughlin 1989, Mad-dock 1992, Camacho et al. 1995).

The localization of pseudotachylite on and near planar sur-faces suggests that it is generated by sliding (hence these sur-faces are called generation planes). The rounding of the frag-ments may be due to thermal spalling. A link between slick-ensides, slickenlines and pseudotachylite has been suggestedby Spray (1989) from observing mechanical excavation ofsandstone, which created thin layers of melt with shiny andstriated surfaces.

Pseudotachylite has also been successfully created in fric-tional sliding experiments (e.g. Spray 1987, 1995). Fila-ments of frozen melt drawn out across fractures “like warmmozarella” are a delicate feature of experimentally-producedpseudotachylite which has not so far been observed in tec-tonic pseudotachylite, but could be a useful diagnostic fea-


ture. The experiments resolve the controversy about the ori-gin of pseudotachylites: both crushing and frictional melt-ing are likely to be involved, depending on the strain rate(Spray 1995). Crushing occurs in experiments at strain ratesof and melting at strain rates from to

The passage of seismic or shock waves, with a rapid in-crease in strain rate, may involve three sequential processes:initial fracture, comminution, and melting when heat cannot be dissipated fast enough. Together with results fromthermal modelling, experiments show that pseudotachylitesshould 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 isgenerated to melt the rock (e.g. Spray 1987, 1988). On theother hand, Magloughlin (1989) has argued that pseudota-chylites were generated from a hydrous cataclasite with a lowmelting temperature. The permeability of the wall rocks isan important factor in determining whether excess fluid pres-sures can build up during faulting and suppress melting (Maseand Smith 1985). Experiments confirm the widely held viewthat pseudotachylites can form during seismic slip at strainrates greater than

Some pseudotachylites are associated with features suchas craters, shatter cones and shock metamorphism whichshow that they have been produced during meteorite impacts,for example the Vredefort Dome in South Africa, from whereShand (1916) first coined the term pseudotachylite, and

the Sudbury structure (Spray and Thompson 1994). Otherpseudotachylites have been described from the bases of largelandslides (e.g. Masch et al. 1985).

2.11.3 Misidentification

Veins of tourmaline, chlorite, ultracataclasite and fine-grainedigneous dykes can easily be mistaken for pseudotachyl-ite. Passchier and Trouw (1996) suggest some microstruc-tures that allow the distinction of pseudotachylite from othertypes of veins. The diagnostic microstructures that identifypseudotachylite as a melt in distinction from hydrothermalveins are the quench and devitrification textures. The distinct-ive chemistry, reflecting the wall rocks, is also diagnostic anddistinguishes pseudotachylite from ultracataclasites. Distinc-tion from igneous dykes relies on evidence for faulting associ-ated with the generation of pseudotachylite, and the differentchemistry of the dykes. Pseudotachylites should be distin-guishable from ultramylonites by their lack of evidence forextensive intracrystalline plasticity.

There are nevertheless many cases in which detailed evid-ence must be sought to prove the origin of a pseudotachyl-ite by frictional melting. It is a relatively uncommon rocktype in the field. The requirement of formation in hot anhyd-rous conditions effectively limits the host rocks to igneous ormetamorphic 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 distor-tion form by diffusive mass transfer (DMT). It is often diffi-cult to identify positively the nature of the diffusive flux, andthe evidence for the role of a solution may be indirect. Nev-ertheless, it is a safe generalization that up to amphibolite fa-cies, by far the majority of DMT occurs via solution becausesolid-state diffusion occurs very slowly at these temperatures(e.g. Rutter 1983).

3.2 Fundamental deformation mech-anisms of diffusive mass transferby solution

Diffusion of material occurs in response to gradients in chem-ical potential, as summarized by Fick’s law which states thatthe flux of the material (J) is proportional to the chemicalpotential gradient

the concentration. The chemical potential of a componentis given by:

where U is the molar internal energy, T the temperature, Sthe molar entropy, is the normal stress, is the porefluid pressure, and is the molar volume (e.g. Green 1980).Therefore variations in normal stress and pore fluid pressurecan establish chemical potential gradients necessary for DMTto occur; internal strain energy may also play a role (Wintschand Dunning 1985, Bell and Cluff 1989). The diffusion willbe from sites of high to low normal stress; a gradient innormal stress is essential to cause material flux (Raj 1982).Hence deformation accommodated by DMT through an in-tergranular solute is often called pressure solution, despitethe fact that solubility itself is not significant to the thermo-dynamics of DMT. However, recent experiments that demon-strate a crystallographic dependence of solubility may posesome problems for current pressure solution models basedon this thermodynamic approach (Becker 1995, Den Brock1996).

The exact geometry of the diffusion path is still unclear.Rutter (1976, 1983) has suggested that a water film on theorder of nm thick must be present along grain boundaries toallow DMT to occur. This concept is supported by observa-tions of arrays of fluid inclusions along grain boundaries innaturally deformed halite, interpreted as evidence of a syn-tectonic continuous film of fluid along grain boundaries thatsubsequently healed into discrete fluid inclusions (Urai et al.1986), and by observations of grain boundary pores whichprobably formed in equilibrium with a fluid in quartz (Hip-pert 1994). Alternatively, islands of solid material may sup-port normal stresses and allow diffusion to occur in a solutionwithin intervening channels (Raj and Chyung 1981, Raj 1982,Spiers and Schutjens 1990). It may be necessary for differentphases to be adjacent to each other for a continuous fluid filmto be present (Hickman and Evans 1991).

The undercutting mechanism proposed for pressure solu-tion is completely different and does not require diffusionin a solute film between surfaces under high normal stressgradients (e.g. Bathurst 1958, Tada and Siever 1986). In thismechanism, grain contact deformation occurs by intracrystal-line plasticity or cataclasis, while dissolution of free surfacesaround the contacts maintains small contact areas and there-fore the high stresses necessary for the solid-state grain de-formation. The relative importance of solute film diffusionversus undercutting depends on temperature, grain size andfree grain surface area. Calculations by Tada et al. (1987)suggest that undercutting will provide faster strain rates un-der most circumstances, but that solute film diffusion will be-come important for very small free grain surface areas, as inthe final stages of compaction, and in dense aggregates.

The distribution of fluid in a rock, and therefore the effect-iveness of pressure solution, is dependent on surface energy(measured by the dihedral angle, ) and deformation. Fluidswith greater than 60° will occur as isolated pores at equi-librium, as interconnected channels at triple grain junctionsfor and along all grain surfaces for(e.g. Watson and Brennan 1987). The first two cases clearlyconstrain the potential for pressure solution because the dif-fusion pathways are limited, and probably apply to fluids inmost static geological situations. However, experiments andtheory show that fluids can wet grain boundaries completelyunder deformation, leading to enhancement of diffusion creepby orders of magnitude (e.g. Urai 1983, Copper et al. 1989,Heidug 1991, Tullis et al. 1996). Moreover, dihedral anglesmay be a function of pressure and temperature, as established

24

B is the particle velocity per unit potential gradient and c is

CHAPTER 3. DIFFUSIVE MASS TRANSFER BY SOLUTION 25

by experiments on brines in halite, which show that the di-hedral angle reduces below the 60° threshold at elevated pres-sures and temperatures (Lewis and Holness 1996). Unfortu-nately experiments also suggest that evidence for fluids alonggrain boundaries may be rapidly removed after deformationceases.

DMT is often inferred from microstructures that show asystematic relation to surfaces which may have been underhigher normal stress (e.g. surfaces parallel to a shape fab-ric). However, it is the principal strains which are apparentfrom most microstructures, and thus there is always an ambi-guity about interpreting DMT from the such microstructures(cf. Groshong 1988). The above inference relies on an as-sumption 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 emphas-ized because they can be inferred from the microstructures.Any links made between DMT microstructures and principalstress orientations or gradients in normal stress must involvesome assumptions about the nature of the deformation, suchas coaxiality.

3.3 Grain surface solution textures

Solution creates distinctive grain surface textures which areclearly observed in the scanning electron microscope. Thetypical texture consists of pits on the grain surface, giving acorroded appearance. Irregular pores with channel and cavemorphologies 1 to long and deep have beendescribed in a naturally deformed micaceous quartzite (Hip-pert 1994); similar channel morphologies have been producedduring experimental deformation of quartz by DMT (DenBrock and Spiers 1991). Atomic-scale details of the solu-tion process can be seen using atomic force microscopy (e g.Gratz et al. 1991).This technique shows that calcite dissolu-tion and precipitation occurs in layers, and that solution gen-erates crystallographically-controlled etch pinholes less than5 nm deep, and rhombic etch cores more than 90 nm deep(Hillner et al. 1992). Quartz solution occurs along nm-sizedpits and ledges parallel to the and directions(Gratz et al. 1991).

3.4 Indenting, truncating and inter-penetrating grain contacts

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

confidence in sediments with rounded grain shapes that al-low the pre-solution grain shape to be accurately delineated.Truncated contacts may appear superficially like flat contactsformed by mutual concordance between overgrowths (Sec-tion 3.9), but can be distinguished by the evidence for mater-ial removal.

The formation of indenting or truncating grain contacts canbe understood from theoretical and photoelastic treatments ofthe stress distribution in contacting grains. The values of theprincipal stresses, and the mean and differential stresses, in-crease with the radius of curvature, suggesting that DMT willpreferentially occur in grains with larger radii of curvature,creating indenting grain contacts (e.g. McEwen 1981). Trun-cating grain contacts are expected from approximately equalrates of DMT between two grains of equal curvature.

Interpenetrating grain contacts are mutually interlockingprotrusions of grains into each other, which have a similarmorphology to microstylolites (Section 3.6), and indicate ma-terial removal along the contacts (Figs. 3.1c, 3.4). Su-tured grain contacts may look superficially like interpenet-rating grain contacts, but these are formed by intracrystal-line plasticity (Section 4.8). Sutured contacts can usually bedistinguished from interpenetrating contacts by the presenceof subgrains, which may form promontories along a suturedcontact that are slightly misoriented with respect to the latticeof the rest of the grain (Fig. 3.1d). By contrast, interpenet-rating contacts formed by DMT usually occur between grainswith uniform lattice orientations.

26 CHAPTER3. DIFFUSIVE MASS TRANSFER BY SOLUTION


3.5 Strain caps

Strain Caps are strongly foliated domains enriched in micasor less soluble minerals around opposite surfaces of relativelyrigid objects (Passchier and Trouw 1996). The foliation tracein the strain cap is concordant with the margins of the object,and grades into continuity with the bulk rock foliation withincreasing distance from the object, so that it appears to bedeflected around the object (Plate 13). The concentration ofless- soluble minerals such as micas in strain caps suggeststhat they form by removal of the more soluble matrix aroundthe rigid object, which can be predicted from the thermody-namic considerations above, since the object will concentratenormal stress on some surfaces due to its greater rigidity.

3.6 Microstylolites


Microstylolites are microscale discontinuities with offset orwave-like shapes that may truncate grains or other mark-ers 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 teethor columns, which may have flat crowns at the top andsteep walls along their sides (Fig. 3.5). The offset or wave-like geometry is observed in all sections, showing that thethree dimensional shape is a forest of columns and match-ing pits approximately perpendicular to the plane of the mi-

crostylolite. Microslickolites are distinguished from micro-stylolites by having teeth that are oblique to the plane. Theoblique teeth create a lineation on the microslickolite surface.Microstylolites may follow grain boundaries, especially of in-soluble grains. Microstylolites can be conspicuous becauseof coatings or fillings of iron oxides, hydroxides, phyllosilic-ates (which may have a strong grain shape fabric) or organicmatter, which distinguish them from the wall rock. The thick-ness of the filling is commonly highly variable. The morpho-logical classification of cross-sections through stylolites byGuzetta (1984) into sharp peak, rectangular, wave, smoothand composite types is useful for microstylolites (Fig. 3.6a,Plate 15). Parameters for describing microstylolites quant-itavely are shown in Fig. 3.6b (from Andrews and Railsback1997).

3.6.2 Formation and propagation

The geometrical similarities between stylolites and micro-stylolites suggests that they have the same genesis. The dis-cussion in this section refers to sylolites, because most previ-ous work has focussed on the mesoscale features.

Truncation of markers is direct evidence that material hasbeen removed along a stylolitic surface. The minimum thick-ness of material removed along a stylolite is approximatelytwice the amplitude, or the height of the teeth (Tada andSiever 1989). The walls of the teeth are parallel to the direc-tion of movement between the opposite sides of the stylolite.Railsback and Andrews (1995) have shown that the orienta-tion of teeth is more constant than the orientation of the plane

28 CHAPTER. 3 DIFFUSIVE MASS TRANSFER BY SOLUTION

of the stylolite, and is useful for kinematic analysis becauseit shows the direction of maximum finite shortening. The ob-liquity of the teeth of slickolites shows that there has beena component of shortening parallel to the slickolite surface(Fig. 3.6c).

The coatings or fillings on stylolite surfaces are usually in-terpreted as insoluble residues that did not diffuse and wereconcentrated at the sites of solution. The relative insolubil-ity of these coatings is good evidence that diffusion was viasolution. This is further supported by a correlation betweenthe thickness of the coating and the concentration of this ma-terial in the adjacent layers of rock (e.g. Borradaile et al.1982), and by the influence of heterogeneities on the stylolitetrace, which show that solution was concentrated in areasof higher inferred normal stress, and was impeded by insol-uble material. The concentration of insoluble material on astylolite surface compared to its concentration in the host rockcan be used to give an estimate of the amount of shortening(Railsback and Andrews 1995). However, in some cases thestylolite filling material is not found in the host rock, and istherefore the product of metamorphic reactions (Beach 1979).Reaction products can also mark grain boundaries to indicatewhere fluid transport has occurred (McCaig 1987).

A stylolite can be regarded as an anticrack, or an ellips-oidal volume removed from the rock (Fletcher and Pollard1981). If the resulting hole is closed up by elastic deform-ation, stresses are induced around the anticrack tip, whichcause the stylolite to propagate in its own plane. Thereforea stress concentration under the appropriate conditions is suf-

ficient to cause a stylolite to develop and propagate under itsown stresses. Anticrack tip stresses may be dissipated in azone at the stylolite tip which can be referred to as a processzone, analogous to the non-linear zone ahead of a microcracktip (Section 2.2.1.1). Carrio-Schaffhauser et al. (1992) foundevidence for a process zone in the form of enhanced poros-ity at stylolite tips in limestones, and suggested that materialremoved from the process zone was deposited in a zone ofreduced 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 localizedalong authigenic clay layers and around pre-existing phyllo-silicates. The proposal that contacts between different phasesare necessary for continuous fluid films (Hickman and Evans1991) suggests a good reason for enhanced DMT aroundphyllosilicates, and may also apply to the formation of straincaps.

3.7 Diffusive mass transfer and cleav-age

3.7.1 Classification

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


cleavage), which hinders discussion about deformation mech-anisms, but Powell (1979) and Borradaile et al. (1982) give anexcellent alternative, non-genetic classification, which is fol-lowed here. The main subdivision is made between spacedcleavage, consisting of cleavage surfaces that are separatedby tabular bodies of rock called microlithons, and continu-ous cleavage, in which cleavage is penetrative throughout thewhole body of the rock. Since most cleavages are concen-trated into domains at some scale, this distinction is evidentlyscale-dependent. Powell (1979) suggests that any cleavagewith domains spaced less than 0.01 mm (i.e. at the limit ofoptical microscopic resolution) should be called continuousat a microscopic scale.

3.7.2 Spaced cleavages

Disjunctive cleavage

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

the rock, forming a zonal cleavage. Another important as-pect of description is the transition from cleavage domainsto microlithons, which may be sharp or gradational. Manycleavages referred to as solution cleavages belong in this cat-egory. Disjunctive cleavage domains are often composition-ally distinct from microlithons, having a higher proportion ofopaque minerals and phyllosilicates. Zonal cleavages withmarked compositional contrasts between cleavage domainsand microlithons constitute a new compositional layering,which has been called solution striping (e.g. Williams 1972).Truncation of markers in cleavage domains suggests that ma-terial loss by DMT has occurred (Fig. 3.8), although sheardisplacements on fractures may create a similar effect. Ma-terial removal can be distinguished from the latter effect insome cases by inconsistent offsets of a folded layer on sev-eral cleavage surfaces, and large variations in thickness of amarker in adjacent microlithons (Fig. 3.8). The amount ofmaterial removal can be estimated from truncation of markersand other effects such as imbrication of chert pebbles and off-set of bedding. The volume loss may vary from a few percentfor a weak cleavage to over 35% for very strong cleavage incherts (e.g. Alvarez et al. 1976), and 30-40% for disjunctivecleavage in coarse siltstones and sandstones (Murphy 1990).Material transfer by diffusion can also be inferred from thecompositional contrasts between the cleavage domains andthe microlithons. A solution can be inferred as the diffusingphase because of the relatively low grades of cleavage form-ation.

Crenulation cleavage

This is a spaced cleavage in which microlithons contain anearlier fabric that is systematically related to the fabric ofthe cleavage domains. The cleavage domains are usuallyaxial planar to folds (crenulations) in the earlier fabric (Plates16, 17). The crenulations may be symmetric or asymmet-ric, and rounded or angular, and the transition from cleav-age to microlithon may be sharp or gradational. The crenu-lation wavelength may be determined by the thickness of fol-ded layers as predicted by buckling theory (cf. Price andCosgrove 1990), or the grain size may dictate a minimumwavelength. The cleavage domains are commonly localizedalong the limbs of the folds, especially those associated withasymmetrical folds, in which the cleavage domains are usu-ally along only one of the two unequal limbs. Most crenu-lation cleavages are zonal because the cleavage domains arecompositionally quite distinct from the microlithons. Phyl-losilicates and opaques are concentrated in cleavage domainsand quartz, calcite and feldspar are concentrated in hinges.

While the essential role of buckling in crenulation cleav-age formation is evident, the importance of DMT is clearfrom the compositional zoning. Cleavage domains in typ-ical crenulation cleavages are depleted in Si, Ca, Na,P, Mg, Fe, and F, and enriched in Al, K, Rb, Y, Ce, Sc, andBa, and there is a net volume loss from the cleavage domainsthat is approximately balanced by volume gain in the micro-lithons (Manktelow 1994). The zoning can be explained in asimplistic way by the normal-stress dependence of diffusion:crenulation hinges, where the primary foliation lies at a highangle to the maximum principal stress, may be sites of low

30 CHAPTER 3. DIFFUSIVE MASS TRANSFER BY SOLUTION

normal stress. It is also possible that anisotropy in diffusionpathways may play an important role in the differentiationduring crenulation (Marlow and Etheridge 1977).

Other lines of evidence that can be used to show the import-ance of DMT are the narrow widths of quartz and feldspargrains in cleavage domains relative to microlithons, trunca-tion of grains, the association of crenulation cleavage withovergrowths and mica beards, and the lack of other deforma-tion mechanisms (Gray 1977, Borradaile et al. 1992).

3.7.3 Continuous cleavage

The most characteristic continuous cleavage is slaty cleav-age, defined as a preferred orientation of phyllosilicates toofine to be seen with the unaided eye (Fig. 3.9). Slaty cleavagecommonly has cleavage domains with spacings of 0.1 mmor less, which are easily visible under the optical microscopeor the SEM. Mechanical rotation of grains (cataclasis), kink-ing (intracrystalline plasticity) as well as DMT are involvedin the formation of slaty cleavage (e.g. Wood 1974). Elec-tron microscope studies show that phyllosilicates and quartzare mobilized by DMT in the development of cleavage do-mains (Knipe and White 1977,1979, White and Knipe 1978,Knipe 1979). Structural studies using a variety of strain mark-ers including reduction spots, trace fossils and grain bound-aries suggest that slaty cleavage formation is accompaniedby volume loss (e.g. Ramsay and Wood 1973: 20%, Wrightand Platt 1982: 50%, Wright and Henderson 1992: 40-60%),

which must have occurred by DMT via solution at these meta-morphic grades. However, bulk volume loss is not permit-ted by several geochemical studies on the same rocks, as dis-cussed 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, carbon-ate, quartz and feldspar (Borradaile et al. 1982). The depletedelements 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 un-likely at low grades because average shale and slate compos-itions are very similar (Erslev and Ward 1994). The reasonsfor the contradiction between this conclusion and the largevolume losses suggested by the structural studies is presentlyunclear. Whether or not large scale volume changes accom-pany cleavage formation, the pronounced local compositionaldifferentiation is clear evidence that DMT processes are es-sential to slaty cleavage development.

3.8 Grain surface deposition textures

Crystal growth from solution occurs by two different mech-anisms with different microstructures (Bennema and van derEerden 1987). Atomically flat crystal faces first nucleatesteps to which atoms can attach, and growth then occurs byaccretion along the steps in crystallographically controlledplanes. Dislocations (Section 4.2) may provide important


nucleation sites for both growth and solution on crystal sur-faces (e.g. Casey 1995). Above a critical temperature knownas the roughening transition, growth rates are isotropic, andsub-spherical crystals may be produced. Deposition by thefirst mechanism can be recognized from the euhedral natureof grain surfaces seen under the microscope; the second maybe responsible for botryoidal textures.

3.9 Overgrowths, porosity reduction,pressure shadows and fringes, andmica beards


Overgrowths are mineral deposits surrounding a grain in arim. The rim and grain are usually the same mineral. The rimis commonly in crystallographic continuity with the grain,and may therefore be difficult to detect. The overgrowthmay, however, be separated from the original grain by a zoneof inclusions (Fig. 3.10), and CL can distinguish subtle fea-tures of overgrowths (Plate 18). A distinctive microstructureformed by overgrowths in well cemented rocks are relativelystraight grain boundaries and 120° triple junctions betweenthree grains (Fig. 3.11a). This texture arises from mutual im-pingement of overgrowths growing at equal rates from adja-cent grains. It can easily be confused with granoblastic poly-gonal textures (Section 5.4) unless the overgrowth can be dis-tinguished 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 strainmarkers and the precipitation of cements in pores in the formof overgrowths (e.g. Carrio-Schaffhauser and Gaviglio 1990).Crystal growth in pore spaces typically results in a convolutedpore-grain interface which is a fractal curve. Models suggestthat the fractal dimension of the curve increases with the ra-tio between precipitation and dissolution rates, and also in-creases from about 2.5 to about 2.75 in natural sandstones asporosity reduces during diagenesis (Aharonov et al. 1997).

Pressure shadows and fringes are domains of secondarymineral growth adjacent to grains (Plate 19). Pressure shad-ows lack distinctive internal crystal forms, while pressurefringes have fibrous mineral fillings. Mica beards are a dis-tinctive type of pressure fringe, consisting of fibrous mica de-fining a good shape fabric. Passchier and Trouw (1996) preferto use strain instead of “pressure” on the basis that it is non-genetic; however, they point out that the term strain shadow isalso potentially misleading as it incorrectly implies that strainis low in the shadow zone. Analysis of steady-state pure sheararound a rigid sphere shows that there are two volumes of lowpressure at the ends of the object perpendicular to the max-imum applied stress, but that differential stress and strain arehigh in these areas (Masuda and Mizuno 1995). With this in-sight, and since pressure shadow is so firmly entrenched inthe literature, it is probably the better choice of terminology,despite its genetic connotations.

The mineralogy in a pressure shadow or fringe may be thesame as the grain, the same as the matrix, or different from


both. Crystallographic continuity may be maintained withthe grain, host, or neither. Fibres may be perpendicular tothe grain boundary or oblique. They may be found on morethan two sides of polyhedral grains. Fibres may be straight,curved, deformed or undeformed. All of these basic aspectsof pressure shadow and fringe description are important ininterpreting the mechanism of formation.

3.9.2 Mechanisms

Pressure shadows and fringes are considered to form by DMTbecause they are a new mineral growth under low grade con-ditions. Diffusion towards sites of low mean stress is pre-dicted by the thermodynamic approach of Section 3.2, whichexplains why pressure shadows and fringes form at the endsof grains perpendicular to the inferred maximum principalstress.

However, the geometry of the pressure shadows and fringesis controlled by a number of other factors including strain,grain shape, whether growth occurs at the grain interface (an-titaxial) or the matrix interface (syntaxial), whether the ori-entation of fibres is controlled by the grain surface (face con-trolled) or the incremental extension direction (displacementcontrolled), and whether the shadows or fringes are deformedor not. Some of these factors are illustrated for a coaxialdeformation in Fig. 3.12. Given due consideration of thesefactors, fibres can be used to deduce the extensional strain,the deformation path, and the type of flow, as discussed be-

low for fibrous microveins (e.g. Durney and Ramsay 1973,Elliot 1973, Ramsay and Huber 1983, Passchier and Trouw1996). The use of fibres as shear sense indicators is discussedin 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 dir-ection (e.g. indenting, truncating or interpenetrating grainboundaries), and by material addition (e.g. overgrowths)in the extension direction (Fig. 3.11b). Grain shape fabricshave been produced experimentally by DMT (e.g. Schutjens1991). Lack of microfractures distinguish DMT grain shapefabrics from cataclastic grain shape fabrics, and the lackof undulatory extinction, sub-grains or recrystallized grains(Chapter 4) distinguishes them from intracrystalline plasticgrain shape fabrics.

3.11 Fluid inclusion planes

Fluid inclusions consist of nm to mm sized cavities filled byfluids, which may also contain solid material. They occuras isolated inclusions, in clusters, and in planes. They arecommon in thin section, where they may appear as dark in-clusions which reveal their fluid contents on examination athigher 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 in-clusions that form during growth of the original minerals ofa rock, and secondary inclusions that form subsequently (e.g.Roedder 1984). Primary inclusions can be recognized be-cause they are disposed on euhedral crystal forms, whereassecondary inclusions cut across crystal growth features.

Fluid inclusion planes (FIPs) are planes of fluid inclusionsthat often have a strong preferred orientation on a microscopicscale, and may also have regionally consistent orientations(Fig. 3.13, Fig. 2.5).

The fluids in any set of FIPs are generally composition-ally homogeneous, and may be distinct from fluids in othersets of FIPs that occur in the same rock. FIPs form by trap-ping of a fluid during precipitation of microcrack fillings.The microcracks are usually extension microcracks formedby the mechanisms described in Section 2.3. Healing of mi-crocracks in quartz can occur as rapidly as micrometres/day(Smith and Evans 1984). The healing rate depends on tem-perature, concentration of the fluid, and microcrack di-mensions (e.g. Brantley 1992). Healing leaves a plane ofcylindrical or spherical fluid inclusions along the former mi-crocrack, firstly by forming cylindrical tubes of fluid paral-lel to the microcrack tip (necking down), and pinching off orovulation to form approximately spherical or negative crys-tal shapes (Fig. 3.14). Isolation of spheres occurs becausegrain boundary migration rates depend on the thickness ofthe fluid phase (the microcrack can be considered as a typeof grain boundary): slower migration rates occur in thickerfluid films. A local thickening of the microcrack will slow


the grain boundary migration rate at that point, isolating afluid inclusion behind the rest of the more rapidly migratingboundary (e.g. Urai 1983). FIPs are primary evidence forthe importance of fluids and DMT during deformation. Theyform perpendicular to and can be very useful in kinematicanalysis, because they can constrain both the orientation of

and the fluid pressure (e.g. Lespinasse and Cathelineau1995). They can also give evidence for fluctuating fluid pres-sures associated with stress cycling and therefore probablywith earthquake faulting (e.g. Robert et al. 1995).

3.12 Microveins

Microveins are microscopic tabular zones of secondary min-eral growth. The rich variety of microvein textures observableunder the microscope can allow very detailed interpretationsof their formation. CL is valuable for analysis of microveins,because variations in fluid chemistry and temperature duringmicrovein filling can be reflected in the luminescence, delin-eating delicate growth features (e.g. Dietrich and Grant 1986,Urai et al. 1991).

One of the most important features of a microvein is theorientation of the opening vector, the vector which connectspoints that were originally joined before microvein opening.The opening vector may be perpendicular to the wall (anextension microvein) or have components of displacementboth perpendicular and parallel to the walls (a shear micro-vein). The opening vector can be determined from displacedmarkers, or from some features of microvein fillings, as de-scribed below. The opening vector can be considered for thenet opening history of the microvein (the cumulative open-ing vector) or for individual increments of opening (the in-cremental opening vector).

Common microvein fillings are carbonates, quartz, chlor-ite and epidote. The texture of the filling may be massive,equant (blocky; Plate 20), fibrous, laminated, euhedral (idio-morphic), or botryoidal. Filling textures are a function of nuc-leation and growth kinetics, the rate of microvein opening, thegeometry of the opening vector and the geometry of the walls.Euhedral crystal terminations and botryoidal textures are dia-gnostic of growth into open space, and are useful in ore par-agenetic studies for discriminating epithermal environments.Euhedral crystal faces can be recognized by growth zoning inthe 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 microveinsalso contain inclusions of the same mineralogy as the wallrock. Lines of inclusions parallel to the microvein marginsare known as inclusion bands, while those at higher angles tothe 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 wallrock followed by incorporation of fragments into the filling.The latter mechanism may occur where irregularities on mi-crovein walls obstruct opening, and must be broken off foropening to occur (e.g. Urai et al. 1991). The presence ofnumerous inclusion bands is evidence for a cyclic process ofmicrocrack opening followed by filling, a process known ascrack-seal (Ramsay 1980). Each inclusion band representsone cycle. The width between inclusion bands isfor many rock types, and there may be up to thousands ofbands in a microvein (Ramsay and Huber 1987). Inclusiontrails are markers of the position of particular points on themicrovein margin at successive opening positions: they aretherefore parallel to the opening vector (Fig. 3.15b).

Paradoxically, microstylolites sub-parallel to microveinmargins have also been described, particularly along inclu-sion bands (e.g. Cox 1987). The shortening demonstratedby such microstylolites could be part of the crack-seal cycleif fluid pressures decreased to less than lithostatic in part ofthe cycle, causing the microvein to close and experience com-pressional stress. These microstylolites may also be due to alater deformation, unrelated to the microvein formation.

Laminated microveins have planar bands, often composedof phyllosilicates, subparallel to the margins. The bands


may form as inclusion trails where the opening vector has alarge component parallel to the margin (Cox 1987). In thissituation there is no distinction between inclusion trails andbands.

Fibrous microveins are particularly rewarding for kin-ematic interpretation. Fibrous fillings are common and formby progressive microvein opening at a rate that can bematched by crystallization of the filling. Fibres can growby at least five different filling sequences: these need to becarefully established before kinematic interpretations can bemade. Syntaxial growth means growth from the wall rocktowards the microvein centre on both sides of the fracture(Fig. 3.17a, Plate 21). The diagnostic features of syntaxialgrowth are two separate bands of fibres on either side of acentral suture. The fibres are not continuous across the suture.The microvein fill is similar to the wall rock and may be incrystallographic continuity with it. Fibre widths generally in-crease in the direction of growth due to competition betweenfibres, which ensures that the faster growing and thereforelarger fibres overgrow and eventually isolate the slower andsmaller fibres from the precipitating solution (e.g. Smith1964). Therefore in syntaxial growth, fibre widths may in-crease towards the suture. By contrast, antitaxial growth oc-curs from the microvein towards the wallrock, and may occursymmetrically on both sides of the microvein (Fig. 3.17b) orasymmetrically on only one side (Fig. 3.17c, 3.18). The dia-gnostic feature of antitaxial growth is fibre continuity acrossthe microvein. Antitaxial growth is characterized by fillingsof different material from the wall rock, and no crystallo-graphic relationship between the filling and wall rock. Sym-

metric antitaxial growth may result in a median line of inclu-sions along the centre of the microvein, unlike asymmetricgrowth. Fibre widths can also be used to distinguish sym-metric growth (fibre widths increase symmetrically in twoopposite directions towards the wall rock) from asymmetricgrowth, in which the widths increase unidirectionally acrossthe microvein. Composite growth means both syntaxial andantitaxial protions in the same microvein (Fig. 3.17d). Non-systematic growth (“ataxial” - Passchier and Trouw 1996)histories may involve fracture at any point in the crystalfibres, which are termed stretched crystal fibres by Durneyand Ramsay (1973). They exhibit none of the systematic fea-tures described for the other four categories of growth above(Fig. 3.17e), and the sides of the fibres have distinctive in-terlocking teeth, dividing then fibres into tablets (Plate 22).The lack of directional growth indicators (e.g. fibre wideningdirection) in these microveins is diagnostic.

In many fibrous microveins, the fibres grow parallel to theincremental opening vector: Such tracking or displacement-controlled fibres can be used to deduce incremental strainhistories with great effect. They can be recognized becausefibres connect markers across the microvein, because theyare parallel to inclusion trails (Urai et al. 1991), or becausethey have a constant orientation between microvein walls ofvariable shape (Plate 23). The strain history deduced fromthe fibres can be plotted on a diagram showing rotation ofthe incremental strain on the horizontal axis against strain onthe vertical axis (a cumulative incremental strain history orcish diagram, e.g. Fisher and Anastasio 1994, Hedlund et al.1994).


Non-tracking fibres do not track the incremental openingvector (e.g. Durney and Ramsay 1973). Non-trackingfibres can be recognized because they do not connectmarkers across the microvein (Fig. 3.19). However, markersmay not be connected even by tracking fibres if a micro-vein has an early history of shear displacement before filling

occurred. Inclusion trails may be used to distinguish track-ing and non-tracking fibres in this case. Non-tracking growthoccurs because fibre growth directions are determined by theorientation of the growth surface, like face-controlled pres-sure fringes. Fibre boundaries are either perpendicular to thegrowth surface (Fig. 3.19), or along the bisector of two ad-jacent growth surfaces. The fibre boundaries have no fixedrelationship 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 andthe growth surface and the shape of the growth surface(Fig. 3.19, Urai et al. 1991). These fibres can only be par-allel to the opening vector when when the trackingefficiency is 1.

Fibres in microveins with any of the above growth histor-ies are often curved. The curvature can be primary (formedduring crystal growth), or secondary (due to subsequent de-formation). Primary curvatures can form in both tracking andnon-tracking fibres due to rotation of the incremental open-ing direction with respect to the previously formed part of thefibre, and can be recognized because the curved fibres showno evidence of strain or recovery.

All low-temperature microvein fillings form by precipita-tion from solution, and are therefore excellent evidence forDMT via solution. Microveins are a conspicuous feature ofgreenschist-facies and lower grade deformation because bothcataclasis (fracture) and DMT via solution are required fortheir formation. At higher grades, other deformation mech-anisms operate, and microvein textures have low preservationpotential because of recrystallization.

Chapter 4

Intracrystalline Plasticity

4.1 Introduction

Permanent distortion of a crystal lattice without fracture oc-curs by intracrystalline plasticity. The definitive feature ofall intracrystalline plastic deformation mechanisms is the in-volvement of dislocation motion (Section 4.2), but solid statediffusion (Chapter 5) is an integral part of some mechanismsconsidered in this chapter. Intracrystalline plasticity has along history of research and a vast literature, particularly inthe materials science field, which is necessarily rather con-densed in this chapter. Useful references that provide moreexplanation and detail are Hull (1975) and Nicolas and Poir-ier (1976).

4.2 Fundamental mechanisms of in-tracrystalline plasticity

A dislocation is one type of imperfection or defect in a crystallattice. A useful way to understand dislocations is to imaginecutting into the lattice, stretching the lattice apart along thecut, and inserting an extra plane of atoms into the cut. Thedislocation is the line along the edge of the extra plane ofatoms (Fig. 4.1). The stretching of the lattice to accommod-ate the extra atoms causes a distortion, the orientation andsize of which is measured by the Burgers vector whichcan be perpendicular to the dislocation line (an edge disloca-tion), parallel to the line (a screw dislocation) or oblique (amixed dislocation) (Fig. 4.1). If has a magnitude of onelattice unit, the dislocation is described as perfect, but par-tial dislocations have with magnitudes of fractions of lat-tice units. The lattice distortion causes a stress field aroundthe dislocation, and is one way of storing strain energy in thecrystal. A direct image of dislocations can be obtained in thetransmission electron microscope (TEM) where they usuallyappear as dark lines (Fig. 4.2). They can also be revealedby the technique of etching, in which a flat crystal surfaceis exposed to acid, creating a small depression (etch pit) atthe intersection of a dislocation with the surface of the crys-tal. The etching occurs due to dissolution which is enhancedby the lattice distortion and strain energy of the dislocation.The density of dislocations can be used to measure past stressfields (Section 9.8.4). Movement of a dislocation is accom-plished by breaking bonds in the intact lattice ahead of thedislocation 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 slipplanes, in a direction within the plane known as the slip dir-ection; this process is called dislocation glide. The combin-ation of the slip plane and slip direction is known as the slipsystem, which is specified by the crystallographic orientationsof the slip planes and directions. For example, slip along theprism planes of quartz in the <a> direction is annotated by

Breaking bonds during glide requires en-ergy, which can be provided by heat, so the stress necessary tocause glide decreases with increasing temperature. Slip sys-tems in some directions are more active than in others becauseof anisotropy in crystal properties. The relative ease of glidebetween different slip systems changes with temperature.

Glide is impeded by impurities in the crystal lattice, andby the stress fields associated with other dislocations. Theseobstacles may be overcome by dislocations changing theirslip plane. Screw dislocations can accomplish this by a pro-cess called cross-slip (Fig. 4.4). Edge dislocations can changetheir slip plane if a lattice vacancy replaces the last atom in thehalf-plane of the dislocation (Fig. 4.4). This process of dis-location climb therefore involves diffusion, although it is anintracrystalline plastic deformation mechanism because dis-location motion occurs. Deformation by a combination ofdislocation glide and climb is called dislocation creep, andoccurs at higher temperatures, and lower strain rates than dis-location glide (Nicolas and Poirier 1976).

4.3 Deformation twins

A twin is a region of a crystal that is rotated or reflected withrespect to the rest of the crystal (the host). Twins may formduring crystal growth or deformation: the latter are known asdeformation or mechanical twins. Growth twins are generallystraight and of constant thickness, but deformation twins havevariable thickness (thinning and branching towards the edgeof a crystal), and are commonly bent (Plate 24, Fig. 4.8). De-formation twins are common in carbonates (where their mor-phology is temperature-dependent; Section 9.9.2), and feld-spars.

Deformation twins form by shear of the crystal lattice withrespect to the host lattice along the twin plane, together withminor rearrangements of the twinned lattice points. Thetwin plane is a mirror plane comprising an array of par-tial or twinning dislocations. The strain due to twinningcan be measured, and deformation twins can also be usedto measure stress and temperature during their formation

39

40 CHAPTER 4. INTRACRYSTALLINE PLASTICITY

CHAPTER 4. INTRACRYSTALLINE PLASTICITY 41

(Sections 9.8.5, 9.8.9, 9.9.2).

4.4 Undulatory extinction

Undulatory extinction is visible in cross-polarized light whena single crystal has variable extinction positions (Fig. 4.5).The extinction position may change consistently from oneend of a crystal to the other, producing a continuous sweepof extinction as the stage is rotated. Irregular patches of dis-tinct extinction positions can occur, and sectors of differentextinction positions may radiate from the centre of the crystal.Undulatory extinction is caused by distortion of the crystallattice by dislocations of a consistent orientation (Fig. 4.6a).It forms at low strains during intracrystalline plastic deform-ation, and as such it is a sensitive indicator of intracrystallineplastic deformation.

4.5 Intracrystalline deformationbands, kink bands and subgrains:Recovery

Intracrystalline deformation bands are tabular low-strain do-mains within crystals, separated from other parts of the crys-tal along approximately planar boundaries across which thereis a slight change in lattice orientation. “Intracrystalline” isused to distinguish these smaller scale features from cata-clastic deformation bands (Section 2.5). Kink bands are sim-

ilar to intracrystalline deformation bands, but have sharper,more planar boundaries. The change in lattice orientationacross an intracrystalline deformation or kink band may bevisible from changes in extinction position (Fig. 4.7) orchanges in the orientation of twin or exsolution lamellae infeldspars or calcite (Fig. 4.8). Kinks are different from twinsbecause the change in orientation across a kink plane is not afixed amount, and the kink plane is not a mirror plane.

Subgrains are seen under the microscope as areas withingrains of slightly different crystallographic orientation, separ-ated by boundaries subparallel to crystal planes. The variationin crystallographic orientation is visible in quartz as slightchanges in extinction position which affect discrete areas,in contrast to the progressive change in extinction positioncharacteristic of undulatory extinction. Subgrains in quartzare usually elongate parallel to the prism planes, which formthe subgrain boundaries (Fig. 4.9). Tabular subgrains thushave a similar geometry to kink bands. An alternative pat-tern in quartz consists of approximately square subgrains withboundaries parallel to both prism and basal planes, known asa chessboard pattern (Fig. 4.10). Basal subgrain boundariesare only visible in grains with their c-axes subparallel to theplane of the section, in contrast to prismatic subgrain bound-aries. Lattices with high dislocation densities possess a largeinternal strain energy (Section 4.2), which can be loweredby dislocation movement into surfaces surrounding relativelydislocation-free volumes (Fig. 4.6). This process of recoveryresults in a microstructure of low-energy volumes (subgrains)surrounded by walls across which there is a slightly differentlattice orientation: hence the walls are sometimes called low-angle boundaries. Undulose extinction, deformation and kinkbands, and subgrains are a sequence of microstructures thatform with progressive strain, by generation and movementof dislocations into low-angle boundaries (Fig. 4.6; White1976). The change in lattice orientation across the bound-aries, and the dislocation density in the boundaries, increaseswith strain.

Subgrain boundaries are orientated approximately perpen-dicular to the glide direction of the moving dislocationsand therefore the orientations of the subgrain boundaries


may change with temperature. Most natural quartz subgrainboundaries formed in the stability field are pris-matic, while in the stability field, both prismatic andbasal subgrain boundaries (chessboard pattern) form (Kruhl1996). These observations do not fully agree with exper-imental results, or with the suggestion that basal subgrainsare indicative of water-present deformation (Mainprice et al.1986), but the observations are robust empirical evidence, al-lowing the occurrence of chessboard patterns to be used as ageothermobarometer (Section 9.9.4). The size of subgrainscan be used as a palaeopiezometer (Section 9.8.3).

4.6 Deformation lamellae

Deformation lamellae are crystallographically orientatedplanar features wide (Drury 1993). Deforma-tion lamellae can be seen under the optical microscope us-ing a high magnification and a narrow diaphragm to en-hance relief contrasts. In quartz, they have a slight extinc-tion or refractive index contrast with the adjacent host grain(Fig. 4.11), and they commonly have a sub-basal orienta-tion, sub-perpendicular to prismatic deformation bands (e.g.Spang and Van der Lee 1975, Drury 1993). Deformationlamellae are commonly slightly curved, and have a single,consistent orientation within a grain. They are known fromboth experimentally and naturally deformed rocks, especiallyin quartz but also in olivine and plagioclase (e.g. Den Brockand Spiers 1991). The nature of deformation lamellae is enig-matic: observations include slip bands, walls of tangled dis-locations, twin boundaries and planes of glass (e.g. McLarenet al. 1967, Christie and Ardell 1974, Twiss 1974). However,TEM studies have shown that typical deformation lamellaein quartz consist of elongate subgrains wide with asub-basal orientation bounded by curved dislocation

walls made up of well-ordered arrays of two to three sets ofdislocations (Blenkinsop and Drury 1988, McLaren 1991).The change in orientation across the subgrain walls is lessthan 2°. The lamellae may have high fluid inclusion densit-ies, and variable dislocation densities which suggest a highlyrecovered structure (Blenkinsop and Drury 1988). These de-formation lamellae probably formed by recovery of disloca-tion slip bands, which leads to a variable decrease in dislo-cation density and precipitation of water in fluid inclusions(Drury 1993). Deformation lamellae can be used as a crudepaleopiezometer (Section 9.8.6). Planar deformation features(PDFs) are different type of planar feature caused by shockmetamorphism (Section 8.4, 8.11).

4.7 Grain shape fabrics and ribbongrains

One of the most characteristic microstructures formed by in-tracrystalline plasticity are flattened or elongated grains witha preferred orientation, resulting in a grain shape fabric, orshape preferred orientation (Fig. 4.12). The effect of move-ment of dislocations through a crystal is to change its shapetowards that of the strain ellipsoid. Apart from the distortionof individual grains, intracrystalline plastic grain shape fab-rics can also form by coalescence of grains of similar orient-ation during dynamic recrystallization (Section 4.8; Meansand Dhong 1982). Extreme strain can result in monocrys-talline grains with very large aspect ratios known as ribbongrains (Fig. 4.13); they may have intracrystalline plastic de-formation features such as undulatory extinction or subgrains.Grain shape fabrics can also be produced by both cataclasisand DMT. Intracrystalline plastic grain shape fabrics can bedistinguished by their association with other intracrystallineplastic microstructures.

4.8 New grains, core and mantle struc-ture: Dynamic recrystallization

New grains, usually equant and strain free, are commonaround and within larger original grains in moderately orhighly strained rocks. The new grains may have a strong crys-tallographic fabric which is often systematically related to ad-jacent older grains. The new grains may be concentrated ina mantle that partly or completely surrounds the older grains,described as core-and-mantle structure (Fig. 4.14), and theymay have similar sizes and orientations to adjacent subgrains.These features are characteristic of dynamic recrystallization,or recrystallization that occurs syntectonically. The lack ofstrain in the new grains shows that they are at an advancedstage of recovery. The new grains form by two main mech-anisms: subgrain rotation (SGR), and grain boundary migra-tion (GBM). SGR and GBM are structural transformationswhich do not per se accommodate deformation, and thereforethey are not strictly speaking deformation mechanisms (Uraiet al. 1986), but merely structural rearrangements.

Movement of dislocations into subgrain walls during re-covery causes progressive rotation of the subgrains (see


previous section) leading to the formation of a new grain(Fig. 4.15). The collection of dislocations into the sub-grain walls may occur by subgrain walls sweeping throughthe crystal, or by movement of the dislocations towards thewalls, and some GBM probably accompanies SGR (Urai etal. 1986). There is a continuum between subgrains and newgrains formed by this process, so that the definition of a newgrain as distinct from a subgrain is somewhat arbitrary. Agenerally used criterion is that the new grain lattice orient-ation differs from its host by more than 10°, but other val-ues have been used (e.g. Urai et al. 1986). SGR is read-ily inferred from the coexistence of new, smaller grains withsubgrains of similar size within the older grains (Fig. 4.16).Other diagnostic microstructures for SGR include clusters ofnew grains with similar orientations, which they inherit froma single parent grain (Urai et al. 1986), and the coincidenceof one or more lattice directions in adjacent subgrains. GBMoccurs by the movement of a grain boundary from one graininto another, and finally by the closure of the boundary toisolate a new grain in a different lattice orientation from thehost (Fig. 4.17).

The grain boundary migrates in response to an internalstrain energy gradient from a less-deformed grain with alower dislocation density to a more highly strained one. Thefinal closure of the boundary may be achieved by an in-termediate stage consisting of a bridging subgrain bound-ary that accommodates progressively more misorientation(Means 1981). GBM can be recognized by the presence of


new grains along grain boundaries, together with the serratedand lobate shapes of grain boundaries (Fig. 4.18, Plate 25).The characteristic shape of these grain boundaries, also calledsutured grain boundaries, is caused by the fact that theboundaries migrate away from their centre of curvature (e.g.Hirth and Tullis 1992). A bimodal distribution of grain sizesmay develop between the old and new grains (e.g. Lloydand Freeman 1994). New grains often develop where thereis large lattice distortion and changes in lattice orientation,such as deformation bands and kink planes (Fig. 4.5). Thisoccurs because grain boundaries migrate more rapidly at amisorientation of about 5°. Drury et al. (1985) emphasizethe importance of these special sites by proposing a third cat-egory of dynamic recrystallization: sub-boundary migration,which is defined as subgrain growth in areas with large gradi-ents of strain and orientation. New grains also form in highstress sites around porphyroclasts. The size of recrystallizedgrains can be used as a paleopiezometer (Section 9.8.2).

Dynamic recrystallization in naturally-deformed quartz oc-curs by either a mixture of 50% SGR and GBM (Fig. 4.19),or by SGR alone (Drury et al. 1985). Both mechanisms maybe observed because they occur sequentially and cyclically(Lloyd and Freeman 199la, b, 1994). Three regimes of dy-namic recrystallization in quartz, depending on deformationconditions, have been indentified from experiments (Hirthand Tullis 1992). The Hirth and Tullis classification of dy-

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


A crystallographic fabric, or lattice or crystal preferred ori-entation (L or CPO) is a concentration of lattice orientationsin one or a limited number of directions. Crystallographicfabrics can be recognized quickly under a optical microscopeby observing that many grains have the same extinction pos-ition or interference colour. A useful way to detect the pres-ence of a crystallographic fabric in more detail is to insertthe sensitive tint plate or a quarter-wavelength plate, whichdistinguish the different optical axes and therefore provides amore precise picture of the lattice orientations than extinctionposition or interference colour. This is particularly useful forquartz because of its low order interference colours. Detailedmeasurements of crystallographic orientation are made withthe universal stage under the optical microscope, by electrondiffraction, channelling, and backscattering techniques underthe electron microscope, and by X-ray and neutron diffrac-tion. Crystallographic fabrics can be produced in rocks byat least four distinct mechanisms (e.g. Hobbs et al. 1976,Mainprice and Nicolas 1989). The first two mechanisms, an-isotropic crystal growth (Chapter 5.3) and rigid-body rotation


of grains in a flowing matrix, can be recognised because in-dividual grains are not deformed. A simple way in whichintracrystalline plasticity can produce a crystallographic fab-ric is shown in Fig. 4.20, in which a grain is deformed by theoperation of a single slip system. As the grain rotates, theslip plane becomes orientated normal to the finite shorteningdirection and the slip direction rotates towards a bulk shearplane (cf. Allison and LaTour 1977). Although this is a sim-plified case, it illustrates that lattice reorientation can occurby dislocation glide. Fourthly, crystallographic fabrics maybe created by recrystallization. The exact mechanism is un-clear, but experiments demonstrate that the mechanism of dy-namic recrystallization determines the type of LPO (Gleasonet al. 1993).

A great deal of work has gone into theoretical attempts topredict crystallographic fabrics, and to compare them withexperimental results and natural LPOs (see reviews by Law1990, Wenk and Christie 1991). The general developmentof a crystallographic fabric by intracrystalline plasticity can

be understood through relatively complex models that allowfor multiple slip systems and make additional assumptions,such as strain compatibility between grains, uniform stress,or viscoplastic self consistent theory, which minimizes stressand strain differences from an average value (e.g. Wenk andChristie 1991). The work of Jessell (1988a, b, Jessell andLister 1990) for quartz incorporates effects of both the intra-crystalline plastic mechanisms referred to above, lattice rota-tion and recrystallization. These models show that the dom-inant slip plane does not necessarily align with the bulk shearplane, and the dominant slip direction is not necessarily paral-lel to the shear direction, in contrast to the single slip systemmodel above. The simulations also show that the LPO is af-fected by the type of flow during deformation, the finite strainmagnitude and type, and the temperature, which determineswhat slip systems are active. Crystallographic fabrics carrylarge amounts of structural information, and have been usedto analyze shear sense (Section 7.9), finite strain, and gradeof deformation.

Chapter 5

Diffusive Mass Transfer and PhaseTransformations in the Solid State

5.1 Introduction

Material removal, transport and deposition without fractur-ing, lattice distortion or melting at metamorphic grades atand above amphibolite facies suggest diffusive mass trans-fer (DMT) in the solid state. The major evidence for solidstate as opposed to fluid assisted DMT is provided by phe-nomena that can be explained in terms of known solid statediffusion coefficients, including many metamorphic textures.Recent experimental evidence suggests that a variety of trans-formations, some involving DMT, may also be important de-formation mechanisms in the solid state (Section 5.9). Therehas been considerable recent discussion about the potentialtectonic importance of superplasticity, which involves a com-posite solid-state deformation mechanism: these microstruc-tures and mechanisms are described in the last section, 5.10.

5.2 Fundamental deformation mech-anisms of solid state diffusive masstransfer and phase transforma-tions

DMT in the solid state may occur though the crystal lat-tice by the movement of lattice defects. This is known asvolume diffusion, and the resulting deformation is Nabarro-Herring creep (e.g. Nicolas and Poirier 1976). Coble creepis a different type of deformation resulting from solid stateDMT through grain boundaries, which have different diffu-sion characteristics from the grain interiors (e.g. White andWhite 1981). Deformation by either or both types of creepis known by the general term diffusion creep. Diffusion oc-curs in response to gradients in chemical potential (Fick’slaw) 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 fourtypes (Green 1985, Kirby and Stern 1993), which may be im-portant in deformation because they are sensitive to deviatoricstress. The transformed polymorph is related to the originalby crystallographic rules in the first three types, which areknown as coherent transformations.

1. Displacive transformations are changes in crystal struc-

2.

3.

4.

ture without bond breaking and with minor shapechange. An example is the phase transition inquartz.

Martensitic-like transformations are coherent transform-ations involving dominantly shear strain. Importantgeological examples include:

and

Coherent exsolution may involve both dislocation move-ment and DMT to effect crystallographic and chemicalrearrangements. Exsolution of clinopyroxene from or-thoenstatite is one of the best studied examples.

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

5.3 Grain shape fabrics and ribbongrains

A grain shape fabric of unstrained grains is an important mi-crostructure in metamorphic rocks at higher grades (Plate 26).The grain shape fabric can be produced by anisotropic grainboundary migration recrystallization, although the details ofthe mechanism are not clear (e.g. Jessell 1987), and by aniso-tropic crystal growth (e.g. Shelley 1989a, b). Both of theseprocesses are important DMT mechanisms.

Quart-mica rocks at higher metamorphic grades commonlyhave a distinctive microstructure of flat grain boundariesbetween quartz and mica parallel to the mica basal plane,and quartz-quartz grain boundaries approximately perpendic-ular to the mica flakes (Plate 27). A related effect is pin-ning of quartz-quartz grain boundaries at the end of micaflakes. These microstructures attest to impeded grain bound-ary movement, and can be attributed to a greater surface en-ergy between quartz and mica than between quartz and quartzgrains. Where one phase has a strong preferred orientation,the impeding effect on grain boundary growth may create agrain shape fabric in the other phase, which grows parallel tothe shape fabric (Plates 26, 27).

Monocrystalline quartz ribbon grains (Fig. 5.1) in highgrade gneisses may form by the above mechanism, andare characteristically free of internal structures (by contrast

52

CHAPTER 5. DIFFUSIVE MASS TRANSFER AND PHASE TRANSFORMATIONS IN THE SOLID STATE 53

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

to lower grade ribbons; Section 4.7, Fig. 4.13). Gowerand Simpson (1992) proposed that the geometry of quartz-feldspar grain boundaries in ribbon grains is largely con-trolled by a combination of dislocation creep and diffusionthat results in a microstructure of straight quartz-feldsparboundaries perpendicular to the shortening, and cusps point-ing in the extension direction. However, recently MacKinnonet al. (1997) have proposed that the textures of some ribbongrains may form by filling of microfractures, based partly onthe evidence that grains adjacent to the ribbons appear to betruncated by the ribbon boundaries, and that the ribbons con-tain wall rock particles that have very similar geometries toinclusion bands (Section 3.12). Some of these features areillustrated in Fig. 5.1. Both models account for many of theobserved features of high-temperature ribbon grains. A pos-sible approach to distinguishing the two mechansims may liein careful examination of the ribbon tips, which could yieldevidence for incipient microcracking or diffusion processes.

5.4 Foam texture, static and second-ary recrystallization

Monomineralic metamorphic rocks, particularly at highermetamorphic grades, may have a distinctive microstructureconsisting of approximately hexagonal grain sections withstraight or curved grain boundaries and strain-free grain in-teriors, known as foam or granoblastic polygonal texture(Fig. 5.2). This microstructure is readily interpreted in termsof grain boundary migration and DMT in the solid statedriven by surface energy. The process of achieving a min-imum energy configuration under hydrostatic stress is staticrecrystallization. The lowest surface energy will occur for thesmallest surface area to volume ratio. The minimum surfaceenergy configuration for equal-sized, space-filling polyhedraare rhomb dodecahedra (truncated octahedra). However, thisfigure does not have an isotropic distribution of surface en-ergies, because angles between grain edges around four grainjunctions are unequal. Grain edges and boundaries may curveto allow all four grain edges to intersect in the same angleof 109.5°, resulting in a structure of polyhedra with curvedfaces, similar to liquid films and as observed in annealed al-loys and chromitites in ultramafic intrusions (Smith 1964). Intwo dimensions, 120° triple junctions are common betweenaggregates of the same phase. The surface area to volumeratio can also be reduced by an increase in grain size, whichoften accompanies static recrystallization, and is known asOstwald ripening or exaggerated grain growth. This micro-structure can be useful in revealing a post-tectonic period ofrelatively high temperatures.

Grain shape fabrics can be produced even under static re-crystallization by at least two processes. An older fabriccan control grain boundary mobility by surface energy effects(Section 5.3) or a new mineral growth may overgrow a preex-isting fabric: this is known as mimetic crystallization.

Relatively large, irregular crystals which commonly con-tain inclusions of secondary phases are formed by the pro-cess of secondary recrystallization. This occurs when secondphases become incorporated into the growing crystal, rather

than pinning the crystal boundaries as described in Sec-tion 5.3, which are then free to grow at a faster rate. Thishas been distinguished as a separate type of grain boundarymigration mechanism called fast or free grain boundary mi-gration by Urai et al. (1986).

5.5 Decussate texture

Decussate texture consists of randomly orientated, interlock-ing elongate crystals (Plate 28). It is especially common inamphiboles and micas, and arises because of unequal growthrates in different crystallographic directions. The growthanisotropy modifies the ideal foam texture to favour grainboundaries parallel to the faster growth directions, and maylead to randomly orientated acicular crystals for particularlyhigh anisotropies.

5.6 Porphyroblasts and inclusiontrails


Porphyroblasts are single crystals grown during metamorph-ism with a larger size than the adjacent grains in the matrix.They are only widespread in rocks that have been at uppergreenschist facies or higher, and are most common in meta-pelites or metabasites. Chlorite, chloritoid, biotite, garnet,cordierite, sillimanite, kyanite, andalusite, and staurolite arecommon porphyroblastic minerals.

Porphyroblasts commonly contain inclusions, the mostcommon of which are opaque minerals, aluminium-richphases such as sillimanite or spinel, quartz, zircon, apatite, ru-tile, and sphene. A porphyroblast with a very high density ofinclusions is a poikiloblast. Inclusions may have a variety oftextures that contain important microstructural information,and need to be described carefully. The shape of individualinclusions may be euhedral, platy, linear, or rounded. Inclu-sion trails are aligned inclusions that define a fabric which isgiven 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 par-allel to the length of the porphyroblast. may be straightor curved. It may be continuous and curve smoothly fromthe core of the porphyroblast to the rim; it may be sharplydeflected or cut off along deflection and truncation surfacesrespectively, which are sub-parallel to more external parts ofthe inclusion trail.

Special types of inclusion trails include snowball textures(Plate 29), which are spiral-shaped inclusion trails that curvethrough more than 180°, with the total curvature decreasingfrom the centre towards either end of the rotation axis (e.g.Powell and Treagus 1970, Busa and Gray 1992). Millipedetexture consists of a straight, parallel which is deflectedinto at the porphyroblast margin. Helicitic texture consistsof folds defined by Inclusions may be crystallographicallycontrolled by the porphyroblast: this can give rise to a numberof distinctive textural zoning patterns such as sector zoningand re-entrant zones. The relation between and is very


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

5.6.2 Growth mechanisms

Porphyroblasts are a mineral growth and therefore form byDMT. In many cases there is no evidence for a fluid phaseand it is assumed that diffusion was in a solid state. Thiscan sometimes be demonstrated by the observation that themost mobile components are those with the highest solid statediffusion coefficients. Porphyroblasts grow in response tothermodynamic considerations including composition, tem-perature, internal strain, and surface energy. Unusually largeporphyroblasts, especially those which contain high densitiesof inclusions, may have grown by secondary recrystallization(Section 5.4). The size and distribution of porphyroblasts re-flects a balance between the energy required for nucleationand that for growth. High ratios of nucleation to growth ener-gies will favour the formation of fewer and larger porphyro-blasts.

5.6.3 Relationship to deformation

Porphyroblast textures can yield detailed information aboutthe relation between porphyroblast growth (and thus P-T con-ditions) and deformation. This is best analyzed by lookingat the relationship between inclusions and which can beclassified into four non-interpretive categories (Fig. 5.3):

1.

2.

3.

4.

Inclusions are randomly orientated, is commonly alsocurved around the porphyroblast (Fig. 5.3a).

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

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

is continuous with and and have similarshapes and orientations (Fig. 5.3d).

These categories are objective descriptions, but also allowstraightforward interpretations based on the original ideas ofZwart (1960, 1962), and the recent comprehensive treatmentof Passchier and Trouw (1996). The first category suggestsgrowth of a porphyroblast before deformation i.e. pretectonic(e.g. Borradaile et al 1982, p. 438-441), but exceptions arenoted by Passchier and Trouw (1996). The curvature of isdue to deformation of the matrix around the more rigid por-phyroblast (Plate 30).

The second category is called intertectonic by Passchierand Trouw (1996), and indicates porphyroblast growth fol-lowing the deformation event that generated but preced-ing another event that formed (Fig. 5.4). The two deform-ations that created and may be separated by other de-formations that are not recorded by or (e.g. Johnsonand Vernon 1995). An intertectonic interpretation is reliablydemonstrated when is folded and is straight (Plate 31,e.g. Borradaile et al. 1982, p. 441, 453).

The third category indicates syntectonic porphyroblastgrowth (Plate 32). A common example of different geomet-ries assumed by and is a coarser grain size in than

which can be interpreted to show that the matrix was over-grown by the porphyroblast relatively early and subsequentlyunderwent grain size reduction. Coarsening of towards themargin of a porphyroblast indicates growth during progradeconditions. A particularly complex pattern may be producedby overgrowth of one type of in a pressure shadow and an-other type of in a strain cap, both Si being formed in thesame event (Shoneveld 1977).

The fourth category indicates post-tectonic growth (Plate33). However, even in the case of concordant and withno obvious differences in geometry, may be curved aroundthe porphyroblast, suggesting either that the last incrementof deformation outlasted porphyroblast growth, or that lowstrain deformation occurred subsequently.

Detailed analyses shows that some porphyroblasts can notbe accommodated into the above simplified scheme. Trunca-tion of within the porphyroblast may indicate two phasesof porphyroblast growth and foliation development. Trunca-tion surfaces can also be formed by post-tectonic overgrowthof an early foliation which has been truncated by a strain cap


during a later deformation, or a later stage of a progressive de-formation (Passchier et al. 1992). Bell et al. (1992) have ar-gued that truncation surfaces form in the strain shadow at theend of a porphyroblast during progressive deformation. Thesame interpretations may apply to deflection surfaces. Prob-lems of interpreting curved inclusion trails are discussed inSection 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 sincemetamorphic mineral growth requires mobility of compon-ents. Relic minerals are mineral grains that have been mostlyreplaced by reaction rims. Reaction rims that completely sur-round a grain are coronas (Plate 34), which are more commonin high grade rocks. Symplectites are lamellar or vermicularintergrowths that are common in reaction rims (Plate 35). Theintimate mixture of the phases involved, the fine grain sizeand the disequilibrium grain shape shows that diffusion wasnot possible over long distances.

5.8 Chemical zoning

Metamorphic minerals, especially garnets, amphiboles,pyroxenes and feldspars, commonly show systematic com-positional variations from core to rim, or a relatively con-stant core composition and a quite different rim composition.These variations can be seen in the colours of some min-erals (e.g. blue-green colour variations in amphiboles), butare most accurately revealed by electron microprobe studies.Zoning is clear evidence for DMT, and may form in two dif-ferent ways.

Growth zoning. Preferential partitioning of an element intoa mineral during growth can cause growth zoning by de-pletion of that element within an equilibrium volume.This describes how growth zoning may occur at con-stant temperature and pressure: however, the P-T con-ditions may change during mineral growth, which willalso change the distribution coefficients for elementsbetween mineral and matrix. This is probably the mostcommon way in which growth zoning is produced.

Reaction zoning. After a crystal has formed, DMT may oc-cur with neighbouring minerals in response to a changein P-T conditions from the formation of the originalmetamorphic rock. Commonly the rims of minerals havecompositions that reflect lower metamorphic grades thanthe cores: the rims have equilibriated during the coolingof a rock. Such zoning is known as retrograde zoning.

5.9 Solid state phase transformationmicrostructures

Experiments suggest that phase transformations under devi-atoric stress could have distinctive microstructures, but nat-

ural examples are so far virtually undescribed. Thequartz phase transition can be viewed as Dauphiné twinningon the scale of the unit cell, so that relict Dauphiné twins aswell as microfractures (Section 2.3.11) could constitute mi-crostructural evidence of this deformation mechanism (Kirbyand Stern 1993).

Martensitic-like transformations can occur during cool-ing under hydrostatic stress, but in the case of the

transformation, deformation-induced trans-formation can be distinguished by lack of twinning in theclinoenstatite (e.g. Kirby and Stern 1993). The best knownmicrostructures of Martensitic-like transformation are thoseof the olivine-spinel system. “Fingers” or lobes of spinelgrow into olivine parallel to in experiments on(Vaughan et al. 1984), and a similar phase transformation mi-crostructure of aragonite crystals has been observed growinginto calcite parallel to (Hacker and Kirby 1993). Thesemicrostructures are consistent with the theoretical analysisof Green (1985), which suggest that ellipsoids of the stablephase should grow with long axes parallel to possibly co-alescing into ellipsoids separated by cusps. However, laterexperiments on olivine showed lens shaped spinel inclusionsperpendicular to These can be accounted for by stress re-distribution at a microscopic scale between stronger olivineand weaker spinel (Green and Burnley 1989).

Exsolution of Ca-clinopyroxene from orthoenstatite is sim-ilar to the transformation describedabove, with the addition of Mg, Fe and Ca diffusion. Coolingexsolution may be distinguished from deformation exsolutionby the orientation of the exsolved phase (Champness and Lor-imer 1974).

5.10 Superplasticity

Superplasticity was first used to describe the behaviour ofmetals deformed in extension to large strains without failure(e.g. Langdon 1982). The dominant deformation mechan-ism in these experiments is grain boundary sliding, with in-compatibilities between grains mainly relieved by solid stateDMT around grain boundaries. Unfortunately the term hasbeen used in at least three different ways in a geological con-text (Gilotti and Hull 1990):

1.

2.

3.

As a deformation mechanism consisting of grain bound-ary sliding accommodated by DMT and/or intracrystal-line plasticity.

As a set of deformation conditions and mechanical re-sponses in rock. In particular, dependence of strain rateon an inverse power of grain size, high temperatures, andproportion between strain rate and a low power of stresshave been regarded as characteristic (Section 9.4.3).

As a description of strain e.g. “continuous, homogen-eous deformation to very large strain”. This “phenomen-ological” definition proposed by Gilotti and Hull (1990)is not in widespread use, and the more common first geo-logical definition will be followed here.

Microstructures characteristic of rocks considered to havebeen deformed by Superplasticity have been summarized by


Fliervoet and White (1995). They include:

1.

2.

3.

4.

5.

6.

7.

8.

9.

A very fine grain size (less than in quartzites,more in carbonates) (e.g. Behrmann 1985).

Equiaxed grains with diamond or blocky shapes (e.g.Drury and Humphreys 1988).

Alignment of grain boundaries over several grains (e.gWhite 1977).

An inverse correlation between finite strain and grainsize (e.g. Evans et al. 1980).

A weak crystallographic fabric (Rutter et al. 1994).

An inverse correlation between dislocation density andgrain size (e.g. Behrmann 1985).

High dislocation densities and voids at grain triple junc-tions (e.g. White 1977).

Large mismatches between the crystallographic orienta-tions of adjacent grains (e.g. Fliervoet and White 1995).

Rotations between grains that can not be explained byrotation axes perpendicular to known Burgers vectors(e.g. Fliervoet and White 1995).

Several of these microstructures are ambivalent in their in-terpretation. For example, Fliervoet and White (1996) de-scribe an exceedingly fine-grained quartz mylonite which de-formed exclusively by dislocation creep with no grain bound-ary sliding. A fine grain size is apparently a necessary butnot sufficient condition for superplasticity. Crystallographicfabrics may be developed during superplastic flow of fine-grained calcite rocks, with the implication that they could inprinciple become strong after sufficient strain (Schmid et al.1987, Rutter et al. 1994). Fig. 5.5 shows an ultramylonitethat shows some of the characteristic features of superplas-ticity. The clear definition and recognition of superplasti-city in rocks is still a difficult matter. An important exampleof superplasticity related to phase transformation may occurwhen olivine transforms to spinel. This is known as the anti-crack theory of phase transformation faulting, which suggeststhat olivine transforms into spinel in anticracks which local-ize into faults within which grain boundary sliding occurs onfine grained spinel (Green and Burnley 1989, Burnley et al.1991). The theory accounts well for earthquakes that occur atdepths too great for frictional behaviour, and is well suppor-ted by experimental evidence (e.g. Tingle et al. 1993).

Chapter 6

Magmatic and Sub-magmatic Deformation

6.1 Introduction

Identification of deformation microstructures and mechan-isms in rocks containing melt has several important tec-tonic implications, notably for the problem of melt extractionand in the interpretation of pluton ascent and emplacementmechanisms. However, microstructural criteria to distinguishmagmatic, sub-magmatic and non-magmatic deformation mi-crostructures and mechanisms are not well established, andthe distinction is best made using a combination of meso-scopic and microscopic evidence. The mesoscopic evidenceis briefly described in Section 6.4.

6.2 Fundamental deformation mech-anisms and microstructures inrocks containing melt

6.2.1 Magmatic flow

Flow of magma (i.e. melt and crystal phases) by transportof rigid crystals is often regarded as the typical deformationmechanism in melt-bearing rocks. Magmatic flow has beendefined as flow by displacement of melt and rigid-body rota-tion of crystals without sufficient interaction to cause crystalplastic deformation (Paterson et al. 1989). As shown be-low, crystal interaction in melts may lead to deformation bycataclasis and diffusive mass transfer (DMT) as well as in-tracrystalline plasticity. A more useful, general definition formagmatic flow is flow of melt and crystals without crystal de-formation; this definition allows for the possibility that crys-tals may be deformed by processes other than crystal plas-ticity, and also describes the flow of a suspension. It leadsnaturally to a definition of magmatic microstructures as thosewhich indicate melt-present deformation without crystal de-formation.

Some experimental studies suggest that viscosities of melt-laden systems reduce abruptly by orders of magnitude whenthe proportion of melt increases beyond a value known as thecritical melt fraction, CMF (Arzi 1978, Van der Molen andPaterson 1979). The CMF is commonly taken to be 30%, butmay be as much as 50% (Vernon et al. 1988) or as little as10-20% for gabbroic rocks (Nicolas et al. 1988). Experi-ments on two silicate melts reported an increase in viscosityof three orders of magnitude, and a change from Newtonianto non-Newtonian behaviour, as melt fraction increased from

40 to 60% (Lejeune and Richet 1995). The importance of theCMF is further suggested by the observation that the max-imum proportion of phenocrysts in volcanic rocks is 55-65%:volcanic rocks with larger proportions of phenocrysts may notbe able to erupt because their viscosities are too high (Marsh1981, Wickham l987).

A minimum melt proportion for magmatic flow can alsobe deduced from the critical packing density of crystals tobring them into a coherent mass. The critical packing densitydepends on crystal shape, size, size distribution, packing ar-rangement, and amount of compaction. Table 6.1 summarizesporosity at critical packing density, which is equal to the min-imum melt proportion, for some combinations of these factorsin geometrical models, and estimates for magmas. Surfaceenergies of the crystal and liquid phases may be important inflow because they can affect melt distribution (e.g. Jarewiczand Watson 1984, 1985), and indeed an explicit relationshipbetween the CMF and surface energy can be formulated (Ri-ley 1990). A more fundamental parameter than melt fractionfor determining the rheology of melt-laden systems may bethe contiguity (the fraction of grain surface area in contactwith other grains), because contiguity affects surface energyand the resistance to shearing at the average surface. A loadbearing framework of crystals breaks down at contiguities ofless than 0.15-0.2 (Miller et al. 1988), which correspond tovariable equilibrium melt fractions depending on surface en-ergies and grain size and shape distributions (German 1985).These factors may explain the variation in estimates of theCMF.

Viscosities of suspensions depend on fluid (melt) com-position, pressure, temperature, and proportion, size andshape distribution of solids. The viscosity of melts contain-ing spherical crystals is often approximated by the Einstein-Roscoe equation (Roscoe 1952):

where is the viscosity of the pure melt, and is thecrystal fraction for coherent packing. Values of 0.6 forseem to represent silicate systems well. n has a theoreticalvalue of 2.5 for variably-sized spheres, which is also a goodfit to data for silicate systems, but the viscosity-melt fractionrelationship has a slight temperature dependence which canbe allowed for by letting n vary between 2.0 and 2.5 withtemperature (Lejeune and Richet 1995). Another relationshipbetween viscosity and crystal fraction includes the effect ofgrain 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 themechanical properties of the solid phases or by the factors af-fecting diffusive mass transfer (DMT) through the melt phase(see below).

While some experiments and observations suggest a signi-ficant mechanical change at the CMF, other experiments castdoubt on the existence of a large change in viscosity over anarrow range of melt fraction (e.g. Rushmer 1995), and ithas been argued that the abrupt change in viscosity observedin some previous experiments was due to a combination oftemperature effects and increase in water contents of melts(Rutter and Neumann 1995). There may be a continuous de-crease in magma viscosity as melt fraction increases, and noCMF, if the change in water content is allowed for. Yet otherexperiments suggest that the change in strength at the CMFis due to dilatancy hardening (cf. Brace and Martin 1968) atlow melt fractions (Renner et al. 1999).

6.2.2 Sub-magmatic flow

Sub-magmatic flow can be defined as deformation involvingflow of melt and crystals with crystal deformation. Sub-magmatic microstructures are accordingly those indicatingmelt-present deformation with crystal deformation, and cor-respond to the pre-full crystallization fabric of Hutton (1988).Crystal deformation by DMT through the melt phase is prob-ably the dominant deformation mechanism at low melt frac-tions and lower strain rates. The thermodynamic consider-ations for DMT through melt are the similar to those de-scribed for DMT through solution in Section 3.2. Experi-ments show that strength decreases by an order of magnitude,and intracrystalline plasticity changes to melt-enhanced diffu-sion creep, when melt proportion increases from zero to only3-5% (Cooper and Kohlstedt 1984, Dell’Angelo and Tullis1988, Dell’Angelo et al. 1987). The weakening occurs dueto enhanced grain boundary diffusion though the melt. Ex-periments on rock analogues by Park and Means (1996) alsodemonstrate the importance of this process in some systems atlow melt fractions. Surface energy considerations may be im-portant in sub-magmatic deformation. Most diffusion creepmodels assume isotropic surface energy, but measurementssuggest an order of magnitude anisotropy in olivine, for ex-ample (Cooper and Kohlstedt 1982). This may affect strengthby allowing a continuous film of melt along two-grain bound-aries instead of limiting melt to three or more grain junctions,as usually assumed (Hirth and Kohlstedt 1995).

6.2.3 Magmatic and sub-magmatic flow andrheology

It appears that two fundamental changes in deformationmechanisms occur as melt fraction is increased from thesolid to liquid states, and these correspond to reductions instrength or viscosity. An order of magnitude decrease instrength occurs as melt fraction rises from zero to only afew percent melt, and solid-state deformation mechanismschange to melt-enhanced diffusion creep, with other deform-ation mechanisms (cataclasis, grain boundary sliding, and in-tracrystalline plasticity) also possibly accommodating crys-tal deformation. As melt fraction increases still further, sus-pension flow can occur, reducing strength by three ordersof magnitude, and crystal deformation ceases. Figure 6.1shows these two transitions in mechanism schematically, butit should be emphasized that the values and range of melt frac-tions over which the transitions occur, and the magnitudes ofthe changes in viscosity, vary with magma composition, andare presently the subject of some debate.

6.3 Mesoscopic evidence for magmaticand sub-magmatic flow

It is important to use mesoscopic evidence in conjunctionwith microscopic observations to distinguish magmatic, sub-magmatic and non-magmatic flow. Fabrics in magmatic rockscan be defined by phenocryst alignment, matrix mineral shapefabrics, compositional banding, enclave or schlieren shapefabrics, or magnetic anisotropy. Phenocrysts may be tiled orimbricated (e.g. Blumenfeld 1983, Blumenfeld and Bouchez1988). The fabrics can have any shape from prolate to ob-late. The mere existence of any of these fabrics is not dia-gnostic of melt-present flow, but detailed examination on themesoscopic scale may provide some strong indicators to dis-tinguish magmatic/sub-magmatic from non-magmatic flow.The following criteria are proposed as diagnostic of magmaticor sub-magmatic flow because they suggest extremely large,non-systematic strain gradients in an homogeneous rock ona meter scale, which are mechanically unlikely in the non-magmatic state. Microscopic studies should be used to backup these lines of evidence.

1.

2.

3.

Abrupt and non-systematic changes in fabric orientation.

Abrupt and non-systematic changes in fabric intensity.

Irregular, polyclinal, disharmonic or rootless folds in

CHAPTER 6. MAGMATIC AND SUB-MAGMATIC DEFORMATION 61

massive rocks that have no sign of mechanical aniso-tropy or shear surfaces (e.g. McLellan 1984).

The relationship between igneous layering and fabrics canalso be used to suggest magmatic/sub-magmatic flow. Forexample, Pons et al. (1995) describe centimetre-scale cycleswith sharp bases consisting of a layer of fine-grained am-phiboles to a medium grained amphibole-plagioclase layer toa slightly porphyritic plagioclase-K-feldspar-quartz layer inalkali granites. The layers are parallel to the preferred orient-ation in any of these minerals, and the fabric never cross-cutsthe layering. The layering and fabric can be explained bytrapping of a mafic cumulate layer from a slowly convectingmagma, leaving an increasingly felsic rich magma to crys-tallize as the upper layers. By contrast, Paterson and Ver-non (1995) discuss several examples of an apparently mag-matic foliation that cross-cuts both gradational compositionalchanges and contacts between phases of different magmatic

compositions. These foliations may have formed in a shortinterval of cooling after the melt has been emplaced but be-fore final solidification.

Phenocryst density and size distributions may be affectedby magmatic flow. Interaction between phenocrysts in con-centrations greater than 8% in a moving fluid creates a “graindispersive pressure” which is proportional to the velocitygradient in the magma and is therefore greatest at the marginsof an intrusion (Bagnold 1954, Komar 1972a, b). Phenocrystsare concentrated into the centre of the intrusion, and may alsocoarsen in this direction. The velocity distribution and pheno-cryst concentration should be plug-shaped with a central re-gion of high density and an abrupt decrease in density towardsthe sides of an intrusion, due to the effect of phenocryst con-centration on viscosity, and this is indeed observed in manynatural examples (e.g. Ryan 1995).

Discordance between fabrics in a xenolith and a surround-ing igneous rock is often interpreted as evidence for mag-


matic flow, but this criteria is not diagnostic: fabrics maybe preserved in xenoliths during non-magmatic deformationdue to competence contrasts between the xenolith and sub-solidus matrix. Sub- or non-magmatic fabrics on this scaleare demonstrated by deformation of individual crystals suchas feldspar megacrysts. However, concordance between fab-rics within enclaves and the host rock can be evidence formagmatic/sub-magmatic flow if the enclaves can be inter-preted as partially molten during deformation (e.g. Vernonand Paterson 1993).

Shear zones may develop during intrusion by magmatic,sub-magmatic and non-magmatic mechanisms (e.g. Guine-berteau et al. 1987, Pons et al. 1995). Magmatic shearzones can be identified by fabrics defined by unstrained ig-neous minerals (e.g. Miller and Paterson 1994) or reorient-ation of a magmatic fabric in the shear zones (e.g. Pons etal. 1995). Non-magmatic shear zones in the latter study hadpressure shadows filled by epidote around megacrysts, andribbon grains formed by intracrystalline plasticity in quartz.Magmatic shear zones were identified in the experiments ofPark and Means (1996) by analyzing movements of solid in-clusions, but there was very little direct microstructural evid-ence to distinguish the shear zones from the unsheared walls.

Criteria which are sometimes misused to demonstrate mag-matic flow include imbrication, which may occur in mylon-ites (Section 7.11.3), magnetic anisotropy, which is wellknown in metamorphic rocks, and S-C fabrics, which mayform in either sub-magmatic or non-magmatic deformation(Blumenfeld and Bouchez 1988).

6.4 Magmatic microstructures

6.4.1 Grain shape fabrics

The typical microstructure of magmatic flow consists of apreferred orientation of euhedral phenocrysts, in an isotropicmatrix that shows igneous textures (Plate 36). Grain shapefabrics are commonly defined by feldspars and micas in felsicrocks, and may be defined by feldspar, olivine and pyroxenein mafic rocks (e.g. Benn and Allard 1989). This is diagnosticof magmatic flow if there is no sign of any other deforma-tion mechanism. Since intracrystalline plastic deformationmicrostructures are readily visible in quartz after only smallstrains, the lack of deformation in quartz (no undulose ex-tinction, subgrains, or kink bands) is a reliable criterion formagmatic flow. Lack of sub-solidus deformation can also bechecked from inclusions (e.g. rutile) in quartz: the inclusionsshould be undeformed themselves and randomly orientated(e.g. Mitra 1976, Stel 1991). Other primary, undeformedigneous features such as strongly euhedral crystals, igneous(e.g. idiomorphic) zoning, growth twins, and ophitic texturecan be used to demonstrate the absence of intracrystalline de-formation. A fabric in igneous rocks can only be confirmed asa magmatic microstructure by lack of any evidence for otherdeformation microstructures or mechanisms, including any ofthe microstructures described in previous Chapters.

It may be difficult to distinguish magmatic deformationfrom non-magmatic deformation followed by static recrystal-lization. Post-deformational annealing can be revealed by de-

formed inclusions in annealed minerals, and by a granoblastictexture in mineral aggregates (especially quartz) which them-selves define a shape fabric. Primary igneous fabrics can alsobe recognized by random spatial relationships between dif-ferent phases. By contrast, solid state deformation produceshigher frequencies of contacts between like grains becausenew grains preferentially nucleate on the same phase (e.g.Ashworth and McLellan 1985).

6.4.2 Crystallographic fabrics

Magmatic crystallographic fabrics are due to orientation ofeuhedral, inequant phenocrysts during flow. In mafic rocks,(010) crystal faces are parallel to the magmatic foliation inolivines, pyroxenes and feldspars (Benn and Allard 1989).The [001] direction of olivine and clinopyroxene is parallel tothe magmatic lineation, and [100] has been observed parallelto the magmatic lineation in feldspars within gabbros and ton-alites (Benn and Allard 1989). Such crystallographic fabricscan be distinguished from fabrics due to intracrystalline plas-ticity by complete lack of evidence for intracrystalline plasticmicrostructures (Chapter 4), and in the case of olivine, by thefact that high-temperature non-magmatic deformation leadsto [100] parallel to the lineation rather than [001], as observedfor magmatic fabrics (Benn and Allard 1989).

6.5 Sub-magmatic microstructures

6.5.1 Grain shape fabrics

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

6.5.2 Intracrystalline plasticity

Intracrystalline plasticity in quartz may be expected in thepresence of melt under appropriate differential stresses (>1MPa) and temperatures (700 to 800°C) on the basis of graniteand quartz flow laws (Rutter and Neumann 1995). The verycommon appearance of undulose extinction in quartz withingranites suggests that intracrystalline plasticity occurs duringsub-magmatic deformation (Plate 37). Subgrain formation orrecrystallization of quartz has been taken as an indicator ofsub-magmatic deformation where an overall igneous textureis preserved and no other non-magmatic deformation event isknown (e.g. Bouchez and Gleizes 1995). Deformation twinsand bent twins in feldspar may also hint at sub-magmaticintracrystalline plasticity (Plate 38). However, these micro-structures on their own do not demonstrate that melt waspresent during deformation.

CHAPTER 6. MAGMATIC AND SUB-MAGMATIC DEFORMATION 63

6.5.3 Diffusive mass transfer

Despite the experimental evidence for the importance of melt-aided diffusion creep, microstructural evidence has remainedelusive, probably because melt films have almost no poten-tial for preservation in the geological environment. However,experiments suggest some features that could be used: trun-cated, embayed, scalloped or overgrown grain boundaries inigneous rocks may be analogous to some of the features dis-cussed in Chapter 3 that indicate DMT through a fluid phase(e.g. Dell’Angelo et al. 1987, Park and Means 1996). Thepotential importance of surface energy anisotropy in melt dis-tribution needs to be investigated and may have some micro-scopic expression in the form of differences between crystalfaces.

6.5.4 Cataclasis

Paradoxically, some of the clearest sub-magmatic microstruc-tures involve cataclasis of the crystals, such as microfractureswhich are healed by melt (e.g. Hibbard 1987, Bouchez et al1992, Karlstrom et al. 1993). Microfractures in plagioclasecan be demonstrated to have been filled by melt from the fol-lowing criteria:

1.

2.

3.

4.

The microfractures are intragranular. This allows thatthe plagioclase crystals were in contact with melt.

The microfracture filling is compositionally and crystal-lographically continuous with the same phase in the ig-neous groundmass of the rock.

The composition of the microfracture filling is compat-ible with the later stages of the igneous petrographic his-tory of the rock. Plagioclase microfracture fillings mayhave lower anorthite contents than their host crystals,consistent with progressive evolution towards a min-imum melt composition (Bouchez et al. 1992). The rela-tion between quartz and plagioclase in the microfracturefillings also suggests a residual melt: feldspars are onthe walls or tips of the microfractures.

Early crystals (e.g. biotite, sphene) are trapped withinthe microfracture fillings.

Cataclastic microstructures observed in experiments con-firm the potential importance of cataclasis in sub-magmaticdeformation in the experiments of Rutter and Neumann(1995). Up to 10% melt, axially-orientated cracks formedand filled with melt, and the sample was faulted. Between10 and 45% melt, cataclastic flow occurred with pore col-lapse. Axially orientated microcracks formed by high meltpressures have been observed in granitic aggregates contain-ing 2-15% melt (Dell’Angelo and Tullis 1988). Connollyet al. (1997) have demonstrated that microcracking causedby volume increase during melting is a viable way to createpermeable fracture and melt-pool networks in a muscovite-bearing quartzite. The syn-kinematic experiments by Parkand Means (1996) also recorded fracture, localized along akink band boundary.

On a larger scale, several cataclastic features are commonlyassociated with melt in migmatites. Metatextites often consist

of a competent body sub-divided by fractures that are filled bymelt. Melt may form in pressure shadows at the ends of boud-ins, and fill faults (e.g. Quick et al. 1992). Quartz-feldsparneosomes have been described accumulated under imper-meable refractory layers such as amphibolite sheets which areboudinaged, allowing the neosomes to rise into boudin necks,and forming a geopetal structure (Burg 1991). Segregationsin shear zones and along axial planes of folds in migmatitesare common. Localization of the melt in these structures in-dicates that the melt was syntectonic. This sort of evidencehas great relevance to the problem of extracting melt to formplutonic 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 observedin their experiments, and suggested that indented boundariesand truncation of growth zoning might be indicative of theprocess in nature. Another suggestion from their experimentsis that filter pressing and expulsion of melt might be recog-nizable from layers characterized by intracrystalline plasti-city adjacent to layers of less-deformed or undeformed rockformed by crystallization of the expressed melt. Overgrowthsof feldspar along low-stress boundaries have been consideredas indicators of magmatic or sub-magmatic deformation (Hi-bbard 1987), but they may equally be be formed by sub-solidus DMT processes (Paterson et al. 1989).

6.7 Non-magmatic deformation

Non-magmatic microstructures reflect deformation withoutany melt present, which has also previously been known assub-solidus or solid state deformation. Non-magmatic is pre-ferred to “solid state” because it allows for the presence offluids other than melt. Two sort of evidence can be used tosuggest that non-magmatic deformation has continued as partof the same deformation event shown by magmatic or sub-magmatic features. The first sort is evidence for high temper-ature deformation, e.g. prism <c> slip in quartz (recognizedby quartz c-axes with an orientation close to the lineation e.g.Law et al. 1992, Lagarde et al. 1994), basal subgrain bound-aries (Section 4.6), or albite exsolution lamellae indicatingexsolution above the alkali feldspar solidus. A distinctivegrain boundary shape in quartz consisting of reticular grainboundaries, resulting in a mosaic pattern, indicates crystallo-graphic control of grain boundaries, extreme grain boundarymobility and therefore high temperatures (Gapais and Bar-barin 1986). However, a difficulty with this sort of evidenceis that the high temperature deformation may be a later eventwhich is entirely unrelated to the magmatic deformation (Pa-terson et al. 1989).

The second line of evidence is kinematic; if magmatic/sub-magmatic microstructures are kinematically compatible withnon-magmatic microstructures, they can plausibly be relatedto the same event. This type of evidence is probably morereliable. One example is non-magmatic S-C fabrics record-ing the same shear direction and sense as magmatic/sub-


magmatic flow (Blumenfeld and Bouchez 1988, Miller andPaterson 1994). Another case is the formation of quartz rods,folds and boudins in a non-magmatic state along an intrus-ive interface, while the bulk of the intrusion was still mag-matic, as suggested by the fact that all these structures dieout away from the interface (Stel 1991). The non-magmaticsimple shear zones described by Ramsay (1989) have a tan-gential extension direction around an intrusion: these can berelated to ballooning strains during successive phases of ex-pansion of the pluton.

Myrmekite is sometimes taken as evidence for non-magmatic deformation, but this is unreliable as myrmekitecan form by direct crystallization (Paterson et al. 1989). Hi-bbard (1987) has suggested that magmatic myrmekite can bedistinguished by growth in dilational sites (“pressure shad-ows”) around phenocrysts, in contrast with tectonically in-duced myrmekite which forms in volumes perpendicular tothe 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 withrespect to the directions of maximum and minimum stretch-ing rate, the instantaneous stretching axes, ISA (e.g. Hanmerand Passchier 1991). This rotation is also known as the shear-induced vorticity or internal vorticity (Means et al. 1980, Wil-liams et al. 1994). Features such as pressure shadows, por-phyroclast inclusion trails and fibrous vein fillings may pre-serve some information about the orientation of the ISA, andcan be used to determine internal vorticity. However, shearsense is commonly used in a much looser way to indicate thedisplacement sense of a fault or shear zone, usually specifiedby one or a combination of the usual terms for fault or shearzone movement i.e. normal, reverse, sinistral, dextral. In thiswider usage, shear sense is defined with respect to a shearplane, which is a useful external reference frame. Rotationrelative to this frame of reference is external vorticity, con-sisting of the sum of internal vorticity and any rotation of theISA with respect to the external frame of reference, or spin,which should be accounted for in a rigorous shear sense eval-uation.

Shear sense should only be analyzed in a section that isperpendicular to the shear plane and parallel to the shear dir-ection, which is known as the shear sense observation planeor the vorticity profile plane (Robin and Cruden 1994). Theshear direction is often assumed to be parallel to mineralstretching lineations or slickenlines, and the shear plane tofoliations or fault planes; hence the shear sense observationplane would be parallel to the lineation and perpendicularto the foliation (Fig. 7.1). Other sections can give incor-rect shear senses when using marker displacements as shearsense indicators. However, the assumption that mineral foli-ation and lineation are parallel to the shear plane and sheardirection breaks down in two circumstances: when the totalshear strain is low, and in complex types of three-dimensionalstrain. Transpression with a high ratio of pure shear to simpleshear is one example of the latter. In this case, the maximumfinite stretching axis (tracked by the mineral lineation) maycome to lie perpendicular to the shear direction (Fig. 7.2;Tikoff and Fossen 1993). This may explain some puzzlingcases where shear sense indicators are not seen in the planeparallel to the lineation and perpendicular to the foliation, butin planes perpendicular to the lineation (e g. the Larder Lake-Cadillac deformation zone, Canada; Robert 1989). The vor-ticity vector in these cases must be near to the maximum fi-

nite stretching axis, and the vorticity profile plane must beperpendicular to the lineation. The vorticity vector in gen-eral may lie at any angle to the maximum ISA, depending onthe amount of transpression, and this relationship can changeacross a single shear zone (Robin and Cruden 1994, Tikoffand Greene 1997). In these circumstances, the shear direc-tion can only be established reliably by using shear sense in-dicators other than displaced markers, observed in a variety ofsections. The shear direction is within a plane across whichthe shear sense appears to reverse.

These complications do not apply to any of the examplesused in this chapter, in which all illustrations are in the planeperpendicular to the foliation and parallel to the lineation,which is the correct shear sense observation plane in thesecases. A convention used in many studies is to take the shearplane as perpendicular to the plane of the illustration, and theshear direction as horizontal within the plane of the illustra-tion, so that the shear sense can be specified by dextral (clock-wise) or sinistral (anticlockwise). Unfortunately this conven-tion is not always explicitly recognized, sometimes creatingthe misleading impression of horizontal movement on dip-slip faults or shear zones. The convention is used throughoutthis book, so that dextral or sinistral can be used to describeshear sense. Microscopic observations such as shown in theplates in this chapter are often invaluable aids to shear sensedetermination, especially in fine-grained rocks. This is be-cause of the finer detail visible in thin sections, and becausethe sections can be cut accurately in the shear sense observa-tion plane, which may not be visible in outcrop.

Contradictory shear sense indicators are common in out-crop or thin section. There are good theoretical reasons forthis, including the fact that material lines rotate in oppos-ite directions under general shear (e.g. Passchier and Trouw1996). Therefore it is often necessary to evaluate a large num-ber of shear sense indicators for reliability. This may meanexamining many thin sections, but a useful technique can beto cut a number of hand specimens on a rock saw in the shearsense observation plane. The shear sense may become visibledue to the smooth surface of the saw cut and because it is thecorrect observation plane, and the procedure allows rapid ex-amination of a large number of specimens. Specimens mustbe orientated correctly in the field as discussed by Prior et al.(1987) or Passchier and Trouw (1996). Orientation errors canoccur in the thin section laboratory, but errors can be checkedif all rock fragments created during sectioning are carefullypreserved, so that they can be reconstructed to the pre-sawingconfiguration.

65

66 CHAPTER 7. MICROSTRUCTURAL SHEAR SENSE CRITERIA

7.2 Curved foliation

Many shear zones have a foliation that curves smoothlyacross the shear zone as shown in Fig. 7.1. The angle betweenthe foliation trace and the shear plane decreases from about45° or less to approach zero at the centre of the shear zone(Plate 39). The rotation of the foliation from the outside tothe centre of the shear zone (clockwise in Fig. 7.1) is the sameas the shear sense of the shear zone (dextral). This criterionis one of the most common and reliable for deducing shearsense, and can be applied to shear zones on any scale.

This sort of foliation follows the orientation of the XYplane of the finite strain ellipsoid, and the curvature is dueto increased intensity of finite strain towards the centre ofthe shear zone (Ramsay and Graham 1970). Such foliationsare referred to as strain-sensitive by Hanmer and Passchier(1991) and constitute one of the most common and reliableshear sense indicators. An important caveat to the use of suchfoliations is that the foliation should be formed during shear-

ing, as suggested by the observation that the rock has no fab-ric outside the shear zone. A pre-shear foliation may not ex-hibit this characteristic curvature, and is a less reliable shearsense indicator.

7.3 Oblique foliations and shape pre-ferred orientations

Oblique foliations are grain shape fabrics that maintain an ap-proximately constant, acute angle to the shear plane across ashear zone. The shear plane in this case is often represen-ted by a compositional banding, so that the fabric consists ofcompositional bands within which there is an oblique foli-ation. The shear sense is given by the acute rotation from theoblique foliation to the compositional banding (clockwise ordextral in Fig. 7.3). Oblique foliations and the bands paral-lel to the shear plane have been called and respectively(Law et al. 1984). The geometry of oblique fabrics has strongsimilarities with some S-C fabrics (see Section 7.5). The ob-lique foliations described above are examples of shape pre-ferred orientations, which can be used as shear sense indicat-ors by a different method (Shelley 1995). The aspect ratio ofthe grain long axes is measured and plotted against their anglewith respect to an arbitrary reference orientation. Frequencycurves can be plotted for the highest and the mean aspect ra-tios, in class intervals of 5-10°. The angular distribution ofaspect ratios is typically skewed, and the shear sense is givenby two criteria: the rotation sense from the longer part of thecurve towards the highest aspect ratio orientation, and fromthe mean orientation of all grains towards the orientation ofthe highest aspect ratio orientation.

Shape preferred orientations can result from passive be-haviour of grains that behave as strain markers (i.e. theyflatten and rotate towards the shear plane as described inSection 7.2), from anisotropic mineral growth (Chapter 5.3),or from rotation of grains behaving as rigid objects in afluid (e.g. Shelley 1989b). Development of shape preferredorientations may be counteracted by dynamic recrystalliza-

CHAPTER 7. MICROSTRUCTURAL SHEAR SENSE CRITERIA 67

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

7.4 Porphyroclast systems


A porphyroclast system comprises a porphyroclast with amantle and tails (or wings) of grains attached and derivedfrom the porphyroclast (Passchier and Simpson 1986). Themineralogy of the mantle and tails may be the same as the por-phyroclast (a “mantled porphyroclast”; Passchier and Trouw1996), or may be related to the porphyroclast by a meta-morphic reaction, such as a feldspar porphyroclast breakingdown to tails of quartz and mica.

Two tails are usually developed around a porphyroclast,symmetrically related to each other by a 180° rotation aboutan axis parallel to the shear plane and perpendicular tothe shear direction. Fig. 7.4 shows four categories of por-phyroclast system geometry, three of which are named fromtheir similarity to the Greek letters: and (Passchierand Simpson 1986, Passchier 1994). and havemonoclinic symmetry (Figs. 7.5-7.7), while haveorthorhombic symmetry. Complex types (Fig. 7.8) consistof more than one tail on each side of the porphyroclast(Passchier and Simpson 1986). The geometry of porphyro-


clast systems can be described with respect to a referenceplane that passes through the centre of the porphyroclast par-allel to the planar part of the tail far from the porphyroclast,and that contains a symmetry axis (Fig. 7.4; Passchier andSimpson 1986). An important distinction is made betweenin-plane tails that both lie on the reference plane, and stair-stepping tails that lie on opposite sides of the reference plane(Fig. 7.3). types have been subdivided into whichare found around isolated porphyroclasts in an homogeneousmatrix, and which are are found in S-C ormylonites, with short tails that curve into C- or(Passchier and Simpson 1986). tails were origin-ally defined by the criterion that an individual tail crossedthe reference plane near the porphyroclast, and therefore thetwo tails stair-stepped (Passchier and Simpson 1986). How-ever,the characteristic feature of a is commonly takento be a deflection near the porphyroclast which produces afeature known as an embayment (Fig. 7.9), so that systemsdo not necessarily stair step. There is a great deal of variationin the detailed morphology of porphyroclast systems, and allgradations exist between and (Fig. 7.10).

7.4.2 Mechanisms of formation

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

sidered to develop by dynamic recrystallization of the por-phyroclast to form a mantle of finer grains around the clast,which subsequently become entrained in the flow of the mat-rix. The geometry of tails depends on the rate of recrystalliz-ation relative to the rate of rotation of the porphyroclast, theshape of the porphyroclast, the rheology of the tails and mat-rix, the shear strain, and the degree to which the tails and mat-rix adhere to the porphyroclast (the coupling). Many of thesefactors can be simplified by analyzing the relation betweenthe recrystallizing mantle and the “separatrix”, which is asurface around the porphyroclast that separates closed fromopen flow lines. The separatrix is either eye-shaped or bow-tie shaped in simple shear, depending on a variety of factorsincluding the rheology of the matrix. The shape of the tailsis governed by the shape of the separatrix and the location ofthe separatrix relative to the mantle. A few generalizationscan be made from the experimental work of Ten Brink andPasschier (1995):

A mantle entirely within the separatrix will not develop tails.

A mantle that intersects the separatrix will first formtails which then develop into tails, thus explainingwhy there are transitional types between the two end-members.

A mantle that encloses the separatrix will first formtails, which develop into type tails for an eye-

shaped separatrix, but remain as tails for a bow-tie shaped separatrix.


tails will stair step in a bow-tie shaped separatrix, butnot in an eye-shaped one.

tails only develop with high coupling.

The evolution of tails is incompletely understood at present,but these approaches offer the possibility that analysis of tailswill yield considerable information on rheology and deform-ation conditions.

There are three simple guidelines for using porphyroclastsystems as shear sense indicators. It should be emphasizedthat shear sense can only be determined from tails with mono-clinic symmetry, i.e. and tails.

7.4.3 Stair-step direction: and tails

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

7.4.4 Faces of a tail

The asymmetry of an individual tail can be used todefine shear sense (Simpson and Schmid 1983). An indi-vidual tail has two faces in contact with the matrixaround the porphyroclast, one of which is approximately par-allel to the shear plane, and the other which is curved andoblique (Fig. 7.9a). Provided the face parallel to the shearplane can be identified, the curvature and obliquity of theother face can be used to give the shear sense: it is convexin the direction of movement of the adjacent matrix, and theacute rotation of the curved face to the shear plane is in thesame sense as an oblique foliation. It is usually possible tocheck this criterion by stair-step direction, but it is sometimesuseful if only one tail is well-developed (Fig. 7.5).

7.4.5 Deflection and embayments oftails

The deflection of a tail from the distal end towardsthe porphyroclast can be used to give the rotation sense ofthe porphyroclast (Passchier and Simpson 1986). The de-flection of the tail is in the opposite sense to the rotation ofthe porphyroclast. For example, the deflection of the tails inFig. 7.4 is to the left as the porphyroclast is approached, giv-ing a clockwise or dextral rotation and sense of shear. An em-bayment is the term for the approximately triangular regionof matrix between a tail and the adjacent porphyro-clast (Fig. 7.9b; Passchier and Simpson 1986). The interfacebetween the porphyroclast system and the matrix in the em-


bayment has the shape of a fold: the direction of fold clos-ure is the same direction as the rotation of the porphyroclast(Figs. 7.7, 7.9b).

Ideal conditions for the use of porphyroclast systems asshear sense indicators are that the grain size of the matrix issmall with respect to the porphyroclast, the fabric of the mat-rix is homogeneous, the porphyroclast systems were formedin one phase of deformation, the tails are long enough todefine a reference plane, and observations are made in theshear sense observation plane (Passchier and Simpson 1986).Porphyroclasts should be spaced widely enough not to inter-fere with one another. Structures similar to tails can developin a passive matrix around a rotating porphyroclast, but thesehave formed in a different manner from the porphyroclast sys-tems described above and should not be used in the same wayfor shear sense determination (Section 7.10).

7.5 S-, C- and


Many shear zones contain two or three planar fabrics thatcan be divided into a sigmoidal penetrative foliation (an S-foliation), and discrete planar shears that displace the fo-liation, which are known as C- or (Berthé etal. 1979), extensional crenulation cleavage or ecc (Plattand Vissers 1980), shear bands or shear planes (White etal. 1980, Simpson 1986), or normal-slip crenulations (Den-

nis and Secor 1987, 1990). Figure 7.11 shows the essen-tial geometry of these fabrics and the angular relationshipsbetween them, which are useful for description and classific-ation. S-surfaces intersect C- or in a line perpen-dicular to, or at a high angle to, the shear direction, as definedby the lineation on C- or surfaces. Angle is definedas that between the shear plane and the local orientation ofS-surfaces between C or and angle is thatbetween the shear plane and which are inclinedto 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- oris equal to between S- and C-surfaces, and

between S- and S-, C- and havea wide variation in morphology (Fig. 7.12, Plates 40, 41).Two major subdivisions of S-C mylonites were proposed byLister and Snoke (1984), but problems with this classifica-tion have been pointed out (Passchier and Trouw 1996). Analternative approach to description and classification (Blen-kinsop and Treloar 1985) uses:

1.

2.

3.

4.

The angles and

The structures that define the S-surfaces,

The spacing of the C- or and

The relative strengths of S- and C- or

This classification is descriptive, yet uses features that havepossible genetic significance. Three end members (porphyro-clastic, megacrystic and banded types) identified using the


scheme are shown in Fig. 7.13 and the relations betweenand 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 strainand microstructure, and can be used to distinguish betweenmodels for the formation of S-, C- and A com-mon feature of several models is that S-surfaces are parallelto the XY plane of the local finite strain ellipsoid (Ramsay1967, 1980, Berthé et al. 1979, Lister and Snoke 1984, Blen-kinsop and Treloar 1995). C- and are zones of

localized shear analogous to Y-shears and Riedel shears re-spectively, as observed in simple shear experiments and faultgouges (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 failuresurface. It is likely that anisotropy controls the formation andthe orientation of i.e. the value of (Platt andVissers 1980, White et al. 1980, Platt 1984, Passchier 1984).

If S-surfaces are parallel to the axes of the local finite strainellipsoid, then the value of should decrease with progress-ive strain in these domains. However, C- andprobably do not not rotate with increasing strain. This issuggested by the planar nature of C- and if sig-nificant rotations occurred in a material with discontinuousC- and strain incompatibilities would distort ini-tially planar surfaces (Shimamoto 1989). If decreases and

remains constant with progressive strain, then should alsodecrease, which has been reported in many studies (e.g. Zeeet 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. Dennisand Secor 1987, 1990). Volume changes and “stretching”shear zones that extend parallel to their length may be import-ant in the formation of some S-, C-, and (Behrmann1984, Passchier 1991).

Complications in the use of S-, C-, and as kin-ematic 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 seenparticularly clearly in thin section.


7.5.3 Curvature of S-foliation

Many S-foliations curve into like the curved foliation de-scribed in Section 7.2, and the rotation sense from the S-foliation to the C- or gives the sense of shear re-liably (Fig. 7.12, Plates 39, 40).

7.5.4 Shear on C- or

The sense of shear on C- or is always the sameas the overall sense of shear. It can sometimes be determinedindependently from the curvature of S-foliations by identify-ing features such as veins that are offset by shear on C- or

7.6 Pressure shadows and fringes

7.6.1 Kinematics of pressure shadows andfringes in shear zones

The formation and geometry of pressure shadows and fringeshas been discussed in Chapter 3.9, and here it needs to bere-emphasized that pressure shadows and fringes grow in thevolume of low mean stress around the inclusion, which ap-proximately coincides with the direction of the ISA. There-fore in the simplest case of undeformed pressure shadows,the pressure shadow or fringe will be oblique to the shearplane, and the shear sense can be deduced from the acuterotation from the pressure shadow/fringe to the shear plane(Fig. 7.15). Some pressure shadows/fringes have a stair-step

geometry that can be used in the same way as porphyroclasts(Section 7.4.3). It may be difficult to distinguish pressureshadows from porphyroclasts (Plate 44). Rotation of the por-phyroclast relative to the ISA and deformation of the pressureshadow or fringe complicate this simple geometry. Despitethe large variety of intricate structures that may be produced(e.g. Etchecopar and Malavieille 1987, Aerden 1996) thereare two reliable methods of determining shear sense frommore complex pressure shadows and fringes.

7.6.2 Geometry of the last increment of growth

The last increment of growth may be virtually undeformedand can therefore indicate the direction of the last maximumISA. In the case of antitaxial growth, the last increment willbe closest to the inclusion, while the opposite will be true forsyntaxial growth. The sense of rotation from the last max-imum ISA to the shear plane give the shear sense. Fibre ori-entation in pressure fringes can be used to indicate the ISAonly if the fibres are displacement-controlled.

7.6.3 Shape

Fortunately a simple empirical rule for shear sense determ-ination is clear from computer simulations and natural ex-amples, even without detailed considerations of the mechan-ism of pressure shadow or fringe formation. The shape of thewhole pressure shadow or fringe, defined either by its medianline or by its enveloping surfaces, forms a doubly-inflectedcurve which is either an S or a Z shape. S shapes imply dex-tral rotation of the core objects and therefore dextral shear


sense, and vice versa (Figs. 7.16, 7.17). One complicationthat may occur in the use of pressure shadows or fringes asshear sense indicators is that the shadow/fringe may be ro-tated with the inclusion (White and Wilson 1978). Pressureshadows formed in the extension field may be rotated into theshortening field and therefore give the incorrect shear sense.Longer pressure shadows can come to resemble por-phyroclast tails by this process (Hanmer and Passchier 1991).

7.7 Mica fish

Mica fish are a distinctive type of porphyroclast consistingof elongated single crystals of mica, often associated withtails (e.g. Lister and Snoke 1984, Hanmer 1986, Passchierand Trouw 1996). Stair-stepping of the tails can be used as ashear sense indicator following Section 7.4, and the sense ofthe acute rotation from the long axis of the fish to the shearplane also gives the shear sense directly (Plate 42). The form-ation of mica fish is not well understood, but may involvedevelopment of tails by dynamic recrystallization as well ascataclasis, and rotation of the fish. Other minerals (e.g. horn-


blende, Plate 43) may show the same geometrical features asmica fish.

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 porphyroblastswith respect to the matrix foliation and the flow axes dur-ing non-coaxial flow (Fig. 7.18a). However, Bell and cowork-ers (e.g. Bell 1985, Bell and Johnson 1989, Bell et al. 1992,Aerden 1995) have argued that many form by porphyro-blasts overgrowing a new foliation that formed in a differ-ent orientation relative to an external frame of reference, andthat the porphyroblasts themselves do not have an externalvorticity. The latter interpretation leads to the opposite senseof shear to the former (Fig. 7.18). A third possibility for inter-preting curved is that a pre-existing foliation rotates duringcoaxial flow (Fig. 7.18b; Ramsay 1962). A fourth possibilityfor generating curved inclusions with continuous and isby post-tectonic overgrowth of folds (helicitic structure). Thelast case can be distinguished by the presence of folds witha similar geometry to the inclusion trail in the matrix awayfrom the porphyroblast.

The above ambiguities are sufficient to preclude the useof porphyroblast inclusion trails as a shear sense indicatorunless the mechanism by which they formed is well known.It is important to know the regional kinematic framework to

understand the porphyroblast growth mechanism, especiallywhether the porphyroblast is in a domain of coaxial or non-coaxial flow.


Crystallographic fabrics are usually represented by lowerhemisphere, equal area stereographic projections of Crystal-lographic directions of individual grains or sub-grains. It isstandard practice to use the shear sense observation planeas the projection plane, with the foliation as a vertical east-west plane, and the lineation as a horizontal east-west line.This section focuses on the use of quartz c-axis Crystallo-graphic fabrics, which are readily obtained from the universalstage following the procedures in Turner and Weiss (1963) orPasschier and Trouw (1996).

Quartz c-axes in non-coaxial shear are known to have anasymmetrical distribution relative to foliation and lineationfrom experiments, numerical models and natural examples.The asymmetry has been demonstrated as a reliable shearsense criterion (e.g. Behrmann and Platt 1982, Bouchez etal. 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 usedfor shear sense determination when deformation is the res-ult of intracrystalline plasticity, the foliation and lineation areclearly defined and related to the finite strain, the deformationis homogeneous, and the mineral used for the determination is


dominant in volume (Bouchez et al. 1983). Crystallographicfabrics should always be used in conjunction with other shearsense indicators.

The crystallographic fabric asymmetry is best defined bythe fabric skeleton, which are the lines that connect the max-imum concentrations of c-axes (Behrmann and Platt 1982,Law 1987). Two types of asymmetry are distinguished: ex-ternal asymmetry is defined by the obliquity of the centralgirdle to the foliation (angle in Fig. 7.19; angle nomen-clature after Law 1987), and by the difference between theangles and (Fig. 7.19). Internal asymmetry is definedby the angles and between the arms of the crossedgirdle and the central girdle (Fig. 7.19). The most importantshear sense criterion is the obliquity of the central girdle tothe foliation (Fig. 7.19). The shear sense is given by the dir-ection of the acute rotation from the girdle to the shear plane.The difference between the values of and or between

and can also be used (Fig. 7.19, cf. Law 1987, 1990,Law et al. 1994). At high temperatures where prism <c> slipoperates, a completely different type of fabric forms, consist-ing of a point maximum of c-axes near the lineation. Theacute rotation from the point maximum of the c-axes to thelineation is in the opposite direction to the shear sense (Main-price et al. 1986).

Quartz a-axes can also be used for shear sense determin-ation, but they require X-ray measurement. Calcite c-axesdevelop a girdle in non-coaxial shear that is inclined to thefoliation in the opposite direction to the quartz c-axis girdle(e.g. Schmid et al. 1987, Wenk et al. 1987). Asymmet-ric girdles in calcite determined by X-ray diffractionhave proven to be reliable shear sense indicators (Erskine etal. 1993). Plagioclase, olivine and orthopyroxene asymmetricfabrics have also been used to determine shear sense (Mercier1985, Gapais and Brun 1981, Passchier and Trouw 1996).

7.10 Asymmetric microboudins

Microboudins can be observed in thin section and share thesame characteristics as the mesoscale examples described inoutcrop or experiments. Boudins in non-coaxial deformationare asymmetric, due either to modification of initially sym-metrical boudins (types I and II in Fig. 7.20; Hanmer 1986),or to development of asymmetric boudins ab initio (type IIIin Fig. 7.19, Goldstein 1988). The three types of asymmetricboudin can be used as shear sense criteria according to thefollowing rules: Type 1 asymmetric boudins have two facesparallel to the shear plane, and two inclined to the shear planesuch that an acute rotation from the inclined faces to the shearplane gives the shear sense. Types II and III boudins are sep-arated on shears that are inclined to the shear plane, and dis-place the boudins in the same sense as the overall shear zone(analogous to and Riedel shears, Section 7.5, Plate 6).

However, the geometry of boudins in non-coaxial shear de-pends on the initial shape and orientation of the boudinagedlayer (Goldstein 1988). Boudins can be separated by shears


antithetic to the dominant sense of shear, even when layerslie in the extensional field (Fig. 7.21). “Bookshelf sliding”also involves antithetic shearing. Therefore microboudins canonly be reliably used for shear sense determination if theirinitial shape and orientation are known: this may be possiblefrom observations outside shear zones.

7.11 Asymmetric microfolds androlling structures

The vergence of asymmetric folds in shear zones is often thesame as the dominant shear sense (e.g. Fossen and Hoist1995), but fold vergence opposed to the shear sense is alsowell documented (e.g. Krabbendam and Leslie 1996). Op-positely verging folds can be generated by shear of bucklefolds in a layer inclined in the shortening direction of non-coaxial flow (e.g. Ramsay et al. 1983, Little et al. 1994) orby heterogeneous simple shear (Fig. 7.22; Krabbendam andLeslie 1996). Folds in simple shear zones may be highly non-cylindrical, and this also complicates the use of fold vergencefor shear sense determination. Asymmetric microfolds, likeasymmetric microboudins, should not be used for shear sensedetermination unless the way in which they formed is known.Reliability can be enhanced when the vergence of several lay-ers in different orientations can be observed, or when the ini-tial orientation of the folded layer can be determined, for ex-ample by observations outside the shear zone.

A special type of asymmetric fold is developed by the in-teraction between rigid inclusions and layering. This type offold is localized around the inclusion, and has the same ver-gence as the sense of shear. Such structures developed at highshear strains by interaction between porphyroclast and matrixare referred to as rolling structures (Van Den Driessche andBrun 1987).

7.12 Shear sense criteria in rocks con-taining melt

7.12.1 Magmatic shear zones

Shear planes and shear directions that existed during deform-ation of melt-bearing rocks may be difficult to identify be-cause they leave no microstructural imprint (e.g. Park andMeans 1996). Compositional layering may function as ashear plane because of its rheological anisotropy (Benn andAllard 1989). It has been suggested that magmatic flowplanes and directions may correspond to the average ori-entation of planar and linear shape fabrics respectively (e.g.Guineberteau et al. 1987). However, magmatic foliationdefined by megacrysts is not generally parallel to the shearplane (e.g. Paterson and Vernon 1995, Yoshinobu and Pa-terson 1996) because megacrysts rotate at variable rates de-pending on their aspect ratios and the rheology of the matrix(e.g. Tikoff and Teyssier 1995). Indeed the first shear sensecriterion suggested below is based on the obliquity betweengrain shape fabrics and the shear plane. Possible dangers ofthe interpretation of magmatic or sub-magmatic foliations as


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

7.12.2 Oblique grain shape fabrics

Grain shape fabrics of olivine, pyroxene and feldspar in gab-bros have been used to determine sense of shear (Benn andAllard 1989) on the basis that they form a fabric at an acuteangle to the shear plane. The shear sense is given by the acuterotation from the fabric to the shear plane. However, sinceelongate grains may rotate past the shear plane, this criterionneeds to be used with caution, and a large number of obser-vations should be collected.

7.12.3 Tiling and imbrication

Tiled and imbricated clasts have been proposed as shearsense indicators in rocks containing melt (e.g. Blumenfeld1983, Paterson 1989), and as diagnostic of melt-present, non-coaxial deformation. Imbrication has been comprehensivelymodelled by Tikoff and Teyssier (1994), who suggested threepossibilities corresponding to magmatic, sub-magmatic andsolid state deformation respectively:

1.

2.

Rigid clasts in a fluid, which form imbricated trains ofclasts that rotate as a unit when they have coalesced(“Jeffrey rotating train” model, after Jeffrey’s (1922)analysis of rigid objects in a fluid).

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

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

In all cases, the imbrication of clast trains gives the shearsense according to the rule that the acute rotation from thetrain long axes to the shear plane is the same as the shearsense. Individual clast long axes can not be used reliably asshear sense indicators because they may rotate past the shearplane in the first two models, even in simple shear. The resultsof numerical modelling show that clast interaction dependsstrongly on clast density, and that the March models (2 and3), possibly corresponding to sub-magmatic and solid statedeformation, are more effective at imbricating clasts. There-fore clast imbrication is not diagnostic of melt-present de-formation.

7.12.4 S-C fabrics

S-C fabrics have been described in rocks that have no evid-ence of solid state flow (Blumenfeld and Bouchez 1988, Bennand Allard 1989), and these fabrics can be used as a shearsense indicator as described in Section 7.5. Euhedral, littledeformed plagioclase grains in S- and C-orientations havebeen described by Miller and Paterson (1994), suggesting thatthe S-C fabric formed in a sub-magmatic state.

7.12.5 Sub-magmatic microfractures

Sub-magmatic microfractures can be used as a shear sense in-dicator. They are inclined to the shear plane in the same ori-entation as T fractures (Fig. 2.17); the acute rotation from thefractures to the shear plane opposes the shear sense (Bouchezet al. 1992).


7.13 Shear sense criteria for faults

7.13.1 Shear sense observations on faults

Determining the shear sense (sensu lato) of faults requiresobservations in the shear sense observation plane which canonly be identified if the net slip vector of the fault is known.In practice this almost always requires the presence of slick-enlines to identify the shear direction within the fault plane.Incohesive fault rocks are particularly difficult to sample andsection. It may be possible to collect intact and orientatedspecimens by removing them in a metal casing, and to pre-pare them for thin sectioning by impregnation with adhesive.

7.13.2 Displaced grain fragments

Small displacements on faults can be measured by matchingfragments of distinctive grains in the cataclastic fault matrixto their parents in the walls of the fault, or by matching indi-vidual grains on either side of the fault (Fig. 7.23, Plate 45).The sensitive tint plate is useful for this task. However, thesense of shear determined from fracturing of porphyroclastsor boudins is ambiguous, because both synthetic and anti-thetic fractures can form (Section 7.10).

7.13.3 Risers and slickenfibres

Risers (Section 2.9) on fault surfaces have been used histor-ically to determine shear sense, making the assumption thatthey are congruous, but the occurrence of incongruous stepsis well known. A variety of other fault surface features have

been suggested as shear sense indicators by observations onfaults with known shear senses in specific conditions (e.g.Petit 1987, Doblas et al. 1997). It is not clear how widespreadthese features are, and many of them could have ambiguousinterpretations on faults with unknown shear sense. Until fur-ther work is reported, the only generally reliable fault surfacekinematic indicator at present are risers created by slicken-fibre terminations (or accretion steps), which are always con-gruous (Fig. 7.23). When using this criterion, it is useful toexamine fault surfaces in more detail under a binocular mi-croscope such as those in common use for micropalaeonto-logy, or under the SEM. This technique is particularly goodfor dealing with incohesive specimens.

7.13.4 Gouges

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

7.13.5 Jogs and bends

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

Chapter 8

Shock-induced microstructures and shockmetamorphism

8.1 Introduction

A number of distinctive natural microstructures have recentlybeen accepted as characteristic of shock metamorphism onthe basis of field observations, detailed microstructural stud-ies, and comparisons with impact experiments. The productsand processes of shock deformation are distinct from othergeological microstructures and mechanisms, which justifiesa separate Chapter for their description. Shock effects havewell-established links with meteorite impact structures onearth, which form under pressures ranging from <5 to > 100GPa, temperatures up to 10 000 °C, and typical strain rates of

to (Fig. 8.1; Stöffler and Langenhorst 1994, Reimold1995).

These pressures and strain rates are far greater than valuesfor tectonic crustal deformation (using the term tectonic toimply endogenous processes). There has been vigorous de-bate about which microstructures are diagnostic of meteoriteimpact: this is discussed in Section 8.11.

8.2 Shock mechanisms

Four fundamental shock mechanisms have been identifiedfrom experimental work that simulates shock conditions:cataclasis, intracrystalline plasticity, solid state phase trans-formation, and melting (e.g. Huffman et al. 1993, Stöffler andLangenhorst 1994). Cataclasis results in irregular, randomlyorientated microfractures, and sub-parallel sets of planar mi-crofractures (Section 8.3). Intracrystalline plasticity is shownby twinning and, possibly, shock mosaicism (Section 8.4), butthe formation of the latter is not well-understood.

The most distinctive shock-induced microstructures andmechanisms are associated with phase transformation at highcompressive stress and with pressure-release melting. Quartzundergoes a phase transformation to an amorphous state inwhich Si coordination with O changes from four-fold to six-fold under pressures of approximately 5-35 GPa (Stöffler andLangenhorst 1994). The phase transformation is heterogen-eous at relatively low temperatures and high strain rates, andis localized on crystallographic planes leading to the form-ation of one type of planar deformation feature (PDF, Sec-tion 8.4). At higher temperatures and strain rates, the amorph-ization is homogeneous. Quenching following melting forms

a distinctive type of glass called diaplectic glass at moder-ate pressures (25-50 GPa; 8.6), and a fused glass called le-chatelierite at higher pressures (Section 8.8). Coesite andstishovite, the high pressure polymorphs of silica, may beformed during shock metamorphism by solid state transform-ation or direct crystallization from melt (Section 8.7). Theseshock microstructures can be classified into high and lowpressure types (Table 8.1).

8.3 Microfractures

Non-planar, unorientated microfractures are common inshocked materials. Planar, open microfractures (termedplanar fractures) parallel to low index crystallographic planesin quartz are a distinctive shock microstructure. They oc-cur parallel to (0001) and planes, with a spacing of

(Stöffler and Langenhorst 1994). In experiment-ally shocked dunite, intragranular microfractures formed insets of variable orientation with separations of less than

and more planar fractures with a spacing offormed in subregions of single crystals (Reimold and Stöffler1978). Their orientation is crystallographically controlledand a function of the direction of shock wave propagation,and their density increases linearly with shock pressure in a

80

CHAPTER 8. SHOCK-INDUCED MICROSTRUCTURES AND SHOCK METAMORPHISM 81

given rock type. PDFs (next Section) do not cross the mi-crofractures, which has been interpreted to indicate that thefractures formed earlier than the PDFs during the shock event(Stöffler and Langenhorst 1994).

8.4 Planar Deformation Features(PDFs)

Planar Deformation Features are intracrystalline zones of op-tical and/or crystallographic contrast with the host crystal,typically on the order of micrometres wide and with spacingsof 2 to as seen under the optical microscope (Fig. 8.2,Plate 46). Quartz is the mineral in which shock characteristicdeformation is best developed and has been most extensivelystudied. PDFs in quartz occur in sets parallel to the crys-tallographic directions given in Table 8.2 Other orientationsinclude and As many as18 different sets may form in one grain. The orientations ofPDFs in quartz can be related to the shock pressure as dis-cussed in Section 8.10. Planar features with the dimensionsof PDFs in shocked quartz are spectacularly revealed in SEM-CL images (Seyedolali et al. 1997), though it is not known yetcertain that they correspond to PDFs. A fundamental distinc-tion can be made under the optical microscope between dec-orated 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 moredetail. About ten times more PDFs are visible in the TEMthan under the optical microscope (Goltrant et al. 1991),

and PDFs have spacings from 1 to and thicknessesof under the TEM. Four different types of PDFs canbe distinguished in the TEM (Goltrant et al. 1991):

1.

2.

3.

4.

Brazil twins in the basal plane.

Bands of high dislocation density, with concentrationsof voids or bubbles in rhombohedral planes.

Bands of variable proportions of glass and crystallinequartz in the form of microcrystallites ~ 10 nm in sizein rhombohedral planes.

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 onrhombohedral planes, and their association with dislocationsshows that they are mechanical twins (Leroux et al. 1994).Types 2 and 3 are thought to be secondary features dueto post-shock annealing of type 4 PDFs, which are recog-nized, together with basal Brazil twins, as characteristic offresh shock microstructures, and are the only type of lamellaefound in experimentally shocked material. The glass lamel-lae can be subdivided into narrow transformation lamellae (<10 nm wide) which are only visible in the TEM, and widetransformation lamellae (50-500 nm wide) which are the op-tically visible lamellae, and which may contain a finer scalesub-lamellar structure of pillars at high angles to the lamellaeboundaries, which is clearly revealed by etching (Gratz et al.1996).

82 CHAPTER 8. SHOCK-INDUCED MICROSTRUCTURES AND SHOCK METAMORPHISM

A number of models have been put forward for PDF form-ation in quartz, which have largely been consolidated in thework of Huffman et al. (1993), Langenhorst (1994), andHuffman and Reimold (1996). Huffman et al. (1993) pro-posed that PDFs formed by heterogeneous solid-state trans-formation to amorphous silica along planes of progressivelylower compressibility, from basal to m, z, and a ori-entations, with increasing pressure. Optically visible PDFsare coalesced TEM-scale PDFs in domains between micro-faults. Langenhorst (1994) considered the temperature pro-files associated with shock fronts to show that, at relativelylow pressures (<25 GPa), melts were unlikely to form andPDFs should consist of short-range order (quasi-diaplectic)silica glass formed by solid-state transformation. At interme-diate pressures (25-35 GPa), thin bands of melt may form inaddition to the solid state transformation, and at 35-50 GPa,large-scale melting during compression may form diaplecticglass. At >50 GPa, melting continues after shock compres-sion to form lechatelierite by quenching under low pressure.This model successfully accounts for the presence of bothamorphous silica in PDFs at temperatures too low for melting,and diaplectic glass quenched from a melt. Fresh, impact-induced PDFs are either Brazil twins or amorphous silicalamellae formed by solid state transformation or quenchingof melt.

Post-shock annealing may have several effects on PDFs.The bands of high dislocation density in quartz (type 2, dec-orated PDFs) are likely to be due to post-shock annealing(Fig. 8.2). Annealing may be recognized by a distinct typeof fluid inclusion, which is typically much smaller (tens ofnm) than inclusions trapped during crystal growth (Goltrantet al. 1991), although annealed PDFs may also contain in-clusions up to diameter (Leroux et al. 1995, Lerouxand Doukham 1996). Fluid inclusions may form by diffusionalong dislocations during recrystallization of PDFs, becausewater is much more soluble in amorphous silica than in crys-talline quartz. The tiny crystallites in the interior regions ofPDFs (type 3 of Goltrant et al. 1991) are a characteristic an-nealing texture (e.g. Langenhorst and Clymer 1996).

Zircon, like quartz, is a favorable mineral to investigateshock effects because of its transparency, lack of importanttwinning or cleavage, and uniaxial optical character. Fur-thermore, it is a very refractory mineral which is resistantto alteration and thermal overprint. Shock features in zirconcan be seen after etching in the SEM (Bohor et al. 1993).Planar features continuous across the whole grain occur witha spacing of or less (Fig. 8.4, Kamo et al. 1996). It

is uncertain whether these features are analogous to PDFsin quartz: recent TEM analyses of experimentally shockedzircon single crystals suggests that these features are micro-cleavage, not PDF-type lamellar defects (Leroux et al. 1998).Zircon grains may also have a granular, polycrystalline tex-ture of zircon crystals (“strawberry texture”), which hasbeen observed in zircon from laboratory experiments and atseveral confirmed impact sites, including the Vredefort dome,but never in other circumstances: this texture may thereforebe diagnostic of impact (Figs. 8.5, 8.6).

8.5 Mosaicism

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


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 shockevents: therefore a direct comparison with subgrain forma-tion by recovery may not be valid (Stöffler and Langenhorst1994). Remnants of planar features associated with mosa-icism in quartz suggest that the mosaic domains may haveformed by the intersection of two sets of PDFs (Leroux andDoukham 1996). Mosaicism has also been described fromother minerals, for example feldspars, olivine and pyroxene(Hörz and Quaide 1973). A patchy optical extinction patternsimilar to mosaicism in feldspar can be formed by growthzoning (Sharpton and Schuraytz 1989).

8.6 Diaplectic glass

Diaplectic glass is a distinctive type of silica glass that oc-curs in homogeneous patches and associated with some typesof PDF. It can be distinguished from either synthetic quartzglass or lechatelierite (Section 8.8) by optical properties and

density (Table 8.3). Langenhorst’s (1994) model for the form-ation of diaplectic glass suggests that it forms by quenchingof a melt before shock pressure is completely released. Itsdistinctive properties are due to some long range order whichis inherited from the solid state. Diaplectic glass in naturalimpact rocks recrystallizes during annealing to characteristictextures which depend on the shock pressure reached. Forshock pressures less than 35 GPa, the glass becomes brown.Microstructures of spherical single crystals, all inthe same orientation, called ballen, form at higher pressures.At still higher pressures, the ballen have different crystallo-graphic orientations (Grieve et al. 1996, Plates 47, 48).

8.7 High pressure polymorphs ofquartz - Coesite and stishovite

Coesite in some shocked samples appears as colourless-brown, sized aggregates of individual grains lessthan in size, typically twinned parallel to (010), oftenenclosed by diaplectic glass or isotropic quartz with abundantremnants of PDFs (Grieve et al. 1996). The aggregates maybe aligned on crystallographic planes. Coesite associatedwith pseudotachylite veins from the Vredefort Dome is col-ourless, with high relief, and very low birefringence (Fig. 8.7;Martini 1991). It is either massive or forms radiating needlesup to long. Coesite has been observed at impactcraters and artificial explosion craters, but never in shock ex-periments: this is attributed to the much shorter duration ofthe pressure pulse in experiments compared to natural events(Stöffler and Langenhorst 1994). Stishovite, the higher pres-sure polymorph of silica, can be distinguished from coesite bya 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 ag-gregates parallel to PDFs in experimentally shocked quartz(Ashworth and Schneider 1985). The difficulty of produ-cing stishovite in shock experiments suggests that relativelylong shock durations may be required, and that it may crys-tallize from a high-pressure melt (Stöffler and Langenhorst1994). The preservation of stishovite appears to require rapidquenching because it is transformed to amorphous quartz atrelatively low temperatures.

8.8 Lechatelierite

Lechatelierite is a silica glass which may contain flow struc-tures and vesicles, and may occur as veins within quartz withPDFs. The optical properties of lechatelierite are summarizedin Table 8.3. The vesicles and flow structures demonstratethat lechatelierite forms as a quench product from a melt un-der low pressure, in distinction from diaplectic glass. Meltingto form lechatelierite requires pressures in excess of those forwhole rock melting, and therefore it is only preserved in veinswhere pressures were locally increased by shock reverbera-tion, or as inclusions in whole rock melts or tectites (Stöfflerand Langenhorst 1994).


8.9 Tectites, microtectites and spher-ules

Tectites are rounded silicate glass bodies usually less than afew centimetres in diameter (Glass 1990). They are usuallyblack, but may be translucent, brown or green. Three majortypes are recognized:

1.

2.

3.

Splash forms: spheres, disks, dumbell and teardropshapes.

Ablated forms, which are similar to splash forms butwith an additional flange.

Muong Nong types, which are larger, layered chunks oftectite glass.

Tectite glass has flow structures and abundant inclusions oflechatelierite. Baddeleyite, coesite, metallic spherules, andcorundum may occur as inclusions, as well as quartz, zir-con, rutile, chromite and monazite in the Muong Nong types(Glass 1990). The chemistry of tectites is similar to felsic vol-canic glasses, with contents of >65%, but with higher

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

Tectites occur over large areas of the earth’s surface calledstrewn fields, demonstrating that they were deposited fromthe atmosphere. Terrestrial and lunar volcanism, and met-eorite impact, have been proposed as possible origins fortectites, but a number of arguments favour the impact hypo-thesis (Glass 1990). Flow structures demonstrate that tectiteswere molten: the flanges of ablated forms were formed dur-ing a second period of melting. Lechatelierite inclusions sug-gest temperatures of 2000°C, far greater than volcanic tem-peratures. The chemical composition of tectites shows thatthey were derived by melting of sediments and not from ter-restrial or lunar igneous rocks. Furthermore, the tectites intwo strewn fields can be linked to proven impact craters: forexample, tectites in Czechoslovakia came from the Ries craterin Germany, while the Ivory Coast strewn field is derivedfrom the Bosumtwi crater in Ghana.

Microscopic tectites (microtectites) have been describedfrom several sites at the Cretaceous-Tertiary boundary, andassociated with impact at the Chicxulub site in Mexico (e.g.Smit et al. 1992, Olsson et al. 1997). Many of these mi-


crotectites are spherical, with diameters of 0.2 to 5 mm, andcomposed of clays or calcite. Some of these spherules containa core of glass, suggesting that the clays formed by devitrific-ation. Flow structures in the glass, and associated shockedgrains, demonstrate the impact origin of these spherule beds.

8.10 Shock barometry and thermo-metry

Establishing shock pressures has fundamental implicationsfor establishing the origin of shock microstructures and isuseful to estimate the rate of shock pressure attenuation formodelling impact processes. Shock pressures have beencalibrated experimentally using microstructures, the densityof quartz, and the optical, X-ray diffraction, spectroscopicand thermoluminescent characteristics of quartz (Stöffler andLangenhorst 1994, Grieve et al. 1996). Shock temperaturescan be calculated in non-porous rocks from shock pressures(e.g. Langenhorst 1994). In the following Sections, twomethods of shock wave barometry using optical microscopy(microstructures and optical properties of quartz) are brieflydescribed.

8.11 Calibration of shock pressuresfrom microstructures

Microstructures form with increasing shock pressure in asequence that has been clearly established by experimentalwork. In quartz, from the lowest pressure, fracturing evolves

to mosaicism and PDFs, followed by diaplectic glass, coes-ite and stishovite, and lechatelierite at the highest pressure.However, the accurate pressure calibration of this general se-quence is subject to the effects of a number of other variables(8. 10.3) which are not yet quantified with the exception of theinitial temperature of the target rocks and porosity.

Figure 8.9 shows onset pressures for shock microstructuresin quartz and feldspar based on experiments on single crystalsof quartz or non-porous crystalline rocks. The lines showingdecrease in onset pressure with temperature were determinedby experiments on granite and quartzite (Reimold and Hörz1986, Huffman et al. 1993, Huffman and Reimold 1996).Microfracturing occurs at the lowest pressures: planar micro-fractures form at pressures greater than 5 to 10 GPa (Huffmanet al. 1993). Mosaicism in quartz is first seen from pressuresof 6 to 10 GPa. The onset pressures for PDFs in these experi-ments fall between 15 and 18 GPa, but this is well above fig-ures given by other workers, who suggest formation of non-basal PDFs (i.e. not Brazil twins) at pressures as low as 10GPa. These lower pressures are shown on Fig. 8.2 by labelsgiving the orientations of the dominant PDFs (after Stöfflerand Langenhorst 1994, Grieve et al. 1996). The nature andorientations of PDFs in quartz changes as a function of pres-sure. Brazil twins form at the lowest pressures, followed byPDFs. PDFs are generally considered to form at pressuresof ~20 GPa, and become the dominant orientation at ~25GPa (e.g. Langenhorst and Deutsch 1994). Widths of experi-mentally created PDFs at the TEM scale change in thicknessfrom <100 nm at <25GPa to ~200 nm at >25 GPa. Ini-tial (i.e. target) temperature has an effect on PDF develop-ment. Approximately equal numbers of grains show singleand multiple sets of PDFs in experiments at 18 GPa and 250C

86 CHAPTER 8. SHOCK-INDUCED MICROSTRUCTURES AND SHOCK METAMORPH1SM


(Huffman and Reimold 1996). The number of grains withPDFs decreases as temperature increases, and multiple setsare preferentially eliminated; at 750°C, there are very fewgrains with single sets of PDFs in samples shocked at thispressure. The shape of the PDFs also changes from narrowwith sharp boundaries at low temperatures to broad with wavyboundaries at higher temperatures. There is some evidencefor a change from dominantly to orientations as temper-ature is increased from 25°C to 440°C (Huffman et al. 1993).

Quartz phase may affect the character of PDFs: samplesshocked at 20-25 GPa and 540°C showed dominantlyPDFs, compared to a broad range of PDF orientation at630°C, corresponding to the change from to sta-bility in the target rocks (Langenhorst et al. 1992, Langen-horst and Deutsch 1994). Experimentally shocked and

are also distinguished by the observation that positiveand negative planes of the same form never coexist in samplesshocked at less than 573°C, the transition tem-perature, as expected from the trigonal symmetry ofThis is in distinction from PDFs parallel to the hexagonal pyr-amids of samples shocked in the stability field, andhas been used as a geothermometer for the pre-shock tem-perature of the target rocks: for example the temperature ofthe Ries crater and Eocene-Oligocene impact targets can beconstrained to less than 573°C (Engelhardt and Bertsch 1969,Langenhorst and Clymer 1996).

A pressure calibration for porous rocks (Table 8.4) hasbeen established from work on the Barringer Crater in Ari-zona (Kieffer 1975, 1981, Kieffer et al. 1976). In general,the onset pressure for any microstructure decreases with in-creasing porosity, because stress concentrations can occur onthe grain scale. PDFs seem to be less abundant in poroustargets than crystalline ones, and there is some evidence thatdifferent orientations occur in porous rocks:and may be important (e.g. French et al. 1974). Asequence of microstructures similar to those in quartz andfeldspar with increasing shock pressures has been establishedin olivine and pyroxene. Planar fractures and mosaicism de-velop in olivine at 10 to 15 GPa, intergranular brecciationat 30 GPa, and glass, annealing and melting occur above 50GPa. In pyroxenes, twinning occurs at 5-10 GPa, mosaicismand PDFs at 20-25 GPa, and glass, melting and recrystalliza-tion above 75 GPa (Huffman and Reimold 1996).

Details of the techniques for measuring PDF orientationsare described in Engelhardt and Bertsch (1969), Stöffler andLangenhorst (1994) and Grieve et al. (1996). There are vari-ous methods of assigning an average shock pressure to ori-entation measurements from a sample, which should consistof as many grains as possible. According to the method ofRobertson (1975), individual grains are assigned the value ofthe highest pressure PDF observed within the grain accord-ing to some experimental calibration. The shock pressure istaken as the average of individual grain pressures. Variationson this method which account, respectively, for grains withno PDFs and grains converted to diaplectic glass, have beenproposed (Fel’dman 1994).

Mosaicism, the first appearance of diaplectic glass, 100%diaplectic glass, and lechatelierite, as well as the existenceof coesite and stishovite can all be used as quartz barometers

from Fig. 8.7. The effect of temperature on the high pres-sure phases is not well known, however. Microstructures infeldspar show a similar pattern with the notable differencethat mosaicism occurs at considerably lower pressures. Thebest pressure estimates are given by the coexistence of two ormore microstructures.

8.11.1 Calibration of shock pressures from op-tical properties of quartz

The refractive indices and birefringence of quartz at roomtemperature decrease continuously from samples shocked at25 GPa to those shocked at 35 GPa, which corresponds to thetransition to diaplectic glass (Fig. 8.10). The refractive indexof diaplectic glass decreases slightly as pressure increases fur-ther to 50 GPa. The effect of increasing pre-shock temperat-ure is to restrict the pressure interval over which the decreasesoccur: 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 variablesthat should be considered in an accurate calibration of shockpressure (e.g. Huffman et al. 1993, Huffman and Reimold1996). Different calibrations between experiments with dif-ferent pulse lengths suggest that this is a factor which has notbeen adequately explored, and raises the problem of extra-polation from experiment to nature (Huffman and Reimold1996). Grain size, grain size distribution, porosity and modalmineralogy may also have effects (e.g. Robertson and Grieve1977, Reimold and Stöffler 1978, Grieve et al. 1996).

Basal PDFs appear to be more resistant to post-shock an-nealing than other types of PDF (Leroux et al. 1994), render-ing shock pressure calibrations from PDFs unreliable wherethere has been post-shock annealing. Other barometers maybe questionable in this situation: for example, diaplectic glassis a metastable phase and may not be preserved. Stishovite


inverts in only 3 days at 250°C in the presence of water, andthus has limited preservation potential (Gigl and Dachville1968).

8.12 Diagnostic impact microstruc-tures

The identification of microstructures that are diagnostic ofimpact has profound implications for the understanding of theevolution of the earth. The subject has been very controver-sial, with much debate on the extent to which shock-inducedmicrostructures could be produced by volcanic eruptions orother endogenous processes as alternatives to meteorite im-pact.

PDFs have been associated with shock metamorphism onthe basis of experiments and observations of their distribu-tion in suspected impact target rocks (e.g. McIntyre 1962,Carter 1965, Hörz 1968, Robertson and Grieve 1977). How-ever, PDFs can be superficially similar to deformation lamel-lae formed by tectonic processes as described in Section 4.6,and this raises the controversial questions of whether PDFscan be diagnostic of shock metamorphism, and whether shockfeatures can be produced in volcanic eruptions.

Commonly used criteria for identifying impact-inducedPDFs include the following (e.g. Alexopoulos et al. 1988):

1.

2.

3.

4.

Well-defined, sharp features.

Straight.

Often occur in multiple sets per grain.

90% are in the and directions.

However, single sets of PDFs are developed in shock experi-ments under higher temperature conditions and are observed

in impact rocks, so that the third criterion is not general. Mul-tiple, intersecting sets of tectonic lamellae are also known(e.g. Lyons et al. 1993), but these lamellae are distinctly dif-ferent to PDFs, mainly because they are non-planar (Reimold1994). Several studies of PDF orientations in shocked targetsshow that criterion 4 is false: in particular, shock-inducedbasal PDFs (Brazil twins) are common at low shock pres-sures. Straightness (planarity), and sharpness or definition re-main important characteristics for identifying PDFs, but havenot been quantitavely defined: this lack of objectivity hascaused considerable controversy in the past. Optical studiesalone appear to be incapable of fully distinguishing impactfrom tectonic lamellae.

SEM and TEM-scale features do allow a clear distinctionbetween fresh PDFs and tectonic deformation lamellae. Etch-ing with HF liquid or vapour followed by SEM examinationis a more widely available technique than the TEM (Gratz etal. 1996). Relatively wide, glass-filled PDFs from impactsand shock experiments are revealed as sharp, planar featuresless than wide after etching, often containing sublamel-lar structure (see Section 8.4). By contrast, etched tectoniclamellae are much wider than and often sinuous, similarto their appearance under the optical microscope (e.g. Joreauet al. 1997; Fig. 4.11). Table 8.5 summarizes the differencesbetween shock-induced PDFs and deformation lamellae inquartz. Post-shock annealing reduces PDFs to a dislocationstructure similar to deformation lamellae, making the distinc-tion between the two difficult (cf. Goltrant et al. 1992). BasalBrazil twins have only ever been observed in quartz fromimpact structures and in experiments with deviatoric stressesover ~4 GPa at room temperature (Leroux et al. 1994). Theseobservations were used by Leroux et al. (1994) to argue thatbasal Brazil twins are strong evidence for shock.

Stöffler (1972) suggested that mosaicism is the most com-mon manifestation of shock. However, mosaicism is notdiagnostic of shock effects, and may also be caused by tec-


tonic deformation or even chemical effects, especially in feld-spars (e.g. Sharpton and Schuraytz 1989, Officer and Carter1991). The occurrence of coesite is known from inclusionsin high-pressure metamorphic rocks and kimberlites, but thevery fine-grained, polycrystalline nature of coesite in impactrelated rocks is diagnostic (Grieve et al. 1996). Diaplecticglass can be formed by static compression at 25-35 GPa, butthis is much higher than any pressures recorded by rocks atthe surface of the earth (Fig. 8.1): diaplectic glass remains adiagnostic indicator for shock pressures.

Explosive volcanism and deep-seated explosive activityare possible alternatives to meteorite impact as a source ofshock waves. Controversy exists over what pressures may beobtained in volcanic explosions; estimates range from values

within the field of crustal metamorphism (0.5 GPa) to valuesas much as 10 GPa. Planar microstructures, mosaicism andBrazil twins in volcanic rocks have been claimed as shockmicrostructures (e.g. Carter et al. 1986, 1990, Huffman andReimold 1993), but much of this evidence has been stronglyrefuted (Sharpton and Schuraytz 1989). The case for shockfeatures related to volcanism or deep-seated explosive activ-ity is not strong. In summary, Brazil twins, glass-filled PDFs,diaplectic glass, the distinctive very fine grained, polycrystal-line occurrence of coesite, and shock features in zircon arediagnostic of impact shock. As stressed by Reimold (1994),a thorough evaluation of the case for shock metamorphismshould examine a variety of minerals and evidence on macro-scopic to microscopic scales.

Chapter 9

From Microstructures to Mountains:Deformation Microstructures, Mechanismsand Tectonics

9.1 Introduction

Deformation microstructures and mechanisms can be used todeduce conditions of deformation, because they imply partic-ular relations between stress, strain rate (and possibly strain),temperature, fluids, microstructure, and mineralogy. Theserelations are derived from laboratory experiments and the-ories of material deformation. There are three types of re-lationship: failure criteria (Section 9.2), which give stressconditions for faulting, frictional sliding laws (Section 9.3),which give stress or stress-strain rate conditions for frictionalsliding, and flow laws (Section 9.4), which give stress-strainrate relationships for flow accommodated by intracrystallineplasticity, diffusive mass transfer (DMT), or composite de-formation mechanisms. Relationships between flow laws areusefully illustrated on a deformation mechanism map (Sec-tion 9.6) which shows fields of stress, temperature and grainsize in which particular mechanisms give the highest strainrate. The rheology of the lithosphere is often shown by a plotof stress against depth, or a lithospheric strength envelope(Section 9.7).

Although failure criteria and flow laws are essential toolsin tectonic interpretation, they are based on laboratory exper-iments which have serious limitations when extrapolated tonature (e.g. Carter and Tsenn 1987, Paterson 1987). Two ofthe most important problems are the limited scale of triaxialtests, which means that mesoscopic heterogeneities in rocksare not accounted for, and the fast strain rates of experiments(typically to ) compared to most tectonic de-formation ( to ), which mean that the deform-ation mechanisms may differ between nature and experiment.Microstructures offer a potential solution to these problems:if deformation microstructures and mechanisms can be iden-tified in nature that correspond to experiments, then laborat-ory results can be extrapolated with more confidence. ThisChapter provides some basic failure criteria, frictional slidinglaws, and flow laws to be used in conjunction with micro-structural observations to make tectonic interpretations. Theprimary intention of the Chapter is to provide quantitativedata based on the concepts in the previous Chapters.

9.2 Failure criteria

9.2.1 Coulomb and Mohr failure criteria

The simplest failure criterion, known as the Coulomb cri-terion, is that failure occurs when the shear stress on a plane

equals a material property called the cohesion andthe normal stress multiplied by the coefficient of internalfriction,

The criterion in terms of the principal stresses is:

where is the uniaxial compressive strength. The value ofis around 0.58 for many rocks, while varies widely for

different rock types. A useful generalization of rock typesin terms of their uniaxial compressive strengths is given byPaterson (1978):

1.

2.

3.

4.

Igneous and high-grade metamorphic rocks.

Low-porosity sedimentary and low-medium grade meta-morphic rocks.

High porosity sedimentary and some low grade meta-morphic rocks.

Low porosity dolomites and quartzites.

However, the relationship between shear and normal stressof real rocks is not linear as suggested by the Coulomb cri-terion. The Mohr criterion describes the relationship by anempirical fit to data from rock mechanics experiments, andso is limited to materials that have been extensively tested.The Griffith criterion (next Section) is a better approximationto rock behaviour in most situations.

9.2.2 Griffith failure criteria

Griffith’s thermodynamic approach to failure outlined in Sec-tion 2.2.1 leads to a failure criterion of the form:

90

CHAPTER 9. FROM MICROSTRUCTURES TO MOUNTAINS 91

is the uniaxial tensile strength. This simple criterion iswidely applicable, but limitations are the difficulty of meas-uring (which is dependent on the type of test used) andthe observation that in most rocks, the uniaxial compressivestrength is 10 to 12 times the value of (Jaeger andCook 1979). The Griffith criterion predicts thatValues of for use in the Griffith criterion can be calculatedfrom the generalized values of given in Section 9.2.1.

The simple Griffith criterion does not predict any effect ofon failure: this was accommodated in the Extended Grif-

fith criterion by Murrell (1963), giving the failure criterion:

This criterion implies that in better agreementwith observed values. A more rigorous treatment of triaxialstress acting on a body containing randomly orientated ellips-oidal cavities leads to a failure criterion which is independentof and implies lower values of than the Griffith fail-ure criterion (Murrell and Digby 1970). The effect of fric-tional sliding on closed Griffith flaws with a coefficient offriction and stress required for closure was dealt withby McClintock and Walsh (1962) in the Modified Griffith cri-terion:

is a function of Poisson’s ratio and the axial ratios of theflaws: for penny-shaped flaws andfor biaxial stress. The value of is (Murrell 1965);a value of 100 MPa was used for in lithospheric failuremodelling by Kusznir and Park (1984, 1987). If is takento be negligible, this criterion becomes equivalent to the Cou-lomb criterion and implies thatwhich generally has little relationship to the observed valuesof (Jaeger and Cook 1976). However, this criterion doesprovide a better fit to stresses measured at compressive failurethan either the original or extended Griffith criteria.

9.3 Pore fluid pressure and faulting

The principle of effective stress (Terzaghi 1943) describes theimportant mechanical effect of pore fluid pressure onfailure in many experiments (e.g. Paterson 1978). The prin-ciple states that normal stresses are reduced by the value ofthe pore fluid pressure, so that effective stresses can be sub-stituted in any of the failure criterion above. The effectivestresses (“conventional effective stresses” of Paterson 1978)are calculated from:

More detailed experiments and poroelasticity theory indicatethat the normal stress is not reduced by the exact value of thepore fluid pressure. The effective stresses are given by

where is known as the effective stress coefficient and hasvalues less than or equal to 1 for elastic properties. de-creases with permeability, porosity and confining pressure,

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

9.4 Fracture mechanics and failurecriteria

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

9.5 Frictional sliding laws

9.5.1 Byerlee’s law

Amonton’s Law (Section 2.2.2) describes frictional behaviourof rocks well at low normal stresses, but at higher normalstress, a threshold shear stress must be applied before anymovement occurs: this is called the cohesion, so that fric-tional sliding in rocks is described by:

The form of this equation is the same as used for the Coulombcriterion; however, the cohesion and coefficient of friction ofintact rock (the coefficient of internal friction) are differentfrom those of frictional sliding. Byerlee (1978) discoveredthat values of in most rocks are in the range of 0.5-1, withan average of approximately 0.75, and that frictional slidingin 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, roughnessof the sliding surface and temperature. Byerlee’s Law is of-ten assumed to describe the state of stress in the upper crust(e.g. Burov and Diament 1996) which implies that failure inthe upper crust is governed by frictional sliding of previouslyformed faults in ideal orientations for slip.

Marone et al. (1992) have proposed that distributed de-formation in a gouge zone has a separate frictional slidinglaw, because stress is re-orientated within the sheared layer

92 CHAPTER 9. FROM MICROSTRUCTURES TO MOUNTAINS

so that the maximum possible shear stress is supported. Thisresults in a sliding law of the form:

where is the coefficient of sliding on a discretesurface.

9.5.2 Rate and state dependent frictional slid-ing

Dynamic effects during frictional sliding, such as stick slipbehaviour, are not described by Amonton’s and Byerlee’sLaws. Stick-slip behaviour involves a steady rise in stresswith no movement followed by a sudden rapid episode ofstress drop and sliding, and is analogous to earthquake fault-ing. Stick slip behaviour is favoured over stable sliding inrocks with higher quartz contents, by sliding along roughersurfaces, by the lack of gouge, by high normal stresses, andlow temperatures (e.g. Paterson 1978). Dynamic frictionalsliding can be modelled by making the coefficient of frictionin Amonton’s Law depend on the rate of frictional slidingV:

where is the coefficient of friction at a reference velocityV*, a and b are empirically determined constants, and is astate variable (e.g. Rice and Gu 1983). L is the critical slipdistance for sliding to return to a stable rate after a change insliding velocity. It is possible to expand the above expressionwith more than one state variable. The value of a expressesthe dependence of on a change in velocity, and b expressesa change in that decays with time. This behaviour is foundto describe laboratory experiments very well. The stabilityof a sliding process depends on the value of if thisis negative, increasing velocity will lead to a reduction inand the fault is said to show velocity weakening, and hence becapable of nucleating earthquakes. On the other hand, ifis positive, velocity changes will lead to increased friction andthe sliding will be stable.

The rate dependent friction law can be applied to the nuc-leation of earthquakes in the crust (Tse and Rice 1986, Scholz1990). is expected to be positive in the upper few km ofwell-developed faults due to the presence of a layer of gouge(which is characterized by positive values of ). Thereforeearthquakes are not expected to nucleate in the top part of thecrust except in areas lacking well-developed faults. Below thelevel of gouge, there is an upper stability transition to negativevalues of which are the conditions for earthquake nucle-ation (Scholz 1990). becomes positive with increasingtemperature, so that there is also a lower stability limit be-low which earthquakes will not nucleate: this is probably ap-proximately at the 300°C isotherm in continental crust, whichexplains the observation that earthquakes generally nucleatein the top 15 km of the crust (e.g. Sibson 1982, Meisnerand Strehlau 1982). However, this picture is complicated bythe relation between velocity dependence and displacement:positive velocity dependence at small displacements changesto negative dependence with increasing displacement, due to

localization on shear surfaces, which always has a negativevelocity dependence (Marone et al. 1992).

Seismological observations support the above picture, butthere is a lack of corresponding knowledge about naturalmicrostructures produced by seismic slip. Pseudotachylites(Section 2.11) are the only direct evidence for seismic slip(e.g. Sibson 1977), but seismic slip can also be tentativelysuggested by textures such as crack-seal vein fillings whichindicate repetitive fracture followed by fracture sealing (Sec-tion 3.12). Particle size distributions of gouge may give an in-dication of its mechanical behaviour: velocity strengtheningbehaviour changed to velocity weakening when the fractal di-mension of the particle size distribution of the gouge attaineda value of 2.6 in experiments (Biegel et al. 1989). Anotherpotential microstructural indicator of seismic slip is the evid-ence of continuous deformation alternating with discontinu-ous deformation in slickenside surface material (Power andTullis 1989).

9.6 Flow laws

9.6.1 Diffusive mass transfer: Grain size sens-itive creep

The fundamental flow laws for deformation accommodatedby diffusive mass transfer can be derived from theoreticalprinciples for three cases: diffusion through a grain bound-ary fluid (pressure solution, Section 3.2), diffusion throughthe grains (Nabarro-Herring creep, Section 5.2), and diffu-sion through grain boundaries (Coble creep, Section 5.2). Allflow laws for diffusion-accommodated deformation have onedistinguishing characteristic: strain rate is a function of grainsize.

Pressure solution creep

Diffusive mass transfer through a solution requires threesteps: solution, diffusion and precipitation. The strain ratemay be controlled by the rate of diffusion in the fluid, the rateof solution or precipitation at the source or sink, or the rate ofdiffusion at the source or sink (Paterson 1995), and the flowlaw depends on which process is rate-limiting. The follow-ing 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 para-meter varying from 40 to 140 that depends on a geomet-rical model for the arrangement of the grains, is the molarvolume 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 coefficientin the grain boundary fluid is the activation en-thalpy for pressure solution R is the gas constant

T is the temperature (K), and is thedifferential stress (Pa). The reference state for and is

It is conventional to quote flow laws for the uniaxial


shortening strain rate under axisymmetric stress, and the dif-ferential stress relations between other componentsof the stress and strain rate tensors can be obtained from ap-propriate geometrical modifications.

The terms are sometimesgathered together as a phenomenological coefficient repres-enting effective grain boundary diffusivity (e.g. Spiers et al.1990). Diffusion control is relevant for salt in which solutionrates are very fast (e.g. Spiers et al. 1990), and has beenreported for experiments on quartz (e.g. Gratier and Guiget1986). 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 andGuegen 1996). Paterson (1995) presents a model for pressuresolution creep in porous materials based on an analogy withparticulate flow in soil mechanics, leading to equations foreach of the three possible rate limiting processes.

Diffusion-limited pressure solution creep in quartz hasbeen described by the following flow law (modified by An-gevine and Turcotte 1983 after Rutter 1976):

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 inA flow law for the solution/deposition-controlled case is

given by :

is the reference state diffusion coefficient for grainboundary diffusion, is the grain boundary width, andis the activation enthalpy for grain boundary diffusion. Thespecification of is difficult because it depends on thediffusing species. The strain rate will however be limited bythe diffusion rate of the slowest species, which is normallyquoted. In olivine, the values of the diffusion coefficients in-crease in the order Si < O < Fe; the value for Si is rate limit-ing. Values for Coble creep flow laws are given in Table 9.1.

Nabarro-Herring creep

The flow law for Nabarro-Herring creep is similar to the pre-vious law, with the following important differences: the grain

where is the reference state volume diffusion coefficient,and is the activation enthalpy for volume diffusion. Para-meters for the flow law are given in Table 9.1.

A flow law for diffusion creep that combines both Cobleand Nabarro-Herring creep (Ashby 1970, Raj and Ashby1971) is:

The value of A may differ in all the above expressions sinceit depends on the preferred geometrical model of the grains.

9.6.2 Intracrystalline plasticity

Two different flow laws can be formulated from theory to de-scribe flow controlled by dislocation glide (abbreviated to dis-location glide; Section 4.2) and flow controlled by dislocationclimb, known as dislocation creep (Section 4.2).

Dislocation glide

Theoretical flow laws for dislocation glide are characterizedby 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, andF is a constant (e.g. Dorn 1954). Dislocation glide is notconsidered to be important in most minerals under geologicalconditions because of the relatively high stresses required forthe process.

Dislocation creep

At lower stresses, higher temperatures and slower strain rates,a power law relationship describes the strain rate-stress rela-tionship:

where A is a pre-exponential constant, is the activation

ample, “Dry” quartz, with H contents of less thanSi, is an order of magnitude stronger than “wet” quartz (the

where kC is the velocity of the deforming crystal facee.g. Raj 1982). Solution/deposition controlled pressure solu-tion has been described in quartz (Gratier and Jenatton 1984),but seems to be less common than diffusion control. Pa-terson’s (1995) model suggests that for quartz and rocksaltaggregates, source/sink diffusion control is likely to be ratelimiting under geological conditions, but in high-temperaturelaboratory conditions, precipitation control is effective forquartz. Experimentally determined values for the relevantparameters are given in Table 9.1.

Coble creep

The flow law for Coble creep has a very similar form to thatfor pressure solution creep since both processes involve diffu-sion around grain boundaries. The major differences are thelack of a term for solubility, and differences in the values ofthe diffusion coefficient and activation enthalpy:

size dependence is to the square of d there is no grain bound-ary thickness to take into account, and the diffusion coeffi-cient and activation enthalpies are different:

enthalpy for dislocation creep, and n is the stress exponent(Weertman 1968). A depends on the volume diffusion coeffi-cient and the Burgers vector of the active dislocation system.

Unfortunately there are severe problems in using thissimple flow law. One of the most important variables that isnot explicitly taken into account is the effect of water. For ex-


considerable amount of literature on this subject is summar-ized in Paterson and Luan 1990). The effect is probably due inpart to replacement of Si-O-Si bonds by weaker Si-OH-OH-Si (hydroxyl) bonds, known as hydrolytic weakening (Griggsand Blacic 1965, Griggs 1967), which is similar to the mech-anism that controls sub-critical microcrack propagation (Sec-tion 2.2.1.1). Hydrolytic weakening enables easier disloca-tion glide and creep, although other mechanisms may alsoplay a role in the weakening effect of water. This suggeststhat the flow law should incorporate a function of the wateractivity (e.g. Kronenberg and Tullis 1984, Ord and Hobbs1985). Experiments show that the power law exponent de-pends on water content, at least for low water contents (e.g.Blacic and Christie 1984). In any event, it is difficult to de-termine water contents at the time of deformation, althoughKronenberg et al. (1990) have shown how quartz water con-tents measured across a natural shear zone vary with shearstrain, possibly reflecting a variation during deformation.

Mean stress is another variable that should appear in theflow law, which may have important rheological effects byincreasing water diffusion rates, increasing water solubility,and increasing activation enthalpy. Quartz phase appears toaffect both the activation enthalpy ( is higher for than

(e.g. Linker and Kirby 1981). The quartz phase af-fects diffusion coefficients in a remarkably similar way to ac-tivation enthalpies of power law creep, showing that the creepenthalpies are essentially the same as the enthalpies of diffu-sion (Gilletti and Yund 1984, Dennis 1984). Trace amounts ofimpurities in quartz may also cause considerable weakening,probably through their control on glide velocity (e.g. Jaoul etal. 1984).

9.6.3 Empirical flow laws from experimentaldata

Experimental data can be fitted to empirical laws of the fol-lowing general form for low stress levels:

where A is a pre-exponential constant ordepending on the units of ), Q is an activ-

ation energy P is the mean stress (Pa), V is an ac-tivation volume often ignored due to its relatively

oretical relationships: and indicates dislocationcreep, and indicates pressure solution or Coblecreep, and and indicates Nabarro-Herring creep.Superplasticity is characterized by and (e.g.Boullier and Guegen 1975, Schmid et al. 1977). Tables 9.2-9.4 are a compilation of some results for minerals and rocks,divided into grain-size sensitive Table 9.2), andgrain-size insensitive flow laws ( Table 9.3 for min-erals and Table 9.4 for rocks). When using these tables orany such rock mechanics data, it is advisable to consult theoriginal reference to find any important specific conditions ofthe experiments which may limit the applicability of the flowlaws.

At high stress levels, the power law relationship does notfit the experimental data (“power law breakdown”, Carter andTsenn 1987). Two alternative empirical relationships may de-scribe the data well:

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

and the power law exponent n which is lower for

values of m and n suggest interpretations of the deformationmechanism from experimental data by analogy with the the-

small contribution to the exponential term), and m and n aregeneral dimensionless grain size and stress exponents. The

e.g. Heard (1963), with C, and n as the empirically-derivedconstants.

9.7 Polymineralic deformation

While flow laws such as those in Section 9.4.3 can describethe bulk rheology of rocks, microstructural evidence showsthat strain is distributed in a complex manner between miner-als with different rheologies, which should be accounted forin a detailed understanding of deformation. An empirical ap-proach was taken by Tullis et al. (1991), who derived a flow


law for a bimineralic rock that weighted the flow law para-meters of each mineral by the volume fractions of theweak and strong minerals as follows:

of the composite is given by the sum of the shearstrengths of each mineral multiplied by their volume frac-tions:

where and are the shear strengths and volumefractions of the weak and strong minerals respectively. Forthe IWL case,

The distinction between LBF and IWL behaviour occurs atthe conditions where both give the same bulk strength. Oneither side of this equality, the more stable microstructureis the one that dissipates less strain energy. For a quartz-feldspar composite with dislocation creep flow laws, the LBFmicrostructure is only stable for very low volume fractions ofquartz.

This model shows that strength of the composite alwayslies between the limiting strengths of the strong and weakphases, but the relative strengths of the composite comparedto the end members are a complex function of their strengthcontrasts and their volume fractions. In a granite, 10-20%volume fraction of quartz reduces the composite strengthby one half compared to the strength of pure feldspar, as-suming IWL behaviour, and that the deformation occurs inthe conditions appropriate for dislocation creep flow laws.Quartz effectively governs the strength of the composite atvolume fractions over 10%. Feldspars in gabbros signific-antly weaken the composite strength from a pure pyroxenerock up to 900°C: at higher temperatures, pyroxene is weakerthan feldspar, and the composite has an intermediate strength.

Both the Tullis and Handy approaches outlined above arelimited to two phase rocks. A generalization to any numberof phases is given by Ji and Zhao (1993) for two extremeassumptions. If all phases are loaded equally (isostress cri-terion), the flow law is given by taking the volume average ofthe strain rates of individual phases:

An average flow law is considered to lie between thephysically unrealisitic extremes of the isostress and isostraincases. The average is determined by minimising the dif-ference between two composite flow laws, that each allowfor variable, complementary proportions of the isostress andisostrain flow law parameters. The results achieve reasonablefits to some experimental data and the method is advantage-ous when the effects of more than two phases need to be con-sidered.

In another approach to the deformation of polymineralicrocks, Ji and Zhao (1994) have used a fibre loading model toderive an expression for the elastic solution to the problemof rigid fibres in a weaker matrix, and generalized from theelastic solution to a viscoplastic case. The equation they de-rive includes an expression for the aspect ratio of the fibres.Their model describes some experimental results better thanthe Tullis model, but they point out that it assumes uniformly

where n, Q, and A are the parameters in the dislocation creepflow law (Section 9.4.2.2), and the subscripts and s arefor composite, weak, and strong minerals respectively. Thisgives a reasonable fit to experimental data.

A more fundamental way to treat polymineralic rocks isto model the relationship between the behaviour of differentminerals. Handy (1990, 1994) has suggested that for twominerals of contrasting strength, the microstructure may bedescribed as a load bearing framework (LBF) if the strongermineral supports stress around pockets of a weaker mineral,or as an interconnected weak layer (IWL) if the weaker min-eral forms a matrix around microboudins of the stronger min-eral. The behaviour of the composite depends on the contraston rheological properties of the minerals, and on their volumefraction.

Quantitative observations can be used to distinguish LBFand IWL microstructures. In granite deformed under am-phibolite facies, for example, the LBF case may be char-acterized by similar grain shapes and sizes in all minerals,and strain magnitudes, as estimated approximately by grainshapes, are low (Schulmann et al. 1996). Quartz has largerecrystallized subgrains, and K-feldspars are fractured, whileplagioclase recrystallizes. There is a transition to an IWLwith increasing strain, which is manifest from higher strainin quartz than feldspar, and the appearance of S-C fabrics.The ultimate stage of the process is a transition to anothertype of IWL microstructure in which there is little contrastbetween the quartz and feldspar grains: all minerals have re-crystallized completely in a mylonite.

Two flow laws are necessary to describe the LBF and IWLcases separately, which can be derived from considering thatthe total rate of viscous strain energy dissipation is the sumof the rates in the weak and strong minerals. In the LBF case,both minerals deform at the same rate and the shear strength

where a function x is introduced that describes the sensitivityof the strain rate partitioning in the rock to the strength con-trast of the minerals. The nature of this function can be sug-gested by considering limiting values. One possibility sug-gested by Handy (1994) is:

where n, A, and Q are the flow law parameters, the subscriptsr and i refer to the isostressed aggregate and the individualphases (total number m and volume fractions f) respectively.An alternative assumption is that all phases have the samestrain and strain rate. This isostrain criterion leads to the fol-lowing aggregate properties:


distributed fibres with a constant orientation, which may limitits geological application. Flow laws for polymineralic rocksrequire more theoretical and experimental research.

9.8 Deformation mechanism maps

The rheology predicted by the flow laws in Section 9.4 canbe compared on a deformation mechanism map (Fig. 9.1),which shows the relationships between deformation mech-anism, stress, strain rate, temperature, and grain size (e.g.Stoker and Ashby 1973, Ashby and Verrall 1978). Thesemaps may be shown in plots of stress normalized by theshearmodulus against temperature normalized by melt-ing temperature (homologous temperature, ). The shearmodulus normalization allows for the systematic dependenceof the flow law constant A on and for the temperature de-pendence of the modulus (cf. Tsenn and Carter 1987). Themelting temperature normalization allows for better compar-ison between different materials because transitions betweendifferent rheological regimes depend on These nor-malizations are relevant for materials-science applications,but the direct values of stress and temperature are more usefulfor tectonic studies.

The fundamental assumptions for constructing these mapsare that deformation will occur by the mechanism thatprovides the fastest strain rate, and that steady state conditionsare obtained. Deformation mechanism fields in stress andtemperature space are contoured for strain rate using the flowlaw of the deformation mechanism with the fastest strain rate;boundaries between fields correspond to loci of equal strainrate. Different maps are constructed for different grain sizes;alternatively grain size can be added as a third dimension.The possibility of contributions to strain rate by more thanone mechanism can be allowed for (e.g. Stoker and Ashby1973). For example, both diffusional creep mechanisms canbe combined as in the last equation in Section 9.4.1.3, and dif-fusion and dislocation processes can also combine. However,flow by dislocation glide and dislocation creep are mutuallyexclusive.

Deformation mechanism maps for geological materialshave been produced for only quartz, olivine and calcite (e.g.Rutter 1976, White 1976, Carter and Tsenn 1987), because arange of flow laws for different mechanisms are not availablefor other minerals or rocks. Microstructures are a guidingprinciple to the construction and use of deformation mechan-ism maps: similarity in experimental and natural microstruc-tures and mechanisms allows for the correct choice of flowlaw parameters (e.g. Busch and Van der Pluijm 1995).

9.9 Lithospheric strength envelopes

A Lithospheric Strength Envelope (LSE) shows the relation-ship of lithospheric strength with depth derived from elasti-city, failure criteria, frictional sliding laws and flow laws(Fig. 9.2). The term yield strength envelope is commonlyused (e.g. Goetze and Evans 1979), but since the upper partof the envelope is governed by cataclastic failure, LSE is abetter term. LSEs are a powerful way of describing the be-

haviour of the lithosphere and have important applications tolarge-scale tectonic problems (e.g. Burov and Diament 1995,1996).

The upper part of the LSE is usually constructed usingByerlee’s Law for frictional sliding to represent cataclasticbehaviour. Coulomb or Griffith failure criteria can also beused to model the behaviour of intact lithosphere (cf. Sib-son 1983, Kusznir and Park 1987). A power law for disloca-tion creep is usually assumed for the lower, plastic part of theLSE, justified by observations that suggest that lithosphericstress/temperature/grain size conditions are appropriate forthis deformation mechanism. Dislocation glide can also beincluded (e.g. Burov and Diament 1995).

The LSE is based on a model of the compositional structureof the earth. Gross simplifications are necessary due to thelimited availability of rock mechanics data, and the difficultyof applying them to the earth. The upper crust is commonlymodelled by wet quartzite, the mid to lower crust by feldsparor diorite, and the mantle as olivine. A temperature-depthprofile is chosen for the crustal model. Appropriate steadystate geotherms can be calculated from surface heat flux (e.g.Kusznir and Park 1984, 1987), but Burov and Diament (1995)have modelled an evolving geotherm due to cooling after oro-geny.

The LSE is delineated by the lesser of the cataclastic orplastic strengths. The stresses can either be calculated fora given strain rate, or by applying a constant stress at theboundaries of a model that analyses the stress and strain withtime. The latter approach shows that stress applied uniformlyat the boundaries of the lithosphere is redistributed into themore competent layers which deform elastically, a processreferred to as stress amplification (Kusznir 1982). Eventu-ally the elastic strength of all layers is exceeded, leading towhole lithosphere failure, and deformation continues by fric-tional sliding and intracrystalline plasticity (Kusznir and Bott1977, Kusznir and Park 1982, 1984, 1987). Hence steadystate LSEs have no part that is controlled by elastic strength.

The shape of the simplest LSE (for a lithosphere of homo-geneous composition) consists of a rapid linear increase instrength with depth from the surface of the earth governed bythe high normal-stress dependence of Byerlee’s Law or a fail-ure criterion (e.g. the top part of Fig. 9.2). The cataclasticpart of the LSE depends on the orientation of the failure sur-face: differential stresses for faulting or frictional sliding atany level of the crust increase in the order normal - strike-slip - thrust fault. The strength of the lithosphere reachesa maximum where the cataclastic strength is equal to theflow stress for intracrystalline plasticity. Below this point theflow stress is less than the cataclastic strength, and continuesto decrease with depth as temperature increases. LSEs forlayered models of the earth generally show the above patternrepeated in each layer (Fig. 9.2). There are dramatic strengthcontrasts across layers, especially between the crust and themantle. This suggests that detachment could occur at theselayer boundaries, which are sometimes misleadingly called“brittle-ductile transitions”, but are more accurately referredto as brittle-plastic transitions (Section 1.3, 1.5).

However, saw-tooth LSEs such as those in Fig. 9. 2 canbe criticized because the validity of Byerlee’s law has not


been established at mid to lower crustal pressures, and thehigh stresses predicted by extrapolation of the law to theseconditions are unrealistic (Ord and Hobbs 1989). Moreover,the semibrittle deformation regime is not recognised, prob-ably because of the paucity of published flow laws. Morerealisitic possibilities for the shape of the LSE at mid crustaldepths are either a vertical line of constant stress with a valueequal to the stress at the base of the seismogenic zone (Hobbset al. 1986), or non-linear, pressure-dependent functions de-rived from Mohr-Coulomb constitutive laws for permanentdeformation (Ord 1991). The experimental basis for the lat-ter laws is very limited, and they have the disadvantage ofpredicting 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. Mi-crostructural methods have concentrated mainly on measur-ing differential stress (Sections 9.10.2-9.10.6), and the ori-

entation of the stress tensor (Sections 9.10.7-9.10.8). Pa-laeopiezometers are calibrated by experiments and theory,but there are large variations in calibrations of the same pa-laeopiezometer. Furthermore, different palaeopiezometersyield considerably different values when applied to the samerocks. The latter variations can potentially lead to further tec-tonic insights (Section 9.10.9). Determination of the meanstress is considered in the separate Section on geothermoba-rometry (Section 9.11).

9.10.2 Recrystallized grain size

can be related to differential stress ( MPa) by an expressionof the form:

The size of grains recrystallized during deformation (d, mm)

where m and are constants, which can be determined ex-perimentally or derived theoretically (e.g. White 1979). Val-ues of m and are given in Table 9.5, and the calibrationsfor quartz are plotted in Fig. 9.3. The differential stressesfor quartz predicted by different calibrations differ by up tofour orders of magnitude, for at least seven possible reasons,


which are worth examining in detail because they highlightthe problems of recrystallized grain size palaeopiezometry.

1.

2.

3.

4.

5.

6.

7.

Recrystallized grain size appears to depend on the re-crystallization mechanism: subgrain rotation (regime 2)gives smaller grains than grain boundary migration (re-gime 3) (e.g. Drury et al. 1985, Drury and Urai 1990,Hirth and Tullis 1992).

The importance of the effect of water is obvious fromcomparing the dry and wet calibrations of Ord andChristie (1984). However, it does not appear to be im-portant in olivine (Van der Wal 1993).

The possibility that the recrystallized grain size is sens-itive to temperature is suggested by some data on metalsand olivine (e.g. Mercier et al. 1977, Ross et al. 1980),and predicted by theory (Mercier 1980a).

The possibility that the recrystallized grain size is de-pendent on quartz phase is suggested by the flow lawdependence on phase (9.4.2.2).

Kinetic effects may severely hamper determination ofthe calibration constants if equilibrium grain size is notattained during experiments (Twiss 1977). The lack ofsuch equilibrium for grain sizes on the order of mm isstrongly suggested by the experiments of Kronenbergand Tullis (1984); hence recent experiments have beenon much finer grain sizes (e.g. Post and Tullis 1999).

Equilibrium grain size may not be achieved if grainboundaries are pinned by impurities (e.g. Evans andWhite 1984).

Recrystallized grain size has been measured by varioustechniques and may be specified by different parameters:a standard practice is to measure the mean linear inter-cept grain size and multiply by 1.5 to allow for trunca-tion and sampling effects (e.g. Christie and Ord 1980).However, this will not be correct for grains with a fabric.

9.10.3 Subgrain size

Subgrain size can be related to stress by an expressionof the form:

where and are calibration constants, is the shear mod-ulus, and is the most common Burgers vector. Values of1 and 2 for have been proposed on theoretical grounds(Twiss 1977). Another possible relationship proposed byTwiss (1986) is:

where C and are constants, implying that there is astable subgrain size at zero stress. However, the second rela-tionship does not fit data for olivine and quartz better than thefirst. This may be due to the lack of accurate, low stress meas-urements (Twiss 1986). Until these are available, empiricallyderived constants for the first relationship are probably themost satisfactory calibration of the subgrain size palaeopiezo-meter, as given in Table 9.6. Problems with the applicationof this palaeopiezometer include a possible temperature de-pendence, the sensitivity of subgrain size to water known forolivine (Twiss 1986, Van der Wal 1993) and the problem ofsubgrain size measurement. Subgrain sizes may be measuredunder the optical microscope in reflected light after etching(see Ord and Christie 1984 for the method) or by electron mi-croscopy. Optically determined subgrain sizes are commonlyan order of magnitude larger than subgrain sizes measuredby electron microscopy (Ord and Christie 1984). Availabledata at present may not allow reliable extrapolation outsidethe range of subgrain sizes used to calibrate the palaeopiezo-meter, since alternative calibrations fit the experimental dataalmost equally well, yet give enormous differences when ex-trapolated (Twiss 1986).

9.10.4 Dislocation density

The dislocation density relationship can bewritten as:

where D and are constants, implying a steady statefree dislocation density. However, data for olivine and quartzare not fitted any better by the second relationship. Table 9.7presents calibration constants for quartz and olivine using thefirst expression. Problems with this palaeopiezometer arepossible temperature dependence and the difficulty of meas-uring dislocation density accurately. Invisibility of disloca-tions under the TEM may cause an underestimate in disloca-tion density of 30% (Ord and Christie 1984).

9.10.5 Twinning - differential stress

Three different approaches have been proposed to relate twin-ning in carbonates to differential stress. Two of these (Jam-ison and Spang 1976 and Laurent et al. 1981, 1990) takeno account of the fact that grain size is known to have aneffect on the twinning, and are therefore less suitable than

where k and u are constants, and and are defined as above.A theoretical value of u is 0.5 (Weathers et al. 1979), butmeasured values have a range from 0.45 to 3.33 (Twiss 1986).An alternative form of the relationship sometimes used is:


the approach of Rowe and Rutter (1990), who have calib-rated three twinning palaeopeizometers from laboratory ex-periments, which all appear to be independent of temperature,strain rate and strain.

Twinning incidence,

is the proportion of grains with optically visible twins inany grain size class. It can be related to differential stress

In practice, stresses are calculated from for each grain sizeclass and plotted against grain size: they should lie on linesof constant slope for different values of The standard errorin the experimental calibration was 31 MPa.

Twin density,

is the number of twins per mm. This is calculated by theslope of the relation between number of twins in a grain sizeinterval and grain size. appears to be independent of grainsize, and is related to with a standard error of 43 MPa, by:

Maximum twin volume,

is the maximum volume proportion (%) of twinned ma-terial in a grain size class. In the experiments, the area frac-tion of twins was taken as equal to the volume fraction, andwas measured from a two-dimensional thin section. isrelated to with a standard error of 41 MPa, by :

As with the previous palaeopeizometers, there are a numberof potential problems. Clearly if twinning is the only deform-ation mechanism, the amount of twinning measured by any ofthe three parameters will be a function of strain. The methodthus assumes that after twinning, equilibrium is reached with

a non-twinning deformation mechanism. In principle, grainsize distribution could have an effect on this palaeopiezo-meter because stress distributions in grains are strongly in-fluenced by the size of neighbouring grains (Newman 1994),but this has not yet been evaluated or demonstrated. Thispalaeopiezometer has been calibrated only for calcite (othercarbonates twin at different stress levels), and since the cal-ibration experiments were carried out at 400°C and above, itis probably not appropriate for low temperature deformation(Burkhard 1993). The best results will be obtained in samplesthat have experienced a single, coaxial strain event (the con-ditions of the experimental calibration), otherwise the methodwill overestimate stress (Rowe and Rutter 1990).

There is some experimental evidence that twinning inpyroxene could be used as a palaeopiezometer: Tullis (1980)suggests that the minimum stress needed for twinning inpyroxenes is 100 MPa.

9.10.6 Deformation lamellae

The appearance of deformation lamellae (Section 4.7) in ex-periments is characteristic of flow laws with an exponentialrelationship between stress and strain rate, which occur atrelatively high stresses. This observation leads to the sug-gestion that there may be a critical differential stress for theformation of deformation lamellae, dependent on the type ofbonding and crystal structure of the material (Blenkinsop andDrury 1988). Values of 100-200 MPa have been suggestedfor quartz. However, it is possible that the critical stressmay be inversely temperature-dependent, so that at presentdeformation lamellae in quartz should only be taken as qual-itative indicators of relatively high stress levels (Drury 1993).

The relation derived by Koch and Christie is ( MPa):

(MPa) and grain size (d, mm) by the following equation:

Spacing of deformation lamellae (s, mm) may constitute apalaeopiezometer (Koch and Christie 1981, McLaren 1991).


9.10.7 Principal stress orientations from de-formation lamellae

Techniques of using deformation lamellae in quartz to de-duce principal stress orientations were developed by Carterand Friedman (1965) and Carter and Rayleigh (1969). Ofthe three methods proposed by Carter and coworkers, the “ar-row” method is the easiest and yields consistent results. Theorientations of the c-axis and deformation lamellae are firstmeasured from individual quartz grains. The plane containingthe pole to a lamella and the c-axis is constructed (Fig. 9.4).

and are located within this plane at 45° to the lamellaplane. is distinguished from by the fact that the c-axislies closest to An arrow from the c-axis to the lamellapole will point from to (Fig. 9.4). Unique andorientations can be derived by eigenvector analysis from anumber of such measurements (e.g. Spang and Van der Lee1975). This construction implies that the deformation lamel-lae form in a plane of maximum resolved shear stress with amaximum shear stress direction parallel to the projection ofthe c-axis onto the lamellae plane. The evidence presented inSection 4.7 shows that recovery is an integral part of lamellaedevelopment, so that a simple analogy with a slip system forthe 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 theoreticalbasis remains unclear. The principal stress axes from deform-ation lamellae are in good agreement with those calculatedfrom the carbonate twin method from the same samples (e.g.Spang and Van der Lee 1975). It is possible to consider otherangular relationships between and the lamella plane, butsome results suggest that the stress axes do not change ap-preciably from the 45° constraint (Spang and Van der Lee1975). It is also possible to analyze for the principal stresseswithout using a fixed angular relationship between and thelamella plane, by using the right dihedra technique (cf. An-gelier 1984).

9.10.8 Principal stress orientations and strainsfrom twins

Twins have been used to deduce principal stress orientationsby assuming that the twin plane is a shear plane and the twindirection is parallel to the maximum resolved shear stress

(e.g. Turner 1953, Weiss 1954). Twinning in calcite occurson the e planes in the direction of the plane contain-ing the pole to the twin and the c-axis (Fig. 9.4). andare therefore at 45° to the twin plane in the plane of the c-axisand the pole to the twin - a similar geometry to that used forthe quartz deformation lamellae method. However, in con-trast to that method, is in the opposite quadrant from thec-axis (Fig. 9.4). A similar construction can be applied to ftwins in dolomite, but here is in the same quadrant as thec-axis (Fig. 9.4).

A development of this technique is to use the right dihedramethod to determine the and orientations that are com-patible with the largest numbers of twins. The value of themaximum number of compatible twins (the MAX number)gives a measure of the degree to which the data can be fittedby a single stress orientation (Pfiffner and Burkhard 1987).This method has the advantage that it does not assume a fixedangle between and the twin plane. A further developmentis the determination of principal stress orientations and ab-solute magnitudes by assuming homogeneous stress distribu-tion and a critical resolved shear stress for twinning (Laurentet al. 1981, 1990).

It is possible to calculate a finite strain tensor by incorpor-ating the shear strain necessary for twinning into the orient-ation analysis (e.g. Groshong 1972, 1974, 1984). A com-puter program to carry out the analyses is described by Evansand Groshong (1994). The measurements necessary for thedetermination include the c-axis orientation, twin set orienta-tion, average thickness and number of twins, and grain width.25 grains in two perpendicular thin sections should be meas-ured.

Twins which do not fit the calculated strain or stress fieldsare a problem in all these methods. Two basic philosophiesto deal with this problem have been adopted. The first is torecognize that some twins will inevitably develop in incom-patible orientations due to stress inhomogeneity. The propor-tion of such twins (negative expected values, NEV, Groshong1972) can be used to test the homogeneity of the sample, andthe incompatible grains can be removed from the data set.Alternatively, the incompatible twins are presumed to reflectdifferent superimposed stress fields, and they can be used toseparate out a number of different stress fields (e.g. Lacombeet al. 1990). This method has been criticized by Burkhard(1993) because it has yielded unrealistic results, and the as-


sumption of stress homogeneity is clearly false in a polysizedgrain aggregate.

9.10.9 General problems with palaeopiezomet-ers

A general problem with all palaeopiezometers is the interpret-ation of the results. The significance of the stress level recor-ded can only be evaluated in the context of an evolving stressfield, analogous to the problem of interpreting results fromgeothermobarometry within a P-T path. Interpretations be-come especially problematic when considering results fromthe same sample that differ according to the method used. Forexample, stresses measured by dislocation density, subgrainsize and recrystallized grain size may differ because they havevariable dependence on strain magnitude. Dislocation dens-ities may be reset after less than 1% strain, compared to 5%

for subgrain sizes and perhaps 30% for recrystallized grainsizes; dislocation densities may therefore record later tectonicevents with low strains such as stress during uplift (e.g. White1979a). This problem can potentially be overcome by plot-ting stress levels derived from one palaeopiezometer againstanother. If the stresses are recorded by both palaeopiezomet-ers from the same part of the stress-time path, they will lieon a line with a unit slope, while deviations from this lineindicate 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 (geo-thermobarometry) is traditionally the domain of thermody-namics and mineral chemistry. However, there are a num-ber of promising new developments in the application of mi-crostructures to geothermobarometry, which are particularlyrobust because they are calibrated from naturally deformedsamples.

In addition to these new, quantitative approaches, experi-ments and thermodynamic results allow some generalizationsabout the minimum temperatures for plasticity. Plasticity inthe form of twinning can occur even at room temperature incalcite, but appears to require a temperature of 300°C in dolo-mite. Plasticity (sensu lato) in quartz and mica is generallyrestricted to temperatures greater than 200°C. The onset ofplasticity in feldspars is commonly taken as 450°C, but thisvalue is not well constrained. The common mafic minerals(clinopyroxene, amphibole, orthopyroxene and olivine) maybe plastic from 500°C or above. These numbers must how-ever be treated with great caution, because there are manyother variables involved in the mechanism transitions (not-ably stress/strain rate, see Section 1.3), and because of theproblems of extrapolating from experiments to nature (Sec-tion 9.1).

9.11.2 Calcite twin morphology

Twin morphology in calcite appears to be strongly temperat-ure dependent, and has been proposed as a geothermometer(Burkhard 1993). Four types of twins can be distinguished(Fig. 9.5), together with appropriate temperature ranges fordeformation:

Type I. Thin (1 mm), straight, rational twins form at temper-atures of less than 200°C.

Type II. Thicker, straight and rational twins (> 1 mm),which are slightly lens shaped, form at 150-300°C.

Type III. Thick, curved, irrational twins, which may containtwins within twins, form at temperatures greater than200°C.

Type IV. Thick, irrational and patchy twins, breaking up intotrails of small grains, form at temperatures over 250°C.


9.11.3 Sutured quartz grain boundaries

The geometry of sutured grain boundaries (Section 4.8)formed by grain boundary migration depends on the temper-ature during deformation. With increasing temperature, thelength of segments, serrations or lobes increases (Fig. 9.6,Kruhl and Nega 1996). The geometry of a grain boundary canbe characterized by its fractal dimension, D, which is bestmeasured by the “divider” method (Kruhl and Nega 1996).A grain boundary is divided into linear segments (“strides”)

length. The process is repeated over a range of values of r.

D is derived from the gradient of a log-log plot of L againstr. Kruhl and Nega (1996) established the relation shownin Fig. 9.7 empirically from measurements of natural su-tures with known temperatures of deformation. The relationbetween T (°C) and D is approximately:

D values of sutures produced in experimental deformationappear to be affected by the experimental strain rate as wellas temperature (Takahashi et al. 1997). However, the extentto which strain rate may be important in natural deformationis unknown, especially because the experimental strain rates

were limited to to and maximum strains of31%. Ultimately the validity of a temperature - only depend-ence of D rests on the empirical evidence.

9.11.4 Subgrain boundary orientation inquartz

The orientation of subgrain boundaries in quartz appears tobe controlled by the phase present during deformation (Sec-tion 4.6). The presence of chessboard patterns, indicatingboth basal and prism-parallel sub-grain boundaries, is restric-ted to deformation in the field (Kruhl 1996). Thismeans temperatures above 573°C at 0 MPa, and 825°C at1000 MPa (Gross and Van Heege 1973). Chessboard patternscan only be seen in grains with c-axes subparallel to the planeof the section (high interference colours) and must be distin-guished from three types of pseudochessboard pattern (Kruhl1996).

1.

2.

3.

Two sets of prismatic subgrain boundaries visible ingrains with c-axes at high angles to the section (i.e. lowbirefringence).

Rectangular kink band boundaries that are not crystallo-graphically orientated.

Two sets of rhombohedral subgrain boundaries in grainswith c-axes subparallel to the section, which can be dis-tinguished by their inclined extinction positions.

of individual length r. The length of the grain boundary Lis equal to the product of the number of segments and their

For a fractal grain boundary, L is related to r by:

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Abbreviated references collected bychapter

Italics indicates important general sources for the chaptertopic.

Chapter 1. Beaumont et al. 1996, Byerlee 1968, Carter& Kirby 1978, Chester & Logan 1987, Evans et al. 1990,Griggs & Handin 1960, Kirby & Kronenberg 1984, Knipe1989, 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.

Chapter 2. Agar et al. 1988, Allison & La Tour 1977, Ameen1992, 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, Blen-kinsop & Sibson 1991, Boldt 1995, Borg & Maxwell 1956,Borg et al. 1960, Borradaile 1981, Brace & Bombolakis1963, 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, Gallagher1981, Gallagher et al. 1974, Griffith 1924, Grocott 1981,Hadizadeh 1980, Hadizadeh & Rutter 1983, Hadizadeh &Tullis 1992, Hancock 1985, Hippert 1994, Hirth & Tullis1991, Horii & Nemat-Nassar 1985, 1986, House& Gray1982, Hull 1988, Inglis 1913, Jaeger & Cook 1979, Jamison& Stearns 1982, Kemeny & Cook 1987, Knipe 1986, 1989,

Krantz 1979, 1983, Krantz & Scholz 1977, Kulander &Dean 1995, Labuz et al. 1987, Lawn & Wilshaw 1975a, b,Lespinase & Pêcher 1986, Lin & Williams 1992, Lindquist etal. 1984, Logan et al. 1981, Maddock 1986, 1992, Maddocket al. 1987, Magloughlin 1989, Maltman 1994, Marone& Scholz 1989, Masch et al. 1985, Mase & Smith 1984,McEwen 1981, Means 1987, Menéndez et al. 1996, Meredith1983, 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, Rutteret al. 1986, Sammis & Biegel 1989, Sammis et al. 1987,Schofield & Worth 1968, Scholz 1972,1990, Seyedolali et al.1997, Shand 1916, Shimada 1986, Shimamoto & Nagahama1992, 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 DerEerden 1987, Borradaile et al. 1982, Brantley 1992, Carrio-Schaffhauser & Gaviglio 1990, Carrio-Schaffhauser et al.1992, Casey 1995, Cox 1987, Den Brock 1996, Den Brock& Spiers 1991, Dietrich & Grant 1986, Durney & Ramsay1973, Elliot 1973, Erslev & Ward 1994, Fisher & Anastasio1994, Fletcher & Pollard 1981, Gratz et al. 1991, Gray1977, 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 & Etheridge1977, Masuda & Mizuno 1995, McEwen 1981, McCaig1987, Murphy 1990, Onasch 1994, Passchier & Trouw 1996,Petit & Matthauser 1995, Powell 1979, Price & Cosgrove1990, Railsback & Andrews 1995, Raj 1982, Raj & Chyung1981, Ramsay 1980, Ramsay & Wood 1973, Ramsay &Huber 1983, 1987, Robert et al. 1995, Roedder 1984, Rutter1976, 1983, Schutjens 1991, Smith 1964, Smith & Evans1984, 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 4. Allison & LaTour 1977, Blenkinsop & Drury1988, Christie & Ardell 1974, Den Brock & Spiers 1991,

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Chapter 6. Arzi 1978, Ashworth & McLellan 1985, Bagnold1954, Brace & Martin 1968, Benn & Allard 1989, Blu-menfield 1983, Blumenfield & Bouchez 1988, Bouchez &Gleizes 1995, Bouchez et al. 1992, Burg 1991, Conolly etal. 1997, Copper & Kohlstedt 1982, 1984, Dell’Angelo &Tullis 1988, Dell’Angelo et al. 1987, Finney 1970, Gapais& Barbarin 1986, German 1985, Gray 1968, Guineberteauet al. 1987, Hibbard 1987, Hirth & Kohlstedt 1995, Hutton1988, 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 & Murase1984, 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 & Neumann1995, 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, Behr-mann 1984, 1987, Behrmann & Platt 1982, Bell 1985, Bell &Johnson 1989, Bell et al. 1992, Benn & Allard 1989, Berthéet al. 1979, Bjørnerud & Zhang 1995, Blenkinsop & Treloar1995, 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 & Holst1995, 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 etal. 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,

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, Robert1989, Robin & Cruden 1994, Rutter et al. 1986, Rykkelid& Fossen 1992, Scheuber & Andriessen 1990, Schmid etal. 1987, Shelley 1989b, 1995, Shimamoto 1989, Simpson1986, Simpson & Schmid 1983, Sylvester 1988, Ten Brink& Passchier 1995, Tikoff & Greene 1997, Tikoff & Fossen1993, Tikoff & Teyssier 1995, Turner & Wiess 1963, Van denDriesche & Brun 1987, Wenk et al. 1987, White & Wilson1978, White et al. 1980, Williams et al. 1994, Yoshinobu &Patterson 1996, Zee et al. 1985.

Chapter 8. Alexopoulos et al. 1988, Ashworth & Schneider1985, Bohor et al. 1993, Carter 1965, Carter et al. 1986,1990, Engelhardt & Bertsch 1969, Fel’dman 1994, Frenchet al. 1974, Gigl & Dachville 1968, Glass 1990, Goltrantet 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, Langenhorst1994, Langenhorst & Clymer 1996, Langenhorst & Deutsch1994, 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 etal. 1998, Robertson 1975, Robertson & Grieve 1977, Seye-dolali 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, Burkhard1993, Burov & Diament 1995, 1996, Busch & Van der pluijm1995, Byerlee 1978, Caristan 1982, Carter & Friedman 1965,Carter & Rayleigh 1969, Carter & Tsenn 1987, Carter etal. 1993, Christie & Koch 1982, Christie & Ord 1980, DeBresser 1988, De Bresser & Spiers 1990, Dennis 1984,Detournay et al. 1989, Dorn 1954, Drury 1993, Drury & Urai1990, Drury et al. 1985, Evans & Groshong 1994, Farver &Yund 199la, b, Fowler 1990, Gilletti & Yund 1984, Goetze& Evans 1979, Gratier & Guiget 1986, Gratier & Jenatton1984, Griggs 1967, Griggs & Blacic 1965, Groshong 1972,1974, 1984, Gross & Van Heege 1973, Hallam & Ashby1990, 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 & Blacic1983, Kronenberg & Tullis 1984, Kronenburg et al. 1990,Kruhl 1996, Kruhl & Nega 1996, Kumpel 1991, Kusznir1982, 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,

REFERENCES 125

Mercier 1980a, Mercier et al. 1977, Muller & Briegel 1978,Murrell 1963, 1965, Murrell & Digby 1970, Nemat-Nasser& Horii 1982, Newman 1994, Ord 1991, Ord & Christie1984, Ord & Hobbs 1985, Ord & Hobbs 1989, Paterson1978, 1987, 1995, Paterson & Luan 1990, Pavlis & Bruhn1988, Peng & Johnson 1972, Pfiffner & Burkhard 1987,Post & Tullis 1999, Power & Tullis 1989, Raj 1982, Raj &Ashby 1971, Raterron & Jaoul 1991, Rice & Gu 1983, Rosset al. 1980, Rowe & Rutter 1990, Rutter 1976, Schmid et al.

1977, 1980, Scholz 1990, Schulmann et al. 1996, Shelton etal. 1981, Sibson 1982, 1983, Spang & Van der Lee 1975, Spi-ers et al. 1990, Stoker & Ashby 1973, Takahashi et al. 1997,Terzaghi 1943, Tse & Rice 1986, Tsenn & Carter 1987, Tullis1980, Tullis et al. 1991, Turner 1953, Twiss 1977, 1986, Vander 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 10accretion steps 20activation enthalpy 92-96activation volume 94adhesive strength, wear 10alteration 9, 10, Plate 1Amonton’s law 10, 91amphibolite facies 24, 52, 96amygdale 22annealing 62, 82, 88anticrack 58antitaxial fibre growth 33, 37, 3.17, 3.18arrow method 102, 103asperities, asperity ploughing 10, 2.4, 20asterism 82

ballen 83, Plates 47, 48barrier theory 9bend 79bookshelf sliding 77botryoidal texture 32, 35boudin 19, 2.17, Plate 6Boussinesq configuration 12Brazil twin 81, 8.3, 85, 88, 89breccia 5, 6brittle 1,3brittle-ductile transition 4, 5brittle-plastic transition 3, 97Burger’s vector, 39, 4.1, 58Byerlee’s law 91, 97

calcite twin morphology 104cataclasis 1-3, 5, 7-24, 30, 38, 47,60,63,74cataclasite 6

foliated cataclasite 6, Plate 8protocataclasite 6ultracataclasite 6, 23

cataclastic flow 4, 18, 19, 63cathodoluminescence 2.6, 12, 17, 32,35, 81cement, cementation 17, 32chemical potential 24, 52chemical zoning 2, 57chessboard pattern 41, 4.10, 47Chixculub structure 84Cish diagram 37cleavage 2, 13, 28-30

classification 28, 29continuous 29crenulation 29, 30, Plates 16, 17disjunctive 29, 3.7domain 29, 30slaty 30, 3.9spaced 29, 30zonal 29

coefficient of friction, coefficient of internal friction 90, 91coesite 16, 80, 83, 8.7, 85-88coherent exsolution 52coherent transformation 52cohesion 5, 90, 91composite deformation mechanism 3, 52, 90composite fibre growth 37, 3.17constrained comminution 10contact melting 63contiguity 59core-and-mantle structure 4.14, 42, 49corona 2, 57, Plate 34crack-seal 35, 92crater 23creep 92-94

Coble 52, 92-94diffusion 52, 54, 60, 97dislocation 93-97grain size sensitive 92, 93Nabarro-Herring 52, 92, 94pressure solution 92, 93

critical melt fraction 59critical packing, packing density 59, 60critical slip distance 92cross-slip 39, 4.4crystal plasticity 4, 5, 19, 59, 104crystallites 22crystallographic fabric, preferred orientation 1, 2, 19, 22, 50,

51, 4.20, 58, 62, 75, 76, 7.19curved foliation 66

Dauphine twin 57debris streaking 20decussate texture 2, 54, Plate 28defect 52deflection surface 54, 57deformation bands 2, 5, 7, 18, 41, 4.6, 47, 50deformation lamellae 2, 14, 4.11, 47, 49, 101-103, 9.4

127

128 INDEX

deformation mechanism 1, 2, 3, 5, 90classification 1, 2map 90, 97, 98

deformation microstructure 1-3, 5, 90classification 1, 2

deformation (mechanical) twin 2, 39-41, 4.8, Plate 24, 62,81, 87, 102-104, 9.4, 9.5

deformation zone rocks 5deformation

continuity 4, 5, 18, 19,57distribution 4, 5, 18mechanism and mode 4, 5scale 4

dendritic crystals 22diaplectic glass 80-83, 85, 87, 89diffusion 24, 29, 53, 57, 82, 94-94

coefficient 52, 92-94creep - see creepgrain boundary 93volume 93

diffusive mass transfer 1-3, 47, 63, 90, 92by solution 3, 24-38in melts 60solid state 52-58

dihedral angle 24dilatancy hardening 60dislocation 1, 30, 39, 4.1, 4.2, 4.3, 4.4, 47

climb 39, 93creep 39, 5.1, 54, 93-97density 39, 41, 47-49, 58, 81, 82glide 39, 51,93-95

displaced grain fragments 79, 7.23, Plate 45displacement control 33displacive transformation 52divider method 105Dorn law 93ductile stringer 19, 2.17ductility 4, 5dynamic recrystallization 47-50, 4.19

effective stress 91Einstein-Roscoe equation 59enclave 60erosional sheltering 20etching, etch pit 39, 88exaggerated grain growth 54exponential law 93-97extensional crenulation cleavage 70external asymmetry 76, 7.19

fabric skeleton 76face control 33, 38, Plate 19failure criteria 90, 91, 97

Coulomb and Mohr failure criteria, 90, 97, 98and fracture mechanics 91Griffith criteria 91,97

fault gouge - see gougefault breccia - see brecciafibre 32, 33, 37, 65

fibre strain 17growth histories: antitaxial, ataxial, composite, non-

systematic, syntaxial 37, 3.17, 3.19non-tracking 38, 3.19tracking 37, 38, Plate 23tracking efficiency, 38, 3.19

Fick’s law 24, 52finite shortening, extension 25flow laws 90, 92-97fluid inclusions planes 2, 12, 24, 2.5, 33-35, 3.13, 3.14, 47foam texture 2, 54fold 4, 19, 2.17, 60, 63, 70foliation 54-57, 65, 75

curved 66, Plate 39external foliation 54, 55, 5.3, Plates 29-33internal foliation, 54, 55, 5.3, Plates 29, 31-33oblique foliation, 66, 67, 7.3steady-state, strain insensitive 61strain sensitive 66

fractal dimension 10, 22, 92, Plates 1-4fracture 1fracture toughness 9fragmentation 10frictional sliding 7, 10, 2.4, 90-92, 97

generation plane 22geotherm 97geothermobarometry 104, 105gouge and gouge zone microstructures 5, 6, 10, 17, 2.17,

Plate 6, 72grain boundary migration 4.12, 47, 4.17, 48-50, 4.18, 4.19,

95, 99, 100grain boundary width - fast or free 54grain dispersive pressure 61grain shape fabric 2

cataclastic 19, 33diffusive mass transfer 33, 3.11, 52, Plates 26, 27, 54intracrystalline plastic 33, 4.12, 47magmatic 62, Plate 36sub-magmatic 62

grain size sensitive creep 92, 93grain surface deposition textures 2, 30grain surface solution textures 2, 25granoblastic polygonal texture 32, 54, 62growth twins 39, 62, 81growth zoning 57, 83, Plate 18

helicitic texture 54Hertzian configuration 12homologous temperature 97hydrolytic weakening 9, 94

igneous zoning 62imbrication 60, 62, 78inclusion 35, 3.15, 3.16, 62, 84

band 35, 3.15, 3.16trail 35, 3.15, 3.16, 37, 38, 54, 65

indenting grain contacts 2, 25, 3.1, 3.2, 33, 63independent particulate flow 2, 22

INDEX 129

instantaneous stretching axis, axes 65, 73, 7.15, 7.17interconnected weak layer 96internal asymmetry 76, 7.19internal strain energy 4, 48, 52, 55interpenetrating grain contacts 2, 25, 3.1, 3.4, 33intracrystalline plasticity 1, 3, 4, 14, 19, 22-24, 30, 39-51, 57,

60, 62, 63, 75, 90, 93, 97intracrystalline deformation bands 2, 41island and channels 24isostrain criterion 96isostress criterion 96

jog 79

kink, band 2, 14, 30, 41, 4.6, 4.7, 4.8, 50, 62

landslide 23lattice preferred orientation 58, see crystallographic fabriclechatelierite 80-85lineation 66, 7.1, 7.2, 75lithospheric strength envelope 90, 97, 98, 9.2load bearing framework 96low angle boundaries 41lower stability limit 92

magmatic flow 3, 59-63, Plate 36magmatic microstructures 59, 62magmatic shear zones 62, 77, 78magnetic anisotropy 60, 62Martensitic-like transformation 52, 57mean stress 94megacryst 62metatextite 63mica beard 32mica fish 74, Plate 42microboudins, asymmetric 76, 7.20, 7.21, Plate 6microcrack 2, 7, 2.1, 2.5, 2.6, 10-16, 33

axial 12, 17, 2.14, 19, 63,91characteristics and observation 10-12circumgranular 10-12classification 10-12cleavage 13, 19cone 12dynamic propagation 7elastic mismatch 13, 14, 2.11, Plate 5en-echelon 17, 2.15en passant 17, 2.15extension 12, Plate 2extension force 7flaw-induced 2.6, 13, 2.10grain boundary 13impingement 2.6, 12, 13, 2.7, 2.8, 2.9intragranular 10, 2.5, 12, 13, 2.9, 18, 19microfault- induced 14, 19microscopic feather fracture 14, 15, 2.12modes I, II, III 7, 2.2, 12phase transformation-induced 16plastic mismatch 14, 2.12, 2.13refracturing 13, 2.11

sub-critical propagation 9, 94thermally-induced 15, 16transgranular 10, 2.5, 12, 13, 16wing 13, 2.10

microcrystallite 81, 82microfault 2, 7, 17, 18, 2.14-2.16, 7.23, Plates 3, 45microfold 77, 7.22microfracture 1, 2, 4, 7, 80, 85, 86, Plate 4

mirror 20mist 20surface features 2, 19, 20velocity hackle 20Wallner line 20

microlithon 29, 30microphenocryst 22microslickolite, 27, 28, 3.6microstylolite 1, 2, 27, 28, 3.5, Plate 14, 3.6

characteristics 27filling 24-28formation and propagation 27, 28teeth, walls, crown 27, 3.5, 28

microvein 1, 2, 7, 35-38free-face growth 35fibrous 33, 35, 37, 38laminated 35opening vector 35

migmatite 63millipede texture 54mimetic crystallization 54mode of failure 4molar entropy 24molar internal energy 24molar volume 24, 92-94mosaicism 81-83, Plate 48, 85-89mylonite 5, 6, 62

protomylonite 6ultramylonite 6

necking down 33, 35, 3.14neosomes 63new grains 47, 50non-magmatic deformation 63, 64non-systematic fibre growth 37, 3.17, Plate 22normal slip crenulation 70

oblique grain shape fabric 78ophitic texture 62order-disorder transformation 52Ostwald ripening 54overgrowth 1, 2, 4, 13, 32, 33, 3.10, Plate 18, 3.11, 4.11

P-foliation 19, Plate 6, 79P-shear, fracture 19, 20paleopiezometry, palaeopeizometer, 98-101

deformation lamellae 101-103dislocation density 100, 101general problems 104maximum twin volume, 101principal stress and strains 103, 104, 9.4

130 INDEX

recrystallized grain size 98-100, 9.3subgrain size 100, 101twinning - differential stress 100, 101twinning density, 101twinning incidence, 101

particle size distributions 10, 22, 92perlitic texture Plates 47, 48permeability 91phase transformation microstructures 2, 57, 58, 80-92phenocryst 60-62phenomenological coefficient 93phyllonite 6pinching off 33, 3.14planar deformation feature 47, 8.2, 80-89

decorated 81,82non-decorated 81, 82sub-lamellar structure 81, 88

plastic, plasticity 3, 14poikiloblast 54Poisson’s ratio 97polymineralic deformation 94-97pore fluid, pressure 3, 22-24, 91porosity 12, 13, 18, 87, 88, 93

reduction 19, 32porphyroblasts 1, 2, 54-57

characteristics 54, 55growth mechanisms 55internal, external foliations 54, 75, 7.18intertectonic 55-57, 5.3, 5.4, Plate 31plate 33posttectonic 55-57, 5.3pretectonic 55-57, 5.3, Plate 30relationship to deformation 55-57, 5.3shear sense indications 75, 7.18syntectonic 55-57, 5.3, Plate 32

porphyroclast, porphyroclast systems 5, 14, 67-70characteristics and classification 67, 68, 7.4complex type 67, 7.4, 7.8

67-70, 7.4, 7.7, 7.9, 7.10, 74deflection, embayments 68-70, 7.7, 7.9faces of a tail 69, 7.9in-plane 67, 68

67, 7.4mantle 67mechanisms of formation 68, 69

67-69, 7.4, 7.5, 7.6, 7.9stair-step 67, 68tail, wing 67

power law 93-97breakdown 94

pre-exponential constant 93-97pre-lithification deformation 2, 22pressure shadows and fringes 1, 2, 4, 32, Plate 19, 33, 3.12,

63, 65, 73,74last increment of growth 73,kinematics in shear zones 73, 7.15shape 73, 7.16, 7.17

process zone 9, 28prod mark 20, 2.18

pseudotachylites 2, 5-7, 22, 23, Plates 9-11, 83, 92characteristics 22misidentification 23origin 22, 23

rate and state dependent frictional sliding 92reaction rims 2, 57reaction zoning 57recovery 2, 41, 47, 82recrystallized grain size 49, 50reentrant zoning 54relict minerals 2, 57rheology 90rhomb dodecahedra 54ribbon grain 2, 4.13, 52-54, 5.1ridges and grooves 20Riedel, conjugate Riedel shear 2.17, 19, Plates 6, 7, 20, 72,

79fractures 15

risers 20, 2.19, 79congruous/incongruous 20, 2.19

rolling structures 77roughening transition 32

S-, C- and 5, 20, 62, 63, 66, 70-73, Plates 40, 41characteristic and classification 70-72, 7.11, 7.13, 7.14curvature of S-foliation 73formation and evolution 72shear on C-or 19, 73shear sense from 73

scaly clays 22scanning electron microscope 2.6, 19, 2.18, 2.19, 25, 30, 50,

79, 82, 88schlieren 60secondary recrystallization 54, 55sector zoning 54semibrittle 3, 14, 19sensitive tint plate 2.6, 12, 17, 50, Plates 1-4, 45, 79separatrix 68, 69shape preferred orientation 47, 66shatter cones 23shear band 70shear modulus 97shear sense 65-79

criteria 65-79for faults 79, 7.23, Plate 45in rocks containing melt 77-79observation plane 65, 71

shock-induced microstructures 3, 80-89, 8.1shock mechanisms and metamorphism 23, 80shock wave barometry and thermometry 85-88

calibration 85-87problems 87, 88

slickenfibre 20, 2.19, 79slickenline 20slickenside 20, 92slip 39

direction 39, 51plane 39, 4.1, 51

INDEX 131

systemsnowball texture 54, Plate 29solid state transformation 52-58, 80-82solubility 24, 92-94spherulites 22, 83, 84spin 65state variable in frictional sliding law 92static recrystallization 54, 62stick-slip behaviour 92stishovite 80, 83, 8.8, 85-87strain cap 2, 27, Plate 13strain rate 90, 92-97strawberry texture 82, 8.5, 8.6stress 1

amplification 97corosion 9exponent 93-97intensity factor 7, 9

stretched crystal fibre 37, 3.17strewn field 84sub-boundary migration 50sub-magmatic flow 3

microstructures 3, 59-63, Plates 37, 38, 78, 79subgrain 1, 2, 14, 19, 32, 41, 47, 4.9, 4.13, 47, 4.16, 50, 62

boundary orientation in quartz 41, 63, 105rotation 47, 4.15, 4.16, 48, 49, 4.19, 50, 99, 100size 1, 47

superplasticity 3, 52, 57, 58, Fig. 5.5, 94surface tension force, energy 7, Plate 27, 54, 55sutured grain boundary 25, 50, 3.1, 4.19, 104, 105, 9.6, 9.7symplectite 2, 57, Plate 35syntaxial fibre growth 33, 37, 3.17, Plate 21

T fracture 19, 2.17, 20, 79tectite, microtectite 83, 84

tiling 60, 78tool track or mark 20, 2.18transmission electron microscope 19, 39, 4.2, 81-83, 88, 101transpression 65, 7.2triple grain junction 54, 5.2truncating grain contact 2, 25, 3.1, 3.3, Plate 12, 30, 33truncation surface 30, 3.7, 54, 55twin, twinning 47, 57, see also deformation (mechanical)

twin, growth twin, paleopiezometry

undercutting mechanism 24undulatory extinction 1, 2, 33, 41, 4.5, 4.6, 4.13, 49, 62, 82uniaxial compressive strength 90, 91uniaxial tensile strength 90, 91upper stability transition 92

velocity weakening, strengthening 92vesicle 84viscosity 59-61volume loss 29, 3.8, 30vorticity 65

external 65internal 65profile plane 65, 7.1shear-induced 65vector 65, 7.1

wear groove 20, 2.18whole lithosphere failure 97

Y-shear Plate 7, 19, 2.17, 72Young’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 rotationor displacement of the fragments. Alteration of the feldspar to more sodic compositions (and ultimately to laumontite) occurs along the microcracks, showingthat 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, demonstratingextension in the plane of the section. The microcrack is filled by angular fragments of quartz, biotite and feldspar in random orientations which can not bematched 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 beenrotated 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 constrainedcomminution 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 largefeldspar 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 inextinction, compared to the birefringent host grain elsewhere. The strained area has the same geometry as the area of high stress predicted from modellingthe 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 P-foliation which trends in the opposite direction. The two features surround an asymmetrical boudin. All three features give a sinistral shear sense. Comparewith Figs. 2.17, 7.20. (Chapters 2.7, 7.13.4).PPL, 2 mm, Tarskavaig thrust zone, Tarskavaig, Isle of Skye, Scotland, U.K.


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 Riedelshears 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 microfracturesat 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 materialparallel 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 ofsizes. All these features are typical of pseudotachylite. See Plates 10, 11. (Chapter 2.11).PPL, 2 mm, Pseudotachylite in grey gneiss mylonite, Buhwa, Zimbabwe.


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 finegrained, moderately birefringent phyllosilicate picks out a weak foliation in bands which are concave towards the apex of the veinlet. These bands suggest aprimary 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 weakfoliation 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 aflow 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 planarsurfaces. 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.


Plate 13. Strain cap. The porphyroclast of feldspar at centre has foliae of muscovite wrapped around the top and bottom surfaces. Flakes of muscovitedistributed sparsely through the matrix around the porphyroclast define a weak foliation. The presence of muscovite in the matrix suggests that the muscovitewas 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 carbonatesalong the stylolites. Shell fragments are truncated by the stylolites. The amplitude of the sinusoidal stylolite suggests approximately 0.3 mm of shorteningperpendicular to the stylolite trace. (Chapter 3.6).PPL, 2 mm, Carboniferous limestone, Somerset, U.K.


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 thefold limbs, defined by muscovite and chlorite alignment parallel to the axial surfaces of the folds, and by concentrations of opaque minerals which are foundelsewhere throughout the rock, suggesting that diffusive mass transfer has contributed to the formation of this cleavage by removal of the more solublephases 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 alignedbiotite 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 biotitelayers, 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 ofthe 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 beenalmost 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 concentricgrowth zones shown by different colours disrupted by solution. The complex zonation patterns show variations in fluid chemistry during growth. (Chapter3.9).SEM-CL, 0.9 mm. Hydrothermal quartz vein cross-cutting the Ventersdorp Contact Reef, Klerksdorp goldfield, Witwatersrand, South Africa.


Plate 19. Pressure fringe. Fibrous quartz overgrows euhedral faces of pyrite crystals. The quartz fibres are perpendicular to several faces, showing a typicalface-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 quartzlines 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. Thecalcite 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 meetsthe 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 grainsize 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 butundeformed nature of the fibres shows that the curvature is a primary growth feature. Since the fibres grow towards a median suture, wall rock control oftheir orientation is unlikely, and the fibres probably tracked the incremental opening vector, which rotated with respect to the vein. See Fig. 3.17a. (Chapter3.12).XPL, 4.3 mm, Quartz-calcite vein, locality unknown.


Plate 22. Non-systematic (stretched) fibres. Individual quartz fibres stretch all the way across the microvein. There is no symmetry or suture, and the sidesof the fibres have interlocking teeth. They are subdivided into tablets perpendicular to their length, each of which represents a growth increment. These arethe 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 marginorientation 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.


Plate 25. Grain boundary migration. Advanced stages of grain boundary migration. As well as highly convoluted grain boundaries, blue grains can be seenentirely 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 diffusivemass 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 quartzgrains are perpendicular to muscovite basal planes, and often pinned at the ends of the muscovite grains (centre). Both these effects are due to the highersurface energies between quartz and mica than between quartz grains. The resultant quartz grains are elongate parallel to the muscovite fabric. Also seePlate 26. (Chapter 5.3).XP, 2 mm, Mica-garnet schist, Miami district, Zimbabwe.


Plate 28. Decussate texture. Randomly orientated, interlocking, elongate actinolite crystals show a good example of this texture, produced by diffusive masstransfer 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 isdefined 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 biotitewraps 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.


Plate 31. Intertectonic porphyroblasts. The two biotite porphyroblasts contain an internal fabric of relatively large quartz grains. The external fabricis 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 diagnosticfeatures 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 and theexternal 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.


Plate 34. Corona. The central hornblende grain is surrounded by an intergrowth of orthopyroxene, plagioclase and magnetite, a typical texture produced byprograde 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 occursbecause diffusion distances during the reaction were too small to allow an equilibrium microstructureto 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 fromupper right to lower left. The phenocrysts are completely unstrained and set in a matrix of pyroxene. Undeformed igneous zoning can be seen in the largestphenocryst. No evidence for deformation is seen in the hand specimen, which also shows the alignment of euhedral phenocrysts. The complete lack ofmicrostructural evidence for strain demonstrates that the fabric is magmatic. (Chapter 6.4).XPL, 2 mm, Basalt, Bembezi river, Zimbabwe.


Plate 37. Magmatic/submagmatic grain shape fabric. Euhedral plagioclase and biotite crystals define a grain shape fabric parallel to the top/bottom of theplate. The only other evidence for deformation is slight undulatory extinction in quartz. Primary, undeformed zoning is visible in the dark plagioclase crystalon 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 thequartz, 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 feldsparmegacryst alignment is a conspicuous feature of the outcrop from which this specimen was collected. The feldspar megacryst is recrystallized along itsmargins, 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 fabricand 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 tothe 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.


Plate 40. fabric or extensional crenulation cleavage. A penetrative S-foliation from upper left to lower right is defined by hornblende grains. S isdeflected 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 quartzgrain shapes and compositional banding. Discrete shears parallel to the top and bottom edges are C-surfaces. A few discrete shear surfaces are inclined tothe 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 (SeeFigs. 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 appearto have formed by microfracture along basal cleavage planes. The basal cleavage is parallel to the long axis of the fish. Recrystallization occurs along theleft 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 shearplane (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.


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 fromthe 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 arepressure 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, suggestingrotation 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


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 typicalplanar, 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. Theright 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 individualballen are filled by a birefringent phase, probably a clay mineral, suggesting that the ballen structure formed in a similar way to perlitic texture by hydrationof 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 theevidence 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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