Deformation Micro Structures and Mechanisms in Minerals and Rocks

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Deformation Microstructures and Mechanisms

in Minerals and Rocks

by

Tom Blenkinsop

Department of Geology,

University of Zimbabwe, Harane Zimbabwe

KLUWER ACADEMIC PUBLISHERS

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eBook ISBN: 0-306-47543-X Print ISBN: 0-412-73480-X

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Acknowledgements

Most of the photomicrographs were developed and printed by Cuthbert Banda, whose assistance, with that of other members of the staff of the Geology Department, University of Zimbabwe, was invaluable. James Preen guided the preparation of the TEX version of the text. Faith Samkange and Maxwell Matongo were able research assistants, supported by the University of Zimbabwe Research Board. The following are thanked for contributing photomicrographs or thin sections: P. Dirks (Plates 11, 30, 31, 32, 33, 35, 44), R. Fernandes (Fig. 2.9), H. Frimmel (Plate 18), S. Kamo (Figs. 8.4 - 8.6), H. Leroux (Figs. 8.2, 8.3), J.E.J. Martini (Figs. 8.7, 8.8), U. Reimold (Plate 46), J. Stowe (Plate 22), D. Van Der Wal and M. Drury (Fig. 4.2). Plates 1 - 4 and Figs. 2.5, 2.18, 2.19, 3.18 of core material from the Cajon Pass drillhole were made at the Institute for Crustal Studies, University of California, Santa Barbara, as part of research on deformation mechanisms with R. Sibson, supported by the National Science Foundation, U.S.A., under grant DAR-84-10924. The assistance of the technical staff at U.C.S.B. is gratefully acknowledged. Some research for this book was supported by the IUGS Commission on Tectonics, COMTEC.

Detailed reviews of chapters from the following are greatly appreciated: P. Dirks, R. Hanson, H. Jelsma, W. Means, A. Ord, M. Paterson, U. Reimold, E.H. Rutter, A. Schmid Mumm.

Fig. 8.2 was reproduced from Leroux et al. (1994), and Fig. 8.7, 8.8 from Martini (1991), all with kind permission of Elsevier Science - NL Sara Burgerhartstraat 25, 1055 KV Amsterdam, The Netherlands. Fig. 8.4 was reproduced from Krogh et al. (1996) with kind permission from the American Geophysical Union. Fig. 9.5 was reproduced from Burkhard (1993), and Fig. 7.20 from Goldstein (1988), both with kind permission from Elsevier Science Ltd. The Boulevard, Langford Lane, Kidlington OX5 1GB, U.K. Figs. 9.6, 9.7 were reproduced from Kruhl and Nega (1996), with kind permission from Springer Verlag. Full details of these publications are given in the references.

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Symbols, Abbreviations and Units

Numbers in brackets give the chapters and sections where the symbols are used. Units are given if they are referred to in the text. a A B c CL CMF d D DMT E GBM G ISA J kC l L LPO LSE m n P PDF

Microcrack long axis (2.2.1), Rate and state variable friction law constant (9.3.2)

Flow law constant, (9.4.1), or (9.4.3)

S to C-surface angle (7.5), Flow law constant (9.4.3), Effective stress coefficient (9.2.3) Microcrack short axis (2.2.1), Flow law constant (9.4.3)

Burger’s vector (4.2, 9.8)

Particle velocity/unit potential gradient (3.2), Flow law constant (9.4.3) Microcrack or flaw length (2.2, 2.3, 2.4.2), concentration of particles (3.2) Cohesion, MPa (9.2, 9.3), Reference state solubility, mole fraction (9.4.1) Angles defining external asymmetry of an LPO (7.9)

Cathodoluminescence (2, 3, 8) Critical melt fraction (6.2)

Particle size (2.2.2), Flaw spacing (2.4.2), Grain size, m (9.4), mm (9.8) Subgrain size, mm (9.8.3)

Fractal dimension (2, 9.9.3)

Reference state diffusion coefficient for Pressure solution creep, (9.4.1) Reference state diffusion coefficient for Nabarro-Herring, Coble creep, (9.4.1) S to C- or angle (7.5), Grain boundary width, m (9.4)

Diffusive mass transfer (1, 3, 5, 9) Twinning density, (9.8.5) Young’s modulus (2.2)

Strain rate, strain rate at 0 K in the Dorn Law, (9.4) Volume fraction of phase in polymineralic flow law (9.5) Crystal fraction, fraction at critical packing (6.2.1) Microcrack extension force, Critical value (2.2) Grain boundary migration (4.8, 9.4, 9.8)

Surface tension force (2.2), Angle between flaw and (2.3.4) Fracture toughness (2.2)

Viscosity of suspension, Viscosity of pure melt (6.2.1) Instantaneous stretching axes (7.1, 7.7)

Twinning incidence (9.8.5) Diffusive flux (3.2)

Angle between flaw array and (2.4.2)

Dislocation density palaeopiezometer constant (9.8.4) Stress intensity factors for mode I, II, and III opening (2.2) Threshold, Critical stress intensity factors (2.2)

Velocity of deforming crystal face, (9.4.1)

Model microcrack length (2.3, 2.4.2), Dislocation density palaeopiezometer constant (9.8.4) Critical slip distance (9.3.2), Length of grain boundary (9.9)

Lattice preferred orientation (4.9, 5.10) Lithospheric strength envelope (9.7)

Grain size exponent in flow law (9.4, 9.5), Grain size palaeopiezometer constant (9.8.2) Coefficient of friction (2, 9.2), Chemical potential (3.2), Shear modulus, GPa (9.8.3, 9.8.4) Coefficient of internal friction (9.2)

Power law exponent for flow laws (9.4, 9.5) Mean stress, Pa (3.2, 9.4)

Planar deformation feature (7)

xi

b

f

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xii SYMBOLS, ABBREVIATIONS AND UNITS PPL r R RF S S SEM ST TEM u U V () { } < >

Pore fluid pressure (3.2, 9.2.3) Plane polarized light

Flaw to microcrack angle (2.3.4, 2.4.2)

Activation enthalpy, pressure solution and grain boundary diffusion, (9.4) Activation enthalpy for volume diffusion, and diffusion creep, (9.4) Length of stride in divider method (9.9.3)

Gas constant, (9.4)

Reflected light

Density, (9.4.1), Dislocation density, (9.8.4) Deformation lamellae spacing, mm (9.8.6)

Molar entropy (3.2) Cohesion, MPa (9.2)

Oblique foliation, Bands parallel to shear plane (7.3) External, Internal foliations (5.6, 7.8)

Scanning electron microscope Sensitive tint plate

Differential stress, Pa or MPa (9.4)

Maximum, intermediate, minimum principal stresses, Compression positive (2,3,5) Remote stress for microcrack closure (9.2.2), Remote applied stress (2.2.1) Normal stress (2,3, 9.2), Stress at 0 K in Dorn Law (9.4)

Temperature,0C and K, Melting temperature Uniaxial tensile strength, MPa (9.2.2) Transmission electron microscope Shear stress, MPa (2.2.2, 9.2, 9.3)

Dislocation density exponent in dislocation density palaeopiezometer (9.8.4) Molar internal energy (3.2)

Subgrain size exponent in subgrain size palaeopiezometer (9.8.3)

Microcrack velocity (2.2.1), Sliding velocity (9.3.2), Activation volume, (9.4.3) Molar volume, (3.2, 9.4)

Maximum twin volume, % (9.8.5)

Thickness of a grain boundary fluid (9.4.1), Grain size palaeopiezometer exponent (9.8.2) Angles defining internal asymmetry of an LPO (7.9)

Crossed polarized light

State variable in dynamic frictional sliding law (9.3.2) Miller-Bravais indices of a crystal face

Miller-Bravais indices of a representative face of a form Miller-Bravais indices of a crystal direction

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Introduction and Terminology

1.1 Introduction

Deformation microstructures in rocks and minerals are

mi-croscopic features created by deformation. A deformation

mechanism is a process on one scale that accommodates an

imposed deformation on a larger scale. This book describes microstructures 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. The book aims to be a guide to the recognition and interpretation of deformation microstructures and mechanisms, and should be used in conjunction with a petrographic microscope.

Why is the study of deformation microstructures and mechanisms useful ? Deformation mechanisms are determ-ined by temperature, stress (both hydrostatic and deviatoric components), 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. For example, microstructures called subgrains are formed by the deformation mechanism of intracrystalline plasticity, which indicates 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 of information for tectonic analysis and for understanding the behaviour 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 of materials science to minerals and rocks has provided much of the new impetus. However, the literature is scattered through journals in a large number of disciplines, and most structural geology textbooks have limited coverage of the field. At least some of this diverse literature is reviewed in this book, which contains a comprehensive reference list. Hopefully the book is 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 to elastic (recoverable) deformation which is not seen directly in the geological record. Deformation microstructures can be divided into three major categories:

1.

2.

3.

Microfractures, displacement, and/or rotation of rigid particles 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, transport and deposition without fracturing, permanent lattice dis-tortion or melting. Examples of microstructures at sites

of material removal are distinctive types of grain con-tacts and microstylolites. Microstructures indicating material deposition include microveins, overgrowths, pressure shadows and porphyroblasts. Many meta-morphic textures associated with deformation fall into this 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 basis of optical microscope observations. Table 1.1 summarizes the scheme, gives examples of specific microstructures, and refers to the relevant chapters and sections of the book.

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

1.

2.

3.

Cataclasis - deformation by microfracturing, sliding

and/or rotation of rigid particles (Chapter 2). Brittle de-formation is often used as a synonym for cataclasis, but is more accurately defined as deformation by fracture or microfracture.

Diffusive mass transfer (DMT) - deformation by

diffu-sion, the movement of lattice defects, ions, atoms or molecules 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 a crystal lattice (Chapter 4).

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A general term for both categories 2 and 3 is useful because they are often closely associated with each other: crystal plasticity or simply plasticity means deformation by either or both DMT and intracrystalline plasticity. DMT can be split into two major sub-divisions: diffusion via a solution (pressure solution, Chapter 3), and diffusion in the solid state (Chapter 5). Solid state phase transformations may occur during deformation, some involving DMT: these are included in Chapter 5.

Two or more mechanisms may act simultaneously together within a single mineral. An example is the combination of sliding on grain boundaries, and mass transfer of material by diffusion to fill space created by sliding. This is is one type of superplasticity, which is a composite deformation

mech-anism; one in which two or more mechanisms are coupled

together in the same mineral under the same conditions. An important composite deformation mechanism involves the in-teraction of microfractures and intracrystalline plasticity: this is called semibrittle deformation. The different components of a composite mechanism may have variable strain rates: the mechanism with the slowest rate determines the composite strain rate, and is said to be rate-limiting.

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

Three chapters deal with relatively new developments in structural geology, which can involve all three categories of deformation microstructures and mechanisms. Magmatic

and sub-magmatic deformation (Chapter 6) describes the

de-formation of rocks that contain melt. Important dede-formation mechanisms 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 to

the displacement of rock masses during deformation on all scales: this is one of the most important types of tectonic information. Shock-induced microstructures and shock

meta-morphism are produced by meteorite impacts (Chapter 8), and

are the focus of much current interest, because microstruc-tural studies have a central role to play in the debates about mass extinction and other possible effects of large meteorite impacts on the earth’s evolution.

1.3 Deformation

microstructures and

mechanisms 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 to allow 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 minmin-eral within a polyminmin-eralic rock: a common ex-ample is the intracrystalline plasticity of quartz in a shear zone 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 other mechanisms are substituted or added during the same de-formation: for example, faulting (cataclasis) may relieve high stress 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. Mechanisms that can only accommodate deformation in a single direction, for example slip on a single fault set, are especially restrictive and invariably require supplementary mechanisms.

The variation in conditions in the earth, particularly of tem-perature and pressure, causes a corresponding variation in de-formation mechanisms. Two systematic changes occur with increasing depth: temperature increases by the geothermal gradient, and pressure increases due to the effect of gravity. Cataclasis is relatively insensitive to the variation in temper-ature, but is suppressed by pressure. The plastic deformation mechanisms (intracrystalline plasticity and solid-state DMT) behave in the opposite way: they are strongly promoted by temperature but relatively insensitive to pressure. As a res-ult, cataclasis is generally restricted to the upper crust (where pressures are low), and crystal plasticity occurs in the lower crust and the rest of the lithosphere (where temperatures are high). The transition from cataclasis to plasticity is some-times called the brittle-plastic transition. Experiements have identified an important intermediate regime of semibrittle be-haviour where microfractures interact with intracrystalline plasticity (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 the semibrittle - plastic transition in mechanisms at mid-lower

crustal levels.

These generalizations need to be heavily qualified because of the other variables that affect deformation mechanisms. Stress, strain rate and pore fluids are some of the most im-portant additional variables to be considered. For example, high stresses or strain rates may cause cataclasis at greater depths or higher temperatures, which would otherwise be as-sociated with plasticity, as in the above example. Pore flu-ids are essential for solution-assisted DMT, and may promote cataclasis 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 to be considered as an independent variable (Chapter 9). These qualifications, particularly the variation in properties of dif-ferent minerals, mean that no unique depth or temperature can be given for the brittle-semibrittle and semibrittle-plastic transitions. However, the transitions are important concepts

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4 CHAPTER 1. INTRODUCTION AND TERMINOLOGY

for understanding the behaviour of the earth, and the condi-tions for the transicondi-tions can be predicted for specific models of the earth, as described in Chapter 9.

1.4 The

description

of

deformation:

Scale, continuity, distribution,

mechanism and mode

Deformation of minerals or rocks should be described in terms of three fundamental attributes. Continuity is the con-nectivity of material points through the deforming body; de-formation can be characterized as continuous if points remain connected, or discontinuous if not. The distribution of de-formation can be described as localized (e.g. a shear zone) or distributed. The deformation mechanism can be described in one of the three categories given in Section 1.2. All three at-tributes depend on the scale of observation. Scales are loosely defined in this book as macro (greater than outcrop, i.e. > 10 m), meso (outcrop, i.e. 1 cm - 10 m) or micro (microscopic, i.e. < 1 cm).

Permanent deformation of minerals or rocks involves breaking atomic bonds and is therefore is discontinuous at the atomic scale. As the scale of observation is increased, these atomic discontinuities can not be discerned, and deformation appears to be continuous. The top surface of Fig. 1.1 shows schematically how deformation continuity is a function of the scale of observation. An example of the discontinuous to con-tinuous transition with increasing scale is the accommodation of a fold by cataclastic flow (Chapter 2.6). At micro to meso scales, the deformation is by discontinuous fracture, but at a

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

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

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

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

Figure 1.1 makes some important links between 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 themes of this book: deformation microstructures can be used, al-beit with care and a certain amount of ambiguity, to identify deformation mechanisms. This link is possible by induction and deduction from the microstructures and theoretical un-derstanding of the mechanisms, and by comparison with ex-periments.

1.5 Ductility and the “brittle-ductile

transition”

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

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

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

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

1.6 Character

and

classification of

de-formation zone rocks

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

Some definitions and classifications of deformation zone rocks have advocated the use of deformation mechanisms as a classificatory tool. For example, Wise et al. (1984) pro-posed to include crystal plasticity as an essential element of the definition of mylonites. As for any other classifica-tion of observaclassifica-tional data, the use of mechanisms should be avoided because subjective interpretations are required, that may change in the light of new knowledge. Nomenclature and classification of deformation zone rocks is extensively discussed in Snoke et al. (1998).

The non-genetic classification scheme of Sibson (1977) is

the most widely used today. The scheme is based on the dis-tinction between random fabric and foliated deformation zone rocks, as well as the cohesion of the rock (the degree to which it behaves as a continuous body during deformation) and pro-portion of matrix to porphyroclasts. The original scheme is slightly modified to allow for a range of foliation intensities in fault rocks in Table 1.2, by including the additional category of foliated cataclasites (e.g. Chester and Logan 1987), and by extending the gouge, breccia and pseudotachylite categor-ies to the foliated category. Random fabric has been changed to Unfoliated to allow for deformation zone rocks that have some order to their fabric, for example fragments with a jig-saw texture, and Foliated has been changed to Strongly fo-liated in order to contrast with Unfofo-liated. Many deforma-tion zone rocks have two or even three distinct foliadeforma-tions. All foliations should be considered when assessing foliation in-tensity, and the existence of different foliations can be used to further classify deformation zone rocks (e.g. S-C mylon-ites, Lister and Snoke 1984). The crush breccia series of the original 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 larger fragments in a finer grained matrix is a fundamental character of most deformation zone rocks. However, as noted by Sibson (1977), any pidgeon-hole classification such as Table 1.2 suf-fers from the problem that the classification criteria may show a continuous range of variation. This is especially problem-atic in the assessment of foliation development, which may be difficult to quantify, or to judge objectively on a qualitative basis. Another problem with the classification is the defini-tion of cohesion, which is specified in the original classific-ation as primary cohesion, i.e. cohesion during deformclassific-ation. Post-tectonic processes may decrease cohesion, for example by weathering, or increase cohesion by cementation. These processes may not be recognized easily or allowed for when assessing primary cohesion.

1.7 Format and use of this book

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

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6 CHAPTER 1. INTRODUCTION AND TERMINOLOGY

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

A general list of symbols, abbreviations and units prefaces this chapter. Abbreviations and conventions for all 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 of the image in mm. Shear sense (see Chapter 7) is given as sinistral or dextral (assuming that the shear or fault plane is vertical); all illustrations with shear senses given are perpen-dicular to the shear plane and parallel to the shear direction, and the shear plane is approximately parallel to the horizontal edge of the image.

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Cataclasis

2.1 Introduction

Microfractures, displacements and rotations of rigid particles with no permanent lattice distortion are microstructures formed by cataclasis. There are two fundamental cataclastic deformation mechanisms: microcracking (Section 2.2.1) and

frictional sliding (Section 2.2.2). The single most

distinct-ive cataclastic microstructure is the microfracture, defined as a tabular or planar microscopic discontinuity. The term thus includes microfaults, microscopic deformation bands, microjoints, microcracks, microveins, and microscopic slip surfaces. 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 are

briefly discussed in Section 2.11 because they are associated with cataclastic mechanisms.

2.2 Fundamental

cataclastic

deforma-tion mechanisms

2.2.1 Microcracking

Microcracking involves microcrack nucleation followed by propagation. Nucleation is irrelevant in a geological context because 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 assumes an elliptical microcrack (Inglis 1913, Griffith 1924). For a uniaxial remote applied stress parallel to the microcrack axis, the tangential stress, on the microcrack surface var-ies from a negative (tensile stress) equal to the value of at the long axis of the microcrack, to a compressive stress at the short axis (Fig. 2.la). In a biaxial stress field is equal to at the long axis of the microcrack (Fig. 2.1a; Jaeger and Cook 1979). The stress state around an elliptical microcrack illustrates the essential concepts that large tensile stresses can develop at the tip of a microcrack surface in com-pression, and that the maximum tensile stress will develop in the direction of the maximum applied stress. These two concepts explain why extension microcracking perpendicular to the least principal stress is a widespread and fundamental cataclastic mechanism.

The second model treats the microcrack as flat with a sharp tip (e.g. Lawn and Wilshaw 1975a), and gives a versatile solution for the stress field around the microcrack, in the polar coordinates 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 factor and a function of which depends on the propagation mode. The three microcrack propagation modes shown in Figure 2.2 are tensile or opening mode, (mode I), in which displace-ments are perpendicular to the fracture plane, and two shear modes, in which displacements are parallel to the fracture plane, sometimes called sliding and tearing modes (modes II and III, Fig. 2.2). The stress intensity factor depends on the propagation mode, the microcrack length and the applied stress

Given these descriptions of the stress around the flat mi-crocrack, it is possible to deduce the failure criteria for brittle solids from Griffith’s postulate that a microcrack will extend when the total energy change with the propagation of the mi-crocrack is negative or constant. The energy terms involved in microcrack propagation are release of mechanical energy, which must be equal to the energy required for creation of surface area.In terms of force, the microcrack extension force,

G, must be greater than or equal to twice the surface tension force, (the doubling factor accounts for the two sides of the crack). This leads to the classical Griffith result for fail-ure under a tensile remote stress in a solid with Young’s modulus E:

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

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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 by postulating the existence of a small zone ahead of the micro-crack tip in which non-linear forces are expended in breaking bonds. The additional energy involved is incorporated in the energy balance in the form of a new parameter the fracture

toughness, to replace the surface energy term The energy balance condition is now:

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

This analysis is valid provided that the size of the non-linear zone is much smaller than the length of the microcrack i.e. it does not affect the elastic stress system as a whole: this assumption is called the small scale yielding (Rice 1968) or the small scale zone approximation (Lawn and Wilshaw 1975a). More detailed analyses can relate to the dis-placement on the microcrack if some function of the cohesive forces ahead of the tip with distance is postulated: these are the cohesive force models (e.g. Rudnicki 1980). The exist-ence of such a non-linear zone ahead of the microcrack tip is known experimentally from ceramics and metals, where it is referred to as a process zone, and it has been detected from acoustic emission and microcracking in geological materials (e.g. Swanson 1981, Peck 1983, Labuz et al. 1987).

Sub-critical microcrack propagation

The above analysis is restricted to microcracks that 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 conditions well below the critical stress intensity factor, at rates that depend on temperature and chemical environment as well as stress intensity. This phenomenon is known as sub-critical

microcrack growth and is potentially of enormous geological

importance (e.g. Anderson and Grew 1977, Das and Scholz 1981, Atkinson 1982, Meredith 1983). A diagram showing microcrack velocity (V) as a function of mode I stress intens-ity factor illustrates the main features of sub-critical mi-crocrack growth (Fig. 2.3), which have been demonstrated for a range of geological materials (e.g. quartz, granite, andesite, basalt, calcite, oil shale, sapphire and glass). The

relationship falls into three parts:

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

Region 2. Velocity is dependent on water concentration and

temperature, but not

Region 3. Velocity increases extremely rapidly with until

microcrack growth becomes dynamic at

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

Other types of chemical reaction occur with sub-critical microcracking. An example is the reaction of plagioclase to increasingly sodic compositions and ultimately to laumontite, which creates a 60% volume increase (Blenkinsop and Sibson 1991). This alteration occurs along cleavage microcracks, resulting in a texture of expanded, matching fragments, 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 into some models of crustal processes, for example to explain the difference between creep (stress intensity factor between and and seismic faulting (stress intensity factor equal to

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

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10 CHAPTER 2. CATACLASIS

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 the

coefficient 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 stress until sufficient area of contact is established to support the normal load (Fig. 2.4). The contact area is considered to have an adhesive strength which must be exceeded over the entire area for sliding to occur. This model successfully predicts the normal-stress dependence of friction, but it underestim-ates the value of because there are other contributions to the shear strength in addition to the adhesive strength of the contacts (Scholz 1990). These include:

1.

2.

3.

Interlocking of asperities. Oblique surfaces of contact are created between two asperities that come into contact after sliding (Fig. 2.4). Interlocking may be relieved by shearing through asperities (adhesive wear), or by slid-ing on the oblique contacts, which causes dilatancy. Increasing area of contact due to asperity shearing. As sliding continues, asperities fail and the contact area in-creases.

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

Adhesive strength, asperity interlocking and increasing contact area have been combined in an elegant and simple model by Wang and Scholz (1994, 1995), which accounts for experimental results very well.

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

where N (d) is the number of particles greater than size d, and

D is the fractal dimension. D for particles in cataclastic rocks

ranges from 1.88 to 3.08 (e.g. Sammis et al. 1987, Sammis and Biegel 1989, Olgaard and Brace 1983, Wang 1987), with many results for natural and experimental gouges around the value of 2.6 (Marone and Scholz 1989, Biegel et al. 1989). Sammis et al. (1987) showed that a distribution of spheres of unequal sizes reduces impingement stresses on individual fragments, and proposed that microcracking in a gouge will proceed in order to minimize the probability of neighbour-ing grains havneighbour-ing equal sizes. This constrained comminution model 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 by at least two other processes of particle size reduction. Alter-ation, for example the laumontization of feldspar described above, may lead to lower values of D (Blenkinsop and Sib-son 1991). At advanced stages of gouge formation, selective microfracture of larger particles takes place, creating PSDs with fractal dimensions greater than 2.58 (Blenkinsop 1991). Plates 1 to 4 show a sequence of cataclastic textures arranged in order of increasing fractal dimension of PSDs, which is the evolutionary sequence of textures with strain.

A simple theory of wear can predict that the volume of gouge created by sliding, and hence the thickness of the gouge, will increase linearly with displacement (Scholz 1990), and some experimental results confirm this relation-ship (e.g. Teufel 1981). This relationrelation-ship has also been claimed for faults (e.g. Robertson 1983, Hull 1988). How-ever, there are at least two reasons the experimental data should not be applied directly to faults. Firstly, the roughness of laboratory prepared surfaces is not comparable to natural fault surfaces (Brown and Scholz 1985), and secondly, once a gouge layer has accumulated sufficient thickness to prevent interaction between the sliding surfaces, wear will no longer occur according to the simple model of surfaces in contact with each other. The reported data on fault gouge 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 single grains), transgranular (across two or more grains) and

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cir-12 CHAPTER 2. CATACLASIS

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-cemencemen-ted, low poros-ity rocks have transgranular microcracks. This classification is useful in discriminating various microcrack mechanisms described in Sections 2.3.3 to 2.3.11.

Tectonic trans- or intragranular microcracks commonly link contact points between adjacent grains, and are kinked or curved. Despite grain-scale irregularities in fracture 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 their importance may be underestimated or even completely over-looked. The only manifestation of a fracture left in the rock may be a line of fluid inclusions which are the trace of a fluid inclusion plane (Fig. 2.5, Section 3.11), and careful obser-vations of well-prepared thin sections at high magnifications under the optical microscope are necessary to detect such mi-crocracks if they contain small inclusions. Mimi-crocracks can have a wide range of aspect ratios and densities. Extension microcracks may have regionally systematic orientations be-cause they form perpendicular to (Section 2.2.1.1), and have been used very effectively to deduce regional stress sys-tems (e.g. Lespinasse and Pêcher 1986).

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

2.3.2 Microstructures and mechanisms

Nine microcrack “mechanisms” can be distinguished, mainly from experiments, where extension microcracks, usually known as axial microcracks, form from about half the peak strength through to post-failure (e.g. Tapponier and Brace 1976). These are secondary mechanisms compared to the fundamental physics of tensile microcracking described in Section 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 link several pairs of contact points around grains. Four basic patterns are shown in Fig. 2.7 (Gallagher et al. 1974), which depend on the boundary loads, packing arrangement,

size, sorting, and grain shape. Impingement microfracturing can be understood from an analysis of the stress field cre-ated on loading a plane surface by a pointed (Boussinesq configuration) or spherical (Hertzian configuration) object (Fig. 2.8). The point load should produce radial microcracks sub-perpendicular to the indenter, and such microcracks have been observed in indenter experiments (e.g. Lindquist et al. 1984). The spherical indenter (possibly more geologically realistic) will contact the plane surface along a spherical sur-face, inducing a region of compressive stresses immediately below the indenter, surrounded by tensile stresses near the edge of the contact surface. An extension microcrack in the form 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 via a 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 case of two spheres loading each other: in this case, an extension microcrack initiates at the edge of the contact plane between the two spheres at a critical load that depends on porosity, grain size, elastic moduli and fracture toughness, providing that there are pre-existing flaws present in the loaded grains (Fig. 2.8, Zhang et al. 1990). The critical load measured in

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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 pressure solution (Section 3.4) show that tensile microcracks should form 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 impingement mechanism is linking of contact points by intragranular mi-crocracks (Fig. 2.9, Table 2.1), although the contact points may not be visible in the plane of the section. Impingement microcracking is suppressed in well-cemented or low poros-ity crystalline rocks because impingement contacts are lack-ing and tensile stress concentrations are dramatically reduced by 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 other microcracks, pores and grain boundaries. They form because of the tensile stresses that develop on the flaw surface when remote stresses are imposed. They are recognized in experi-ments on analogues and on rocks (e.g. Brace and Bombola-kis 1963, Tapponnier and Brace 1976). Analytical solutions show that the microcracks will grow along curved trajectories from both ends of a flaw to produce the well-known “wing crack” geometry (Fig. 2.10, Horii and Nemat-Nasser, 1985, 1986, Kemeny and Cook 1987, Baud et al. 1996). For isol-ated microcrack growth, a flaw length 2c, orientisol-ated at to is assumed to have both cohesive and frictional resistance to shear, and tensile microcracks grow from both ends with

a critical length (approximately 1.0) in even the slightest tensile value of microcrack growth becomes unstable, but remains stable in compression. Flaw-induced microcracking can be recognized by the connection of the microcrack to a

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 dominimport-ant 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 in controlling the overall microcracking process, and Atkinson (1982) asserts that it dominates the upper 20 km of the crust. The opening of grain boundaries (a type of pre-existing flaw) is well known to contribute to experimental cataclasis (e.g. Dunn et al. 1973, Hadizadeh 1980, Tapponnier and Brace 1976). The weakness of some natural grain boundaries, even when these are overgrown by optically continuous quartz ce-ments, is apparent from the observation that overgrowths are a major component in the matrix of faults (Pittman 1981). Microfracturing of pre-existing flaws can be identified by mi-crofracture of a cement, and may involve all three types of microcrack (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 recognized because 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 grains at contacts with micas in experimental studies (Tapponier and Brace 1976, Wong and Biegel 1985), and in a naturally deformed quartzite (Hippert 1994). These microcracks are length l (Fig. 2.10), making an angle with the flaw. The

res-ults show that microcracks will grow by tensile failure from the edges of the flaw along paths which fit experimental ob-servations very well (Horii and Nemat-Nasser 1985). After

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

2.3.8 Plastic mismatch microcracks

Where intracrystalline plasticity is localized in one area (e.g. in twins, deformation lamellae and kinks, Chapter 4), mi-crocracks may be initiated due to the strain incompatibility between the area of plastic deformation and the adjacent un-deformed area (e.g. Olson and Peng 1976, Wong and Biegel 1985; Fig. 2.12). Microcracks have been observed along kink bands in naturally deformed enstatite, and normal to the kink bands in experimentally deformed quartz (Carter and Kirby 1978). Plastic mismatchs may account for the common microcracking of feldspar porphyroclasts 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 due to stress concentrations created by intracrystalline plasticity (e.g. Lawn and Wilshaw 1975a). Plastic mismatch-induced

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

2.3.9 Microfault-induced microcracks:

Micro-scopic feather fractures (mffs)

Mffs are intragranular microcracks found only adjacent to faults. 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 of shear faults, and parallel to They did not occur adja-cent to an incipient shear, and therefore formed in response to 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 or parallel to the applied direction of Microcrack density and length increase with displacement and normal stress (confin-ing pressure) (Conrad and Friedman 1976, Teufel 1981). Tec-tonic analogues of mffs have been observed associated with shear surfaces between pebbles in contact with each other (McEwen 1981). T fractures as described by Petit (1987)

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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 formation of mffs due to contact stresses on the sliding surface. Mffs are intra- or transgranular microcracks that can be recognized by their localization adjacent to the fault plane, their inclinations of 20-50° to the fault plane, and their wedge-shape opening towards the fault plane (Table 2.1). Mffs can be distinguished from Riedel microfractures, which may also be associated with fault planes (e.g. Petit 1987), by the shear offset along the latter.

2.3.10 Thermally-induced microcracks

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

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

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

2.3.11 Phase

transformation-induced

micro-cracks

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

transformation involves a volume increase of 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 has been used as evidence for the former presence of coesite (e.g. Wang and Liou 1991). Although it is clear that the transition has occurred in these rocks be-cause relict coesite can be found in some inclusions, the ques-tion arises whether such microcracks could be due to elastic mismatches between the silica phase and the host. The key evidence for phase transition microcracking is the observation that there is no microcracking around other types of inclusion, including rutile. The extensional nature of the microcracks and their origin from tips of inclusions are consistent with the mechanism. Fracture surface energy measurements in the conditions of the quartz phase transition imply that microfractures could also be associated with this transforma-tion, and other minerals undergoing similar phase transform-ations (Kirby and Stern 1993). Radial microcracks have also been observed around calcite inclusions which have replaced aragonite, a transformation that involves a 8.5% volume in-crease (Wang and Liou 1991). The distinctive features of phase transformation microcracking are the association of in-tragranular microcracks with evidence for phase transform-ation. In the case of the transition, radial microcracks around inclusions of quartz after coesite are dis-tinctive.