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Current issues and new developments in deformation mechanisms, rheology and tectonics

S. DE MEER, M. R. DRURY, J. H. P. DE BRESSER & G. M. PENNOCK

Vening Meinesz Research School of Geodynamics, Faculty of Earth Sciences, Utrecht University, P.O. Box 80.021, 3508 TA Utrecht, The Netherlands (e-mail. [email protected])

Abstract: We present a selective overview of current issues and outstanding problems in the field of deformation mechanisms, rheology and tectonics. A large part of present-day research activities can be grouped into four broad themes. First, the effect of fluids on defor- mation is the subject of many field and laboratory studies. Fundamental aspects of grain boundary structure and the diffusive properties of fluid-filled grain contacts are currently being investigated, applying modern techniques of light photomicrography, electrical conductivity measurement and Fourier Transform Infrared (FTIR) microanalysis. Second, the interpretation of microstructures and textures is a topic of continuous attention. An improved understanding of the evolution of recrystallization microstructures, boundary mis- orientations and crystallographic preferred orientations has resulted from the systematic application of new, quantitative analysis and modelling techniques. Third, investigation of the theology of crust and mantle minerals remains an essential scientific goal. There is a focus on improving the accuracy of flow laws, in order to extrapolate these to nature. Aspects of strain and phase changes are now being taken into account. Fourth, crust and lithosphere tectonics form a subject of research focused on large-scale problems, where the use of analogue models has been particularly successful. However, there still exists a major lack of understanding regarding the microphysical basis of crust- and lithosphere-scale localiza- tion of deformation.

The motion and deformation of rocks are pro- mechanisms, rheology and tectonics. We have cesses of fundamental importance in shaping subdivided our review into four broad themes the Earth, from the outer crustal layers to the that reflect a large part of present-day research deep mantle. Reconstructions of the evolution activities: (1) the effect of fluids on deformation; of the Earth therefore require detailed knowledge (2) the interpretation of microstructures and of the geometry of deformation structures and textures; (3) deformation mechanisms and their relative timing, of the motions leading to rheology of crust and upper mantle minerals; deformation structures and of the mechanisms and (4) crust and lithosphere tectonics. This governing these motions. These problems con- introductory paper also serves to introduce the cern structures on all scales, from grain scale or papers presented in this volume. smaller to regional or global scale. Earth scien- tists in the early years of rock deformation studies focused strongly on extensive, detailed The effect of fluids on deformation descriptions of structures. Since the 1960s, the emphasis has been more on the mechanisms Fluids influence virtually all aspects of deforma- behind structure development and on the role tion mechanisms and rheology in the Earth on of the rheological or flow properties of rocks scales ranging from grain to plate boundaries during deformation within the framework of (Carter et al. 1990). Deformation in turn has an large-scale tectonics. Integration of laboratory important influence on fluid distributions in research, theoretical work on microphysical rocks (Daines& Kohlstedt 1997) and on rock processes, microstructural and outcrop-scale transport properties (Fischer & Paterson 1989). studies, and modelling of tectonics has become The involvement of water in deformation has more widespread, but at the same time the field been demonstrated in numerous field and labora- has broadened enormously. Consequently, the tory studies. One of the principal effects on need for dialogue between researchers from rheology results from the presence of water in different disciplines is ever increasing. grain boundaries. Grain boundary water sup- The objective of this paper is to present a ports a fast intergranular diffusion path, which selective overview of some current issues and allows stress-driven mass transport, resulting recent developments in the field of deformation in permanent, time-dependent deformation

From: DE MEER, S., DRURY, M. R., DE BRESSER, J. H. P. & PENNOCK, G. M. (eds) 2002. Deformation Mechanisms, Rheology and Tectonics: Current Status and Future Perspectives. Geological Society, London, Special Publications, 200, 1-27. 0305-8719/02/$15 © The Geological Society of London. Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

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(Paterson 1973, 1995; Rutter 1976, 1983; Green fault zones, fault creep, and the propagation, 1984; Lehner 1990, 1995; De Meer & Spiers arrest, and recurrence of earthquake ruptures. 1995, 1999). This process of dissolution-precipi- Besides the physical role of fluid pressures con- tation creep (or pressure solution) is an impor- trolling rock strength in crustal faults, it is also tant mechanism for: compaction in sedimentary clear that fluids can exert mechanical influence rocks (Tada et al. 1987); healing, sealing and through a variety of chemical effects. In recent strength recovery in active fault zones (Sleep & years, much attention has been focused on the Blanpied 1992; Hickman et al. 1995; Bos & role of pressure solution in the strength recovery, Spiers 2000; Bos et al. 2000; Imber et al. 2001); healing and sealing of faults and the role of deformation under low temperature meta- phyllosilicates therein (e.g. Gratier et al. 1994; morphic conditions (Elliot 1973; St6ckhert et al. Bos & Spiers 2000; Bos et al. 2000). It is generally 1999); and evaporite flow (Spiers et al. 1990; believed that pressure solution and subcritical Spiers & Carter 1998). crack growth have a significant weakening Despite the large amount of work already effect on fault zone rheology. However, Bos done on pressure-solution creep, many unre- and co-workers found that in their experiments solved problems remain. At present, the elemen- on halite-clay mixtures, pressure solution only tary diffusive and interfacial processes remain resulted in weakening of fault gouges when clay poorly understood. In particular, the structure was added. In the monomineralic halite fault and diffusive properties of water-bearing grain gouge, pressure-solution compaction and healing boundaries are the subject of ongoing debate. effects dominated, leading to frictional behav- Experimental studies of pressure-solution creep iour. In halite-clay mixtures, the presence of in crustal rocks have largely focused on compac- phyllosilicates at grain boundaries prevented tion of granular quartz or quartz-phyllosilicate grain contact healing, leading to a mechanism mixtures (Schutjens 1991; Mullis 1993; Dewers of frictional sliding along clay foliae, accommo- & Hajash 1995; Renard & Ortoleva 1997). How- dated by pressure solution of asperities. ever, compared with time scales accessible in the Recent results on dissolution-precipitation laboratory, pressure solution is slow in these creep, the properties of water-bearing grain materials, hampering reliable determination of boundaries, and the role of fluids in faulting (as bulk kinetics or the rate-controlling mechanism. well as vein formation) are discussed below. The involvement of fluids in faulting processes and shear zone development is widely recog- nized. A comprehensive review on the mechani- Dissolution-precipitation creep cal involvement of fluids in faulting is given by Hickman et al. (1995). Fluids are linked to a Dissolution-precipitation creep involves three variety of faulting processes, including long- serial steps (Fig. 1): (1) dissolution of material term structural and compositional evolution of at grain boundaries under high normal stress;

ii Fig. 1. Schematic illustration of pressure-solution creep. (a) Uniaxial compaction of a granular aggregate in the presence of saturated solution (saturated with respect to the stressed solid) at fluid pressure Pf. (b) Enlargement of grain contact area showing the three serial steps of pressure-solution creep: 1 - dissolution within the stressed grain boundary; 2 - diffusion through the grain boundary fluid; 3 - precipitation on the pore walls, an is the effective mean normal stress across the contact, Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 3

(2) diffusion through the grain boundary fluid caused by a thick fluid film supported between phase; and (3) precipitation at grain contacts clay minerals and, for example, quartz. Clay under low normal stress or on free pore walls minerals have a relatively large surface charge, (Raj 1982; Rutter 1983; Lehner 1990; De Meer leading to large hydration forces. Therefore, & Spiers 1997). Since interfacial reactions and clay-quartz boundaries are expected to have a diffusion occur serially, either the grain bound- thicker film than quartz-quartz boundaries ary diffusive properties or the interface reaction (Israelachvili 1992; Heidug 1995), promoting kinetics control the rate of the process. water film diffusion. When dissolution or precipi- Zhang et al. (2002) present the first systematic tation is rate controlling, cations in the solution investigation into the effect of applied stress, released by the dissolution of clays cause changes grain size and Mg 2+ content on the compaction in the solubility, dissolution and precipitation of wet calcite powder at room temperature. Well- rates of quartz (e.g. Dove & Rimstidt 1994; controlled starting aggregates were prepared by Renard et al. 1997; Dove 1999), leading to dry compacting granular samples before wet acceleration or deceleration of pressure-solution compaction. Dry compaction was carried out at creep rates. Furthermore, clay coating of avail- applied stresses higher than those for wet com- able precipitation sites (pore walls) will also paction. This minimized grain rearrangement lead to a decrease in pressure-solution creep and sliding during the subsequent wet compac- rates. Apparently, the presence of clays can tion stage of the tests. An acoustic emission result in both an increase and a decrease of transducer receiver was incorporated in the intergranular pressure-solution creep rates in experimental set-up in order to detect any grain quartz. cracking or brittle deformation. As no acoustic In addition to experimental work, the effect of emissions were detected and as the experimental clay on pressure-solution creep has also been conditions did not favour solid-state plastic modelled in numerical studies. Gundersen et al. deformation, intergranular pressure solution is (2002) present the results of a numerical study proposed to be the most likely mechanism for on the effects of clay on pressure-solution creep compaction" of the calcite aggregates.2+By system- in sandstone. They studied the effect of clay on atically increasing the amount of Mg added to the dissolution, grain boundary diffusion and the solution phase, compaction was drastically precipitation steps as well as on global transport inhibited. This agrees with the crystal growth (centimetre to decimetre scale). More specifically, literature in which it is known that Mg 2+ inhibits Gundersen et al. studied the effect of clays on: (1) precipitation of calcite (Reddy & Wang 1980; the increase of the kinetic coefficient of dissolu- Mucci & Morse 1983). It is therefore inferred tion or increased solubility; (2) the coating of that pressure solution in the calcite aggregates pore walls that inhibits precipitation; and (3) was rate limited by the precipitation step. the increase of the grain boundary water film In nature, clastic sediments often contain a thickness that enhances diffusion rates. The significant amount of clay, which can strongly results of the modelling work show that dissolu- influence the rate of pressure solution in a tion is a purely local process that governs the number of ways. For example, clay coatings on amount of mass transport which dissolves at pore walls may have an inhibiting effect on preci- the grain-to-grain contact. The model also pre- pitation, thus slowing down pressure solution dicts that diffusion affects both local and global creep (Baker et al. 1980; Tada & Siever 1989; (cm to dm scale) processes as it governs the rate Mullis 1991). On the other hand, small amounts of mass transport into the pore volume. Finally, of clay within grain boundaries may promote the model predicts that precipitation controls relatively rapid and ongoing pressure-solution global mass transport by limiting fluid super- creep (e.g. Weyl 1959; Dewers & Ortoleva saturation. 1991). Experimental evidence for such an effect The suggestion that clay has an accelerating has been reported by Hickman & Evans (1995). effect on pressure-solution creep in sandstone The accelerating effect of clay on intergranular (Gundersen et al.), can be assessed using the pressure-solution creep has been attributed to results of compaction creep experiments on different mechanisms. If diffusion is rate control- quartz-muscovite mixtures from Niemeijer & ling, a clay film on grain boundaries may enhance Spiers (2002). These experiments (at 500°C) grain boundary diffusion as it consists of a show that the compaction rate is not accelerated collection of clay ptatelets separated by thin by the addition of muscovite; rather, a modest water films that provide 'easy' diffusion paths decrease in compaction rate is observed. The (Weyl 1959; Hickman & Evans 1995). Alterna- results of previous work (Niemeijer et al. 2002) tively, Renard & Ortoleva (1997) put forward implied that, under similar experimental con- the hypothesis that enhanced diffusion can be ditions in muscovite-free samples, pressure Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

4 S. DE MEER ET AL. solution in quartz is rate limited by the dissolu- as different grain boundary structures may lead tion step. The decrease in compaction rate by to orders of magnitude difference in the predicted the addition'" of muscovite observed by Niemeijer3+ pressure-solution creep rates. & Spiers may be caused by dissolved A1 The observations by Den Brok et al. (2002), on dominating any accelerating effects of alkali- the development of stress-induced solid/fluid metal cations. This is expected to decrease the interface roughness, are consistent with the solubility, dissolution and precipitation rates of dynamically stable island-channel structure quartz. above. From theoretical considerations, it is known that the flat surface of an elastically strained solid is morphologically unstable (e.g. Structure and properties of water-bearing Leroy & Heidug 1994). Furthermore, the elastic grain boundaries strain energy of a flat surface can be relaxed by the formation of a rough surface. This roughness Recently, attention has been focused on in situ can develop, for example, by diffusion through observation of grain boundary structure and an aqueous solution. Den Brok et al. performed measurement of the diffusive properties of in situ experiments on (K-) alum single crystals actively dissolving fluid-filled grain contacts in order to study this process. In the experiments, using a variety of new techniques such as: grooves developed in orientations perpendicular reflected light interferometry and transmitted to the maximum compressive stress. These light photo-micrography (Hickman & Evans grooves were dynamic in nature, and moved as 1991, 1995); electrical conductivity (De Meer the local stress field changed. et al. 2002); and Fourier Transform Infrared (FTIR) microanalysis (De Meer pers. comm.). The aim of these experiments was to elucidate Channelized and pervasive fluid flow the structure and diffusive properties of wetted grain boundaries. Water may be present (Fig. Regional to crustal scale shear zones are often 2) in the form of: (1) strongly adsorbed thin described as channels of fluid flow (e.g. Etheridge films (Rutter 1976, 1983; Hickman & Evans et al. 1984; Pili et al. 1997). Le H~bel et al. (2002) 1991, 1995; Renard & Ortoleva 1997); (2) non- investigated the role of fluids in deformation equilibrium or dynamically stable island- processes in a regional shear zone in Southern channel networks or films (Raj & Chyung 1981; Brittany. They used geochemical methods to Raj 1982; Lehner 1990; Spiers & Schutjens assess the amount of fluids involved and the 1990); or (3) isolated inclusions or connected scale of mass transfer. The microstructures pre- crack arrays (Gratz 1991; Den Brok 1992). It is served and extensive vein development were of major importance to determine the structure consistent with solution-precipitation creep as of actively 'pressure dissolving' grain boundaries, the predominant deformation mechanism. Dis- solution of quartz and feldspars resulted in local enrichment of the rock in micas. Quartz and feldspar had been precipitated in numerous veins of which the composition reflected the local rock composition. Their geochemical data suggested that the extent of local volume loss of quartz and feldspar was in the range of 45- a b 75%. Oxygen isotope data were consistent with local control of fluid compositions and Le H6bel et al. show that the amount of fluid Grain A involved during deformation was limited and Pf P, fluid transfer occurred on a restricted scale with no significant fluid flow across lithological Grain B layers. Le H6bel et al. conclude that the regional C shear zone they studied in Southern Brittany acted as a trap for early fluids rather than provid- Fig. 2. Structure of water-bearing grain boundaries. ing a channel for extensive fluid migration. (a) Adsorbed thin film (thickness up to c. 20 nm). (b) Dynamically stable island-channel structure with In nature, vein-growth mechanisms can be channel thickness up to ,~250 nm. An adsorbed thin divided into two main classes: (1) vein formation film might be present at grain-to-grain contacts dominated by advective fluid flow (e.g. McCaig points. (c) Array of connected cracks (Gratz' model; et al. 1995); and (2) vein formation dominated Gratz 1991). Pf is the pore fluid pressure. by diffusional transport (e.g. Jamtveit & Yardly Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 5

1997; Bons 2000). In the first class of vein-growth experiments on sandstones, permeability mechanisms, material that precipitates in the showed negligible stress-induced anisotropy vein (nutrients) can be transported over long dis- before the onset of shear-enhanced compaction. tances to a vein-growth site. In the second class During the initiation of cataclastic flow, the per- of vein-growth mechanisms, it is generally meability tensor showed significant anisotropy believed that nutrients are derived locally (cm which diminished again with progressive devel- to dm scale) due to the limited distance of diffu- opment of cataclastic flow. Permeability aniso- sional transport. Diffusional transport along tropy is thus transient in nature under the concentration gradients is normally cited as applied experimental conditions. driving the nutrient transport towards growing antitaxial fibrous veins (e.g. Bons& Jessell 1997; Means & Li, 2001). This suggestion has The interpretation of mierostruetures and been tested by Elburg et al. (2002), using textures constraints from geochemistry, on carbonaceous shale-hosted fibrous calcite veins from the The development and application of new northern Flinders Ranges, South Australia. techniques of microstructure and texture charac- From major and trace element data they infer terization, such as automated electron back that, besides locally-derived nutrients, material scattered diffraction (EBSD) and similar electron was transported over distances of at least microscopy methods (Randle & Engler 2000; decimetres to over 100 m. The fibrous texture of Kunze et al. 1994; Leiss et al. 2000; Prior et al. a vein is thus no proof of local derivation of 1999), as well as computer-integrated polariza- nutrients. Furthermore, they found that fluid tion microscopy (CIP) and related light micro- flow was pervasive, as there was no evidence scopy techniques (Panozzo-Heilbronner & Pauli for preferential channelized fluid flow through 1993; Fueten & Goodchild 2001) have the the veins. potential to revolutionize our understanding of In order to understand the effects of fluids in microstructural processes in materials. With active tectonic environments, the fluid transport EBSD, the complete orientation from regions properties of rock need to be evaluated, particu- as small as 0.5 gm can be measured and displayed larly permeability. However, permeability is dif- as an orientation map. CIP involves computer- ficult to estimate in many geological processes assisted collection and analysis of crystallo- because of its sensitivity to pressure, temperature graphic orientations in optically uniaxial and stress. Laboratory experiments on perme- minerals using light microscopy. Maps showing ability development as a function of pressure, the orientation, misorientations and orientation temperature and stress provide useful constraints gradients of the c-axis can be produced with the on fluid transport properties of rock. Experi- CIP method. mental observations have showed that perme- Over the last twenty-five years a tremendous ability can be modified significantly under both amount of work on microstructures in rocks hydrostatic and non-hydrostatic stresses. More- and minerals has allowed a qualitative under- over, permeability depends sensitively on the standing of microstructural development to be anisotropic development of damage. In most established (Passchier & Trouw 1996; Blenkin- theoretical models to date, permeability is sop 2000). New techniques offer the possibility commonly prescribed as either a constant or a of providing a complete characterization of function of the effective mean stress (e.g. Ingeb- microstructures. Combining these microstruc- ritsen & Sanford 1998), rather than a more com- tural studies with experimental and theoretical plete second-rank tensor description. Because it studies should lead to an improved quantitative is very difficult to measure permeability simulta- understanding of the microstructural develop- neously in several different directions at elevated ment in minerals and rocks. pressures, knowledge about permeability aniso- tropy is largely lacking. For this reason, Zhu et al. (2002) have developed a so-called 'hybrid Dynamic recrystallization triaxial compression test'. When combined with conventional triaxial compression experiments, Recrystallization during deformation has a this new type of testing makes it possible to strong effect on microstructure and texture devel- measure permeability along both the minimum opment (Av6 Lallemant 1975; Karato 1988; and maximum principal stress directions. The Wenk et al. 1997; Herwegh et al. 1997) and tests provide quantitative estimates of the devel- may have an important influence on rheology opment of permeability anisotropy as a function of all types of materials (White 1977; White of effective mean and differential stress. In their et al. 1980; Urai et al. 1986; Peach et al. 2001). Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

6 S. DE MEER ET AL.

The basic mechanisms of dynamic recrystalliza- Geometric dynamic recrystallisation tion are reasonably well understood from a qua- litative viewpoint (e.g. Poirier & Nicolas 1975; White 1977; Poirier & Guillop6 1979), but quan- titative general theories of recrystallization are lacking in areas that account for the development of microstructure with strain, the dependence of recrystallized grain size on deformation condi- tions, the variations of recrystaUization mechan- isms with deformation conditions, and the effect of recrystallization on texture. Dynamic recrystallization has been studied (a) Initial microstructure extensively in metals. Metallurgists recognize three types of recrystallization, termed conven- tional dynamic recrystallization, continuous dynamic recrystallization and geometric dynamic recrystallization (Humphreys & Hatherly 1995; Doherty et al. 1997). The first type of recrystalli- zation involves the formation of new dislocation- free grains in the deformed or recovered struc- ture. The dislocation-free grains then grow at (b) High strain equi-dimensional microstructure the expense of the old deformed grains. In contin- uous dynamic recrystallization a new grain Fig. 3. Geometric dynamic recrystallization. structure evolves gradually and homogeneously (a) Elongated grains form at intermediate strains, during deformation with no distinct nucleation showing cusps along grain boundaries. (b) At high strains, geometric dynamic recrystallization occurs; an and growth of strain free grains (Humphreys & equi-dimensional grain structure is produced by the Hatherly 1995, 167-171; Doherty et al. 1997, impingement of irregular grain boundaries once the 248). Mechanisms proposed in continuous grains are strained to a width that is smaller than the recrystallization include subgrain growth, the amplitude of the grain boundary irregularities. development of new high-angle boundaries by merging of lower-angle boundaries, and the increase of boundary misorientation through 1979; Drury et al. 1985; Urai et al. 1986; Drury & the accumulation of dislocations into subgrain Urai 1990; Hirth & Tullis 1992). In many cases boundaries. Continuous recrystallization occurs recrystallization in minerals involves hybrid in metal alloys in which grain boundary migra- mechanisms (Drury & Urai 1990) which occur tion is inhibited by a high second phase and by a combination of migration and rotation pro- solute content (Doherty et al. 1997). Geometric cesses. The different terminology reflects differ- dynamic recrystallization is important in metals ences between recrystallization in metals and which do not undergo conventional discontinu- minerals and also, in some cases, a different clas- ous dynamic recrystallization, like aluminium sification of similar mechanisms (Humphreys & (Humphreys 1982). The mechanism (Fig. 3) Hatherly 1995; Drury & Urai 1990). Geometric involves the formation of new grains from the dynamic recrystallization has not yet been recog- original high-angle grain boundaries once the nized in minerals. However, high-strain micro- original grains have flattened to approximately structures of marble deformed in torsion (Pieri the diameter of the subgrain size, without invol- et al. 2001a, b) resemble geometric recrystalliza- vement of rotation recrystallization. Kassner (in tion structures (cf. Fig. 3), suggesting that this Doherty et al. 1997) suggests that geometric mechanism may be operative in geological dynamic recrystallization may generally occur materials. in high stacking fault materials at high strains, In order to estimate temperature and strain resulting in substantial grain size reduction rate conditions in naturally deformed rocks, a (Humphreys & Hatherley 1995; Drury & Hum- link is needed between experimentally and natu- phreys 1986). rally deformed samples. One way of making such The recrystallization terminology used by a link is to construct a recrystallization mechan- geologists concentrates on the role of grain ism map for dynamically recrystallized rock boundary migration (migration recrystallization) covering laboratory as well as natural conditions. versus the formation of new high-angle grain Stipp et al. (2002b) present such a map for dyna- boundaries from sub-boundaries (rotation mically recrystallized quartz. Natural data are recrystallization) (White 1977; Poirier & Guillop6 obtained at the Tonale fault zone (Italian Alps) Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 7 where a temperature gradient was determined and textural changes is likely not only to improve across a single shear zone using mineral assem- our understanding of these processes, but may blages in the metasedimentary rocks in the lead to routine use of new microstructural shear zone (Stipp et al. 2002a). Three distinct parameters. dynamic recrystallization regimes were defined An important development in misorientation within increasing temperature ranges, namely, studies of subgrains in materials deformed at regimes of bulging recrystallization, subgrain elevated temperature is the observed power law rotation recrystallization and grain boundary relationship between the average misorientation migration recrystallization. Here, bulging is a and strain (Hughes et al. 1997; Pennock et al. combination of migration and rotation pro- 2002). The theoretical subgrain misorientations cesses. This subdivision corresponds to regimes values predicted at higher strains by Mika & recognized in experimentally deformed quart- Dawson (1999) using finite element analysis zites (Hirth & Tullis 1992) of known deformation modelling are in good agreement with the experi- conditions. An estimate of the differential stress mental values found by Hughes et al., who also during deformation is obtained by applying a found that misorientation distributions from recrystallized grain size piezometric relation to transmission electron microscopy studies could the observed natural quartz microstructures. be scaled to a single curve. Misorientation distri- Stress estimation, however, relies on the theoreti- butions could, therefore, be useful in determining cal approach of Twiss (1977), since experimental the amount of strain accommodated by disloca- calibration of recrystallized grain size and stress tion creep in natural rocks. is not only very limited for quartz (see Hacker The distributions of the minimum angle of et al. 1990), but also fails to take into account misorientation (disorientation) from EBSD the proposed presence of different recrystalliza- studies are often presented, especially for sub- tion regimes. grain boundaries (Mainprice et al. 1993; Faul & Fitz Gerald 1999; Fliervoet et al. 1999). During rotation recrystallization, the misorientation Boundary misorientations angle distribution shifts to higher angles at higher strains (in aluminium, Hughes et al. During recrystallization, large misorientations 1997; in NaC1, Trimby et al. 2000; Pennock are a necessary precursor to the development of et al. 2002). Grain boundary migration recrystal- new high-angle grain boundaries. Grain bound- lization reduces the frequency of intermediate aries with certain misorientations have a higher misorientation angles in halite and in quartz at mobility and may affect grain growth and texture low strains (Trimby et al. 2000, 1998), although evolution (Humphreys & Hatherly 1995). For- these misorientations increase at higher strains mulating the development of misorientations, in quartz (Neumann 2000; Trimby et al. 1998). as a function of strain and temperature and in Problems in accurately defining the transition terms of microstructural processes, is therefore from a low-angle boundary to a high-angle an important step towards improving our under- boundary make it difficult to define the 'grain standing of the role of grain boundaries in recrys- size' in rocks (White 1977; Drury & Urai 1990; tallization and texture formation. The increasing Trimby et al. 1998). Trimby et al. (1998) intro- volume of data from automated EBSD studies, duced the 'grain boundary hierarchy' concept coupled with statistical analysis, is improving that accounts for the distribution of grain sizes our understanding of misorientation (Randle and boundary misorientations. The grain bound- et al. 2001; Wheeler et al. 2001; Prior 1999; ary hierarchy is defined by the variation of Lloyd 2000, Humphreys 2001). Fig. 4a shows domain size with the minimum misorientation deformed halite, mapped using EBSD. Misorien- angle (Fig. 4d). Such data can be readily tations, subgrain and grain scale textures can be obtained using the EBSD technique in the related to the microstructure, and displayed as a SEM. The information embodied in the grain misorientation angle distribution (Fig. 4b). boundary hierarchy is essential to the proper Large-step scan mapping (step size greater than characterization of grain and subgrain size distri- the average subgrain size, but less than the aver- butions. Usually an arbitrary angle is used to age grain size) is suitable for determining textures separate sub-boundaries from grain boundaries. (Fig. 4c) and can give a reasonable estimate of Most geologists use an angle of 10 ° (White average misorientations in materials with long- 1977) while material scientists use 15 ° . In range gradients (Pennock et al. 2002). The infor- materials with subgrains, the grain size will mation available from misorientation data from depend on which definition is used. This problem EBSD studies is a rapidly developing field and is illustrated in Figure 4d which shows the interpreting the data in terms of microstructural variation in measured grain size as a function Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

8 S. DE MEER ET AL.

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111

e 1Ol Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 9 of minimum misorientation angle in a sample of Leiss et al. also found that two end-member halite deformed at high temperature. If a geo- types of texture developed in amphibole and logical definition of a grain boundary is used, plagioclase from amphibolites in the Ingales the grain size is 40 gm, while a grain size of complex, Northern Cascades, USA. They sug- 60 gm is obtained if the metallurgical definition gest that the textures developed by crystal plastic of grain boundaries is used. This discrepancy slip, with texture reflecting different strain states arises because of the type of misorientation in different units of the Ingalis nappe complex. angle distribution developed during recovery Neutron diffraction techniques were used by and subgrain rotation (Fig. 4b). If the grain size Zucali et al. (2002) to study a deformed horn- is measured by light microscopy, then the distinc- blendite from the Sesia-Lanz0 zone (Italian tion between subgrain boundaries and grain Alps). This eclogite facies rock consists almost boundaries can only be made qualitatively. entirely of amphiboles with grain sizes ranging While there may be a relatively sharp struc- from 0.1-0.8mm. Conventional X-ray diffrac- tural transition from subgrain to grain bound- tion methods are not particularly suitable for aries at an angle ranging between 10-25 ° , the quantitative texture analysis of this type of properties of the boundaries may vary continu- rock, since the technique only allows sampling ously with misorientation. Boundaries should of a relatively small part of the aggregate, result- be considered on the basis of their properties ing in poor statistics if the grain size is large (e.g. rather than structure (Lloyd & Freeman 1994; Kocks et al. 1998). Part of the problem can be Lloyd et al. 1997). Consequently, the minimum solved if results of several parallel samples are misorientation used to define a grain boundary summed (e.g. Schwerdtner et al. 1971). In an could be different when considering different attempt to further improve quantitative analysis, processes. For recrystallization, boundary mobi- Zucali et al. compared their results from neutron lity is the key parameter and the transition from diffraction with spectra of X-ray diffraction low mobility to high mobility may occur at low summed from at least three slabs of the specimen. misorientations (Humphreys & Hatherly 1995). The analyses resulted in textures that agree with It is noted in this respect that the orientation of earlier work on amphibolite facies rocks (Gapais grain boundary might also be of importance & Brun 1981), and showed that the X-ray tech- (Randle 1998). nique alone provides enough information to obtain a reliable quantitative texture analysis. The presence of a lattice preferred orientation Textures in a rock usually indicates that dislocation creep processes were active during deformation; grain Crystallographic textures, or lattice preferred microstructures in the same rock are then often orientations, in metamorphic rocks can provide implicitly assumed to be associated with the information on the mechanisms, kinematics and deformation process. Heilbronner & Tullis conditions of flow in the Earth (e.g. Law 1990, (2002) performed texture analysis on experi- Wenk & Christie 1991, Schmid 1994, Bunge mentally deformed quartzites which were et al. 1994 and Leiss et al. 2000). subsequently annealed at the deformation tem- Leiss et al. (2002) used neutron diffraction to perature. The analysis was confined to c-axis analyse the textures of polyphase amphibolites pole figures constructed from orientation containing amphibole, plagioclase and occa- images obtained by light microscopy methods. sional quartz. They found that care was needed These images were calculated using the CIP in quantitative texture analysis and that more method (Heilbronner 2000). Heilbronner & than three individual pole figures must be mea- Tullis found that the microstructures of the sured to obtain accurate results for plagioclase. samples changed substantially during annealing,

Fig. 4. Microstructural and textural information from EBSD mapping of halite, deformed (in compression) to a strain of 50%, containing 6 ppm water. (a) EBSD map: a dark line is drawn for differences in orientation >10 ° between neighbouring pixels. The majority of boundaries >10 ° surround grains but many also occur within grains. The difference in shading within the grains shows the deviation of (110) poles from the compression axis (vertical). 4 ~tm step scan map after replacing 25% unindexed pixels, mostly along etched boundaries. (b) Misorientation angle distribution, showing the frequency of misorientations between pixels; low angle misorientations dominate the distribution. (e) Boundary hierarchy, showing that the average domain size (the diameter of a circle, with the equivalent area as the domain) depends on the minimum misorientation angle used to define the domain. (d) { 110} pole figure, where Y represents the compression axis, and inverse pole figure of compression axis. Equal area projection, upper hemisphere, contoured in 0.5 steps of mean uniform density, c. 3000 grains, 40 gm step scan size. Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021 l0 S. DE MEER ET AL. with a static recrystallized grain size that with experimentally and naturally deformed increased by a factor of 2 to 5. In contrast, the quartzites (Stipp et al. 2002b; Hirth & Tullis textures remained more or less unchanged. 1992). The simulations show that the microstruc- Grain microstructure and texture thus document ture and grain size at high strain depends on the different events from the history of the rock. relative rates of rotation recrystallization, Peridotites from the Ronda massif, southern nucleation recrystallization and grain boundary Spain, form a natural example of this (Van der migration, with the fastest process dominating Wal & Vissers 1996). These upper mantle rocks a microstructure. The rates of these processes show olivine crystallographic preferred orienta- depend mainly on temperature and strain rate. tions that developed during mylonitization, but Three recrystallization regimes have been recog- the granular microstructure of the rock is attrib- nized in experimentally deformed quartz (Hirth uted to static recrystallization. These findings and Tullis 1992). The simulations of Piazolo pose problems on the interpretation of natural et al. show that a transition from rotation domi- deformed rocks in general, and quartzites in nated recrystallization (quartz regime II - Hirth particular (e.g. St6ckhert et al. 1999; Hirth et al. & Tullis terminology) to migration dominated 2001). In such studies, grain size is used to con- recrystallization (quartz regime III) can be strain flow stress or strain rate, yet annealed associated with an increase in grain boundary recrystallized grains cannot be used to estimate mobility, which may be related to increasing palaeostress (Twiss 1977). The results of Heil- temperature and/or water content. An increase bronner & Tullis show that crystallographic of temperature, however, will also change the preferred orientations or textures cannot be driving force for grain boundary migration and used to distinguish annealed grains from this effect needs to be incorporated in future deformed grains. However, grain-shape criteria recrystallization models. are expected to be of use.

Deformation mechanisms and rheology of Computer simulation of microstructure crust and upper mantle minerals development Experimental deformation studies can provide Computer simulation modelling of microstruc- direct information on the deformation mechan- ture development provides a means of improving isms and rheology of minerals and rocks. The our understanding of the interplay of the main limitation of such data is the problem of processes involved in deformation, including extrapolation to slow natural strain rate and dynamic recrystallization. Furthermore, such from laboratory to larger scales (Paterson 1987, models can be used to pinpoint key areas for 2001). In contrast, studies of microstructures in future research and for predicting microstruc- naturally deformed minerals can potentially tures based on deformation and temperature provide information on the deformation regimes that are geologically important but mechanisms that actually operate in the Earth. which are not attainable experimentally. Jessell The limitations on studies of natural microstruc- & Bons (2002) have reviewed the current status tures are that the deformation conditions are of numerical modelling of all scales of micro- unknown, and the processes which are dominant structure evolution in rocks, paying particular in microstructure development may not be the attention to simulation and prediction of texture processes which control the rheology (Bos & development, grain boundary geometries, crystal Spiers 2000). growth and deformation in two-phase systems. The combination of EBSD studies of materials before and after deformation is a powerful tech- Flow laws and extrapolation to nature nique for studying detailed changes in orienta- tion, microstructure and texture, which can Creep laws for crust and upper mantle rocks are then be compared to numerical models (Bhatta- usually thought to follow the Dorn-type 'power charyya et al. 2001). Finite element models law' equation. Renner & Evans (2002) show have been used to investigate inhomogeneous that data for low strain deformation of calcite deformation and lattice rotations on the scale do not fit this type of law, and that exponential of individual grains (Zhang et al. 1994; Zhang creep laws (e.g. Goetze 1978; Tsenn & Carter & Wilson 1997; Mika & Dawson 1999). 1987) are more appropriate. The strength of Piazolo et al. (2002) used a numerical model different calcite rocks in the dislocation creep (Jessell et al. 2001) to simulate dynamic recrystal- regime varies with grain size, similar to the lization in quartz. The results can be compared Hall-Petch relationship found in metals. Renner Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS Ii

& Evans suggest that an extra variable dependent modated a substantial fraction of the deforma- on grain size - or subgrain size - is needed in tion. These results agree well with observations calcite flow laws and they show that flow laws made by Zhang et al. (2000) in simple shear for several dislocation creep mechanisms can be deformation of synthetic olivine aggregates. modified to include this effect. An alternative In calcite, current results show softening at approach is the internal state variable approach high strains associated with dynamic recrystalli- (e.g. Covey-Crump 1998). Renner & Evans zation (Schmid et al. 1987; Rutter 1999; Pieri show that the internal state variable model et al. 2001a, b) although softening is moderate proposed by Stone (1991) provides a good in most studies (maximum 10% in the torsion description of calcite flow strength. In the Stone tests ofPieri et al. 2001a, at 7 = 10) and not asso- model, deformation occurs by a combination of ciated with obvious shear localization. Grain size subgrain-size dependent glide in addition to dif- reduction was common in all tested calcite fusion-controlled subgrain boundary migration, materials, but no evidence was found for a with subgrain size and distribution as the internal switch in mechanism from dislocation creep to state variables. grain-size-sensitive (diffusion) creep (cf. Rutter The extrapolation of calcite flow laws to nat- & Brodie 1988), although diffusion processes ural conditions is discussed by De Bresser et al. might have contributed to deformation at high (2002) who note that there is a discrepancy strain according to Pieri et al. A detailed charac- between nature and experiment. Natural myloni- terization of the change in grain size distribution tic microstructures suggest that dislocation creep with increasing strain in experimental, uniaxial is the dominant mechanism, while the extra- deformation of Carrara marble is given by Ter polation of experimental flow laws predict that Heege et al. (2002). In this work, a bimodal dis- grain-size-sensitive creep should be dominant. tribution at the start evolved into a grain size A variety of flow laws has been obtained for distribution close to log normal (cf. Ranalli marbles, and extrapolation of these flow laws 1984) at strains of ~35%. The median and aver- produces large variations in strength predictions age grain size decreased with increasing strain, for calcite under natural conditions. De Bresser but showed a complex dependence on tempera- et al. suggest that there may be a 'missing link' ture. Associated with the evolution of grain size in current constitutive equations for creep in with increasing strain, a general weakening of calcite (see also Renner & Evans). As an alterna- 10-25% was observed by Ter Heege et al. This tive approach, De Bresser et al. derive a hypothe- weakening can be accounted for if the change tical flow law based on estimates of deformation in grain size distribution with increasing strain conditions in natural calcite mylonites. is included in composite flow laws comprising dislocation and diffusion processes. A gradual shift in distribution towards smaller grain sizes Rheological and microstructural evolution then results in an increased contribution to the towards high strain deformation of the relatively weak diffusion- creep mechanisms. Current flow laws for most minerals are based on In anhydrite, strain weakening up to 50% has low-strain experiments, yet natural deformation been observed in torsion test to shear strains of often involves enormous strain. The recent ~8 (see Heidelbach et al. 2001). The weakening expansion of methods for high-strain experi- was accompanied by a reduction in grain size ments (direct shear in a saw cut assembly, (from 12 to 6 gm). A switch in dominant defor- Zhang et al. 2000; and torsion testing, Casey mation mechanism as a function of strain has et al. 1998; Paterson & Olgaard 2000) may lead been proposed for this material, from dislocation to a new generation of high-strain flow laws. creep to diffusion creep, but full details have not Exciting results have been published on the yet been published. Textures of the deformed mechanical, microstructural and textural evolu- anhydrite were not weakened, but rather tion towards high strain of olivine (Bystricky increased in strength towards higher strain et al. 2000), calcite (Pieri et al. 2001a, b) and (Heidelbach et al. 2001). This seems to contradict anhydrite (Heidelbach et al. 2001). the general belief that diffusion creep does not In olivine, Bystricky et al. (2000) found that result in a strong texture development. the stress decreased by 15-20% during large- All three materials discussed above showed strain torsion deformation of olivine up to grain size reduction (by dynamic recrystalliza- shear strains of 5 (Fig. 5). This softening was tion) and mechanical weakening towards high associated with a grain size decrease from strain. However, weakening was limited (at least 20 gm to 3-6 lain. The development of a strong in case of calcite and olivine), and was substan- texture indicates that dislocation creep accom- tially less than would be expected if a complete Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

12 S. DE MEER ET AL.

250

~_j~ 200150 : ...... ~- =t~uv'''''~- -- - ~'-"

P,

5o ra~ 1,.. ~ K86 411|1.0% 3.0% 5.~ 70% axialshortening 0 I " | I HI I I I 0 1 2 3 4 5 Shear strain 7 Fig. 5. Selection of stress-strain curves for olivine materials at T = 1200-1300 °C, P = 300 MPa. BOO: synthetic aggregates (grain size D ,,~ 20 ~tm) of hot-pressed San Carlos olivine powders deformed in torsion at T = 1200 °C and shear strain rate 6 x 10-5 s -t (Bystricky et al. 2000), ZOO: synthetic aggregates (D ~ 35 ~tm) of hot-pressed San Carlos olivine powders deformed in direct shear in a saw cut assembly at T = 1200°C (top curve) and 1300 °C (bottom curve), shear strain rate 105 s-1 (Zhang et al. 2000). CP81: natural Aheim dunite (D ~ 900 lain; top curve) and Anita Bay dunite (D ~ 1001.tm; bottom curve) deformed in axial compression at T = 1200 oC, strain rate 10 s s t (Chopra & Paterson 1981). K86: synthetic aggregates (grain size c. 65 lain) of hot-pressed San Carlos olivine powders deformed dry (top curve) and wet (bottom curve) in axial compression at T = 1300 °C, strain rate 10 -5 s -1 (Karato et al. 1986). Note that the compressional stresses (or) measured for CPS1 and K86 have been converted into shear stresses (T) applying T = (1/V/3)a. switch in deformation mechanism had taken (Rudnick 1992). Thus, depending on local com- place. De Bresser et al. (1998, 2001) have argued position, the strength of the lower crust may be that rather than producing a switch in mechan- influenced by several minerals including quartz, ism, dynamic recrystallization might lead to a mica, feldspar, amphibole, pyroxene and balance between grain size reduction and grain garnet. Rutter & Brodie (1992) have provided a growth processes set up in the neighbourhood comprehensive review on the rheology of the of the boundary between the dislocation-creep lower crust based on experimental studies and field and the diffusion-creep field on a deforma- naturally deformed rocks from exhumed lower tion mechanism map. If this model holds, only crustal terrains. The strength of the lower crust minor rheological weakening can result from depends on the timing of deformation and meta- dynamic recrystallization accompanying defor- morphism (Rutter & Brodie 1992). The rheology mation. Further, the mechanical behaviour at of lower crust undergoing prograde metamor- high strain should be described by composite phism is strongly influenced by dehydration flow laws encompassing dislocation as well as reactions which generate high pore fluid pres- diffusion creep rather than by a single constitutive sures. Deformation mechanisms in this case rate equation. It seems worthwhile to test this may be dominated by solution-transfer and model against the high-strain experimental data. brittle deformation at low effective stress (Ether- idge et al. 1984; Rutter & Brodie 1992). Localized shear zones may develop by the concentration of The lower crust deformation into transiently fine-grained zones produced by dehydration reactions (Brodie & Experimental constraints on the rheology of Rutter 1987). The onset of melting may also the lower crust are limited compared to upper result in significant weakening of the lower crustal and upper mantle rocks (Kohlstedt et al. crust (Burg & Vigneresse 2002). 1995). The composition of the lower crust may After metamorphism, lower crustal rocks are vary considerably from mafic to intermediate expected to be relatively dry. Subsequent Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 13 deformation should be controlled by currently very little known about the rheology quartz+feldspar in intermediate lower crust of ultra-high pressure crust made up of minerals (White & Bretan 1985) and feldspar + pyroxene such as omphacite, garnet, jadeite and coesite. in mafic lower crust. Structures in lower crustal The rheology of high pressure crust has an terranes show that deformation can be concen- important influence on processes such as oro- trated into fine-grained shear zones formed by genesis, subduction and exhumation (Dewey metamorphic reactions (Rutter & Brodie 1992; et al. 1993; Austreim 1997). Kruse & Stfinitz 1999). These lower crustal St6ckhert (2002) reviews the information from fine-grained shear zones may deform by grain- field-based and experimental studies on defor- size-sensitive creep processes (Bouillier & Gue- mation mechanisms and stress levels in high guen 1975). pressure (HP) and ultra-high pressure (UHP) Deep seismic reflection profiles often show a metamorphic rocks with the aim of constraining highly reflective lower crust. Seismic reflectors the physical conditions along subduction zones in the lower crust have been interpreted as to depths of 100 kin. HP and UHP rocks are shear zones formed during crustal extension either undeformed or deform by dissolution- (Reston 1990). The lower crust is often consid- precipitation creep suggesting very low stress ered to be relatively weak and to act as a decoup- levels during (U)HP metamorphism in felsic ling zone between the higher strength upper crust rocks. Available flow laws for dislocation creep and upper mantle (Reston 1990). This view has also provide an upper bound to stress levels been questioned by Schmid et al. (1996) and along subduction zones within the uncertainties Handy et al. (2001), who point out that the of the extrapolation of experimental flow laws geometry of the lower crust revealed in seismic to natural conditions. St6ckhert concludes that profiles of the Alps is consistent with a strong deformation along subduction zones with a sub- lower crust and detachment of the upper crust duction channel filled with crustal material is: (1) from the lower crust and upper mantle. Consid- highly localized; and (2) occurs predominantly ering likely variations in lower crust composi- by dissolution-precipitationcreep with a Newto- tion, a large variation in lower crust strength nian theology at very low stress levels. might be expected depending on grain size and St6ckhert notes that eclogites often show water content. different behaviour, as many studies (Van Roer- Early work on the theology of feldspar has mund & Boland 1981; Lardeaux et al. 1986; been reviewed by Tullis (1983). Recent experi- Piepenbreier & St6ckhert 2001) have found evi- mental studies on anorthite have established dence for dislocation creep even at temperatures flow laws for diffusion creep (Wang et al. 1996; as low as 400-500 °C. Deformation by disloca- Dimanov et al. 1999) and dislocation creep tion creep at such low temperatures is incompati- (Rybacki & Dresen 2000) under both wet and ble with extrapolated flow laws for diopside dry conditions. As with quartz and olivine, the (Boland & Tullis 1986; Bystricky & Mackwell water content has a strong influence on the 2001) which implies that Na-pyroxenes jadeite strength and flow parameters of anorthite. The and omphacite have a much lower dislocation- effect of melt on anorthite strength has been creep strength than diopside. Preliminary creep investigated by Dimanov et al. (1998, 2000). data on synthetic jadeite (Orzol et al. 2001) The rheology of dry diabase has been studied indeed suggest a much lower strength than by Mackwell et al. (1998) who show that dry diopisde. However, recent creep data on eclogites lower crust may be much stronger than expected (Jin et al. 2001) indicate that dry eclogite has a from earlier experimental studies (Shelton & similar flow strength to harzburgite. A low Tullis 1981). In feldspar-poor regions of the creep strength in naturally deformed eclogites lower crust the strength may be controlled may be explained by a strong water weakening by pyroxenes. Bystricky & Mackwell (2001) effect in Na-pyroxenes (Buatier et al. 1991). describe experimental flow laws for clino- Clearly, further experimental studies on these pyroxenite which indicate that a pyroxene-rich materials are needed to resolve the 'eclogite lower crust could be stronger than the upper rheology problem'. mantle. Discoveries of ultra-high pressure mineral assemblages (Chopin 1984) have shown that con- Effects of melts on rheology tinental crust may be subducted or thickened, so that familiar crustal minerals are replaced by Melting during deformation occurs in the conti- high pressure polymorphs. While there are nental crust during orogenesis and in a wide large uncertainties concerning the rheology of range of tectonic situations in the upper mantle. normal lower crust (Handy et al. 2001), there is Intuitively, melts are expected to drastically Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

14 S. DE MEER ET AL. weaken rocks, but the effects of melt on rheology Further, microstructures provide constraints on can be quite variable and complex. Recent the flow mechanisms operative during deforma- reviews on this topic include Kohlstedt et al. tion and on the initiation of strain localization (2000) for mantle rocks, Nicolas & Idelfonse (Vissers et al. 1997; Jin et al. 1998). It has (1996) for oceanic crust and Rosenberg (2001) become clear that the rheology of the crust and for granitic compositions. Burg & Vigneresse upper mantle has an important influence on a (2002) discuss the rheology of partially molten wide range of tectonic processes, from the felsic rocks, extending the analysis presented by development of orogenic belts to the evolution Vigneresse et al. (1996) and Vigneresse & Tikoff of sedimentary basins. (1999) to include non-linear effects of melting The regional analysis of fault surface slip data and crystallization. The rheology of partially is a powerful method of tectonic analysis (e.g. molten rocks is often considered in terms of a B6nard et al. 1990). Wiesmayr et al. (2002) 'rheologically critical melt percentage' which apply this method to a part of the exhumed crus- marks the transition at which most of the viscos- tal slab of the High Himalaya in northwestern ity drop from full solid to full liquid occurs (Arzi Bhutan. Two sets of faults were distinguished, 1978; Van der Molen & Paterson 1979; Rosen- which differed in strike and age. Older faults in berg 2001). Vigneresse et al. (1996) suggested one set were found to be more steeply dipping, that the rheological transition between melt and while the second set of faults, with different solid occurs at different melt fractions during strike, showed more shallow dips. This might heating or cooling. For the case of melting mean that the principal stresses rotated with during heating, the solid loses cohesion at a time or, alternatively, that the entire rock mass melt content of between 20-25%. In contrast, rotated while the stress field remained constant. during cooling and crystallization a magma In the latter case, deep ramp structures on a gains cohesion once the melt content has crustal-scale thrust system could be the cause of decreased to 25-30% melt. The term 'melt the rotation. escape threshold' is proposed to describe the loss of cohesion of a melting rock and the term 'particle locking threshold' to describe the onset Softening and localization of cohesion in crystallizing magma. Burg & Vigneresse discuss how positive feedback effects One aspect of lithosphere tectonics which puzzles tend to localize deformation during melting and many researchers is the localization of strain in how negative feedback effects during crystalliza- deformation (shear) zones. Although localized tion result in a more distributed deformation. In zones are ubiquitous in natural settings, their consequence, during melting, the viscosity of the initiation remains a matter of debate (e.g. solid-melt mixture can be described by the Braun et al. 1999; Rutter 1999). In pressure- geometric average of solid and melt viscosity at insensitive viscous materials, persistent strain constant stress, while during crystallization the localization may only occur if weak zones are arithmetic average applies. Based on this present or are created by some softening process approach Burg & Vigneresse derive temperature, (Bowden 1970; Cobbold 1977; Poirier et al. 1979; melt content and viscosity relationships for par- Rutter 1999). However, localization can also tially molten rocks that are completely different involve hardening if deformation is dilatant for the cases of melting and crystallization. (Rudnicki & Rice 1975; Hobbs et al. 1990; Burg 1999). High-strain experiments on single- phase rocks show that some strain softening Crust and lithosphere tectonics occurs in many materials at high strain (Rutter 1999; Bystricky et al. 2000; Pieri et al. 2001a, b) The tectonic structure and evolution of the crust but the amount of softening is limited and and lithosphere can be analysed from different shear localization has not been observed (see viewpoints. Regional field studies generally pro- above). More significant strain softening has vide insight into the geometry and kinematics been reported in anhydrite (Heidelbach et al. of deformed parts of the crust and upper 2001) and feldspar aggregates (Tullis et al. mantle (Schmid et al. 1987). Microstructural 1989), but the concentration of deformation analysis focuses on key structures within the has only been reported in feldspar. large-scale deformation zones, such as faults or Most current experimental data are consistent ductile shear zones (e.g. Imber et al. 2001). with the suggestion of De Bresser et al. (1998, Inferences can be made on the dynamics of 2001) that dynamic recrystallization does not deformation, given that reliable palaeostress lead to drastic softening by inducing a change indicators are present (Blenkinsop 2000). to dominant grain size sensitive deformation, Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 15 which implies that other softening processes may Microphysical and numerical models for be important. Deformation mechanism maps dynamic recrystallization (Wenk et al. 1997; constructed for olivine from low strain experi- Shimizu 1998; Piazolo et al. 2002) are likely to mental data (Drury & Fitz Gerald 1998; Jin play an important role in resolving how much et al. 1998) imply that dynamic recrystallization softening can be produced by recrystallization. may induce significant softening and localization Shear zones are commonly fine grained com- by inducing a switch from normal dislocation pared to the adjacent rocks, and thus grain size creep to a hybrid mechanism (Fig. 6) involving reduction is often suspected of acting as a soften- deformation by a combination of dislocation ing mechanism. Cataclasis and metamorphic creep on the weak slip system and grain bound- reactions are important processes of grain size ary sliding (Hirth & Kohlstedt 1995). Dynamic reduction in natural shear zones (Stfinitz & Fitz recrystallization can result in other types of Gerald 1993; Vissers et al. 1997; St/,initz & softening including, geometric softening (Ion Tullis 2001; Imber et al. 2001). Handy & Stiinitz et al. 1982; White et al. 1985) and structural (2002) studied strain localization caused by frac- softening (Urai et al. 1986; Peach et al. 2001). turing and reaction weakening in spinel lherzolite from an exhumed passive continental margin. They observed two types of shear zones, with Dry Olivine T = 800°C different mineral assemblages and deformation 1000 microstructures. In the first type of shear zone, ultra-fine grained products of fracturing and retrograde reactions under high temperature 100 conditions (accompanied by fluid infiltration) allowed grain boundary sliding and diffusion processes to operate, probably lowering the 10 strength of the rock by more than an order of magnitude (see Fig. 6). The second type of shear zone developed under low temperature 1 conditions, and showed weakening associated with dilatancy and retrograde hydration, which was further enhanced once these type 2 zones 0.1 coalesced subparallel to the lithosphere-scale extensional shearing plane. Extensional shear zones beneath rifted margins thus might nucleate 0.01 as cracks in the initially strong upper mantle 0.001 0.01 0.1 1 10 100 rock, subsequently evolving into trans-litho- sphere weak zones as grain size is reduced due Grainsize mm to retrograde reactions associated with fluid Fig. 6. Deformation mechanism map for dry olivine infiltration. after Drury & Fitz Gerald (1998). Grain-size- insensitive dislocation creep controlled by slip on the hard [c] slip systems is dominant at high stress and Tectonic models large grain size. Diffusion creep and grain boundary sliding are dominant at low stress and small grain Insights into crust and lithosphere tectonics have size. At intermediate stress and grain size the benefited substantially from a scale modelling dominant mechanism is grain-size-sensitive dislocation creep where deformation occurs by a approach in which the outer shell of the Earth combination of slip on the weak slip system and grain is represented on laboratory scale by single- boundary sliding (Hirth & Kohlstedt 1995). The layer sandbox models or sand-silicone multi- shaded area marked 'rx grain size' shows the range of layers. Sandbox models usually consist of sand grain sizes expected for dynamic recrystallization in layers positioned above a rigid substratum, dry olivine. Under these conditions, grain size representing crust and lithospheric mantle, reduction by dynamic recrystallization may result in a respectively. Sand layers and substratum are switch from [c] dislocation creep to [a] dislocation separated from each other by means of a plastic creep. The shaded box marked as 'reaction grain size' sheet. With this approach, characteristic struc- shows typical grain sizes for polyphase bands that have been produced by metamorphic reaction. Grain tures of the crust have been reproduced and size reduction by metamorphic reaction can result in analysed in terms of geometry and evolution, change in deformation mechanism to diffusion creep such as: thrust wedge systems (Davis et al. in the fine-grained reaction productions (Newman 1983; Colletta et al., 1991); horst and graben sys- et al. 1999; Handy & Stfinitz 2002). tems (Koopman et al. 1986); and strike slip fault Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

16 S. DE MEER ET AL. zones (Schreurs 1994; Richard et al. 1995). In (e.g. Beaumont et al. 1994; Braun et al. 1999; these models, the crust is considered as being Gueydan et al. 2001). made up of frictional, Mohr-Coulomb materials which are decoupled from the lithospheric mantle. In contrast, multilayer sand-silicone Outstanding problems and goals for future models incorporate frictional as well as viscous work layers, which couple brittle and ductile rheolo- gies. In this way, the strength profiles of a layered We end this paper by outlining some of the out- lithosphere (e.g. Goetze & Evans 1979; Kohlstedt standing problems related to topics reviewed in et al. 1995) can be simulated. Essential in this the previous sections. approach is a correct scaling of relative strengths of the sand and silicone layers in the model with respect to brittle and ductile rheologies in Fluids and grain boundaries nature. Most analogue models are essentially two- Although during recent decades an enormous dimensional, although three-dimensional amount of field, laboratory and experimental models of, for example, continental indentation work has been directed towards understanding (e.g. Davy & Cobbold 1988) and oblique rifting the role of fluids in deformation processes, (Tron & Brun 1991) have also been produced. many questions remain unresolved. For example, The models almost invariably demonstrate that the involvement of fluids in faulting remains lithosphere-scale deformation patterns are uncertain. Questions that need to be answered highly dependent on the characteristics of the include: what are fluid pressures at depth, what theological layering as shown by Brun (2002). is the chemical role of fluids in fault and shear In particular, the rheology of the mantle and zones, and how can porosity and permeability the brittle-ductile coupling determine whether be altered? The primary goal of future studies deformation is distributed at lithosphere scale should be to identify and quantify the processes or localized in narrow zones. A high-strength and parameters that are most important in upper mantle, represented by a sand layer in controlling fault zone rheology (Hickman et al. the analogue model, tends to localize deforma- 1995) and reactivation (Imber et al. 2001). tion in zones of necking (in extension) or thrust- Despite extensive study and accumulation of a ing (in compression) which cross-cut the whole large amount of experimental data gathered on lithosphere (Davy et al. 1990, 1995). In contrast, dissolution-precipitation creep (or pressure- models with a ductile upper mantle and lower solution creep), several important questions are crust below a brittle top layer, result in more- still not fully answered. For example, the influ- or-less homogeneous thinning or thickening, ence of clays on intergranular pressure solution even if the ductile mantle is stronger than the in important rock-forming minerals such as ductile lower crust. Not only the relative quartz and calcite remains unclear from both strengths of the layers are of importance, but the experimental and theoretical points of view. also the rates of deformation. At low strain Furthermore, the structure and diffusive proper- rates, the lower crust may act as d~collement ties of grain boundaries during active pressure between upper crust and mantle (Brun 2002) solution are still not well understood, although therefore favouring localization. At high strain the different suggested grain boundary structures rates, the increased strength of the lower crust may lead to orders of magnitude difference in couples upper crust and mantle, resulting in predicted strain rate (Den Brok 1998). deformation to become more distributed. Pat- While the importance of fluids is widely recog- terns of distributed deformation appear to be in nized, current numerical models of geodynamic disagreement with structures in recent orogens processes seldom include fluid effects. The rheol- (Butler 1986), but might apply to Archean or ogy of the upper crust in the standard models is Proterozoic tectonics (e.g. Shackleton 1993; usually considered as a frictional or Mohr- Brun 2002). Coulomb material (e.g. Beaumont et al. 1994). While analogue models provide important Current strength profiles for the crust only contributions to our understanding of litho- include the physical effect of fluids on effective sphere tectonics, temporal theological changes, pressure and frictional sliding (e.g. Kohlstedt for example resulting from cooling or phase or et al. 1995; Evans & Kohlstedt 1995). Models grain size changes, cannot be incorporated. which include the chemical effect of fluids, such Also, measurement of stress or strain rate at as pressure solution are needed. For example, any point in the model is not possible. This Bos & Spiers (2002) describe a new model requires the application of numerical models for upper crustal deformation which involves Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 17 frictional-viscous deformation by a combination conditions. An assessment of environmental con- of sliding along phyllosilicates accommodated by ditions during deformation is then attainable, pressure solution of quartz. thereby improving boundary conditions for geodynamic modelling. Microphysical models for dynamic recrystalli- Microstructure development zation (e.g. Shimizu 1998) suggest that the grain size should be dependent on temperature as well Quantitative methods of microstructure and as stress. If this is generally valid, then large texture characterization have provided extensive errors in stress estimates may occur (De Bresser data on grain and subgrain sizes and shapes and et al. 2001). Another problem is the lack of on misorientation distributions. Future numeri- reliable experimental data for dynamic recrystal- cal models should investigate the full range of lization scaling laws. The application of early microstructural characteristics. The development calibrations of scaling laws in olivine (Ross of improved models for microstructure and et al. 1980) and quartz (Mercier et al. 1977; texture development requires information on Koch 1983) will result in overestimates of natural basic parameters as input for the models. This stresses (Green & Botch 1987). The calibration concerns information on diffusion coefficients, problems for quartz are discussed by Stipp et al. mobility and energies of subgrain and grain (2002a) who suggest that the theoretical model of boundaries. Many of these parameters for miner- Twiss (1977) is the best relationship to use. For als are unknown or badly constrained. There is a olivine, Van der Wal et al. (1993) reported that need for further experimental and micro- the dynamically recrystallized grain size was structural studies of these basic properties (e.g. independent of temperature and water content. Duyster & St6ckert 2001). However, later studies have shown that the Scaling laws between dynamic recrystallized grain size-stress relationship does vary with grain size and stress are often used to estimate melt content (Hirth & Kohlstedt 1995), water the flow stress during natural deformation content (Jung & Karato 2001) and flow behav- (Mercier et al. 1977). If temperatures can be esti- iour (Van der Wal et al. 1993, Fliervoet et al. mated from metamorphic mineral assemblages, 1999, Zhang et al. 2000) (see Fig. 7). In conse- then strain-rate estimates can be obtained from quence, parameters such as the water content experimental flow laws, given that these flow and the rate-controlling deformation mechanism laws are suitable for extrapolation to natural need to be identified, so that the correct scaling

1000 .(1) -,...... %. .,3,e(2) C~ 12.

cD 100 o - ",'e:ee,~,e/.o~'~. C'OO,e^ t O0 % ",,.%

10 I I 10 -6 1 0 -5 0.0001 0.001 Grainsize (m) Fig. 7. Plot of dynamic recrystallized grain size versus differential stress in experimentally deformed olivine rocks. 1 - data on dry single crystals at 1650 °C and wet and dry dunites at 1200-1300°C (Karato et al. 1982; Van der Wal et al. 1993). 2 - data from dry, fine grained synthetic olivine rocks (Zhang et al. 2000). 3 - data from olivine with high water content (Jung & Karato 2001). Plot shows that the scaling law between stress and grain size varies depending upon water content and which deformation mechanism dominates. Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

18 S. DE MEER ET AL. law can be used for stress estimates. A link Geodynamic models: significance of between flow laws and scaling laws (Fig. 7) localization of deformation implies that self-consistent relationships must be used in the estimation of stresses and strain It is important to obtain new experimental data rates from natural microstructures. and to develop new models for the grain size The problems with calibration and application formed by dynamic phase transformations. of grain size scaling laws impose limitations on Microstructures in natural shear zones (Stfinitz the applicability of quantitative 'grain size & Fitz Gerald 1993; Fliervoet et al. 1997; Kruse palaeo-piezometery', although studies of grain & Stfinitz 1999; Newman et al. 1999) show that size variation in naturally deformed rocks can the grain size of reaction products is generally give useful information on the qualitative a factor of 2-10 times smaller than the grain variations of stress levels. Further studies on size in single-phase domains. Dynamic phase the controls of grain size during large-strain transformations might be of more importance deformation are required. in grain size reduction and associated rheological weakening (cf. Rutter & Brodie 1988) than conventional, mono-phase dynamic recrystalli- High-strain flow laws zation. Geodynamic modelling studies show that So far, published flow laws for geological crustal and lithosphere scale localization has an materials (Kohlstedt et al. 1995; Evans & Kohl- important influence on the architecture and stedt 1995) have all been derived on the basis of width of compressional orogenic belts (Ellis low-strain deformation experiments. Recent et al. 2001) and is a key process in continental high-strain experiments, however, have clearly break up (Regenauer-Lieb & Yuen 2000) and shown that 'steady state' was probably not the initiation of subduction at passive margins achieved at the strains for which the flow laws (Branlund et al. 2000). Ductile shear localization are calibrated. This is illustrated in Figure 5 for may also be a crucial process that must be the case of olivine. Even if the amount of soften- included in mantle convection models in order ing found under experimental conditions is rela- for the models to reproduce plate tectonics (e.g. tively small, softening under natural conditions Tackley 2000; Bercovici et al. 2000). To date, could be larger if the temperature and stress strain softening processes have usually been dependence of high-strain deformation are differ- included in geodynamic models in an empirical ent to low-strain flow laws. Obviously, one pri- way (Govers & Wortel 1995; Braun et al. 1999; mary goal of future experimental studies should Tackley 2000; Gueydan et al. 2001). There is an be to establish high-strain flow laws. This may urgent need for materials based, microstructure not be easy, because high-strain deformation of and strain dependent rheological models that rocks might be governed by some combination include the microphysics of the processes of processes rather than a single mechanism. involved in deformation. Special attention should thus be devoted to designing experiments that allow testing of end- member mechanisms, for example by fabricating well-controlled starting materials. Special References emphasis must lie on grain size as being a funda- ARZI, A. A. 1978. Critical phenomena in the rheology mental parameter in the deformation behaviour of partially melted rocks. Tectonophysics, 44, of materials. Grain size is usually described by 173-184. an average, one-dimensional value (i.e. the dia- AUSTREIM, H. 1997. Influence of fluid and deformation meter) but other aspects may be important such on metamorphism in the deep crust and conse- as the distribution of grain sizes (Ter Heege quences for the geodynamics of collision zones. et al. 2002) and the distribution of grain bound- In: HACKER, B. R. & LIOU, J. G. (eds) When Con- ary misorientations (Trimby et al. 1998). tinents Collide: Geodynamics and Geochemistry of Considering the materials already studied in Ultrahigh Pressure Rocks. Petrology and Structural high-strain deformation, a new focus on lower Geology, 10, 297-323, Kluwer, Dordrecht. AvI~ LALLEMENT, H. G. 1975. Mechanism of preferred crustal rocks is needed. For example, feldspars orientations of olivine in tectonic peridotite, have been found to exhibit strong weakening at Geology, 3, 653-656. high strain (Tullis & Yund 1985) and this weak- BAKER, P. A., KASTNER, M., BYERLEE, J. D. & LOCK- ening can lead to shear localization (Tullis et al. N~, D. A. 1980. Pressure solution and hydro- 1989). Further information might allow more thermal recrystallization of carbonate sediments reliable interpretation of seismic reflectors in - an experimental study. Marine Geology, 38, the lower crust. 185-203. Downloaded from http://sp.lyellcollection.org/ by guest on October 1, 2021

DEFORMATION MECHANISMS, RHEOLOGY & TECTONICS 19

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