Geology of the Earthquake Source: an Introduction

Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

Geology of the earthquake source: an introduction

A˚ KE FAGERENG1* & VIRGINIA G. TOY2 1Department of Geological Sciences, University of Cape Town, Private Bag X3, Rondebosch 7701, South Africa 2Department of Geology, University of Otago, PO Box 56, Dunedin 9054, New Zealand *Corresponding author (e-mail: [email protected])

Abstract: Earthquakes arise from frictional ‘stick–slip’ instabilities as elastic strain is released by shear failure, almost always on a pre-existing fault. How the faulted rock responds to applied shear stress depends on its composition, environmental conditions (such as temperature and pressure), ﬂuid presence and strain rate. These geological and physical variables determine the shear strength and frictional stability of a fault, and the dominant mineral deformation mechanism. To differing degrees, these effects ultimately control the partitioning between seismic and aseismic deformation, and are recorded by fault-rock textures. The scale-invariance of earthquake slip allows for extrapolation of geological and geophysical observations of earthquake-related deformation. Here we emphasize that the seismological character of a fault is highly dependent on fault geology, and that the high frequency of earthquakes observed by geophysical monitoring demands consider- ation of seismic slip as a major mechanism of ﬁnite fault displacement in the geological record.

Rick Sibson has, throughout his career, pointed out the geological and physical parameters likely to that earthquakes occur in rocks (e.g. Sibson 1975, control their prevalence discussed. The mechanisms 1977, 1984, 1986, 1989, 2002, 2003). This simple by which rocks were deformed can be inferred from fact implies that fault rocks exert a critical control their textures (Knipe 1989); these relationships for on earthquake nucleation and propagation, and the typical fault rocks encountered in exhumed should contain an integrated record of earthquakes. fault zones are reviewed in this introduction. A con- Reading this record, Sibson has added significantly ceptual fault-zone model building on previous to our understanding of how faults work. His intui- review works (e.g. Sibson 1983; Scholz 2002) that tive ability to relate structures in exhumed rocks to is based on observations of the depth-distribution seismological observations of active deformation of fault rocks, inferred variability in deformation has contributed significantly to bridging the gap mechanism, fault strength and frictional stability between how geologists and geophysicists think with depth is also described. It is explained how about fault mechanics, and has led to the develop- this model provides basic controls on the depth ment of earthquake science as a truly interdisciplin- limits of the crustal seismogenic zone. Earthquake ary field. This volume celebrates Rick Sibson’s size and scaling parameters, in relation to interpret- career by bringing together a collection of papers ing data from different scales and methods, are also that present innovative studies spanning a range briefly reviewed; and, finally, the chapters in this of topics within the field of earthquake geology. Special Publication are outlined. Included are descriptions of fault-zone rocks from active and ancient faults, rock deformation experiments at seismic slip rates, and theoretical treat- Earthquake mechanism and frictional ments of fault initiation and slip. stability In this introduction the mechanism of earthquake faulting – generally accepted to be the release of Faults slip at a range of speeds, including slow creep tectonic elastic strain energy stored in a rock at plate tectonic rates (order of 1–10 cm year21), volume – by fast slip on a pre-existing or new fault fast frictional sliding (c.1ms21) during earthquake (Brace & Byerlee 1966) is discussed. Also included slip and displacement at a range of recently discov- is a short discussion on how rate- and state-variable ered, intermediate slip rates (such as slow slip and frictional stability may affect the rate at which fault low-frequency earthquakes: see the review by slip occurs (Dieterich 1972, 1979). The deformation Peng & Gomberg 2010 and references therein). mechanisms likely to be active at different con- The mechanisms of fault creep and slip at intermedi- ditions within a fault zone, which may lead to ate velocities are not yet obvious, but earthquake seismic or aseismic deformation, are outlined, and slip is generally thought to arise from ‘stick–slip’

From:Fagereng,A˚ ., Toy,V.G.&Rowland, J. V. (eds) Geology of the Earthquake Source: A Volume in Honour of Rick Sibson. Geological Society, London, Special Publications, 359, 1–16. DOI: 10.1144/SP359.1 0305-8719/11/$15.00 # The Geological Society of London 2011. Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

2 A˚ . FAGERENG & V. G. TOY frictional instability on a pre-existing fault (Brace & (Dieterich 1972, 1979; Scholz et al. 1972; Ruina Byerlee 1966) or, less commonly, by failure on a 1983; Marone 1998; Scholz 1998). Rate- and state- new fault surface. In this model, based on the variable effects on fault friction include an increase elastic rebound theory first suggested by Reid in ms with loading time, t, related to an increase (1911), elastic strain accumulates in rock surround- in real contact area with time under an applied ing the existing or potential fault surface (Fig. 1a), normal stress (Dieterich 1972). In addition, the until the resultant shear stress on the plane exceeds dynamic friction (mk), the sliding resistance experi- its frictional strength and failure occurs. Thus, earth- enced on a fault after slip has initiated, appears to quake faulting is a frictional process, where rupture vary with slip velocity, v (Fig. 1b) (Scholz et al. initiation depends on frictional resistance on the 1972). Although there are several forms of a consti- existing or potential fault plane. The shear stress tutive equation fitting laboratory observations of at failure, tf, is commonly approximated by the rock friction, a commonly used description of the empirical Amontons’ law: time and velocity dependence of friction is the Dieterich–Ruina law (Beeler et al. 1994): = + − tf C0 ms(sn Pf ) (1) = + + m m0 a ln(v/v0) b ln(v0u/D); where C is cohesive strength, m is the static coeffi- (2) 0 s du/dt = 1 − uv/D cient of friction, sn is normal stress and Pf is fluid pressure. Equation (1) indicates the shear stress conditions at which a fault will slip; however, it does not where m0 is the coefficient of friction at a reference define at what rate this slip occurs. slip velocity (v0), D is the critical slip distance, inter- Experimental studies of fault friction, generally preted as the slip increment needed for renewal of in homogeneous rock samples, have indicated that surface contacts (Dieterich 1979), a and b are empiri- a rate- and state-variable friction law can describe cal material constants, and u is a state variable observed sliding behaviour on laboratory faults that evolves over time. Geological effects, such as

time

Elastic strain (a)

Start of new cycle Interseismic strain Post-seismic accommodation (b) Slow Fast 0.07

0.06 T2 b2 0.05

0.04 ° T1 = quartz at 150 C, velocity weakening Δm 0.03 ° T2 = quartz at 600 C, velocity strengthening a2 0.02

0.01 a1 T b1 0.0 1

-0.01 0 2 4 6 8 10 Slip distance/D

Fig. 1. (a) Fault movement by elastic rebound. Starting from a relaxed state at the beginning of a seismic cycle, the rock surrounding the fault accumulates elastic strain. When this strain reaches a value where the resultant shear stress on the fault exceeds its frictional strength, displacement occurs on the fault, releasing stored elastic strain in the wall rock. (b) Example of the velocity dependence of friction, using material properties (a and b) for quartz from Blanpied et al. (1995) and Equation (2), at temperatures of 150 8C (solid line) and 600 8C (dashed line). Note the different evolution of the frictional coefﬁcient (m) predicted for velocity-weakening and velocity-strengthening conditions. Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

GEOLOGY OF THE EARTHQUAKE SOURCE 3 mineralogy, fluid pressure, dominant deformation by fluid presence between grains (Carter 1975). mechanism and rock structure, as well as environ- Because of high porosity and, commonly, high mental variables including temperature and pressure, fluid content, diagenetic chemical processes such are here included in the empirical material constants as cementation in pore spaces, recrystallization and a, b and D, and in the time-dependent state variable, u. fluid-assisted diffusional mass transfer into sites of Stick–slip instabilities, proposed by Brace & low stress are also likely to contribute to this sort Byerlee (1966) to result in earthquake slip, may of deformation (Maltman 1994). These chemical only occur where the fault material exhibits unstable processes are unlikely to accommodate much dis- or conditional frictional stability (Scholz 2002). placement but, rather, lead to an increase in cohe- Frictional stability depends on the empirical con- sion and rigidity, suppressing granular flow and stants a, b and D; where, if the difference (a 2 b) increasing the ability of the rock to store elastic is positive, the material is said to be ‘velocity strain. A shallow transition from continuous granu- strengthening’ and m increases with slip (Fig. 1b). lar flow to frictional, discontinuous deformation of Such material is inherently stable and cannot slip cohesive rock capable of elastic rebound may there- at seismic velocities (Dieterich 1979). However, if fore occur in response to increasing confining (a 2 b) is negative, friction decreases with slip pressure, and to mechanical and chemical cementa- (Fig. 1b) and the material is ‘velocity weakening’. tion (Moore & Saffer 2001). In this case, the material is either unstable or conditionally stable; earthquakes may only nucleate in Grain-size-sensitive creep the unstable regime but can propagate into conditionally unstable regions (Scholz 1998). In a general sense, deformation mechanisms The rate- and state-friction formulations imply comprising a mixture of grain-boundary sliding that the seismogenic behaviour of a fault depends and material transfer by solid-state or solution- purely on its frictional stability, not on mechanical accommodated diffusion are termed grain-size- strength. Although Equation (2) is a good empirical sensitive creep because they have an inverse depen- fit to experimental data, geological controls on dence on grain size (Ashby & Verrall 1973; Zeuch the empirical constants a, b and D, and the physi- 1983). Diffusive mass transfer occurs when a cal meaning of the state variable, u, are largely chemical potential gradient is created by a gradient unknown. Also, because rock friction experiments in normal stress along a grain boundary subject to have commonly been performed using homo- a local stress field (Gibbs 1877; Paterson 1973). geneous samples in dry conditions at room tempera- Diffusion of material occurs from relatively high ture, it is uncertain how well these friction laws normal stress regions to zones of low or negative (e.g. Equation 2) approximate the frictional behav- compressive stress. At high temperatures solid-state iour of natural heterogeneous fault rocks at seismo- diffusion may occur along grain boundaries or genic depths in the presence of a fluid phase and through the crystal lattice, but diffusivity is gener- ongoing chemical reactions. The value of shear ally too low for this mechanism to operate at stress at failure (Equation 1) is also likely to be crustal conditions (particularly ,350 8C: McClay affected by fluid-pressure-driven fluctuations in 1977). Despite this, rock textures in low-grade effective normal stress and time-dependent changes metamorphic rocks commonly indicate strain in cohesion by mineral precipitation and cementa- accommodation by diffusive mass transfer. Such tion. The slip behaviour of natural faults is therefore fabrics, which generally involve truncated objects likely to be a complex function of several geological and ubiquitous fibrous growths, are interpreted as parameters. Mineral deformation mechanisms other evidence for fluid-assisted diffusive mass transfer than frictional sliding may be of particular impor- (e.g. Sorby 1853; Durney 1972). This is commonly tance during crustal deformation. known as ‘pressure solution’ but as that technically only refers to dissolution under non-hydrostatic stress (Durney 1972), this phenomenon is here Fault-zone deformation mechanisms termed dissolution–precipitation creep. Granular flow In the presence of a reactive fluid phase in fine- grained rocks, dissolution–precipitation creep can At shallow depths, ductile shear, a macroscopically occur at low temperatures (down to ,150 8C in sili- continuous deformation, can occur by granular flow. cilastic rocks: e.g. McClay 1977; Duebendorfer Frictional grain-boundary sliding occurs, with little et al. 1998). Commonly described in low-grade or no breaking or modification of single particles metamorphic rocks, it is recognized as a major (Borradaile 1981). Rigid grains must be able to flow mechanism competing with frictional sliding rotate and flow past each other, that is, the aggregate and cataclasis in the mid- to upper crust (Gratier must be able to dilate. This mechanism is therefore & Gueydan 2007 and references therein). How- suppressed by high confining pressure and assisted ever, these mechanisms operate at different rates, Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

4 A˚ . FAGERENG & V. G. TOY timescales and stress conditions. Whereas frictional this review focuses on rock deformation within the sliding can occur at rates of up to m s21, dissol- seismogenic zone, the depth at which dislocation ution–precipitation creep generally occurs at slow creep becomes more efficient than frictional failure shear strain rates in the range of 10215 –10211 s21 is important as it may control the depth to the base (Pfiffner & Ramsay 1982). Frictional sliding also of the seismogenic crust (Sibson 1984; Scholz occurs episodically at relatively high differential 1988). The flow law for dislocation creep may be stress, and between episodes of failure a relatively expressed as (Hirth et al. 2001): long period of quiescence is required to build up 1˙ = Af m (s − s )ne(−Q/RT) (3) the necessary shear stress for reactivation (Equation H2O 1 3 1). Dissolution–precipitation creep, on the other where 1˙ is pure shear strain rate, A is a material con- hand, can be active continuously for tens to stant, fH O is water fugacity, m is a water fugacity thousands of years under low differential stress 2 exponent, n is the stress exponent, (s 2 s )isdiffer- (Gratier & Gueydan 2007). Experimentally, it 1 3 ential stress, R is the universal gas constant, T is has also been shown that frictional sliding and temperature and Q is activation energy. dissolution–precipitation creep can operate simul- The variables A, n and Q are solely functions of taneously, in a ‘frictional–viscous flow’ (Bos & the deforming material, so the primary factors that Spiers 2001). In this case, frictional sliding on affect the rate of dislocation creep in the absence weak planes results in cavitation that is accommo- of lithological variation are differential stress, temp- dated by dissolution–precipitation of intervening erature and composition. It is commonly inferred and surrounding material. This appears to be a par- that temperature exerts the primary control, ticularly efficient mechanism in gouge containing leading to a temperature-dependent transition from interconnected planes of phyllosilicates (Niemeijer frictional to viscous deformation in the crust, & Spiers 2006). which occurs at T 350 8C for quartz (Kohlstedt The dissolution–precipitation creep process can et al. 1995; Scholz 1998; Sto¨ckhert et al. 1999) be divided into dissolution, transport and precipi- and T 450 8C for feldspar (Tullis & Yund tation phases, of which the slowest determines the 1991). Mafic rocks are brittle to greater tempera- rate of the process as a whole (Rutter 1976). All tures, with a brittle–viscous transition in olivine at these processes have a strong dependence on grain temperatures in excess of 650–700 8C (Karato & size, where strain rate is inversely proportional to Wu 1993; Passchier & Trouw 2005). Boettcher grain size, or to grain size cubed in the case of et al. (2007) observed a change from velocity- diffusion-limited creep (Rutter 1976). In highly weakening to velocity-strengthening behaviour of fractured rocks fracture spacing may be at least as olivine in laboratory friction experiments, and important as grain size (Gratier et al. 2009). Conse- extrapolated the temperature dependence of this quently, microfracturing, leading to both decreased transition to approximately 600 8C at geological grain size and decreased fracture spacing, signifi- strain rates. They interpret the physical reasoning cantly increases the rate of dissolution–precipi- for this transition in frictional stability to be the tation creep (Gratier & Gueydan 2007; Gratier onset of dislocation glide, allowing for asperity et al. 2009). Thus, dissolution–precipitation creep deformation and slip-hardening (Boettcher et al. is likely to be an important deformation mechanism 2007). Thus, if dislocation creep is the major in fine-grained, fractured, fluid-saturated rocks, viscous deformation mechanism in the mid- to which are commonly present within fault gouge of lower crust, thermal structure and mineralogy deter- the crustal earthquake source. mine the maximum depth of frictional failure and unstable frictional behaviour, possibly also con- Dislocation creep straining the depth of the seismogenic zone (Scholz 1988). However, if other mechanisms (e.g. The third macroscopically ductile deformation mech- grain-size-sensitive creep) are more important, anism rocks may experience is the movement of temperature may not provide the primary control. atomic-scale line defects by dislocation glide and climb, which, with recovery, is known as dislocation creep. While brittle failure of intact rock, frictional sliding and dissolution–precipitation creep are gen- Fault-rock assemblages and fault-zone erally considered the dominant rock deformation structure mechanisms at low temperature (subgreenschist- facies conditions), viscous flow by dislocation The deformation mechanisms outlined in the pre- creep is inferred to be the most important mineral vious section produce distinctive textures, so the deformation mechanism at higher temperatures fault-rock assemblages may be used to assess the (greenschist-, amphibolite- and lower-granulite- dominant mechanisms active during deformation facies conditions: Passchier & Trouw 2005). As (e.g. Sibson 1977; Knipe 1989). However, fault-rock Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

GEOLOGY OF THE EARTHQUAKE SOURCE 5 microstructures result from a complex combina- Fig. 2d) provides a definite record of seismic slip. tion of protolith, the pressure–temperature–fluid However, pseudotachylyte is rare in fault zones pressure conditions during initial deformation and exhumed from seismogenic depths, indicating that the exhumation path (Snoke et al. 1998; Smith friction melt is either rarely produced or rarely pre- et al. 2011), thus requiring careful assessment of served relative to the frequency of earthquake slip cross-cutting relationships and post-deformation events observed on active faults (Sibson & Toy overprinting. 2006). Transfer of seismic slip across a dilational step-over may lead to a significant fluid pressure Rock textures produced in the drop in the dilatant region (Sibson 1985), causing frictional regime formation of an implosion breccia comprising wall- rock fragments cemented by hydrothermal precipi- In a monolithological fault zone, predominant tates (Fig. 2a) (Sibson 1986; Pavlis et al. 1993), fault-rock type can be expected to change with but it cannot be said with certainty that the slip depth (Sibson 1977; Woodcock & Mort 2008). leading to dilatancy happened at seismic rates At shallow levels, incohesive (unless post- (Cowan 1999). Other structures that may be associ- kinematically cemented) rocks with random fabrics ated with seismic slip include rounded spherical are observed (Fig. 2a, b), classified as fault breccia aggregates in clayey gouge, observed in laboratory (clast/matrix ratio .0.3; Fig. 2a) and gouge shear experiments at fast slip rates (Boutareaud (clast/matrix ratio ,0.3; Fig. 2b). These rocks et al. 2008; Ujiie et al. 2011) and in samples from probably form by granular flow at low confining natural, seismogenic faults (Boullier et al. 2009; pressures, such as in the top few kilometres of Boullier 2011), and incrementally developed hydro- active fault zones where low levels of microseismic thermal veins (Boullier & Robert 1992; Fagereng activity are commonly observed. They experience 2011). Unfortunately, these structures are either dif- dilatancy hardening at relatively low shear strains ficult to recognize in exhumed rocks, which have (Marone et al. 1990), so should be velocity strength- experienced overprinting by further deformation at ening, and contain clays, such as montmorillonite, depths shallower or deeper than the seismogenic which have a low frictional coefficient (Morrow regime, or may also be interpreted as resulting et al. 1992). They are therefore probably records from slower slip rates. of aseismic slip, or were produced around the shal- Cataclastic fault rocks may also form in the pres- low termination of upward-propagating ruptures. ence of dissolution–precipitation creep (Gratier & Cohesive rocks formed by cataclastic processes, Gamond 1990), as testified by ubiquitous pressure- defined as fracturing and rotation of mineral grains solution selvage seams in many fault-rock assem- and grain fragments accompanied by dilatancy and blages (e.g. Fig. 2c). Distinct microstructures frictional sliding along grain boundaries, are resulting from dissolution–precipitation creep or referred to as cataclasites (Sibson 1977; Woodcock cataclasis have been observed in simulated fault & Mort 2008). Although Sibson (1977) classified rocks experimentally deformed either at much less cataclasites as containing only random fabrics, it than, or close to, seismic slip rates, which display, has been demonstrated that cataclasites may also respectively, velocity-strengthening and velocity- develop a foliation defined by alignment of elongate weakening behaviour (e.g. Niemeijer & Spiers grains (Chester et al. 1985). Cataclasites comprise 2007). By correlation to these experiments, it is angular grains in a fine matrix (Fig. 2b, c), with reasonable to infer that natural microstructures greater matrix fraction and decreasing total grain result from aseismic creep or seismic slip (e.g. van size accompanying increased finite fault displace- Diggelen et al. 2010; Rowe et al. 2011). However, ment. Within a fault zone, one or more principal as it is likely that frictional sliding and dissol- slip zones (psz) are commonly defined by thin ution–precipitation creep coexist as the main defor- layers of ultracataclasite (with .90% matrix) mation mechanisms in the seismogenic crust, and, (Sibson 2003). Some fault zones may have one dis- as discussed above, operate at different timescales tinct principal slip zone surrounded by a commonly within the earthquake cycle, discontinuous defor- asymmetric, substantially thicker, highly fractured, mation textures produced by frictional slip and damage zone (Chester & Chester 1998), while continuous flow by dissolution–precipitation creep other faults comprise numerous anastomosing, thin may cross-cut and destroy each other forming slip zones surrounding lenses of fractured protolith complex resultant textures (Fig. 2c) (e.g. Knipe (Faulkner et al. 2003). 1989, 1990; Gratier & Gueydan 2007). A better Within cataclastic fault rocks, it is notoriously understanding of the interplay between frictional difficult to assess whether deformation textures sliding and dissolution–precipitation creep is formed during displacement at seismic, aseismic therefore likely to be important for understanding or intermediate slip rates. Cowan (1999) proposed the mechanics of earthquake nucleation and propa- that only pseudotachylyte (lithified friction melt; gation, as it may be possible that the diffusive Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

6 A˚ . FAGERENG & V. G. TOY

a b

c 500 μm d

500 μm

e f 54/080

500 μm 200 μm

Fig. 2. Examples of natural fault-rock textures. (a) Calcite-cemented dolomite fault breccia from the Naukluft Thrust, Namibia. (b) Two distinct packages of cataclasite (mint and dark green) separated by a zone of uncemented, very ﬁne-grained fault gouge running diagonally from top left to bottom right. The compass in (a) and (b) is 8 cm long. (c) Photomicrograph of ultracataclasite, with cross-cutting pressure-solution selvages (subhorizontal) truncating hydrothermal quartz veins (subvertical). (d) Photomicrograph with crossed polars of a pseudotachylyte fault vein (horizontal) with two injection veins (subvertical) into protocataclasite derived from quartzofeldspathic mylonite. The pseudotachylyte injection veins have been altered to chlorite towards their ends. (e) Mylonite, with fabric indicating that quartz and mica behaved viscously, deforming into elongate polygranular masses, while individual feldspars formed porphyroclasts that underwent microfracturing. (b)–(e) are all from the hanging wall of New Zealand’s Alpine Fault Zone. (f) Quartzofeldspathic mylonite with fabric diagnostic of continuous deformation in all phases – quartz, feldspar and mica – from the Grebe Shear Zone, Fiordland, New Zealand (courtesy of J. Scott). Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

GEOLOGY OF THE EARTHQUAKE SOURCE 7 mechanism is at least as important as frictional (T . 300 8C) (Voll 1976; Sibson 1984; Sto¨ckhert faulting at seismogenic depths (Rutter & Mainprice et al. 1999; Mariani et al. 2006). Fluid–rock reac- 1979, Gratier & Gueydan 2007). tion and hydrolytic weakening mechanisms may Seismic coupling, x, refers to the ratio of fault also allow creep to occur at shallower levels displacement accommodated by seismic slip (cumu- where a reactive fluid phase is present (Wintsch late over multiple earthquake cycles) over total plate et al. 1995; Hippertt & Egydio-Silva 1996). tectonic slip rate. For most crustal faults, x appears Conversely, in feldspars dislocation creep only to be indistinguishable from 1, that is, the vast becomes the dominant mode of deformation above majority of slip occurs seismically (e.g. Scholz approximately 450 8C, when deformation is com- 2002). Exceptions to this, shown by faults accom- monly accompanied by mineralogical change (Pryer modating significant displacement at aseismic or 1993; Passchier & Trouw 2005). Consequently, intermediate slip rates, are found in the San quartzofeldspathic mylonites developed at 300 , Andreas Fault System (Hill et al. 1990), and along T , 450 8C tend to have semi-discontinuous tecto- several subduction thrust interfaces (Ruff & nite fabrics developed by viscous creep of quartz Kanamori 1980). The implication of x ¼ 1 is that and mica combined with cataclastic deformation fault rocks deformed by slip at seismogenic depths of feldspar (Fig. 2e). In this mixed brittle–viscous are more likely to have experienced seismic slip regime, bulk rheology depends on the ratio of than not. It may therefore be appropriate to reconsi- high- to low-viscosity minerals, and the geometry der the proposal of Cowan (1999), who suggested of the phase distribution in the rock (Handy 1990). that fault slip is most likely to have been aseismic Consequent variations in bulk rheology may also and that faults should be considered to have slipped be reflected in the seismic style of actively deform- slowly until proven fast. Instead, we propose that the ing, heterogeneous shear zones (Fagereng & fact that earthquakes appear to be the main mode of Sibson 2010). discontinuous deformation in the frictional regime At temperatures exceeding 500 8C, quartzofelds- (e.g. Sibson 1989) should lead us to assume that pathic rocks tend to deform by continuous viscous most shallow crustal fault rocks were deformed seis- shearing flow (Fig. 2f), commonly distributed over mically. Hence, we should also look for ways of shear zones several hundreds of metres thick. defining rocks deformed by slower slip styles. Within such viscous shear zones, pseudotachylytes are sometimes observed (Sibson 1980; Clarke & Rock textures produced in the viscous regime Norman 1993), indicating that frictional sliding at seismic velocity can occur within a regime where The frictional process required for earthquake viscous mechanisms accommodate deformation faulting is pressure-dependent because high confin- most of the time. In some cases, pseudotachylytes ing pressure prevents frictional sliding and pro- have formed in the deep crust in relation to dehy- motes ductile flow (Byerlee 1968). Viscous creep, dration reactions in a cold, subducting slab at on the other hand, is primarily temperature- eclogite-facies conditions (e.g. Austrheim & Boundy dependent, and becomes more efficient as tempera- 1994), while in other instances, ruptures nucleated ture increases with depth (Equation 3). From these within the frictional regime have propagated down- relationships arise commonly presented crustal wards into normally aseismic, viscously deforming strength curves, which predict a frictional–viscous crust (Lin et al. 2005; Moecher & Steltenpohl transition at the cross-over point between the fric- 2011). Despite different mechanisms, these deep tional and viscous strength curves as calculated occurrences of pseudotachylyte signify that dynamic based on Equations 1 and 3 (Fig. 3) (Sibson 1977; changes in physical parameters, such as fluid Brace & Kohlstedt 1980). The conditions at which pressure or shear strain rate, can lead to transient this transition occurs depend on the behaviour of shear instabilities (or velocity weakening) and the mineral constituents of the fault rocks, primarily earthquake slip at depths normally considered quartz, feldspar and mica in quartzofeldspathic con- deeper that the earthquake source. tinental crust, their relative proportions, and phase arrangement. At depths where temperatures are high enough Conceptual fault model for efficient recovery, and where high confining pressures prevent easy frictional sliding, rocks So far, we have reviewed the concepts of elasticity develop mylonitic textures by dislocation creep, and frictional stability that govern brittle faulting, with or without dissolution–precipitation creep and have discussed the main deformation mechan- (White 1973; Sibson 1977; Hippertt & Egydio-Silva isms inferred for the mid- to upper crust, and 1996). In quartz and mica, these textures are com- typical fault-rock assemblages. Likely variations in monly developed at pressures and temperatures geological features, dominant deformation mechan- above the onset of greenschist-facies conditions isms, frictional stability, seismic behaviour and Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

8 A˚ . FAGERENG & V. G. TOY

Geologic Deformation Rate-and-state Seismic Crustal features mechanisms frictional behaviour behaviour strength profile

Soft sediment shear strength granular flow

+ Stable Stable but Pressure-dependent frictional faulting Unstable Dissolution- stable ruptures may propagate precipitation creep Conditionally through

Porosity + Frictional faulting DISCONTINUOUS

Diagenetic zone Clay-rich gouge Clay-rich DEFORMATION 150 ºC 6 km pseudo-

- b) tachylyte

(a Seismogenic zone Dissolution-

ºC/km geotherm

ºC) precipitation creep (Nucleation of earthquakes) Cataclasite +

metamorphism Sub-greenschist Frictional faulting Dissolution- precipitation ? creep 300 ºC

Temperature ( Temperature 12 km Seismic/aseismic transition zone

and depth (km) based on 25 Dislocation creep + Dissolution-

Aseismic CONTINUOUS Mylonite

Lower greenschistLower precipitation creep DEFORMATION creep Temperature-dependent viscous flow mechanisms 450 ºC 18 km Upper greenschist

Fig. 3. Conceptual fault-zone model, modified from Sibson (1983) and Scholz (1988, 2002). crustal strength with depth are summarized in seismogenic zone, within which earthquake Figure 3. Major transition zones appear to occur at ruptures nucleate and can easily propagate. Very temperatures of approximately 150 and 300 8C, low-grade (subgreenschist) metamorphic reactions corresponding to depths of approximately 6 and occur throughout this temperature range. The domi- 12 km, respectively, assuming a typical continental nant deformation mechanisms are frictional sliding geothermal gradient of 25 8Ckm21. and dissolution–precipitation creep, leading to The upper (colder) transition represents a tran- mixed frictional–viscous behaviour (Gratier & sition from shallow, predominantly aseismic, defor- Gamond 1990; Niemeijer & Spiers 2007). Assum- mation into the top of the seismogenic zone. Above ing a quartz-dominant lithology, frictional proper- this transition, rocks are poorly lithified and a fault ties are predicted to lie in the unstable field zone predominantly consists of clay-rich gouge. (Blanpied et al. 1995), and temperatures are too Diagenetic chemical reactions occur, causing pro- low for deformation by viscous creep to accommo- gressive cementation and decreasing porosity with date tectonic strain rates (Hirth et al. 2001). depth, accompanied by increasing frictional shear Shear strength is highly dependent on the fluid strength as sn increases with increasing overburden pressure state, and on the dominant deformation pressure. The dominant flow mechanisms are gran- mechanism. ular flow and dissolution–precipitation creep, At greenschist-facies conditions (.300– accompanied by brittle failure in more cemented 350 8C) dislocation creep and diffusive mass trans- rocks. Frictional properties at shallow depths are fer are inferred to dominate in continental crustal also presumed to be stable to conditionally stable; material. However, in mafic rocks, these mechan- thus, generally aseismic frictional sliding is pre- isms only become dominant over frictional failure dicted on pre-existing planes (Scholz 1998), at significantly higher temperatures (Karato & Wu although it is likely that large ruptures that nucleate 1993). At depths where viscous flow mechanisms at greater depths will propagate to the surface dominate, brittle or frictional failure occurs locally, through this regime. in the event of local elevation of strain rate, fluid The temperature range from approximately 150 pressure, or in the presence of different material to about 300 8C represents the predicted crustal properties such as composition or grain size. As Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

GEOLOGY OF THE EARTHQUAKE SOURCE 9 earthquakes occur by frictional instability, seismic by viscous mechanisms at greater depths (e.g. events are not expected to nucleate in these creeping Sibson 1983, 1984; Scholz 1988). If this inter- materials, although ruptures may propagate into pretation is correct, then the intercept between them because of elevated strain rates at the rupture Equations (1) and (3) defines the greatest depth at tip. At high temperatures (.350 8C), frictional which earthquakes can nucleate. The corollary is properties of a quartz-dominant lithology are that the peak strength that the crust can sustain expected to lie in the field of frictional stability then defines the depth to the base of the seismogenic (Blanpied et al. 1995), suggesting aseismic fric- zone, and this then depends strongly on strain tional sliding will occur and that propagating rup- rate, the composition and the viscosity of the tures will be rapidly arrested. dominant mineral phase, and the frictional resist- The transition from frictional sliding to viscous ance on pre-existing faults, as defined by their dislocation creep depends on material properties, frictional coefficient and the effective normal temperature, effective normal stress and strain rate stress (sn 2 Pf). (Equation 3), and therefore probably occurs at In addition to variations in dominant defor- different depths in different parts of a natural fault mation mechanism, the frictional stability along zone. Similarly, the frictional stability depends on faults may also determine seismogenic behaviour. material properties. Thus, because most natural As earthquake nucleation requires frictional insta- faults contain a range of geological materials and bility, seismic slip is confined to velocity- experience a range of conditions, the transition weakening materials at conditions favouring from stable to unstable sliding is likely to be irregu- frictional sliding and elastic strain in surrounding lar in space (and time?). rock. As quartz is velocity-strengthening above 300–350 8C and below approximately 150 8C (Blanpied et al. 1995), seismogenesis in quartz- The crustal seismogenic zone dominated continental crust is constrained to T , 350 8C. The base of the seismogenic zone is The fault model in Figure 3 may be translated thus likely to be controlled by a range of parameters to different heat flow regimes by changing the including composition, strain rate, temperature and thermal gradient, and hence the depths on the verti- effective stress, and is therefore complicated by cal axis. The temperature and depth variability of the probable heterogeneity in both composition various factors affecting slip style emphasizes the and fluid pressure distribution within the continental complex interplay of physical parameters determin- crust. ing the up- and down-dip limits of the crustal seismogenic zone. In theory, the seismogenic zone is defined by Earthquake size and scaling relationships where earthquakes can nucleate, and in practice it is defined by the depth-distribution of microseismic Within the heterogeneous rocks of the earthquake activity and mainshock–aftershock sequences. The source, earthquakes occur over a large range of latter is problematic because there are many major length scales (Gutenberg & Richter 1944) but crustal fault zones that have not experienced signifi- follow a consistent self-similar pattern (Kanamori cant earthquakes since seismometers have been & Anderson 1975; Hanks & Kanamori 1979): that available to record earthquake arrivals (e.g. New is, 1024 , u/L , 1025, where u is average slip Zealand’s central Alpine Fault: Evison 1971), so and L is fault length (Wells & Coppersmith 1994). there is little control on the depth to which ruptures This implies that irrespective of earthquake size, may propagate. Furthermore, although commonly slip distance scales with the size of the fault on defined as a range of crustal depths, the extent of which sliding occurs. the seismogenic zone may vary along strike of a Earthquake size is traditionally measured by the single structure, and comprise regions of strong earthquake magnitude. A variety of magnitude and weak coupling on a fault plane (Bilek & Lay scales, based on measuring the amplitude of a 2002). However, it is unclear how interseismic specific seismic wave at a specific frequency, have coupling relates to the seismogenic zone (e.g. been used through history. This approach is proble- Kaneko et al. 2010). matic because of the difficulties encountered in The depth of the seismogenic zone varies signifi- correcting for distance and instrument response, cantly between tectonic settings, but the tempera- and the different source spectra for different magni- ture at its base in generally around 350 8Cin tude events. continental crust (Scholz 2002). This is approxi- While earthquake magnitude is a logarithmic mately coincident with the expected frictional– scale based on instrument response to seismic radi- viscous transition, so may be defined by a transition ation, the seismic moment (Aki 1967) provides a from the shallow elastic–frictional regime to creep physically meaningful measure of earthquake size. Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

10 A˚ . FAGERENG & V. G. TOY

Seismic moment (in N m) is deﬁned as a scalar by ~1000 km the relationship: ~ 300 km Mw 9 Mw 8 ~30 km M0 = GuA (4) ~150 km

25km ~ ‘typical’ geological study area where G is shear modulus and A is rupture area. Hanks & Kanamori (1979) designed a magnitude scale based on the seismic moment, the moment ~300 km magnitude scale, where magnitude, Mw, is pro- Mw 6 portional to log10 M0. A consequence of this Mw 8 M 7

~30 km w relationship is that Mw is proportional to rupture area (Aki 1981). Dynamic rupture characteristics, that is, slip acceleration, distance and velocity, vary linearly ~20 km with stress drop, Dt (Scholz 2002). A number of studies have found that Dt is relatively constant Mw 4 Mw 6 over a large range of magnitudes and tectonic set- Mw 5 200 m tings (Kanamori & Anderson 1975; McGarr 1999; ~5 km rock outcrop Allmann & Shearer 2009). The mean stress drop is approximately 3 MPa, although variation over the range 0.03 , Dt , 30 MPa is evident in published ~2 km data sets (Scholz 2002; Choy et al. 2007; Allmann & Shearer 2009). Mw 2 Mw 4

A consequence of constant Dt is that it is poss- Mw 3 20 m ible to estimate the relationships between u, L and ~500 m rock outcrop Mw (Hanks & Kanamori 1979; Wells & Copper- smith 1994; Sibson 2011). The effect of these Fig. 4. Relative size of the rupture areas from scaling relationships on earthquake rupture area is earthquakes of Mw 9 down to Mw 2, based on shown graphically in Figure 4. As well as illustrat- relationships determined by Kanamori & Anderson ing the variation in rupture size with earthquake (1975) and Hanks & Kanamori (1979). Dimensions have magnitude, Figure 4 emphasizes the difficulty of been adapted to elliptical rupture areas for ease of studying earthquake parameters at outcrop scale, presentation and for more realistic shapes of the larger which is important when using geological obser- earthquakes, where the down-dip dimension is limited by vations to make inferences about earthquake the thickness of the seismogenic zone. mechanics – as discussed in many contributions within this volume. For example, 25 km, a typical map-scale study area, is equivalent to the rupture Davis 1993), implying a 10-fold increase in fre- length of an elliptical Mw 6 rupture. A typical quency for each unit decrease in magnitude. As single outcrop of about 20 m, on the other hand, the world experiences one Mw 8 earthquake a is only visible on an Mw 4 sized rupture plane in year, on average, this implies that 10 Mw 7 events, Figure 4, and anything larger than an Mw 2is 100 Mw 6, 1000 Mw 5, and so on, occur every year. inferred to rupture more than the entire outcrop. It If applied to a single fault, the frequency– is important to be mindful of these scaling relation- magnitude relationship tends to significantly under- ships when comparing geological and geophysical estimate the size of the largest earthquake possible earthquake observations, as different processes are on the fault (Wesnousky et al. 1983; Schwartz & visible (and invisible) at different scales. Coppersmith 1984). This may be a consequence of The empirical Gutenberg–Richter relationship the different scaling of small and large earthquakes (Ishimoto & Iida 1939; Gutenberg & Richter (Scholz 1997), but is also an effect of the basis for 1944) is another fundamental earthquake scaling the frequency–magnitude relationship. It is com- relationship. This frequency–magnitude relation- monly inferred that the Gutenberg–Richter relationship is a consequence of the self-similar scaling of ship derives from a power-law distribution of faults; earthquakes over many magnitudes, and implies each fault failing in a characteristic earthquake that, in any region, for each unit decrease in magni- during which an entire fault segment fails – in tude, the earthquake frequency increases by a factor other words, the characteristic earthquake is depen- of 10b. The relationship appears to hold to magni- dent on the fault size (King 1983; Schwartz & tudes down to Mw21 in scientific borehole exper- Coppersmith 1984; Wells & Coppersmith 1994; iments (Abercrombie 1995). The value of b is Wesnousky 1999). The geological implication is approximately 1.0 in most regions (Frohlich & that there is likely to be some parameter limiting Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

GEOLOGY OF THE EARTHQUAKE SOURCE 11 the earthquake size on any given fault, which may temperature and geological setting during defor- be recognized as a geological or geometrical mation must be interpreted. In the second group of feature determining the end points of a rupture papers, Fagereng discusses observations in an (Sibson 1989). exhumed me´lange shear zone, interpreted as under- These scaling relationships imply a scale- thrust sedimentary rocks and scattered basalts invariance in the dynamics of earthquakes, sug- deformed below a subduction dećollement. He high- gesting that information about the geological lights how a subduction megathrust comprises a effects of earthquakes at a range of scales, derived wide zone of heterogeneous rock, within which from microscopic to outcrop scale studies, can be slip occurs by both localized and distributed defor- extrapolated to the mechanics of larger events. mation, and speculates on how observations in the However, we must always be mindful that most exhumed shear zone can be interpreted in terms of geological studies are performed at scales very the fault-slip behaviour within modern analogues. different from those considered by geophysicists Rowe et al. also present geological observations observing active faults (Fig. 4). The chapters in from a fossil subduction thrust interface. They find this volume, outlined in the next and final section three mutually cross-cutting fault-rock textures, of this introductory chapter, therefore span a range and interpret that these were produced during slip of scales from microscopic observations of natural at seismic, aseismic and intermediate slip rates and laboratory faults, through to field studies of during different parts of the earthquake cycle. fault rocks from a range of inferred depths and set- Field and microstructural observations from a tings, to larger-scale geophysical and theoretical different tectonic setting, a low-angle normal fault, contributions. are presented by Smith et al., who also review the extensive literature available on the Zuccale Fault. They identify a range of deformation mechanisms Chapters in this volume that were active broadly contemporaneously within the fault core. The relative importance of The first two original research papers describe these mechanisms appears to have varied over observations from drilling projects in active fault time as a function of temperature, fluid presence, zones. These contributions – based on direct obser- fault-zone structure and mineralogy, resulting in vations of rocks actively deforming within the seis- fundamental changes in fault-slip behaviour. mogenic zone – set the scene for studies involving Overall, these three chapters present new ideas on interpretations of exhumed fault rocks, laboratory how to interpret fault-rock assemblages in relation experiments and theoretical models, which all rely to the range of seismic styles observed in active on understanding what an active fault zone looks faults, and address questions of the deformation like. In the first of these chapters, Boullier reviews mechanisms responsible for different slip modes geological observations from drilling projects in within the seismogenic zone. the Nojima and Chelungpu faults, both responsible Toy et al. then describe pseudotachylyte occur- for recent large earthquakes. Her review highlights rences in the transpressional Alpine Fault Zone, how both fault-zone geology and fluid pressure New Zealand, which, although ubiquitous, are state vary considerably between these two faults, insufficiently abundant to account for all increments so that no unique solution can be applied to the of earthquake slip on this fault. They propose that geology of all crustal fault zones. Ellsworth & many of the pseudotachylyte veins were formed Malin present seismological observations from the during moderate magnitude events near the base San Andreas Fault Observatory at Depth, where of the seismogenic zone and that some represent they identify a new type of fault-guided seismic events that propagated downwards into normally wave (Fw). They argue that fault-guided seismic aseismic regimes. Voluminous pseudotachylytes waves in this location propagate within an approxi- are mostly restricted to exhumed slivers of the foot- mately 30 m-wide damage zone, which extends to a wall of the fault. A distinct group of pseudotachy- depth of 7 km, that is, into the seismogenic zone. lytes are interpreted to have formed at shallower They suggest that this damage zone arises either depths, in the middle seismogenic zone, probably from fracturing around the principal slip zone of in the presence of fluid patches on heterogeneous large earthquakes or is generated by aseismic creep. fault planes; this observation challenges the view While deep drilling projects have the advantage that thermal pressurization can significantly lower of direct sampling from known depths in well- fault strength at asperity contacts. characterized active faults, deformation structures This paper provides a transition to a third group developed within the seismogenic zone may also of papers that discuss rupture processes at the base be studied in exhumed fault-rock assemblages. Out- of the seismogenic zone and into the mid- to lower crops provide natural laboratories at much larger crust. Allen & Shaw present a study of a several observational scale than a drill core, but depth, hundreds of metres wide, high-strain shear zone, Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

12 A˚ . FAGERENG & V. G. TOY which they interpret as an anisotropically weakened fracture healing. Fault-related calcite veins are zone generated in the mid-crust. The shear zone also described by Nuriel et al., who investigate hosts pseudotachylyte veins, which they interpret the structure, petrography, geochemistry and geo- as evidence of first-generation ruptures in intact chronology of calcite precipitates along the East rock. Some of the pseudotachylytes were generated Anatolian and Dead Sea fault zones. The nature of in a regime where creep could still occur, possibly the calcite precipitates and their relationships to by down-dip propagation of shallower-level rup- faulting vary significantly, and Nuriel et al. con- tures. They suggest that the anisotropy of the clude that geochemical and structural analysis, com- sheared rocks constrained rupture geometry as bined with U–Th geochronology, can be used to they were exhumed into the seismogenic zone. A provide age constraints on the timing of brittle petrographical study by Altenberger et al. also deformation. Upton et al. also present a study of illustrates that earthquakes may nucleate below fluid flow in active settings as they compare and expected seismogenic zone depths. These authors contrast the distribution of heat flow, hot springs determine the pressure–temperature conditions of and hydrothermal veins in Taiwan and New formation for a pseudotachylyte vein, from which Zealand – two young, transpressive orogens. Both they calculate a formation depth of 21–23 km orogens have two fluid-flow systems centred at the with an ambient temperature of 800 8C. The textures drainage divide: a shallow topographically driven observed in the vein suggest that it formed by initial system involving dominantly meteoric water viscous deformation followed by brittle failure, where steep veins develop in an extensional perhaps as increasing strain rate resulted in grain- regime; and a deeper flow of mineralizing, rock- size reduction and an increased surface area pro- exchanged fluids in which subhorizontal veins moted frictional melting. Moecher & Steltenpohl form during vertical stretching. In Taiwan, veins also describe pseudotachylytes formed at warmer were formed locally during embrittlement of conditions than normally expected for earthquake initially weak slates; in New Zealand, a more nucleation. They report host rock temperatures uniform rheology leads to distributed vein for- exceeding 600 8C for multiple generations of fric- mation. In the last chapter in this section, Ujiie tion melt in extensional shear zones developed et al. present high-velocity friction experiments in during late- to post-Caledonian crustal extension. which clay-rich fault gouge is subjected to thermal They suggest that frictional melting of rocks nor- pressurization and fluidization. In both wet and mally deforming by slow creep in the aseismic dry tests, fault weakening occurs as fluid pressure regime occurred as earthquake ruptures, nucleated increases (and therefore effective normal stress at shallower depths, propagated at high strain decreases) within the fault plane. In the dry tests, rates, into the viscous regime below the base of this is caused by thermal pressurization as water is the seismogenic zone. To complement the descrip- released from clay dehydration driven by frictional tions of rocks representing seismic slip at and heating. In wet tests, fluid pressure is increased by below the seismogenic zone, Nu¨chter & Ellis both frictional heating and shear-enhanced compac- present numerical models of coseismic loading tion of the gouge. In addition to concluding that around major faults in a range of tectonic settings. earthquake ruptures may be enhanced by fluid press- Their models predict a coseismic stress drop in the urization in both wet and dry experiments, these upper crust, accompanied by loading to high stress authors show the microstructures that developed at depths below the base of the seismogenic zone. during this process, which can be compared to In all of these papers it is clear that, although it is observed fault rocks that may have accommodated rare for earthquakes to nucleate below the seismo- similar effects. genic zone, earthquakes significantly change the Earthquakes regularly occur on pre-existing surrounding stress field. This, combined with high faults, particularly if weakened by fluid presence. strain rates and strength anisotropy, may lead to However, it is important to consider when a fault slip propagation into the normally viscous regime, may be reactivated as opposed to when it is prefer- and significant loading of the mid- to lower crust. able for new faults to initiate. Nortje et al. analyse Many active and ancient fault zones contain relative timing and orientation data for mineralized clear evidence of the effects of fluids on faulting. faults in the Mount Gordon Fault Zone, Mt Isa, Aus- The fourth set of papers focuses on fluid flow in tralia. Here, deformation was accommodated by the crustal fault zones. Barker & Cox studied a initiation of multiple new orientations of strike-slip swarm of hydrothermal quartz–calcite veins devel- faults, rather than by reactivation of pre-existing oped incrementally during progressive fold growth structures. This implies that the existing faults had and associated reverse faulting. They conclude high cohesion and/or insufficient fluid pressure for that, in this setting, fluid pathways changed dynami- reactivation. They also conclude that pre-existing cally as permeability varied from high to low during faults with high cohesion may have acted as barriers cycles of episodic slip, fracture development and to fluid flow, so that intersections between new and Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

GEOLOGY OF THE EARTHQUAKE SOURCE 13 old faults host mineralization as a result of the inter- References action between newly formed, high-permeability Abercrombie faults, and cohesive, relatively impermeable, pre- , R. E. 1995. Earthquake source scaling existing structures. For more general cases, Scholz relationships from 21 to 5 ML using seismograms recorded at 2.5-km depth. Journal of Geophysical presents a theoretical treatment of how splay faults Research, 100, 24 015–24 036. form at acute angles to the primary fault when the Aki, K. 1967. Scaling law of seismic spectrum. Bulletin of main fault becomes misaligned for reactivation. the Seismological Society of America, 72, 1217–1231. He gives compelling examples from literature on Aki, K. 1981. A probabilistic synthesis of precursory active faults of how, in the dip–slip case, splay phenomena. In: Simpson,D.&Richards, P. (eds) faults commonly develop when the main fault orien- Earthquake Prediction. American Geophysical tation is guided by a pre-existing weak plane, which Union, Washington, DC, 566–574. Allmann Shearer becomes frictionally misaligned with the greatest ,B.P.& , P. M. 2009. Global vari- principal compressive stress. ations of stress drop for moderate to large earthquakes. Journal of Geophysical Research, 114, B01310, As outlined above, the 16 original papers in this doi:10.1029/2008JB005821. volume cover a range of topics on the slip behaviour Ashby,M.F.&Verrall, R. A. 1973. Diffusion accom- of faults in various tectonic settings. The variety in modated flow and superplasticity. Acta Metallurgica, approach and methodology in the chapters, and the 21, 149–163. range of geological features they describe, is testa- Austrheim,H.&Boundy, T. M. 1994. Pseudotachylyte ment to the wide scope of earthquake geology, as generated during seismic faulting and eclogitization of Sibson describes in the final chapter of this book. the deep crust. Science, 265, 82–83. Beeler Tullis Weeks In his chapter, Rick outlines his thoughts on the , N. M., ,T.E.& , J. D. 1994. The future directions of the field of earthquake roles of time and displacement in the evolution effect in rock friction. Geophysical Research Letters, 21, geology. He points out that earthquakes influence 1987–1990. a wide range of geological processes at a variety Bilek,S.L.&Lay, T. 2002. Tsunami earthquakes of distances from the rupture source. While palaeo- possibly widespread manifestations of frictional seismology has traditionally focused on fault-zone conditional stability. Geophysical Research Letters, surface processes, Sibson highlights the need to 29, 1673, doi:10.1029/2002GL015215. expand the concept to near-field and far-field Blanpied, M. L., Lockner,D.A.&Byerlee,J.D. effects, both on the surface and in the subsurface. 1995. Frictional slip of granite at hydrothermal His conclusion, that many processes commonly conditions. Journal of Geophysical Research, 100, considered smooth and progressive are rather inter- 13 045–13 064. Boettcher, M. S., Hirth,G.& Evans, B. 2007. Olivine mittent and tied to the seismic cycle, should be kept friction at the base of oceanic seismogenic zones. in mind by all geologists studying deformation of Journal of Geophysical Research, 112, B01205, the Earth’s lithosphere. doi:10.1029/2006JB004301. Borradaile, G. J. 1981. Particulate flow of rock and the formation of cleavage. Tectonophysics, 72, Summary 305–321. Bos,B.&Spiers, C. 2001. Experimental investigation into The uniformitarian approach of considering Earth the microstructural and mechanical evolution of deformation as slow and constant on geological phyllosilicate-bearing fault rock under conditions timescales has become engrained in geological favouring pressure solution. Journal of Structural Geology, 23, 1187–1202. thought. In this volume, rocks deformed by earth- Boullier, A.-M. 2011. Fault-zone geology: lessons from quake slip are identified in settings of crustal short- drilling through the Nojima and Chelungpu faults. ening, extension and transpression. Experiment and In: Fagereng,A˚ ., Toy,V.G.&Rowland,J.V. theory suggest that seismic slip is common. It is time (eds) Geology of the Earthquake Source: A Volume to reconsider the common approach of assuming in Honour of Rick Sibson. Geological Society, exhumed faults were aseismic until proven guilty London, Special Publications, 359, 17–38. of seismic slip, especially given the high frequency Boullier, A.-M. & Robert, F. 1992. Palaeoseismic of earthquakes recorded by seismic networks around events recorded in Archaean gold-quartz vein net- the globe. Future studies need to identify the mech- works, Val d’Or, Abitibi, Quebec, Canada. Journal of Structural Geology, 14, 161–179. anisms behind aseismic slip, the factors governing Boullier, A.-M., Yeh, E.-C., Boutareaud, S., Song, the partitioning between seismic and aseismic S.-R. & Tsai, C.-H. 2009. Microscale anatomy of the deformation, and identify what characterizes an 1999 Chi–Chi earthquake fault zone. Geochemistry earthquake nucleation site. Geophysics Geosystems, 10, Q03016, doi:10.1029/ 2008GC002252. A˚ . Fagereng acknowledges funding from the UCT Boutareaud, S., Calugaru, D.-G., Han, R., Fabbri, O., Research Development Fund. The authors thank R. Law Mizoguchi, K., Tsutsumi,A.&Shimamoto,T. for review comments that improved the final manuscript. 2008. Clay-clast aggregates: a new structural evidence Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

14 A˚ . FAGERENG & V. G. TOY

for seismic sliding? Geophysical Research Letters, Gibbs, J. W. 1877. On the equilibrium of heterogeneous 35, L05302, doi:10.1029/2007GL032554. substances. Transactions of the Connecticut Academy, Brace,W.&Byerlee, J. D. 1966. Stick–slip as a mech- 3, 108–248, 343–524. anism for earthquakes. Science, 153, 990–992. Gratier,J.P.&Gamond, J. F. 1990. Transition between Brace,W.F.&Kohlstedt, D. L. 1980. Limits on litho- seismic and aseismic deformation in the upper crust. spheric stress imposed by laboratory experiments. In: Knipe,R.J.&Rutter, E. H. (eds) Deformation Journal of Geophysical Research, 85, 6248–6252. Mechanisms, Rheology and Tectonics. Geological Byerlee, J. D. 1968. Brittle–ductile transition in rocks. Society, London, Special Publications, 54, 461–473. Journal of Geophysical Research, 73, 4741–4750. Gratier,J.P.&Gueydan, F. 2007. Deformation in the Carter, R. M. 1975. A discussion and classification of presence of fluids and mineral reactions. In: Handy, subaqueous mass-transport with particular reference M. R., Hirth,G.&Hovius, N. (eds) Tectonic to grain-flow, slurry-flow and fluxoturbidites. Earth- Faults: Agents of Change on a Dynamic Earth. MIT Science Reviews, 11, 145–177. Press, Cambridge, MA, 319–356. Chester,F.M.&Chester, J. S. 1998. Ultracataclasite Gratier, J. P., Guiguet, R., Renard, F., Jenatton,L.& structure and friction processes on the Punchbowl Bernard, D. 2009. A pressure solution creep law for fault, San Andreas system, California. Tectonophysics, quartz from indentation experiments. Journal of 295, 199–221. Geophysical Research, 114, B03403, doi:10.1029/ Chester, F. M., Friedman,M.&Logan, J. M. 1985. 2008JB005652. Foliated cataclasites. Tectonophysics, 111, 139–146. Gutenberg,B.&Richter, C. F. 1944. Frequency of Choy, G. L., McGarr, A., Kirby,S.H.&Boatwright, earthquakes in California. Bulletin of the Seismological J. 2007. An overview of the global variability in Society of America, 34, 185–188. radiated energy and apparent stress. In: McGarr, A., Handy, M. R. 1990. The solid-state flow of polyminer- Abercrombie,R.&Di Toro, G. (eds) Earthquakes: alic rocks. Journal of Geophysical Research, 95, Radiated Energy and the Physics of Faulting. Ameri- 8647–8661. can Geophysical Union, Geophysical Monographs, Hanks,T.C.&Kanamori, H. 1979. A moment magni- 170, 15–24. tude scale. Journal of Geophysical Research, 84, Clarke,G.L.&Norman, R. 1993. Generation of pseudo- 2348–2350. tachylyte under granulite facies conditions, and its Hill, D. P., Eaton,J.P.&Jones, L. M. 1990. Seismicity, preservation during cooling. Journal of Metamorphic 1980–86. In: Wallace, R. E.(ed.) The San Andreas Geology, 11, 319–335. Fault System, California. US Geological Survey, Cowan, D. S. 1999. Do faults preserve a record of seismic Professional Paper, 1515, 115–151. slip? A field geologist’s opinion. Journal of Structural Hippertt,J.&Egydio-Silva, M. 1996. New polygonal Geology, 21, 995–1001. grains formed by dissolution–redeposition in quartz Dieterich, J. H. 1972. Time-dependence of rock friction. mylonite. Journal of Structural Geology, 18, Journal of Geophysical Research, 77, 3690–3679. 1345–1352. Dieterich, J. H. 1979. Modeling of rock friction 1. Exper- Hirth, G., Teyssier,C.&Dunlap, W. 2001. An evalu- imental results and constitutive equations. Journal of ation of quartzite flow laws based on comparisons Geophysical Research, 84, 2161–2168. between experimentally and naturally deformed Duebendorfer, E. M., Vermilye, J., Beiser,P.A.& rocks. International Journal of Earth Sciences, 90, Davis, T. L. 1998. Evidence for aseismic deformation 77–87. in the western Transverse Ranges, southern California: Ishimoto,M.&Iida, K. 1939. Observations sur les seisms implications for seismic risk assessment. Geology, 26, enregistre par le microseismograph construite dernier- 271–274. ment (I). Bulletin of the Earthquake Research Institute, Durney, D. W. 1972. Solution-transfer, an important Tokyo University, 17, 443–478. geological deformation mechanism. Nature, 235, Kanamori,H.&Anderson, D. L. 1975. Theoretical basis 315–317. of some empirical relations in seismology. Bulletin of Evison, F. F. 1971. Seismicity of the Alpine Fault, New the Seismological Society of America, 65, 1073–1095. Zealand. Royal Society of New Zealand Bulletin, 9, Kaneko, Y., Avouac, J.-P. & Lapusta, N. 2010. Towards 161–165. inferring earthquake patterns from geodetic obser- Fagereng,A˚ . 2011. Geology of the seismogenic subduc- vations of interseismic coupling. Nature Geoscience, tion thrust interface. In: Fagereng,A˚ ., Toy,V.G.& 3, 363–369. Rowland, J. V. (eds) Geology of the Earthquake Karato,S.&Wu, P. 1993. Rheology of the upper mantle Source: A Volume in Honour of Rick Sibson. Geologi- – A synthesis. Science, 260, 771–778. cal Society, London, Special Publications, 359, 55–76. King, G. C. P. 1983. The accommodation of large strains Fagereng,A˚ .&Sibson, R. H. 2010. Melange rheology in the upper lithosphere of the Earth and other solids and seismic style. Geology, 38, 751–754, doi: by self-similar fault systems: the geometrical origin 10.1130/G30868.1. of b-value. Pure and Applied Geophysics, 121, Faulkner, D. R., Lewis,A.C.&Rutter, E. H. 2003. On 761–815. the internal structure and mechanics of large strike-slip Knipe, R. J. 1989. Deformation mechanisms – recognition fault zones: field observations of the Carboneras fault from natural tectonites. Journal of Structural Geology, in southeastern Spain. Tectonophysics, 367, 235–251. 11, 127–146. Frohlich,C.&Davis, S. D. 1993. Teleseismic b values; Knipe, R. J. 1990. Microstructural analysis and tectonic or, much ado about 1.0. Journal of Geophysical evolution in thrust systems: examples from the Research, 98, 631–644. Assynt region of the Moine Thrust Zone, Scotland. Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

GEOLOGY OF THE EARTHQUAKE SOURCE 15

In: Barber,D.J.&Meredith, P. G. (eds) Defor- Pavlis, T. L., Serpa,L.F.&Keener, C. 1993. Role of mation Processes in Minerals, Ceramics and Rocks. seismogenic processes in fault-rock development – Mineralogical Society of Great Britain and Ireland. an example from Death Valley, California. Geology, Unwin Hyman, London, 228–261. 21, 267–270. Kohlstedt, D. L., Evans,B.&Mackwell, S. J. 1995. Peng,Z.&Gomberg, J. 2010. An integrated perspective Strength of the lithosphere: constraints imposed by of the continuum between earthquakes and slow-slip laboratory experiments. Journal of Geophysical phenomena. Nature Geoscience, 3, 599–607. Research, 100, 17 587–17 602. Pﬁffner,O.A.&Ramsay, J. G. 1982. Constraints on Lin, A., Maruyama, T., Stallard, A., Michibayashi, geological strain rates: arguments from ﬁnite strain K., Camacho,A.&Kano, K. 2005. Propagation of states of naturally deformed rocks. Journal of Geophy- seismic slip from brittle to ductile crust: evidence sical Research, 87, 311–321. from pseudotachylyte Woodroffe thrust, central Pryer, L. L. 1993. Microstructures of feldspars from Australia. Tectonophysics, 402, 221–235. a major crustal thrust zone: the Grenville Front, Maltman, A. (ed.). 1994. The Geological Deformation of Ontario, Canada. Journal of Structural Geology, 15, Sediments. Chapman & Hall, London. 21–36. Mariani, E., Brodie,K.H.&Rutter, E. H. 2006. Exper- Reid, H. F. 1911. The Elastic-Rebound Theory of imental deformation of muscovite shear zones at high Earthquakes. University of California Publica- temperatures under hydrothermal conditions and the tions, Bulletin of the Department of Geology, 6, strength of phyllosilicate-bearing faults in nature. 413–444. Journal of Structural Geology, 28, 1569–1587. Rowe, C. D., Meneghini,F.&Moore, J. C. 2011. Tex- Marone, C. 1998. Laboratory-derived friction laws and tural record of the seismic cycle: strain-rate variation their application to seismic faulting. Annual Review of in an ancient subduction thrust. In: Fagereng,A˚ ., Earth and Planetary Sciences, 26, 643–696. Toy,V.G.&Rowland, J. V. (eds) Geology of the Marone, C., Raleigh,C.B.&Scholz, C. H. 1990. Fric- Earthquake Source: A Volume in Honour of Rick tional behavior and constitutive modeling of simulated Sibson. Geological Society, London, Special Publi- fault gouge. Journal of Geophysical Research, 95, cations, 359, 77–96. 7007–7025. Ruff,L.&Kanamori, H. 1980. Seismicity and the sub- McClay, K. R. 1977. Pressure solution and Coble creep in duction process. Physics of the Earth and Planetary rocks and minerals. Journal of the Geological Society, Interiors, 23, 240–252. London, 134, 57–70. Ruina, A. 1983. Slip instability and state variable fric- McGarr, A. 1999. On relating apparent stress to the stress tion laws. Journal of Geophysical Research, 88, causing earthquake fault slip. Journal of Geophysical 10 359–10 370. Research, 104, 3003–3011. Rutter, E. H. 1976. The kinetics of rock deformation by Moecher,D.P.&Steltenpohl, M. G. 2011. Petrologi- pressure solution. Philosophical Transactions of the cal evidence for co-seismic slip in extending middle– Royal Society of London, 283, 203–219. lower continental crust: Heier’s zone of pseudotachy- Rutter,E.H.&Mainprice, D. H. 1979. On the possi- lyte, north Norway. In: Fagereng,A˚ ., Toy,V.G.& bility of slow fault slip controlled by a diffusive mass Rowland, J. V. (eds) Geology of the Earthquake transfer process. Gerlands Beitrage zur Geophysik, Source: A Volume in Honour of Rick Sibson. Geo- 88, 154–162. logical Society, London, Special Publications, 359, Scholz, C. H. 1988. The brittle–plastic transition and the 169–186. depth of seismic faulting. Geologische Rundschau, 77, Moore,J.C.&Saffer, D. 2001. Updip limit of the seis- 319–328. mogenic zone beneath the accretionary prism of south- Scholz, C. H. 1997. Size distributions for large and small west Japan: an effect of diagenetic to low-grade earthquakes. Bulletin of the Seismological Society of metamorphic processes and increasing effective America, 87 , 1074–1077. stress. Geology, 29, 183–186. Scholz, C. H. 1998. Earthquakes and friction laws. Morrow, C., Radney,B.&Byerlee, J. 1992. Frictional Nature, 391, 37–42. strength and the effective pressure law of montmorillo- Scholz, C. H. 2002. The Mechanics of Earthquakes and nite and illite clays. In: Evans,B.&Wong, T. (eds) Faulting, 2nd edn. Cambridge University Press, Fault Mechanics and Transport Properties of Rocks. Cambridge. Academic Press, New York, 69–88. Scholz, C. H., Molnar,P.&Johnson, T. 1972. Detailed Niemeijer,A.R.&Spiers, C. J. 2006. Velocity studies of frictional sliding of granite and implications dependence of strength and healing behaviour in for the earthquake mechanism. Journal of Geophysical phyllosilicate-bearing fault gouge. Tectonophysics, Research, 77, 6392–6406. 427, 231–253. Schwartz,D.P.&Coppersmith, K. J. 1984. Fault Niemeijer,A.R.&Spiers, C. J. 2007. A microphysical behavior and characteristic earthquakes: examples model for strong velocity weakening in phyllosilicate- from the Wasatch and San Andreas Fault Zones. bearing fault gouges. Journal of Geophysical Journal of Geophysical Research, 89, 5681–5698. Research, 112, B10405, doi:10.1029/2007JB005008. Sibson, R. H. 1975. Generation of pseudotachylyte by Passchier,C.W.&Trouw, R. A. J. 2005. Microtec- ancient seismic faulting. Geophysical Journal of the tonics. Springer, Berlin. Royal Astronomical Society (now Geophysical Paterson, M. S. 1973. Nonhydrostatic thermodynamics Journal International), 43, 775–794. and its geological applications. Reviews of Geophysics Sibson, R. H. 1977. Fault rocks and fault mechanisms. and Space Physics, 11, 355–389. Journal of the Geological Society, 133, 191–213. Downloaded from http://sp.lyellcollection.org/ by guest on September 27, 2021

16 A˚ . FAGERENG & V. G. TOY

Sibson, R. H. 1980. Transient discontinuities in duc- flow laws and predictions on the temperature of the tile shear zones. Journal of Structural Geology, 2, brittle–plastic transition. Journal of Structural 165–171. Geology, 21, 351–369. Sibson, R. H. 1983. Continental fault structure and the Tullis,J.&Yund, R. A. 1991. Diffusion creep in feldspar shallow earthquake source. Journal of the Geological aggregates: experimental evidence. Journal of Struc- Society, 140, 741–767. tural Geology, 13, 987–1000. Sibson, R. H. 1984. Roughness at the base of the seismo- Ujiie, K., Tsutsumi,A.&Kameda, J. 2011. Reproduction genic zone: Contributing factors. Journal of Geophysi- of thermal pressurization and fluidization of clay-rich cal Research, 89, 5791–5799. fault gouges by high-velocity friction experiments Sibson, R. H. 1985. Stopping of earthquake ruptures at and implications for seismic slip in natural faults. In: dilational fault jogs. Nature, 316, 248–251. Fagereng,A˚ ., Toy,V.G.&Rowland, J. V. (eds) Sibson, R. H. 1986. Earthquakes and rock deformation in Geology of the Earthquake Source: A Volume in crustal fault zones. Annual Review of Earth and Plane- Honour of Rick Sibson. Geological Society, London, tary Sciences, 14, 149–175. Special Publications, 359, 267–286. Sibson, R. H. 1989. Earthquake faulting as a structural Van Diggelen, E., Holdsworth, R. E., De Bresser, process. Journal of Structural Geology, 11, 1–14. J. H. P., Spiers, C., Smith, S. A., Walker,R.J.& Sibson, R. H. 2002. Geology of the crustal earthquake Bowen, L. 2010. The microstructural character and source. In: Lee, W. H. K., Kanamori, H., Jennings, evolution of fault rocks from SAFOD and potential P. C. & Kisslinger, C. (eds) International Handbook weakening mechanisms along the San Andreas Fault. of Earthquake and Engineering Seismology – Part A. In: 2010 American Geophysical Union Fall Meeting. Academic Press, New York, 455–473. American Geophysical Union, San Francisco, CA, Sibson, R. H. 2003. Thickness of the seismic slip zone. Abstract T52B-06. Bulletin of the Seismological Society of America, 93, Voll, G. 1976. Recrystallization of quartz, biotite and 1169–1178. feldspars from Erstfeld to the Leventina Nappe, Sibson, R. H. 2011. The scope of earthquake geology. In: Swiss Alps, and its geological significance. Schweizer- Fagereng,A˚ ., Toy,V.G.&Rowland, J. V. (eds) ische Mineralogische und Petrographische Mitteilun- Geology of the Earthquake Source: A Volume in gen, 56, 641–647. Honour of Rick Sibson. Geological Society, London, Wells,D.L.&Coppersmith, K. J. 1994. New empirical Special Publications, 359, 319–332. relationships among magnitude, rupture length, rupture Sibson,R.H.&Toy, V. G. 2006. The habitat of fault- width, rupture area, and surface displacement. Bulletin generated pseudotachylyte: presence v. absence of of the Seismological Society of America, 84, friction melt. In: McGarr, A., Abercrombie,R.& 974–1002. Di Toro, G. (eds) Earthquakes: Radiated Energy and Wesnousky, S. G. 1999. Crustal deformation and the the Physics of Faulting. American Geophysical stability of the Gutenberg–Richter relationship. Bulle- Union, Geophysical Monographs, 170, 153–166. tin of the Seismological Society of America, 89, Smith, S. A. F., Holdsworth, R. E., Collettini,C.& 1131–1137. Pearce, M. A. 2011. The microstructural character Wesnousky, S. G., Scholz, C. H., Shimazaki,K.& and mechanical significance of fault rocks associated Matsuda, T. 1983. Earthquake frequency distribution with a continental low-angle normal fault: the and the mechanics of faulting. Journal of Geophysical Zuccale Fault, Elba Island, Italy. In: Fagereng,A˚ ., Research, 88, 9331–9340. Toy,V.G.&Rowland, J. V. (eds) Geology of the White, S. 1973. Syntectonic recrystallisation and texture Earthquake Source: A Volume in Honour of Rick development in quartz. Nature, 244, 276–278. Sibson. Geological Society, London, Special Publi- Wintsch, R. P., Christoffersen,R.&Kronenberg, cations, 359, 97–114. A. K. 1995. Fluid-rock reaction weakening of Snoke, A. W., Tullis,J.&Todd, V. R. (eds). 1998. fault zones. Journal of Geophysical Research, 100, Fault-related Rocks: A Photographic Atlas. Princeton 13 021–13 032. University Press, Princeton, NJ. Woodcock,N.H.&Mort, K. 2008. Classification of fault Sorby, H. C. 1853. On the origin of slaty cleavage. Edin- breccias and related fault rocks. Geological Magazine burgh New Philosophical Journal, 55, 137–148. 145, 435–440. Sto¨ckhert,B.,Brix,M.F.,Kleinschrodt,R.,Hurford, Zeuch, D. H. 1983. On the inter-relationship between A. J. & Wirth, R. 1999. Thermochronometry and micro- grain-size-sensitive creep and dynamic recrystalliza- structures of quartz – a comparison with experimental tion of olivine. Tectonophysics, 93, 151–168.